CHAD EDWARD WOLAK - Home - GETD

CHAD EDWARD WOLAKMesozoic structure, stratigraphy, and magmatism in the eastern Pueblo Mountains,

southeast Oregon and northwest Nevada: A record of an allochthonous arc terrane.(Under the direction of SANDRA J. WYLD)

Mesozoic rocks in the eastern Pueblo Mountains, along the Oregon-Nevada

border, represent a Middle Jurassic stratovolcano complex. New mapping, stratigraphic,

magmatic, and structural studies indicate these rocks constitute a distinct terrane (Pueblo

terrane), whose relationship to other Mesozoic terranes, in the U.S. Cordillera, is

uncertain.

The Pueblo terrane contains a thick sequence of volcanogenic strata intruded by

shallow crustal-level comagmatic plutons (~179-176 Ma). Volcanogenic strata consist of

dacitic-andesitic lava, rhyolitic-andesitic volcaniclastics, and rhyolitic-dacitic tuffs. These

rocks reflect an arc constructed on transitional, but perhaps continental, crust and

deposition in a subaerial environment.

A major ductile shear zone (Pueblo Mountains shear zone) deforms the

southeastern Pueblo terrane, and is characterized by greenschist grade mylonites with

northeast-striking, moderately southeast-dipping foliations and down-dip mineral-

stretching lineations. Along its northwest boundary, strain quickly fades, and the

northwestern Pueblo terrane is undeformed by the shear zone. In it, shear sense indicators

include abundant S-C fabrics and mantled porphyroclasts that indicate top-to-the-

northwest directed shear and reverse-sense compressive tectonism. Preliminary 40Ar/39Ar

geochronologic data suggest regional metamorphism and the shear zone are Albian (mid-

Cretaceous).

INDEX WORDS: Oregon, Nevada, Mesozoic, Pueblo Mountains, shear zone, arc,

stratovolcano.

MESOZOIC STRUCTURE, STRATIGRAPHY, AND MAGMATISM IN THE

EASTERN PUEBLO MOUNTAINS, SOUTHEAST OREGON AND NORTHWEST

NEVADA; A RECORD OF AN ALLOCHTHONOUS ARC TERRANE

by

CHAD EDWARD WOLAK

B.A., The University of the South, 1994

A Thesis Submitted to the Graduate Faculty of The University of Georgia in Partial

Fulfillment of the Requirements for the Degree

MASTER OF SCIENCE

ATHENS, GEORGIA

2001

2001

Chad Edward Wolak

All Rights Reserved

MESOZOIC STRUCTURE, STRATIGRAPHY, AND MAGMATISM IN THE

EASTERN PUEBLO MOUNTAINS, SOUTHEAST OREGON AND NORTHWEST

NEVADA; A RECORD OF AN ALLOCHTHONOUS ARC TERRANE

by

CHAD EDWARD WOLAK

Approved:

Major Professor: Sandra J. Wyld

Committee: Steven M. HollandMichael F. RodenJames E. Wright

Electronic Version Approved:

Gordhan L. PatelDean of the Graduate SchoolThe University of GeorgiaDecember 2001

iv

ACKNOWLEDGMENTS

This project was supported by the National Science Foundation grant 10-21-

RR176-239, and would not have been possible without the help and support of S.J.

Wyld, J.E. Wright, S.M. Holland, and M.F. Roden. I would also like to thank the

University of Georgia Geology Department for an exceptional graduate experience, Alison

Mote for her assistance with my fieldwork, and my friends in Denio, Nevada for their

hospitality. To my family and friends in Athens (and elsewhere in the United States),

thank you for your love, generous hearts, and patience.

v

TABLE OF CONTENTS

Page

ACKNOWLEDGMENTS..................................................................................................iv

CHAPTER

1 INTRODUCTION, GEOLOGIC FRAMEWORK, AND PREVIOUS

WORK.............................................................................................................1

Geologic Framework and Previous Work................................................9

2 PRE-CENOZOIC GEOLOGIC UNITS OF THE EASTERN PUEBLO

MOUTAINS..................................................................................................19

Stratigraphic Units of the Pueblo Terrane.............................................19

Intrusive Rocks of the eastern Pueblo Mountains................................65

Geochemistry of Igneous Rocks and Tectonic Setting of Magmatism.75

Relationship Between Intrusive and Volcanic Units.............................87

Stratigraphic and Magmatic Summary and Interpretation of the Pueblo

Terrane..................................................................................................89

3 STRUCTURAL GEOLOGY AND METAMORPHISM IN THE

EASTERN PUEBLO MOUNTAINS...........................................................96

Mesozoic Deformation and Metamorphism.........................................96

Cenozoic Deformation........................................................................148

4 SIGNIFICANCE OF THE PUEBLO TERRANE IN THE U.S.

CORDILLERA............................................................................................161

REFERENCES.................................................................................................................173

1

CHAPTER 1

INTRODUCTION, GEOLOGIC FRAMEWORK, AND PREVIOUS WORK

During the Paleozoic and Mesozoic, periods of prolonged subduction occurred

along the western margin of the North American continent. Plate convergence and related

subduction of oceanic crust resulted in the accretion of several outboard terranes to the

western edge of the continent, produced widespread and diverse arc magmatism, and

resulted in major periods of orogenesis and shortening deformation, primarily in the

Mesozoic. These processes lead to the development of the Cordilleran orogenic belt of

the western U.S. and the collage of tectonostratigraphic terranes that form the margin of

the continent (Fig. 1.1).

Initial synthesis of the development of the U.S. Cordillera emphasized two-

dimensional models involving “accordion” tectonics wherein subduction, extension and

shortening were largely orthogonal to the continental margin, and accreted terranes may

have originated far out in the proto-Pacific ocean basin (Hamilton, 1969; Burchfiel and

Davis, 1972; Burchfiel and Davis, 1975; Schweickert and Cowan, 1975; Speed, 1979;

Harper and Wright, 1984; Miller and others, 1984; Bruekner and Snyder, 1985; Edelman,

1991). More recently, Cordilleran geologists have begun to emphasize the importance of

oblique subduction and the component of margin-parallel strike-slip displacement that

this can generate in the upper plate (Beck, 1986; Ave Lallemant and Oldow, 1988;

Tickoff and Saint Blanquat, 1997). It is increasingly recognized that terranes in the U.S.

Cordillera may have been translated north or south along the margin from their original

point of origin (Lund and Snee, 1988; Lahren and Schweickert, 1989; Busby-Spera and

2

Figure 1.1. Simplified map of the U.S. Cordillera (modified from Silberling and others,

1984) emphasizing the distribution and characteristics of Paleozoic and Mesozoic

tectonostratigraphic terranes, and showing the location of the Pueblo terrane (PT) in

southeastern Oregon. Heavy solid and dashed lines show the known and inferred

locations of the Mojave Snow Lake Fault (MSLF), Western Nevada Suture (WNS), and

Salmon River Suture (SRS) (locations from Wyld and Wright, 2001). G and R are the

towns of Gerlach and Reno. Boxed area in northwest Nevada and southeast Oregon

shows location of Figure 1.2.

N

0 100 200 km

46o

44o

42o

40o

WNS

MSLF

PT

Nort

h A

mer

ican

Mio

geo

clin

e

NorthAmerican

Miogeocline

120 o 116 o

G

KlamathMtns.

Blue Mtn.province

R

SRS

Mezozoic batholith

Mesozoic forearcbasin

Mesozoic arc built ontransitional crust

Mesozoic arc built oncontinental crust

Late Mesozoicsubduction complex

E. Mesozoicbasinal terrane

E. Mesozoicshelf terrane

Paleozoic and Mesozoicoceanic arc rocks

Paleozoic and Mesozoicophoilitic rocks

Paleozoic and Mesozoicsubduction complex

Paleozoic deepmarine allochthons

Key totectonostratigraphic

terranes

E. MesozoicL. Mesozoic Paleozoic

4

Saleeby, 1990; Schweickert and Lahren, 1990; Wyld and others, 1996; Wyld and Wright,

2000).

Determining the point of origin and displacement history of individual terranes is

important to deciphering the tectonic evolution of the U.S. Cordillera. As described by

Coney and others (1980), a tectonostratigraphic terrane is characterized by internal

homogeneity and continuity of stratigraphy, tectonic style and history. Boundaries

between terranes are fundamental discontinuities in stratigraphy that cannot be explained

easily by conventional facies changes or unconformity. Most boundaries separate totally

distinct temporal or physical rock sequences, and are known or suspected faults that

usually display a complex structural history (Monger, 1977; Berg and others, 1978). In

order to evaluate the origin and role of a terrane in an orogenic belt, its geologic history

must be defined and compared with that of adjacent terranes. Key information that must

be known is: age, stratigraphy, lithology and depositional environment of rocks in the

terrane; tectonic setting of magmatism, if present; and timing and character of deformation

and/or metamorphism, if present. Once this information is known, the geologic history of

the terrane can be evaluated and compared with that of adjacent terranes to determine

what kind of displacement may have occurred on the bounding faults between them.

In this study, I focus on the Pueblo terrane in southeastern Oregon (Figs. 1.1 and

1.2), whose role in the U.S. Cordilleran orogen is unknown. This terrane, which consists

of pre-Cenozoic volcanic and plutonic rocks, is juxtaposed against the Black Rock terrane

to the southeast across a major ductile shear zone, the Pueblo Mountains shear zone, and

surrounded on all other sides by Cenozoic strata (Figs. 1.1, 1.2, and 1.3). The Black Rock

terrane consists of deep marine sedimentary and volcanic rocks of middle Paleozoic to

Triassic age cut by Early Jurassic to Cretaceous plutons (Russell, 1984; Wyld, 1990;

1996; 2000; Quinn and others, 1997). The Black Rock terrane is considered to be little

displaced from its original latitudinal position relative to the continental margin, based on

5

Figure 1.2. Simplified geologic map of northwest Nevada and southeast Oregon (modified

after Wilden, 1964; Roback and others, 1987; Quinn and others, 1997; Wyld, 2000; Wyld

and Wright, 2001). D and W are the towns of Denio and Winnemucca. Box shows area of

Figure 1.3 and Plate 1.

Jurassic felsic tointermediatevolcanogenic strata

Upper Triassic deepmarine sedimentary strata

Pz-Mz strata ofuncertain affinity

Cretaceousplutons

Quaternary

Tertiary

Cretaceousalluvial strata

Triassic to LowerJurassic volcanic andsedimentary strata

mid to upper Paleozoicsedimentary andvolcanic strata

Jurassicplutons

PUEBLO TERRANE BLACK ROCK TERRANE

BASINAL TERRANE

Jurassicplutons

41°

42°

119°

Bla

ck R

ock

Des

ert

BLACKROCK

RANGE

PINEFORESTRANGE

OREGON

NEVADA

W

JACKSONMTNS.

BILKCREEKMTNS.

D

0 10 20 30 km

N

118°

BasinalTerrane

Pueblo Terrane

BlackRock

Terrane

Figure 1.3

WesternNevadaShearZone

PUEBLOMTNS.

Approximatedboundary betweenthe Black Rock andBasinal terranes

7

Figure 1.3. Simplified geology of the eastern Pueblo Mountains, northern Pine Forest

Range, and Lone Mountain (location shown in Figure 1.2). PMSZ is the Pueblo

Mountains shear zone. Distribution of Cenozoic rocks and some Cenozoic bedding data

in the Pueblo Mountains from Roback and others (1987). All other geologic data in the

Pueblo Mountains from this study. Geology of the northern Pine Forest Range and Lone

Mountain is from Wyld and Wright (2001). See Plate 1 for more detail.

5‘ 00“

42o

NORTHERNPINE

FORESTRANGE

0 1 kmN

A'

A

B

B'

C '

D '

D

C

JCD

JSB

JSB

JSB

JSB

JSBJSB

JPM

JDI

JAU

JTU

JTU

JTU

JTU

JT

JAU

JAU

Qal

Qls

Qal

Qal

QlsQls

Qls

Qal

Qal

JCT

PUEBLOMOUNTAINS

Strawberry Butte

Denio Basin

Van Horn Basin

Pueblo Mtn.

Denio, NV.

22

19

22

34

2421

23

26

34

23

25

14

30

23

20

21

14

13

22

4142

40

26

49

3931

25

85

35

50

3547

69

49

49

49594669

63

32

6557

69

54

60

52

4364

4555

21

50

5439

21 80

66

30

43

65

44

7035

43

32

36

42 55

53

41

64

78 51 45

41

43

76

74

39

79

80

60

7880

48

59 574555

53

58

46

54

75

54

49

6855

61

59

75

46 30

39

55

6276

5151

45

35

30

5573

40

40

4439

37

3540

7054

2623

25

21

11

61

A’

55‘ 00“

42o

5‘ 00“

Denio SummitShear Zone

AntelopeValley Shear

Zone

55‘ 00“

E“

E’

E

*

*

*

*

*

*

*

*

*

*

*

*

*

*

*

*

denotes zircon sample localitystrike and dip of bedding

strike and dip of S1 foliation

strike and dip of compaction fabric Cenozoic normal fault - ballson downthrown side

ductile fault - teeth on upthrown side

Northwest boundary of CretaceousPMSZ (Pueblo Mountains shear zone)- teeth on shear zone side

Cretaceous shear zones

denotes geochemistry sample locality*denotes argon sample locality

SYMBOLS

LONEMOUNTAIN

PMSZ

PMSZ

trend and plunge of L1 lineation

58

39

55

6166

76

34

65

30

41 Tertiary

CretaceousGranodiorite plutons(108 - 115 Ma)

Cowden Creek tonaliteto quartz diorite (188+/- 2 Ma)

Baltazor pluton(182 Ma)

Alluvium

Landslide

Volcanic andsedimentary strata

Jura

ssic

Qua

tern

ary

LEGEND

JBZ

JCD

Pueblo Terrane

Andesite unit (176Ma +/- 2 ma)

JAU

Tuff unit (179+/- 2 Ma)

JTU

JSB

JDI

JPM

JCT

Strawberry Butte quartzdiorite (176 +/- 3 Ma)

Diamond Inn hypabyssalrhyolite (160 +/- 2 Ma)

Catlow Creek quartz-monzodiorite (179 +/- 2 Ma)

Jura

ssic

Stratigraphic units

Plutonic units

Pueblo Mountainhypabyssal andesite

Early Jurassicplutons (200-185 Ma)

Paleozoic amphiboliteand biotite schist

Triassic pluton(~235 Ma)

Black Rock Terrane

23

Qal

Qls

JCD

JBZ

= structure section lineC C’

Qal

JCD

Qls

JSB

Qal

Qal

Qal

Qls

Qls

JPM

Qal

Qls Qls

9

stratigraphic and detrital zircon studies that link it to the adjacent continental margin in

the Paleozoic and Mesozoic (Darby and others, 2000; Wyld, 2000).

Wyld and Wright (1999, 2001) have argued that the Pueblo Mountains Shear Zone

is part of a larger structural boundary in the U.S. Cordillera along which outboard

terranes, including the Pueblo terrane, were displaced several hundred kilometers north

from their original point of origin. This boundary, shown in Figure 1.1, extends from the

southern Sierra through western Idaho. According to these studies, the strike-slip

boundary was subsequently reactivated during a younger compressional event, resulting

in the dip-slip shear sense indicators now found along parts of its length, including the

Pueblo Mountains Shear Zone.

In this thesis, I present the results of a detailed study of the geology of pre-

Cenozoic rocks of the Pueblo Mountains. In this chapter, I provide information on the

geologic framework of the Pueblo Mountains, and summarize the results of previous

studies of the pre-Cenozoic rocks of this area. In chapter two, the focus is on describing

and interpreting the pre-Cenozoic rock units of the Pueblo Mountains including both

stratigraphic and intrusive rocks. In chapter three, I discuss the Mesozoic and Cenozoic

structural evolution of the Pueblo Mountains. In chapter four, I evaluate the data

presented in chapters one through three in terms of the regional geology of this part of the

Cordillera.

GEOLOGIC FRAMEWORK AND PREVIOUS WORK

The Pueblo terrane is located in the Pueblo Mountains, a narrow northeast-

striking mountain range along the Nevada-Oregon border (Fig. 1.2) that was uplifted

during Cenozoic Basin and Range extension. Much of the Pueblo Mountains is overlain

by Cenozoic volcanic strata (Fig. 1.2) erupted from a caldera system on the northeast side

of the range (Roback and others, 1987). Only on the eastern side of the range, adjacent to

10

the steep, southeast-dipping, range-bounding normal fault, are pre-Cenozoic rocks

exposed (Fig. 1.2). These rocks are unconformably overlain by the Cenozoic strata.

The Cenozoic and Mesozoic geology of the Pueblo Mountains was originally

described in four masters theses (Burnham, 1971; Rowe, 1971; Tower, 1972; and Harold,

1973). Each of these workers mapped one-quarter of the range, and together they

produced a preliminary geologic map of the Cenozoic and older rocks. Their mapping

indicated that the pre-Cenozoic rocks of the Pueblo Mountains consisted of schist,

phyllite, greenstone, quartzite, metaconglomerate, metabreccia, metasandstone, aplite,

quartz diorite, biotite granodiorite, and quartz monzonite. K-Ar dates by Harold (1972)

indicated that the pre-Cenozoic rocks of the Pueblo Mountains were likely Mesozoic in

age: two whole rock analyses of dioritic plutonic rocks yielded ages of 100 +/- 2.0 Ma and

92.4 +/- 1.3 Ma, one analysis of biotite from a quartz monzonite yielded an age of 91.3

+/- 1.3 Ma, and one analysis of plagioclase from an andesitic rock yielded an age of 108+/-

1.5 Ma.

Roback and others (1987) also mapped the Pueblo Mountains for a U.S.

Geological Survey mineral resource study. This study provided more detail on the

distribution and lithology of pre-Cenozoic igneous and metamorphic rocks of the range,

concluded that the protoliths of metamorphic rocks were felsic to intermediate plutons,

tuffs, lavas, and volcaniclastic rocks, and emphasized that a strong northeast-striking and

southeast-dipping foliation cuts these rocks on the eastern side of the range. Unpublished

geochronology cited by Roback (1988) suggested that the pre-Cenozoic volcanic rocks of

the Pueblo Mountains were Jurassic, rather than the Cretaceous age implied by prior K-

Ar analysis: zircon from a volcanic tuff yielded a discordant U-Pb age of ~155 Ma, and

Rb/Sr whole rock isochrons from tuffs yielded an age of ~160 Ma.

More recent work in the Pueblo Mountains was by Brown (1996) who mapped

the pre-Cenozoic rocks as part of an M.S. thesis. Her geologic units did not differ

11

substantially in lithology or distribution from those of Roback and others (1987), but she

provided new details on the petrography of these rocks, their deformational history, and

their age. One important contribution of Brown’s study was a structural analysis of the

deformed rocks on the east side of the range, indicating that these rocks are mylonites

defining a shear zone that records top-to-the-northwest shear, and suggesting that the

shear zone graded into non-deformed rocks to the northwest. In addition, Brown (1996)

provided new geochronologic data. U-Pb zircon analyses from five plutonic bodies (two

of which were found to actually be volcanic rocks; see later section) suggested late-Early

to early-Middle Jurassic ages for all (~189-176 Ma), although all analyses were

discordant. These data indicated an older age for both plutons and volcanogenic strata

than was suggested by prior radiometric analyses. In addition, Brown (1996) presented

one 40Ar/39Ar step heating analysis of metamorphic biotite from the shear zone which

indicated that deformation and syntectonic metamorphism in the Pueblo Mountains

occurred in the mid-Cretaceous at ~90 Ma (Brown, 1996). These data suggest that the

Cretaceous K-Ar dates of Harold (1972) from the plutonic and volcanic rocks of the

Pueblo Mountains likely reflect post-emplacement Ar-loss during younger

metamorphism.

In this study, I mapped the pre-Cenozoic rocks in the eastern Pueblo Mountains

at a scale of 1:24,000. Results of this mapping are shown in Figure 1.3 and Plate 1;

structure cross sections across the mapped area are shown in Figure 1.4. Previous

mapping in the eastern Pueblo Mountains had primarily focused on distinguishing

deformed rocks of the shear zone from undeformed rocks, mapping out plutonic bodies,

and delineating zones of lithologic variation in the pre-Cenozoic rocks, with little attempt

to define any stratigraphy in the region. In contrast, my study focused in part on defining

a stratigraphy within the pre-Cenozoic volcanogenic rocks. Two major units were defined

12

Figure 1.4a-d. Northwest to southeast structure sections through the eastern Pueblo

Mountains and northern Pine Forest Range along lines A-A’, B-B’, C-C’, and D-D’ (see

Figure 3 for locations). PMSZ is Pueblo Mountains Shear Zone: shear sense is top-to-

the-northwest. Curved contact between Andesite and Tuff units in sections C-C’, and D-

D’ is inferential and based on evidence discussed in the text indicating that strain in the

Pueblo Mountains Shear Zone dies out near the contact between the two units and that

the contact is stratigraphic not structural.

A A'(NW) (SE)

?

2000

1500

1000

500

2500

2000

1500

1000

500

2500PMSZ

?

TU

JAU

JTU

JTU

JCDJSB

Qu JBZ

B B '(NW) (SE)

0 1000 m

no vertical exaggeration

2500

2000

1500

1000

500

2500

2000

1500

1000

500

JTU

Qu

PMSZ

JDI

JAU

TUJSB

JSB

JTU Qu

JCD

elev

atio

n in

met

ers

elev

atio

n in

met

ers

KG

(a)

(b)


0 1000 m


Strawberry Butte quartz

diorite (176 +/- 3 Ma)

Diamond Inn hypabyssal

rhyolite (160 +/- 2 Ma)

Pueblo Mountain

hypabyssal andesite

Catlow Creek quartz

monzodiorite (~179 +/- 3 Ma)

JDI

JSB

JCT

JPM

INTRUSIVE UNITS

Andesite unit (176 +/- 2 Ma) - mostly

andesite flows with lesser dacitic

andesite flows, felsic tuff, tuff-breccia,

and volcanic sandstones

Tuff unit (179 +/- 2 Ma) - mostly rhyolitic to

dacitic pumice, crystal, lithic and ash tuff with

lesser felsic volcanic sandstone, slitstone, and

breccias, and felsic to intermediate flows

STRATIGRAPHIC UNITS

JAU

mylonitic foliation - closerspacing and thicker linesdenote greater strain

JURA

SSIC

Pueblo Terrane

2500

2000

1500

1000

500

(SE)

Qu

(NW) C C '

2500

2000

1500

1000

500elev

atio

n in

met

ers

0 1000 m

Qu

Tu

JTUQu

JAU

PMSZ

2000

1500

1000

500

2500

(SE)(NW) D D '

elev

atio

n in

met

ers

2000

1500

1000

500

2500Qu

Qu

JAUTu

Qu

JPMJTU

PMSZ

(d)no vertical exaggeration

JCD

KG

JBZ

Jurassic Cowden Creek quartz diorite

with minor tonaite (188 +/- 2 Ma)

Cretaceous Granodiorite

pluton (108-115 Ma)

Jurassic Baltazor

pluton (182 Ma)

Tertiary volcanic and

sedimentary strata

Quaternary landslides

and alluviumQu

Tu

LEGEND

(c)

0 1000 m


JTU

15

in my study, the informally named Tuff unit and Andesite unit (Figs. 1.3, 1.4, and 1.5).

The Tuff unit consists of a variety of felsic tuffaceous rocks with lesser andesite flows

and felsic to intermediate volcaniclastic strata (Fig. 1.5). The Andesite unit contains

similar rock types, but andesitic flow rocks predominate.

These two volcanogenic units are intruded by a pluton and two small stocks (Fig.

1.3). Names for these intrusions, the Strawberry Butte pluton (JSB), Catlow Creek stock

(JCT), and Pueblo Mountain stock (JPM), were originally proposed by Brown (1996) and

are retained here. A deformed felsic quartz porphyry intrusion (Fig. 1.3) was not

recognized in previous studies and is herein named the Diamond Inn hypabyssal rhyolite

(JDI). Three small stocks shown on previous maps of the area by Roback and others

(1987) and Brown (1996) were not found to exist in this study (one is volcanogenic, the

others are altered portions of the Strawberry Butte pluton). In fault contact with the tuff

unit in the southeasternmost part of the Pueblo Mountains is another intrusion, the

Cowden Creek pluton (JCD) (Fig. 1.3), originally named by Brown (1996).

The southeast side of the Pueblo Mountains is deformed in the Pueblo Mountains

shear zone (Fig. 1.3) whose strain is most pronounced to the southeast but dies out to the

northwest near the boundary between the Tuff and Andesite units (Fig. 1.4). The Pueblo

Mountains shear zone is characterized by greenschist facies mylonites with a northeast-

striking, moderately southeast-dipping foliation and a prominent down-dip mineral and

stretching lineation (Fig. 1.3). Outcrop-scale shear sense indicators are abundant and

consistently indicate top-to-the-northwest shear sense. This shear zone thus records

northwest-southeast-directed reverse-sense tectonism.

In this thesis, I propose the term Pueblo terrane for all pre-Cenozoic rocks of the

Pueblo Mountains excluding the Cowden Creek pluton. The basis for this terminology is

the following, and is explained in more detail in later sections. (1) The Tuff and Andesite

units, and the plutons and stocks that intrude them, form related parts of a single

16

Figure 1.5. Stratigraphic units and intrusions of the Pueblo Mountains. Thickness of

interlayering shown schematically. Ages of plutons and volcanic units are based on U-Pb

zircon analyses, and are all from this study except for the age of the Catlow Creek stock,

which is from Brown (1996).

