Geology of the Blue Mountains Region of 

Oregon, Idaho, and Washington:

Petrology and Tectonic Evolution 

of Pre-Tertiary Rocks of thte 

Blue Mountains Region

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Geology of the Blue Mountains Region of 

Oregon, Idaho, and Washington:

Petrology and Tectonic Evolution 

of Pre-Tertiary Rocks of the 

Blue Mountains Region

TRACY L. VALUER and HOWARD C. BROOKS, editors

U.S. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1438

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON : 1995
U.S. DEPARTMENT OF THE INTERIOR 

BRUCE BABBITT, Secretary

U.S. GEOLOGICAL SURVEY 

Gordon P. Eaton, Director

Any use of trade, product, or firm names 

in this publication is for descriptive purposes only and 

does not imply endorsement by the U.S. Government

Text and illustrations edited by 

Helen Gibbons and Andrea C. Eddy

Library of Congress Cataloging-in-Publication Data

Petrology and tectonic evolution of pre-Tertiary rocks of the Blue Mountains region / Tracy L. Vallier and Howard C. 

Brooks, editors.

p. cm. (Geology of the Blue Mountains region of Oregon, Idaho, and Washington) (U.S. Geological Survey 

professional paper ; 1438) 

Includes bibliographical references. 

Supt. of Docs, no.: 119.16: 1438

1. Geology, Structural Blue Mountains Region (Or. and Wash.) 2. Island arcs Blue Mountains Region (Or. and 

Wash.) 3. Rocks Blue Mountains Region (Or. and Wash.) I. Vallier, T. L. (Tracy L.) II. Brooks, Howard C. 

III. Series. IV Series: U.S. Geological Survey professional paper ; 1438. 

QE747.B58P48 1995

551.7'009795 dc20 94-48560 

________________________________________________________________________CIP

For sale by U.S. Geological Survey, Information Services 

Box 25286, Federal Center, Denver, CO 80225
PREFACE

U.S. Geological Survey Professional Paper 1438 is 

one volume of a five-volume series on the geology, 

paleontology, and mineral resources of the Blue 

Mountains region in eastern Oregon, western Idaho, 

and southeastern Washington. This professional paper deals specifically with petrology and tectonic evolution. Other professional papers in the series include 

Professional Paper 1435 on paleontology and biostratigraphy, Professional Paper 1436 on the Idaho 

Batholith and its border zone, Professional Paper 

1437 on Cenozoic geology, and Professional Paper 

1439 on stratigraphy, physiography, and mineral resources. The purpose of these volumes is to familiarize readers with work that has been completed in the 

Blue Mountains region and to emphasize the region's 

importance for understanding island-arc processes 

and the accretion of an allochthonous terrane to a 

continent. These professional papers provide current 

interpretations of a complex island-arc terrane that 

was accreted to ancient North America in the late 

Mesozoic Era, of a large batholith that was emplaced 

after accretion had occurred, and of overlying Cenozoic volcanic rocks that were subsequently uplifted and 

partly stripped off the older rocks by erosion.

Modern island arcs are not well understood, and 

even less so are ancient arcs that have been deformed, metamorphosed, and subsequently accreted 

to continents. We have learned that characteristics of 

modern arcs change significantly both along and 

across the arcs' axes and that studies of arc fragments are less than satisfactory because they generally do not characterize an entire arc. For example, 

the landward trench slopes of arcs can differ greatly, 

depending on whether materials from the descending 

slab are being accreted or the slope is being tectonically eroded; which process dominates apparently is 

related to the volume of sediment in the adjacent 

trench and the vector of plate convergence. In addition, some arcs (Aleutian) have broad, long, and sediment-filled fore-arc basins, whereas in others (TongaKermadec) the fore-arc insular slopes descend precipitously to trench depths and are only interrupted in

places by narrow fault-bounded terraces. Moreover, 

some arcs (Tonga-Kermadec) have erupted primarily 

tholeiitic igneous products throughout their histories, 

and others (Aleutian) have a long history of both calcalkaline and tholeiitic eruptive activity. Ridge axes of 

island arcs may be narrow or broad, and in some arcs 

(Solomons and Vanuatu), the axial regions have extended to form deep bathymetric and sedimentary basins. Even back-arc basins have different origins and 

histories of development. Some (Mariana Trough and 

Lau Basin) have active spreading ridges, whereas 

others (Aleutian basin) are floored by ancient oceanic 

crust that was trapped behind the arc.

Because our knowledge of the diverse processes 

within modern arcs is limited, it becomes even more 

important to study ancient analogs. Just by the nature of their on-land exposures, ancient arcs can provide insights into sedimentary facies, magmatic 

evolution, and deep crustal processes that can only be 

studied in modern arcs by geophysical methods, 

dredging, and drilling. Few ancient island arcs have 

exposures as well developed and as extensive as 

those in the Blue Mountains region. Particularly 

spectacular and helpful are outcrops provided by intensive stream erosion, which has left some canyon 

walls more than 2 km deep (Snake River canyon west 

of the Seven Devils Mountains).

Most earth scientists who have worked in the Blue 

Mountains region agree that pre-Tertiary rocks there 

form one or more allochthonous terranes. The importance of such terranes in the evolution of circum-Pacific continental margins has been recognized for 

more than a decade, but many complex questions remain. For example, how, when, and where did most 

of the circum-Pacific allochthonous terranes form? 

How did they accrete to continents? What are the 

mechanisms of amalgamation processes during terrane. formation and transport? And, perhaps most 

importantly, what are the effects of these processes 

on mineral and hydrocarbon resources? While these 

volumes provide some answers, the data and interpretations contained in them will no doubt raise new 

and equally intriguing questions for future generations of earth scientists.

in

CONTENTS

[Numbers indicate chapters]

Page

1. Petrology of the Canyon Mountain Complex, eastern Oregon 1

William P. Leeman, Hans G. Ave Lallemant, David C. Gerlach, John F. Sutter, and 

Richard J. Arculus

2. A closer look at the Bald Mountain batholith, Elkhorn Mountains, and some comparisons with the Wallowa batholith, Wallowa Mountains, northeastern Oregon 45 

William H. Taubeneck

3. Petrology of pre-Tertiary igneous rocks in the Blue Mountains region of Oregon, 

Idaho, and Washington: Implications for the geologic evolution of a complex 

island arc 125

Tracy L. Vallier

4. High-pressure, low-temperature schistose rocks of the Baker terrane, northeastern

Oregon 211

Ellen M. Bishop

5. Mafic and ultramafic rocks of the Baker terrane, eastern Oregon, and their implications for terrane origin 221

Ellen M. Bishop

6. Tectonic implications of U-Pb zircon ages of the Canyon Mountain Complex,

Sparta complex, and related metaplutonic rocks of the Baker terrane, northeastern 

Oregon------------------ - - - 247

Nicholas W Walker

7. Pre-Cretaceous tectonic evolution of the Blue Mountains province, northeastern

Oregon 271

Hans G. Ave Lallemant

8. Pre-Tertiary deformation in the Desolation Butte quadrangle, northeastern

Oregon --------- -- -- - ------- 305

James G. Evans

9. The Bourne and Greenhorn subterranes of the Baker terrane, northeastern Oregon: 

Implications for the evolution of the Blue Mountains island-arc system 331 

Mark L. Ferns and Howard C. Brooks
VI CONTENTS

10. An 40Ar/39Ar chronicle of the tectonic development of the Salmon River suture

zone, western Idaho 359

Lawrence W. Snee, Karen Lund, John F. Sutter, David E. Balcer, and Karl V. Evans

11. Geology of the northern part of the Ironside Mountain inlier, northeastern Oregon 415

Peter R. Hooper, Michel D. Houseman, John E. Beane, Gregory M. Caffrey, Kenneth R. Engh, 

James V. Scrivner, and A. John Watkinson

12. Petrology and deformation history of the Burnt River Schist and associated

plutonic rocks in the Burnt River Canyon area, northeastern Oregon 457 

Roger P. Ashley

13. Gravity studies of an island-arc/continent suture zone in west-central Idaho and

southeastern Washington 497

Gregory B. Mohl and Richard L. Thiessen

14. Metamorphic and structural development of island-arc rocks in the Slate CreekJohn Day Creek area, west-central Idaho 517

Karen Lund
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX,

EASTERN OREGON

By WILLIAM P. LEEMAN/ HANS G. AVE LALLEMANT, 1 DAVID C. GERiACH,2

JOHN F. SUTTER,3 and RICHARD J. ARCULUS4

CONTENTS

Page

Abstract________________________________ 1

Introduction 1

Acknowledgments 2

Regional geologic setting 2

Geochronology 5

Canyon Mountain Complex 5

Sparta complex 6

Geology of the Canyon Mountain Complex 6

Previous investigations 6

Geologic descriptions of units 9

Ultramafic member 9

Gabbroic member 10

Sheeted (sill-like) member 10

Structure _____________________________ 11

Petrography and mineralogy 11

Paleothermometry and paleobarometry 12

Geochemistry and petrology 12

Analytical methods 13

General geochemical characteristics 13

Comparative geochemistry and evaluation of

tectonic setting 20

Petrogenetic models 30

Plagiogranite suite 30

T^Tnfic rocks -------________ «- - . _ _-._.__« ^2

Summary 37

Tectonic implications 38

References cited 39

ABSTRACT

The Permian Canyon Mountain Complex of northeastern Oregon 

has many attributes of ophiolites, including a stratified structure 

and characteristic lithologic succession. The stratigraphic sequence

Department of Geology and Geophysics, Rice University, Houston, TX 77251 

2Charles Evans and Associates, 301 Chesapeake Dr., Redwood City, CA 94063 

3U.S. Geological Survey, Reston, VA 22092

4Present address: Department of Geology, Australian National University, 

Canberra ACT 0200, Australia

Manuscript approved for publication February 18, 1994.

consists of a basal harzburgite tectonite overlain successively by an 

ultramafic-to-mafic cumulate member (penetratively deformed in 

the eastern part of the complex and undeformed in the western 

part), and an upper sheeted member (plagiogranites, mafic dikes 

and sills, and keratophyres). The majority of the keratophyres appear to have formed prior to the sheeted member and may represent 

part of a Permian island arc. The tectonite and metamorphosed 

cumulate rocks are cut by undeformed isotropic gabbro, which may 

be cogenetic with the layered cumulate gabbros and possibly some 

of the mafic dikes. However, because clinopyroxenes from basaltic 

rocks of the sheeted member differ (lower Ca contents and higher 

Fe/Mg ratios) from those in the gabbros, it is probable that most of 

the mafic dikes were emplaced during a separate event. Field relations and 40Ar/39Ar age determinations demonstrate that the plagiogranites, mafic dikes, and some keratophyres were emplaced in 

the sheeted member more or less contemporaneously at about 260 

Ma. U-Pb zircon geochronology for a single high-level cumulate gabbro (Walker, 1986) suggests a time interval of perhaps 15 m.y. between formation of the gabbro and emplacement of the sheeted 

member. The mafic dikes, sills, and minor flows apparently formed 

by partial melting of a source that was strongly depleted with 

respect to incompatible trace elements. A source resembling that for 

midocean-ridge basalt (MORE) from which melts had been previously extracted is plausible. The plagiogranite suite appears to have 

been derived by hydrous partial melting of a mafic source rock (such 

as amphibolite), possibly representing subducted oceanic crust. 

Thus, the Canyon Mountain Complex appears to be a composite 

body whose members formed over a significant period of time, possibly in different tectonic settings; there is no compelling reason to 

propose a cogenetic relation between magmas of the sheeted member and other rocks of the complex. For the mafic magmas, discriminant diagrams indicate transitional affinities between ocean-floor 

and primitive island-arc basalts. However, striking depletions in Ta 

and Hf, relative to other incompatible elements, suggest a subduction-related volcanic-arc (or possibly fore-arc or marginal-basin) setting for their origin.

INTRODUCTION

Participants at a Penrose Field Conference concluded that ophiolites are pseudostratified rock assemblages (Anonymous, 1972) comprising four diagnostic 

lithologic units or members from bottom to top, an 

ultramafic member, a gabbroic member, a mafic sheeted-dike complex, and a mafic volcanic complex as
PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

well as associated minor rock types that commonly 

are found in ophiolites. These associated rock types 

include chromitites, sodic felsic intrusive and extrusive rocks (such as trondhjemites, keratophyres), and 

sedimentary rocks (such as chert, shale, limestone). 

Although no genetic significance was intended in the 

Penrose Conference definition, many geologists consider ophiolites to have formed at midocean spreading 

centers (see Coleman, 1981). In recent years, the recognition that ophiolites have quite varied characteristics has led to alternative interpretations of their 

tectonic settings and processes of formation.

One or more of the four key lithologic members 

may be absent in surface exposures of some ophiolites as a result of tectonic disruption. However, 

many apparently intact ophiolite complexes lack one 

or more of these units, or have units that are similar 

but occur in uncharacteristic style. For example, the 

Point Sal ophiolite (southern California) has a sill 

member but no dike member (Hopson and Frano, 

1977). The (informal) Sparta (ophiolite) complex of 

eastern Oregon is characterized by an unusually 

large proportion of silicic intrusive rocks (Prostka 

and Bateman, 1962; Phelps, 1978, 1979); in fact, nontrivial volumes of felsic igneous rocks are associated 

with most so-called ophiolites in Oregon and California (Bailey and Blake, 1974). Coleman and Peterman 

(1975) designated such silicic intrusive rocks as "oceanic plagiogranites" and proposed that they formed 

by extensive fractional crystallization of basaltic 

magmas at midocean ridges.

Alternative tectonic environments proposed to account for the now-recognized ophiolite variants include actively spreading back-arc basins such as the 

Marianas (Karig, 1971), leaky transform faults such 

as the Cayman Trough (Thompson and others, 1980; 

Holcombe and Sharman, 1983), and island arcs such 

as Tonga (Ewart and Bryan, 1972). The uncertainty 

over tectonic setting is well illustrated by the various 

interpretations of the Troodos ophiolite (Cyprus); an 

island-arc origin is suggested on the basis of geochemical and petrologic data (see Miyashiro, 1973; McCulloch and Cameron, 1983; Robinson and others, 1983), 

but this model has been widely contested (for example, Hynes, 1975; Moores, 1975; Church and Coish, 

1976). Multiple magmatic suites that may have originated in different tectonic settings and (or) at different times have been recognized in some cases (see 

Coleman, 1977; Thayer, 1977b), including the Betts 

Cove (Coish and others, 1982), New Caledonia (Dupuy 

and others, 1981), Troodos (Robinson and others, 

1983; Schminke and others, 1983), Vourinos (Noiret 

and others, 1981; Beccaluva and others, 1979), Pindos 

(Capedri and others, 1980), and Oman ophiolites (Alabaster and others, 1982). Juxtaposition of coherent 

stratal sequences formed in distinct settings may be 

more the rule than an exception in convergent-margin 

tectonic zones (Moores and others, 1984).

The Canyon Mountain Complex, Sparta complex, 

and several other possible ophiolite fragments occur 

in the Blue Mountains province of northeastern Oregon (see Mullen, 1985). This province consists of five 

tectonostratigraphic terranes (Silberling and Jones, 

1984) believed to have formed at indeterminate distances from the North American craton and then to 

have been accreted to the craton at later times (see 

Coney and others, 1980). In order to reconstruct the 

plate tectonic evolution of these terranes, it is necessary to understand the original tectonic settings of 

ophiolite remnants preserved in suture zones separating the terranes.

The Canyon Mountain Complex is the best exposed 

and most completely preserved ophiolite in the province. Herein, we review published petrologic, geochemical, and structural data, and present new 

geochemical data for this body, especially for the basaltic components. We specifically address the petrogenesis of felsic and mafic magmas associated with 

the sheeted member and late-stage intrusions, and 

the geochemical constraints on the type of tectonic 

setting in which these rocks formed. To further elucidate ages of formation, we also present new radiometric ages for samples from the Canyon Mountain 

Complex and Sparta complex.

ACKNOWLEDGMENTS

We are grateful to Bruce Chappell, Mike Rhodes, 

Ben Powell, Doug Blanchard, and Roman Schmitt for 

access to analytical facilities. Lisanne Pearcy and 

Tracy Vallier provided thorough and constructive 

reviews. This work was supported by the National 

Science Foundation through grants GA-25894, EAR76-13372, EAR-79-04829, and EAR-82-06531.

REGIONAL GEOLOGIC SETTING

Postulated remnants of ophiolites occur throughout 

the Blue Mountains province of northeastern Oregon 

and adjacent areas of Idaho. Their bearing on regional tectonic interpretations is discussed within the 

context of four of the five tectonostratigraphic terranes of the province. From north to south, these are 

the Wallowa, Baker, Izee, and Olds Ferry terranes 

(Silberling and Jones, 1984; the fifth terrane, the 

Grindstone, will not be discussed in this chapter). An
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON

oceanic affinity is inferred for these terranes on the 

basis of their internal geology and the fact that the 

igneous rocks within each terrane are characterized 

by initial Sr isotopic compositions (87Sr/86Sr) at or 

below 0.704 (Armstrong and others, 1977). There is 

no compelling evidence for the presence of old continental crust within the province.

The Wallowa terrane (formerly called the Seven 

Devils terrane by Dickinson, 1979) consists of Permian and Upper Triassic volcaniclastic and volcanic 

rocks with affinity to low-I^O basalt-andesite-daciterhyolite suites typical of present-day island arcs (Vallier and Batiza, 1978). In three areas in the Snake 

River canyon (Vallier, 1974), the volcanic rocks overlie 

gabbro, plagiogranite, and diabase of Permian to Triassic age (Balcer, 1980; Walker, 1981, 1982); contacts 

between the plutonic and volcanic rocks are intrusive 

or tectonic. The supposed basement exposures in the 

Snake River canyon are lithologically similar to the 

Late Triassic Sparta complex (fig. 1.1), which is usually included in the Baker terrane. Phelps (1978, 

1979) and Ave Lallemant and others (1980) concluded 

that the Sparta complex is the plutonic equivalent of 

a low-K2O island-arc volcanic suite; they suggested 

that it is a part of the root of the Wallowa arc. Volcanic rocks of the Wallowa terrane are overlain by Upper 

Triassic limestone and Upper Triassic to lower Upper 

Jurassic clastic sedimentary rocks (Morrison, 1964; 

Nolf, 1966; Vallier, 1974, 1977).

The Baker terrane (formerly called the Central Melange terrane by Dickinson, 1979) consists of disrupted fragments of oceanic crust, including local melange 

deposits, large coherent tectonic slices of ophiolitic 

and volcanic assemblages, and Triassic to possibly 

Lower Jurassic chert-argillites (Dickinson and Vigrass, 1965; Blome and others, 1983; Nestell, 1983; 

Blome and others, 1986). Analyses of basaltic and 

gabbroic blocks (Mullen, 1985; Bishop, chap. 5, this 

volume) indicate that they are varied in composition 

and include tholeiitic (that is, midocean-ridge), calcalkaline, and rare alkalic basalt types. Late Triassic 

blueschists occur near Mitchell (fig. 1.1; Hotz and others, 1977), and somewhat lower pressure schistose 

rocks occur in serpentine-matrix melange elsewhere 

in the Baker terrane (Bishop, chap. 5, this volume). 

The Canyon Mountain Complex is by far the largest 

and most complete ophiolite fragment in this terrane.

The Olds Ferry terrane (formerly called the Huntington terrane by Dickinson, 1979) is a poorly exposed Late Triassic volcanic arc (Brooks and Vallier,

1978). It consists of Upper Triassic volcanic (basalt, 

andesite, and rhyolite) and volcaniclastic rocks and 

rare limestone (Brooks and Vallier, 1978; Brooks,

1979). Gabbro and plagiogranite plutons were emplaced locally during the Late Triassic (Henricksen, 

1975). Lower Jurassic volcanogenic flysch-like deposits, technically overlying the Triassic rocks (Ave Lallemant, 1984a), were previously included in the 

Mesozoic Clastic terrane of Dickinson (1979).

The Izee terrane (formerly called the Mesozoic 

Clastic terrane by Dickinson, 1979) comprises Upper 

Triassic to Upper Jurassic turbidites deposited on the 

Baker terrane. Volcanic tuff beds occur mainly in its 

lower parts. Dickinson (1979) proposed that the turbidites were deposited in a fore-arc basin associated 

with the Olds Ferry arc to the south; however, the 

upper units of Jurassic age may represent an overlap 

sequence. Fault-bounded lenses and blocks of chertargillite, gabbro, and serpentinite occur throughout 

this terrane.

Two major erogenic events apparently affected the 

rocks in the Blue Mountains province (Ave Lallemant 

and others, 1980; Ave Lallemant, chap. 7, this volume). First, during Triassic and Early Jurassic time, 

south to southeastward subduction of an oceanic plate 

beneath the Olds Ferry volcanic arc caused the formation of the melange and blueschists of the Baker terrane, and resulted in tectonic imbrication and folding 

of the melange into Upper Triassic and earliest Jurassic forearc-basin turbidites (Brown and Thayer, 1966; 

Dickinson and Thayer, 1978; Ave Lallemant, chap. 7, 

this volume). Exposed rocks of the Olds Ferry terrane 

do not appear to be affected by the Late Triassic deformation. However, a Late Triassic to Early Jurassic 

depositional hiatus signifies uplift of the terrane at 

that time (Brooks and Vallier, 1978). Neither unconformities nor folds of this age are apparent in the 

Wallowa terrane; Late Triassic deformation there 

seems to be confined to three major northeast-trending left-lateral shear zones (Vallier, 1974; Ave Lallemant, 1983; Ave Lallemant and others, 1985).

The second erogenic event was manifested by shortlived north-south shortening that affected all terranes 

(Ave Lallemant, 1984a, chap. 7, this volume). Timing of 

this event is well constrained in the Wallowa terrane, 

where the youngest deformed sediments are of early 

Oxfordian age (Morrison, 1964) and the oldest posttectonic granitic intrusion is 160 Ma or late Oxfordian 

in age (Armstrong and others, 1977). This deformation 

may correspond to the collision of the Wallowa (or Blue 

Mountains) arc with the other four terranes. Subsequently, between 160 and 120 Ma (Armstrong and others, 1977), a number of granitic intrusions were 

emplaced in all four terranes, and locally these Jurassic and Cretaceous plutons crosscut contacts between 

terranes. No penetrative structures younger in age 

than Late Jurassic have been recognized in the Blue 

Mountains province of Oregon, but they are developed
46°

120° 119°

45°

25 50 KILOMETERS

rF

__________________118°______

EXPLANATION

Granite

Metasedimentary and metavolcanic rocks

Ophiolitic rocks

117°

ASHINGTONI rv^F--Si U

( OREGON IDAHO J

AREA OF MAP

Mitchell

OLDS

FERRY

TERRANE

o f 

o

D 

O

O 

I

a

o 

o

*T3

§o

s

O i i

O

FIGURE 1.1. Simplified geology of Blue Mountains province (modified from Walker, 

1977), showing four tectonostratigraphic terranes of Silberling and Jones (1984): 

Wallowa, Baker, Izee, and Olds Ferry terranes. Ophiolitic rocks include the Canyon Mountain Complex (CMC) and the (informal) Sparta complex (SC) of Phelps

(1979). Other basement rock types include Permian to Jurassic metasedimentary 

and metavolcanic rocks and predominantly Cretaceous granitoid intrusions. Dotted lines indicate terrane boundaries.
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON

TABLE 1.1. 40Ar/ 39Ar age determinations using hornblende separates from hornblende diorites of the Canyon Mountain Complex

and from hornblende gabbro of the Sparta complex

[Error estimates in Age column reflect analytical precision and are calculated following methods of Cox and Dalrymple (1967) and Dalrymple and others (1981). Monitor 

mineral was MMhb-1, which has an accepted age of 519.4 Ma. Samples were irradiated in U.S. Geological Survey's TRIGA (Training Research Isotope General Atomic) 

reactor in Denver, Colo. Corrections for Ca- and K-derived Ar isotopes were made using constants suggested by Dalrymple and others (1981); decay constants used in 

age calculations are from Steiger and Jager (1977). J, irradiation parameter; pet, percent; T, temperature]

40Ar/39Ar 37Ar/39Ar 36Ar/39Ar 39Ar 

(pet)

40Ar 

(pet)

39Ar 

(moles)

K/Ca 

(molar)

Age

(Ma)

Sample CMG-190A, hornblende diorite, Canyon Mountain Complex

[44°18'59" N., 118°51'44" W.; J=0.006618±0.000033; sample weight=0.9740 g]

800 352.7

950 77.17

1150 43.87

Fuse 109.8

rTV*4-ol CT£1G OO*O

Plateau age (950 and

64.00

41.59

60.93

61.22

1 150 °C steps) 

1.125

.1916

.0849

.3096

3.0

15.2

62.4

19.4

7.1

30.8

53.6

21.0

0.294

1.50

6".08

1.89

0.00777

.0122

.00818

.00814

276.5±12.8

263.5 ±1.8

260.6±1.3

255.7±2.6

260.6

261.2±1.6

Sample CMG-190B, hornblende diorite, Canyon Mountain Complex

[44°18'59" N., 118°51'44" W.; J=0.006349±0.000032; sample weight=1.0376 g]

950 116.2

1150 94.80

Fuse 315.0

Preferred age (1150 °C

41.92

71.97

69.80

0.3188

.2563

.9922

20.1

58.2

21.6

21.7

26.0

8.7

1.90

5.39

2.01

0.0121

.00688

.00710

267.7±1.5

261.9±3.3

287.6±21.4

268.7

261.9±3.3

Sample SP-96A, hornblende gabbro, Sparta complex

[44°48'35" N., 117°22'45" W.; J=0.006614±0.000033; sample weight=1.0039 g]

550 221.8

650 157.2

750 72.72

850 26.68

950 23.30

1050 22.62

1150 28.05

Fuse 115.3

Tn f- o 1 rr a G £1 cro

PI d t".A £in £1 cro ( Q ^\ O a n t\ 1 ICLbCdU. dgC \*y*J\J dllU

26.02

24.33

16.09

12.02

11.72

11.93

12.56

16.91

1 150 °C steps) 

0.7090

.4971

.1891

.02433

.01288

.01140

.02980

.3281

0.9

.1

.7

12.8

28.2

44.5

9.8

3.0

6.4

7.8

24.9

76.5

87.6

89.2

72.1

17.1

0.448

.0073

.249

6.68

14.7

23.1

5.09

1.57

0.0196

.0210

.0320

.0429

.0440

.0432

.0410

.0304

162.5±13.3

140.4±67.3

203.6±10.5

228.5±1.2

228.3±1.2

225.9+1.2

226.3±1.5

220.7±4.6

oo^i n ^b.U

OOT H-l-1 O ^/.Uxl.o

in Idaho along the North America-Blue Mountains suture between Orofino and McCall (Sutter and others, 

1984). However, post-Cretaceous 65° clockwise rotation 

of the Blue Mountains province (Wilson and Cox, 1980; 

Hillhouse and others, 1982) may be responsible for 

many faults and long-wavelength, low-amplitude folds 

that affect the Canyon Mountain Complex and overlying Eocene rocks (Thayer, 1956).

GEOCHRONOLOGY

A major question about the origin of the Oregon 

ophiolites concerns relationships between the associated mafic rocks (gabbros, basalts) and silicic rocks

(plagiogranites, keratophyres). Thus, absolute ages 

for these different rock types constitute an important 

constraint on their formation. Because reliable age 

determinations for ophiolites of the Blue Mountains 

province are sparse, we have determined three new 

40Ar/39Ar ages using hornblende separates from 

rocks of the Canyon Mountain Complex and Sparta 

complex (table 1.1). These are discussed within the 

context of other available data.

CANYON MOUNTAIN COMPLEX

Previous age determinations include concordant zircon U-Pb ages of 278 Ma for an undeformed intrusive
PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

gabbro from Canyon Mountain, and 278 Ma for a plagiogranite that cuts cumulate gabbro; U-rich zircon (with 

5-10 times more U than zircons from the plagiogranite) from a quartz diorite yielded slightly discordant UPb ages of about 268 Ma (Walker and Mattinson, 1980; 

Walker, 1981, 1986). Hornblende from a late-stage 

hornblende pegmatite yielded a 40Ar/39Ar plateau age 

of 262+14 Ma (Ave Lallemant and others, 1980).

New 4°Ar/39Ar hornblende ages are reported herein 

(table 1.1) for two relatively late-stage hornblende diorite intrusions. One of these samples yields a 40Ar/39Ar 

plateau age of 261.2+1.6 Ma (for 78 percent of the 

argon released). The other sample may be slightly disturbed as it did not yield a plateau age. However, more 

than 58 percent of the argon released (the 1,150°C 

step) yielded an age of 261.9+3.3 Ma; the total-gas age 

obtained is about 269 Ma. The ages for these samples 

are analytically indistinguishable.

From the available geochronological data, it can be 

argued that all intrusive phases of the Canyon 

Mountain Complex are older than 260 Ma but probably no older than about 280 Ma. The data appear to 

resolve a small but significant difference in age between the gabbroic and sheeted members of the complex. Considering the uncertainties of these ages, we 

cannot preclude that the magmas of the gabbroic and 

sheeted members of the complex were coeval and 

possibly cogenetic; at the least they may have formed 

in the same general tectonic environment.

SPARTA COMPLEX

Although some investigators previously included 

both the Sparta complex and the Canyon Mountain 

Complex in the Baker terrane, this interpretation is 

questionable because of the similarity between rocks 

in the Sparta complex and those in the Wallowa terrane. Also, the Sparta complex differs from the Canyon 

Mountain Complex in that it lacks an ultramafic member, owing to faulting (Prostka and Bateman, 1962). 

The structurally lowest member of the Sparta complex 

consists mainly of cumulate pyroxenites, which grade 

upward into cumulate gabbros; the highest gabbros 

are massive and contain primary amphibole and 

quartz (Phelps, 1978). Overlying and intruding the 

gabbros is a composite trondhjemite-quartz diorite 

pluton. It also intrudes volcanic rocks that were considered to be part of the Wallowa terrane by Vallier 

(1977). The apparently large proportion of felsic magmas supports the idea that the Sparta complex is part 

of the Wallowa terrane (Phelps, 1978, 1979; Phelps 

and Ave Lallemant, 1980). Phelps (1979) interpreted 

the gabbro-plagiogranite association to be cogenetic.

Published ages for the quartz diorite include (1) a 

K-Ar biotite age of 213+5 Ma (Marvin and Cole, 

1978), (2) hornblende 40Ar/39Ar plateau ages of 

223.0+3.3 and 223.2+5.5 Ma and a 40Ar/39Ar total-gas 

biotite age of 212.1 Ma (Ave Lallemant and others, 

1980), and (3) a concordant U-Pb zircon age of 223 

Ma (Walker, 1983, 1986). The reasonably close agreement among these ages suggests that the plagiogranite unit was emplaced about 223 Ma. Walker (1986) 

also reported a concordant U-Pb zircon age of 253 Ma 

for an albite granite unit within the plagiogranite 

member. These data indicate that silicic magmatism 

occurred over a relatively wide range in time.

Because no age information was available to ascertain the relationship of the cumulate gabbro member 

to the silicic units, we dated a hornblende separate 

from the massive hornblende-quartz gabbro unit; a 

40Ar/39Ar plateau age of 227.0+1.3 Ma was obtained 

(table 1.1). Thus, the formation of the gabbro was 

nearly coeval with that of most(?) of the plagiogranite. The sole Permian plagiogranite must be part of 

an older basal complex, possibly a part of the Permian segment of the Wallowa arc or perhaps an allochthonous fragment of older basement (see Walker, 

1982; T.L. Vallier, oral commun., 1987). It seems unlikely that the gabbro is as old as Permian. The exact 

relationship between the albite granite and the gabbro-plagiogranite assemblage remains unresolved.

GEOLOGY OF THE CANYON MOUNTAIN COMPLEX

PREVIOUS INVESTIGATIONS

The first detailed map of the Canyon Mountain 

Complex (Thayer, 1956) showed it as an almost rectangular east-trending body about 22 km in length and 8 

km in width (fig. 1.2). An ultramafic member in the 

northern part is overlain successively southward by an 

ultramafic-mafic layered zone, a gabbroic member, a 

sheeted sill-like (multiply intruded) zone, and a volcanic sequence. Plagiogranite intrusions occur locally 

throughout the complex but are concentrated near the 

contact between the gabbroic and sill-like members. 

Thayer (1963a) interpreted the Upper Permian and 

(or) Lower Triassic volcanic rocks in the southern part 

of the complex (keratophyre, quartz keratophyre, and 

spilite) to be hosts to the younger ultramafic and gabbroic intrusions of the complex. Apparently, emplacement of late-stage plagiogranite (quartz diorite to 

albite granite) magmas in the upper part of the complex caused metasomatism of the gabbroic and volcanic wallrocks. The ultramafic and gabbroic rocks were 

believed to have formed at upper-mantle and lower-
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON

crustal depths, respectively, in a large stratiform complex (analogous to the Stillwater intrusion) and to 

have been remobilized as a crystal mush and emplaced 

into higher structural levels (Thayer, 1963a, b; Thayer 

and Himmelberg, 1968; Thayer, 1969a). According to 

Thayer (1963b), layering in these rocks formed due to 

flow differentiation; primary structures were largely 

destroyed except for nodular structures in the podiform chromitite deposits (Thayer, 1969b).

On the basis of detailed structural and petrofabric 

studies, it has been demonstrated that the Canyon 

Mountain Complex consists of two distinct mantle diapirs that were emplaced at relatively high temperatures (Ave Lallemant, 1976; Misseri and Boudier, 

1985). To account for the disproportionately large volumes of plagiogranite and silicic volcanic rocks, Ave 

Lallemant (1976) suggested that these diapirs ascended either beneath a volcanic island arc or at 

least proximal to such an arc in a marginal basin. 

Thayer (1977a) briefly accepted this model but later 

in the same year stated that the ultramafic-mafic 

rocks of the complex formed in a stratiform intrusion 

beneath a midoceanic ridge, that these rocks were

118° 55' 118° SO- 45'

44C 

20'

5 KILOMETERS

EXPLANATION

Surficial deposits (Quaternary)

Canyon Mountain Complex (Permian) 

Consists of:

Keratophyre 

Plagiogranite 

Gabbro with minor pyroxenite

Volcanic rocks (Tertiary)

Harzburgite 

Serpentinite

FIGURE 1.2. Simplified geology of the Canyon Mountain Complex 

(modified from Thayer, 1977b). Heavy lines represent faults, 

dashed where covered. Area of map shown in figure 1.1.

deformed during ascent as a crystal mush, and that 

they were subsequently intruded by diabase and plagiogranite magmas and then covered by equivalent 

volcanic rocks (Thayer, 1977b).

On the basis of petrologic and geochemical studies 

of the ultramafic and gabbroic rocks, Himmelberg and 

Loney (1980) concluded that (1) the major harzburgitic component of the ultramafic member represents residual upper-mantle material that had been subjected 

to a previous partial-melting event unrelated to the 

formation of the remainder of the ophiolite, (2) the 

podiform chromitite found only in dunite bodies represents cumulate material in the basal harzburgite unit, 

(3) the gabbroic member formed by fractional crystallization of basaltic magma that had been emplaced 

into depleted mantle wallrocks, and (4) the large volume of silicic rocks implies an island-arc setting for 

the formation of the Canyon Mountain Complex.

Detailed geologic mapping of the sheeted member 

(Gerlach, 1980) confirmed the abundance of silicic 

magmas (fig. 1.3) and provided information on relative timing for the emplacement of its respective 

magma types. On the basis of petrologic and geochemical studies, Gerlach, Ave Lallemant, and Leeman (1981) and Gerlach, Leeman, and Ave Lallemant 

(1981) interpreted the plagiogranites as the products 

of partial melting of hydrated mafic rocks at relatively shallow depths within an island arc; emplacement 

of basaltic magmas was invoked to provide the necessary heat.

Further detailed mapping and petrologic studies of 

the ultramafic and gabbroic rocks led E.M. Bishop 

(Mullen, 1983a, b; Bishop, chap. 5, this volume) to 

conclude that (1) several generations of gabbros (some 

of which are penetratively deformed) intruded the depleted harzburgite tectonite member, and (2) the most 

primitive gabbros were tholeiitic whereas the more 

evolved gabbros were calc-alkalic. Considering the regional geology, she proposed that the Canyon Mountain Complex formed in a fore-arc setting. On the 

basis of regional tectonic constraints, Ave Lallemant 

(1984b) suggested that it formed in an intra-arc extensional basin produced by en echelon strike-slip 

faulting related to oblique subduction.

Despite uncertainties concerning details of the 

original setting for the Canyon Mountain Complex, 

the overall evidence indicates that the ultramafic 

rocks (plus some of the mafic rocks in the eastern 

part of the complex) ascended diapirically, that early 

mafic magmas were intruded syntectonically or pretectonically, and that post-tectonic mafic and felsic 

intrusions cut the ultramafic rocks and led to formation of high-level magma chambers, dike and sill 

complexes, and possibly lava flows as well.
00

188° 55' 118° 52'30"

1 + X * XiLrlilv X * :^^^ii^lll|pllliiF;PPf f*Wf PHi!ii!!^-< > ;-, - s-,s r ,-^> i^illi^^-^ ^*.

;Ifj ;^ ?£££££*$£ i; i >^ ;I{i ;^r ;£ ;£ ;^e >;f;

-Tv/Xi" r ^1" r ^"' r ,V r M" r ^" ^"'r i'""^ ^V*r ^V*r I'/'r ^* r ^/r X^r ^**^T J J/T?!" r ^ii^%j^-7 ?

\* r \" r ** ~r\" r \" r~'^£^^S^T^ ^1" r \" 7*"'^ T ^S~t~*~S "r-i," r ?/ r -7 .^-*7 ^^.--L^ r ^'t" r ^i" r \" r '"'t" t ^'i." r ^ L" r ^t" r \" r ^t" r \* r ^<./*^~ii.'"?'"7 <.A \^-^:^:-;^^- ^-'^'''-^ " J 7 " J 7 ".^.^.^^J A r

Explanation

Alluvium and colluvium 

(Quaternary)

Canyon Mountain Complex (Permian) - Consists of: 

Diabase, thick sills 

Basaltic sills, dikes 

Keratophyre and minor spilite 

Hornblende diorite

Clamo Formation 

(Tertiary)

Plagiogranite 

Gabbro 

Pyroxenite 

Faults

44° 

17' 

30"

O

r 

o

H 

O 

H

i

o

§

3

H 

H 

H

I 

S

O

w f 

a

H

S 

O

01

» 

H

2

i

FIGURE 1.3. Geology of sheeted member and keratophyre unit of the Canyon Mountain Complex showing major dikes and sill-like intrusive lenses (modified from Gerlach,

1980). Contacts, faults, and sills or dikes are dashed where covered. Area of map shown in figure 1.2.
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON

GEOLOGIC DESCRIPTIONS OF UNITS

In general, the lithologic units (or members) of the 

Canyon Mountain Complex (fig. 1.4) are characteristic of many ophiolites; they are described individually 

later. E.M. Bishop (Mullen, 1983a, b; Bishop, chap. 5, 

this volume) presents additional details of the geology and mineralogy of the ultramafic and mafic members of the complex. Contacts between lithologic 

units trend approximately east-southeast to westnorthwest and are near vertical. On the east, north, 

and west the ophiolite is fault bounded; on the south 

it is overlain by Eocene conglomerates and volcanic 

rocks (Thayer, 1956).

ULTRAMAFIC MEMBER

The ultramafic member consists mostly of harzburgite with rare Iherzolite. In a few areas, the peridotites

show faint layering due to orthopyroxene concentrations along foliation planes. Large irregular dunite 

bodies occur throughout, and these commonly include 

layers of nodular and disseminated chromitite. Pretectonic or syntectonic (diapiric) gabbro, gabbro pegmatite, and pyroxenite dikes crosscut the peridotites. 

Clearly post-tectonic dunite occurs in veinlike layers 

and as dikes in the peridotites. Interstitial plagioclase 

occurs in rocks toward the top of the ultramafic member and locally near dikes of gabbro.

All rock types in the ultramafic member are penetratively deformed and have well-developed foliation 

and lineation (Ave Lallemant, 1976). The chromitites 

have relict cumulate textures (Thayer, 1969b); Himmelberg and Loney (1980) suggested that the chromitites and the enclosing dunite bodies together 

represent the basal part of the overlying gabbroic 

member, which was folded into the tectonite harzburgite. The dunite dikes and veins are not folded,

Point Sal

California

UJ

O

X

Trinity

California

top not 

exposed

Del Puerto Fidalgo Semail

California Washington Oman

Karmoy

Norway

not exposed

Canyon Mountain 

Complex

Oregon

UJ

. o

X

Bay of 

Islands

Newfoundland

Papua

New Guinea

P\7 V V W T

VV t X]

11 'III

% nn

% nn

&li

l^

itf

r°

-i

(0 

-3 (/) 

UJ

Z

-40

X

1-

-5

-6

Explanation

|::::tnl Keratophyre -

|vvvvvvi| Pillow lava and spilite ^\^\

iKMijrsi Dikes and sills of diabase 

H1^ -! or basalt ___.=^-

|yX%| Plagiogranite

[^v^ffjl Isotropic gabbro

Layered gabbro

Unconformity

Fault Arrows 

indicate relative 

movement

FIGURE 1.4. Comparative stratigraphic sections of upper parts of selected ophiolites worldwide (from Gerlach, Ave Lallemant, and

Leeman, 1981). TV, Tertiary volcanic rocks.
10 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

but have a tectonite fabric; they appear to have 

formed by metasomatic transformation of harzburgite 

(Dungan and Ave Lallemant, 1977). Subsequent to 

dunite formation, gabbro, pyroxenite, rare diabase, 

and very rare plagiogranite dikes and sills intruded 

the ultramafic rocks. Most of the peridotites, particularly those exposed near the bounding fault along the 

north margin of the complex, locally are serpentinized and sheared to various extents.

GABBROIC MEMBER

Unlike the petrographically and structurally homogeneous harzburgite tectonites of the ultramafic 

member, the gabbroic member is very heterogeneous. 

Most primary igneous textures in the eastern part of 

the complex have been obliterated by intense penetrative deformation, whereas in the west primary 

textures are preserved.

Cyclical igneous layering in the western massif is 

consistent with crystal fractionation by gravitational 

settling within a periodically refilled basaltic-magma 

chamber. The lowermost units are dunites and 

chromitites, which are overlain by olivine pyroxenites. Relict cumulate textures are ubiquitous, and 

rocks are not penetratively deformed. The olivine pyroxenites in turn are overlain by layered gabbroic 

rocks that locally contain lenses of wehrlite or olivine 

pyroxenite. The gabbroic rocks comprise various proportions of plagioclase, clinopyroxene, olivine, and 

rare orthopyroxene; corresponding rock types include 

gabbro, troctolite, norite, and anorthosite. Depositional structures such as slumps and trough banding are 

rare. Diabase dikes occur throughout this unit but 

are more numerous southward toward the contact 

with the sheeted member. At Canyon Mountain, they 

form a northeast-trending swarm and locally form 

coarse-grained stocks (the ophitic gabbros of Ave Lallemant, 1976).

In contrast, the gabbroic member of the eastern 

massif represents a number of small independent intrusive bodies. The main rock type is coarse-grained 

gabbro containing relict cumulate structures. This 

gabbro in turn contains inclusions of fine-grained 

gabbro and is cut by gabbro pegmatite, olivine pyroxenite, and several other gabbroic intrusions (Ave Lallemant, 1976; Mullen, 1983a, b; Bishop, chap. 5, this 

volume). Penetrative deformation overprinted all of 

these rocks as they were partly folded into the tectonite harzburgites of the ultramafic member. This 

deformation was followed by intrusion of isotropic 

gabbro, pyroxenite, diabase, hornblende pegmatite, 

and plagiogranite bodies.

SHEETED (SILL-LIKE) MEMBER

The sheeted member consists of east-west-striking, 

steeply dipping layers of gabbro, keratophyre, basalt, 

basaltic andesite, diabase, and plagiogranite. The 

contact with the underlying gabbroic member is gradational. Gabbro, diabase, and plagiogranite are 

mostly concentrated in the structurally lower northern part, and keratophyre, basalt, and basaltic andesite are more abundant at structurally higher levels 

to the south. The largest plagiogranite bodies occur 

near the contact between the sheeted and gabbroic 

members. Field relations indicate that gabbros, 

keratophyres, and basalts are the oldest rocks in the 

sheeted member and that they were intruded by diabase and plagiogranite. Rocks of the plagiogranite 

suite were emplaced in the general sequence hornblende diorite, quartz diorite, and trondhjemite; 

these rocks were intruded by later basalts and diabases. Rare banded keratophyre lenses apparently 

represent one of the latest magmatic phases. Relative 

timing and duration of intrusive events as deduced 

from crosscutting relations in the sheeted member 

are shown in figure 1.5.

Because neither pillow basalts nor cherts were 

found in the sheeted member, Thayer (1977b) interpreted it to be a dike complex. However, some keratophyres display fine lamination, which Ave Lallemant 

(1976) previously interpreted as pyroclastic layering. 

Also, the presence of vesicles in some basalt and basaltic andesite units suggest that they are lava flows. 

If these rocks are indeed extrusive, then conformable 

attitudes between them and the diabase and plagioDecreasing relative ageGabbro

Hornblende diorite 

Quartz diorite 

Trondhjemite 

Diabase and basalt 

Keratophyre

FIGURE 1.5. Relative timing and duration of magmatic pulses that 

produced sheeted member in upper part of the Canyon Mountain Complex. Evidence is based on field relations including intrusive contacts, crosscutting dikes and lenses, and presence of 

xenoliths and intrusion breccias. Segments are dashed where 

uncertain.
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON 11

granite lenses imply that the sheeted member is a 

sill rather than a dike complex.

STRUCTURE

The penetrative-deformation structure in the ultramafic member and in the eastern part of the gabbroic 

member is manifested by well-developed foliation, 

lineation, and strong preferred orientation of all constituent minerals (Ave Lallemant, 1976). As deduced 

from the preferred orientations, the major principal 

extension direction (X) is vertical. This inference and 

the overall structural style led to the conclusion that 

the Canyon Mountain Complex was emplaced as two 

separate massifs by diapiric flow in the upper mantle 

(Ave Lallemant, 1976). In the western massif, the extension axis (X) is vertical in the north, but rotated 

and even overturned to the south near the contact 

with the gabbroic member, where the cumulate layering is vertical to slightly overturned. These data indicate that the ophiolite is folded into a large 

asymmetric east-west-trending anticline (Ave Lallemant, 1976; Thayer and others, 1981). The contact 

between the gabbroic and sheeted members is steep, 

but aeromagnetic data (Thayer and others, 1981) 

suggest that at depth the contact rotates toward a 

more southerly dip of about 45° and that the ophiolite extends southward for some distance beneath 

surface cover. If this anticline is unfolded, the cumulate layering in the gabbroic member, layering in the 

keratophyre units, and the intrusive layers of diabase and plagiogranite all become subhorizontal. Rotation of the sheeted member is poorly constrained in 

time; it may have occurred during Late Triassic subduction, during Triassic subduction, during Jurassic 

collision, or possibly in Eocene or later time. Because 

Eocene conglomerates overlying the southern part of 

the ophiolite also have a steep southward dip (average is about 45°, but locally as high as 70°), folding 

of the Canyon Mountain Complex cannot be dated 

closely (Thayer, 1956). Although we consider the 

sheeted member to be a stack of rotated sills and 

minor flows (Gerlach, Ave Lallemant, and Leeman, 

1981), others (for example, Thayer, 1963a; Misseri 

and Boudier, 1985) favor emplacement of this member as steeply dipping dikes because elsewhere in the 

complex most dikes are steeply dipping.

PETROGRAPHY AND MINERALOGY

The peridotite of the ultramafic member is generally harzburgite of uniform composition consisting of

olivine (Fos8-9i) with 10 to 25 percent orthopyroxene 

(Engo_92), 3 to 5 percent clinopyroxene (0.8-1.0 percent Cr2O3), and 1 to 2 percent chrome spinel (Himmelberg and Loney, 1980). Pargasitic amphiboles 

commonly border or enclose pyroxene grains. Only 

rarely does the peridotite contain enough clinopyroxene (as much as 10 percent) to make it a true Iherzolite. The texture is xenoblastic-granular; most grains 

are slightly flattened and elongate parallel to foliation and lineation, respectively. Olivines and orthopyroxenes in the dunites in which chromitite deposits 

occur are slightly enriched in Mg (Fo92_96, 

and rare clinopyroxenes contain 1 to 2 percent 

(Mullen, 1983a, b); textures in the dunites are similar to those in the harzburgites, except for the chromitite layers, which contain subhedral to euhedral 

chromite grains. Dunite veins and dikes have textures similar to those in the harzburgite, although 

the olivines are coarser. All peridotites and dunites 

are somewhat serpentinized.

Ultramafic and gabbroic rocks in the eastern part of 

the gabbroic member are mostly tectonites that display the same penetrative-deformation structures as 

found in harzburgites of the ultramafic member. Although these gabbroic rocks have relict primary 

banding, their overall texture is xenoblastic-granular; 

only rarely are cumulate textures preserved (Ave Lallemant, 1976). Himmelberg and Loney (1980) found 

evidence of a slight Fe-enrichment trend in going 

from dunites to gabbros, on the basis of compositions 

of olivine (Fogs in dunites to Fogs in gabbros), orthopyroxenes (Engs to Envs), and clinopyroxenes 

(Ca49Mg4gFe2 to Ca46Mg42Fei2); plagioclases in the 

same rocks range in composition from Angg in dunites 

to Anvo in gabbros. Pearcy (1987) observed that such 

compositional trends are not necessarily regular.

The mafic cumulates of the western massif are 

folded but not penetratively deformed. Cumulate textures are generally preserved except where the rocks 

are altered near diabase and plagiogranite intrusions 

and near brecciated zones. Mineral compositions for 

these rocks are similar to those given for the tectonite metacumulates in the eastern part of the complex 

(Himmelberg and Loney, 1980).

Both the ultramafic and gabbroic members contain 

gabbro, gabbro pegmatite, and pyroxenite dikes that 

were emplaced prior to penetrative deformation. Their 

textures are similar to those described above for compositionally equivalent rocks. Post-tectonic dikes and 

veins have igneous textures. For example, in some 

gabbro dikes, euhedral plagioclase crystals are aligned 

parallel to the dike walls; this texture may reflect flow 

alignment of crystals parallel to conduit walls. Pyroxenite, gabbro pegmatite, and hornblende pegmatite
12 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

dikes are very coarse grained and have allotriomorphic to hypidiomorphic textures. Diabase dikes display 

subophitic textures but vary in grain size.

The gabbroic rocks have undergone subsolidus alteration of at least two different types. Proximal to 

intrusions of hornblende pegmatite or plagiogranite 

or near mylonite zones, clinopyroxenes are partly to 

completely amphibolitized. In contrast, gabbroic 

rocks in permeable zones are commonly altered to 

greenschist-facies assemblages, including hydrogrossular, chlorite, prehnite, calcite, and serpentine.

The sheeted member consists of screens of cumulate gabbros in the lower (more northerly) part and 

keratophyres and basaltic rocks in the upper (more 

southerly) part; these rocks are intruded by many 

sills, dikes, and stocks of plagiogranite, diabase, and 

minor basalt. The keratophyres are porphyritic, bearing phenocrysts of quartz and plagioclase in a finegrained felty, spherulitic, or trachytic groundmass; 

quartz phenocrysts tend to be embayed but commonly display high-temperature bipyramidal morphology. 

Basalts are fine grained with intergranular to intersertal textures, and consist of plagioclase, clinopyroxene, and opaque minerals. A few vesicular basaltic 

andesites contain elongate plagioclase phenocrysts in 

a groundmass containing abundant microlites of plagioclase and oxides. Diabases contain diopsidic to 

subcalcic augite, plagioclase, magnetite, and ilmenite 

and have characteristic ophitic texture. Plagiogranites consist of plagioclase, quartz, hornblende, and 

small amounts of biotite and opaque minerals; their 

textures include hypidiomorphic-granular, granophyric, and porphyritic. Plagioclase phenocrysts display 

oscillatory zoning in some samples. Modal green 

hornblende varies from less than 10 percent in trondhjemites to more than 50 percent in some diorites; it 

occurs as an interstitial phase, as euhedra, and rarely as poikilitic grains with inclusions of plagioclase 

and quartz. Severe local alteration of rocks in the 

sheeted member produced secondary albite, chlorite, 

epidote, and calcite.

Electron-microprobe analyses of fresh clinopyroxenes in selected mafic rocks mainly from the sheeted 

member (fig. 1.6) display a marked difference in composition from those of clinopyroxenes in the cumulate 

and tectonite members. These data suggest that the 

basaltic magmas that intruded different parts of the 

complex do not represent cogenetic liquids.

PALEOTHERMOMETRY AND PALEOBAROMETRY

Information on pressures and temperatures of formation for rocks of the Canyon Mountain Complex is

sparse. Himmelberg and Loney (1980) estimated temperatures of about 880 to 975 °C for the harzburgite 

tectonites and cumulate pyroxenites using analyses 

of coexisting orthopyroxenes and clinopyroxenes and 

the geothermometer of Wells (1977). These temperature estimates are unreasonably low for near-liquidus conditions, and it is probable that the mineral 

compositions have readjusted to subsolidus conditions due to low cooling rates or to recrystallization.

Mercier (1980; oral commun., 1982) applied his 

own geothermometer-barometer to seven harzburgites and obtained apparent equilibration temperatures of 1,000 to 1,100 °C and rather high pressures 

of 15 to 25 kb for samples from the deformed eastern 

part of the complex, and 900 to 1,000 °C and 8 to 15 

kb for the western part. Because Mercier's method is 

based largely on empirical relationships, the calculated temperatures and pressures may not be correct in 

an absolute sense; however, they are internally consistent and suggest at least relative pressure-temperature conditions for recrystallization of the two 

structurally distinct parts of the complex.

GEOCHEMISTRY AND PETROLOGY

Reconnaissance geochemical investigations were undertaken to characterize all common rock types of the

Explanation

This study: Himmelberg and 

Loney (1980):

m Gabbro d Metacumulates

^ rx. i * ^ i x 

 Diabase A Cumulates

O Basalt $ Harzburgite

FIGURE 1.6. Compositions of clinopyroxenes from mafic rocks of 

the Canyon Mountain Complex projected onto the pyroxene 

quadrilateral. Data from Himmelberg and Loney (1980) are representative of clinopyroxenes from the tectonite and cumulate 

members (see also Mullen, 1983a; Bishop, chap. 5, this volume). 

Note distinction between pyroxenes from cumulate rocks and 

those from gabbro, diabase, and basaltic rocks that have nearliquidus or "magmatic" compositions. Di, diopside; En, enstatite; 

Fs, ferrosilite; Hd, hedenbergite.
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON 13

Canyon Mountain Complex. Representative samples 

have been analyzed for major and (or) trace elements 

using a variety of methods (see section "Analytical 

Methods," below). These samples include the plagiogranite and keratophyre suites previously described in 

detail by Gerlach, Leeman, and Ave Lallemant (1981). 

Of primary concern here are the volcanic and hypabyssal rocks that most closely represent magmatic liquids; relatively little emphasis has been placed on the 

cumulate gabbros or the various ultramafic tectonites, 

the chemical and (or) mineralogical compositions of 

which have been described by others (Thayer and 

Himmelberg, 1968; Himmelberg and Loney, 1980; 

Mullen, 1983a, b; Bishop, chap. 5, this volume). Analytical results as well as brief descriptions of the analyzed samples are given in tables 1.2 through 1.6. 

Specific sample locality information is available from 

the authors.

Most of the rocks selected for geochemical study are 

relatively fresh and have retained their original mineral compositions and textures. However, because many 

of our samples exhibit at least some alteration or lowgrade metamorphic recrystallization, it is necessary to 

carefully assess those geochemical parameters that 

supposedly reflect primary petrogenetic processes. The 

problems of evaluating the effects of element mobility 

in altered rocks and of deciphering tectonic setting and 

petrogenetic history of ophiolitic rocks have been summarized by Pearce and Cann (1973), Winchester and 

Floyd (1977), Pearce (1982, 1983), and others. Most 

attention has focused on geochemical comparisons 

with modern basaltic associations such as midoceanridge basalt (MORE), oceanic-island or within-plate 

basalt (WPB), and island-arc basalt (IAB) suites, which 

have diagnostic geochemical features. To facilitate 

such a comparison, it is first necessary to identify those 

samples that retain or closely approach their original 

magmatic compositions.

ANALYTICAL METHODS

Representative samples were analyzed in two 

batches as follows: Group I (CM-prefix samples) and 

Group II (CMG- and DC-prefix samples). Major-element analyses were determined for Group II samples 

by combined electron microprobe analysis (EMPA) of 

fused glasses and atomic absorption spectrophotometry (AAS) (Gerlach, Leeman, and Ave Lallemant, 

1981), and for Group I samples by X-ray fluorescence 

(XRF) analysis of flux-fused samples (see Jacques 

and Chappell, 1980). Trace-element analyses were 

obtained on Group II samples by instrumental neutron activation analysis (INAA) at Oregon State University and at the National Aeronautic and Space 

Administration (NASA) Johnson Space Center 

(Gerlach, Leeman, and Ave Lallemant, 1981), and by 

XRF analysis of pressed powder disks at the University of Massachusetts (Frey and others, 1984). Group 

I samples were analyzed for a few trace elements by 

the latter method at the Australian National University. For selected Group II samples, Sr, Ba, Sc, V, Cr, 

Ni, Zn, and Zr were determined by inductively coupled plasma (ICP) spectroscopy at Rice University.

Analytical uncertainties are generally within 1 to 5 

percent of the amount present for the major elements 

and 5 to 10 percent for most trace elements. Uncertainties may be somewhat larger for elements with 

low abundance or low sensitivity, such as Th, Ta, Ba, 

and Nb; the results for these elements (and especially for Nb) are reported to indicate the extremely low 

abundances in most samples. Dashes in tables indicate that abundances either were not determined or 

were below detection limits.

Analysts included the following: XRF, B.C. Gerlach 

at the University of Massachusetts and Bruce Chappell at the Australian National University; INAA, 

W.P. Leeman at Oregon State University and B.C. 

Gerlach at NASA; AAS, B.C. Gerlach at NASA; 

EMPA, B.C. Gerlach at Rice University; ICP, W.P. 

Leeman at Rice University.

Bata are presented in separate tables (tables 1.2 - 

1.6) according to rock type. However, we have arbitrarily included gabbroic samples with greater than 

0.7 weight percent TiO£ in table 1.4 with diabase, dolerite, and basaltic rocks because all of these rocks are 

believed to approximate original magmatic liquids in 

composition. Throughout this paper, all samples in 

table 1.4 have been plotted together as "magmatic" 

samples. Inspection of figures 1.7 and 1.8, for example, shows that the compositions of these samples do 

indeed resemble basaltic compositions and define coherent geochemical trends. All major-element data 

shown in diagrams are based on analyses recalculated to 100 percent totals on an anhydrous basis.

GENERAL GEOCHEMICAL CHARACTERISTICS

Major-element characteristics of the rocks of the 

Canyon Mountain Complex generally resemble those 

of most ophiolites. The main lithologic groups are 

easily distinguished in SiO£ variation diagrams (fig. 

1.7). The ultramafic rocks define a restricted grouping (low SiO£, alkalis, TiO£, P^O^, and A^Oa, and 

high MgO) consistent with their dominantly olivine 

and pyroxene mineral assemblages. With exception of 

the cumulate gabbros, which have relatively high
14 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 1.2. Analyses of major-element oxides and selected trace elements in ultramafic rocks (peridotite, harzburgite, pyroxenite) of the

Canyon Mountain Complex

Sample -

Rock type - 

CM65

Serpentinized 

peridotite

CM37

Tectonite 

harzburgite

CM38

Tectonite 

harzburgite

CM60

Pyroxenite

CM54

Pyroxenite

CM64

Pyroxenite

CM61

Cumulate olivine 

pyroxenite

CM62

Plagioclasebearing pyroxenite

CM63

Olivine-plagioclase 

pyroxenite

Major-element oxides, in weight percent

SiO2 - 

Tin

ALA--

T?«. Fe2f\O3 

TJI,-/-V

MnO 

MgO

CaO 

Na2O 

V f\

P2Uf)5 

H n+

H

/-Vpn

Total 

37.43 

0.09

5.70 

A qq

4.85

n IR

qo 70

q 1 1

0.04

0.08

0.01

11.09

0.30

0

1 R

100.09

42.71 

0.01

1.19

O.OD

0 1 q

42.25

1.10

0.02

0.02

0.01

4.44

n qi

n q9

100.17

43.78 

0.01

0.80 

1.45

0.14

49 Q7

\J* 1 \J

0.01

0.01

0.01

0.01

n qi

3 01

99.72

47.32 

0.04

1.19 

9 qs.

O QC

0.10

OC K.C

1 q OR

0.08

0.08

0.01

R c\n

0.30

0

1 Q

99.72

51.27

0

1 £

3.33 

1 9q

4.47

U. J-U

17.77

0.17

0.03

0.01

1 1Q

0.20

n 91

99.63

49.77 

n 1*;

2.82 

1 1Q

A *Q

0.14

20.51

18.49

0.14

0.03

0.01

1 ^9

0.17

0.14

99.61

42.30 

0.11

2.73 

4 98

5

£*£*

0

1 a

CiC QK

Q 4.9

0.10

0.03

0.01

7.41

0.30

n 9R

99.64

41.24 

0.08

6.23 

9 qn

01^

oc ni

Q 79

0

1 £

0.03

0.01

fi QR

n qo

n 9q

99.67

41.33 

0.07

6.21

9 T^

c 77

0

1C

OK 7fi

Q 49

n IQ

0.03

0.01

7.41

O Qfl

0.20

99.63

Trace elements, in parts per million

V - . CO

4,325

41 

2,960

38 

3,300

70 

2,985

262 

2,255

228 

4,360

125 

1,495

93 

1,895

84 

1,785

MgO and Al^Os, the remaining rocks define what appear to be coherent trends more or less typical of igneous liquid lines of descent. A plot of A^Os versus 

SiC>2 (fig. 1.7) is particularly effective in discriminating those rocks with compositions appropriate for 

magmatic liquids (that is, non-cumulate gabbro, diabase, basaltic rocks, and the more siliceous plagiogranite and keratophyre suites) from rocks having 

high proportions of cumulate minerals (cumulate 

gabbro and pyroxenite) and rocks from which melt 

may have been extracted (harzburgitic peridotites). 

Solely on the basis of petrography or field relations, 

it is often difficult to differentiate coarse-grained, 

dominantly cumulate gabbroic rocks from gabbros 

that represent magmatic liquids. For example, many 

dolerites and coarse-grained diabases have higher 

contents of incompatible elements (such as K, Ti, P) 

than expected for purely cumulate gabbros. Also, because the cumulate rocks represent concentrates of 

plagioclase, olivine, and pyroxenes, they are relatively enriched in Al, Mg, and Ca. A plot of the relatively 

immobile components A^Os versus TiC>2 (fig. 1.8) 

clearly discriminates a group of mafic rocks that resemble modern oceanic basalts (see Melson and others, 1976; Bryan and others, 1981) from the other 

lithologic groups. In this chapter we rely mainly on 

these presumably magmatic or near-magmatic samples to understand petrogenetic processes.

The basaltic to felsic rocks of the intrusive complex 

and the extrusive keratophyres are characterized by a 

relatively strong covariation between SiC>2 and the oxides CaO, MgO, FeO* (total Fe as FeO), Na2O, and 

TiO2 (fig. 1.7) not unlike that expected for magmatic 

liquid lines of descent. However, considerable scatter 

for contents of K^O and possibly ?2Os may reflect alteration. Alteration effects are more clearly seen in 

plots of various elements against TiO2 (fig. 1.8), which 

tends to be one of the least mobile constituents of 

igneous rocks under conditions of alteration (Pearce, 

1983).

The possible effects of crystal-liquid fractionation 

must be considered in evaluating overall compositional variations for both the felsic and mafic rocks. The 

mafic rocks alone display strong positive correlations 

between TiO2 and P2O5 , Ce, and Hf (fig. 1.8), as well 

as between TiO2 and Zr (see fig. 1.14). Weaker posi-
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON 15

TABLE 1.3. Analyses of major-element oxides and selected trace elements in gabbroic rocks (with less than 0.7 weight percent

Canyon Mountain Complex

[---, not determined or below detection limit]

of the

Rock type 

CM55

Anorthositic 

gabbro

CM56

Foliated 

anorthositic 

gabbro

CM57

Foliated 

gabbroic 

anorthosite

CM58

Foliated 

gabbroic 

anorthosite

CM28

Cumulate 

gabbro

CM34

Cumulate 

anorthositic 

gabbro

CM35

Cumulate 

anorthositic 

gabbro

OC277

Olivineplagioclase 

cumulate 

(troctolite)

Major-element oxides, in weight percent

Si02 

Tin

AI n.

Fe2 03 ~

FeO _____

MnO ___

MgO

CaO ------ -

Na2 0 

K c\

P

r\ 2O5 

H

rvH

r\-

co

Total - 

4.C K.A

0.09

90 ^7

r\ ci

9 M

0.06

11.90

1 °. OS

1 89

0.06

0.01

2 87

n IQ

0.30

QQ 61

44 31

0.09

14.86

C nQ

0.11

i a nn

11.60

1.47

0.04

0.01

*r.O J019

0 IQ

QQ fiO

44 07

0.05

OC f\f\

O cc

1 8Q

0.04

16.47

1 99

0.06

0.01

2

rjc

0.30

0.14

QQ QQ

40 no

0.08

99 91

n cc

o no

0.07

O QK

15.08

1.17

0.09

0.01

0 QO

0.30

0.15

QQ ^Q

47.05

0

1 Q

18.00

0.96

q ci

0.10

10.06

17.11

0.90

0.06

0.01

9 O3

0.17

0.01

100.15

45.19

0

1 £

oo na

1 OQ

2

/>o

0.07

a nf\

i a QO

1.00

0.05

0.01

9 O3

0 21

0.06

QQ 7Q

45.82

0 97

90 fiQ

9 Oft

o. no

0.09

7 SO

1^ 4Q

1.14

0.03

0.01

9 07

0 9^

0.01

QQ 7°.

47.50

0.01

24.40

c 70

0.08

Q ^4

11.90

0.02

100.61

Trace elements, in parts per million

Sc -- 

V 47 79 50 87 146 114

Cr 1,130 1,085 496 357 429 354

CoNi 

Th ______

Sm ~

TTn

Tb ____

Yb 

Lu 

9 1

150

Qd R

47.4

CO

0.03 

0.11 

0.05 

0.17

0.012

O

nCO

0.007

tive correlations between TiO2 and Th, Ba, and possibly K2O in the mafic rocks (fig. 1.8) may reflect 

alteration of some samples, because normally these 

incompatible elements are coherently enriched in 

evolving mafic liquids. Negative correlations between 

TiO2 and Mg number (fig. 1.8) and between TiO2 and 

abundances of Sc, Co, Cr, and Ni (table 1.4) and the 

positive correlations just mentioned are consistent 

with differentiation involving removal of Fe-Ti-oxidepoor mineral assemblages from the evolving mafic 

liquids. Under such conditions Ti behaves as an essentially incompatible element, and its abundance

correlates with that of other incompatible trace elements (see Cann, 1970; Saunders and others, 1980).

On some of the TiO2 variation diagrams (see fig. 

1.8), the plagiogranites and keratophyres show trends 

that preclude their derivation by simple fractionation 

from mafic precursor magmas. In the felsic rocks, 

TiO2 and Mg number decrease together along with 

other transition metals (tables 1,5 and 1.6) in accord 

with fractionation of Fe-Ti oxides and ferromagnesian 

minerals. Fractionation of such minerals is consistent 

with the small degree of Fe enrichment (fig. 1.9) and 

the strong alkali-enrichment trend (fig. 1.10) observed
16 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 1.4. Analyses of major-element oxides and selected trace elements in mafic rocks (diabase, dolerite,

[ , not determined or

Sample »- 

Rock type -

CM26

Olivine 

gabbro

CM27

Olivine 

gabbro

CM19

Hornblendeplagioclase 

dolerite

CM20

Gabbro

CM21

Cumulate 

gabbro

CM39

Gabbro

CM40

Gabbro

CMG124A

Porphyritic 

plagioclaseclinopyroxene 

diabase

CMG49

Quartz 

diabase

OC244 OC256

Diabase Diabase

OC259

Diabase

Major-element oxides,

SiO2 

TiO2 

ALA --

Fe2O3 -

FeO 

MnO -----

MgO ~

PnO

Na2O --

P

/~i 2^)5 ------

H20-~~

CO2 

Total -

48.56 

1.29

15.19

2.22

7.81

0.19

8.04 

11.16 

2.45 

0.15 

0.06 

2.43

0.17

0.08

99.80

49.52 

1.22

15.52

2.03

7.27

0.18

7.76 

10.81 

3.07 

0.13 

0.09 

1.96

0.24

0.01

99.81

49.75 

1.89

14.87

0.75

10.53

0.20

5.72 

10.05 

3.91 

0.20 

0.16 

1.92

0.08

0.07

100.10

49.54 

2.14

14.62

0.63

11.54

0.22

4.85 

9.10 

4.24 

0.19 

0.20 

1.78

0.09

0.53

99.67

50.83 

2.34

14.59

1.16

10.93

0.23

4.67 

7.36 

4.89 

0.29 

0.21 

2.31

0.16

0.04

100.01

49.99 

1.20

15.17

2.58

6.94

0.18

6.58 

9.03 

4.21 

0.19 

0.07 

3.12

0.28

0.09

99.63

47.29 

2.12

15.13

3.64

7.92

0.22

6.94 

8.13 

3.98 

0.40 

0.17 

3.34

0.35

0.29

99.92

50.80 

1.44

15.10

10.20

0.21

7.78 

10.10 

2.70 

0.22

...

...

98.55

56.30 

1.73

14.90

9.59

0.24

5.79 

6.09 

3.79 

0.29

...

...

98.72

53.90 52.20 

0.84 1.61

16.70 15.50

8.68 9.57

0.15 0.19

6.25 8.81 

9.74 7.21 

3.49 3.98 

0.03 0.13

 .

...

99.78 99.20

54.30 

1.05

15.10

10.20

0.16

6.93 

7.99 

4.49 

0.32

...

...

100.54

Trace elements,

Sr 

Ba

Sc ------ 

y ____

CrCo

Ni 

Zn - -

Ta - 

Nb -

TIf

Th

ZrLa -

Ce ~ 

Nd -

Sm -

EU - 

Tb -

YhJ. U ~~

Lu 

15

329 

272

::

_

7

_

30

278 

262

9

30

370 

86

...

18

35

396 

38

18

20

430 

26

...

16

15

113

...

7

...

30

356 

133

11

132

12

36.5 

236 

173

41.6

106

102

0.1 

2.99 

0.54

102

2.62

9.0 

7.3 

3.40

1.32

0.97

3.69

0.53

109

27

31.1 

174 

124

26.6

47

121

0.06 

0.2 

3.79 

0.60

138

3.06

11.2 

10.7 

4.39

1.48

0.96

4.41

0.68

72 125

20

40.5 35.9 

60 220

30.4 36.8

22 96

88 114

0.1 

1.19 2.41

44 102

1.15 2.83

3.7 9.0 

4.0 9.8 

1.77 3.56

0.74 1.26

0.52 0.83

1.92 2.88

0.30 0.52

83

31

41.1 

71

37.1

36

104

0.2 

2.04

74

2.00

6.4 

7.2 

2.76

0.97

0.72

2.85

0.42

in these rocks. A small number of samples described 

as basaltic andesites and spilites, and probably some 

so-called epidiorites, display significant alkali enrichment that in part is due to alteration; these rocks 

have been lumped in with the plagiogranite suite in 

fig. 1.10.

The positive correlation between TiC>2 and P2C>5 in 

the mafic rocks is consistent with concurrent fractionation of apatite and Fe-Ti oxides (see fig. 1.8). Negative 

correlations of TiC>2 with Ce, Hf, and Zr simply indicate that Ti is not incompatible in the more evolved 

magmas, whereas Ce, Hf, and Zr are (tables 1.4-1.6).
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON 17

basaltic rocks, and gabbroic rocks with more than 0.7 weight percent "HO^) of the Canyon Mountain Complex

below detection limit]

CMG162

Diabase

CMG166

Diabase

CMG211

Diabase

CMG207

Diabase

CMG43C

Hornblende 

epidiorite

CMS

Gabbro

CM6

Plagioclase 

phyric 

basalt

CM25 CMG32

Basalt Basalt

CMG81

Basalt

CMG227

Spilite

OC216 CMG18A

Spilite Porphyritic 

basaltic 

andesite

CMG146B

Porphyritic 

basaltic 

andesite

in weight percent

51.30

1.00

15.50

...

9.57

0.19

7.75

10.20

2.68

0.10

_.

.__

...

98.29

52.40

1.23

16.30

10.10

0.19

7.42

8.63

3.33

0.12

...

...

...

...

99.72

52.50

1.86

14.90

11.50

0.18

6.95

8.29

3.33

0.60

...

.__

100.11

55.90

1.32

14.80

...

10.00

0.18

6.15

7.55

4.36

0.20

...

...

...

100.46

51.90

0.72

15.90

9.64

0.18

8.46

9.15

2.81

0.68

...

...

...

...

99.44

48.65

1.49

15.97

1.87

8.18

0.20

7.91

8.41

3.21

0.55

0.12

3.08

0.24

0.21

100.09

50.91

1.55

14.84

1.84

8.53

0.30

7.33

5.89

4.19

0.46

0.10

3.23

0.34

0.10

99.61

50.07 53.00

1.21 1.39

15.56 15.40

1.18

8.71 10.90

0.18 0.19

6.69 7.02

9.01 7.64

3.60 4.43

0.77 0.20

0.07

2.56

0.09

0.14

99.84 100.17

54.90

1.57

15.10

...

11.90

0.18

5.97

5.65

5.18

0.09

...

100.54

61.00

1.04

15.10

...

9.71

0.18

3.48

4.19

5.58

0.04

...

...

...

...

100.32

61.50 57.80

1.23 1.30

15.70 15.20

...

7.68 8.38

0.23 0.13

2.83 5.62

2.12 5.16

6.64 5.45

0.15 0.89

...

...

._

...

98.08 99.93

59.10

1.00

15.1

...

10.00

0.14

6.43

2.59

4.92

0.42

...

...

...

...

99.70

in parts per million

167

7

42.8

289

162

37.6

66

82

.__

0.2

1.50

0.28

55

1.57

5.2

5.3

2.18

0.80

0.50

2.15

0.34

197

9

40.1

288

124

38.2

49

91

0.09

0.4

2.09

0.28

77

1.77

7.5

6.4

2.79

0.97

0.71

2.65

0.43

41.3

...

80

40.3

0.09

3.28

0.68

...

2.86

10.2

8.8

4.03

1.40

1.03

3.88

0.60

79

22

37.4

214

138

31.4

44

109

0.8

2.40

0.21

85

2.86

8.3

8.2

3.41

1.02

0.82

3.10

0.49

...

42.4

...

121

38.9

...

...

1.31

0.37

...

1.10

2.5

3.4

0.96

0.43

0.26

1.25

0.18

50

...

332

267

...

...

8

50

...

325

240

...

...

...

...

7

...

...

129

40 19

39.8

321 303

123 162

36.3

67

98

...

...

2.65

0.63

90

2.19

9 7.9

6.9

3.77

1.47

0.87

3.16

0.50

78

1

40.4

349

15

33.0

17

117

...

...

2.01

0.45

69

1.71

5.0

6.8

2.85

1.18

0.73

2.74

0.44

45

...

29.4

...

2

21.1

8

106

...

2.00

0.38

68

1.75

5.4

6.0

2.45

0.98

0.57

2.37

0.40

91

141

20.4 34.9

7

4 250

8.8 27.6

32

76

0.09

0.2

4.04 2.27

0.19 0.55

84

3.62 2.09

14.2 6.6

17.1 7.0

6.16 3.08

1.78 1.09

1.43 0.72

4.95 2.81

0.74 0.42

83

33

33.5

248

26

26.9

17

95

...

...

1.58

0.54

51

1.34

5.4

8.4

4.08

1.84

0.99

3.12

0.45

The scatter observed for some elements or oxides may 

reflect alteration (Ba, K^O) and possibly analytical uncertainties (Th) at low concentration levels. More importantly, the significant decrease in contents of many 

incompatible elements (Ce, Hf, Zr, Th) and the retention of comparable levels for others (K^O, Th, Ba) observed in going from the basaltic rocks to the felsic 

ones is inconsistent with a comagmatic relationship 

between these two groups of rocks. Further comparisons of covariations between other expectedly incompatible elements (fig. 1.11) reveal that, among the 

mafic rocks, La (like Ce) correlates strongly with Ti,
18 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 1.5. Analyses of major-element oxides and selected trace

[ , not determined or

Sample CM17 CM15 CM51 CM41 CM42 CM11 CM10 CM18 CM44

Rock type Feldspathic Diorite with Diorite Hornblende 

diorite cumulate diorite 

clinopyroxene pegmatite 

plagioclase

Hornblende Hornblende Diorite Quartz diorite Hornblende 

diorite diorite (xenolith in (chloritized) plagiogranite 

keratophyre)

Major-element oxides,

SiO2 - 

Tin

AI n~

fl U

Fe2O3 - 

peQ ____

MnO 

MgO 

POO

Na2O - fi

v r»

P

r\ 2U5 

HO+

H2cr -

CO2

Total -----

53.64

0.69

n

oq

1.38 

5.80 

0.15

5.33 

600

A QC

0.84

0.07

9 7^

0.25 

0.11

99.58

KK ni

0

£Q

4.56 

4.80 

0.17 

4.45

3 81

n 31

0.08

3

QQ

0.31 

0.05

100.16

KK 44

n Q4

1 C KK

2.14 

6.29 

0.16 

4.10

A O£?

n 98

n iq

3 m

0.37 

0.38

99.75

KO vn

n *w

i a an

2.36 

4.59 

0.17 

5.04 

10.14

3 CO

3 no

0.62 

0.34

100.09

C O "I "1

0.80

14.86

1.57 

6.20 

0.18 

4.45

n 79

4

rjf*

0.43

0.05

2

qc

0.26 

0.15

99.89

54.15

1.17

15.49

2.74 

8.19 

0.19 

4.12

C OQ

3 87

O Q8

0.08

2

f*rj

0.27 

0.09

99.74

59.58

1.10

1 c OQ

1.97 

6.38 

0.15 

3.00

fi 1Q

3 Q4

n °»9

0.10

i an

0.19 

0.01

99.98

71.46

0.30

14.61

0.50 

0.45 

0.02 

0.56

o 79

6 or

n 9^

0.05

1.12

0.15 

0.05

99.59

7C QC

0.14

n

an

0.73 

1.63 

0.07 

0.73

f\ QC

fi 94

0.07

0.01

n QI

0.13 

0.18

99.59

Trace elements,

Sr 

Ba 

V 

OC

155

17

on

185 

IP.

oc

216

7

1 ^

223 

iia

40

166

7fi

KK

371 

?;

KH

164 

a

on

6 

1

9fl

5 

9

NbHfCe 5276 8 9 6 4

MJ

Sm -

Eu 

Tb 

Yb 

13

Hf, and Zr, and slightly less so with Th. Thus, it seems 

reasonable that all of these elements and probably Ta 

and Nb as well (see Wood, Joron, and Treuil, 1979) are 

controlled mainly by magmatic processes and are not 

significantly affected by alteration or weathering processes (see Menzies and others, 1977). On the other 

hand, the alkalis (K, Na, Ba) usually display poor correlations with the other determined incompatible elements, and it appears that they may be unreliable as 

petrogenetic indicators (see Pearce, 1983). Despite the 

wide range in bulk composition and the apparent influence of alkali metasomatism, it is notable that K^O 

contents are generally low (<1 weight percent) in all of 

the rocks studied.
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON 19

elements in rocks from plagiogranite suite of the Canyon Mountain Complex

below detection limit]

CM 12 CM43 CMG174C CMG190B OC314 CMG137 CMG2 CMG55A CMG191 OC114

Plagiogranite Hornblende Epidiorite Hornblende Hornblende Tonalite Tonalite Trondhjemite Trondhjemite Trondhjemite 

plagiogranite (altered diorite diorite 

gabbro)

in weight percent

74.45 77.42 59.40 62.30 59.40

0.30 0.10 0.44 0.65 0.90

12.45 11.72 15.10 14.20 14.90

1.27 0.85

1.96 1.06 6.06 7.67 7.26

0.07 0.04 0.13 0.16 0.18

0.85 0.34 7.91 4.90 4.84

2.37 1.16 6.62 5.47 6.28

5.49 6.01 3.62 4.54 4.71

0.17 0.08 0.11 0.56 0.40

0.04 0.01

1.30 0.68

0.15 0.12

0.01 0.08

100.88 99.67 99.39 100.45 98.87

61.10 67.30 69.70 71.90 76.00

0.65 0.55 0.50 0.36 0.21

16.30 16.10 15.80 15.50 13.80

8.30 5.54 4.22 3.31 1.56

0.18 0.18 0.10 0.07 0.06

3.71 0.94 1.12 0.60 0.38

5.31 3.30 3.29 3.65 1.04

4.41 5.00 4.40 3.91 5.99

0.26 0.40 0.98 0.48 0.15

...

100.22 99.31 100.11 99.78 99.19

in parts per million

146

15 20 28

22.0 33.8 31.3

16 3 164

1 1 206 73 72

29.7 23.8 21.8

32 47

87

0.2

1.57 2.29 2.20

0.63 0.49 0.12

63

3.44 1.88 1.95

10 20 9.2 7.0 6.2

3.9 6.5 8.9

1.25 2.56 2.82

0.47 0.84 0.83

0.20 0.60 0.63

0.84 2.68 2.87

0.15 0.44 0.49

Considering the more reliable elements, it is clear

that the least fractionated felsic rocks with the highest TiO2 contents have lower incompatible-element

contents than many of the basaltic rocks. This point

is well illustrated in comparing chondrite-normalized

rare-earth-element (REE) profiles for rocks of the

Canyon Mountain Complex (fig. 1.12). This figure

107 148 100 49

57 63 46 23

28.4 15.8 12.5 4.7 4.5

8 45 28 6

10 7.4 13 13 3

20.7 6.2 7.5 4.8 1.8

 1 4 4 10

76 57 38 19

0.09 0.11 0.10

0.6

1.71 3.12 2.99 4.65 5.83

0.46 0.35 0.42 0.42 0.29

100 108 153 198

1.74 3.42 3.07 2.34 4.21

5.7 10.4 9.3 6.7 14.0

5.2 9.6 8.3 4.9 11.1

2.14 3.39 2.68 1.80 3.84

0.91 0.83 0.80 0.59 0.97

0.65 0.86 0.70 0.53 0.91

2.26 4.03 2.90 2.68 4.74

0.36 0.63 0.49 0.41 0.77

contrasts REE contents for a variety of mafic to felsic

rocks, including the basaltic andesites and spilites.

Despite evidence for selective alteration of some of

these rocks, REE contents and relative abundances

are broadly similar for the mafic rocks. From this observation, it appears that the REE have not been significantly mobilized and they should be useful for
20 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 1.6. Analyses of major-element oxides and selected trace elements

[ , not determined or

Sample- - CMG15 CMG157A OC214

Rock type Quartz Plagioclase Plagioclase 

keratophyre keratophyre keratophyre

SiO2 - 75.80 75.00 73.20 

TiO2 0.14 0.27 0.59 

A124 - 14.20 12.60 12.60

FeaOg 

FeO 2.70 3.94 5.75 

MnO 0.10 0.07 0.15

MgO 0.25 0.49 2.56 

CaO 0.56 0.83 0.77

Na2O - 6.23 6.19 5.09

K2O - 0.11 0.06 0.03 

P205 

Hn+

H2cr 

PO

Total 100.09 99.45 100.74

gr ______ 72 49 23 

Ba 20 10 8 

Sc - 9.7 13.3 18.4

V V _ -- ^j 9 -- 79| ^

Cr 19 15 44

Co 0.8 1.6 12.0 

jiji _____ 4 7 22

7,n _______ AG 70 107

Ta 0.12 0.08

Nb ------ 0.2 

Hf 6.47 3.04 3.41

TVi -- ft 1fi 018 ft 1fi

Zr 225 98 123

T a - 4 9^ 9 V7 911

Ce 14.3 7.8 8.2 

Nd 14.9 7.9 8.6

Sm 5.07 2.99 3.42

TTii f\ *7O f\ QO f\ nn

TK 1 V7 1 fin O Qft

vv* a of\ o oc A f\ A

Lu 1.02 0.58 0.64

petrogenetic purposes. Another striking feature of the 

data is the strong similarity between REE profiles for 

the mafic and the felsic rocks, with the exception 

that the latter have marginally higher La/Sm ratios 

and conspicuous Eu anomalies. The most unusual 

REE pattern obtained is for an altered gabbroic rock 

(the so-called epidiorite, sample CMG-174C) that displays significant depletion of heavy REEs.

OC224 OC241 OC243 CM1

Quartz Keratophyre Quartz Keratophyre 

keratophyre keratophyre

Major-element oxides,

79.40 76.60 76.10 77.92 

0.16 0.18 0.18 0.14 

11.00 12.80 12.70 11.47 

1 17

1.73 3.21 3.82 0.60 

0.04 0.04 0.10 0.02

1.01 0.24 0.52 0.31 

0.97 1.50 0.52 0.82

5.10 5.49 6.07 5.85

0.09 0.09 0.09 0.10

0.02 

0 78

0.23

0 91

99.50 100.15 100.10 99.64

Trace elements,

83 27 

38 7 10 

8.7 11.9 15.0 

8 11 6 

4 3 22 2

1.9 1.4 3.1 

 jo q

oc 70

n OQ o 07

0.1 0.1 

3.71 3.65 3.30 

0.39 0.16 0.21

120 108

1 QO 0 (\A Q 9Q

7.8 11.5 10.4 10

6 Q 11 q

2

71 °. ^8 °i fi7

A OQ A QC 1 f\Q

ft so n ss n Q8

A f\O O A O A O^7

f\ CA f\ CA f\ CQ

COMPARATIVE GEOCHEMISTRY AND 

EVALUATION OF TECTONIC SETTING

Having established that most of the analyzed incompatible trace elements are likely to be representative of original magmatic abundances, it is of 

interest to determine the affinity of the mafic rocks of 

the Canyon Mountain Complex to other present-day
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON 21

in keratophyres and quartz keratophyres of the Canyon Mountain Complex

below detection limit]

CM3

Keratophyre

CM7

Keratophyre

CM9

Keratophyre

CM45

Keratophyre

CM47

Quartz 

keratophyre

CM49

Quartz 

keratophyre

CM50

Quartz 

keratophyre

CM46

Quartz 

keratophyre

in weight percent

75.98

0.14

11.53

1.12

2.27

0.07

0.53

0.74

5.58

0.43

0.05

0.94

0.17

0.07

99.62

73.62

0.22

12.94

1.77

1.52

0.04

0.68

1.64

5.89

0.32

0.05

0.78

0.15

0.04

99.66

75.98

0.15

12.12

1.36

1.19

0.03

0.40

1.49

5.63

0.25

0.01

0.85

0.16

0.23

99.85

74.82

0.17

12.44

1.14

1.98

0.08

0.45

1.64

5.81

0.06

0.01

0.98

0.14

0.18

99.90

78.52

0.10

11.13

0.74

0.89

0.02

0.52

0.61

6.07

0.06

0.01

0.68

0.14

0.23

99.72

75.72

0.16

11.91

0.80

1.73

0.06

1.02

0.32

6.44

0.09

0.01

1.01

0.20

0.14

99.61

74.32

0.21

12.18

1.36

2.05

0.08

0.50

1.63

5.38

0.08

0.02

1.36

0.19

0.24

99.60

77.22

0.10

11.83

0.65

1.08

0.04

0.64

0.29

6.39

0.07

0.01

0.72

0.17

0.36

99.57

in parts per million

10

12

2

10

9

3

30

3

1

3

3

10

4

2

10

15

2

15

5

2

15

4

1

19 36 10 13

basaltic associations. Unfortunately, a similar comparison cannot yet be attempted either for the cumulate and ultramanc members of the complex, owing 

to insufficient data and because these rocks do not 

approximate magmatic liquids in composition, or for 

the felsic rocks, because of limited data and ambiguity of the available petrogenetic discriminators for 

such rock types (see Pearce and others, 1984).

Direct comparison of representative diabases and a 

basaltic andesite of the Canyon Mountain Complex 

with basaltic rocks from other environments is illustrated in figure 1.13, which shows incompatibleelement abundances normalized to Wood's (1979) estimated primordial-mantle composition. Similar profiles 

are obtained by normalizing to estimated midocean 

ridge basalt (MORE) abundances (see Sun and Nesbitt,
22 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

20

15

10

CaO

XX

*M +H- °

*%*

30

20

10

AI2O3

X

1-

LU

0 

CE

LU 

Q_

hX

O

LU

50

40

30

20

10

n

-MgO

m

_

** M

M w 

X

v

i

Explanation

 Keratophyre

n Plagiogranite

A Andesite

+ Basalt

* Cumulate gabbro

M Ultramafic rocks

-

_

~

*Jt fth iMj n

XX* ^yfrjfrtft £ n

. i . i . n n m riffinj

12

10

8

6

4

2

0

FeO

+ B <^

f a*

^ MM 

x * 

x

- l.U

£

£ 0.8

Z

O

U 0.6 

LU

Q

X 0.4

0.2

0.0 

0.25

0.20

0.15

0.10

0.05

K20

+

A

+ D

_l_ £3 

+ a *a a

+ + A ° ' -

. D a a

**+ + n

++ . n A D a

IJ^^ ,MM+. ,A . , " *

P O

+ +

a

a

* a°

a a m m -

o a

m m 

XMNX , . , . , . "dOi

°'0°40 50 60 70 8(

0 

3.0

2.5 

2.0 

1.5 

1.0

0.5 

0.0.

Na2O

TiO2

+a

S A

D° 

D

++r +

A a a 

a

'40 50 

Si02 CONTENT, IN WEIGHT PERCENT

60

aJ

70 80

FIGURE 1.7. SiO2 variation diagrams for all analyzed rock samples of the Canyon Mountain Complex. FeO*, total

iron as FeO.
*J

INCOMPATIBLE-ELEME

£|gl EXCEPT Mg# ' IN PARTS PER M

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DD

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24 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

10

MgO

Island arcs

4 6 8 10 12 14 16 18

FeO* CONTENT, IN WEIGHT PERCENT

0

DC LU . 

Q_

1-

X 75

CD

LiJ

£

Z

Z

LiJ 1-

o cc O 55

LiJ

Q

X

O

45 Q n

O.U

2.5

2.0

1.5

1.0

0.5

i i i

Si02 B

A

/ A

 CA / 

/TH

/ A n

/ D

A/A A n °

W+ n n

4i-L.

TSr /^

Ti02 C 

Jl n Explanation

jjj a Galapagos Ridge

)ii/ a n MORE averages

Jl|f A Troodos

" Hlf A + Cany°n Mountain -

"§A A "

-f^MOPB ^ ;

0.0 

0 5 10 15

FeOVMgO RATIO

FIGURE 1.9. Fe-Mg variations in basaltic rocks of the Canyon 

Mountain Complex compared to trends for average midoceanridge basalts (MORE; Melson and others, 1976), lavas from 

the Galapagos Ridge (Perfit and others, 1983), and glasses 

from upper pillow lavas of the Troodos ophiolite (Robinson 

and others, 1983). FeO*, total iron as FeO. A, MgO versus 

FeO*. Canyon Mountain Complex and Troodos ophiolites 

both show only moderate Fe enrichment relative to Galapagos and Skaergaard rocks. Island-arc data field (stippled) and 

Skaergaard trend (arrow) are from Robinson and others 

(1983). B, SiO2 versus FeO*/MgO ratio. Samples from all 

suites display only limited increase in FeO*/MgO ratio with 

increasing SiO2; Canyon Mountain Complex samples straddle the calc-alkalic (CA)/tholeiitic (TH) boundary of Miyashiro (1973). C, TiO2 versus FeO*/MgO ratio. Mafic samples of 

all suites, including the Canyon Mountain Complex, plot 

within the MORE data field (from Casey and others, 1985).

1977; Sun and others, 1979; Pearce, 1983). Other than 

anomalous spikes in Ba and K, the Canyon Mountain 

Complex basaltic andesite has a profile similar to those 

of the diabases (fig. 1.13A). Furthermore, excluding the 

epidiorite, all Canyon Mountain Complex mafic rocks 

display striking relative depletions in the light REEs, 

Nb, and Ta relative to the heavy REEs. All analyzed 

samples have relatively high ratios of K, Th, or Ba to 

Nb (or Ta), and the Hf-Zr-Sm segments of their profiles 

are slightly convex upward. Except for the Ba-Th-K-Nb 

segments, these profiles are similar to those of many 

MORBs as exemplified by a sample from the Ecuador 

Rift in the East Pacific (fig. 1.135). For comparison, a 

sample of the Geotimes basaltic unit of the well-studied Semail ophiolite (Oman) closely resembles the 

MORE profile shown (except for a K spike). This particular basaltic unit has been interpreted as a fragment of 

oceanic crust; its K spike is explained by extensive

FeO*+MnO

Na2O+K2O MgO

Explanation

Basalt (O), diabase ( )

Cumulate gabbro 

Ultramafic rock

Plagiogranite suite

Keratophyre

Spilite trend (alkali enrichment)

FIGURE 1.10. A (Na2O+K2O)-F (FeO*+MnO)-M (MgO) 

ternary diagram for samples of the Canyon Mountain 

Complex. Basaltic and equivalent intrusive rocks (diabases, dolerites) overlap in composition and display 

only moderate Fe enrichment. Relatively evolved rocks 

(including basaltic andesite) and slightly altered rocks 

(spilite, epidiorite) plot in plagiogranite field; alkali enrichment trend (arrow) indicates effect of alteration.
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON 25

alteration (Alabaster and others, 1982). The relative 

depletions of Nb (and Ta) seen in the Canyon Mountain 

Complex samples are uncharacteristic of MORE samples (Noiret and others, 1981) and suggest that the 

Oregon ophiolite perhaps formed in a setting other 

than near a midocean ridge.

Oceanic-island basaltic rocks (such as those found in 

Hawaii or Iceland) or intracontinental (within-plate)

basaltic rocks generally are quite distinct in composition from ophiolitic basalts (see fig. 1.13.D). The Salahi 

basalt unit, also part of the Semail ophiolite, is a possible exception (fig. 1.13B). This unit is interpreted as 

having formed at a seamount by Alabaster and others 

(1982), who viewed the Semail ophiolite as a composite 

body formed as tectonic conditions changed with time. 

Commonly characterized by low relative abundances of

(

6

5

4

o

5

5 2

CC 

LU

0_ .

Hf ' 

D

-

D

A

 T

M- + D "°

 £^-

A 4. D

T- | i.i.

^DU

200

150

H

LUb 100

^~ "^

81

1- H 50

z ^

LU ^

^ rr

Zr - -

D

D

+

D .+ + n 

* A+ +

: . +*°* : +

i.i.i .

H 12,000

CC

0- 10,000

Z

K

_ 8,000

LUIZ 6,000

0

o

I- 4,000

Z

LU

ii 2,000

_J

LU

LU 0

_J

CO

< °'7

Q_ 

5 0.6

O 

? 0.5

0.4

0.3

0.2

 | . | i | 

 Ti + + -

+

~

" + A + A "

. A + A +

D

A

DD

" Qn

D

n .

1 . I.I

 Th +

+ a

A A + 

D

-P

D D 

A A "

D

n -

.

- + *

a.

Qj ^ 80

CD 5

< °- 60

Q. ?

^>

O

y 40

^

20

0

1.0

IZ 08

- LU 

}f| LU 0.6

O ^

O ^ 0.4

O rr]

^ >

> 0.2

r\ r\

, , . i .

.Ba

D

D

.

D

-

Q +

+ D

+ + -

+

+ "*" "

. 4- 1 . 1 . 1

. | . | . | 

'K2O n Explanation "

A + Basalt 

A Andesite

A D Plagiogranite 

+ " Keratophyre _ n ' '

D

- A n n

+

a +

+ ^" +

+ + n A D.

* 

°' 1 1 2 3 4 5 "'"1 2 3 4 5

La CONTENT, IN PARTS PER MILLION

FIGURE 1.11. La-variation diagrams for incompatible elements in magmatic rocks of the Canyon Mountain Complex.
26 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

102

10°

: Keratophyre

0 0

102

101

100 I I

: Plagiogranite

-+- CMG190B 

-*- OC314 

-o- CMG137

-n- CMG55A 

 CMG191 

-*- OC114

-CD- CMG2 

- - CMG174C

102

LLI

o 

o

LU 

LU 

DC

LJJ

E 

O

O

§o 

o 10°

: Basaltic andesite

102

10°

\ I i i r i i i i i

t Basalt

102

101

100

Spilite

102

101

10°

Diabase

-o- CMC 124 A

-+- OC244

-o- OC256

-4- OC259

-n- CMG162

-*- CMC 166

-* CMG211

-*- CMG207

La Ce Nd Sm Eu Tb Yb Lu La Ce Nd

RARE-EARTH ELEMENTS (REE)

Sm Eu Tb Yb Lu

FIGURE 1.12. Plots of rare-earth element (REE) contents of magmatic rocks of the Canyon Mountain Complex normalized to 

REE contents of chondrite meteorites; normalization values are 

from Sun and Nesbitt (1977). Note similarity of profiles for all 

mafic rock types (including basalt, diabase, basaltic andesite,

and spilite). Most felsic rocks (plagiogranite and keratophyre) display significant negative Eu anomalies, and plagiogranites are 

slightly enriched in light REEs compared to the mafic rocks. Sample numbers are listed in explanation at lower right corner of each 

graph.
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON 27

Nb and Ta, basalts from volcanic arcs associated with 

subduction processes appear to be the strongest analogs for basaltic magmas of the Canyon Mountain Complex; these are exemplified (fig. 1.13C) by samples from 

the Andean (Laguna Maule, Chile) and Aeolian (Salina, 

Italy) volcanic arcs, and similar rocks occur in some 

potassic volcanic provinces such as Vulsini (Italy).

Various other discriminant diagrams (figs. 1.14- 

1.16) have been proposed to elucidate the origin of 

basaltic magmas from different tectonic settings (see 

Pearce and Cann, 1973; Beccaluva and others, 1979; 

Wood, Joron, and Treuil, 1979; Noiret and others, 

1981; Pearce, 1982; Shervais, 1982; Mullen, 1983b); 

these emphasize some of the details of the normalized abundance profiles seen in figure 1.13. Although 

these diagrams have little applicability to silicic magmas, data for both the mafic and felsic rocks of the 

Canyon Mountain Complex are plotted for comparison. In figure 1.14, the mafic rocks of the Canyon 

Mountain Complex plot within the island-arc basalt 

field of the Th/Yb-Ta/Yb plot, overlap both the 

MORE and island-arc basalt fields in the Ti-Zr plot, 

and lie largely within the MORE field of the V-Ti 

plot. Similar ambiguity is seen in figure 1.15, in 

which Canyon Mountain Complex samples resemble 

island-arc basalts in the Hf/Th-Ta/Th plot, but are 

transitional between island-arc basalts and MORE in 

the Cr-Yb and Ti/Cr-Ni plots.

LLJ 

LU H

4° <?

103

102

101

100

= A

Mafic rocks of 

Canyon Mountain 

Complex

103

102

100

Subduction-related ; 

basalts i

ni i B -D- Ecuador Rift 

~*~ Oman Geotimes 

~**~ Oman Salahi

Oman ophiolite 

and MORB

-; 101

- 10°

10-1

Within-plate basalts

Ba Th K Nb Ta La Ce Sr Nd Hf Zr Sm Eu Ti Tb Yb Ba Th K Nb Ta La Ce Sr Nd Hf Zr Sm Eu Ti Tb Yb

INCOMPATIBLE ELEMENTS

FIGURE 1.13. Plots of incompatible-element contents normalized 

to incompatible-element contents estimated for primordial mantle (Wood, 1979). A, Mafic rocks from the Canyon Mountain 

Complex, including two diabases (CMG49, CMG166) and a basaltic andesite (CMG18A). Samples plotted for other tectonic 

settings: B, Oceanic basalts from Ecuador Rift midocean-ridge 

basalt (MORB) (Perfit and others, 1983) and Geotimes and

Salahi basalts from Semail ophiolite, Oman (Alabaster and others, 

1982). C, Subduction-related basalts from Vulsini (Rogers and others, 1985) and Salina Island (Ellam, 1986), Italy, and from Laguna 

del Maule (L. Maule), Chile (Frey and others, 1984). D, Withinplate basalts from Kilauea, Hawaii, and Dish Hill, California 

(Thompson and others, 1984), and from Vadalda, Iceland (Wood, 

Joron, Treuil, Norry, and Tarney, 1979).
28 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

or

I

\- 10-1 r

Explanation 

X Cumulate gabbro 

+ Basalt 

A Andesite 

n Plagiogranite 

 Keratophyre

10-210-2 10-1 100 

Ta/Yb RATIO

UJ

O

12,000

Z 10,000

II

^ 8,000

6,000 

{£ 4,000 

 2,000

0 50 100 150 200 250

Zr CONTENT, IN PARTS PER MILLION

UJ

or

UJ

o_

500

400

300

> W 200 

CC

^ 100

0

c

0 5,000 10,000 15,000 20,000

Ti CONTENT, IN PARTS PER MILLION

FIGURE 1.14. Discriminant diagrams showing compositions 

of magmatic rocks of Canyon Mountain Complex relative to 

present-day volcanic suites: IAB, island-arc basalts; WPB, 

within-plate basalts; MORE, midocean-ridge basalts. A, 

Th/Yb versus Ta/Yb plot (from Pearce, 1982) shows estimated primordial-mantle composition (asterisk) and composition field for non-arc basalts (stippled band). Arrow labelled 

"WPB" indicates mantle enrichment trend typical of oceanisland (within-plate) magmatism; arrow labelled "IAB" indicates enrichment of Th relative to Ta characteristic of 

subduction-zone (island-arc) magmatism. B, Ti versus Zr 

plot (from Pearce and Cann, 1973); fields discriminate basalts from various tectonic settings: 1, island-arc tholeiites; 

2, island-arc basalts and MORE; 3, calc-alkalic basalts; 4, 

MORE. C, V versus Ti plot (from Shervais, 1982).

The ternary discriminant diagrams shown in figure 

1.16 again suggest that the basalts of the Canyon 

Mountain Complex are transitional in character. They

103

102

10°

IAB il MORE + WPB

10° 101 102 

Yb CONTENT, IN PARTS PER MILLION

103

102

O

100

IAB

102

Ni CONTENT, IN PARTS PER MILLION

103

1.0

Ta/Th RATIO

FIGURE 1.15. Discriminant diagrams showing compositions of magmatic rocks of Canyon Mountain Complex 

relative to present-day volcanic suites. IAB, island-arc 

basalt; MORE, midocean-ridge basalt; WPB, withinplate basalt. A, Cr versus Yb plot adapted from Cr-Y 

diagram of Pearce (1982) assuming Y/Yb=10, which is 

typical of most basalts. Stippled area corresponds to 

overlap between IAB and MORB+WPB fields. B, Ti/Cr 

versus Ni plot (from Beccaluva and others, 1979). C, 

Hf/Th versus Ta/Th plot (from Noiret and others, 1981).
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON 29

lie within the island-arc basalt field in the Th-Ta-Hf 

diagram but also partly overlap with MORE fields in 

the Zr-Yb-Ti and MnO-P2O5-TiO2 diagrams. Despite 

the uncertainty of assigning specific tectonic settings 

solely on geochemical grounds, the distinct Nb-Ta depletion characteristic of rocks of the Canyon Mountain 

Complex is analogous to that documented in volcanicarc rocks. Fortunately, this apparent affinity is compatible with geologic criteria, particularly with the 

large relative abundance of silicic hypabyssal rocks.

Hf/3 Ti/100

Th 

A

Ta Zr 

B

30Yb

TiO2

Explanation

[ | Felsic rocks

A Keratophyre 

n Plagiogranite

Mafic rocks 

 Basalt

^^| Cumulate gabbro 

KxxXj Ultramafic rocks

10MnO 

C

10P205

FIGURE 1.16. Ternary discriminant diagrams for Canyon 

Mountain Complex magmatic rocks compared to basaltic 

rocks from other tectonic settings. CAB, calc-alkalic basalt; IAB, island-arc basalt; IAT, island-arc tholeiite; 

MORE, midocean-ridge basalt; WPB, within-plate basalt; 

OIT, ocean-island tholeiite; OIA, ocean-island alkalic basalt. A, Th-Ta-Hf diagram (from Wood, Joron, and Treuil,

1979). JB, Zr-Ti-Yb diagram, modified from Pearce and Cann's 

(1973) similar diagram that used Y instead of Yb (see fig. 

1.15). C, TiO2-MnO-P2O5 diagram (from Mullen, 1983b). 

Ocean-island basalts are subdivided into OIT and OIA fields. 

Note that data fields for Canyon Mountain Complex cumulate gabbros and ultramafic rocks, used alone, can lead to 

confusion in assigning tectonic setting.
30 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

PETROGENETIC MODELS

Further insights into the origin and history of the 

Canyon Mountain Complex may be gleaned from consideration of the petrogenetic processes that produced its lithologic members. Very little detail is 

known about the petrogenesis of the tectonite and 

metacumulate members, but published information 

on mineral compositions (Himmelberg and Loney, 

1980) and the general characteristics of these units 

suggest that they may well represent a partial section of oceanic lithosphere. There is also relatively 

little information on the cumulate and intrusive gabbros. They clearly represent a phase of magmatism 

that occurred after deformation of the structurally 

lower parts of the complex. On the basis of clinopyroxene analyses (Himmelberg and Loney, 1980) and 

mineral-melt partition coefficients (see Elthon and 

Casey, 1985), the magmas that precipitated the cumulates apparently had Na2O and TiO2 contents not 

unlike those of some of the mafic dikes in the sheeted 

member. Nevertheless, as pointed out earlier, the 

phenocrysts in these dikes differ substantially in 

composition from the minerals in the cumulate rocks. 

Thus, it is probable that these two members are unrelated, in which case the basaltic magmas of the 

sheeted member must represent a distinct phase of 

magmatism.

Finally, there is the problem of the plagiogranitekeratophyre suite. The majority of keratophyres 

appear to have formed as a hypabyssal sequence, 

probably in an island arc, prior to emplacement of the 

Canyon Mountain Complex (Gerlach, Ave Lallemant, 

and Leeman, 1981). However, published 4°Ar/39Ar 

age determinations (Ave Lallemant and others, 1980; 

this chapter) and our field studies suggest that the 

diverse magmas of the sheeted member formed as 

part of the same magmatic phase at approximately 

262 Ma. In contrast, the maximum U-Pb zircon age of 

about 278 Ma for an intrusive gabbro high in the cumulate-gabbro sequence (Walker, 1986) suggests that 

they formed during an earlier magmatic phase. Thus, 

available data imply that the Canyon Mountain Complex consists of at least two mafic-ultramafic magmatic suites separated by a significant time interval 

(perhaps as much as 15 m.y). Furthermore, the geochemical data summarized earlier demonstrate that 

the plagiogranites (and minor keratophyres) are unlikely to be strictly cogenetic with the contemporaneous basaltic magmas despite their mutually intrusive 

relations in the sheeted member. Examination of the 

petrogeneses of these two roughly coeval groups of 

magmas provides insight into the conditions under 

which they formed.

PLAGIOGRANITE SUITE

The origin of the plagiogranites and keratophyres 

of the Canyon Mountain Complex was discussed by 

Gerlach, Leeman, and Ave Lallemant (1981), so only 

their essential arguments are repeated here. Silicic 

lavas (low-K2O dacites to rhyodacites) are found at 

certain oceanic spreading centers; in some cases 

these appear to be extreme differentiates of MORBlike magmas (Sinton and Byerly, 1980; Clague and 

others, 1981; Perfit and others, 1983). It has been 

proposed that plagiogranites in ophiolites represent 

hypabyssal equivalents of such magmas (Coleman 

and Peterman, 1975; Kay and Senechal, 1976; Pallister and Knight, 1981). Whereas this may be a reasonable model in many cases, the fact that mafic and 

felsic rocks of the Canyon Mountain Complex have 

overlapping REE profiles and contents of several 

other incompatible elements such as Hf, Th, Ta, and 

Nb (figs. 1.8, 1.11, 1.12, 1.17; tables 1.4-1.6) effectively precludes such a simple cogenetic link between the 

two magma types. Figure 1.17 illustrates REE variations that result from fractionation of basaltic magmas to rhyodacitic differentiated liquids at the 

Galapagos spreading center (Clague and others, 

1981; Perfit and others, 1983) and in the Troodos 

ophiolite (Kay and Senechal, 1976). In contrast, in 

the Canyon Mountain Complex (La/Yb)n and Ybn values are similar in the basaltic and felsic rocks.

Compositional variations for relatively immobile 

major and trace elements appear to be coherent within both the mafic and silicic suites (figs. 1.7, 1.8), and 

are somewhat consistent with the operation of fractional-crystallization processes. However, as discussed earlier in this chapter, inconsistencies in 

incompatible-trace-element abundances and compositional trends for the two suites indicate that they 

cannot have been derived from a common parental 

magma. Alternatively, on the basis of experimental 

studies, Dixon and Rutherford (1979) proposed that 

some plagiogranites form by means of liquid immiscibility. We consider this process to be insignificant for 

the Canyon Mountain Complex because the required 

element partitioning between mafic and felsic liquids 

is inconsistent with that actually observed (Gerlach, 

Leeman, and Ave Lallemant, 1981). Furthermore, the 

dominance of felsic rocks in the Canyon Mountain 

Complex is inconsistent with either fractional crystallization or liquid immiscibility.

It was proposed instead that the felsic magmas 

formed as a result of extensive partial melting of hydrated (that is, amphibole-bearing) mafic rocks in the 

underlying crust (Gerlach, Leeman, and Ave Lallemant, 1981). Although this model cannot be uniquely
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON 31

quantified, it is consistent with available experimental studies on melting of various basaltic compositions in the presence of high water contents 

(Holloway and Burnham, 1972; Helz, 1976; Spulber 

and Rutherford, 1983; Rapp and others, 1991) in that 

compositions of the experimentally produced liquids 

approximate those of natural plagiogranites. Assuming single-stage batch melting at low to moderate 

pressure (<5 kb) and at partial water pressure less 

than or equal to total pressure, Gerlach, Leeman, 

and Ave Lallemant (1981) demonstrated by inverse 

modeling that REE profiles in the required plagiogranite sources are likely to resemble those in the

mafic rocks of the Canyon Mountain Complex and in 

many other ophiolites. This process presumably is 

driven by convective heating associated with the emplacement of basaltic magmas. The availability of 

aqueous fluids is a critical feature of the model and 

can be accommodated either by hydrothermal convection in the shallow crust in areas of crustal extension 

(Gregory and Taylor, 1979) or by subduction processes that provide access of fluids or water-bearing magmas (see Pearce, 1982). The depth of partial melting 

is uncertain but, judging from the exposed stratigraphic section, the main plagiogranite bodies were 

emplaced at a minimum depth of 5 km. Although

A

1.0

0.5

0.0

85

75

Z O 

LJJ DC

O I- 65

O I

Q 

CO 55

45

Canyon Mountain Complex

Explanation

+ Basalt 

A Andesite 

n Plagiogranite 

' Keratophyre

10 20 30 40

B

1.5 

1.0 

0.5

0.0

75

70

65

60

55

50

45

Galapagos and Troodos

Explanation

n GSC-Perfit 

 GSC-Clague 

O Galapagos MORE 

A Troodos

20 40 60 80 100

ROCK/CHONDRITE Yb CONTENT (Ybn)

FIGURE 1.17. Rare-earth element (REE) fractionation in A, magmatic rocks of the Canyon Mountain Complex and B, Troodos ophiolite 

(Kay and Senechal, 1976) and midocean-ridge basalt (MORE) and 

differentiated lavas from Galapagos spreading center (GSC) 

(Clague and others, 1981; Perfit and others, 1983). Plots of chondrite-normalized La/Yb and Yb abundances illustrate REE fractionation; values of (La/Yb)n=l and Ybn=l correspond to chondritic

relative abundances of REE (that is, flat profiles). Variation of SiO2 

with Ybn illustrates the range in bulk rock compositions. Ybn may be 

variably enriched depending upon details of crystallization process. 

Because La is more incompatible than Yb in such magma suites, 

(La/Yb)n ratio should increase significantly during fractionation, as is 

the case in both Troodos and Galapagos rocks. Note difference in scale 

for both (La/Yb)n and SiO2 in parts A and B of figure.
32 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

there is field evidence for hydrothermal alteration of 

wallrocks near these bodies and even for limited in 

situ melting (for instance, leucocratic segregations in 

hornblende-rich rocks that may be altered basalt or 

gabbro), the primary partial-melting zone must have 

been at a greater structural depth than is presently 

exposed. An attractive feature of this model is that it 

can account for the contemporaneity of the mafic and 

felsic magmas and the common observance of plagiogranite intruding gabbro and basalt in the upper 

part of the complex.

The thick section of keratophyres in the uppermost 

part of the Canyon Mountain Complex is believed to 

represent a slightly older or possibly coeval subvolcanic pile into which the plagiogranites and associated 

mafic magmas were emplaced (Gerlach, Ave Lallemant, and Leeman, 1981). Because the keratophyres 

are compositionally and mineralogically similar to the 

most evolved rocks of the plagiogranite suite (trondhjemites), it is possible that keratophyre magmas (or 

their precursors) had a similar origin. Allowing for 

metasomatic effects (Na gain or K loss), the geochemical features of these rocks (such as low Ta and Nb 

despite their evolved compositions), their porphyritic 

texture, and the common occurrence of plagioclase 

and quartz phenocrysts (indicative of originally dacitic to rhyodacitic magmas) suggest that the keratophyres formed in an island-arc environment.

MAFIC ROCKS

Mafic rocks of the sheeted member are dominantly 

basaltic. Their major-element compositions may be 

compared in some detail with those of volcanic rocks 

from the Galapagos Ridge (Perfit and others, 1983), 

average MORE compositions (Melson and others, 

1976), and fresh glasses from the Troodos ophiolite 

(Robinson and others, 1983) by juxtaposing SiC>2 and 

MgO variation diagrams (figs. 1.18, 1.19) for these 

rocks. Although there is some scatter in the Canyon 

Mountain Complex analyses, many clearly resemble 

MORE analyses with respect to MgO, P2C>5, TiO2, 

FeO* (total Fe as FeO), and A12O3. The relative depletion in CaO and variable enrichment in Na2O and K^O 

compared to MORE averages are attributed to posternplacement alteration. Such patterns of alteration are 

not unusual for ophiolitic basalts (see Hopson and 

Frano, 1977; Vallier and Batiza, 1978; Schminke and 

others, 1983). Unfortunately, even minor modification 

of major-element compositions precludes use of detailed mass-balance methods of petrogenetic modeling. 

We can, however, use data for the more immobile trace 

elements to evaluate the relative importance of fractional crystallization and partial-melting processes in 

producing these magmas.

Two fundamental questions concern (1) whether 

the various mafic magmas are cogenetic, and (2) 

whether they are related to the lower parts of the 

complex (such as the undeformed isotropic or layered 

gabbros). The first question can be assessed through 

inspection of figures 1.8, 1.11, and 1.20. In general, 

the basaltic rocks define strong coherent trends 

among such elements as Ti, P, Zr, Hf, LREEs, and to 

a lesser extent Th. For some of the analyzed samples, 

even Ba and K correlate well with these other incompatible elements. This correlation is notably poorer 

for some of the basaltic andesites and spilites, which 

are presumably altered to varying degrees. Although 

not shown in detail, there is antithetic depletion of 

compatible transition elements with increasing incompatible-element contents throughout the suite of 

rocks composing the Canyon Mountain Complex (see 

figs. 1.18 and 1.19).

A simple test of coherent fractionation involves comparison of relative enrichment factors for the immobile trace elements (Saunders and others, 1980), 

which should correlate systematically with degree of 

fractionation in a cogenetic magma suite. Because 

phenocryst-magma distribution coefficients (D) for the 

incompatible trace elements are very low in magnitude (much less than 0.1), during crystallization in 

closed magma systems these elements tend to be enriched in the residual liquids in inverse proportion to 

the fraction of original liquid remaining at any given 

stage. Thus, for a suite of cogenetic derivative liquids, 

if the incompatible-trace-element abundances are normalized to those in a relatively primitive liquid, such 

normalized ratios should be highly correlated for 

equally incompatible elements (that is, those with 

similar D values). Such relationships for representative elements in the Canyon Mountain Complex mafic 

rocks (normalized to elemental abundances in a relatively unfractionated diabase, sample CMG162) are 

shown in figure 1.20. Normalized Ce values (Cen) are 

comparable to those for other light REEs, Zr, and Hf. 

Ce appears to be slightly enriched relative to TiO2, 

Eu, and Yb; this behavior is consistent with the relative D values characteristic of these elements (see 

Frey and others, 1978). Similar plots are shown for Ba 

and K^O to illustrate their anomalously large (alteration-related) enrichments relative to those expected 

for closed-system fractionation. Finally, the surprisingly unsystematic behavior of Th may partly reflect 

analytical uncertainties. In general, the behavior of 

the immobile incompatible elements implies cogenetic 

relations among most of the mafic rocks. In detail, 

however, there are inconsistencies and irregularities
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON 33

12

10

8

6

4

2

8

/; I- 6

LU 

ODC 4

LU

QL

£ 2

O

LU 

^ 0 

1 A ' ' ' '

-CaO - A

T-? -ft*-

-H*- A

++±D

++ A

 n

+ + D A

a

AA

 i-^-. i i 

- MgO "^j. Explanation

+^[^ n Galapagos Ridge .

° 1f+ MORE averages _

, A Troodos 

A + Canyon Mountain

D A A

n A

n n

A

D

, AA.

z

z

LLJ

I- 1.0

O

o

LU

g

0 0.5

0.0

0.8

0.6

0.4

0.2

0.0 w v

1 A ' K2o ; -

+

 ++ A .

i ^\

^4P -r^A %^4^

, D +.

 i i i 

P2O5

D

D

D

D

-1^"° °

 * j. A A A

17

15

13

i 1 i 1 . 1 i

AI2O3

, +

.-^A

~4 t /A

|W+;fcA A

D

D

A

D

^ ,

11 

16

14

12

10

8

4 

6

5

4

3

2

1

3

2

1

r\

 _ FeO* D J

n n

D

n

" A^A" D "

IpF "T- "

-i- QA^A

AA i.i.i.

,.,.,.

 Na2O . -p "

-kH^-+

"*"++ , D A^

+ + a

-H. ++ cP A A

J. D A

"^B A

AA

A

i.i.i.

. i . i . i .

TiO2

_i_ D D

+_!_

D

^-+ n D

°4f +

Jg+++AA

++ ^

AA A

A AA

i.i.i.

40 50 60 70 80 40 50 60 70 8(

Si02 CONTENT, IN WEIGHT PERCENT

FIGURE 1.18. Silica variation diagrams for basaltic rocks from 1976), glasses from the Troodos ophiolite (Robinson and others,

the Canyon Mountain Complex contrasted with average mid- 1983), and selected lavas from the Galapagos Ridge spreading

ocean-ridge basalt (MORE) compositions (Melson and others, center (Perfit and others, 1983). FeO*, total Fe as FeO.
34 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

11-

12

10

8

6

4

2

75

H

Z

LLJ

O

cc oc UJ 65 

CL

^

O

HI 55

Z" ~

h"** A^

,.,.,.,.

CaO A

^3

1 Q 4>

A + + 

D D . ++ +

-A t -

DA + + + .

D T

" D A +

-

4

i . i . i . i

SiO2

A 

A

0 A

D 

A

n n A A +

° ° +++ ll + "

+ + D jTuPkh

1 . 1 . 1 . 1

1 /

16

15

14

13

12

11

18

16

14

12

10

8

6

A

i i i + i 

; AI203 A A ++ 

A + +A-K + +

A ++ \ ?ft

A + 

D

D D

 A

D

A D

n n

D 

D

i . i . i . i

 FeO* ;

n

a a

a

a

a .a

4 + m

D J- M_ '

A A ++ ± *|f +

A + + 

A ^

-

A

i . i . i . i

z a

LLJ

H

o

0 1.0

UJ

o

X

o

0.5

0.0 ,

K2o ;

.

+

A . a + +

a +

A ° a A+ + +

.a A 5ftj-+ +

' .,.,.% ° r* .'

5

4

3

2

4

- Na2O + +

+ ++» +

-* D ^ + + -

n * +

/P D A + + + +

+

ff D

 _0t

A n n n- +

A A

A

. 1 . 1 . t . 1 

.0

0.6

0.4

0.2

n n

,.,.,.,.

 '2^5 Explanation

n

n n Galapagos Ridge

n MORE averages

A Troodos

D + Canyon Mountain "

n n ++ D J.

A + " mm

A "^ A +

 1 . 1 . 1 . 1 .

H3

2

1

n

TiO2

-

n n a +

H" n

 A + ,T n , ^-

- D A A + + "+

A A

4 A

i . i . i . i

6 8 10 024

MgO CONTENT, IN WEIGHT PERCENT

10

FIGURE 1.19. MgO variation diagrams for basaltic rocks from the Canyon Mountain Complex; other samples as given in 

previous figure. FeO*, total Fe as FeO.
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON 35

o: 

o

LJL

Iz

LLJ

S 

I

m

H 

LLJ

§ 1

LLJi 

LLJ

CD n

O

o

a: 

o

fc

LJJ

X 

O

EC

Z 

LLJ

LJJ 

Q 

X

O

Eun

Ybn

(Ti02)n (K20)n

n

D 1 2 3-0 1 2 3

Ce ENRICHMENT FACTOR (Cen)

FIGURE 1.20. Covariation of incompatible elements in mafic tionation of liquidus phases (olivine, pyroxenes, plagioclase) from

rocks of the Canyon Mountain Complex; incompatible- basaltic magmas causes equivalent enrichment of these elements

element contents of each sample are normalized to their such that normalized abundances will theoretically follow the

abundances in diabase sample CMG162. Closed-system frac- equiline in each plot. (See section "Mafic Rocks" for discussion.)
36 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

between the behavior of these and the more compatible (such as transition metal) elements. It is not clear 

to what extent these reflect varied metasomatic and 

(or) alteration effects as opposed to the emplacement 

of distinct and unrelated batches of magma within the 

sheeted member.

The second question, concerning possible relations 

between Canyon Mountain Complex mafic magmas 

and gabbroic rocks, cannot be resolved definitively in 

this study due to insufficient data. Trace-element 

data for a single cumulate troctolitic gabbro (approximate mode is plag75ol25OpX5; sample OC277, table 

1.3), show that it is characterized by very low REE 

abundances except for a striking positive Eu anomaly 

(fig. 1.21). Assuming that this rock is a cumulate cognate to typical basaltic rocks of the sheeted member, 

we used the method of Shaw (1970) to model REE 

profiles of hypothetical parental liquids required for 

compositional mass balance between the troctolite 

and several basaltic rocks chosen as possible fractionated liquids. The choice of liquid composition is not 

very critical because REE profiles are rather uniform 

for the basalts. Representative model profiles are 

shown in figure 1.21. Despite the latitude in choice of 

model parameters, we conclude that the parental liquids had significant positive Eu anomalies if the troctolite precipitated from the sheeted member basaltic 

magmas. On the other hand, Bishop (chap. 4, this 

volume) analyzed several intrusive noncumulate gabbros that have REE profiles roughly subparallel to, 

but less enriched than, those of the basalts. The gabbros she analyzed clearly have no cumulate relation 

to the basaltic rocks. They could possibly represent 

more primitive precursor magmas, but to test this 

hypothesis requires a combination of data not yet 

available to us. For example, the discrepancy in pyroxene compositions between analyzed basalts and 

gabbros must somehow be resolved before cogenetic 

links can be demonstrated.

Another question related to the petrology of the 

mafic rocks (or their parental magmas) concerns 

their ultimate source. We consider an origin by partial melting within the mantle to be quite reasonable, 

but inverse modeling of the REE data can provide 

more specific constraints on the required characteristics of the mantle source rocks. Using both equilibrium- and fractional-melting equations (Shaw, 1970) 

and a reasonable range in partition coefficients (Leeman and others, 1977; Frey and others, 1978), we 

have estimated the compositions of a variety of possible source rocks. We considered three bulk lithologies: spinel Iherzolite (mode: ol+opx75_g5cpx25_i5), 

feldspathic Iherzolite (ol+opx65_75cpx15_2oplagio_2o), 

and garnet Iherzolite (ol+opX75CpX2o-22gt5-3)- Other

aspects of the calculations follow Leeman (1976) and 

Leeman and others (1977).

A representative set of models is illustrated in figure 1.22. In these models it was assumed that the 

analyzed basaltic rock had undergone as much as 25 

percent fractionation of olivine between the melt-extraction stage and emplacement in the sheeted member. Corrections for this effect lower the REE contents 

of the hypothetical parental magma, but have no effect on its relative REE abundances. From this assumed range of liquid compositions, we calculated

LU

Sz io1

Q 111

0°= 

CC

10

10'.-1

La Ce Sm Eu Tb Yb Lu 

RARE-EARTH ELEMENTS (REE)

Explanation

Calculated residual liquids

O 10 percent crystallization (F=0.9) 

- - 50 percent crystallization (F=0.5) 

Canyon Mountain Complex 

n Troctolite sample OC277

Other gabbros (Mullen, 1983a) 

O Sample Ml 10G 

A Sample M666 

A Sample M707

FIGURE 1.21. Chondrite-normalized rare-earth element (REE) plot 

showing patterns for calculated parental magmas that can produce troctolite cumulate (such as sample OC277) by either 10 

percent or 50 percent fractional crystallization. Except for their 

positive Eu anomalies, these patterns are similar to the range 

for basaltic rocks of the Canyon Mountain Complex (stippled 

band). Shown for comparison are three gabbro samples analyzed 

by Mullen (1983a). F, fraction of original liquid remaining.
L PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON 37

REE profiles required for the sources and residues 

corresponding to formation of between 2 and 25 percent melt. Ranges in these compositions, shown in figure 1.22, illustrate that the source for the basaltic 

magmas was strongly depleted in light REEs, regardless of which model peridotite lithology is adopted; 

details of the modeling have only second-order influence on this conclusion. It appears that this source 

essentially resembles that which produces MORE and

10'

10 1

10"

10

z 1 °"

LJJ

H

O 1C

O

LJJ

LJJ

CC Hri- Q 10

Z

o

O 10"

o

O «n- cc 10

 2.

10

101

10"

10"

10"

CMG32-

Spinel Iherzolite source

CMG32-

Plagioclase Iherzolite source :

Garnet Iherzolite source :

La Ce Sm Eu Tb Yb Lu 

RARE-EARTH ELEMENTS (REE)

that this source has heavy-REE contents of about 1 to 

5 times chondritic abundances.

SUMMARY

The principal results and conclusions derived from 

our petrogenetic and geochemical studies are listed 

below. They should be considered with the caveats that 

we have focused our studies primarily on the upper 

parts of the Canyon Mountain Complex and that much 

remains to be learned about the petrogenesis of the 

ultramafic and gabbroic parts of the complex.

(1) The effects of chemical alteration in Canyon 

Mountain Complex rocks are sufficiently minor 

(when care is taken to obtain good samples) that for 

many elements significant information is retained 

concerning original magmatic compositions.

(2) The basalts underwent different degrees of fractionation and were derived from distinct batches of 

unrelated magma emplaced at different times. For 

the Canyon Mountain Complex as a whole, a minimum of three phases of basaltic magmatism can be 

recognized on the basis of field and (or) geochemical 

criteria; these magmatic phases separately produced 

the deformed metacumulates (earliest phase), the undeformed (or only partly deformed) layered cumulate 

gabbros, and the intrusive isotropic gabbros. The 

sheeted member may also represent distinct phases 

of magmatism.

(3) A cogenetic relationship between sheeted-member basaltic rocks and the gabbros of the undeformed 

part of the complex is tenuous and unresolved at this 

time because of limited mineralogical and geochemical constraints.

(4) Basalts in the upper part of the complex crystallized from magmas formed by partial melting of a 

light-REE-depleted source like that which produces

FIGURE 1.22. Chondrite-normalized rare-earth element (REE) 

plot (heavy solid line) for representative basalt sample 

(CMG32) from the Canyon Mountain Complex, and results 

(shaded areas) of partial melting calculations. Reasonable 

ranges are shown for calculated hypothetical source rocks 

(Source) and ultramafic residue (Residue) after melt extraction (see Shaw, 1970). Model spinel, plagioclase, and garnet 

Iherzolites were used to represent source-rock mineralogy at 

a range of depths appropriate to melting. Shaded areas represent range in calculated REE contents resulting from varied degrees of melting (2-25 percent) and a range in partition 

coefficients (see Leeman and others, 1977; Frey and others, 

1978). As degree of melting increases, source composition 

approaches that of the liquid. Despite wide latitude in these 

calculations, a source strongly depleted in light REEs is 

required to produce the REE profiles characteristic of mafic 

rocks from the Canyon Mountain complex.
38 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

MORE, except for the notable depletion of both Nb 

and Ta (whose concentrations approach analytical detection limits). Such depletion is very unusual for 

MORE and suggests, instead, an island-arc petrogenesis. Light-REE-depleted arclike rocks are indeed 

found in other ophiolites (such as Vourinos, Troodos) 

and in fore-arc regions (for example Japan, Marianas).

(5) Separate petrogeneses are required for the basaltic rocks and for the plagiogranite suite (diorite to 

trondhjemite) to explain the absence of normal fractionation trends and specifically the similar abundances of many immobile incompatible elements.

(6) The plagiogranites most likely represent partial 

melts of hydrated mafic rocks with compositions similar to those of gabbroic rocks in many ophiolites, such 

as those that are analogous to parts of oceanic or marginal-basin crust. Coexistence with the upper basalts 

suggests that the basaltic magmas may have provided 

the necessary heat source to drive crustal-level melting. Thus, magma probably was produced concurrently from two or more, different-depth sources.

(7) The keratophyres into which the Canyon Mountain Complex was emplaced indicate an island-arc 

association judging from the mineralogy and geochemistry of these rocks. They so closely resemble 

the trondhjemites of the plagiogranite suite that a 

similar model of crustal anatexis is considered reasonable. Presumably the keratophyre parent magmas 

formed by melting in the deep-crustal root of a volcanic arc.

TECTONIC IMPLICATIONS

Studies of the Canyon Mountain Complex offer 

some interesting perspectives on the formation of 

ophiolites and on the nature of convergent tectonic 

processes in general. The complex appears to comprise a variety of lithotectonic units that are not 

strictly related to one another in a cogenetic sense. 

Specifically, the tectonite harzburgite and metacumulate members together have characteristics similar to 

those envisaged for oceanic lithosphere. Diapiric ascent of these rocks in one or more stages (Ave Lallemant, 1976; Misseri and Boudier, 1985) is possibly 

analogous to diapirism inferred at spreading centers 

(Casey and Dewey, 1984; Casey and others, 1985; 

Rabinowitz and others, 1987), or in fore-arc regions 

(Bloomer, 1983). The deformation in these rocks could 

be related to primary flowage in an extensional setting or to tectonic factors associated with convergence 

and accretion in a fore-arc region. The undeformed 

isotropic and layered gabbros represent a second 

magmatic phase that did not necessarily occur in the

same place or manner as the first. Primarily on the 

basis of their respective mineral compositions, Mullen 

(1983a, 1985) suggested that these rocks have calcalkaline affinities, whereas the earlier gabbros have 

tholeiitic affinities. These distinctions are rather tenuous, and further work would certainly be useful to 

elucidate the associated tectonic history. Nevertheless, magmas of the sheeted member have definite island-arc affinities, although the basalts are relatively 

primitive. The combined geochemical characteristics 

of these rocks suggest that they may form by fusion of 

fundamentally MORE source material but that the 

distinctive Ta and Nb depletion is related to hydrous 

conditions of melting proximal to a subduction zone 

(Saunders and others, 1980; Wood and others, 1980; 

Pearce, 1982, 1983). The close association of these 

transitional basalts with plagiogranites and keratophyres also suggests proximity to an island arc 

(Gerlach, Ave Lallemant, and Leeman, 1981) onto 

which the remainder of the complex was accreted. 

Another notable feature of the Canyon Mountain 

Complex is the absence of a well-developed sheeteddike complex, a pillow lava sequence, and pelagic sedimentary rocks. Although such units could have been 

removed tectonically, it is possible that they were 

never present. Our interpretation of the sheeted 

member as a sill-like unit (Gerlach, Ave Lallemant, 

and Leeman, 1981), if correct, is consistent with development of the latest magmatic phase under conditions of tectonic convergence, which would have 

favored the stagnation of magma at crustal levels and 

the formation of sills (see Leeman, 1983).

Based on the foregoing discussion, our general impression is that the Baker terrane, of which the Canyon Mountain Complex is a part, resembles a forearc region to which a variety of unrelated oceanic 

lithotectonic blocks accreted during Triassic to Early 

Jurassic convergence. This accretionary phase resulted in telescoping part of an oceanic or marginal basin 

and juxtaposition of unrelated fragments of oceanic 

lithosphere, seamounts, and remnants of island arcs.

This situation is reminiscent of many other ophiolite complexes in which the mafic rocks define two or 

more distinct magma series. Examples include the 

Troodos (McCulloch and Cameron, 1983; Robinson 

and others, 1983; Schminke and others, 1983; Dilek 

and others, 1990), Vourinos (Noiret and others, 1981; 

Beccaluva and others, 1984), Pindos (Capedri and others, 1980), New Caledonia (Dupuy and others, 1981), 

Betts Cove (Coish and others, 1982), Bay of Islands 

(Casey and others, 1985), Papua (Jacques and Chappell, 1980), Oman (McCuUoch and others, 1981; Alabaster and others, 1982), and southern Chile (Stern, 

1980) ophiolites. Multiple stratigraphic sequences
1. PETROLOGY OF THE CANYON MOUNTAIN COMPLEX, EASTERN OREGON 39

that represent distinct accreted lithotectonic blocks 

(including seamounts, oceanic crust, island-arc edifices) are recognized in some of the better studied complexes (such as Troodos and Oman). It has been 

proposed that some of these sequences formed in diverse tectonic settings and were emplaced adjacent to 

one another during accretion of different ophiolites (or 

fragments thereof) in a common collision zone (see 

Capedri and others, 1980). In Papua, New Guinea, 

local geologic evolution was complicated by primitive 

island-arc magmatism that postdated and is unrelated to the main ophiolite complex. Good geochronologic 

control is essential to fully understand the accretionary history of such compound ophiolites.

Despite the inferred polyphase magmatic and tectonic evolution of the Canyon Mountain Complex, exclusive of the basal ultramafic unit, the overall bulk 

composition of the complex is similar to basaltic andesite (Pearcy and others, 1990). This relatively 

mafic composition is consistent with our interpretation that the complex represents a section of immature oceanic to island-arc crust.

The application of geochemical studies to ophiolite 

petrogenesis provides another tool that can help in 

unraveling some of the problems of ophiolite formation. Nd or Hf isotopic studies may be particularly 

useful because these elements are resistant to mobilization during weathering and alteration. Interestingly, Nd data suggest non-MORB affinities for many 

ophiolites: £Nd<+8 for Troodos (McCulloch and Cameron, 1983), Oman (McCulloch and others, 1981), 

Vourinos (Noiret and others, 1981), and other ophiolite complexes. As we have learned in our work on 

the Canyon Mountain Complex, geochemical studies 

need to be integrated with comprehensive geologic 

and geochronologic studies to appreciate the complexities of ophiolite evolution.

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2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH,

ELKHORN MOUNTAINS, AND SOME COMPARISONS WITH THE

WALLOWA BATHOLITH, WALLOWA MOUNTAINS,

NORTHEASTERN OREGON

By WILLIAM H. TAUBENECK1

CONTENTS

Abstract 46 

Introduction 47 

Acknowledgments 52 

Age relations and distribution of rocks near Bald Mountain 

batholith - -- ------ 53

Geographic names 54

Metamorphic aureole of Bald Mountain batholith 54

Ribbon chert __________________________ 54

Cherty argillite 54

In the Belleview wedge 54

South of the Elkhorn pluton 56

Argillite and tuffaceous argillite 57

Limestone____________________________ 58

Metagabbro 59

Ultramafic rocks 61

Structure of Elkhorn pluton - . ....... ._ 62

Peripheral contacts and associated foliation in margins

of pluton ___________________________ 62

Internal foliation 64

Intraplutonic contact 68

Rocks of Bald Mountain batholith 70

Norite of Badger Butte ... 70

Setting and general description 70 

Microscopic description 70 

Tonalite of Bald Mountain and granodiorite of

Anthony Lake 71 

Setting and general description 71 

Comparative modal compositions and

crystallization of augite 72 

Assimilation of graphite-bearing metasedimentary

rocks in Elkhorn pluton 82 

Distinctive petrographic features of rocks of Elkhorn

Specific gravity of rocks of Elkhorn pluton 84 

Compositions of cores of oscillatory-zoned plagioclase 

in Elkhorn pluton and other granitic intrusions of 

northeastern Oregon 84

1Department of Geosciences, Oregon State University, Corvallis, OR 97331-5506

Rocks of Bald Mountain batholith Continued

Tonalite of Bald Mountain and granodiorite of 

Anthony Lake Continued 

Volume percent and origin of calcic cores in

plagioclase of Elkhorn pluton 85 

Interludes: A newly described feature of

oscillatory-zoned plagioclase in granitic rocks 85

Pegmatites and aplites of Elkhorn pluton 87 

Xenoliths of country rocks and inclusions of unknown

origin in Elkhorn pluton 87

Quartz diorite of Limber Creek 91

Setting and general description 91

Microscopic description 92

Leucogranodiorite of Trail Creek 94

Setting and general description 94

Microscopic description 94

Granite of Clear Creek 94

Setting and general description 94

Microscopic description 95

Quartz diorite of Wolf Creek 95

Setting and general description 95

Microscopic description 95

Distinguishing features 96

Boundary quartz diorite unit 96

Setting and general description 96

Microscopic description 96

Distinguishing features 98

Granodiorite of Indiana Mine Road 98

Setting and general description 98

Microscopic description 98

Distinguishing features 98

Granodiorite of Beaver Meadow 99

Setting and general description 99

Microscopic description 99

Distinguishing features 100

Granite of Anthony Butte 100

Setting and general description 100

Microscopic description 101

Distinguishing features 102

Leucogranite of Dutch Creek 102

Setting and general description 102

Microscopic description 102

Distinguishing features 102

Satellitic bodies east of granite of Anthony Butte 103

Tonalite of North Fork 103

45
46 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Satellitic bodies east of granite of Anthony Butte Continued 

Tonalite of North Fork Continued

Setting and general description 103

Microscopic description 103

Distinguishing features 104

Granite of Isham Spring 104

Setting and general description 104

Microscopic description 104

Distinguishing features 104

Granitic rocks of the Guard Station inlier 105

Overview of granitic rocks 105

Method of emplacement of Elkhorn pluton 108

Volatiles contents of granitic magmas in relation to

interstitial zeolites and potassium feldspar in granitic

rocks of the Bald Mountain and Wallowa batholiths - 113

Major-element compositions 118

Conclusions 121

References cited 121

ABSTRACT

The Late Jurassic Bald Mountain batholith, about 450 km2 in 

area, is in the Elkhorn Mountains of northeastern Oregon. It intrudes a melange terrane (Baker terrane) that consists of the Devonian to Triassic Elkhorn Ridge Argillite which is a heterogeneous 

assortment of metasedimentary rocks as well as scattered masses 

of metagabbro and serpentinized ultramafic rocks of uncertain age. 

Contact-metamorphic effects, evident as a wide thermal aureole 

that includes high-temperature minerals characteristic of the pyroxene hornfels facies, overprint a regional (greenschist-facies) metamorphism, characterized by green biotite. The most widespread 

pyroxene hornfels-facies rocks are argillites and metagabbros with 

two-pyroxene assemblages, and cherty argillites with metapelitic 

assemblages characterized by cordierite and orthoclase, with or 

without one or more other minerals such as sillimanite and garnet. 

One notable high-temperature aureole rock is a rare pargasitic 

hornblende-peridotite hornfels that contains small amounts of 

green spinel and plagioclase. The grain size of recrystallized quartz 

in bedded chert near the batholith is a good approximation of the 

intensity of thermal metamorphism.

Forceful intrusion apparently was the dominant process responsible for the emplacement of the batholith. Much wallrock yielded 

by plastic flow. Segregation banding in intensely deformed metagabbro resembles the bedding in some metasedimentary rocks. Intensely deformed salients of country rock that project far into the 

batholith are commonly gneissic and include the Archean gneiss 

unit of Lindgren (1901). Scattered boudins in the gneiss indicate 

extreme plastic deformation. Much of the gneiss apparently is underlain by the batholith.

An assessment of stoping as a significant emplacement mechanism is severely hampered by the absence of critical field evidence 

within the batholith. However, the size differential between xenoliths in the margin and in the interior of the major unit of the 

batholith combined with size relations in the upper parts of small 

cupolas of granitic rocks alongside the batholith imply that stoping 

played a more important role than is commonly recognized in mesozonal plutons.

In the area south of lat 45° N. at least nine discrete intrusive 

units of the Bald Mountain batholith were emplaced in a mafic-tofelsic sequence that commenced with three small gabbroic bodies 

and a small unit of quartz diorite, continued with the very large 

compositionally zoned Elkhorn pluton (tonalite and granodiorite) 

that constitutes most of the batholith, and ended with four small 

felsic units. The tonalite and granodiorite of the compositionally

zoned Elkhorn pluton are separated for about 9 km on the north 

and 4 km on the northwest by an intraplutonic contact, concealed 

for about 4 km between the north and northwest exposures by Tertiary volcanic rocks. Microscopic evidence indicates that augite 

crystallized initially throughout much of the tonalite but only in 

some of the granodiorite in the Elkhorn pluton. The small volume 

percent of calcic cores in oscillatory-zoned plagioclase of the 

Elkhorn pluton indicates that the granitic rocks crystallized from 

a melt that contained only a comparatively minor amount of solid 

constituents. An interlude is the name herein applied to concentrations of one or several mafic minerals of small size in a distinct 

layer of plagioclase between consecutive oscillations in oscillatoryzoned plagioclase. Heretofore, interludes apparently have not been 

reported in the plagioclase of granitic rocks.

Good to excellent planar structure characterizes most margins 

of the Elkhorn pluton except near much of the southern contact. 

Faint to fairly good planar structure that generally is difficult to 

detect occurs sporadically in parts of the interior of the pluton. 

New procedures that take into account realistic (rather than idealized) field relations facilitate recognition of planar structure on 

smooth joint surfaces, exposures controlled by steep diagonal 

joints, and exposures shaped by exfoliation.

Xenoliths of metagabbroic and metasedimentary rocks throughout the central and southwestern parts of the Elkhorn pluton are 

more abundant than xenoliths of country rocks in most batholiths. 

The scarcity of xenoliths of argillaceous rocks as compared to chert 

is attributed to a more rapid rate of assimilation. Granitic rocks in 

the Elkhorn Mountains with reddish-brown to red biotite and little 

or no iron oxide are attributed to assimilation of graphite-bearing 

metasedimentary rocks.

Far more abundant than xenoliths of country rocks are mafic 

inclusions of uncertain origin (see, for example, Pabst, 1928). No 

mafic synplutonic dikes occur in the Bald Mountain batholith at 

the present level of erosion. On the basis of field observations and 

petrographic data, most of the derivations postulated for mafic inclusions in circum-Pacific granitic intrusions can be questioned or 

discounted for mafic inclusions in the Bald Mountain batholith.

A distinct zone in which microcline twinning characterizes the 

potassium feldspar extends for 19 km across the Elkhorn pluton 

and is related to deformation rather than to the concentration of 

volatiles.

Much of this report concerns the part of the Bald Mountain batholith that lies north of lat 45° N. in the Anthony Butte 7^-minute 

quadrangle. Here many small intrusive units of the batholith occur, 

and deformation features of the rocks are more pronounced than in 

the southern part of the batholith. The plutons commonly are separated in part by dismembered screens. One granodiorite body contains small amounts of orthopyroxene and is notable for reaction 

relations between orthopyroxene, cummingtonite, hornblende, and 

biotite. This body is one of two intrusions that contain red-brown 

biotite and little or no iron oxide. Another pluton in the study area 

is the largest granite body in northeastern Oregon and contains 

traces of orthopyroxene and garnet. Several intrusions are composed of silica-rich granitic rocks. Both satellitic intrusions in metasedimentary rocks near this part of the batholith contain either 

red-brown or red biotite and almost no iron oxide.

The mineralogy of the granitic rocks coupled with the pyroxene 

hornfels-facies thermal metamorphism of the wallrocks indicates 

that much of the batholith crystallized from relatively hot and waterundersaturated magmas.

Rocks of the Bald Mountain batholith range in composition 

from norite that contains about 52 weight percent SiO2 to leucogranite that contains about 77 weight percent SiO2- The rocks are 

metaluminous to weakly peraluminous. The compositionally zoned 

Elkhorn pluton ranges inward from border rocks that contain
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 47

about 61 weight percent SiO2 and 0.5 weight percent ^O to core 

rocks that contain about 69 weight percent SiO2 and 2.1 weight 

percent ^O.

Rocks of the Bald Mountain batholith show a typical calc-alkaline compositional trend on an Al2O3-FeO-MgO (AFM) diagram. 

They also show marked late-stage iron enrichment on diagrams 

used to distinguish calc-alkaline from tholeiitic rocks.

INTRODUCTION

The name "Bald Mountain batholith" originated, according to Hewett (1931), with Lindgren (1901), who 

derived it at the turn of the century from the local 

designation "Bald Mountain" for a prominent peak of 

tonalite in the Elkhorn Mountains of northeastern Oregon. In 1917 the U.S. Board on Geographic Names 

officially designated the peak "Ireland Mountain," and 

the local usage of the term "Bald Mountain" gradually 

was discontinued. The new name "Ireland Mountain" 

was soon changed to "Mount Ireland" (fig. 2.1). However, the name that Lindgren informally used for the 

batholith was retained by Taubeneck (1957) mostly in 

recognition of Lindgren's many pioneering contributions, including a remarkable geologic map of much of 

northeastern Oregon and closely adjacent parts of 

Idaho. Lindgren's (1901) classic reconnaissance map,

118° 30' 118° 00' 117° 30' 117° 00'

45' 

30'

45' 

00'

44' 

30'

Approximate

boundary of

Eagle Cap

Wilderness

Area

Of JO 

Cornucopia &J ^9 a? si

AREA OF MAP

IZEE TERRANE

OLDS 

FERRY 

TERRANE -

I_____

10 20 30 40 50 KILOMETERS

FIGURE 2.1. Index map showing major geographic features of 

northeastern Oregon. Dashed lines are terrane boundaries 

(from Vallier, chap. 3, this volume), queried where uncertain.

based on approximately three months of field work, 

encompasses about 17,550 km2 , has a contour interval 

of 500 ft, accurately indicates the peak then known as 

Bald Mountain, and shows much of the associated 

batholith. The accomplishments of Lindgren in the 

Blue Mountains region strongly suggest that he is 

among the most perceptive of the great American reconnaissance geologists.

The 450-km2 Bald Mountain batholith (figs. 2.2 

and 2.3) is mostly in the central part of the Elkhorn 

Mountains. About 78 percent of the batholith is south 

of lat 45° N.; here exposures generally are good to 

excellent except in some of the westernmost and 

easternmost parts of the batholith. Contacts of the 

batholith north and east of Mount Ireland (fig. 2.2) 

commonly are concealed by glacial deposits. The 

batholith is mountainous south of lat 45° N. with elevations commonly greater than 2,375 m; total relief 

is 1,555 m. Prior to 1965, no map was available for 

the area north of lat 45° N. and only a 1901 map (of 

the Sumpter 30-minute quadrangle) at a scale of 

1:125,000, was available for the area south of lat 45° 

N. Modern l:24,000-scale topographic maps for the 

Elkhorn Mountains and adjacent areas were not 

available until the 1960's and 1970's when the U.S. 

Geological Survey published the Anthony Butte 

(1965), Limber Jim Creek (1965), Little Beaver Creek 

(1965), Marley Creek (1965), Tucker Flat (1965), Anthony Lakes (1972), Bourne (1972), Crawfish Lake 

(1972), Elkhorn Peak (1972), Granite (1972), Mt. Ireland (1972), Rock Creek (1972), and Trout Meadows 

(1972) 7/2-minute quadrangles (fig. 2.2A).

South of lat 45° N., the composite Bald Mountain 

batholith consists of at least nine discrete intrusive 

units emplaced in a mafic-to-felsic sequence. Intrusive activity commenced with the emplacement of 

three small gabbroic bodies and a small unit of 

quartz diorite, continued with the intrusion of a tonalite and granodiorite body herein called the 

Elkhorn pluton, which constitutes most of the batholith, and ended with the injection of four small felsic 

units. The three gabbroic plutons consist of the norite 

of Willow Lake, the norite of Badger Butte, and the 

quartz gabbro of Black Bear; the small unit of quartz 

diorite is the quartz diorite of Limber Creek (fig. 2.4). 

The Elkhorn pluton is compositionally zoned from a 

mafic margin (tonalite of Bald Mountain) to a felsic 

core (granodiorite of Anthony Lake). The four felsic 

units are the leucogranodiorite of Mount Ruth, the 

leucogranodiorite of Red Mountain, the leucogranite 

of Elk Peak, and the leucogranodiorite of Trail Creek. 

Seven of the units listed above were identified in an 

earlier report (Taubeneck, 1957), and the names used 

here are modified from that report. Two of the
48 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

units the quartz diorite of Limber Creek and the 

leucogranodiorite of Trail Creek are newly mapped

and named. The rock name for the Elk Peak unit has 

been changed from leucocratic quartz monzonite

118° 22' 30" 118° 15' 00"

AREA OF MAP

118°00'

118° 00' 00"
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 49

(Taubeneck, 1957) to leucogranite in accordance with 

the classification of Streckeisen (1973).

This chapter describes the metamorphic aureole 

and structure of the Elkhorn pluton, then updates

EXPLANATION

Alluvium (Quaternary) 

Volcanic rocks (Tertiary) 

Jep "M Leucogranite of Elk Peak (Late Jurassic)

Jccl Granite of Clear Creek (Late Jurassic)

.i i i_i

r^Jabjn Granite of Anthony Butte (Late Jurassic)

PI iiin

f^\JmrXfj Leucogranodiorite of Mount Ruth (Late Jurassic) 

Leucogranodiorite of Red Mountain (Late Jurassic) 

Leucogranodiorite of Trail Creek (Late Jurassic) 

Jbm Tonalite of Bald Mountain (Late Jurassic) 

Jal I Granodiorite of Anthony Lake (Late Jurassic)

Norite of Willow Lake (Late Jurassic) 

I Jbb_ N>] Norite of Badger Butte (Late Jurassic)

jb ;.;.;| Quartz gabbro of Black Bear (Late Jurassic)

A iiniiiiiiuiii B

D IIIIIIIIIIMIII E

Metagabbro and Elkhorn Ridge Argillite, undivided

(Triassic and Paleozoic)

118° 22'30" 118° 15'00" 118° 07'30" 118°00'00"

Contact-Dashed where intraplutonic 

Gradational igneous - facies boundary

Boundary Divides rocks of same lithology but 

differing mineralogy; possible igneous 

intrusive contact

Southwest edge of zone of distinct eastward and 

northeastward decrease in xenolith population 

within Monumental salient

Strike and dip of bedding 

Inclined 

Vertical

Strike and dip of foliation 

Inclined 

Vertical

FIGURE 2.2. Bald Mountain batholith. A, Generalized 

geology of all but northeastern part of Bald Mountain batholith. Except in area south and southwest 

of Hunt Mountain, contacts of the batholith with 

country rocks were mapped anew (since 1950's) on 

7>2-minute quadrangles shown in upper right corner. Within batholith, contacts of small intrusions 

south and southeast of Anthony Lake were mapped 

in 1950's on Sumpter 30-minute quadrangle (1901) 

and are only approximate. B, Topographic and cultural features of Elkhorn Mountains and extent of 

Bald Mountain batholith (heavy line is contact between batholith and country rocks). Also shown are 

named salients and wedges in southern part of Bald 

Mountain batholith.

02468 KILOMETERS 

I I I I I I I I I
50 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

118° 15'00" 118° 07'30"

3 KILOMETERS

EXPLANATION

Volcanic rocks (Tertiary) SATELLITIC ROCKS OF BALD MOUNTAIN BATHOLTTH

ROCKS OF CENTRAL PART OF BALD MOUNTAIN BATHOLITH 

Tonalite of Bald Mountain (Late Jurassic)

ROCKS OF PERIPHERAL PART OF BALD MOUNTAIN BATHOLITH

Granite of Isham Spring (Late Jurassic) 

i Jnf ^ Tonalite of North Fork (Late Jurassic)

PREBATHOLJTHIC ROCKS

Metagabbro and Elkhorn Ridge Argillite, 

undivided (Triassic and Paleozoic)

Contact-Dotted where concealed 

Strike and dip of bedding

Leucogranite of Dutch Creek (Late Jurassic)

Granite of Anthony Butte (Late Jurassic) 

 Locally contains large quartz diorite 

inclusions (qd)

Granodiorite of Beaver Meadow (Late Jurassic)

Granodiorite of Indiana Mine Road (Late 

Jurassic)

Boundary quartz diorite unit (Late Jurassic) 

Quartz diorite of Wolf Creek (Late Jurassic)

FIGURE 2.3. Geology of northeastern part of Bald Mountain batholith.
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 51

petrographic data for the norite of Badger Butte, the 

tonalite of Bald Mountain, and the granodiorite of 

Anthony Lake before describing the newly identified 

quartz diorite of Limber Creek and leucogranodiorite 

of Trail Creek.

The batholith north of lat 45° N. was not suitable 

for field studies until the 1960's because of dense 

timber that concealed the limited exposures, poor access, and the lack of topographic maps. About 80 km2 

of forest in the Anthony Butte quadrangle (fig. 2.2A) 

burned during 1960 in a wildfire known as the Anthony Burn. Moss and lichens that originally covered 

most exposures could not survive direct sunlight and 

reduced moisture. Formerly concealed exposures 

within the burn became apparent. In addition, much 

new bedrock was exposed along more than 180 km of

118° 17'30' 118° 15'00"

3 -".I ^ **^^ ec-N v *t ^

2 KILOMETERS

. Og .

EXPLANATION

Alluvium (Quaternary)

Glacial deposits (Quaternary)

Tonalite of Bald Mountain (Late Jurassic)

Quartz diorite of Limber Creek (Late Jurassic)

Metagabbro and Elkhorn Ridge Argillite, 

undivided (Triassic and Paleozoic)

Contact Dashed where approximately 

located, dotted where inferred

Strike and dip of beds 

Inclined 

Vertical

Strike and dip of foliation 

Inclined 

Vertical

Strike and dip of joints 

Inclined 

Vertical

FIGURE 2.4. Geology of area between Monumental salient (see fig. 2.2B) and central part of Bald Mountain batholith. Arrows pointing 

toward breccia indicate east and west extent of brecciated quartz diorite of Limber Creek.
52 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

roads that were constructed during and after the fire 

for purposes of firefighting, reforestation, erosion control, and timber salvage. Areas outside the Anthony 

Burn remain in timber and are poorly exposed, especially the southeastern part of the Limber Jim Creek 

quadrangle (fig. 2.2A), where at least 95 percent of 

the scattered exposures in an area of 20 km2 are covered with moss and lichen. Thus, the fire and new 

roads were critical factors in creating adequate exposures to determine many geologic relations.

Much of this chapter concerns an area (fig. 2.3) in 

the Anthony Butte quadrangle where many heretofore-unreported intrusive units of the Bald Mountain 

batholith occur within and near the Anthony Burn. 

These units are mostly felsic; some have a high quartz 

content. The largest pluton consists of the hereinnamed granite of Anthony Butte, which is by far the 

most extensive granite in northeastern Oregon.

A major purpose of this chapter is to increase the 

data base for rocks of the batholith and to provide a 

reference for future studies. Ten new major-element 

analyses are included, but petrogenetic interpretations for the array of rocks in the batholith are not 

possible without additional data from trace elements 

and isotope geochemistry.

Mineral proportions (by volume) were determined 

by the point-counter method (Chayes, 1949), tallying 

at least 2,400 points for each of two or three large 

thin sections (generally 800-960 mm2) of each analyzed rock. Two thin sections each were used for 

modal analysis of norite, quartz diorite, the granodiorite of Anthony Lake, and all tonalite except in the 

inlier 1.5 to 2 km southeast of the Grande Ronde 

Guard Station; three thin sections each were used for 

all other rocks. Classification of the rocks is based on 

the IUGS system (Streckeisen, 1973), except that a 

distinction between tonalite and quartz diorite was 

not made in the border rocks of the tonalite of Bald 

Mountain.

Specific-gravity determinations for rocks of the tonalite of Bald Mountain and the granodiorite of Anthony Lake were made on a beam balance by 

balancing the specimen in air and in water and 

measuring the change in length of the lever arm. 

Measurements are reproducible with ±0.01 precision.

Refractive indices of minerals in the metamorphic 

aureole were determined on suitably oriented grains 

by the immersion method using a sodium vapor lamp. 

After each determination, the immersion liquid was 

tested on a refractometer. The accuracy of refractiveindex measurements is regarded as ±0.002.

Optic-axial angles for potassium feldspar were obtained (except where stated as estimates) with a fourring universal stage by direct rotation (from one optic

axis to the other) of crystals showing nearly centered 

acute-bisectrix figures.

Granitic rocks of the Bald Mountain batholith commonly are compared in this chapter with those of the 

Wallowa batholith. The Wallowa batholith is within 

the Wallowa Mountains (fig. 2.1), called in Oregon 

the "Swiss Alps of America." The peaks of the Wallowa Mountains and intervening glaciated canyons as 

much as 1,375 m in depth contribute to many superior exposures of granitic and metasedimentary rocks. 

The two batholiths and the Cornucopia stock, the 

major satellite of the Wallowa batholith, dominate 

the Late Jurassic granitic rocks of northeastern Oregon. Almost all of the Wallowa batholith and most of 

the Cornucopia stock are within the approximately 

1,450-km2 Eagle Cap Wilderness Area (fig. 2.1).

Fieldwork totaling 14.5 months in the Elkhorn 

Mountains was done in 1955, 1956, 1966, 1970, 1971, 

and 1983 through 1987. Fieldwork totaling 24 months 

in the Wallowa Mountains was done from 1956 

through 1964 and from 1966 through 1969. The work 

in both mountain ranges was done alone.

ACKNOWLEDGMENTS

The writer thanks the Department of Geosciences 

of Oregon State University for office space and clerical support.

Tracy L. Vallier, Howard C. Brooks, Ellen M. Bishop, and Keith F. Oles provided many helpful suggestions for improving the manuscript.

A.J. Piwinskii provided indispensable analytical assistance at the University of Chicago in 1967 in microprobe determination of the composition of cores of 

oscillatory-zoned plagioclase feldspars. Additional 

core compositions were determined by Ellen M. Bishop at Washington University in St. Louis in 1983.

The writer also gratefully acknowledges the many 

courtesies of Howard C. Brooks and Mark L. Ferns of 

the Baker City office of the Oregon Department of 

Geology and Mineral Industries. During recent years 

these courtesies have ranged from the temporary 

storage of rock specimens and field equipment to discussions of the nineteenth- and early twentieth-century history of the Elkhorn Mountains and adjoining 

areas of northeastern Oregon.

Support for studies in the Elkhorn Mountains during 1970 and 1971 from National Science Foundation 

Research Grant GA-31852 is acknowledged with special gratitude. Studies of granitic rocks in the Wallowa Mountains were supported during the period 

1956 through 1969 by National Science Foundation 

Grants G2415, G14553, GP-472, and GA-427.
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 53

AGE RELATIONS AND DISTRIBUTION OF ROCKS 

NEAR BALD MOUNTAIN BATHOLITH

The Bald Mountain batholith intrudes a widespread and heterogeneous assortment of metasedimentary rocks that compose the Elkhorn Ridge 

Argillite (Gilluly, 1937) as well as scattered masses of 

metagabbro with small amounts of serpentinized ultramafic rocks.

The Elkhorn Ridge Argillite consists of predominant 

argillite, tuffaceous argillite, cherty argillite, and ribbon chert with small amounts of tuff, graywacke, massive chert, limestone, and rare conglomerate. Fossils 

of Pennsylvanian, Permian, and Triassic age occur 

within limestone pods in the Elkhorn Ridge Argillite 

(Brooks and others, 1982) and indicate that the different limestones are parts of a complex lithologic mix of 

similar-appearing rocks. The Elkhorn Ridge Argillite 

has persistently defied attempts to stratigraphically 

subdivide it. Its most northerly exposures, heretofore 

unmapped, are located in the northern part of the

inlier of pre-Tertiary rocks (fig. 2.5) in the Little Beaver Creek quadrangle (fig. 2.2A) and are part of the 

Baker terrane (Silberling and others, 1984).

The ages of the metagabbro and associated ultramafic rocks within the aureole are uncertain. Metagabbro 

exposed 50 km southeast of the batholith is at least 

241 Ma (Brooks and others, 1982), whereas gabbroic 

rocks about 70 km southwest of the batholith are at 

least 278 Ma (Walker and Mattinson, 1980).

Small masses of previously unreported ultramafic 

rocks are associated with metagabbro and schistose 

amphibolite north of lat 45° N. in the Tucker Flat 

and Anthony Butte quadrangles (fig. 2.2A). These 

masses, mostly in country rocks near the southeast 

part of the granite of Anthony Butte (fig. 2.3), are 

among the northernmost exposures of ultramafic 

rocks in Oregon and are part of the Baker terrane.

K-Ar and Rb-Sr ages and Sr isotopic compositions 

(Armstrong and others, 1977) of rocks from the Bald 

Mountain batholith suggest emplacement of the batholith about 160 Ma in a volcanic-arc environment.

118° 22' 30' 118° 20' 00"

Grande Ronde Guard Station

2 KILOMETERS

Jn

EXPLANATION

Alluvium (Quaternary) 

Volcanic rocks (Tertiary)

GRANITIC ROCKS OF THE GUARD STATION INLIER

Northern biotite granodiorite (Late Jurassic)

> T Jc

-L

Central quartz diorite (Late Jurassic)

Southeastern biotite granodiorite and biotite 

tonalite (Late Jurassic)

Southern quartz gabbro and quartz diorite 

(Late Jurassic)-Divided into:

Quartz diorite 

Quartz gabbro

PRE-JURASSIC ROCKS

Elkhorn Ridge Argillite (Triassic and 

Paleozoic)

Metagabbro

 Contact Dotted where concealed or inferred 

Z£ Strike and dip of bedding

FIGURE 2.5. Reconnaissance geology of Guard Station inlier. Many very small granitic bodies, including at least two consisting of

granite, are not shown on map.
54 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

The northern part of the batholith is overlain by 

Miocene Grande Ronde Basalt in the Tucker Flat and 

Anthony Butte quadrangles (fig. 2.2A) and by older 

Tertiary volcanic rocks (Walker, 1973) to the west in 

the Limber Jim Creek quadrangle. The older volcanic 

rocks extend south of lat 45° N. into the Crawfish 

Lake quadrangle, where they conceal the west contact 

of the batholith with country rocks for about 6.5 km.

GEOGRAPHIC NAMES

For reference purposes certain parts of the batholith 

and adjacent country rocks have been assigned geographic names (fig. 2.2B). The westward-projecting 

mass of tonalite in the southwestern part of the batholith is known as the Monumental salient (Hewett, 

1931). The wedge of country rocks that separates the 

north side of the Monumental salient from the central 

part of the batholith herein is called the Belleview 

wedge. The eastward-projecting mass of tonalite in 

the southeastern part of the batholith is designated 

the Hunt Mountain salient (Taubeneck, 1957). The 

wedge of country rocks that separates the north side 

of the Hunt Mountain salient from the central mass of 

the batholith is called the Muddy Creek wedge 

(Taubeneck, 1957).

METAMORPHIC AUREOLE OF BALD MOUNTAIN 

BATHOLITH

The well-developed contact aureole of the Bald 

Mountain batholith includes thermally metamorphosed ultramafic rocks, metagabbro, metadiorite, 

greenstone, graywacke, limestone, and chert. Argillaceous rocks of the Elkhorn Ridge Argillite metamorphosed to pyroxene hornfels facies and including 

argillite, cherty argillite, and tuffaceous argillite 

(Pardee and Hewett, 1914) are the predominant 

type of country rock. Ribbon chert and cherty argillite are the main rock types in the Belleview wedge, 

whereas metagabbro composes nearly all of the Muddy Creek wedge (fig. 2.2B). The widespread high-temperature assemblages in argillaceous rocks, impure 

limestone, metagabbro, and serpentinite near the 

contact of the batholith are attributed to the emplacement of hot and relatively dry tonalitic magmas.

Aureole rocks are considered in the order of ribbon 

chert, cherty argillite, argillite and tuffaceous argillite, 

limestone, metagabbro, and ultramafic rocks. The metamorphic mineralogy of cherty argillite is characteristic of pelitic assemblages, whereas the mineralogy of 

argillite and tuffaceous argillite is representative of

basic assemblages. Field and laboratory data are 

mostly a summary of work done in the 1950's.

RIBBON CHERT

The grain size of recrystallized quartz in ribbon chert 

near the batholith generally is a good field approximation of the intensity of thermal metamorphism. Ribbon 

chert is moderately common throughout much of the 

aureole and locally is the dominant rock type. The 

rhythmic layers, mostly less than 7 cm thick, are separated by thin argillaceous partings. Quartz crystals 

in chert within 150 m of the contact are glassy and 

commonly 1 to 3 mm across in hand specimens. Grain 

size progressively decreases away from the batholith 

and generally is less than 0.5 mm in rocks more than 

350 m from the contact.

Rocks of the Belleview wedge provide an example 

of the correlation of grain size in ribbon chert with 

intensity of thermal metamorphism. Ribbon chert 

and cherty argillite that are plastically deformed and 

intensively recrystallized characterize all but the 

southwestern part of the Belleview wedge. The grain 

size of most of the chert exceeds 0.5 mm and suggests metamorphism at temperatures of and above 

hornblende hornfels facies.

The correlation of grain size in ribbon chert and 

intensity of thermal metamorphism is evident elsewhere in northeastern Oregon in rocks of the Baker 

terrane. For example, the grain size of recrystallized 

quartz in ribbon chert is a useful guide to the proximity of the west and southwest contacts of the Big 

Lookout Mountain stock, which is about 68 km eastsoutheast of Hunt Mountain, and of the south and 

southeast contacts of the Pedro Mountain stock, 

which is about 64 km southeast of Hunt Mountain 

(figs. 2.1 and 2.2).

CHERTY ARGILLITE

High-temperature mineral assemblages in metamorphosed cherty argillite in the Belleview wedge differ in mineralogy from assemblages in cherty argillite 

south of the Elkhorn pluton. Moreover, gneiss derived 

primarily from cherty argillite extends throughout 

most of the Belleview wedge, whereas the distribution 

of metasedimentary gneiss is restricted elsewhere in 

the aureole. Accordingly, metamorphosed cherty argillite of the Belleview wedge and of the aureole south of 

the batholith are described separately.

IN THE BELLEVIEW WEDGE

Metasedimentary rocks in an 18-km2 area of intense 

metamorphism comprising most of the Belleview
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 55

wedge were justifiably regarded as Archean gneiss by 

Lindgren (1901, p. 577), who correctly reported that 

the rocks have no counterparts in northeastern Oregon. Lindgren (1901, p. 594) rejected "the supposition 

of contact-metamorphic origin" for the gneiss because 

"it is completely recrystallized over a large area." The 

rock type and age assignment of Lindgren (1901) underscore the unique features of the metasedimentary 

rocks. The gneiss is mostly an intensely metamorphosed cherty argillite but includes some very argillaceous ribbon chert that also was converted into gneissic 

rocks. Metamorphosed ribbon chert typically is composed of more than 90 to 95 percent quartz and therefore occurs either as small and distinct quartz-rich 

areas within the gneiss or as much larger areas, many 

meters in width, that alternate with gneiss. The visible 

difference in quartz content between areas of metamorphosed ribbon chert and gneiss is accentuated by the 

common retention of very crude bedding (the original 

argillaceous partings) in the ribbon chert, in contrast 

to the obliteration of sedimentary bedding in the 

gneiss. Foliation in the gneiss is mostly defined by 

wispy streaks and lenticular concentrations of biotite 

as much as 3 cm long. Assorted boudins in the gneiss 

commonly intensify the foliation. The gneiss is not 

migmatitic or cut by leucocratic veins, except locally by 

small pegmatites that contain tourmaline.

Modal analyses of ten rock specimens of gneiss gave 

the following range in volume percent for the five 

major minerals: quartz, 41.1 to 74.5; plagioclase, 7.1 to 

29.6; biotite, 8.0 to 23.4; orthoclase, 0.3 to 8.8; and 

cordierite, 2.1 to 7.3. Significant minerals present in 

some but not all rock specimens are sillimanite, andalusite, and garnet. Muscovite (from 0.3 to 2.4 percent 

by volume) occurs in eight rocks and is mostly a replacement of cordierite and andalusite, as well as of 

orthoclase and plagioclase. Some muscovite in three 

rocks could be primary. Accessory minerals generally 

include fairly constant amounts of apatite, tourmaline, 

monazite, and iron oxides, but widely variable quantities of graphite (from 0.1 to 4.2 percent by volume).

The overall grain size of the gneiss is approximately 

1 mm; quartz and cordierite have maximum dimensions of about 3 mm, whereas biotite and plagioclase 

have maximum dimensions of only about 1.5 mm. Biotite and cordierite in strongly deformed rocks exhibit 

marked parallelism to foliation, as do most of the 

small aggregates of graphite (as much as 0.25 mm in 

length) that typically occur along grain boundaries. 

The most extreme elongation of cordierite is about 10 

to 1. In strongly deformed rocks some quartz that is 2 

to 3 mm in greatest dimension and parallel to foliation has length-to-width ratios of as much as 4 to 1 

and rarely 5 to 1.

All 21 rock specimens of gneiss were collected from 

about 9 to 360 m from the batholith. The diagnostic 

high-temperature pair of cordierite and orthoclase occurs with quartz, plagioclase, and red-brown biotite 

in each of the 21 rock specimens. The potassium feldspar is orthoclase, as determined on the flat stage by 

estimated optic-axial angles (2V) of about 55° for 

nearly centered acute-bisectrix figures. Zoning in the 

plagioclase is a clear indicator of disequilibrium. 

Rather strong normal zoning predominates, but locally as many as four oscillations occur.

Sillimanite, fibrolite, garnet, and andalusite occur 

in two or more of the 21 rock specimens of gneiss. 

Sillimanite is the name applied here to prisms that 

are approximately 0.01 mm or more in width, whereas fibrolite denotes much thinner prisms that occur 

mostly as fibrous bundles, sprays, and mats. Sillimanite is present in two rocks, 9 and 30 m from the 

batholith. Garnet and fibrolite each occur in six rocks 

from 9 to about 360 m from the batholith; two of the 

rocks contain both minerals. Garnets are mostly from 

0.05 to 0.5 mm across; a few are as much as 1.0 mm 

across. Crystals less than about 0.1 mm are commonly euhedral to subhedral, whereas larger garnets are 

mostly irregular, commonly perforated, and frequently characterized by a core of opaque inclusions. Andalusite, present in eight rocks from 20 to about 345 m 

from the batholith, is a relict mineral generally composing not more than 0.2 percent by volume of each 

rock. Simultaneous extinction of relicts in three thin 

sections indicates that andalusite originally occurred 

as crystals of at least 0.5 to 1.0 mm in size. The 

relicts also suggest that andalusite was of wide occurrence during the early history of the gneiss. Significantly, in the extreme southwestern part of the 

Belleview wedge, beyond the area of gneiss, euhedral 

porphyroblasts of andalusite occur about 1 km from 

the batholith.

No specimens of gneiss were collected farther than 

360 m from the contact with the batholith, so no report 

of metapelitic assemblages in the interior of the Belleview wedge is possible in this chapter. Some specimens 

of basic metasedimentary rocks, however, provide 

information about metamorphic intensities in the interior of the wedge. Basic metasedimentary rocks containing amphibole and (or) pyroxene are rare in the 

Belleview wedge, but they do occur very near the center 

of the wedge, approximately 1.9 km from both the south 

and north contacts of the batholith. The rocks are well 

foliated and contain hornblende commonly from 0.5 to 

1.0 mm in greatest dimension. Among five rock specimens that were collected, one contains brown hornblende, plagioclase, diopside, and quartz, whereas 

three contain brown hornblende and plagioclase with
56 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

small amounts of both biotite and quartz. The fifth rock 

specimen contains cummingtonite, plagioclase, and biotite with minor quartz and brown hornblende. The 

five rock specimens from the center of the Belleview 

wedge are very similar to cummingtonite and brown 

hornblende schists found within 100 m of the contact 

on the south side of the batholith where they are in the 

upper part of the hornblende hornfels facies.

SOUTH OF THE ELKHORN PLUTON

Contact metamorphism of cherty argillite south of 

the Elkhorn pluton is superimposed on a regional 

greenschist-facies metamorphism characterized by 

green biotite. The aureole is exposed in its entirety 

only on the south side where country rocks extend 

outward for 14 km from the contact of the pluton 

without being concealed by Cenozoic volcanic rocks or 

wide alluvial valleys. The outer limit of the aureole is 

difficult to determine in argillaceous rocks because 

the distinction between green and brown biotite is 

not optically possible in very fine grained metasedimentary rocks. Brown biotite is much easier to identify in metagabbro near the aureole boundary; it 

occurs in chlorite in metagabbro 6.5 km from the 

Elkhorn pluton.

Proceeding toward the batholith in the zone of 

brown biotite, garnet is present about 2,500 m from 

the contact but is not common within the aureole, 

whereas cordierite is common and occurs at least 

1,320 m from the contact. Orthoclase is found at 

least 1,270 m from the contact. Inadequate rock specimens of cherty argillite prevent a more precise determination of the first appearance in the aureole of 

garnet, cordierite, and orthoclase.

Grain size in metamorphosed cherty argillite south 

of the batholith varies considerably even within the 

same thin section. Within a given thin section, grain 

size is determined by the ratio between quartz and 

argillaceous material, including graphite. For example, a quartz-rich layer or lens in a thin section commonly will have a grain size several times larger 

than that of associated argillaceous layers and lenses. Grain size decreases as the amount of biotite and 

graphite increases. Thus grain size in the metapelitic 

cherty argillite cannot be correlated as closely with 

distance from the batholith as can the size of quartz 

crystals in metamorphosed ribbon chert. The general 

range in grain size for most rock specimens of cherty 

argillite within 15 m of the contact is 1 to 2 mm for 

quartz, 0.2 to 0.5 mm for biotite, and 0.4 to 1.3 mm 

for orthoclase, plagioclase, and cordierite.

Petrographic data in this paragraph apply to 71 

aureole specimens of cherty argillite, all collected approximately S. 10° E. of Summit Lake (fig. 2.2) along 

a traverse of 1,090 m that commenced at the contact 

of the batholith. Orthoclase, cordierite, quartz, and 

biotite are the major rock components. Plagioclase is 

present only in small amounts, if at all, and exhibits 

good normal zoning as far from the intrusive contact 

as 125 m. The potassium feldspar is orthoclase as determined (for samples from contact outward for 410 

m) on the flat stage by estimated 2V from about 50° 

to 60° for nearly centered acute-bisectrix figures. Cordierite is elongated parallel to foliation; elongation 

ratios are most pronounced in schists between 240 

and 620 m from the batholith. Four very strongly foliated schists contain cordierite with elongation ratios ranging from 5 to 1 to as much as 10 to 1; ratios 

of as much as 12 to 1 occur in two other schists. 

Within about 575 m of the batholith, muscovite occurs only in small amounts, .but it is much more 

abundant farther from the intrusive contact. Most 

muscovite within half a kilometer of the batholith is 

an alteration of cordierite, whereas many flakes of 

muscovite farther than 575 m from the contact are 

independent of cordierite. Muscovite not associated 

with cordierite commonly crosscuts foliation; this relation implies a retrograde origin for the muscovite.

Biotite is widely distributed throughout the aureole 

and allows investigation of the possibility that the 

mineral composition of biotite changes progressively 

with increasing metamorphism. Pitcher and Sinha 

(1957) showed that changes in composition of biotite 

in the aureole of the Ardara pluton, Northern Ireland, coincide with systematic variations in refractive 

index. Biotite from the aureole of the Bald Mountain 

batholith was concentrated from 72 rocks that were 

collected south of the batholith from within 950 m of 

the contact. The P index for biotite from rocks within 

230 m of the contact is generally between 1.637 and 

1.655, whereas the index is mostly between 1.607 

and 1.637 for rocks farther from the contact. The index for biotite from 12 of the 72 rocks is inconsistent 

with these values. Therefore, the data show only an 

approximate correlation between refractive index of 

biotite and metamorphic intensity. The higher indices 

that characterize most biotite within 230 m of the 

batholith presumably denote increases in ferrous iron 

and titanium in the biotite (Pitcher and Sinha, 1957).

Graphite is much more abundant in the aureole 

rocks than is apparent from the examination of thin 

sections because it occurs as small flakes along grain 

boundaries. The mineral commonly is rather inconspicuous partly because the thin sections are made at 

right angles to the planes of schistosity and bedding 

in which the flakes lie. Accordingly, the small crystals 

in general show the minimum cross-sectional area
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 57

and are easily overlooked in many rocks that contain 

comparatively small amounts of graphite. An estimated 20 percent of the argillaceous rocks (argillite 

as well as cherty argillite) collected south and southeast of Summit Lake contain more than 2 percent 

graphite. The presence of graphite in argillaceous 

rocks can be verified by placing -100-mesh powder of 

crushed rock in a large glass beaker and vigorously 

introducing tap water, allowing the larger grains to 

settle from the suspended minute particles, and decanting and repeating the process. Where graphite is 

present, a grayish-black scum of the mineral will collect on the top of the water in the beaker.

Rutile also is more abundant than apparent in thin 

sections, in which it is discerned as small crystals 

that mostly appear opaque in ordinary light and commonly are masked by flakes of associated biotite. The 

crystals of rutile are best seen petrographically in 

convergent light. The abundance of rutile became evident after the heavy minerals were extracted from 

the 72 rocks that were crushed for the separation of 

biotite. Knee-shaped twins of rutile are common 

among the heavy minerals; rutile probably is more 

widespread in metamorphosed argillaceous rocks 

than has been realized.

Sporadic tourmaline is seen in thin sections of 

rocks from near the contact of the Elkhorn pluton to 

beyond the aureole. Much of the tourmaline is in biotite-rich horizons within cherty argillite. The mineral 

was seen neither in microfractures (in thin sections) 

nor as megascopic crystals in veins and along joints. 

Concentrations of tourmaline were obtained from 21 

rocks within 20 to 2,400 m of the contact, but no pattern was recognized in the indices of refraction, distribution, or color of tourmaline throughout this part 

of the aureole. The distribution and occurrence of 

tourmaline suggest that it is not a metasomatic byproduct of granitic emanations. Presumably boron in 

the argillaceous sediments contributed to the growth 

of tourmaline.

Remarkable concentrations of apatite occur in a 

thin section of a cherty argillite collected southeast of 

Summit Lake and about 20 m from the near-vertical 

contact of the batholith. The largest concentration of 

apatite is confined to a quartz-rich area of 24 mm2 

that contains 451 apatite crystals mostly 0.01 to 0.20 

mm across. Apatite constitutes 16.2 modal percent of 

the rock within the quartz-rich area. Two hundred 

ninety-three crystals of apatite occur along three 

chainlike stringers that represent annealed fractures. 

The other 158 crystals are restricted to grain boundaries between quartz. The second eye-catching occurrence of apatite is a lens-shaped pod packed with 93 

crystals; the pod is about 1.0 mm long and as much

as 0.35 mm wide. Two more concentrations of apatite 

are associated with continuous sequences of biotite 

that define recrystallized argillaceous material. One 

occurrence consists of 23 closely spaced crystals within 2.0 mm, whereas the other has 9 chainlike crystals within a distance of 0.5 mm.

Two additional concentrations of apatite are present 

in a thin section of a cherty argillite also collected 

southeast of Summit Lake but about 35 m from the 

batholith. Both occurrences are veinlike and parallel 

to the foliation, and they extend mostly along grain 

boundaries of quartz. Locally, however, one or more 

comparatively large crystals of quartz are crossed by 

the veins. One vein containing about 560 small crystals of apatite extends from the edge of the thin section for about 12 mm into the interior of the section. 

The second vein contains an estimated 700 crystals of 

apatite and extends across the width of the thin section for about 16 mm. All concentrations of apatite in 

the two rock specimens are interpreted as of pneumatolytic origin.

Index determinations for apatite from 26 aureole 

rocks within 1,000 m of the batholith range from 

1.633 to 1.640 for ne and from 1.637 to 1.643 for nw; 

these values are essentially the same as indices for 

apatite extracted from four widely scattered tonalite 

specimens from the batholith. All indices are characteristic of fluorapatite.

Garnet, from about 0.02 mm to rarely as much as 

0.5 mm across, occurs in 12 rock specimens from 3 to 

about 2,500 m from the batholith. Much of the garnet 

is irregular, but nearly all crystals less than about 0.1 

mm in four rock specimens are subhedral. Crystals in 

half of the rock specimens are commonly perforated 

with voids as much as 0.02 mm across; many crystals 

in two rock specimens contain opaque inclusions.

Andalusite as partly resorbed porphyroblasts is 

present in two rock specimens from an area about 

780 m from the batholith. The embayed and scalloped crystals, from 0.5 to 0.9 mm in greatest dimension, are separated from the graphitic groundmass by 

a clear rim of quartz as much as 0.2 mm wide. Andalusite, which also was recovered as a heavy mineral 

from one rock specimen that was crushed for the separation of biotite, clearly was unstable during the 

late crystallization history of the aureole rocks southeast of Summit Lake.

ARGILLITE AND TUFFACEOUS ARGILLITE

Argillite and tuffaceous argillite within 670 m of 

the south contact of the batholith occur mostly as 

rocks of the hornblende hornfels facies, but a narrow
58 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

zone of pyroxene hornfels-facies rocks is present near 

the batholith. The outward extent of hornblende 

hornfels-facies rocks beyond 670 m is unknown.

Rocks of both facies are well exposed for almost 8 km 

near the contact west of Hunt Mountain (fig. 2.2). The 

rocks are contact schists rather than hornfelses. Orthopyroxene was found as far as 38 m from the batholith, 

but common replacement by cummingtonite suggests 

that the original outward distribution of orthopyroxene 

was greater. Typical assemblages containing orthopyroxene are (1) orthopyroxene-plagioclase-quartz, (2) 

orthopyroxene-plagioclase-biotite-quartz, and (3) orthopyroxene-plagioclase-diopside-quartz with or without hornblende. Refractive-index determinations of a 

and Y for orthophyroxene separated from six specimens 

gave compositions of En48 to En52 . Determinations of 

a and Y indices for plagioclase separated from the same 

six specimens gave compositions of An34 to An44.

Both cummingtonite and hornblende are common 

in the hornblende hornfels-facies rocks. The two amphiboles occur together in some rocks but separately 

in other rocks. Cummingtonite is common from the 

contact outward for about 200 m but occurs as far 

outward as 650 m. Both brown and green hornblende 

are present; most brown hornblende occurs within 

about 100 m of the batholith. Some brown hornblende contains cores of cummingtonite. The typical 

assemblage containing either cummingtonite or hornblende as the only amphibole includes plagioclase 

and quartz with or without biotite. Diopside may occur as an additional mineral in rocks that contain 

hornblende as the only amphibole.

Eight rock samples of argillite were collected from 

the dismembered screen between the herein-named 

granodiorite of Indiana Mine Road (fig. 2.3) and the 

boundary quartz diorite unit for comparison with 

metamorphic assemblages in argillaceous rocks near 

the south contact of the Elkhorn pluton. Five of these 

samples are pyroxene hornfels-facies rocks that contain variable amounts of orthopyroxene, diopside, 

hornblende, plagioclase, and quartz; biotite is an additional mineral in two of the five samples. The remaining three samples are hornblende hornfels-facies 

rocks that contain hornblende, plagioclase, and 

quartz, with or without diopside.

LIMESTONE

Scattered xenoliths of calc-silicate rocks occur 

throughout much of the batholith, but wallrock contacts with metacarbonate rocks are rare at the 

present levels of exposure. Metamorphosed limestone 

near the batholith occurs either as marble, as bedded

calc-silicate rocks, or locally as garnet skarns. Nearly 

all exposures of these rocks within 30 m of the contact are poor and discontinuous. Wollastonite is the 

highest temperature index mineral in the metacarbonate rocks. The best and most accessible examples 

of both a garnet skarn and bedded calc-silicate rocks 

containing wollastonite are about 1.2 km S. 55° E. of 

Antelope Peak (fig. 2.2).

Abandoned prospect pits southeast of Antelope 

Peak aid in following the garnet skarn along the contact of the batholith. Minor quartz and epidote occur 

with the garnet. A narrow zone of scheelite as much 

as 1.5 cm wide commonly extends along the garnetmarble contact; smaller amounts of scheelite occur 

within the massive garnet. The garnet-marble contact, several meters or more from the edge of the 

batholith, is sharp and irregular and has extensions 

of garnet projecting into the marble in a manner consistent with metasomatism of a pure limestone.

Tonalite as much as 4 m from the contact of the 

garnet skarn on the southeast slopes of Antelope Peak 

contains diopside instead of hornblende and biotite. 

Considerable sphene is characteristic of diopside tonalite. Thin sections of 2 of 11 rocks show associations 

of diopside and hornblende that suggest replacement 

of amphibole by pyroxene. Formation of diopside is 

attributed to reaction of the solidifying magma with 

limestone and enrichment of CaO with concomitant 

loss of A12O3, (Mg,Fe)O, and SiO2 to the wallrocks. In 

the tonalite, TiO2 freed by cessation of crystallization 

of hornblende and biotite combined with CaO to form 

sphene. However, the volume of garnet skarn requires 

the transfer of much more material than could have 

been derived from the apparent volume of diopside 

tonalite as exposed at the surface.

Bedded calc-silicate rocks near the contact southeast 

of Antelope Peak occur as fine-layered sequences mostly from 5 to 20 mm in width but also in layered sequences more than 20 mm wide. Only the wide layers 

were sampled. Assemblages in seven rock specimens 

are wollastonite-diopside-plagioclase and wollastonitegarnet-diopside. Plagioclase separated from three 

rocks is An93_9g. The four minerals in the two assemblages have random orientations except where wollastonite is aligned between two or more closely spaced 

crystals of garnet. Preferential deformation of plagioclase is evident by recrystallized shear zones, bent 

twinning lamellae, and microfaults with offset twin 

lamellae.

Wollastonite in the Belleview wedge occurs in four 

small lenses of bedded calc-silicate rocks at distances 

of about 100, 250, 310, and 1,300 m from the batholith. 

The lenses are from 0.5 to 3.0 m in maximum width 

and from about 4 to 7 m in length. The most conspic-
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 59

uous feature of the four metamorphosed carbonate 

lenses is repetitious sequences of apparently monomineralic zones of red-brown garnet, green diopside, and 

white wollastonite. Some layers are as much as 20 mm 

across, but most are less than 8 mm; many layers are 

less than 3 mm, and some are less than 1 mm. Sawed 

surfaces (45 to 65 cm2) of hand specimens clearly reveal that the calc-silicate layers interfinger. Layers 

commonly are less than about 0.3 mm wide where they 

terminate. Coarsely crystalline marble occurs with the 

calc-silicate sequences in two of the metacarbonate 

lenses. Zones of pure marble contain no calc-silicate 

minerals, whereas marble zones with minor impurities contain either scattered silicate crystals or crystals concentrated in laminae that commonly are very 

thin. For example, some layers of garnet in marble do 

not exceed 1 mm in width and 25 mm in length.

Two thin sections of fine-layered calc-silicate rocks 

from each of the four metacarbonate lenses disclose 

that the layering consists primarily of alternating sequences of wollastonite-diopside, garnet-diopside, and 

plagioclase-diopside. Two or more layers that contain 

the same minerals commonly differ in modal proportions, but no true monomineralic zones were detected, except for short distances of not more than about 

8 mm. However, zones of wollastonite with only 

about 4 volume percent diopside extend across the 

width (24 mm) of thin sections, as do both zones of 

garnet with only about 10 volume percent diopside 

and zones of diopside with only about 10 volume percent plagioclase. The metacarbonate lens farthest 

(1,300 m) from the batholith differs from the other 

lenses in the presence of minor quartz in some layers 

of both diopside-plagioclase and garnet-diopside.

Foliation in the calc-silicate layers is mostly defined 

by elongated crystals of wollastonite; the other minerals are more or less equigranular and generally less 

than 0.5 mm across. Many crystals of wollastonite are 

almost 1 mm in length and rarely as much as 3 mm. 

In strongly foliated rocks many wollastonite crystals 

have elongation ratios from 7:1 to as much as 9:1. 

Polysynthetic twinning in plagioclase in strongly deformed rocks commonly is parallel to bedding. Twinning lamellae in plagioclase are deformed both by 

plastic deformation and by rupture. Maximum deformation of plagioclase occurs in crystals more than 1 

mm across that exhibit common differences in directions of twinning lamellae from 15° to 20°.

METAGABBRO

Thermal metamorphism of metagabbro without a 

dynamic component is characterized by the random 

crystallization of amphibole needles that penetrate

original plagioclase in a decussate fashion. The gabbroic texture is somewhat modified but never obliterated. 

Most metagabbro within 1,000 m of the batholith, however, is dark amphibolite that commonly shows only 

traces of the original gabbroic texture. Many of the 

metagabbros are schistose. Intensely deformed metagabbros commonly show gneissic layering with alternating light and dark layers that resembles the 

bedding in some metasedimentary rocks (fig. 2.6). The 

difficulty in recognizing intensely metamorphosed metagabbro is underscored by a small inlier of pre-Tertiary 

rocks about 2.5 km due south of Chicken Hill (fig. 2.2). 

Throughout this inlier, metagabbro with gneissic layering is shown as "argillite series" on an otherwise 

excellent map by Pardee (1941). Good exposures of the 

misinterpreted metagabbro are on either side of the 

North Fork of Trail Creek (fig. 2.2JB). The light layers 

are plagioclase, whereas the dark layers are mostly 

hornblende and diopside. In some metagabbroic rocks 

near the batholith the light layers are elongated lenses. Much less commonly the metagabbros contain so 

little plagioclase that on weathered surfaces they are 

a uniform greenish-black to coal black. Some metagabbro at the contact of the batholith at an elevation of 

about 2,578 m and 0.59 km S. 15° E. of Summit Lake 

(fig. 2.2) shows contorted layering that suggests plastic 

deformation.

Strongly foliated amphibolite occurs mostly within 

200 m of the batholith, throughout much of the Muddy Creek and Belleview wedges (fig. 2.2B), and as 

screens and inclusions within the batholith. Foliation 

generally dips within 20° of vertical. At least a few 

unequivocal vestiges of gabbroic texture (fig. 2.7) survive in all bodies of metagabbro outside of the batholith except one small mass in the Belleview wedge. 

Although most gabbroic relics in the amphibolites are 

lens-shaped (fig. 2.7), some are small knotlike masses, commonly from 0.5 to 2.5 cm across; locally, two 

or more closely spaced knots resemble boudins or 

even large rolled garnets.

The typical foliated amphibolite is characterized by 

a simple assemblage of plagioclase, hornblende, iron 

oxide, and apatite. Plagioclase compositions range 

from An62 to An86, as determined by refractive indices 

of crystals separated from nine samples of amphibolite. 

Minor quantities of quartz and green diopside are 

present in some amphibolites. Plagioclase in many amphibolites has been completely recrystallized to equant 

xenoblastic grains that are water-clear and commonly 

untwinned. Rounded and lens-shaped relics of the original plagioclase persist in some amphibolites. Recrystallization is accompanied by reduction in average 

grain size from considerably more than 1.0 mm to 

slightly less than 0.1 mm. The average grain size of
60 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

hornblende in rock specimens of strongly foliated amphibolite ranges from 0.35 to 0.10 mm. The foliated 

amphibolite is indicative of directed pressure associated with the forceful injection of granitic magma.

A notable feature of some schistose amphibolites 

north of lat 45° N. is the presence of so-called segregation veins and lenses that are leucocratic except 

for central concentrations of garnet. Locally, where 

extreme deformation and recrystallization have occurred, the garnet-filled veins and lenses provide evidence of the metagabbroic origin of the amphibolite.

Pyroxene hornfels-facies amphibolites occur in the 

northern and western parts of the batholith as judged 

by samples collected near (1) Gorham Butte (fig. 2.2), 

(2) from a dismembered screen along the southwest 

contact of the herein-named granite of Clear Creek 

(fig. 2.2) with the Elkhorn pluton, (3) from a dismembered screen between the herein-named leucogranite

FIGURE 2.6. Gneissic layering in intensely deformed metagabbro. Such layering can be mistaken for metasedimentary bedding. Lower rock specimen is from small inlier of 

metagabbro about 2.5 km south of Chicken Hill (fig. 2.2).

Upper specimen was collected about 125 m from Bald Mountain batholith near westernmost part of Muddy Creek wedge 

(fig. 2.2B). Coin is about 18 mm in diameter.
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 61

of Dutch Creek and the boundary quartz diorite unit 

(fig. 2.3), and (4) from a large inclusion at least 3 m 

long in the Elkhorn pluton. This inclusion is 1.65 km 

N. 31° E. of Badger Butte, about 60 m due west of 

Trail Creek (fig. 2.2B), on the west side of a dirt road 

that closely parallels Trail Creek; it is just north of a 

conspicuous xenolith of layered calc-silicate rock at 

least 12 m long. The consistent assemblage for the 

pyroxene hornfels-facies amphibolites from all four localities is orthopyroxene-diopside-plagioclase, with or 

without hornblende.

ULTRAMAFIC ROCKS

Serpentinized ultramafic rocks are present in the 

aureole as small lenses commonly less than 25 m 

wide that constitute about 1.5 percent of the country 

rock within 2 km of the batholith. According to 

Evans (1977), the following minerals in serpentinite 

indicate progressive zones of increasingly intense 

thermal metamorphism: (1) tremolite, (2) talc, (3) anthophyllite, (4) enstatite, (5) spinel, (6) pargasitic 

hornblende, and (7) plagioclase. The first mineral in 

the series, tremolite, marks the transition from 

greenschist to hornblende hornfels facies, and the 

next-to-last mineral, pargasitic hornblende, denotes 

the transformation to pyroxene hornfels facies. Tremolite occurs in nearly all ultramafic rocks within 1.6 

km of the Bald Mountain batholith. Green spinel appears before enstatite, but otherwise the progressive 

sequence of index minerals closely parallels the 

parageneses of Evans (1977).

Green spinel (n=1.756 to 1.767) is present in most 

ultramafic rock specimens collected within 200 m of

the batholith. Specimens from near the contact commonly contain 4.0 to 11.3 volume percent spinel. 

Grain size increases and crystals become more nearly 

euhedral as the batholith is approached. Subhedral 

to euhedral spinel crystals between 0.3 and 0.5 mm 

across are common in thermally metamorphosed serpentinite within about 30 m of the contact. Some 

green spinel is enclosed or adjacent to chlorite and 

less commonly phlogopite. The chlorite apparently is 

clinochlore, as suggested by birefringence of about 

0.08, distinct polysynthetic twinning, biaxial + figure, 

and estimated 2V of 25°. Chlorite in the aureole 

rocks overlaps the stability field of enstatite.

The highest temperature metamorphic assemblage 

in ultramafic rocks in the Elkhorn Mountains is near 

the apex of the Belleview wedge (fig. 2.2B), where 

metaperidotite contains olivine, clinopyroxene, enstatite, pargasitic hornblende, and commonly plagioclase 

with small amounts of phlogopite, green spinel, and 

uncommon chlorite. The assemblage, which indicates 

pyroxene hornfels facies, is rare in high-grade contactmetamorphic aureoles; plagioclase peridotite hornfels 

were first reported by Frost (1976) near the contact of 

a gabbro that is an early unit of the Mount Stuart 

batholith, Washington. In the Elkhorn Mountains the 

high-grade plagioclase peridotite hornfels are exposed 

at elevations between 2,225 and 2,270 m in a lens as 

much as 30 m wide near the crest of the ridge about 

4.3 km N. 31° E. from the summit of Mount Ireland. 

These exceedingly tough rocks can only be adequately 

sampled with a long-handled rock hammer such as 

those available in England and Germany or with a 

somewhat heavier American sledge hammer.

Pargasitic hornblende, olivine, and clinopyroxene 

are the most abundant of the major minerals in the

FIGURE 2.7. Lens-shaped relic of metagabbro (arrows point to top and bottom margins) in sheared and recrystallized amphibolite. Coin is about 18 mm in diameter. Specimen from small inlier south of Chicken Hill.
62 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

pyroxene hornfels-facies rocks northeast of Mount Ireland. The hornblende is either almost colorless, pale 

brown, or light reddish brown. Very small amounts of 

phlogopite occur in all rock specimens. Chlorite is rare 

and apparently has been removed by reactions (Fawcett and Yoder, 1966) that formed forsterite, enstatite, 

spinel, and H2O. Most rock specimens contain green 

spinel but in much less abundance than in lower temperature assemblages. Much of the spinel may have 

reacted with one or more Ca-Mg-silicate minerals to 

form calcic plagioclase and forsterite (Evans, 1977). 

Plagioclase is present in small amounts in most rock 

specimens but does not exceed 3 percent by volume in 

contrast to from 20 to 25 percent by volume in rocks 

described by Frost (1976).

STRUCTURE OF ELKHORN PLUTON

PERIPHERAL CONTACTS AND ASSOCIATED FOLIATION IN 

MARGINS OF PLUTON

The Elkhorn pluton constitutes about 93 percent of 

the batholith and is a single intrusion that ranges 

outward from a felsic core composed of the potassium-poor granodiorite of Anthony Lake to a margin 

composed of the mafic tonalite of Bald Mountain. The 

complexity of contacts of the tonalite of Bald Mountain with country rocks differs from south to north as 

does the intensity and distribution of foliation in the 

tonalite.

The south contact of the tonalite of Bald Mountain 

is best exposed near the serrated crest of a glaciated 

ridge south and southeast of Summit Lake (fig. 2.2); 

here, along about 4 km of its length, the contact is 

sharp and simple with few inclusions of country rock. 

The mafic inclusions (Pabst, 1928) of uncertain origin 

that characterize most circum-Pacific plutons of tonalite and granodiorite compose about 0.26 volume 

percent (range, 0.18-0.35 volume percent) of the tonalite of Bald Mountain as determined at eight localities near the contact by measurement of inclusions 

exposed across areas ranging in size from 56 to 70 

m2. In comparison, mafic inclusions of unknown origin in a 12-km2 area near Anthony Lake in the interior of the pluton (fig. 2.8) average 0.35 volume 

percent (range, 0.17-0.56 volume percent) of the rock 

as determined at 58 localities by measuring inclusions across areas ranging in size from 47 to 187 m2. 

Thus, mafic inclusions of unknown origin are not as 

abundant in mafic tonalite south and southeast of 

Summit Lake as they are in granodiorite near Anthony Lake in the interior of the pluton. Near some peripheral contacts south and southeast of Summit 

Lake, the mafic inclusions are randomly oriented.

However, they commonly are elongated (fig. 2.9) 

within 350 m of the contact, with some elongation 

ratios as much as 5 to 1. Planar structure in two areas southeast of Summit Lake extends at least as far 

as 425 m from the contact. Planar structure roughly 

parallels the contact and dips either north or south, 

mostly within 10° of vertical.

Foliation in tonalite in the south-central and southeastern parts of the batholith is mostly localized near 

contacts. In the Monumental salient, to the southwest (fig. 2.25), foliation is absent in tonalite near 

many contacts. The absence of planar structure near 

many contacts of the Monumental salient and the restricted occurrence of planar structure near most other contacts in the southern parts of the batholith are 

anomalous in a mesozonal pluton and contrast with 

widespread foliation in the tonalite in the northern 

part of the batholith.

The zone of foliation along the east contact of the 

batholith near Antelope Peak (fig. 2.2) is wider and 

more pronounced than the zone along the southern 

contact of the batholith. Elongated mafic inclusions 

of unknown origin occur as far inward from the contact as 500 m before the shapes become rounded and 

irregular. Elongation ratios commonly are 6 to 1 and 

many are 8 to 1. Inclusions are much more abundant 

here than along the south contact of the batholith. 

Mafic inclusions of unknown origin in most exposures 

constitute between 0.5 and 2.2 volume percent of the 

tonalite, but only a few recognizable inclusions of 

country rocks occur. Near Antelope Peak the contact 

of the Elkhorn pluton is sharp and relatively simple. 

The foliated border zone of the batholith southeast of 

Antelope Peak is well exposed in the minichasm of 

the Bulger Ditch (fig. 2.25). Since the turn of the 

century the irrigation waters have cut an almost vertical trench as deep as 12 m through most of the border-zone rocks.

Contact relations northwest of Antelope Peak are 

more complex near Gorham Butte (fig. 2.2). Here the 

contact consists of a zone as much as 60 m wide of 

foliated and crudely interlayered country rocks and 

tonalite, commonly alternating in a large-scale litpar-lit fashion. Planar structure extends inward as 

far as 900 m from the zone of interlayered rocks. Foliation parallels the contact and generally dips 60°- 

75° NE. Recognizable inclusions of country rocks occur inward from the zone of interlayered rocks, but 

inclusions of all types generally constitute less than 2 

percent of the tonalite. Numerous elongated mafic inclusions of unknown origin are present in tonalite as 

far as 500 m from the zone of interlayered rocks; 

some inclusions near the contact have elongation ratios exceeding 25 to 1 (fig. 2.10). One extreme exam-
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 63

pie is 132 cm long and only 4 cm wide. Locally, small 

lensoid concentrations of mafic inclusions make up as 

much as 20 percent of the tonalite.

From Gorham Butte the contact of the batholith 

continues northwestward for about 3 km as a mixed 

zone of tonalite and foliated country rocks as much 

as 80 m wide. Foliated tonalite strikes about N. 60° 

W. and dips mostly 50°-70° NE. Farther west, where 

the contact zone is intruded by the granite of Anthony Butte (fig. 2.3), a dismembered screen represented 

by large inclusions (as long as 0.5 km) and numerous 

smaller country-rock xenoliths occurs along and near 

the tonalite-granite contact. From Gorham Butte 

northwestward along the tonalite contact for about 5 

km, small inclusions of country rocks and mafic inclusions of unknown origin make up as much as 5 

percent of the tonalite. The country rock inclusions

contain many calc-silicate xenoliths, which suggest 

that emplacement of the tonalite in this part of the 

batholith proceeded alongside a large limestone lens. 

Throughout an area slightly larger than 45 km2 

that lies west, southwest, and northwest of the granite of Anthony Butte (figs. 2.2 and 2.3), nearly all tonalite is foliated, in marked contrast to the restricted 

distribution of peripheral foliation in tonalite south of 

lat 45° N. Foliated tonalite west of the granite of Anthony Butte extends south to an intraplutonic contact 

with granodiorite (fig. 2.2), located just north of lat 

45° N. and described in a later part of this chapter. 

Elongation ratios of mafic inclusions commonly are 

between 6 to 1 and 17 to 1. Concentrations of mafic 

inclusions occur in lens-shaped zones as much as 2 m 

wide and 35 m long; the mafic inclusions constitute as 

much as 40 percent of the rock in several zones.

FIGURE 2.8. Granodiorite of Anthony Lake near center of 

Elkhorn pluton. View to west. Anthony Lake is 2 km north of 

the two peaks on central skyline. Granodiorite in this part of 

Elkhorn pluton averages 0.35 percent inclusions, mostly of unknown origin but including some of recognizable country rocks

as much as 0.9 m across. Some granodiorite is characterized by a 

faint to fairly good planar structure. Northernmost sampled granodiorite showing evidence of early crystallization of augite was 

collected on ridge extending northeastward from prominent unnamed peak on left horizon.
64 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

In general, the complexity of contacts of tonalite 

with country rocks correlates with the dip of the contact. Most vertical to near-vertical plutonic contacts 

in northeastern Oregon are relatively simple in comparison with contacts that dip outward from the pluton at about 60° or less. Magma commonly crosscuts 

and (or) intrudes along several or more bedding 

planes where it ascends beneath outward-dipping 

contacts but not where contacts are essentially vertical. The lens-shaped concentrations of mafic inclusions north of lat 45° N. are interpreted as relatively 

large slabs of wallrock that were gradually fragmented into many closely spaced inclusions as they were 

incorporated into the magma.

In northeastern Oregon the most accessible outward-dipping contact with well-exposed sequences of 

alternating tonalite and country rocks is along the 

north contact of the Cornucopia stock (Taubeneck, 

1964, fig. 2). The pertinent exposures, visible from 

distances of several hundred meters, are in open terrain about 5.3 km N. 38° W. of Cornucopia (fig. 2.1). 

Ready access from Cornucopia is by a trail to Pine 

Lakes, which lie nearly 6 km northwest of Cornucopia. Within the stock the trail crosses a small stream 

at an elevation of about 1,850 m and roughly parallels the intrusive contact at higher elevations to the 

north. West-northwest of the stream crossing, at an 

elevation of about 2,160 m, are alternating sequences 

of bedded metasedimentary rocks and tonalite that 

extend northward to an elevation of about 2,260 m. 

Most of the alternating sequences of metasedimen-

'". - »J!

FIGURE 2.9. Cluster of five elongated mafic inclusions (arrows) of 

unknown origin about 90 m from contact of Elkhorn pluton 

southeast of Summit Lake (fig. 2.2). Elongation ratios from 2:1 

to 5:1 are common for inclusions near south contact of Elkhorn 

pluton but are much lower than ratios for mafic inclusions along 

north contact. Knife is 10 cm long.

tary rocks are less than 4.5 m wide. Commonly the 

relative volumes of metasedimentary rocks and tonalite in the larger exposures are about equal. Bedding dips mostly between 50° to 65° outward (north) 

but locally as shallowly as 30°. The metasedimentary 

rocks consist of a variety of rock types including calcsilicate rocks, mafic rocks of amphibolitic composition, and rare cherty beds (as much as 15 cm wide). 

Many of the small-scale features commonly associated with contacts of granitic intrusions, as reported in 

the literature worldwide, occur in this vicinity.

INTERNAL FOLIATION

The Elkhorn pluton is distinctive among the zoned 

intrusions of tonalite and granodiorite in northeastern Oregon in the occurrence of sporadic planar 

structure in the granodiorite of Anthony Lake that 

forms the core of the pluton and also in the tonalite 

of Bald Mountain that composes the interior of the 

Monumental salient. The structural relations generally are elusive but locally are clear for distances of

FIGURE 2.10. Extreme elongation of mafic inclusions (arrows) in 

tonalite near Gorham Butte (fig. 2.2). Elongation ratios of all inclusions here exceed 25:1. Knife is 10 cm long.
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 65

100 m and more. The sporadic structures are indicative of numerous magma surges and pulses during 

the consolidation phase of emplacement of the granodiorite core and the interior of the tonalite in the 

Monumental salient.