160 +/- 2 MaJDI

179 +/- 2 Ma

base not exposed

188 +/- 2 MaJCD

176 +/- 3 MaJSB

~179 +/- 2 Ma

176 +/- 2 Ma

JCT

min

imun

str

uctu

ral t

hick

ness

in m

eter

s


JPM

Strawberry Buttequartz diorite

Diamond Innhypabyssal rhyolite

Catlow Creekquartz monzodiorite

JSB

JDI

INTRUSIVE UNITS

clast supported felsicvolcaniclastic sandstones

dacitic to rhyolitic lithic tuff-breccia,crystal-pumice tuff, and ash-flow tuff

mostly andesite flows with lesserdacitic andesite flows, felsic tuff,tuff-breccia, and volcanic sandstones

Andesite unit

mostly rhyolitic to dacitic pumice,crystal, lithic and ash tuff

rhyodacite to daciticandesite flows

matrix-supported rhyolitic todacitic volcanic sandstones,siltstones, and breccias

Tuff unit

STRATIGRAPHIC UNITS

JCT

SYMBOLS

faultcontact

PUEBLO TERRANE ( JURASS IC )

JCDJurassic Cowden Creek quartzdiorite with lesser tonalite

LEGEND

Tertiary volcanic andsedimentary strata

unconformity

Tu

500

2000

0

1000

1500

Tu

JPM

18

magmatic terrane based on stratigraphic, petrologic, geochemical, and geochronological

data. (2) This magmatic terrane differs significantly in terms of depositional environment,

lithology, age, and deformational history from the Black Rock terrane, which lies

southeast of the Pueblo Mountains across a narrow Cenozoic valley (Fig. 1.2). (3) The

Pueblo terrane and Black Rock terrane are separated from one another by a system of

northeast-striking ductile shear zones, including the Pueblo Mountains shear zone and

two other shear zones in the northernmost Pine Forest Range and Lone Mountain (Figs.

1.3 and 1.4a; Wyld and Wright, 2001). The Cowden Creek pluton is not considered part

of the Pueblo terrane because it is older than, and in fault contact with, the adjacent Tuff

unit, and because it is similar to another pluton (the Baltazor pluton) located southeast of

the Pueblo Mountains shear zone in the northernmost Pine Forest Range (Figs. 1.3 and

Plate 1).

19

CHAPTER 2

PRE-CENOZOIC GEOLOGIC UNITSOF THE EASTERN PUEBLO MOUNTAINS

All pre-Cenozoic rocks in the Pueblo Mountains have been partially recrystallized

to greenschist facies assemblages during regional metamorphism that accompanied

development of the Pueblo Mountains Shear Zone. Because the focus of this chapter is

on protolith geology, I omit the prefix “meta -” from the descriptions in the following

sections and I refer to the pre-Cenozoic rocks of the Pueblo Mountains in terms of their

protoliths. Metamorphic features of pre-Cenozoic Pueblo Mountains rocks are described

in detail in chapter 3.

STRATIGRAPHIC UNITS OF THE PUEBLO TERRANE

Tuff Unit

The Tuff unit is located on the southeast side of the Pueblo terrane and contained

entirely within the Pueblo Mountains shear zone (Figs. 1.3 and 1.4). Although all rocks

in the tuff unit are deformed by the shear zone fabric and recrystallized by syntectonic

metamorphism, protolith lithologies in most rocks can still be recognized on the basis of

relict pre-deformational textures. Textural clues that provide insight into protolith

lithologies are euhedral porphyroclasts, broken crystal fragments, rounded sand grains,

and flattened pumice clasts.

Approximately 80% of the Tuff unit consists of massive felsic tuffaceous rocks,

including pumice, crystal, lithic and ash tuffs (Fig. 1.5). These rocks are interbedded

locally with felsic to intermediate flows, and with volcaniclastic strata, including volcanic

sandstones, siltstones, and breccias (Fig. 1.5). Interlayering between different rock types

20

in the Tuff unit is generally on the scale of 0.25 m to 10’s of meters and strikes northeast

parallel to the unit as a whole (Fig. 1.3). Scarce bedding measurements (Fig. 1.3 and Fig.

2.1; Plate 1) indicate a consistent north-northeast to northeast strike but variable dip from

290 – 410 southeast to 330 - 550 northwest. These data suggest that bedding in the Tuff

unit is folded on the megascopic scale by northeast trending folds with steeply southeast-

dipping axial planes. Because of the extensive deformation in the Tuff unit and scarcity of

bedding data, it is not possible to calculate stratigraphic thickness for the unit. A

structural thickness was calculated using the average strike and dip of foliation in the unit.

This thickness is ~2250 m.

Four facies of volcanogenic strata are defined in the Tuff unit and described below.

These are the massive tuff facies, the plane-laminated tuff facies, the lava flow facies, and

the volcanic breccia and sandstone facies.

Massive Tuff Facies

This facies constitutes most of the Tuff unit. Rocks within this facies consist

entirely of tuffs containing an abundance of fine grained matrix in which are variable

amounts of scattered, poorly sorted, and angular crystal fragments, flattened pumice, and

volcanic lithics. Absence of any internal stratification features is characteristic of

tuffaceous rocks in this facies.

Clasts in the tuffs of this facies include up to 50% pumice, up to 25% crystal

fragments, and up to 10% lithic fragments. These are set in a fine-grained recrystallized

matrix, inferred to be derived from volcanic ash, that makes up 50-90% of the rock. The

matrix is composed of very fine-grained crystals (< 0.5 mm in diameter) and generally

consists of subequal amounts of quartz and metamorphic muscovite, with minor amounts

of plagioclase (Fig. 2.2). Rarely, potassium feldspar may be present; however, due to

alteration in thin sections that appear to contain potassium feldspar, its presence is

21

Figure 2.1. Lower-hemisphere equal-area net projections of bedding data from Tuff unit

volcaniclastic strata. Solid circles are poles to bedding planes (So). Star is the calculated

fold axis and line is the calculated axial plane.

*

Tuff unit bedding

N21oE, 68oSE

S15oW, 13o

23

Figure 2.2. Photomicrograph of felsic tuff from Tuff unit massive tuff facies. Large

crystals are quartz porphyroclasts. Matrix composed mostly of fine-grained quartz

(white to grey and equant) and metamorphic muscovite (elongate high birefringent

crystals). Foliation, defined by muscovite, trends from upper left to lower right. Width of

view is 3.2 mm. Cross polarized light.

24

uncertain. This matrix is inferred, on the basis of composition and fine-grain size, to have

been derived from recrystallization of felsic ash, specifically ash rich in silica, potassium,

and aluminum. Locally, some tuffs also contain 35-55% mafics in the matrix, including 25-

50% metamorphic biotite, 0-30% metamorphic epidote and 0-10% metamorphic chlorite,

in addition to quartz (0-30%), plagioclase (0-5%) and muscovite (0-25%). This

more mafic matrix is inferred to have been derived from recrystallization of ash rich in

silica with lesser potassium, aluminum, iron, magnesium, and calcium.

Crystal fragments in the tuffs consist exclusively of quartz and plagioclase. These

are present in variable amounts, from dominantly quartz (Fig. 2.2) or dominantly

plagioclase to some combination of both. These crystals range in size from 0.25-2 mm.

Lithic clasts are lapilli-sized (2-64 mm), and consist of quartz-rich aphyric and plagioclase

phyric volcanic rock fragments (Fig. 2.3). Rarely, these volcanic lithics contain some very

fine-grained potassium feldspar. Pumice clasts in these tuffs (Fig. 2.4a-b) are recognized

on the basis of three features. (1) They have a highly flattened shape. Their thickness is

typically < 0.5 mm while their long dimension averages 5 mm and is up to 30 mm. The

amount of flattening strain indicated by these dimensions is large, much more than is seen

in clasts that are obviously volcanic lithics (compare Figs. 2.3 and 2.4a-b). This is

consistent with the fact that pumice is rheologically weak and can be easily strained both

during deposition and in later deformation. (2) Their composition, which is almost

entirely fine-grained metamorphic muscovite (Fig. 2.4b), differs with the composition of

matrix and lithic clasts in the tuffs which both contain much more quartz and/or feldspar,

including some obvious phenocrysts and/or crystal fragments (Figs. 2.2; 2.3; and 2.4b).

(3) Their ragged and diffuse edges, which contrast with the sharp boundaries of volcanic

lithic clasts (compare Figs. 2.3 and 2.4a-b), is consistent with the greater ease of

recrystallization likely with volcanic glass.

25

Figure 2.3. Photomicrograph of lithic tuff from the Tuff unit. Dark rock fragment in upper

right corner is a plagioclase phyric volcanic lithic. Plagioclase laths are replaced by mostly

fine-grained metamorphic muscovite. Rock fragment in the white circle is an aphyric

quartz-rich lithic. Width of view is 6.5 mm. Cross polarized light.

26

Figure 2.4a. Outcrop of pumice tuff from Tuff unit, showing flattened and elongate

pumice clasts (white color) in fine-grained groundmass (grey color). Photograph is looking

normal to foliation plane; long axes of pumice clasts define stretching lineation. Portion of

lens cap (upper left of photograph) is 4 cm.

27

Figure 2.4b. Photomicrograph of pumice tuff from Tuff unit showing pumice clasts (high

birefringent, mostly yellow-orange colored, wispy features: locations indicated by white

arrows) replaced by metamorphic muscovite. Foliation trends from upper left to lower

right. Width of view is 6.5 mm. Cross polarized light.

28

Tuffs in this facies can be divided into six types based of percentage and size of

different clasts. Pumice tuff and lapilli tuff contain 20-50% flattened pumice fragments

(Figs. 2.4a-b). Crystal tuff contains 10-25% crystal fragments (Fig. 2.2). Crystal pumice

and pumice crystal tuffs contain a combination of these clast types in varying proportion.

All of these tuffs also locally contain up to 10% volcanic lithic fragments. Ash tuffs are

similar in mineralogy, texture and grain size to the matrix of pumice and crystal tuffs, but

they contain no lithic or pumice clasts, and less abundant (<10%) quartz and plagioclase

crystal fragments. Ash tuffs are the least common tuff type, while other tuff types are

the most common.

Most of the tuffs are felsic, probably of dacitic to rhyolitic composition, based on

the mineralogy of their crystal clasts (quartz and plagioclase) and matrix (quartz,

plagioclase, muscovite, +/- potassium feldspar). Some more mafic tuffs are present also;

these differ from the felsic tuffs primarily in that their matrix consists mainly of

metamorphic biotite and epidote in the matrix, instead of quartz and metamorphic

muscovite. Also, more mafic tuffs contain fewer plagioclase crystal clasts than the felsic

tuffs. Considering the silica content of minerals in these slightly more mafic tuffs, coupled

with the percent abundance of the different minerals, these tuffs are not much more mafic

than the more common felsic tuffs, and are most likely dacitic in composition.

Two broad processes of deposition can be considered for the volcaniclastic rocks

of the massive tuff facies: (1) epiclastic reworking of older volcanogenic rocks, or (2)

deposition in a primary pyroclastic process. Epiclastic deposition appears unlikely

because the abundance of fine-grained matrix and common presence of pumice in the rocks

of this facies, coupled with the poor sorting and angularity of lithic and crystal clasts, is

inconsistent with the abrasion and size winnowing expected during epiclastic transport.

The lack any internal stratification features or any non-volcanic debris in these rocks also

argues against epiclastic deposition.

29

Deposition from primary pyroclastic eruptions is more consistent with the

various features of this facies. Three types of eruptions could produce voluminous

tuffaceous deposits: pyroclastic fall (deposition via fallout from an ash cloud),

pyroclastic flow (deposition, generally into topographically low areas, of pyroclastic

debris by gravity controlled surface flows containing high particle concentrations), or

pyroclastic surge (deposition, typically mantling topography but thickest in depressed

areas, of pyroclastic debris by low particle concentration surface flows). The principal

distinguishing features between felsic to intermediate deposits from these three different

types of eruptions are as follows. Pyroclastic falls are generally homogenous nearest the

vent, consist of abundant pumice and ash, nearest the vent they can be up to 25 m thick

but are normally much thinner, can have blocks up 10’s of cm in diameter, and may or

may not have internal stratification. Pyroclastic flow deposits are composed of crystals,

glass shards and pumice, and lithic fragments in highly variable proportions. Poor sorting

is characteristic throughout the entire flow sheet but can vary vertically. Pyroclastic surge

deposits form localized accumulations of well stratified volcaniclastic debris, can contain

plane laminated or low-angle cross stratification, are better sorted than flows and falls,

and are composed mostly of lithics and crystals. In contrast, falls and flows contain more

ash and pumice (Fisher and Schmincke, 1984; Cas and Wright, 1987).

The massive tuff facies rocks are poorly sorted throughout, are massive with no

observable internal stratification, and have variable amounts of ash and pumice lapilli,

with lesser crystals. Based on the predominance of pumice and ash in massive tuff facies

rocks, it is unlikely these rocks are pyroclastic surge deposits. Massive tuff facies rocks

have characteristics indicative of both pyroclastic falls and flows (i.e., they contain

variable amounts of pumice, crystals, ash, and lithics, and they are poorly sorted and have

no internal stratification). However, heterogeneity in these rocks and their highly variable

proportions of clast types suggest these rocks most likely represent pyroclastic flows

30

deposits. Therefore, the tuffaceous rocks of the massive tuff facies are interpreted as

pyroclastic pumice and ash flows. Because these tuffs are all significantly recrystallized

and deformed in the Pueblo Mountains shear zone, it is not possible to tell whether they

are welded or not. Pyroclastic falls, surges, and flows are often intimately associated with

each other (Fisher and Schmincke, 1984; Cas and Wright, 1987) and it is possible that all

three depositional processes may be represented in these rocks.

Plane Laminated Tuff Facies

A distinctive, but only locally present, facies in the Tuff unit consists of lithic

tuffs with prominent plane laminations (Fig. 2.5). These laminated rocks occur on the

first ridge south of Van Horn Creek, and on the north wall of the west end of Denio Creek

Canyon, as laterally discontinuous horizons up to 10 m thick (Fig. 2.5; Plate 1). Plane

laminae in this facies are 1-4 mm thick and are defined by grain size variations (Fig. 2.6).

These laminations are generally planar, and rarely show low angle cross bedding.

The laminated tuffs consist of 30-40% moderately sorted clasts, 0.5-2 mm in size,

set in a fine-grained (< 0.25 mm) recrystallized matrix, inferred to be derived from

volcanic ash, that constitutes 60-70% of the rock. Clasts are predominately aphyric felsic

volcanic lithics, with minor pumice, and rare quartz crystal fragments. Matrix consists of

muscovite (5-60%), quartz (0-30%), and epidote (up to 20%). This matrix is broadly

similar in composition to matrix in massive tuffs, and is likewise inferred to be derived

from recrystallization of felsic volcanic ash.

Presence of plane laminations in this facies suggests the possibility of deposition

by epiclastic processes; however, planar laminated stratification dominates in this facies,

even over outcrop thicknesses of 10 m. The only epiclastic depositional environment

likely to be associated with planar laminated rocks of this thickness is the foreshore or

backshore beach environment (Reading, 1996). However, this environment is inconsistent

31

Figure 2.5. Photograph of plane-laminated tuffs facies of the Tuff unit. Outcrop located

along the north wall of the west end of Denio Creek Canyon. Light laminae are composed

mostly of pumice and rhyolitic lithic fragments. Metamorphic epidote is primarily

responsible for the darker colored layers. Lens cap in center of photograph is 5.5 cm

wide.

32

Figure 2.6. Photograph of plane laminations in plane laminated tuff facies of the Tuff unit.

Laminations are 1-4 mm thick and are defined by grain size variations and compositional

variations in clast types. Pencil is 14 cm.

33

with the abundance of fine-grained matrix in these rocks, as well as the lack of any other

features supportive of a shallow marine setting of deposition elsewhere in the Tuff unit

such as cross-bedded stratification, or clast-supported, well sorted, and rounded clastic

rocks. I therefore conclude that the plane laminated facies was not deposited by epiclastic

processes.

Alternatively, certain pyroclastic processes can result in planar laminated

stratification, including surges and fallout from ash clouds. In contrast to other types of

pyroclastic flows that are high particle concentration flows, surges are low concentration

flows that move as expanded, turbulent dispersions (Fisher and Schmincke, 1984; Cas and

Wright, 1987). Pyroclastic surge deposits associated with intermediate to felsic volcanic

systems form localized accumulations of well stratified volcaniclastic debris, typically at

the base or top of pyroclastic ash flows, that are dominated by low angle cross-bedding

near the vent but may be characterized entirely by planer laminated stratification further

from the vent. Surge deposits are also similar in composition to associated pyroclastic ash

flows, but are better sorted and tend to be more enriched in lithic and crystal fragments

with proportionally less ash and pumice. These features are all consistent with the plane

laminated tuff facies of the Tuff unit.

Pyroclastic air fall deposits also may form planar laminated sequences of

volcaniclastic debris that are similar in composition to ash flow tuffs but better sorted;

these air fall deposits may thus be difficult to distinguish from planar laminated surge

deposits (Fisher and Schmincke, 1984; Cas and Wright, 1987). However, fall deposits

typically occur as laterally extensive sheets and their internal stratification features are

commonly crude and vague rather than distinct. In addition, they are more likely to be

composed mostly of pumice or ash. These features are less consistent with the plane

laminated tuff facies of the Tuff unit than are features associated with surge deposits. I

therefore conclude this facies most likely represents surge deposits.

34

Lava Flow Facies

Lava flows in the Tuff unit are recognized on the basis of a clearly defined

porphyritic texture of euhedral plagioclase phenocrysts set in a fine-grained groundmass.

They are up to 150 meters thick and are most common in the northern part of the Tuff

unit. Typical flows (Fig. 2.7) are relatively felsic and contain 15-25% quartz. These rocks

contain 35 to 40% phenocrysts, 1-2 mm in size, of which 30% are euhedral plagioclase

and 0-10% quartz. The remaining 60-65% of these rocks is groundmass comprised of

minerals less than 1 mm in size, including 15% quartz, 20% plagioclase microlites, 5-10%

metamorphic epidote, 0-10% metamorphic biotite, and 15% metamorphic muscovite. No

potassium feldspar was observed in stained thin-sections, and there is no textural

evidence that metamorphic mafic minerals replaced igneous mafic phenocrysts. The

quartz-rich and mafic poor mineralogy of these flow rocks, coupled with the presence of

quartz and plagioclase phenocrysts and the lack of mafic phenocrysts, suggest they are

relatively felsic in composition. Based on the percentage of different minerals in these

flows and the typical compositions of these minerals (Deer and others, 1983), they are

most likely dacitic in composition.

Locally, more mafic flows are present and contain more common and abundant

metamorphic biotite and epidote in the groundmass, as compared to the more felsic flows

described above. The more mafic flows contain 15-25% euhedral plagioclase phenocrysts,

1-3 mm in size, in a groundmass composed of 25-50% metamorphic biotite, 25%

metamorphic epidote, 25% quartz, and minor amounts of plagioclase, chlorite, opaques,

and metamorphic muscovite. No quartz phenocrysts are present nor is there any evidence

that metamorphic minerals are replacing mafic phenocrysts. No potassium feldspar was

observed in stained thin-sections. Based on the percentage of minerals present in these

flows, and the lower silica content of mafic minerals versus felsic minerals, these rocks are

more mafic than the dacitic flows described above, and are best interpreted as andesites.

35

Figure 2.7. Photomicrograph of porphyritic lava flow from the Tuff unit containing

plagioclase phenocrysts (long white crystals). Groundmass consists of very fine-grained

quartz and plagioclase (dark grey equant crystals) and metamorphic muscovite (acicular

high birefringent crystals – red to pink). Foliation is defined by aligned metamorphic

muscovite and trends from lower left to upper right. Width of view is 6.5 mm. Cross

polarized light.

36

Volcanic Breccia and Sandstone Facies

The volcanic breccia and sandstone facies of the Tuff unit consist of ~20% breccia,

~75% sandstone, and ~5% siltstone that, in few places, show bedding ranging from a few

cm to 10’s of cm in thickness (Fig. 2.8 and 2.9). However, in most outcrops this facies is

massive. The volcanic breccia and sandstone facies is only locally present in the Tuff unit,

mostly in the northern part of the unit, and forms laterally discontinuous successions that

are 10’s of meters thick. In general, no internal stratification features are seen in these

volcaniclastic rocks. However, a few outcrops show crude graded bedding (Fig. 2.8).

Siltstones in this facies form thin porcelaneous beds (Fig. 2.9) that are siliceous and

recrystallized.

Volcaniclastic sandstones contain 25-60% sand grains set in a fine-grained (<0.2

mm) recrystallized matrix that comprises 40-75% of the rock. These rocks show an

obvious clastic texture and are moderately sorted with subangular to subrounded clasts

(Fig. 2.8, 2.9, and 2.10). Sand-sized clasts in these rocks generally include plagioclase-

phyric volcanic lithics and aphanitic quartz-rich volcanic lithics, with some plagioclase

and quartz crystal fragments. These clasts are set in a matrix that consists mostly of

metamorphic muscovite, with minor metamorphic epidote, quartz and opaque minerals.

Locally, clasts in the sandstones are dominated by plagioclase crystal fragments, with

lesser lithics and no quartz. In these rocks, biotite rather then muscovite, forms the

metamorphic mica phase in the matrix.

Volcaniclastic breccias (Fig. 2.8 and 2.11) contain 35-70% clasts set in a fine-

grained recrystallized matrix that comprises 30-50% of the rock. Clasts in these rocks are

subrounded to angular and include up to 15% plagioclase-phyric volcanic lithics, and up

to 55% aphanitic quartz-rich volcanic lithics. Clasts range from 5-30 mm in diameter. The

matrix of these rocks consists mainly of metamorphic muscovite, with lesser quartz and

minor opaques.

37

Figure 2.8. Outcrop of the volcanic breccia and sandstone facies of the Tuff unit, showing

graded bedding between volcanic breccia (below) and volcanic sandstone (above). Outcrop

located in Denio Basin. Hammer head is 20 cm.

38

Figure 2.9. Interbedded volcanic siltstone beds (black fine-grained layers) and sandstone

(grey coarser layers) in volcanic breccia and sandstone facies of the Tuff unit. Outcrop

located in Denio Basin. Knife in bottom right is 9 cm long.

39

Figure 2.10. Photomicrograph showing clastic texture of volcanic sandstone from the Tuff

unit volcanic breccia and sandstone facies. Larger white and grey crystals (0.25 – 1 mm in

diameter) are plagioclase crystal clasts and have subrounded to subangular grain shapes.

Matrix consists of biotite (high birefringent acicular crystals) and very fine-grained,

equant quartz (grey to white), epidote (to fine-grained to see in photograph), and opaques

(black). Width of view is 6.5 mm. Cross polarized light.

40

Figure 2.11. Photograph, looking normal to foliation plane, of volcanic breccia from the

volcanic breccia and sandstone facies of the Tuff unit. Clasts are volcanic lithics: dark grey

clasts are plagioclase phyric lithics and white clasts are aphanitic quartz-rich lithics.

Portion of lens cap shown is 4 cm.

41

When evaluating the deposition environment of the volcanic breccia and sandstone

facies, two broad depositional processes are possible. These rocks were either deposited

by pyroclastic or epiclastic processes. Interlayering between volcanic breccia, sandstone,

and siltstone (from few to 10’s of cm in thickness; Figs. 2.8 and 2.9), graded bedding (Fig.

2.8), and rounded volcanic sandstone and siltstone clasts, are features that appear most

consistent with an epiclastic (rather than pyroclastic) mode of deposition.

The exact mode and environment of deposition of this facies is unclear; however,

several lines of evidence exist, from which conclusions about deposition of this facies can

be drawn. (1) The abundance of matrix implies limited transport or reworking and

winnowing by air or water. (2) The presence of breccias requires either gravity-driven

mass movement or transport in a high-shear stress water environment. (3) Clastic rocks of

the volcanic breccia and sandstone facies are similar in composition to clasts and matrix of

tuffaceous rocks, but they have better sorted and rounded clasts, and they contain less

matrix and no pumice. (4) This facies is only locally present in the Tuff unit, and it forms

discontinuous lenses. Collectively, these features appear most consistent with an

interpretation that volcanic breccia and sandstone facies clastic rocks were deposited in a

proximal alluvial environment, on the flanks of a volcanic system, in which primary

pyroclastic debris derived from the massive tuff facies was transported and resedimented

a relatively short distance downslope.

Geochronology of the Tuff Unit

A volcaniclastic sandstone from the volcanic breccia and sandstone facies was

sampled from the Tuff unit for U-Pb geochronology. The location of this sample is

shown in Figure 1.3; latitude and longitude are 42o 00’ 38” N and 118o 40’ 16” W. In thin

section, this sample consists of 50% clasts and 50% matrix. Clasts include of 25%

aphyric felsic volcanic lithics and 25% unidentifiable clasts replaced partly or totally by

42

opaques. The matrix consists mostly of muscovite, with lesser opaque minerals and

epidote.

Zircon grains were extracted from this sample using standard rock grinding and

mineral separation techniques. The sample was crushed in a jaw crusher and disc grinder,

then heavy minerals were separated from the sample via a density separation gemini table.

Heavy liquids and a Franz magnetic barrier laboratory separator were used to isolate

zircon grains from the sample. Several zircon grains (10-15) were spot analyzed on the

Super High Resolution Ion Microprobe (SHRIMP) at Stanford University by J.E.

Wright. According to J.E. Wright the sample yields an age of 179 +/- 2 Ma. This age is

Aalenian (earliest Middle Jurassic) according to the time scale of Gradstein and others

(1994).

Andesite Unit

The Andesite unit is located on the northwest side of the Pueblo terrane and is

largely unaffected by deformation in the Pueblo Mountains shear zone (Figs. 1.3 and 1.4).

Recrystallization is consequently less advanced than in the Tuff unit, and protolith

textures and mineralogy in Andesite unit rocks are more easily recognized.

The Andesite unit consists mostly of lavas (95% of the unit) that are interbedded

locally with felsic tuffaceous and volcaniclastic rocks (Fig. 1.5). Lavas are massive and

lack any bedding features indicating individual flows. Interlayered tuffaceous and

volcaniclastic rocks average 10’s of meters of thickness, but can be as thick as 300 m, and

usually occur as isolated horizons within thicker accumulations of lava. Scare bedding

measurements in the Andesite unit, coupled with map distribution of tuffaceous and

clastic horizons indicates layering is generally north to northeast-striking and that map

scale folding is present in the northernmost part of the unit (Fig. 1.3; Plate 1). Stereonet

analysis of the scarce bedding measurements suggests that layering in the Andesite unit is

43

folded by megascopic folds with a gently to moderately northeast plunging fold axis and a

steeply east-southeast-dipping axial plane (Fig. 2.12). Because of the variable strike and

dip of the scarce bedding measurements in the Andesite unit and the evidence for

megascopic folding, it is difficult to calculate the thickness of the unit. A minimum

thickness is based on the maximum vertical distribution over which the unit is exposed at

Pueblo Mountain (Plate 1), which is ~800 m. A maximum thickness is based on the

maximum horizontal distance over which the unit is exposed, as measured perpendicular

to the dominant north-northeast trend of layering (Plate 1), which is ~4400 m. These are

structural, not stratigraphic, thicknesses due to the evidence for megascopic folding in the

Andesite unit. Regardless of uncertainty, it is safe to conclude that the original

stratigraphic thickness of the Andesite unit was at least a few hundred meters, if not

many hundreds of meters.

Flows

All flow rocks in the Andesite unit are porphyritic, with phenocrysts 1-3 mm in

size including 5-40% plagioclase and up to 8% hornblende (partially replaced by some

combination of metamorphic actinolite, chlorite, biotite, and epidote) (Figs. 2.13 and

2.14). The remaining 60-95% of the flow rock is groundmass comprised of crystals less

than 1 mm, including 25-65% plagioclase microlites, 0-5% quartz, 0-25% mafics

(typically hornblende), 2-20% opaque minerals, and 0-20% interstitial chlorite (perhaps

replacing mafic glass). Rarely, minor amounts of metamorphic biotite and muscovite are

present in the groundmass. No potassium feldspar was observed in stained thin sections.

Based on the abundance of plagioclase in these rocks, coupled with the scarcity of

quartz and the moderate amount of mafics, it appears that the flows are most likely

andesitic in composition. The composition of these flows is thus slightly more mafic than

that of flows in the Tuff unit.

44

Figure 2.12. Lower-hemisphere equal-area net projections of bedding data from Andesite

unit volcaniclastic strata. Solid circles are poles to bedding planes (So). Star is the

calculated fold axis and line is the calculated axial plane.

Andesite unit bedding

*

N50oE, 79oSE

N55oE, 23o

46

Figure 2.13. Typical flow rock from the Andesite unit, showing porphyritic texture:

white phenocrysts are plagioclase and smaller dark phenocrysts are hornblende. Knife is 9

cm long.

47

Figure 2.14. Photomicrograph of lava from the Andesite unit showing typical texture and

mineralogy. Phenocrysts include plagioclase (large white crystal laths) and less common

hornblende (yellow - brown subhedral, acicular crystal shape in lower left corner that is

replaced by an aggregate of very fine-grained metamorphic minerals). Groundmass

consists mostly of plagioclase microlites (white flecks) and metamorphic mafic minerals

(black). Width of view is 3.2 mm. Cross polarized light.

48

Tuffaceous and Volcaniclastic Rocks

A spectrum of tuffaceous and volcaniclastic rocks are found as local horizons

within the Andesite unit. All consist entirely of volcanogenic material (andesitic and/or

rhyolitic volcanic lithics, feldspar and/or quartz crystal fragments, and recrystallized fine-

grained felsic matrix inferred to be derived from volcanic ash), and nearly all contain an

abundance of fine-grained matrix, enclosing angular and poorly sorted clasts. These rocks

are thus all interpreted as either primary pyroclastic deposits or deposits derived from

minor downslope remobilization of primary pyroclastic debris. Three principal

tuffaceous and volcaniclastic assemblages have been found, generally in different parts of

the Andesite unit, and each of these is described and interpreted below.

In Arizona Canyon (see Plate 1, northern part of Andesite unit, for location), the

volcaniclastic assemblage consists of felsic ash tuff with a prominent discontinuous flow

banded fabric (Fig. 2.15), and volcanic sandstone with a well developed and pervasive

fine-scale planar laminated stratification (Fig. 2.16). The tuff contains ~15% angular clasts

in a fine-grained matrix. Clasts include up to 10% plagioclase and quartz crystal

fragments, 0.25-1 mm in size, and scarce volcanic lithics, rarely up to 3 cm in size,

including andesitic clasts with abundant plagioclase microlites, and felsic clasts rich in

very fine-grained quartz. Matrix consists of fine-grained plagioclase, quartz, and minor

potassium feldspar, with common metamorphic biotite. Based on the composition of

clasts and matrix, the Arizona Canyon tuff is probably of a broadly dacitic andesite

composition. Associated volcaniclastic sandstones in Arizona Canyon are of similar

composition to the nearby tuffs. These sandstones contain abundant (85%) subangular

clasts, including felsic volcanic lithics (55%), plagioclase (25%) and quartz crystal

fragments (5%), in a recrystallized matrix consisting of epidote, biotite and minor

potassium feldspar, quartz and plagioclase.

49

Figure 2.15. Photomicrograph of Andesite unit tuff, sampled from the north wall of

Arizona Canyon, with alternating light and dark bands defining a well developed flow

banded fabric. Width of view is 6.5 mm. Plane polarized light.

50

Figure 2.16. Arizona canyon planar-laminated deposit from the Andesite unit.

Laminations are very thin dark horizons that trend from lower left to upper right. Lens

cap is 5.5 cm.

51

A broadly similar assemblage of volcaniclastic rocks occurs on top of Pueblo

Mountain (Fig. 1.3 and Plate 1). Here, felsic ash tuff with a prominent flow banded fabric

is also present and is interlayered with thinly-bedded volcanic breccia (Fig. 2.17). The tuff

on Pueblo Mountain is probably rhyolitic in composition: clasts, 0.25-2 mm in size,

include potassium feldspar, quartz and lesser plagioclase crystal fragments, while the

matrix contains abundant potassium feldspar, lesser quartz and plagioclase, and minor

epidote. On Pueblo Mountain, the volcanic breccias contain 50-70% subangular to angular

volcanic lithic clasts, mostly 1-2 mm in size but locally up to 25 cm, that include

potassium feldspar and/or quartz-rich rhyolitic lithics that are variably homogeneous or

flow-banded. Matrix in these breccias consist mostly of fine-grained quartz, potassium

feldspar and plagioclase, with some scattered and sand-sized quartz crystal fragments,

and minor chlorite and epidote. Planar bedding, 5-8 cm in thickness, is common.

The discontinuous compaction fabric in the tuffs in Arizona Canyon and on top of

Pueblo Mountain appears very similar to the fabric defined by flattened pumice fiamme

in welded ash flow tuffs (Fisher and Schmincke, 1984; Cas and Wright, 1987). In addition,

the overall texture of these tuffs, with scarce, poorly sorted, and angular crystal and

lithics clasts in a fine-grained matrix, is also similar to that found in ash flow tuffs (Fisher

and Schmincke, 1984; Cas and Wright, 1987). Based on these features, I interpret the

Arizona Canyon and Pueblo Mountain tuffs to be welded pyroclastic ash flow deposits.

In both areas, the associated volcaniclastic rocks are similar in composition to the tuffs

and they are, likely closely related to the tuffs, either as products from the same eruptive

phase as the tuffs or as resedimented material derived from the tuffs. The plane laminated

volcanic sandstone in Arizona Canyon is similar in composition, texture, and stratigraphic

features to the plane-laminated facies of the Tuff unit, and is likewise interpreted to be

pyroclastic surge deposits. The volcanic breccias on Pueblo Mountain could either be

near-vent microbreccia surge deposits, or possibly volcaniclastic material redistributed by

52

Figure 2.17. Pueblo Mountain volcanic breccias of the Andesite unit. Finger points to

bedding plane between coarser beds (top) and a finer beds (bottom). Scattered white

clasts (in top beds) are aphanitic rhyolitic lithic fragments.

53

fluvial processes (Fisher and Schmincke, 1984; Cas and Wright, 1987).

A second volcaniclastic assemblage is found in Colony Creek and on the top of

Pueblo Mountain. Rocks here are distinctive volcanic breccias whose clasts are

everywhere flattened in shape (Fig. 2.18a-b) although no tectonic foliation is evident in

the area and the breccia matrix is undeformed. Clasts vary in abundance from 35-55% and

are up to 15 cm long and 3 cm thick. They include minor plagioclase crystal fragments and

abundant felsic to andesitic volcanic lithics, most of which are aphanitic with a felty

groundmass composed of variable amounts of plagioclase, quartz, and potassium feldspar.

In thin section, the clasts are very similar to and difficult to distinguish from the matrix,

which consists of fine-grained plagioclase, quartz, potassium feldspar, and minor

metamorphic biotite and epidote. This matrix and is inferred to be derived from felsic

volcanic ash. The breccias are poorly sorted and massive with no internal stratification.

They appear best interpreted as pyroclastic block and ash flow deposits, based on the

following (Fisher and Schmincke, 1984; Cas and Wright, 1987): abundance of fine-grained

matrix inferred to be recrystallized volcanic ash; massive unstratified character;

predominance of compositionally homogenous volcanic lithics in the clast population; and

similarity in composition of the lithic clasts and the matrix. In addition, and most

importantly, blocks in block and ash flows are primarily juvenile, non-vesiculated cognate

lithics, derived from the eruptive magma and still hot when deposited (Cas and Wright,

1987). Hot clasts such as this could easily become flattened during deposition due to their

rheological weakness relative to that of cold lithics.

The final volcaniclastic rock assemblage in the Andesite unit consists of variably

matrix to clast-supported volcanic breccias and microbreccias. This assemblage forms two

thick horizons, one in Arizona Canyon, and another found locally elsewhere in the

Andesite unit (Plate 1). Clast abundance ranges from 20-80% and clast sizes vary from 5

mm to 60 cm. They are everywhere poorly sorted and vary from subrounded to angular

54

Figure 2.18a-b. Both photographs are of flattened Andesite unit volcanic breccias located

in Colony Creek Canyon. Light colored elongate ellipses are flattened, aphanitic, felsic to

intermediate volcanic lithic clasts suspended in a fine-grained matrix (darker color). Long

axes of flattened clasts define a compaction fabric that trends left to right. Lens cap is 5.5

cm wide.

56

(Fig. 2.19a-d). The clasts consist entirely of andesitic and rhyolitic volcanic lithics.

Andesitic lithics consists of abundant plagioclase microlites and some plagioclase

phenocrysts, and are very similar in texture and composition to andesite lavas elsewhere

in the Andesite unit. Rhyolitic lithics vary from internally structureless to flow banded

(Fig. 2.20), and are rich in quartz and potassium feldspar. Matrix of these breccias varies

in abundance from 20-80% and consists of scattered plagioclase and minor quartz crystal

fragments up to 1 mm in size, surrounded by a finer-grained assemblage of quartz,

potassium feldspar, plagioclase, and minor metamorphic chlorite, epidote, and calcite.

This matrix is inferred to be derived from the recrystallization of felsic ash.

No bedding or internal stratification features are seen in the volcanic breccia and

microbreccia assemblage, although there is variation from outcrop to outcrop in the

abundance and size of clasts, and in the relative abundance of andesitic versus rhyolitic

lithics. The exact mode of deposition of these breccias is unclear. Absence of pumice, and

abundance of volcanic lithics relative to ash matrix, argues against deposition by

pyroclastic ash or pumice and ash flows. The poly-lithologic character of the lithics and

the lack of any evidence that they were hot when deposited also argues against deposition

from pyroclastic block and ash flows. Alternatively, the lack of internal stratification in

the breccias, and their commonly matrix-supported character, suggests deposition by

some sort of mass flow. The best interpretation seems to be that they were deposited

from debris flows remobilizing primary volcanic debris on the flanks of the volcano. Most

likely this material mixed with fragments of previously deposited andesite lava as the

debris flow moved downslope.

57

Figure 2.19a. Volcanic breccia of the Andesite unit, showing volcanic lithics set in a fine-

grained felsic matrix (light tan color). Several rhyolitic lithics (white clasts) are visible and

range ~1-14 cm in diameter. Directly above the right end of the knife is an andesitic lithic

clast (grey color with white dots). Dots are plagioclase phenocrysts. Photograph is from

the north wall of Arizona canyon. Knife is 9 cm long.

58

Figure 2.19b. Volcanic breccias of the Andesite unit, showing volcanic lithics set in a fine-

grained felsic matrix (light tan). White clasts are aphyric rhyolitic lithics, large dark grey

clasts (bottom center and center of photograph) are flow banded rhyolite clasts, and

medium grey clasts with white dots (adjacent to the rhyolite clast in the center) are

plagioclase phyric andesitic lithics. Photograph is from the south wall of Arizona

Canyon. Hammer (right side of photograph) is 28 cm long.

59

Figure 2.19c. Volcanic breccia of the Andesite unit, showing volcanic lithics set in a fine-

grained felsic matrix. Highly angular grey, tan and white clasts are rhyolitic lithics.

Photograph is from the south wall of Arizona Canyon. Hammer (lying horizontal in

upper part of photograph) is 28 cm long.

60

Figure 2.19d. Volcanic microbreccias of the Andesite unit, showing volcanic lithics set in a

fine-grained felsic matrix (light grey speckled pattern). Photograph is from the south wall

of Arizona canyon. White clasts are rhyolitic lithics and dark grey angular clasts are fine-

grained andesitic lithics. Knife is 9 cm long.

61

Figure 2.20. Photomicrograph of part of a large rhyolitic flow-banded clast taken from an

Andesite unit clast-supported volcanic breccia (hand specimen from north wall of

Arizona Canyon). White, grey, and tan bands define flow-banded fabric that trends

vertically across photograph. White crystal in bottom center is plagioclase. Bottom right

corner shows a rock fragment (~2 mm in diameter) composed of darker and lighter lithic

fragments. Width of view is 6.5 mm. Plane polarized light.

62

Geochronology of the Andesite unit

Age data for the Andesite unit are from U-Pb analysis of zircon separates taken

from the plane-laminated volcanic sandstone (surge deposit) in Arizona Canyon. This

sample was collected at 42o 07’ 47” N and 118o 39’ 33” W (Fig. 1.3 and Plate 1). Zircon

grains were extracted from this sample using techniques described above. The zircon

population was homogeneous, consistent with derivation of the sandstone from a primary

pyroclastic deposit. 10-15 zircon grains were spot analyzed on the Super High

Resolution Ion Microprobe (SHRIMP) at Stanford University by J.E. Wright and yield

an age of 176 +/- 2 Ma (J.E. Wright). According to the time scale of Gradstein and others

(1994) this age is Bajocian (Middle Jurassic).

Relation between the Tuff and Andesite Units

Structure contouring of the mapped contact between the Tuff an Andesite units

(Plate 1; Fig. 1.3) indicates that it strikes northeast and dips 330 southeast. The Andesite

unit therefore dips beneath the Tuff unit, suggesting that the Andesite unit is the older of

the two. The contact is contained, however, within the northwest edge of the Pueblo

Mountains shear zone (Plate 1; Fig. 1.3), suggesting that it is better viewed as a structural

rather than a stratigraphic contact. In addition, there is clear evidence for megascopic folds

of bedding and layering in both the Tuff and Andesite units, as well as a lack of any

macroscopic features that would indicate the “up” direction, such as graded beds. There is

therefore no basis for assuming, with any supporting evidence, that the structurally

overlying Tuff unit is younger than the Andesite unit. U-Pb zircon ages from the two

units do not resolve this question because their ages overlap within error.

To evaluate this problem, the following observations are useful to consider. (1)

The Tuff and Andesite units are similar in terms of rock types, facies and age. Both

contain a similar suite of rock types (andesitic to dacitic lava, felsic pyroclastic tuff and

63

surge deposits, and intermediate to felsic volcanic breccia and sandstone), although the

relative abundance of rock types varies between the units. Facies in both units reflect

active, near vent volcanism, both explosive and effusive, with only limited evidence of

epiclastic reworking of volcanogenic debris. Neither unit contains any non-volcanic

sedimentary rocks. Both units have similar ages, identical within error. (2) The contact

between the Tuff and Andesite units is gradational and cannot be mapped as an abrupt

lithologic boundary. The contact was instead mapped where andesite lavas began to

dominate over felsic tuffs. (3) Strain in the Pueblo Mountains shear zone dies out

gradually to the northwest (Plate 1; Fig. 1.3), near but not exactly at the boundary

between the Tuff and Andesite units. At distances of greater than 200 m from the Tuff

unit–Andesite unit contact, a mylonitic foliation characterizes rocks of the Tuff unit.

Closer to the contact, Tuff unit strata are more variably deformed and range from

mylonites to schists. Moving to the northwest in the Andesite unit, strata near the

Tuff–Andesite unit contact are also deformed variably, but less strongly; within 200–400

m of the contact Andesite unit strata vary from mylonitic to schistose to only weakly

deformed. Farther northwest in the Andesite unit, the strata are undeformed (Plate 1; Fig.

1.3). There is thus no structural evidence that the Tuff–Andesite contact is characterized

by increased strain as would be expected if the contact was purely structural. Rather, it

appears that strain in the Pueblo Mountains shear zone coincidentally dies out close to

the contact between the two units. Based on these observations, it is unlikely that there is

substantial displacement between the two units along their contact, although the presence

of higher strain rocks in the Tuff unit requires that there must be some displacement

between the two units. (4) Structures within the Pueblo Mountains shear zone clearly

indicate top-to-the-northwest shear sense (see Chapter 3). If these kinematics motions are

removed, the Tuff unit restores to a position some distance down to the southeast relative

to the Andesite unit. (5) Shear deformation, which occurred in the mid-Cretaceous, is the

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first and only deformation to leave an imprint on the Tuff and Andesite units. It is

therefore reasonable to conclude that layering in the volcanogenic units was originally

subhorizontal prior to shear zone deformation. Stratigraphic contacts between units

would have also been subhorizontal prior to shear zone deformation.

Based on these observations, it appears that the Tuff and Andesite units form

temporally and spatially related accumulations from a single volcanic-plutonic complex,

and that their mutual contact likely represents a deformed stratigraphic contact rather that

a structural boundary across which significant fault displacement occurred. In addition,

considering the displacement history of the Pueblo Mountains shear zone and the fact

that this is the only deformation to affect the volcanic units (points 4 and 5 above), the

Tuff unit could originally have been stratigraphically below the Andesite unit, but it could

not have been above it. If the Tuff unit was originally below the Andesite unit, then the

contact is now overturned due to displacement within the Pueblo Mountains shear zone.

This, however, is not consistent with evidence that the Tuff – Andesite unit contact is a

stratigraphic boundary that is not substantially deformed. Alternatively, the Tuff and

Andesite units originally interfingered laterally, and the gradational contact currently

observed between them is just the moderately deformed manifestation of this laterally

interfingered boundary. This interpretation is consistent with all the observations listed

above and is the preferred interpretation of the relation between the Tuff and Andesite

units.

Depositional Setting of the Tuff and Andesite Units

The depositional setting of the Tuff and Andesite units is interpreted to be

subaerial. This conclusion is based on the following observations and relations. (1)

Accumulations of marine strata typically show obvious evidence of marine deposition,

such as presence of limestone, chert, abundant shale, Bouma sequences, pillow lavas,

65

and/or marine fossils (Reading, 1996). The Tuff and Andesite units of the Pueblo terrane

contain none of these. (2) The Andesite and Tuff units both contain felsic to intermediate

volcaniclastic rocks and pyroclastic flow deposits, including pumice and ash flows, block

and ash flows, surge deposits, and welded ash flows. According to Cas and Wright

(1987), these types of pyroclastic deposits are generally, if not entirely, restricted to the

subaerial environment. (3) With rare exception, tuffaceous and volcaniclastic rocks in the

Andesite and Tuff units show little sorting or rounding of clasts, and contain a large

abundance (up to 90%) of fine-grained matrix inferred to be recrystallized volcanic ash.

These features are not consistent with the winnowing and transport expected in a marine

environment (Reading, 1996). (4) Internal stratification is rare in clastic rocks of the Tuff

and Andesite units; most clastic rocks are massive and show neither internal stratification

or bedding. In contrast, typical accumulations of clastic marine strata show abundant

bedding and sometimes sedimentary structures within individual beds (Reading, 1996). (5)

Finally, the assemblage of intermediate to felsic lavas and pyroclastic flows in the Tuff

and Andesite units, with few strata that contain evidence of epiclastic transport

processes, is typical of the facies found around intermediate to felsic subaerial

stratovolcanoes (Cas and Wright, 1987; Reading, 1996).

INTRUSIVE ROCKS OF THE EASTERN PUEBLO MOUNTAINS

Four plutonic bodies intrude the stratigraphic units of the Pueblo terrane. These

include (Fig. 1.3 and Plate 1): the Strawberry Butte quartz diorite pluton, which is a large,

elongate body that intrudes the Tuff unit; the Pueblo Mountain hypabyssal andesite

stock, which intrudes the andesite unit and underlies Pueblo Mountain; the Catlow Creek

quartz monzodiorite stock, a small body intruding the northern part of the Andesite unit;

and the Diamond Inn hypabyssal rhyolite stock, which intrudes both the Strawberry

Butte pluton and part of the Tuff unit. An additional pluton, the Cowden Creek quartz

66

diorite-tonalite, is present in the southeastern part of the Pueblo Mountains, and is in

fault contact with the Tuff unit (see chapter 3 for detail). Compositions for the Catlow

Creek stock, Strawberry Butte and Cowden Creek plutons were determined by plotting

modal mineralogy on the I.U.G.S. classification diagram for phaneritic rocks (Raymond,

1995). The Diamond Inn and Pueblo Mountain stocks were also assigned compositions

following Raymond (1995). However, instead of assigning these rocks plutonic names,

they are given the volcanic equivalent name because they show hypabyssal textures and

field relations (discussed further in sections below). The following sections provide more

detailed descriptions of each plutonic body.

Strawberry Butte Quartz Diorite

The Strawberry Butte quartz diorite is located entirely within the Pueblo

Mountains Shear Zone (Fig. 1.3), but is heterogeneously deformed, so that rocks of this

pluton vary from non-foliated to strongly-foliated. Regardless of macroscopic fabric

development, all of the pluton exhibits significant dynamic recrystallization at the

microscopic scale. Despite this, the original texture and mineralogy of the pluton is still

reasonably recognizable.

The typical texture and mineralogy of the Strawberry Butte quartz diorite is

shown in Figure 2.21. The pluton is everywhere porphyritic with phenocrysts

constituting 40-50% of the rock. Plagioclase is the primary phenocryst phase, and occurs

as euhedral laths ranging in length from 1 to 3 mm. Minor phenocrystic phases are

hornblende (5% of the rock) and quartz (<1% of the rock), both of which average 1 mm in

diameter. Hornblende is sometimes fresh, but commonly is partially or totally replaced

by biotite and/or epidote. When replaced, typical amphibole crystal shapes are still

preserved. The groundmass of the Strawberry Butte quartz diorite is fine grained (< 1

mm) and consists of 35-40% plagioclase, up to 5% quartz, 10-15% metamorphic biotite

67

Figure 2.21. Photomicrograph showing plutonic porphyritic texture of Strawberry Butte

quartz diorite. Phenocrysts are plagioclase laths and blocks (large white and light grey

crystals). Groundmass consists of very fine-grained plagioclase (white and nearly equant),

lesser metamorphic biotite (brown patches), and minor quartz (light grey and equant).

Width of view is 6.5 mm. Cross polarized light.

68

and epidote, and up to 3% oxides. No potassium feldspar was observed in stained thin-

sections. Mafic metamorphic minerals are assumed to have been derived from

recrystallization of original igneous mafics, but the composition of these cannot be

determined.

All of the groundmass crystals are anhedral suggesting that they likely reflect

dynamic recrystallization of the original igneous groundmass. Some grain size reduction

likely accompanied dynamic recrystallization (see chapter 3 for detail) but it is clear,

based on the obvious porphyritic texture of the pluton and the small size of the

phenocrysts, that the original groundmass must have been fine-grained. The Strawberry

Butte pluton was therefore likely intruded at a shallow crustal level.

A sample of the Strawberry Butte pluton was collected for U-Pb zircon dating at

41o 58’ 03” N and 118o 42’ 14” W (location shown in Figure 1.3 and Plate 1). This sample

contains 40% phenocrysts, including 30% plagioclase and 10% igneous amphibole (now

replaced by metamorphic biotite and epidote). The groundmass consists of 40%

plagioclase, 10% quartz, 10% metamorphic biotite and epidote. Zircons were extracted

from this sample using techniques described above in a previous section. J.E. Wright spot

analyzed several zircon grains (10-15) on the Super High Resolution Ion Microprobe

(SHRIMP) at Stanford University. This sample yields an age of 176 +/- 3 Ma. Using the

ages provided by the time scale of Gradstein and others (1994), this sample is Bajocian

(Middle Jurassic).

Pueblo Mountain Hypabyssal Andesite

The Pueblo Mountain hypabyssal andesite is a nearly circular body north of the

Pueblo Mountains shear zone boundary (Fig. 1.3) that is unaffected by shear zone-related

deformation. Rocks of the Pueblo Mountain hypabyssal andesite show two types of

textures: some are strongly porphyritic in texture while others have a slightly porphyritic

69

to nearly equigranular texture. Despite these textural differences, all samples from the

Pueblo Mountain hypabyssal andesite share a common mineralogy.

Porphyritic samples contain from 45 to 75% phenocrysts, including euhedral

plagioclase (most common) that ranges in length from 1-10 mm, euhedral to subhedral

amphibole (now replaced largely or entirely by opaques) that average 1 mm in size, and

subhedral clinopyroxene, 1-2.5 mm in size (now replaced partially or entirely by

actinolite and epidote). Groundmass of the porphyritic rocks consists of abundant small

plagioclase laths, with minor quartz, potassium feldspar, oxides, and metamorphic

chlorite and epidote. The nearly equigranular rocks of the Pueblo Mountain hypabyssal

andesite have a similar overall mineralogy to the porphyritic rocks, but they have a fine-

grained subhedral granular texture with few phenocrysts and no plagioclase microlites.

Catlow Creek Quartz Monzodiorite Stock

The Catlow Creek stock is located north of the Pueblo Mountains shear zone (Fig.

1.3) and is unaffected by shear zone-related deformation. Petrographic analysis of the

Catlow Creek stock shows that this intrusion displays a spectrum of textures ranging

from medium-grained and equigranular to a porphyritic texture with a fine-grained

groundmass. In outcrop, these textural variants typically grade and/or dike into each

other. Based on this observation, they appear to be part of a single magmatic body.