Examination of the granodiorite core was mostly 

confined to a 12-km2 area south to east of Anthony 

Lake; the greatest number of continuous exposures of 

the granodiorite core occur here. The area is characterized by sharp peaks and cirque basins. The southeastern part of the core, deeply dissected by a major 

glaciated valley that includes Dutch Flat (fig. 2.2), remains to be studied.

Textbooks, field manuals, and the nonetheless 

highly commendable memoir of Balk (1937) provide 

little guidance for the recognition of planar structure 

throughout much of the interior of the Elkhorn pluton. Therefore, field procedures and observations are 

detailed herein because the Bald Mountain batholith 

is not unique in its elusive planar structure, rocks 

with relatively smooth joint surfaces, and common 

exposures either controlled by steep diagonal joints 

or shaped by exfoliation.

Diagonal joints warrant primary consideration. 

Serrate ridges as well as notched or V-shaped peaks 

are common physiographic expressions of diagonal 

joints in recently glaciated granitic terrains. Gunsight Mountain (fig. 2.2jB) is merely one example of a 

notched peak south of Anthony Lake. Much of the 

Wallowa batholith also is dominated by diagonal 

joints. Excellent examples of sawtooth ridges occur in 

unit 3 of the Wallowa batholith about 9.8 km S. 25° 

W. of the south end of Wallowa Lake (fig. 2.11).

Most exposures south to east of Anthony Lake are 

controlled by two prominent northeast-striking diagonal-joint sets that dip 35°-45° NW and 60° SE, respectively. Faintly to even moderately delineated 

planar structure is difficult to detect because (1) the 

relatively smooth joint surfaces obscure the structure, (2) the diagonal joints restrict the interval for 

recognition across outcrops, and (3) the glaring reflection of the sun from the dominant diagonal surfaces 

(the northwest-dipping set) makes further structural 

elucidation nearly impossible.

The character of the exposed rock surface is an important factor in the recognition of planar structure; 

with increasing roughness of the rock surface, planar 

structure is increasingly discernible. The most favorable surfaces are those created by dynamite blasting 

done by the U.S. Forest Service in making trails. 

Weak planar structure that is visible in massive rock 

on large surfaces exposed by blasting generally cannot 

be seen on adjacent smooth-jointed natural exposures. 

Large talus blocks at the base of cliffs, especially

blocks 1 m or larger in size, as well as natural exposures with rough surfaces commonly reveal structure 

that is obscure on relatively smooth surfaces.

Diagonal-joint surfaces south to east of Anthony 

Lake are among the most difficult in the pluton to 

examine with any degree of success. On these surfaces, traces of planar structures mostly trend at angles 

within 45° to 75° of joint strikes. Even fairly well delineated planar structure is not easy to recognize looking downdip along surfaces that are inclined at 35° to 

45°, because the observer is looking at an acute angle 

at a surface that is rapidly becoming too distant for 

inspection. Therefore the surface area that can be 

viewed at the proper distance or proper angle is inadequate. Any structure in the rocks is much easier to 

detect if the observer turns 180° and then scans a 

surface that is at a roughly uniform radial distance. 

Viewing an exposure from several different angles and 

different distances is essential to the recognition of 

any poorly defined planar structure. The angle of recognition of elusive planar structure on diagonal-joint 

surfaces commonly is as low as 20°.

The diagonal-joint set that dips 35°-45° NW. exposes 

the largest continuous surface areas amenable to the 

identification of planar structure. As noted previously, 

however, even fairly well defined planar structure is 

generally difficult to recognize looking downdip at angles of 35° to 45°. But looking updip is also difficult 

because in this area it necessitates facing southeast, 

toward the sun. The constant reflection of sunlight 

from the white surfaces of granodiorite is an almost 

insuperable obstacle; however, in October, before the 

arrival of the winter snows, many areas of northwestdipping diagonal joints (especially where angles are 

near 45°) are in shade for as many as four consecutive 

hours and some throughout the day.

Taking advantage of smoke from wildfires and using 

artificial-shading devices are two means of facilitating 

the recognition of planar structure. Major wildfires 

are fairly common during August and September in 

northeastern Oregon and generally burn for 15 to 25 

days. Smoke from wildfires as much as 130 km away 

can reduce midday visibility to a distance of 6 to 10 

km. Planar structure is more easily detected in the 

Elkhorn pluton on any day when smoke reduces visibility to less than about 25 km. Thus very good to 

excellent field observations are possible throughout 

the day when visibility is only 6 to 10 km.

Artificial-shading devices permit the detection of 

fairly to moderately well delineated planar structure. 

The devices are almost always practical on south-facing slopes, and they also are useful during July and 

August for most north-facing slopes. Examples of effective artificial-shading devices are pieces of cardboard
66 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

and a standard collapsible umbrella. An advantage of 

cardboard is its availability in a wide range of sizes.

An unopened umbrella's compact size means that it is 

small enough to fit inside a day pack, but recognition

117° 30' 117° 15'
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 67

of weak structure commonly requires a larger shadow 

than that provided by an umbrella (depending in part 

on the hour of day, month of year, and latitude of the 

area). For example, even at high noon on October 15 

at lat 45° N., a 6-ft (~1.8-m) person standing on a 

horizontal surface will cast a 2-m shadow that rapidly 

lengthens within 90 minutes. The shadow at noon on 

June 15 is only about 0.7 m, but it is 4 m long at noon 

on December 15. Therefore, the effectiveness of the 

same artificial-shading device for revealing weak 

structures varies according to time and place.

Examination of most exposures south and southeast of Anthony Lake indicates respectively that east 

and northeast structural trends and west-northwest 

trends dominate (fig. 2.2). About 4 to 5 km east and 

east-southeast of the lake, trends have changed to N. 

0°-25° W. Overall, including areas not previously 

mentioned to the north and southwest of the lake, 

the apparent pattern of structural trends in the interior of the Elkhorn pluton crudely defines an arc of 

about 130°. The arcuate pattern is consistent with 

upward surges of magma during the progressive inward consolidation of the core.

Planar structure where present within the interior 

of the Monumental salient varies from faint to well 

defined. Tonalite in the western one-fourth of the salient occurs as exfoliation knobs. Such rounded surfaces complicate the recognition of planar structures 

and the determination of their orientations, but the 

extent of the complications depends on many factors. 

For example, structure in exfoliation knobs north of

EXPLANATION

Qs

Tor! 

Jgr 

Jgb 

JEs

Surficial deposits (Quaternary)-Shown 

only near Tallotre fault

Columbia River Basalt Group (Miocene) 

Granitic intrusive rocks (Jurassic)

Gabbroic intrusive rocks (Jurassic)

Sedimentary rocks (Jurassic and Triassic) 

iiinnii -Includes some mafic volcanic rocks

 Contact Short dashed There approximately 

located

_ _ _ Gradational unit boundary

f High-angle fault Dashed Trhere approximately 

"~^~~ located; dotted There concealed. Bar and 

ball on downthrown side

X Aplite body

FIGURE 2.11. Generalized geologic map showing 

units 1 through 4 of Wallowa batholith; these 

four compositionally zoned plutons, composed of 

tonalite and granodiorite, constitute most of the 

Wallowa batholith. Locations of three large aplite bodies are also shown.

the intraplutonic contact (fig. 2.2) is comparatively 

easy to recognize partly because the structure is accentuated by the large size, pronounced elongation, 

and relative abundance of mafic inclusions. In contrast, the relative scarcity, smaller size, and absence 

of pronounced elongation of mafic inclusions in the 

western part of the Monumental salient contribute to 

the difficulty of recognizing planar structures there. 

The contrast is intensified by weaker fabrics in rocks 

from the salient due, for instance, to their containing 

less than half as much hornblende, which is important in defining the fabric. Structural definition in 

the western part of the salient is, at best, only fair to 

fairly good.

Exfoliation knobs in the salient commonly are surrounded at ground level by spalled slabs and rinds. 

The most instructive slabs are more than 0.6 m2 , at 

least 7 cm thick on two adjoining sides, and comparatively fresh; have originated from the central and lower parts of knobs that are more than 2.5 m high; and 

provide rock surfaces for inspection in three different 

directions. Planar structure can be determined more 

readily in such slabs than in the knobs themselves. 

Rotating a slab in the air is essential in searching for 

unknown structure but is more important in confirming and documenting the attitude of suspected structures. The confirmed attitude is marked on each slab. 

A larger slab's original position on the knob generally 

is apparent and may be verified by lifting and then 

moving a slab until the respective curved surfaces of 

knob and slab coincide precisely.

Many knobs do not yield even one suitable exfoliation slab for the possible detection and measurement 

of planar structures. Furthermore, some appropriate 

slabs cannot be convincingly restored to their original 

position. Examination of the knobs commonly discloses incipient to well-developed exfoliation cracks, many 

of which permit the breaking of large pieces of rock for 

inspection. Massive exfoliation surfaces are best broken with a heavy, long-handled rock hammer such as 

a 2-kg German hammerhead on a hickory handle or a 

6-lb American sledge. Large specimens broken from 

exfoliation knobs facilitate the recognition of fair to 

moderately well defined planar structure, but the surface areas are inadequate for the recognition of weak 

structures.

Joint surfaces rather than exfoliation surfaces characterize the well-exposed rock at higher elevations in 

the central and eastern parts of the Monumental salient. The tonalite throughout most of the salient is 

distinctive compared to tonalite elsewhere in northeastern Oregon in its relatively large number of xenoliths composed of country rocks. Oriented xenoliths are 

the best indicators of otherwise weak fabrics. Xenoliths
68 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

of foliated country rocks, mostly metagabbro, commonly occur as elongated inclusions many of which are 

small slivers less than 5 cm long. The slender dark 

slivers that contain abundant hornblende are readily 

visible and, where aligned, accentuate or confirm a 

weak planar structure. Some narrow inclusions of ribbon chert also are arranged with their longest axes 

parallel to enclosing foliation planes. It is notable that 

aligned xenoliths are not everywhere accompanied by 

foliation planes. However, aligned xenoliths without 

associated foliation planes generally are parallel to 

planar structure in nearby areas.

A representative axis of planar structure in the interior of the Monumental salient trends west-northwest at an angle across the salient (fig. 2.2) rather 

than parallel to the length of the salient, which runs 

east-west. This structural axis closely parallels a line 

drawn through the narrowest part of the neck of the 

salient and coincides with the overall west-northwest 

trends of country rocks on either side of the neck. 

These gross relations suggest that the salient initially 

was an independent, satellitic intrusion that subsequently coalesced with the main part of the batholith.

Planar structure within the neck of the salient varies from faint to well defined but is absent in many 

exposures. Trends in the neck are much more variable 

than in the interior of the salient and locally arc as 

much as 45° within 245 m. In the northwest part of 

the neck, structures generally trend north-northwest 

to north-northeast (fig. 2.2A). In the southeast part of 

the neck, structures trend northeast. However, in the 

restricted exposures of the northern part of the border 

area between the neck and the salient, structures are 

mostly west-northwest, parallel to both the long direction of the neck and structures to the southwest in the 

interior of the salient. These west-northwest trends 

are contrary to the structural pattern that would be 

expected as a result of magma passing through the 

neck between the batholith and the salient, and support the interpretation that the salient originally was 

a satellite of the batholith. The lack of structural symbols on figure 2.2 in the important 1-km-wide zone 

just south-southwest of the neck is caused by a glacial 

valley (fig. 2.4) that contains only a small number of 

iron-stained roches moutonnees, each without discernible planar structure.

Extensive limonitic alteration along closely spaced 

joints associated with gold-quartz mineralization 

(Lindgren, 1901; Ferns and others, 1982) greatly complicates the recognition of planar structure throughout 

much of the neck. The principal joint set includes one 

major and many minor zones of atypical joints that are 

very closely spaced (2.5 to 20 cm apart), hydrothermally altered, and accompanied by gold-quartz veins that

mostly parallel the jointing. In areas outside of the 

zones of atypical jointing, joints of the principal joint 

set are less regular, mostly spaced from 0.3 to 2.4 m 

apart, and generally parallel to the trend of the nearest zone of atypical joints. The principal joint set commences about 5 km S. 78° E. of Mount Ireland (fig. 

2.2), strikes about N. 10° W. for almost two km, and 

then gradually arcs to the northeast through the neck. 

In the area of figure 2.4, the major zone of atypical 

joints extends northward from the southernmost vertical joint symbol (striking N. 20° E.), then arcs 

through Cable Cove and the two roches moutonnees on 

the north side of Silver Creek before continuing northeast to the vicinity of the spring east of the North Fork 

John Day River. Rock near the spring is a limoniticstained grus peppered with many century-old prospect 

pits that assist in confirming the northeast continuation of the mineralization and closely spaced joints. 

The major zone of atypical joints has a maximum 

width of about 1,200 m within 0.5 km north and northeast of Cable Cove. Minor zones of atypical joints, to 

either side of the major zone, are mostly from 2 to 15 

m wide. Of the 64 joint symbols shown on figure 2.4, 

51 represent the atypical closely spaced joints. Associated quartz veins are mostly 3 to 50 mm wide; a few 

are as much as 60 cm wide.

The gaping joints and gold-quartz veins are the result of pronounced upward expansion and arching of 

the granitic rocks on both sides of the neck. Whatever the earlier emplacement history of the tonalite, 

the dominant joint set indicates that the neck and 

salient reacted as a coherent unit during the final episode in the emplacement of the Elkhorn pluton.

INTRAPLUTONIC CONTACT

The gradational zoning within the Elkhorn pluton 

from tonalite in the margin to granodiorite in the 

core is disrupted 6.5 km north to 8 km northwest of 

Anthony Lake and again about 6 km west of Anthony 

Lake by an intraplutonic contact (fig. 2.2) where core 

magma on the south and east sides intruded already 

solidified tonalite on the north and west sides. Tonalite north and west of the intraplutonic contact has a 

color index of 17 to 30, whereas younger core rocks 

south and east of the intraplutonic contact have an 

index of 9 to 16. Rocks on either side of the intraplutonic contact in most areas also can be distinguished 

by the presence or absence of planar structure. Older 

tonalite north and west of the contact is strongly foliated, whereas younger core rocks south and east of 

the contact are generally unfoliated. The foliation 

that occurs in core rocks near the intraplutonic con-
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 69

tact (mostly about 8 km N. 50° W. of Anthony Lake) 

is faint to fairly good.

The intraplutonic contact invariably separates 

rocks of different color index and fabric, but the 

younger core rocks are not everywhere of uniform 

composition. Nearly all core rocks on the north are 

granodiorite, whereas those northwest and west of 

Anthony Lake are either granodiorite or rocks transitional between granodiorite and tonalite; core rocks 

near the intraplutonic contact about S. 67° W. of Anthony Lake are tonalite. Locally, the intraplutonic 

contact north to northwest of Anthony Lake is accentuated by small intrusions of leucogranodiorite that, 

in outcrop, are elongate parallel to the contact. The 

best exposed and most accessible intrusion of leucogranodiorite, about N. 9° E. of Anthony Lake, is approximately 600 m long and as much as 80 m wide.

Relations near the intraplutonic contact north and 

northwest of Anthony Lake can be readily determined 

for only about 2.5 km in the general vicinity of where 

the contact changes in strike from almost east-west to 

about N. 75° E. Tonalite in this area has a color index 

of about 25, whereas the granodiorite (core) has an 

index of about 12. Just east of the directional change 

in the contact, foliated tonalite (on the north-northwest side) with elongated inclusions is within 6 m of 

nonfoliated granodiorite (on the south-southeast side) 

containing a few rounded inclusions of strongly foliated tonalite. Inclusions of tonalite are rare or absent 

elsewhere in core rocks near the intraplutonic contact.

From due north of Anthony Lake the intraplutonic 

contact extends roughly westward for about 7 km before 

disappearing beneath Cenozoic volcanic rocks. Elongated inclusions and faint to fairly good planar structure 

are found in core rocks near the contact for about 1.5 

km east of the Cenozoic volcanic rocks in notable contrast to their absence in core rocks farther east.

A fundamental consideration is whether the intraplutonic contact turns southward beneath the Cenozoic 

volcanic rocks and reappears somewhere north of the 

norite of Badger Butte (fig. 2.2). Planar structure and 

associated elongated inclusions in core rocks curve 

from about N. 45° E. on the north at lat 45° N. to S. 

20° E. about 2.5 km to the south. Virtually all foliation 

in core rocks near the intraplutonic contact occurs in 

a relatively small area near and just south of lat 45° 

N. and apparently formed only where the core magma 

was most closely constricted during emplacement. The 

curving foliation in the core rocks suggests that the 

intraplutonic contact continues, concealed, beneath the 

Cenozoic volcanic rocks and reappears to the south.

Relations near the north end of the elongated leucogranodiorite of Trail Creek (fig. 2.2) indicate that 

the intraplutonic contact reappears in a poorly exposed area about 1 km north-northwest of the elongated leucogranodiorite. Mafic tonalite, strongly 

foliated, outcrops for about 500 m northwest of the 

leucogranodiorite and extends eastward at least as 

far as the intraplutonic contact as mapped on figure 

2.2. The nearest exposures of nonfoliated granodiorite are about 75 m farther east. Almost due south, 

near the east contact of the leucogranodiorite of Trail 

Creek, nonfoliated granodiorite with a color index of 

10 to 14 is exposed intermittently southward for 

about two-thirds of the length of the leucocratic intrusion. Strongly foliated tonalite with a color index 

of 19 to 30 is on both the west and south sides of the 

leucogranodiorite. The relations indicate that the intraplutonic contact extends southward along the east 

margin of the leucogranodiorite (fig. 2.2).

Most bedrock between 300 to 900 m south-southeast of the leucogranodiorite of Trail Creek is concealed by widespread glacial deposits, but a solitary 

exposure of the intraplutonic contact occurs in a 

driftless area almost midway between the southern 

extremity of the leucogranodiorite and the northeast 

corner of the norite of Badger Butte. At this locality 

the intraplutonic contact separates a strongly foliated 

mafic tonalite on the west with a color index in a 

representative rock specimen of 23.8 and a quartz 

content of 15.3 from a nonfoliated felsic tonalite on 

the east with a color index in a typical rock specimen 

of 13.2 and a quartz content of 27.8.

Metasedimentary xenoliths in the driftless area 

occur in core tonalite near the exposure of the intraplutonic contact. Relict bedding in the xenoliths, where 

preserved, strikes about N. 15° W. and dips vertically. 

The largest xenolith is almost 1 m long and as much as 

0.35 m wide. Much larger metasedimentary xenoliths 

occur in core rocks about 300 m to the north-northwest, 

including a banded calc-silicate xenolith more than 12 

m long that strikes roughly N. 15° W. and dips 80° NE.; 

weak planar structure in the core rocks strikes roughly 

north and dips about 75° E. The source of the metasedimentary xenoliths (which include pyroxene hornfelsfacies rocks with two-pyroxene assemblages) is unknown, but their restricted distribution near the south 

end of the intraplutonic contact suggests that in this 

general vicinity an earlier wallrock contact of the batholith with metasedimentary rocks was obliterated during the emplacement of core magma. No trace of this 

probable former contact occurs to the south at the 

present level of erosion. Core tonalite to the south is in 

sharp contact with the norite of Badger Butte with no 

intervening metasedimentary rocks.

The intraplutonic contact, concealed within 450 m of 

the norite of Badger Butte, terminates at the northeast 

corner of the norite as determined by the contrasting
70 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

fabric and modal compositions of granitic rocks on the 

north (border rocks) and east (core rocks) sides of the 

norite. Strongly foliated tonalite north of the norite has 

in three specimens a color index of 24.9 to 30.6, a 

quartz content of 12.6 to 15.3 volume percent, and only 

trace amounts of potassium feldspar, whereas nonfoliated or weakly foliated (one locality) tonalite within 90 

m of the east side of the norite has in three specimens 

a color index of 13.9 to 16.0, a quartz content of 23.2 

to 27.5 volume percent, and a potassium feldspar content of 0.6 to 1.9 volume percent.

The dip of the intraplutonic contact, whether outward as in classical cauldron subsidence or inward, 

can only be indirectly approximated from attitudes of 

limited planar structure in core rocks and even more 

restricted bedding in local metasedimentary inclusions near the norite of Badger Butte. A nearly vertical dip for the intraplutonic contact is suggested by 

planar structure in core rocks near the concave side 

of the contact in the vicinity of lat 45° N. (fig. 2.2). 

Near the southern extremity of the intraplutonic contact, bedding in metasedimentary xenoliths in core 

rocks is vertical or dips about 80° E. Poor planar 

structure in associated tonalite also dips steeply eastward. In review, field evidence for the outward or inward dip of the intraplutonic contact is inconclusive.

ROCKS OF BALD MOUNTAIN BATHOLITH

For convenience, rocks of the Bald Mountain batholith are divided into those of the central area that 

constitutes most of the batholith (fig. 2.2) and those of a 

northeastern peripheral area (fig. 2.3) that is small but 

encompasses diverse rock types and includes most of 

the heretofore-unmapped intrusive units. Rocks of the 

two areas are described separately in two mafic-to-felsic 

sequences commencing with the intrusive-rock units 

(fig. 2.2) that have been restudied or discovered since 

1957. The reexamined units are the norite of Badger 

Butte and the two units of the compositionally zoned 

Elkhorn pluton, the tonalite of Bald Mountain and the 

granodiorite of Anthony Lake (names modified from 

Taubeneck, 1957). Herein-named additional units within the Elkhorn pluton are the quartz diorite of Limber 

Creek, leucogranodiorite of Trail Creek, and granite of 

Clear Creek. The mafic-to-felsic intrusive sequences 

were established on the basis of field relations.

The newly named units in the northeastern peripheral area (fig. 2.3) are the quartz diorite of Wolf 

Creek, boundary quartz diorite unit, granodiorite of 

Indiana Mine Road, granodiorite of Beaver Meadow, 

granite of Anthony Butte, and leucogranite of Dutch 

Creek. Granitic rock units mapped in figure 2.3 are 

distinctly different from other granitic intrusive rocks

in northeastern Oregon. Accordingly, the format used 

in describing each rock unit concludes with a section 

in which the distinguishing features of the unit are 

summarized.

NORITE OF BADGER BUTTE

SETTING AND GENERAL DESCRIPTION

The norite of Badger Butte (Taubeneck, 1957) is the 

largest of the three small gabbroic bodies that are the 

earliest intrusive units in the batholith. Limited access, poor exposures in most areas, and contacts that 

were almost entirely concealed by extensive deposits 

of glacial drift restricted fieldwork during the 1950's 

(Taubeneck, 1957). Modern roads with deep roadcuts, 

and clear-cut logging of dense forests of lodgepole pine 

have disclosed much bedrock since about 1970. A new 

study of the norite of Badger Butte was motivated by 

the desire to more accurately determine its distribution, to collect specimens from a larger area than was 

originally possible, and to obtain specimens that contain few or no metamorphic minerals.

The dark color of the norite of Badger Butte is 

caused primarily by the grayish plagioclase, which is 

the dominant mineral. The rock is medium grained 

with plagioclase and pyroxene as much as 4.0 mm in 

length. Poikiloblastic biotite is commonly 10 cm 

across and rarely 20 cm. The norite generally exhibits a parallelism of plagioclase and a less pronounced 

alignment of pyroxene, but the dark color of the rock 

tends to obscure the foliation. Seventeen specimens 

were collected for petrographic study. Modal data given in table 2.1 are for the five rocks with the lowest 

amounts of postconsolidation biotite and amphiboles; 

the new data constitute a much closer approximation 

of the composition of the original norite than did previously published modes (Taubeneck, 1957).

The norite of Badger Butte, about 6 km2 in areal 

distribution, is surrounded by the tonalite of Bald 

Mountain (fig. 2.2). Deep roadcuts that expose the tonalite-norite contact both to the northeast and to the 

southeast of Badger Butte reveal that the norite was 

intruded by the tonalite. Also, dikes of tonalite within the norite contain inclusions of norite.

MICROSCOPIC DESCRIPTION

The general petrographic features of the norite of 

Badger Butte were given by Taubeneck (1957). Additional descriptive data recorded herein pertain exclusively to deformation and crystallization histories.

Deformation features in plagioclase are more pronounced in the norite of Badger Butte than in any
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 71

TABLE 2.1. Modes of the norite of Badger Butte

[Values in volume percent. Each analysis represents at least 2,400 point counts for each of two thin sections; , not present]

Specimen 

number

o o

84 

a R

:? O 

QQ 

Quartz

- 0

- 4

- 4

g 

6 

3 

6

 5

Plagioclase

65.5 

67.5 

69.6 

66.5

C~7 Q

Hypersthene

18.5 

19.7 

11.1 

10.8

T 1 "7

Augite

9.1 

9.4 

3.9 

7.2

Q 1

Hornblende

3. 

1. 

5. 

4. 

1

5 

0 

8 

5

T

Biotite

0.7

4.2 

4.4

Opaque 

minerals

1.4 

1.5 

.8 

1.4

T A

Nonopaque 

accessory minerals

0 4 

3 

3 

6

A

Average 2.1 67.4 15.6 3.6 1. 1.3 .4

other unit of the batholith. Plagioclase is recrystallized into groups of small and nearly equant crystals 

at pressure points along margins of crystals where 

plagioclase contacts other plagioclase or pyroxene. 

Twinning lamellae of plagioclase commonly are bent 

as much as 10°, less commonly as much as 20°, and 

rarely more than 25°. Some plagioclase crystals in 3 

of the 17 rocks studied are deformed into S shapes in 

which bending generally is most extreme near the 

ends of crystals. More intense plastic deformation occurs in a few crystals that show three curve reversals 

in the trends of twin lamellae. Whereas plagioclase 

shows the effects of strong plastic deformation, pyroxene commonly is fractured; an exception is one 

2.1-mm-long prism of hypersthene that is bent 24°.

Hypersthene crystallized early; many small crystals (0.1 to 0.5 mm) are enclosed within plagioclase. 

Augite surrounds or partly surrounds hypersthene, 

but inclusions of augite in plagioclase are not common. Hypersthene commonly is replaced partly by 

cummingtonite, whereas augite is replaced partly by 

actinolite. Some rocks contain no fibrous amphibole, 

but others contain only relics of pyroxene.

Hornblende and biotite are present as undeformed, 

poikiloblastic crystals that commonly continue as 

narrow extensions along grain boundaries between 

plagioclase (original or recrystallized), pyroxene, 

quartz, iron oxide, and apatite. In some thin sections, 

either hornblende or biotite along a microfracture is 

in optical continuity respectively with hornblende or 

biotite that is intersected by the same microfracture. 

Relations along microfractures support a postconsolidation origin for hornblende and biotite.

TONALITE OF BALD MOUNTAIN AND GRANODIORITE 

OF ANTHONY LAKE

SETTING AND GENERAL DESCRIPTION

Rocks in the margin and the core of the compositionally zoned Elkhorn pluton were named the tonalite of Bald Mountain and the granodiorite of Anthony 

Lake, respectively (Taubeneck, 1957). Except along 

the intraplutonic contact, the two rock units represent gradational facies of the same intrusion (fig. 

2.2), and the only major distinction between the rock 

types is a difference in mineral proportions. Accordingly, one general description of both lithologies is 

given.

The tonalite and granodiorite are medium-grained 

pale-gray rocks containing from 9 to 30 modal percent biotite and hornblende. Granodiorite and much 

tonalite in the field are almost white in the glare of 

the sun. Granodiorite near Anthony Lake has a color 

index of 11, whereas the two ferromagnesian minerals commonly constitute as much as 25 modal percent of the tonalite in the north and near contacts of 

the Elkhorn pluton.

The crystal habit and maximum size of the major 

rock-forming minerals were recorded in the field by 

examining more than 100 large exposures of both the 

tonalite and granodiorite. Most hornblende occurs in 

euhedral prisms that commonly are as long as 8 mm 

and rarely 14 mm. The euhedral form of the hornblende indicates crystallization from a melt. Only in 

some tonalite in the margin of the pluton does the 

hornblende have megascopic features that might suggest an origin as restite (White and Chappell, 1977), 

that is, unmelted solid residue in granitic rocks that 

originated as partial melts in the deep crust. Much 

biotite occurs as euhedral books commonly as much 

as 5 mm across and rarely 8 mm across. Tabular plagioclase commonly is as long as 6 mm with some 

larger crystals present in nearly all exposures. Plagioclase crystals as long as 10 mm occur in many outcrops; the largest one measured was 14 mm. Quartz 

is interstitial with common maximum dimensions of 

5 mm and rarely 8 mm. Potassium feldspar is interstitial with poikilitic crystals in granodiorite commonly 10 mm across and rarely 16 mm. Detailed 

petrographic descriptions of the tonalite of Bald
72 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Mountain and granodiorite of Anthony Lake are 

given in Taubeneck (1957).

COMPARATIVE MODAL COMPOSITIONS AND 

CRYSTALLIZATION OF AUGITE

New modal analyses for the tonalite of Bald Mountain and granodiorite of Anthony Lake have provided 

a basis for comparing rock units or subgroups of 

rocks throughout the Elkhorn pluton in terms of 

their petrographic characteristics. Much of the new 

data is for tonalite north and west of the intraplutonic contact where relatively few specimens were collected during the early 1950's (Taubeneck, 1957). 

Special attention is paid to augite, which is generally 

rare or uncommon in tonalite and granodiorite in 

western North America.

Tonalite north of the intraplutonic contact can be 

geographically divided into a distinct subgroup of 

northern rocks that generally contain a small amount 

of augite and a southern subgroup without augite. 

Rocks with augite occur in an 18-km2 area in the 

northwesternmost part of the tonalite of Bald Mountain (north of line A-B, fig. 2.2), whereas rocks without 

augite are to the south. Modal analyses of 18 rocks of 

the northern subgroup and 12 rocks of the southern 

subgroup are given in table 2.2.

The most notable mineralogical feature of the 

northern tonalite (table 2.2) is the widespread occurrence of augite cores in hornblende, whereas augite 

is uncommon or absent in most tonalite elsewhere in 

the batholith. The augite cores within hornblende are 

generally 0.3 to 1.5 mm across, but two cores exceed 

3.0 mm in greatest dimension. The augite is sufficiently abundant to indicate that it was the dominant early mafic mineral to crystallize. Accordingly, 

the tonalitic magma must have been water-undersaturated during the early crystallization history of the 

rocks.

Augite originally was much more abundant in the 

crystallizing tonalite than the small modal percentage shown in table 2.2 implies. The formerly larger 

amounts of augite within hornblende in each of the 

18 rocks is indicated by hornblende with bleached 

cores that enclose minute crystals of quartz (fig. 2.12; 

Taubeneck, 1964). The bleached cores in hornblende 

represent an advanced stage in the replacement of 

augite by hornblende (Taubeneck, 1967), and the 

common quartz inclusions represent silica liberated 

during the conversion of pyroxene to amphibole. The 

small inclusions of quartz in some cores are almost 

vermicular in shape rather than mostly oval, elliptical, or elongated as shown in figure 2.12. Generally 

the sieve-like areas of quartz in hornblende are distinctly bleached; some are only faintly bleached; a 

few are not associated with a change in color of the 

amphibole.

The original crystal habit of augite is best seen in 

longitudinal sections rather than in nearly equidimensional cross sections, in which the crystal faces 

are blurred by even comparatively minor growth of 

hornblende. The 36 thin sections used for modal 

analysis of rocks of the northern subgroup (2 thin 

sections of each of the 18 rock samples) contain six 

prismatic crystals of augite ranging from 2.3 to 3.2 

mm in length, surrounded by 0.03- to 0.15-mm rims 

of hornblende. The thinnest rims of hornblende invariably are along the prismatic faces; the rims are 

sufficiently thin to conclude that four of the augite 

crystals were initially bounded by euhedral prisms. 

None of the original faces that terminated the six 

prismatic crystals are clearly defined, but the approximate location of the faces that bounded two of the 

crystals is apparent. The two augite crystals that 

most nearly retain their original shape were either 

euhedral or almost euhedral before replacement by 

hornblende commenced.

Nine small inclusions of augite in plagioclase in the 

36 thin sections also verify the early crystallization of 

augite. The largest inclusions are about 0.3 mm 

across; all inclusions are either subhedral or euhedral.

FIGURE 2.12. Sketch of hornblende 

crystal with bleached core enclosing minute inclusions of quartz 

(white); black areas represent magnetite inclusions. Width of hornblende crystal is 0.8 mm.
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 73

TABLE 2.2. Modes of the tonalite of Bald Mountain north of the intraplutonic contact

[Values in volume percent. Each analysis represents at least 2,400 point counts for each of two thin sections; 0.0, <0.05; , not present]

Specimen Quartz 

number

Potassium 

feldspar

Plagioclase Biotite Hornblende Augite Opaque Nonopaque 

minerals accessory minerals

Tonalite of northern subgroup

±\ j. j 

R126 - 

R138 - 

R139 - 

R140 ' - 

R141 - 

R148 

R152 

R155 

R156 

R157 - 

R158 - 

R164 - 

R165 

R166 - 

R167 - 

R170 

R760 

Average

17.7 

17.5 

20.6 

21.7 

18.1 

23.4 

22.8 

20.6 

23.5 

24.7 

17.3 

21.2 

20.1 

23.2 

20.7 

20.4 

20.5 

16.4

20.5

1.3 

.0 

.9 

2.5 

1.0 

2.3 

3.1 

1.1 

2.6 

.4 

.2 

2.4 

2.2 

2.1 

1.1 

.3 

2.4 

.3

1.5

54.8 

58.4 

57.1 

56.2 

54.3 

52.8 

50.0 

59.5 

50.2 

50.4 

56.2 

53.4 

50.6 

52.6 

52.4 

53.8 

53.4 

58.4

54.1

11.9 

10.4 

10.6 

10.0 

12.0 

10.7 

10.1 

12.9 

11.7 

13.3 

11.4 

11.3 

12.0 

9.6 

11.9 

11.3 

10.8 

8.8

11.2

13.1 

13.0 

10.1 

9.1 

13.7 

10.2 

13.5 

5.0 

11.0 

10.4 

14.1 

11.2 

14.7 

12.2 

13.3 

13.5 

12.3 

15.0

12.0

0.3 

.1 

.2 

.0 

.3 

.2 

.3 

.0 

.6 

.0 

.4 

.2 

.1 

.1 

.0 

.3 

.1 

.5

.2

0.5 

.5 

.2 

.2 

.3 

.2 

.1 

.7 

.2 

.5 

.3 

.1 

.1 

.1 

.3 

.3 

.3 

.4

.3

0.4 

.1 

.3 

.3 

.3 

.2 

.1 

.2 

.2 

.3 

.1 

.2 

.2 

.1 

.3 

.1 

.2 

.2

.2

Tonalite of southern subgroup

PI i _____

±\ j y

R40 

R41 

R43 

122 

133 

134 

137 

220 

Average

24.8 

18.9 

22.3 

17.6 

19.8 

22.2 

18.9 

22.9 

22.2 

22.6 

20.3 

23.8

21.4

2.0 

.3 

3.7 

.9 

.2 

3.2 

1.4 

1.0 

3.4 

1.6 

2.2 

.7

1.7

55.8 

55.3 

54.7 

56.7 

55.2 

52.3 

53.1 

54.4 

54.6 

50.3 

56.5 

59.9

54.9

10.5 

8.9 

8.9 

9.4 

9.0 

10.5 

17.2 

10.5 

8.5 

12.5 

8.1 

9.3

10.3

6.4 

15.8 

10.1 

15.0 

15.2 

11.3 

8.9 

10.5 

10.9 

12.4 

12.4 

6.0

11.2

 

 

 

 

:::

0.3 

.6 

.2 

.2 

.5 

.4 

.4 

.5 

.3 

.5 

.4 

.2

.4

0.2 

.2 

.1 

.2 

.1 

.1 

.1 

.2 

.1 

.1 

.1 

.1

.1

Tonalite near line A-B

R150 - 

R151 - 

R171 -

Average

14.8 

14.4 

13.5

14.2

0.0 

.0 

.1

.0

55.2 

55.3 

54.7

55.1

12.1 

11.1 

11.4

11.5

17.3 

18.6 

19.7

18.5

0.1 

.0

.1 

.1

0.3 

.4 

.4

.4

0.2 

.2

.1

.2

The volume percentage of augite in the crystallizing magma can be crudely approximated by point 

counting the areas of bleached cores, the local sieved 

areas that are not bleached, the areas of hornblende 

between two or more augite relicts that extinguish 

simultaneously, and all relict augite as well as inclusions of augite in plagioclase. This method indicates 

the presence of from 0.4 to 2.3 volume percent augite

during the early crystallization history of the northernmost subgroup of tonalites in the Elkhorn pluton. 

Unit 1 of the Wallowa batholith (fig. 2.11) provides 

the most complete documentation for comparative purposes of the early crystallization of augite in the tonalitic magmas of northeastern Oregon. Augite relicts in 

hornblende are common in 78 thin sections of 39 rock 

specimens, as are small inclusions of euhedral to
74 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

subhedral augite in plagioclase. Inclusions of augite in 

plagioclase in unit 1 commonly are larger than any in 

the Elkhorn pluton; 11 inclusions are from 0.3 to 0.7 

mm in greatest dimension, and one is 0.9 mm across. 

The procedure used to determine the original content 

of augite in the northern tonalites of the Elkhorn pluton indicates that six tonalites of unit 1 of the Wallowa 

batholith initially contained from 1.9 to 3.1 volume 

percent augite.

Augite cores in hornblende were not seen in thin 

sections of the southern subgroup of tonalites (table 

2.2), but the former presence of augite is disclosed in 

most rocks by hornblende with bleached cores that 

enclose minute crystals of quartz. Accordingly, augite 

was the first mafic mineral to crystallize throughout 

the southern area of tonalite as well as in the northern area. Lack of augite cores in the southern rocks 

could be a function of earlier water saturation than 

in the northern rocks and consequent replacement of 

augite by hornblende.

The average modal compositions of the northern 

and southern subgroups of tonalites (table 2.2) show 

an abnormally small increase in the felsic components 

of the rocks southward toward the intraplutonic contact (or toward the central area of the Elkhorn pluton). An explanation for the small change in modal 

composition inward is complicated by overlying Cenozoic volcanic rocks on the north and west that conceal 

both interior and peripheral tonalites. Border tonalites on the east were mostly removed during intrusion of the granite of Anthony Butte. In addition, an 

explanation for the small change in modal composition inward is hindered by poor exposures throughout 

the area of tonalite.

The northern subgroup of tonalites may consist of 

an earlier pulse of magma that slightly preceded the 

emplacement of the southern subgroup of tonalites. 

This possibility requires an internal intrusive contact 

(approximating line A-B, fig. 2.2) and derives modal 

support from low contents of quartz and potassium 

feldspar and high contents of hornblende in the three 

rock samples (R150, R151, and R171, table 2.2) collected nearest the postulated contact. The northern 

subgroup of tonalites, in addition to containing augite, is characterized by plagioclase that is antiperthitic as seen in at least a few crystals in nearly 

all thin sections. Moreover, the plagioclase in many 

thin sections is clouded (Poldervaart and Gilkey, 

1954). Distinctions between the plagioclase in the 

northern and southern tonalites are consistent with 

the probability of two tonalitic intrusions separated 

by an intrusive contact near the line A-B.

The systematic changes in modal mineralogy that 

typically occur inward from wallrock contacts of the

Elkhorn pluton are well defined in the triangular 

area of tonalite west of the intraplutonic contact and 

north of the norite of Badger Butte (fig. 2.2). Tonalites in this area can be divided, from south to 

north, into three subgroups (table 2.3) that are at 

distances from the norite contact of 0.3 to 0.5 km, 1.0 

to 2.0 km, and 2.2 to 3.2 km. The most apparent systematic modal changes (table 2.3) northward in the 

three subgroups are an increase in potassium feldspar from trace amounts to several volume percent, 

an increase in quartz from about 14 to 21 volume 

percent, and a decrease in hornblende from about 

18.5 to 10.5 volume percent. Similar changes typically occur inward from the margins of other tonalitic 

intrusions in northeastern Oregon, although the 

changes are not necessarily as pronounced for an 

equivalent horizontal distance as those north of the 

norite of Badger Butte. Moreover, in the Cornucopia 

tonalite unit in the Wallowa Mountains, the inward 

increase in quartz and decrease in hornblende is accompanied by almost no increase in potassium feldspar (Taubeneck, 1967).

Outcrops of tonalite near the south contact of the 

Elkhorn pluton were sampled where the contact is 

best exposed to determine the modal composition of 

border tonalite in contrast to other tonalite and also 

in contrast to granodiorite in the core of the pluton. 

Eight specimens from within 40 m of the contact 

were collected westward for 12 km from Hunt Mountain (fig. 2.2). Previously published modal analyses 

(Taubeneck, 1957) for tonalite sampled throughout 

most of the Elkhorn pluton indicate that the eight 

border-rock samples (table 2.4) contain minimum 

amounts of quartz and potassium feldspar and high 

amounts of hornblende. Augite cores in hornblende 

occur in three of the eight rock samples (table 2.4); 

four of the five remaining rock samples contain hornblende with bleached cores that enclose minute crystals of quartz.

The augite cores (in three samples) and the indirect 

evidence of former augite (in four samples) prompted 

an attempt to more nearly determine the original distribution of augite in the crystallizing magma south of 

the intraplutonic contact. Reexamination of thin sections of rocks collected during the early 1950's (Taubeneck, 1957) disclosed small cores of augite in 3 rock 

samples, as well as bleached cores (indicative of 

former augite) in the hornblende of 21 other rock samples. The 24 rock samples are confined to the southern 

and eastern parts of the batholith. Forty-six additional rock samples were collected to increase sampling 

density in and near this part of the batholith. Four 

rock specimens in this group have augite cores in 

hornblende, whereas 16 other rock specimens contain
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 75

TABLE 2.3. Modes of the tonalite of Bald Mountain west of the intraplutonic contact and north of the norite of Badger Butte 

[Values in volume percent. Each analysis represents at least 2,400 point counts for each of two thin sections; 0.0, <0.05]

Specimen Quartz 

number

Potassium Plagioclase 

feldspar

Biotite Hornblende Augite Opaque Nonopaque 

minerals accessory minerals

Tonalite 0.3 to 0.5 km north of norite

77 ______ 

79 

Average

74 

75 

1 c _

Average

15. 

12. 

13.

13.

17. 

18. 

17.

17.

3 

6

7

9

9 

3 

2

8

0.1 58.6 

.1 55.8 

.0 55.8

.1 56.7

Tonalite 1

.0 53.4 

.2 55.5

6

10 

10

9

.2 

.6 

.2

.0

17. 

19. 

18.

18.

6 

7 

3

5

1.1 

.3 

1.2

.9

0.8 

.6

.5

.6

0.3 

.3 

.3

.3

.0 to 2.0 km north of norite

11 

12 

11

.2

.4 

.5

12. 

16. 

14.

8 

2 

3

0.1 

.3 

.2

.2

0.2 

.3

.4

.3

0.1 

.3 

.2

.2

Tonalite 2.2 to 3.2 km north of norite

/- o

70 

72 _

Average

19. 

20. 

23. 

20.

21.

4

1 

1 

8 

6

2

1.3 53.9 

2.0 53.2 

3.5 51.4 

3.3 56.6 

2.4 51.1

2.5 53.2

12 

11 

10 

11 

12

11

.1 

.4 

.9 

.9 

.5

.8

11. 

12. 

10. 

7. 

10.

10.

9 

7 

4 

0 

6

5

1.1 

.0 

.2 

.0 

.3

.3

0.2 

.2 

.2 

.2 

.3

.2

0.1 

.4 

.3 

.2 

.2

.3

TABLE 2.4. Modes of the tonalite of Bald Mountain near the south contact of the Elkhorn pluton 

[Values in volume percent. Each analysis represents at least 2,400 point counts for each of two thin sections; , not present]

Specimen Quartz Potassium Plagioclase Biotite Hornblende 

number feldspar

Bl 7.6 65.4 9.5 16.4

B2 19.4 62.6 11.2 7.5

B3 17.2 62.4 7.9 12.2

B4 18.4 63.7 9.1 8.1

B5 20.1 65.3 7.0 6.8

B6 15.2 0.2 65.8 5.6 12.5

B7 19.5 61.7 10.4 8.0

B8 18.7 1.2 55.8 11.3 11.4 

Average 17.0 .2 62.8 9.0 10.3

Augite Opaque Nonopaque 

minerals accessory minerals

0.3 0.7 0.1 

 .2 .1

.1 .2 .1

 .6 .1

 .6 .2

.1 .4 .2

the diagnostic bleached cores in hornblende that denote the former presence of augite.

Tonalite in which augite is known to have crystallized extends west from Hunt Mountain and southwest from Antelope Peak to about 5 km N. 80° W. of 

Summit Lake. Early augite crystallization is also evident in rocks that extend northward from near Summit Lake to within about 2 km of Anthony Lake, 

including the southeastern part (about 35 km2) of the 

granodiorite of Anthony Lake (fig. 2.2). Thus, all tonalite south and southeast of the granodiorite (except 

in and very near the Monumental salient) is within 

the area of augite crystallization. Additional sampling 

of tonalite would surely enlarge the area of known
76 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

augite crystallization throughout the general area 

from 5 to 13 km N. 50°-75° W. of Summit Lake. However, sampling density in the granodiorite is adequate 

to conclude that any appreciable extension of the area 

of augite crystallization farther northward in granodiorite is unlikely.

Tonalites of the Monumental salient (table 2.5) are 

distinctly more siliceous than tonalites elsewhere in 

the Elkhorn pluton at equivalent distances from peripheral contacts (tables 2.2-2.4). A precise comparison of modal data for different subgroups of border 

tonalites would necessitate that each group be sampled at identical distances from wallrocks. Accordingly, no exact comparison of modal data is feasible; 

some contacts of the Elkhorn pluton are overlain by 

Cenozoic volcanic rocks, whereas other contacts are 

not exposed within distances of 15 to 60 m. General 

comparisons of modal data are adequate, however, in 

substantiating the conspicuous compositional differences between tonalite in the salient and tonalite 

elsewhere in the Elkhorn pluton. The average modal 

content of quartz in border rocks (0.0 to 0.5 km from 

contact) of the salient is roughly between 4 and 11 

volume percent more than in other parts of the pluton. The average modal content of hornblende in border rocks of the salient is between 4 and 12 volume 

percent less than in border rocks elsewhere in the 

pluton. Interior tonalite (0.5 to 3.0 km from nearest 

contact) of the salient contains about 3 volume percent more quartz than tonalite from 1.5 to 3.5 km 

north of the south contact of the pluton in the general area north of Summit Lake.

The high modal quartz and low modal hornblende 

contents of border tonalite of the salient could have 

originated from the peripheral crystallization of a 

separate tonalitic magma that was more siliceous 

than tonalitic magmas elsewhere in the Elkhorn pluton. This possibility is supported indirectly by the interpretation from structural data that the salient 

originally was a satellite of the Elkhorn pluton.

Direct support for the crystallization of a more siliceous magma in the Monumental salient is derived 

from the absence of any evidence in hornblende for 

the former presence of augite, in contrast to evidence 

for the presence of augite in tonalite from nearby 

parts of the Elkhorn pluton. For example, in tonalite 

west of the Muddy Creek wedge and at least 3 km 

north of the south contact of the Elkhorn pluton, 

bleached cores containing minute crystals of quartz 

occur in hornblende in eight rocks containing 23.1 to 

25.6 volume percent quartz. One specimen containing 

23.6 volume percent quartz also contains augite 

relicts. Farther west, specimen 482 (table 2.5, fig. 

2.13) containing 24.1 volume percent quartz is the

nearest tonalite to the Monumental salient showing 

evidence in hornblende for the former presence of augite. The salient constitutes the largest area of tonalite in the Elkhorn pluton showing no evidence for the 

former presence of augite. At least a few rocks of the 

salient containing less than 25.6 volume percent 

quartz should show such evidence if the salient crystallized from the same magma as the tonalite to the 

north and northeast in the main part of the pluton.

Mafic tonalite along the southwest, south, east, and 

northeast contacts of the Elkhorn pluton grades inward, across a wide transitional zone, to granodiorite 

in the central part of the exposed pluton. On the north 

and northwest the intraplutonic contact abruptly separates tonalite on the north from a more felsic rock on 

the south that generally is granodiorite. Pertinent relations are defined in figure 2.13 by the content of 

potassium feldspar in specimens of tonalite, granodiorite, and rocks transitional between typical tonalite 

and typical granodiorite. Corresponding modal analyses are given in tables 2.5, 2.6, and 2.7.

An arbitrary unit boundary between the tonalite of 

Bald Mountain and the granodiorite of Anthony Lake 

is shown in figures 2.2 and 2.13; a more precise 

boundary would require the collection of additional 

samples, especially from north and northeast of Nip 

and Tuck Pass (fig. 2.2), east and northeast of Chicken Hill, and west of Antelope Peak. Moreover, the 

present study reveals that modal analyses of only 

two thin sections of specimens that are borderline in 

composition between tonalite and granodiorite will 

introduce significant errors for the purposes of nomenclature in perhaps as many as a third of the 

specimens. Potassium feldspar contents in two thin 

sections of the same rock differ by less than 15 percent of the smaller value in about half of the rock 

samples but by more than 45 percent in a quarter of 

the rock samples. Trial and error shows that variations by even 45 to 95 percent in potassium feldspar 

content from one thin section to another are of little 

consequence in a mode that is the average of tabulations from five thin sections. Therefore, a rigorous determination of the geographic distribution of the 

tonalite of Bald Mountain and the granodiorite of 

Anthony Lake would require perhaps as many as five 

thin sections for all specimens in the gradational 

zone between tonalite and granodiorite. Other considerations pertaining to modal variations in thin sections of granitic rocks were reviewed by Emerson 

(1964).

Distinct mineralogic and textural differences occur 

in rocks along the south side of the intraplutonic contact. Granodiorite predominates, and most textural 

differences occur within 250 m of the contact. Some
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 77

TABLE 2.5. Modes of the tonalite of Bald Mountain in the border and interior of the Monumental salient, in 

the area of the neck between the salient and the main part of the Elkhorn pluton, and in the remaining parts 

of the Elkhorn pluton

[Values in volume percent. Each analysis represents at least 2,400 point counts for each of two thin sections; 0.0, <0.05; , not present]

Specimen Quartz 

number

Potassium 

feldspar

Plagioclase Biotite Hornblende Opaque Nonopaque 

minerals accessory minerals

Border tonalite (0.0 to 0.5 km from contact) of the Monumental salient

404 

411 -----

413 

414 

415 

416 

417 

418 

419 

A £Lf\ ____

469 

AQH ««. 

491 

493 

Average

21.5 

25.5 

25.5 

23.9 

27.9 

25.4 

26.0 

26.7 

24.7 

27.3 

19.6 

20.6 

24.9 

28.0 

24.3

24.8

0.8 

1.3 

.1

.4 

.2 

.8 

.8 

1.9

.1

.4

62.8 

60.6 

62.8 

61.3 

54.7 

56.0 

56.9 

59.2 

55.7 

57.6 

62.5 

61.2 

60.4 

55.7 

57.3

59.0

9.9 

7.0 

5.7 

4.5 

11.5 

11.7 

10.5 

8.3 

11.2 

8.3 

8.9 

8.8 

9.1 

9.7 

10.6

9.1

4.8 

5.4 

5.2 

10.0 

5.7 

6.4 

6.0 

4.6 

7.4 

4.2 

8.7 

9.2 

5.3 

6.3 

7.5

6.4

0.1 

.1 

.5 

.2 

.0

.3 

.3 

.1 

.5 

.2 

.1 

.1 

.1 

.1

.2

0.1 

.1 

.2 

.1 

.2 

.1 

.1 

.1 

.1 

.2 

.1 

.1 

.2 

.1 

.2

.1

Interior tonalite (more than 0.5 km from contact) of the Monumental salient

400 

401 

403 

405 

406 

407 

408 

409 

410 

423 

426 

428 

429 

463 

A £L A

467 

468 

483 

492 

Average

25.7 

26.6 

27.4 

28.5 

28.1 

27.5 

27.2 

28.0 

24.9 

22.7 

25.5 

28.4 

27.2 

25.3 

24.4 

29.0 

27.3 

27.3 

23.2 

29.8 

28.7 

26.1 

26.9

26.8

0.2 

1.5 

1.0 

1.4 

.2 

1.2 

1.2 

.8 

1.9 

.1 

.4 

.7 

1.0 

1.6 

.8 

.5 

.6 

.5 

1.1 

.7 

3.1 

2.7 

.6

1.1

59.9 

59.6 

57.6 

56.8 

58.2 

61.0 

59.8 

60.1 

61.5 

58.8 

60.6 

60.3 

60.8 

59.3 

60.5 

59.4 

58.7 

58.9 

58.1 

56.4 

57.5 

54.6 

57.5

58.9

8.9 

8.7 

9.4 

9.6 

8.9 

6.5 

7.2 

7.2 

7.6 

10.1 

7.5 

5.0 

7.5 

7.4 

9.1 

7.3 

8.1 

8.2 

10.0 

6.7 

8.4 

9.2 

10.0

8.2

5.0 

3.1 

3.6 

3.3 

4.1 

3.2 

3.8 

3.4 

3.7 

7.7 

5.2 

5.2 

3.3 

5.8 

4.7 

3.6 

5.2 

5.0 

7.5 

5.5 

2.1 

7.3 

4.7

4.6

0.2 

.4 

.7 

.3 

.4 

.5 

.6 

.3 

.3 

.4 

.7 

.3 

.1 

.5 

.4 

.1 

.0 

.0

.7 

.1

.1 

.3

0.1 

.1 

.3 

.1 

.1 

.1 

.2 

.2 

.1 

.2 

.1 

.1 

.1 

.1 

.1 

.1 

.1 

.1 

.1 

.2 

.1 

.1 

.2

.1

Tonalite in the area of the neck between the salient and the main part of the Elkhorn pluton

442 

444 

446 

449 

462 

475 

Average

24.5 

24.0 

27.3 

25.1 

24.3 

20.7 

24.2 

23.9

24.2

1.5 

.9 

2.0 

1.4 

1.8 

1.5 

.1 

2.1

1.4

61.2 

60.2 

59.4 

60.2 

56.1 

59.9 

61.2 

61.6

60.0

8.1 

8.9 

7.5 

8.2 

7.9 

7.7 

6.5 

7.9

7.8

4.7 

6.0 

3.7 

5.1 

9.7 

10.0 

7.9 

3.8

6.4

.0

.1 

.1 

.0 

.6

.1

0.0 

.0 

.1 

.0 

.1 

.1 

. 1 

.2

.1

Table continued on next page.
78 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 2.5. Modes of the tonalite of Bald Mountain in the border and interior of the Monumental salient, in 

the area of the neck between the salient and the main part of the Elkhorn pluton, and in the remaining parts 

of the Elkhorn pluton Continued

[Values in volume percent. Each analysis represents at least 2,400 point counts for each of two thin sections; 0.0, <0.05; , not present]

Specimen Quartz 

number

Potassium 

feldspar

Plagioclase Biotite Hornblende Opaque Nonopaque 

minerals accessory minerals

Tonalite in the remaining parts of the Elkhorn pluton

31 

1 O .-

*} A _ 

35 ______

37 

^Q 

45 ______

46 

62 

*7Q -»

O 17

100 

158 

167 

169 

174 

220 

246 

251 

252 

255 

256 

257 

258 

259 

264 

265 

271 

470 

471 

472 

473 

482 

530 

Average

27.1 

20.1 

23.7 

25.4 

26.0 

24.4 

19.8 

23.9 

22.9 

27.3 

23.5 

23.2 

27.5 

25.6 

19.7 

24.4 

26.7 

26.4 

10.6 

23.8 

21.3 

20.3 

26.3 

21.3 

21.6 

22.5 

19.2 

19.8 

19.6 

20.5 

20.6 

21.5 

21.2 

17.8 

20.9 

25.9 

26.1 

19.4 

18.1 

24.1 

24.1

22.5

3.2 

.8 

3.4 

2.4 

3.5 

3.2 

1.8 

2.3 

2.8 

1.4 

3.1 

.6 

1.9 

1.4 

.6 

1.8 

2.2 

3.1

.7 

1.2 

1.3 

1.4 

.8 

3.6 

1.6 

1.8 

.4 

.2 

.7 

.7 

1.7 

1.5 

1.1 

.3 

2.3 

1.6 

.5

3.7 

4.1

1.7

53.6 

57.4 

58.3 

60.7 

56.6 

56.3 

58.6 

55.4 

55.9 

57.4 

57.4 

60.6 

53.3 

58.2 

55.9 

54.6 

55.4 

57.5 

59.8 

59.9 

61.2 

56.7 

59.2 

62.1 

55.8 

59.7 

58.9 

57.2 

57.0 

59.7 

58.6 

57.0 

58.7 

60.2 

57.1 

56.5 

57.4 

58.4 

50.0 

59.2 

59.8

57.6

9.6 

5.7 

8.6 

6.7 

6.8 

8.3 

7.5 

8.4 

11.0 

10.4 

9.1 

8.3 

9.2 

7.5 

9.2 

9.5 

8.7 

7.9 

10.9 

9.3 

6.7 

10.1 

7.2 

8.3 

10.1 

7.9 

8.2 

10.4 

9.4 

9.3 

10.7 

9.4 

8.1 

9.0 

9.5 

8.4 

7.2 

9. 1 

11.9 

7.8 

5.9

8.7

6.0 

15.3 

5.2 

4.1 

6.2 

7.1 

11.4 

9.0 

6.4 

3.1 

6.0 

6.7 

6.6 

6.4 

14.3 

9.2 

6.3 

4.5 

17.9 

6.0 

9.2 

10.9 

5.2 

7.4 

8.0 

8.0 

11.2 

11.4 

13.0 

9.4 

8.9 

9.9 

9.9 

11.5 

11.2 

6.3 

7.3 

12.0 

19.5 

4.1 

4.9

8.8

0.4 

.5 

.5 

.5 

.6 

.5 

.6 

.7 

.7 

.2 

.7 

.4 

1..0 

.5 

.2 

.4 

.6 

.4 

.6 

.2 

.3 

.6 

.5 

.0 

.8 

.2 

.6 

.7 

.6 

.2 

.4 

.3 

.5 

.3 

.8 

.4 

.3 

.4 

.4 

.9 

.9

.5

0.1 

.2 

.3 

.2 

.3 

.2 

.3 

.3 

.3 

.2 

.2 

.2 

.5 

.4 

.1 

.1 

.1 

.2 

.2 

.1 

.1 

.1 

.2 

.1 

.1 

.1 

.1 

.1 

.2 

.2 

.1 

.2 

.1 

.1 

.2 

.2 

.1 

.2 

.1 

.2 

.3

.2

rocks are considerably finer in grain size than normal granodiorite, whereas others are slightly coarser. 

Mineralogic differences are less obvious, but well-defined megascopic variations in color index and, especially, hornblende content occur at least as far as 450 

m from the contact, and variations in mineral proportions not recognized in the field are apparent by 

modal analysis. For example, 3 of 10 specimens from 

a 4-km area just south of the intraplutonic contact 

are tonalites containing 2.0, 2.4, and 3.0 volume percent potassium feldspar, respectively (specimens 160,

111, 213, table 2.7, fig. 2.13). Specimen 111 was collected only 15 m from specimen 110 (7.3 volume percent potassium feldspar) and could not be separately 

shown at the scale of figure 2.13. Five thin sections 

were used to obtain modal data for each of the three 

aberrant tonalite samples as well as for specimen 

110. Therefore, the low values for potassium feldspar 

in the three tonalite samples were adequately verified, as was the absolute difference of about 5 percent 

modal potassium feldspar between specimens 110 

and 111 collected within 15 m of each other.
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 79

A disproportionately large amount of quartz is a 

further mineralogic variation in some specimens of 

granodiorite from within 1.5 km of the intraplutonic 

contact. Modal percentages for quartz that are an average of analyses of five thin sections for each of three 

rock samples are 32.5, 31.6, and 32.6 volume percent, 

respectively (specimens 110, 159, 212, table 2.7, fig. 

2.13). The very high values for modal percentages of

118° 15' 00" 118° 07'30"

210. 

(8.9)

14 

(6.1,

'-.. 53 .196 

(8.5) ( 74)»-54 (7.0)

13 . ' < 9' 8 > 

(6.7) A10 

.A7<11 ' 3 > 

»11 

(6.7) .A11

EXPLANATION

Granodiorite of Anthony Lake 

Tonalite of Bald Mountain 

Norite of Badger Butte , 194 

(4.3)

Contact-Dashed where intraplutonic

Approximate boundary between granodiorite of 

Anthony Lake and tonalite of Bald Mountain

Sample locality-Showing sample number and 

(in parentheses) potassium feldspar 

content in volume percent

FIGURE 2.13. Sample localities and potassium feldspar contents for rock specimens of the granodiorite of Anthony Lake and the

tonalite of Bald Mountain.
80 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 2.6. Modes of rocks transitional between the tonalite of Bald Mountain and the granodiorite of

Anthony Lake

[Values in volume percent. Bach analysis represents at least 2,400 point counts for each of two thin sections]

Specimen Quartz 

number

30 

33 

o /:

39 ______

40 

41 

57 

f^Q 

59 ______ 

60 

63 

64 

a c

101 

120 ----- 

121 

135 ~

162 

163 

1 £ A

T a c

166 _____

170 

173 

1 7Q _

1 QA _- _ _

195 _____

198 

199 _____

215 ----- 

217 

269 -----

Average

23.6 

22.8 

24.1 

24.6 

29.1 

29.1 

26.5 

27.0 

26.7 

27.5 

29.2 

22.2 

28.0 

28.0 

28.3 

25.5 

27.7 

29.7 

26.2 

26.3 

26.3 

29.2 

30.0 

29.4 

28.2 

28.8 

28.8 

26.1 

28.7 

30.6 

24.7 

27.2 

27.6 

27.3 

27.9

27.2

Potassium 

feldspar

4.8 

5.3 

4.5 

6.3 

3.4 

5.7 

5.1 

5.3 

4.8 

5.3 

5.1 

2.8 

5.9 

6.0 

5.3 

2.9 

5.3 

5.5 

4.6 

4.7 

4.0 

4.8 

3.1 

5.9 

5.1 

2.9 

6.7 

5.7 

4.3 

3.4 

4.1 

3.4 

6.1 

6.8 

4.4

4.8

Plagioclase

51.4 

55.2 

55.9 

54.8 

54.7 

53.7 

57.5 

57.1 

54.9 

50.4 

51.1 

62.0 

48.6 

53.5 

54.2 

59.2 

54.9 

52.5 

53.4 

54.7 

53.8 

52.7 

51.8 

50.2 

48.3 

51.4 

53.0 

52.2 

51.9 

51.9 

55.8 

52.3 

52.4 

50.3 

53.5

53.5

Biotite

9.8 

8.3 

9.7 

10.3 

7.4 

6.8 

6.9 

5.8 

7.4 

7.9 

11.0 

6.7 

8.9 

7.9 

6.8 

7.3 

7.8 

5.9 

9.9 

8.5 

9.9 

8.6 

9.7 

7.9 

10.0 

9.4 

7.2 

9.8 

8.6 

7.5 

8.3 

8.1 

7.1 

8.1 

8.9

8.3

Hornblende Opaque Nonopaque 

minerals accessory minerals

9.4 0 

8.0 

5.0 

3.2 

4.4 

3.9 

3.4 

4.1 

5.9 

8.3 

2.9 

5.2 

8.1 

4.1 

4.7 

4.5 

3.7 

5.5 

4.9 

5.2 

5.6 

4.0 

4.7 

6.2 

7.5 

6.4 

3.8 

5.6 

6.0 

5.9 

6.2 

8.3 

6.2 

7.1 

4.8

5.5

7 

3 

6 

6 

7 

6 

4 

5 

2 

4 

5 

7 

4 

4 

5 

4 

4 

7 

7 

4 

3 

6 

4 

3 

6 

9 

4 

5 

4 

5 

6 

5 

5 

3 

3

5

0.3 

.1 

.2 

.2 

.3 

.2 

.2 

.2 

.1 

.2 

.2 

.4 

.1 

.1 

.2 

.2 

.2 

.2 

.3 

.2 

.1 

.1 

.3 

.1 

.3 

.2 

.1 

.1 

.1 

.2 

.3 

.2 

.1 

.1 

.2

.2

quartz in rock specimens such as these and the low 

values for potassium feldspar in other specimens are 

not understood, but general proximity to the intraplutonic contact suggests that the variations are related 

to this intrusive contact. Careful sampling and detailed field studies near the contact are essential before relations can be adequately appraised. The 

variations may result from perturbations in magma 

chemistry that resulted from surging and mixing prior 

to crystallization.

The most felsic rocks of the granodiorite core of the 

Elkhorn pluton occur in an area of about 2 km2 just 

west and southwest of Anthony Lake. Modal analyses 

of five rock samples from this area are given in table 

2.8. The average content of both potassium feldspar 

and quartz for these five samples is generally from 1 

to 3 volume percent higher than in rocks from other 

areas near Anthony Lake, whereas the content of 

hornblende is from 0.5 to 2.5 volume percent lower.

Many crystals of hornblende throughout most of the 

Elkhorn pluton south of the intraplutonic contact are 

strongly corroded where they are in contact with either quartz or potassium feldspar; resorption is greatest where hornblende is completely surrounded by one 

or both minerals. Obviously, hornblende was not stable in the rest magma of rocks that contain a relatively high content of quartz and potassium feldspar. The 

low average content of hornblende for the five specimens (table 2.8) near Anthony Lake together with the 

high average contents of quartz and potassium feldspar suggests that the low content of hornblende may 

be partly a function of greater resorption in this area 

rather than less crystallization of hornblende.

The most felsic rocks (table 2.8) of the granodiorite 

core contain more quartz than equivalent parts of the 

four similarly zoned units that grade from mafic tonalite to granodiorite in the Wallowa batholith (fig. 2.11), 

which is about 70 km to the northeast of the northern
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 81

TABLE 2.7. Modes of the granodiorite of Anthony Lake and the three aberrant rock specimens 

oftonalite (111, 160, 213) just south of the intraplutonic contact

[Values in volume percent. Table includes all rock specimens collected inside approximate boundary between granodiorite 

of Anthony Lake and tonalite of Bald Mountain (fig. 2.13). Rock specimen 111 was collected 15 m from specimen 110

and 213 represent at least 2,400 point counts for each of five thin sections; all other analyses represent at least 2,400 

point counts for each of two thin sections]

Specimen Quartz

4 _______

11 

13 

14 ______

15 

16 

17 

19 - 

20 

2i ______

23 ______ 

24 ______

25 

26 

27 

44 ______

47 ______

48 

49 ______

50 

52 _

53 

54 

71 ______

107 

113 

159 

161 

171 

197 

202 

204 

205 

207 

209 

2io 

2ii _____

212 

214 

A9 

AID 

All 

Average

27.5 

26.2 

24.2 

25.1 

28.9 

29.2 

25.5 

26.4 

29.2 

29.4 

27.0 

26.5 

26.7 

27.1 

26. 3 

24.1 

26.2 

25.8 

26.8 

28.7 

24.6 

28.8 

26.9 

27.6 

27.8 

29.4 

25.6 

24.7 

26.6 

28.6 

27.3 

27.6 

28.7 

28.9 

29.7 

23.6 

29.0 

27.5 

25.1 

29.8 

32.5 

29.7 

28.3 

26.9 

31.5 

27.7 

31.1 

28.8 

29.8 

27.5 

27.9 

28.4 

31.8 

27.8 

26.5 

29.2 

30.3 

32.6 

30.0 

31.0 

28. 1 

30.4 

28.2 

31.2

28.0

Potassium 

feldspar

7.3 

7.1 

4.8 

5. 3 

7.0 

5. 3 

6.8 

6.8 

6.7 

6.7 

5.1 

8.5 

7.6 

7. 2 

5.1 

7.2 

6.8 

8.4 

9.7 

6.1 

5.9 

6.6 

5.7 

8.1 

6.2 

9.7 

5.2 

7.1 

5.6 

7.1 

6.9 

7. 1 

7.4 

9.8 

6.8 

8.1 

7.0 

6.7 

5.9 

10.0 

7.3 

7.7 

5.3 

8.1 

9.4 

8.6 

8.4 

7.0 

6.6 

4.2 

6.9 

6.7 

7.3 

8.2 

7 .5 

6.9 

4.7 

12.7 

8.5 

8.2 

9.3 

7.9 

11.3 

7 .0

7.2

Plagioclase

50.4 

53.8 

55.7 

55.6 

52.1 

52.9 

51.4 

54. 1 

53.5 

53.6 

56.3 

53.9 

51.8 

52.8 

55.2 

55.3 

53.9 

55.3 

53.2 

53.3 

59.5 

52.3 

56. 6 

53.8 

56.4 

47.2 

55.1 

56.1 

54.3 

51.8 

49.3 

49.8 

50.6 

48.8 

52.4 

56.7 

51.8 

55.6 

55.0 

51.2 

49.2 

51.3 

55.3 

52.5 

48.5 

50.6 

47.6 

47.9 

51. 3 

55.6 

50.4 

51.4 

47.6 

50.1 

54.9 

51.2 

54.5 

45.3 

49.8 

49.7 

51.3 

49.1 

49.5 

50.2

52.4

Biotite

9.8 

7.6 

9.5 

9.5 

6.9 

7.5 

9.6 

7.2 

6.1 

7. 1 

6. 1 

7.4 

7.9 

8. 1 

7.4 

8.3 

8.7 

6.9 

7.3 

8.9 

6.4 

6.6 

6.9 

7.3 

4.8 

7. 1 

7.2 

7. 2 

7.5 

6.4 

9.8 

8.6 

5.9 

7.8 

6.3 

7.1 

7.7 

6.0 

7.2 

5.5 

8.0 

7.9 

5.6 

7.3 

7.7 

7.4 

8.6 

9.1 

8.0 

6.0 

8.9 

7 .8 

6.9 

8. 1 

6.5 

6.6 

6.4 

6.5 

7.4 

7.7 

8.1 

7.8 

6.3 

7. 3

7.4

Hornblende

3.9 

4.7 

4.8 

3.5 

4.2 

4.3 

5.7 

4.8 

3.5 

2.7 

4.7 

3. 1 

4.7 

4.1 

4.8 

4.4 

3.8 

2.9 

2.4 

2.4 

2.7 

4.7 

3.1 

2 . 1 

4.1 

5.7 

5.7 

3.9 

5. 3 

5.3 

5.7 

6. 1 

6.6 

3.7 

4.1 

3.9 

3.8 

3.5 

5.8 

3.0 

2.4 

2.3 

4.7 

4.5 

2.4 

5.0 

3.8 

6.4 

3.6 

5.6 

5.0 

5.0 

5.7 

4.9 

4.1 

5.8 

3.6 

2.4 

3.6 

2.7 

2.6 

4.2 

4.2 

3.2

4. 2

Opaque Nonopaque 

minerals accessory minerals

0.8 

.4 

.8 

.8 

.5 

.5 

.8 

.6 

.7 

.3 

.5 

. 5 

1.1 

.5 

.8 

.4 

.4 

.6 

.4 

.4 

.7 

.7 

.5 

.9 

.6 

.6 

.8 

.7 

.5 

.6 

.7 

.6 

.6 

.7 

.5 

.4 

. 5 

.5 

.8 

.3 

.5 

.8 

.5 

.5 

.4 

.5 

.4 

. 6 

.5 

.8 

.7 

.5 

.5 

.6 

.3 

.2 

.4 

.3 

.5 

.5 

.5 

.5 

.4 

.8

.6

0.3 

.2 

.2 

.2 

.4 

.3 

.2 

.1 

.3 

.2 

.3 

.1 

. 2 

.2 

.4 

. 3 

.2 

. 1 

.2 

.2 

. 2 

. 3 

.3 

.2 

. 1 

. 3 

.4 

.3 

.2 

.2 

. 3 

.2 

. 2 

.3 

.2 

. 2 

.2 

. 2 

.2 

. 2 

.1 

.3 

.3 

.2 

. 1 

.2 

. 1 

.2 

.2 

.3 

.2 

.2 

.2 

. 3 

.2 

. 1 

. 1 

. 2 

.2 

. 2 

. 1 

.1 

.1 

. 3

.2

Aberrant specimens of tonalite just south of the intraplutonic contact

111 

160 - 

213 

25.7 

28.5 

26.7

2.4 

2.0 

3.0

59.2 

52.5 

58.3

7.6 

9.6 

7.6

4.7 

6.7 

3.8

0.3 

.5 

.4

0.1 

.2 

.2

Average 26.9 2.5 56.6 1.3 5.1
82 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 2.8. Modes of the granodiorite of Anthony Lake in the Bald Mountain batholith near Anthony Lake 

compared with modes of core granodiorite of units 2 and 4 of the Wallowa batholith, northeastern Oregon

[Values in volume percent. Each analysis represents at least 2,400 point counts for each of two thin sections]

Specimen Quartz 

number

Potassium 

feldspar

Plagioclase Biotite Hornblende Opaque 

minerals

Nonopaque 

accessory minerals

Average 25.0

Core granodiorite (near Anthony Lake), Bald Mountain batholith

A7 31.0 8.2 49.7 7.7 2.7 0.5 0

A8 28.1 9.3 51.3 8.1 2.6 .5

f\J ~*iy.*i / . j *ij7.X / O 1 £. ~J

A10 28.2 11.3 49.5 6.3 4.2 .4 

All 31.2 7.0 50.2 7.3 3.2 .8

Average 29.8 8.7 50.0 7.4 3.4 .5

2

1

1

1 

3

2

Core granodiorite of unit 2, Wallowa batholith

116 -- 

272 -----

273 ~ 

289 ~

24.1 

25.5 

26.4

1/1 C.

9.4

9.1 

12.2 

8.4

T 1 1

49.4

48.9 

47.9

. y

9.0 

8.3 

8.9

4.6 

4.5 

7.6

c c

7 0

3 

4 

6

/

3

2 

2 

o

10.1 49.5 8.9 5.8 .5 .2

Average 27.5

Core granodiorite of unit 4, Wallowa batholith

765 

766 

767 

768 

779 

29.6 

o o n <io . U 

27.7

27.0

25.0

15.0 

TO n -Lo U

14.6

10.9

10.7

45.2

41.7 

45.1

47.5

49.5

6.5

6 Q

. y 

7.5

8.4

6.1

2.8 0

4 . 7 

4.5

5.7

7.6

6

4 

3

3

6

0.3

.3 

.3

.2

.5

13.8 45.8 7.1 5.1 .4 .3

part of the Elkhorn pluton. For example, five rock 

samples from the granodiorite core near Anthony Lake 

contain an average of 29.8 volume percent quartz, 

whereas five samples from the core of unit 2 of the 

Wallowa batholith (Taubeneck, 1987) contain an average of 25.0 volume percent quartz, and five samples 

from the core of unit 4 contain an average of 27.5 

volume percent quartz (table 2.8). The comparatively 

high content of quartz in both the granodiorite of Anthony Lake and the younger felsic intrusions of the 

Bald Mountain batholith suggests a genetic relationship between these silica-rich granitic units.

ASSIMILATION OF GRAPHITE-BEARING METASEDIMENTARY 

ROCKS IN ELKHORN PLUTON

The higher content of modal quartz in tonalite of the 

Monumental salient (table 2.5) than in tonalite elsewhere in the Elkhorn pluton apparently resulted from 

the intrusion of a separate and more siliceous magma

(see previous section, "Comparative Modal Compositions and Crystallization of Augite"). Assimilation of 

siliceous metasedimentary rocks may also have contributed to the high content of modal quartz in rocks 

of the Monumental salient. Xenoliths of siliceous metasedimentary rocks are relatively common throughout 

most of the salient, and the rounded shapes of many 

xenoliths indicate assimilation.

Thin sections of tonalite containing red-brown biotite and low amounts of modal iron oxide support the 

interpretation that metasedimentary rocks were assimilated into the Monumental salient, as well as into 

many other granitic intrusions in the Elkhorn Mountains. Both the red-brown biotite and the low modal 

iron oxide content indicate that the rocks are reduced; 

assimilation of graphite-bearing metasedimentary 

wallrocks is the logical way to explain the reduction 

(Flood and Shaw, 1975, p. 162). Notably, the most continuous graphite-bearing sequences of metasedimentary wallrocks in contact with the Elkhorn pluton
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 83

occur west and southwest of Summit Lake in the vicinity of the Monumental salient.

Much of the tonalite in and near the Monumental 

salient contains evidence of reduction: Including rock 

specimens collected during the 1950's (Taubeneck, 

1957), 14 of 23 tonalite samples collected within 400 

m of the contact of the salient contain red-brown biotite and have an average of 0.09 volume percent iron 

oxide, and 10 of 15 tonalites from within the area of 

the neck contain red-brown biotite and have an average of 0.06 volume percent iron oxide. In contrast, 

modal iron oxide is much more abundant constituting 0.26 to 1.10 volume percent in nearly all tonalite bodies in northeastern Oregon that do not contain 

red-brown or red biotite.

Further evidence of reduced rocks in the Elkhorn 

pluton is provided by an analysis of a border tonalite 

containing red-brown biotite and collected southwest 

of Summit Lake. This tonalite was analyzed for selected oxides using classical wet methods (C.O. Ingamells, 

analyst, 1966, Pennsylvania State University); the 

analysis yielded values for Fe2O3 and FeO of 0.39 and 

3.39 weight percent, respectively. These iron oxide values and the red-brown color of biotite are typical of Stype granitic rocks (Hine and others, 1978; White and 

Chappell, 1983; Whalen and Chappell, 1988), which 

are derived from the melting of sedimentary source 

rocks.

The granitic rocks of northeastern Oregon, even the 

cordierite-bearing trondhjemites of the Wallowa 

Mountains, are I-type granitoids on the basis of their 

low initial 87Sr/86Sr ratios (Armstrong and others, 

1977), low K2O/Na2O ratios, high Fe2O3 /FeO ratios, 

and the unrefutable evidence for the magmatic crystallization of cordierite in the trondhjemites (Taubeneck, 1987). Accordingly, granitic rocks in the Elkhorn 

pluton with red-brown biotite and minor amounts of 

iron oxide crystallized from I-type magmas that acquired their low Fe2O3 /FeO ratios by assimilation of 

graphite-bearing metasedimentary rocks.

Tonalite with red-brown biotite that is, reduced tonalite is present not only within the Monumental salient, but also, sporadically, eastward from the salient 

to the contact south of Hunt Mountain and northward 

to tonalite west of the norite of Badger Butte (fig. 2.2). 

Some of the tonalite in the Monumental salient lacks 

red-brown biotite and contains moderate amounts of 

iron oxide. The distribution of xenoliths in the salient 

indicates that even this unreduced tonalite assimilated metasedimentary rocks; presumably tonalite that 

is not reduced assimilated metasedimentary rocks 

that contained little or no graphite. No assessment of 

the importance of assimilation during the emplacement and crystallization of the Elkhorn pluton is possible without trace-element data and oxygen-isotope 

compositions for both metasedimentary rocks and associated border and interior tonalite.

DISTINCTIVE PETROGRAPHIC FEATURES OF ROCKS OF 

ELKHORN PLUTON

Three petrographic features that distinguish rocks 

of the Elkhorn pluton from many other granitic rocks 

are (1) very strong oscillatory zoning in plagioclase, 

(2) magnetite concentrations as much as 2 mm across, 

and (3) general absence of rutile needles in quartz.

Zoning in plagioclase in plutonic rocks south of the 

intraplutonic contact is sharp and strongly oscillatory: Except in the two salients, at least one crystal in 

every two thin sections of each rock sample generally 

exhibits 48 or more distinct oscillations. As many as 

82 oscillations occur, but more than 65 is uncommon. 

Delicate rhythmic oscillations especially characterize 

the most strongly zoned crystals. Zoning in the Monumental salient is slightly less intense than throughout the main part of the pluton, although nearly all 

rocks contain some plagioclase with 44 or more oscillations and one crystal has 72 oscillations. Zoning 

throughout the Hunt Mountain salient is less uniform than elsewhere in the pluton, but about onehalf of the rocks contain some plagioclase with at 

least 40 oscillations.

The very strong oscillatory zoning in plagioclase of 

the Elkhorn pluton is not confined to rocks well within 

the batholith but commonly persists to within less 

than 100 m of peripheral contacts with country rocks. 

For example, the maximum of 22 oscillations in plagioclase in specimen 250 (fig. 2.13), collected 6.5 m from 

the contact at Antelope Peak (fig. 2.2), is greater than 

the number of oscillations in crystals throughout 

many plutons of batholithic dimensions.

The character of the oscillatory zoning in plagioclase is a significant feature in distinguishing between different plutonic environments as well as 

between individual plutons (Hutchison, 1970). The 

four major units of the Wallowa batholith (fig. 2.11) 

provide a convenient basis for contrasting oscillatory 

zoning in four compositionally zoned plutons that 

grade inward from peripheral tonalite to core granodiorite. Each of the four units is large enough to 

qualify as a small batholith. Oscillatory zoning in 

plagioclase of 76 rock specimens from unit 1 ranges 

from less than 20 oscillations in rocks in the east to 

as many as 45 in rocks in the extreme west. Zoning 

in different parts of unit 2, as studied in 70 rock 

samples, ranges from 10 oscillations to as many as 

36. Seventeen or fewer oscillations occur in plagioclase in 84 rock samples from unit 3 and in most of
84 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

90 rock samples of unit 4. Thus, oscillatory zoning in 

plagioclase from the Elkhorn pluton is much more 

pronounced and uniform than in other compositionally zoned plutons in northeastern Oregon; it is also 

more conspicuous than in other granitic bodies such 

as those of the Prince Rupert region of British Columbia (Hutchison, 1970).

Concentrations of magnetite in large clots that contain many crystals is a second distinctive feature of 

most rocks of the Elkhorn pluton. Clots are absent or 

rare in tonalite that has assimilated adequate graphite in metasedimentary rocks to reduce the modal iron 

oxide content to less than about 0.2 volume percent. 

Clots of magnetite in most thin sections are about 1 

mm across, but many are about 2 mm across. Individual crystals of magnetite mostly are less than 0.1 mm 

across. Therefore, some clots contain more than 100 

crystals. Many crystals retain their euhedral outline, 

but others commonly have coalesced, obliterating 

most faces. Open spaces in compact aggregates of 

magnetite fairly commonly are filled with sphene.

Clots of magnetite are not numerous enough to occur in every thin section, but they are present in 

most rocks throughout the Elkhorn pluton. In contrast, other granitic plutons in northeastern Oregon 

contain few or no clusters of magnetite. Some concentrations similar to those in the Elkhorn pluton are 

present in units 3 and 4 of the Wallowa batholith 

(fig. 2.11) and in cordierite trondhjemite of the Cornucopia stock (Taubeneck, 1964).

A third distinctive feature of rocks of the Elkhorn 

pluton is the general absence of needles of rutile in 

quartz in comparison to other intrusions of tonalite 

and hornblende-bearing granodiorite in northeastern 

Oregon. (Data are insufficient to make a comparison 

with plutons outside of northeastern Oregon). Rutile 

needles in quartz in the Elkhorn pluton are estimated to occur in only about 3 percent of the rock samples. The slender needles generally are not more 

than 0.2 mm long, lack birefringence, and were not 

mentioned in previous reports (see Taubeneck, 1957). 

In contrast, long hairlike needles of rutile generally 

are common in quartz in granitic rocks elsewhere in 

northeastern Oregon. Quartz in some of these rocks 

contains needles as much as 1 mm long, and a few 

needles are considerably longer; birefringence is commonly discernible in the longer needles; scores of 

hairlike needles commonly occur in a single crystal of 

quartz; and many groups of needles are oriented.

SPECIFIC GRAVITY OF ROCKS OF ELKHORN PLUTON

The specific gravity of rocks of the Elkhorn pluton 

decreases inward in accordance with the compositional

variations indicated by modal analyses. Specific-gravity determinations for 109 specimens yielded values of 

2.78 to 2.67. Border rocks near the southern contact of 

the Elkhorn pluton and most mafic tonalite north of 

the intraplutonic contact have densities of 2.75 to 2.78, 

whereas those of other tonalites are 2.70 to 2.75. Densities of granodiorite are 2.67 to 2.70.

COMPOSITIONS OF CORES OF OSCILLATORY-ZONED

PLAGIOCLASE IN ELKHORN PLUTON AND OTHER 

GRANITIC INTRUSIONS OF NORTHEASTERN OREGON

Limited compositional data for calcic cores in oscillatory-zoned plagioclase in the Elkhorn pluton, Cornucopia stock (Taubeneck, 1964), and Wallowa batholith 

(Taubeneck, 1987) were obtained with the assistance 

of A.J. Piwinskii at the University of Chicago in 1967. 

The focus of that study was to determine the maximum anorthite content of cores of oscillatory-zoned 

plagioclase. Compositions of plagioclase feldspar were 

measured with an Applied Research Laboratory (ARL) 

electron microprobe analyzer, following the methods 

described by Smith (1965) and by Ribbe and Smith 

(1966). The pertinent plagioclase compositional data 

are recorded in the following six paragraphs, together 

with data for the cordierite trondhjemite of Crater 

Lake (Taubeneck, 1964) obtained by E.M. Bishop in 

1983 at Washington University in St. Louis.

In specimens of five border rocks of the tonalite of 

Bald Mountain, the anorthite content in cores of eight 

crystals ranged from An4y 2 to An^iQ. The most calcic 

spot compositions were found in three other crystals 

from the same group of rocks: An51 5Ab47 ^Or^ Q> 

An53 5Ab44 8Or! o, and An^ 2Ab41 0Or! 3 .

In four rock samples (six crystals) of the granodiorite of Anthony Lake, the anorthite content ranged 

from An4g 9 to An50 i. The most calcic composition 

found was An50 iAb48 5Ori 2-

The anorthite content ranged from An49 Q to An53 5 

for six crystals in five rock specimens from the interior of the Cornucopia tonalite unit (Taubeneck, 1967, 

p. 6). Analysis of a seventh crystal yielded an even 

more calcic (spot) composition: An5g 4Ab39 gOr0 7.

Core compositions for 10 crystals from five border 

rocks (Taubeneck, 1967, p. 6) of the Cornucopia tonalite 

unit ranged from An49 2 to Ang0 0. The most calcic compositions found, local spots in each of three other crystals, were An63 .iAb353Or07, An64 oAb34 2Or0. 8, and

Eight crystals in two border-rock samples of unit 4 

of the Wallowa batholith (Taubeneck, 1987) ranged in 

composition from An496 to An55 5; five crystals in 

three core-rock samples yielded a similar range of 

composition, from An49 0 to An54 8, indicating an in-
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 85

significant difference between core- and border-rock 

plagioclase.

The most calcic composition detected in four crystals in two specimens of the Tramway trondhjemite 

unit (Taubeneck, 1964), the oldest trondhjemite of the 

Cornucopia stock, was An54 2Ab45 0Or0 6- An39 7 was 

the most calcic composition detected by E.M. Bishop 

in nine crystals from six specimens of the cordierite 

trondhjemite of Crater Lake (Taubeneck, 1964), the 

youngest trondhjemite of the Cornucopia stock.

VOLUME PERCENT AND ORIGIN OF CALCIC CORES IN 

PLAGIOCLASE OF ELKHORN PLUTON

In the restite model, granitic magmas are visualized as crystal mushes that incorporate as much as 

about 60 percent of restite as single crystals (rather 

than refractory rock inclusions) of cordierite, sillimanite, garnet, zircon, and calcic cores of plagioclase 

(Chappell and others, 1987). Only zircon and calcic 

cores of plagioclase occur in the granitic rocks of the 

Elkhorn pluton. The euhedral habit of the zircon 

(Taubeneck, 1957, p. 216) indicates that the crystals 

are of magmatic origin, except for possible zircon 

cores too small to be seen with a petrographic microscope. A major linchpin of the restite model is that 

calcic cores of plagioclase in granitic rocks that contain less than 70 weight percent SiC>2 (such as granitic rocks of the Elkhorn pluton) are residuals of 

restite that have been overgrown by magmatic, oscillatory-zoned plagioclase of more sodic compositions 

(Chappell and others, 1987).

Seventeen rock specimens of the granodiorite of Anthony Lake and 11 specimens of the tonalite of Bald 

Mountain (from near the contact south and east of 

Summit Lake) were used for this exploratory study of 

the volume percent and origin of calcic cores in plagioclase of the Elkhorn pluton. The volume percent of 

calcic core was determined for crystals that show a 

large number of oscillations and apparently were cut 

through the center. To select appropriate crystals for 

examination, one must view many random slices of 

plagioclase in thin section (see Dowty, 1980). Some of 

the most favorable orientations occur in one or more 

crystals of synneusis aggregates (see Vance, 1969, figs. 

10-14), but aggregates of two or more crystals were 

excluded from consideration in this study because of 

interference during growth. Commonly not more than 

a few suitable crystals occur in a thin section. Point 

counts for each plagioclase were tabulated for the core 

and the remainder of the crystal. In each case, the 

remainder of the crystal consists of oscillatory-zoned 

andesine bordered by rim oligoclase. The rim oligoclase 

is restricted volumetrically in peripheral tonalite but

is as abundant as 35 percent by volume for some crystals in the granodiorite of Anthony Lake. For all volume 

percent determinations, a mechanical stage was adjusted to record 20,491 points for an area of 920 mm2.

Calcic cores in 36 crystals of the granodiorite of 

Anthony Lake range from 0.8 to 5.9 volume percent 

of the plagioclase. The average volume of core is 2.4 

volume percent. The crystals examined are from 0.9 

to 5.2 mm in maximum dimensions. The average core 

volume of 23 plagioclase crystals from the tonalite of 

Bald Mountain is 3.7 volume percent in a range from 

1.7 to 7.8 volume percent. The crystals examined are 

from 1.1 to 4.9 mm in maximum dimensions.

Biotite did not commence to crystallize in the 

Elkhorn pluton until just before growth of the plagioclase rim of oligoclase; quartz and potassium feldspar 

are interstitial. Therefore, irrespective of the origin of 

the cores of plagioclase in the granodiorite of Anthony Lake and the tonalite of Bald Mountain, the small 

core volumes indicate a magma that was emplaced as 

a near-liquid rather than as a crystal mush.

The following considerations were instrumental in 

the conclusion that calcic cores in the plagioclase of 

the Elkhorn pluton are of magmatic origin: The cores 

are not separated by higher birefringence from the 

surrounding oscillatory-zoned plagioclase or by an encircling corrosion rim. Polysynthetic twinning, where 

present, extends through oscillatory-zoned plagioclase 

into cores. The two-dimensional shape (mostly rectangular or squarish) of the cores closely approximates in 

gross outline the outer configuration of the plagioclase 

crystals. Furthermore, in a simple twin that exhibits 

considerable disparity in size, that part of the core 

within the longer individual is proportionally longer 

than the part within the shorter individual.

INTERLUDES: A NEWLY DESCRIBED FEATURE OF OSCILLATORYZONED PLAGIOCLASE IN GRANITIC ROCKS

Significant features of the oscillatory zoning in plagioclase of the Elkhorn pluton are interruptions in the 

rhythmic crystallization that herein are called "interludes." An interlude in oscillatory-zoned plagioclase is 

a concentration of either one or several different mafic 

minerals of small size in a concentric zone of plagioclase between consecutive oscillations. No descriptions 

of interludes were found during a literature search 

focused on oscillatory zoning in plagioclase of granitic 

rocks. A hint that interludes may occur in the Tokuwa 

batholith, central Japan, is the statement of Shimizu 

(1986, p. 63) that some plagioclase crystals have "lines 

of tiny inclusions along the crystal growth lines." In 

northeastern Oregon, examples of interludes occur in 

plagioclase of the Wallowa batholith and Cornucopia
86 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

stock, as well as in plagioclase of the Elkhorn pluton. 

It is illogical to assume that interludes are restricted 

to the plagioclase of granitic intrusions in northeastern 

Oregon, especially considering the common worldwide 

occurrence of oscillatory zoning in the plagioclase of 

granitic rocks containing less than about 70 weight 

percent SiO2 .

Interludes are subdivided for discussion into complete interludes, which are continuous and extend between oscillations around the circumference of a 

plagioclase crystal, and incomplete interludes, which 

are discontinuous and do not extend around the circumference. Incomplete interludes pinch out parallel 

to either one or two crystal faces. In this study of 20 

rock specimens (40 thin sections) of the granodiorite of 

Anthony Lake, only six interludes are complete. In two 

of the complete interludes, the only mafic mineral is 

an iron oxide, whereas in four of the complete interludes, both iron oxide and hornblende are present. The 

complete interludes contain from 41 to 168 mafic crystals. Twenty-one incomplete interludes are present; 

four contain inclusions of iron oxide only. The proportions of iron oxide and hornblende inclusions in incomplete interludes range from entirely iron oxide in some 

interludes to almost as many crystals of hornblende as 

of iron oxide in other interludes. The 21 incomplete 

interludes contain from 6 to 69 mafic crystals.

Concentrations of iron oxide and hornblende in plagioclase also occur in the cores of 121 crystals in the 

40 thin sections; 10 additional cores contain concentrations of iron oxide only. Cores that contain both mafic 

minerals have from 6 to 149 mafic crystals; cores that 

contain only iron oxide have from 7 to 21 mafic crystals. Almost invariably the cores of plagioclase that 

contain mafic inclusions are mottled, in contrast to 

cores that contain no mafic inclusions. Mottling consists of spots and blotches that exhibit different intensities of light under crossed nicols as the plagioclase is 

rotated on the stage of a microscope. A few mottled 

cores also contain small areas (square and rectangular 

in shape) of a more sodic plagioclase that is sufficiently 

different in composition to produce a distinct Becke 

line. Six of the 121 cores that contain both hornblende 

and iron oxide locally show resorption along as much 

as about 35 percent of the core periphery. Resorption 

was not detected, however, around the circumference 

of any core with concentrated mafic inclusions or any 

core without mafic inclusions.

Complete interludes differ from incomplete interludes in greater overall thickness, larger size of mafic 

inclusions, conspicuous mottling of plagioclase, and 

the common occurrence of microcracks localized within or adjacent to the interlude. The four complete interludes containing inclusions of both hornblende and

iron oxide are from about 0.3 to 0.5 mm in width, 

whereas the two complete interludes with inclusions 

only of iron oxide are about 0.05 and 0.10 mm in 

width. In comparison, each incomplete interlude is 

less than 0.05 mm in maximum width; some have a 

maximum width of less than about 0.02 mm.

Most hornblende crystals in the complete interludes are subhedral to euhedral prisms of stubby 

habit with maximum dimensions from about 0.02 to 

0.08 mm. Crystals larger than 0.12 mm are rare. 

Overall, the hornblende crystals are slightly larger 

than very similar inclusions of hornblende in the 121 

cores of plagioclase that contain concentrated inclusions of iron oxide and hornblende. Cubic crystals of 

iron oxide in the complete interludes commonly are 

about 0.02 mm on an edge. Some crystals are 0.03 

mm; a few are as much as about 0.04 mm. Similarities in size and habit between the mafic minerals in 

the plagioclase cores and those in the complete interludes suggest a genetic relationship that is much less 

apparent in the incomplete interludes, where crystallization was relatively short-lived. Crystals of both 

iron oxide and hornblende in the incomplete interludes are an estimated one-third to two-fifths as 

large as those in the complete interludes.

Point counts (20,491 points for an area of 920 mm2) 

for hornblende confirm that collectively, the smallest 

crystals occur in the incomplete interludes rather 

than in the plagioclase cores or any other part of the 

plagioclase. One hundred sixteen counts were tabulated for the 1,511 hornblende crystals in the 121 plagioclase cores, whereas only 4 points were recorded for 

the 126 crystals of hornblende in the incomplete interludes. Much larger crystals of hornblende occur in plagioclase outside of the interludes and between the 

calcic cores and outer margins of the oligoclase rims. 

The largest inclusion of hornblende between a core 

and an incomplete interlude is a euhedral prism about 

0.65 mm long and 0.15 mm wide. Larger crystals of 

hornblende (including resorbed crystals) occur between the incomplete interludes and the oligoclase 

margins. In review, the small overall size of hornblende in the incomplete interludes is explained by a 

more restricted period of growth after nucleation than 

in either the complete interludes, the calcic cores, or 

elsewhere within the plagioclase.

Each complete interlude is characterized by mottled 

plagioclase that commonly extends in an irregular 

manner beyond the interlude for a short distance into 

adjoining oscillations. The mottling obscures detailed 

examination of adjacent oscillations and therefore prevents a close determination of possible fine-scale resorption along the borders of interludes. However, no 

wholesale resorption is apparent. Locally, the mottled
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 87

interludes are further distinguished by small areas 

(square and rectangular in shape) of a more sodic plagioclase, as is also observed in some of the mottled 

cores of plagioclase that are characterized by concentrations of mafic inclusions. Mottling occurs in the incomplete interludes only where they are widest and 

hence generally does not obscure detailed relations 

along the margins of the interludes. No resorption was 

detected along the margin of any incomplete interlude.

Mottling in both complete interludes and the cores 

of plagioclase that contain concentrated mafic inclusions undermines the suggestion of Presnall and Bateman (1973) that mottled cores of plagioclase are 

evidence of restite. The cause for the local mottling of 

plagioclase in the granodiorite of Anthony Lake remains unknown, but the common occurrence of a mafic 

inclusion centered in a light blotched area repeatedly 

suggests that mottling is actuated by the presence of 

mafic inclusions. However, this interpretation is not 

entirely satisfactory because many blotches contain no 

inclusion, at least not in the plane of the thin section.

Three of the six complete interludes exhibit microcracks that are confined mostly to the interlude but 

also extend for short distances into adjacent oscillations. The microcracks possibly represent differential 

strain induced during cooling of layers of different 

plagioclase composition, enhanced perhaps by the numerous mafic inclusions within each interlude.

About three-quarters of the interludes occur within 

15 percent of the midpoint of the total number of oscillations in each plagioclase. The most consistent position of interludes relative to oscillations occurs 

about 4 km east-southeast of Anthony Lake where 

interludes in specimens 18, 19, and 20 (fig. 2.13) are 

within about 10 percent of the midpoint of the total 

oscillations in each plagioclase, with each interlude 

closer to the rim than to the core. The lateral extent 

of the interludes beyond the locations of these three 

rock specimens is unknown, but the distribution of 

interludes is sufficiently widespread to imply a significant event, or events, in the solidification of the 

magma before crystallization of quartz, biotite, potassium feldspar, and much of the plagioclase.

This survey of interludes in the granodiorite of Anthony Lake also suggests that interludes might be 

useful in studies of the progressive compositional 

changes in mafic minerals during crystallization of 

granitic rocks, as well as the crystallization histories 

of plagioclase, both locally and throughout a pluton.

PEGMATITES AND APLITES OF ELKHORN PLUTON

Pegmatites are rare except in the interior of the 

Elkhorn pluton, where they occur as small dikes less

than 0.7 m wide. Some pegmatites are less than 1.0 m 

long and not more than 10 cm wide. South and southeast of Anthony Lake a few of the larger aplite dikes, 

from 20 to 60 cm wide, contain either a continuous or 

discontinuous narrow core of pegmatite. Almost all 

pegmatites have a simple mineralogy: feldspar, quartz, 

and minor biotite. Scattered crystals of black tourmaline occur in two pegmatites.

Twenty-three pegmatites containing tourmaline, 

mostly in the Belleview wedge, were noted in country 

rocks near the contact of the Elkhorn pluton. With one 

exception, these pegmatites are small or very small.

Aplite dikes are not common in the tonalite of Bald 

Mountain except in the northeast; they are much 

more abundant in the granodiorite core of the pluton. 

The dikes generally are 2 to 10 cm wide in the tonalite 

of Bald Mountain and commonly as much as 17 cm 

near the center of the exposed pluton. Some aplite 

dikes in the granodiorite core are as much as 70 cm 

wide. The dikes mostly are flesh-colored and composed 

of feldspar, quartz, and a small amount of biotite.

The overall scarcity of pegmatite and aplite dikes 

in the Elkhorn pluton is consistent with the crystallization of a relatively dry magma. Moreover, almost 

no quartz veins occur in the core area, where aplites 

and pegmatites are most abundant. Only one small 

group of closely associated quartz veins was found in 

a 5-km2 area south and southeast of Anthony Lake. 

The small veins contain as much as about 2 volume 

percent chalcopyrite and are confined within a zone 

less than 20 m in length.

XENOLITHS OF COUNTRY ROCKS AND INCLUSIONS OF 

UNKNOWN ORIGIN IN ELKHORN PLUTON

The scarcity of metasedimentary xenoliths except 

near contacts with metasedimentary wallrocks is characteristic of many batholiths in the Western United 

States, ranging from the small Snoqualmie batholith 

(Erikson, 1969) to the larger Boulder batholith (Grout 

and Balk, 1934; Becraft and others, 1963) to the 

immense Sierra Nevada batholith (Bateman, 1983; 

Noyes and others, 1983).

The Elkhorn pluton is unique among the zoned intrusions in northeastern Oregon in the comparative 

abundance of xenoliths of country rocks in the interior 

of the pluton. Elsewhere in northeastern Oregon xenoliths of country rocks are relatively rare except near 

contacts such as in the Wallowa batholith (Taubeneck, 

1987) and Cornucopia stock (Taubeneck, 1964).

Field studies of the xenoliths in the Elkhorn pluton 

commenced in 1987 with the goal of determining their 

character, relative abundance, and distribution. The 

studies ultimately should provide insight concerning
88 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

emplacement mechanisms, extent of assimilation, and 

the origin of the comparatively numerous mafic inclusions of unknown origin that cannot be visibly correlated with known country rocks. Mafic inclusions of 

unknown origin in granitic rocks commonly have been 

interpreted as restite of deep-seated origin, carried 

upward from its source by the granitic magma. General field studies are well advanced in the Monumental 

salient and in an area of about 12 km2 south and east 

of Anthony Lake. The ensuing discussion of field observations refers to these areas, respectively, as the 

"salient" and the "lakes area."

Xenoliths of country rock in the two areas consist of 

(in order of increasing abundance) rare calc-silicate 

rocks, cherty argillite, ribbon chert, and mafic foliated 

rocks (mostly foliated metagabbro). Xenoliths of country rock constitute substantially less than 0.01 volume 

percent of the granodiorite in the lakes area and an 

estimated 0.02 volume percent of the tonalite in the 

salient. The xenoliths are more easily seen and sampled in the many accumulations of large talus blocks 

in cirque basins than in massive outcrops. Counts of 

metasedimentary xenoliths observed each day during 

the examination of talus are as follows: calc-silicate 

rocks, 0 to 1 in the lakes area and 0 to 2 in the salient; 

cherty argillite, 0 to 3 in the lakes area and 6 to 13 in 

the salient; ribbon chert, 3 to 8 in the lakes area and 

19 to 34 in the salient.

Xenoliths of mafic foliated rocks in the lakes area 

are three to four times as abundant as all other xenoliths of country rock as judged mostly by counts in 

four large accumulations of talus. Initial attempts to 

count the number of mafic foliated xenoliths were 

abandoned because of their great abundance, especially in much of the salient area where many thousands of very small foliated-rock slivers occur. Counts 

in the salient exceeded 300 such xenoliths per day 

before tabulations were terminated.

The greatest dimensions of xenoliths for each of 

the four types are calc-silicate rocks, as much as 22 

cm in the lakes area and 15 cm in the salient; cherty 

argillite, as much as 45 cm in the lakes area and 51 

cm in the salient; ribbon chert, as much as 70 cm in 

the lakes area and 89 cm in the salient; and mafic 

foliated rocks, as much as 29 cm in the lakes area 

and 31 cm in the salient. The smaller size of the mafic foliated xenoliths is attributed to fragmentation 

along the sharply defined planes of foliation in contrast to much less well defined foliation in chert and 

most cherty argillite.