Catlow Creek samples displaying an equigranular texture typically contain 65-

70% euhedral plagioclase, 15% anhedral potassium feldspar, 10% interstitial quartz,

about 3% subhedral clinopyroxene (now mostly replaced by metamorphic actinolite +/-

epidote), about 3% euhedral biotite (locally replaced by metamorphic chlorite), and 2%

opaques. Crystals in these rocks vary in size from 1-6 mm. Porphyritic Catlow Creek

rocks contain 10-15% phenocrysts (1-5 mm in size) in a microlitic to subhedral granular

groundmass. Phenocrysts include plagioclase and clinopyroxene (locally replaced, partly

70

to completely, by metamorphic actinolite +/- oxides +/- epidote) plus some combination

of potassium feldspar +/- biotite, and a locally present unidentified phase now replaced

by metamorphic mica and actinolite. Groundmass is these porphyritic rocks consists of

40-50% plagioclase laths, 15-25% potassium feldspar, 5-10% quartz, up to 10%

clinopyroxene (partly to totally replaced actinolite) and biotite (fresh), 2% oxides, and up

to 25% metamorphic chlorite, epidote, and actinolite.

The Catlow Creek stock was dated by Brown (1996) using U-Pb zircon analysis.

Her data yielded a discordant age of 179+/-2 Ma. This age is Aalenian (early Middle

Jurassic) according to the time scale of Gradstein and others (1994).

Diamond Inn Hypabyssal Rhyolite

The Diamond Inn hypabyssal rhyolite intrudes the Strawberry Butte pluton and

the Tuff unit (Fig. 1.3). The map pattern of this hypabyssal intrusion is highly irregular

(Fig. 1.3), and dikes of it locally cut across the Strawberry Butte pluton. Rocks

surrounding the Diamond Inn stock are affected by a large zone of alteration characterized

by abundant silicification and replacement of phenocryst voids by iron and / or

manganese oxides. The Diamond Inn hypabyssal rhyolite is contained entirely within the

Pueblo Mountains shear zone (Fig. 1.3), and is generally foliated and recrystallized.

This intrusion is porphyritic and contains 15-20% phenocrysts of quartz, 1-2 mm

in size, in a much finer-grained groundmass. Although obviously originally phenocrysts,

the large quartz grains are now polycrystalline reflecting dynamic recrystallization during

shear zone deformation. The groundmass of the stock consists of crystals < 0.5 mm,

including 35-50% quartz, 30-40% metamorphic muscovite, up to 5% plagioclase, and up

to 1% opaques. In assigning this rock a composition, modal mineral percentages were

calculated under the assumption that metamorphic muscovite is likely derived from the

recrystallization of igneous potassium feldspar.

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A U-Pb geochronology sample was collected from the center of this small stock at

42o 00’ 37” N and 118o 39’ 47” W (Fig. 1.3 and Plate 1). This sample consists of 20%

quartz in a groundmass of 35% quartz, 40% metamorphic muscovite, up to 4%

plagioclase, and up to 1% opaques. Zircon crystals were extracted from this sample using

techniques described above and were spot analyzed on the Super High Resolution Ion

Microprobe (SHRIMP) at Stanford University by J.E. Wright. According to J.E. Wright

the sample yields an age of 160 +/- 2 Ma. This age is Callovian (late Middle Jurassic)

according to the time scale of Gradstein and others (1994).

Cowden Creek Quartz Diorite - Tonalite

The Cowden Creek pluton is heterogeneously deformed in the Pueblo Mountains

shear zone, with the most prominent foliation near the faulted northwest boundary with

the Tuff unit (Fig. 1.3). Mapping shows that the Cowden Creek pluton is characterized

by two intermingled magmatic phases. Approximately 10% of the intrusion is a fine

grained mafic phase and the remaining 90% of the intrusion is a medium grained more

felsic phase.

The felsic phase likely originally had a subhedral granular texture, with crystals 1-

5 mm in size, but is now generally porphyroclastic with a fine-grained matrix due to

dynamic recrystallization. The composition of most samples of this phase is 50-60%

porphyroclastic plagioclase, 10-15% fine-grained quartz, ~25-30% biotite and 10-15%

metamorphic epidote. Much of the biotite in these rocks occurs as fine-grained aggregates

of crystals that are aligned with the foliation: this biotite is the product of dynamic and

metamorphic recrystallization and is not igneous. However, some biotite (usually 5-10%)

occurs as large (1-2 mm), randomly oriented, single crystals that are inferred to be of

igneous origin. Further clues to the original mafic content of the felsic phase comes from a

nearly undeformed sample, which contains, in addition to 10% quartz and 50%

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plagioclase, 15% large (1-2 mm) igneous hornblende crystals, only ~5% large igneous

biotite, and the remainder 10% epidote and 10% fine-grained recrystallized biotite. This

suggests that the felsic phase was originally a biotite hornblende quartz diorite.

The mafic phase consists of crystals mostly < 1 mm in size, although locally

phenocrysts of plagioclase up to 1-2 mm are present. The overall composition of these

rocks is 10-15% quartz, 0-30% plagioclase, 30-50% biotite and 25-30% epidote. It is the

abundance of epidote and biotite that give this phase its distinctive darker color. Biotite is

everywhere fine-grained and it is difficult to determine on textural grounds whether it is of

igneous or metamorphic origin, or some combination. However, epidote can be seen in

some samples to be completely replacing relict euhedral plagioclase. This suggests that

much of the epidote in the mafic phase is after plagioclase and that the original plagioclase

content of this phase was higher that the currently observed 0-30%. A best estimate

based on thin section petrography is that plagioclase abundance was likely originally ~30-

40%. This makes the mafic phase tonalitic with an unknown original mafic mineralogy,

but possibly hornblende and biotite as in the felsic phase.

In outcrop, evidence of mutual diking between the mafic and felsic phases is

common. These dikes are typically < 1 m wide. Chilled margins along distinct edges of

these dikes show that locally the mafic phase intrudes the felsic phase whereas the felsic

phase elsewhere intrudes the mafic phase. In places, the edges of these dikes are more

diffuse and clearly show an intermingled comagmatic relationship (Fig. 2.22a-b). These

field relations suggest that the mafic and felsic phases reflect magma mingling.

A sample was taken from the felsic phase for U-Pb zircon geochronology. This

sample consist of 50% phenocrysts, including 45% plagioclase, 4% biotite (distinguished

from metamorphic biotite by its large and equidimensional crystal shape) and 1% quartz.

The groundmass consists of 25% quartz, 15% metamorphic biotite and 7% epidote, and

3% oxides. Sample location (Fig. 3 and Plate 1) is 41o 57’ 46” N and 118o 40’ 37” W.

73

Figure 2.22a. Outcrop photograph showing intermingling of Cowden Creek felsic phases

(white) with the mafic phase (dark grey). Knife is 9 cm long.

74

Figure 2.22b. Outcrop photograph of dikes of the mafic phase (dark grey) intruding the

felsic phase (white and light grey) of the Cowden Creek pluton. Mafic dikes have chilled

margins (1-2 cm wide). Hammer in bottom of photograph is 28 cm long.

75

Zircon crystals were extracted from this sample using techniques described in a

previous section. Zircon crystals (10-15) were spot analyzed on the Super High

Resolution Ion Microprobe (SHRIMP) at Stanford University by J.E. Wright and yielded

an concordant age of 188 +/- 2 Ma. This age is Toarcian (late Early Jurassic) according to

the time scale of Gradstein and others (1994).

GEOCHEMISTRY OF IGNEOUS ROCKS AND TECTONIC SETTING OF

MAGMATISM

Samples of lava, plutonic rocks, and felsic tuff were collected from the Pueblo

Mountains in order to establish the compositions of igneous rocks and to evaluate the

paleotectonic setting of magmatism. Sample sites are shown in Figure 1.3 and Plate 1.

Fourteen samples were collected from rocks of the Pueblo terrane; six from andesite lava

of the Andesite unit, four from the Strawberry Butte quartz diorite, two from the

Diamond Inn hypabyssal rhyolite stock, and two from felsic tuffs (one crystal and one

ash tuff – both muscovite-rich) of the Tuff unit. The andesites and Strawberry Butte

pluton were analyzed with the goal of using their geochemistry to define the paleotectonic

setting of magmatism in the Pueblo terrane, and were chosen because Andesite unit lavas

are the most widespread lavas and the Strawberry Butte is the largest pluton in the

terrane. Samples from the Diamond Inn stock and the Tuff unit were analyzed to better

define the original igneous composition (rhyolitic or dacitic) of recrystallized felsic rocks

in the terrane. One sample from the felsic phase of the Cowden Creek pluton was also

analyzed in order to evaluate the origin of this fault bounded pluton.

When collecting geochemistry samples, care was taken in choosing fresh

(unweathered) rock with no evidence of alteration and showing minimal strain. All

samples were powdered in a steel shatterbox and analyzed by Activation Laboratories

Limited in Ontario, Canada. Major elements were analyzed by lithium metaborate /

76

tertraborate fusion inductively coupled plasma technique. Trace elements were analyzed

by fusion inductively coupled plasma-mass spectrometry technique. Data from these

analyses is shown in Table 2.1.

Typical silica values (Table 2.1) for Andesite unit lavas range from 50-56%.

According to Raymond (1995), these values are characteristic of basalts and basaltic

andesites; however, those compositions are not consistent with the mineralogy and

phenocryst assemblage of Andesite unit lavas. This suggests some mobility of silica

related to later metamorphism. Likewise, the Strawberry Butte and Cowden Creek

samples typically have silica values ranging from 63-66% (Table 2.1), which is slightly

higher than expected based on the mineralogy of these intrusions; typical silica content for

quartz diorite and tonalite is roughly 58-62% (Raymond, 1995). Silica values of the two

samples from the Diamond Inn stock are 68% and 75%. This is nearly consistent with the

mineralogy of the Diamond Inn stock because rhyolitic rocks have a 69 – 77% silica range

(Raymond, 1995). The two felsic tuff samples have silica values of 67 and 70%. These

values are consistent with their rhyolitic to dacitic mineralogy.

Based on the above analysis, it appears that silica has been mobile in rocks of the

Pueblo Mountains. Some other major elements also have slightly anomalous values given

the mineralogy of the various analyzed rocks. These include unusually high Mg and Fe

values in andesites and unusually high Al concentrations and slightly lower Na and K

concentrations in Diamond Inn rhyolite samples.

Because major elements in the samples show some anomalous values, and because

major elements are commonly mobile during metamorphism (Cann, 1970; Humphris and

Thompson, 1978; Hanson, 1980; Shervais, 1982), these elements cannot

be reliably used to evaluate petrogenesis of the metamorphosed igneous rocks in the

Pueblo Mountains. However, most trace elements and rare earth elements (REE) are

considered immobile during greenschist facies metamorphism (Cann, 1970; Humphris and

TABLE 2.1. Major, trace, and rare earth element analysis from andesitic, dioritic and rhyolitic rocks, Pueblo terrane

Strawberry ButteAndesite unit lavas quartz diorite

sample PMW -12 PMW-19 PMW-123 PMW-232 PMW-246 PMW-271 PMW-90

SiO2 51.06 51.44 52.07 51.95 49.84 55.75 65.35TiO2 1.15 1.07 1.09 1.09 1.01 0.91 0.38Al2O3 20.17 17.58 17.79 17.74 17.82 17.25 17.01Fe2O3 7.88 10.93 9.38 11.38 10.15 7 4.19MnO 0.04 0.06 0.19 0.05 0.11 0.16 0.02MgO 3.96 4.36 4.01 4.22 5.45 2.73 1.34CaO 5.17 4.1 4.57 4.93 8.72 5.45 1.59Na2O 4.54 5.24 4.73 4.84 3.29 4.34 6.49K20 3.99 1.71 3.78 1.13 1.59 3.14 2.95P205 0.57 0.55 0.44 0.34 0.38 2.99 0.18LOI 1.9 2.5 2.25 2.63 1.88 2.99 0.99Total 100.43 99.53 100.3 100.29 100.24 100.11 100.49

Trace elements in ppmSr 822 675 640 587 1010 898 436K 33100 14100 31300 93700 13100 26000 24400Rb 207 57 75 28 77 49 31Ba 743 283 1,320 442 465 710 896Th 10.85 4.02 4.53 1.55 1.94 12.08 2.65Ta 0.62 0.34 0.34 0.19 0.19 0.53 0.21Nb 9.62 4.13 4.85 2.75 2.43 7.38 1.72Ce 34.82 40.25 41.52 25.97 27.64 49.09 29.97P 2488 2401 1921 1484 1659 1305 785Zr 189 91 125 82 70 208 87Hf 5.44 2.37 3.45 2.35 1.98 5.3 2.59Sm 5 4.85 5.04 3.76 3.65 5.01 2.55Ti 6889 6398 6554 6560 6068 5468 2290Y 28.76 22.8 24.41 20.57 17.88 22.55 7.79Yb 2.83 1.93 2.29 1.86 1.58 1.98 0.72Sc 25 22 26 27 27 18 6Cr 40 25 43 46 108 23 27La 16.56 18.34 19.66 12.03 12.9 23.91 17.23Pr 5.15 5.51 6.04 3.97 4.18 6.98 3.88Nd 20.74 22.62 23.11 16.25 16.12 25.59 13.79Eu 1.12 1.42 1.56 1.34 1.39 1.46 0.84Gd 4.54 4.72 4.6 3.49 3.22 4.24 1.96Tb 0.78 0.64 0.74 0.6 0.53 0.65 0.27Dy 4.8 3.78 4.4 3.51 3.13 3.79 1.43Ho 1 0.7 0.86 0.74 0.64 0.75 0.28Er 2.92 2.09 2.51 2.13 1.78 2.17 0.78Tm 0.44 0.29 0.36 0.3 0.25 0.31 0.11Lu 0.43 0.28 0.34 0.29 0.24 0.31 0.1

Total Fe as Fe2O3.

TABLE 2.1. Continued.

Cowden Creek Diamond InnStrawberry Butte quartz diorite quartz diorite Tuff unit felsic tuff hypabyssal rhyolite

PMW-145 PMW-198 PMW-SBPZr PMW-CDPZr PMW-35 PMW-76 PMW-273 PMW-FSZr

63.46 66.04 58.1 63.12 70.73 67.09 68.07 74.740.47 0.44 0.93 0.47 0.45 0.38 0.4 0.41

16.59 16.03 17.25 16.62 19.04 15.34 15.54 17.014.74 3.84 6.99 4.13 0.67 3.41 3.48 0.610.13 0.05 0.12 0.09 0 0.08 0.01 01.65 2.16 2.89 1.91 0.6 1.59 5.35 0.223.17 2.59 5.66 3.44 0.02 2.29 0.32 0.034.67 4.16 4.46 3.65 0.38 3.85 0.45 1.083.89 3.58 2.28 4.69 5.98 3.01 2.94 3.940.21 0.19 0.35 0.18 0.02 0.2 0.14 0.121.49 1.31 1.39 1.79 2.53 3.16 3.55 2.28

100.46 100.38 100.43 100.07 100.41 100.4 100.24 100.43

605 496 580 692 - - - -23200 29700 18900 38900 - - - -

65 57 107 74 - - - -1,150 1,010 647 1,180 - - - -4.83 7.48 25.31 4.58 - - - -0.34 0.46 0.92 0.35 - - - -2.92 3.54 16.41 2.5 - - - -

23.18 25.95 52.13 32 - - - -916 829 1528 786 - - - -102 121 507 106 - - - -2.73 3.53 12.71 2.84 - - - -2.4 2.57 5.58 2.78 - - - -

2794 2638 5600 2842 - - - -11.18 12.8 28.43 11.9 - - - -1.13 1.35 2.86 1.21 - - - -10 11 18 11 - - - -50 156 55 31 - - - -

11.99 13.39 24.99 17 - - - -3.22 3.56 7.36 4.2 - - - -

11.87 12.73 26.83 14.5 - - - -0.79 0.81 1.44 0.949 - - - -2.15 2.17 4.86 2.36 - - - -0.32 0.35 0.78 0.35 - - - -1.82 2.12 4.78 2.02 - - - -0.39 0.46 0.99 0.41 - - - -1.18 1.34 2.86 1.2 - - - -0.17 0.2 0.43 0.175 - - - -0.18 0.22 0.45 0.194 - - - -

79

Thompson, 1978; Hanson, 1980; Sun, 1980; Wood, 1980; Pearce, 1982; Shervais, 1982),

such as has affected the rocks of the Pueblo Mountains. They can therefore effectively be

used to evaluate petrology and tectonic setting of magmatism for these rocks.

Trace element and REE data from selected plutonic and volcanic samples of the

Pueblo Mountains are listed in Table 2.1 and plotted on tectonic discrimination diagrams

(La vs Nb; La vs Th; Th – Hf/3 – Ta; Ti vs Zr; and Th/Yb vs Ta/Yb) in Figure 2.23, a

spider diagram (following Pearce, 1983) in Figure 2.24, and a REE plot in Figure 2.25.

Only andesitic and dioritic samples were used to plot on these diagrams because crystal

fractionation processes in more felsic rocks may substantially influence trace and REE

element values, thereby obscuring information about the source. According to Pearce

(1982), the various elements plotted in Figures 2.23-2.25 are unlikely to be significantly

affected by crystal fractionation processes for intermediate composition igneous rocks.

Discrimination diagrams (Fig. 2.23) are useful for establishing the tectonic setting

of magmatism. According to these diagrams, igneous rocks from the Pueblo terrane plot

entirely or largely in the volcanic arc field, indicating that they formed at a convergent

plate boundary due to subduction-related magmatism. Several of the plots also

discriminate between the tholeiitic (low-K), calc-alkaline (med-K), and shoshonitic (high-

K) series of arc rocks. According to these diagrams (Fig. 2.23d-e), igneous rocks from the

Pueblo terrane are calc-alkaline to shoshonitic.

Spider diagrams also indicate that the volcanic and plutonic rocks were generated

at a convergent margin (Pearce, 1982; Gill, 1987; Wilson, 1989). Specifically, they show

enrichment in Sr, Rb, Ba and Th, display a strong Nb-Ta trough, and are relatively

depleted in the less compatible elements Ti, Y, and Yb (Fig. 2.24a). For comparison with

Pueblo terrane rocks, typical calc-alkaline to shoshonitic oceanic and continental arc lavas,

and island arc tholeiites are plotted on Figure 2.24b (data from Pearce, 1983 and Wilson,

1989). The principal differences between the oceanic and continental arc rocks on spider

80

Figure 2.23. Trace element discrimination diagrams of Pueblo Mountain lavas and

plutons. Dark circles = Andesite unit lavas, dark squares = Strawberry Butte pluton,

open squares = Cowden Creek pluton. MORB: mid-ocean ridge basalt field, locally

divided into enriched (E-MORB) and tholeiitic or normal (N-MORB) subfields. Volcanic

arc field locally divide into tholeiitic, calc-alkaline, and shoshonitic subfields. Plots (a) and

(b) are La-Nb and La-Th plots, respectively, after Gill (1987). (c) Ti-Zr plot after Pearce

(1982). (d) Th-Hf/3-Ta plot after Wood (1980), values in ppm. (e) Discrimination

diagram (after Pearce, 1982) based on Th - Ta covariations using Yb as a normalizing

factor. TH = tholeiitic, CA = calc-alkaline, TR = transitional, SHO = shoshonitic, ALK =

alkaline.

TaTh

Hf/3

prim

ativ

e ar

cth

olei

ites

calc

-alkal

ine

arc

rock

s E-MORB

WPB

N-MORB

(d)

orogenic andesite

E-MORBN-MORB

high-K

Th (ppm)

30

20

10

2 4 6 8 10 12 14 16 18 20 22 24

low-K

med-K

(b)

La(ppm)

La(ppm)

Nb (ppm)

30

20

10

2 4 6 8 10 12 14 16

orogenic

andesite

high-K

med-K

N-MORB

E-MORBlow-K

(a)

SHO

CA

CA

TH TH

TR

ALK

TR

10

1.0

0.1

0.01

0.01 0.1 1.0 10

(e)

MORB

WPB

Vo

lca

nic

a r

c

10000

100010 100

Ti(ppm)

Zr (ppm)

within platelavas

arc lavas

MORB

(c)

Th/Yb(ppm)

Ta/Yb(ppm)

82

Figure 2.24. MORB normalized trace element spider diagrams for: (a) Pueblo Mountain

Andesite unit lavas (dashed box pattern), Strawberry Butte pluton (light grey), and

Cowden Creek pluton (solid black line) (normalizing values from Pearce, 1983). (b)

Typical calc-alkaline to shoshonitic continental arc lavas (horizontal lines), oceanic arc

lavas (vertical lines), and tholeiitic island arc lavas (open pattern). Data from Pearce

(1983) and Wilson (1989).

1

100

10

.1

Ba

Th Nb

Zr Sm

Ti

YCeTa

P Hf YbRb

K

Sr

Andesiteunit lavas

StrawberryButte pluton

CowdenCreekpluton

1

100

10

.1

Ba

Th Nb

Zr Sm

Ti

YCeTa

P Hf YbSr Rb

calc-alkaline toshoshonitic continental

arc lavas

calc-alkalineto shoshonitic

ocean arclavas

Rock

/MO

RB

(a)

(b)

Rock

/MO

RB

(felsic phase)

island arctholeiites

84

Figure 2.25. Chondrite normalized REE diagrams (normalization values from Nakamura,

1974). (a) Pueblo Mountains geochemistry samples: Andesite unit lavas (within dashed

pattern), Strawberry Butte pluton (light grey) and Cowden Creek pluton (solid black

line). (b) Some typical medium to high-K continental arc lavas (horizontal lines) and

medium to high-K island arc lavas (vertical lines). Data from Gill (1987) and Wilson

(1989).

100

10

1La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb

Rock

/Cho

ndrite

s

Andesiteunit lavas

StrawberryButte pluton

Cowden Creekpluton

(a)

100

10

1La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb

Rock

/Cho

ndrite

s

med to high-Kisland arc lavas

med to high-Kcontinental arc

lavas

(b)

(felsic phase)

86

diagrams are that the oceanic arc rocks have more pronounced negative Ta, Nb, Zr and Hf

troughs and somewhat lower values of Ce, P, Sm, Ti, Y and Yb. Also, tholeiitic island arc

rocks generally have lower values for all trace elements as compared to calc-alkaline to

shoshonitic oceanic and continental arc lavas, and they have no pronounced troughs (Fig.

2.24b). Pueblo terrane samples display trends that are clearly not tholeiitic. Instead they

are transitional between the trends for calc-alkaline to shoshonitic oceanic and continental

arc rocks (compare Figs. 2.24a-b). This possibly suggests that Pueblo terrane igneous

rocks represent a transitional arc environment, in which, the arc was constructed on crust

transitional in thickness and composition between oceanic and continental crust.

REE data from Pueblo terrane rocks are shown in Figure 2.25a and compared with

REE data for typical medium to high-K island arc lavas and continental arc lavas in Figure

2.25b (data from Gill, 1987 and Wilson, 1989). The Pueblo terrane rocks are enriched in

light REE, have heavy REE roughly 3 to 15 times chondrites and show no Eu anomaly

(Fig. 2.25a). According to Gill (1987) and Wilson (1989), these characteristics are typical

of arc lavas, consistent with the conclusions based on other trace element data. It is not

possible to clearly distinguish continental versus oceanic medium to high-K arc lavas

based on their REE signature (Fig. 2.25b), and the Pueblo terrane rocks have a REE

signature compatible with either setting.

Geochemical data from Pueblo terrane igneous rocks lead to the following

conclusions. First, the Pueblo terrane lavas and plutonic rocks are calc-alkaline to

shoshonitic, and formed above a subduction zone in a magmatic arc (Fig. 2.23a-e). Second,

based on the spider diagram trends in Figure 2.24a-b, these rocks appear transitional in

geochemistry between continental arc and oceanic arc rocks. Finally, Andesite unit lavas

and the Strawberry Butte pluton have very similar geochemical signatures, consistent

with prior conclusions based on petrography and age data that they are related

components of a single volcanic-plutonic complex (Figs. 2.23, 2.24, and 2.25).

87

The Cowden Creek pluton has a similar geochemical signature to the Pueblo

terrane igneous rocks (Figs. 2.23, 2.24, and 2.25). Based on discrimination plots and

spider diagrams (Figs. 2.23-2.24), it also formed in a magmatic arc environment and is

calc-alkaline to shoshonitic.

RELATIONSHIP BETWEEN INTRUSIVE AND VOLCANIC UNITS

Of the five intrusive bodies found in this study, two (the Cowden Creek pluton

and Diamond Inn stock) have ages that are significantly older and younger, respectively,

than the ages of the Tuff and Andesite units. The Cowden Creek pluton is 188 +/- 2 Ma

and the Diamond Inn stock is 160 +/- 2 Ma, versus 179 +/- 2 Ma and 176 +/- 2 Ma for

the Tuff and Andesite units, respectively. These two intrusions are therefore not directly

related to the volcanic units and will not be further discussed in this section.

Of the remaining intrusions (the Strawberry Butte pluton, Pueblo Mountain and

Catlow Creek stocks) the Strawberry Butte has been dated at 176 +/- 3 Ma and the

Catlow Creek was dated by Brown (1996) at approximately 179 Ma (discordant zircon

age data). The age of the Strawberry Butte is identical within error to the age of the two

volcanic units, and overlaps the most with the Andesite unit (176 +/- 2 Ma). This

suggests that the Strawberry Butte pluton may be comagmatic with the volcanic units,

specifically the Andesite unit. The Catlow Creek ages (although discordant) overlap

within error the ages of the volcanic units suggesting that it also may be comagmatic with

the volcanic units. Further analysis of these possibilities, and the possible relation

between the Pueblo Mountain stock and the volcanic units can be evaluated by

considering intrusive relations and composition.

In terms of intrusive relations, the Strawberry Butte pluton intrudes the Tuff unit,

and thus could be comagmatic with this unit or the Andesite unit, as suggested above. The

88

Catlow Creek and Pueblo Mountain stocks intrude only the Andesite unit. They could be

comagmatic with the Andesite unit, but not with the Tuff unit.

The Strawberry Butte quartz diorite consists primarily of plagioclase, lesser

hornblende and rare quartz phenocrysts, in a groundmass of plagioclase, quartz (<5%)

and metamorphic epidote and biotite. This composition and phenocryst assemblage is

similar to the dominant rock type in the Andesite unit (andesite lava). This suggests these

units are comagmatic, which agrees with the age data for both units, and with the map and

intrusive relations between the two units. The Strawberry Butte pluton is not very

similar in composition to the felsic tuffs that dominate the Tuff unit. This pluton contains

hornblende, less quartz, and is overall more mafic in composition than Tuff unit felsic

tuffs. Therefore, it is unlikely that this pluton is comagmatic with the Tuff unit. This

conclusion is consistent with the age relations and intrusive map relations between the

two units.

The Pueblo Mountain hypabyssal andesite (equivalent to shallow level diorite)

consists of plagioclase, lesser hornblende and minor clinopyroxene phenocrysts set in a

groundmass of abundant plagioclase, minor quartz, metamorphic chlorite and epidote, and

rare potassium feldspar. This composition has some similarities with Andesite unit lavas.

Both rock types are clearly porphyritic, have abundant plagioclase and lesser hornblende

phenocrysts, and minor quartz in the groundmass. Also, hand samples of both rock types

have a similar appearance with the only difference being that the plagioclase phenocrysts

are larger in the Pueblo Mountain stock than in Andesite unit lavas. The main difference

in these two rock types is that the Pueblo Mountains stock contains phenocrystic

clinopyroxene and minor potassium feldspar in the groundmass. Based on the above

similarities and differences, the Pueblo Mountain stock could be comagmatic with the

Andesite unit, but this can not be concluded for certain.

89

The Catlow Creek stock is a quartz monzodiorite varying in texture from

porphyritic to equigranular and its phenocrysts assemblage consists of plagioclase with

lesser hornblende, and sometimes minor clinopyroxene and biotite. Variable amounts of

plagioclase and quartz are present in the groundmass along with lesser metamorphic

epidote, biotite, chlorite and actinolite. The Catlow Creek stock bears no similarity with

Andesite unit lavas, the Strawberry Butte pluton, or the Pueblo Mountains stock.

However, this stock does have some similarities with Andesite unit volcaniclastic rocks.

Both rock types contain similar amounts of potassium feldspar and quartz. Preliminary

age data (179 +/-2 Ma) from Brown (1996) suggests that the Catlow Creek stock is either

the same age, or slightly younger, than the Andesite unit. Based on these similarities in

composition and age, the Catlow Creek stock could be comagmatic with the source of

rhyolitic volcaniclastics in the Andesite unit. However, this conclusion is speculative.

STRATIGRAPHIC AND MAGMATIC SUMMARY AND INTERPRETATION OF

THE PUEBLO TERRANE

Data presented in the previous sections indicate that pre-Cenozoic rocks in the

eastern Pueblo Mountains consist of a large Middle Jurassic (Aalenian to Callovian; 196-

160 Ma) volcanic-plutonic complex (the Pueblo terrane) and an Early Jurassic (Toarcian;

188 Ma) pluton ( the Cowden Creek pluton) that is in fault contact with the Pueblo

terrane (Fig. 1.3 and Plate 1).

The Pueblo terrane consists of two volcanogenic units, the Tuff unit and the

Andesite unit, that are intruded by one large pluton (Strawberry Butte) and three small

stocks. The Tuff unit has been dated at 179 +/- 2 Ma and consists of rhyolitic to dacitic

pyroclastic ash flows and pyroclastic surge deposits, lavas flows, and immature

epiclastics (volcanic sandstones and breccias). The Andesite unit has an age of 176 +/- 2

Ma and consists mostly of andesitic lava flows, with much less common volcaniclastic

90

rocks that range in composition from andesitic to rhyolitic. The latter include welded

pyroclastic ash flow tuffs, block and ash flow deposits, volcanic breccias probably

deposited from debris flows mobilizing primary volcanic debris, and pyroclastic surge

deposits. These two middle Jurassic volcanic units are in gradational contact with one

another and likely interfinger laterally, although precise stratigraphic relations are

obscured by deformation in the Pueblo Mountains shear zone. The original stratigraphic

thickness of the volcanic units cannot be determined because of later deformation and lack

of exposure of either the base or top of the units. Calculated structural thicknesses for

each unit are at least 800 m for the Andesite unit and 2000 m for the Tuff unit.

The Strawberry Butte quartz diorite pluton has been dated at 176 +/- 3 Ma and is

likely comagmatic with the Andesite unit based on similarities in age, mineralogy and trace

and REE geochemistry. The small Catlow Creek quartz monzodiorite stock and Pueblo

Mountain hypabyssal andesite stock could be comagmatic with the Andesite unit, based

on mineralogy and preliminary age data (Brown, 1996), but this conclusion cannot be

confirmed without additional age control. The Diamond Inn hypabyssal rhyolite stock

has been dated at 160 +/- 2 Ma and is significantly younger than other rocks in the Pueblo

terrane.

Geochemical data from andesite lavas of the Andesite unit and plutonic rocks of

the Strawberry Butte pluton indicate that Pueblo terrane magmatism took place in a

volcanic arc setting. Several relations argue that this arc was continental (i.e., built above

sea level on thickened crust of continental or transitional character; Gill, 1987). (1) Pueblo

terrane igneous rocks range from andesitic to rhyolitic, with no basaltic compositions

represented. (2) Pueblo terrane igneous rocks have a calc-alkaline to shoshonitic

composition, with no tholeiitic rocks present. (3) Spider diagrams indicate that Pueblo

terrane igneous rocks have a geochemical signature transitional between that of oceanic

and continental arcs. (4) Stratigraphic and facies relations indicate that the volcanic strata

91

of the Pueblo terrane were deposited in a subaerial setting, and pyroclastic deposits are

common in the volcanic stratigraphy. According to Pearce (1982), Gill (1987), and Wilson

(1989), these features are most characteristic of continental arc assemblages. Oceanic arcs,

in contrast, are typically dominated by basaltic to basaltic composition igneous rocks,

with a greater predominance of tholeiitic rocks relative to calc-alkaline and shoshonitic

rocks, a greater predominance of flows relative to pyroclastic deposits, and a largely

marine environment of deposition for volcanogenic strata.

In addition to the geochemical relationships discussed above, the distribution of

volcanic facies in the Pueblo terrane provides further insight concerning the tectonic and

depositional setting of magmatism in the Pueblo terrane. Figure 2.26 is a generalized facies

model for continental stratovolcanoes (from Cas and Wright, 1987) that shows the spatial

relationship of proximal and distal volcanogenics to their volcanic center. Distal ring plain

rocks are thick successions of immature volcanic detritus deposited in either, or both,

alluvial or shallow marine settings. These epiclastic deposits are contemporaneous with

proximal near-vent cone facies rocks. Proximal facies are dominated by lavas, domes, and

shallow intrusives commonly dacite to basaltic andesite. Epiclastic or pyroclastic volcanic

breccias are intimately associated with lava flows in the near-vent cone facies. The cone is

flanked by various pyroclastic and epiclastic deposits that can, but may not, have a wide

compositional range and eruptive style. According to Reading (1996), lavas and welded

tuffs will represent much of the sedimentary and volcanic rock record in proximal (near-

vent) settings.

In the northwestern Pueblo terrane (the Andesite unit), andesite lavas dominate

and are interbedded with minor pyroclastic deposits (block and ash flows, pyroclastic ash

flows, and pyroclastic surges). In the southeastern Pueblo terrane (the Tuff unit),

pyroclastic pumice and ash flows (massive tuff facies) are the most abundant rock types.

These are interlayered with lesser pyroclastic surge deposits (plane-laminated tuff facies),

92

Figure 2.26. Generalized facies model for continental stratovolcanoes (modified from Cas

and Wright, 1987). Lavas and pyroclastics dominate the near vent cone facies and

epiclastic volcanic debris dominates the distal ring plain environment.

0 5 10 km

proximal near vent cone faciesdistal ring plain environment

lavas, domes, and intrusive feeder bodies

epiclastic sediments contemporaneous with pyroclastics

pyroclastic flow and volcaniclastic mass-flow deposits

basement

subsurface intrusives

94

dacitic lavas (lava flow facies), and volcaniclastic rocks (volcanic breccia and sandstone

facies).

When these Pueblo terrane rocks are compared with facies descriptions in the

continental stratovolcano model (Fig. 2.26; Cas and Wright, 1987), Andesite unit

volcanogenic rocks match well with the proximal near-vent cone facies assemblages (the

solid black pattern in Figure 2.26). Tuff unit tuffaceous and volcaniclastic rocks match

well with proximal facies rocks deposited on the flanks of the volcano (white in Figure

2.26). Notice also, in Figure 2.26, that pyroclastic flow rocks (white color) are

interfingered with near-vent lavas (solid black). In the Pueblo terrane a similar pattern is

seen: Andesite unit lavas contain minor tuffaceous and volcaniclastic layers, and Tuff unit

pyroclastic flows contain some lava flows. In summary, the distribution of volcanic facies

in the Pueblo terrane matches well with that of a continental stratovolcano model (Fig.

2.26; Cas and Wright, 1987), consistent with a conclusion that the Pueblo terrane is a

continental arc whose volcanic stratigraphy was deposited in a subaerial setting.

According to Cas and Wright (1987), rapid and inconsistent lithologic and compositional

changes in vertical succession, and somewhat rapid lateral lithologic changes, are typical

features of continental stratovolcano deposits. In the Pueblo terrane, near-vent facies

Andesite unit rocks (mainly flows and lesser coarse pyroclastics) quickly give way to

pumice-rich tuffs in the northern Tuff unit (Fig. 1.3; Plate 1). Within the Tuff unit,

volcaniclastic rocks and lava flows are most common in the northern Tuff unit and are less

common in the southern Tuff unit (Plate 1). In the Tuff unit, its four facies (the massive

tuff, plane-laminated tuff, lava flow, and volcanic breccia and sandstone facies) are

randomly interlayered and, in vertical succession, any one facies may change abruptly

into any other facies. These facies changes are common in the Tuff unit and no discernable

pattern of facies changes is present. The Pueblo terrane volcanic stratigraphy appears to

95

display the rapid, inconsistent vertical and lateral changes in lithology that is expected in

continental stratovolcanoes.

96

CHAPTER 3

STRUCTURAL GEOLOGY AND METAMORPHISM IN THE EASTERN PUEBLO

MOUNTAINS

Two broad periods of deformation have the affected rocks of the Pueblo

Mountains. The earliest period occurred in the Mesozoic and was associated with

development of the Pueblo Mountains shear zone (Fig. 3.1; Plate 1). This deformation

was accompanied by regional metamorphism that affected, to varying degrees, the

Mesozoic volcanic and plutonic rocks of the range. Subsequently, these rocks were

uplifted, eroded, and overlain along a major angular unconformity, by Tertiary strata. The

second broad period of deformation occurred in the Cenozoic and was associated with

localized brittle deformation within Tertiary and Mesozoic rocks, and with normal

faulting along the eastern range front and westward tilting of Tertiary and older rocks. The

main focus of this chapter is the Mesozoic deformation and metamorphism, but the

Cenozoic deformation is also described.

MESOZOIC DEFORMATION AND METAMORPHISM

Mesozoic deformation in the Pueblo terrane produced a ductile shear zone in the

southeastern part of the Pueblo Mountains (Pueblo Mountains shear zone; Fig. 3.1; Plate

1). The shear zone is defined by strongly foliated to mylonitic rocks. Regional

metamorphism accompanied shear zone deformation and also affected the undeformed

rocks northwest of the shear zone.

97

Figure 3.1. Simplified geology of the eastern Pueblo Mountains, northern Pine Forest

Range, and Lone Mountain (location shown in Figure 1.2). PMSZ is the Pueblo

Mountains shear zone. Distribution of Cenozoic rocks and some Cenozoic bedding data

in the Pueblo Mountains from Roback and others (1987). All other geologic data in the

Pueblo Mountains from this study. Geology of the northern Pine Forest Range and Lone

Mountain is from Wyld and Wright (2001). See Plate 1 for more detail.

5‘ 00“

42o

NORTHERNPINE

FORESTRANGE

0 1 kmN

A'

A

B

B'

C '

D '

D

C

JCD

JSB

JSB

JSB

JSB

JSBJSB

JPM

JDI

JAU

JTU

JTU

JTU

JTU

JT

JAU

JAU

Qal

Qls

Qal

Qal

QlsQls

Qls

Qal

Qal

JCT

PUEBLOMOUNTAINS

Strawberry Butte

Denio Basin

Van Horn Basin

Pueblo Mtn.

Denio, NV.

22

19

22

34

2421

23

26

34

23

25

14

30

23

20

21

14

13

22

4142

40

26

49

3931

25

85

35

50

3547

69

49

49

49594669

63

32

6557

69

54

60

52

4364

4555

21

50

5439

21 80

66

30

43

65

44

7035

43

32

36

42 55

53

41

64

78 51 45

41

43

76

74

39

79

80

60

7880

48

59 574555

53

58

46

54

75

54

49

6855

61

59

75

46 30

39

55

6276

5151

45

35

30

5573

40

40

4439

37

3540

7054

2623

25

21

11

61

A’

55‘ 00“

42o

5‘ 00“


AntelopeValley Shear

Zone

55‘ 00“

E“

E’

E

*

*

*

*

*

*

*

*

*

*

*

*

*

*

*

*

denotes zircon sample localitystrike and dip of bedding

strike and dip of S1 foliation

strike and dip of compaction fabric Cenozoic normal fault - ballson downthrown side

ductile fault - teeth on upthrown side

Northwest boundary of CretaceousPMSZ (Pueblo Mountains shear zone)- teeth on shear zone side

Cretaceous shear zones

denotes geochemistry sample locality*denotes argon sample locality

SYMBOLS

LONEMOUNTAIN

PMSZ

PMSZ


58

39

55

6166

76

34

65

30

41 Tertiary

CretaceousGranodiorite plutons(108 - 115 Ma)

Cowden Creek tonaliteto quartz diorite (188+/- 2 Ma)


Alluvium

Landslide


Jura

ssic

Qua

tern

ary

LEGEND

JBZ

JCD

Pueblo Terrane

Andesite unit (176Ma +/- 2 ma)

JAU

Tuff unit (179+/- 2 Ma)

JTU

JSB

JDI

JPM

JCT

Strawberry Butte quartzdiorite (176 +/- 3 Ma)

Diamond Inn hypabyssalrhyolite (160 +/- 2 Ma)

Catlow Creek quartz-monzodiorite (179 +/- 2 Ma)

Jura

ssic

Stratigraphic units

Plutonic units


Early Jurassicplutons (200-185 Ma)

Paleozoic amphiboliteand biotite schist

Triassic pluton(~235 Ma)

Black Rock Terrane

23

Qal

Qls

JCD

JBZ

= structure section lineC C’

Qal

JCD

Qls

JSB

Qal

Qal

Qal

Qls

Qls

JPM

Qal

Qls Qls

99

Pueblo Mountains Shear Zone

The Pueblo Mountains shear zone affects the Tuff unit, Cowden Creek and

Strawberry Butte plutons, and Diamond Inn stock, and the southeasternmost edge of the

Andesite unit (Fig. 3.1; Plate 1; see chapter 2). Strain in the shear zone is heterogeneous

and dies out to the northwest near the Tuff-Andesite unit contact (Fig. 3.1; Plate 1).

Three phases of ductile deformation are recognized in the Pueblo Mountains shear zone.

D1 occurs throughout and is manifested primarily by a mylonitic foliation and lineation.

D2 occurs only locally and is manifested by folds of the D1 foliation with a locally

developed axial planar foliation. D3 is a rarely developed foliation that locally cuts across

D1 and D2 structures. All three phases of deformation are inferred to reflect progressive

deformation in the shear zone rather than separate, unrelated events, because they are all

restricted to the shear zone region (Fig. 3.2) and because they all reflect similar

kinematics.

D1 Structures

The main structures in the Pueblo Mountains shear zone are a well-developed

foliation (S1) and a stretching or mineral lineation (L1). The S1 foliation is generally

mylonitic, and less commonly schistose (non-mylonitic). Where mylonitic, the S1 foliation

can be locally differentiated into an S-C (locations shown in Figure 3.2; Plate 1), or rarely

an S-C’ fabric (terminology of Passchier and Trouw, 1998). Overall, the S1 and L1 fabrics

are defined by the preferred alignment of metamorphic minerals, flattened clasts and

stretched grains (in appropriate rock types), and pressure shadows around

porphyroclasts.

Where only a single S1 plane of foliation is present, it is generally pervasive and

continuous (Figs. 3.3 and 3.4). Where S-C fabrics are developed, both strike and dip in a

similar direction, but the S planes are cross-cut and deformed into a sigmoidal shape by

100

Figure 3.2. Simplified geologic map of the eastern Pueblo Mountains showing the location

of all structures excluding S1 foliation. See lower right of figure for key to symbols and

structures. Light, medium, and dark grey shading correspond to tuffs, intrusives, and

lavas, respectively.

Plutons andhypabyssal rocks

Jura

ssic

Quaternary andTertiary

Key to map units

Andesite Unit

Tuff Unit

**

0 1 km

N

54

51

*

Denotes location of observedS-C fabric (D1)

Denotes location of F2 folds

Denotes location of S3 foliation

SYMBOLS PUE

BLO M

NTS

. SHEA

R ZO

NE

102

Figure 3.3. Planar, penetrative, and well-developed S1 foliation in Tuff unit pumice tuff

outcrop. View of photograph is normal to foliation plane. Lens cap is 5.5 cm.

103

Figure 3.4. Photograph of well-developed S1 foliation in the Cowden Creek pluton. View

is northeast and parallel to foliation plane. Portion of lens cap shown is 4 cm.

104

the C planes, which are nearly planar (Fig. 3.5). The rare S-C’ fabrics are antithetic, with

respect to the shear zone, and C’ planes are discontinuous. C and C’ planes, where

present, are planar to slightly undulating fabrics that range from poorly to well-

developed. The C planes have a spacing of 1-40 mm, and generally average 1-15mm. C’

planes have only been noted in thin section and have a spacing of 1-2 mm. According to

Passchier and Trouw (1998), C and C’ foliations develop during the same deformational

phase that produces the S foliation. The C foliations form soon after the development of

the S foliation and C’ foliations develop during the waning stages of the deformation that

produces S foliations.

Figure 3.6 shows the orientation of the S1, S, and C foliations in the Pueblo

Mountains shear zone. Data in this figure shows that all three foliations have a similar

northeast-southwest strike and moderate southeast dip, consistent with the conclusion

that they all formed as progressive elements of a single deformational event. The

orientation of the L1 lineation is also shown in Figure 3.6. This linear fabric element

consistently plots in the down-dip direction with respect to the S1, S, and C foliations,

and has an average trend and plunge of S73°E and 50°, respectively. Collectively, the

orientations of the D1 foliations and lineations indicate that D1 deformation reflects

northwest-southeast shortening, with extension in the down-dip direction.

No D1 folds of bedding have been observed at the outcrop scale in the Tuff unit;

however, scare bedding data from this unit suggests that it has been folded on the

megascopic scale (Fig. 3.1; Plate 1). Poles to bedding in the Tuff unit are shown on a

stereonet in Figure 3.7. These poles plot in two broad clusters, reflecting the fact that

bedding dips either northwest or southeast (Plate 1), and the clusters define a great circle

whose axis trends S15oW, 13o (Fig. 3.7). These data suggest that bedding in the Tuff unit

has been folded on the megascopic scale about folds with an average axis that gently

105

Figure 3.5. View looking northeast parallel to well-developed S-C fabric in a Tuff unit

tuffaceous rock. S plane foliations are penetrative and are the more vertically oriented of

the foliations. C plane foliations are spaced a few cm apart, and are oriented closer to

horizontal than are S plane foliations. White arrows are aligned with C planes, curved

white lines represent S planes, arrows indicate sense of shear, and NW is northwest.

NW

106

Figure 3.6. Lower-hemisphere equal-area net projections of structural data for D1

deformation from the Pueblo Mountains shear zone in the eastern Pueblo Mountains.

Solid circles are poles to S1 foliation planes and open circles are L1 lineations; Bulls-eye

patterns are kamb contouring of these data sets. Stars and solid squares are poles to S and

C plane foliations, respectively. Great circles are averaged foliation planes. The bottom

right two plots compare S and C plane foliation data with kamb contours of poles to S1

foliation planes.

N3

5E,

54

SE

S1 f

olia

tions

L 1 li

neat

ions

S p

lane

folia

tions

C p

lane

folia

tions

S p

lane

folia

tions

and

S1 c

ont

our

sC p

lane

folia

tions

and

S1 c

ont

our

sS7

3E,

50

N1

8E,

27

SE

N3

0E,

47

SE

***

* ** *

** * ***

*** *

**

** *

**

***

* ** *

** * ***

** * *

**

** *

**

108

Figure 3.7. Lower-hemisphere equal-area net projections of bedding data from Tuff unit

volcaniclastic strata. Solid circles are poles to bedding planes (So). Star is the calculated

fold axis and line is the calculated axial plane.

*

Tuff unit bedding

N21oE, 68oSE

S15oW, 13o

110

plunges southwest (individual axes plunge gently both northeast and southwest) and

steeply southeast-dipping axial planes. I interpret this folding to be a manifestation of D1

deformation because it deforms bedding, not S1 (compare stereonet of Figure 3.7 with

Figure 3.6) and because it reflects the same direction of shortening (northwest-southeast)

as the S1 foliation.

The fault contact between the Cowden Creek pluton and the Tuff unit is also

interpreted to have formed during D1 deformation. Rocks near adjacent to, and along both

sides of this contact, which juxtaposes the 188 +/- 2 Ma Cowden creek pluton against the

179 +/- 2 Ma Tuff unit, is characterized by a strong mylonitic foliation and well-

developed lineation. The orientation of these structures is identical to S1 and L1 elsewhere

in the shear zone.

D1 Shear Sense Indicators

Indicators of shear sense associated with D1 deformation are abundant. These

include mantled porphyroclasts (sigma and delta structures; Figs. 3.8, 3.9, and 3.10), S-C

and S-C’ fabrics (Figs. 3.5, 3.11, and 3.12). According to Marshak and van der Pluijm

(1997) and Passchier and Trouw (1998), these features are useful in determining shear

sense within a ductile shear zone. To establish shear sense accurately in both outcrop and

in geographically oriented thin sections, these features were viewed in a plane that is

normal to the S1 foliation and parallel to the L1 lineation. Mantled porphyroclasts, and S-

C fabrics, consistently indicate top-to-the-northwest shear sense (Figs. 3.5, 3.9, 3.10,

3.11, 3.12).

M 1 Metamorphism and Microtextures Associated with D1 Deformation

In the Pueblo Mountains shear zone, syn-tectonic metamorphism accompanied D1

deformation. Original igneous and volcaniclastic protolith rocks (discussed in chapter 2)

111

Figure 3.8. Photograph of mantled plagioclase porphyroclasts (white crystals) in the

Cowden Creek pluton. The most obvious porphyroclast is circled and forms a sigma

structure. Arrows show shear sense and NW is northwest.

NW

112

Figure 3.9. Photomicrograph of polygonal quartz neoclasts (larger white and grey crystals

in center of photograph) in the Diamond Inn stock. They represent a relict quartz

phenocryst that forms a somewhat poorly defined delta structure. Top left of thin section

photograph is up direction. Arrows show shear sense and NW is northwest. Width of

view is 4.5 mm. Cross polarized light.

NW

113

Figure 3.10. Photomicrograph of rotated plagioclase crystal, and quartz pressure

shadows, that form a delta structure in a Tuff unit crystal-ash tuff. NW is northwest and

arrows show top-to-the northwest shear sense. Width of view is 2 mm. Cross polarized

light.

NW

114

Figure 3.11. Photomicrograph of Tuff unit crystal tuff containing an S-C fabric defined by

aligned metamorphic muscovite (elongate high birefringent crystals). Schematic drawing

shows S plane foliations trend nearly north to south, and C plane foliations trend nearly

east to west. Large arrows show shear sense direction. NW is northwest. Width of view is

4.5 mm. Cross polarized light.

NW

S

C

115

Figure 3.12. Photomicrograph from an altered portion of the Strawberry Butte pluton

containing a moderately developed S-C fabric. S plane foliations are oriented from top left

to bottom right of photograph, and are defined by long axes of plagioclase laths (grey and

white). In the plagioclase crystal in the right-center part of the photograph, albite twins

are bent and S shaped. C plane foliations trend from lower left to the upper right of

photograph. Three obvious, distinct C planes are visible and are defined by aligned

metamorphic muscovite (acicular high birefringent mineral). Arrows show shear sense

direction along C plane (thin white line). NW is northwest and top left is up direction.

Width of view is 10mm. Cross polarized light.

NW

116

presently contain a variety of metamorphic minerals. In order to establish the

metamorphic grade of these rocks, modal analysis of 71 thin sections was used to

determine their metamorphic mineral assemblages. The following paragraphs describe the

metamorphic mineral assemblages of shear zone rocks, interpret their metamorphic grade,

describe D1-related microstructures in these rocks, and interpret of temperature

conditions of deformation within the shear zone.

M1 Metamorphism

For this discussion, metamorphic rocks in the shear zone are divided into two

broad protolith compositional categories. These are rhyolitic to dacitic rocks, and

andesitic or dioritic rocks. In these rocks, aligned metamorphic minerals define the S1

foliation and L1 lineation, indicating that these metamorphic minerals are syn-tectonic

with D1 deformation.

Metamorphosed rhyolitic to dacitic rocks in the shear zone include the Diamond

Inn hypabyssal stock and most Tuff unit rocks. The Diamond Inn hypabyssal rhyolite

has a unique metamorphic mineral assemblage, distinct form other shear zone rocks, of

only quartz and muscovite. Rhyolitic to dacitic volcanogenic rocks in the Tuff unit

contain the metamorphic assemblage muscovite + quartz + plagioclase, + epidote +/-

calcite, biotite, potassium feldspar, and chlorite. Andesitic to dioritic rocks in the shear

zone include some lavas in the Tuff unit, the Strawberry Butte and the Cowden Creek

plutons. The M1 mineral assemblage in the rocks is plagioclase + quartz, +/- chlorite,

biotite, epidote and muscovite, +/- rare calcite and actinolite.