Xenoliths of chert range in shape and size from 

rare blocks and slabs to very small circular relicts. 

Some xenoliths more than about 14 cm in smallest 

dimensions lack bedding and are derivatives either of

very coarse ribbon chert or conceivably of massive 

chert which constitutes less than about 0.03 volume 

percent of country rocks within 3 km of the batholith. 

A few xenoliths in the salient area (but not in the 

lakes area) are slender ribbons up to 15 cm long. The 

irregular to subrounded shape of many xenoliths indicates magmatic resorption. Some xenoliths less 

than 20 cm in greatest dimension are almost elliptical, whereas a few less than 3 cm are circular. The 

xenoliths are not rimmed by a mafic corona as in the 

Ballachulish granodiorite, Argyllshire, Scotland 

(Muir, 1953). The shapes of the smaller xenoliths in 

particular indicate assimilation.

The grain size of recrystallized quartz in the xenoliths is no larger than in bedded chert near contacts 

of the Elkhorn pluton or in much chert throughout 

the Belleview wedge. Therefore, either the grain size 

in high-temperature aureole rocks does not continue 

to increase at magmatic temperatures, or the temperatures of aureole rocks and crystallizing granitic 

magma were too close to produce any additional increase in grain size.

The light-colored xenoliths of chert are difficult to 

detect in the granitic rocks. Consequently, the search 

for inclusions is more successful from a stationary 

position than while moving. The inclusions of chert 

are best seen in the Monumental salient near the 

headwall in the cirque basin just north of Mount Ireland (fig. 2.2) and along the crests of the three ridges 

that enclose this basin.

Cherty argillite composes the large number of rocks 

in the Elkhorn Mountains that are gradational between ribbon chert and argillite (Pardee and Hewett, 

1914). Most xenoliths of cherty argillite are surrounded by an eye-catching halo of iron oxide that is visible 

for appreciable distances. Cherty argillite is common 

in the aureole but is disproportionately uncommon 

among xenoliths of country rocks. Assimilation of 

cherty argillite is much more prevalent than assimilation of ribbon chert, as is apparent along and near 

contacts with wallrocks. Accordingly, the relative scarcity of xenoliths of cherty argillite is consistent with 

the comparatively rapid rate of assimilation of pelitic 

rocks in contrast to that of ribbon chert.

Mafic foliated xenoliths, generally less than 18 cm 

long, are mostly metagabbro that is moderately fine 

grained and strongly foliated. Some mafic schistose 

xenoliths could be meta-argillite; no unequivocal megascopic distinction is possible. Many moderate to large 

xenoliths are transected by finer grained shear zones 

such as are conspicuous in metagabbro in the aureole. 

Xenoliths vary from uncommon, almost-black rocks 

with sparse visible plagioclase to common, lighter 

rocks with approximately equal amounts of horn-
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 89

blende and plagioclase. Many xenoliths are angular or 

somewhat rectangular, whereas some are subrounded. 

Substantially less than 1 percent of the xenoliths have 

reacted with the magma to produce a halo of recrystallized but nonfoliated rock surrounding a foliatedrock core. Otherwise, the xenoliths exhibit no or only 

rare indication of reaction with the magma.

A distinct decrease in the number of xenoliths occurs in the Monumental salient east and northeast of 

a boundary shown as line D-E in figure 2.2. Xenoliths increase again farther east, commencing slightly 

more than 1 km from the east contact of the salient. 

The distribution of xenoliths must be related to 

source regions and magmatic-flow directions. However, the xenoliths are distributed randomly and cannot 

be followed along zones of inclusions to a source in 

peripheral wallrocks, as in the Donegal granite in 

Northern Ireland (Pitcher and others, 1959).

The Monumental salient is mostly in contact with 

metagabbro for about 2.5 km along its east side. 

Elsewhere metagabbro constitutes less than 2 percent of wallrock, which is almost entirely ribbon 

chert and cherty argillite with small amounts of argillite. Xenoliths of ribbon chert and cherty argillite 

occur throughout the salient, although less commonly 

east of the line D-E in figure 2.2. The distribution of 

mafic foliated xenoliths, however, shows a closer relation to the line D-E and the distance to metagabbroic 

wallrock. Most of the mafic foliated xenoliths east of 

line D-E are of larger size and can be correlated with 

metagabbro along the east side of the salient. The 

thousands of small slivers of foliated metagabbro are 

mostly west of line D-E and presumably were derived 

largely from a western source. The scarcity of metagabbro in the aureole west of line D-E suggests a 

source for the slivers at a different level or levels 

from those now exposed by erosion. Whatever the 

source or sources, a thorough mixing of magma is required to explain the distribution of xenoliths.

Mafic inclusions of uncertain origin constitute more 

than 99 percent of all inclusions in the lakes area 

and an estimated 95 percent in the salient. In the 

lakes area, the mafic inclusions average 0.35 volume 

percent of the granodiorite; in the salient, they average 0.32 volume percent of tonalite in the interior of 

the salient to the west of line D-E (fig. 2.2). Nearly 

all inclusions are less than 0.6 m in greatest dimension; a few in the lakes area are slightly more than 1 

m across, and one is 2.1 by 1.6 m.

The mafic inclusions exhibit notable variations in 

color, grain size, texture, mineral proportions, and 

shape. Color ranges from grayish black to gray, grain 

size from fine to medium grained, and texture from 

equigranular to heterogranular. The heterogranular

inclusions contain large crystals of hornblende and 

(or) plagioclase. It is not known whether many of 

these large crystals are phenocrysts or porphyroblasts. Many mafic inclusions of unknown origin contain few (if any) large crystals of hornblende or few 

(if any) of plagioclase; some contain no large crystals 

of either mineral. Modal contents of hornblende and 

plagioclase vary significantly. Haloes of hornblende 

rim some inclusions but not more than 5 percent of 

the inclusions in any area. Most inclusions are subrounded, some have one or more angular corners, 

some are angular with no rounded surfaces, and a 

few appear almost spherical. Scalloped margins are 

rare. According to Didier (1987, p. 47), inclusions 

with angular shapes "prove the proximity" of their 

source, whereas inclusions with ellipsoidal shapes 

originated "much farther away."

The inclusions are sufficiently scattered that two or 

more of them generally cannot be closely examined 

and compared from one vantage point; swarms of inclusions are extremely rare. For comparative purposes, specimens of some inclusions can be carried 

throughout the day, but most inclusions are in massive rock that defies sampling even with a heavy 

sledge. In the ensuing discussion, one of only three 

known clusters containing more than 10 inclusions is 

described.

The swarm occurs at an elevation of 2,225 m, approximately 0.5 km N. 2° W. of the summit of Mount 

Ireland. The inclusions are in an area of about 1.4 

m2 on a large and conspicuous diagonal joint face 

that dips 25° N. The cluster is 2 m west of an aplite 

dike that strikes about N. 20° E; the aplite is extremely wide (about 20 cm) compared to relatively 

few others within 500 m.

The cluster contains four xenoliths of ribbon chert, 

one xenolith of cherty argillite, two metagabbro xenoliths and two probable metagabbro xenoliths (one 7 

cm long, spear-shaped; the other, about 1 by 2 cm, 

rectangular), and 23 mafic inclusions of uncertain origin from 2 to 30 cm in greatest dimension. Three of 

the 23 inclusions are angular to very angular, whereas six have one or more pointed corners but are 

mostly subrounded. The remaining 14 inclusions are 

subrounded. Of the 23 inclusions, 21 contain large 

crystals of hornblende, but some of these 21 have almost no large crystals of plagioclase. The two other 

inclusions contain no large crystals of either hornblende or plagioclase. The inclusions range considerably in color from almost light gray to dark gray, 

reflecting differences in the amount of hornblende 

and plagioclase. Many of the 21 inclusions with large 

crystals of hornblende differ from their neighbors, 

but some appear almost identical.
90 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Exceptional numbers of closely associated xenoliths 

of metasedimentary wallrocks together with mafic inclusions of unknown origin occur near the contact of 

the Monumental salient near elevation 2,089 m about 

5.2 km S. 77° E. of the summit of Mount Ireland (near 

summit 6,853 on the Mount Ireland 7/^-minute quadrangle). Inclusions of all the types commonly total 

about 35 percent of the rock by volume. Roughly onehalf of all inclusions are metasedimentary xenoliths, 

and the remainder are mafic inclusions of uncertain 

origin. The derivation of the inclusions of uncertain 

origin cannot be determined by megascopic observations, but observations do trigger a few thoughts. If 

the mafic inclusions of uncertain origin are restite, 

why are they 50 times more abundant near a wallrock 

contact than throughout the interior of the salient? 

What is the significance of the positive correlation between excessive numbers of inclusions of unknown origin and excessive numbers of xenoliths of country 

rocks?

Exceptional numbers of closely associated xenoliths 

of country rocks together with mafic inclusions of unknown origin occur in a small satellitic intrusion about 

5.1 km S. 65° E. of the summit of Mount Ireland and 

roughly 0.7 km south of the south contact of the salient. The inclusions are in an offshoot of tonalite (at 

elevations from 1,920 to 1,972 m) that intrudes metasedimentary rocks of the Elkhorn Ridge Argillite on 

either side of the boundary between T. 8 S. and T. 9 S, 

R. 36 E. At least one-half of the outcrops within 60 m 

of the west contact of the offshoot contain 20 to 35 

percent inclusions. An estimated 12 to 15 percent of 

the inclusions are xenoliths of country rocks, including 

metagabbro that does not crop out within 3 km of the 

offshoot of tonalite. Some mafic inclusions of unknown 

origin are elongated and angular, but most are subrounded. Some contain no large crystals of either hornblende or plagioclase. Why is the concentration of 

supposed restite in a small satellitic offshoot of tonalite between 50 and 100 times greater than throughout 

the main part of the Monumental salient? Why are so 

many xenoliths of country rocks concentrated with the 

supposed restite? Have excessive numbers of xenoliths 

of country rocks been mixed in one small offshoot with 

an excessive amount of supposed restite? Or have all 

inclusions in the offshoot been derived from mesozonal 

levels near the south side of the Monumental salient?

In the Elkhorn Mountains there are too many disquieting field relations, including the overall diversity among the inclusions of uncertain origin, to allow 

acceptance of the restite model without question.

Four other explanations for the origin of the mafic 

inclusions are (1) disaggregated and dispersed synplutonic mafic dikes, (2) genetically related plutonic

rocks, (3) xenoliths of reconstituted mafic country 

rocks, and (4) globules of either basic or intermediate 

magmas that intruded granitic magma in the lower 

crust. Assessment of all explanations is not feasible 

merely from field observations and examination of 

thin sections of three norite and six metagabbro xenoliths. However, sufficient overall information is 

available to question all but the fourth explanation 

for the derivation of the mafic inclusions.

No synplutonic dikes occur at the present level of 

erosion in any granitic intrusion in the Elkhorn Mountains. Accordingly, the possibility that mafic to intermediate inclusions of unknown origin in the Elkhorn 

pluton are dismembered fragments of synplutonic 

dikes is remote. The absence of synplutonic dikes is 

another feature that distinguishes the Bald Mountain 

batholith from most circum-Pacific batholiths.

Three inclusions of genetically related noritic rocks 

were found in an area about 12 km2 south to east of 

Anthony Lake. The norite xenoliths possess the unmistakable features of the norite of Badger Butte, 

such as local recrystallization of plagioclase into 

small grains at pressure points and the postconsolidation extension of hornblende along grain boundaries of plagioclase and quartz. Thin sections reveal 

no further reduction in grain size beyond that which 

occurred at plagioclase pressure points during emplacement of the norite. The xenoliths have maintained their integrity, despite traveling at least 

several kilometers from their point of origin. Accordingly, evidence in the Elkhorn pluton does not support the possibility that the mafic inclusions of 

uncertain origin were derived from genetically related plutonic rocks.

The possibility that the mafic inclusions of unknown origin represent reconstituted mafic country 

rocks is considered mostly from field observations, 

supplemented by thin sections of six xenoliths of 

metagabbro. The largest xenolith of metagabbro seen 

in the lakes area is 29 cm in maximum dimension; 

the xenoliths generally are less than 18 cm long. In 

comparison, many mafic inclusions of unknown origin 

in the lakes area exceed 29 cm in greatest dimension; 

a few exceed even 100 cm. Clearly, the larger inclusions of unknown origin were not derived from the 

smaller xenoliths of metagabbro. Field observations 

of xenoliths of metagabbro reveal little or no mineralogical or textural reaction with the magma, other 

than the uncommon hornblende rims that border a 

small number of the xenoliths. Metagabbro xenoliths 

are foliated in contrast to the nonfoliated mafic inclusions of unknown origin. The grain size of plagioclase 

and hornblende in the mafic inclusions of unknown 

origin is quite variable; most inclusions contain large
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 91

crystals of hornblende; many also contain large crystals of plagioclase. In contrast, nearly all plagioclase 

and hornblende in the six xenoliths of metagabbro 

are restricted to a grain size from 0.05 to 0.5 mm. No 

large crystals of plagioclase or hornblende that qualify as either phenocrysts or porphyroblasts occur in 

the six xenoliths. Accordingly, reconstituted xenoliths 

of metagabbro are an improbable source for the mafic 

inclusions of uncertain origin.

In summary, the diversity of textures and compositions as observed during field observations of the mafic inclusions of uncertain origin suggests a multiple 

source that possibly includes the intrusion in granitic 

magma at depth of basic to intermediate magmas.

QUARTZ DIORITE OF LIMBER CREEK

SETTING AND GENERAL DESCRIPTION

The quartz diorite of Limber Creek, represented by 

four small intrusions near the neck of the Monumental salient, is herein named for rocks that crop out in 

the headwaters of Limber Creek (fig. 2.4). Rocks of the 

four intrusions are medium grained and have a color 

index of about 18 to 46, in sharp contrast to the nearby tonalite of Bald Mountain which has an index of 10 

to 18. The quartz diorite in the upper drainage of Limber Creek is part of the largest intrusion, exposed as 

an elongated east-west mass of 0.4 km2 and designated for location purposes as the western quartz diorite. 

This elongated body is concealed on the west by glacial deposits and does not reappear west of the map 

area (fig. 2.4). The western quartz diorite is sharply 

truncated on the east by the tonalite of Bald Mountain and is partly separated from the tonalite on the 

south by a dismembered screen of country rocks.

Planar structure in the western quartz diorite, absent in numerous exposures, is fair to fairly good at 

some locations; strikes are approximately east-west 

and dips are mostly vertical to 70° S. Favorable sites 

for viewing the planar structure are the numerous 

small cliffs from 1 to 5 m high that are 150 to 400 m 

west of Limber Creek. Only three elongated mafic inclusions of unknown origin were noted in the area of 

small cliffs during a search of 90 minutes. Elongation 

ratios are from 2.5:1 to 4.0:1. Good planar structure in 

the quartz diorite is confined to a small area of perhaps 1,000 m2 in the extreme west, about 75 m south 

of the contact with country rocks and near the glacial 

deposits (fig. 2.4). Here, mafic inclusions of uncertain 

origin have elongation ratios of as much as 8:1.

Xenoliths of country rocks are common in the western quartz diorite. West of Limber Creek they are concentrated within 50 m of both the north and the south

contacts of the quartz diorite; the xenoliths occur 

mostly as slices of metasedimentary rocks as much as 

3 m wide and 7 m long. In contrast, xenoliths are 

much smaller and less abundant in the interior of the 

western quartz diorite, where maximum dimensions 

of xenoliths rarely exceed 0.7 m. The relatively flat 

crest of a north-northwest-trending ridge roughly 0.5 

km west of Limber Creek provides a convenient cross 

section for observing the distribution and size variations of the xenoliths within the quartz diorite, the 

features of the intruded metasedimentary rocks of the 

screen, and, just to the south, the peripheral exposures of the tonalite of Bald Mountain. Metasedimentary xenoliths in the tonalite are comparatively few in 

comparison with those within the western quartz diorite; they also are relatively small with one conspicuous exception. The largest metasedimentary xenolith 

found (18 m long and as much as 2 m wide) in the 

tonalite of the Monumental salient is just south of the 

screen and about 0.3 km west of Limber Creek.

A second of the four intrusions of quartz diorite in 

the neck area is a very small exposure of about 0.003 

km2 , approximately 0.5 km east of Limber Creek and 

80 m north of the western quartz diorite (fig. 2.4). 

The rocks are well exposed, except on the east, and 

probably connect at depth to the western quartz diorite. This small intrusion was not sampled and will 

receive no further consideration.

A third body of the quartz diorite of Limber Creek, 

on the east side of the neck, is exposed in a 0.13-km2 

area (fig. 2.4) about 5 km S. 75° E. from the headwaters of Limber Creek. This mass (designated for location purposes as the eastern quartz diorite) is intruded 

on the northeast by a small body of tonalite containing 

red-brown biotite and minor iron oxide, separated on 

the northwest from the batholith by a narrow screen 

of ribbon chert, and concealed on the south by glacial 

deposits. The quartz diorite is well exposed in the east 

half of the intrusion but poorly exposed throughout 

much of the west half. Many xenoliths of metasedimentary rocks, as much as 6.5 m in maximum dimension, occur in the east-central part of the eastern 

quartz diorite, together with xenoliths of metagabbro 

that do not exceed 1 m across. Distance from source is 

a plausible explanation for the disparity in size between xenoliths of the two rock types; the eastern 

quartz diorite is in contact with metasedimentary 

wallrocks (fig. 2.4), whereas the nearest surface exposures of metagabbro are 0.55 km to the southeast.

The eastern quartz diorite is cut by leucocratic granitic dikes (as much as 8 m wide) that contain no 

hornblende and only minor biotite. Dikes on the west 

intrude the screen of ribbon chert that separates the 

quartz diorite from the tonalite of the batholith. Two
92 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 2.9. Modes of the quartz diorite of Limber Creek

[Values in volume percent. Each analysis represents at least 2,400 point counts for each of two thin sections; 0.0, <0.05]

Specimen Quartz 

number

Plagioclase Biotite Hornblende Augite Opaque Nonopaque 

minerals accessory minerals

Western quartz diorite

234 -

236 --

441 

476 

Average

6

14

21

14

14

23

15

.2

.2

.2

.4

.2

.6 

.6

45.

48.

46.

46.

51.

47. 

47.

0

7

0

2

6

3 

5

2

11

11

9

.5

.4

.6

.6

.9

.1 

.9

45

19

27

22

16 

26

u

.7

.1

.6 

.6

£t \J

*±

3

1

1

2 

2

1 0

£t

1

1

0

1

1

2

2

1

1

1

1 

1

Eastern quartz diorite

447 11.9 45.1 13.4 29.3 0.2 0.0 0.1

448 15.8 43.4 7.4 32.6 .6 .1 .1

470 11.3 54.2 11.5 22.7 .0 .0 .3

Average 13.0 47.5 10.8 28.2 .3 .0 .2

Brecciated quartz diorite

445 14 . 9 45.6 6>1 32.5 Q.6 0.1 0.2

472 11.1 46.3 4.6 37.5 .2 .1 .2

Average 13.0 45.9 5.4 35.0 .4 .1 .2

excellent examples of partial assimilation of ribbon 

chert occur in the granitic dikes near the east contact 

of the screen. As much as 3 m of ribbon chert in a 

thicker sequence has been partly assimilated in a 5- 

m-wide dike at an elevation near 2,160 m, about 20 

to 30 m east of a prominent gulch. The location is 

roughly N. 40° E. of a conspicuous prospect dump at 

an elevation of about 2,100 m. Dikes to the northeast, at an elevation of about 2,260 m, exhibit many 

small-scale examples of the partial assimilation of 

pieces of ribbon chert as much as 0.5 m in greatest 

dimension. Thus even leucocratic granitic dikes can 

assimilate some refractory chert; assimilation commenced along the argillaceous partings. More extensive assimilation of refractory ribbon chert occurs 

within the western quartz diorite, as well as in those 

parts of the screen south of the western quartz diorite that were intruded by the quartz diorite in a 

crude lit-par-lit fashion.

A fourth occurrence of the quartz diorite of Limber 

Creek is along a restricted part of the east contact of 

the neck of the Monumental salient where the quartz 

diorite in contact with the tonalite of Bald Mountain 

is mostly a giant breccia. The quartz diorite (including breccia) is not more than about 30 m wide and is 

merely a remnant of a larger mass that was intruded, brecciated, and dispersed during emplacement of 

the tonalite. The breccia is as much as 20 m wide 

and contains blocks of quartz diorite as large as 6.4 

by 4.6 m. For identification purposes this small body 

is designated as the brecciated quartz diorite. As is 

true of both the western and eastern quartz diorites, 

the brecciated quartz diorite contains more xenoliths 

of metasedimentary rocks than the tonalite of Bald 

Mountain. Modal analyses of six rock specimens of 

the western quartz diorite, three rock specimens of 

the eastern quartz diorite, and two rock specimens of 

the brecciated quartz diorite are given in table 2.9.

MICROSCOPIC DESCRIPTION

Rocks of the western quartz diorite, eastern quartz 

diorite, and brecciated quartz diorite (table 2.9) crystallized from magmas sufficiently alike to warrant a 

single petrographic description. In addition to the 

overall petrographic description, a discussion of several late-stage features of the brecciated quartz diorite will be presented.
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 93

Augite occurs in hornblende as relict crystals as 

much as 2.1 mm in greatest dimension. Replacement 

of augite generally is advanced, and commonly only 

bleached areas in hornblende with small grains of 

quartz signify the former presence of augite. Nearly all 

larger relicts of augite are composed of many isolated 

parts that extinguish simultaneously. For example, 

eight cores of augite that range in greatest dimension 

from 1.1 to 1.9 mm are composed of from 5 to 12 

detached (in plane of thin section) parts that extinguish in unison. No augite relicts contain inclusions of 

plagioclase, but one large relict is in reaction relation 

to a core of hypersthene about 0.1 by 0.3 mm. Orthopyroxene and subsequently clinopyroxene apparently 

were the earliest major minerals to crystallize.

Primary hornblende is brown and generally between 0.5 and 5.0 mm in greatest dimension, whereas recrystallized hornblende is light green and 

mostly about 0.1 mm across. Many smaller crystals 

of brown hornblende, mostly less than 1.0 mm in 

greatest dimension, are subhedral, but larger crystals are commonly irregular and even poikilitic. From 

some massive crystals, narrow fingers of hornblende 

extend outward along grain boundaries and, rarely, 

along fractures in plagioclase. One poikilitic hornblende roughly 2 by 4 mm encloses nine small laths 

of plagioclase; a few poikilitic crystals surround tabular plagioclase as much as 0.7 mm in length. However, during the overlapping crystallization histories of 

the two minerals, several hornblende crystals as 

much as 0.8 mm by 0.4 mm were enclosed in plagioclase. Late-stage deformation and recrystallization of 

hornblende is much more prevalent than in any other quartz diorite or tonalite in northeastern Oregon. 

Hornblende in the quartz diorite of Limber Creek 

commonly has recrystallized into a mosaic of nearly 

equidimensional grains. For example, after recrystallization, one original crystal of hornblende about 1.1 

by 0.8 mm consists of 97 grains, each about 0.1 mm 

across.

Tabular plagioclase is mostly between 0.3 and 4.0 

mm in greatest dimension. The bulk of the plagioclase 

is andesine, but core compositions are as calcic as 

about An53. Some plagioclase shows oscillatory zoning, generally from three to seven broad and rather 

indistinct bands; a few crystals exhibit as many as 12 

oscillations. Polysynthetic twinning fairly commonly 

is bent as much as 7° and infrequently from 8° to 16°. 

Locally, at least a few crystals in every rock specimen 

except 476 (table 2.9) have recrystallized along contacts with adjacent plagioclase to form a mosaic of 

xenoblastic grains mostly 0.1 to 0.04 mm across; recrystallization of plagioclase is not as extensive as recrystallization of hornblende. Slender needles of rutile

(mostly less than 0.1 mm long) that are aligned crystallographically in plagioclase are a distinctive feature 

of the quartz diorite. The interior of a few of the larger 

plagioclase crystals is mottled in some thin sections. 

The mottling is characterized either by imperfect 

square or rectangular areas that are interconnected 

and extinguish simultaneously or by discontinuous 

patches. Mild clouding of the plagioclase is attributed 

to thermal metamorphism (MacGregor, 1931). Clouding in some thin sections consists of very fine particles 

that are concentrated toward the central part of plagioclase, whereas in other thin sections larger particles 

are aligned along twinning planes.

Quartz is interstitial. Some crystals in every thin 

section show wavy extinction, but they extinguish 

within a relatively small rotation of 3°. Quartz alongside bent plagioclase or recrystallized plagioclase 

commonly shows almost no wavy extinction. Apparently quartz crystallized after the mechanical stress 

that was responsible for the recrystallization of hornblende and plagioclase. Two examples were noted of 

narrow quartz veins that cross plagioclase but cannot 

be traced into areas of interstitial quartz. Rutile needles mostly less than 0.15 mm long are characteristic 

of the quartz. Orientation of the needles generally is 

evident where more than 12 to 15 needles occur in a 

relatively small area.

Biotite crystallized late, as is demonstrated by a 

common reaction relation to hornblende, by general 

occurrence along grain boundaries, by some poikilitic 

crystals, and by thin stringers of biotite that extend 

along fractures that cross plagioclase. The striking 

red-brown to red color of the biotite as seen microscopically is also megascopically visible in thin sections. 

Biotite fairly commonly is bent from 7° to 12°, infrequently from 13° to 17°, and rarely as much as 26°.

Accessory minerals are allanite, zircon (mostly euhedral), apatite, and iron oxide. The allanite replaces 

hornblende, as is commonly true in northeastern Oregon in granitic rocks such as quartz diorite, tonalite, 

and granodiorite. Much of the apatite occurs as relatively large crystals more than 0.08 mm in cross section. The rocks are notable for the small amount of 

iron oxide.

The brecciated quartz diorite is distinguished petrographically from the western and eastern quartz 

diorites by resorption of hornblende and resorption of 

plagioclase that is not recrystallized. Resorption is 

minor in rock specimen 445 but strong in specimen 

472 (table 2.9).

Resorption of hornblende is denoted by irregular and 

embayed contacts with quartz, rounded appendages of 

hornblende that extend into quartz, and isolated pieces 

of hornblende in quartz that extinguish simultaneously
94 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 2.10. Modes of the leucogranodiorite of Trail Creek

[Values in volume percent. Each analysis represents at least 2,400 point counts for each of three thin sections; 0.0, <0.05]

Specimen Quartz Potassium Plagioclase Biotite 

number feldspar

Opaque Nonopaque 

minerals accessory 

minerals

93 39.5 i4.Q 42.1 4.3 Q.I Q.Q

97 34.6 13 . 8 46 . 3 4 . 8 .5 . Q

104 36 .6 13.3 48.5 1.5 .1 .0

Average 36.9 13.7 45.6 3.6 .2 .0

with neighboring crystals of hornblende. Resorbed plagioclase generally occurs as an irregular core separated 

by a sharp Becke line from a clear rim that constitutes 

as much as 60 percent by volume of some crystals. 

Cores commonly show higher birefringence than rims 

and always are darker from clouding than the rims, 

which never display clouding. Fractures in the cores as 

much as 0.01 mm wide do not extend into either the 

rims or the quartz adjacent to the plagioclase. Many of 

the larger fractures contain chlorite.

Perhaps heat from the younger tonalite of Bald 

Mountain generated an interstitial melt phase in the 

brecciated quartz diorite to produce the unique features of the plagioclase and the resorption of hornblende. This possibility cannot be assessed without a 

systematic sampling of several large blocks of breccia 

(from margin to center), as well as the sampling of 

the brecciated quartz diorite farthest from the tonalite of Bald Mountain.

LEUCOGRANODIORITE OF TRAIL CREEK

SETTING AND GENERAL DESCRIPTION

The leucogranodiorite of Trail Creek (fig. 2.2) is 

herein named for exposures in the headwaters area 

of Trail Creek, along the crest of a north-south-trending ridge extending from about 1.7 to about 3.4 km 

north of Badger Butte. It underlies an area about 1 

km2 and is the largest leucogranodiorite body in the 

batholith. The leucogranodiorite intrudes the tonalite 

of Bald Mountain on the northwest, but elsewhere its 

contacts with the tonalite and the granodiorite of Anthony Lake are concealed.

The leucogranodiorite is medium grained and contains widely scattered crystals of dark biotite from 3 to 

5 mm across. Fresh specimens are light gray, but nearly all outcrops are weathered to various shades of yellow. The rock disintegrates readily into a sandy grus; 

most good exposures are restricted to steep slopes near

the south and north extremities of the Trail Creek 

body. Modal analyses of three rock specimens of the 

leucogranodiorite are given in table 2.10.

Notable quantities of aplite are associated with the 

leucogranodiorite. The dimensions of the largest aplite 

dikes 15 to 18 m wide greatly exceed the dimensions 

of aplite dikes in other parts of the batholith. Some 

large aplite dikes extend or occur outside the leucogranodiorite; one is as far as 0.6 km from the nearest 

surface exposure of leucogranodiorite.

MICROSCOPIC DESCRIPTION

Plagioclase (An16_42) occurs as subhedral to anhedral crystals throughout most of the leucogranodiorite of Trail Creek. Oscillatory zoning in the core is 

surrounded by a wide mantle of progressively more 

sodic oligoclase. Potassium feldspar is orthoclase and 

mostly interstitial; some anhedral to subhedral crystals as much as 5 mm show exsolution lamellae and 

Carlsbad twinning. Orthoclase that is not interstitial 

may contain euhedral crystals of both quartz and 

plagioclase. Myrmekite occurs in limited amounts as 

narrow fringes about 0.02 mm wide between plagioclase and orthoclase. Strong undulatory extinction 

and local recrystallization characterize the interstitial quartz. Biotite is subhedral to anhedral except 

for small crystals in orthoclase. Most biotite occurs as 

flakes, less than 0.7 mm across, that commonly are 

along grain boundaries. The leucogranodiorite shows 

considerable alteration of biotite to chlorite and minor amounts of epidote. Accessory minerals are iron 

oxide, apatite, and zircon.

GRANITE OF CLEAR CREEK

SETTING AND GENERAL DESCRIPTION

The granite of Clear Creek (fig. 2.2) is herein 

named for rocks exposed in an area about 0.67 km
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 95

TABLE 2.11. Modes of the granite of Clear Creek

[Values in volume percent. Each analysis represents at least 2,400 point counts for each of three thin sections; 0.0, <0.05]

Specimen 

number

R147 

R148 - 

Quartz

31.7 

33.5

Potassium 

feldspar

21.9 

20.9

Plagioclase

39.7 

39.8

Biotite

6.6 

5.3

Hornblende

0.1 

.5

Opaque 

minerals

0.0 

.0

Nonopaque 

accessory minerals

0.0 

.0

Average 32.6 21.4 39.7 6.0 .3 .0 .0

near the junction of Clear Creek and the Grande 

Ronde River. The unit intrudes the tonalite of Bald 

Mountain but is overlain on the west by Cenozoic 

volcanic rocks. A dismembered screen of country 

rocks extends along the southwest contact between 

the granite and the tonalite.

The granite is medium grained and light gray to 

slightly pinkish where fresh; it contains scattered 

dark biotite. Many outcrops are covered with lichens; 

some of the best exposures form bluffs on the north 

side of the Grande Ronde River opposite Clear Creek. 

Foliation is shown by aligned biotite and varies from 

faint to strong with mostly east-northeast trends. The 

modes of two rock specimens are given in table 2.11.

MICROSCOPIC DESCRIPTION

Most plagioclase is anhedral, but small subhedral 

crystals, not more than 0.8 mm long, occur enclosed 

within quartz and potassium feldspar. The plagioclase 

is antiperthitic oligoclase with cores, in larger crystals, 

of sodic andesine. In most crystals zoning is normal, 

but in some large crystals normal zoning encloses several faint oscillatory zones. Twinning lamellae in a few 

crystals are bent as much as 5°; fractured crystals also 

occur with twinning planes in adjoining pieces at angles as much as 14°. The potassium feldspar is anhedral and mostly a fine microperthitic orthoclase; some 

crystals show faint microcline gridiron twinning. Myrmekite is common as a narrow fringe between plagioclase and potassium feldspar, but the intergrowths are 

not prominent. Quartz shows strong undulatory extinction and local recrystallization. Biotite occurs as 

subhedral crystals as much as 1 mm across in potassium feldspar, as irregular flakes along grain boundaries, and as very small stringers along a few 

fractures. Some biotite along grain boundaries is bent. 

Hornblende is present as strongly corroded relicts that 

commonly are surrounded by quartz.

Accessory minerals are apatite, zircon, monazite, 

allanite, and iron oxide. The iron oxide occurs as 

small cubic crystals that commonly are in aggregates

of two or more intergrown cubes. Alteration products 

are restricted to uncommon chloritization of biotite 

and a slight kaolinitization of potassium feldspar.

QUARTZ DIORTTE OF WOLF CREEK

SETTING AND GENERAL DESCRIPTION

The quartz diorite of Wolf Creek is herein named for 

rocks that crop out on the north side of the canyon of 

Wolf Creek (fig. 2.2B). The rock is exposed in an eastwest-trending zone about 3.25 km2 (fig. 2.3) that extends westward into the headwaters of Beaver Creek, 

where the unit is divided by a wedge-shaped body 

composed of the presumably younger granodiorite of 

Beaver Meadow and a dismembered screen of country 

rock about 0.1 km wide. Not all contacts are exposed. 

The quartz diorite is intruded along its south contact 

by the granite of Anthony Butte and elsewhere is 

overlain by the Columbia River Basalt Group.

The quartz diorite is a gray, medium-grained rock 

with a color index of about 27. An alignment of tabular plagioclase is typical. Conspicuous crystals of poikilitic biotite, commonly between 10 and 15 mm in 

greatest dimension, contribute to a foliation that generally is good to excellent. Mafic inclusions help to 

define the foliation but commonly constitute less 

than 1 percent of the rock. Foliation trends from 

about east-west to west-northwest with northward, 

mostly steep dips. Modal analyses of six quartz diorite rock specimens are given in table 2.12.

MICROSCOPIC DESCRIPTION

The plagioclase consists of subhedral laths of slightly 

zoned calcic andesine. The margins of many crystals 

show slight normal zoning, whereas the interiors of 

some crystals have as many as eight oscillatory zones. 

Plagioclase in rock specimens from near the contact 

with the granite of Anthony Butte is moderately clouded by dustlike particles, probably as a result of thermal 

metamorphism (MacGregor, 1931). Some plagioclase
96 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 2.12. Modes of the quartz diorite of Wolf Creek

[Values in volume percent. Each analysis represents at least 2,400 point counts for each of two thin sections; 0.0, <0.05; , not present]

Specimen Quartz 

number

Average 15.4

Potassium Plagioclase 

feldpsar

Biotite Hypersthene Augite Hornblende

2.1 54.5 15.2 1.6 4.3 6.5

Opaque Nonopaque 

minerals accessory minerals

PI "7 

POT 

R122 -

R125 -

15.3

16.3

5.1

19.4 

18.6

17.8

3.5

4.9

0.1

1.6

59.3

56.1

61.3

4 o o

51.1

50.4

14

14

4

16 

22

17

9

6

6

6

6

1.2

1.7

6.1

0.0 

.2

.3

5.3

5.9

1.3

3.7 

6.1

3.7

0

20

1

8

.3 0

.5 1

.7 

.1

.4

2 0

0

1 

1

1

0

1

1

1 

1

1

.3 .1

encloses small crystals of hypersthene and augite (generally less than 0.1 mm across), but crystals of hypersthene as large as 0.9 mm also occur in plagioclase. 

Twinning lamellae in plagioclase commonly are bent 

8° and rarely as much as 12°. Bent twinning lamellae 

in plagioclase adjacent to quartz in which strain shadows are the maximum deformation effects suggest latestage deformation of plagioclase before crystallization 

of quartz. Recrystallization of plagioclase to a fine mosaic occurs along two microfractures as well as at grain 

boundaries where plagioclase crystals impinge on one 

another. Twinning lamellae on opposite sides of microfractures commonly are offset or trend in slightly different directions. Microfractures that do not extend 

into adjacent quartz suggest protoclastic deformation 

of plagioclase before crystallization of quartz.

Hypersthene is subhedral and generally not more 

than 2.0 mm in length. Discrete crystals of augite are 

of similar size, but much of the augite either surrounds or partly rims hypersthene. Small crystals of 

hypersthene and augite enclosed in plagioclase are 

evidence that pyroxene began to appear in the magma at an early stage of crystallization. Hornblende 

and biotite are in reaction relation to both pyroxenes 

but also occur as discrete crystals. Hornblende and 

biotite crystallized very late and conceivably are poikiloblastic. For example, narrow fringes of biotite 

and hornblende extend along boundaries between 

plagioclase crystals, between plagioclase and quartz, 

and between crystals of quartz. Biotite also extends 

along several microfractures that offset twinning lamellae of plagioclase.

Quartz and potassium feldspar are interstitial. 

Some quartz shows strong undulatory extinction, and 

local recrystallization has occurred in two rock samples. The potassium feldspar is orthoclase as indicated by (flat-stage) estimates of 2V of less than 60° for 

nearly centered acute-bisectrix figures. Iron oxide, 

apatite, and stubby crystals of zircon are the main 

accessory minerals.

DISTINGUISHING FEATURES

The quartz diorite is unique among the granitic intrusions of northeastern Oregon in containing hypersthene and augite. Poikiloblastic(?) texture of biotite 

and hornblende also is distinctive, as well as the deformational features of plagioclase.

BOUNDARY QUARTZ DIORITE UNIT

SETTING AND GENERAL DESCRIPTION

A quartz diorite that occurs in a timbered area about 

0.8 km2 along the boundary between the Anthony Butte 

quadrangle and the Tucker Flat quadrangle (fig. 2.2A) 

is herein informally designated the boundary quartz 

diorite unit (fig. 2.3). Most of the unit's better outcrops 

are located in the eastern third of the body. Farther 

west, much of the quartz diorite has disintegrated into 

a greenish-gray grus, which is well exposed along several abandoned logging roads. The quartz diorite is 

locally intruded on the north by the leucogranite of 

Dutch Creek, but the two rock types are mostly separated by a dismembered screen of metasedimentary rocks. 

A fragmented screen on the south separates the quartz 

diorite from the granodiorite of Indiana Mine Road. The 

relationship between these two granitic rock units is 

unknown, but presumably the quartz diorite is older 

than the granodiorite. The Columbia River Basalt 

Group overlies the quartz diorite on the west and east.

The quartz diorite is a gray, medium-grained rock 

ranging from massive to slightly foliated. In nearly 

all outcrops, a few crystals of plagioclase, hornblende, 

and biotite are as much as 10 mm across, but most 

crystals are 1.5 to 4.0 mm across. Modal analyses of 

two rock specimens are given in table 2.13.

MICROSCOPIC DESCRIPTION

The plagioclase is subhedral andesine that shows 

as many as 16 oscillatory zones surrounded by a nar-
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 97

TABLE 2.13. Modes of granitic rocks of the northeastern (peripheral) part of the Bald Mountain batholith

[Values in volume percent. Each analysis of the boundary quartz diorite unit represents at least 2,400 point counts for each of two thin sections; all other analyses represent at least

2,400 point counts for each of three thin sections; 0.0, <0.05; , not present]

Specimen Quartz Potassium Plagioclase Biotite Muscovite Orthopyroxene Augite Hornblende Cummingtonite Opaque Nonopaque 

number feldspar minerals accessory

minerals

Average 34.0

Boundary quartz diorite unit

R115 -

R116 

Average

18.2

13.0

15.6

1.8

.0

.9

52.

56.

54.

,0

,0

.0

13,

11.

12.

. 1 

.8

.4 

0.2

.3

 .2

14.

18.

16.

,6

.6

,6 

0.1

.2

.2

0.1

.1

.1

Granodiorite of Indiana Mine Road

R187

R188

R189

R190

 33.

 34.

 36.

 31.

,3

,5

.9

,1

5.

11.

6.

11.

,5

,7

.7

,9

52.

44.

49.

50.

.9

.9

.5

.2

8.

8.

6.

6.

3

9

6 

1 

___

 

0.2

 .7 

0.0

.0

.0

.0

0.0

.0

.1

.0

49.4 7.5 .2 .0 .0

Average 32.5

Granodiorite of Beaver Meadow

R118 -

R119 -

R127 -

31.

32,

33.

.1

.8

.8

20,

11,

19.

.6

.2

.1

42.4

46.5

40.1

5,

7

6,

.7 

.8

.2 

0.4

.2

0.2

.6

.3

___

0.6

.3

 0.0

_ __ 1

.0

17.0 43.0 6.6 .2 .4 .3

Granite of Anthony Butte

R120

R121

R123

R128

R129

 32.7

 32.9

 32.8

 31.4

 31.2

28.2

23.4

23.6

28.4

25.4

36.4

38.0

37.2

36.0

37.5

2.6 

5.7 

4.3 0.2 

4 _1 

5.9

0.0

 .0

1.8 .0

 

 

0.1

.0

.1

.1

.0

Average 32.3 25.8 37.0 4.5 .4 .0 .1

Average 37.9

Leucogranite of Dutch Creek

R22 38.4 29.1 30.5 1.6 0.4 

R24 38.2 26.8 33.6 1.1 .1 

DOR tl O O Q C. Q 1 Q O 1 1 

0.0 0.0

.2 .0

i n

28.2 32.0 1.6 .2 .1 .0

row mantle that is progressively more sodic toward 

the periphery. The cores of some plagioclase crystals 

contain small inclusions (<0.1 mm) of augite and 

hornblende. The margins of a few adjacent plagioclase crystals are recrystallized at pressure points to 

a fine-grained mosaic of nearly equant crystals less 

than 0.1 mm across. Twinning lamellae in plagioclase 

are bent as much as 12°. Bent twinning lamellae are
98 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

common at the ends of plagioclase crystals where 

they impinge on one another. Quartz is interstitial 

and commonly shows strong undulatory extinction; at 

least a few sutured contacts indicative of recrystallization under stress are present in every thin section. 

Hornblende occurs as subhedral and anhedral crystals; some that are poikilitic are of late origin. Relicts 

of augite as large as 1.6 mm occur within some hornblende. The former presence of augite in other crystals of hornblende is indicated by bleached cores that 

enclose minute crystals of quartz (Taubeneck, 1964). 

Biotite occurs in reaction relation to hornblende and 

commonly as irregular crystals >2 mm across. Late 

crystallization of biotite is indicated commonly by biotite along grain boundaries and rarely by biotite as 

small flakes along fractures in plagioclase. Interstitial potassium feldspar is orthoclase, as verified on 

the flat stage by estimated 2V of 50°-55°for nearly 

centered acute-bisectrix figures. Accessory minerals 

are iron oxide, apatite, zircon, and sphene.

DISTINGUISHING FEATURES

The boundary quartz diorite unit is characterized 

by small crystals of augite within plagioclase and by 

millimeter-size cores of augite in hornblende. Both 

occurrences of augite confirm the early crystallization 

of clinopyroxene. The deformation features of the plagioclase distinguish this quartz diorite from other 

quartz diorites in the Bald Mountain batholith, with 

the exception of the quartz diorite of Wolf Creek.

GRANODIORITE OF INDIANA MINE ROAD

SETTING AND GENERAL DESCRIPTION

The granodiorite of Indiana Mine Road is herein 

named for the Indiana Mine Road (fig. 2.2B) that extends across the western part of the intrusion. The 

unit consists of an elongated mass about 3 km2 (fig. 

2.3) that extends westward for about 3 km from 

Dutch Creek (fig. 2.25). The granodiorite intrudes 

country rocks along its southwestern margin and 

contains several sizable inclusions of wallrocks in its 

western part. By far the largest inclusion extends 

nearly east-west for about 1 km and dips steeply 

north. The granodiorite's west contact with the granite of Anthony Butte and its north contact with the 

boundary quartz diorite unit are concealed. A dismembered screen on the north partly separates the 

granodiorite from the boundary quartz diorite unit. 

The normal mafic-to-felsic sequence occurring in the 

composite batholiths and stocks of northeastern Oregon suggests that the granodiorite is younger than 

the boundary quartz diorite unit but older than the 

granite of Anthony Butte. The eastern and southeastern parts of the granodiorite are overlain by the Columbia River Basalt Group.

The granodiorite is a light-colored, medium-grained 

rock with scattered relatively small crystals of biotite 

that generally define a strong east-west-striking, 

steeply north-dipping foliation. The rock tends to 

crumble and is poorly exposed; much of it has disintegrated into a distinctive yellowish grus, which is 

best seen along several abandoned logging roads. 

Modal analyses of four rock specimens of the granodiorite are given in table 2.13.

MICROSCOPIC DESCRIPTION

Plagioclase (An25_34) is subhedral to anhedral and 

only slightly zoned. Resorption of plagioclase is common where it contacts potassium feldspar; some plagioclase in quartz also has irregular borders that 

appear corroded. Much of the plagioclase is antiperthitic. Deformation of the crystals is pronounced 

and twinning lamellae commonly are bent as much 

as 8°. The quartz is strongly strained and sutured 

boundaries indicate considerable recrystallization. 

The potassium feldspar is orthoclase microperthite; 

at least a few Carlsbad twins occur in almost every 

thin section. Biotite, generally less than 1.0 mm in 

largest dimension, is mostly anhedral and restricted 

to grain boundaries. Cleavage traces of some biotite 

are bent. Late crystallization of biotite is verified by 

small crystals along microfractures in plagioclase. 

Small amounts of resorbed hornblende in two rock 

samples (R189 and R190 in table 2.13) indicate that 

hornblende originally was more abundant. Accessory 

minerals are apatite, zircon, allanite, monazite, and 

iron oxide, but they generally are not abundant 

enough (<0.05 volume percent) to be recorded in modal analyses.

DISTINGUISHING FEATURES

The granodiorite of Indiana Mine Road is distinguished mineralogically by a very high content of 

quartz, almost no hornblende, and very little iron 

oxide (0.01 volume percent). The high content of 

quartz and paucity of iron oxide clearly suggest assimilation of graphite-bearing cherty argillite and 

ribbon chert. However, the biotite in each of the four 

rock specimens (table 2.13) exhibits no trace of the 

red-brown (typical) to reddish (atypical) color of biotite in other contaminated granitic rocks in this part
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 99

of the Bald Mountain batholith. The deformation fabric of this rock unit is almost unique among the 

granodiorite bodies of northeastern Oregon in the 

comparatively abundant deformed plagioclase, bent 

biotite, and recrystallized quartz.

GRANODIORITE OF BEAVER MEADOW

SETTING AND GENERAL DESCRIPTION

The name "granodiorite of Beaver Meadow" is herein applied to a body of granodiorite that is exposed in 

a 0.35-km2 area in the Anthony Burn east of Beaver 

Meadow (fig. 2.25). The granodiorite is a wedgeshaped mass situated between the quartz diorite of 

Wolf Creek to the north (beyond country rock screen) 

and south and overlain to the west by alluvium in 

Beaver Meadow and farther west by the Columbia 

River Basalt Group (fig. 2.3). Contacts with the quartz 

diorite are concealed, but the granodiorite presumably 

is younger as judged from the typical mafic to felsic 

intrusive sequence in composite batholiths.

The granodiorite is medium grained, light gray 

where fresh, and mottled with dark biotite. Scattered 

megacrysts of potassium feldspar are common and 

are as much as 2.5 cm across. These megacrysts generally constitute less than 2 percent of the rock by 

volume. Foliation is well developed and shown by 

aligned biotite and commonly also by plagioclase and 

megacrysts of potassium feldspar. The foliation is 

somewhat variable but strikes approximately N. 70° 

W. and has a steep northeastward dip. Modes of 

three rock specimens are given in table 2.13.

MICROSCOPIC DESCRIPTION

The plagioclase ranges from sodic andesine to calcic 

oligoclase and commonly shows slight zoning. The mineral occurs as anhedral crystals mostly 0.5 to 4.0 mm 

long and as smaller subhedral to euhedral crystals 

within potassium feldspar. A small number of resorbed 

crystals are present within quartz. Anhedral plagioclase is antiperthitic, and twin planes are commonly 

bent 5° to 8°. Some microfractures offset twinning lamellae that differ in trend on opposite sides of the 

microfracture. Almost every thin section contains at 

least one example of the recrystallization of plagioclase 

to a fine-grained (0.05-0.20 mm) mosaic along contacts 

where two crystals directly impinge on each other.

Potassium feldspar is present as small interstitial 

crystals and also as subhedral, commonly Carlsbadtwinned crystals more than 1 mm across. Many of 

the larger crystals are microperthitic. The potassium

feldspar is exclusively orthoclase in rock specimen 

R119, but some microcline also is present in specimen R127; microcline is more abundant than orthoclase in rock specimen R118 (table 2.13).

Myrmekite is mostly inconspicuous and nearly always occurs as a very narrow fringe between plagioclase and potassium feldspar. Only a few of the typical 

cauliflower-like protuberances are present in any thin 

section, and they are absent in some sections. Quartz 

occurs generally as interstitial crystals up to 4 mm 

across, but some euhedral crystals, mostly less than 

0.5 mm across, are present in potassium feldspar. The 

quartz shows weak to strong undulatory extinction as 

well as some recrystallization.

Orthopyroxene has been mostly replaced by fibrous 

cummingtonite and typically is present only as small 

cores within aggregates of cummingtonite. Pseudomorphs after orthopyroxene suggest that it originally 

occurred as stout subhedral prisms mostly 0.3 to 0.6 

mm long; the largest prism was 0.9 mm long. Relatively early crystallization of orthopyroxene is indicated 

by three unreplaced crystals that occur inside plagioclase. A slight to moderate alteration to a reddish iron 

oxide is associated with nearly all orthopyroxene.

Cummingtonite is characterized by polysynthetic 

twinning and extinction angles between 14° to 18°. 

Small quantities of green hornblende occur in reaction relation to cummingtonite and also as independent crystals along grain boundaries. Hornblende 

associated with cummingtonite commonly occurs as a 

very narrow fringe around the cummingtonite. This 

replacement is always peripheral and never extensive. The interstitial hornblende consists of thin 

stringers and fingers, commonly less than 0.03 mm 

wide, that locally are poikilitic.

Biotite occurs as (1) poikilitic crystals as much as 

2.5 mm across, (2) small flakes along grain boundaries, (3) small subhedral crystals as much as 0.5 

mm long within quartz and potassium feldspar, (4) 

reaction rims on both cummingtonite and hornblende, and (5) very small flakes, mostly less than 

0.2 mm long, within microfractures in plagioclase 

and potassium feldspar. All occurrences except the 

subhedral crystals in quartz and potassium feldspar 

indicate late to very late crystallization. Cleavage 

traces of biotite commonly are bent between 5° to 10° 

and, rarely, more than 30°. Deformation of biotite 

generally is most intense at the corners of plagioclase 

and between closely adjacent crystals of plagioclase 

that are at high angles to each other. Biotite is redbrown in rock specimens R119 and R127 but not in 

specimen R118 (table 2.13).

Alteration products are relatively uncommon. Some 

biotite shows chloritization, and potassium feldspar
100 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

commonly is dusted by kaolin. Flakes of muscovite as 

much as 0.1 mm long are visible on some potassium 

feldspar crystals and also along a few grain boundaries. Trace amounts of calcite occur interstitially and 

also in association with hornblende. Accessory minerals are apatite, zircon, allanite, and very small 

amounts of iron oxide. The virtual absence of iron oxide 

indicates that the granodiorite of Beaver Meadow crystallized under more reducing conditions than prevailed 

in most granitic plutons in northeastern Oregon.

DISTINGUISHING FEATURES

Foliation and megacrysts of potassium feldspar are 

the two distinguishing megascopic features of the 

granodiorite of Beaver Meadow. Nearly all other granodiorite bodies in northeastern Oregon lack foliation, 

and no other granodiorite contains megacrysts of potassium feldspar.

The presence of orthopyroxene is an even more notable distinction. Furthermore, the reaction relations 

between orthopyroxene, cummingtonite, hornblende, 

and biotite are exceptional. A literature search failed 

to disclose any reference to a similar reaction sequence in granitic rocks.

Another distinction of the granodiorite is the scarcity of iron oxide. Iron oxide occurs in such small 

amounts that it was not recorded during the modal 

analysis of three samples (26,727 points, across nine 

thin sections) of the granodiorite (R118, R119, and 

R127 in table 2.13). In contrast, iron oxide constitutes, respectively, 0.23, 0.23, 0.27, and 0.30 volume 

percent of the four small intrusions of granodiorite in 

the batholith south of lat 45° N., and 0.27, 0.30, 0.31, 

and 0.36 volume percent of the four small intrusions 

of granodiorite in the Wallowa batholith extending 

from about 45 to about 70 km to the northeast. Each 

of the other eight intrusions of granodiorite contains 

more than 60 times as much iron oxide as the Beaver 

Meadow unit.

Red-brown biotite in rock specimens R119 and 

R127 (table 2.13) is petrographic evidence of reduction, here attributed to the assimilation of graphitebearing metasedimentary rocks. However, a chemical 

analysis (Ken-ichiro Aoki, 1973, Tohoku University, 

Japan) of rock specimen R118 (table 2.13) which does 

not contain red-brown biotite yielded values for 

Fe2O3 and FeO of 0.26 and 2.13 weight percent, respectively, indicating that R118 also is reduced. The 

absence of red-brown biotite in R118 is difficult to interpret without additional whole-rock chemical analyses, each accompanied by a mineral analysis of 

biotite, including biotite from R118.

GRANITE OF ANTHONY BUTTE

SETTING AND GENERAL DESCRIPTION

The granite on either side of the major ridge in the 

northern part of the batholith is herein called the 

granite of Anthony Butte. This rock unit is exposed 

over an irregular area approximately 14 km2 (fig. 

2.3); in size it greatly exceeds any other granite body 

in northeastern Oregon. The granite notably contains 

two very large inclusions of quartz diorite and country rock; xenoliths probably constitute at least 20 

percent of the volume of the intrusion, as indicated 

by surface exposures. The granite of Anthony Butte 

intrudes the tonalite of Bald Mountain on the south 

and west, the quartz diorite of Wolf Creek on the 

north, and country rocks on the east. The granite's 

contact with the granodiorite of Indiana Mine Road 

is concealed, but the granite is probably younger 

than the granodiorite according to the typical maficto-felsic sequence. A fragmented screen on the south 

and southwest partly separates the granite from the 

tonalite of Bald Mountain.

Fresh specimens of the granite of Anthony Butte 

have about 5 percent dark biotite scattered throughout a white matrix of feldspar and quartz. Weathered 

specimens are discolored and dominated by various 

shades of yellow. Exposures are generally poor, partly 

because the rock tends to disintegrate. Some relatively fresh outcrops occur locally as rounded knobs that 

stand 2 to 5 m above the surrounding ground level.

The granite is a predominantly medium-grained 

rock; most crystals of plagioclase and quartz are 2 to 

6 mm across. Twinned megacrysts of potassium feldspar occur in many parts of the intrusion and constitute as much as 2.3 percent of the rock by volume. 

The volume percentage of megacrysts was determined by placing a transparent grid across large exposures and counting the grid squares and parts of 

squares that cover each megacryst. Megacrysts commonly are 2.0 to 2.5 cm long, but some are as much 

as 3.5 cm. The megacrysts apparently crystallized 

from the magma and thus are phenocrysts rather 

than porphyroblasts.

Foliation is well developed in many parts of the intrusion and is shown both by biotite and by aligned 

megacrysts of potassium feldspar. Some foliation is 

almost gneissic and is locally accentuated by elongated inclusions and schlieren of nearly assimilated 

metasedimentary rocks. Intricate swirled foliation occurs locally. Foliation generally closely parallels external contacts as well as internal contacts with 

large inclusions and dips more than 65° N. to NE. 

Modal analyses of five rock specimens are given in 

table 2.13.
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 101

MICROSCOPIC DESCRIPTION

Plagioclase is present as anhedral crystals about 1 

to 4 mm long and as much smaller euhedral to subhedral crystals in potassium feldspar. The plagioclase 

is oligoclase that shows slight normal zoning. Some 

highly resorbed crystals of plagioclase occur in both 

quartz and potassium feldspar. Antiperthite is common in rock specimens R121 and R123 (table 2.13). 

Twinning lamellae in many larger crystals are bent 5° 

and rarely 12°. Microfractures are fairly common in 

most thin sections, and some plagioclase shows minor 

crushing. Trends of twinning planes in adjacent pieces 

of broken plagioclase differ by as much as 13°. Local 

recrystallization of plagioclase to a mosaic of small 

grains, commonly less than 0.1 mm across, occurs 

along some microfractures and also at some pressure 

points where crystals impinge on each other.

Potassium feldspar is microperthitic, and many of 

the larger crystals show Carlsbad twinning. Excellent microcline gridiron twinning is common in rock 

specimens R120 and R123, (table 2.13), but much of 

the potassium feldspar in both rocks is orthoclase, as 

revealed by the relatively small 2V of many acutebisectrix figures. Microcline occurs locally in rock 

specimen R121, but potassium feldspar in rock specimens R128 and R129 is exclusively orthoclase. Alteration of potassium feldspar to kaolin is ubiquitous in 

all rock samples; flakes of muscovite as much as 0.5 

mm across occur within the orthoclase of rock specimen R128.

Myrmekite in small amounts is present in all the 

rock samples, mostly as a very narrow fringe between plagioclase and potassium feldspar. Conspicuous myrmekite is restricted to rock specimen R123 

(table 2.13), in which myrmekite occurs as cauliflower-like protuberances that project as far as 0.4 mm 

into adjacent potassium feldspar.

Quartz occurs as large crystals from 3 to 8 mm 

across, as smaller interstitial crystals, and as much 

smaller euhedral to subhedral crystals in potassium 

feldspar. The large areas of quartz generally are devoid of other minerals and are believed to have crystallized relatively early. Some are egg- or lens-shaped. 

Large crystals generally show weak to strong undulatory extinction, but sutured boundaries indicate that 

some recrystallization has occurred. However, unstrained crystals as much as 4 to 6 mm across occur in 

the same thin section with strongly strained quartz 

and recrystallized quartz.

In rock specimen R123 (table 2.13), orthopyroxene 

occurs with iron oxide and cummingtonite as small 

relicts and as remnant cores of subhedral crystals of 

orthopyroxene; in the partially replaced crystals, the

pyroxene is surrounded by cummingtonite. Orthopyroxene is rarely found in true granites; the small 

(about 0.4 km2) leucogranite of Mt. Rubidoux (Larsen, 1948) in southern California is the best known 

orthopyroxene-bearing granite in North America. The 

orthopyroxene crystals in specimen R123 are 0.15 to 

0.9 mm long and are characterized by pale-pink to 

greenish pleochroism and by alteration to dark-red, 

apparently nearly opaque iron oxide. However, in 

thin section most of the iron oxide appears translucent when viewed with the upper condenser inserted. 

Patches of iron oxide as large as 0.5 mm across compose as much as 60 volume percent of a few pseudomorphs. Orthopyroxene in felsic granitic rocks 

typically is iron-rich and commonly is ferrohypersthene. The iron oxide associated with the orthopyroxene in the granite of Anthony Butte strengthens 

the probability of an iron-rich composition for this 

pyroxene. Moreover, the small amount of iron oxide 

in the rock (table 2.13) and consequent implication of 

low oxygen fugacity during crystallization also 

strengthens the probability of an iron-rich orthopyroxene. Most of the orthopyroxene typically is replaced by fibrous cummingtonite that generally 

occurs as compact aggregates and clusters of small 

crystals. In contrast to the cummingtonite that is associated with iron oxide in obvious pseudomorphs after orthopyroxene, some cummingtonite in interstitial 

areas without iron oxide appears to have formed as a 

primary mineral of magmatic crystallization. Of possible significance in this respect is the report by 

Sams and Saleeby (1988, p. 876) of cummingtonite in 

the southernmost Sierra Nevada in a hornblende 

quartz diorite that contains neither orthopyroxene 

nor clinopyroxene. Irrespective of origin, cummingtonite in the granite of Anthony Butte commonly is 

fringed by biotite in reaction relation.

Most biotite is anhedral and generally occurs along 

grain boundaries as small crystals 0.1 to 0.8 mm long. 

Cleavage in some crystals is bent. Small flakes of biotite also occur along a few microfractures in plagioclase. Some biotite in rocks that show comparatively 

little foliation occurs as clusters of as many as a dozen 

flakes as much as 1.5 mm across. Biotite also is 

present as subhedral crystals enclosed within either 

potassium feldspar or quartz. Biotite crystallized before some potassium feldspar and quartz, but much of 

the biotite along grain boundaries is very late, as is 

the biotite along microfractures in plagioclase.

Biotite in rock specimens R123 and R128 is redbrown, but this indicator color, typical of biotite in 

reduced rocks, is absent in specimens R120, R121, and 

R129 (table 2.13). Chemical analyses (Ken-ichiro Aoki, 

1973, Tohoku University, Japan) for rock specimens
102 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

R121, R123, and R128 (table 2.13) yielded, respectively, the following values in weight percent: Fe2O3 =0.74, 

FeO=1.80; Fe2O3 =0.52, FeO=1.93; and Fe2O3 =0.17, 

FeO=1.91. Ratios of the iron values for the three rocks 

correlate with the absence of red-brown biotite in rock 

specimen R121, a moderate red-brown color in the biotite of specimen R123, and a strong red-brown color 

in the biotite of specimen R128.

Accessory minerals are garnet, apatite, allanite, 

zircon, monazite, and iron oxide. The garnet occurs 

only in rock specimen R128 (table 2.13) in which it 

constitutes 0.1 volume percent of the rock. Garnet is 

present as subhedral to euhedral crystals 0.05 to 0.80 

mm in diameter, is faint pink, and locally is fringed 

by biotite. The garnet either is a primary magmatic 

mineral or occurs as xenocrysts from assimilated 

metasedimentary rocks. More than 99 percent of the 

garnets in 21 metapelitic rocks from the aureole are 

0.03 to 0.50 mm across; garnets showing well-developed crystal habit commonly occur as crystals as 

much as about 0.1 mm in diameter. Crystals larger 

than about 0.2 mm invariably are irregular, commonly contain a core of opaque inclusions, and generally 

are perforated. In contrast, the numerous garnets 

larger than 0.2 mm in the granite contain no inclusions, are not perforated, and have subhedral to euhedral borders. Therefore, features of the garnet in 

the granite suggest crystallization from a melt.

DISTINGUISHING FEATURES

The granite body is distinct in the common strong 

foliation, abundance of large igneous and metasedimentary inclusions, and relative size which exceeds 

that of other granite bodies in northeastern Oregon 

by at least 20 times. Mineralogical distinctions are 

common megacrysts of potassium feldspar, low content of iron oxide (0.02 volume percent), red-brown 

biotite in two specimens, and local occurrence of garnet, orthopyroxene, and cummingtonite.

LEUCOGRANITE OF DUTCH CREEK

SETTING AND GENERAL DESCRIPTION

A poorly exposed leucogranite in a heavily wooded 

area about 0.25 km2 in the headwaters of Dutch 

Creek (fig. 2.2B) is herein called the leucogranite of 

Dutch Creek (fig. 2.3). It intrudes the boundary quartz 

diorite unit on the south and east, is overlain on the 

north and west by rocks of the Columbia River Basalt 

Group, and is in contact with a dismembered screen of 

country rocks on the southwest. Throughout much of 

the area, the leucogranite has disintegrated into a

grus; rock specimens can be obtained from only a few 

outcrops. However, it is unexpectedly common as float, 

which typically shows faint to moderate foliation.

Fresh leucogranite is pale gray with sparsely scattered dark biotite. Most rock specimens, however, are 

discolored by weathering and are dominated by shades 

of yellow and yellow orange. The grain size might best 

be characterized as fine to medium because most crystals are 0.5 to 2.0 mm. Modes of three rock specimens 

of the leucogranite are given in table 2.13.

MICROSCOPIC DESCRIPTION

Plagioclase in the xenomorphic-granular leucogranite is a weakly zoned oligoclase that is partly altered 

to small flakes of white mica. The potassium feldspar 

is mostly orthoclase, as verified in many crystals by 

2V of less than 60°. Indistinct microcline gridiron 

twinning occurs in a few crystals. The orthoclase 

shows faint exsolution and minor development of 

small (0.05-0.1 mm) secondary albite granules at the 

margins of larger potassium feldspar crystals. Myrmekite in minor amounts occurs both as bulbous protuberances and as narrow fringes along contacts 

between plagioclase and potassium feldspar.

Quartz is characterized by strong undulatory extinction and considerable recrystallization that ranges from strong in rock specimen R22 to moderate in 

specimen R24 to minor in specimen R25 (table 2.13). 

The grain size of crystals in rock specimen R22 commonly has been reduced by recrystallization from 

grains more than 1 mm across to about 0.5 mm and 

locally to less than 0.1 mm.

Biotite is present mostly along grain boundaries as 

irregular flakes less than 0.5 mm long. A few flakes are 

found along microfractures in plagioclase and potassium feldspar. Muscovite occurs mostly associated with 

potassium feldspar as flakes as much as 0.8 mm across, 

but also in small quantities intergrown with biotite. 

Accessory minerals are iron oxide, apatite, and zircon.

No field or petrographic evidence was noted to suggest that the character of the leucogranite of Dutch 

Creek has been modified by assimilation of country 

rocks. Rock specimens R22, R24, and R25 (table 2.13) 

contain, respectively, 0.04, 0.18, and 0.07 volume percent iron oxide. These iron oxide contents would be 

considered low for tonalites and granodiorites, but 

not for leucogranite, such as the leucogranite of 

Dutch Creek.

DISTINGUISHING FEATURES

The leucogranite is distinct in having a grain size 

that is intermediate between aplite and normal medi-
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 103

TABLE 2.14. Modes ofsatellitic granitic rocks northeast of the Bald Mountain batholith

[Values in volume percent. Each analysis represents at least 2,400 point counts for each of two thin sections of the tonalite of North Fork or for 

each of three thin sections of the granite of Isham Spring; 0.0, trace amount; , not present]

Specimen Quartz 

number

Potassium 

feldspar

Plagioclase Biotite Hornblende Opaque Nonopaque 

minerals accessory minerals

Tonalite of North Fork

R133 29.0 0.3 56.5 11.8 2.3 0.0 0.1

R134 30.8 1.5 55.4 11.5 .7 .0 .1

Average 29.9 .9 55.9 11.7 1.5 .0 .1

Granite of Isham Spring

R136 36.3 25.7 33.6 4.4

R137 36.0 26.5 33.3 4.2

Average 36.1 26.1 33.5 4.3

0.0 

.0

.0

0.0 

.0

.0

um-grained granitic rocks. Moreover, the leucogranite 

has the highest average content of both quartz and 

potassium feldspar of any unit in the Bald Mountain 

batholith. The modal composition of the leucogranite 

closely resembles that of aplite, except for the leucogranite's higher content of biotite. Recrystallization of 

quartz is notable, especially in rock specimen R22 

(table 2.13).

SATELLITIC BODIES EAST OF GRANITE OF 

ANTHONY BUTTE

TONALITE OF NORTH FORK

SETTING AND GENERAL DESCRIPTION

A somewhat irregular body of tonalite that occupies a 0.4-km area surrounded by country rocks just 

north of the North Fork of Anthony Creek is called 

the tonalite of North Fork (fig. 2.3). The country 

rocks are mostly schistose metasedimentary rocks 

but include amphibolites derived from metagabbro. 

The tonalite is medium grained with a color index of 

about 12. Biotite commonly is concentrated within 

moderately distinct planes that accentuate a well-defined foliation that strikes from due west to N. 60° W. 

and dips 35°-60° N. Modal analyses of specimens of 

tonalite from two widely separated outcrops are given in table 2.14.

MICROSCOPIC DESCRIPTION

The plagioclase is andesine and occurs mostly as 

subhedral crystals that are only weakly zoned. The

crystals are antiperthitic, and some are markedly embayed by quartz. Twinning lamellae are commonly 

bent 6° and rarely as much as 14°. Local recrystallization has occurred along the contact of some plagioclase where it is in contact with another plagioclase 

crystal. Quartz is characterized by moderate to strong 

undulatory extinction, and about half of the quartz 

shows one or more sutured contacts. Features of 

quartz are best viewed with low-power magnification.

Biotite ranges from fairly well-formed plates to 

ragged grains and aggregates. It generally is independent of hornblende, and only a few examples of a 

reaction relation were observed. The biotite in both 

specimens (table 2.14) is pleochroic from grayish yellow (X-axis direction) to either reddish brown or, 

more generally, red (Z axis), in contrast to the brownish color of biotite (Z) in most other granitic rocks of 

northeastern Oregon. The pronounced reddish coloration of the biotite in thin sections is readily visible 

megascopically. Both megascopically and microscopically, colors of biotite in the tonalite of North Fork 

(as in all reduced rocks in the Elkhorn Mountains) 

are best seen in crystals more than 0.5 mm in maximum dimension. In the tonalite of North Fork the 

occurrence of much biotite along grain boundaries, 

including those between adjacent crystals of quartz, 

indicates late crystallization. The cleavage of some 

biotite is bent. A conspicuous feature of the biotite is 

numerous enclosed crystals of apatite. For example, 

one flake about 0.6 mm long contains 11 small crystals of apatite.

Hornblende occurs mostly as euhedral to subhedral 

crystals less than 2.5 mm long, but much hornblende
104 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

that is enclosed within quartz is partly resorbed. 

Early crystallization of some hornblende is implied 

by crystals as large as 0.3 mm enclosed within plagioclase. Interstitial potassium feldspar is orthoclase as 

judged from estimated 2V (flat stage) of about 55° for 

nearly centered acute-bisectrix figures. Accessory 

minerals are ubiquitous apatite and zircon as well as 

uncommon allanite and iron oxide. Local alteration 

products are white mica (on plagioclase) and chlorite 

(after biotite).

DISTINGUISHING FEATURES

The tonalite of North Fork is distinctive in modal 

composition, texture, and mineralogic features and 

has no known counterpart in northeastern Oregon. 

Nearly all other granitic rocks in northeastern Oregon 

with 30 percent modal quartz contain less biotite, 

more hornblende, and considerably more orthoclase 

than this tonalite. Trondhjemites in northeastern Oregon with 30 percent quartz contain less biotite and 

hornblende and commonly slightly more orthoclase 

than this tonalite. The core of the Cornucopia tonalite 

unit in the Wallowa Mountains (Taubeneck, 1967, fig. 

1) most nearly approximates the tonalite of North 

Fork, but the Cornucopia core rocks contain only 

about 60 percent as much biotite and essentially no 

hornblende.

The tonalite of North Fork has a distinctive fabric 

characterized by deformed plagioclase, strained and 

recrystallized quartz, and bent biotite. Other distinctive features are the reddish biotite and the numerous enclosed crystals of apatite. The virtual absence 

of iron oxide (0.01 volume percent) in the tonalite 

also is noteworthy. The reddish biotite and dearth of 

iron oxide in the tonalite of North Fork are evidence 

of the assimilation of graphite-bearing metasedimentary rocks.

GRANITE OF ISHAM SPRING

SETTING AND GENERAL DESCRIPTION

A small (about 0.04 km2) mass of granite that crops 

out just south of the North Fork of Anthony Creek and 

about 1.1 km north of Isham Spring is herein called 

the granite of Isham Spring (fig. 2.3). The granite is a 

tabular body that was emplaced concordantly between 

schistose metasedimentary rocks that strike roughly 

east-west and dip 35°-50° north. The granite is a medium-grained light-gray rock with scattered, small 

crystals of biotite that define a distinct foliation. The 

granite is well exposed and contains no megacrysts of

potassium feldspar. Modal analyses of two rock specimens are given in table 2.14.

MICROSCOPIC DESCRIPTION

The plagioclase is weakly zoned oligoclase that 

shows some corrosion by both quartz and potassium 

feldspar. Plastic deformation of plagioclase is not as 

common as fracturing, but twinning lamellae in one 

crystal are bent 14°. Microfractures are common, and 

a few crystals are composed of annealed pieces of plagioclase in which the twinning planes trend in different directions by about 10°. The potassium feldspar is 

orthoclase microperthite (2V<60°). The commonly 

Carlsbad-twinned crystals are as much as 8 mm 

across and anhedral; some enclose subhedral to euhedral quartz. Conspicuous myrmekite occurs predominantly as cauliflower-like protuberances that extend 

as far as 0.3 mm inside the orthoclase. Quartz is interstitial and exhibits very strong extinction and considerable recrystallization. Nearly all areas of interstitial 

quartz reveal some recrystallization, especially as 

viewed with low-power magnification. Red-brown biotite (both rock specimens, table 2.14) generally occurs 

along grain boundaries as small and irregular flakes 

that commonly show deformation effects. Cleavage 

planes in biotite are bent as much as 21°. Biotite locally extends along fractures in plagioclase. Accessory 

minerals are zircon, monazite, allanite, apatite, and a 

very small amount of iron oxide. Much of the zircon 

and monazite occurs in clusters of two or more crystals. The largest cluster consists of 27 crystals of monazite. Alteration products include small amounts of 

chlorite associated with biotite and minor kaolinitic 

alteration of orthoclase.

DISTINGUISHING FEATURES

Recrystallized quartz is the most distinctive feature of the major minerals of the granite. The only 

other granitic rock body in the Elkhorn Mountains 

with as much recrystallized quartz is the leucogranite of Dutch Creek. The granite of Isham Spring also 

is characterized by conspicuous myrmekite as bulbous and wartlike masses. The scarcity of iron oxide, 

only 0.005 volume percent, and the presence of redbrown biotite are other distinctive features of this 

rock. The clusters of zircon and monazite are unique 

among the granitic rocks of northeastern Oregon. Intrusions of cordierite trondhjemite in the Cornucopia 

stock (Taubeneck, 1964) and Wallowa batholith are 

characterized by concentrations of zircon but not of 

monazite.
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 105

GRANITIC ROCKS OF THE 

GUARD STATION INLIER

The northern part of the Bald Mountain batholith is 

concealed by Cenozoic volcanic rocks, and only two 

small exposures of tonalite occur near the bottoms of 

canyons incised in rocks of the Columbia River Basalt 

Group about 3 to 4 km north and northeast of the 

leucogranite of Dutch Creek (fig. 2.3). More widespread and significant exposures of granitic rocks are 

present about 18 km west on either side of the Grande 

Ronde River (fig. 2.2B) in an area about 6.5 km2 (fig. 

2.5) that is partly in the northwest corner of the Limber Jim Creek quadrangle but mostly in the southwest corner of the Little Beaver Creek quadrangle 

(fig. 2.2A). For purposes of reference the area of preTertiary rocks shown in figure 2.5 is called the Guard 

Station inlier after a U.S. Forest Service building, the 

Grande Ronde Guard Station, that is located well 

within the area. Most and possibly all granitic rocks 

in the inlier are satellitic to the batholith rather than 

an integral part of the batholith. Biotite granodiorite 

and biotite tonalite in the southeastern part of the 

inlier could be peripheral components of the batholith.

Country rocks in the Guard Station inlier are metasedimentary rocks and metagabbro that have their 

counterparts in the aureole of the batholith south of 

lat 45° N. The metasedimentary rocks are mostly ribbon chert, cherty argillite, and argillite that are correlative with the Elkhorn Ridge Argillite and are a 

northward extension of the Baker terrane. Strata in 

the east half of the inlier strike approximately eastwest and dip steeply southward, whereas most metasedimentary rocks to the west strike about N. 70° E. 

and dip >70° SE. or (less commonly) NW. Whether 

the contact between metagabbro on the south and 

metasedimentary rocks on the north (fig. 2.5) is tectonic or intrusive is debatable.

OVERVIEW OF GRANITIC ROCKS

For purposes of discussion the granitic rocks of the 

Guard Station inlier have been divided geographically 

and lithologically into four units (fig. 2.5): (1) a northern biotite granodiorite, (2) a central quartz diorite, 

(3) a southeastern biotite granodiorite and biotite tonalite, and (4) a southern quartz gabbro and quartz 

diorite. Many very small granitic intrusions, including 

at least two consisting of granite, are not shown in 

figure 2.5; several larger intrusions could not be sampled because of ubiquitous grus. The Guard Station 

inlier is a unique area in northeastern Oregon in the 

large number and great diversity of its small granitic 

intrusions. Moreover, petrographic evidence indicates

that the granitic magma of nearly every intrusion assimilated graphite-bearing metasedimentary rocks. 

The rock types are considered briefly by geographic 

distribution commencing in the north.

The northern biotite granodiorite (fig. 2.5) crops out 

for about 250 m along each of two small tributaries of 

the Grande Ronde River. The granodiorite intrudes 

metasedimentary rocks on the southwest and is overlain by Cenozoic volcanic rocks elsewhere. Modal analyses of three rock specimens are given in table 2.15.

Discrete crystals of muscovite as much as 1 mm 

long are the most distinctive feature of the northern 

biotite granodiorite. Some muscovite is intergrown 

with biotite, but most occurs as irregular crystals on 

either plagioclase or orthoclase. Muscovite is rare in 

granitic rocks in northeastern Oregon and is not 

known to occur in granodiorite bodies except within 

the Guard Station inlier. Another notable feature in 

each rock specimen (table 2.15) of the northern biotite granodiorite is the distinctive reddish-brown biotite. Other significant features of the northern biotite 

granodiorite are scarcity of iron oxide (only 0.016 

modal percent), traces of garnet (in rock specimen 

R592, table 2.15), and a relatively high content of 

quartz. Zircon commonly is zoned.

Many small and poorly exposed granitic intrusions 

occur in the central part of the Guard Station inlier near 

the contact between metasedimentary rocks and metagabbro. Most of the intrusions are in metasedimentary 

rocks, are elongated in an approximate east-west direction, are medium grained, and contain hornblende. The 

largest of the elongated bodies is an unmapped mass of 

grus extending from about 0.5 to 1.5 km east of the 

Guard Station. No specimens could be obtained because 

of poor exposures and deep weathering. The best exposed intrusion into the metasedimentary rocks is the 

central quartz diorite, which crops out in an area about 

0.04 km2 approximately 200 m south of the Grand 

Ronde Guard Station. Modal analyses of two rock specimens are given in table 2.15.

The central quartz diorite is characterized by crystals mostly less than 1 mm across, by cloudy plagioclase, and by evidence of early crystallization of 

abundant augite. Approximately one-half of the hornblende crystals contain bleached cores enclosing 

minute grains of quartz (Taubeneck, 1964), and some 

of the original augite is preserved as cores in hornblende. The largest augite is a core about 0.7 mm 

across surrounded by a rim of hornblende between 

0.1 and 0.2 mm wide. The abundance of augite as the 

early mafic phase in the fine-grained quartz diorite 

supports the general conclusion that granitic rocks in 

the Elkhorn Mountains crystallized from relatively 

hot and water-undersaturated magmas.
106 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 2.15. Modes ofplutonic rocks of the Guard Station inlier

[Values in volume percent. Each analysis represents at least 2,400 point counts for each of three thin sections; 0.0, trace amount; , not present]

Specimen Quartz Potassium 

number feldspar

R590 36.2 10.1 

R591 36.9 7.3 

R592 33.9 16.8

Average 35.7 11.4

R56 8.3 2.8 

Average 8.2 2.9

R54 33.6 27.2

R594 39.1 16.3 

R595 38.0 12.7

Average 38.5 14.5

R596 39.6 1.5 

R597 35.7 1.9

Average 36.7 1.7

R46 11.0 

Average 11 . 4 

R49 9.7 

Average 11.8 

D/1 t; Tit; o/lc

Plagioclase Biotite Muscovite Augite Hornblende Opaque 

minerals a

Northern biotite granodiorite

47.0 6.2 0.5 0.0 

48.4 7.3 .1 .0 

44.5 3.6 1.2 .0

46.6 5.7 .6 .0

Central quartz diorite

65.5 14.5 0.6 8.2 0.1

64.7 12.3 1.0 10.6 .3 

65.1 13.4 .8 9.4 .2

Leucogranite from central part of inlier

34.9 3.8 0.5 

Southeastern biotite granodiorite

40.3 4.2 0.1 0.0 

44.0 4.9 .4 .0

42.1 4.6 .3 .0

Southeastern biotite tonalite

52.9 6.0 0.0 

53.4 9.0 .0

53.1 7.5 .0

Southern quartz gabbro

25.4 0.6 62.3 0.2

28.6 1.0 58.3 .2

Southern quartz diorite

52.7 5.8 31.4 0.1

49.2 5.6 30.2 .9 

50.9 5.7 30.8 .5

Biotite tonalite dike from southern part of inlier

58.6 7.1 0.0 0.0

Granite from southern part of inlier

Nonopaque 

iccessory minerals

0.0 

.0 

.0

.0

0.0

.0 

.0

0.0

0.0 

.0

.0

0.0 

.0

.0

0.5

.5 

.5

0.3

.3 

.3

0.0

n n
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 107

Granitic intrusions in the central part of the Guard 

Station inlier also occur in metagabbro near the contact with the metasedimentary rocks. The intrusions 

are less abundant than those into the metasedimentary rocks, smaller in size, and generally contain no 

hornblende. Only one of these intrusions was sampled: a medium-grained leucogranite (specimen number R54, table 2.15) about 12 m wide on the east side 

of the Grande Ronde River about 900 m south of the 

Guard Station. The leucogranite is distinguished by 

muscovite in irregular crystals as much as 1.5 mm 

across, bent cleavage traces in some of the red-brown 

biotite, quartz showing strong undulatory extinction 

and some recrystallization, and near absence of iron 

oxide.

The southeastern biotite granodiorite and biotite 

tonalite make up a rather poorly exposed mass that 

extends for about 1.4 km along the southeast contact 

of the inlier. These granitic rocks intrude both the 

metagabbro and the metasedimentary rocks and are 

overlain on the east, southeast, and south sides by 

Cenozoic volcanic rocks (fig. 2.5). The foliation strikes 

mostly N. 20° E. to N. 60° E., approximately parallel 

to the contact with country rocks. The northern part 

of the mass is a biotite granodiorite, whereas the 

southern part is a biotite tonalite. Relations between 

the tonalite and granodiorite are unknown; presumably the granodiorite is younger. Modal analyses of 

two rock samples of the biotite granodiorite and two 

of the biotite tonalite are given in table 2.15.

The most distinctive feature of the biotite granodiorite is the unusual amount of strongly corroded plagioclase. In contrast, only minor corrosion of plagioclase 

occurs in the biotite tonalite. The granodiorite also is 

distinguished by muscovite in flakes as much as 0.8 mm 

across, by a very high content of quartz, and by such 

small amounts of iron oxide that none was tabulated 

during the point counting of six thin sections (16,462 

points). Both rock specimens of biotite granodiorite 

(table 2.15) are characterized by red-brown biotite.

Many features of the biotite tonalite duplicate characteristics of other granitic rocks in the Elkhorn 

Mountains north of lat 45° N. but are distinct from 

most granitic rocks elsewhere in northeastern Oregon. 

These features are local recrystallization of plagioclase 

at pressure points, mild bending of twinning lamellae 

in plagioclase, stringers of biotite and orthoclase along 

fractures in plagioclase, high content of quartz, and a 

dearth of iron oxide (0.02 volume percent). Biotite is 

red-brown in rock specimen R596 (table 2.15) but apparently not in specimen R597.

The southern quartz gabbro and quartz diorite (fig. 

2.5) are the dominant intrusions into metagabbro in 

this part of the inlier. Presumably the quartz gabbro

is older than the quartz diorite, according to the common mafic-to-felsic intrusive sequence. Near the north 

contact of the quartz diorite with the metagabbro is a 

banded gneiss with sharply defined layers of plagioclase and hornblende. The gneiss is more than 12 m 

wide and is interpreted as a regional east-west-trending shear zone that controlled the emplacement of the 

quartz diorite. Modal analyses of the southern quartz 

gabbro and quartz diorite are given in table 2.15.

The quartz gabbro is a medium-grained melanocratic rock dominated by many nearly equant crystals of 

hornblende mostly 3 to 6 mm across. Slender extensions of amphibole that project into adjacent plagioclase from well-defined crystal faces of hornblende are 

a notable microscopic feature of the gabbro. The extensions are in optical continuity with the large primary crystals of hornblende and are regarded as a 

product of thermal metamorphism. Another characteristic of the quartz gabbro is the preservation of most 

of the original augite within plagioclase rather than 

within hornblende. Subhedral to euhedral augite 0.1 

to 0.5 mm across is fairly common within plagioclase. 

In contrast, only small, irregular, and scattered relicts 

of augite remain enclosed within hornblende. Features of deformation are widespread. Twinning lamellae of plagioclase are bent as much as 12°; locally, 

plagioclase is crushed and, along most shears, annealed or recrystallized. Some hornblende has crystallized along fractures within plagioclase and also along 

a few crush zones. Quartz shows strong strain shadows and local recrystallization. Clouded crystals of apatite are another earmark of the quartz gabbro.

The quartz diorite is almost fine grained; most 

crystals are either slightly larger or smaller than 1 

mm. Poikilitic crystals of biotite as much as 5 mm 

across distinguish the rock megascopically by contrasting markedly with the relatively small crystal 

size of the remainder of the rock. The quartz diorite 

is characterized by a very large amount of hornblende but a comparatively low content of biotite. 

The absence of augite is surprising in a mafic quartz 

diorite having a color index of 37 and a quartz content of only about 10 volume percent.

A medium-grained dike of biotite tonalite (table 

2.15) without chilled contacts cuts the quartz diorite 

and locally confirms the assumption that the quartz 

diorite is older than the more felsic granitic rocks in 

the Guard Station inlier. Felsic granitic rocks north 

of lat 45° N. typically show deformation features, and 

the dike is no exception. Twinning lamellae in plagioclase are bent as much as 8°, and sutured contacts in 

quartz are relatively common. In addition, biotite extends along fractures in plagioclase. The dike contains small amounts of apatite, zircon, iron oxide,
108 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

and unresorbed hornblende. The low content of iron 

oxide (0.03 volume percent) and abundant quartz 

(table 2.15) suggest assimilation of siliceous components of the Elkhorn Ridge Argillite, but the brown 

of the larger biotite crystals has only a slight, if any, 

reddish tint.

The quartz diorite in the extreme southwestern 

part of the inlier is intruded by a medium-grained 

granite (table 2.15); this small body, exposed over an 

area only about 10 m across, is concealed to the 

south by alluvium and Cenozoic volcanic rocks. The 

granite is either an east-west-striking dike or the 

northernmost part of a concealed pluton. Deformation features in the granite are twinning lamellae 

bent as much as 5° and quartz showing strong wavy 

extinction as well as some recrystallized contacts. 

The red-brown color of biotite is sufficiently intense 

to recognize readily even in most crystals no more 

than 0.1 mm across. Accessory minerals (not recorded 

in 8,431 point counts) are apatite, zircon, monazite, 

and iron oxide.

METHOD OF EMPLACEMENT OF 

ELKHORN PLUTON

Studies of the emplacement of even small batholiths such as the Bald Mountain batholith involve 

formidable structural problems that invite an integrated team investigation. Only prolonged field studies in areas of favorable exposures can provide 

evidence of long-distance effects of emplacement 

(Balk, 1937). Unfortunately, most of the area surrounding the Bald Mountain batholith is not suitable 

for studies of emplacement mechanisms. Many pertinent rock exposures in the east are concealed by valley alluvium (fig. 2.2). Cenozoic volcanic rocks are 

within 1 km of the batholith on the southwest, and 

they overlie the granitic rocks on the north and 

northwest. Furthermore, major structural features 

have not been recognized in the Elkhorn Ridge Argillite (Gilluly, 1937; Pardee, 1941). The absence of established fold axes and well-defined marker horizons 

restricts the discussion of regional structural features 

to strike-and-dip patterns displayed by hundreds of 

individual beds.

In the Elkhorn Mountains the regional trend is essentially east-west and the strata commonly dip 

south at angles steeper than 50°. As intrusive contacts are approached, the regional trend becomes progressively disrupted. In most parts of the world, 

including the Elkhorn Mountains, structural patterns 

that surround small stocks are more straightforward 

than the pattern around a nearby batholith. Located

3 to 8 km south of the Bald Mountain batholith, the 

satellitic Lake Creek and Grays Peak stocks mostly 

are concordant, with strikes of surrounding strata 

approximately parallel to intrusive contacts (Taubeneck, 1957, figs. 4 and 5). The strong deflection of the 

regional trend around the two stocks implies that 

emplacement was by forceful intrusion.

The earliest intrusive units of the batholith are the 

three small gabbroic bodies (fig. 2.2). The forceful intrusion of the tonalite of Bald Mountain, which surrounds the norite of Badger Butte, left almost no 

trace of the norite's host rock. Therefore, little proof 

remains regarding the emplacement mechanism of 

the norite. Adequate metasedimentary rocks occur 

within and near the norite of Willow Lake (Taubeneck and Poldervaart, 1960, fig. 3) to suggest that 

this body was emplaced by forceful intrusion. Unlike 

the two norite bodies, only part of the quartz gabbro 

of Black Bear (fig. 2.2) is in contact with the tonalite 

of Bald Mountain. Accordingly, considerable evidence 

pertaining to the emplacement of this mass survives 

in country rocks. A dismembered screen between the 

quartz gabbro and the tonalite provides additional 

evidence. The strike of ribbon chert changes progressively in a clockwise direction from east-west at the 

southeast corner of the quartz gabbro (fig. 2.2) to N. 

25° W. on the west side. The strike of ribbon chert 

further changes within the dismembered screen from 

north-south on the southwest side to east-west on the 

east side. The eastern part of the quartz gabbro is 

concealed by alluvium. This 180° change in strike of 

the ribbon chert around the exposed contacts of the 

quartz gabbro of Black Bear is consistent with emplacement by forceful intrusion.

Gabbroic units in composite batholiths commonly 

occur as comagmatic mafic forerunners around large 

granitic plutons. The mafic rocks of the Bald Mountain batholith mostly were the products of high-alumina basaltic magma that was fundamental in 

transferring heat and mass into the crust. Accordingly, subsequent emplacement of the Elkhorn pluton 

was facilitated by intrusion of the three peripheral 

masses of gabbroic rocks, as well as by the intrusion 

of the quartz diorite of Limber Creek.

The zoned Elkhorn pluton that constitutes the largest part of the batholith is broadly concordant along 

about 85 percent of the exposed contacts (fig. 2.2). The 

two major areas of discordance are at the neck of the 

Monumental salient and along the northwest contact 

of the salient (fig. 2.2). Some major discordances are 

almost inevitable in a forcefully emplaced intrusion 

with the dimensions of the Elkhorn pluton. Observations of salt domes (Trusheim, 1960) and experiments 

on diapirism (Ramberg, 1981) show that mature dia-
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 109

pirs commonly crosscut the structures that they induced during their forceful intrusion.

Tertiary volcanic rocks west and northwest of Anthony Lake overlie the contact of the Elkhorn pluton 

and prevent an evaluation of emplacement mechanisms for this part of the pluton. However, limited 

exposures of metagabbro shown in figure 2.2 about 

2.5 km south of Chicken Hill indicate that the general concordancy of the pluton continues from the vicinity of the norite of Badger Butte northward under 

the Tertiary volcanic rocks. Nearly vertical gneissic 

layering in the metagabbro south of Chicken Hill 

strikes about N. 10° W. as is consistent with a northstriking concealed contact. The intensity of metamorphism in the metagabbro implies that the concealed 

contact of the Elkhorn pluton is not more than a few 

hundred meters to the east.

The area encompassed by figure 2.2 does not extend far enough to the west and south of the Elkhorn 

pluton to reveal the prevailing east-west regional 

trend that occurs in areas not concealed by Tertiary 

volcanic rocks. Maps by Evans (1986) and Elfrink 

(1987) show that trends are essentially east-west in 

country rocks more than 6 km west of the Monumental salient. The east-west trend in most of the area 

south of the Elkhorn pluton is shown on maps by 

Pardee (1941) and by Ferns and others (1987). Conspicuous disruption of the regional trend near the 

Elkhorn pluton is compatible with forceful intrusion 

of the largest unit of the batholith.

Emplacement of the Elkhorn pluton along the 

southeast side of the Belleview wedge and through 

the neck of the Monumental salient was preceded by 

intrusion of the quartz diorite of Limber Creek. The 

dike-like western quartz diorite (see section "Quartz 

Diorite of Limber Creek") extends eastward along the 

southeast side of the Belleview wedge where it is 

truncated at the neck by the tonalite of Bald Mountain (fig. 2.4). The original eastward dimensions of 

the western quartz diorite are unknown; speculation 

suggests that this quartz diorite extended continuously or discontinuously eastward through much of 

the neck and perhaps to the general location of the 

brecciated quartz diorite. The basis for this speculation is the occurrence of scattered xenoliths of the 

quartz diorite of Limber Creek in the tonalite. The 

largest xenolith, exposed on a flat surface, is a slab 

that tapers almost to a point at one end; dimensions 

are 1.7 m long with a maximum width of 0.6 m near 

the wide end but only 0.04 m at the narrow end.

Intrusions of the quartz diorite of Limber Creek 

are localized on either side of the neck of the Monumental salient. The brecciated quartz diorite is clearly a remnant of a larger intrusion that was mostly

removed during emplacement of the tonalite of Bald 

Mountain. The western quartz diorite unmistakably 

exerted a structural control on the emplacement of 

the Monumental salient, initially or ultimately, by 

confining tonalitic magma to the south (fig. 2.4). The 

tonalite of the salient probably also was confined 

south of the neck, at least in part, by extensions of 

the western quartz diorite and the brecciated quartz 

diorite. Country rock originally within the area of the 

neck is envisioned as being squeezed and dragged 

upward during extreme plastic deformation as is 

characteristic of rocks of the Belleview wedge. Subsequently, or very late in the severing of the country 

rock partition between rocks of the salient and those 

of the main batholith, most wallrock very near the 

contacts of the neck was bent aside into a crude concordance (fig. 2.4).

Two major discordances in metasedimentary rocks 

occur within about 75 m of the east contact of the neck 

and one on the west side of the neck. One marked 

discordance along the east contact of the neck is southwest of the brecciated quartz diorite where cherty argillite on the ridge crest at 2,347 m dips 75° NW. and 

strikes S. 65° W. (fig. 2.4) at right angles to the contact. 

The second discordance along the east contact occurs 

farther south at the corner of the neck where an almost rectangular area of ribbon chert is surrounded on 

the northeast and southeast by glacial deposits (fig. 

2.4). Most strikes in this small area vary from N. 65° 

E. to N. 85° W. with dips from vertical to 50° south. On 

the west side of the neck the only highly discordant 

strata are just north of where the western quartz diorite is cut off by the tonalite of Bald Mountain. For 

slightly more than 200 m (mostly north of the vertical 

attitude with east-west strike on fig. 2.4) the metasedimentary rocks strike into the contact at high angles.

Isolated boudins of chert and less common boudins 

of ribbon chert in Lindgren's (1901) Archean gneiss 

unit demonstrate extreme plastic deformation of country rock in the Belleview wedge during forceful emplacement of the Elkhorn pluton. Nearly all boudins 

in the gneiss occur as randomly scattered lens-shaped 

to ellipsoidal masses without stratigraphic continuity. 

Each dismembered boudin originally was part of one 

or more continuous beds. Most boudins are relicts of 

lamellae of ribbon chert originally interbedded with 

more extensive sequences of cherty argillite. Less 

commonly the boudins are relicts of layers of massive 

chert. Some boudins are pieces of ribbon chert that 

contain two or more chert lamellae. Lenticular to 

blocklike boudins of ribbon chert rarely exceed 1 m in 

length. The dimensions of most single-ribbon-wide 

boudins are less than 20 by 10 cm; many are less than 

8 cm in greatest dimension.
110 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

The boudins are revealing vestiges of extreme plastic deformation. Rocks of the wedge flowed upward as 

they were confined by intruding magma. The extreme 

plastic deformation and recrystallization require an 

exceptional heat source and thus imply that granitic 

magma must have underlain much of the wedge. Some 

planar structure in tonalite east of the Belleview 

wedge dips westward at angles of no more than 30° 

(fig. 2.2). The low dips, if projected westward, would 

extend under the eastern part of the wedge. The topography is such that all planar structure near the wedge 

may decrease in dip at depths of less than 300 m. 

Extreme metamorphism of the Belleview wedge rocks 

suggests that the three-dimensional form of much of 

the wedge very crudely resembles an almost flat-bottomed boat. Such a configuration for the wedge at 

depth is consistent with deformation and recrystallization features of the metasedimentary rocks, although 

it is probably an oversimplified proposal for this dilemma, which was faced first by Lindgren (1901).

Folding in metasedimentary rocks in the Belleview 

wedge represents the compressional stresses on wallrock by incoming magma. Many folds with planar 

limbs and very angular hinges are chevron folds. The 

folds are tightly appressed (fig. 2.14A), and some are 

almost isoclinal. Layers are attenuated along limbs 

and thickened at hinges. One of the best exposures is 

of a 48-m-thick sequence of tightly folded ribbon 

chert (fig. 2.14B) on a ridge at an elevation of 2,246 

m almost 3.6 km due south of Badger Butte and 1.2 

km from the nearest exposed contact of the Elkhorn 

pluton. The strongest folding in the wedge mostly occurs in the interior rather than along contacts with 

the tonalite of Bald Mountain. Concentration of the 

most-intense folding well within the wedge adds 

credibility to the idea that Belleview wedge rocks 

were confined by rising magma everywhere but on 

the west side. Country rocks unable to flow downward and compressed laterally could explain the internal folding.

Other structural features of rocks near and within 

the Elkhorn pluton suggest that forceful intrusion 

was the dominant process responsible for the emplacement of the pluton.

Peripheral joints and upthrusts occur in granitic 

rocks near many contacts of the Elkhorn pluton. The 

joints and upthrusts dip inward, generally at angles of 

20° to 40°. Narrow aplite dikes fairly commonly fill 

and accentuate the peripheral joints and upthrusts. 

Two accessible localities for observing peripheral 

joints are near the summit of Gorham Butte and in 

the minichasm of the Bulger Ditch (fig. 2.2.B). Good 

examples of the much less common peripheral upthrusts are near the contact southeast of Summit

Lake. The peripheral joints represent the tension produced by the upward flow of semisolid magma past 

steep confining walls. The Elkhorn pluton made additional room for itself along the peripheral upthrusts 

by crowding the walls outward as well as upward.

Pulled-apart lenses of ribbon chert in an argillaceous matrix denote the plastic deformation that occurred in wallrock south of the Elkhorn pluton to 

accommodate intruding magma. The shape and regularity of the lenses are related to the amount of argillaceous material in the rocks. Excellent examples of 

these lenses occur as much as 600 m south of the 

Elkhorn pluton, but they are well developed only in 

rocks that contain at least 20 percent argillaceous 

matrix. In strongly deformed rocks with considerable 

argillaceous material, most lenses range from 2.5 to 

18 cm in length and commonly are 0.3 to 1.4 cm 

wide; evidence of bedding survives because biotite 

shows minimum deflection around the lenses. Bedded 

chert with all lamellae 5 cm or more in width commonly includes only a small amount of argillaceous 

material. Plastically deformed lamellae are more oblong than lens shaped; some are sausage shaped. 

Dismembered ribbon chert with very little argillaceous matrix consists of tightly packed cylindrical 

chunks of chert totally unlike angular pieces of sedimentary or igneous breccias. Biotite conspicuously 

winds between the dismembered cylinders of chert. 

No bedding survives.

Another point worthy of consideration is that argillaceous wallrocks of the Elkhorn pluton are schists, 

whereas ribbon cherts and cherty argillites of the 

Belleview wedge are gneissic rather than massive 

hornfelses. Foliated wallrocks along and near contacts of the pluton are consistent with the outward 

compression and upward drag of a forcefully intruded 

magma.

A review of the field evidence indicates that forceful intrusion was the dominant process in the emplacement of the Elkhorn pluton. Xenoliths entrained 

throughout the granitic rocks of the pluton at the 

present level of erosion suggest that stoping of country rocks was of minor overall importance. The space 

taken by pieces of wallrock represents only a fraction 

of 1 percent of the total volume of granitic rocks in 

nearly all surface exposures of the pluton. However, 

a field observer sees in the granitic rocks evidence of 

only the final stages of intrusion at the present level 

of erosion; stoping may have taken place on a larger 

scale during the early stages of intrusion. Evidence 

for such stoping is discussed below.

Structures within the Elkhorn pluton indicate that 

major stoping occurred along at least part of the intraplutonic contact. In the vicinity of the leucogranodior-
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 111

ite of Trail Creek, the intraplutonic contact is 

essentially at right angles to the structure of the foliated tonalite on the west (fig. 2.2). Most exposures near 

the intraplutonic contact are poor, but the only reasonable conclusion is that the foliated tonalite is abruptly 

truncated by the intraplutonic contact. Much foliated

tonalite to the east must have been removed by stoping 

along a semicircular fracture that now is the intraplutonic contact. On a much smaller scale, the brecciated 

quartz diorite of Limber Creek (fig. 2.4) is a convincing 

example of stoping. The quartz diorite was emplaced 

in country rocks, but the only remaining contact with

B

FIGURE 2.14. Folding in metasedimentary rocks in Belleview 

wedge. A, Typical tightly appressed fold in bedded chert. 

Similar small-scale folds are present in metasedimentary 

rocks in most other areas near Elkhorn pluton, but nowhere 

is folding as widespread as within Belleview wedge. Coin is

about 18 cm in diameter. B, Rock specimen from 48-m thick 

sequence of folded chert in Crawfish Lake quadrangle (fig. 

2.2A) about 1.3 km from contact of Elkhorn pluton. Attenuation 

and thickening of layers is common. Coin is about 24 mm in 

diameter.
112 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

wallrocks is along the southeast side. Stoping must 

have removed the other wallrock contacts, as well as 

an unknown volume of the shattered quartz diorite.

The removal by stoping of a large volume of foliated tonalite east of the leucogranodiorite of Trail 

Creek, together with the field relations between the 

brecciated quartz diorite of Limber Creek and adjacent rocks, prompts a closer consideration of stoping 

as a significant emplacement mechanism throughout 

the Elkhorn pluton. A rigorous and direct assessment 

of stoping is possible only by comparing the sizes of 

xenoliths near wallrocks with the sizes of xenoliths 

in the interior of the pluton and by examining the 

upper parts of cupolalike extensions of granitic rocks 

near the margin of the Elkhorn pluton.

The largest xenolith along the contact south to east 

of Summit Lake (fig. 2.2) is a slice of ribbon chert 60 

m long and as much as 9 m wide. This inclusion, 

about 40 m from the contact of the pluton, verifies 

that some xenoliths near their source are much larger than any xenoliths within the central part of the 

pluton, which do not exceed 1 m in greatest dimension. Three larger xenoliths, all of metagabbro, occur 

in peripheral tonalite near the contact with the granite of Anthony Butte. The larger of the two xenoliths 

shown in figure 2.3 is nearly 300 m long and as 

much as 80 m wide. The notable discrepancy in size 

between peripheral xenoliths and interior xenoliths 

suggests that larger xenoliths originally within the 

central parts of the pluton sank rapidly enough in 

the region of slower cooling to leave only small xenoliths behind as an incomplete testimony of piecemeal 

stoping.

The roof and upper part of the Elkhorn pluton 

were removed by erosion before Albian time (Taubeneck, 1960). Accordingly, roof rocks can be examined 

only in the upper part of cupolalike extensions of tonalite near the margin of the pluton. If several of the 

dissected cupolas in country rocks near the contact of 

the Elkhorn pluton provide accurate small-scale models of the former roof of the pluton, stoping of as 

much as 15 percent of the roof rocks was in progress 

during consolidation of the magma.

A readily accessible cupola occurs at an elevation of 

about 2,225 m approximately 0.45 km N. 35° W. of the 

Upper Baisley-Elkhorn Mine (fig. 2.2B) and approximately 1.7 km S. 45° W. from the summit of Hunt 

Mountain. At this locality a precipitous stream has cut 

down through overlying metagabbro to reveal the top 

of a small cupola. The contact of the metagabbro and 

tonalite, exposed for about 12 m along its length, dips 

30° to 35° SW. Southwest of the stream in the immediate vicinity of the contact, xenoliths of metagabbro 

range from 2 cm to almost 2 m across. An inclusion of

metagabbro about 4 m in greatest dimension occurs on 

the northeast side of the stream. The xenoliths constitute 10 to 15 volume percent of the cupola.