Several authors (Mason, 1978, Winkler, 1979; Turner, 1981; Yardley, 1989;

Philpotts, 1990; Miyashiro, 1994) have documented the metamorphic minerals

assemblages that define the various metamorphic facies. Figure 3.13 (modified from

Philpotts, 1990) is a plot showing the pressure and temperature boundaries for each

117

Figure 3.13. Diagram of approximate pressure and temperatures under which various

metamorphic mineral facies form. Boundaries between facies are broad zones in which a

number of important relations take place. ACF plots are of common quartz-bearing

mineral assemblages in the metamorphic facies. Minerals plotted include andalusite (A),

kyanite (k), sillimanite (sill), muscovite (M), grossularite (Gr), almandine (Alm), pyrope

(Py), orthopyroxene (Opx), clinopyroxene (Cpx), calcic plagioclase (Pl), epidote (Ep),

lawsonite (Lw), Laumonite (Lm), pumpellyite (Pm), prehnite (Pr), calcite (C), aragonite

(Ar), dolomite (D), actinolite (Ac), hornblende (Hb), glaucophane (Gl), anthophyllite

(Ath), talc (Tc), biotite (B), chlorite (Ch), pyrophyllite (Pph), and stilpnomelane (Stn).

Tem

per

atur

e (

oC)

Pressure (GPa)

100

200

300

400

500

600

700

800

900

0

0.0

0.10.2

0.30.40.5

0.6

0.7

0.8

0.9

1.01.1

1.2

Garnet melts

Cla

y

Ch

C

PrLm

Ch T

cA

cC

Ep

MEp

M,K

CA

cTc

Ch(

Alm

)(B

)

M,P

ph

EpPm Pr

CD

Ch

Stn

M(S

ill)K

C

Pl

Gr

Cpx

Hb

Ath

Alm

Cpx

Hb

C

Pl

Gr

M,A

B

M

Gl

Lw

Ar

Gr

Ar

K

Cpx

Opx

Alm

-Py

Zeolit

ePr

ehni

tepum

pel

lyit

e

Gre

ensc

hist

Blu

esch

ist

Eclo

git

e

Am

phi

bolit

e

Gra

nulit

e

San

idin

ite

119

facies, and the metamorphic mineral assemblages that define each facies. For the purpose

of establishing metamorphic grade in the Pueblo Mountains shear zone, the M1 minerals

of these rocks are compared to the metamorphic mineral assemblages in Figure 3.13.

The following is a synthesis of M1 mineral assemblages for all shear zone

protoliths: +/- muscovite +/- quartz, +/- plagioclase +/- biotite +/- epidote +/- potassium

feldspar +/- chlorite +/- calcite +/- actinolite. According to Mason (1978), Winkler (1979),

Turner (1981), Yardley (1989), Philpotts (1990), and Miyashiro (1994) actinolite +

chlorite + biotite + epidote + albite + quartz + muscovite + sphene + calcite + opaque

minerals define greenschist facies rocks for felsic to intermediate volcanics, and dioritic

protolith compositions. Other facies, whose pressures and temperatures bound the

greenschist facies zone, are the zeolite, prehnite-pumpellyite, blueschist, eclogite, and

amphibolite facies (Fig, 3.13). None of the minerals that are diagnostic of these facies are

observed in the shear zone. It is clear that rocks in the Pueblo Mountains shear zone are

greenschist grade.

Greenschist facies mineral assemblages approximately form between 300 and 500

°C (Fig. 3.13; Philpotts, 1990). As shown in this figure, increases in pressure gradually

shifts this temperature range towards higher temperature. This study is not concerned

with pressure in the shear zone; however, the temperature range, above, is important

because the following paragraphs analyze temperature data from microstructures and

compares that data with the temperature range indicated by M1 minerals.

Microstructures in Quartz and Plagioclase

According to Passchier and Trouw (1998), individual microstructures, and sets of

microstructures, form in specific minerals at specific temperature conditions. Therefore,

microstructures may be used to constrain the temperatures conditions of the deformation

that formed them. This information is useful because, microstructures in rocks from the

120

Pueblo Mountains shear zone can be used to estimate temperature conditions of

deformation during formation of this shear zone.

Passchier and Trouw (1998) describe various microstructures, and the temperature

conditions at which they form, for several minerals. However, the only minerals abundant

enough in the shear zone to be important to this study, are quartz and plagioclase. Of

these, quartz is a less reliable indicator of temperature conditions than plagioclase. This is

because strain-rate, differential stress, and the presence of water in the lattice and along

grain boundaries, are other factors in addition to temperature that can govern deformation

in quartz (Passchier and Trouw, 1998). The following paragraphs describe

microstructures in the Pueblo Mountains shear zone and provide an estimate of

temperature conditions during D1 deformation by comparing microstructures, from

within the shear zone, with the temperature data (in Passchier and Trouw, 1998) at which

they form.

In nearly all Pueblo Mountains shear zone rocks, some amount of quartz is

present, and is generally abundant. In places, it is igneous crystals that are much larger

than metamorphic (dynamically recrystallized) quartz and, as indicated by

microstructures, display lesser amounts of recrystallization and recovery than

metamorphic quartz. However, in the shear zone, the majority of quartz is very fine-

grained recrystallized quartz that is the product of deformation and metamorphism.

Figure 3.14a (from Passchier and Trouw, 1998) illustrates the recovery process,

which is the ordering mechanism that competes against disordering (i.e., deformation) in

strained crystals. Dislocations, caused by deformation, first produce undulose extinction.

As deformation wanes and recovery becomes more dominant, these dislocations become

concentrated along deformation bands and eventually they form individual subgrains.

Further along the recovery road, deformed crystals become smaller crystal aggregates

characterized by equigranular-polygonal grain boundaries (top left rectangle in Figure

121

Figure 3.14 (from Passchier and Trouw, 1998). (a) Schematic illustration of the recovery

process. See text for detailed description. (b) Schematic illustration of shapes of grain

aggregates and grain boundaries. These shapes are the product of recrystallization in

deformed rocks.

Und

ulose

ext

inct

ion

Def

orm

atio

n ban

d

Sub

gra

in b

oun

dar

y

dis

loca

tions

recovery

equi

gra

nula

r-poly

gona

lin

equi

gra

nula

r-poly

gona

lse

riat

e-poly

gona

l

seriat

e-in

terlobat

ein

equi

gra

nula

r-in

terlobat

eeq

uigra

nula

r-in

terlobat

e

equi

gra

nula

r-am

oeb

oid

ineq

uigra

nula

r-am

oeb

oid

seriat

e-am

oeb

oid

Sha

pe

of

gra

in a

ggre

gat

es

a

b

123

3.14b). These crystals no longer show evidence of strain (i.e., undulose extinction or

subgrains), and are termed “strain-free” and dynamically recrystallized.

In shear zone rocks, both phenocrystic quartz and quartz crystal clasts, commonly

display undulose extinction and well developed subgrains boundaries (Figs. 3.15 and

3.16). The majority of the quartz, in the shear zone, is in the groundmass and matrix of

these rocks and is very fine-grained recrystallized quartz neoclasts. This quartz lacks

undulose extinction and subgrains (i.e., they are strain-free; Figs. 3.9 and 3.17) indicating

that high levels of recovery (as defined above) are represented in these crystals. In most

rocks, recrystallized quartz has nearly equigranular and polygonal grain boundaries (Figs.

3.15 - ignore large quartz crystal with subgrains, and 3.17; compare to shapes of grain

aggregates in Figure 3.14b). In fewer rocks, grain boundaries are inequigranular-interlobate

to seriate-interlobate (Fig. 3.9; compare to Figure 3.14b). These recrystallized quartz

crystals are “strain-free” indicating that these rocks have been greatly affected by the

recovery process (Passchier and Trouw, 1998).

All of the microstructures described above, form in quartz at temperatures ranging

from 400 to700o C (Passchier and Trouw, 1998). This temperature range is a preliminary

estimate of temperature conditions, in the shear zone, during D1 deformation. Further

temperature data is obtained by analysis of microstructures in plagioclase crystals, and is

discussed below.

Like quartz, plagioclase is nearly always present in shear zone rocks, is often

abundant, and displays a variety of microstructures. These include mantled

porphyroclasts (Figs. 3.8 and 3.10), poorly developed core-mantle structures,

recrystallized plagioclase neoclasts (Fig. 3.18), bent albite twinning (Figs. 3.12 and 3.19),

undulose extinction, and fractured and boudinaged crystals. Fractured plagioclase crystals

and bent albite twins (Figs. 3.12 and 3.19) are the most common microstructures observed

in shear zone plagioclase. Fractured crystals are often boudinaged and, when boudinaged,

124

Figure 3.15. Photomicrograph of large quartz crystal (bottom center) displaying undulose

extinction and well-developed subgrain boundaries (dark shades of grey to black). Smaller

quartz crystals (white and grey and somewhat equant) are recrystallized and have

inequigranular-interlobate to polygonal grain boundaries as defined by Passchier and

Trouw (1998; Fig. 3.13b). Note – these quartz crystals are “strain-free”. Width of view is

4.5mm. Cross polarized light.

125

Figure 3.16. Photomicrograph of large quartz crystal with well-developed undulose

extinction and subgrains (dark shades of grey to black). Thin-section is from a Tuff unit

ash-crystal tuff. Width of view is 4.5 mm. Cross polarized light.

126

Figure 3.17. Photomicrograph of “strain-free” recrystallized quartz neoclasts (white to

grey equant crystals) in the matrix of a Tuff unit tuffaceous rock. Gain boundaries are

equigranular-interlobate to equigranular-polygonal (Passchier and Trouw, 1998; Fig.

3.14b). Width of view is 1.3 mm. Cross polarized light.

127

Figure 3.18. Photomicrograph of recrystallized plagioclase neoclasts that form a thin

semi-continuos rim (white color) developed along the edges of a plagioclase lath partly

replaced by metamorphic epidote (very fine-grained high birefringent equant crystal

aggregates). Subhedral recrystallized plagioclase laths (light grey) are scattered throughout

the groundmass. Thin section from the Strawberry Butte pluton. Width of view is 1.3

mm. Cross polarized light.

128

Figure 3.19. Photomicrograph of plagioclase lath with bent albite twins in altered thin

section of the Strawberry Butte quartz diorite. Width of view is 4.5 mm. Cross polarized

light.

129

are extended in a direction parallel to the L1 stretching lineation. Also common are rotated

plagioclase crystals that form sigma and delta structures (Fig. 3.10). Moderate amounts of

recrystallized plagioclase neoclasts are present in these rocks, both in the groundmass and

as small neoclasts along the edges of larger plagioclase phenocrysts (Fig. 3.18) and crystal

clasts. In some places, recrystallized plagioclase around crystal edges form poorly defined

core-mantle structures. Undulose extinction is rarely observed in plagioclase in shear zone

rocks.

In plagioclase, fracturing, bent albite twins, recrystallization of plagioclase

neoclasts, undulose extinction, and core-mantle structures develop at low-grade

temperature conditions ranging from 300 to 400o C. At low to medium-grade conditions

(400 to 500o C) recrystallization of plagioclase along grain boundaries becomes a more

dominant process. Above 500o C fracturing in plagioclase in less common, subgrains

become well-developed, and core-mantle structures are still present, although the

boundary between core and mantle is less pronounced that at low-grade temperature

conditions (temperature data from Passchier and Trouw, 1998). Collectively, the

abundance of fracturing and bent albite twins in plagioclase, lesser core-mantle structures,

and rare undulose extinction, initially indicate that temperature conditions of shear zone

deformation (D1) ranges from 300 to 400o C.

Considering the moderate amount of recrystallized plagioclase in the shear zone,

which begins to develop around 400o C, it is more likely that temperature conditions were

closer to 400o C. At temperatures above 500o C, fracturing in plagioclase is rare and

subgrains should develop; therefore, temperature conditions of D1 deformation in the

shear zone probably did not exceed 500o C. In summary, the data above approximates

temperature conditions of D1 deformation at 400o C (all temperature data for plagioclase

from Passchier and Trouw, 1998).

130

By combining temperature conditions of shear zone deformation, as indicated by

microstructures in both quartz and plagioclase, a reasonable estimate of temperature

conditions can be established. Microstructures in quartz, although a less reliable indicator

of temperature than plagioclase (Passchier and Trouw, 1998), suggests temperatures of

400 to 700o C. Microstructures in plagioclase indicates that shear zone deformation

occurred at temperatures around 400o C. Because these temperatures overlap, they

together indicate that temperature conditions of deformation in the Pueblo Mountains

shear zone is closely approximated at, or perhaps slightly above, 400o C.

Summary of D1 Temperature Conditions

Greenschist facies metamorphic mineral assemblages, such as those in the Pueblo

Mountains shear zone, form at temperatures ranging from approximately 300-500o C (Fig.

3.13; Philpotts, 1990), and microstructures in the shear zone indicate that temperatures

conditions for D1 deformation is ~400o C. This overlap in temperatures, indicated by

microstructures and metamorphic mineral assemblages, indicates peak recorded

metamorphic temperatures reached approximately 400o C.

Age Constraints for D1 Deformation and Metamorphism

In the Pueblo Mountains, field relations provide an initial age constraint for D1

deformation and metamorphism. The youngest rock unit affected by D1 deformation is

the 160 +/- 2 Ma Diamond Inn hypabyssal rhyolite stock. The oldest rocks unaffected by

D1 deformation are Tertiary volcanogenic strata dated at 17 Ma (Hart and others, 1989).

Therefore, D1 deformation and metamorphism must have occurred sometime between 160

and 17 Ma.

131

In addition to the these cross-cutting relationships, the age of D1 deformation and

metamorphism can be more tightly constrained by age data from two workers. The first

worker (Harrold, 1972), dated two pre-Cenozoic dioritic plutonic rocks in the Pueblo

Mountains shear zone using K-Ar techniques. Whole rocks analysis of these rocks

yielded ages of 100 +/- 2.0 Ma and 92.4 +/- 1.3 Ma. These ages were interpreted by

Harrold (1972) as original crystallization ages (mid-Cretaceous) for these minerals.

However, by comparing the sample localities of these rocks (Harrold, 1972) with the

detailed map produced by this study, the rocks Harrold (1972) dated are mapped by this

study as the Strawberry Butte quartz diorite, which has been dated by this study (J.E.

Wright). Based on U-Pb zircon age data (chapter two of this study), the rocks Harrold

(1972) dated as mid-Cretaceous are in fact Jurassic (~179-176 Ma; chapter 2). Based on

these new age data in this study, I suggest that K-Ar ages from Harrold (1972) may

represent post-emplacement Ar-loss during younger metamorphism of these rocks.

Therefore, these mid-Cretaceous ages are instead, ages of metamorphism, not original

crystallization of igneous minerals, and they provide preliminary age data for shear zone

deformation and metamorphism.

The conclusion above is supported by data from Brown (1996). She presented

one 40Ar/39Ar step heating analysis biotite from the Cowden Creek pluton. The sample

was from a portion of the pluton deformed into an S-C mylonite and the biotite in the

sample is aligned and grew as a metamorphic phase during D1 deformation (Brown, 1996).

The age spectra from this sample defines an irregular plateau at ~95 Ma, suggesting that

these rocks cooled through the closure temperature of biotite to Ar loss (300-350°C;

McDougall and Harrison, 1999) in the mid-Cretaceous.

To better constrain the age of D1 deformation and metamorphism in the Pueblo

Mountains, I collected eight rocks samples containing M1 minerals, from within the shear

zone for 40Ar/39Ar analysis. Metamorphic muscovite and biotite mineral separates from

132

these rocks have been analyzed by Alexander Iriondo. The exact data from this analysis is

not yet available; however, it is known that metamorphic ages for these minerals are

Albian (112-99 Ma: mid-Cretaceous; Alexander Iriondo, personal communication).

D2 Deformation

Locally present in the shear zone are folds of the S1 foliation. These folds are

interpreted as D2 structures and their locations are shown in Figure 3.2 and Plate 1. D2

folds are only observed in the Tuff unit.

Two distinct styles of D2 folds are present. (1) The most common are parallel,

asymmetric, folds with amplitudes of 0.25-0.5 m and rounded to subrounded fold hinges.

Rarely, an axial planar S2 cleavage is associated with these folds. (2) The less common D2

folds are disharmonic (typically a few centimeters in length), with subangular to kinked

fold hinges, and an amplitude of has a 0.25-0.5 cm.

Fold axes and axial planes of both fold types have similar orientations. F2 fold axes

trend northeast to southwest and are gently plunging, and axial planes of these folds are

northeast-striking and dip steeply to the southeast (Fig. 3.20a). All asymmetric D2 folds

verge northwest indicating a top-to-the-northwest shear sense. This is the only shear

sense indicator associated with D2 structures.

Axial planar cleavage, rarely observed in the first fold type, is the only evidence of

possible M2 metamorphism related to D2 deformation. Although, formation of this

cleavage could be in response to other processes other than metamorphism. Thin sections

of rocks, from shear zone areas containing D2 deformation, were analyzed for D2

structures and any metamorphic minerals that may have developed during this

deformation; however, no D2 structures or M2 minerals were observed.

133

Figure 3.20a-b. Lower-hemisphere equal-area net projections of structural data from D2

and D3 deformation in the eastern Pueblo Mountains. (a) D2 deformation: Solid circles are

poles to F2 axial planes and stars are F2 fold axes. Great circle is average strike and dip of

F2 axial planes. (b) D3 deformation: Open circles are poles to S3 foliation planes and great

circle is average strike and dip of S3 foliation planes.

*

*

**

****

*

*** ***

**

***

N35E, 65SE

N36E, 23SE

F2 fold axes and

axial planes

S3 foliation

(a)

(b)

135

Very little data is available to constrain the timing of these folds. They are

observed in only the Tuff unit (179 +/- 2 Ma) and must be younger than the S1 foliation

(~95 Ma; Brown, 1996) they fold. The oldest overlying Tertiary strata, which are

undeformed by D2 deformation are dated at ~17 Ma by Hart and others (1989).

Therefore, D2 folds must have developed between ~95 and 17 Ma. No other age data is

available to further constrain the timing of D2 folding.

D3 Deformation

In the Pueblo Mountains shear zone, D1 and D2 structures are locally cross-cut by

a spaced S3 foliation. This foliation is observed in only two locations, both of which are in

Tuff unit (Fig. 3.2; Plate 1). It is a planar, discontinuous foliation with spacing that ranges

from 1-10 cm. It strikes to the northeast and dips gently southeast (Fig. 3.20b).

S3 was not observed in thin sections of rocks collected from areas containing D3

structures. No shear sense indicators, associated with the S3 foliation, were observed in

outcrops of these rocks.

The only available data, to constrain the timing of D3 deformation, is that D3

structures must be younger than the structures they cross-cut and older than the youngest

rocks they do not affect. Therefore, D3 deformation is younger than D2 and D1

deformation (dated at ~95 Ma; Brown, 1996), and older than undeformed overlying

Tertiary strata (dated at ~17 Ma; Hart and others, 1989). No other age data is available to

further constrain the timing of D3 deformation.

Strain Heterogeneity in the Pueblo Mountains Shear Zone

Strain related to all three phases of deformation, D1, D2 and D3, is heterogeneous

in the Pueblo Mountains shear zone. In an effort to evaluate the significance of this, the

136

relative, amount of strain in the shear zone, was analyzed according to spatial distribution

and rock type. A description of this analysis is presented in the following four

paragraphs are its results are illustrated in Figure 3.21.

For D1 deformation, the degree of strain affecting different rocks and areas can be

divided into three categories. (1) highest strain is manifested by a well-developed and

pervasive mylonitic, or schistose, S1 foliation and strong L1 stretching lineation; areas and

rocks with S-C fabrics are also included in this category. (2) Intermediate amounts of

strain are manifested by more moderately-developed, non-pervasive, mylonitic or

schistose foliation. (3) Lowest amounts of strain are represented by a non-pervasive,

weakly-developed schistose foliation.

Dark grey shading depicts the locations of highest strain areas (category 1) in the

shear zone (Fig. 3.21). Nearly the entire Tuff unit is part of this category, as this unit

generally displays a pervasive, well-developed foliation and lineation. Other high strain

areas include the northwest side of the Cowden Creek pluton (near its contact with the

Tuff unit), all of the Diamond Inn stock, and localized areas within the Strawberry Butte

pluton.

Parts of the shear zone characterized by intermediate amounts of strain (category

2) are depicted by medium grey shading (Fig. 3.21). In the Tuff unit, some parts of Denio

Canyon, the east end of Van Horn Canyon, and the southern tip of the map area, are

zones of intermediate strain. Other areas, from north to south across the study area,

include the southeasternmost part of the Andesite unit (along the Tuff-Andesite unit

contact), the northern part of the Strawberry Butte pluton (just north of Denio Canyon),

scattered areas in southwestern part of the Strawberry Butte pluton, and much of the

Cowden Creek pluton.

Areas of lowest strain amounts (category 3), are represented by the lightest grey

shading, and areas in the shear zone without shading (excluding Quaternary units) are

137

Figure 3.21. Simplified geologic map of the eastern Pueblo Mountains that uses light,

medium, and darker grey shading to show distribution and strength of strain within the

Pueblo Mountains shear zone. Key in lower right of figure summarizes significance of

shading. Note- important topographic features are shown.

non-pervasive weakly-

developed schistose foliation

SHEAR ZONE FABRICS

intense, pervasive mylonitic to

schistose foliation and well-

developed stretching lineation

incr

easi

ng s

trai

n

moderately-developed, non-pervasive

mylonitic to schistose foliation

5451

PUE

BLO M

NTS.

SHE

AR

ZONE

JCD

JCT

JSB

JAU

JTU

JPM

JPM

TU

QU

QU

QU QU

QU

QU

QU

QU

QU

QU

QU

QU

QU

QU

QU

QU

QU

QU

JAU

JAU

JAU

JAU

JTU

JTU

JTU

JTU

JTU

JSBJSB

JCD

JCD

JSB

JSB

JDI

JTU

TU

TU

TU

TU

0 1 km

N

Andesite Unit flows and lesservolcaniclastics and tuffs

JTU

JAU

Tuff Unit tuffaceous rocks andlesser volcaniclastics and flows

Stratigraphic units

JSB

JCT

JPM

JDI


Catlow Creek quartzmonzodiorite


Diamond Innhypabyssal rhyolite

Intrusive units

PUEBLO TERRANE(Jurassic)

Key to Map Units

QU

TU

JCD

Tertiaryundivided

Quaternaryundivided

Cowden Creekquartz diorite

SYMBOLS

Northwest boundary of

Cretaceous PMSZ

Ductile thrust fault -

teeth on upthrown side

Denotes location

of S-C fabric

JSB

JSB

DenioCreek

DenioBasin

VanHornBasin

Van HornCreek

*

*

*

Denotes location

of F2 folds

Denotes location

of S3 foliation

139

rocks with no macroscopic S1 foliation (Fig. 3.21). These areas are generally restricted to

parts of the Strawberry Butte pluton (specially along its northwestern contact with

Quaternary and Tertiary deposits) and parts of the Cowden Creek pluton.

Rock type and position of the rock within the shear zone appear to act as controls

for the amount, and distribution, of strain in shear zone. Rock type is a control because,

owing to their mineralogy and textures, different rock types have different competencies.

Competence is defined by Twiss and Moores (1992) as the relative rate at which a

material ductily deforms in response to stress. Under the same amount of stress, a more

competent rock will deform less than a less competent rock. These less competent rocks

are generally termed incompetent. In general, rocks with more mafic compositions are

more competent than rocks with more felsic compositions. Also, rocks textures influence

the competency of a rock. Rocks with interlocking crystalline textures are more

competent than rocks with non-crystalline ash or matrix-rich textures. For example, a

gabbro will be more competent than a rhyolitic tuffaceous rock. Several examples of this

exist in the Pueblo Mountains. Tuff unit tuffaceous rocks nearly always display highest

strain (Fig. 3.21). This is likely because, owing to their felsic composition and ash-rich

matrix, they are incompetent. Intermediate strain parts of the Tuff unit generally coincide

with the location of more competent rocks, such as volcaniclastics and flows, which are

more mafic in composition than the tuffs and have stronger textures (compare areas of

intermediate strain in Figure 3.21 with the location of Tuff unit volcaniclastics and flows

in Plate 1).

Greater competency is most likely the reason why intrusives (Strawberry Butte

and Cowden Creek plutons) generally display intermediate to no strain amounts (Fig.

3.21). An exception is the Diamond Inn stock, which displays high amounts of strain

(Fig. 3.21). This is most likely because it is of lesser competency owing to its rhyolitic

composition.

140

The other control, governing magnitude and distribution of strain, is the position

of the strained rock in the shear zone. In general, the northwestern rocks of the Tuff unit

and Strawberry Butte plutons show less strain than shear zone rocks located to the

southeast, and an overall trend of decreasing strain, from southeast to northwest, is visible

across the Pueblo Mountains shear zone (observe mylonitic foliation in structure sections

- Fig. 1.4a-d; and Fig. 3.21). This is trend is also seen along the Tuff-Andesite unit

contact, where strain dies out to the northwest (Fig. 3.21).

An interesting feature, regarding strain heterogeneity in the shear zone, is that S-C

fabrics, and D2 and D3 structures are observed only in higher strain areas of the Tuff unit

(Figs. 3.2 and 3.21) and not in any other rocks in the shear zone. Based on the spatial

distribution of these structures, it appears that a competency difference between

tuffaceous rocks and intrusives may control the location of where these structures form.

It is probable that strain was not strong enough to produce S-C, D2 or D3 structures in

non-tuffaceous shear zone rocks.

In summary, strain related to D1 deformation is heterogeneously distributed in the

Pueblo Mountains shear zone (Fig. 3.21) and two factors, competency of the rock type

and position of the rock within the shear zone, control this distribution. Competency

contrasts between tuffaceous and intrusive shear zone rocks appears to have controlled

the distribution of S-C fabrics, and D2 and D3 structures. Analysis of the distribution of

D1 strain reveals a general trend, of strain decreasing towards the northwest, across the

Pueblo Mountains shear zone. Position of the rocks, and in part competency contrasts

between different rock types, likely controlled the development of this pattern.

Rocks Northwest of the Pueblo Mountains Shear Zone

Strongly deformed rocks in the Pueblo Mountains shear zone extend only ~500 m

into the Andesite unit, otherwise they are only found south of the Tuff-Andesite unit

141

contact (Figs. 3.1, 3.2, and 3.21; Plate 1). Moving from this contact towards the

northwest, strain dies out rapidly. Thus most of the Andesite unit, macroscopically,

appears unaffected by Mesozoic ductile deformation. The two stocks that intrude this

unit, the Pueblo Mountain and Catlow Creek stocks (Fig. 3.1 and Plate 1), are also

unaffected. However, the orientations of bedding in the Andesite unit and the map pattern

of clastic horizons north of Arizona Canyon (Plate 1 and Fig. 3.1) suggest that layering in

the Andesite unit is folded. The limited bedding data in the Andesite unit is plotted in

Figure 3.22. Poles to bedding define two clusters with a calculated axial plane that dips

steeply to the southeast and gently plunging fold axis that trend northeast (Fig. 3.22).