Roof rocks in the small cupola southwest of Hunt 

Mountain probably are a small-scale replica of the 

long-vanished roof of the Elkhorn pluton. Xenoliths in 

the cupola are as much as four times as large as any 

xenolith in the interior of the Elkhorn pluton. Furthermore, the size differential between xenoliths in the 

margin and those in the interior of the pluton suggests 

that large xenoliths with comparatively small surface 

areas and rapid sinking velocities have disappeared 

into the interior parts of the pluton. Therefore, the size 

differential of xenoliths within the Elkhorn pluton 

combined with relations exhibited in small cupolas implies that stoping was a significant emplacement 

mechanism, but the magnitude of the stoping process 

can never be closely appraised in mesozonal batholiths 

such as those of northeastern Oregon because the critical field evidence is gone.

Northeast of the Elkhorn pluton and north of lat 

45° N. is a concentrated group of six intrusions within the batholith (fig. 2.3). The units are the quartz 

diorite of Wolf Creek, boundary quartz diorite unit, 

granodiorite of Indiana Mine Road, granodiorite of 

Beaver Meadow, granite of Anthony Butte, and leucogranite of Dutch Creek. The problem concerning the 

mechanism of emplacement is to determine how six 

different magmas successively acquired space for 

themselves to build up the composite batholith. No 

conclusive answer is possible partly because of the 

widespread distribution of overlying Tertiary volcanic 

rocks (fig. 2.3).

Significantly, each of these six units is elongated in 

an approximate east-west direction subparallel to the 

Monumental and Hunt Mountain salients (figs. 2.2B) 

of the Elkhorn pluton. By analogy with the two salients of the Elkhorn pluton, these six units probably 

were emplaced mostly by forceful intrusion. The dismembered screens in figure 2.3 reveal that a preferred 

route for intrusion of each successive unit was near a 

contact between an earlier intrusion and country rock. 

The pattern of the dismembered screens as well as the 

common strong planar structure and deformation features of the granitic rocks imply that emplacement 

was by forceful intrusion.

Two very large inclusions in the granite of Anthony 

Butte suggest that the emplacement of this unit was 

more complex than that of other units of the batholith. 

The southeastern part of the granite of Anthony Butte 

tapers to a narrow wedge (fig. 2.3) between the tonalite of Bald Mountain and country rock. The granite is 

strongly foliated as if forcefully intruded into a restricted space. Within the granite to the northwest are
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 113

two very large northwest-trending inclusions composed mostly of earlier foliated quartz diorite, some 

country rocks, and local hybrid rocks that nearly all 

occur as part of the large northern inclusion. Each 

inclusion is approximately 1.5 km long. Foliation and 

bedding in the inclusions strike dominantly northwestward. The two large inclusions are in the western 

and southwestern parts of the granite body; field relations suggest that magma in this area initially was 

intruded along three elongated conduits that subsequently coalesced. The characteristic mafic-to-felsic intrusive sequence in composite batholiths suggests that 

except for the leucogranite of Dutch Creek, the granite 

of Anthony Butte was the last intrusive unit emplaced 

within the area of figure 2.3. Earlier granitic rocks 

that surround the granite of Anthony Butte on nearly 

all sides should have restricted the forceful emplacement of yet another magma, whereas the two large 

inclusions of quartz diorite (relatively massive) might 

have favored multiple intrusion along paralleling 

channels rather than major intrusion along a solitary 

channel. In any event the features of the granite of 

Anthony Butte and associated rocks seem best interpreted as resulting from the forceful emplacement of 

tightly constricted magma that initially intruded 

along three more or less parallel passageways.

Likewise, the elongated, east-northeast-trending inclusion of country rock almost 1 km long in the granodiorite of Indiana Mine Road (fig. 2.3) almost surely 

is a clue to the growth of the magma body. The inclusion suggests that early in the emplacement history 

the granodioritic magma intruded along at least two 

east-west-trending channels before coalescing.

Thus these six units' overall strong planar structure and deformation features, their general eastwest configuration controlled by so-called basement 

rocks, and the features of their dismembered screens 

and large inclusions are consistent with forceful intrusion as the dominant emplacement mechanism for 

this part of the batholith.

The five other units of the batholith are small intrusions in the Elkhorn pluton. Any statement regarding 

the emplacement mechanism of the leucogranodiorite 

of Red Mountain and the leucogranite of Elk Peak (fig. 

2.2) is necessarily inconclusive, but emplacement of 

the granite of Clear Creek and the leucogranodiorite 

of Trail Creek was definitely controlled by structure. 

The granite of Clear Creek was emplaced alongside an 

inclusion of country rock in the tonalite of Bald Mountain, whereas the leucogranodiorite of Trail Creek intruded along the intraplutonic contact.

The leucogranodiorite of Mount Ruth (Taubeneck, 

1957) is the smallest and best exposed of these five 

small units that intrude the Elkhorn pluton. The leucogranodiorite, which crops out over a triangular area 

about 0.15 km2 at the summit of Mount Ruth (fig. 

2.15), is a sheetlike intrusion; its subhorizontal contact with the underlying tonalite of Bald Mountain 

dips 10° to 15° NE. The leucogranodiorite is moderately fine grained (0.5 to 1.0 mm) with scattered phenocrysts (about 4.0 mm across) of biotite and plagioclase. 

The resemblance of the texture of the leucogranodiorite to that of some epizonal rocks suggests that emplacement occurred after the batholith was deroofed 

or partly deroofed. In any event, presumably the subhorizontal walls of the intrusion moved apart because 

the hydrostatic pressure exerted by the leucogranodiorite magma exceeded the lithostatic pressure associated with the tonalite of Bald Mountain.

VOLATILES CONTENTS OF GRANITIC MAGMAS

IN RELATION TO INTERSTITIAL ZEOLITES AND

POTASSIUM FELDSPAR IN GRANITIC ROCKS OF

THE BALD MOUNTAIN AND

WALLOWA BATHOLITHS

Interstitial zeolites in granitic rocks were reported 

in the Cornucopia tonalite unit (Taubeneck, 1967) 

and now are known to occur in many granitic plutons 

in northeastern Oregon, westernmost Idaho, and 

northwest Nevada. The maximum known content of 

interstitial zeolite in any specimen of these granitic 

rocks is 0.63 volume percent but few contain more 

than 0.15 volume percent. Heulandite is by far the 

dominant interstitial zeolite; chabazite is comparatively rare.

Interstitial heulandite in the Cornucopia tonalite 

unit characteristically contains many small apatite 

crystals. However, the association of apatite with 

heulandite is much less common in other granitic 

intrusions, including the Elkhorn pluton. The heulandite-apatite association represents hydrothermal 

crystallization of a residual fluid rich in volatiles.

Interstitial heulandite and chabazite are easily 

overlooked and superficially resemble very small 

amounts of interstitial orthoclase such as may occur 

locally between plagioclase crystals in tonalites and 

granodiorites. Interstitial zeolites in the tonalite of 

Bald Mountain and granodiorite of Anthony Lake 

were incorrectly tabulated as orthoclase in earlier 

modal analyses (Taubeneck, 1957).

Heulandite in granitic rocks also occurs along fractures and as a selective replacement of one or more 

zones of oscillatory-zoned plagioclase. Both occurrences in the Elkhorn pluton are relatively unimportant. 

For example, only two fractures containing heulandite were seen during a close examination of 437 thin 

sections.
114 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

The following summary of zeolite distribution in 

granitic rocks of northeastern Oregon is restricted to 

interstitial zeolite in the Elkhorn pluton and in the 

four major units of the Wallowa batholith (fig. 2.11). 

The Elkhorn pluton contains the smallest amount of 

interstitial zeolite of these five compositionally zoned 

intrusions of tonalite and granodiorite. On the average, the Elkhorn pluton (number of samples (n)=193) 

contains 0.0054 volume percent interstitial zeolite. In 

contrast, unit I (n=76) of the Wallowa batholith (fig. 

2.11) contains 0.011 volume percent interstitial zeolite, unit 2 (rc=70) contains 0.0073 volume percent, 

unit 3 (n=124) contains 0.0294 volume percent, and 

unit 4 (n=90) contains 0.0087 volume percent. All 

four major units of the Wallowa batholith contain 

more interstitial zeolite than the Elkhorn pluton, but 

unit 3 contains in excess of 400 percent more zeolite.

Paucity of interstitial zeolite in the Elkhorn pluton 

is consistent with crystallization of a relatively dry 

magma. Because aplites are generally believed to 

represent late-stage residual liquids remaining after 

extensive crystallization of a granitic magma, the 

general abundance of aplites and pegmatites associated with a granitic intrusion may correlate with the 

relative abundance of interstitial zeolite in some plutons. The possibility of such a correlation is strengthened in northeastern Oregon by the occurrence of 

several times as much aplite and pegmatite in unit 3 

of the Wallowa batholith as in either the Elkhorn 

pluton or the other Wallowa units.

Interstitial zeolites are present in such small 

amounts in granitic rocks that their possible use as a 

general guide to the volatiles content of a magma 

seems very limited. A more common mineral with diFiGURE 2.15. View across Dutch Flat (lower left foreground) toward Mount Ruth, highest peak on left skyline. Solid line near 

summit of Mount Ruth denotes subhorizontal contact between 

overlying leucogranodiorite of Mount Ruth and underlying tonalite of Bald Mountain, as viewed from well within granodiorite of Anthony Lake (fig. 2.2). Contact dips toward observer, but 

sheetlike intrusion does not intersect any peaks in foreground 

nor higher peaks to left of photograph. View looking S. 25° W. 

from a point 3.6 km S. 65° E. of Anthony Lake.
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 115

agnostic properties that vary systematically such as 

potassium feldspar would be much more suitable in 

attempting to gauge the relative volatiles contents of 

granitic magmas.

Emeleus and Smith (1959) and many subsequent 

authors suggested that volatiles have a catalytic effect 

on the ordering of potassium feldspar. If so, potassium 

feldspar that crystallized from relatively dry magmas 

presumably should be characterized by little or no evidence of exsolution, a relatively small 2V, and no microcline twinning. As both orthoclase and microcline 

occur in different parts of the Elkhorn pluton, changes 

in properties such as exsolution textures, size of 2V, 

and distribution of crosshatch twinning might provide 

useful data regarding the relative water content of the 

crystallizing magma.

Exsolution lamellae are absent from potassium 

feldspar in the Hunt Mountain salient and rare in 

the Monumental salient but relatively common elsewhere south of the intraplutonic contact. The lamellae occur as straight and very narrow stringlike lines 

about 0.001 mm wide and as much as 0.6 mm long. 

The microperthite threads are so narrow that in thin 

section (plane-polarized light) they resemble cleavage 

traces and can scarcely be detected at low magnifications. The stringers commonly are confined to the interior of larger crystals. Individual members of a 

swarm of exsolution lamellae are generally rigidly 

parallel to each other, and the stringers are estimated to constitute as much as 15 volume percent of 

some crystals. Stringers do not cross grain boundaries but may extend to within less than 0.1 mm of 

boundaries. The scarcity of aplite and pegmatite 

dikes in the Hunt Mountain and Monumental salients is consistent with the absence or rarity of exsolution lamellae in both salients. Elsewhere in the 

Elkhorn pluton, however, the relative abundance of 

aplite and pegmatite dikes, such as near Anthony 

Lake, does not correlate locally with a greater number or width of exsolution lamellae.

Significant differences in the volatiles content of 

crystallizing magma in the Elkhorn pluton might be 

expressed by inward-increasing values of 2V for orthoclase and orthoclase microperthite. Average values of 

2V for potassium feldspar vary from about 49° in each 

of the two salients to about 52° for most of the remainder of the pluton. Higher values of 2V occur in a diagonal zone that crosses the pluton approximately 

midway between Summit Lake and Anthony Lake 

(fig. 2.16). In the Hunt Mountain salient, determinations of 2V for 12 crystals in 7 rock samples range 

from 47.0° to 51.0° and average 49.2°; in the Monumental salient, determinations of 2V for 14 crystals in 

6 rock samples range from 43.0° to 51.5° and average

49.3°; in the remainder of the pluton, excluding the 

diagonal microcline zone that crosses the batholith, 

determinations of 2V for 27 crystals in 14 rock samples range from 50.5° to 53.5° and average 52.2°.

The restricted range in values of 2V implies that 

potassium feldspar has a nearly uniform structural 

state and bulk chemical composition throughout most 

of the pluton. Average values of 2V for samples from 

an area near Anthony Lake where aplite and pegmatite dikes are concentrated are only about 3° higher 

than in the two salients where aplite and pegmatite 

dikes are uncommon. Accordingly, only a tenuous correlation is evident between the value of 2V in potassium feldspar and the abundance of aplite and 

pegmatite dikes in the Elkhorn pluton.

The occurrence of much microcline in a diagonal 

zone (fig. 2.16) that crosses the Elkhorn pluton suggests that the structural state of the crosshatched 

feldspar is not directly linked to the volatiles content 

of the Elkhorn magma. Crosshatched twinning occurs 

in most rock samples from within the zone but in 

only 3 of 197 rock samples from outside the zone. An 

assessment of the significance of the distribution of 

microcline requires consideration of the general textural features of rock types that contain this distinctive feldspar in contrast to rock types that lack 

microcline.

Rock types that lack microcline almost invariably 

contain quartz with moderate to strong undulatory 

extinction but little or no recrystallization. Parts of a 

single crystal of quartz commonly extinguish at angles that differ by several degrees and rarely by as 

much as 8°. Potassium feldspar, however, generally 

extinguishes simultaneously, and nearly all biotite is 

undeformed. In contrast, significantly different properties of quartz, potassium feldspar, and biotite characterize the various rock types within the zone of 

microcline (fig. 2.16). Quartz commonly shows very 

strong undulatory extinction, with parts of one crystal extinguishing at angles that differ by 5° to 12° 

and, rarely, as much as 16°. Sutured contacts are relatively common, and some zones of fine-grained 

(0.03-0.05 mm) recrystallized quartz extend through 

what originally was a single crystal. In a few rock 

types such narrow zones of fine-grained quartz may 

subdivide an original crystal into as many as seven 

or eight parts.

Much potassium feldspar shows wavy extinction 

with as much as an 11° difference in the extinction 

position of parts of one crystal. Many crystals show a 

dramatic change in 2V from core to rim. Angle changes 

in three examples are from 51.5° near the center of one 

large crystal to 68° near the edge, from 53.0° near one 

side of a second crystal to 72.0° near the other side,
116 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

and from 52.5° near the center of a third crystal to 

67.0° near a corner of a plagioclase that projects into 

the potassium feldspar. In each example the highest 

value of 2V is close to the crystal margin and near a 

pressure point associated with a corner of plagioclase. 

The three examples of widely varying optic-axial 

angles are in different rock samples from within the 

outer part of the zone of microcline (fig. 2.16). Near 

the center of the zone many, commonly small, crystals show pervasive crosshatched twinning. Some

large crystals, however, show microcline twinning everywhere except in their cores.

Deformation effects are less pronounced in biotite 

than in quartz or potassium feldspar, but crystals of 

biotite in rock samples from the microcline zone are 

much more likely to show bent cleavage traces than 

those in rock samples from outside the zone. Fracturing in some biotite accompanied plastic deformation. 

For example, one large biotite lath is crossed nearly at 

right angles to cleavage by slightly irregular fractures

EXPLANATION

+ Rocks with microcline twinning in potassium feldspar Numbers 

refer to samples that show anomalous microcline twinning

* Rocks without microcline twinning in potassium feldspar 

- Approximate boundary of microcline zone within Elkhorn pluton

FIGURE 2.16. Location of rocks of Elkhorn pluton with and without microcline twinning in potassium feldspar. Solid-line contacts surround

the norite of Badger Butte and divide the Elkhorn pluton from wallrocks.
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 117

that contain small (0.02-0.04 mm) fragments of biotite with cleavage directions rotated as much as 23°.

Exsolution lamellae within the zone of microcline 

(fig. 2.16) are significantly different from lamellae in 

rock samples outside the zone. Lamellae within the 

zone commonly are coarser, lens shaped, slightly 

curved, continuous to crystal boundaries, and only 

approximately parallel. In addition, some en echelon 

sequences occur. A few lens-shaped lamellae are as 

much as 0.03 mm wide near a midpoint and gradually taper in both directions to a sharp termination. 

Lens-shaped lamellae are mostly 0.05 to 0.30 mm 

long. Some lamellae gradually curve as much as 10° 

or rarely even 15° throughout their length, whereas 

others change direction only in a restricted part of 

the lamellae near one termination. Lamellae that extend to a crystal boundary commonly are noticeably 

coarser at the boundary than within the crystal. The 

corner of a plagioclase crystal that is surrounded by 

potassium feldspar is a favored locality for one or 

more lamellae to extend into the potassium feldspar. 

Most en echelon sequences involve three or more lamellae each less than 0.1 mm long. In summary, features of lamellae within the zone of microcline 

support the conclusion that the occurrence of microcline is a byproduct of deformation rather than a 

function of the volatiles content of the crystallizing 

magma.

Crosshatched twinning in three rock samples (fig. 

2.16) outside the diagonal zone of microcline occurs 

in a small part (0.03 mm2) of one crystal in specimen 

71, in two small crystals in specimen 49, and in 

many crystals of different size in specimen 29. Petrographic features indicate limited deformation in rock 

specimen 71, moderate deformation in specimen 49, 

and strong deformation in rock specimen 29. A second zone of microcline-bearing rocks may occur in the 

vicinity of rock specimen 29, but such a zone, if 

present, must be comparatively minor in extent. Additional minor zones of microcline are conceivable 

near rock specimens 71 and 49 (fig. 2.16).

The leucogranodiorite of Trail Creek has the most 

aplite and pegmatite dikes of any unit that intrudes 

the Elkhorn pluton. Therefore, the leucogranodiorite 

has high potential for a possible occurrence of microcline that can be related to a relatively high volatiles 

content. However, no crosshatched twinning occurs in 

any thin sections of the leucogranodiorite. Determinations of 2V for four crystals in three rock samples 

(table 2.10) range from 47° to 56°. Some exsolution 

lamellae that are several times as wide as those in 

undeformed orthoclase in the Elkhorn pluton provide 

the only strong suggestion of a magma with a higher 

volatiles content.

Aplite dikes are generally regarded as having crystallized from late-stage residual liquids of granitic 

magmas. Each of four aplite dikes from widely separated parts of the Elkhorn pluton has exsolution lamellae in potassium feldspar. Lamellae are approximately 

0.001 mm or less in width. The presence of exsolution 

lamellae in four small dikes of aplite, each less than 

15 cm wide, is consistent with an increased volatiles 

content in the aplitic melts. In contrast, no lamellae 

occur in the tonalite of Bald Mountain throughout the 

31-km2 Hunt Mountain salient and only rarely 

throughout the larger Monumental salient.

Microcline in two of the aplites, however, is not unequivocal evidence of increased volatiles because of the 

uncertainty of the magnitude of deformation effects. A 

small percentage of the microcline clearly is associated 

with healed microfractures that cross crystals of potassium feldspar. Narrow zones of delicate and incipient 

crosshatched twinning faithfully follow the microfractures. In addition, crystals of quartz with extinction 

positions that differ by 5° and, rarely, 6° occur in both 

aplite dikes and suggest that all microcline is related 

to deformation. A definitive study of the occurrence of 

microcline in aplite dikes of the Elkhorn pluton would 

require many additional specimens.

The Wallowa batholith (fig. 2.11) is composed of a 

more diverse array of granitic intrusions, including 

sizable bodies of aplite, than the Bald Mountain batholith and hence has greater possibilities for rocks in 

which microcline can be attributed to a volatile-rich 

magma. Microcline occurs in parts of each of the four 

major units of the batholith as well as in nearly all 

small units (not shown on fig. 2.11) of granodiorite and 

granite. Some microcline in the various intrusions 

may be unrelated to tectonism, but the evidence is 

inconclusive. Accordingly, the focus turns to the three 

largest known aplite bodies in northeastern Oregon. 

One of these aplites is within unit 1 of the Wallowa 

batholith, about 0.55 km north of Hobo Lake (fig. 2.11); 

this intrusion is quadrilateral in plan view, with longest and shortest boundaries about 125 m and 55 m, 

respectively. The second of the three aplites is in unit 

3, about 1.75 km northeast of Traverse Lake. Estimated dimensions are 250 m long and as much as 125 m 

wide. The third aplite, in unit 4 and slightly northwest 

of Olive Lake, is irregular with an estimated maximum dimension of 300 m. Each of the three aplitic 

units contains less than 1.5 volume percent biotite; 

has an average grain size that ranges in most rock 

specimens from 0.5 to 1.0 mm; is composed of nearly 

equal amounts of quartz, plagioclase, and potassium 

feldspar; and contains about 77 weight percent SiO2 

(C.O. Ingamells, analyst, 1962, Pennsylvania State 

University).
118 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Microcline is the only potassium feldspar in the aplite near Hobo Lake, whereas orthoclase constitutes 

at least 60 percent of the potassium feldspar in each 

of the other aplites. Petrographic characteristics of 

the aplite near Hobo Lake support the conclusion 

that the microcline crystallized from a volatile-rich 

magma. Of the three intrusions, the aplite near Hobo 

Lake has by far the strongest alteration of potassium 

feldspar to kaolinite, the most extensive distribution 

of flakes of muscovite (as much as 0.5 mm long) associated with plagioclase, and the coarsest exsolution 

lamellae (as much as about 0.025 mm wide). In addition, the aplite near Hobo Lake contains the only micropegmatite in thin sections of the three intrusions. 

Perhaps the relative abundance of needles of rutile in 

quartz is another indication of a high volatiles concentration. A crude correlation has been recognized 

in northeastern Oregon between the occurrence of rutile needles and the presence of hydrous alteration 

minerals. In any event, the aplite near Hobo Lake 

seemingly is an unequivocal example of a rock whose 

potassium feldspar is exclusively microcline and a 

product of crystallization of a magma with a high 

volatiles content. In contrast to their presence (locally) in the other two large aplites, neither annealed 

microfractures in potassium feldspar nor pressure 

points are present to suggest that any of the microcline is pressure induced.

MAJOR-ELEMENT COMPOSITIONS

Major-element compositions of 10 rock specimens 

from units of the Bald Mountain batholith are presented in table 2.16 to augment 11 analyses reported 

earlier (Taubeneck, 1957), which are not repeated 

here. The following discussion incorporates all analyses given earlier except that of the rhyolite porphyry 

dike (Taubeneck, 1957).

Most specimens of the norite of Badger Butte contain biotite and amphiboles that are regarded as 

postconsolidation. The specimen of the norite analyzed earlier (Taubeneck, 1957) contained only 1.9 

volume percent pyroxene but 9 volume percent biotite and 22.7 volume percent amphibole. A new analysis (specimen 99, table 2.16) is a much closer 

approximation of the composition of the original magma, and the analyzed specimen contains 27 volume 

percent pyroxene, no biotite, and 3.2 volume percent 

amphibole. This analysis plots in the field of highalumina basalt on both an alkali-silica diagram and 

an alumina-alkali diagram (Kuno, 1968), as does the 

single analysis of the norite of Willow Lake (Taubeneck, 1957).

The role of basaltic magmas in the petrogenesis of 

the granitic rocks of the Bald Mountain batholith 

cannot be assessed within the framework of available 

data, but the noritic components of the batholith suggest the involvement of magmas that had the composition of high-alumina basalt.

The major unit of the batholith ranges inward from 

mafic border rocks of quartz diorite and tonalite to a 

core of granodiorite. The chemical change inward, 

however, was not documented earlier (Taubeneck, 

1957) in the absence of an analysis of a sample of the 

border rocks. Rock specimen B6 (table 2.16) is a representative border-rock sample collected about 40 m 

from the south contact of the batholith, whereas rock 

specimens 15 and 27 (table 2.16) are granodiorites 

from the interior of the compositionally zoned intrusion. Analyses recalculated to 100 percent H2O-free 

show that silica ranges from about 61 weight percent 

in the border rock to about 69 weight percent in the 

core rocks. I^O, the only other major oxide to increase inward, ranges from 0.54 weight percent in 

the border rock to about 2.1 weight percent in the 

core rocks.

The leucogranite of Dutch Creek (rock specimen 

R22) is the most felsic unit in the batholith. Its silica 

content is 77.92 weight percent after recalculating 

the analysis to 100 percent, in comparison to 76.4 

weight percent silica in a 7.5-cm-wide aplite dike 

(Taubeneck, 1957) in the granodiorite of Anthony 

Lake. Normative quartz is 39.75 percent in the leucogranite but only 34.0 percent in the aplite. The differentiation index for the aplite, however, is 94.36, 

essentially the same as the value 94.22 for the leucogranite of Dutch Creek.

The analysis of the leucogranite of Dutch Creek 

closely approximates the compositions of the three 

large bodies of aplite (see previous section) in units 1, 

3, and 4 of the Wallowa batholith (fig. 2.11). Silica 

contents for the aplite bodies are 77.15, 77.05, and 

77.04 weight percent; normative quartz percentages 

are, respectively, 38.39, 36.71, and 37.26. Differentiation indexes are, respectively, 94.81, 94.42, and 92.95.

The leucogranite of Dutch Creek as well as the 

three large aplite intrusions indicate that the compositions of the last intrusive units within both the 

Bald Mountain and the Wallowa batholiths are those 

of minimum melts, irrespective of the manner in 

which these melts may have originated.

One important criterion for classification of granitic 

rocks involves the molecular ratio of Al2O3/(CaO+ 

Na2O+K2O), abbreviated A/CNK. Tonalites of the batholith are metaluminous, whereas most granodiorites 

and all granites are weakly peraluminous (A/CNK = 

1.0 to 1.1). The values of A/CNK imply that the granitic
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 119

TABLE 2.16. Chemical analyses and corresponding norms and modes of selected specimens from the Bald Mountain batholith

[Specimens B6, 15, and 27 analyzed at Pennsylvania State University in 1966 by C.O. Ingamells using wet chemical methods. Specimens R118, R121, R123, and 

R128 analyzed at Tohoku University, Japan, in 1973 by Ken-ichiro Aoki using wet chemical methods. Specimens 99, R22, and R125 analyzed in 1985 in 

laboratories of U.S. Geological Survey using X-ray fluorescence. LOI, loss on ignition; n.d., not determined; 0.0, <0.05; , not present]

Specimen No. 1 - 99 B6 R125 15 27 R118 R121 R123 R128 R22

Major-element oxides (weight percent)

SiO2 - - 

A1203 - 

Ti02 - - 

Fe203 - 

FeO

MnO

McrO IVlg \J 

CaO - 

SrO - _ - _ 

BaO 

Na20 

KoO

HoO+ 1M.£\J _ -

H20~ 

P205 

LOI - 

Total ------

Quartz 

f-\ ii i

Albite 

Anorthite 

Corundum 

Wollastonite 

Enstatite 

Ferrosilite - 

Magnetite 

Ilmenite - 

A r> of iff* _ _ ~.,_^. __ _ _

 - 51.3

 19.1

 1.07

 3.32

*» 18

__ 5.91

 9.74

 - n.d.

 - n.d.

 3.28

- n.d.

 n.d.

 .31

.42

 99.95

 2.40

1 AA

 27.91

 27.37

 _

 7.36

 12.15

 4.29

- 4.82

__ 1 94

74

60.70

17.83

.75

2.11

o 01

O.Ol

2.98

6.58

.09

.05

4.01

.53

.40

.13

.15

n.d.

99.71

CIPW

16.46

o 1 c o.lb 

34.26

29.36

1.09

7.49

3.29

3.09

1.44

.36

62.0

16.3

.63

.45

A CO 4.D-6

,\J*s OQ

3.10

6.53

n.d.

n.d.

3.34

1.72

n.d.

n.d.

.20

.51

99.49

68.59

1523

.35

1.46

1 Q(^

AQ .Uo

1.71

3.83

.05

.10

3.82

2.08

.04

.19

.13

n.d.

99.61

norms (weight

16.50

10.16

24.69

28.58

5.57

6.30

5.61

0.95

1.21

.43

Modes

27.35

-tfy OQ

32.57

18.29

.04

4.29

1.96

2.13

.67

.31

68.79

16.09

.32

1.38

1 -L.\AJ fiO

.07

1.26

3.77

.06

.11

3.89

2.14

.50

.00

.07

n.d.

100.05

74.51

12.71

.24

.26

9 1°. £t. 1O

.07

.49

1.84

n.d.

n.d.

3.01

3.77

.58

.09

.06

n.d.

99.76

73.06

13.44

.27

.74

1.80

AC .Ub 

.67

2.04

n.d.

n.d.

3.22

3.76

.63

.10

.07

n.d.

99.86

72.95

13.76

.27

.52

i qo. i.yo

AC .(Jo

.43

1.62

n.d.

n.d.

3.74

3.69

.51

.08

.04

n.d.

99.60

74.39

12.81

.23

.17

1 91 J..*7-L

.39

1.14

n.d.

n.d.

3.67

4.12

.58

.09

.05

n.d.

99.64

77.0

12.4

.07

.39

18

09

.15

.78

n.d.

n.d.

2.90

4.95

n.d.

n.d.

<.05

.44

99.33

percent, water free)

27.75

19 1^4. 79I £t

33.12

18.36

.69

3.16

1.41

2.01

.61

.17

36.85

0940 £i£t,*Vj

25.70

8.82

.48

_

1.23

3.46

.38

.46

.14

33.99

22.42

27.48

9.75

.53

1.68

2.38

1.08

.52

.17

32.13

rto no

31.96

7.85

.77

1.08

2.81

.76

.52

.10

33.30

£f±*\J\J 2460

31.38

5.38

.36

.98

3.19

.25

.44

.12

39.75

oq cr\

24.86

3.92

.81

.38

.01

.53

.14

(volume percent)

[0.0, <0.05]

iqjuartz Potassium feldspar

Plagioclase 

Biotite - 

Muscovite 

Hornblende 4 -j.

Hypersthene - 

Cummingtonite 

Opaque oxides -

02

_ 

 67.8

 _

 _

 3.2

QO.

- 17.7

 _ ._

 - 1.4

 IE n

.2

65.8

5.6

. 

12.5

.1

.4

17 Q l/.O

1.6

50.4

17.6

_

8.4

o. 7 O. i

.3

.1

9fi ^

8.5

53.9

7.4

_

3.1

_

.5

8.1

53.8

7.3

2.1

_

.9

01 1

O-L.l

20.6

42.4

5.7

.2

_

._

°.9 Q

23.4

38.0

5.7

_

.0

QO Q o^.o

23.6

37.2

4.3

_

.2

1.8

.0

O1 A ol.4

28.4

36.0

4.1

_

OQ A oo.4

29.1

30.5

1.6

.4

_.

...

.0

Nonopaque accessory

minerals 

Specimens

99

B6

R125

15

27

.4 .2 .1 .1 .2 .0 .0 .1 .1 .0

are from the following units:

Norite of Badger Butte

Tonalite of Bald Mountain

Quartz diorite of Wolf Creek

Granodiorite of Anthony Lake

Granodiorite of Anthony Lake

R118

R121

R123

R128

R22

Granodiorite of Beaver Meadow

Granite

Granite

Granite

of Anthony Butte

of Anthony

of Anthony

Butte

Butte

Leucogranite of Dutch Creek
120 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

rocks of the batholith are I-type in the classification of 

Chappell and White (1974). I-type geochemistry is verified by all analyses' plotting in the Na2O field on a 

diagram of Na2O versus K2O (White and Chappell, 

1983).

The alkali-lime index of Peacock (1931) is the silica 

value at which curves on variation diagrams intersect 

for CaO and total alkalies (Na2O+K2O). Peacock arbitrarily distinguished calc-alkaline rocks as those with 

an index between 56 and 61, and calcic rocks as those 

with an index greater than 61. The alkali-lime index 

for rocks of the Bald Mountain batholith is about 63.5 

in comparison to about 62 for the Mount Stuart batholith in Washington (Erikson, 1977), about 63 for 

widely scattered granitic rocks in the Klamath Mountains of Oregon and California (Hotz, 1971), about 

61.5 for the Wooley Creek batholith in northern California (Barnes, 1983), about 63 for the western Sierra 

Nevada (Ross, 1972), and about 64.5 for the southern 

California batholith (Larsen, 1948). By the chemical 

classification of Peacock, nearly all granitic suites in 

the Western United States west of the quartz diorite 

line (Moore, 1959) are calcic rather than calc-alkaline. 

Moreover, many tholeiitic suites, whether on continental margins or in oceanic islands, are calc-alkaline by 

the alkali-lime index. The Peacock index is useful in 

the comparison of regional data but does not define 

calc-alkaline suites as this term has been increasingly 

used since the work of Wager and Deer (1939) on the 

Skaergaard intrusion. In contemporary petrogenetic 

studies, the (Na2O+K2O)-FeO-MgO (AFM) plot has 

become widely adopted as a discriminant diagram to 

identify calc-alkaline trends.

An AFM diagram (fig. 2.17) for rocks of the Bald 

Mountain batholith has the typical calc-alkaline trend 

of little or no iron enrichment for intermediate compositions. Also, a Na2O-K2O-CaO diagram (Taubeneck, 

1957, fig. 18) indicates that the rocks have a trend 

similar to the common calc-alkaline trend found by 

Nockolds and Alien (1953).

The calc-alkaline trend also is clearly distinguished 

in a diagram showing the rate of iron enrichment relative to silica (fig. 2.18). Consistent with calc-alkaline 

suites, rocks of the Bald Mountain batholith show essentially no iron enrichment for compositions with 

less than 70 weight percent SiO2. However, marked 

enrichment occurs at higher percentages of silica because the relative percent decrease in MgO as it approaches zero on the abscissa of a Harker diagram 

progressively exceeds any corresponding absolute decrease in FeO* (total iron as FeO). The sharp increase shown in figure 2.18 in the FeO*/(FeO*+MgO) 

ratio for compositions with more than 70 percent 

SiO2 is graphically visualized from general Harker

variation diagrams for contents of FeO and MgO 

throughout the spectrum of granitic rocks.

Relative iron enrichment in late calc-alkaline plutonic units is responsible for many analyses of granites plotting in the tholeiitic field of Miyashiro (1974). 

For example, at least 32 of the 497 chemical analyses 

of batholithic rock samples of southern California 

(Baird and Miesch, 1984) are tholeiitic on Miyashiro's 

diagram. Most of these 32 rock samples have more 

than 74 weight percent SiO2, and many are undoubtedly leucogranites or alaskites.

FeO

Na2 O+K2O MgO

FIGURE 2.17. Alkali-iron-magnesium (AFM) ternary diagram for 

rocks of Bald Mountain batholith. Boundary between tholeiitic 

and calc-alkaline fields is from Kuno (1968).

O 

o>

0.5

45 50 55 60 65 70 75 

SILICA CONTENT. IN WEIGHT PERCENT

FIGURE 2.18. Graph showing rate of iron enrichment, as FeO*/ 

(FeO*+MgO), with increasing silica for rocks of Bald Mountain 

batholith. FeO* is total iron as FeO. Boundary between tholeiitic and calc-alkaline fields is modified from McBirney (1984, 

p. 307) by extending the termination of the original line from 

about 63 percent SiO2 to 70 percent SiO2.
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 121

CONCLUSIONS

The principal conclusions of this study are as follows:

(1) All units of the Bald Mountain batholith except 

some small felsic bodies mostly south of lat 45° N. 

have the overall features of mesozonal (Buddington, 

1959) plutons. Contacts with country rocks are mostly concordant in the broad sense with much of the 

discordance on the scale of outcrops. Relations with 

country rocks along the contact of the batholith 

range from relatively simple in the south to relatively complex in the north.

(2) The batholith is characterized by an extensive 

metamorphic aureole that includes widespread pyroxene hornfels-facies rocks. Salients of country rock 

that project far into the batholith are intensely deformed and commonly gneissic; they include the Archean gneiss unit of Lindgren (1901). Intense contact 

metamorphism of metagabbro produced a rock with 

gneissic layering that resembles some well-bedded 

metasedimentary rocks. The relative grain size of recrystallized quartz in ribbon chert near the batholith 

is a good approximation of the intensity of thermal 

metamorphism.

(3) The general concordance of structural trends in 

country rocks adjacent to the contact of the batholith 

and especially in dismembered screens between units 

of the batholith is strong evidence for forceful emplacement. Much wall rock yielded by plastic flow. 

Planar structure and deformation textures in granitic 

rocks are more pronounced in the northern onefourth of the batholith than elsewhere. Some tonalite 

in the north is gneissoid; the granite of Anthony 

Butte locally is almost gneissic. Structural style in 

the north is characterized by dismembered screens 

between intrusive units. The Elkhorn pluton is distinctive among the zoned intrusions of tonalite and 

granodiorite composition in northeastern Oregon in 

the occurrence of planar structure and inclusions of 

country rocks in the granodiorite core. The major 

structural feature within the Elkhorn pluton is an intraplutonic contact where core magma, mostly of 

granodioritic composition, intruded outer solidified 

tonalite. The intraplutonic contact provided the 

structural control for the emplacement of leucogranodiorite as elongated intrusions.

(4) Augite was the dominant early mafic mineral 

to crystallize throughout much of the tonalite and 

some of the granodiorite of the Elkhorn pluton. The 

small volume percent of calcic core in oscillatoryzoned plagioclase indicates that the granitic rocks 

crystallized from a melt that contained only a comparatively minor amount of solid constituents. An interlude is a concentration of one or several different

mafic minerals of small size in a distinct layer of plagioclase between consecutive oscillations in oscillatory-zoned plagioclase. The larger interludes are 

characterized by mottled plagioclase; similar mottling 

in cores of plagioclase cannot be cited as an unequivocal petrographic criterion for a restite origin of plagioclase cores in the Elkhorn pluton or elsewhere.

(5) Granitic rocks in the Elkhorn Mountains with 

reddish-brown to red biotite and minor or no iron oxide 

crystallized from magmas that assimilated graphitebearing metasedimentary rocks. These rocks are characterized by a much lower Fe2O3 /FeO ratio than the 

normal I-type granitic rocks of northeastern Oregon.

(6) The granodiorite of Beaver Meadow is notable 

for a reaction relation between orthopyroxene, cummingtonite, hornblende, and biotite.

(7) Much of the batholith crystallized from hot and 

relatively dry magmas as judged by early crystallization of pyroxene and by aureole rocks that include 

pyroxene hornfels facies. The general scarcity of pegmatite and aplite dikes in the Elkhorn pluton is consistent with the crystallization of a relatively dry 

magma, as is the paucity of interstitial zeolites in the 

tonalite and granodiorite.

(8) The zone of microcline that extends for 19 km 

across the Elkhorn pluton is attributed to deformation during final crystallization of the magma rather 

than to increased volatiles content.

(9) Early intrusive units in both the Bald Mountain batholith and the nearby Wallowa batholith are 

very similar norites with the composition of high-alumina basalt. However, compositionally zoned plutons 

of tonalite and granodiorite of the two batholiths differ in the decidedly higher content of quartz in rocks 

of the Bald Mountain batholith. The compositions of 

the last intrusive units of the two batholiths are 

those of minimum melts.

(10) Rocks of the Bald Mountain batholith show an 

exceptionally well defined calc-alkaline trend on an 

AFM diagram. The Peacock index does not identify 

the calc-alkaline trend of most granitic suites west of 

the quartz diorite line in Washington, Oregon, and 

California.

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McBirney, A.R., 1984, Igneous petrology: San Francisco, Freeman, 

Cooper and Co., 509 p.

Miyashiro, Akiho, 1974, Volcanic rock series in island arcs and active continental margins: American Journal of Science, v. 274, 

p. 321-355.

Moore, J.G., 1959, The quartz diorite boundary line in the Western 

United States: Journal of Geology, v. 67, p. 198-210.

Muir, I.D., 1953, Quartzite xenoliths from the Ballachulish Granodiorite: Geological Magazine, v. 40, p. 409-427.

Nockolds, S.R., and Alien, R., 1953, The geochemistry of some igneous 

rock series: Geochimica et Cosmochimica Acta, v. 4, p. 105-142.

Noyes, H.J., Wones, D.R., and Frey, F.A., 1983, A tale of two plutons Petrographic and mineralogic constraints on the petrogenesis of the Red Lake and Eagle Peak plutons, central Sierra 

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Pabst, Adolf, 1928, Observations on inclusions in the granitic rocks 

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Geological Sciences, v. 17, p. 325-368.

Pardee, J.T., 1941, Preliminary geologic map of the Sumpter quadrangle, Oregon: State Department of Geology and Mineral Industries, scale 1:96,000.

Pardee, J.T., and Hewett, D.F., 1914, Geology and mineral resources 

of the Sumpter quadrangle, Oregon: Oregon Bureau of Mines 

and Geology, Mineral Resources of Oregon, v. 1, no. 6, p. 3-128.

Peacock, M.A., 1931, Classification of igneous rock series: Journal 

of Geology, v. 39, p. 54-67.

Pitcher, W.S., Read, H.H., Cheesman, R.L., Pande, I.C., and Tozer, 

C.F., 1959, The main Donegal granite: Quarterly Journal of 

Geological Society of London, v. 114, p. 259-305.
2. A CLOSER LOOK AT THE BALD MOUNTAIN BATHOLITH, NORTHEASTERN OREGON 123

Pitcher, W.S., and Sinha, R.C., 1957, The petrochemistry of the 

Ardara aureole: Quarterly Journal of Geological Society of 

London, v. 113, p. 393^408.

Poldervaart, Arie, and Gilkey, A.K., 1954, On clouded plagioclase: 

American Mineralogist, v. 39, p. 75-91.

Presnall, D.C., and Bateman, PC., 1973, Fusion relationships in the 

system NaAlSi3O8-CaAl2Si2O8-KAlSi3O8-SiO2-H2O and generation of granitic magmas in the Sierra Nevada batholith: Geological Society of America Bulletin, v. 84, p. 3181-3202.

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Ross, D.C., 1972, Petrographic and chemical reconnaissance study 

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 1964, Cornucopia stock, Wallowa Mountains, northeastern Oregon: Geological Society of America Bulletin, v. 75, p. 1093-1116.

 1967, Petrology of the Cornucopia tonalite unit, Cornucopia 

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America Centennial Field Guide, v. 1, p. 327-332.

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3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE 

BLUE MOUNTAINS REGION OF OREGON, IDAHO, AND

WASHINGTON: IMPLICATIONS FOR THE 

GEOLOGIC EVOLUTION OF A COMPLEX ISLAND ARC

By TRACY L. VALUER

CONTENTS

Page

Abstract________________________________ 125

Tfif"Y*o(1iipi^0ii ________________________ __ _ _ _ __.__._ ______..____. 126

Acknowledgments 126

Methods 127

Field and laboratory methods 127

Rock classification and data presentation 130

Interpretations of petrogenesis and tectonic setting 131

Geologic framework 13 2

RflJrAl* tpT*T*3Tlf* ____-___«_____________________________________________.____ 1 38

(rl*iTlQStonf* tGT*T*3.TlG __-.__-____ -««-- -. _ -.____________..____ 138

Wallowa terrane 138

O1 Q§ J(1 6I*T*V tGTT*3.riG~~~~---_____~ "~~~------______._.________________ 139

IZGG tcrrfliiG __ _____ ___._.__________________________« 139

Age and field characteristics of pre-Tertiary igneous rocks

in the Blue Mountains region 140 

Late Paleozoic and Triassic igneous rocks of the Baker

terrane 140

Plutonic rocks of the Baker terrane 141

Canyon Mountain Complex 141

Other plutonic rocks of the Baker terrane 143

Petrogenesis of plutonic rocks of the Baker terrane ~ 145

Volcanic rocks of the Baker terrane 145

Late Paleozoic and Triassic igneous rocks of the Wallowa

terrane 149 

Plutonic and associated hypabyssal rocks of the Wallowa

terrane 150 

Sparta complex 150 

Oxbow Complex 161 

Cougar Creek Complex 163 

Wolf Creek-Deep Creek and Imnaha plutonic rocks -- 165 

Sheep Creek plutonic rocks 166 

Plutonic rocks of the southern Seven Devils

Mountains 166 

Plutonic rocks of the Salmon River canyon 167 

Plutonic rocks associated with the Blue Mountains

anticline - 167 

Wildhorse River dike 167

Petrogenesis of plutonic rocks of the Wallowa

terrane__________________________ 168 

Volcanic (flow) and associated hypabyssal intrusive

rocks of the Wallowa terrane 170

Late Paleozoic and Triassic igneous rocks of the Wallowa

terrane Continued 

Volcanic (flow) and associated hypabyssal intrusive

rocks of the Wallowa terrane Continued 

Permian volcanic (flow) and

associated hypabyssal intrusive rocks 

Triassic volcanic (flow) and associated hypabyssal

intrusive rocks 175 

Petrogenesis of Permian and Triassic volcanic (flow)

and associated hypabyssal intrusive rocks 178 

Triassic and Early Jurassic(?) igneous rocks of the Olds Ferry

170

iOCK.S ------ --- ----- ____________ ________ ________ j_oJVolcanic (flow) and associated hypabyssal intrusive

rocks ______________________________ 183 

Petrogenesis of igneous rocks of the Olds Ferry terrane 185 

Jurassic igneous rocks associated with the Coon Hollow

Formation of the Wallowa terrane 186 

Description of the igneous rocks 186 

Petrogenesis 187 

Jurassic to Cretaceous plutonic rocks of Blue Mountains

190

193

193

194

Description of the plutonic rocks 

Petrogenesis of the Jurassic to Cretaceous plutonic rocks -

Tectonic implications of petrologic and other data 

Tectonic setting of igneous rocks of the Baker terrane 

Tectonic setting of igneous rocks of the Wallowa terrane 194 

Tectonic setting of igneous rocks of the Olds Ferry terrane 195 

Tectonic setting of Jurassic to Cretaceous igneous rocks 196

Plate tectonic evolution of Blue Mountains island arc 197 

Paleomagnetic and faunal evidence for latitudinal

Plate movements and terrane tectonism 198 

Seven-stage model for plate tectonic evolution of Blue

Mountains island arc 199

£\6l6i 611 C6 S

ABSTRACT

Pre-Tertiary rocks in the Blue Mountains region of Oregon, Idaho, 

and Washington constitute a complex multi-generation intraoceanic 

volcanic arc, the Blue Mountains island arc, consisting of five

125
126 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

terranes: the Baker, Grindstone, Izee, Olds Ferry, and Wallowa 

terranes. The known igneous history of the Blue Mountains island 

arc extends from Pennsylvanian(?) to Early Cretaceous time. The 

oldest isotopically dated igneous rocks, Pennsylvanian(?) in age, are 

exposed in the Wallowa and Baker terranes. The largest volumes of 

igneous rocks formed in the Early Permian in the Wallowa and 

Baker terranes, in the Middle and Late Triassic in the Wallowa and 

Baker terranes, in the Late Triassic and Early Jurassic in the Olds 

Ferry terrane, and in the Late Jurassic and Early Cretaceous in all 

terranes except the Grindstone terrane.

Chemical data are used to assign the igneous rocks to tholeiiticand calc-alkaline-magma series and to interpret plate tectonic settings, magmatic-source rocks, and processes of melting, magma 

mixing, assimilation of older rocks, differentiation, and metamorphism. Erupted in an intraoceanic-arc setting, almost all late Paleozoic and Triassic igneous rocks of the Blue Mountains island arc 

are tholeiitic. The Jurassic to Cretaceous intrusive rocks are calcalkaline. The change from tholeiitic- to calc-alkaline-magma series 

had occurred by the Middle Jurassic.

The Baker terrane is petrologically and tectonically distinct. 

The Paleozoic and Triassic igneous rocks are tholeiitic, alkaline, 

and calc-alkaline, and crystallized in a large variety of oceanic tectonic settings (spreading ridge, ocean island or seamount, and volcanic arc). Most occur in fault-bounded blocks and as blocks or 

knockers in serpentinite-matrix melange. Many probably were accreted along the inner wall of one or more trenches and subsequently were deformed and tectonically shuffled extensively within 

the fore-arc regions of the Blue Mountains island arc.

Igneous rocks of the Wallowa terrane are related to magmaticaxis (volcanic-front) activity of an island arc. Most of the Permian 

and Triassic rocks belong to the island-arc tholeiite magma series. 

Siliceous flows are dominant volcanic rocks in the Permian sequence, whereas the Triassic flows are mafic, intermediate, and 

siliceous. Calc-alkaline hypabyssal rocks intrude the Middle and 

Upper Jurassic Coon Hollow Formation. The plutonic and volcanic 

(flow) rocks of the Olds Ferry terrane are mostly tholeiitic. The flow 

rocks are chemically similar to tholeiitic rocks of the Tonga and 

Kermadec island arcs of the southwest Pacific, except for higher 

concentrations of K-group elements (K, Rb, Cs, Ba, and Sr) in rocks 

from the Olds Ferry terrane. Calc-alkaline plutonic bodies of Late 

Jurassic and Early Cretaceous age occur in all terranes of the Blue 

Mountains island arc except the Grindstone. Some Jurassic to Cretaceous plutonic bodies were emplaced along terrane boundaries.

Paleomagnetic data indicate that the Wallowa terrane and, by 

inference, the other terranes in the Blue Mountains island arc 

were always in the northern hemisphere. Southward movement occurred in the Late Permian and Early Triassic and northward 

movement occurred in the Late Triassic to Late Jurassic. Final 

accretion of the arc to the North American craton occurred in the 

Early Cretaceous. The volumes and compositions of igneous rocks 

reflect major changes in plate motions.

Characteristics of the igneous rocks are strongly related to tectonic setting. Changes in tectonic setting and associated igneous 

rock characteristics probably occurred in the (1) late Paleozoic, especially during voluminous Early Permian siliceous volcanism, (2) 

Late Permian to Early Triassic during relative igneous quiescence 

and probable southward transport of the island arc, (3) Middle and 

early Late Triassic during igneous activity in the Wallowa and 

Baker(?) terranes, (4) Late Triassic through the Early Jurassic 

during igneous activity in the Olds Ferry terrane, (5) Middle and 

Late Jurassic during relative quiescence and northward transport 

of the island arc, (6) Late Jurassic to Early Cretaceous when calcalkaline plutonic bodies intruded at least four terranes, and (7) 

mid- and Late Cretaceous during accretion of the Blue Mountains 

island arc to North America.

INTRODUCTION

The origins of metamorphosed pre-Tertiary plutonic 

and volcanic rocks in the Blue Mountains region (fig. 

3.1) of eastern Oregon, western Idaho, and southeastern Washington have long puzzled and intrigued geologists. The close affinity of these rocks to a volcanic arc 

was first suggested by Hamilton (1963), who compared 

the chemistry of volcanic rocks from the Seven Devils 

Mountains and closely adjacent areas with those from 

the Aleutian island arc of Alaska. Most geologists who 

have worked in the Blue Mountains region now agree 

that it contains fragments of a pre-Tertiary island arc. 

In this chapter I refer to that arc as the Blue Mountains island arc. In general, all late Paleozoic and Triassic igneous rocks of the Blue Mountains island arc 

either formed directly within the arc or were accreted 

to it along subduction zones. Interpretation of the 

plate tectonic setting of Jurassic to Cretaceous plutonic rocks is less constrained, and I refer to those rocks 

within the broader context of the Blue Mountains region. The term "Blue Mountains region," as used in 

this chapter, includes parts of the Blue Mountains, 

Columbia Intermontane, Deschutes-Umatilla Plateau, 

and Joseph Upland provinces (fig. 3.2; Walker, 1990).

Volcanic arcs may lie within an ocean basin (intraoceanic island arc), be built on a continent (continental arc), or be a combination of the two. Studies of 

modern volcanic arcs show that there are significant 

petrologic and tectonic differences both within and between arcs. Some of the factors contributing to the 

differences include (1) age of the arc (maturity), (2) 

characteristics of subducting lithosphere, including 

age and evolution, (3) physical and chemical characteristics of the mantle and crust in which the magmas 

form and through which they ascend, (4) fractionation 

processes, (5) relative movements of the converging 

plates, (6) subduction angles, and (7) processes such 

as collisions of volcanic arcs with oceanic spreading 

ridges, oceanic plateaus, continents, and other arcs.

In this chapter, I (1) describe the pre-Tertiary igneous rocks of the Blue Mountains region using field, 

petrographic, and chemical data, (2) interpret the petrology of the igneous rocks, (3) discuss the implications of the petrology for specific tectonic settings 

within the Blue Mountains island arc, and (4) propose a plate tectonic model for the arc's geologic evolution on the basis of igneous petrology, rock ages, 

and paleomagnetic data.

ACKNOWLEDGMENTS

I am indebted to many scientific colleagues for help 

with this and associated studies. I particularly appre-
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 127

ciate the intellectual stimulation and friendship of 

H.C. Brooks, D.L. White, and D.W. Scholl, who have 

been patient mentors and associates. I thank W.H. 

Taubeneck for introducing me to many of the geologic 

problems in the Blue Mountains region and for his 

continued encouragement. I thank E.M. Bishop, H.C. 

Brooks, M.L. Ferns, and N.W. Walker for providing 

thoughtful reviews of an early draft of this manuscript. I am very grateful to the chemists within the 

U.S. Geological Survey for their painstaking efforts 

of providing high-quality data. I thank Jim Pearl and 

Leda Beth Gray for providing previously unreported 

radiometric ages using K-Ar methods. Without the 

willing and capable assistance of my wife and field 

companion, Trudy Vallier, I would not have completed the research for this chapter.

METHODS

FIELD AND LABORATORY METHODS

Beginning in 1963, I spent several complete summer field seasons and parts of many others in the 

Blue Mountains region, mostly in the Snake River 

canyon and adjacent areas of the Wallowa and Seven 

Devils Mountains (figs. 3.1-3.3). Thin sections of 

more than 300 igneous rocks collected throughout the 

region were studied to select appropriate samples for 

chemical analyses. Vallier (1967a) and Vallier and 

Batiza (1978) gave petrographic descriptions of some 

of the rocks that were chemically analyzed for this 

chapter.

The chemical data presented here are results of 

analyses completed during the 1967-1987 interval at

119° 118° 117° 116°

Walla Walla 

<£___WASHINGTON _

Enterprise

WALLOWA

MOUNTAINS SALMON 

RIVER

Homestead 

Cornucopia

MOUNTAINS

'New Meadows .

Sparta BROWNLEE J 

 DAM GREENHORN ELKHORN 

MTS

CANYON MTN 

STRAWBERRY

A IRONSIDE 

MTN

45° -

FIGURE 3.1. Map of Blue Mountains region of Oregon, Idaho, and Washington showing major features mentioned in text. For more

detail refer to figure 3.3 and to quadrangles shown in figure 3.2.
128 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

the U.S. Geological Survey. During that time, changes in procedures and instrumentation increased the 

precision and accuracy of some techniques. Because 

some already published major- and minor-element 

oxide analyses (Vallier and Batiza, 1978) were not

completed in U.S. Geological Survey laboratories, the 

rocks were reanalyzed to provide consistency in the 

data set; the new analyses are included in the tabulated data. The analyzed samples generally are typical of a given outcrop. Volcanic-rock samples were

118° 117° 116°

45° 

30'

45° 

00'

44° 

30'

River

^ /^ Lewiston

 y\ r^ \P.\O ^\w>

WASHINGTON

;? 1

O / CACTUS 

<( MTN

BOLES

WAPSHILLA \ CREEK 

\ \

OREGON

JIM CREEK BUTTE 

DEADHORSE RIDGE -

Pendleton

Enterprise

o La 

Grande

'ARBUCKLE 

MOUNTAIN SAWTOOTH 

RIDGE \

40 MILES 

I

I I I I

50 KILOMETERS _ , _ OXMAN\ Baker City Q \

JOHN 

DAY

BATES

BLUE 

CANYON

BOWEN ' 

VALLEY

SPARTA

DURKEE

ENCINA

LOST' 

BASIN

CORNUCOPIA

HALFWAY

MIN- , 

ERAL _/.

OLDS 

/ FERRY

IMNAHA

HARL 

BUTTE

HOMESTEAD

bow7

\

\

WOLF 

REEK

KERNANl 

POINT

HE 

DEVIL

CUPRUM

HORNET

KEUTERVILLE

Grangeville

O

GOODWIN MEADOW

N WHITE BIRD 

- GRAVE POINT

-LUCILE

v KIRKWOOD 

CREEK

'Area of 

figure 3.3

COPPERFIELD

FIGURE 3.2. Map of Blue Mountains region showing quadrangle province, whereas Olds Ferry terrane is split between Blue Mounmaps referred to in tables. Inset map shows physiographic prov- tains and Columbia Intermontane provinces. Wallowa terrane is in

inces (modified from Walker, 1990). Grindstone, Baker, and Izee Blue Mountains, Columbia Intermontane, and Joseph Upland

terranes (see fig. 3.8) are almost entirely within Blue Mountains provinces.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 129

117°00' 

46°00'

116°45' 116°00'

45°45' 

45°30' 

3.3. Major rivers, creeks, and towns of Snake River and Salmon River areas that are referred to in tables and 

text. For specific locations, see appropriate quadrangle map (fig. 3.2).
130 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 3.1. Classification scheme used for most igneous rocks in the Blue Mountains island arc

[Names granodiorite and granite are used only for unaltered high-K20, high-Si02 Jurassic to Cretaceous plutonic rocks. Quartz keratophyre is

high-Na20, low-I^O rhyolite]

SiC>2 content 

(weight percent)

Plutonic rock Volcanic rock

<52 Gabbro or norite; diabase

52-56 Meladiorite

56-62 Diorite

62-70 Quartz diorite; tonalite; granodiorite

>70 Trondhjemite; granite

Basalt

Basaltic andesite

Andesite

Dacite

Rhyolite; quartz keratophyre

selected mainly from aphyric flows; rocks with abundant phenocrysts were avoided.

Methods used for chemical analyses were X-ray fluorescence spectrometry, both wavelength-dispersive 

(Taggart and others, 1981,1987) and energy-dispersive 

(Johnson and King, 1987), for eight major- and two 

minor-element oxides and selected trace elements (particularly Rb, Sr, Ba, Y, Zr, and Nb); wet chemical and 

gravimetric analysis for ferrous iron, water (H2O+ and 

H2O), and carbon dioxide (Peck, 1964; Kolthoff and 

others, 1969; Tillman, 1977); direct-current arc-emission and optical-emission spectrometry (Golightly and 

others, 1987; Wilson and others, 1987) for a few trace 

elements; and instrumental neutron activation analysis for the remaining trace elements including the 

rare-earth elements (Baedecker, 1979; Baedecker and 

McKown, 1987). Some trace elements were determined 

by more than one method; the value used here depends 

on the presumed precision and accuracy of the data. 

Initial strontium isotope ratios were determined using 

methods that are discussed by Fleck and Criss (1985). 

Previously unreported K-Ar radiometric ages were determined using standard methods.

ROCK CLASSIFICATION AND DATA PRESENTATION

Most igneous rocks described in this chapter are 

classified by texture and chemical composition. The 

use of mineral modes for classification is hampered 

because of metamorphic effects. For example, regional 

metamorphism to the greenschist or amphibolite facies and contact metamorphism to the albite-epidote 

or hornblende hornfels facies caused extensive replacement of primary minerals in rocks older than 

Late Jurassic. Plagioclase feldspars have been partly 

to wholly replaced by albite. Mafic minerals are extensively replaced by chlorite and epidote. Primary 

textures in shear zones of some plutonic complexes 

were obliterated during deformation and metamorphism; some of the resultant foliated rocks are schist, 

amphibolite, mylonite, and gneissic mylonite.

The lava flows are both phyric and aphyric; the 

groundmasses of mafic and intermediate volcanic rocks 

(basalt, basaltic andesite, and some andesite) mostly 

have intersertal, intergranular, and hyalopilitic textures, whereas siliceous volcanic rocks (dacite and rhyolite) have felty groundmasses. Most rocks selected 

from large intrusive bodies have hypidiomorphicgranular textures.

A simple igneous-rock classification (table 3.1), 

based on texture and silica content, is used throughout this chapter, except for the Jurassic to Cretaceous plutons; the classification of rocks from them 

depends in part on the amount of K2O (for example, 

K2O content is used to distinguish granodiorite from 

granite). Igneous rocks older than the Late Jurassic 

are variably metamorphosed, and the prefix "meta" 

should be added to the rock name, but "meta" is 

omitted in this chapter in order to simplify descriptions and discussions. Some massive mafic-rock bodies in the Baker terrane have been so deformed and 

metamorphosed that original colors, textures, and 

minerals are destroyed. These mostly green rocks are 

referred to as greenstones.

Classification of metamorphosed siliceous igneous 

rocks is a problem. Many low-K2O trondhjemites and 

quartz diorites are Na2O rich and have been referred to 

by others as plagiogranite, soda granite, metatrondhjemite, and albite granite. I generally refer to those 

rocks as Na2O-rich trondhjemite and Na2O-rich quartz 

diorite. Low-K2O siliceous volcanic rocks containing 

abundant Na2O have been referred to as quartz keratophyre, soda rhyolite, and plagiorhyolite. In this chapter, the name quartz keratophyre is used because it is 

recognized by most igneous petrologists. The name spilite is not used, although metamorphosed basalt and 

basaltic andesite with Na2O contents greater than 4.5 

weight percent occur in Permian and Triassic mafic 

rocks throughout the Blue Mountains region.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 131

Igneous rocks are separated into tholeiitic-, alkalic-, 

and calc-alkaline-magma series. Gill (1981) explained 

the history of this classification and commented on the 

difficulties in assigning volcanic rocks to a specific 

magma series. Particular problems arose in separating the igneous rocks of the Blue Mountains island arc 

into magma series because of the metamorphic effects. 

Selected major-element plots, such as AFM ([Na2O + 

K2O]-FeOtotai-MgO) and FeOtotal/MgO ratio versus 

SiO2 content diagrams, and certain trace-element 

plots, such as rare-earth element (REE), Ti-Zr-Y and 

Ti-Zr, and FeOtotal/MgO versus Cen/Ybn ratio diagrams, serve to distinguish calc-alkaline from tholeiitic 

rocks; and some trace-element plots, such as Ti-Zr-Y, 

Ti-Zr, and REE diagrams, distinguish alkalic basalt. 

The FeOtotal/MgO ratio versus SiO2 content diagram 

is helpful for distinguishing tholeiitic and calc-alkaline 

rocks for SiO2 contents between about 53 and 70 percent. The FeOtotal/MgO versus Cen/Ybn ratio diagram 

is used to distinguish calc-alkaline from tholeiitic 

rocks because it shows both the iron enrichment and 

the flatness of the chondrite-normalized-REE diagram. Ce (rather than La) and Yb (instead of Lu) are 

used in the diagram because of suspected greater mobility of La during weathering and metamorphism and 

the better precision of Yb measurements. An arbitrary 

line at the Cen/Ybn ratio of 2.5 is used to divide tholeiitic and calc-alkaline rocks in the FeOtotal/MgO versus Cen/Ybn ratio diagrams.

There are some difficulties in using the diagrams 

mentioned above for distinguishing tholeiitic from 

calc-alkaline rocks. Some rocks show neither flat 

REE patterns nor conclusive iron enrichment; for example, rocks whose REE patterns show LREE (light 

REE) contents as 15 to 20 times chondrite abundances and HREE (heavy REE) contents as 8 to 12 times 

chondrite values and whose FeOtotal/MgO ratio versus SiC>2 content diagram is not diagnostic are difficult to assign to a specific magma series. Herein, 

those rocks are referred to as transitional between 

the tholeiitic and calc-alkaline series. High-A^Oa basalts (greater than 18 weight percent A^Os) are rare 

in the analyzed rock suites, in part because of the 

mostly aphyric nature of the rocks selected for chemical analyses.

Metasomatism has affected most pre-Jurassic 

rocks. In particular, sodium metasomatism was extensive. Some of the Permian volcanic rocks would be 

classified as alkalic if only major-element diagrams, 

such as the Na2O versus SiO2 content diagram (MacDonald and Katsura, 1964), were used. REE diagrams are not entirely conclusive because the LREE's 

may have been mobilized. The HREE's, on the other 

hand, probably were immobile during metamorphism. Reviews of suspected REE mobility in metamorphosed basalt are given by Frey and others 

(1974), Hermann and others (1974), Hellman and 

Henderson (1977), Menzies and others (1977, 1980), 

and Humphris and others (1978).

Most chemical data discussed in this chapter are 

new. A few of the analyses have also been reported by 

Ferns and others (1983, 1987), Walker (1986), and 

Ferns and Brooks (chap. 9, this volume); the rocks 

analyzed by Vallier and Batiza (1978) were reanalyzed 

for consistency and the new analyses are presented 

here. White and Vallier (1994) briefly discuss the 

chemistry of rocks in the Pittsburg Landing area of 

Oregon and Idaho. In general, except for the Baker 

terrane rocks, my samples seem to adequately characterize the pre-Tertiary igneous rocks of the Blue 

Mountains region. Additional chemical analyses and 

(or) comments on the igneous petrology are given in 

this volume (see Ashley, chap. 12; Bishop, chaps. 4 

and 5; Leeman and others, chap. 1; Hooper and others, chap. 11; Taubeneck, chap. 2; and Walker, chap. 

6). Dissertations, theses, and journal articles that contain chemical analyses of igneous rocks and (or) discuss important field relations and petrography of 

igneous rocks from the Blue Mountains region are 

those by Huntting (1942), Taubeneck (1957, 1967); 

White (1973); Skurla (1974); Trauba (1975); Shorey 

(1976); Almy (1977); Armstrong and others (1977); 

Phelps (1978, 1979); Mullen (1978, 1983, 1985); 

Gerlach (1980); Himmelberg and Loney (1980); Phelps 

and Av6 Lallemant (1980); Gerlach and others (1981); 

Brooks, Ferns, Coward and others (1982); Brooks, 

Ferns, and Mullen (1982); Sarewitz (1982, 1983); 

Scheffler (1983); Brooks and others (1983, 1984); 

Ferns and others (1983, 1987); Hunt (1985); Goldstrand (1987); and LaPierre and others (1988). The 

isotope ratios and isotopic ages determined and compiled by Hendricksen and others (1972), Armstrong 

and others (1977), Walker (1986), and Criss and Fleck 

(1987) are helpful for interpretations of the region's 

petrologic and tectonic evolution. For comparisons of 

plutonic rocks from the Blue Mountains region with 

Cretaceous plutonic rocks of the Idaho batholith, refer 

to Hyndman (1983) and to individual chapters in the 

volume by Vallier and Brooks (1987). Chemical and 

petrographic data for Eocene plutonic rocks of the Idaho batholith were reported by Motzer (1986).

INTERPRETATIONS OF PETROGENESIS 

AND TECTONIC SETTING

It is not easy to interpret igneous-rock data from 

modern volcanic arcs, let alone those from ancient arcs 

that have undergone several thermal and tectonic
132 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCICS, BLUE MOUNTAINS REGION

events. Furthermore, volcanic arcs are notorious for 

showing large differences in rock chemistry both 

across and along magmatic axes and within individual 

volcanoes themselves (Gill, 1981; Kay and Kay, 1994). 

In this chapter, comparison between igneous rocks of 

the Blue Mountains island arc and those of modern 

island arcs use as models the calc-alkaline and tholeiitic rocks from the Aleutian island arc (Kay and others, 

1990; Kay and Kay, 1994) and tholeiitic rocks from the 

Tonga and Kermadec island arcs of the southwestern 

Pacific Ocean (Ewart and others, 1973; Vallier and 

others, 1985; Ewart and Hawkesworth, 1987; Vallier 

and others, 1991).

Source rocks of volcanic-arc parent magmas are a 

subject of major debate (see reviews by Gill, 1981; 

Kay, 1984). Data presented in this paper are not adequate for a detailed treatment of source rocks and 

source areas.

Differentiation processes in arc magmas are thoroughly discussed by Grove and Baker (1983, 1984) 

and Grove and Kinzler (1986), who used experimentally determined phase equilibria of mafic melts at 

low pressure to study processes of fractionation, assimilation, and magma mixing. They showed that 

both calc-alkaline- and tholeiitic-series magmas can 

be generated from a common parent and that composition of the parent magma, pressure of fractionation, 

fugacity of water, extent of magma mixing, and crustal contamination are all important in determining 

differentiation paths.

Grove and Baker (1984) concluded that the proportions of olivine, plagioclase, and pyroxene crystallizing from a parent magma during fractionation are 

important compositional controls in determining 

whether a crystallizing magma follows a tholeiitic or 

calc-alkaline trend. They did not discuss the roles of 

magnetite and amphibole fractionation, although they 

believed that in some cases both are important. The 

fractionation of plagioclase, olivine or orthopyroxene, 

clinopyroxene (generally augite), and magnetite has 

the greatest effect on the final rock compositions, according to Gill (1981). Kay and Kay (1985a, b, 1994) 

discussed the important role of amphibole fractionation in the generation of calc-alkaline volcanic rocks. 

Garnet crystallization may also play a role in the evolution of calc-alkaline magmas, and particularly may 

account for the prominent HREE depletion.

The tholeiitic trend, according to Grove and Baker 

(1984), is produced by low-pressure fractionation of a 

basalt magma, characterized by the sequential crystallization of olivine, plagioclase, and augite, with 

plagioclase continuously dominating the assemblage. 

Crystallization continues until the remaining magma 

reaches a reaction point, where olivine reacts with

the liquid to form augite, plagioclase, and pigeonite. 

A dramatic increase in total Fe and a mild decrease 

in SiO2 occur in the liquid at the reaction point.

The calc-alkaline trend (Grove and Baker, 1984), in 

contrast, develops when olivine, calcic plagioclase, and 

augite crystallize in nearly equal mass proportions. 

This phase assemblage precipitates under conditions 

of moderate pressure and water undersaturation in 

the middle to upper crust. The crystallization process 

is characterized by a small increase in total Fe, an 

increase in SiO2, and a decrease of MgO in derivative 

liquids. Assimilation of a crustal component and mixing of basaltic liquids with siliceous residual melts are 

additional processes in the calc-alkaline-crystallization sequence. Relative proportions and types of crystallizing phases change in response to an increase in 

total pressure and an increase in Pu2O- Crystallization can occur at elevated pressure if basalt magma is 

slowed during its ascent or held in a crystallizing 

magma chamber.

I use three diagrams to separate rocks into magma 

series or trends and to compare rocks from different 

arcs: (1) the FeOt0ta\/MgO ratio versus SiO2 content 

diagram of Miyashiro (1974) that shows iron-enrichment trends (fig. 3.4) to distinguish tholeiitic from 

calc-alkaline fields; (2) the FeOtotal/MgfO versus Cen/ 

Ybn ratio diagram (fig. 3.5), which is used not only to 

distinguish calc-alkaline and tholeiitic series but also 

to show differences between the tholeiitic suites, such 

as Tongan versus Aleutian (Okmok Caldera) tholeiitic 

suites; and (3) REE diagrams for both plutonic rocks 

(figs. 3.6A, 3.6B) and lavas (fig. 3.7), which show relative LREE enrichment and HREE depletion for calcalkaline igneous rocks and relatively flat patterns for 

tholeiitic igneous rocks. The FeO^^i/MgO versus 

Cen/Ybn ratio diagram is a new method of portraying 

these magma series and may be applicable in other 

petrologic investigations.

Chemical and other geological data can be combined to interpret petrogenesis and tectonic settings 

of the igneous rocks (table 3.2). Almost all pre-Tertiary rocks in the Blue Mountains region, except for a 

few alkalic flows in the Baker terrane, are interpreted using the tholeiitic/calc-alkaline classification 

scheme. Rocks that do not fit solidly within specific 

fields of the diagrams or that have REE patterns between the flat tholeiitic trend and the sloping calcalkaline trend are called transitional. I follow many 

of Gill's (1981) interpretations of tectonic settings.

GEOLOGIC FRAMEWORK

Important to understanding the geologic framework 

of pre-Tertiary (particularly the Permian and Triassic)
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 133

rocks in the Blue Mountains region are the concepts 

that (1) they formed in a volcanic arc (Hamilton, 

1963), and (2) major tectonostratigraphic terranes 

compose the arc (Brooks and others, 1976; Vallier and

80

75

hZ

LLJ o

£ 70

Q.

hX

o 65

60

z

LLJ

Iz 

o 

o

55

50

45

III

2345678 

FeOtota |/MgO RATIO

EXPLANATION

ALEUTIAN ISLAND ARC 

Volcanic rocks 

High-MgO basalt 

High-Al2O3 basalt 

Tholeiitic rocks of Okmok Caldera 

| D | Calc-alkaline lavas

Plutonic rocks

: E : Mafic plutonic rocks 

JJFj Intermediate plutonic rocks

j_G | Siliceous plutonic rocks Dashed where 

uncertain

TONGA ISLAND ARC 

JLXJ Tholeiitic rocks of Ata Island

BLUE MOUNTAINS REGION 

JKJ| Cretaceous to Jurassic plutonic rocks

FIGURE 3.4. FeOtotal/MgO ratio versus SiO2 content for volcanic 

and plutonic rocks of Aleutian island arc (Kay and others, 1990; 

Kay and Kay, 1994) and volcanic rocks of Ata Island of Tonga 

island arc (Vallier and others, 1985). Area KJ, representing plutonic rocks of Blue Mountains region, was constructed from 

analyses of rocks of the Wallowa and Bald Mountain batholiths 

(W.H. Taubeneck, written commun., 1987). Calc-alkaline-tholeiitic compositional boundary from Miyashiro (1974).

others, 1977; Brooks and Vallier, 1978; Dickinson, 

1979; Brooks, 1979a; Mullen and Sarewitz, 1983; 

Silberling and others, 1984). The tectonostratigraphicterrane concept, used to explain complex geologic relations in western North America (see Davis and others, 

1978; Coney and others, 1980; Howell and others, 

1985), has been particularly helpful in understanding 

the evolution of the Blue Mountains island arc. The 

early-proposed names for terranes (oceanic crust terrane, arc terrane, and so on) became unacceptable because they had genetic connotations that did not hold 

up as processes responsible for the terranes' geologic 

evolution became better understood. Silberling and 

others (1984) renamed the terranes for localities rather than for interpreted origins. Vallier and Brooks 

(1986) recommended that those terrane names be 

used when discussing the pre-Tertiary rocks of the 

Blue Mountains region. Therefore, the terrane names 

of Baker, Grindstone, Izee, Olds Ferry, and Wallowa 

(Silberling and others, 1984) are used in this chapter. 

Some boundaries between the terranes (fig. 3.8) have 

been slightly modified and should be compared with 

the boundaries proposed by Silberling and others 

(1984).

In describing the terranes, I suggest a few new interpretations, some of which are dealt with in more 

detail elsewhere (see Ferns and Brooks, chap. 9, this 

volume). The basic premises are the following:

(1) All terranes in the Blue Mountains region are 

kindred. Since becoming coherent terranes, they have 

always been part of the same complex island arc.

(2) The Grindstone and Baker terranes are parts 

of the same terrane (or tectonic region) that formed 

within a late middle Paleozoic to Late Triassic island 

arc (Morris and Wardlaw, 1986). Their differences 

partly reflect multiple deformations within the forearc region or regions.

(3) Rocks within the inliers that occur in the Blue 

Mountains anticline of north-central Oregon and 

along the Tucannon River tributaries in Washington 

are more similar to rocks of the Baker terrane than 

to those of the Wallowa terrane. Igneous rocks from 

those areas, however, are discussed in sections concerning the Wallowa terrane. The small number of 

samples preclude a definitive terrane assignment.

(4) The highly metamorphosed and deformed rocks 

near New Meadows, Idaho (Bonnichsen, 1987), have 

protoliths of chert and ultramafic rocks that are more 

similar to rocks in the Baker terrane than to those in 

any other terrane of the Blue Mountains region.

(5) The Wallowa and Olds Ferry terranes are 

parts of volcanic edifices and subjacent intrusions 

built on, and within, a more ancient arc framework 

(or basement).
134 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

(6) The (informal) Sparta complex is part of the 

Wallowa terrane rather than the Baker terrane as 

mapped by Silberling and others (1984).

(7) The Izee terrane is part of an extensive belt that 

includes the Weatherby Formation (Brooks, 1979a) 

and presumably parts of the Squaw Creek Schist 

(Brooks and Vallier, 1978). This belt extends from

east-central Oregon into western Idaho, where it turns 

northward and merges with the border zone of the 

Idaho batholith.

(8) Rocks from both the Izee and Olds Ferry terranes compose parts of the Squaw Creek Schist 

(Hamilton, 1963) in the Riggins, Idaho, area and probably some of the rocks (for example, amphibolites and

/ / 

/ /

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/ / .

O

0>

o 

O

0)

4 5 

Ce n/Yb n RATIO

EXPLANATION

[Bj

l£l

[p]

i E: 

 Gl

ALEUTIAN ISLAND ARC 

Volcanic rocks

High-Al2C>3 basalt

Tholeiitic rocks of Okmok Caldera

Calc-alkaline lavas 

Plutonic rocks

Mafic plutonic rocks

Intermediate plutonic rocks

Siliceous plutonic rocks

TONGA ISLAND ARC

jjKJ Tholeiitic rocks of Ata Island, Tonga 

! Z j Tonga-Kermadec island arc lavas

BLUE MOUNTAINS REGION 

IBM! Bald Mountain batholith

FIGURE 3.5. FeOtotal/MgO versus Cen/Ybn ratios for igneous 

rocks of the Aleutian and Tonga-Kermadec island arcs. 

Data fields correspond to similarly designated fields in figure 3.4, with addition of Tonga-Kermadec island arc lavas. 

In this plot, Jurassic to Cretaceous plutonic rocks of Blue 

Mountains region are represented by data field for Late 

Jurassic Bald Mountain batholith (BM). Line, dashed and

queried where uncertain, divides tholeiitic from calc-alkaline 

fields. Tholeiitic fields show increasing iron content with relatively constant chondrite-normalized rare-earth-element ratios. Calc-alkaline fields show relatively constant iron content 

with increasing chondrite-normalized rare-earth-element ratios, n, chondrite normalized.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 135

La Ce Nd Sm Eu Tb

RARE-EARTH ELEMENTS

Yb Lu

EXPLANATION

ALEUTIAN ISLAND ARC 

lEll! Mafic plutonic rocks (n=6) (Oligocene and younger)

Til Mafic plutonic rocks of Terrible Mountain, Attu Island 

(n=5) (Eocene?)

FIGURE 3.6. Rare-earth element (REE) diagrams for Aleutian island arc plutonic rocks and Bald Mountain batholith. Normalizing values (Masuda and others, 1973) used for this diagram and 

for other REE diagrams in this chapter: La, 0.378; Ce, 0.976; 

Nd, 0.716; Sm, 0.230; Eu, 0.0866; Tb, 0.589; Yb, 0.249; Lu, 

0.0387. A, Mafic plutonic rocks of Aleutian island arc, including 

five rocks from Terrible Mountain on Attu Island (Vallier, unLa Ce Nd Sm Eu Gd Tb Yb Lu 

RARE-EARTH ELEMENTS

EXPLANATION

ALEUTIAN ISLAND ARC 

F . | Intermediate plutonic rocks (n=8)

Siliceous plutonic rocks (n=7)

BLUE MOUNTAINS REGION 

Bald Mountain batholith

pub. data, 1984) and plutonic rocks from Adak and other islands 

(Kay and others, 1990). Missing data points for Nd and Sm indicate elements that were below detection limit of the analytical instrument, were not determined, or had large coefficient of 

variation. B, Aleutian island arc intermediate and siliceous plutonic rocks and Late Jurassic Bald Mountain batholith of Blue Mountains region.

Ce Nd Sm Eu Tb

RARE-EARTH ELEMENTS

EXPLANATION

ALEUTIAN ISLAND ARC 

Tholeiitic rocks of Okmok Caldera

Calc-alkaline lavas

Yb Lu

FIGURE 3.7. Rare-earth element (REE) diagram for Aleutian calcalkaline (n=8) and tholeiitic (n=6) lavas and Tongan (Ata Island) 

tholeiite (rc=12). Note spread of Okmok Caldera data, reflecting 

broad range of differentiated lavas, and relative increase in 

light rare-earth element contents among Aleutian island arc 

calc-alkaline lavas (n=8). For normalizing values used, see figure 3.6. Missing Sm and Nd data points indicate that the elements were below detection limit of the analytical instrument, 

were not determined, or had large coefficient of variation.

TONGA ISLAND ARC 

Tholeiitic rocks of Ata Island
136 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 3.2. Chemical characteristics, petrogenetic interpretations, and tectonic implications of tholeiitic- and calc-alkaline-magma series

in the Blue Mountains island arc

[amph, amphibole; cpx, clinopyroxene; gt, garnet; mt, magnetite; ol, olivine; opx, orthopyroxene; plag, plagioclase]

Tholeiitic series Calc-alkaline series

Chemical characteristics

K-group elements 1 Low to intermediate Intermediate to high.

Ti-group elements2 Variable (low to high) Low.

FeOt0ta/MgO ratio High (1 to 6+; increases with Low (1 to 2; remains about constant

increasing SiO2). with increasing SiO2).

Cen/Ybn ratio Low (remains about constant with High (increases with increasing

increasing SiO2).

87Sr/86Sr ratio Low Low.

Petrogenetic interpretations

Sources and melting Partial melting of mantle wedge ±

material from subducting slab ± 

crustal material.

Parental magma Island-arc tholeiite 

Dominant form of differentiation Fractionation 

Type of fractionation Plag, ol, ±opx, cpx, mt 

Degree of mixing Low 

Water content Low 

Crustal assimilation Low 

Pressure for fractionation Low 

Temperature of eruption High 

Partial melting of mantle wedge ± 

material from subducting slab ± 

crustal material.

High-A^Os basalt.

Fractionation.

Plag, ol, ±opx, cpx, mt, ±amph, ±gt.

Low to intermediate.

High.

Intermediate.

High.

Low.

Tectonic implications

Setting Island arc3 Island and continental arcs.4

Crustal thickness Thin (< 20 km) Thick (>30 km).

Relative age of subducting slab Old Old to young.

Convergence rates >7 cm/yr <7 cm/yr.

Angle of Benioff zone High Intermediate to low.

K-group includes K, Rb, Cs, Ba, and Sr. 

2Ti-group includes Ti, Zr, Hf, Nb, and Ta.

Intraoceanic volcanic arc; volcanic activity occurs mostly along the magmatic axis or volcanic front. 

Volcanic activity is not necessarily along the magmatic axis or volcanic front.

FIGURE 3.8. Terranes of Blue Mountains region showing some differences from terranes designated by Silberling and others (1984). 

Major differences and suggested changes are (1) continuation of 

Izee terrane as narrow belt into western Idaho to include Weatherby Formation and Squaw Creek Schist, (2) change in contact 

between Baker and Wallowa terranes to include (informal) Sparta 

complex of Phelps (1979) in Wallowa terrane, (3) continuation of

Baker terrane into Cuddy Mountains area of western Idaho (see 

fig. 3.1 for geographic locations), (4) assignment of border zone near 

McCall, Idaho, to Olds Ferry terrane in this figure, although it may 

belong to Baker terrane, (5) pre-Tertiary rock inliers in Blue Mountains anticline of northern Oregon are assigned to Baker terrane, 

and (6) small inliers in Tucannon River area of southern Washington 

can be assigned to either Wallowa terrane or Baker terrane.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 137

120° 119° 118 116°

46° -

Walla 

VWalla WASHINGTON

River , OREGON /./

45

EXPLANATION

TERRANES OF THE BLUE MOUNTAINS REGION ___ ROCKS NOT ASSOCIATED WITH TERRANES

[.':'IZ-.':'| Izee terrane Cenozoic volcanic and sedimentary

r~OF=] Olds Ferry terrane 

JVVA^ Wallowa terrane

Baker terrane 

GR Grindstone terrane

Boundary between terranes Dashed 

where covered or uncertain

rocks

 ip: ;| Late Cretaceous Idaho batholith and 

related plutonic rocks

[ :}'P^;j Cretaceous to Jurassic plutonic rocks 

of the Blue Mountains region

Paleozoic and Precambrian North 

American cratonal rocks
138 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

gneisses) that occur within the border zone of the Idaho 

batholith south of Riggins (Aliberti, 1988; Manduca, 

1988).

BAKER TERRANE

The Baker and Grindstone terranes are structurally the most complex of the five recognized terranes 

in the Blue Mountains region. The Baker terrane is 

by far the most extensive (fig. 3.8). It had been called 

fragmented oceanic crust (Brooks and others, 1976), 

the oceanic crust terrane (Vallier and others, 1977), 

the dismembered oceanic crust terrane (Brooks and 

Vallier, 1978; Brooks, 1979a), and the central melange terrane (Dickinson and Thayer, 1978; Dickinson, 1979) before being named the Baker terrane by 

Silberling and others (1984).

The Baker terrane is a seemingly chaotic assemblage of Devonian through Late Triassic igneous and 

sedimentary rocks that have undergone varying, but 

pervasive, metamorphism, mostly to greenschist facies. Small amounts of amphibolite- and blueschistfacies rocks are present. Locally, hornfels-facies rocks 

surround Jurassic to Cretaceous intrusions. The Canyon Mountain Complex has many characteristics of 

an island-arc ophiolite (Bishop, chap. 5, this volume; 

Leeman and others, chap. 1, this volume).

Large parts of the Baker terrane probably formed in 

the fore-arc region, or regions, of the Blue Mountains 

island arc (Mullen, 1985; Bishop, chaps. 4 and 5, this 

volume; Leeman and others, chap. 1, this volume; Ave 

Lallemant, chap. 7, this volume). Most structural deformation apparently occurred in the Late Triassic 

and Early Jurassic(?). A complicating factor, however, 

is that volcanic and plutonic rocks in the Burnt River 

Schist, in the Canyon Mountain Complex, and in some 

parts of the Baker terrane (Mullen, 1985; Ferns and 

Brooks, chap. 9, this volume; Ashley, chap. 12, this 

volume; Bishop, chap. 5, this volume; Leeman and 

others, chap. 1, this volume; Peter Hooper, oral commun., 1987) have many chemical characteristics similar to those of rocks that form along the magmatic 

axis of a volcanic arc; this similarity suggests that the 

tectonic setting changed over time (for example, from 

magmatic axis to fore-arc or from fore-arc to magmatic axis). Another possibility is that some magmatism 

occurred within the fore-arc region itself (for example, 

see Marlow and others, 1988).

GRINDSTONE TERRANE

Reviews of structure, paleontology, and evolution of 

the Grindstone terrane were given by Buddenhagen

(1967), Blome and others (1986), Morris and Wardlaw 

(1986), Miller (1987), and Blome and Nestell (1991). 

Blome and Nestell (1991) reported differences in faunas, lithology, and metamorphic grade that suggest 

that the Grindstone terrane is a tectonic block that is 

separate from other terranes in the Blue Mountains 

region; the terranes probably were juxtaposed in the 

latest Triassic or Early Jurassic. Ferns and Brooks 

(chap. 9, this volume), however, showed that the 

southern part of the Baker terrane has Paleozoic faunal affinities with the Grindstone terrane; thus, at 

least in the late Paleozoic, the Baker and Grindstone 

terranes may have been kindred (see Morris and 

Wardlaw, 1986). Miller (1987) suggested that the 

Grindstone terrane is the continuation of similar terranes in the northern Klamath Mountains.

Blome and Nestell (1991) adequately explained some 

of the complex structural relationships. Briefly, they 

hypothesized that Paleozoic blocks slid and slumped 

into Permian and Lower Triassic base-of-slope and basinal sediments from shelf and other relatively shallow-water environments; these blocks and basinal 

sediments subsequently were covered by Upper Triassic and Lower Jurassic debris- and turbidity-flow deposits. Reshuffling along some of the faults mapped by 

Buddenhagen (1967) also must have occurred.

Long-distance transport of an exotic terrane-sized 

block probably is not necessary to explain either the 

geologic relationships in the Grindstone terrane or 

the differences between the Baker and Grindstone 

terranes. Lithological and chronological heterogeneities reported from fore-arc regions of the Tongan 

(Bloomer and Fisher, 1987), Marianas (Bloomer, 

1983; Fryer and others, 1985), and Aleutian (Vallier 

and others, 1994) island arcs indicate that fore-arc 

regions are very complex entities, particularly the 

parts that form along the inner trench wall. The 

Grindstone terrane probably is related to the Baker 

terrane in both time and space; upper Mesozoic sedimentary rocks were deposited in local basins that 

formed during and after deformation of the outer 

fore-arc region (including the inner trench wall) of 

the Blue Mountains island arc during the Late Triassic and Early Jurassic.

WALLOWA TERRANE

The Wallowa terrane (Silberling and others, 1984) 

was previously called the volcanic arc terrane (Vallier 

and others, 1977), the Wallowa Mountains-Seven Devils Mountains volcanic arc terrane (Brooks and Vallier, 1978), and the Seven Devils terrane (Dickinson 

and Thayer, 1978). The terrane includes the Wallowa
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 139

and Seven Devils Mountains in addition to the Snake 

River canyon north of Brownlee Dam (figs. 3.1, 3.8).

Wallowa terrane rocks are Pennsylvanian(?) 

through Early Cretaceous in age. A Permian stratified sequence of siliceous volcanic rocks (flows and 

associated volcaniclastic rocks) seemingly overlies an 

upper Paleozoic basement that probably was part of 

an ancient arc terrane. Permian stratified and intrusive rocks of the Wallowa terrane are overlain unconformably by Middle and Upper Triassic (Ladinian 

and Karnian) mafic to intermediate volcanic flow 

rocks and related volcaniclastic rocks; both the Permian and Triassic rocks were cut by intrusions of 

Middle and Late Triassic age. Massive carbonate and 

sandstone-mudstone flysch units of Late Triassic 

(mostly Norian) and Early Jurassic age unconformably cap the upper Paleozoic and Triassic (Ladinian 

and Karnian) sequences. Middle Jurassic conglomerate, sandstone, shale, and tuff and Middle and Upper 

Jurassic flysch-type turbidites occur in fault-bounded 

inliers in the Snake River canyon. Late Jurassic and 

Early Cretaceous intrusions cut the older rocks.

The Wallowa terrane is believed by some workers 

to be part of the Wrangellia superterrane (hereafter 

called Wrangellia for brevity), which extends from 

the Blue Mountains of Oregon to southeastern Alaska (Jones and others, 1977; Hillhouse and others, 

1982). This interpretation was questioned by Mullen 

and Sarewitz (1983), Sarewitz (1983), and Scheffler 

(1983) on the basis of geochemical data for Triassic 

volcanic rocks of the Wallowa terrane; these rocks 

are very different from lavas of the same age in undisputed parts of Wrangellia. Vallier (1986) suggested 

that Triassic volcanic rocks in the Wrangellia superterrane (specifically in the Wallowa terrane and in 

the Vancouver Island area of British Columbia) may 

have evolved in different parts of the same island 

arc, the Wallowa terrane rocks forming along the 

magmatic axis and the Vancouver Island rocks forming in the back-arc region during a rifting event in 

the Middle and Late Triassic. The chemistry of the 

Triassic volcanic rocks on Vancouver Island, however, 

do not support an island-arc origin.

OLDS FERRY TERRANE

The Olds Ferry terrane, as mapped by Silberling 

and others (1984), is part of the volcanic arc terrane 

of Vallier and others (1977) and the Juniper Mountain-Cuddy Mountains volcanic arc terrane of Brooks 

and Vallier (1978). Although Silberling and others 

(1984) included the Weatherby Formation (Brooks, 

1979a) in the Olds Ferry terrane, it is interpreted in

this chapter to be part of the extensive clastic belt 

that includes the Izee terrane.

The Olds Ferry terrane is in part lithologically 

similar to, but chronologically distinct from, the Wallowa terrane. Like the rocks in the Wallowa terrane, 

those in the Olds Ferry terrane evolved along the 

magmatic axis of an island arc. The Olds Ferry terrane volcanic rocks are Late Triassic (late Karnian 

and Norian) in age, and the intrusive rocks, exclusive 

of the Jurassic to Cretaceous plutons, range in age 

from Middle Triassic to Early Jurassic(?). In contrast, 

volcanic rocks of the Wallowa terrane range in age 

from Early Permian to Late Triassic (early Karnian), 

and intrusive rocks, exclusive of the Jurassic to Cretaceous plutons, are late Paleozoic (mostly Permian) 

to Late Triassic. Additional radiometric ages of plutonic rocks are needed from both terranes, particularly the Olds Ferry terrane.

IZEE TERRANE

The Izee terrane contains a thick coherent sequence of clastic sedimentary rocks with subordinate 

limestone and volcanic rocks of Late Triassic to Middle Jurassic age (Dickinson, 1979; Blome and others, 

1986). It is the flysch terrane of Brooks and Vallier 

(1978).

The Izee terrane (fig. 3.8) is here considered to be 

part of the long belt of clastic rocks that includes the 

Weatherby Formation of eastern Oregon and western 

Idaho (Brooks, 1979a, b; Mann, 1988) and parts of the 

Squaw Creek Schist (Hamilton, 1963) of the Riggins 

Group. As previously defined (Silberling and others, 

1984), the Izee terrane was more limited in extent. 

Rocks of the Izee terrane along the west boundary as 

defined by Silberling and others (1984) overlie an unconformity that was cut across rocks of the Grindstone and Baker terranes. Rocks of the Izee terrane 

that are exposed along the east boundary as defined 

by Silberling and others (1984) lie beneath Tertiary 

volcanic rocks. Rocks lithologically similar to those of 

the Izee terrane of Silberling and others (1984) crop 

out a short distance east of that east boundary, in the 

Ironside Mountain area, and form a wide but coherent 

belt that can be traced eastward in Oregon and thence 

northeastward across the Snake River into Idaho. 

These flyschlike rocks, called the Weatherby Formation in Oregon (Brooks, 1979a, b), are Jurassic in age 

(Imlay, 1986), in part the same age as most rocks in 

the Izee terrane of Silberling and others (1984).

As pointed out by Ave Lallemant (chap. 7, this volume), the Weatherby Formation is generally separated from the Olds Ferry terrane by a thrust fault
140 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

(Bruce, 1971; Hendricksen, 1974; Skurla, 1974; Mann, 

1988). On the north, the Baker terrane is faulted over 

the Weatherby Formation along the Connor Creek 

fault in eastern Oregon and western Idaho, but near 

Riggins, Idaho, the Squaw Creek Schist (probably in 

part correlative with the Weatherby Formation) is 

faulted over the Wallowa terrane, the probable result 

of thrusting during the Early Cretaceous accretion of 

the Blue Mountains island arc to North America. It is 

evident that sediments in the Weatherby Formation 

were derived in large part from the erosion of the Olds 

Ferry terrane (Brooks, 1979a, b).

The Izee terrane, as denned here, includes the 

Weatherby Formation and Squaw Creek Schist and 

forms a coherent belt of clastic rocks that extends 

into easternmost Oregon and westernmost Idaho. 

This belt narrows along the northwest side of the 

Cuddy Mountains and disappears to the northeast 

beneath the Columbia River Basalt Group; it reappears along the Little Salmon River south of Riggins, 

Idaho (fig. 3.8). Rocks with characteristics of the 

Squaw Creek Schist have been mapped north of Riggins along the Idaho batholith border zone (Myers, 

1982; Lund, 1984; Lund and Snee, 1988; Lund, chap. 

14, this volume; Snee and others, chap. 15, this volume), but because of the map scale in figure 3.8, the 

belt is not shown north of Riggins.

AGE AND FIELD CHARACTERISTICS OF

PRE-TERTIARY IGNEOUS ROCKS IN THE

BLUE MOUNTAINS REGION

Intrusive bodies in the Blue mountains region can 

be separated by age into three major groups: (1) late 

Paleozoic and Triassic (mostly in the Wallowa and 

Baker terranes), (2) Middle Triassic to Early Jurassic(?) (mostly in the Olds Ferry terrane), and (3) Late 

Jurassic to Early Cretaceous (in all terranes except 

the Grindstone).

Plutonic rocks of the oldest group are generally 

metamorphosed to the greenschist facies; some have 

no primary mafic minerals remaining, whereas others have pyroxene (commonly two pyroxenes) and 

blue-green actinolitic amphibole as major primary 

minerals. Plagioclase feldspars have been replaced 

wholly, or in part, by albite. Almost all of these plutonic rocks are K-poor and have tholeiitic chemical 

affinities. The youngest plutonic rocks, in contrast, 

are not metamorphosed, have pyroxene, amphibole, 

and biotite as primary mafic minerals in their modes, 

are K-rich compared to rocks of the oldest group, and 

have calc-alkaline chemical affinities. The Middle Triassic to Early Jurassic(?) plutonic rocks, mostly of

the Olds Ferry terrane, are generally tholeiitic and 

have field relations and petrographic characteristics 

that are similar to both other groups.

The geologic field relations of the group of Jurassic 

to Cretaceous plutonic bodies differ from those of the 

two older groups. The Jurassic to Cretaceous plutonic 

bodies, which include the large Bald Mountain and 

Wallowa batholiths, intrude all terranes except the 

Grindstone. Some of these younger intrusive masses 

were emplaced along terrane boundaries. Such intrusion between established terranes is apparent along 

the border between the Baker and Izee terranes. In 

contrast, the oldest (late Paleozoic and Triassic) intrusions do not intrude terrane boundaries, although 

parts of them have been incorporated into some terrane boundaries by faulting.

Late Paleozoic and Triassic plutonic rocks (group 1) 

are widespread in the Blue Mountains region, particularly in the Baker and Wallowa terranes. Although 

several papers have been written about the Canyon 

Mountain Complex (see reviews by Mullen, 1983; 

Leeman and others, chap. 1, this volume) and the 

(informal) Sparta complex (Gilluly, 1933; Prostka, 

1962, 1963; Almy, 1977; Phelps, 1978; Ave Lallemant 

and others, 1980), those two well-described igneous 

complexes make up only a small percentage of the 

exposed late Paleozoic and Triassic plutonic rocks. 

Many in the Baker terrane occur in fault blocks and 

irregular fault slivers and locally as knockers within 

serpentinite-matrix melange (see maps by Brooks 

and others, 1976; 1983; Brooks, Ferns, Coward, and 

others, 1982; Brooks, Ferns, and Mullen, 1982; Ferns 

and others, 1982, 1983, 1987).

Volcanic (flow) rocks of late Paleozoic and Triassic 

age in the Baker terrane generally are exposed in 

fault-bounded and discontinuous outcrops. Pillow 

lavas occur in places. In contrast, flows in the Wallowa and Olds Ferry terranes are stratigraphically 

more continuous. Pillowed flows are abundant in 

some areas along the Snake River and in the Seven 

Devils Mountains of the Wallowa terrane. Flows and 

associated volcaniclastic rocks are stratigraphically 

stacked to thicknesses greater than 1,000 m in places 

along the walls of Snake River canyon (Vallier, 1977).

LATE PALEOZOIC AND TRIASSIC IGNEOUS ROCKS 

OF THE BAKER TERRANE

Late Paleozoic and Triassic igneous rocks of the 

Baker terrane constitute only a small percentage of 

its total rock volume. These igneous rocks are fragments representing several tectonic regimes. The wide 

range in crystallization ages for these rocks, from the
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 141

TABLE 3.3. Lithology, sample locality, and field description of late Paleozoic and Triassic plutonic rocks of the Baker terrane

[Rocks within each group arranged by increasing SiO2 content to correspond to chemical analyses in tables 3.4 and 3.5]

Sample Rock type Sample locality Field description and remarks

Plutonic rocks of the Canyon Mountain Complex

CM-79-3 Quartz diorite

CM-79-1 Trondhjemite

CM-79-4 Trondhjemite

CM-79-2 Trondhjemite

Southeast ridge of Canyon Mountain; 

elevation 2,274 m (7,460 ft); NW1/4 

NE1/4NW1/4 sec. 28, T. 14 S., R. 32 

E.; John Day quadrangle; Oreg.

North side of Berry Creek; elevation 

1,417 m (4,650 ft); NE1/4SW1/4 sec. 

31, T. 14 S., R. 32 E.; John Day 

quadrangle; Oreg.

On ridge north of Sheep Gulch; 

elevation 1,661 m (5,450 ft); NE1/4 

SE1/4 sec. 24, T. 14 S., R. 31 E.; John 

Day quadrangle; Oreg.

On ridge about 200 m from gravel pit 

that is north of Berry Creek; 

elevation 1,567 m (5,140 ft); SE1/4 

SE1/4SE1/4 sec. 25, T. 14 S., R. 31 E.; 

John Day quadrangle; Oreg.

Quartz-rich cumulate layer within 

gabbroic rocks. Permian U-Pb 

radiometric age (Walker, 1986).

Coarse-grained dike in dike complex of 

varied lithology (diabase, basalt, 

quartz keratophyre, quartz diorite).

Large dike, one of small quartz-rich 

intrusions that cut gabbro.

Large dike, one of small quartz-rich 

intrusions that cut gabbro.

Other plutonic rocks of the Baker terrane

BR-79-1

V-l-79

Diorite

Trondhjemite

About 5 km up Burnt River canyon 

(near "Y" split in canyon); along 

Burnt River road, elevation 900 m 

(2,950 ft); unsurveyed; Lost Basin 

quadrangle; Oreg.

Along U.S. Highway 30 in Pleasant 

Valley; elevation 1,161 m (3,810 ft); 

NW1/4NW1/4 sec. 30, T. 10 S., R. 42 

E.; Oxman quadrangle; Oreg.

Small irregular intrusion associated 

with volcanic breccia. Intrudes Burnt 

River Schist.

Dike rock in classic association with 

serpentinite and gabbro.

Pennsylvanian to the Late Triassic, combined with an 

extraordinary diversity of igneous rock types indicates 

a long and complex pre-Jurassic evolution.

Several metamorphic facies are represented in 

rocks of the Baker terrane (Kays and others, 1988; 

Evans, chap. 8, this volume; Bishop, chap. 4, this volume). The metamorphic-rock diversity is related to a 

broad range of pressure and temperature (PIT) conditions. Both low-P/T greenschist and actinolite facies and high-P amphibolite and blueschist facies 

have been recognized. In addition, Jurassic to Cretaceous plutons produced wide contact aureoles that 

overprinted older metamorphic facies.

PLUTONIC ROCKS OF THE BAKER TERRANE

The six plutonic rocks described in this section (tables 

3.3-3.5) include (1) quartz diorite (sample CM-79-3)

andtrondhjemite (samples CM-79-1, CM-79-2, and CM79-4) from the Canyon Mountain Complex and (2) 

diorite (sample BR-79-1) and trondhjemite (sample V-l79) from other parts of the Baker terrane.

CANYON MOUNTAIN COMPLEX

The Canyon Mountain Complex is a dismembered 

ophiolitic body composed of peridotite, gabbro, and 

subordinate amounts of siliceous, intermediate, and 

diabasic intrusive bodies (Thayer, 1956, 1963, 1977; 

Thayer and Himmelberg, 1968; Ave Lallemant, 1976; 

Dungan and Ave Lallemant, 1977; Himmelberg and 

Loney, 1980; Gerlach, 1980; Gerlach and others, 1981; 

Mullen, 1983; Walker, 1986 and chap. 6, this volume; 

Pearcy, 1991; Bishop, chap. 5, this volume; and Leeman and others, chap. 1, this volume).
142 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 3.4. Major- and minor-element oxides of late Paleozoic and Triassic plutonic rocks of the Baker terrane

[Oxides in weight percent, normalized to 100 percent; prenonnalization volatile contents included to indicate amount of alteration. , not determined or below

instrumental detection limit]

Plutonic rocks of the Canyon Mountain Complex

Other plutonic rocks of 

the Baker terrane

Sample CM-79-3 CM-79-1 CM-79-4 CM-79-2 BR-79-1 V-l-79

SiO2 

Tin

A12O3 -

^ o

Fe2O3 - 

^ O

FpO _____ _

MgO ------

MnO

CaO -----

Na2O -----

j^If 2Ci (j 

Pr\ 2O5 -

TT f-\ +

H

(-\ 2U 

rr»

FeOtotalMgO -

69.10

CO

13.56

I

CO

9 49

2 nQ

.07

6.56

2 DO

nc.

n^

1.35

.31

nc.

2.35

nc on

1 C

1 o 01

1 97

9 04 A.^Jri 

01 .O -L

.04

1.84

A 8.9

00

0.94

.16

1 C

2.20

70 i c

on ,£*\J 

1 O 1 P.

70

1.17

.59

09

1.24

^ <^n

1 c

0.76

.17

n^

3.17

79 17

no

U Q7

1 99

r*c

OQ

1 19

5

CC

.04

0.73

.14

.07

6.02

^7 91

.74

19.51

1 1 %

3

QQ

2 7Q

np.

5 1 ^

6 OQ

79

91

2.54

.14

99

1.80

HA. J^

OQ

1 9 ftfl

9 49

.42

en

no

A OC

4 CO

.01

02

0.59

.10

.04

5.20

The Canyon Mountain Complex is of Early Permian (intrusive) age, as shown by concordant 206Pb/ 

238U ages on zircon of 276 and 268 Ma (Walker, 1986 

and chap. 6, this volume) and 40Ar/39Ar ages as 

young as 262 Ma (Ave Lallemant and others, 1980; 

Mullen, 1983). On the basis of the concordant 206Pb/ 

238U ages, Walker concluded that at least two separate crystallization events are represented in the 

complex, one at or before 276 Ma and the other about 

268 Ma. One sample analyzed by Walker (1986 and 

chap. 6, this volume) yielded discordant ages on zircon of 314±29 Ma (206Pb/207Pb) and 278 Ma (206Pb/ 

238U), leading Walker to speculate that 278 Ma may 

be a minimum age and that part of the Canyon 

Mountain Complex may be as old as 314±29 Ma, or 

Pennsylvanian(?). However, Walker did not extend 

the formal age of the Canyon Mountain Complex to 

Pennsylvanian(?) on the basis of these data.

The siliceous intrusive bodies sampled for this 

study (table 3.3) include a quartz diorite (sample 

CM-79-3) within the large gabbro body, two small 

trondhjemite intrusions (samples CM-79-2 and CM79-4) that cut upper parts of the same gabbro, and a 

trondhjemite dike (sample CM-79-1) in the sheeted 

member of the ophiolitic body. All rock samples have 

hypidiomorphic-granular textures with cloudy plagioclase, quartz, blue-green amphibole, and rare clinopyroxene as primary minerals. Chlorite and epidote are 

common secondary minerals.

The siliceous rocks have low K^O contents (table 

3.4). Three of the rocks have high Na2O contents and 

low amounts of CaO, suggesting extensive soda metasomatism. High Na2O contents also occur in other siliceous intrusive bodies in the Canyon Mountain 

Complex (Leeman and others, chap. 1, this volume). 

A FeOtotal/MgO versus SiO2 diagram (fig. 3.9) shows 

the large range in silica content for Baker terrane 

plutons (mostly the Canyon Mountain Complex; 

Thayer, 1977; Leeman and others, chap. 1, this volume) and the trend towards Fe-enrichment, particularly if compared to the calc-alkaline Jurassic to 

Cretaceous plutons.

Trace-element data (table 3.5) are similar to those 

from other siliceous intrusive bodies in the Canyon 

Mountain Complex (Gerlach and others, 1981; Leeman and others, chap. 1, this volume). Samples CM79-3 (quartz diorite) and CM-79-1 (trondhjemite dike 

rock) have similar trace-element concentrations with 

relatively high Sr contents and depleted REE contents (for example, subtotals of seven REE's are 

21.61 and 23.96 parts per million for samples CM-79- 

3 and CM-79-1, respectively, compared to subtotals of 

46.28 parts per million for sample CM-79-4 and 43.69 

parts per million for sample CM-79-2).
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 143

TABLE 3.5. Trace elements of late Paleozoic and Triassic plutonic rocks of the Baker terrane

[Results in parts per million. Analysis by neutron activation unless otherwise noted. , not determined or below detection limit or discarded because of high

coefficient of variation]

Plutonic rocks of the Canyon Mountain Complex

Other plutonic rocks of 

the Baker terrane

Sample CM-79-3 CM-79-1 CM-79-4 CM-79-2 BR-79-1 V-l-79

Rb -----

Gli*

Ba -

rni.

U~ -

La ----- -

Ce» 

Sm -

TTi-i

/~t J

TVU

Tm ----- -

Yb ----- -----

Lu -----

Y - .

7-r

Tfl

Nb -

Ni . .

Co -

rv

c?«

Zn -

1 7fi

1 a

2

0

8 7

3

er

1 9fi

6

r\ A .04

no

3

QQ

cn

2CK

1 CO

5 no

fi7Q

220

1 3

42

1 ft

1 OC

91

.17

2 a

8

0

4.1

1.15

1

1 a

^ 7^

90

277

1 3Q

S CO

.4

o

1 3

IqnU\J

l3fi

^K.

29

7 2

94. ^ITT. 1J.

5 a

 90U\J

1 9 E>

6

OC

97

27 c

1QO \-\j£t

^ ^4

9 9

0

0

77

1 97

ac\

27

6

1

22 0 £i£t.\J

5.4

.84

1.40

6.90

1 n^

280

200

^ ^8

.14

1 ^

1.4

11

OQC

301

1 a

19

K

137

9 9

70

94

an

.11

26

9 "'GOnr\

1 OQ

.07

2 10

13 f\

19 4

11

2 140

2 160

an

q

.7

0

91

3 7

1.11

0 K

C1

2

0

.42

297

120

4.5

93

1 R

1

10

Analysis by X-ray fluorescence.

Analysis by emission and optical spectroscopy.

An FeOtotal/MgO versus Cen/Ybn diagram (fig. 3.10) 

shows that most rocks of the Canyon Mountain Complex (Leeman and others, chap. 1, this volume) plot 

in a tholeiitic field (strong FeOtotal enrichment and 

flat REE pattern) near the Tonga-Kermadec field. 

Only the diorites (Leeman and others, chap. 1, this 

volume) show a LREE-enrichment trend. Plots of 

REE's (fig. 3.11), normalized against the chondrite 

values of Masuda and others (1973), show patterns 

that are characteristic of other rocks sampled from 

the Canyon Mountain Complex (for example, Leeman 

and others, chap. 1, this volume). Samples CM-79-2 

and CM-79-4 show a fractionated-tholeiitic pattern 

with a large negative Eu anomaly, indicating extensive plagioclase fractionation.

OTHER PLUTONIC ROCKS OF THE BAKER TERRANE

Two other intrusive rock bodies in the Baker terrane were sampled (tables 3.3-3.5). Both are from 

more easterly outcrops in the terrane and include diorite (sample BR-79-1) associated with the Burnt 

River Schist and trondhjemite (sample V-l-79) from 

serpentinite-gabbro-trondhjemite outcrops near 

Pleasant Valley, Oreg. (table 3.3). Zircons from sample BR-79-1 have slightly discordant 206pb/238u ages 

of 233 and 230 Ma; zircons from sample V-l-79 yielded concordant ages of 244 Ma (Walker, 1986, p. 80). 

Another trondhjemite intrusive body from the Greenhorn Mountains area yielded a U-Pb age of 243 Ma 

(Brooks, Ferns, Coward, and others, 1982; Walker, 

1986). All of these intrusive bodies are younger than
144 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

80

75

HZ

LLJ 

O

£70

a.

IX

LU 65 

£

z

l-'60

z

O

o

CN

O

50

45

FIGURE 3.9. FeOtota/MgO ratio versus SiO2 content for plutonic 

rocks of Canyon Mountain Complex (CM) and other plutonic 

rocks of Baker terrane. Many mafic plutonic rocks show tholeiitic affinities, whereas some siliceous plutons have decidedly calcalkaline characteristics. Dots, data from Leeman and others 

(chap. 1, this volume); squares, data from Thayer (1977); crosses, new data for the Canyon Mountain Complex (this chapter); 

triangles, data for other plutonic rocks of Baker terrane (this 

chapter). Compare with area (KJ) representing Jurassic to Cretaceous plutonic rocks of Blue Mountains region (see fig. 3.4). 

Calc-alkaline-tholeiitic compositional boundary from Miyashiro 

(1974).

23456 

FeOtota |/MgO RATIO

cc 4 

O

2 3 

O

CIV

I C

' )

0 } 23456789

Ce n/Yb n RATIO

FIGURE 3.10. FeOtotai/MgO versus CeIffbIi ratios for plutonic rocks of Baker terrane. Data field 

for Canyon Mountain Complex (CM) from Leeman and others (chap. 1, this volume). Squares, 

diorite of Canyon Mountain Complex (Leeman and others, chap. 1, this volume); open circles, 

Canyon Mountain Complex (this chapter); solid circles, other plutonic rocks (this chapter). Data 

fields for Tonga-Kermadec island arc lavas (Z), tholeiitic rocks of Okmok Caldera of Aleutian 

island arc (C), and Bald Mountain batholith of Blue Mountains region (BM; see fig. 3.5) are 

shown for comparison. Line, dashed and queried where uncertain, divides tholeiitic and calcalkaline fields, n, chondrite normalized.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 145

La Ce Sm Eu Gd Tb Yb Lu

RARE-EARTH ELEMENTS

EXPLANATION

Trondhjemite, Baker terrane 

Sample V-l-79 

Sample CM-79-4 

Sample CM-79-2

Diorite, Baker terrane 

Sample BR--79-1

Quartz diorite, Baker terrane 

Sample CM--79-3

FIGURE 3.11. Rare-earth element (REE) diagram for plutonic rocks 

of Baker terrane. Sample numbers correspond to those in tables 

3.3 through 3.5; for normalizing values used, see figure 3.6.

the isotopically dated rocks of the Canyon Mountain 

Complex. These isotopic ages indicate an Early and 

Middle Triassic phase of igneous activity in the Baker terrane.

The diorite and trondhjemite samples (BR-79-1 and 

V-l-79) show important differences from samples of 

the Canyon Mountain Complex on the basis of traceelement contents (table 3.5) and diagrams (figs. 3.10 

and 3.11). For example, both samples show LREE enrichment; sample V-l-79 plots in the Aleutian Okmok 

Caldera tholeiitic field and sample BR-79-1 plots in 

the Bald Mountain batholith calc-alkaline field on 

the FeOtotai/MgO versus Cen/Ybn plots (fig. 3.10). 

REE plots of the two samples (fig. 3.11) also show 

relative enrichment in LREE's compared to granitoids of the Canyon Mountain Complex; the HREE's 

are extremely depleted in sample BR-79-1 (Yb, about 

3 times chondrite; Cen/Ybn=6), which indicates a calcalkaline fractionation history. The geologic evolution 

of the Burnt River Schist, from which this sample 

was collected, is problematical; most igneous rocks 

from that unit are of volcanic-arc (magmatic-axis) origin (P.R. Hooper, oral commun., 1987) and therefore 

may be related to the igneous evolution of either the 

Wallowa terrane or the Olds Ferry terrane.

PETROGENESIS OF PLUTONIC ROCKS OF THE BAKER TERRANE

The diversity of plutonic rocks in the Baker terrane 

suggests more than one petrogenetic (and tectonic) environment. The plutonic rocks of the Canyon Mountain 

Complex definitely are different, both in age and composition, from many other plutonic rocks of the terrane.

The siliceous rocks of the Canyon Mountain Complex are classified as metamorphosed I-type granitoids (White and Chappell, 1983) mostly on the basis 

of their low K2O/Na2O ratios and Cr contents. An average analysis of I-type granitoids, however, shows 

higher concentrations of Sr, Ba, Rb, Th, U, Zr, and 

REE's (White and Chappell, 1983, p. 32) relative to 

granitoid plutonic rocks of the Canyon Mountain 

Complex. Melts that fractionated to yield siliceous 

plutonic rocks of the Canyon Mountain Complex had 

an igneous source and must have been derived from a 

very depleted mantle similar to mantle that produces 

plagiogranites under midocean spreading ridges and 

plagiogranites associated with many onland ophiolites (Coleman and Peterman, 1975; Coleman, 1977).

On the basis of chemical analyses of ultramafic and 

gabbroic rocks in the Canyon Mountain Complex, 

Mullen (1983) concluded that although most gabbros 

and so-called transitional rocks are tholeiitic, some 

calc-alkaline gabbros also occur, and that parent 

magma (or magmas) for the calc-alkaline gabbros 

originated by partial melting of a hydrous source. Also, she speculated that the origin of felsic plutons 

(plagiogranites) was by partial fusion of altered oceanic gabbro upon intrusion by the younger gabbro 

magma. Mullen (1983) related the origin of the Canyon Mountain Complex to early fore-arc magmatism 

generated by hydration and melting of depleted mantle adjacent to the leading edge of a subducting slab.

VOLCANIC ROCKS OF THE BAKER TERRANE

The upper Paleozoic and Triassic volcanic rocks of 

the Baker terrane are more diverse than those in any 

other terrane of the Blue Mountains region. Work by 

E.M. Bishop (Mullen, 1978, 1985; Bishop, chap. 5, 

this volume), for example, clearly pointed out the 

petrologic diversity of the volcanic rocks and led to 

the conclusion that the rocks have volcanic-arc, within-plate, and midocean-ridge petrotectonic affinities 

and are part of a complex fore-arc terrane. My data 

in part duplicate the work by Mullen (1985) and 

Ferns and Brooks (chap. 9, this volume) and provide 

additional analyses (tables 3.6-3.8).

In the Baker terrane, all volcanic rocks that were 

sampled for this study are metamorphosed to the 

greenschist or amphibolite facies and are deeply
146 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 3.6. Lithology, sample locality, and field description of Paleozoic and Triassic volcanic (flow) and hypabyssal (dike and

sill) rocks of the Baker terrane

[Rocks arranged by increasing SiO2 content to correspond to chemical analyses in tables 3.7 and 3.8. Do. and do., ditto]

Sample Rock type Sample locality Field description and remarks

V-3-82 Basalt

V-2-82 - do 

B133A1 - do 

BB135B - do ~ 

V-l-82 - do 

VB-31 Amphibolite

VB-30 - do 

BM8001 Basaltic andesite -----

Secondary pinnacle on Vinegar Hill; 

elevation 2,301 m (7,550 ft); NE1/4 

SE1/4NW1/4 sec. 12, T. 9 S., R. 35 

E.; Bates quadrangle, Oreg.

Main pinnacle of Vinegar Hill; elevation 

2,335 m (7,660 ft); NE1/4SE1/4NW1/4 

sec. 12, T. 9 S., R. 35 E.; Bates quadrangle, 

Oreg.

North side of State Highway 220; elevation 

1,189 m (3,900 ft); SE1/4NW1/4 sec. 29, T. 

10 S., R. 39 E.; Blue Canyon quadrangle, 

Oreg.

North side of State Highway 220; elevation 

1,229 m (4,030 ft); SE1/4SE1/4 sec. 24, T. 

10 S., R. 38 E.; Blue Canyon quadrangle, 

Oreg.

Vinegar Hill; elevation 2,274 m (7,460 ft); 

NE1/4NW1/4 sec. 12, T. 9 S., R. 35 E.; 

Bates quadrangle, Oreg.

Above and west of road along Lodge Creek; 

elevation 1,219 m (4,000 ft); SW1/4NE1/4 

sec. 20, T. 10 S., R. 41 E.; Encina 

quadrangle, Oreg.

Above and west of road along Lodge Creek; 

elevation 1,204 m (3,950 ft); SW1/4NE1/4 

sec. 20, T. 10 S., R. 41 E.; Encina 

quadrangle, Oreg.

South side of road between Richland, 

Oreg., and Snake River, north of cattle 

guard; elevation 960 m (3,150 ft); 

SW1/4SE1/4 sec. 29, T. 11 S., R. 46 E.; 

Mineral quadrangle, Oreg.

Flow rock.

Pillow lava.

Flow rock.

Flow rock, probably pillow 

lava.

Sill or flow within chert-clast 

sandstone unit.

Probable metamorphosed flow 

cutting breccia, chert, and 

argillite.

Do.

Pillow lava.

weathered, as shown by high water contents (table 

3.7). Several are pillow lavas (table 3.6). Where the 

flows are intercalated with metasedimentary rocks 

they can be traced laterally for only a few tens of 

meters; most exposures are bounded by faults. 

Greenstone derived from volcanic rocks occurs in 

places as blocks or knockers in a serpentinite-matrix 

melange.

A (Na2O+K2O)-FeOtotal-MgO diagram (fig. 3.12) 

shows the characteristic Fe-enrichment and Na- and 

K-depletion of most Baker terrane volcanic rocks. No 

coherent fields are apparent on the diagram, in part 

because of the suspected mobility of the alkali elements and the small number of samples. However,

the rocks plot in the Fe-enriched (tholeiitic) and MgFe-enriched (alkalic) fields. Plots on an FeOtotal/MgO 

ratio versus SiO2 content diagram (fig. 3.13) also indicate that most are Fe-enriched. High TiC>2 contents 

are characteristic of alkalic rocks (see samples V-l-82 

and BM8001 in table 3.7), and the low TiO2 concentrations in samples V-2-82, B133A1, and BB135B 

suggest volcanic-arc affinities.

A REE diagram (fig. 3.14) shows a wide range of 

trends similar to those discussed by Mullen (1985); it 

strengthens her conclusion that the basaltic rocks 

crystallized in different tectonic settings. Samples V1-82 and BM8001 have alkalic basalt trends (high 

LREE abundances with HREE contents about 10
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 147 

TABLE 3.7. Major- and minor-element oxides of Paleozoic and Triassic volcanic (flow) and hypabyssal (dike and sill) rocks of the Baker terrane

[Oxides in weight percent, normalized to 100 percent; prenormalization volatiles contents included to indicate amount of alteration. , not determined]

S*

Si

A 

Fi 

F< 

M 

M 

C, 

N 

K

P;

H 

H

rji

ample 

i02 -- --

1O __________

& o

,2O3 

:>O .................

:gO------

;nO 

qO __ _ _

a2Q... ............ 2 o ____

O - __.___

u O

2Q+. ..............

20---- 

O~

FeOtota/MgOV q QO

46.07 

1.78 

15.73 

4.56 

6.67 

9.09 

.20 

10.52 

3.50 

1.60 

.28

4.10 

.27 

4.79 

1.19

V-2-82

47.76 

1.03 

15.41 

1.99 

10.77 

8.64 

.21 

11.29 

2.37 

.36 

.17

2.23 

.17 

.08 

1.45

B133A1 BB135B V-l-82 VB-31 VB-30 BM8001

50.05 50.16 

1.26 1.39 

16.02 17.13 

3.06 2.39 

5.28 6.95 

6.48 6.64 

.15 .15 

14.24 11.66 

3.16 3.10 

.17 .25 

.13 .18

2.13 2 

.12 

1.51 

1.24 1

.35 

.38 

.17 

.37

50.27 

2.45 

15.21 

1.61 

9.45 

8.38 

.21 

9.11 

2.57 

.35 

.39

2.25 

.38 

.01 

1.30

52.03 52.30 52, 

1.71 1.68 2, 

15.43 14.85 18, 

2.17 3, 

8.67 11.22 9, 

7.14 7.24 2, 

.32 .19 

8.33 7.90 2, 

3.75 4.11 5, 

.34 .37 1, 

.11 .14

3.13 

.34 

.20 

1.49

3, 

1.55 4,

,71

,53 

.76 

.39 

.30 

.87 

,29 

.67 

.14 

.65 

,69

.98 

.50 

.02 

.30

TABLE 3.8. Trace elements of selected Paleozoic and Triassic volcanic (flow) and hypabyssal (dike and sill) rocks of the Baker 

terrane

[Results in parts per million. Analysis by neutron activation unless otherwise noted. , not determined or below detection limit or discarded because of high coefficient 

of variation]

Sample 

Rb1 

Sr 1 

Ba1 

TVU

u .... 

La 

Ce 

Sm 

/"i j

rrvL

Tin 

Yb 

Lu 

Y1 

TJf

Ta

Nb1 -----

Ni1 ~

Co 

fr

Sc - 

7n ___ - ......

V-3-82

21 

188 

99 

.96 

.59 

10.4 

21.6 

4.0 

1.31

.70

2.41 

.37 

26 

133 

2.89 

.72 

16

66 

679 

33

IfU

V-2-82

10 

172 

189 

.50 

.45 

3.54 

8.0

.89 

.70

3.42 

.52 

30 

60 

1.54

9 

96 

57 

320 

62

15U

B133A1

2 

223 

32

.5

7 

15

1.04 

3.2 

.71

2.5 

 .41 

23 

95 

2.3 

.48 

7 

73 

37 

310 

35

79

BB135B

4 

162 

56 

.9

11 

20

1.17 

3.9 

.65

2.1 

.32 

22 

97 

2.2 

.84 

13

36 

121 

34

7R

V-l-82

13 

373 

677 

2.61 

.73 

22.5 

44.7 

6.3 

1.7 

5.5 

.85 

.42 

2.26 

.33 

27 

212 

4.74 

1.75 

25

47 

295 

21 

inn

VB-31

4 

142 

126

4 

12 

3.4 

1.23 

4.6 

.99 

.53 

4.3 

.60 

34 

82 

2.9

26 

39 

165 

37

130

BM8001

38 

125 

646 

1.7 

.4 

20 

39 

5.6 

2.05

1.22

4.1 

.66 

44 

165 

4.0 

1.58 

23

49 

56 

36

937

Analysis by X-ray fluorescence.
148 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Fedtotal

Na2 0+K2 0 50 MgO

FIGURE 3.12. (Na2O+K2O)-FeOtotal-MgO (AFM) diagram for volcanic (flow) rocks of Baker terrane, showing iron enrichment in 

most rocks. Dots, data from Mullen (1978, 1985); crosses, data 

from this chapter. Curved line (from Kuno, 1968) approximately 

separates calc-alkaline field from tholeiitic field. This diagram 

generally is not useful for distinguishing alkalic from tholeiitic 

basaltic rocks.

70

in 65 

O 

cc iii

0-

(-

X

CD 60

Z

Z 55

O 

O

CN

O 50 

C/3

45 J I

0 1 23456 

FeOtota ,/MgO RATIO

FIGURE 3.13. FeOtotal/MgO ratio versus SiO2 content for volcanic 

rocks of Baker terrane. Dots, data from E.M. Bishop (Mullen, 

1978, 1985; Bishop, chap. 5, this volume); crosses, new data 

(this chapter). Data fields A through D represent volcanic rocks 

of Aleutian island arc and correspond to those in figure 3.4. 

Note strong Fe enrichment in most Baker terrane volcanic 

rocks. Calc-alkaline-tholeiitic compositional boundary from Miyashiro (1974).

times chondrite values). The other samples have either flat tholeiitic or up ward-sloping transitional (between tholeiitic and calc-alkaline) REE trends.

Discrimination diagrams (Pearce and Cann, 1973) 

using Y, Ti, and Zr data (figs. 3.15 and 3.16) also show 

the chemical diversity of basaltic rocks in the Baker 

terrane. Two rocks (samples V-l-28 and BM8001) are 

alkalic on the basis of high contents of Ti and closely 

associated elements (Zr, Hf, and Ta), relatively high 

quantities of alkali and alkaline-earth elements, high 

concentrations of Th, and enriched-LREE patterns. In 

addition, two of Mullen's (1985) samples of basaltic 

rocks of the Baker terrane (BRS and V27) are alkalic. 

Three basalts (samples V-2-82, B133A1, and VB-31) 

are tholeiitic (midocean-ridge basalt or island-arc basalt affinities) on the basis of their relatively flat and

La Ce Sm Eu Gd Tb Yb Lu

RARE-EARTH ELEMENTS

EXPLANATION

Basalt, Baker terrane 

Sample V-l-82 

Sample BB135B 

Sample V-3-82 

Sample B133A1 

Sample V-2-82

Basaltic andesite, Baker terrane 

Sample BM8001

Amphibolite, Baker terrane 

Sample VB-31

Tholeiitic rocks of Ata Island, 

Tonga island arc

FIGURE 3.14. Rare-earth element (REE) diagram for volcanic 

rocks of Baker terrane. Sample numbers correspond to those in 

tables 3.6 to 3.8; for normalizing values used, see figure 3.6. Ata 

Island field (Vallier and others, 1985) shows island-arc tholeiite 

trend for mafic lavas. Samples V-l-82 and BM8001 show alkalic 

basalt trends; samples V-3-82, V-2-82, B133A1, BB135B, and 

VB- 31 are tholeiitic. Some light rare-earth element enrichment 

may have occurred in samples V-3-82 and BB135B. Nd, Sm, and 

Gd data are missing for some samples either because element 

was not analyzed, there was a high coefficient of variation, or 

content was below detection limit of instrument.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 149

(or) depleted-LREE patterns (fig. 3.14), relatively low 

Ti contents, and plotted positions (within "OFB+LKT 

+CAB" field) on the discrimination diagrams (figs. 

3.15 and 3.16). Sample B133A1, especially, has very 

low amounts of Rb and Ba, which are compatible with 

a midocean-ridge basalt setting, although ion mobility 

during metamorphism may have affected some elemental concentrations. The relatively high quantities 

of Ti in these rocks are more similar to midoceanridge basalt than to island-arc basalt. Ocean island 

(seamount) tholeiites are difficult to distinguish, but 

sample V-3-82 is the best candidate, although it could 

represent a high-K island-arc transitional basalt. The 

samples from Vinegar Hill (V-l-82, V-2-82, and V-3-82; 

table 3.6) in the Greenhorn Mountains, selected from 

knockers within serpentinite-matrix melange, are 

both alkalic and tholeiitic; their great compositional 

diversity and present positions must represent a mixing of rocks during melange formation, probably in 

the fore-arc region of the Blue Mountains island arc 

(Mullen, 1985; Bishop, chap. 5, this volume).

The petrogenesis of volcanic rocks in the Baker terrane has been discussed in more detail by E.M. Bishop (Mullen, 1985; Bishop, chap. 5, this volume). 

According to her interpretation, all of the major oceanic tectonic settings (midocean ridge, seamount or 

oceanic island, and volcanic arc) are represented. The 

Baker terrane apparently has been involved in several tectonic episodes, not only during the convergence 

of oceanic plates, but also during, and subsequent to, 

the accretion of the Blue Mountains island arc to 

North America.

Ti/100

Zr 50 Yx3

FIGURE 3.15. (Zr)-(Ti/100)-(Yx3) discrimination diagram (Pearce 

and Cann, 1973). CAB, calc-alkaline basalt; LKT, low-potassium 

tholeiite; OFB, ocean-floor basalt; and WPB, within-plate basalt. 

Dots represent lavas of Baker terrane.

LATE PALEOZOIC AND TRIASSIC IGNEOUS ROCKS 

OF THE WALLOWA TERRANE

The lithologies and chemistries of igneous rocks of 

the Wallowa terrane suggest that it is part of the late 

Paleozoic (mainly Early Permian) and Triassic magmatic axis (volcanic front) of the Blue Mountains island arc. The terrane is very small (fig. 3.8) when 

compared with modern arcs. The magmatic axes of 

modern island arcs have volcanoes (counting submarine edifices) spaced at distances of 50 to 70 km (Aleutian island arc) or 20 to 30 km (Tonga island arc). The 

Aleutian and Tonga arcs are long and wide; volcanoes 

along the Aleutian Islands, for example, extend for 

about 1,200 km, and the entire arc, from the trench 

inner wall to the abyssal depths in the back-arc region, 

is generally about 150 to 200 km wide. If the Alaska 

Peninsula is included in the Aleutian arc, then the 

length of the magmatic axis is greater than 2,000 km. 

The Wallowa terrane, as mapped, is about 150 by 200 

km in size. As such, that terrane alone would have 

supported only a few volcanoes at any one time during 

its geologic evolution. In the Middle (Ladinian) and 

Late (Karnian) Triassic, for example, the Wallowa terrane would have supported no more than ten (analogy 

to the Tonga arc), and probably five or fewer (analogy 

with the Aleutian arc), volcanoes. Therefore, it is evident that a large part (probably at least 50 percent 

and possibly as much as 70 percent) of the Wallowa 

terrane is missing from the immediate area. The missing parts of the Wallowa terrane were either translated northward to become part of western Canada and

1.5

1.0

OFB + LKT 

+ CAB

0 50 100 150 200 

Zr CONTENT, IN PARTS PER MILLION

FIGURE 3.16. Ti versus Zr content (Pearce and Cann, 1973), showing ocean-floor affinity of volcanic rocks of Baker terrane. CAB, 

calc-alkaline basalt; LKT, low-potassium tholeiite; and OFB, 

ocean-floor basalt. High Ti contents in upper part of OFB field 

suggest seamount alkalic basalt.
150 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

southeastern Alaska or were overridden during tectonic convergence. Processes of tectonic convergence were 

(1) subduction during the times when the Wallowa 

terrane was part of an intraoceanic island arc and (2) 

accretion (overthrusting and wrench faulting) when 

the North American craton was in direct contact with 

the terranes.

Although most of the shallow crustal rocks of the 

Wallowa terrane are composed of Permian and Triassic volcano-root plutons, volcanic flow rocks, and volcaniclastic rocks, there is an overlying carapace of 

Upper Triassic limestone and Upper Triassic to Middle Jurassic siliciclastic rocks (Vallier, 1977). Furthermore, Jurassic to Cretaceous plutons intrude the 

older rocks. The Permian and Triassic volcanic strata 

comprise the Seven Devils Group (Vallier, 1977). A 

composite geologic column for the Wallowa terrane 

and local variants across that terrane are shown 

schematically in figure 3.17. This schematic diagram, 

in contrast to a stratigraphic column by Vallier 

(1977), shows the Windy Ridge Formation as a lithofacies equivalent of the lower part of the Hunsaker 

Creek Formation and shows the Doyle Creek Formation as a lithofacies equivalent of the upper part of 

the Wild Sheep Creek Formation.

PLUTONIC AND ASSOCIATED HYPABYSSAL ROCKS OF THE 

WALLOWA TERRANE

Plutonic rocks of the Wallowa terrane were emplaced mostly in late Paleozoic (mostly Permian), 

Middle and Late Triassic, and Jurassic to Cretaceous 

time intervals. Only the late Paleozoic and Triassic 

plutonic rocks are discussed in this section. A total of 

39 samples of plutonic and hypabyssal rocks from the 

Wallowa terrane were analyzed (table 3.9), all 39 for 

major- and minor-element oxides (table 3.10), and 36 

for trace elements (table 3.11). These igneous rocks of 

the Wallowa terrane are separated into specific rock 

masses and (or) grouped by geographic area (table 

3.9; figs. 3.1, 3.2, 3.3): the (informal) Sparta complex; 

Oxbow Complex; Cougar Creek Complex; Wolf CreekDeep Creek and Imnaha plutonic rocks; Sheep Creek 

plutonic rocks; plutonic rocks of the southern Seven 

Devils Mountains; plutonic rocks of the Salmon River 

canyon; plutonic rocks associated with the Blue 

Mountains anticline; and Wildhorse River dike.

The compositional range of coarse-grained plutonic 

rocks of late Paleozoic (mostly Permian) and Triassic 

age includes norite and gabbro, meladiorite and diorite, and tonalite (quartz diorite and trondhjemite). 

K-rich plutonic rocks (granodiorite and granite) of Permian and Triassic age are extremely rare. In addition, 

dikes of quartz keratophyre and keratophyre porphyries, diabase, and aplite were collected and analyzed 

where field relations indicated that they are offshoots 

of larger plutonic bodies. One orthogneiss associated 

with the Blue Mountains anticline was analyzed.

Na2O-rich siliceous plutonic bodies are common in 

the Wallowa terrane. They are characterized by 

abundant blue-colored bipyramidal quartz phenocrysts, high contents of albite (and hence Na2O), and 

low amounts of K^O-bearing minerals. Quartz keratophyre dikes are common in the Oxbow Complex and 

the (informal) Sparta complex and generally are 

closely associated with mafic dikes. In the Oxbow 

Complex, the felsic and mafic dikes occur together in 

sheetlike sequences.

The FeOtotai/MgO ratios versus SiO2 contents (figs. 

3. ISA, 3.1&B), FeOtotal/MgO versus Cen/Ybn ratios 

(fig. 3.19), and REE diagrams (figs. 3.20A-3.20£) 

show the broad compositional diversity of plutonic 

rocks in the Wallowa terrane. Tholeiitic, transitional, 

and calc-alkaline late Paleozoic and Triassic plutonic 

rocks all occur within the terrane.

SPARTA COMPLEX

The (informal) Sparta complex was thoroughly studied by Prostka (1962), Almy (1977), Phelps (1978, 

1979), and Phelps and Ave Lallemant (1980) and will 

not be discussed in detail here. Chemical analyses given in tables 3.10 and 3.11 complement the previously 

published work. Phelps (1979) presented analyses of 

34 rocks from the Sparta complex. He separated out 

the quartz diorite unit, which, according to Prostka 

(1962, 1963) is the youngest plutonic unit from the 

trondhjemite unit (older plutonic sequence), and speculated that they might be co-magmatic. On the basis 

of recent isotopic age data (Walker, 1986 and chap. 6, 

this volume), however, it is apparent that the two plutonic masses represent separate events.

Walker (1986 and chap. 6, this volume) reviewed 

previous isotopic age determinations for the younger

FIGURE 3.17. Schematic geologic columns for Wallowa terrane. A, Four localities (see fig. 3.1) within Wallowa terrane. Stratigraphy interpreted mostly from Nolf (1966), 

Vallier (1977), Folio (1994), and White and Vallier (1994). 

Lines joining columns indicate time correlations, queried 

where uncertain. Queries below "fcws in right-hand columns indicate that the stratigraphic unit beneath the 

Wild Sheep Creek Formation is unknown. B, Composite 

column for Wallowa terrane. Circled "1" labels area of undivided Permian and Triassic Cougar Creek Complex of 

Vallier (1968), Oxbow Complex of Vallier (1967a, b), and 

(informal) Sparta complex of Phelps (1979).
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 151

Homestead 

region, Oregon

Northern Wallowa 

Mountains, Oregon

Pittsburg Landing, 

Idaho and Oregon

Grande Ronde River, 

Washington and Idaho

Tcr

 fide

Oxbowx x 

x Complex x

X X K X X 

X X X X

	EXPLANATION

Tcr Columbia River Basalt Group (Miocene)

Jch Coon Hollow Formation (Upper and Middle Jurassic)

J~fih Hurwal Formation (Lower Jurassic and Upper Triassic)

"fimb Martin Bridge Limestone (Upper Triassic)

"fils "Lower Sedimentary Series" of Prostka (1962) (Upper

	Triassic) 

"fide Doyle Creek Formation (Upper and Middle

	Triassic) Includes: 

"fidck Kurry unit (Upper Triassic)

"fiws Wild Sheep Creek Formation (Upper and Middle

	Triassic)

"RPcc Clover Creek Greenstone (Triassic and Permian)

Phc Hunsaker Creek Formation (Lower Permian)

Pwr Windy Ridge Formation (Lower Permian)

11111 [ [ I Volcanic flows

[ " I Sandstone, siltstone, and shale

pcPl Conglomerate

I ! . '. | Sandstone

[ i ' i 'I Limestone

|A A| Breccia

I:' A'. I Breccia and sandstone

PCX] Pillow lava

~~ ^" Interfingering contact

S Intrusive contact

 ^ ^ Unconformity
152 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 3.9. Lithology, sample locality, and field description of late Paleozoic and Triassic plutonic and hypabyssal (dike) rocks of the

Wallowa terrane

[Rocks within each group arranged by increasing SiO2 content to correspond to chemical analyses in tables 3.10 and 3.11. Do. and do., ditto]

Sample Rock type Sample locality Field description and remarks

Sparta complex

SP-79-1 Trondhjemite

V-8-79

V-100-79 Quartz keratophyre --

Bishop Springs, along State Highway 86; 

elevation 777 m (2,550 ft); NW1/4NW1/4 

sec. 6, T. 9 S., R. 44 E.; Sparta 

quadrangle; Oreg.

Ridge west of road to Sparta; elevation 

1,036 m (3,400 ft); SE1/4NW1/4 sec. 23, 

T. 8 S., R. 45 E.; Sparta quadrangle; 

Oreg.

Ridge east of Clover Creek; elevation 982 m 

(3,220 ft); NW1/4NW1/4 sec. 1, T. 8 S., R. 

42 W.; Sawtooth Ridge quadrangle; Oreg.

Youngest pluton (stock) that cuts 

older suite (gabbro) of the Sparta 

complex. Late Triassic in age 

(Walker, 1986).

Dike rock from sequence of 

trondhjemite and diabasedikes, 

probably associated with the older 

suite (gabbro) of the Sparta 

complex.

Dike rock from sequence of quartz 

keratophyre, trondhjemite, and 

diabase dikes, probably associated 

with older suite (gabbro).

Oxbow Complex

V-4-73 Gabbro North side of small tributary near prospect; Small intrusion; several textures,

elevation 1,000 m (3,280 ft); SE1/4NW1/4 mostly massive, 

sec. 2, T. 19 N., R. 4 W.; Homestead 

quadrangle; Idaho.

V-67-33 Diabase porphyry Along old road that parallels Snake River; Mafic dike crosscut by trondhjemite.

elevation 537 m (1,760 ft); NW1/4SW1/4 Some fluxion structure, 

sec. 3, T. 7 S., R. 48 E.; Copperfield 

quadrangle, Oreg.

VC-184 Trondhjemite Below major switchback along old road Cataclastic rock from wide zone of

west of Indian Creek; elevation 774 m dikes. 

(2,540 ft); NW1/4SE1/4 sec. 9, T. 19 N., 

R. 4 W.; Copperfield quadrangle; Idaho.

VC-45 do North side of Scorpion Creek; elevation 817 Dike cutting gabbroic body.

m (2,600 ft); NW1/4SE1/4 sec. 21, T. 19 

N., R. 4 W.; Copperfield quadrangle; 

Idaho.

LC-48 do North side of Scorpion Creek; elevation 677 Dike rock from sequence of gabbro,

m (2,200 ft); SE1/4NW1/4 sec. 21, T. 19 diabase, and trondhjemite dikes. 

N., R. 4 W.; Copperfield quadrangle; 

Idaho.

CO-79-1 do West side of Oxbow dam; elevation 560 m Dike rock in sequence of diabase and

(1,840 ft); SE1/4NW1/4 sec. 10, T. 7 S., trondhjemite dikes; Late Permian

R. 49 E.; Copperfield quadrangle; Oreg. U-Pb age (Walker, 1986).

Cougar Creek Complex

T-68-14 Gabbro - South edge of Cougar Creek complex along Small pluton that cuts metamorphosed

trail that parallels Snake River; and sheared sedimentary and 

elevation 396 m (1,300 ft); SE1/4 NW1/4 volcanic rocks, 

sec. 36, T. 26 N., R. 1 W.; Kernan Point 

quadrangle; Idaho.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 153

TABLE 3.9. Lithology, sample locality, and field description of late Paleozoic and Triassic plutonic and hypabyssal (dike) rocks of the

Wallowa terrane Continued

[Rocks within each group arranged by increasing Si02 content to correspond to chemical analyses in tables 3.10 and 3.11. Do. and do., ditto]

Sample Rock type Sample locality Field description and remarks

Cougar Creek Complex Continued

V-61-79

V-5-73

T-68-26

V-12-73

V-6-73

V-37-79

KC-79-1

T-68-11

T-68-23

V-66-79

V-35-79

.... do 

Diorite -

.... do 

Quartz dioriteTrondhjemite

.... do 

.... do - - 

South edge of Cougar Creek Complex along 

trail that parallels Snake River about 60 

m south of sample locality T-68-14; 

elevation 396 m (1,300 ft).

North side of Klopton Creek along ridge; 

elevation 1,479 m (4,850 ft); SE1/4SE1/4 

sec. 26, T. 27 N., R. 1 W.; Grave Point 

quadrangle; Idaho.

Along trail that parallels Snake River 

about 400 m south of Corral Creek; 

elevation 369 m (1,210 ft); NW1/4SE1/4 

sec. 9, T. 26 N., R. 1 W.; Kirkwood Creek 

quadrangle; Idaho.

Along road between White Bird and 

Pittsburg Landing, Idaho; elevation 

1,037m (3,400 ft); extreme southwest 

corner of SW1/4SE1/4 sec. 24, T. 27 N., 

R. 1 W.; Grave Point quadrangle; Idaho.

North side of Klopton Creek along ridge; 

elevation 1,256 m (4,120 ft); SE1/4SW1/4 

sec. 26, T. 27 N., R. 1 W.; Grave Point 

quadrangle; Idaho.

North of White Bird Creek-Salmon River 

confluence; elevation 550 m (1,800 ft); 

NE1/4SE1/4 sec. 34, T. 29 N., R. 1 E.; 

White Bird quadrangle; Idaho.

Ridge south of Corral Creek; elevation 885 

m (2,900 ft); NE1/4SW1/4NE1/4 sec. 16, 

T. 26 N., R. 1 W.; Kirkwood Creek 

quadrangle; Idaho.

Along trail that parallels Snake River 

about 400 m north of Cougar Creek; 

elevation 397 m (1,300 ft); T. 1 N., R. 50 

E. (unsurveyed); Kirkwood Creek 

quadrangle; Oreg.

Along trail that parallels Snake River; 

elevation 390 m (1,280 ft); NW1/4 NE1/4 

sec. 36, T. 26 N., R. 2 W.; Kernan Point 

quadrangle; Idaho.

Along trail that parallels Snake River; 

elevation 385 m (1,260 ft); NE1/4 sec. 30, 

T.26 N., R. 1 W.; Kirkwood Creek 

quadrangle; Idaho.

North of White Bird Creek-Salmon River 

confluence; elevation 543 m (1,780 ft); 

NE1/4SE1/4 sec. 34, T. 29 N., R. 1 E.; 

White Bird quadrangle; Idaho.

Do.

Massive outcrops of pluton cut by 

Klopton Creek fault. Forms the 

northern rock unit of the Cougar 

Creek Complex.

Dike rock from sequence of mafic dikes 

striking about N. 65° E.

Greatly altered and sheared; prehnite 

and calcite zones indicate extreme 

calcium mobilization.

Associated with gabbro sampled at 

same locality as V-5-73. Minor part 

of mostly gabbroic pluton.

From series of dikes and small stocks 

(trondhjemite, quartz diorite, and 

diabase).

Pluton that cuts across mafic sequence. 

Probably among youngest rocks in 

Cougar Creek Complex. Permian UPb age(Walker, 1986).

Dike rock from sequence of 

trondhjemite and mafic dikes.

Dike rock from sequence of dikes that 

range from gabbro to trondhjemite in 

composition.

Sheared, partly mylonitized dike rock 

from sequence of dikes striking 

about N. 50° E.

Dike rock from series of dikes and 

small stocks (trondhjemite, quartz 

diorite, and diabase).

Table continued on next page.
154 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 3.9. Lithology, sample locality, and field description of late Paleozoic and Triassic plutonic and hypabyssal (dike) rocks of the

Wallowa terrane Continued

[Rocks within each group arranged by increasing SiO2 content to correspond to chemical analyses in tables 3.10 and 3.11. Do. and do., ditto]

Sample Rock type Sample locality Field description and remarks

Cougar Creek Complex Continued

TP-13 Diorite Pittsburg Landing area; elevation 565 m Clast typical of well-rounded igneous

(1,850 ft); SW1/4SE1/4 sec. 28, T. 27 N., rock clasts in conglomerate of the 

R. 1 W.; Grave Point quadrangle; Idaho. Coon Hollow Formation; all 29

igneous rock clasts collected at this 

locality apparently were eroded from 

the nearby Cougar Creek Complex.

TP-15 Quartz diorite - do -- Do.

TP-18 do ------ - do - - -- ------ Do.

Wolf Creek-Deep Creek and Imnaha plutonic rocks

V-70-79 Gabbro - Near mouth of Wolf Creek; elevation 320 m Part of basement intrusive complex.

(1,050 ft); SE1/4 sec. 9, T. 28 N., R. 2 W.; Most outcrops are massive; sampled

Wolf Creek quadrangle; Idaho. gabbro shows rare rough layering.

DR-SR-79-1 Trondhjemite Along Snake River near confluence with From outcrops that are mostly quartz

Imnaha River 100 m upstream from "Mile diorite and diorite in composition. 

192" and near old mine; elevation 305 m Associated with rugged outcrops of 

(1,000 ft); unsurveyed part of canyon; gabbro upstream. 

Deadhorse Ridge quadrangle; Oreg.

V-72-79 Aplite - Near mouth of Deep Creek; elevation 310 Dike rock from abundant mostly

m (1,020 ft); near south boundary of T. 4 vertical white to buff dikes that cut 

N., R. 49 E. (unsurveyed); Cactus gabbro. 

Mountain quadrangle; Oreg.

Sheep Creek pluton

T-68-2 Meladiorite

HD-79-1 Quartz dioriteT-68-3 - do 

T-68-1 Trondhjemite

North side of Bills Creek in Snake River 

Canyon; elevation 442 m (1,450 ft); 

NE1/4NW1/4 sec. 16, T. 24 N., R. 2 W.; 

He Devil quadrangle; Idaho.

North side of Sheep Creek near trail; 

elevation 500 m (1,640 ft); NE1/4SE1/4 

sec. 35, T. 25 N, R. 2 W.; He Devil 

quadrangle; Idaho.

North side of Sheep Creek near trail 200 m 

downstream from sample site HD-79-1; 

elevation 482 m (1,580 ft); NE1/4SE1/4 

sec. 35, T. 25 N., R. 2 W.; He Devil 

quadrangle; Idaho.

North side of Bills Creek; elevation 450 m 

(1,480 ft); NE1/4NW1/4 sec. 16, T. 24 N., 

R. 2 W.; He Devil quadrangle; Idaho.

Mafic part of Sheep Creek pluton. 

Contact zone with country rock is 

sheared.

Permian U-Pb age (Walker, 1986). 

Light gray rocks contain blue quartz. 

Coarse grained.

Probably same age as dated sample 

HD-79-1. Similar texture and 

mineralogy.

Silicic part of Sheep Creek pluton at 

Bills Creek. Coarse grained.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 155

TABLE 3.9. Lithology, sample locality, and field description of late Paleozoic and Triassic plutonic and hypabyssal (dike) rocks of the

Wallowa terrane Continued

[Rocks within each group arranged by increasing SiO2 content to correspond to chemical analyses in tables 3.10 and 3.11. Do. and do., ditto]

Sample Rock type Sample locality Field description and 

remarks

Plutons of southern Seven Devils Mountains

V-2-84

V-4-84

V-3-84

V-l-84

Gabbro

Diorite -

Quartz dioriteAlong trail near Joe's Gap; elevation 2,440 

m (8,000 ft); SW1/4NW1/4 sec. 4, T. 21 

N., R. 2 W.; Cuprum quadrangle; Idaho.

Along trail near Black Lake Fork; elevation 

1,982 m (6,500 ft); NW1/4NE1/4 sec. 2, T. 

21 N., R. 2 W.; Cuprum quadrangle; 

Idaho.

Along trail east of Joe's Gap; elevation 

2,395 m (7,850 ft); extreme northwest 

corner NW1/4NW1/4SW1/4 sec. 3, T. 21 

N., R. 2 W.; Cuprum quadrangle; Idaho.

Along trail northwest of Joe's Gap; 

elevation 2,250 m (7,380 ft); NE1/4NE1/4 

sec. 5, T. 21 N., R. 2 W.; Cuprum 

quadrangle; Idaho.

Permian(?) or Early Triassic(?) White 

Mountain pluton of White (1968). 

Probably related to Cougar Creek 

and Oxbow Complexes.

Pactolian pluton of White (1968). 

Probably Permian or Early Triassic 

in age and related to Cougar Creek 

and Oxbow Complexes.

Horse Pasture pluton of White (1968). 

Probably Permian or Early Triassic 

in age and related to Cougar Creek 

and Oxbow Complexes.

Big Lake pluton of White (1968). 

Probably Permian or Early Triassic 

in age and related to Cougar Creek 

and Oxbow Complexes.

Salmon River canyon plutonic rocks

V-39-79

V-57-79

Diorite -

Trondhjemite

North side of Salmon River; elevation 780 

m (2,560 ft); SW1/4NW1/4 sec. 3, T. 31 

N., R. 2 W.; Boles quadrangle; Idaho.

In Box Canyon along U.S. Highway 95 and 

the Salmon River; elevation 512 m (1,680 

ft); SW1/4SE1/4 sec. 1, T. 26 N., R. 1 E.; 

Lucile quadrangle; Idaho.

Relatively massive diorite and quartz 

diorite pluton; altered and 

mineralized, particularly along 

shear zones.

Dike rock from sequence of silicic and 

mafic dikes. Permian U-Pb age 

(Walker, 1986).

Plutons within the Blue Mountains anticline

V-20-79 Hornblende gneiss 

V-17-79 Trondhjemite

Along road that parallels Stanley Creek; 

elevation 1,073 m (3,520 ft); NE1/4NE1/4 

sec. 32, T. 25 N., R. 32 E.; Granite 

Meadows quadrangle; Oreg.

At major turn along road that parallels 

Pearson Creek; elevation 1,073 m (3,520 

ft); SE1/4NE1/4NE1/4 sec. 9, T. 3 S., R. 

33 E.; Tamarack Gulch quadrangle; 

Oreg.

Within Blue Mountains anticline; 

probably part of metagabbro unit 

mapped by Trauba (1975).

Small pluton that cuts volcanic 

sequence (silicic and mafic flows and 

volcaniclastic rocks).

Table continued on next page.
156 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 3.9. Lithology, sample locality, and field description of late Paleozoic and Triassic plutonic and hypabyssal (dike) rocks of the

Wallowa terrane Continued

[Rocks within each group arranged by increasing SiO2 content to correspond to chemical analyses in tables 3.10 and 3.11. Do. and do., ditto]

Sample Rock type Sample locality Field description and 

remarks

Wildhorse River dike

V-5-84 Trondhjemite North side of bridge that crosses Wildhorse Dike rock along major fault separating

River (northwest of Cuddy Mountains); Wallowa and Baker terranes. 

elevation 738 m (2,420 ft); 

SE1/4NE1/4NE1/4 sec. 33, T. 18 N., R/ 4 

W.; Brownlee quadrangle, Idaho.

TABLE 3.10. Major- and minor-element oxides of late Paleozoic and

[Oxides in weight percent, normalized to 100 percent; prenormalization volatile contents included

Sparta complex Oxbow Complex

Si02 - - 

Tif^

AloOo 

£t O "C1 ^ f~\ Fe2O3 - 

FeO 

MgO 

MnO - 

CaO ----- -

Na2O 

v r\

P2Ur\5 

H O+

H20--~-~ 

r*f\

FeOtotal/MgO -----

S-79-1

70.07 

.59 

14.40 

2.36 

2.48 

1.22 

.06 

3.89 

3.56 

1.25 

.12

0.84 

.09 

.04

3.77

V-8-79

74.54 

.36 

12.97 

1.67 

1.74 

.88 

.05 

3.26 

3.67 

.79 

.07

0.75 

.09 

.07 

3.70

V- 100-79

79.85 

.29 

12.67 

.13 

.11 

.29

.39 

5.73 

.54

0.52 

.26 

.04

.78

V-4-73

51.49 

1.35 

16.68 

5.88 

6.00 

4.74 

.19 

9.14 

3.61 

.72 

.20

3.10 

.84 

.02 

2.38

V-67-33

53.60 

1.23 

18.20 

2.37 

7.61 

4.12 

.23 

8.33 

3.29 

.79 

.23

2.60 

.17 

.02 

2.37

VC-184

71.40 

.58 

14.44 

2.12 

1.62 

1.21 

.12 

2.63 

4.14 

1.62 

.12

1.00 

.12 

.32 

2.92

VC-45

73.63 

.43 

14.67 

1.62 

.93 

.84 

.07 

1.11 

6.57 

.05 

.08

0.85 

.15 

.03 

2.84

LC-48

74.73 

.39 

13.99 

1.42 

1.12 

.81 

.06 

1.42 

4.56 

1.42 

.08

1.00 

.14 

.04 

2.96

CO-79-1

75.56 

.33 

13.27 

.95 

1.15 

.64 

.03 

1.90 

4.41 

1.70 

.06

0.85 

.11 

.21 

3.13

Cougar Creek Complex Continued Wolf Creek-Deep Creek and 

Imnaha plutonic rocks

Sheep Creek pluton

SiO2 

TiO2 - 

Al2Og -

Fe2O3 -

MgO 

MnO 

CaO

Na2O - 

P

P» 2O5 

H n+

H

/-vr*r\

FeOtotal/MgO -----

V qc 70

76.47 

.30 

12.86

 f A *)

CO

no

3.96

73

O

rjc

.04

3.01

TP 13

56.66 

.80 

22.07 

1.10

1.44

9 99

03

8

1 1

5.76

1 ^O

31

2.00

/JO

09

1.10

TP-15

65.78 

1.44 

14.31 

1.74

2 nn

1 19

19

6 KG

6.03

3Q

.44

1 3O

3 1O

3.00

TT3 1 Q

69.07 

1.18 

17.48 

1.07

9 QQ

1 3^

.07

9 10

3.69

31

1.70

.44

CO.

2.93

V-70-79

50.65

.78 

18.40

2

CA

.17

11 42

2.64

A£?

1 (\(\

.11

.07

1.33

FID GD

79-1

72.31 

.40 

14.64

.77

1 97

1.34

4.75

4.09

O Q3

1 /?

.04

1.47

V-79-7Q

78.68 

.17 

11.79

QQ

1Q

91

4

c\n

3.96 

.05

0.42

.04

4.41

T

/?Q O

53.80

.75 

17.87

9 Q7

5 O9

5.04

.14

10.81

3.04

OQ

1 31

.10

.10

1.51

TTFI 7Q 1

68.22 

.56 

14.88

9 47

2 45

1.95

.10

4.13 

.49

.10

1 19

.11

.04

2.40
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 157

plutonic rocks. Earlier work gave K-Ar and 40Ar/39Ar 

ages of 223 to 212 Ma. U-Pb studies by Walker (1986) 

yielded a 206pj-)/238u age Of 215 Ma for the younger 

plutonic suite, in good agreement with the K-Ar and 

40Ar/39Ar age determinations. Zircons in trondhjemite 

from the older plutonic suite have a 206pjy238u age Of 

252 Ma. On the basis of these new data, field relations, and a regional understanding of similar plutonic complexes, it is apparent that there are at least two 

separate plutonic suites in the Sparta complex: the 

older plutonic rocks are Late Permian in age and the 

younger plutonic rocks crystallized in the Late Triassic. Good outcrops of both suites are along State Highway 86 in the Powder River canyon west of Richland, 

Oreg.; particularly good outcrops of the younger pluTriassic plutonic and hypabyssal (dike) rocks of the Wallowa terrane

to indicate amount of alteration. , not determined or below instrumental detection limit]

tonic rocks are near Bishop Springs along State Highway 86 in the Sparta quadrangle (table 3.9; fig. 3.2).

Some obvious compositional differences between the 

two plutonic suites are shown in tables 3.10 and 3.11 

and in figure 3.18A. The older suite is composed mostly of gabbro, trondhjemite, quartz diorite, quartz keratophyre, and diabase and is very similar to parts of 

the Canyon Mountain, Oxbow, and Cougar Creek 

Complexes; rare ultramafic rocks occur along the 

southern border of the Sparta complex and rocks of 

intermediate compositions are uncommon. In places, 

the rocks are variably sheared and metamorphosed. 

Included in the older suite are abundant Na-rich felsic 

plutonic bodies, particularly albite-rich trondhjemite 

(albite granite of Gilluly, 1933). The younger suite

Cougar Creek Complex

T-68-14 V-61-79 V-5-73 T-68-26 V-12-73 V-6-73 V-l-37-79 KC-79-1 T-68-11 T-68-23 V-66-79

49.35

1.03

21.50

2.46

5.89

5.05

.15

11.73

2.51

.23

.10

0.13

.91

.08

1.61

49.98

1.10

18.21

3.37

6.81

6.58

.19

11.13

2.38

.20

.05

1.02

.10

.03

1.50

Sheep Creek

50.62

.83

19.51

3.30

7.41

4.92

.23

10.11

2.81

.18

.08

0.68

.28

.02

2.11

51.76

.51

9.67

2.81

7.40

13.64

.21

12.19

1.29

.45

.07

1.40

.30

.20

.73

53.25

.98

18.50

3.08

6.89

4.77

.23

8.19

3.27

.60

.24

2.30

.32

.16

2.03

54.51

.45

24.21

2.06

2.30

1.46

.07

10.80

3.66

.36

.12

0.21

.32

.03

2.85

Plutons of southern Seven Devils Mountains

pluton Continued

62.81 69.36

.70 .36

15.22 15.80

3.37 1.93

3.39 1.53

3.80 1.16

.13 .12

6.41 4.19

3.81 4.05

.25 1.36

.11 .14

2.48 0.63

.16 .11

.01 .05

1.69 2.82

Salmon River

canyon plutonic

rocks

71.67

.41

14.16

2.01

2.02

1.22

.08

4.95

3.15

.23

.10

0.97

.17

.02

3.14

74.90

.28

13.18

1.84

1.42

.89

.03

2.87

4.35

.18

.06

0.67

.30

.20

3.46

Plutons within Blue

Mountains anticline

76.44

.28

13.11

1.09

1.16

.76

3.15

3.78

.23

0.83

.14

.02

2.82

Wildhorse

River dike

T-68-3 T-68-1 V-2-84 V-4-84 V-3-84 V-l-84 V-39-79 V-57-79 V-20-79 V-17-79 V-5-84

68.93

.53

14.54

2.53

2.01

1.67

.09

5.05

4.08

.52

.05

1.12

.18

.25

2.57

70.74

.45

14.94

.63

1.66

1.49

.03

4.30

4.98

.71

.07

1.28

.20

.04

1.50

51.66

.50

19.89

1.78

4.57

6.66

.10

10.97

2.67

1.20

2.33

.16

.20

.93

59.31

1.27

16.42

3.95

4.55

2.67

.18

5.88

4.05

1.38

.34

1.40

.06

.06

3.04

59.39

1.20

15.71

3.57

4.61

2.92

.19

6.41

4.62

1.07

.31

2.01

.04

2.00

2.68

69.60

.57

14.72

1.93

2.43

1.30

.09

2.80

4.11

2.32

.13

0.860

.07

.07

3.20

57.34

.93

17.39

2.07

5.91

4.92

.18

5.39

4.55

1.18

.14

2.48

.16

.01

1.58

74.60

.29

13.65

1.70

.55

.49

2.85

5.32

.50

.05

0.63

.04

.10

4.24

44.22

1.03

21.07

3.03

9.65

6.90

.23

12.30

1.34

.17

.06

3.51

.21

.28

1.79

77.69

.38

12.09

1.06

.91

.53

.03

2.94

4.19

.12

.06

0.40

.14

.03

3.52

77.08

.26

12.16

.86

.70

.55

.04

1.57

4.53

2.20

.05

0.68

.06

1.01

2.66
158 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 3.11. Trace elements of selected late Paleozoic and Triassic

[Results in parts per million. Analysis by neutron activation unless otherwise noted. , not

Sparta complex Oxbow Complex Cougar Creek Complex

Sample ----- S-79-1 V-8-79 V-100-79 V-4-73 CO-79-1 T-68-14 V-61-79

33

99

35

283

V-5-73

1.7 

4

1.3 

.61

.29

.22

1.49

.21

.4

9

29

34

298

Rb 129 1 15 21 X 3

Sr ----- 172 469 194 ^ll 127 2410 1367

Ba 26 *349 117 443 440 94 1 69

Th 2.94 1.06 .38 2.13

U 2.04 .65 1.02

La 9.6 5.3 5.1 11.4 3 1

Ce 19.6 11.3 12.4 25.5 7 5

Sm - 3.1 1.4 3.4 2 1.2

Eu .92 .63 1.11 .81 .76 .64

Gd 3.42 3.4 1.8 2.6

Tb .56 .63 .61 .43 .33

Tm 

Yb 2.04 1.53 2.17 3.21 3.0

Lu .33 .29 .33 .53 .48 .19

Y 234 245 2 8 125 228 222 40

Zr 234 470 187 184 164 U

Hf - - - 7.11 5.57 1.76 4.63 1.0 .6

Ta .254 .41 .35

Nb - 19

Ni-~----- 26 235

Co 7.6 26 .4 30 3 26

Cr~ ------ 8.2 24 13 2 49

Sc 9 2 17 5 31 7 27

Zn 230 91 274

	Cougar Creek Wolf Creek-Deep Creek and 

	Complex Continued Imnaha plutonic rocks

Sample- TP-15 TP-18 V-70-79 DR-SR- V-72-79 T-68-2 HD-79-1 T-68-3 

___________________________________79-1_________

Rb ... 49 18 *3

Sr 454 1 238 X 357 266 *211

Th - 2.48 2.63 .88 1.0 .90

U ... 1.19 .90 -- .36

La 13.1 9.5 2 2.9 3 7

Ce 31.5 22.8 4 .4 6 22

Sm - 8.4 1.2 2.0

Eu 1.85 1.51 .65 .37 .43 1.11

Gd ----- 2.1 6.4

Tb 1.20 .34 .37 1.25

Tm - .99 .60

Yb 6.95 4.79 1.4 1.65 1.00 4.70

Lu .89 .66 .20 .26 .14 .69

Y *54 *41 40 2 12 1 8 254

Zr ^06 487 X 32 107 412

Hf 7.04 6.19 .6 3.13 4.6 1.50

Ta - .47 .38 .22 .30 .24

Nb - X5 X8

Ni ... 7 ... 226

Co 6 8 32 2.6 .7 22

Cr 2 5 117 3.7 7.4 56

Sc 15 21 38 3 4 32

Zn 37 127 78 83

Sheep Creek pluton

T-68-1

7

2360

140

198

154

1.68

.77

8.9

21.1

3.4

.90

3.70

.73

3.51

.57

239

124

3.81

.38

26

10

12

18

69

6

2 150

210

1.50

.80

8

19

3.1

.85

.55

3.0

.48

231

150

4.3

.33

26

6

7

13

35

10

2 160

150

1.4

.7

5

15

3.1

.72

3.5

.52

2.5

.38

227

130

3.9

.32

2 8

3

9

12

Analysis by X-ray fluorescence. 

Analysis by emission spectroscopy.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 159

plutonic and hypabyssal (dike) rocks of the Wallowa terrane

determined or below detection limit or discarded because of high coefficient of variation]

Cougar Creek Complex Continued

T-68-26

417

1 63

.42

.20

4.1

10.6

1.9

.39

...

.31

.31 IIY

*39

.73

~.

8

45

964

85

2 125

V-12-73

130S

492

.29

5

12.1

.83

.63

.30

2.56

.39

125

*41

1.0

8

27

7

31

2 115

V-6-73

1 4T8

473

.24

. 

2.9

6.6

1.7

.93

.26

1.04

.16

1^3

X 51

.97

6

9

2

12

239

V-37-79 KC-79-1 T-68-11 T-68-23 V-66-79 V-35-79

:233 377 ^06 479 *239 ^04

*99 230 *71 424 *215 497

0.6 1.57 .34 .79

.97 .36

7 15.4 3.4 5.1

17 34.9 6.5 9.5

3.3 4.0 1.4 1.6

1.05 1.11 .6 .6

3.5 2.0

74

...

3.0 2.5 1.36 1.38

.45 .41 .22 .21

loo 24 ^16 47 13

403 110 :98 416 1 255 1 255

2.7 3.1 2.55 3.15

.12 .24 .14

16 1 6

20 5 7 6.1 23

23 1 3 4 29 4

23 5 13 9 210

264 2 150 248 229

Plutons of southern Seven Devils Mountains Salmon River canyon Rocks within Blue

V-2-84

33

250

194

.15

...

1.27

4.35

1.37

.72

...

.30

.19

1.18

.18

23

.48

.03

_

22

240

37

V-4-84

26

260

350

1.11

.52

8.62

23.60

5.20

1.66

6.03

.96

.63

3.99

.61

35

217

6.33

.24

'.'.'.

16

139

25

91

V-3-84

24

150

208

1.86

.94

10.50

29.50

5.89

1.69

7.10

1.12

.71

4.49

.67

40

215

6.03

.31

".'.

17

14

27

94

plutonic rocks Mountains anticline

V-l-84 V-39-79 V-57-79 V-20-79 V-17-79

42 *24 11 :4

200 483 *220 817 471

552 ^le 497 35 443

2.87 .40 1.5

1.14 .6

15.40 7 17

36.90 21 4 14

4.78 4.2 1.7 2.5

1.14 .85 .77 1.12

5.28 6.1 3.3

.86 1.04 .48 .66

.60

3.88 3.99

.57 .70 .37 .58

35 *37 249 12 *26

166 *65 470 38 4 14

4.97 2.6 .6 3.3

.39 .61 .27

23

7 23 24 36 1.5

4 44 24 13 1.4

12 29 27 47 13

51 2 115 104 2

TP-13

^l?

*231

.52

.48

4.7

11.9

2.8

1.02

...

.50

.30

2.35

.34

IOA

ll QC

3.56

.21

*

5

4

18

26

Wildhorse

River dike

V-5-84

33

2 85

386

3.12

.81

9.40

20.30

2.12

.48

2.00

.35

.30

1.96

.30

220

73

2.73

.36

_

3

14

5

33
160 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

o

DC

80

75

70

O 

O

60

55

50

45

"x, G

1 23456 

FeOtota |/MgO RATIO

EXPLANATION

(Informal) Sparta complex of Phelps (1979)

Bishop Springs plutonic rocks 

If Gabbro, diorite, tonalite Data field

represents compositional range;

diamonds, separate samples

Older plutonic rocks Heavy line encloses

range of samples for each composition 

Gabbro

Trondhjemite

Quartz keratophyre 

Oxbow Complex of Vallier {1967a, b) 

^ Trondhjemite Data field represents

compositional range; crosses, separate 

samples 

Amphibolite

o

DC

80

75

70

- 65

UJ

£

O 

O

CM

O 

to

60

55

50

45

i r r r

B

1 234567 

FeOtota |/MgO RATIO

EXPLANATION

Cougar Creek Complex of Vallier (1968) Data 

field (CC) represents compositional range; 

dots, separate samples

Sheep Creek plutonic rocks

Wolf Creek-Deep Creek and Imnaha plutonic rocks

Southern Seven Devils Mountains plutonic rocks

Salmon River canyon plutonic rocks

Wildhorse River dike

FIGURE 3.18. FeOtotal/MgO ratio versus SiO2 content for plutonic 

rocks of Wallowa terrane. Calc-alkaline tholeiitic compositional 

boundary from Miyashiro (1974). A, (Informal) Sparta complex of 

Phelps (1979) and Oxbow Complex of Vallier (1967a, b). New data 

(this chapter) supplemented by data from Almy (1977) and Phelps 

(1978). Note Fe-enriched compositions of almost all plutonic rocks. 

Quartz keratophyre samples (data field Q) show calc-alkaline affinity on the basis of this diagram only. Strongly fractionated 

igneous rocks (SiOg contents above approximately 70 percent)

generally show calc-alkaline affinities on this diagram as explained 

in text. For comparison, areas representing Aleutian plutonic rocks 

(E, F, G) and Jurassic to Cretaceous plutonic rocks of Blue Mountains 

region (KJ) are shown (see fig. 3.4). B, Cougar Creek Complex of 

Vallier (1968) and other Permian and Triassic plutonic rocks of 

Wallowa terrane. Note calc-alkaline affinities of many plutonic rocks. 

For comparison, area (KJ) representing Jurassic to Cretaceous plutons (see fig. 3.18A) is shown. Also, compare with plots of plutonic 

rocks of Oxbow Complex and (informal) Sparta complex (fig. 3.18A).
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 161

consists of gabbro, diorite, and quartz diorite (tonalite); highly siliceous trondhjemite is rare or absent. 

These younger rocks, compared to rocks in the older 

suite, are less felsic, are less metamorphosed, and 

have undergone less shearing. Biotite is common in 

some of the younger plutonic rocks, whereas it is absent in the older plutonic rocks.

The chemical differences between the older and 

younger suites are well shown in figure 3. ISA. Highly fractionated trondhjemite and quartz keratophyre, 

mostly of the older group, occupy data fields in the 

upper part of the figure and overlap both Aleutian 

siliceous plutonic rocks and the more siliceous Jurassic to Cretaceous plutons of the Blue Mountains region. The data field for the (undivided) Sparta 

complex and Oxbow Complex on the FeOtoia\/MgO 

versus Cen/Ybn ratio diagram (fig. 3.19) overlaps 

parts of the field for the Aleutian Okmok Caldera 

tholeiite. Rocks of these two complexes are more enriched in LREE's than most rocks of the Canyon 

Mountain Complex.

OXBOW COMPLEX

The Oxbow Complex was described by Vallier 

(1967a, b, 1974) and named for rocks exposed near 

the Oxbow Dam and in the Indian Creek canyon 

northeast of the dam. Geology of the Oxbow region 

also has been studied by Phelps (1978), Balcer 

(1980), Schmidt (1980), Walker (1986), and Ave Lallemant (chap. 7, this volume).

Rocks in the Oxbow Complex are similar to many 

in the Sparta complex and in the Cougar Creek Complex (figs. 3.ISA, 3.18B, 3.19, 3.20A, 3.20B). Rock 

types are amphibolite, mylonite, gneissic mylonite, 

gneiss, diabase, quartz keratophyre, gabbro, diorite, 

quartz diorite, and trondhjemite. Rocks in a wide 

shear zone consist of mylonite, gneissic mylonite, and 

amphibolite; protomylonite and ultramylonite also occur. The shear-zone unit is extremely foliated. The 

foliation planes strike approximately parallel to the 

strike of the dikes (about N. 35°-60° E.); dips on foliation planes vary from about 60° NW to vertical to

cc 

O

O)

O

O 

u_

CM

EXPLANATION

 Cougar Creek Complex of Vallier (1968)

O Oxbow Complex of Vallier (1967a, b) and (informal) 

Sparta complex, undivided

A Sheep Creek plutonic rocks

4 Wolf Creek-Deep Creek and Imnaha 

plutonic rocks

 Southern Seven Devils Mountains plutonic rocks 

Wildhorse River dike

4 5 

Ce n /Yb n RATIO

FIGURE 3.19. FeOtotal/MgO versus Ce^fbn ratios for late 

Paleozoic and Triassic plutonic rocks of Wallowa terrane. 

Data fields for Bald Mountain batholith (BM), tholeiitic 

rocks of Okmok Caldera (C), Canyon Mountain Complex 

(CM), and Aleutian island arc plutonic rocks (E and G) 

are shown for comparison. Diagram also includes plutonic rocks of Cougar Creek Complex of Vallier (1968) (data 

field CC) and the undivided Oxbow Complex of Vallier 

(1967a, b) and (informal) Sparta complex of Phelps (1979) 

(data field OS). Line, dashed and queried where uncertain, 

separates tholeiitic from calc-alkaline fields, n, chondrite 

normalized.
162 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

RARE-EARTH ELEMENTS 

Sm Eu Gd Tb Yb Lu

EXPLANATION

Sparta complex Oxbow Complex of 

(informal) Vallier (1967a, b)

Sample V-100-79 A Sample V-4-73 

Sample SP-79-1 Sample CO-79-1

Ce

RARE-EARTH ELEMENTS 

Sm Eu Gd Tb Yb Lu

EXPLANATION

Sheep Creek plutonic rocks

A Sample T-68-1 

 Sample T-68-3

0 Sample HD-79-1 

 Sample T-68-2

La Ce

RARE-EARTH ELEMENTS 

Sm Eu Gd Tb

La Ce

RARE-EARTH ELEMENTS 

Sm Eu Gd Tb Yb Lu

O 50

La Ce

EXPLANATION

Cougar Creek Complex of Vallier (1968) 

Sample V-5-73 A Sample V-37-79 

Sample V-6-73 Sample TP-18 

Sample T-68-23 D Sample KC-79-1 

Sample V-12-73

RARE-EARTH ELEMENTS 

Sm Eu Gd Tb

La Ce

EXPLANATION

Southern Seven Devils plutonic rocks 

A Sample V-2-84 Sample V-4-84 

 Sample V-l-84 o Sample V-3-84

RARE-EARTH ELEMENTS 

Sm Eu Gd Tb Yb Lu

30

20

EXPLANATION EXPLANATION

Wolf Creek-Deep Creek and 

Imnaha plutonic rocks

 Sample V-72-79

A Sample V-70-79

 Sample DRSR-79-1

Aleutian island arc

1 Mafic plutonic rocks 

1 of Terrible Mountain,

Attu Island (n=5)

(Eocene?)

Wildhorse River dike 

A Sample V-5-84

Salmon River canyon

plutonic rocks 

0 Sample V-39-79

Blue Mountains anticline 

plutonic rocks

 Sample V-20-79

 Sample V-17-79
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 163

60° SE. Lineations of stretched quartz and amphibole 

crystals are nearly horizontal; some have shallow 

plunges.

Mapped field relations show that the oldest units 

in the Oxbow Complex are gabbro, trondhjemite, 

meladiorite, and diabase that crop out in Idaho near 

the mouth of Indian Creek. Dikes whose predeformation compositions include quartz keratophyre, diabase, gabbro, and basalt occur in the wide shear 

zone, which subsequently was intruded by quartz diorite, trondhjemite, and diabase. All gradations of 

mechanical deformation are evident in the Oxbow 

Complex: unsheared plutonic and hypabyssal rocks 

through cataclasites, mylonites, mylonite gneisses, 

and ultramylonites. Rare tachylite in some shear 

zones suggests postmetamorphic movement. The history of tectonic movement, however, is not simple. 

There is good evidence for Triassic sinistral movement (Ave Lallemant, chap. 7, this volume).

U-Pb radiometric ages (Walker, 1986) of rocks in 

the Oxbow Complex are 249 Ma (206Pb/238u on mylo. 

nite sample CO-80-1) and 225 to 222 Ma (206Pb/238u 

on two zircon fractions from trondhjemite sample CO79-1). An 40Ar/39Ar age of 214 Ma on hornblende from 

an amphibolite in the shear zone is interpreted as the 

time of Ar retention in the hornblende after metamorphism (Phelps, 1978). Ave Lallemant and others 

(1980) reported an 40Ar/39Ar age of 220 Ma for an 

amphibolite of the Oxbow Complex. Balcer (1980) reported weighted-average 40Ar/39Ar plateau ages of

FIGURE 3.20. Rare-earth element (REE) diagrams of late 

Paleozoic and Triassic plutonic and hypabyssal rocks of 

Wallowa terrane. Sample numbers correspond to those in 

tables 3.9 to 3.11; for normalizing values used, see figure 

3.6. A, Oxbow Complex of Vallier (1967a, b) and (informal) Sparta complex of Phelps (1979). B, Cougar Creek 

Complex of Vallier (1968). All samples yield nearly flat 

REE patterns (excluding Eu), except sample KC-79-1, 

whose trend is transitional to calc-alkaline. C, Wolf 

Creek-Deep Creek and Imnaha plutonic rocks. These plutonic rocks are strongly tholeiitic. Note negative europium anomaly for sample DRSR-79-1. Data field (n=5) for 

Terrible Mountain (Attu Island of Aleutian island arc; see 

fig. 3.6) is shown for comparison. D, Sheep Creek plutonic 

rocks. This diagram shows that Sheep Creek plutonic 

rocks are derived from tholeiitic magma by extreme fractionation with some light rare-earth element (LREE) enrichment. E, Southern Seven Devils Mountains plutonic 

rocks. Big Lake pluton (sample V-l-84) has so-called transitional pattern shown by some LREE enrichment. F, 

Miscellaneous plutonic rocks. Note strong LREE depletion of hornblende orthogneiss (sample V-20-79) from 

Blue Mountains anticline. Sample V-5-84 trend is transitional to calc-alkaline.

227.8±5.3 Ma and 219.2±7.6 Ma for two hornblende 

cooling ages and interpreted them as "a minimum age 

for deformation and peak regional metamorphism in 

this terrane" (p. 49). Events interpreted from these 

isotopic ages are (1) Late Permian (about 249 Ma) 

dike injections, probably into an older (or comagmatic) gabbro-trondhjemite complex, (2) Late Triassic 

(225-222 Ma) injection of trondhjemite, and (3) Late 

Triassic (228-219 Ma) shearing and metamorphism 

that occurred during sinistral movement (Ave Lallemant, chap. 7, this volume). An unsheared lamprophyre dike with an 40Ar/39Ar age of 197.4 Ma cuts the 

mylonites (Ave Lallemant and others, 1985), thereby 

indicating that most deformation in the Oxbow Complex had ceased by middle Early Jurassic time.

The petrography of most rock types in the Oxbow 

Complex was discussed by Vallier (1967a). In general, the older protoliths are metamorphosed to the 

amphibolite (and hornblende hornfels) facies and to 

the greenschist (and albite-epidote hornfels) facies. 

Mafic rocks were more susceptible to metamorphism 

than the felsic rocks. Most siliceous plutonic rocks 

have an assemblage of albite, chlorite, blue-green 

hornblende, epidote and, locally, secondary quartz, 

sphene, leucoxene, white mica, and (or) pyrite as the 

metamorphic phases. Amphibolite-grade mafic rocks 

have feldspar (andesine and oligoclase) and bluegreen amphibole as major minerals, with chlorite and 

epidote and possibly also sphene, pyrite, and (or) 

quartz as minor minerals.

Most rocks selected for chemical analyses (tables 

3.9-3.11) are siliceous, although mafic and intermediate rocks compose as much as 50 percent of the exposures. The gabbro (sample V-4-73) lies close to the 

contact between the Hunsaker Creek Formation and 

the Oxbow Complex and may cut both rock bodies; 

field relations are not clear. Except for sample VC-45, 

K^O values are higher than those from many other 

trondhjemitic plutonic bodies in the Wallowa terrane 

(table 3.10). Sample VC-45 also has a high Na2O value. The FeOtotal/MgO ratios versus SiO2 contents 

and FeOtotal/MgO versus Cen/Ybn ratios (figs. 3.18A, 

3.19) show an Fe-enrichment trend in both the mafic 

and silicic rocks. The isotopically dated siliceous sample (CO-79-1) has relatively high REE contents (figs. 

3.20A, 3.20B) compared to rocks from the Cougar 

Creek Complex and the (informal) Sparta complex.

COUGAR CREEK COMPLEX

The Cougar Creek Complex (Vallier, 1968, 1974; 

Balcer, 1980; Walker, 1986) is best exposed for about 

10 km along the Snake River canyon between Temperance Creek and Pittsburg Landing (fig. 3.3), but it
164 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

apparently extends northeasterly beneath Tertiary 

basalt and crops out again in the Salmon River canyon north of Whitebird, Idaho. Furthermore, exposures along the South Fork of the Clearwater River 

east of Grangeville, Idaho, are strikingly similar to, 

and essentially on structural trend with, the Cougar 

Creek Complex. The area between the Salmon River 

near Whitebird and the South Fork of the Clearwater 

River east of Grangeville has a Tertiary basalt cover. 

If the Cougar Creek Complex does extend from Temperance Creek on the Snake River northeasterly 

about 65 km to the South Fork of the Clearwater 

River, then the complex is a major structural feature 

within the Wallowa terrane. The Cougar Creek Complex in its type area near Pittsburg Landing has 

been faulted against sheared Permian and Triassic 

stratified rocks along the southeast side and is faulted over rocks of the Jurassic Coon Hollow Formation 

on the northwest.

The Cougar Creek Complex is structurally and petrologically diverse. Most shear zones composed of mylonite and amphibolite are deformed dike (or sill) 

sequences. The dike/shear zones trend about N. 60°- 

70° E. and dip steeply. In places the mylonitic foliation itself is folded, indicating more than one deformation. Individual dikes range in width from a few 

centimeters to as much as 10 m. A marked foliation, 

resulting from mylonitization, parallels the preexisting dike (or sill) trends in most outcrops. A distinct 

subhorizontal lineation is formed by stretched quartz 

porphyroblasts and long axes of amphibole crystals.

The wide range of igneous rock compositions is 

shown in tables 3.9 to 3.11. The coarse-grained plutonic rocks comprise gabbro, norite, meladiorite, diorite, quartz diorite, and trondhjemite; finer grained 

diabase, quartz keratophyre, basalt, and andesite 

also occur, particularly in dike (or sill) zones. Protoliths of the mylonite, gneissic mylonite, and amphibolite in dike zones probably include most of the rock 

types just mentioned; in places plagioclase-rich dikes 

are similar to plagioclase cumulates.

Chlorite schist and hornblende schist are locally 

abundant, particularly along the Oregon side of the 

Snake River canyon near Salt Creek. Metasedimentary rocks are faulted into the complex; metasedimentary rock outcrops along the trail just south of 

Suicide Point are composed of stretched-clast conglomerate and breccia, sandstone, and argillite.

The effects of metamorphism are variable in rocks 

of the Cougar Creek Complex; most dike rocks and 

rocks within wide zones of shearing are in the greenschist and amphibolite facies. Many plutonic rocks 

(samples T-68-14, V-61-79, V-5-73, V-6-73, and KC-79- 

1), however, are not intensely altered. Primary minerals (calcic plagioclase, orthopyroxene, clinopyroxene, hornblende, and biotite) are relatively unaffected 

in some samples selected several meters from fractures and shears. In the amphibolite bodies, however, 

high-temperature metamorphic minerals such as 

hornblende and andesine show retrograde metamorphic effects, indicated by rims of albite surrounding 

the andesine and rims of chlorite and (or) epidote 

surrounding hornblende.

Plutonic rocks of the Cougar Creek Complex are 

mildly Fe-enriched, as shown on the FeOtotal/MgO 

ratio versus SiC>2 content diagram (fig. 3.18B), and 

somewhat more calc-alkaline or transitional than 

most other pre-Jurassic plutonic rocks of the Blue 

Mountains island arc. Most have low K^O contents 

(table 3.10). Trace-element concentrations (table 3.11) 

indicate a wide range of rock types and diverse crystallization histories. Plots of FeOt0tal/MgC) versus 

Cen/Ybn ratios and chondrite-normalized REE diagrams (figs. 3.19 and 3.20B) indicate that most plutonic rocks of the Cougar Creek Complex are 

tholeiitic, marked by relatively flat REE patterns. 

Only sample KC-79-1 (fig. 3.20B) has a trend that 

appears to be transitional between calc-alkaline and 

tholeiitic.

Balcer (1980) determined 40Ar/39Ar ages on amphiboles from a deformed gneissic unit; the ages, based 

on Ar-release spectra, range from 239.2 to 225.8 Ma 

and average 233.6 Ma. Balcer (1980, p. 48) suggested 

that the average age "is the time at which this area 

of the Cougar Creek shear zone cooled below 500±50° 

C and as such represents an upper limit to the time 

of peak metamorphism and deformation." Further 

work by Walker (1986) on rocks from a similar 

gneissic unit, using two zircon fractions and a TeraWasserburg diagram, suggests a late Carboniferous 

(Middle Pennsylvanian) age of 309(?) Ma for magmatic crystallization. Younger siliceous plutonic bodies cut the older gneissic and sheared-dike units. One 

of these younger bodies, a hornblende-biotite quartz 

diorite (sample KC-79-1), contained two zircon fractions that yielded the concordant latest Permian 

206Pb/238u age of 246 Ma. Walker (1986) subsequently sampled a similar intrusive body of the Cougar 

Creek Complex and determined a Late Permian 

206Pb/238U age of 256 Ma. A siliceous protomylonite 

collected from the dike zone near the mouth of Two 

Corral Creek yielded a U-Pb age 262 Ma (Walker, 

1986).

The Cougar Creek Complex had a long history of 

igneous activity that may have started as early as 

309 Ma. The 262-Ma age of the mylonitized silicic 

dike analyzed by Walker (1986) is about the same as 

the age of the Lower Permian tuffaceous rocks of the
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 165

Hunsaker Creek Formation. This similarity in age 

suggests that the rocks are related. The younger, and 

relatively unmetamorphosed, plutonic bodies with 

ages of 256 and 246 Ma probably mark the end of 

Permian igneous activity. These Late Permian rocks 

are about the same age as the older plutonic rocks in 

the Sparta complex and the Oxbow Complex. The average 233.6-Ma amphibole age that was determined 

by Baker (1980) is strikingly close to the fossil age of 

the Middle Triassic (Ladinian) volcanic rocks of the 

Wild Sheep Creek Formation in the Wallowa terrane.

The 309(?)-Ma rocks of the Cougar Creek Complex 

may be the oldest rocks in the Wallowa terrane. If 

some of the rocks of the Canyon Mountain Complex 

are as old as 314 Ma, as suspected by Walker (chap. 

6, this volume), then the two plutonic units may have 

been spatially linked during the late Carboniferous 

and Permian. These old plutonic rocks may represent 

parts of an ancient arc terrane that later hosted Permian and Triassic igneous intrusions and volcanic 

eruptions.

Recent mapping (T.L. Vallier, unpub. data, 1992) of 

the Cougar Creek Complex suggests that the rock 

body is a tilted fragment of island-arc crust that 

formed at deep to intermediate depths (probably 5- 

15 km). The rocks occur in distinct bands or layered 

units. Mylonite zones both occur within and separate 

some of the layered units. The older rocks to the 

north consist of gabbro, norite, and diabase that have 

been deformed and recrystallized into amphibolite 

and gneissic gabbro. The igneous rocks become more 

fractionated and siliceous to the south (shallower 

depth), ending in mostly trondhjemite, quartz diorite, 

rhyolite, diabase, and rare gabbro intrusive bodies. 

The igneous rocks of the Cougar Creek Complex are 

overlain by volcanic flow rocks, conglomerate, breccia, and sandstone.

WOLF CREEK-DEEP CREEK AND IMNAHA PLUTONIC ROCKS

A plutonic-rock complex crops out in places beneath 

Tertiary basalt in the Snake River canyon east of the 

mouth of the Imnaha River (Morrison, 1963; Vallier, 

1974). These rocks are herein informally referred to 

as the Wolf Creek-Deep Creek plutonic rocks in the 

eastern outcrops and the Imnaha plutonic rocks in 

the west near the mouth of the Imnaha River (figs. 

3.2, 3.3). The two areas are separated by an inlier of 

the Columbia River Basalt Group that reaches river 

level and covers the older plutonic rock bodies (Vallier, 1974). The plutonic rock bodies are bordered on 

the northeast by a northwest-trending shear zone 

(Morrison, 1963; Chen, 1985), which is well exposed

in the lower parts of Divide Creek. The metamorphosed and structurally deformed rocks in the shear 

zone are similar to many rocks of the Cougar Creek 

Complex.

An initial 87Sr/86Sr ratio of 0.70311+0.00001 was 

obtained for a gabbro (sample V-70-79, table 3.9) 

near the mouth of Wolf Creek. The rock is relatively 

unaltered (tables 3.10 and 3.11) and has very low 

REE contents (fig. 3.20C) that show a flat tholeiitic 

trend, a positive Eu anomaly, and depletion in 

LREE's. The initial Sr-isotope ratio is low even for 

midocean-ridge basalt and island-arc basalt (Gill, 

1981); chemically, it is the most primitive gabbro yet 

analyzed from the Wallowa terrane.

A striking feature of the Wolf Creek-Deep Creek 

plutonic rocks is the abundance of pegmatitic and aplitic dikes. The pegmatite dikes are composed predominantly of large feldspar and quartz crystals 

(some have lengths of 5 cm or more), with or without 

muscovite. The pegmatite dikes crop out mostly in 

the eastern parts of the inlier along the Snake River; 

they range from a few centimeters to tens of meters 

in width and anastomose irregularly within the host 

rocks. The aplite dikes crop out mostly in the western part of the inlier; these steeply dipping white to 

buff dikes are thin (all are less than 50 cm wide, and 

most are 10 to 20 cm wide) and cut sharply across a 

gabbro host. The aplite dikes protrude like slivers 

from the gabbro host; from a distance, the sheer 

black canyon walls appear to be broken by thin, 

nearly vertical white lines.

An aplite selected for chemical analyses (sample V72-79) contains mostly plagioclase and quartz and, as 

expected, has relatively high amounts of SiO2, A^Oa, 

CaO, and Na2O. K^O content is extremely low (table 

3.10). Trace-element analyses (table 3.11) show low 

amounts of REE's and Y, and relatively high concentrations of Zr and Hf. The Cen/Ybn ratio is about 1.5 

(fig. 3.19). From field relations and chemistry, it appears that the aplite is a late-stage differentiation 

product of the host gabbro. As the gabbro cooled, extension-induced fractures developed, thereby allowing 

volatile-charged fluids to be injected into the fractures, where they crystallized as dikes.

Walker (1986) selected a trondhjemite sample from 

a plutonic body near the mouth of Deep Creek for 

isotopic dating using the U-Pb method. Three zircon 

fractions had concordant 206pb/238u ages of 231 Ma 

(late Ladinian on the time scale of Palmer, 1983).

A sample selected for chemistry from the Imnaha 

plutonic bodies (DR-SR-79-1) is petrographically and 

chemically similar to many of the other Permian and 

Triassic trondhjemites in the Wallowa and Baker terranes. It has notably high SiO2 and low K2O contents.
166 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Its tholeiitic affinity is indicated by both a low Cen/ 

Ybn ratio (fig. 3.19) and the REE plot (fig. 3.20C), 

which is relatively flat with a strong negative Eu 

anomaly. This sample gave slightly discordant 206Pb/ 

238U ages (Walker, 1986) of 228 and 225 Ma (Karnian). Balcer (1980) used the 40Ar/39Ar dating technique 

on hornblende from four samples of plutonic bodies 

exposed near the mouth of the Imnaha River and determined ages of 225.0+2.3 Ma, 223.3+5.9 Ma, 

215.6+3.6 Ma, and 213.0+2.0 Ma (Karnian and Norian 

Stages of the Triassic).

Rocks from the Imnaha and Wolf Creek-Deep 

Creek plutonic bodies apparently crystallized in the 

Middle and Late Triassic (late Ladinian and Karnian, with some crystallization occurring possibly as 

late as the Norian). The ages are the same as, and 

slightly younger than, those of the major Triassic volcanic units in the Wallowa terrane (Ladinian and 

Karnian Wild Sheep Creek and Karnian Doyle Creek 

Formations). The plutonic and volcanic events, therefore, are related in time and space. The plutonic bodies probably are the crystallized magma chambers 

that fed one or more Middle and Late Triassic volcanoes in the Wallowa terrane.

Sheared rock bodies (Morrison, 1963; Chen, 1985; 

Ave Lallemant, chap. 7, this volume) in the Divide 

Creek area north of the Imnaha plutonic bodies probably are composed of Permian and Triassic rocks that 

were deformed prior to and during intrusion of the 

plutons. Foliated rocks in the Imnaha plutonic bodies 

(Morrison, 1963) imply that some deformation occurred during crystallization.

SHEEP CREEK PLUTONIC ROCKS

Plutonic bodies in the Sheep Creek area of the 

Snake River canyon (fig 3.3) have a broad compositional range, but the contacts between the bodies 

have not been well mapped. A gabbro and diorite 

body crosses the river near the mouth of Bills Creek 

and parallels the river for about 12 km in Idaho along 

the hanging wall of a thrust fault. In the Sheep Creek 

canyon, the intrusive bodies are mostly trondhjemite 

and quartz diorite. Four rocks were sampled for 

chemical analyses (tables 3.9-3.11); one, a quartz diorite sample (HD-79-1), was dated isotopically using 

the U-Pb method (Walker, 1986). This sample is similar, both petrographically and in outcrop characteristics, to many of the Permian and Triassic plutonic 

rocks in other parts of the Wallowa terrane. Where 

mapped north of Sheep Creek, the quartz diorite body 

is both in structural contact with and overlain unconformably by Triassic rocks of the Wild Sheep Creek 

Formation. Trondhjemite and quartz diorite clasts in

overlying conglomerate of the Wild Sheep Creek Formation indicate erosion of the plutonic rocks during 

the Middle Triassic (Vallier, 1977).

The Sheep Creek plutonic rocks are calc-alkaline 

on the basis of the FeO^0^ai/MgO ratios versus SiC>2 

contents (fig. 3.185), but are definitely tholeiitic on 

the basis of the FeOtotal/MgO versus Cen/Ybn ratios 

(fig. 3.19) and REE patterns (fig. 3.20Z>). The REE 

patterns are nearly flat. Cen/Ybn ratios range from 

1.3 to 1.6. All samples have negative Eu anomalies. 

Sr contents are lower than in many other Permian 

and Triassic plutonic rocks in the Wallowa terrane; 

Hf, Zr, Ba, and Ta values are within the range of 

those for other Permian and Triassic plutonic rocks 

in the terrane.

Two zircon fractions from sample HD-79-1 yielded 

concordant Early Permian U-Pb ages of 263 Ma 

(Walker, 1986). Clasts of Sheep Creek plutonic rocks 

in overlying strata of the Wild Sheep Creek Formation indicate significant tectonic uplift and erosion 

from about 263 to 235 Ma (the latter is the probable 

oldest age of the Wild Sheep Creek Formation).

PLUTONIC ROCKS OF THE SOUTHERN 

SEVEN DEVILS MOUNTAINS

Plutonic rock bodies in the southern Seven Devils 

Mountains are well described by White (1968) who 

mapped six plutons in an "old mafic suite," three in 

an "older granitic suite," and two in a "younger granitic suite." I selected samples from four separate 

plutonic bodies for chemical analysis (tables 3.9- 

3.11): gabbro (V-2-84) from the White Mountain pluton; diorite (V-4-84) from the Pactolian pluton; diorite 

(V-3-84) from the Horse Pasture pluton; and quartz 

diorite (V- 1-84) from the Big Lake pluton. The first 

three belong to White's (1968) "old mafic suite," and 

sample V-l-84 belongs to his "older granitic suite."

Three of the four samples are Fe-enriched (fig. 

3.18B). Their K2O contents are higher than those of 

most other late Paleozoic and Triassic plutonic rocks 

in the Wallowa terrane; P2Os and Rb also are enriched 

(tables 3.10 and 3.11). In the more siliceous samples 

(diorite and quartz diorite), Th, U, Zr, and Hf abundances also are greater than in most other Permian 

and Triassic plutonic rocks of the Wallowa terrane. 

The FeOtotal/MgO versus Cen/Ybn ratios (fig. 3.19) and 

REE patterns (fig. 3.20.E) show that the plutonic rocks 

are tholeiitic to transitional in composition.

Crystallization ages are unknown. White (1968) 

suggested that the two older suites are most likely 

Late Triassic to Late Jurassic in age. Some of the 

plutonic bodies are in fault contact with Permian 

stratified rocks on the west and are on trend with
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 167

the Oxbow Complex, which is Late Permian and Trias sic in age. On the basis of chemistry, mineralogy 

(mostly greenschist-facies minerals), and field relations all of which are similar to those of known Permian and Trias sic plutonic rocks in the Wallowa 

terrane I infer that the "old mafic suite" and "older 

granitic suite" of White (1968) are also Permian and 

Trias sic in age.

PLUTONIC ROCKS OF THE SALMON RIVER CANYON

There is a broad assortment of plutonic bodies in 

the Salmon River canyon between Riggins, Idaho, 

and the mouth of the Salmon River (figs. 3.1, 3.3). 

Some are dike units and are very similar to Permian 

dike rocks in the Cougar Creek Complex. Other plutonic bodies, from field relations and mineralogy, are 

similar to known Permian and Triassic plutonic bodies in other parts of the Wallowa terrane. Still others 

are nearly unaltered and may be as young as Late 

Cretaceous in age.

Atrondhjemite sample (V-57-79, table 3.9) was taken 

from one of several plutonic bodies exposed in the Salmon River canyon north of Riggins, Idaho. This sample 

is from the same outcrop as Walker's (1986) sample LU 

8-1, which was radiometrically dated. Sample V-57-79 

was chemically analyzed (tables 3.10-3.11), and an 

initial Sr-isotope ratio of 0.70425+0.00003 was determined. Two zircon fractions yielded concordant 206Pb/ 

238U ages of 260 Ma (sample LU 8-1; Walker, 1986). 

The U-Pb age of 260 Ma (Early Permian) suggests that 

the pluton may have crystallized from a magma that 

had served as a source for some of the siliceous flows 

and tuffs of the Hunsaker Creek Formation.

PLUTONIC ROCKS ASSOCIATED WITH THE 

BLUE MOUNTAINS ANTICLINE

Along the Blue Mountains anticline (fig. 3.1; Trauba, 1975; Shorey, 1976), igneous rocks have similarities to rocks from both the Baker and Wallowa 

terranes. For example, ultramafic rocks are abundant 

along the Blue Mountains anticline (Trauba, 1975); 

this abundance is more characteristic of the Baker 

terrane than of the Wallowa terrane. In the Tucannon River area of southern Washington (fig. 3.1), four 

inliers expose pre-Tertiary rocks (Huntting, 1942) 

that are similar to those found in both terranes. The 

largest and most central inlier (Huntting, 1942, fig. 

4) has argillite, chert, and greenstone that are very 

similar to many rocks composing the Elkhorn Ridge 

Argillite of the Baker terrane. The southernmost exposures along Panjab Creek (a tributary of the Tucannon River), however, have volcanic (flow) and 

volcaniclastic rocks similar to those of the Seven 

Devils Group (Vallier, 1977) in the Wallowa terrane. 

Huntting (1942) mapped a small amount of supposed 

diorite that is similar to metagabbro common to both 

terranes. Therefore, rocks in those inliers along the 

Tucannon River may have been part of either the 

Baker or Wallowa terrane.

The rocks analyzed for this study have been placed 

in the Wallowa terrane. My most recent work suggests they should instead be assigned to the Baker 

terrane; however, sufficient justification for such a reassignment is beyond the purpose and scope of this 

chapter.

Hornblende gneiss and trondhjemite of probable 

Permian and possible Triassic age from the inlier 

south of Pendleton, Oreg. (Trauba, 1975), were sampled for chemical analyses (V-20-79 and V-17- 79, tables 3.9-3.11) and strontium isotopic analysis (sample 

V-21-79). Major-, minor-, and trace-element chemistry 

of sample V-20-79, a hornblende gneiss, suggests that 

it was derived from a primitive gabbro. Its extremely 

low REE contents, especially the extremely depleted 

LREE's (fig. 3.20F), and a Lan/Lun ratio of about 0.3 

(Yb was not determined) indicate its tholeiitic affinity. 

Another hornblende gneiss sample, V-21-79, has an 

initial 87Sr/86Sr ratio of 0.70367±0.00006, which suggests a depleted mantle source.

Sample V-17-79, a Na2O-rich trondhjemite (albite 

granite), contains quartz (light blue) and albite as the 

major minerals, with some epidote, chlorite, and 

sphene as secondary minerals. High amounts of Na2O 

and low K^O contents (table 3.10) show its similarity 

to some of the older trondhjemitic plutonic rocks in 

the Wallowa and Baker terranes. The FeOtotal/MgO 

versus Cen/Ybn ratios (fig. 3.19) and REE patterns 

(fig 3.20F) strongly suggest a tholeiitic classification. 

This trondhjemite is very similar to siliceous plutonic 

rocks of the Canyon Mountain Complex and to some 

in the (informal) Sparta complex. Although trondhjemite sample V-17-79 has a higher initial 87Sr/86Sr 

ratio (0.70418+0.00005) than sample V-21-79, described above, its ratio is well within the range expected of igneous rocks from an intraoceanic island 

arc (Gill, 1981).

WILDHORSE RIVER DIKE

A trondhjemite dike sample (V-5-84) was collected in 

the Wildhorse River canyon (table 3.9). This metamorphosed dike is near, and may be within, a major shear 

zone that separates the Baker and Wallowa terranes 

(Mann, 1988). The dike most likely was emplaced
168 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

during Permian to Late Triassic time. Compared to 

most other Permian and Triassic intrusive bodies in 

the Wallowa terrane, this dike rock has unusually high 

amounts of K^O and Rb, a relatively high amount of 

Ba, and a low Sr content (tables 3.10 and 3.11). The 

FeOtotal/MgO versus Cen/Ybn ratios (fig. 3.19) and REE 

patterns (fig. 3.20F) show characteristics that are transitional between tholeiitic and calc-alkaline. It has one 

of the highest Cen/Ybn ratios (about 2.7) of all Wallowa 

terrane Permian and Triassic plutonic rocks (fig. 3.19). 

High contents of LREE's, Rb, and Ba parallel the K2O

content. Metasomatism may have been important in 

producing the dike rock's unusual chemistry.

PETROGENESIS OF PLUTONIC ROCKS 

OF THE WALLOWA TERRANE

Plutonism was active throughout most of the Permian and Triassic in the Wallowa terrane (fig. 3.21). 

Isotopic ages indicate Permian and Middle to Late 

Triassic plutonic pulses. The Permian ages of plutonic rocks in the Wallowa terrane are about the same'

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180 200 

Jurassic

220 240

Triassic

260 280 

Permian

300 320 

Pennsylvanian

GEOLOGIC AGE (Ma) AND PERIOD

FIGURE 3.21. Isotopic ages ranging from Pennsylvanian(?) to Early Jurassic for plutonic rock 

samples from Baker, Wallowa, and Olds Ferry terranes of Blue Mountains island arc. Shading indicates ages determined by K-Ar methods; other ages were determined by U-Pb methods (Walker, 1986). Where a large range is reported in K-Ar ages for a particular sample or 

group of samples from the same plutonic body, the oldest age was used in compiling the 

histograms. Palmer's (1983) time scale-is followed in this diagram. See text for specific age 

determinations and data sources. A, Baker terrane. AM, melange of Aldrich Mountain; B, 

dike or sill cutting Burnt River Schist; CM, samples from Canyon Mountain Complex; G, 

trondhjemite of Granite Creek; M, blueschist near Mitchell, Oreg., which is about 50 km 

west-northwest of John Day, Oreg. (fig. 3.1); PV, trondhjemite at Pleasant Valley in Oxman 

quadrangle, Oreg. B, Wallowa terrane. CC, Cougar Creek Complex of Vallier (1968); I, plutonic rocks of Imnaha-Deep Creek area; O, Oxbow Complex of Vallier (1967a, b) (youngest 

age is for crosscutting lamprophyre dike reported by Ave Lallemant, chap. 7, this volume); 

S, (informal) Sparta complex of Phelps (1979); SC, plutonic rocks of Sheep Creek; SR, Salmon River plutonic rocks; C, Olds Ferry terrane. C, Cuddy Mountains plutonic rocks; H, 

Brownlee Reservoir plutonic body exposed near Huntington, Oreg.; MI, plutonic rocks of 

Mineral-Iron Mountain mining district.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 169

as the youngest ages of dated rocks in the Canyon 

Mountain Complex in the Baker terrane (fig. 3.21A); 

these age relations suggest a comagmatic series. The 

Middle to Late Trias sic plutonic pulse is particularly 

distinct in the Wallowa terrane (fig. 3.2IB).

Plutonic rocks of the Permian and Trias sic suites in 

the Wallowa terrane are similar. Na2O-rich trondhjemites occur in both suites, but they are more common in the Permian plutonic rocks. Almost all of the 

larger complexes have gabbro, diorite, quartz diorite, 

and trondhjemite plutonic bodies with diabase-quartz 

keratophyre-trondhjemite dike sequences. I presume 

that the Permian plutonic bodies served as zones of 

weakness for the intrusion of Middle and Late Triassic 

plutons; a good example is the Late Triassic dioritequartz diorite unit (exposed along State Highway 86 

near Bishop Springs), which apparently was emplaced 

into older Permian rocks of the (informal) Sparta complex. Quartz keratophyre dikes and associated trondhjemite plutonic bodies of Permian age are no doubt 

related to the Na2O-rich and K^O-poor Permian volcanic rocks (flows and tuffs) of the Wallowa terrane.