These are folds of bedding (So) and are D1 deformation. As seen in the Tuff unit, the

orientation of these folds indicates northwest-southeast directed compression.

Evidence of metamorphism is present in the Andesite unit and the stocks that

intrude it. In this section, I evaluate whether or not microstructural features are present in

these rocks, and I describe metamorphic assemblages in these rocks.

The Pueblo Mountain and Catlow Creek stocks contain sufficient amounts of

plagioclase, potassium feldspar, and quartz, and Andesite unit rocks contain enough

plagioclase, that if these rocks are deformed, then these minerals could contain

microstructures. However, subgrains, recrystallized (strain-free) neoclasts, bent albite

twins, fractured plagioclase crystals, and the preferred alignment of metamorphic minerals

are not observed in these rocks. The only microstructure observed is faint undulose

extinction in quartz, which is rare and restricted to the Catlow Creek stock.

To establish metamorphic grade in the Andesite unit, and Pueblo Mountain and

Catlow Creek stocks, metamorphic assemblages of these rocks were determined by modal

analysis of 38 thin sections. These assemblages are compared to known metamorphic

facies assemblages illustrated in Figure 3.13. Rocks north of the shear zone are broadly

andesitic/dioritic in composition and, based on these compositions, they act as more

142

Figure 3.22. Lower-hemisphere equal-area net projections of bedding data from Andesite

unit volcaniclastic strata. Solid circles are poles to bedding planes (So). Star is the

calculated fold axis and line is the calculated axial plane.

Andesite unit bedding

*

N50oE, 79oSE

N55oE, 23o

144

reliable indicators of metamorphic grade than do more felsic rocks, and it is appropriate to

compare their metamorphic assemblages to those for metabasites. This is following the

convention of Mason (1978), Winkler (1979), Turner (1981), Yardley (1989), and

Miyashiro (1994).

The metamorphic minerals assemblages for rocks north of the shear zone is as

follows: the Pueblo Mountain stock consists of plagioclase + epidote +/- quartz,

actinolite, potassium feldspar, and chlorite, the Catlow Creek stock is comprised of

quartz + plagioclase + epidote, +/- actinolite, potassium feldspar, chlorite, and calcite.

Andesite unit andesites contain plagioclase + epidote, +/- chlorite, actinolite, biotite,

quartz, muscovite, and potassium feldspar (only present in one sample), and rhyolitic to

andesitic Andesite unit volcaniclastic and tuffaceous rocks consists of +/- quartz,

plagioclase, +/- lesser potassium feldspar, epidote, biotite, chlorite, and calcite.

For metabasic rocks, prehnite and pumpellyite define the subgreenschist facies,

chlorite + actinolite + epidote + albite + quartz is typical of greenschist facies

metabasites, and more calcic plagioclase and the coexistence of actinolite and metamorphic

hornblende define the amphibolite facies (Miyashiro, 1994; Philpotts, 1990). According

to these authors, the metamorphic mineral assemblages of rocks northwest of the shear

zone, as described above, are greenschist facies.

Little direct evidence is available to constrain the timing of greenschist

metamorphism northwest of the Pueblo Mountains shear zone. Andesite unit rocks, from

the structurally highest part of the Pueblo Mountains stratigraphic section, have an

igneous crystallization age of 176 +/- 2 Ma (J.E. Wright, this study). Metamorphism in

these rocks must be younger than this age and must be older that the unmetamorphosed

~17 Ma Tertiary strata (Hart and others, 1989) that overlie these rocks.

One indirect line of evidence, provides some insight into the timing of

metamorphism. Harrold (1972) dated two rocks north of the shear zone using K-Ar

145

techniques (these rocks are in addition to the previously discussed shear zone rocks he

dated). Biotite from his quartz monzonite stock (Catlow Creek stock; this study) yielded

an age of 91.3 +/- 1.3 Ma, and plagioclase from an andesitic rock (Andesite unit lava; this

study) yielded an age of 108+/- 1.5 Ma. Harrold (1972) interpreted these ages as original

crystallization ages (mid-Cretaceous) for these minerals. However, the Catlow Creek

stock was dated by Brown (1996) at ~179 Ma, and an Andesite unit rock from the top of

the Pueblo terrane stratigraphic section (the sample Harrold dated was from a lower part

of the section), at 176 +/- 2 Ma (J.E. Wright, this study). Again, I suggest these mid-

Cretaceous ages represent metamorphic ages, not igneous crystallization ages.

Two important conclusions, concerning rocks northwest of the shear zone, can be

drawn from the deformational and metamorphic data presented above. The first is that,

moving from the Tuff-Andesite unit contact northwest into the Andesite unit, strain

quickly dies out. This is supported by the fact that all rocks northwest of the shear zone

show no microscopic evidence of Mesozoic ductile deformation. The second conclusion is

that these rocks are the same metamorphic grade (greenschist) as rocks within the shear

zone, and that metamorphism both northwest of the shear zone, and within it, is roughly

the same age (mid-Cretaceous). This conclusion is significant because it implies that shear

zone rocks, and rocks northwest of the shear zone, are from approximately the same

structural levels within the crust. If these packages of rocks were from different crustal

levels, this would be reflected by their having different metamorphic grades. Also, both

packages of rocks are affected by similar map-scale folding. These imply that only limited

dip-slip displacement occurred along the northwestern boundary of the shear zone,

consistent with conclusions presented in chapter two (Discussion of the Tuff Unit –

Andesite Unit Contact), that there was little structural displacement along the Tuff-

Andesite unit contact.

146

Summary and Interpretation of Mesozoic Deformation

The Pueblo terrane consists of a thick sequence of Middle Jurassic volcanogenic

strata intruded by three shallow-level rock units, also Middle Jurassic. One other small

intrusion is present, it is Late Jurassic (Fig. 3.1; Plate 1).

Following the deposition and crystallization of these rocks, the southern half of

this terrane was deformed by the Pueblo Mountains shear zone. This shear zone is at

least 3 km wide, dips southeast (Fig. 3.1; Plate 1), and contains one major (D1), and two

lesser (D2 and D3), phases of deformation. The first (D1), produced a foliation that is

often mylonitic and intensely deforms these rocks (Fig. 3.21) and large scale folds. In

other rocks, the magnitude of this deformation is less. The second and third, D2 and D3,

phases of deformation, are manifested by F2 folds and an S3 foliation, respectively, that

deform younger structures. Shear sense indicators from D1 and D2 deformation indicate

that reverse-sense compressional tectonism displaced shear zone rocks upwards and to

the northwest. Strain in the shear zone is strongest in the southeast, fades to the

northwest (Fig. 3.21), and is absent in the northern third of the Pueblo terrane (Figs. 3.1

and 3.21; Plate 1).

Four interesting similarities exist, between D2 folding and S1 structures, that

provide further insight into the timing of D2 deformation. (1) The average S1 foliation and

D2 axial plane, respectively, have identical northeast strikes and similar southeast dips

(Fig. 3.6 and 3.20a). (2) Both D1 and D2 shear sense indicators reveal top-to-the-

northwest shear sense. (3) The locations of D2 folds, like D1 structures, are restricted to

within the shear zone (Figure 3.2 and Plate 1). (4) Both D1 and D2 structures, as indicated

by their orientations and shear sense indicators, were produced by northwest-southeast

directed compressive forces. These similarities lead to the conclusion that D1 and D2

structures were not produced by discrete deformational events. Instead, both structures

147

were produced by a single protracted deformational event, with D2 structures forming

soon after D1 structures.

The localized nature of D2 structures is somewhat enigmatic. D2 structures may be

locally developed because this protracted deformational event included only localized

pulses of greater strain, which produced D2 structures in addition to D1 structures, in

only a few areas. Another possibility is that D2 deformation was not strong enough to

deform non-Tuff unit rocks, all of which are intrusive are more competent than the tuffs.

In addition, D3 deformation likely formed in response to the same protracted

deformation that produced D1 and D2 structures. The argument for this conclusion is

similar to the argument presented above and is summarized as follows. The average S1

foliation, F2 axial plane, and S3 foliation all have similar northeast strikes and southeast

dips (Figs. 3.6 and 3.20a-b). These orientations indicate northwest-southeast compressive

tectonism produced all three phases of deformation, and the locations of all deformational

phases are restricted to within the shear zone. Based on these relations, these three

phases likely formed one after the other during a single protracted deformational event.

Rocks in the Pueblo Mountains shear zone are greenschist grade and, as indicated

by microstructures and metamorphic minerals, temperature conditions for deformation in

the shear zone reached approximately 400° C. Field relations and preliminary

geochronologic data indicate deformation and metamorphism in these rocks is mid-

Cretaceous (~95 Ma; Brown, 1996). A mid-Cretaceous northeast-southeast compressive

tectonic force is the likely cause for deformation and syn-tectonic metamorphism.

These observations above allow brief speculation into the possible role of the

Pueblo Mountains shear zone in northwest Nevada and southeast Oregon. Because strain

within the Pueblo Mountains shear zone is greatest along its southeast side, and wanes to

the northwest, it is possible this shear zone represents only a less deformed, northwest

148

portion, of a larger shear zone. This larger shear zone would have to be located to the

southeast, and its southeast portion would contain rocks that are more deformed and of

higher metamorphic grade than those of the Pueblo Mountains shear zone. This

speculation is consistent with the fact that other shear zones, containing amphibolite

grade metamorphic rocks, and are located southeast of, the Pueblo Mountains (Fig. 3.1) as

documented by Wyld and Wright (2001). These shear zones are roughly the same age as

the Pueblo Mountains shear zone.

CENOZOIC DEFORMATION

According to Burnham (1971), Rowe (1971), Harrold (1972), and Tower (1973),

Tertiary rocks in the Pueblo Mountains are predominately basalt flows and dikes. Lesser

rock types include poorly consolidated boulder pebble conglomerate (basaltic in

composition), ash-flow tuff, welded tuffs, porphyritic and equigranular gabbro,

tuffaceous sandstones and conglomerates, rhyolite and latite flows, lithic-rich volcanic

sandstones, and andesite flows. More recently, Hart and others (1989) similarly describe

these rocks as mostly basalt flows, dikes, shallow intrusives, and cumulates, with lesser

silicic flows and pyroclastic units. Hart and others (1989) produced a 1 km thick

composite section of the Pueblo Mountains basalts that contains, in places, up to 64

individual units.

Little age data is available for Tertiary rocks in the Pueblo Mountains. Carlson and

Hart (1985) determined that ages of these rocks to range from 11-16 Ma in age. Hart and

others (1989), in a more detailed study, established a more precise age for these rocks

(15.2-17.3 Ma: mid-Miocene).

149

Unconformity in the Pueblo Mountains

The Tertiary strata of the western Pueblo Mountains are separated from

Mesozoic rocks by a major angular unconformity (Fig. 3.23). Directly west of this

unconformity, several bedding measurements from Tertiary basalt flows were obtained by

this study and previous studies. Based on these measurements, Tertiary strata strike

north-south and dip 23o to the west (Fig. 3.24). In contrast, the underlying Mesozoic

Andesite unit strikes east-west and dips 21o to the north, and the Tuff unit strikes north-

south and dips 41o to the west (Fig. 3.1; Plate 1). Unlike Mesozoic rocks, Tertiary strata

are unmetamorphosed and are not ductily deformed. These relations indicate that

sometime after deposition of Jurassic volcanic strata and after Cretaceous deformation

and metamorphism of these rocks, there was a period of exhumation and erosion prior to

deposition of Tertiary strata.

Tertiary Tilting

All Tertiary strata in the Pueblo Mountains strike north and dip to the west (Figs.

3.1 and 3.24; Plate 1). As shown in Figure 3.24, the average bedding dip of the Tertiary

strata is 23° to the west, with very little variation in direction or dip value. Because all the

mapped Tertiary strata have this orientation, the age of tilting must post-date the age of

the youngest Tertiary rocks, which is 11 Ma (Carlson and Hart, 1985).

All Mesozoic rocks and structures underlying the unconformity must have

likewise been tilted during this event. Removing the effects of this younger tilting, the

moderately southeast-dipping Tuff – Andesite unit contact must have been closer to

vertical and the moderately southeast-dipping Mesozoic foliations must originally have

been more steeply dipping to the southeast. In particular, the shear zone foliation (S1),

which presently dips on average 54° to the southeast (Fig. 3.6), originally dipped 77° to

the southeast, prior to late Cenozoic tilting.

150

Figure 3.23. View to the north of the Pueblo Mountains unconformity. White line is its

approximate location. Tilted Tertiary strata are on the left side of the photograph, and

Mesozoic rocks are on the right. The west flank of Pueblo Mountain is in the upper right

corner of the photograph.

151

Figure 3.24. Lower-hemisphere equal-area net projection of bedding data from Cenozoic

strata in the Pueblo Mountains. Solid dots are poles to bedding data (most data from

Roback and others, 1987; some data from this study). Line is strike and dip of average

bedding plane.

N3W, 23SW

Bedding data fromCenozoic strata

153

Cenozoic Brittle Deformation

Northwest of the Pueblo Mountains shear zone, Pueblo Terrane rocks and

overlying Tertiary strata are locally affected by a brittle deformational fabric (Fig. 3.25;

Plate 1). This deformation is not observed in the Tuff unit, or in any intrusives located in

the shear zone. The following paragraphs describe this deformation, evaluate its cause,

and establish whether or not the brittle fabric, in the northern Pueblo terrane and Tertiary

strata, is the same.

In Tertiary strata, this fabric is located slightly north of Cottonwood Canyon, at a

small area roughly one km northwest of Pueblo Mountain, and on the east side of the

range at Red Point (Fig. 3.25; Plate 1). In the Andesite unit, this fabric is most common in

its northern half, although it is observed in one location immediately south of Colony

Creek (Fig. 3.25; Plate 1).

In Tertiary or Mesozoic rocks, brittle deformation is not pervasive in the affected

rocks. Instead it only locally deforms portions of outcrop. This fabric is always planar to

very slightly undulating, Three categories of fabric spacing exist: 1-8 mm is most

common, and 4-5 cm and 30-75 cm spacings are rarer. The significance of these spacings,

if any, is uncertain. Figure 3.26 shows the chaotic character of this fabric in Andesite unit

lava, and how this fabric affects these rocks to varying degrees.

Since this fabric is present in two different packages of rocks (Cenozoic and

Mesozoic), the question arises if this is one fabric (younger than both packages of rocks)

or two discrete fabrics of different ages. Based on field observations of the deformational

character of this fabric, in both Tertiary and Mesozoic rocks, it appears that the fabric in

these rocks is the same fabric, and not two distinct fabrics. This observation is important

because it indicates that it is younger than the Cenozoic strata it affects. In an attempt to

further evaluate the nature of this fabric, its orientations measured in Tertiary strata and

154

Figure 3.25. Simplified geology of the northern portion of the eastern Pueblo Mountains

(northern half of mapped study area) emphasizing rocks affected by Cenozoic brittle

fabric (dark grey shaded areas) and location of the Cenozoic Pueblo Mountains caldera

(from Roback and others, 1987). All other geologic data from this study.

JCT

JAU

JTU

JPM

JPM

TU

QU

QU

QU

QU

QU

QU

QU

QU

QU

QU

QU

JAU

JAU

JAU

JAU

JSB

JTU

JTU

TU

TU

TU

TU

0 1 km N

Andesite Unit flows and lesservolcaniclastics and tuffs

JTU

JAU

Tuff Unit tuffaceous rocks andlesser volcaniclastics and flows

Stratigraphic units

JSB

JCT

JPM


Catlow Creek quartzmonzodiorite


Intrusive units

PUEBLO TERRANE (Jurassic)

Keyto

mapunits

QU

TUTertiaryundivided

Quaternaryundivided

Approximate margin ofPueblo Mountains caldera

Denotes areas affected byCenozoic spaced brittle fabricSymbolsNorthwest boundary of CretaceousPueblo Mountains shear zone

Range front normal fault -balls on down side

TU

QU

CottonwoodCreek

Colony Creek

Red PointPueblo

Mountain

156

Figure 3.26. Locally Andesite unit rocks are variably affected by brittle deformation.

Rocks in the left half of the photograph (covered by yellow lichen) display at least two

prominent, widely spaced, fabrics. Rocks on the right side of the outcrop are affected by

a more tightly spaced, chaotic fabric. Photograph taken near the west end of Colony

Creek Canyon.

157

in the Andesite unit (Fig. 3.27). Data in this plot is scattered and no useful pattern is

evident.

This brittle fabric cross-cuts Tertiary basaltic rocks that are 15.2-17.03 Ma (Hart

and others, 1989) and therefore must be of late Cenozoic age. The processes responsible

for this young brittle deformation cannot fully be addressed by data collected in this

study. It could be related to Cenozoic Basin and Range extension or to normal faulting

associated with the development of the Pueblo Mountain caldera (Roback, 1987). The

location of the caldera rim (Roback, 1987) and of the observed brittle fabrics are shown in

Figure 3.25. A sufficiently close spatial relationship exists in the northeastern Andesite

unit, between these features that suggests, faulting related to caldera formation, may be

the cause of brittle deformation.

Range Front Faulting and Quaternary Landslides

The entire eastern side of the Pueblo Mountains is bounded by a range front

normal fault that separates Mesozoic and Cenozoic rocks from Quaternary valley fill

(Fig. 3.1; Plate 1; location from Roback, 1987; Brown, 1996; and this study). This fault is

part of a larger fault, the Steens-Pueblo fault, that begins at the southern tip of the Pueblo

Mountains, and extends northward 130 km into southeastern Oregon (Hart and others,

1989). Along the eastern side of the Pueblo Mountains, this fault is poorly exposed in

most places and is commonly covered by Quaternary landslides. An attempt was made

by this study to structure contour the range front fault to determine its exact dip;

however, because the fault location is poorly exposed and its location, in some places, is

approximated, structure contouring along the fault line yielded no reliable information.

Large and numerous landslides are present along the eastern side of the range (Fig.

3.1; Plate 1). These landslides were evidently facilitated by the southeast dip of the

foliation in the Mesozoic rocks which is subparallel to the southeast-dipping topographic

158

Figure 3.27. Lower-hemisphere equal-area net projection of Cenozoic brittle fabric data.

Solid dots and open squares are poles to fabric in Andesite unit rocks and Tertiary strata,

respectively.

Cenozoic brittledeformation

160

slope of the eastern Pueblo Mountains. Many smaller landslides are present than are

shown in Figure 3.1 and Plate 1.

161

CHAPTER 4

SIGNIFICANCE OF THE PUEBLO TERRANE IN THE U.S. CORDILLERA

The first three chapters of this thesis document the stratigraphic, magmatic,

structural, and metamorphic Mesozoic geologic record of the eastern Pueblo Mountains.

With this knowledge, it is now possible to evaluate the significance of the Pueblo terrane

in the U.S. Cordillera. First, this chapter compares the stratigraphic and magmatic record

of the Pueblo terrane with that of the nearby Black Rock terrane (locations shown in

Figure 4.1) and discusses structural relations between these terranes. Second, this chapter

compares Pueblo terrane geology with some other geologic regions in the Cordillera.

For the purpose of these comparisons, a summary of the geologic record of the

Pueblo terrane is provided here. The Pueblo terrane is a thick sequence of Middle Jurassic

(179-176 Ma) rhyolitic to andesite pyroclastic rocks, volcaniclastics, and flows that are

intruded by shallow-level plutons (Fig. 4.2). These intrusions are contemporaneous with

volcanism, except for one early Late Jurassic (160 Ma) stock. Pueblo terrane rocks, based

on their compositions, depositional environments, and geochemical character, represent a

volcanic-plutonic stratovolcano arc complex that was likely constructed on continental,

but possibly, transitional crust.

In the mid-Cretaceous (Albian; 112-99 Ma), the eastern Pueblo Mountains were

regionally metamorphosed to greenschist grade and their southern two-thirds cross-cut by

the Pueblo Mountains shear zone. This period of compression was followed by some

combination of uplift, erosion, and non-deposition that spanned, either all or part of the

time, until the late Middle Miocene when basaltic strata (date from Hart and others,

1989) covered the Mesozoic rocks and structures of the Pueblo terrane.

162

Figure 4.1. Simplified geologic map of northwest Nevada and southeast Oregon (modified

after Wilden, 1964; Roback and others, 1987; Quinn and others, 1997; Wyld, 2000; Wyld

and Wright, 2001). D and W are the towns of Denio and Winnemucca.

Jurassic felsic tointermediatevolcanogenic strata

Upper Triassic deepmarine sedimentary strata

Pz-Mz strata ofuncertain affinity

Cretaceousplutons

Quaternary

Tertiary

Cretaceousalluvial strata

Triassic to LowerJurassic volcanic andsedimentary strata

mid to upper Paleozoicsedimentary andvolcanic strata

Jurassicplutons

PUEBLO TERRANE BLACK ROCK TERRANE

BASINAL TERRANE

Jurassicplutons

41°

42°

119°

Bla

ck R

ock

Des

ert

BLACKROCK

RANGE

PINEFORESTRANGE

OREGON

NEVADA

W

JACKSONMTNS.

BILKCREEKMTNS.

D

0 10 20 30 km

N

118°

BasinalTerrane

Pueblo Terrane

BlackRock

Terrane

WesternNevadaShearZone

PUEBLOMTNS.

Approximatedboundary betweenthe Black Rock andBasinal terranes

164

Figure 4.2. Stratigraphic summary of the Pueblo terrane represented on a Mesozoic time

scale showing timing of intrusions and deposition of stratigraphic units. Thickness of

Andesite and Tuff units shown schematically and is not actual.

JUR

AS

SIC

TR

IAS

SIC

MID

DLE

LAT

EE

AR

LY

227 m.y.

206 m.y.

180 m.y.

CONTINENTAL ARC

INTERMEDIATE TO FELSIC LAVA, TUFF,AND VOLCANICLASTIC STRATA,

AND RELATED PLUTONS

159 m.y.

andesite lava withlesser felsic

volcanic breccia,sandstone and tuff

felsic tuff withlesser andesitelava and felsic

volcanic sandstoneand breccia

shallowlevel

plutons TUFF UNITANDESITE UNIT

166

To understand the role of the Pueblo terrane in the tectonic history of the

Cordillera, the following paragraphs compare it with the nearby Black Rock terrane. This

comparison is useful because the Black Rock terrane has been linked to the North

American continent, by detrital zircon studies (Darby and others, 2000) and stratigraphic

studies (Wyld, 2000), and is considered little displaced. Therefore, establishing the

relation of the Pueblo terrane to the Black Rock terrane provide insight as to whether the

current location of Pueblo terrane is its original location, or is it an allochthonous arc

terrane.

The Pueblo terrane is different from the nearby Black Rock terrane (locations

shown in Figure 4.1). The Black Rock terrane consists of a thick sequence of Paleozoic

marine strata, mostly sedimentary, overlain by ~3 km of middle Triassic to earliest

Jurassic volcanic and sedimentary strata that reflect construction of a deep marine,

basaltic to andesitic, volcanic arc (Fig. 4.3). These rocks are intruded by several Early

Jurassic plutons (Russell, 1984; Wyld, 1990; Quinn and others, 1997), but Middle to

Late Jurassic plutonic rocks are extremely rare, and no Middle or Late Jurassic

supracrustal strata are preserved.

The Pueblo terrane is separated from the Black Rock terrane by a series of ductile

shear zones, one of which is the mid-Cretaceous Pueblo Mountains shear zone (Fig. 3.1;

Wyld and Wright, 2001). This shear zone trends northeast, dips southeast and is ~3 km

wide. Greenschist grade mylonites characterize this shear zone and the average strike and

dip of their foliations is N35oE, 54oSE. This foliation is accompanied by a well-developed

mineral and stretching lineation whose average trend and plunge is S73oE, 50o. This

orientation is nearly directly down the dip of the foliation indicating this is a dip-slip

shear zone. Numerous shear sense indicators indicate reverse sense top-to-the-northwest

directed shear and no evidence is exists for strike-slip displacement. The age and

kinematics of the Pueblo Mountains shear zone is similar to that of the other shear zones

167

Figure 4.3. Composite stratigraphic history of the Black Rock terrane shown on a

Mesozoic time scale. Data from Quinn and others (1997) and Wyld (2000).

basaltic to andesitic lavaand volcaniclastic brecciaand sandstone

JUR

AS

SIC

TR

IAS

SIC

MID

DLE

LAT

EE

AR

LY

227 m.y.

206 m.y.

180 m.y.

159 m.y.

majorunconformity

3700

m

DEEP MARINEARC

ASSEMBLAGE

BASALTIC TOANDESITIC

VOLCANOGENICSTRATA, ANDDEEP MARINESEDIMENTARY

ROCKS

limestone;siliciclastic conglomerate,sandstone, and shale

chert

basaltic lava andvolcaniclasticsandstone and breccia

andesitic lava andvolcaniclastic sandstone,breccia , and shale

shortening deformation

uplift anderosion

plutons

169

located to the southeast, in the northern Black Rock terrane (Fig. 3.1).

Based on the differences between the geologic records of the Pueblo terrane and

the Black Rock terrane, these are distinct terranes and the Pueblo terrane is not merely a

northward extension of the Black Rock terrane. In fact, the Pueblo terrane matches well

with the definition of a terrane as defined in Chapter 1 of this thesis. A terrane, as

described by Coney and others (1980), Monger (1977), and Berg and others (1978), is

characterized by internal homogeneity and has its own age, stratigraphy, lithology,

depositional environment, tectonic setting of magmatism, and timing and character of

deformation and/or metamorphism. Boundaries between terranes are fundamental

discontinuities in stratigraphy that cannot be explained easily by conventional facies

changes or unconformity. Most boundaries separate totally distinct temporal or physical

rock sequences, and are known or suspected faults that usually display complex

structural histories. So, based on this definition and the fact that the Pueblo terrane is

distinct from the Black Rock terrane and is separated from it by multiple shear zones, it is

logical to refer to the Pueblo terrane as its own terrane.