In the following numbered paragraphs, interpretations summarized in table 3.2 and conclusions made 

by Phelps (1979), Phelps and Ave Lallemant (1980), 

and Leeman and others (chap. 1, this volume) are 

used to discuss the petrogenesis and possible tectonic 

significance of Permian and Triassic plutonism in the 

Wallowa terrane.

(1) The plutonic rocks are mostly tholeiitic in composition, although a few, such as those from the 

Sheep Creek plutons, are transitional between calcalkaline and tholeiitic. Most have low Cen/Ybn ratios, 

low initial Sr-isotope ratios, and low quantities of K 

and Ti. Fractional crystallization of plagioclase, olivine, orthopyroxene (not always present), augite, and 

magnetite (with or without amphibole removal) at 

relatively low pressures and high temperatures (Gill, 

1981; Grove and Baker, 1984) was most likely the 

dominant differentiation process.

(2) Regional geologic relations, isotopic ages, and 

rock chemistries suggest that all late Paleozoic and 

Triassic plutonic bodies were emplaced in a volcanic 

arc, most likely along the magmatic axis.

(3) A Pennsylvanian(?) U-Pb age for a sample 

from the Cougar Creek Complex suggests pre-Permian magmatism. The presence of plutonic rock clasts 

in Lower Permian conglomerates indicates plutonism 

in the Early Permian, but does not rule out the possibility of older igneous activity. The amount of prePermian igneous activity is not known, but it may 

have been extensive.

(4) The oldest well-documented major plutonic 

suite crystallized in the interval from 265 to 245 Ma.

Rocks from three of the plutonic bodies have U-Pb 

isotopic ages in the 265- to 260-Ma range, approximately synchronous with the Early Permian eruptions 

of tuffs and lavas that formed the Hunsaker Creek 

Formation. Some of the plutons may be crystallized 

magma chambers that fed the volcanic eruptions.

(5) The plutonic rock samples from the Canyon 

Mountain Complex of the Baker terrane and the Cougar Creek Complex of the Wallowa terrane that yielded 

Pennsylvanian(?) radiometric ages may be part of the 

magmatic axis of an ancient arc associated with the 

early evolution of the Blue Mountains island arc.

(6) It is clear that trondhjemite with characteristics 

of oceanic plagiogranite can form in tectonic settings 

other than midocean spreading centers (Hawkins, 

1988). Wherever depleted mantle is partially melted 

and the resultant melt is not mixed with abundant 

sediment and (or) liquids derived from the subducting 

slab, then similar magmas should be expected.

(7) The excess Na2O in the Na2O-rich trondhjemites (such as the so-called albite granite of the older 

plutonic suite of the Sparta complex) may have an 

entirely igneous origin (Phelps, 1978). If so, the magma chambers probably maintained open systems 

with free seawater interchange. Spooner and Fyfe 

(1973) and Coleman (1977) discussed the characteristics of hydrothermal alteration in subsea geothermal 

(spreading-center) systems; similar effects probably 

occur in island-arc settings. In an extensional setting, such as that at spreading centers (and in subsea parts of island arcs), circulation of hot seawater 

is probably voluminous. Tholeiitic magmas are hot 

and crystallize at relatively shallow levels in the 

crust; their high temperature and shallow depth can 

account for both the driving thermal energy and seawater accessibility. Seawater, by directly mixing with 

the magma during late stages of crystallization 

through a system of fractures in the upper parts of 

the chamber, probably has a significant role in Na2<3 

metasomatism. Furthermore, in volcanic arcs, seawater brine probably can be trapped in the thick piles of 

subducted sediments; this process would increase the 

quantity of Na2O available for metasomatic exchange. In many places within the Na2O-rich trondhjemite plutons, Na2O contents of rocks increase in 

fractures and shear zones; the fractures probably 

served as conduits for hot seawater. A double diffusive model, such as that proposed by J.L. Bischoff 

(oral commun., 1989), may be necessary to account 

for a brine that contains sufficient Na2O for the extensive metasomatism evident in these altered rocks.

(8) Quartz keratophyre dikes are genetically related to the Na2O-rich trondhjemites; both rock types 

are late-stage differentiation products.
170 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

(9) Diorite and quartz diorite (younger Triassic plutonic suite) and trondhjemite (older Permian plutonic 

suite) in the (informal) Sparta complex may be cogenetic, but are not related through the process of magmatic 

differentiation, as proposed by Phelps (1978,1979). On 

the basis of available age and field data, it is evident 

that separate melting and crystallization took place. 

The source rocks were either the same or very similar, 

and the same processes of differentiation occurred.

(10) Liquids with REE concentrations similar to 

those of the trondhjemite and quartz diorite samples 

could be generated by partial melting of an oceanic 

tholeiite (Phelps, 1978; Phelps and Ave Lallemant, 

1980). Just as possible, and probably more likely, is that 

the magmas resulted from partial melting of a depleted 

mantle with some assimilation of crustal rocks; influences from the downgoing slab probably were minimal. 

This problem, however, is difficult to sort out in modern 

arcs (Gill, 1981; Kay, 1984) let alone in ancient arcs, 

where even the arc polarities are difficult to determine.

(11) Similarities in ages and compositions of many 

plutonic bodies within the Wallowa terrane suggest a 

cogenetic relationship. On the basis of the high volume of these rocks, the plutonic rocks appear to underlie and to have intruded a large part of the 

Wallowa terrane. The older plutons with associated 

dike sequences (Cougar Creek Complex) probably are 

either basement or (and) volcano-root magma chambers for the Lower Permian Hunsaker Creek and 

Windy Ridge Formations; the younger Triassic plutons probably are crystallized magma chambers that 

evolved beneath Triassic volcanoes.

VOLCANIC (FLOW) AND ASSOCIATED HYPABYSSAL 

INTRUSIVE ROCKS OF THE WALLOWA TERRANE

Almost all volcanic (flow) and hypabyssal intrusive 

rocks in the Wallowa terrane are of Permian and Triassic age. Lava flows of Permian age are relatively 

easy to separate from Triassic flows, both in the field 

and in the laboratory. The Permian flows are more 

siliceous, have a near-equilibrium greenschist-facies 

mineralogy, are extremely Na2O-rich and K^O-poor, 

and are associated with thick sequences of siliceous 

tuffs. Reflecting both their original petrologic settings 

and subsequent metamorphism, the Triassic rocks 

are more mafic, are metamorphosed to different intensities (although mostly in the greenschist facies), 

and contain variable amounts of Na2O and

PERMIAN VOLCANIC (FLOW) AND 

ASSOCIATED HYPABYSSAL INTRUSIVE ROCKS

Permian rocks are particularly well exposed in the 

southern Wallowa Mountains and the Homestead,

Oreg., area of the Snake River canyon between Oxbow, Oreg., and the Kleinschmidt grade (Lindgren, 

1901; Livingston and Laney, 1920; Anderson, 1930; 

Gilluly, 1937; Cook, 1954, Wetherell, 1960; Bostwick 

and Koch, 1962; Vallier, 1967a, b, 1974, 1977; Vallier 

and others, 1977; Scheffler, 1983; La Pierre and others, 1988). Some reconnaissance studies farther 

north along the Snake and Salmon River canyons 

and in the Seven Devils Mountains indicate that Permian rock outcrops are much more extensive than 

previously realized. Volcanic (flow) rocks and associated hypabyssal intrusive rocks are relatively rare, 

however, and compose no more than five percent of 

the Permian sequences.

Vallier (1977) divided the Permian rocks into the 

Windy Ridge and Hunsaker Creek Formations (fig. 

3.17). Best exposures of the Windy Ridge Formation 

are along the Idaho Power and Light Company road 

in Idaho immediately north of The Oxbow of the 

Snake River (Vallier, 1974). The Windy Ridge Formation is composed almost entirely of stratigraphically 

coherent quartz keratophyre (and keratophyre) pyroclastic rocks and rare lava flows, all of which are cut 

by mafic dikes. In the siliceous volcanic flows, phenocrysts of quartz and plagioclase feldspar (albite) are 

set in a felty quartz-albite groundmass. Quartz phenocrysts generally are either bipyramidal or rounded 

by magmatic resorption. Primary magnetite is uncommon. Penninite and epidote replace the other, 

originally very rare primary mafic minerals, which 

were ortho- and clinopyroxenes. In addition to quartz 

and albite, penninite, sphene, white mica, epidote, 

calcite, hematite, and pyrite occur in the groundmasses. The mafic dikes are basalt and basaltic andesite porphyries, gabbro, and diabase; in these dikes 

most mafic minerals have been replaced by epidote 

and chlorite and the primary plagioclase is replaced 

by albite. Pumice and other glassy debris in the pyroclastic rocks are entirely recrystallized to quartz and 

albite; vague outlines of the original fragments can 

be observed on some slabbed surfaces and in a few 

thin sections.

The Windy Ridge Formation, originally designated 

as Early Permian(?) in age (Vallier, 1977) is now considered to be a lithofacies equivalent of the lower 

part of the Hunsaker Creek Formation (fig. 3.17). 

Such lithofacies equivalence supports an Early Permian age for the Windy Ridge Formation. The lithologic similarity between this formation and the 

tuffaceous lower part of the Hunsaker Creek Formation, as well as the poorly exposed contact between 

the units, makes it reasonable to correlate the units. 

However, the presence of abundant mafic dikes, the 

dominance of pyroclastic debris, an absence of epiclastic sedimentary rocks, and the siliceous composi-
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 171

tion of the Windy Ridge Formation distinguish it 

from most parts of the Hunsaker Creek Formation.

The rock types of the Hunsaker Creek Formation 

(Vallier, 1967a, 1977) are more diverse than those of 

the Windy Ridge Formation. Most are fragmental (pyroclastic and epiclastic) rocks, many of which are siliceous tuffs that were deposited in a submarine 

environment. Locally abundant and thick conglomerate and sandstone units indicate close proximity to one 

or more major landmasses, probably islands. The rocks 

are dominantly pyroclastic breccia and tuff and epiclastic conglomerate, breccia, sandstone, and siltstone. 

Although flows are rare, in places dikes are abundant; 

concentrations of siliceous dikes near Homestead, 

Oreg., and in the lower parts of Deep Creek canyon in 

Idaho near the Hells Canyon Dam may mark proximity 

to volcanic vents. These dikes and sills are part of the 

Hunsaker Creek Formation. The presence of trondhjemite, quartz diorite, and gabbro clasts in conglomerate units (Wetherell, 1960; Vallier, 1967a, b, 1977) of 

the Hunsaker Creek Formation indicates not only that 

the Permian source area was lithologically diverse, but 

also that the islands were high; erosion of the older 

plutonic rocks indicates substantial uplift. Furthermore, the islands were eroding contemporaneously 

with the outpourings of siliceous lavas and tuffs.

Three samples from the Windy Ridge Formation 

and nine from the Hunsaker Creek Formation were 

selected for chemical analysis (tables 3.12-3.14). The 

mineralogy and chemistry of siliceous rocks from the 

two units are similar. High H2O and CC>2 contents 

(table 3.13) reflect both extensive weathering and 

large amounts of secondary calcite.

Sample VB-79 is from a dike that cuts the Windy 

Ridge Formation, and samples VS5-1 and VS5-7 are 

from rhyolite (quartz keratophyre) flows from the 

Windy Ridge Formation. Sample VB-79 is a low-K basaltic andesite (classification of Gill, 1981). Compared 

to a representative low-K tholeiitic basaltic andesite 

from modern arcs as defined by Gill (1981, p. 100), 

sample VB-79 has relatively high amounts of TiO2 (2.00 

percent for VB-79 versus 0.57 percent for modern arcs), 

Na2O (6.73 percent versus 1.8 percent), ?2O5 (0.41 

percent versus 0.08 percent), MnO (O.42 percent versus 

0.19 percent), Y (40 parts per million versus 17 parts 

per million), Zr (140 versus 28 parts per million), Hf 

(3.4 versus 0.8 parts per million), Nb (7 versus 0.6 parts 

per million), Zn (1,070 versus 89 parts per million), and 

relatively high REE contents (tables 18.13 and 18.14). 

However, it has significantly lower amounts of K2O 

(0.04 percent for sample VB-79 versus 0.40 percent for 

basaltic andesite of modern arcs), CaO (4.27 percent 

versus 11.0 percent), and Sr (106 parts per million 

versus 235 parts per million). The abundance of the 

high-field-strength elements Ti, Zr, Hf, and Nb is more

similar to that of alkalic basalt than of island-arc low-K 

basalt or basaltic andesite. Samples VS5-1 and VS5-7 

are quartz keratophyres (low-K rhyolites). These siliceous rocks have abundant Na2O and low CaO and K2O 

contents (table 3.13). Furthermore, sample VS5-7 is 

enriched in LREE's with a Cen/Ybn ratio of about 2; Sr 

and Ba contents are depleted and Nb and Y are enriched compared to most modern arc rhyolites.

Volcanic (flow) rocks and associated hypabyssal intrusive rocks in the Hunsaker Creek Formation include rare basalt, basaltic andesite, and andesite 

with more abundant dacite and rhyolite (quartz keratophyre). No basalt flow was sampled although 

some of the intraformational dikes and sills are microgabbro, diabase, or basalt porphyry. All rocks have 

near-equilibrium greenschist-facies mineralogies. Albite replaced most calcic plagioclase, and chlorite and 

epidote replaced almost all mafic minerals (predominantly pyroxene); the secondary groundmass minerals epidote, chlorite, sphene, calcite, and white mica 

are ubiquitous. Rare relict clinopyroxene phenocrysts 

occur (for example in sample VC-197).

Basalt sample VB-88 (tables 3.12 and 3.13) from 

the Hunsaker Creek Formation has high amounts of 

TiC>2, Na2O, and ¥<£>§ and a low K2O content compared to other island-arc basalt (Ewart and others, 

1973; Vallier and others, 1985; Kay and Kay, 1991). 

Ba, Zr, Y, Hf, and the REE contents (table 3.14) also 

are relatively high. Basaltic andesite samples (VC131 and VC-197) have relatively high Na2O contents 

compared to lavas from modern island arcs. K2O contents are extremely low. Andesite samples (VB-91 

and LS7-32b) have some chemical variability in the 

amounts of TiO2 and K2O, but are consistently high 

in Na2O. Rhyolite (quartz keratophyre) samples (VC225 and VC-15) also have high Na2O and variable 

K2O contents; trace-element contents of sample VC225 (table 3.14) are similar to those of sample VS5-7 

from the Windy Ridge Formation.

Relatively high Na2O, TiO2 , P2O5 , Hf, Zr, Y, Ba, and 

REE contents, compared to those of modern island-arc 

rocks (for example, see Ewart and others, 1973; Gill, 

1981; Vallier and others, 1985; Ewart and Hawkesworth, 1987; Kay and Kay, 1994), and high FeOtotal/ 

MgO ratios along with low K2O contents and low Cen/ 

Ybn ratios suggest that the petrogenesis of these Permian rocks is different from that of most igneous 

rocks of present-day island arcs. According to FeOtotal/ 

MgO ratios versus SiO2 contents (fig. 3.22), the rocks 

are both tholeiitic and calc-alkaline. The Cen/Ybn ratios (fig. 3.23) indicate a tholeiitic trend between those 

of the Tonga arc and Okmok Caldera in the Aleutians. 

The REE patterns indicate fractionated-tholeiitic 

trends and slight LREE enrichment (fig. 3.24). The 

petrogenesis of these rocks is problematical.
172 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIAEY ROCKS, BLUE MOUNTAINS REGION

TABLE 3.12. Lithology, sample locality, and field description of volcanic (flow) and associated hypabyssal intrusive rocks from the Lower

Permian Windy Ridge and Hunsaker Creek Formations

[Samples arranged by increasing SiO2 content to correspond to chemical analyses in tables 3.13 and 3.14. Do. and do., ditto]

Sample Rock type 1 Sample locality Field description and remarks

VB-88 Basalt - Along road south of Ballard Creek; Sill in bedded-tuff unit; Hunsaker

elevation 543 m (1,780 ft); SE1/4SW1/4 Creek Formation, 

sec. 11, T. 6 S., R. 48 E.; Homestead 

quadrangle; Oreg.

VC-131 Basaltic andesite Ridge between Homestead Creek and Lava flow about 4 m thick in tuff

Whiskey Gulch; elevation 975 m (3,200 breccia sequence; Hunsaker 

ft); center of SE1/4 sec. 16, T. 6 S., R. Creek Formation. 

48 E.; Homestead quadrangle; Oreg.

VC-197 do - Ridge between Irondyke and Holbrook Lava flow about 5 to 6 m thick in

Creeks; elevation 598 m (1,960 ft); mafic tuff breccia; Hunsaker 

SE1/4SW1/4 sec. 21, T. 6 S., R. 48 E.; Creek Formation. 

Homestead quadrangle; Oreg.

VB-79 do Idaho Power and Light Company road; Mafic dike that cuts the Windy

elevation 550 m (1,820 ft); NE1/4 Ridge Formation. 

NW1/4 sec. 17, T. 20 N., R. 4 W.; 

Homestead quadrangle; Idaho.

VB-91 Andesite - North side of Ballard Creek; elevation Lava flow about 3 m thick in tuff732 m (2,400 ft); SW1/4NW1/4 sec. 11, and-sandstone sequence; 

T. 6 S., R. 48 E.; Homestead Hunsaker Creek Formation, 

quadrangle; Oreg.

LS7-32b do North side of Hunsaker Creek; elevation Large clast about 70 cm

869 m (2,850 ft); east edge of SE1/4 indiameter from breccia; 

NE1/4 sec. 6, T. 7 S., R. 48 E.; Hunsaker Creek Formation. 

Homestead quadrangle; Oreg.

LS9-12 Dacite North side of Ballard Creek; elevation Clast about 15 cm in diameter

747 m (2,450 ft); SW1/4NS1/4 sec. 11, from breccia;Hunsaker Creek 

T. 6 S.; R. 48 E.; Homestead Formation, 

quadrangle; Oreg.

LS7-36 do North side of Hunsaker Creek; elevation Lava flow in flow-and-breccia

900 m (2,950 ft); SE1/4NE1/4 sec. 6, sequence; Hunsaker Creek 

T. 7 S., R. 48 E.; Homestead Formation, 

quadrangle; Oreg.

VC-225 Rhyolite Near top of Grassy Ridge about 1.6 km Sill(?) or flow(?); poor outcrop;

west of Kinney Point; elevation 1,905 Hunsaker Creek(?) Formation, 

m (6,250 ft); NW1/4SE1/4 sec. 10, T. 21 

N., R. 3 W.; Cuprum quadrangle; 

Idaho.

VS5-7 do - Idaho Power and Light Company road; Flow in thick sequence of tuff

elevation 550 m (1,820 ft); NW1/4 breccia; Windy Ridge Formation. 

SW1/4 sec. 17, T. 20 N., R. 4 W.; 

Homestead quadrangle; Idaho.

VS5-1 do Idaho Power and Light Company road; Do.

elevation 550 m (1,820 ft); NW1/4 

SW1/4 sec. 17, T. 20 N., R. 4 W.; 

Homestead quadrangle; Idaho.

VC-15 do Ridge between forks of Homestead Flow(?) in tuff-and-sandstone

Creek; elevation 1,310 m (4,300 ft); sequence; Hunsaker Creek 

NE1/4SE1/4 NW1/4 sec. 17, T. 6 S., R. Formation. 

48 E.; Homestead quadrangle; Oreg.

Rocks are metamorphosed to the greenschist facies; Na2O-rich rock can be classified either as spilite (metabasalt and metabasaltic andesite), 

keratophyre (meta-andesite), or quartz keratophyre (metadacite and metarhyolite).
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 173

TABLE 3.13. Major- and minor-element oxides of Permian volcanic (flow) and associated hypabyssal intrusive rocks from the Lower

Permian Windy Ridge and Hunsaker Creek Formations

[Oxides in weight percent, normalized to 100 percent; prenormalization volatile contents included to indicate amount of alteration. , not determined]

Sample- VB-88 VC-131 VC-197 VB-79 VB-91 LS7-32b LS9-12 LS7-36 VC-225 VS5-7 VS5-1 VC-15

SiO2 

Ti02

Fe2O3~ -- 

FeO ------

MgO -

MnO ------

CaO - 

Na2O 

K20-~-~ 

P205 -----

51.95

2.55

16.21

5.81

7.92

5.79

.23

3.13

5.90

.08

.43

54.36

1.89

16.65

4.50

5.34

3.47

.21

10.05

3.14

.02

.37

54.49

.93

17.39

2.97

5.30

7.85

.17

6.37

4.24

.08

.21

54.74

2.00

15.73

5.52

6.21

3.93

.42

4.27

6.73

.04

.41

59.81

1.77

16.44

3.92

4.20

2.74

.25

3.58

6.84

.13

.32

60.04

1.22

16.40

8.18

1.35

1.96

.05

2.83

7.02

.74

.21

62.38

.47

13.92

7.46

4.43

.13

4.40

5.01

.10

1.20

62.94

1.31

14.47

7.95

2.58

.22

3.46

5.93

.73

.41

68.18

.85

15.00

4.98

.85

.51

1.61

6.99

.83

.20

76.42

.35

12.78

.61

1.65

2.15

.08

.45

4.71

.80

76.98

.36

13.21

.55

1.63

1.12

.12

.00

5.80

.18

.05

79.97

.24

12.09

1.02

.16

.24

.16

.00

5.89

.15

.08

H2O+ 

H20-~ -----

C02 ------

FeOtotal/MgO

3.79

.64

1.46

2.27

3.10

.36

.05

2.71

4.10

.79

.08

1.02

2.20 

.27 

.08

2.84

2.49

.31

2.24

2.82

1.71

.65

1.20

4.45

2.77

1.51

1.68 3.08

0.59

.26

1.04

10.45

1.67

.27

.18

1.02

1.10

.11

.02

1.90

0.57 

.15 

.06

4.49

TABLE 3.14. Trace elements of selected volcanic (flow) and associated hypabyssal intrusive rocks 

from the Lower Permian Windy Ridge and Hunsaker Creek Formations

[Results in parts per million. Analysis by neutron activation unless otherwise noted. , not determined or below instrumental 

detection limit or discarded because of high coefficient of variation]

Sample VB-88 VB-79 VB-91 LS7-32b VC-225 VS5-7

Rb1 -- 3 2 4 14 19

Sr1 68 106 116 130 147 40

Ba1 - 989 160 136 168 304 112

Th - .7 1.7 1.2 

U -

La-- - 9 9 9 5 13 13

Ce - 26 12 33 33

Sm 5.2 4.8 4.5 2.6 6.2 5.1

Eu 1.80 1.62 1.48 .43 1.66 1.21

Gd - - - 6.7

Tb - 1.30 1.07 1.06 .45 1.10 .92

Tm .66 .47 .56 .53

Yb - 4.4 4.1 4.6 1.9 4.6 4.4

Lu - .67 .60 .66 .30 .65 .65

Y1 - 43 40 44 21 45 46

Zr1 223 140 217 68 192 176

Hf - 4.6 3.4 5.2 2.2 4.5 4.6

Ta - .56 .94 3.45 2.54

Nb1 - 7 15 19 

Ni -

Co 31 25 15 13 8 6

Cr - 11 21

Sc 35 32 25 36 16 10

Zn------- 168 1,070 168 82 39 88

Analysis by X-ray fluorescence.
174 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

O 

cc

I

80

75

70

O 

O

CM 

O

60

55

50

45

0 1 23456 

FeOtota |/MgO RATIO

FIGURE 3.22. FeOtotal/MgO ratio versus SiO2 content for Permian 

and Triassic lavas and dikes of Wallowa terrane. Dots, Permian 

rocks; triangles, Triassic rocks. (Two samples, one with an FeOtotaj/ 

MgO ratio of 1.71 and an SiO2 content of 48.2 weight percent and 

the other with an FeOtotai/MgO ratio of 2.37 and an SiO2 content 

of 78.29 weight percent, are plotted as Triassic but may be otherwise.) Main data field for Triassic rocks, WA, shows that they are 

strongly tholeiitic. Data fields A through D represent Aleutian 

island arc volcanic rocks (fig. 3.4), X represents tholeiitic rocks of 

Ata Island (Tonga island arc; fig. 3.4), and Q represents quartz 

keratophyre of (informal) Sparta complex of Phelps (1979) (fig. 

3. ISA). Calc-alkaline tholeiitic compositional boundary from 

Miyashiro (1974).

FIGURE 3.24. Rare-earth element (REE) diagram for Early Permian lavas and dikes. Note relatively flat patterns (excluding Eu 

anomaly) for all samples. Sample numbers correspond to those 

in tables 3.12 through 3.14; for normalizing values used, see 

figure 3.6.

123456 

Ce n /Yb n RATIO

FIGURE 3.23. FeOtotal/MgO versus Cen/Ybn ratios for Permian 

(dots) and Triassic (triangles) lavas and dikes of Wallowa terrane. 

~E, data field for Triassic lavas, extent uncertain where queried. 

Three data points plot off scale, as indicated by arrows: Values of 

FeOtota]/MgO ratio for two Permian lavas are both 10.4, and for 

one Triassic lava, 15.2. Two samples (basalt, FeOtotal/MgO=1.71; 

rhyolite, FeOtotal/MgO=2.37) are plotted as Triassic lavas, although they may not be Triassic in age (see text). Data fields B, 

C, and D represent Aleutian island arc volcanic rocks, and fields 

X and Z, Tonga island arc volcanic rocks (fig. 3.5). Line, dashed 

and queried where uncertain, separates tholeiitic and calc-alkaline fields, n, chondrite normalized.

La Ce Sm Eu Gd Tb 

RARE-EARTH ELEMENTS

EXPLANATION

Tm Yb Lu

Sample LS7-32b 

Sample VB-79 

Sample VB-91 

Sample VC-225
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 175

TRIASSIC VOLCANIC (FLOW) AND 

ASSOCIATED HYPABYSSAL INTRUSIVE ROCKS

The Triassic volcanic (flow) and associated hypabyssal intrusive rocks in the Wallowa terrane occur 

in the upper part of the Clover Creek Greenstone 

(Gilluly, 1937; Nolf, 1966) and in the Wild Sheep 

Creek and Doyle Creek Formations of the Seven Devils Group (Vallier, 1977). There are striking differences between the Permian and Triassic sequences in 

outcrop characteristics, field relations, petrography, 

and chemistry (Vallier, 1967a, 1977; Vallier and Batiza, 1978; Scheffler, 1983; La Pierre and others, 1988). 

Compared to Permian lavas, compositions of Triassic 

lavas are more similar to those from present-day island arcs.

Middle (Ladinian) and Late (Karnian) Triassic flows 

and dikes of the Wallowa terrane range in composition 

from basalt to rhyolite (tables 3.15-3.17). Except for 

samples V-26-79, V-51-79, and V-10-79, the samples 

were collected from outcrops in the Snake and Imnaha River canyons and closely adjacent areas.

Sample V-26-79, a basalt, was collected from outcrops along the road paralleling the Tucannon River 

in Washington. Its age could be other than Triassic, 

and its association with chert, argillite, gabbro, and 

greenstone (Huntting, 1942) suggests that those outcrops could easily be assigned to the Baker terrane 

(fig. 3.8) and not to the Wallowa terrane. The chemistry supports assignment to the Baker terrane. For 

consistency with previous terrane assignments (Silberling and others, 1984), however, I have tabulated 

data for this basalt with other Wallowa terrane lavas, although it is not included in the diagrams and 

discussions dealing with the Triassic rocks of the 

Wallowa terrane.

Stratigraphic relations between the mostly marine 

Wild Sheep Creek Formation (green or greenish, pillow lavas, thick debris-flow breccia units, carbonate 

lenses) and the partly subaerial and partly marine 

Doyle Creek Formation (red or-reddish, abundant silicic tuff units, few pillow lavas, large vesicles in 

flows) need further clarification (fig. 3.17). In many 

places the two formations are difficult to separate. 

On the basis of work completed subsequent to the 

initial stratigraphic synthesis (Vallier, 1977), it is apparent that parts of the Doyle Creek Formation are 

subaerial or shallow-marine-lithofacies equivalents of 

the Wild Sheep Creek Formation. However, in many 

places it is evident that strata of the Doyle Creek 

Formation overlie the Wild Sheep Creek Formation. 

The Doyle Creek Formation is considered to be, in 

total, a lithofacies equivalent of both the Wild Sheep 

Creek Formation and the overlying unit. On the 

basis of composition and depositional character (Doyle

Creek rocks are more siliceous and tuffaceous), the 

more oxidized rocks in the Doyle Creek Formation 

may merely represent the final products of volcano 

maturation, growth above sea level, and erosion. Ladinian fossils have been recovered from red marine 

beds of the formation. No fossils, however, have been 

found in the younger and mainly subaerial parts of 

the Doyle Creek Formation near its type area. At 

Pittsburg Landing, Karnian Halobia sp. were recovered in beds (newly named Kurry unit) that overlie 

the Wild Sheep Creek Formation; these strata of the 

Kurry unit have been recently reassigned to the 

Doyle Creek Formation (White and Vallier, 1994). To 

further confuse age relationships, Karnian Halobia 

sp. were recovered from the Wild Sheep Creek Formation from northern parts of the Snake River canyon (Grant, 1980). Thus, the ages of both the Wild 

Sheep Creek and the Doyle Creek Formations are 

now considered to be Middle (Ladinian) and Late 

(Karnian) Triassic; this is a change in age assignment for the Doyle Creek Formation, which was previously thought to be no older than Late Triassic 

(Vallier, 1977).

Metamorphic grades and the extent of metamorphism are lower in the Triassic sequences than in the 

Permian rocks of the Wallowa terrane. Triassic volcanic rocks range in metamorphic facie s from lower 

greenschist to zeolite, whereas Permian rocks are in 

the lower to middle greenschist fades. Higher temperature fades occur in contact aureoles near the 

Late Triassic and the Jurassic to Cretaceous plutonic 

bodies.

Samples of 19 Triassic igneous rocks of the Wallowa terrane were analyzed for major and minor 

elements (table 3.16), and 14 were analyzed for trace 

elements (table 3.17). The analyses complement 

those of Scheffler (1983) and Goldstrand (1987, 

1994).

Discounting samples V-26-79 (problem with age and 

terrane assignments) and V-10-79 (may be Permian 

in age), the SiO2 contents (n=lT) range from 48.29 

percent to 75.67 percent. TiO2 content is relatively 

high in two basalts (samples VC-265 and VB-85), 

three basaltic andesites (samples VC-135, VS6-5, and 

VC-224), and one andesite (sample VB-105) compared 

to mafic rocks of modern island arcs (Gill, 1981; Kay 

and Kay, 1991). High Fe contents are reflected in relatively high FeOtotal/MgO ratios in most of these volcanic and hypabyssal intrusive rocks (table 3.16 and 

fig. 3.22). A12O3 contents vary greatly; three basaltic 

andesites have A^Os concentrations greater than 18 

percent, probably related to a higher volume of plagioclase in the modes. On the basis of the K2O versus 

SiO2 content diagram (Gill, 1981, p. 6), the rocks are
176 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 3.15. Lithology, sample locality, and field description of analyzed volcanic (flow) rocks and associated hypabyssal intrusive rocks, 

mostly from the Middle and Upper Triassic Wild Sheep Creek and Doyle Creek Formations of the Wallowa terrane

[Rocks arranged by increasing SiO2 content to correspond to chemical analyses in tables 3.16 and 3.17. do., ditto; BM, bench mark]

Sample Rock type Sample locality Field description and remarks

V-26-79 Basalt

VC-265 do 

VB-85 do 

VC-413 do 

VC-335 do 

V-51-79 Basaltic andesite 

VC-269 do 

VC-135 do 

VC-302 do 

VS3-1 do 

VC-224 do 

Along road paralleling Tucannon River; 

elevation 960 m (3,150 ft); NW1/4NW1/4 

sec. 15, T. 9 N., R. 41 E.; Panjab Creek 

quadrangle; Wash.

Ridge north of Kibbles Gulch; elevation 975 

m (3,200 ft); SE1/4 sec. 31, T. 21 N., R. 3. 

W.; Cuprum quadrangle; Idaho.

Idaho Power and Light Company road, 

about 0.8 km south of Hells Canyon 

Dam; elevation 580 m (1,900 ft); eastern 

boundary, sec. 21, T. 22 N., R. 3 W.; 

Cuprum quadrangle; Idaho.

North side of Imnaha River canyon, about 

0.8 km N. 30° E. of Indian Crossing; 

elevation 1,440 m (4,720 ft); unsurveyed 

part of T. 5 S., R. 47 E.; Cornucopia 

quadrangle; Oreg.

North side of Imnaha River canyon, 1.6 km 

N. 70° E. of BM 4521; elevation 1,452 m 

(4,900 ft); unsurveyed part of T. 5 S., R. 

47 E.; Homestead quadrangle; Oreg.

Cottonwood Butte, near radio facility; 

elevation 1,738 m (5,700 ft); NE1/4SW1/4 

sec. 33, T. 32 N., R. 1 W.; Keuterville 

quadrangle; Idaho.

North side of unnamed creek between 

Azurite Gulch and Limepoint Creek; 

elevation 884 m (2,900 ft); 

SE1/4SW1/4NW1/4 sec. 13, T. 20 N., R. 3 

W.; Homestead quadrangle; Idaho.

Northeast flank of Grassy Ridge; elevation 

1,006 m (3,300 ft); NE1/4NW4 sec. 16, T. 

22 N., R. 3 W.; Cuprum quadrangle; 

Idaho.

Near Lynes Saddle; elevation 1,555 m 

(5,100 ft); NW1/4NW1/4 sec . 9, T. 20 N., 

R. 3 W.; Cuprum quadrangle; Idaho.

North side of Doyle Creek near Snake 

River; elevation 497 m (1,630 ft), 

unsurveyed area across river from center 

of sec. 17, T. 21 N., R. 3 W. (Idaho); 

Cuprum quadrangle; Oreg.

Southwest side of Black Point; elevation 

1,723 m (5,650 ft); NE1/4NW1/4 sec. 10, 

T. 21 N., R. 3 W.; Cuprum quadrangle; 

Idaho.

Probably northernmost exposure of 

Wallowa terrane (but possibly Baker 

terrane) rocks. Flow within 

sequence of argillite, cherty argillite, 

and greenstone.

Flow about 5 m thick; may be either 

the Wild Sheep Creek or Doyle 

Creek Formation.

Coarse-grained porphyritic flow; Wild 

Sheep Creek Formation.

Flow (6 m thick) in sequence of flows 

and breccias; Wild Sheep Creek 

Formation.

Flow of unknown thickness; Wild 

Sheep Creek Formation.

Porphyritic lava flow, probably 

correlative with the Wild Sheep 

Creek Formation.

Flow unit (about 10 m thick) in 

sequence of flows; Triassic age, most 

likely part of the Wild Sheep Creek 

Formation.

Flow unit about 3 m thick, Wild Sheep 

Creek Formation.

Flow of unknown thickness; Doyle 

Creek Formation.

Flow unit about 5 m thick; lower part 

of the Doyle Creek Formation or 

upper part of the Wild Sheep Creek 

Formation.

Wild Sheep Creek Formation near 

stratigraphic contact with the 

Permian Hunsaker Creek 

Formation. Thick flow unit among 

alternating flows and breccias.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 177

TABLE 3.15. Lithology, sample locality, and field description of analyzed volcanic (flow) rocks and associated hypabyssal intrusive rocks, 

mostly from the Middle and Upper Triassic Wild Sheep Creek and Doyle Creek Formations of the Wallowa terrane Continued

[Rocks arranged by increasing Si02 content to correspond to chemical analyses in tables 3.16 and 3.17. do., ditto; BM, bench mark]

Sample Rock type Sample locality Field description and remarks

VS6-5

VB-87

VB-103

...... do 

Andesite

VB-105

VS3-6

VC-218

VC-262

...... do -- 

Dacite

...... do------

Rhyolite

V-10-79

North side of McGraw Creek, about 50 m 

below contact with the Martin Bridge 

Limestone; elevation 865 m (2,840 ft); 

unsurveyed area directly across Snake 

River from NW1/4NW1/4 sec. 1, T. 20 N., 

R. 4 W. (Idaho); Homestead quadrangle; 

Oreg.

Idaho Power and Light Company road 

about 0.8 km south of Eagle Bar, parallel 

to Snake River; elevation 613 m (2,010 

ft); SW1/4SE1/4 sec. 28, T. 22 N., R. 3. 

W.; Cuprum quadrangle; Idaho,

North side of Kibbles Gulch, along Idaho 

Power and Light Company road parallel 

to Snake River, about 150 m below 

contact with the Martin Bridge 

Limestone, elevation 576 m (1,890 ft); 

center of NW1/4 sec. 31, T. 21 N., R. 3 W.; 

Cuprum quadrangle; Idaho.

South side of Kibbles Gulch, along Idaho 

Power and Light Company road parallel 

to Snake River; elevation 580 m (1,900 

ft); center of NW1/4 sec. 31, T. 21 N., R. 3 

W.; Cuprum quadrangle; Idaho.

North side of Doyle Creek near Snake 

River; elevation 510 m (1,670 ft); 

unsurveyed area across Snake River 

from center of sec. 17, T. 21 N., R. 3 W. 

(Idaho); Cuprum quadrangle; Oreg.

First spur south of McCraw Creek; 

elevation 854 m (2,800 ft); unsurveyed 

area across Snake River from center of 

sec. 1, T. 20 N., R. 4 W. (Idaho) 

Homestead quadrangle; Oreg.

Along trail that parallels Snake River, 

north of Spring Creek; elevation 585 m 

(1,920 ft); unsurveyed area directly 

across Snake River from NE1/4SE1/4 sec. 

36, T. 21 N., R. 4 W. (Idaho); Homestead 

quadrangle; Oreg.

Along road to Sparta, Oreg., parallel to 

Torchlight Creek; elevation 1,250 m 

(4,100 ft); unsurveyed part of T. 7 S., R. 

44 E.; Sparta quadrangle; Oreg.

Red flow of unknown thickness; Doyle 

Creek Formation.

Clast in flow breccia; Wild Sheep 

Creek Formation.

Red flow about 5 m thick; Doyle Creek 

Formation.

Banded flow about 3 m thick; Doyle 

Creek Formation.

Flow about 8 m thick; Doyle Creek 

Formation.

Porphyritic flow about 6 m thick; Doyle 

Creek Formation.

Dark-brown flow; Doyle Creek 

Formation.

Clover Creek Greenstone; age here 

presumed to be Permian on basis of 

surrounding lithologies.

intermediate-K and low-K lavas. The wide range in 

Na2O and I^O contents may in part reflect metamorphic influences on ion mobility rather than just igneous processes. Correlations of K2O content and 

tectonic setting (for example, Mortimore, 1986) are

suspect where metamorphic effects are not well understood (Vallier and Batiza, 1978).

In these Triassic rocks, trace elements (table 3.17) 

are more helpful than major elements (table 3.16) for 

petrologic interpretations. Sr content in some samples
178 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 3.16. Major- and minor-element oxides of Middle and Late Triassic volcanic (flow) rocks and associated hypabyssal intrusive rocks

of the Wallowa terrane

[Oxides in weight percent, normalized to 100 percent; prenormalization volatile contents included to indicate amount of alteration. , not determined

Sample V-26-79 VC-265 VB-85 VC-413 >335 V-51-79 VC-269 VC-135 VC-302 VS3-1

Si02 - 48.17 48.29 49.84 49.87 50.29 53.09 53.16 54.28

T^i"/"* 1 A K O OO O OO 1 AC 1 AC\ Q1 1 HO 1 T7

Al2O3 15.55 17.48 15.99 17.63 16.48 20.15 18.07 16.97

Fe2O3 1.37 3.85 2.98 4.94 4.47 5.67 4.48

FeO 9.50 11.34 9.61 7.45 8.70 4.54 4.06 5.31

MgO 6.28 3.65 5.25 7.61 5.63 4.06 4.05 3.86

MnO - 18 9f) 99 9f» 99 19 90 99

/ t_/^k 1 o C'? 1 A C\C\ O AO C\ CO Q Q 1 1 C\ OO Q O£? £? OC

Na2O ----- 3.56 1.45 4.46 2.92 2.71 2.21 4.36 5.42

K2O .25 .03 .22 .56 .63 .09 1.05 .90

P205 - ----- .12 .34 .31 .21 .10 .13 .10 .44

H2O+ 3.51 4.74 3.35 4.08 1.89 2.11 2.74 2.00

H f\ 01 oo ocr oo 10 OA 1ft

CO2 .28 .34 .08 .19 .11 .01 .21 .05

FeOtota/MgO - 1.71 3.11 2.49 1.33 2.34 2.11 2.26 2.42

54.30 54.97

QO 00

18.52 16.04

4.44 5.24

4.38 4.84

4.65 5.10

oq 17

7 RQ Q f>Q

3 31 319

1 Qf^ 1^1

.31 .10

3 1 A 9 QH

.24 .34

19 9*;

1.80 1.87

Sample VC-224 VS6-5 VB-87 VB-103 VB-105 VS3-6 VC-218 VC-262 V-10-79

Si02 - 

TiO2 

A12O3 

Fe203 

FeO 

MgO 

MnO 

CaO 

Na2O 

K2O 

P205 

H2O+ - - -----

H2O~ 

f^/~\

FeOtota]/MgO 

55.40

1.90

14 79

11.51

5.07

.27

6.05

3.61

1.18

.22

3.04

.14

3.75

55.48

1.91

14.22

10.16

.42

1.88

.13

14.18

1.25

.06

.31

3.26

.24

.27

5.10

55.56

1.20

1 x. 70

8.58

3.95

.19

7.81

5.77

1.15

_

2.17

56.32

1.17

14.70

8.96

2.57

4.62

.17

7.04

4.11

0.24

.10

2.55

.30

.09

2.30

58.69

1.62

14.65

8.08

2.50

3.01

.25

6.49

3.12

1.38

.21

2.62

.42

1.13

3.24

63.65

1.24

15.56

3.61

3.91

2.27

.14

3.09

6.08

.21

.24

1.80

.17

.03

3.15

67.86

.59

1 a 01

3.03

.32

.20

.03

.81

7.85

2.30

.20

0.47

.17

.15

15.23

75.67

.32

1 3 39

1.09

1.11

.20

.05

.58

6.81

.81

.04

0.54

.14

.29

10.45

78.29

.25

12.50

.90

.66

.62

.54

6.12

.12

0.71

.11

.02

2.37

correlates with CaO content. REE contents are similar to those of many other island-arc lavas. The Cen/ 

Ybn ratios have a wide range (fig. 3.23). Basalt, basaltic andesite, and andesite samples have nearly flat 

REE signatures, whereas dacite and rhyolite samples 

show patterns that are more enriched in LREE's (fig. 

3.25); Cen/Ybn ratios in the more siliceous rocks are 

as large as 3.1 (fig. 3.23). None of the lavas has the 

distinct calc-alkaline REE signature of Aleutian lavas 

(fig. 3.25; Kay and Kay, 1994); they are more similar 

to the tholeiitic lavas from Okmok Caldera in the 

Aleutians and to tholeiitic lavas from the Tonga and 

Kermadec island arcs.

PETROGENESIS OF PERMIAN AND TRIASSIC VOLCANIC (FLOW) 

AND ASSOCIATED HYPABYSSAL INTRUSIVE ROCKS

The Permian and Triassic lavas and associated 

hypabyssal intrusive rocks of the Wallowa terrane 

formed within an island arc. On the basis of criteria 

used in this chapter, most are tholeiitic in composition and are assigned to the island-arc tholeiite series (IAT). They formed in both marine and subaerial 

settings on and near island and seamount volcanoes 

above a subduction zone. LaPierre and others (1988) 

reported both tholeiitic and calc-alkaline Triassic lavas, according to their classification.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 179

TABLE 3.17. Trace elements of selected Middle and Upper Triassic volcanic (flow) rocks and associated hypabyssal intrusive rocks of the

Wallowa terrane

[Results in parts per million. Analysis by neutron activation unless otherwise noted. , not determined or below detection limit or discarded because of high coefficient of variation]

Sample- V-26-79 VB-85 VC-413 VC-335 V-51-79 VC-269 VC-302 VS3-1 VS6-5 VB-103 VB-105 VC-218 VC-262 V-10-79

RLl

Ra 1 

rrn

u 

T a

Ce ----- 

Sm - 

Eu -

Gd - 

TV, ___

Tm - 

Yb

Lu --

7V1

TTf

Ta -

Nb -- 

Ni - 

Co - 

Cr -

Sc - -

Zn - 

c

1Q9

QO

0

7

2.5 

.98

.84

3.4

.54

Q1

9 1

.19

153 

41 

400

41 

95

QQ

.7

93

6.0 

1.65

1 49

.63

n q

7O

48

4.1

6 

39

nn

46 

147

QftQ

179

q

n

19

3.0 

.89

ni

1 O

9O

1 C

97

1 9

38 

174

32 

QO

910

4

1 3

3.5 

1.05

n A

3 O

.47

oc

nq

1.7

45 

91

38

inq

q

4Q8

119

q

q

7

1.5 

.63 

2.0

qq

1.4

91

q

40

q

28 

28 

91

26

QQ

17

1Q4

184

5

14

3.0 

.95

an

3 0

.45

10

CO

9 9

31

45

34 

III

34

A QC

QOQ

1.1

q

19

oq

3.5 

1.1

2 a

QQ

7Q

9 3

.5

19 

qn

19

909

200

4

11

2.7 

.76

c.n

2 a

3Q

90

^u

1

0

1.13 

32

33

38 

110

1 39

10

94

6.2 

1.55 

4.8

QQ

3 3

.45

1 O9

4.0

3.4 

15

28 

4Q

24 

QQ

91 3

4

Q

2.8 

.95

cc

3

0

qn

1 O

.77 

33

1 39

42 

Q9

O£?

214

9O1

1 C

94

5.5 

1.49

1 97

.63 

4 Q

79

49

1 OC

4.0

.54 

31

1C

31

1 QC

OC

Q9

9 9

1 9

1 C

4.3 

1.15 

4.0

nc

9 9

qi

99

117

3 n

5.0

8

19

93

8 

41

OQC

1.7

93

48

6.8 

1.82

1 1 C

n n

QO

48

304

fi Q

3.6 

13

on

12

on

Q4

74

2

0

.7

10

qq

2.8 

.61 

3.6

9 7

QQ

99

3 7

.35

10 

2 

o

7

Analysis by X-ray fluorescence.

The petrogenesis of the Triassic lavas and related 

hypabyssal rocks of the Wallowa terrane is very similar to that of Cenozoic and modern island-arc tholeiitic rocks (Gill, 1981). The petrogenesis of the 

Permian lavas and related dikes and sills, however, 

is puzzling. On the basis of several chemical characteristics (FeOtotal/MgO ratios, REE contents and 

trends, and low-K group chemistry), they can be classified as island-arc tholeiites, and interpretations of 

petrology and tectonic settings can be invoked on the 

basis of the compilation in table 3.2. There are several mineralogical and chemical characteristics, however, that distinguish these Permian rocks from most 

arc lavas.

(1) They are very siliceous; dacite and rhyolite are 

the dominant rock types.

(2) The siliceous lava flows have a phenocryst assemblage of bipyramidal quartz and albite, which 

contrasts greatly with the assemblages of most Cenozoic island-arc lavas. Quartz keratophyre flows and 

dikes with similar field, petrographic, and chemical 

characteristics occur in the Eocene sequences of the 

Komandorsky Islands of the Aleutian island arc (Andrei Tsvetkov, oral commun., 1987).

(3) These rocks are characterized by high 

contents, and most are depleted in K^O. Is this 

chemistry the result of igneous processes only or of 

igneous processes in conjunction with extensive seawater interaction? Is it purely a metamorphic effect? 

Is there a tectonic significance?

(4) Most of these Permian rocks have high FeOtotal/ 

MgO ratios and relatively flat to slightly LREE-enriched REE patterns. These characteristics suggest 

that most rocks result from the fractionation of a tholeiitic basalt parent magma.

(5) In comparison with other volcanic-arc rocks, 

these Permian rocks are particularly enriched in highfield-strength elements; such enrichment is more 

characteristic of alkalic seamount and alkalic oceanisland (Hawaiian Islands, for example) rocks than of 

tholeiitic rocks from Cenozoic island arcs. Calc-alkaline rocks are almost always depleted in these elements. Either a locally atypical (for example, highly 

metasomatized) mantle rock served as source rock, 

partial melting took place at great depths where mantle is not as depleted, significant crustal contamination of the magmas occurred, or these elements were 

concentrated by extreme fractionation of magma in
180 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

high-level chambers that were accessible to seawater 

brine.

(6) On the basis of field relations, mineralogy, and 

chemistry, some of the Permian Na2O-rich plutonic 

bodies apparently are the crystallized magma chambers that fed the Permian lavas and hypabyssal intrusive rocks.

La Ce

RARE-EARTH ELEMENTS 

Sm Eu Tb Tm Yb Lu

EXPLANATION

Triassic basalt

 Sample VC-413

 Sample VC-335 

A Sample VB-85

Triassic basaltic andesite 

D Sample V-51-79 

O Sample VC-302 

A Sample VS6-5

La Ce

RARE-EARTH ELEMENTS 

Sm Eu Gd Tb Tm Yb Lu

EXPLANATION

Triassic andesite Triassic dacite and rhyolite

O Sample VB-103 

D Sample VB-105

Sample VC-218 

Sample V-10-79 

Sample VC-262

FIGURE 3.25. Rare-earth element (REE) diagrams for Triassic lavas and dikes of Wallowa terrane. Patterned area (D) is data 

field for Aleutian calc-alkaline lavas (fig. 3.7), shown for comparison. Sample numbers correspond to those in tables 3.15 

through 3.17; for normalizing values used, see figure 3.6. A, Basalt and basaltic andesite. B, Andesite, dacite, and rhyolite. 

(Rhyolite sample V-10-79 may be Permian.) Note strong lightREE enrichment in samples VC-218 and VC-262; heavy-REE 

contents for those samples, however, are not depleted.

Higher grade metamorphic facies in Permian rocks 

than in Triassic rocks can be attributed to one or 

more of the following: (1) deeper burial, (2) a period 

of higher heat flow during eruption and intrusion of 

Triassic lavas and plutonic bodies, (3) higher temperature regimes over a longer period of time (and two 

or more magmatic events), and (or) (4) higher susceptibility of Permian siliceous glassy rocks to the effects of temperature, pressure, and fluids.

The most likely source for magmas of both ages was 

the mantle wedge; fluids from the downgoing slab 

where oceanic crust and associated sediments were 

being subducted also had some influence on the magma's composition (see Kay, 1984). Partial melting of 

the mantle was succeeded by migration of magma upward into chambers at upper-mantle and (or) lowercrust depths. The fractionation of plagioclase, olivine, 

orthopyroxene (not always present), augite, and magnetite was the major differentiation mechanism, although some magma mixing and crustal assimilation 

probably occurred. The fractionating magmas were at 

relatively low pressures and high temperatures.

TRIASSIC AND EARLY JURASSIC (?) IGNEOUS 

ROCKS OF THE OLDS FERRY TERRANE

Exposures of the Olds Ferry terrane, as denned in 

this chapter, extend from Juniper Mountain west of 

Brogan, Oreg., toward the east and northeast through 

the Cuddy Mountains and Peck Mountain, and possibly as far as Riggins, Idaho (figs. 3.2 and 3.8). Some 

of the rocks along the west border of the Idaho batholith north of McCall, Idaho (Aliberti, 1988; Manduca, 

1988), may belong to the Olds Ferry terrane. The 

Weatherby Formation (Brooks, 1979a) was placed in 

the Olds Ferry terrane by Silberling and others 

(1984). However, in this chapter I place that formation in the Izee terrane.

The terrane is composed mostly of volcanic (flow) 

and volcaniclastic (pyroclastic and epiclastic) rocks of 

the Huntington Formation (Brooks, 1979b). Siliceous 

plutonic bodies occur locally. Volcanic (flow) rocks and 

hypabyssal intrusive rocks range in composition from 

basalt to rhyolite with andesite being the most prevalent. Fossils from the Huntington Formation are of 

late Karnian and early Norian ages. Isotopic ages 

from pre-Late Jurassic plutonic bodies include one 

235-Ma U-Pb age (Walker, 1986) and five K-Ar ages 

(Hendricksen and others, 1972) that range from 216 

(middle Norian) to 190 Ma (Toarcian). The 235-Ma 

isotopic age of the small plutonic body exposed along 

the Brownlee Reservoir (see section "Plutonic Rocks," 

below) overlaps isotopic ages of Triassic rocks in the 

Baker and Wallowa terranes.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 181

TABLE 3.18. Lithology, sample locality, and field description ofTriassic and Early Jurassic(?) plutonic rocks of the Olds Ferry terrane 

[Rocks arranged by increasing SiO2 content to correspond to chemical analyses in tables 3.19 and 3.20]

Sample Rock type Sample locality Field description and remarks

Plutons of northern Cuddy Mountains

V-7-84 Diorite

V-6-84 Quartz diorite

Upper reaches of Hornet Creek near Cuddy 

Mine; elevation 2,225 m (7,300 ft); 

SW1/4NW1/4 sec. 17, T. 17 N., R. 3 W.; 

Hornet quadrangle; Idaho.

Johnson Creek Park near U.S. Forest Service 

quarters; elevation 1,845 m (6,050 ft); 

SW1/4NW1/4 sec. 31, T. 17 N., R. 2 W.; 

Hornet quadrangle; Idaho.

Relatively fresh rock; light 

colored in part.

Relatively fresh rock; light 

colored.

Brownlee pluton

HU-79-1 Trondhjemite- In Snake River canyon, along road from Small silicic pluton. Middle 

Huntington to Richland, at Brownlee Triassic age (Walker, 1986). 

Reservoir; elevation 590 m (1,935 ft); extreme 

west edge of NW1/4 sec. 32, T. 13 S., R. 45 E.; 

Olds Ferry quadrangle; Oreg.

PLUTONIC ROCKS

Three plutonic bodies were sampled for chemistry 

(tables 3.18-3.20). Sample HU-79-1 is from a small 

plutonic body exposed along the road that parallels 

the Brownlee Reservoir, which extends along the 

Snake River southward from Brownlee Dam (fig. 3.1); 

the outcrops are about 5 km north of Huntington, 

Oreg. Samples V-6-84 and V-7-84 are from plutonic 

bodies exposed in the Cuddy Mountains of western 

Idaho.

The plutonic body exposed along the Brownlee Reservoir has the minimum U-Pb age of 235 Ma (Walker, 1986) and is the oldest plutonic body thus far 

recognized in the Olds Ferry terrane. Contact relationships are obscure; Brooks (written commun., 

1989) suggested that it may unconformably underlie 

the Huntington Formation. Furthermore, the plutonic body is older than the late Karnian and early Norian Huntington Formation (Brooks, 1979a) and also 

much older than the isotopically dated plutonic rocks 

in the Cuddy Mountains and in the Mineral-Iron 

Mountain mining district (Hendricksen and others, 

1972).

Five pre-Late Jurassic plutonic bodies exposed in 

the Cuddy Mountains and the Mineral-Iron Mountain mining district, gabbro to trondhjemite in composition, range in age (K-Ar methods on hornblende

and biotite) from about 216 to 190 Ma (Hendricksen 

and others, 1972; Cyrus Field, written commun., 

1988). Four of the five plutonic bodies are significantly younger than the pre-Jurassic plutonic bodies in 

the Baker and Wallowa terranes (fig. 3.21). N.W. 

Walker (written commun., 1989), cautioned, however, 

that the K-Ar ages may have been reset by argon release, which would produce younger apparent ages.

Two Jurassic to Cretaceous plutonic bodies in the 

Olds Ferry terrane, one exposed at Peck Mountain 

and the other at Hitt Mountain, have K-Ar ages of 

Late Jurassic (161 Ma) and Early Cretaceous (120 

Ma), respectively (Cyrus Field, written commun., 

1988). These plutonic bodies were not sampled for 

this study.

There are notable petrographic differences among 

plutonic rocks sampled for this study. The rock sampled from the plutonic body exposed along the 

Brownlee Reservoir (sample HU-79-1) is a metamorphosed trondhjemite; both in hand specimen and in 

thin section the rock is similar to trondhjemites from 

the Oxbow Complex and the (informal) Sparta complex of the Wallowa terrane and from the Canyon 

Mountain Complex of the Baker terrane. Feldspars 

are mostly albitized and hornblende is partly replaced by actinolite and chlorite; white mica is common. The other two samples (V-6-84 and V-7-84) are 

not as metamorphosed as sample HU-79-1. They
182 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 3. 19. Major- and minor-element oxides of Triassic and 

Early Jurassic(?) plutonic rocks of the Olds Ferry terrane

[Oxides in weight percent, normalized to 100 percent; prenormalization volatile contents included to indicate amount of alteration]

Plutons of northern Brownlee 

Cuddy Mountains pluton

Sample V-7-84 V-6-84 HU-79-1

CV/"\ KG QO £O QQ 7/i 9Q

T^/"* Qq R7 q^

Al /~\ 17 7O Id Q7 1Q 9Q

Tf,. f) q 17 9 79 9 ftl

FeO 5.15 3.76 1.36

MgO 3.85 2.64 .94

MnO .14 .13 .09 

CaO - 7.99 6.74 2.20

Na2O 2.76 2.94 3.55 

TCO 14ft 114 188

P2O5 - .15 .11 .06

H o+ 1-19 0.70 1.04

HnO- - » - 06 -03 - 08

t~ir\ no n7 1 7

FeOtotal/MgO 2- 08 2- 35 3- 37

TABLE 3.20. Trace elements of Triassic and Early Jurassic(f) 

plutonic rocks of the Olds Ferry terrane

[Results in parts per million. Analysis by neutron activation unless otherwise 

noted. , not determined or below detection limit or discarded because of high 

coefficient of variation]

Plutons of northern Cuddy Brownlee 

Mountains pluton

Sample ~ V-7-84 V-6-84 HU-79-1

Rb 48 41 26

or qf=;n 9on 1190

Ba 664 640 466

Th 2.65 4.47 1.26

__________ QQ 1 99 a A

La 7.41 9.15 5.2

Ce ----- - 18.00 21.00 14.4 

Sm - 3.00 3.59 3.3 

Eu .91 .86 .78

Gd 3.28 3.93 4.05

Yb 1.96 2.77 4.08

Lu .28 .41 .63

Y OA 1 A A f*

7r 7Q QO 1 TA

TJ-F _ 9 Kf) q A9 Q Q

Co 22 15 15

Pr 9fi 1 q 1 Q

Sc 26 23 19

Zn 87 68 110

1Analysis by X-ray fluorescence.

have 

nant 

const 

dote 

Ma 

comp 

MgO 

for be 

to Cr 

regioi 

late 1 

Iowa 

conte 

zoic i 

Bake 

Sr, ai 

enric 

and "\ 

The

80

75 h- z.

LU

u

£ 70 

0-

h0 

LU 65

 z.

H60

LU

1- z.

8 55

CN

0 

V)

50

45 

C

FIGURE 

vole 

rane 

Ferr 

cani 

toni 

plut 

rese 

som 

198* 

yasr

plagioclase, hornblende, and quartz as domiprimary minerals; orthoclase is an uncommon 

ituent. Metamorphic minerals are mostly epiand chlorite; albite is rare, 

jor- and minor-element contents (table 3.19) are 

atible with an island-arc origin. Although FeOtotal/ 

ratios are higher for these three samples than 

)th the Aleutian intrusive rocks and the Jurassic 

etaceous plutonic rocks of the Blue Mountains 

a (fig. 3.26), they plot within the data fields of 

3aleozoic and Triassic plutonic rocks of the Walterrane (fig. 3.18). TiO2 contents are low. K^O 

nts are generally higher than in the late Paleoind Triassic plutonic rocks of the Wallowa and 

r terranes. Quantities of Zr and Y are low; Rb, 

id Ba contents are high (table 3.20). LREE's are 

led in the Cuddy Mountains samples (V-7-84 

f-6-84) with Cen/Ybn ratios of about 2 (fig. 3.27). 

LREE's are not enriched in sample HU-79-1

i i i i i / 1 i

/'"" ~~~/^ < x-<~-~ o ~~~~^/ ~IL~G--^ -

I KJ / jL \ / /y F

lfj/\ WA y EXPLANATION

/ Jj .\ / Plutonic rocks 

\/'\J '' / ° This report i /r-AX-"' / other

/ E rocks and dikes 

/ i ii ii i i

) 1 2345678 

FeOtota |/MgO RATIO

3.26. FeOtotal/MgO ratio versus SiO2 content for plutonic, 

anic (flow), and hypabyssal (dike) rocks of Olds Ferry ter- 

. Heavy line encloses volcanic rock and dike data for Olds 

y terrane. Data field A represents Aleutian island arc volc rocks; fields E, F, and G represent Aleutian island arc plu- 

: rocks; and field KJ represents Jurassic to Cretaceous 

onic rocks of Blue Mountains region (fig. 3.4). Field WA repnts Triassic flows of Wallowa terrane (fig. 3.22). Includes 

e data (squares) compiled by C.W Field (written comm.,

liro (1974).
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 183

from the plutonic body exposed along the Brownlee 

Reservoir, whose Cen/Ybn ratio is less than 1 (fig. 

3.27). REE patterns have a nearly flat tholeiitic 

trend (fig. 3.28).

Four of the six isotopically dated samples of the 

Olds Ferry terrane intrusive bodies are younger than 

pre-Jurassic plutonic rocks of both the Baker and Wallowa terranes (fig. 3.21). The age of the intrusive body 

exposed along the Brownlee Reservoir (sample HU-79- 

1) is similar to ages of several intrusive rocks from the 

Wallowa and Baker terranes. The intrusive body 

along the Brownlee Reservoir may be basement upon 

which strata of the Huntington Formation were deposited; the plutonic bodies exposed in and near the 

Cuddy Mountains are probably volcano-root plutons.

VOLCANIC (FLOW) AND 

ASSOCIATED HYPABYSSAL INTRUSIVE ROCKS

Lava flows from the Upper Triassic (late Karnian 

and early Norian) Huntington Formation are well excc 

O

O)

Z t

0123456 

Ce n /Yb n RATIO

FIGURE 3.27. FeOtotal/MgO versus Cen/Ybn ratios for plutonic, volcanic (flow), and hypabyssal (dike) rocks of Olds Ferry terrane. 

Dots, Late Triassic flows and dikes; triangles, Triassic to Jurassic plutonic rocks. Heavy line encloses field for Late Triassic volcanic (flow) and hypabyssal (dike) rocks of Olds Ferry terrane. 

Data fields B, C, and D represent Aleutian island arc volcanic 

rocks, and X and Z, Tonga island arc volcanic rocks (fig. 3.5). 

Note overlap of data field (heavy outline) that includes all Late 

Triassic flows with data field (X) for tholeiitic rocks of Ata Island. Line, dashed and queried where uncertain, separates tholeiitic and calc-alkaline fields, n, chondrite normalized.

posed along the road that parallels Brownlee Reservoir in Oregon (table 3.21). Interbedded fossiliferous 

marine sedimentary rocks indicate that most of the 

lavas were erupted in a submarine environment. Red 

and maroon tuffaceous rocks (V-17-85 and V-18-85) 

near the upper part of the exposed section suggest 

nearby subaerial or shallow-marine eruptions, and 

reversed grading of pumice clasts in some beds suggests subaqueous deposition.

Most rocks of the Huntington Formation are regionally metamorphosed to the upper greenschist facies. 

The flows have a wide range in textures, including 

pilotaxitic, intersertal, and intergranular. Primary 

minerals are plagioclase, clinopyroxene, and magnetite. Clinopyroxene is abundant in some lavas. The 

replacement of calcic plagioclase by albite is variable 

but ubiquitous; primary orthoclase is rare. In some 

thin sections clinopyroxene is replaced by chlorite and 

epidote. Besides albite, chlorite, and epidote, white 

mica and calcite are common secondary minerals; less 

common accessory and secondary minerals are 

sphene, leucoxene, pyrite, prehnite, and hematite.

Major- and minor-element oxides indicate a wide 

range of rock compositions for the Huntington

La Ce Nd Sm Eu Gd Tb Yb Lu 

RARE-EARTH ELEMENTS

EXPLANATION

Olds Ferry terrane plutonic rocks 

A Sample V-7-84

 Sample V-6-84

 Sample Hu-79-1

FIGURE 3.28. Rare-earth element (REE) diagram for plutonic 

rocks of Olds Ferry terrane. Patterned area (BM) is data field 

for Jurassic Bald Mountain batholith of Blue Mountains region 

(fig. 3.6S), which has a strong calc-alkaline trend. In contrast, 

plutonic rocks of Olds Ferry terrane are tholeiitic. Sample numbers correspond to those in tables 3.18 through 3.20; for normalizing values used, see figure 3.6.
184 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 3.21. Lithology, sample locality, and field description of Late Triassic volcanic (flow) and hypabyssal (dike) rocks of the Olds

Ferry terrane, collected along Brownlee Reservoir road, eastern Oregon

[Rocks arranged by increasing Si02 content to correspond to chemical analyses in tables 3.22 and 3.23. Do. and do., ditto; BM, bench mark]

Sample Rock type Sample locality Field description and remarks

VB-62 Diabase Along road that parallels Brownlee Dike cutting trondhjemite (see

Reservoir on the west side; NW1/4 sec. 32, HU-79-1, tables 3.18-3.20). 

T. 13 S., R. 45 E.; Olds Ferry quadrangle, 

Oreg.

VB-56 Basalt Near bridge that crosses Burnt River east of Flow rock; Huntington

Huntington, Oreg.; NE1/4SE1/4 sec. 8, T. Formation. 

14 S., R. 45 E.; Olds Ferry quadrangle, 

Oreg.

V-13-85 do Near confluence of Burnt and Snake Rivers; Do.

SE1/4NE1/4 sec. 8, T. 14 S., R. 45 E.; Olds 

Ferry quadrangle, Oreg.

V-14-85 Basaltic andesite North of confluence of Burnt and Snake Do.

Rivers; SE1/4NE1/4 sec. 8, T. 14 S., R. 45 

E.; Olds Ferry quadrangle, Oreg.

V-16-85 do Near BM 2017; NE1/4SW1/4 sec. 16, T. 13 Do.

S., R. 45 E.; Olds Ferry quadrangle, Oreg.

V-12-85 Andesite Near confluence of Burnt and Snake Rivers; Do.

SE1/4NE1/4 sec. 8, T. 14 S, R. 45 E.; Olds 

Ferry quadrangle, Oreg.

VB-59 do 0.8 km east of bridge that crosses Burnt Do.

River; NE1/4SE1/4 sec. 8, T. 14 S., R. 45 

E.; Olds Ferry quadrangle, Oreg.

V-15-85 do Near BM 2022; SE1/4SE1/4NW1/4 sec. 21, Do.

T. 13 S., R. 45 E.; Olds Ferry quadrangle, 

Oreg.

V-17-85 Dacite Near BM 2006; NE1/4SW1/4 sec. 9, T. 13 S., Clast in tuff breccia of the

R. 45 E.; Olds Ferry quadrangle, Oreg. Huntington Formation.

V-18-85 Rhyolite Near BM 2006 and about 30 m north of site Do.

for V-17-85; NE1/4SW1/4 sec. 9, T. 13 S., 

R. 45 E.; Olds Ferry quadrangle, Oreg.

Formation (table 3.22), including basalt, basaltic andesite, and andesite as flows, and dacite and rhyolite 

as clasts (in tuff breccias). The flows have low TiO2 

contents; A12O3 concentrations are high, mostly reflecting the amount of plagioclase. Fe enrichment is 

characteristic of the more mafic lavas; the plot of 

FeOtotal/MgO ratios versus SiC>2 contents indicates 

that these rocks are tholeiitic and transitional (fig. 

3.26). K^O contents are high compared to those of 

most island-arc lavas (Gill, 1981). Na2O contents are 

variable and reflect albitization of plagioclase, particularly in the more mafic rocks, in which Na2O ranges 

from 3.19 to 5.41 percent; these are high values compared to those for unaltered mafic lavas from island 

arcs (Gill, 1981).

Trace-element contents may be more useful for interpreting the petrogenesis of these rocks and the extent of ion mobility during metamorphism (table 

3.23). Rb, Sr, Ba, Cs, Th, and U contents can be correlated with K^O content, strengthening the interpretation that high amounts of these trace elements as 

well as the high K^O concentrations are related to 

magmatic rather than metamorphic processes. Zr and 

Y contents are low and correlate with the low abundance of TiO2. The high concentrations of K, Rb, Sr, 

and Ba in sample VB-56 and the high Rb content of 

sample V-13-85, however, are perplexing; thin-section 

analyses suggest that the K^O is in the plagioclase; 

no orthoclase phenocrysts or xenocrysts were observed. The FeOtotal/MgO versus Cen/Ybn ratios and 

REE patterns (figs. 3.27 and 3.29) indicate that most 

of the rocks are tholeiitic. The similarities with plots 

of Ata Island (Tonga) flows are noteworthy, implying a 

somewhat similar island-arc tholeiite petrogenesis.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 185

TABLE 3.22. Major- and minor-element oxides of Late Triassic volcanic (flow) and hypabyssal (dike) rocks of the Olds Ferry terrane

[Oxides in weight percent, normalized to 100 percent; prenormalization volatile contents included to indicate amount of alteration]

CJiO

TiO2 

Fe203 - - 

FeO 

MnO 

PnO

Na2O -

V d

25

H2O+ 

H20- -----

PO

FeOtotal/MgO 

VB-62

49.91 

.86 

17.32 

5.09 

6.63 

6.41 

.27 

8.12 

5.09 

.24 

.06

3.33 

.42 

3.16 

1.75

VB-56 V-13-85 V-14-85 V-16-85 V-12-85

50.10 

.77 

19.78 

3.35 

6.95 

6.35 

.23 

7.04 

3.19 

2.20 

.04

3.99 

.66 

.08 

1.57

50.82 52.65 

.81 .76 

21.90 20.10 

3.82 3.19 

5.32 6.93 

4.35 5.90 

.24 .32 

6.90 3.84 

4.54 5.41 

1.23 .82 

.07 .08

4.37 4.11 

.29 .23 

.29 .06 

2.02 1.66

53.44 

1.01 

19.60 

3.05 

5.71 

5.47 

.16 

4.44 

4.73 

2.23 

.16

3.28 

.52 

.56 

1.54

56.41 

.93 

17.80 

3.19 

7.09 

5.42 

.32 

2.12 

5.75 

.87 

.10

3.58 

.29 

1.12 

1.84

VB-59

57.75 

.93 

17.69 

3.65 

5.96 

4.95 

.22 

1.87 

6.00 

.91 

.07

3.31

.48 

.87 

1.87

V-15-85

59.24 

1.26 

14.47 

5.50 

5.98 

4.16 

.38 

4.12 

4.67 

.48 

.19

2.52 

.27 

.15 

2.53

V-17-85 V- 18-85

65.08 

.87 

16.27 

2.65 

2.82 

1.35 

.15 

2.21 

7.68 

.71 

.21

0.81 

.20 

4.03 

3.85

72.06 

.67 

14.46 

4.18 

.16 

.18 

.04 

.46 

5.33 

2.24 

.22

1.13 

.23 

.09 

21.80

TABLE 3.23. Trace elements of Late Triassic volcanic (flow) and hypabyssal (dike) rocks of the Olds Ferry terrane

[Results in parts per million. Analysis by neutron activation unless otherwise noted. , not determined or below detection limit or discarded because of high coefficient of variation]

Sample - VB-62

TJK Ifi

QY» ll q/l

Ro 17O

p«

Th 

U .

T Q _ A

XJCL *±

Qe ________ g

Sm - 1.9

T?n KQ

f< H _______

Th _____ 4*i

m__

Yb 1.6

Lu .25

Y ^18

Hf i.o

TnJL d. 

Nb 

Co 36

Pr 90

Zn 130

VB-56

X 58 . X 218 

*397

.3

2 

5 

1.4

.32

.21

X 25

.7

28 

20 

36 

114

V-13-85

23 

2 120 

193 

.91 

.22 

.16 

1.74 

4.93 

1.46 

.63

.38

1.46 

.23

37 

.72

23 

19 

34 

125

V-14-85 V-16-85

59 

2 160 2340 

151 689 

2.16 16.90 

.19 1.87 

.20 .63 

1.71 6.63 

5.42 14.70 

1.54 2.65 

.76 .96 

2.07 3.30 

.39 .61 

.27 .40 

1.75 2.50 

.26 .36 

2 10 220 

230 67 

.80 1.78 

.03 .11

26 23 

17 20 

34 34 

182 90

V-12-85

17 

2 120 

214 

1.33 

.28 

.18 

2.51 

6.23 

1.95 

.91 

2.62 

.52 

.34 

2.20 

.35 

2 15 

58 

1.03 

.03

22 

50 

40 

184

VB-59

*9 

X 164

3 

6 

2.0

.59 

.31

.35

X 39 

1.0

21

40

V-15-85

15 

2 150 

254 

1.47 

.65 

.36 

4.55 

11.30 

3.00 

1.20 

4.23

.48 

3.05 

.45 

2 20 

94 

1.75 

.08

23 

4 

42 

216

V-17-85

2 150 

220 

.63 

2.41 

.65 

6.83 

17.10 

3.09 

1.22 

4.71 

.87 

.61 

3.68 

.55 225 

92 

2.42 

.16

7 

24

183

V-18-85

46 2 12 

637 

4.42 

2.91 

.89 

14.30 

34.00 

5.30 

1.59 

6.08 

1.07 

.74 

4.55 

.68 

235 

114 

3.54 

.18

3 

5 

17 

361

Analysis by X-ray fluorescence. 

Analysis by optical spectroscopy.

PETROGENESIS OF IGNEOUS ROCKS 

OF THE OLDS FERRY TERRANE

The chemically analyzed igneous rocks from the 

Olds Ferry terrane formed in an intraoceanic island

arc. The major characteristics of Olds Ferry lavas are 

the ubiquitous presence of plagioclase and clinopyroxene in the modes, Fe enrichment, flat REE patterns, low quantities of high-field-strength elements,
186 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

and intermediate but variable amounts of K-group elements. These rocks have higher contents of the Kgroup elements than most igneous rocks in the Blue 

Mountains island arc older than the Jurassic to Cretaceous intrusive rocks.

K2<D contents are particularly interesting because 

Mortimore (1986) used them to conclude that the 

Olds Ferry rocks were part of a high-K arc that bordered the North American continent. On the basis of 

a plot of K2O versus SiO2 contents (see Gill, 1981, p. 

6), the rocks are high- to intermediate-K basalt and 

basaltic andesite, and intermediate- to low-K andesite. Relatively high K^O content, in comparison with 

that of most Cenozoic tholeiitic island-arc lavas, 

could be explained by one or more of the following: 

(1) It may be related to a large amount of sediments 

that were added to the magma source region above a 

subduction zone, particularly if the sediments had 

been eroded from a continent; (2) it may be related to 

tectonic setting, such as the depth to the dipping 

subduction zone (Gill, 1981); (3) it may merely reflect 

abnormal fractionation of magma beneath a single 

volcano and (or) K mobility during metamorphism.

Parent magmas probably were derived from partial 

melting of a depleted mantle source above a dipping

La Ce Sm Eu Gd Tb 

RARE-EARTH ELEMENTS

Yb Lu

EXPLANATION

Upper Triassic rocks

Sample VB-56 0 Sample V-16-85 

Sample V-13-85 Sample V-15-85 

Sample VB-62 o Sample V-17-85 

Sample VB-59 Sample V-18-85

FIGURE 3.29. Rare-earth element (REE) diagram for volcanic 

(flow) and hypabyssal (dike) rocks of Olds Ferry terrane; these 

rocks are strongly tholeiitic. Patterned area (X) is data field for 

tholeiitic rocks of Ata Island (fig. 3.7), shown for comparison. 

Most samples are from Upper Triassic Huntington Formation of 

Brooks (1979a, b); sample numbers correspond to those in tables 

3.21 through 3.23. For normalizing values used, see figure 3.6. 

Missing data points for Gd indicate that element was not determined, contents were below detection limit, or data were discarded because of high coefficient of variation.

subduction zone. The fractionation of plagioclase, olivine and (or) orthopyroxene, augite, and magnetite 

must have been the dominant differentiation process, 

although assimilation of older arc crust and magma 

mixing cannot be ruled out. The magmas differentiated in relatively high-temperature and low-pressure 

environments.

JURASSIC IGNEOUS ROCKS ASSOCIATED WITH

THE COON HOLLOW FORMATION OF THE

WALLOWA TERRANE

I sampled dikes and sills within the Middle and 

Upper Jurassic Coon Hollow Formation of the Wallowa terrane (Morrison, 1963; White, 1972; White, 

1985; Goldstrand, 1987, 1994; White and Vallier, 

1994; White, 1994) and a partially welded tuff that is 

interbedded with the oldest strata in that formation 

(table 3.24). Middle (Bajocian and Callovian) and 

Late (Oxfordian) Jurassic marine fossils from the 

Coon Hollow Formation were identified by Imlay 

(1986) and Stanley and Beauvais (1990).

DESCRIPTION OF THE IGNEOUS ROCKS

The diabase (sample V-l-85, table 3.24) is metamorphosed to the upper greenschist facies; abundant 

chlorite and epidote have replaced clinopyroxene and 

groundmass glass, and some albite has replaced calcic plagioclase. The microdiorite (sample V-58-79) has 

undergone very low grade metamorphism; most primary minerals are unaltered. Estimated modal mineralogy of sample V-58-79 is plagioclase, 50 percent; 

hornblende, 30 percent; augite, 10 percent; Fe-oxide 

minerals, 1 percent; and epidote+chlorite, 9 percent. 

White mica is rare. In the metamorphosed andesite 

porphyry (sample T-68-16), albite has replaced calcic 

plagioclase, and chlorite and epidote have replaced 

the mafic minerals. The andesite porphyry occurs in 

several places within the Coon Hollow Formation in 

the Pittsburg Landing area.

The rhyolite tuff (sample V-ll-85), described in 

more detail by White and Vallier (1994), crops out 

over a very small area and is associated with abundant alluvial-fan channel deposits of cobble and pebble conglomerate. In the more welded parts of the 

tuff, unaltered feldspar and lithic fragments are set 

in a eutaxitic matrix of altered glass and pumice.

These igneous rocks have low amounts of TiO2 and, 

except for sample V-ll-85, relatively low FeOtotal/ 

MgO ratios (table 3.25). K2O contents are low for 

samples V-l-85, T-68-16, and V-ll-85, and relatively
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 187

TABLE 3.24. Lithology, sample locality, and field description of Jurassic hypabyssal intrusive rocks and a welded tuff interbedded in

lowest part of the Coon Hollow Formation, Wallowa terrane

[Rocks arranged by increasing SiO2 content to correspond to chemical analyses in tables 3.25 and 3.26]

Sample Rock type Sample locality Field description and remarks

V-58-79 Microdiorite

V-l-85 Diabase

T-68-16 Andesite porphyryV-11-85 RhyoliteNorth side Coon Hollow Creek; elevation 

365 ip (1,200 ft); SE1/4SE1/4 sec. 26, T. 

6 N., R. 47 E.; Jim Creek Butte 

quadrangle; Oreg.

East side, north fork of West Creek; 

elevation 870 m (2,850 ft); SW1/4NE1/4 

sec. 21, T. 27 N., R. 1 W; Grave Point 

quadrangle; Idaho.

Along road that parallels Kurry Creek 

near Pittsburg Landing; elevation 518 

m (1,700 ft); SE1/4 sec. 28, T. 27 N., R. 1 

W.; Grave Point quadrangle; Idaho.

Ridge south of West Creek; elevation 457 

m (1,500 ft); extreme NW1/4NW1/4 sec. 

27, T. 27 N., R. 1 W.; Grave Point 

quadrangle; Idaho.

Sill that cuts the Coon Hollow 

Formation.

Large dike that cuts the Coon 

Hollow Formation.

Irregular pluton that cuts the 

Coon Hollow Formation.

Welded tuff; near base of the Coon 

Hollow Formation. Prominent 

ledge.

TABLE 3.25. Major- and minor-element oxides of Jurassic hypabyssal intrusive rocks and a welded tuff interbedded in lowest 

part of the Coon Hollow Formation, Wallowa terrane

[Oxides in weight percent, normalized to 100 percent; prenormalization volatile contents 

included to indicate amount of alteration]

Sample - V-58-79 V-l-85 T-68-16 V-ll-85

SiO2 - 

Al Du

Fe2O3 -

£1 O

TTgO _, __ __ _ 

MtrO

MnO 

CaO - 

Na2O -

P2^O5 ~

H n+

H

Cf~ 2(j- - - 

FeO/MgO

54.74 

1.02 

16.79 

2.13 

5.10 

6.73 

.10 

7.71 

3.90 

1.41 

.37

2.51 

.21 

.12 

1.04

55.02 

.96 

16.54 

1.51 

6.95 

6.66 

.11 

6.38 

5.30 

.35 

.22

3.85 

.27 

3.53 

1.25

60.15 

.72 

17.81 

.96 

4.11 

4.24 

.06 

4.37 

6.79 

.43 

.26

2.20 

.48 

1.80 

1.17

74.14 

.43 

12.20 

1.07 

4.01 

1.53 

.07 

2.58 

3.76 

.15 

.06

2.26 

.42 

2.08 

3.25

high for sample V-58-79. The plot of FeOtotal/MgO ratios versus SiC>2 contents (fig. 3.30) clearly shows the 

similarities between samples V- 58-79, V-l-85, and T68-16 and the large Jurassic to Cretaceous plutons of 

the Blue Mountains region; the tuff sample is more 

Fe enriched and lies outside the data field for Jurassic to Cretaceous plutons.

Trace-element contents show strong enrichments in 

Rb, Sr, and Ba for samples V-58-79 and T-68-16, but 

lower contents were found in the other samples (table 

3.26). Th and U contents are consistently high in all 

four samples. Zr and Y are depleted in parallel with 

the depletion of TiO2- Cen/Ybn ratios of three of the 

four samples are within or close to the data field for 

the Bald Mountain batholith (fig. 3.31); the tuff (sample V-ll-85) has a low ratio and plots within the 

tholeiitic field. The REE diagram (fig. 3.32) shows 

calc-alkaline affinities for all samples except the tuff, 

which has a flat tholeiitic pattern.

PETROGENESIS

The composition, size, degree of metamorphism, 

and field relationships separate the small intrusive 

bodies associated with the Coon Hollow Formation 

from the large Jurassic to Cretaceous plutons of the 

Blue Mountains region. The strongly calc-alkaline 

character of the intrusive rocks has several petrologic 

implications (table 3.2): fractionation of amphibole 

(and garnet?), abundant water in the magma, greater 

mixing in the source region, and a higher pressure 

(deeper) fractionation regime. These possibilities in 

turn have tectonic implications, including slower convergence rates, a shallower dip of the subducting 

slab, and possibly a thicker crust. The important 

point is that there must have been a change in tectonic setting within the Wallowa terrane between the
188 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

80

75

70

60

55

50

45

0 1 234567 

FeOtota |/MgO RATIO

FIGURE 3.30. FeO^^/MgO ratio versus SiO2 content for hypabyssal rocks (dikes and plugs) that cut strata of Coon Hollow 

Formation and for a welded tuff interbedded in lowest part of 

Coon Hollow Formation. Note that dikes and plugs (triangles) 

lie entirely within area (KJ) representing Jurassic to Cretaceous (calc-alkaline) plutons of Blue Mountains region (fig. 3.4) 

and that welded tuff (dot) is more Fe-enriched. Calc-alkalinetholeiitic compositional boundary from Miyashiro (1974).

eruption of the youngest (Karnian) flows of the Doyle 

Creek Formation and the intrusion of the small intrusive bodies (Bajocian or younger in age at Pittsburg Landing and Oxfordian or younger in age in the 

northern Snake River canyon).

The presence of a low-K rhyolite tuff in the lower 

part of the Coon Hollow Formation may signal the 

last gasp of a tholeiitic magmatic regime before a 

change in tectonic setting led not only to the calcalkaline hypabyssal intrusive rocks within the Coon 

Hollow Formation but also to the voluminous Jurassic to Cretaceous calc-alkaline plutons of the entire 

Blue Mountains region.

JURASSIC TO CRETACEOUS PLUTONIC ROCKS 

OF BLUE MOUNTAINS REGION

The Jurassic to Cretaceous plutons are only briefly 

discussed in this chapter. Taubeneck (1957, 1967,

TABLE 3.26. Trace elements of Jurassic hypabyssal intrusive rocks 

and a welded tuff interbedded in lowest part of the Coon Hollow 

Formation, Wallowa terrane

[Results in parts per million. Analysis by neutron activation unless otherwise noted. , not 

determined or below detection limit or discarded because of high coefficient of variation]

Sample V-58-79 V-l-85 T-68-16 V-ll-85

Rh

Sr ----- -

Ba -

Pa

rrn

TT __ .

T a __

Sm -

TTn

/"I j

Tb - 

Yb --

Lu- -

Y - 

Zr -

TJf

To ____

loq

IQGA

Inqc

1.70

en

Ifi

o e

1 91

.41

__ i 0

90

7

9 9

__ 4

q

29QO

oc

1 09

1 19

cc

9

nq

oo en

3 no

1 Oft

3 01

.44

99

1.40

1ft

10

2on

9 09

9R

loq

lOftn

1 90

C1

n

c

91 7

9 7

7q

.10

69

.11

lQn

1 Q

1Q

2 170

inc

qc

1 9Q

nq

4.56

11.10

9 9ft

QC

2 QC.

cc

.40

9 7O

.41

90

114

3 Oft

90

Co -

Pr

Sc- -

Zn -

9Q

ooc

90

7c

qo

9O7

94

no

1ft

1 ^

44

q

94

iq

ei

1Analysis by X-ray fluorescence. 

Analysis by optical spectroscopy.

chap. 2 of this volume and references cited therein) is 

most familiar with the field relations and petrology of 

this large group of rocks. Armstrong and others 

(1977) compiled and discussed a significant number 

of isotopic (mostly K-Ar) ages and Sr-isotope data for 

Jurassic to Cretaceous plutons in the Blue Mountains region. They concluded that most of these plutons crystallized in the interval from 160 to 115 Ma. 

However, on the basis of U-Pb isotopic age dating of 

zircons, Walker (1989; oral commun., 1990) concluded 

that almost all of the plutons were emplaced within 

the interval from 145 to 120 Ma.

In this section, I describe several small Jurassic to 

Cretaceous plutonic bodies that are widely separated 

geographically from the Wallowa and Bald Mountain 

batholiths. Major- and trace-element data, four initial Sr-isotope ratios, and two new K-Ar isotopic ages 

are given. These data certainly are not sufficient for 

thorough study of all Jurassic to Cretaceous plutons 

in the Blue Mountains region, but the data warrant 

some discussion of the petrology and tectonic setting 

of these plutons.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 189

4 6 

Ce n/Yb n RATIO

EXPLANATION

 Jurassic to Cretaceous plutonic rocks (this report)

A Dikes Cut Jurassic Coon Hollow Formation

 Jurassic Coon Hollow Formation (welded tuff)

10

FIGURE 3.31. FeOtotai/MgO versus Cen/Ybn ratios for Jurassic and 

Cretaceous rocks in Blue Mountains region. Line, dashed and 

queried where uncertain, separates tholeiitic and calc-alkaline 

fields. Note that among Jurassic and Cretaceous rocks, only tuff 

in Coon Hollow Formation shows tholeiitic affinity. Data field

KJ corresponds only to those Jurassic to Cretaceous plutonic rocks 

analyzed for this chapter (see table 3.29). For comparison, plot 

also includes Cougar Creek Complex of Vallier (1968) (data field 

CC; fig. 3.19), Canyon Mountain Complex (CM; fig. 3.10), and Bald 

Mountain batholith (BM; fig. 3.5). n, chondrite normalized.

FIGURE 3.32. Rare-earth element (REE) diagram for Jurassic 

welded tuff of Coon Hollow Formation (sample V-11-85) and for 

sill and dike rocks that cut Coon Hollow Formation (samples T16-68, V-58-79, and V-l-85). Patterned area (BM) is data field 

for Late Jurassic Bald Mountain batholith of Blue Mountains 

region (fig. 3.6B), shown for comparison. Note flat pattern for 

sample V-ll-85. Sample numbers correspond to those in tables 

3.24 through 3.26; for normalizing values used, see figure 3.6. 

Missing Tb data point indicates no analysis, content below detection limit, or large coefficient of variation.

La Ce Sm Eu Tb 

RARE-EARTH ELEMENTS

Yb Lu

EXPLANATION

n Sample T-16-68 A Sample V-l-85 

o Sample V-58-79 Sample V-ll-85
190 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 3.27. Lithology, sample locality, and field description of analyzed Jurassic to Cretaceous plutonic rocks of the Blue Mountains region

[Rocks arranged by increasing SiO2 content to correspond to chemical analyses in tables 3.29 and 3.30. Do. and do., ditto]

Sample Rock type Sample locality Field description and remarks

V-12-79 Gabbro - Cutsforth County Park along stream near Near contact with granodiorite

small lake; elevation 1,275 m (4,185 ft); T. 4 (sample V-ll-79). 

S., R. 28 E.; Arbuckle Mountain quadrangle; 

Oreg.

V-24-79 Diorite Carney Butte north of lookout tower; elevation Melanocratic pluton; some

1,503 m (4,930 ft); NE1/4NW1/4 sec. 28, T. 3 cumulate texture. 

S., R. 31 E.; Carney Butte quadrangle; Oreg.

V-23-79 ....do do Do.

V-54B-79 do Along trail, west side of Seven Devils Melanocratic pluton that extends

Mountains south of Baldy Lake; elevation irregularly along west side of 

2,335 m (7,660 ft); unsurveyed part of T. 23 Seven Devils Mountains. 

N., R. 2 W.; He Devil quadrangle; Idaho.

V-46-79 Quartz diorite Outcrop along road near entrance; Winchester Small inlier within the Columbia

State Park; elevation 1,175 m (3,850 ft); T. River Basalt Group. 

33 N., R. 2 W; Winchester East quadrangle; 

Idaho.

V-ll-79 Granodiorite Cutforth County Park along stream near Poor outcrops, near contact with

small lake; elevation 1,280 m (4,200 ft); T. 4 gabbro (sample V-12-79). 

S., R. 28 E.; Arbuckle Mountain quadrangle; 

Oreg.

V-75-79 do Blacktail stock, along road east of Mt. Idaho; Pluton that apparently cuts both

elevation 745 m (2,445 ft); SE1/4SE1/4 sec. the Seven Devils Group and the 

26, T. 30 N., R. 3 E.; Grangeville East Riggins Group, 

quadrangle; Idaho.

VC-295 Granite Small ridge south of Eckels Creek; elevation Small pluton; intrudes rocks of the

1,396 m (4,580 ft); NW1/4NE1/4 sec. 4, T. 20 Hunsaker Creek Formation. 

N., R. 3 W.; Cuprum quadrangle; Idaho.

DESCRIPTION OF THE PLUTONIC ROCKS

The Jurassic to Cretaceous plutonic rocks sampled 

(table 3.27) represent a wide range of rock types. 

Ages of most are unknown, but field relationships 

and compositions, including the small amount of alteration, were used to assign them Jurassic to Cretaceous ages. Sample V-75-79 from the Blacktail stock, 

which is along the South Fork of the Clearwater River in Idaho (fig. 3.1), has been isotopically dated at 

78 Ma by K-Ar methods (Criss and Fleck, 1987), and 

the plutonic body at Winchester State Park in Idaho 

(sample V-46-79) has a K-Ar hornblende age of about 

141 Ma (same as sample 719 of Armstrong and others, 1977). Sample VC-295 is a two-mica granite that 

cuts Permian rocks in the upper reaches of Eckels 

Creek in Idaho south of Hells Canyon Dam (fig. 3.3). 

The absence of granite from known late Paleozoic 

and Trias sic intrusive bodies in the Blue Mountains

region and the lack of appreciable alteration suggest 

a Jurassic to Cretaceous age.

Two previously unreported K-Ar radiometric ages 

(samples V-47-79 and V-3-86, table 3.28), complement 

the age data already available for the Jurassic to 

Cretaceous plutons of eastern Oregon and western 

Idaho (Armstrong and others, 1977; Goldstrand, 

1987). Armstrong and others (1977) cautioned that 

many of the ages are minima; for example, they stated that 136 Ma may be the minimum age for all or 

nearly all of the granitic plutons within the Wallowa 

Mountains (Armstrong and others, 1977, p. 402). In 

contrast, N.W. Walker (oral commun., 1989) stated 

that his U-Pb isotopic age data indicate that 137 Ma 

(for his Pole Bridge unit) is the maximum age determined for the Wallowa batholith.

The consistent K-Ar isotopic ages (about 115 Ma) 

on both biotite and hornblende mineral separates 

from a diorite exposed along Little Granite Creek
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 191

TABLE 3.28. K-Ar ages of selected Jurassic to Cretaceous plutonic rocks of the Blue Mountains region

[Apparent ages, calculated using two or more replicate determinations on each mineral separate, listed as averages of replicate data (weighted hy inverse of variance) or as 

minimum age. Error values for ages are estimates of the standard deviation of analytical precision. Half-lives used in calculations: A£=O.581x10" /yr and Ayrj=4.96x10 /yr. 

Abundance ratio: 40K/K=1.167xlO~4 atom percent, rad, radiogenic; do., ditto. Analysts, L.B. Gray and J.E. Pearl, U.S. Geological Survey]

Apparent age (Ma)

Sample

7-18H

V-47-79

V-3-86

Locality

Intersection of Dry and Cook

Creeks, unsurveyed part,

southeast corner of Jim Creek

Butte quadrangle, Oreg.

Extreme southeast corner of

Washington along Snake

River; SE 1/4SW 1/4 sec. 9,

T. 6 N. R. 47 E.; Lime Kiln

Rapids quadrangle; Wash.

Little Granite Creek, along

trail; elevation 1,160 m (3,800

ft); unsurveyed part, T. 23 N.,

R. 2 W.; He Devil quadrangle;

Idaho.

Rock Mineral 

type

Diorite Biotite 

- do.

Quartz Biotite - 

diorite. do. 

 do _______

 do. Biotite -------

.. do. -----

Hornblende - 

.. do. --- 

KaO

(weight 

percent)

9.00

8.93

7.24 -]

7.16 I

6.86 J

9.31

9.32

.412 1

.401 J

40 Ar^adj 

(moles per 

gram)

1.866xlO'9

1.879xlO'9

1.417xlO'9

(average)

1.599xlO'9

1.601xlO'9

6.967xlO'9

(average)

4!Nd/

40Artotal

0.979

.902

.897

.829

.797

.584

(average)

Replicate 

determination

139.1 1

140.0 J

None -----

115.5 "]

115.6

>

115.3

(average) J

Average or 

minimum age

139.5+2.1

(average)

133.814.0

(minimum)

115.3+2.5

(average)

^ry Creek stock.

TABLE 3.29. Major- and minor-element oxides of Jurassic to Cretaceous plutonic rocks of the Blue Mountains region

[Oxides in weight percent, normalized to 100 percent; prenormalization volatile contents included to indicate amount of alteration. , not determined]

Sample V-12-79 V-24-79 V-54B-79 V-46-79 V-ll-79 V-75-79 VC-295

SiO2 

TiO -

Al20g - 

Fe2Oo 

L* O

FpO

MffCi

MnO - - 

faO

Na2O 

IT" r\

P2(Jg f) 

H f)+

H

c\-

CO2 . ._ _._______.

FeOtotal/MgO ---

47.07

0 AQ

15.69

Q 00

11.09

K 7fi

90

in "¥)

3

0Q

AK

1.30

.13

.01

2.44

CO KQ

01

7.46

1.70

10 94

15.19

00

9 98

1.14

.17

0.57

19

.01

.78

54.41

on

1Q ^.Q

0 O9

413

A. tt

.14

8 7fi

3

QQ

CK

.17

1.53

.10

1.51

64.43

.59

i £ af\

i af\

0 R7

9 fiK.

.07

K I/

4.03

1 Q£

1 C

0.65

09

.01 

1.63

66.54

OK

i a QO

1.76

9 °,°,

1 <V7

1 ^

C CO

3.18

1.45

IK

1.05

.07

.01 

2.49

68.52

.28

17.43

1.43

Qfi

.86

.06

4 ^9

4.45

1 °,7

19

0.55

.11

2.61

75.57

.05

14.89

.58

94

.23

.85

4.35

3.14

.10

0.35

.31

1.02 

3.31

(sample V-3-86, table 3.28) are the youngest thus far 

reported for granitoid plutonic bodies in the Hells 

Canyon-Seven Devils Mountains region. This diorite 

crops out less than 16 km from the 133-Ma Deep 

Creek stock (White, 1973; Armstrong and others, 

1977).

A gabbro and a diorite (samples V-12-79 and V-24- 

79, table 3.27) were collected from layered sequences 

and probably are crystal cumulates. Augite, hypersthene, hornblende, Fe-oxide minerals, and plagioclase are present in both samples. The remaining 

plutonic rocks sampled have hypidiomorphic-granular textures and have pyroxene, hornblende, and biotite as the primary mafic phases and plagioclase, 

quartz, and orthoclase as the dominant felsic phases. 

Muscovite occurs in samples V-75-79 and VC-295.

Some major-element contents (table 3.29) closely reflect mineral modes. For example, high MgO contents
192 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 3.30. Trace elements of Jurassic to Cretaceous plutonic rocks of the Blue Mountains region

[Results in parts per million. Analysis by neutron activation unless otherwise noted. , not determined or below detection limit or discarded

because of high coefficient of variation]

Sample - V-12-79 V-24-79 V-54B-79 V-46-79 V-11-79 V-75-79 VC-295

80

75

70

I 

o

m 65

60

55

50

45

Rb1 

qrl

Rfl 1

Th -

TT _ ..

La 

Ce 

O*Vl

Gd

Tb 

Tm - 

Yh _ _

Lu- -

Yi ............

Zr1 -----

prf ____

Nb1 ----- -

Ni1

Co 

Pr

Sc - - 

7n

9

901

10

1

9

1.1

Q9

O A

.16

1 C

19

oc

q9

qo

CO

qo

qo

c

919

78

.5

q

2.4

CK

CO

.30

.41

1Q

A O

1 9

9.09

cc

qnq

54

1 CQ

9q

0

.6

10

^ -L

o c

1.00

q q

49

1 C

94.

11

45

1.4

1Q

IK

-L«7 

CO

J-»y 

CO

C7

467

6.4 

1 9

14

O A

3 q

.94

CC

1 C

9q

1C

147

4.2

no

qc

14

-iq

37

600

c A n.

4.2 

1.4

9q

2

c

QC

9 ^

.24 

1.7

94

11

1 8

99

7

qq

41

'TOC

1.0

10

1Q

1.7

C1

19

.10

109

2

C

.41

q

q

o

CO

en

90*1

2

ocn

3.77

9 9Q

ICC

q-i c

9 77

cq

9 9

.13

.11

17

9 4

19

1

1

9

17

1Analyses by X-ray fluorescence.

345 

FeOtota |/MgO RATIO

FIGURE 3.33. FeO^^/MgO ratio versus SiO2 content of Jurassic 

to Cretaceous plutonic rocks. Data field KJ represents most of 

the analyzed Jurassic to Cretaceous plutonic rocks of Blue 

Mountains region. Data field G represents Aleutian island arc 

siliceous plutonic rocks (fig. 3.4). Dots, new data. Calc-alkalinetholeiitic compositional boundary from Miyashiro (1974).

correlate directly with pyroxene abundances in sample 

V-24-79. The higher I^O contents reflect abundances 

of orthoclase and (or) muscovite. With the exception of 

sample V-12-79, FeOtotal/MgO ratios increase with increasing SiC>2 content, but most samples still plot 

within the data field of Jurassic to Cretaceous plutonic 

rocks on a FeOt0tal/MgO ratio versus Si(>2 content diagram (fig. 3.33).

Rb, Sr, Ba, Th, and U contents correlate reasonably 

well with the amount of K^O in the samples (tables 

3.29 and 3.30). Ba content is extremely high in sample 

VC-295. The high-field-strength elements are strongly 

depleted in all samples; Zr contents range from 26 to 

147 parts per million, and Y contents range from 5 to 

19 parts per million. Plots of FeOtotal/MgO versus Cen/ 

Ybn (fig. 3.31) and normalized REE (fig. 3.34) satisfactorily show the calc-alkaline character of these samples and their similarities to both the Bald Mountain 

batholith of eastern Oregon and the siliceous plutonic 

rocks of the Aleutian island arc (Kay and others, 

1990).

Four new initial Sr-isotope ratios (table 3.31) range 

from a low of 0.70335±0.00002 to a high of 0.70422± 

0.00004. They are within the range of values reported by Armstrong and others (1977) for other Jurassic 

to Cretaceous plutons in the Blue Mountains region, 

including plutons (such as the Deserette and Council
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 193

TABLE 3.31. 87Sr/86Sr ratios for selected Jurassic to Cretaceous plutonic rocks of the Blue

Mountains region

[Sample localities listed in table 3.27. Sr and Rb contents in parts per million, determined by X-ray fluorescence; analyst, R.J. 

Fleck, U.S. Geological Survey. Measured ratios normalized to 8°Sr/87Sr=0.1194. Mean (n=76) of 87Sr/86Sr ratio obtained on 

standard (NBS SRM-987) is 0.71023±0.00003. Error values for ratios are standard deviations (2cr) and represent analytical 

uncertainty for split]

Sample Rock type Rb Sr Rb/Sr 87Sr/86Sr

V-1 V -L 1-7Q -L 1 \J

V-T^-TQ

V oq 70

V-46-79

Granodiorite - 

Granodiorite ----- 

Diorite 

Quartz diorite -

9^

34

4

54

cr\c

700

Cf\O

AVC

0.057

.046

.007

0.70335±.00002

7fiQRi4- nnnnQ

.70406±.00001

.70422±.00004

Ce Sm Eu Tb Yb Lu

RARE-EARTH ELEMENTS

EXPLANATION

Sample V-75-79 n Sample V-46-79 

Sample VC-295 A Sample V-11-79 

Sample V-54B-79 A Sample V-24-79

FIGURE 3.34. Rare-earth element (REE) diagram showing calc-alkaline trends of Jurassic to Cretaceous plutonic rocks of Blue 

Mountains region. Sample with anomalous trend (number V-24- 

79) has cumulate texture and abundant hornblende. Patterned 

area (BM) is data field for Late Jurassic Bald Mountain batholith (fig. 3.6B), shown for comparison. Sample numbers correspond to those in tables 3.27 through 3.30; for normalizing 

values used, see figure 3.6.

Mountain plutons) that are very near the suture zone 

between the Blue Mountains island arc and the ancient North American craton. The average of 62 Srisotope ratios (table 3.31; Armstrong and others, 

1977), calculated after eliminating the extremely 

high ratios associated with aplites and pegmatites, is 

0.70392. The Bald Mountain batholith (13 samples) 

has the highest average of 0.70441, and the Deep 

Creek stock (10 samples) has the lowest average of 

0.70350.

PETROGENESIS OF THE JURASSIC TO 

CRETACEOUS PLUTONIC ROCKS

The Jurassic to Cretaceous intrusive bodies in the 

Blue Mountains region are I-type plutons (Armstrong 

and others, 1977; White and Chappell, 1983). These 

plutons' characteristics and implied tectonic settings 

are important for understanding the last pre-Cenozoic igneous stage in the evolution of the Blue Mountains region.

The abrupt change in petrologic character of igneous rocks in the Blue Mountains island arc that occurred after about 180 Ma probably can be linked to 

one or more tectonic events. The strongly calc-alkaline affinity of the Jurassic to Cretaceous intrusive 

rocks points to such factors as different depths and 

temperatures of differentiation, fractionation of additional crystallizing phases (or different proportions of 

the same ones), and perhaps the addition of water, 

mixing of magmas, and assimilation of older crust 

(table 3.2). Source magmas probably were derived by 

partial melting of depleted mantle, followed by some 

crustal assimilation. The crust of the region must 

have been thicker than at any time prior to the Middle Jurassic.