The Black Rock terrane has been linked to the continent (Darby and others, 2000;

Wyld, 2000), and since multiple shear zones separate it from the Pueblo terrane, the

questions arises as to what is the nature of the boundary between these terranes. Also, if

the Pueblo terrane is allochthonous with respect to the Black Rock terrane, then from

where did the Pueblo terrane originate. The following paragraphs address these questions.

The Pueblo terrane is different from the early Mesozoic oceanic arc complexes of

the Blue Mountains province, Black Rock terrane, Klamath Mountains, and western and

northern Sierra Nevada. (Fig. 4.4; Wright, 1982; Harper and Wright, 1984; Russell, 1984;

Wright and Fahan, 1988; Wyld, 1990; Hardwood, 1992; Hardwood, 1993; Leeman and

others, 1995; Vallier, 1995; Quinn and others, 1997). However, it is similar in its Jurassic

igneous history and depositional setting to that of the Jurassic continental arc assemblages

170

Figure 4.4. Simplified map of the western U.S. Cordillera (from Wyld and Wright, 2001)

showing Paleozoic to Mesozoic arc rocks, back-arc strata, batholith rocks, the North

American miogeocline, and the locations of the PT (Pueblo terrane), Black Rock terrane

(BRT), Blue Mountain province, Klamath Mountains, and Sierra Nevada. Heavy solid

and dashed lines show the known and inferred locations of the Mojave Snow Lake Fault

(MSLF), Western Nevada Shear Zone (WNS), and Salmon River Suture (SRS). G is the

town of Gerlach. PNR is the Pine Nut Range.

42°

46°

48°

120° 116°

44°

40°

SRS

WNS

MSLF

N

CRETACEOUSBATHOLITHS

PALEOZOIC TOJURASSIC VOLCANICARC AND RELATED

SEDIMENTARY STRATA

continental toshallow marinearc (andesitic

to felsic)

mostlyoceanic/marinearc (basaltic to

andesitic)

EARLY MESOZOICBACK-ARC

BASINAL STRATA

PRECAMBRIANTO PALEOZOIC

MIOGEOCLINALSTRATA

PNR

0 100 200 km

PT

BRT

BLUEMTN.PROVINCE

KLAMATHMTNS.

172

identified from west-central Nevada (Pine Nut Range) and the east-central Sierra south

into Arizona (Fig. 4.4; Busby-Spera and others, 1990; Saleeby and Busby-Spera, 1992;

Wyld and Wright, 1993). It is possible that one of these areas may represent the source

region from which the Pueblo terrane was displaced.

The Western Nevada shear zone (Figs. 4.1 and 4.4) proposed by Wyld and

Wright (2001) is an Early Cretaceous strike-slip crustal discontinuity. Although the

Pueblo Mountains shear zone currently records dip-slip motion, not strike-slip, it is

possible that this shear zone represents a small portion of the larger strike-slip

discontinuity. This is possible if the strike-slip fault along which the Pueblo terrane may

have been displaced northward was reactivated later during mid-Cretaceous shortening

deformation. Thus, the subsequent dip-slip shearing related to this shortening could have

obliterated structural evidence for the earlier strike-slip faulting.

In conclusion, if this speculation is correct, then the Middle Jurassic rocks of the

Pueblo terrane may have originally been situated in west-central Nevada as part of the

large continental to shallow marine arc assemblage that extends into southern California

and Arizona (Fig. 4.4). In the Early Cretaceous, the Pueblo terrane could have been sliced

away from the arc and transported several km north to its current location, thus

juxtaposing this allochthonous arc terrane against the very different Black Rock terrane

(Fig. 4.4). Subsequent compressional deformation in the mid-Cretaceous then could have

overprinted, or completely destroyed, any structural evidence of strike-slip faulting that

may have existed in the Pueblo Mountains shear zone.

173

REFERENCES

Ave Lallemant, H.G., and Oldow, J.S., 1988, Early Mesozoic southward migration of

Cordilleran transpressional terranes: Tectonics, v. 7, p. 1057-1076.

Beck, M.E. Jr., 1986, Model for late Mesozoic-early Tertiary tectonics of coastal

California and western Mexico and speculations on the origin of the San Andreas

Fault: Tectonics, v. 5, no. 1, 1986.

Berg, H.C., Jones, D.L., and Coney, P.J., 1978, Map showing pre-Cenozoic

tectonstratigraphic terranes of southeastern Alaska and adjacent areas – Pagination

2 sheets, U.S. Geological Survey open-file report, 78-1085.

Bohnam, H.F., 1969, Geology and mineral deposits of Washoe and Storey counties,

Nevada: Nevada Bureau of Mines and Geology Bulletin 70, 140 p.

Brown, M.A., 1996, The pre-Tertiary geology, structural evolution, and geochronology

of the Pueblo Mountains, Nevada-Oregon [M.S. thesis]: Houston, Texas, Rice

University, 78p.

Bruekner, H.K., and Snyder, W.S., 1985, Structure of the Havallah Sequence, Golconda

allocthon, Nevada: Evidence for prolonged evolution in an accretionary prism:

Geological Society of America Bulletin, v. 96, p. 1113-1130.

Burchfiel, B.C. and Davis, G.A., 1972, Structural framework and evolution of the

southern part of the Cordilleran orogen, western United States: American Journal

of Science, v. 272, p. 97-118.

Burchfiel, B.C. and Davis, G.A., 1975, Nature and controls of Cordilleran orogenesis,

western United States: American Journal of Science, v. 275-A, p. 363.

Burnham, R., 1971, The geology of the southern part of the Pueblo Mountains,

Humboldt County, Nevada [M.S. thesis]: Corvallis, Oregon, Oregon State

University, 114 p.

Busby-Spera, C.J., 1988, Speculative tectonic model for the early Mesozoic arc of the

southwest Cordilleran United States: Geology, v. 16, p. 1121-1125.

174

Busby-Spera, C.J., Mattinson, J.M., Riggs, N.R., and Schermer, E.R., 1990, The Triassic-

Jurassic magmatic arc in the Mojave-Sonoran Deserts and the Sierran-Klamath

region; Similarities and differences in paleogeographic evolution, in Hardwood,

D.S. and Miller, M.M., eds, Paleozoic and early Mesozoic paleogeographic

relations; Sierra Nevada, Klamath Mountains, and related terranes: Boulder,

Colorado, Geological Society of America Special Paper 255.

Busby-Spera, C.J., and Saleeby, J.B., 1990, Intra-arc strike-slip fault exposed at

batholithic levels in the southern Sierra Nevada, California: Geology, v. 18, p.

255-259.

Cann, J.R., 1970, Rb, Sr, Y, Zr, and Nb in some ocean floor basaltic rocks: Earth and

Planetary Science Letters, v. 19, p. 7-11.

Carlson, R.W., and Hart, W.K., 1985, Regional extent and significance of the Steens

Mountains type basalts in the northern Great Basin: Geological Society of

America – Abstracts with Programs, v. 17, n. 4, p. 212.

Cas, R.A.F. and Wright, J.V., 1987, Volcanic successions – modern and ancient: Allen

and Unwin, London, 528 p.

Coney, P.J., Jones, D.L., and Monger, J.W.H., 1980, Cordilleran suspect terranes:

Nature, v. 288, no. 27, p. 329-333.

Darby, B.J., Wyld, S.J., and Gehrels, G.E., 2000, Provenance and paleogeography of the

Black Rock terrane, northwestern Nevada: Implications of U-Pb detrital zircon

geochronology, in Gehrels, G.E., ed., Paleozoic and Triassic paleogeography and

tectonics of western Nevada and northern California: Geological Society of

America Special Paper 347, p. 77-87.

Deer, W.A., Howie, R.A., and Zussman, J., 1993, An introduction to the rock forming

minerals: Longman Scientific and Technical, Hong Kong, 696 p.

Edelman, S.H., 1991, Relationships between kinematics of arc-continent collision and

kinematics of thrust faults, folds, shear zones, and foliations in the Nevada

Orogen, northern Sierra Nevada, California: Tectonophysics, v. 191, p. 223-236.

Fisher, R.V. and Schmincke, H-U., 1984, Pyroclastic rocks: Springer-Verlag, Berlin, 472

p.

175

Gill, J., 1987, Orogenic andesites and plate tectonics: Springer-Verlag, Berlin, 390 p.

Gradstein, F.M., Agterberg, F.P., Ogg, J. G., Hardenbol, J., van Veen, P, Thierry, J., and

Haung, Z., 1994, A Mesozoic time scale: Journal of Geophysical Research, v.99,

no. 12, p. 24051-24074.

Hamilton, W., 1969, Mesozoic California and the underflow of pacific mantle:


Hanson, G.N., 1980, Rare earth elements in petrogenic studies of igneous systems:

Annual Review of Earth and Planetary Science, v. 8, p. 371-406.

Hardwood, D.S., 1992, Stratigraphy of Paleozoic and lower Mesozoic rocks in the

northern Sierra terrane, California: U.S. Geological Survey Bulletin 1957, 78 p.

Hardwood, D.S., 1993, Mesozoic geology of Mt. Jura, northern Sierra Nevada,

California: A progress report in Dunne, G., and McDougall, K., eds., Mesozoic

Paleogeography of the western United States-II, Pacific Section SEPM, Book 71,

p. 263-274.

Harper, G.D., and Wright, J.E., 1984, Middle to Late Jurassic tectonic evolution of the

Klamath Mountains, California-Oregon: Tectonics, v. 3, p. 759-772.

Harrold, J.L., 1972, K-Ar dates for plutonic rocks, Humboldt County, Nevada, and

Harney County, Oregon: Isochron/West 5: p. 1-5.

Harrold, J.L., 1973, Geology of the north-central Pueblo Mountains, Harney County,

Oregon [M.S. thesis]: Corvallis, Oregon, Oregon State University, 135 p.

Hart, W.K., Carlson, R.W., and Mosher, S.A., 1989, Petrogenesis of the Pueblo

Mountains Basalt, southeastern Oregon and northern Nevada, in Reidel, S.P., and

Hooper, S.R., eds., Volcanism and tectonism in the Columbia River flood-basalt

province: Boulder, Colorado, Geological Society of America Special Paper 239, p.

367-378.

Humphris, S.E., and Thompson, G., 1978, Trace element mobility during hydrothermal

alteration of oceanic basalts: Geochimica et Cosmochimica Acta, v. 42, p. 127-

136.

Jarrard, R.D., 1986, Terrane motion by strike-slip faulting of forearc slivers: Geology, v.

14, p. 783-783.

176

Lahren, M.M. and Schweikert, R.A., 1989, Proterozoic and Lower Cambrian

miogeoclinal rocks of the Snow Lake Pendant, Yosemite-Emigrant Wilderness,

Sierra Nevada, California: Evidence for major Early Cretaceous dextral translation:

Geology, v.17, p. 156-160.

Leeman, W.P., Ave Lallemant, H.G., Gerlach, D.C., Sutter, J.F., and Arculus, R.J., 1995,

Petrology of the Canyon Mountain Complex, eastern Oregon, in Vallier, T.L., and

Brooks., H.C., Geology of the Blue Mountains region of Oregon, Idaho, and

Washington; Petrology and Tectonic Evolution of Pre-Tertiary Rocks of the Blue

Mountains Region: U.S. Geological Survey Professional Paper 1438, p. 1-43.

Lund, K. and Snee, L.W., 1988, Metamorphism, structural development, and age of the

continent-island arc juncture in west-central Idaho, in Ernst, W.G., ed.,

Metamorphism and Crustal evolution of the Western United States, Rubey

Volume VII, p. 296-331.

Marshak, S. and van der Pluijm, B.A., 1997, Earth structure: An introduction to

structural geology and tectonics, McGraw-Hill Companies, Inc., U.S.A., 495 p.

Mason, R., 1978, Petrology of the metamorphic rocks, George Allen and Unwin,

Limited, London, 254 p.

McDougall, I. And Harrison, T.M., 1999, Geochronology and thermogeochronology by

the 40Ar/39Ar method, Oxford University Press, Inc., New York, 269 p.

Miller, E.L., Holdsworth, B.K., Whiteford, W.B., and Rodgers, D., 1984, Stratigraphy

and structure of the Schoonover sequence, northeastern Nevada: Implications for

Paleozoic plate-margin tectonics: Geological Society of America Bulletin, v. 95, p.

1063-1076.

Monger, J.W.H., 1977, Upper Paleozoic rocks of the western Canadian Cordillera and

their bearing on Cordilleran evolution: Canadian Journal of Earth Sciences, 14, no.

8, p. 1832-1859.

Myrashiro, A., 1994, Metamorphic Petrology, Oxford University Press, Inc., New York,

New York, 404 p.

Nakamura, N., 1974, Determination of REE, Ba, Fe, Mg, Na and K in carbonaceous and

ordinary chondrites, Geochimica et Cosmochimica Acta, v. 38, p. 757-775.

177

Passchier, C.W. and Trouw, R.A.J., 1998, Micro-tectonics: Springer-Verlag, Berlin, 289

p.

Pearce, J.A., 1982, Trace element characteristics of lavas from destructive plate

boundaries: in Thorpe, R.S., ed., Andesites, John Wiley & Sons, p. 525-548.

Pearce, J.A., 1983, Role of the sub-continental lithospheric magma genesis at active

continental margins, in Hawkesworth, C.J., and Norry, N.J., eds., Continental

basalts and mantle xenoliths, Shiva Publishing Ltd., Cheshire, p. 230-250.

Philpotts, A.R., 1990, Principles of igneous and metamorphic petrology, Prentice Hall,

Englewood Cliffs, New Jersey, 498 p.

Quinn, M.J., Wright, J.E., and Wyld, S.J., 1997, Happy Creek igneous complex and

tectonic evolution of the early Mesozoic arc in the Jackson Mountains, northwest

Nevada: Geological Society of America Bulletin, v. 109, p. 461-482.

Raymond, L.A., 1995, Petrology: the study of igneous, sedimentary, and metamorphic

rocks, Wm. C. Brown Communications, Inc., Dubuque, Iowa, 742 p.

Reading, H.G., 1996, Sedimentary environments: processes, facies, and stratigraphy,

Blackwell Science Ltd., Oxford, 688 p.

Roback, R.C., Vander Meulen, D.B., King, H.D., Plouff, D., Munts, S.R., and Willett,

S.L., 1987, Mineral resources of the Pueblo Mountains wilderness study area,

Harney County, Oregon, and Humbolt County, Nevada: United States Geological

Survey Bulletin B1-B30 (1740-B).

Roback, R.C., 1988, Deformed Late Jurassic volcanic-plutonic center in the Pueblo

Mountains, Oregon-Nevada: Anonymous Abstracts with Programs – Geological

Society of America, 20, 3, p.224.

Rowe, W., 1971, Geology of the south-central Pueblo Mountains, Harney County,

Oregon [M.S. thesis]: Corvallis, Oregon, Oregon State University, 81 p.

Russell, B.J., 1984, Mesozoic geology of the Jackson Mountains, northwest Nevada:


Saleeby, J.B. and Busby-Spera, C., 1992, Early Mesozoic tectonic evolution of the

western U.S. Cordillera, in Burchfiel, B.C., Lipman, P.W., and Zoback, M.L., eds,

The Cordilleran Orogen: Conterminous U.S.: Boulder, Colorado, Geological

Society of America, The Geology of North America, v. G-3.

178

Schweickert, R.A., and Cowan, D.S., 1975, Early Mesozoic tectonic evolution of the

western Sierra Nevada, California: Geological Society of America Bulletin, v. 86,

p. 1329-1336.

Schweickert, R.A. and Lahren, M.M., 1990, Speculative reconstruction of the Mojave

Snow Lake Fault: Implications for Paleozoic and Mesozoic orogenesis in the

western U.S.: Tectonics, v.9, p. 1609-1629.

Shervais, J.W., 1982, Ti-V plots and the petrogenesis of modern and ophiolitic lavas:

Earth and Planetary Science Letters, v. 59, p. 101-118.

Silberling, N.J., Jones, D.L., Blake, M.C., Jr., and Howell, D.G., 1984, Lithotectonic

terrane map of the western conterminous United States, in Silberling, N.J. and

Jones, D.L., eds., Lithotectonic terrane maps of the North American Cordillera:

U.S. Geological Survey Open-File Report, C1-C43.

Speed, R.C., 1979, Collided Paleozoic microplate in the western U.S.: Journal of

Geology, v.87, p. 279-292.

Sun, S.S., 1980, Lead isotope study of young volcanic rocks from mid-ocean ridges,

oceanic islands and island arcs, in Bailey, D.K., Tarney, J., and Dunham, K., eds.,

The evidence for chemical heterogeneity in the Earth’s mantle: Philosophical

transactions of the Royal Society of London, Series A: Mathematical and Physical

Sciences 297, no. 1431, p. 405-445.

Sun, S.S. and McDonough, W.F., 1989, Chemical and isotope systematics of oceanic

basalts; implications for mantle composition and processes, in Saunders, A.D. and

Norry, M.J., eds., Magmatism in the ocean basins: Geological Society (London)

Special Publications, 42, 313-345.

Tickoff, B., and de Saint Blanquat, M., 1997, Transpressional shearing and strike-slip

partitioning in the Late Cretaceous Sierra Nevada magmatic arc, California:

Tectonics, v. 16, p. 442-459.

Tower, D., 1972, Geology of the central Pueblo Mountains, Harney County, Oregon

[M.S. thesis]: Corvallis, Oregon, Oregon State University, 96 p.

Turner, F.J., 1981, Metamorphic petrology: Mineralogical, field, and tectonic aspects,

Hemisphere Publishing Corporation, U.S.A, 524 p.

179

Twiss, R.J. and Moores, E.M., 1992, Structural geology, W.H. Freeman and Company,

New York, 532 p.

Vallier, T.L., 1995, Petrology of pre-Tertiary igneous rocks in the Blue Mountains region

of Oregon, Idaho, and Washington: Implications for the geologic evolution of a

complex islan arc, in Vallier, T.L., and Brooks., H.C., Geology of the Blue

Mountains region of Oregon, Idaho, and Washington; Petrology and Tectonic

Evolution of Pre-Tertiary Rocks of the Blue Mountains Region: U.S. Geological

Survey Professional Paper 1438, p. 125-210.

Wilson, M., 1989, Igneous petrogenesis: A global tectonic approach, Unwin Hyman Ltd.,

London, 466 p.

Winkler, H.G.F., 1979, Petrogenesis of metamorphic rocks, Springer-Verlag, New York,

348 p.

Wolak, C. and Wyld, S.J., 2000, Mesozoic structure and stratigraphy of the Pueblo

Mtns., SE Oregon: Record of an allochthonous terrane in the U.S. Cordillera:

Abstracts with Programs – Geological Society of America: Cordilleran Section, v.

32, no. 6, p. 76.

Wood, D.A., 1980, The application of a Th-Hf-Ti diagram to problems of

tectonomagmatic classification and to establishing the nature of crustal

contamination of basaltic lavas of the British Tertiary volcanic province: Earth and

Planetary Science letters, v. 50, p. 11-30.

Wright, J.E., 1982, Permo-Triassic subduction complex, southern Klamath Mountains,

northern California: Journal of Geophysical Research, v. 87, p. 3805-3818.

Wright, J.E., and Fahan, M.R., 1988, An expanded view of Jurassic Orogenesis in the

western Unites States Cordillera: Middle Jurassic (pre-Nevadan) regional

metamorphism and thrust faulting within an active arc environment, Klamath

Mountains, California: Geological Society of America Bulletin, v. 100, p. 859-876.

Wyld, S.J., 1990, Paleozoic and Mesozoic rocks of the Pine Forest Range, northwest

Nevada, and their relation to volcanic arc assemblages of the western U.S.

Cordillera, in Hardwood, D.S., and Miller, M.M., eds., Late Paleozoic and early

Mesozoic paleogeographic relations: Klamath Mountains, Sierra Nevada, and

related rocks: Geological Society of America Special Paper 255, p. 219-237.

180

Wyld, S.J., 1996, Early Jurassic deformation in the Pine Forest Range, northwest Nevada,

and implications for Cordilleran tectonics: Tectonics, v.15, p. 566-538.

Wyld, S.J., 2000, Triassic evolution of the arc and backarc of northwestern Nevada, and

evidence for extensional tectonism, in Soreghan, M.J., and Gehrels, G.E., eds.,

Paleozoic and Triassic paleogeography and tectonics of western Nevada and

northern California: Boulder, Colorado, Geological Society of America Special

Paper 347, p. 185-207.

Wyld, S.J., Quinn, M.J., and Wright, J.E., 1996, Anomalous (?) Early Jurassic

deformation in the western U.S. Cordillera: Geology, v.24, no. 11, p. 1037-1040.

Wyld, S.J., and Wright, J.E., 1993, Mesozoic stratigraphy and structural history of the

southern Pine Nut Range, west-central Nevada, in Dunne, G. and McDougall, K.,

eds, Mesozoic paleogeography of the western United States-II: Pacific Section

SEPM, Book 71, p. 289-306.

Wyld, S.J., and Wright, J.E., 1999, Cretaceous crustal discontinuity in northwest Nevada

and southeast Oregon: Manifestation of a major terrane boundary extending from

the Sierra Nevada to western Idaho?: in Geological Association of Canada, Terrane

Paths 99 Circum Pacific Terrane Conference, Abstracts and Programs, p. 73-75.

Wyld, S.J. and Wright, J.E., 2000, Major tectonomagmatic events in the Mesozoic U.S.

Cordillera and implications for plate interactions: Geological Society of America

Abstracts with Programs, v. 32, no. 7, p. A46.

Wyld, S.J., and Wright, J.E., 2001, New evidence for Cretaceous strike-slip faulting in

the U.S. Cordillera; and implications for terrane-displacement, deformation

patterns and plutonism: American Journal of Science.

Yardley, B.W.D., 1989, An introduction to metamorphic petrology, Longman Scientific

and Technical, New York, 248 p.

JSB

altered portion ofthe StrawberryButte quartz diorite

JDI

?

22

46

19

19

74

22

34

24

23 34

23 42

55

51

41

4664

78

76

51

45

66

614143

76

39

57

39

74

70

61

65

36

33

34

3646

2249

46

72

42

42

54

55

47

62

5546

50

50

63 6029

42

4046

35215 66

62

62

356767

5968

5559

41

41

61

1440

40

41

49

41

25

2326

5056

547937

37

11

21

26

40

47

4828

58

49

4439

20

39

3185

50

2135

47 6949

51

49

44

59

46

69 63

32

65

51

69

5460

52

4364

3021

54

50

39

49

65

25

52

55

34

35

530

51

51

19

35

45

48

1850

37

3739

39 29

53

37

32

8385

20

32

1950

30

30

45

55

32

39

61

26

25

14

4555

54

21 6666 65 80

66

30

43

65

44

70

55 35

43 32

35

52

29

49

51

74

76

52

49

73

47

37

41

45

21

30

23

21

13

22

4142

23

2625

7184

49

25

62

45

31

39

25

42

31

1572

80

56

34

39

44

58

45

6450

6570

69

6255

39

35

6633

29

57

65

82

85

86

74

66

78

8756

6638

59

8537

6861

31

5475

21

55

4640

38

58 35

37

30

46

36

5430

86

64 70

72

69

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4643

37

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8880

31

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75

70 62

54

5048

45 60

59

48

8072

34

78

60

80

75

62

35

39

41

39

40

28

GEOLOGY OF THE EASTERNPUEBLO MOUNTAINS

C. E. Wolak, 2001

JTU - Tuff unit (179 +/- 2 Ma) -rhyolitic to dacitic pucime and ashtuffs

JTU

JAU - Andesite unit(176 +/- 2 Ma) - andesite flows

JAUB - rhyolitic to daciticvolcanic breccia

JAUT - rhyolitic ashflow tuff

JAUS - dacitic volcanincsandstone

JAU

JAUB

JAUT

JAUS

JTUL

JTUB

JTUS

JTUB - rhyolitic to andesiticvolcanic breccia

JTUS - dacitic to intermediatevolcanic sandstone

JTUL - dacitic andesite lava

Pueblo Terrane

JSBa

JPM

JCT

Strawberry Buttequartz diorite(176 +/- 3 Ma)

Pueblo Mountainhypabyssalandesite

Catlow Creekquartzmonzodiorite(~179 +/- 2 Ma)

JDI

Diamond Innhypabyssal rhyolite(160 +/- 2 Ma)

AlluviumQal

Tertiary

CretaceousGranodiorite pluton(108 - 115 Ma)

Landslide - arrow showsmovement directionQls

JCDCowden Creek tonaliteto quartz diorite(188 +/- 2 Ma)



JBZ

Qua

tern

ary

ColluviumQcm

Tu

Jura

ssic

KBK

zircon sample locality

40/39 Ar sample locality

geochemistry sample locality

Cenozoic normal fault - ballson downthrown side

ductile fault - teeth onupthrown side

NW boundary of Cretaceous Pueblo Mountainsshear zone - teeth on shear zone side

contact - dashed whereapproximated and dottedwhere concealed

strike and dip of bedding

strike and dip of S1foliation


flattened pumice or vesicles

strike and dip of compaction fabric

SYMBOLS

C plane foliation

S3 foliation

spaced brittle cleavage

S plane foliation

LEGEND

Ju

rassic

Intrusive units Stratigraphic units

D2 fold - arrow indicatestrend and plunge of fold axis

JCD

JBZ

JBZ

KBK

KBK

JCD

JTU

Qal

Qcm

JCD

Qal

JSB

JSB

JCD

Qls

TU

Qls

Qls

JTU

JTU

JSB

JSBJSB

JSBa

JSBa

JSBa

JTU

TU

Qal

JSBa

JSB

QlsQls

Qls

JSB

TU

Qls

JTUQls

JTU

JTU

JTU

JTU

JTU

Qls

Qal

TU

Qal

JTU

Qls

Qls

Qal

Qls

JPM

Qcm

Qcm

JAU

JAU

TU

TU

TU

TU

TU JPM

Qcm

QalQal

Qls

QlsQcm

JAU

JAU

JAU

JAU

JCT

QlsQls

Qal

JAUB

JAUT &

JAUS

JAUB

TU

TU

JAUB

JAUB

JAUB

JAUB

JAUP

JAUP

JAUT

JTUL

JTUB

JTUS

JTUS

JTUS

JTULJTUB

JTUS

JTUB

JTUS

TU

TU

TU

TU

TU

A’

A

B

B’

C’

C

D

D’

TU

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