TECTONIC IMPLICATIONS OF PETROLOGIC AND 

OTHER DATA

Igneous petrology is a helpful tool for interpreting 

the geologic evolution of tectonostratigraphic terranes 

(table 3.2). However, it is the integration of igneous 

petrology with geochronology (both isotopic and paleontologic data), paleomagnetism, stratigraphy, and 

structure that promotes better interpretive endeavors.

In this section I discuss the petrogenetic-tectonic 

association in each terrane and conclude with a plate 

tectonic model for the pre-Cenozoic evolution of the 

Blue Mountains region. The model is preliminary 

and no doubt will be greatly modified as better understanding of Cordilleran tectonics is achieved. For
194 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

example, the model does little to correlate terranes of 

the Blue Mountains island arc with those in Washington (Tabor and others, 1987), British Columbia 

(Monger and others, 1982), and southern Alaska (Silberling and others, 1984). Furthermore, the possible 

correlation of the Blue Mountain island arc terranes 

with those of the Klamath Mountains and the Sierra 

Nevada (Davis and others, 1978; Saleeby, 1983) are 

not integrated into this approach. Rather, I concentrate only on the Blue Mountains region.

TECTONIC SETTING OF IGNEOUS ROCKS 

OF THE BAKER TERRANE

The Baker terrane is structurally and petrologically the most diverse and the most complicated part 

of the Blue Mountains island arc. Previous studies 

(see Davis and others, 1978; Saleeby, 1983) adequately pointed out that the Baker terrane is probably a 

fragment of a much larger tectonic belt that can be 

connected to terranes in British Columbia and southern Alaska and possibly to terranes in the Klamath 

Mountains and the Sierra Nevada of the Western 

United States.

The petrology and ages of igneous rocks within the 

Baker terrane have the following tectonic implications:

(1) The Baker terrane was part of an extensive 

elongate belt within (or bordering) an intraoceanic island arc that was tectonically and magmatically active during the late Paleozoic and Triassic. After 

amalgamation with other terranes in the island arc, 

it was intruded by Late Jurassic to Early Cretaceous 

plutons.

(2) Many igneous rocks of the Baker terrane originated in an ocean basin (or basins) in conjunction 

with sea-floor spreading, rifting, and the growth of 

seamounts and oceanic islands and were subsequently accreted to the fore arc of the Blue Mountains island arc, possibly along the inner wall of a trench 

(Mullen, 1985; Bishop, chap. 5, this volume).

(3) Rocks with volcanic arc characteristics that occur in the Baker terrane probably formed along one 

or more magmatic axes of an island arc. They were 

subsequently faulted, sheared, and shuffled with 

rocks from the oceanic plate during tectonic processes 

that are known to occur in fore-arc regions of longlasting and geologically complex island arcs 

(Bloomer, 1983; Fryer and others, 1985; Bloomer and 

Fisher, 1987). Active fore-arc volcanism in island arcs 

is extremely rare (Marlow and others, 1988); almost 

all isotopically dated igneous rocks from fore-arc regions are older than igneous rocks of an associated 

magmatic axis or volcanic front.

(4) The Canyon Mountain Complex formed within 

an island arc. A possible explanation for its present 

position in the Baker terrane is that it crystallized 

near the magmatic axis of the ancestral (Pennsylvanian to Permian) Blue Mountains island arc and subsequently was incorporated within serpentinitematrix melange during Middle and Late Triassic activity along the inner trench wall (outer-arc high) of 

a subduction zone.

(5) Many of the exotic (not directly associated 

with arc volcanism) igneous rocks occur within serpentinite-matrix melange. Serpentinite-matrix melange is common in the Baker terrane, but the entire 

terrane is not melange. Rather, the terrane is composed of relatively coherent units that are chaotically 

mixed in structural zones, some of which are serpentinite-matrix melange. Serpentinite diapirs, like 

those reported from the Marianas Islands fore arc 

(Fryer and others, 1985), may have been important 

contributors to Baker terrane serpentinite-matrix 

melanges (Mullen, 1983). The regions of serpentinitematrix melange probably mark fault zones (within 

the old fore-arc region of the Blue Mountains island 

arc) where highly pressurized and water-rich mantle 

and crust materials were diapirically emplaced. 

These diapirs could easily incorporate high-pressure 

amphibolites and blueschist-facies rocks from deep 

levels beneath the arc (Bishop, chap. 4, this volume).

(6) The Burnt River Schist is an enigmatic unit in 

the Baker terrane. Many of its somewhat rare volcanic rocks have arc affinities (Ashley, chap. 12, this 

volume; Bishop, chap. 5, this volume; PR. Hooper, 

oral commun., 1987); the age of the Burnt River 

Schist is not well known, but its protolith probably 

formed mostly during the Permian or Triassic (Ashley, 1967; Ashley, chap. 12, this volume).

TECTONIC SETTING OF IGNEOUS ROCKS 

OF THE WALLOWA TERRANE

The Wallowa terrane was the site of an active volcanic front during the Early Permian as well as during the Middle and Late Triassic. There is a 

noteworthy gap in both plutonism and volcanism in 

the Early Triassic, although some rocks of that age 

are present in the Baker terrane. 40Ar/39Ar isotopic 

ages indicate that limited igneous activity occurred as 

late as about 215-210 Ma evident in the (informal) 

Sparta complex, the Oxbow Complex, and in the intrusive body at the mouth of the Imnaha River; U-Pb 

isotopic age data confirmed a Late Triassic (215 Ma) 

age for the youngest pluton of the Sparta complex 

(Walker, 1986). The bulk of the U-Pb ages lie in the
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 195

265- to 245-Ma and 235- to 215-Ma intervals (Walker, 

1986), which overlap the ages of the two major volcanic episodes (Vallier, 1977).

Petrologic and tectonic interpretations for the late 

Paleozoic and Triassic igneous rocks of the Wallowa 

terrane are as follows:

(1) The igneous rocks are mostly tholeiitic; rare 

plutonic rocks have transitional or calc-alkaline affinities. Plutonic bodies are I-type and exhibit no apparent evidence for mixing with continental crust. The 

igneous rocks most likely formed from fractionation 

of a parental tholeiitic basalt magma that had been 

derived from partial melting of the mantle wedge, 

with or without significant input from the subducting 

slab. Magma chambers were relatively shallow (probably less than 10 km), and magma temperatures 

were high. Plagioclase, olivine, orthopyroxene (not always present), augite, and magnetite were the major 

fractionating minerals. Mixing may have occurred, 

but most likely with magmas of similar composition. 

Water content was relatively low.

(2) The Permian and Triassic igneous rocks were 

erupted along the magmatic axis of an intraoceanic 

island arc. This setting implies a separation of 100 to 

150 km, depending on the angle of the subducting 

slab, between the trench and volcanic front. Outcrops 

of high-pressure blueschist (Hotz and others, 1977; 

Bishop, chap. 4, this volume) now lie west of the Wallowa terrane and probably mark the location of the 

ancient trench. A sample from one of these outcrops, 

near Mitchell, Oreg. (about 50 km west-northwest of 

John Day, Oreg.; fig. 3.1), yielded an isotopic age of 

225 Ma (Hotz and others, 1977). Considering the suspected clockwise rotation (Wilson and Cox, 1980), the 

trench was probably southwest of the Wallowa terrane in the Late Triassic.

(3) The Permian volcanic rocks are siliceous and 

dominated by pyroclastic debris. A notable characteristic of the Permian volcanic and plutonic rocks is 

high Na2O content, most likely the result of metasomatism associated with the free interchange of seawater brine during crystallization. The abundance of 

high-field-strength elements suggests a tectonic environment different from those of modern island arcs. 

The high concentrations may reflect tectonic and (or) 

magmatic processes similar to those associated with 

the development of alkalic lavas in extensional (deeply rifted) settings. The Triassic lavas are more mafic 

than the Permian lavas and evolved from predominantly submarine flows and coarse breccias in early 

eruptions to subaerial and more tuffaceous deposits 

during later eruptions.

(4) The Wallowa terrane, as now recognized, may 

have shared its basement rocks with the Baker terrane in the early history (older than late Early Permian) of the Blue Mountains island arc. If this is 

true, then a change in implied plate tectonic relationships promoted volcano growth (formation of Wallowa 

terrane) in the Permian upon a preexisting terrane.

(5) The Jurassic hypabyssal intrusive rocks that 

cut strata of the Middle and Upper Jurassic Coon 

Hollow Formation are calc-alkaline. The petrologic 

change from mostly tholeiitic to calc-alkaline igneous 

rocks implies a change, some time after about 180 

Ma, in factors such as relative lithosphere thickness, 

the tectonic setting with respect to a subduction 

zone, the depth to the subducting slab, and processes 

of differentiation.

(6) On the basis of what we understand about Cenozoic intraoceanic island arcs, a large part of the 

Wallowa terrane (in fact, probably much of the entire 

Blue Mountains island arc) is missing. Particularly 

noteworthy is the missing sea floor that must have 

existed between the arc and the North American continent. Several mechanisms can be called upon to account for the small size of the Wallowa terrane, the 

missing materials, and the present configuration of 

the terranes. These mechanisms include double subduction zones with attendant tectonic erosion, westward subduction and tectonic erosion during part of 

the arc's evolution, tectonic overriding of the Wallowa 

terrane by North America, extensive overthrusting 

during amalgamation of terranes within the Blue 

Mountains island arc and during its accretion to 

North America, and northward translation of large 

blocks into Washington, British Columbia, and southern Alaska (Jones and others, 1977; Saleeby, 1983; 

Wernicke and Klepacki, 1988).

TECTONIC SETTING OF IGNEOUS ROCKS 

OF THE OLDS FERRY TERRANE

Most of the igneous rocks of the Olds Ferry terrane 

are different from those of the Wallowa terrane in 

age and chemistry. From the available data, it appears that the Olds Ferry magmatic front became active at about the same time that igneous activity in 

the Wallowa terrane quieted down. The following 

paragraphs present major petrologic and tectonic interpretations of the igneous rocks of the Olds Ferry 

terrane.

(1) The igneous rocks are tholeiitic. They are more 

similar lithologically to Tongan island-arc tholeiites 

than to igneous rocks from the Wallowa terrane. The 

igneous rocks most likely formed in an intraoceanic 

island arc. The tholeiitic affinity implies a mantle 

wedge source, shallow and hot magma chambers, small 

amounts of magma mixing and crustal assimilation,
196 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

and fractionation predominantly by crystallization of 

plagioclase, olivine, orthopyroxene (may or may not be 

present), augite, and magnetite (table 3.2).

(2) A major difference between Olds Ferry terrane 

and Wallowa terrane volcanic rock chemistries is the 

greater abundance of most incompatible elements 

(particularly K, Rb, Ba, and Sr) in the Olds Ferry 

terrane rocks. This greater abundance of the incompatible elements in the Olds Ferry terrane might be 

related to a greater distance from the subduction zone 

(Gill, 1981) and hence, for example, a shallower dip on 

the subducting slab. A possible explanation for a 

change in subduction-zone geometry would be a 

change in plate-convergence direction and velocity 

with a resultant decrease in subduction angle and the 

eventual formation of a new magmatic axis at a greater distance from the trench. Changes in differentiation processes, however, might also account for the 

higher concentrations of the incompatible elements.

(3) A large part of the Olds Ferry terrane, like the 

other terranes of the Blue Mountains island arc, is 

either missing or covered by Tertiary rocks. A 60-degree counterclockwise rotation of the Olds Ferry terrane, to compensate for the clockwise rotation 

suspected by Wilson and Cox (1980) to have taken 

place since Early Cretaceous time, would orient the 

terrane nearly north-south, parallel and very close to 

the probable cratonal boundary. This position would 

suggest that during accretion much of the terrane 

may have been overridden by the craton. I assume, 

in this reconstruction, that the terrane was not rotated prior to the Early Cretaceous.

(4) The Olds Ferry terrane probably was built on 

rocks lithologically similar to those of the Baker terrane. The lavas probably were erupted onto the older 

terrane after a Late Triassic (late Karnian to early 

Norian) change in convergence direction between the 

oceanic and the Blue Mountains island-arc plates. At 

that time, the Wallowa and Baker terranes became part 

of the Olds Ferry terrane fore-arc region. Mortimore 

(1986) concluded that the Olds Ferry terrane was part 

of an extensive, highly potassic volcanic arc that formed 

along the margin of the North American continent. I 

would modify his conclusion somewhat because lavas in 

the Olds Ferry terrane were erupted in an intraoceanic 

island-arc setting. It is likely that the Olds Ferry terrane was an offshore part of the extensive Late Triassic 

continental-margin volcanic arc of Mortimore (1986).

TECTONIC SETTING OF JURASSIC TO 

CRETACEOUS IGNEOUS ROCKS

The Jurassic dikes and sills intruding the Coon Hollow Formation are significant because they are calcalkaline in character. Different from the older, mostly 

tholeiitic rocks, they are similar in chemistry to the 

Jurassic to Cretaceous plutonic bodies. There may 

have been a significant change in the plate tectonic 

setting of the Wallowa terrane (and probably all of the 

Blue Mountains island arc) after about 180 Ma.

The Jurassic to Cretaceous igneous rocks were emplaced after terranes of the Blue Mountains island 

arc had been amalgamated and, in part, after the arc 

had been accreted to North America. In fact, as 

pointed out by Snee and others (chap. 15, this volume), some of the intrusive bodies along the border 

zone between ancestral North America and the Blue 

Mountains island arc probably formed during the 

final stages of accretion.

The chemistry of these Jurassic to Cretaceous calcalkaline intrusive bodies implies the following tectonic setting (table 3.2):

(1) The crust was thicker in the Jurassic than in 

the Permian and Triassic. The crust of the arc had a 

long period of time in which to thicken, from the 

Pennsylvanian (or earlier) to the Middle Jurassic, or 

as many as 140 to 150 m.y. To put this in perspective, changes in plate motions within the ancestral 

Pacific Ocean during that time interval may have 

been comparable to changes that the present-day 

ocean floors have undergone since the beginning of 

the Early Cretaceous. There were many opportunities for the crust to thicken, not only through magmatic activity, but also through sedimentation, 

tectonic stacking, and underplating. I speculate that 

the crustal thickness of Blue Mountains island arc in 

the Middle Jurassic was at least as great as in the 

modern Aleutian island arc, where it has been measured in places as more than 20 km.

(2) Initial strontium isotope ratios indicate that 

evolved continental crust was not assimilated, nor 

did it serve as a principal source for magmas (Armstrong and others, 1977).

(3) These calc-alkaline magmas crystallized at relatively high pressures and contained more water 

than is typical of tholeiitic magmas. Fractionation 

took place predominantly by crystallization of plagioclase, olivine, orthopyroxene (not always present), 

augite, and magnetite, with amphibole and garnet 

possibly playing minor roles. Magma mixing and assimilation of older arc crust may have been important processes.

(4) Calc-alkaline magmatism in an oceanic setting 

implies converging tectonic plates. Thermal conditions leading to calc-alkaline magmatism may be related to a difference in ages of the converging plates. 

The subducting slab may have been younger and 

therefore much warmer. Young crust being subducted
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 197

would cause a much higher heat flow and probably 

more melting above the subducting slab. Paralleling 

the tectonic setting of the present-day Pacific Ocean 

floor off northern California, Oregon, and Washington, a spreading ridge near the west coast of North 

America during the Late Jurassic and Early Cretaceous may have led to suitable conditions for forming 

large calc-alkaline plutons.

(5) The dip of the subducting slab may have had a 

significant influence on the locations of the Jurassic 

to Cretaceous plutons. A low-angle dip could have 

put the magmatic axis at a greater distance from the 

trench. Some outer-arc terranes, if they existed, may 

have been ripped away from the margin during the 

Early Cretaceous and transported northward by 

wrench faulting for a rendezvous with British Columbia and southern Alaska.

PLATE TECTONIC EVOLUTION OF 

BLUE MOUNTAINS ISLAND ARC

PALEOMAGNETIC AND FAUNAL EVIDENCE 

FOR LATITUDINAL MOVEMENT

Paleomagnetism is a powerful tool for determining 

paleolatitudes of so-called exotic or suspect terranes 

and their positions relative to the North American 

craton through time. Paleomagnetic data have been 

gathered for the Jurassic to Cretaceous plutons of 

the Blue Mountains region (Wilson and Cox, 1980), 

Middle and Upper Triassic lavas of the Wallowa and 

Olds Ferry terranes (Hillhouse and others, 1982), 

and Lower Permian strata of the Wallowa terrane 

(Harbert and others, 1988, 1995). Assuming that terranes of the Blue Mountains island arc are kindred, 

then these paleomagnetic data can be used to describe the movement of the island arc relative to its 

site of accretion to the North American craton. Reconstructing this movement is difficult because the 

North American craton also was moving during the 

Permian through Late Jurassic interval, particularly 

in the Late Jurassic as North America began separating from Europe.

Contrary to earlier speculations (for example, Vallier and Engebretson, 1984), it is now evident that 

parts of the Blue Mountains island arc have been in 

the northern latitudes since the Early Permian (Hillhouse and others, 1982; Harbert and others, 1988, 

1995). Some of the blocks associated with melange 

and disrupted units in the Baker and Grindstone terranes, however, are probably parts of an accretionary 

wedge and may have been scraped off a downgoing 

oceanic plate (or plates) that had formed in southern 

latitudes. Unfortunately, paleomagnetic data cannot

be used to determine paleolongitudes, so they shed 

no light on the arc's distance west of the North 

American craton; The Permian island arc may have 

formed very close to the North American craton or it 

may have formed in the northwestern parts of the 

ancestral Pacific Ocean.

Harbert and others (1988, 1995) concluded that 

Lower Permian strata (the Hunsaker Creek Formation) of the Wallowa terrane were deposited in the 

northern hemisphere at a paleolatitude of about 28° 

N. According to the apparent polar wander path of 

Irving and Irving (1982) and studies of the North 

American craton by Gordon and others (1984) and 

Symons (1990), the site of accretion of the Wallowa 

terrane to the North American craton (that is, the 

site now at 45° N., 116° W) was at a paleolatitude of 

about 11° to 19° N. during the Permian. Thus, during 

Permian time, the island arc on which the Hunsaker 

Creek Formation sediments were deposited lay about 

9° to 17° north of the site of accretion on the North 

American craton.

Hillhouse and others (1982) concluded that Middle 

and Late Triassic volcanic rocks from both the Wallowa and Olds Ferry terranes were at a paleolatitude 

of 20° N. or 20° S. in the Middle and Late Triassic. 

Because the work by Harbert and others (1995) 

shows that the Wallowa terrane was probably in the 

northern hemisphere during the Permian, it follows 

that the Triassic lavas probably erupted in the northern hemisphere as well. According to the apparent 

polar wander path of Irving and Irving (1982), the 

site of accretion was at a paleolatitude of about 15° 

to 20° N. in the Middle and Late Triassic (230-210 

Ma). This paleolatitude implies that by Late Triassic 

time the Blue Mountains island arc was at about the 

same latitude relative to the North American craton 

as when the arc finally accreted to the continent in 

the Early Cretaceous. Thus the island arc moved 

south as much as 1,000 to 2,000 km between the 

Early Permian and the Late Triassic.

In the Late Triassic, both the site of eventual accretion to the North American craton and the Blue 

Mountains island arc itself were at about 20° N. latitude. In other words, the latitude of most of the Blue 

Mountains island arc relative to the site of accretion 

on the North American craton has not changed significantly since the Late Triassic; both the continent 

and the arc moved northward to their present positions. Furthermore, paleomagnetic measurements for 

the Jurassic to Cretaceous plutons in the Blue Mountains region (Wilson and Cox, 1980) indicate very little latitudinal movement, but do suggest clockwise 

rotation of up to 60° since crystallization. Apparently, 

both the Blue Mountains island arc and the site of
198 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

accretion to the North American craton had reached 

their present positions by the Late Jurassic. Faunal 

studies by Pessagno and Blome (1986) showed a 

northward progression of the Izee terrane in the Early and Middle Jurassic from a southerly Tethyan region to a more northerly boreal zone. By Bajocian 

time (approximately 180 Ma) a faunal link had been 

established between the Blue Mountains island arc 

(Wallowa terrane) and the North American craton 

(Stanley and Beauvais, 1990).

These data suggest southward transport of the island arc between the late Early Permian and Late 

Triassic (about 15° of latitude in 30 m.y., which is an 

average of roughly 50 km/m.y.) and northward transport of about the same distance in the Jurassic (the 

island arc now moving with North America, with 

most movement occurring in the Early and Middle 

Jurassic). Because North America also was moving 

northward during the Jurassic and because the 

North American site of accretion and the island arc 

were at about the same latitude by the Late Triassic, 

it is probable that the Blue Mountains island arc has 

not moved a significant distance latitudinally relative 

to the North American craton since the Late Triassic.

PLATE MOVEMENTS AND TERRANE TECTONISM

A comprehensive discussion of the Grindstone terrane has not been attempted here. Buddenhagen 

(1967), Miller (1987), Blome and Nestell (1991), and 

Ferns and Brooks (chap. 9, this volume) have presented and interpreted critical data pertinent to the original tectonic setting and evolution of the Grindstone 

terrane. At least in part its history is probably very 

similar to that of the Baker terrane. There are data, 

however, that tie the Grindstone terrane to the northern part of the Klamath Mountains (Miller, 1987); interpretations of the original tectonic setting of the 

terrane, its relations to both the Klamath and the Blue 

Mountains, and its particular dynamics should remain 

open until additional data are collected.

The Baker terrane is probably kindred to all other 

terranes in the Blue Mountains island arc; for instance, some of the Baker terrane formed in the forearc regions of the Wallowa and Olds Ferry terranes, 

and other parts (including pieces of an ancient arc 

terrane) probably underlie the Wallowa, Olds Ferry, 

and Izee terranes. It is not a so-called exotic terrane 

that formed elsewhere and then collided with the 

other terranes of the arc, although fragments of oceanic crust from the downgoing slab were caught up 

in the arc's subduction zone (or zones). To account for 

the diversity of rock types in the Baker terrane and 

the occurrence of the many distinct structural or melange zones (Evans, chap. 8, this volume; Ferns and 

Brooks, chap. 9, this volume; Bishop, chaps. 4 and 5, 

this volume), the relation between the Blue Mountains island arc and the Baker terrane must be considered in both time and space.

If a long interval of time is combined with the 

changing of tectonic and magmatic locations within an 

island arc, then the mixing of arc and nonarc (for example, oceanic-crust, including seamount and midocean-ridge) igneous rocks should be expected. What 

is now the Baker terrane undoubtedly shifted position 

relative to other parts of the maturing island arc; furthermore, during its long and complex evolution, 

parts of the Baker terrane probably were tectonically 

eroded and either subducted or translated subparallel 

to the arc axis by strike-slip faulting, similar to the 

present translation of fore-arc materials in the western part of the Aleutian island arc Vallier and others, 

1994). For example, what originally had been the 

magmatic axis and fore arc(?) of the Blue Mountains 

island arc in the late Paleozoic may have become 

mostly fore arc in the Middle and Late Triassic. 

Therefore, parts that were related to the late Paleozoic magmatic axis, such as the Canyon Mountain Complex, were reshuffled along and incorporated 

within the Middle and Late Triassic fore-arc region. 

Some of the oldest parts (for example, an ancient prePermian arc terrane) of what is now the Baker terrane could thereby become basement to the Wallowa 

and Olds Ferry terranes. Exotic fragments of oceanfloor rocks, probably from more southerly latitudes, 

that occur as knockers in serpentinite-matrix melange 

and as possible olistoliths within the Elkhorn Ridge 

Argillite probably were scraped off an ancient oceanic 

plate (or plates) during subduction. The high-pressure 

rocks would be brought to the surface from beneath 

the arc framework along major fractures, some of 

which were greased by serpentinite diapirs like those 

discovered in the Marianas fore-arc region (Fryer and 

others, 1985; Bishop, chap. 5, this volume).

A major change in plate-convergence vectors probably occurred between the last major igneous activity 

in the Late Permian (about 260 Ma) and the Middle 

Triassic (about 235 Ma). This is the time of suspected 

southward transport of the Blue Mountains island 

arc. Fragmentation and eastward transport of the 

Melanesian island arc (Kroenke, 1984) in the late Eocene to early Miocene interval (approximately 3,000 

km in 20-25 m.y.) may be an analogous situation.

Differences in ages and petrologic characteristics of 

the Wallowa and Olds Ferry terranes suggest that 

another major change in plate vectors occurred in the 

late Karnian, when the magmatic axis switched from 

the Wallowa terrane to the Olds Ferry terrane. The
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 199

Upper Triassic Huntington Formation formed along a 

magmatic axis in the Olds Ferry terrane. Igneous 

quiescence in the Wallowa terrane and in parts of the 

Baker terrane led to subsidence, the growth of large 

carbonate banks on the subsiding platforms, and the 

formation of rift basins. These conditions accompanied the deposition of the Martin Bridge Limestone 

and the Hurwal Formation in the Wallowa terrane 

and probably some of the Upper Triassic limestone 

bodies in the Baker terrane.

Rocks lithologically similar to those in the Baker 

terrane (chert, serpentinite, limestone, tuffaceous argillite) are faulted into the Weatherby Formation, 

herein considered part of the Izee terrane (see section "Izee Terrane"), near Huntington, Oreg. (Brooks 

and others, 1976; Brooks, 1979a); such rocks also 

crop out on strike with the Olds Ferry terrane near 

New Meadows, Idaho (Bonnichsen, 1987). Their presence suggests that parts of both the Izee and the 

Olds Ferry terranes are underlain by rocks of the 

Baker terrane or a similar (for example, an ancient 

arc) terrane. In addition, a Middle Triassic (235 Ma; 

Walker, 1986) intrusive body beneath (or intruding?) 

strata of the Huntington Formation suggests either 

that an older arc terrane existed prior to deposition 

of Upper Triassic strata in the Olds Ferry terrane or 

that the Huntington Formation includes rocks as old 

as 235 Ma. I prefer the former explanation because 

of field relations. If the Middle Triassic intrusive 

body is part of the Baker terrane or a lithologically 

similar ancient arc terrane, as I suspect, then it follows that the Olds Ferry terrane represents a Late 

Triassic (late Karnian and early Norian) and Early 

Jurassic(?) magmatic axis that grew on the Baker (or 

similar) terrane. A change in converging-plate vectors 

in the Late Triassic can explain the shift in magmatic axis from the Wallowa terrane to the Olds Ferry 

terrane. Such a change may have been analogous to 

the change in arc polarity proposed by Kroenke 

(1984) for the Solomon Islands during the Miocene.

An intra-arc (fore arc to the Olds Ferry terrane) 

Late Triassic to Late Jurassic basin evolved between 

the Olds Ferry and the (composite) Baker and Wallowa terranes and formed what is now the Izee terrane. Strata of the Izee terrane were deposited in 

that intra-arc basin and thus were deposited on parts 

of the Baker and Olds Ferry terranes. Late Cenozoic 

elongate basins in modern arcs, probably analogous 

to the basin that formed the Izee terrane, occur in 

the Solomon Islands and Vanuatu (Kroenke, 1984), 

where subduction polarity changes in the Miocene 

caused profound extension of the older arc frameworks and the formation of intra-arc basins. A Late 

Triassic (late Karnian) change in plate-convergence

vectors between the ancestral oceanic plate (or 

plates) and the Blue Mountains island arc, proposed 

above, could have led to the formation of an intra-arc 

basin into which sediments poured from uplifted 

parts of the fore-arc region (parts of the Baker and 

Wallowa terranes) and the new magmatic axis (Olds 

Ferry terrane). The creation of such a basin would 

account for the formation of the Izee terrane within 

the Blue Mountains island arc. Upper Triassic to 

Upper Jurassic sediments deposited in the Wallowa 

terrane are correlative with those in the Izee terrane; 

this relation indicates that the Wallowa terrane did 

not merely subside, but experienced block faulting, 

subaerial erosion of the uplifted blocks, and marine 

transgression during block subsidence.

A new subduction zone (convergence between oceanic plate and combined Blue Mountains island arc 

and North American craton) apparently was initiated 

by 160 to 150 Ma. The continuity of Jurassic to Cretaceous plutonic bodies that have the same approximate age range and chemistry from California 

through Vancouver Island and into Alaska indicates 

that a volcanic arc was built along the newly formed 

continental margin. Considering the high amount of 

melting that must have occurred to form the plutonic 

bodies, the rarity of equivalent volcanic rocks is puzzling. Either the volcanic rocks were completely eroded or the magmas simply did not reach the surface to 

form volcanoes. Jurassic to Cretaceous plutonic bodies in the Blue Mountains cut across terrane boundaries and across major thrust faults, indicating that 

a significant deformational phase occurred in the 

time interval between deposition of the youngest Jurassic sediments (Callovian in the Izee terrane and 

Oxfordian in the Wallowa terrane, or about 165 to 

155 Ma) and the crystallization of these large plutonic bodies (most are 145 Ma and younger according to 

N.W. Walker, written commun., 1989).

SEVEN-STAGE MODEL FOR PLATE TECTONIC EVOLUTION OF 

BLUE MOUNTAINS ISLAND ARC

A speculative plate tectonic model for the evolution 

of the Blue Mountains island arc from the late Paleozoic to the Late Cretaceous is shown in figure 3.35. 

Major assumptions are that (1) the Blue Mountains 

island arc has always been one complex island arc, 

(2) all terranes of the Blue Mountains island arc are 

kindred, (3) processes occurring in island arcs were 

the same in the late Paleozoic and Mesozoic as they 

are at present, and (4) the petrology of igneous rocks 

can be combined with other geologic data to interpret 

the evolution of a complex island arc. The proposed 

seven stages of arc evolution are as follows:
200 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Stage 1, late Paleozoic and Early Permian, about 

370(?) to 270 Ma (fig: 3.35A). A complex island arc 

began to form in the pre-Early Permian. A magmatic 

axis of some kind had formed by the Pennsylvanian 

(about 310 Ma). The oldest parts (an ancient arc terrane) of what are now the Baker terrane and possibly 

the lower crustal parts of the Wallowa terrane may 

have formed even earlier. Many of these ancient rocks 

presumably have been covered, subducted, and (or) 

ripped away and transported elsewhere. The configuration of this postulated ancient arc, including its polarity, paleolatitude, and paleolongitude, are unknown.

Stage 2, Early Permian, about 270 to 260 Ma (fig. 

3.35E). The Early Permian interval is marked by 

extensive igneous activity and volcano growth. Siliceous arc volcanoes grew on part of the ancient arc 

terrane. This Permian arc (at least the Wallowa terrane) was in the northern hemisphere at about 28° N. 

latitude on the basis of paleomagnetic data (Harbert 

and others, 1988, 1995). The accretion site on the 

North American craton was at about 11° to 19° N. 

latitude on the basis of the apparent polar wander 

path of Irving and Irving (1982). The direction of 

plate convergence (arc polarity) is unknown. Figure 

3.35B shows an ancient oceanic plate being subducted 

northwestward under the arc. Northwestward subduction (sea floor moving southeast to northwest, 

from equatorial to subpolar site) could account, in the 

Baker terrane, for the presence of Tethyan (warm-water) fusulinids in carbonate rocks that presumably 

had been scraped off the ancient oceanic plate during 

subduction to become part of an accretionary prism. 

Warm-water (subtropical) Permian fauna have not 

been reported from the Wallowa terrane; their absence tentatively supports the more northerly paleolatitude (G.D. Stanley, Jr., oral commun., 1990). The 

arc's paleolongitude and geographic relation to other 

terranes and continents are unknown.

FIGURE 3.35 Speculative model to explain geologic evolution of 

Blue Mountains island arc in seven stages, from late Paleozoic to 

mid-Cretaceous, based mainly on petrologic and paleomagnetic 

interpretations. Abbreviations are AA, ancient arc terrane (undivided); AOP, ancient oceanic plate; BA, Baker terrane (includes 

Grindstone terrane); IZ, Izee terrane; NA, North American craton; OF, Olds Ferry terrane; OT, other terranes; PP, Pacific (or 

Farallon) plate; and WA, Wallowa terrane. Schematic cross section (left side) and plan view (right side) are shown for each of 

seven evolutionary stages (A-G). In section and plan view, faults 

are dashed where approximately located, queried where speculative; single-barbed arrow indicates relative movement along fault 

or plate boundary. In section, sedimentary rocks are shown by 

conventional symbols and intrusive rocks by randomly oriented 

dashes. In plan view, paleonorth is up and double-barbed arrow 

indicates relative plate movement; at convergent-plate boundary, 

teeth are on upper plate; *, approximate site of future accretion 

of Wallowa terrane to North American craton. A, Stage 1, Late

Stage 3, Late Permian and Early Triassic, 260 to 

235 Ma (fig. 3.35C). Igneous activity of the arc quieted down during this time interval but did not stop 

altogether. The arc moved southward toward the latitude of the site of eventual accretion to the North 

American craton. During the proposed extensive 

transport, there probably were abundant opportunities for the fragmentation of the arc and the formation of terranes, in addition to the Baker and 

Wallowa terranes, that would ultimately end up along 

different parts of North America (and possibly eastern Siberia). In fact, the Blue Mountains island arc 

during this time interval may have been only a small 

part of an extensive offshore microcontinent. The ancient arc terrane (in stages 1 and 2 above) was changing into what is now the Baker terrane through the 

reworking of its rocks by tectonic and igneous processes. The eastward movement of the Melanesian and 

younger arcs away from Australia (Kroenke, 1984) 

since the early Eocene may be a good model to explain how the Blue Mountains island arc moved 

southward over the Pacific Ocean floor during the 

Late Permian and Early Triassic. Extensive strikeslip faulting, or movement along transform faults, 

would be the simplest method of transporting the arc, 

but no direct evidence for such movement has been 

found in the Blue Mountains rocks. In any scenario, 

the convergence of the ancient oceanic plate (or 

plates) with the Blue Mountains island arc probably 

was in a southeast-to-northwest direction.

Stage 4, late Middle and early Late Triassic, 235 to 

225 Ma (fig. 3.35D). The late Middle and early Late 

Triassic time interval (235-225 Ma) was one of voluminous magmatic activity in the Wallowa terrane and 

of minor, probably related, activity in what is now the 

Baker terrane. This magmatism may have been preceded by a change in converging-plate vectors. There 

must have been a well-developed subduction zone. The

DevonianC?) to Early Permian (3707-270 Ma): Evolution of ancient 

arc terrane; spatial relationships are speculative; presence of landmass on left of plan view is suspected, but relative positions of arc, 

suspected landmass, and North American craton are unknown and 

therefore queried; note that the arc can migrate over the oceanic 

plate. Major topographic and structural features of an island arc are 

shown in cross-section view. B, Stage 2, Early Permian (270-260 

Ma): Formation of siliceous volcanoes (Wallowa terrane) on ancientarc framework; siliceous volcanic rocks erupt at approximately 30° 

N. latitude; fore-arc parts of ancient arc terrane become Baker terrane; question mark in plan view indicates that distance between 

arc and North America is unknown. C, Stage 3, Late Permian and 

Early Triassic (260-235 Ma): Relative igneous quiescence; southward transport of Blue Mountains island arc about 2,000 km relative to North American craton; what remains of ancient arc terrane 

has been incorporated into Baker terrane. Arc shown close to North 

America, but its position in ancestral Pacific is unknown; it could 

have been in western or eastern part.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 201

A Late Devonian(?) to Early Permian (370?-270 Ma)

A'

FORE ARC

Fore-arc basin /Outer high

Accretionary wedge 

Trench

ANCIENT ARC > TERRANE (AA)

SEA LEVEL

Magmatic 

axis

Trench/

Fore arc

AOP

NA

B Early Permian (270-260 Ma)

NW

BLUE MOUNTAINS ISLAND ARC

SE

SEA LEVEL

WA

AOP

Baker terrane

AOP

-25° N

30° N

C Late Permian and Early Triassic (260-235 Ma)

NW

AOP

AOP 

Figure continued on next page.
202 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

presence of blueschist in the western segments of the 

Baker terrane (Hotz and others, 1977; Bishop, chap. 4, 

this volume), and the 60° clockwise rotation of the arc 

since the Early Cretaceous suggest that the arc polarity was associated with south-to-north subduction (assuming that no rotation took place between the Late 

Triassic and the Early Cretaceous). The Blue Mountains island arc was at about 20° N. latitude, approximately directly offshore from its ultimate site of

accretion. Sinistral movement along major shear zones 

in the Cougar Creek and Oxbow Complexes of the 

Wallowa terrane (Ave Lallemant, chap. 7, this volume) 

suggests extensive internal deformation of the arc and 

may account for some of its lateral movement. Formation of a fore-arc basin and irregular basins along the 

inner wall of the trench, olistostromal sliding of large 

blocks, intrusion of serpentinite diapirs, and accretion 

of materials from the oceanic plate probably all ocD Middle and early Late Triassic (235-225 Ma)

D- D'

Back-arc 

rift

Back-arc rift

AOP

SEA LEVEL

20° N

AOP

£ Early and Middle Jurassic (190-155 Ma)

w

AOP

20° N

FIGURE 3.35 Continued. D, Stage 4, Middle and early Late Triassic 

(235-225 Ma): Volcanism in Wallowa terrane; well-developed subduction zone; Triassic volcanic rocks cool at approximately 20° N. 

latitude, about same latitude as accretion site on North American 

craton; left-lateral movement on northeast-striking faults (on section, movement toward reader shown by solid circle, away from 

reader, by circled cross); proposed oceanic plateau or rifted backarc basin may account for extensive outpourings of Middle Triassic

basalt now exposed in the Wrangellia and related terranes. E, Stage 

5, Late Triassic and Early Jurassic (225-190 Ma): Shift of volcanic 

activity to Olds Ferry terrane; inception of intra-arc basin (Izee 

terrane); growth of carbonate banks (Martin Bridge Limestone and 

correlative strata) on volcanically quiescent parts of Wallowa and 

Baker terranes that had become fore arc to Olds Ferry volcanic arc. 

At this time, Blue Mountains island arc and North American continent began moving northward together.
3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 203

curred in the Wallowa terrane fore-arc region, exemplified by parts of what are now called the Grindstone 

and Baker terranes. The Canyon Mountain Complex 

may have been exhumed at this time along the outer 

high (inner trench wall) of the Wallowa terrane's forearc region. Extensive outpourings of Middle Triassic

basalt, now exposed as parts of the Wrangellia terrane 

on Vancouver and Queen Charlotte Islands (Canada) 

and southeast Alaska (Jones and others, 1977), probably formed near the Wallowa terrane, possibly in a 

back-arc environment (Vallier, 1986) or as an oceanic 

plateau (P.R. Hooper, oral commun., 1987).

F Early and Middle Jurassic (190-155 Ma)

sw NE

SEA LEVEL

PP

50° N

NA

G Late Jurassic to Early Cretaceous (155-115)

sw

SEA LEVEL

NA

FIGURE 3.35 Continued. F, Stage 6, Early and Middle Jurassic 

(190-155 Ma): Voluminous sedimentation in Izee terrane as 

intra-arc basin subsided; figure shows configuration of Blue 

Mountains island arc during final stages of northward movement (155 Ma) when arc had already reached site of accretion; 

thrust-faulting in latter parts of interval telescoped terranes. 

Strike-slip faulting active during accretion process; fragments 

of the island arc presumably split off at this stage; Pacific (or

Farallon) plate probable oceanic plate; some igneous activity, 

particularly in Wallowa and Olds Ferry terranes. G, Stage 7, 

Late Jurassic to Early Cretaceous (155-115 Ma): Intrusion of 

large calc-alkaline plutonic bodies throughout the island arc; 

collision of Blue Mountains island arc with North America; 

thrust faults within Blue Mountains island arc not shown in 

plan view because they were not active during this interval; 

strike-slip faults modified from Wernicke and Klepacki (1988).
204 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Stage 5, Late Triassic and Early Jurassic, 225-190 

Ma (fig. 3.35E). The Olds Ferry terrane evolved 

mostly during the 225- to 190-Ma interval. It probably developed on old island-arc crust composed of 

fragments of the Baker terrane. A plate-motion 

change could account for the growth of the Olds Ferry terrane as a magmatic axis; this plate-motion 

change and consequent growth of the Olds Ferry terrane would make the Wallowa and Baker terranes 

parts of the Olds Ferry fore-arc region. An arc-reversal model like that proposed for the Solomon Islands 

arc (Kroenke, 1984, p. 47-61) could explain these relationships, although even a small change in converging-plate vectors could effect a similar result. Thus, 

igneous activity decreased in the Wallowa (and possibly Baker) terrane while it began building the Olds 

Ferry terrane. Carbonate banks (Martin Bridge 

Limestone and correlative strata) grew on slowly 

subsiding parts of the new fore-arc region, and horsts 

and grabens formed islands and basins, which filled 

with sediments such as those in the Hurwal Formation. A large intra-arc basin (between and on the 

Olds Ferry and the composite Baker and Wallowa 

terranes) formed and ultimately became the Izee terrane (herein considered to include the Weatherby 

Formation and Squaw Creek Schist; see section "Izee 

Terrane"). The Olds Ferry terrane probably was an 

oceanic part of the extensive volcanic arc that bordered North America in the Late Triassic and the 

Early Jurassic (Mortimore, 1986), some of which developed as a continental volcanic arc. Northward 

movement of the Blue Mountains island arc along 

with the North American craton (Irving and Irving, 

1982) probably began by at least the latest Triassic 

or earliest Jurassic. It is possible that as early as 225 

Ma, the North American craton and the Blue Mountains island arc were parts of the same tectonic 

(North American) plate.

Stage 6, latest Early to late Middle Jurassic, 190 to 

155 Ma (fig. 3.35Y). Prolific sedimentation occurred 

during the 190- to 155-Ma interval in the intra-arc 

basin that ultimately became the Izee terrane. The 

Blue Mountains island arc continued to migrate 

northward with North America. Proximity to North 

America by about 180 Ma is indicated by an early 

Bajocian fauna in the Wallowa terrane that has 

North American affinities (Stanley and Beauvais, 

1990). Although igneous activity was minor in all terranes during this time interval, a change from tholeiitic to calc-alkaline characteristics of igneous rocks is 

shown by the presence of small calc-alkaline intrusive rocks within the Coon Hollow Formation of the 

Wallowa terrane. Extensive thrust faulting occurred 

in the latter part of this time interval, probably along

an oblique transpressive margin. A bulge in the 

North American craton north of the site of accretion, 

as shown in figure 3.35F, is suspected because of the 

present-day spatial relations. Right lateral strike-slip 

faulting is inferred to have begun the accretion process (Wernicke and Klepacki, 1988). Fragments of the 

complex island arc (and adjacent parts of the Wrangellia terrane) were possibly broken away to be subsequently transported northward along the cratonic 

margin. During the latter parts of this time interval, 

I suspect that either the Pacific plate or Farallon 

plate was the subducted oceanic plate.

Stage 7, Late Jurassic and Early Cretaceous, 155 to 

115 Ma (fig. 3.35G). Large calc-alkaline plutons 

were emplaced during the 155- to 115-Ma interval, 

beginning particularly around 145 Ma (N.W. Walker, 

written commun., 1989). By 155 Ma, the crust apparently had thickened; the onset of plutonism may be 

related to a major change in plate motions. The Blue 

Mountains island arc and the site of accretion on the 

North American craton probably reached their 

present approximate latitude by the earliest Cretaceous. Although most accretion of the island arc to 

the North American craton occurred during the latter 

part of this time interval, especially in the 120- to 

115-Ma interval (Lund and Snee, 1988), the accretion 

process itself probably began somewhat earlier (fig. 

3.35F). Accretionary events such as those outlined by 

Wernicke and Klepacki (1988), which were dominated by strike-slip faults, may have occurred and would 

account for the absence of rocks that must have existed between the North American craton and the 

Blue Mountains island arc. During accretion, the already telescoped and amalgamated terranes were 

probably overridden by the North American craton 

from the east, and extensive wrench faulting may 

have occurred along the suture zone between the arc 

and the North American craton. Following accretion, 

the Blue Mountains island arc was rotated clockwise 

(Wilson and Cox, 1980), covered by Tertiary volcanic 

rocks, and, since the late Miocene, differentially uplifted and eroded.

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3. PETROLOGY OF PRE-TERTIARY IGNEOUS ROCKS IN THE BLUE MOUNTAINS REGION 209

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4. HIGH-PRESSURE, LOW-TEMPERATURE SCHISTOSE ROCKS 

OF THE BAKER TERRANE, NORTHEASTERN OREGON

By ELLEN M. Bisnop1

CONTENTS

Abstract 

Introduction: The Baker terrane 

Acknowledgments - 

High-pressure, low-temperature schists of the Baker terraneBlueschist near Mitchell, Oregon 

Greenhorn Mountains 

Aldrich Mountains (serpentinite-matrix) melange and 

Mount Vernon area 

Schistose rocks of Mine Ridge 

Schistose rocks of Rhea Creek 

Discussion 

Conclusions 

References cited 

Page

211

211

213

213

213

214

216

216

216

217

218

218

ABSTRACT

Schistose rocks are present at widely separated localities within 

the serpentinite-matrix melange of the Baker terrane, northeastern 

Oregon. These schists vary in lithology and metamorphic grade, but 

most contain minerals whose compositions suggest initial metamorphism under conditions at or approaching blueschist-facies conditions. This high-pressure metamorphism was succeeded and 

overprinted by lower pressure metamorphism, usually in greenschist facies. Amphiboles in the high-pressure schists range from 

glaucophane to barroisite and winchite. Microprobe analyses of 

Na/Al relations in amphiboles from blueschists near Mitchell, Oreg. 

suggest initial metamorphism at greater than 7 kilobars (kb), with 

pressure subsequently decreasing to about 5 kb. At Pleasant Hill, 

south of Mount Vernon, blue amphiboles that are near crossite in 

composition suggest initial pressures of 5 kb, followed by pressures 

of 4 to 5 kb. Schists from the Greenhorn Mountains have a similar 

history, but the compositions indicate initial pressures of about 6 to 

5 kb. Blue-green amphiboles from schistose rocks in the Mine Ridge 

area to the southeast suggest initial pressures of about 5 kb. Schists 

from the Rhea Creek area in north-central Oregon near Heppner 

contain no relict high-pressure minerals, but may have been strongly overprinted by contact metamorphism.

This study suggests that some schists of the Baker terrane originated at high pressures, most likely in a subduction zone, and 

were affected by later conditions at decreased pressures and ele-

^2040 Highway 203, Union, OR 97883

vated temperatures. The rarity of true glaucophane blueschists in 

the Baker terrane may indicate that fairly rapid, high shear-stress 

subduction was characteristic of the Blue Mountains island-arc 

subduction zone during the Permian and Early Triassic. Later retrograde greenschist metamorphism may have been related to longterm incorporation in a fore arc under greenschist conditions or a 

shift in the locus of volcanism and the direction of subduction.

INTRODUCTION: THE BAKER TERRANE

The Blue Mountains of eastern Oregon contain a 

complex assemblage of tectonically emplaced and juxtaposed late Paleozoic and early Mesozoic terranes. 

These terranes are considered related to the Blue 

Mountains island arc (Vallier and Brooks, 1986) and 

are named the Wallowa, Baker, Grindstone, Izee, and 

Olds Ferry terranes (Silberling and Jones, 1984; 

fig. 4.1).

The Baker terrane is a disordered mixture of oceanic and island-arc sedimentary, volcanic, and plutonic rocks ranging from Devonian to Triassic in age. 

It has been characterized as an oceanic melange 

(Dickinson, 1979), a relatively coherent fore arc 

(Mullen, 1985), or possibly an ocean-floor environment (Vallier and others, 1977) associated with the 

Blue Mountains island arc. Most rocks in the terrane 

have been metamorphosed to greenschist facies; virtually all are deformed. The location and orientation 

of any subduction zone associated with the Blue 

Mountains island arc is still a matter of conjecture.

The rock types of the Baker terrane are diverse. 

Sedimentary rocks are dominantly chert or siliceous 

argillite. Graywacke and conglomerates, with clasts 

ranging from about 1 cm in fine-grained rocks to nearly 1 m in boulder conglomerates, are distributed 

throughout the terrane. Shales and quartz-rich sandstones are rare. Turbidite sandstones are unknown. 

Disrupted ophiolitic assemblages (greenstone, gabbro, 

layered gabbro, and serpentinized peridotite) and tectonic slices are present throughout the Baker terrane. 

Coherent ophiolitic blocks include the Canyon Moun211
120

Gand Saddle nixie Butte

Celebration Mine

Canyon 

Mountain

Ironside Mountain

FIGURE 4.1. Location map showing region of study and places mentioned in chapters 4 and 5 (Bishop, this volume). Broad gray lines are terrane b

modified from Silberling and Jones (1984) by Vallier (chap. 3, this volume); queried where location uncertain.
4. HIGH-PRESSURE, LOW-TEMPERATURE SCHISTOSE ROCKS OF BAKER TERRANE, NORTHEASTERN OREGON 213

tain Complex (Thayer, 1963; Leeman and others, 

chap. 1, this volume; Bishop, chap. 5, this volume).

The igneous rocks have tholeiitic, calc-alkaline, and 

alkalic compositions. The Canyon Mountain Complex is 

interpreted as island-arc basement (Leeman and others, chap. 1, this volume; Bishop, chap. 5, this volume).

Sheared serpentinite serves as a matrix in melanges or mixed-rock zones that occur from Mitchell at 

the western extremity to the Burnt River on the east 

(fig. 4.1). The nature of these serpentinite-matrix melanges is unresolved. They have been interpreted as 

diapiric upwellings of hydrated peridotite in a forearc environment (Bishop, 1988) and as faulted zones 

of structural weakness (Kays and others, 1988). 

Knockers within the serpentinite-matrix melanges 

include all lithic components of the Baker terrane as 

well as a variety of amphibolites and high-pressure, 

low-temperature schists.

ACKNOWLEDGMENTS

This study was partly funded by Sigma Xi (82-256) 

and the Oregon Department of Geology and Mineral 

Industries. Discussions with Howard Brooks, M. Alan 

Kays, Mark Ferns, Tracy Vallier, and R.G. Coleman 

were helpful. Reviews by Mary Donato and Bernard 

Evans significantly improved the manuscript.

HIGH-PRESSURE, LOW-TEMPERATURE SCHISTS 

OF THE BAKER TERRANE

The high-pressure, low-temperature schistose rocks 

of the Baker terrane are important because they provide insight into the pressure and temperature regimes that affected the Baker terrane, whether fore 

arc or subduction zone. Blueschist and other highpressure, low-temperature metamorphic rocks are associated with melange assemblages, subduction complexes, and accreted exotic terranes elsewhere, 

including the Franciscan Complex (California), the 

North Cascades (Washington), the Klamath Mountains (California and Oregon), the Sanbagawa terrane (Japan), and Otago, New Zealand. Thus, it is 

logical to expect blueschist or related rocks to occur 

in the Baker terrane, which represents a fore-arc or 

subduction complex for the island-arc-derived terranes of northeastern Oregon.

In this study, I examined a number of schistose 

rocks from the Baker terrane that contain blue amphiboles. The goal was to determine whether there 

are any blueschists in the Baker terrane beyond 

those at the already-defined localities near Mitchell 

and Toney Butte (fig. 4.1, Swanson, 1969). I also considered what limits of pressure and temperature 

could be estimated using amphibole compositions. 

This paper briefly describes schistose rocks of the 

Baker terrane and their mineral compositions, and 

speculates on their origins and possible significance.

Estimation of metamorphic pressure and temperature in this study is based upon amphibole chemistry. 

Coleman and Papike (1968) noted that with increasing metamorphic grade the Ca content of glaucophane increases. Brown (1977) found that Ca 

amphibole crystals from high-pressure terranes contain significantly more Na in X(M4) sites (octahedrally coordinated sites occupied by the relatively large 

cations Ca, Na, or K) than amphiboles from low-pressure terranes, and that where critical buffers are 

present, the Na/Ca ratio of amphibole is fixed by 

pressure. Ernst (1968), Liou and others (1975), and 

Brown (1977) observed that A1IV (Al in (Si,Al)O4 tetrahedra within the crystal) increases with temperature in amphiboles. This reaction uses the same 

buffers as the Na/Ca reaction and provides a relative 

geothermometer.

Progressive changes in metamorphic conditions can 

be determined from zonation in amphibole composition. On the basis of these estimates and the composition of other minerals, a metamorphic history is 

suggested for each locality.

BLUESCHIST NEAR MITCHELL, OREGON

The lawsonite blueschist of Meyers Canyon near 

Mitchell in central Oregon (fig. 4.1) was described by 

Swanson (1969). Exposures of Permian to Triassic 

rocks in Meyers Canyon include crystalline marble, 

pelitic schist, quartzite, chert, and mafic metavolcanic rocks, with exposure of serpentinite at the nearby 

blueschist location at Toney Butte. The blueschist at 

Meyers Canyon has been dated by Hotz and others 

(1977) at about 220 Ma using the K-Ar method.

The blueschist of Meyers Canyon contains feathery 

to idioblastic blue amphibole, with pleochroic colors 

ranging from light yellow to pale blue in the alpha 

optical direction and from light purple to pale blue in 

the gamma direction. On the basis of optics, Swanson 

(1969) classified these amphiboles as crocidolite to 

crossite with compositions close to glaucophane (see 

table 4.1). Very fine, slightly corroded prismatic lawsonite is dispersed throughout the rock and also forms 

fine aggregates. White mica (paragonite) also forms 

aggregates and occasional lenticular crystals. Chlorite 

occurs around amphiboles and micas. Quartz and 

albite are groundmass constituents. Small, prismatic 

crystals of Al-rich pumpellyite (see table 4.2) are 

present in mafic schists.
214 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 4.1. Compositions ofamphiboles in schistose rocks of the Baker terrane, northeastern Oregon 

[Values in weight percent. FeO*=FeO+Fe20^; , not determined]

Location 

Analyzed 

mineral --

Si02 - 

TiO2 ----- 

A12Q3 - 

FeO* 

MnO -----

MgO 

CaO--

Na2O - 

K20- 

Total 

Meyers Canyon, 

near Mitchell

Glaucophane 

core

53.6 

.1 

10.2 

14.4 

.2 

9.3 

5.6 

4.9 

.1 

.0 

.0

98.5

Glaucophane rim

54.8 

.1 

9.8 

13.4 

.2 

9.1 

7.4 

4.4 

.1 

.0 

.0

99.3

Mount Vernon area 

(Aldrich Mountains)

Amphibole 

core

53.1 

.0 

4.1 

19.7 

.2 

9.7 

3.5 

5.5 

.1 

.0 

.0

97.9

Amphibole 

rim

52.7 

.2 

4.0 

17.3 

.3 

11.7 

5.2 

4.4 

.1 

.1 

.0

97.9

Bennett Creek, in the 

Greenhorn Mountains

Barroisite 

core

55.0 

.0 

5.5 

17.4 

.3 

12.0 

4.4 

4.7 

.1 

.0 

.0

99.3

Barroisite 

rim

54.5 

.0 

2.4 

12.8 

.3 

16.4 

10.3 

1.6 

.6 

.1 

.0

99.6

Mine Ridge

Amphibole 

in 

amphibole 

schist

47.5 

.6 

10.5 

14.4 

.2 

13.1 

10.1 

2.7 

.6 

.2 

.0

98.3

Amphibole 

in mica 

schist, 

core

47.0 

.3 

11.0 

14.2 

.2 

11.8 

12.0 

1.3 

.6

99.3

Rhea Creek

Hornblende in 

biotite 

schist

44.2 

.2 

14.3 

17.0 

.0 

9.7 

12.5 

1.1 

.0

98.8

Blue 

magnesiogrunerite 

in 

graphite 

schist

53.7 

.0 

.7 

25.1 

.0 

17.8 

.6 

.0

99.7

Large idioblastic amphiboles were selected for microprobe core and rim analysis. Results (table 4.1) show 

that amphiboles from the blueschist of Meyers Canyon 

are zoned from sodic cores to more calcic rims (fig. 4.2). 

Analyses of cores gave an average composition of 

(NaL7,Caa3)(Fe2 .o,MgL8,AlL2)Si8O22(OH)2 (crossite). 

Rim analyses vary widely in the amount of CaO from 

4.5 to 8.6 weight percent which may be a function of 

the precise site analyzed. The average rim composition 

is (NaL3Ca0 7)(Fe! 5MgL7Al0 7)Si8O22(OH)2j which is 

intermediate between actinolite and winchite (Leake, 

1968). A1IV is absent in all analyses.

1.8

1.4

1.0

0.6

0.2

EXPLANATION

Meyers Canyon, near Mitchell

Bennett Creek in Greenhorn Mountains 

Mine Ridge

I_______I_______I

^f

0.2 0.6 1.0 1.4

Ca ATOMS PER FORMULA UNIT

1.8

FIGURE 4.2. Stoichiometric Na/Ca ratios for amphiboles from 

three eastern Oregon localities. Lines connect core (upper) 

and rim (lower) compositions.

The transition from low Ca content in the core to 

moderate Ca content in the rim suggests increased 

miscibility of amphiboles with time, and hence a 

probable increase in temperature (fig. 4.2). Plots of 

amphibole NaM4 (Na in M4 sites) versus A1IV (Al in 

tetrahedral sites) after Brown (1977) indicate that 

pressures during growth of both core and rim exceeded 7 kb. The restriction of Al to six-fold coordination 

suggests that the temperatures of metamorphism 

were relatively low.

GREENHORN MOUNTAINS

The Greenhorn Mountains are midway between 

Mitchell at the western edge of the Baker terrane 

and the Burnt River on the east (fig. 4.1). Permian 

metasedimentary rocks and volcanic greenstones of 

tholeiitic, calc-alkaline, and alkalic compositions are 

juxtaposed with and tectonically intruded by metagabbros and serpentinized peridotites. An oceanic 

component is represented by chert of the Elkhorn 

Ridge Argillite and pillowed greenstones that occur 

as knockers. Alkalic pillowed greenstones that may 

be related to arc rifting are also intercalated with 

cherts (Mullen, 1985). Arc-derived metasedimentary 

rocks include fine-grained graywackes, shales, conglomerates, and tuffaceous pyroclastic rocks. Most 

greenstones associated with or adjacent to these metasedimentary rocks have calc-alkaline geochemical 

and mineralogical characteristics (Mullen, 1985). The 

serpentinite-matrix melange in the Greenhorn Mountains contains clasts of all lithic types and occurs in 

several broad zones.
4. HIGH-PRESSURE, LOW-TEMPERATURE SCHISTOSE ROCKS OF BAKER TERRANE, NORTHEASTERN OREGON 215

TABLE 4.2. Compositions of minerals (excepting amphiboles) in schistose rocks of the Baker terrane,

northeastern Oregon

[Values in weight percent. FeO*=FeO+Fe2&j; , not determined]

Location 

Analyzed 

mineral 

Si02 -----

AloOq 

£t O

TiO2 - 

FeO* -

MnO 

MgO 

CaO 

Na2O -

K20 

Total -----

Meyers Canyon, 

near Mitchell

Pumpellyite 

from blueschist

37.00 

29.03 

.16 

2.57

1.45 

.21.73 

.12 

.01

92.08

Mine Ridge

Epidote

38.82 

24.28 

.15 

11.94 

.01 

.05 

23.64

99.01

Bennett Creek, in the Greenhorn Mountains

Epidote

38.20 

23.31 

.01 

12.40 

.12 

.06 

23.43

97.47

Garnet, rim

38.77 

22.12 

.04 

23.01 

4.76 

.80 

11.02

100.52

Garnet, 

intermediate 

zone

37.81 

21.69 

.05 

22.50 

5.71 

.69 

11.69

100.14

Garnet, core

38.31 

21.36 

.07 

21.06 

7.27 

.59 

11.76

100.15

The best known schists in the Greenhorn Mountains are restricted to an area of about one-half 

square kilometer at the head of Bennett Creek 

(fig. 4.1). There are two lithic types a finely crystalline, highly foliated mica schist and a more massive 

blue-grey greenstone (Mullen, 1979a). The protolith 

of the mica schist was probably a pelitic sediment; 

the more massive rock contains relict pyroclastic 

fragments that suggest origin from a tuff.

The mica schist contains segregation bands of clear, 

granoblastic quartz and clouded, untwinned albite alternating with layers of white mica. The mica contains 

an average of 2 weight percent Na2O (phengite). The 

micaceous bands include fine stringers and veins of iron 

oxides and discontinuous layers of graphite and chlorite. Pretectonic garnet and sphene are broken and 

stretched along axial planes. The garnets are small, 

corroded and embayed by quartz, jacketed by a thin 

rim of chlorite, and restricted to Al-rich quartz-albite 

layers. They are somewhat enriched in MnO (4 to 7 

weight percent; table 4.2). Epidote has a high content 

of A12O3 (table 4.2).

About 20 percent of the Bennett Creek outcrops 

are green tuffaceous schist. The schist contains chlorite, quartz, albite, epidote, and about 10 percent 

blue-green amphibole. Crystalloblasts of epidote commonly cluster around amphiboles. Sphene, paragonite, and apatite are accessory minerals.

The blue-green amphibole has prismatic basal sections that are light straw yellow in the alpha optical 

direction and purplish blue in the beta direction. 

Elongate grains parallel to the (001) plane are light 

greenish yellow to light green-blue in the gamma direction. Amphiboles are commonly bent, with broken 

segments rotated up to 35°.

Microprobe analyses of amphiboles from the Bennett 

Creek locality (table 4.1) indicate an average composition of (NaCa)(Fe2Mg2Al)(Si7Al)O22(OH)2 . This composition is similar to that of barroisites from the Ligurian 

Alps of Italy (Ernst, 1976) and the Sanbagawa II 

terrane of Japan (Toriumi, 1975). Extremes in 

composition are found at the core and rim. Core compositions are highly sodic, containing 4.7 weight percent 

Na2O. The average core composition is (Naj 4Ca0 6) 

(Fe2Mg2 2A10 8)(Si7 8A10 2)O22(OH)2 . This composition 

represents an intermediate solid solution of the glaucophane-hornblende series (Klein, 1969; Brown, 1974; 

Katagas, 1974) or winchite (Leake, 1968). Rim compositions are much less enriched in Na, with only 1.6 

weight percent Na2O. The average rim composition is 

(Na0.4CaL6)(FeL6Mg0.4XSi7 . 8Al02)022(OH)2 . A1IV is 

absent in cores but enriched in rims. The rim composition of amphiboles in the schists of the Bennett Creek 

area resembles that of a sodic actinolite. Amphibole 

compositions suggest formation of cores at high pressures and low temperatures approaching or at blueschist-facies conditions. According to Brown's (1977) 

plot of NaM4 versus A1IV for calcic amphiboles, the 

schists of the Bennett Creek area were initially metamorphosed at 6 to 7 kb at relatively low temperatures 

(possibly 350°C) and later experienced conditions of 4 

to 5 kb at slightly increased temperatures (possibly 

400°C).

The compositions of other minerals support these 

conditions. Garnet is zoned from cores rich in MnO, 

suggesting low temperatures, to rims with less MnO, 

indicating increasing temperatures (Miyashiro, 1953). 

Phengitic micas are characteristic of high-pressure 

terranes in blueschist and lower greenschist facies 

(Velde, 1965).
216 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

ALDRICH MOUNTAINS (SERPENTINITE-MATRIX) MELANGE 

AND MOUNT VERNON AREA

Extensive serpentinite-matrix melanges occur across 

more than 1,000 km2 in east-central Oregon south and 

southwest of the Greenhorn Mountains (fig. 4.1). They 

include serpentinite-matrix melange of Miller Mountain, west of the Canyon Mountain Complex (Brown 

and Thayer, 1966), and melange south of Mount Vernon, collectively and informally termed the Aldrich 

Mountains (serpentinite-matrix) melange by Carpenter and Walker (1992). Similar melange is found near 

the small community of Antone. The components of 

these serpentinite-matrix melanges include chert, 

limestone, greenstone of both tholeiitic and calc-alkaline composition, metagabbro, peridotite, and a variety 

of metamorphic rocks, including amphibolites. Metamorphic clasts in the Aldrich Mountains melange include epidote and garnet amphibolites with schistose 

fabrics, and greenstones in lower greenschist facies 

(Carpenter and Walker, 1992). The amphiboles in the 

amphibolites are blue-green barroisites to winchites or 

brown hornblende. Exchange thermometry data and 

amphibole AlIV/Si ratios indicate that epidote amphibolites with blue-green barroisitic hornblende formed at 

temperatures of 550 to 650°C and pressures just below 

5 kb, whereas garnet amphibolites with brown hornblendes formed at temperatures above 700°C and pressures of about 4 kb (Carpenter and Walker, 1992).

The Aldrich Mountains melange also contains 

schists not reported in Carpenter and Walker's (1992) 

work. These rocks are blue-gray schists exposed 

about 12 km south of Mount Vernon. They contain 

blue amphiboles with typical core compositions of

These amphibole cores are high in Na2O (5.5 weight 

percent) but low in A1IV. They are chemically close to 

riebeckites, but have a pleochroic scheme and overall 

appearance more like those of glaucophane and crossite. Rim compositions of these blue amphiboles, and 

of small, euhedral amphiboles throughout the rock, 

are blue-green barroisite to winchite, with Na2O of 

3.5 to 4.5 weight percent and less A1IV than the cores 

(table 4.1).

SCHISTOSE ROCKS OF MINE RIDGE

Schistose rocks at Mine Ridge (fig. 4.1) were mapped 

and described by Lowry (1968). They are described in 

some detail by Hooper and others (chap. 11, this volume). These rocks comprise a lithologically diverse assemblage of Permian to Triassic pelitic and mafic 

schists and minor chert included in or associated with 

serpentinite of probable Triassic age. The schistose 

rocks occur in serpentinite matrix and are exposed on

the north and west sides of Mine Ridge as well as on 

low ridges about 5 mi south of Hereford (fig. 4.1).

Lowry (1968) obtained K-Ar whole-rock ages of 142 

to 150 Ma for the Mine Ridge rocks. He attributed this 

relatively young age to the contact metamorphic effects of a nearby pluton (Grouse Creek pluton) that he 

believed to be Jurassic in age. However, recent work 

on this intrusion has shown that it is an epizonal body 

33 to 36 Ma in age (Houseman, 1983). No contact effects are evident in the Mine Ridge rocks. Therefore, 

the 142- to 150-Ma age determination must be due to 

a factor other than contact metamorphism, and may 

reflect Ar loss from the Paleozoic to Triassic rocks 

heated to low temperatures by the adjacent intrusion. 

Alternatively, it may reflect a real, relatively young 

age of rocks entrained in a long-lived, high-shearstress subduction-zone environment.

Mica schists contain alternating bands of quartzalbite and lenticular laths of chlorite cut by fine 

strings of graphite. Albite porphyroblasts are twinned 

and deformed. Blue to light-blue-green actinolite is 

also bent and fractured. Epidote is restricted to maficrich bands and is Al-rich (table 4.2). Small aggregates 

of green biotite were the last minerals to form.

Mafic amphibole schists contain coarse amphiboles 

with pleochroic colors ranging from blue to bluegreen in the gamma direction and from blue to 

straw-yellow in the beta direction. These amphiboles 

contain inclusions of clear epidote, have rims altered 

to actinolite and chlorite, and are deformed. White 

mica is rare and phengitic and occurs as small, isolated fibrous crystals. Granular patches of small 

green biotite are found within interstitial chlorite.

Amphiboles from mica schists and mafic schists 

have significantly different compositions (table 4.1). 

Those from mica schists are intermediate between 

actinolite and hornblende, but are best characterized 

as actinolite because of generally low Al content and 

high Na content (fig. 4.3). The average formula is 

(Na0 3CaL7)(Mg2 7Fe! 8A10 5)Si8O22(OH)2 . The composition of these amphiboles is closest to that of barroisites reported from the Sanbagawa III terrane by 

Toriumi (1975)and Ernst and others (1970). On a diagram of NaM4 versus A1IV (after Brown, 1977), they 

plot in the same field as Sanbagawa barroisites, suggesting a pressure range of 5 to 6 kb (fig. 4.3). The 

complete buffering assemblage is present in the 

mafic amphibole schists, and thus this pressure may 

represent a maximum.

SCHISTOSE ROCKS OF RHEA CREEK

Schistose rocks associated with serpentine and 

metagabbro are exposed in a northeast-trending belt
4. HIGH-PRESSURE, LOW-TEMPERATURE SCHISTOSE ROCKS OF BAKER TERRANE, NORTHEASTERN OREGON 217

of pre-Tertiary outcrops approximately 30 km southeast of Heppner, Oreg., along Rhea Creek and near 

Copple Creek (fig. 4.1). These metasedimentary rocks 

include phyllite, pelitic (mica) schists, graphite 

schists, amphibolite, biotite gneiss, marble, and metachert (quartzite) (Shorey, 1976). The schistose rocks 

of Rhea Creek are complexly deformed, containing 

evidence for at least two deformational events. In 

their area of exposure, these rocks seem to be interlayered with marbles and quartzites in coherent 

stratigraphic order. Similar and probably related 

schists are exposed along U.S. Highway 395 approximately 50 km south-southwest of Pendleton.

The mineral assemblages of the Rhea Creek metasedimentary rocks indicate high-grade-greenschist- 

(Shorey, 1976) to amphibolite-facies metamorphism. 

Pelitic schists contain varying amounts of biotite, hornblende, quartz, plagioclase (oligoclase), muscovite, almandine garnet, tourmaline, zircon, and carbonate 

minerals. Amphiboles in the mafic schists and amphibolites are hornblendes with 0.9 to 1.1 weight percent 

Na2O. The average composition is (Na0.3Ca17) 

(Fe2.4Mg1 7Al)(Si6Al2)O22(OH)2. Compositional zoning 

is not evident. A plot of NaM4 versus A1IV suggests 

metamorphism at 3 to 4 kb (fig. 4.3).

FlGUEE 4.3. NaM4 (Na in M4 sites) versus Al (Al in tetrahedrally coordinated sites) for amphiboles from four eastern 

Oregon localities: Meyers Canyon, Bennett Creek, Mine 

Ridge, and Rhea Creek. Shaded areas show fields for amphiboles from the Sanbagawa terrane of Japan (Toriumi, 1975; 

Ernst and others, 1970) and from Otago, New Zealand 

(Brown, 1974). Pressures in kilobars (kb) as determined by 

Brown (1977). Arrow connects core (upper triangle) and rim 

(lower triangle) compositions in single amphibole from Bennett Creek area.

Schists that have a greater percentage of quartz, 

feldspar, and especially graphite contain a blue grunerite rather than hornblende (table 4.1). The grunerite's 

average formula is Cao.i(Fe3Mg4)Si8O22(OH)2. The uncharacteristic pleochroism blue in the alpha direction 

and clear in the gamma direction may be due to the 

extremely reduced character of the iron in the graphite 

schists. The blue color is more intense in amphiboles 

within graphite concentrations. Some of the Rhea Creek 

rocks are blue to steel-blue in outcrop because of the 

combination of graphite, muscovite, and blue grunerite. 

Tourmalines in these rocks are also bluish, primarily 

because of the reduced iron.

Several factors suggest that the schistose rocks of 

Rhea Creek were initially formed by high-pressure, 

low-temperature metamorphism. The rocks occur 

along regional strike of the blueschists near Mitchell. 

Their fabric, deformational history, and lithic types 

are similar to those of the Mitchell rocks. They are 

associated with and probably enclosed by serpentinite-matrix melange. And contact metamorphism and 

metasomatism by adjacent plutons may have produced the mineralogy now observed.

However, despite extensive field, petrographic, and 

microprobe investigation, no relict high-pressure, 

low-temperature (blueschist-facies) minerals have 

been found. Thus, the correlation of the schistose 

rocks along Rhea Creek with the schistose rocks near 

Mitchell is very tentative, and the Rhea Creek rocks 

may have an entirely different origin or late history.

DISCUSSION

The schistose rocks at four localities examined in 

this study (Meyers Canyon near Mitchell, Bennett 

Creek in the Greenhorn Mountains, the Mount Vernon area, and Mine Ridge) have similar metamorphic 

histories: initial high-pressure metamorphism followed by deformation and a second metamorphism at 

higher temperatures and lower pressures. At Mitchell, metamorphic conditions remained at blueschist 

facies. Eastward, the rocks of Bennett Creek, the 

Mount Vernon area (Aldrich Mountains melange), 

and Mine Ridge suggest lower initial pressures and 

higher initial temperatures, with increasing temperatures and decreasing pressures through time. At Bennett Creek, schists underwent metamorphism at 

pressures near 6 kb, followed by greenschist-facies 

metamorphism. The Mine Ridge rocks initially underwent lower-greenschist-facies metamorphism at 

approximately 5 kb, followed by upper-greenschist-facies to amphibolite-facies metamorphism.

The schists most likely formed initially in a subduction-zone environment. The absence of glaucophane in most blue-amphibole-bearing schists of the
218 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Baker terrane may indicate that subduction associated with these rocks was rapid and (or) that shear 

stresses were high. Where shear stresses exceed 100 

megapascals (MPa) and subduction rates are faster 

than 10 cm/yr, amphibolite-facies metamorphism will 

occur (Peacock, 1992). Stability of glaucophane, 

whether epidote or lawsonite subfacies, requires either slow subduction and high shear stress or rapid 

subduction and low shear stress (Peacock, 1992).

A possible model for the formation of the schistose 

rocks of the Baker terrane would associate these 

rocks with subduction-zone metamorphism in a rapidly moving, high-shear-stress environment. Variations in core compositions of blue amphiboles across 

the terrane could reflect variations in subduction 

rates through time and space, variations in shear 

stress within the subduction zone, and (or) variations 

in the depth and conditions of subduction and accretion. Amphibolites of the Aldrich Mountains melange 

assigned to a "dynamothermal aureole" by Carpenter 

and Walker (1992), as well as amphibolites at Mine 

Ridge (Hooper and others, chap. 11, this volume) and 

amphibolite metagabbro in the Greenhorn Mountains 

may have been generated along shear boundaries 

within the same subduction zone by mechanisms elucidated by Peacock (1987). A separate episode of 

thrusting and tectonism may have affected these 

rocks, but is not required for their formation.

The later metamorphism of the schists in the Baker 

terrane at higher temperatures and lower pressures seems linked to the continuing history of the 

Blue Mountains island arc. The greenschist-facies 

rims of amphiboles and the rarity of blueschist-facies 

minerals indicate that higher temperatures and lower 

pressures affected the schistose rocks. Although the 

higher temperature and lower pressure minerals may 

simply indicate that these rocks were "stalled" under 

greenschist conditions, the deformed fabric of these 

rocks, including the later amphiboles, suggests that 

the higher temperature processes were dynamic.

Additional work, including dating of metamorphic 

episodes, fine-tuning of metamorphic petrology, use of 

exchange thermometry to calculate equilibrium temperatures (difficult for the early episodes), and structural studies are needed to further unravel the 

history of these rocks. In turn, this understanding is 

needed for a better understanding of the history of 

the Baker terrane.

CONCLUSIONS

The principal conclusions of this study are as follows: 

(1) Most schistose rocks in the Baker terrane associated with serpentinite-matrix melange underwent

initial blueschist- or near-blueschist-facies metamorphism. Later metamorphism occurred at higher temperatures and lower pressures.

(2) The schists were most likely initially produced 

in a subduction-zone environment and were incorporated in rising serpentinite diapirs.

(3) The lack of true glaucophane amphiboles in the 

schists east of Mitchell may indicate rapid subduction (greater than 10 cm/yr) and (or) high shear 

stress in the subduction zone.

(4) The temperature and intensity of continuing 

metamorphism that affected these rocks generally increased eastward across the Baker terrane.

(5) Metamorphism dated at about 255 to 268 Ma 

by Carpenter and Walker (1992) is possibly related to 

reversal of subduction-zone polarity or inception of a 

new subduction-zone, or to another process with high 

ductile shear. The blueschists. at Mitchell (220 Ma; 

Hotz and others, 1977) did not experience this event 

and may be related instead to Triassic east-facing 

subduction (Vallier, 1992). The relevance of the 142- 

to 150-Ma event recorded in the schistose rocks of 

Mine Ridge (Lowry, 1968) is unclear.

(6) The schists of the Rhea Creek area are not 

clearly related to the other schists. However, factors 

such as their structure, location, and association suggest that they were initially high-pressure schists 

and may have an origin and history similar to those 

of the other schists in the Baker terrane, particularly 

the schists exposed near Mitchell.

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Bishop, E.M., 1988, Igneous petrology of the Baker melange terrane and the Canyon Mountain complex, eastern Oregon: Geological Society of America Abstracts with Programs, v. 20, 

p. 407.

Brown, C.E., and Thayer, T.P., 1966, Geologic map of the Canyon 

City quadrangle, northeastern Oregon: U.S. Geological Survey 

Miscellaneous Geological Investigations Map 1-477, scale 

1:250,000.

Brown, E.H., 1974, Comparison of the mineralogy and phase relations of blueschists from North Cascades, Washington, and 

greenschists from Otago, New Zealand: Geological Society of 

America Bulletin, v. 85, p. 333-344.

 1977, The crossite content of Ca amphibole as a guide to the 

pressure of metamorphism: Journal of Petrology, v. 18, p. 53-72.

Carpenter, PS., and Walker, N.W, 1992, Origin and tectonic significance of the Aldrich Mountains serpentinite matrix melange: 

Tectonics, v. 11, p. 690-708.

Coleman, R.G., and Papike, J.J., 1968, Alkali amphiboles from the 

blueschists of Cazadero, California: Journal of Petrology, v. 9, 

p. 105-122.

Dickinson, W.R., 1979, Mesozoic fore-arc basin in central Oregon: 

Geology, v. 7, p. 166-170.

Ernst, W.G., 1968, Amphiboles: New York, Springer-Verlag, 125 p.

 1976, Mineral chemistry of eclogite and related rocks from 

the Voltri group, western Liguria, Italy: Schweizerische Min-
4. HIGH-PRESSURE, LOW-TEMPERATURE SCHISTOSE ROCKS OF BAKER TERRANE, NORTHEASTERN OREGON 219

eralogische und Petrographische Mitteilungen, v. 56, p. 293- 

343.

Ernst, W.G., Seki, Y., Onuki, H., and Gilbert, M.C., 1970, Comparative study of low-grade metamorphism in the California 

Coast Ranges and the outer metamorphic belt of Japan: Geological Society of America Memoir 124, 276 p.

Hotz, P.E., Lanphere, M.A., and Swanson, D.A., 1977, Triassic 

blueschist from northern California and north-central Oregon: 

Geology, v. 5, p. 659-663.

Houseman, M.D., 1983, Petrology and alteration of the Grouse 

Creek granodiorite porphyry: Pullman, Washington State University, M.S. thesis, 166 p.

Katagas, C., 1974, Alkali amphiboles intermediate between actinolite and riebeckite: Contributions to Mineralogy and Petrology, 

v. 46, p. 257-264.

Kays, M.A., Bishop, E.M., Walker, N.M., and Blackwell, D.L., 

1988, Sodic amphibole-bearing schistose rocks of the Baker 

terrane: Geological Society of America Abstracts with Programs, v. 20, p. 423.

Klein, C., 1969, Two amphibole assemblage in the system actinolite-hornblende-glaucophane: American Mineralogist, v. 54, 

p. 212-237.

Leake, B., 1968, Classification of amphiboles: American Mineralogist, v. 63, p. 1023-1052.

Liou, J.C., Ho, C.O., and Ven, T.P, 1975, Petrology of some glaucophane schists and related rocks from Taiwan: Journal of Petrology, v. 16, p. 80-109.

Lowry, W.D., 1968, Geology of the Ironside, Oregon, quadrangle: 

Oregon Department of Geology and Mineral Industries Open 

File Report.

Miyashiro, A., 1953, Calcium-poor garnet in relation to metamorphism: Geochimica et Cosmochimica Acta, v. 4, p. 179-208.

Mullen, E.D., 1979a, Geology of the Greenhorn Mountains, northeastern Oregon: Corvallis, Oregon State University, M.S. thesis, 372 p.

 1979b, Temperature-pressure progression in high-pressure 

Permo-Triassic rocks of northeast Oregon: Eos (Transactions, 

American Geophysical Union), v. 51, p. 70.

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implications for terrane origin: Geology, v. 13, p. 131-134.

Peacock, S.M., 1987, Creation and preservation of subduction-related inverted metamorphic gradient: Journal of Geophysical 

Research, v. 92, p. 12,763-12,781.

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P-T-t paths in subduction shear zones: Journal of Geophysical 

Research, v. 97, p. 17,693-17,708.

Shorey, E.F., 1976, Geology of part of southern Morrow County, 

northeast Oregon: Corvallis, Oregon State University, M.S. 

thesis, 131 p.

Silberling, N.J., and Jones, D.L., 1984, Lithotectonic terrane maps 

of the North American Cordillera: U.S. Geological Survey 

Open File Report 84-523, 43 p.

Swanson, D.A., 1969, Lawsonite blueschist from north-central Oregon: U.S. Geological Survey Professional Paper 650-B, p. B8- 

Bll.

Thayer, T.P, 1963, The Canyon Mountain Complex, Oregon, and 

the mafic magma stem: U.S. Geological Survey Professional 

Paper 475-C, p. C82-C85.

Toriumi, Mitsuhiro, 1975, Petrological study of the Sambagawa 

metamorphic rocks the Kanto Mountains, central Japan: 

Tokyo, The University Museum, The University of Tokyo, 

99 p.

Vallier, T.L., 1992, Late Triassic change in convergence direction, 

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Vallier, T.L., and Brooks, H.C., 1986, Paleozoic and Mesozoic faunas of the Blue Mountains province A review of their geologic implications and comments on papers in the volume, in 

Vallier, T.L., and Brooks, H.C., eds., Geology of the Blue 

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U.S. Geological Survey Professional Paper 1435, p. 1-7.

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in eastern Oregon and western Idaho, in Stewart, J.H., 

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5. MAFIC AND ULTRAMAFIC ROCKS OF

THE BAKER TERRANE, EASTERN OREGON,

AND THEIR IMPLICATIONS FOR TERRANE ORIGIN

By ELLEN M. BiSHOP1

CONTENTS

Abstract________________________________ 221

Introduction 221

Acknowledgments 223

Canyon Mountain Complex 223

Tectonite peridotite unit 223

Transition zone: Layered peridotites and gabbros unit 224

Gabbro units 227

Relations among gabbro units 229

Late intrusive gabbro: The zone of infiltration 229

Origin of the zone of infiltration 230

Migmatitic and recrystallized gabbro:

Zone of plagiogranite genesis 231 

Origin and significance of the migmatitic zone:

Generation of plagiogranite 233

Geochemistry 236

Origin and tectonic setting 237

Other mafic and ultramafic components of the Baker terrane- 238

Implications for the origin of the Baker terrane 243

References cited 244

ABSTRACT

The Baker terrane is a disordered assemblage of Paleozoic and 

Triassic metaigneous and metasedimentary rocks incorporated in 

serpentinite-matrix melange that occurs across more than 10,000 

km2 of eastern Oregon. Serpentinites, peridotites, gabbros, metagabbros, and volcanic greenstones are substantial components. 

Along with the Canyon Mountain Complex, these rocks provide 

petrologic and geochemical clues to the nature and origin of the 

Baker terrane and its relation to other eastern Oregon terranes of 

similar age.

The Permian Canyon Mountain Complex is the only coherent 

ophiolite associated with the Baker terrane. Its petrology and geochemistry suggest an island-arc (arc-tholeiite to calc-alkaline) affinity. The Canyon Mountain Complex contains two features not 

reported in ophiolite s: gabbros that infiltrated into the tectonite 

peridotite, and well-defined migmatitic zones of partial melting 

and recrystallization of upper, hydrothermally altered gabbros that 

generated plagiogranite. These features suggest that the late- or

1 52040 Highway 203, Union, OR 97883

post-magmatic history of the Canyon Mountain Complex included 

shear, pressure release, minor intrusions, and localized partial 

melting that produced the Canyon Mountain plagiogranites and 

keratophyres.

Mafic volcanic and plutonic rocks of lower greenschist to amphibolite facies occur as knockers and tectonic slivers throughout 

the Baker terrane. The mineralogy and geochemistry of the volcanic rocks indicate formation in a variety of tectonic settings, including midocean ridge, island arc, and rift or transform. Many 

plutonic rocks of the Baker terrane have characteristics of normal 

midocean-ridge basalt (NMORB); others have mineralogy and compositions similar to those of the arc-related Canyon Mountain 

Complex.

A fore arc and (or) erosive high-shear-stress subduction zone are 

the most likely petrotectonic settings for amalgamation of the igneous components of the Baker terrane.

INTRODUCTION

The Paleozoic and Triassic Baker terrane of eastern 

Oregon is a disordered assemblage of basaltic greenstone, silicic greenstone, argillite, chert, graywacke, 

conglomerate, Tethyan and North American limestone 

of Devonian to Triassic age, metadiorite, gabbro and 

metagabbro, and peridotite. These components, as 

well as amphibolites and high- to moderate-pressure 

schists (Bishop, 1988; Bishop, chap. 4, this volume; 

Hooper and others, chap. 11, this volume), are incorporated in serpentinite-matrix melanges throughout 

the terrane. Because of the lithologic and metamorphic heterogeneity of terrane components, the nature 

and origin of the Baker terrane have been major 

problems of eastern Oregon pre-Tertiary geology.

The Baker terrane is associated in space and time 

with the Wallowa terrane, the Olds Ferry terrane, 

the Grindstone terrane, and the Izee terrane, as delineated by Silberling and Jones (1984; fig. 5.1). 

These are generally considered to be arc-related tectonostratigraphic packages, collectively referred to as 

the Blue Mountains island arc (Vallier, chap. 3, this 

volume). The Wallowa terrane is related to Permian 

and Triassic arc volcanism. The Olds Ferry terrane is

221
120° 

Gand Saddle Dixie Butte 

Pine Creek

Strawberry 

Mountain 

Wilderness Baldy Mtn 

Pine Creek Mtn 

Bear Skull Rims

FIGURE 5.1. Location map showing region of study and places mentioned in chapters 4 and 5 (Bishop this volume). Broad gray lines are terrane b

modified from Silberling and Jones (1984) by Vallier (chap. 3, this volume); queried where location uncertain.
5. MAFIC AND ULTRAMAFIC ROCKS OF BAKER TERRANE, EASTERN OREGON 223

a remnant of Late Triassic arc volcanism and sedimentation. The Grindstone terrane may represent an 

accretionary wedge (Blome and Nestell, 1992), and 

the Izee terrane records Triassic and Jurassic postmagmatic sedimentation in a fore-arc or related environment (Dickinson, 1979).

The Baker terrane has been interpreted as an oceanic 

basin (Brooks, 1979), a melange (Dickinson, 1979), and 

a relatively coherent fore arc associated with the magmatically and tectonically active Wallowa arc (Mullen, 

1983a). Its rocks are substantially more deformed and 

older than those of the Izee terrane to the south, and 

can be distinguished from those of the Grindstone terrane by the preponderance in the Baker terrane of 

metaigneous and schistose metamorphic components 

in a largely serpentinite matrix. The Baker terrane 

sedimentary package includes voluminous cherts of the 

Elkhorn Ridge Argillite and sheared siliceous sedimentary rocks of the Burnt River Schist. Both are considered pelagic units, although they contain volcaniclastic 

sedimentary rocks and conglomerates.

Evaluation of the petrology, geochemistry, and character of the ultramafic and mafic rocks (for example, 

peridotite, gabbro, and volcanic greenstone) is one 

route to a better understanding of the Baker terrane. 

This paper discusses the nature of mafic and ultramafic igneous components and their implications for terrane origin, beginning with the largest and most 

complete component the Canyon Mountain Complex.

ACKNOWLEDGMENTS

This work was funded in part by Geological Society 

of America Penrose grants 2271-77 and 2887-81, the 

Bowes Mining Co., and the Oregon Department of 

Geology and Mineral Industries. Discussions with 

H.C. Brooks, J.W. Hawkins, T.P. Thayer, and T.L. 

Vallier were helpful. The manuscript has benefited 

from reviews by Clifford Hopson, Cynthia Evans, 

W.P. Leeman, James Evans, and Lisanne Pearcy

CANYON MOUNTAIN COMPLEX

The Canyon Mountain Complex is a 4-km-thick 

Permian sequence of ultramafic rocks, gabbros, plagiogranites, and keratophyres that is exposed over 

more than 150 km2 in the Strawberry Mountain Wilderness, just south of John Day, Oreg. (figs. 5.1, 5.2). 

It has been dated at 278 to 261 Ma by K-Ar, U-Pb, 

and 40Ar/39Ar methods (Gerlach, and others, 1988; 

Leeman and others, chap. 1, this volume). Isotropic 

and layered gabbros have yielded Nd isochron ages of 

267 to 274 Ma (Gerlach and others, 1988).

The Canyon Mountain Complex is an important 

component of the Baker terrane because it is a nearly 

complete ophiolite sequence and because it contains 

some features rarely reported from similar packages of 

oceanic or arc-related rocks. On the basis of chemistry 

and lithic types, previous workers (Thayer, 1977; Himmelberg and Loney, 1980; Gerlach, and others, 1981; 

Mullen, 1983a) have considered the Canyon Mountain 

Complex to have formed in an island-arc environment. 

Previous investigators examined the economic geology 

(Thayer, 1976), the structure (Ave Lallemant, 1976), 

the origin of the plagiogranite (Gerlach, 1980), the petrology of the peridotite and gabbro (Himmelberg and 

Loney, 1980; Mullen, 1983a), the petrogenesis of the 

gabbro units (Mullen, 1983b), and the origin of the 

layered igneous rocks (Pearcy, 1987, 1991), and have 

made detailed geochemical evaluations (Leeman and 

others, chap. 1, this volume).

TECTONITE PERIDOTITE UNIT

Tectonite peridotite constitutes the lowermost unit 

of the Canyon Mountain Complex. It grades upward 

from glossy green sheared serpentinite at the base to 

relatively fresh rock that is less than 10-percent serpentinized in some locations. Its overall thickness is 

approximately 1.5 km. The principal lithic type (>90 

percent) is harzburgite, with olivine the dominant 

mineral phase (table 5.1). Mineral compositions are 

quite uniform throughout the tectonite peridotite (table 5.2). Elongated chrome spinel defines foliation in 

the harzburgite outcrops.

Lherzolite is rare in the Canyon Mountain Complex 

and constitutes less than 5 percent of the tectonite 

peridotite unit. It is in gradational contact with harzburgite and is most voluminous in the central part of 

the unit where gabbro dikes and veins are abundant.

Dunite most commonly occurs in lenses and elongate bodies from 1 to 100 m in length, but also occurs 

as veins, dikes, and small irregular blobs. Contacts of 

dunite with harzburgite are gradational. Foliation 

and structure defined in outcrop by chromite are parallel with structure in adjacent harzburgite. This 

finding supports a metasomatic origin for the dunite 

of the Canyon Mountain Complex as suggested by 

Dungan and Avci Lallemant (1977).

Dikes and sills of websterite, clinopyroxenite, and 

orthopyroxenite occur throughout the tectonite peridotite unit. Most are less than 20 cm wide. Several generations can be distinguished on the basis of crosscutting 

relations. Early dikes, commonly composed of clinopyroxene-rich websterite, are deformed conformably with 

the host peridotite. These dikes are crosscut by later, 

undeformed websterite or orthopyroxenite dikes. The 

adjacent peridotite is depleted in pyroxene within 2 to
224 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

10 cm of the early dikes. This relationship suggests 

that here, as well as in other ophiolites, the early 

websterite dikes are segregations of melt generated 

locally within the harzburgite (Boudier and Coleman, 

1981; Nicolas and Jackson, 1982).

Podiform chromitite bodies occur throughout the 

tectonite peridotite unit. They are usually oblong to 

lensoidal in shape and range from less than 1 m to 

slightly more that 60 m in length. There are two different spinel compositions (table 5.3). Chrome-rich, 

aluminum-poor spinel occurs in chromitite pods that 

do not contain clinopyroxene and are generally deep 

within the tectonite peridotite unit, whereas more 

aluminum-rich chromite is found in chromitites that

do 'contain clinopyroxene and are near the top of the 

tectonite peridotite unit. The two spinel compositions 

may be produced by fractionation, but the absence of 

intermediate compositions in the Canyon Mountain 

Complex suggests separate sources or interaction of 

Al-rich spinels with gabbroic magma (fig. 5.3).

TRANSITION ZONE: 

LAYERED PERIDOTITES AND GABBROS UNIT

A substantial thickness of interlayered, tectonized 

peridotite and gabbro forms the transition from tectonite peridotite to overlying cumulate and isotropic

118°47'30"

JOHN

-Zone of infiltration

44° 

22' 

00"

Ultramafic rock units 

Harzburgite 

Dunite 

Pyroxenite

Pyroxene-rich peridotite Includes 

wehrtite, olivine-rich gabbro and 

troctolite, rare websterite, Iherzolite

EXPLANATION 

Gabbroic rock units 

Unit G1 Olivine, two-pyroxene gabbro

[;. G2:-| Unit G2 Two-pyroxene gabbro with 

clinopyroxene dominant

[\:G3^| Unit G3 Gabbronorite and norite

I' G4 'I Unit G4 Hydrothermally altered gabbro

KNXXVj Recrystallized gabbros, migmatites, and 

diabase dikes

Plagiogranite rock unit

Plagiogranites Includes trondhjemite 

and quartz diorite

Contact Dashed where inferred, 

dot-dashed where gradational

Fault

Podiform chromitite lens

FIGURE 5.2. Generalized map showing ultramafic, gabbroic, and plagiogranite rock units of the Canyon Mountain Complex. Scale is

approximate.
5 MAFIC AND ULTRAMAFIC ROCKS OF BAKER TERRANE, EASTERN OREGON 

TABLE 5.1. Modes of analyzed samples from the Canyon Mountain Complex 

[tr, trace; , not observed]

225

sample

Lithic type 

Olivine 

Orthopyroxene 

Clinopyroxene 

Plagioclase 

Hornblende 

Actinolite 

PVilnritp ~

Hydrogrossular 

Serpentine 

Opaque minerals 

Accessory minerals 

Total ____ ________

449A

Harzburgite

39

29

2

___

___

9

21

6

1

100

jr/?fj

Dunite

78

2

___

__

__

._

14

6

100

Wehrlite

7

3

87

 _

1

 _

_

1

1

100

C7QA

Olivine

websterite

15

8

61

7

3

_

3

2

1

100

C'TflTD

Olivine

gabbro

_

15

20

47

12

__

6

tr

100

1 1 /"1/-1

Gabbro

dike

1

3

29

18

98

21

tr

tr

100

Unit Gl

(gabbro)

12

13

47

19

3

4

1

1

100

coc

Unit G2

(gabbro)

...

3

32

16

6

3

 to

27

100

707

Unit G3

(gabbro)

...

21

20

52

4

9

tr

1

100

714

Recrystallized

gabbro

_

17

15

61

3

9

tr

tr

100

TABLE 5.2. Compositions of silicate minerals from the tectonite peridotite unit of the Canyon Mountain Complex

[FeO^FeO+Fe20^; , not applicable]

Sample 

Mineral 

Mineral 

grain 

449A (harzburgite)

Olivine

1

Orthopyroxene

2 3

Clinopyroxene

4 5

Olivine

6 7

286 (Iherzolite)

Orthopyroxene

8 9

Clinopyroxene

10 11

Major-element oxides (weight percent)

Si02 

TiO2 - - 

A12Q3 

FeO* 

MnO 

MgO - 

CaO 

Na2O 

Cr203-~ 

NiO -

Total 

42.02 

.00 

.00 

5.11 

.02 

52.17 

.00 

.00 

.04 

.69

100.05

58.30 

.10 

1.19

4.40 

.08 

35.34 

.47 

.00 

.34 

.07

100.26

57.66 

.13 

1.18 

4.15 

.07 

35.33 

.46 

.00 

.46 

.13

99.57

53.15 

.29 

1.99 

1.18 

.06 

17.34 

23.60 

.16 

1.03 

.08

98.88

54.36 

.22 

1.97 

1.96 

.08 

17.38 

22.72 

.34 

1.14 

.05

100.19

40.34 

.00 

.04 

7.21 

.05 

52.20 

.00 

.00 

.00 

.45

100.29

40.33 

.00 

.01 

6.34 

.07 

52.58 

.00 

.00 

.00 

.53

99.92

57.11 

.10 

2.05 

5.51 

.14 

34.47 

.46 

.00 

.56 

.09

100.50

56.24 

.08 

1.75 

5.38 

.07 

35.22 

.25 

.00 

.47 

.13

99.58

53.22 

.10 

2.28 

1.38 

.06 

17.44 

23.59 

.14 

.76 

.11

99.10

53.25 

.11 

2.38 

1.49 

.03 

16.98 

24.05 

.19 

.80 

.02

99.30

Stoichiometry (number of ions per unit cell)1

Ti _______ 

.ru. Al 

Fe 

Mn 

Mg-~ 

Ca 

Na _______

Cr -

Ni

Total 

0.995 

.000 

.000 

.110 

.000 

1.886 

.000 

.000 

.000 

.080

2.991

1.985 

.003 

.048 

.125 

.002 

1.794 

.017 

.000 

.009 

.000

3.981

1.978 

.003 

.098 

.119 

.002 

1.807 

.017 

.000 

.012 

.004

3.987

1.946 

.008 

.086 

.036 

.002 

.947 

.926 

.011 

.030 

.002

3.991

1.963 

.006 

.084 

.059 

.002 

.935 

.879 

.024 

.032 

.001

3.982

0.978 

.000 

.000 

.146 

.000 

1.887 

.000 

.000 

.000 

.050

3.002

0.978 

.000 

.000 

.128 

.000 

1.902 

.000 

.000 

.000 

.050

3.001

1.954 

.002 

.080 

.157 

.004 

1.758 

.016 

.000 

.015 

.002

3.994

1.942 

.002 

.071 

.155 

.001 

1.813 

.009 

.001 

.003 

.003

4.001

1.944 

.002 

.098 

.042 

.001 

.950 

.923 

.010 

.021 

.003

3.998

1.930 

.003 

.103 

.045 

.001 

.935 

.951 

.013 

.023 

.000

3.995

Relative proportions (percent)

Fo 

En2 - 

Fs -

En3 

Wo 

94.5

93.5 93.8

2 

50 

48

3 

50

47

92.8 93.7

93.4 92.1

2 

50 

48

1

50 

49

1 Assuming 4 oxygen atoms per unit cell for olivine, 6 for Orthopyroxene and Clinopyroxene. 

2Enstatite component of Orthopyroxene. 

 ^Enstatite component of Clinopyroxene.
226 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 5.3. Compositions of spinels from podiform chromitites, Canyon Mountain Complex

Sample ----- - 410

Mineral grain 1 2

74

1 2

Rim Core

82 149 169 667

12 12 12 12

Major-element oxides (weight percent)

Si02 -----

A12Q3 - 

Ti02 -----

FeO* ___ ~

MnO -

MgO - 

Cr203 ----- 

NiO

Total 

0.10

1 Q 91

.14

14 97

00

15.01

56.27 

.20

99.43

0.13

1 o en

.14

14.75

10

15.86

56.13 

.22

100.01

0.10 

14.00

09

14 >V7

o/?

15.70

55.24 

.01

99.97

0.20

1 o an

.01

14.28

24

15.95

55.79 

.01

100.15

0.12

1 o 7-1

1C

1 Q 7G

to

15.99

54.81 

.00

98.74

0.13

1 Q QQ

.10

14.82

94

15.02

53.23 

.12

97.04

0.04 

13.84

.16

14.71

00

IE oo

55.72

.08

100.09

0.07 

18 94

7c

17 97

91

14 OQ

48.49 

.18

99.30

0.09 

17 90

69

17.67

91

14.16

48.36 

.19

99.27

0.12 

98 4Q

9Q

16.65

.16

15.71

37.25 

.09

98.76

0.06 

27.13

.38

17.19

.18

15.44

37.39 

.18

97.95

0.07 

28.75

.36

14.46

.20

17.01

38.09 

.21

99.15

0.01 

28.21

.06

12.81

.20

15.76

37.16 

.01

94.22

Relative proportions (percent)

Cr/(Cr+Al) 81.0 

Mg/(Mg+Fe) - 50.4

81.7 

50.5

79.8 

51.9

80.3 

52.8

80.3 

53.7

80.5 

50.3

80.1 

50.9

72.7 

44.9

73.0 

44.5

56.7 

48.5

58.0 

47.3

57.0 

54.1

56.8 

55.2

0.80

0.60

o,

m 

(^

O

0.40

0.20 0.40 0.60 

MgO / (FeO+MgO)

0.80

FIGURE 5.3. Compositions of chrome spinels in chromitite 

bodies, Canyon Mountain Complex. Filled circles, spinels 

from chromitite bodies with clinopyroxene; open circles, 

spinels from chromitite bodies in harzburgite that contain 

no clinopyroxene; filled triangles, accessory spinels of layered peridotites of the transition zone; open triangles, accessory spinels of harzburgite.

gabbro. This transition unit is referred to as the layered peridotites and gabbros unit. The rocks are well 

exposed on sharp, barren ridges and constitute about 

20 percent of the aerial exposure of the Canyon 

Mountain Complex. Their maximum thickness is less 

than 0.5 km.

Similar rocks in most ophiolite sequences are considered to be accumulations of early phases from 

fractional crystallization of basaltic magma. Such sequences are termed "metacumulates" by Coleman 

(1977). They achieve their distinctive textures and 

abrupt changes in composition through hypersolidus 

or subsolidus, metasomatic or metamorphic processes, including double diffusive convection (Irvine, 

1980). Deformation at high-temperature, subsolidus 

conditions accounts for the glide twinning and other 

lattice defects common in minerals of such rocks.

The fabrics of the layered peridotites and gabbros 

of the Canyon Mountain Complex are dominantly deformed adcumulates, similar to other metacumulate 

sequences. Poikilitic textures are present, especially 

in the lower part of the unit. Diopside oikocrysts and 

the enclosed olivine are deformed. Layering is accentuated by deformation, especially in the gabbros.

The layered peridotites and gabbros unit shows no 

regularity in layer thickness or sequence. Layered sequences persist laterally over distances of only several tens of meters and cannot be traced from one ridge 

to the next. Harzburgite is absent. Clinopyroxenerich peridotite constitutes about 70 percent of these 

rocks and is predominantly wehrlite. The wehrlite, 

with varying proportions of clinopyroxene and olivine, occurs in successions of layers whose thickness-
5. MAFIC AND ULTRAMAFIC ROCKS OF BAKER TERRANE, EASTERN OREGON 227

es vary from centimeters to several meters, and it 

constitutes about 70 percent of the layered peridotites and gabbros unit. Websterite, melagabbro, gabbro, troctolite, anorthosite, dunite, and rare Iherzolite 

together constitute about 30 percent of the unit. Contacts between adjacent peridotite units are gradational over several centimeters. Contacts between 

gabbro and peridotite are generally sharp.

Although lithic types vary, mineral compositions 

are monotonously similar throughout the layered 

peridotites and gabbros unit (table 5.3). In the sequences examined along Celebration Ridge (a ridge 

that trends north from Pine Creek Mountain to Celebration Mine, fig. 5.1), there are only slight variations in mineral compositions, so lithic diversity is 

simply a function of different modal abundances.

Minerals are not compositionally zoned and are of 

nearly uniform composition throughout each layer 

(fig. 5.4). Olivine (Fo70_8o) is less Mg-rich than in the 

harzburgite, but exhibits ubiquitous deformation and 

twinning. Bronzitic orthopyroxene (En77_82) commonly 

shows kink-banding as well as other evidence of deformation. Clinopyroxene is Ca-rich diopside, commonly 

with glide twins. Pearcy (1987, 1991) reports compositional ranges of approximately Ca48Mg45_48Fe4_8 in 

wehrlite and websterite, and Ca47_49Mg44_45Fe6_8 in 

gabbros of the layered peridotites and gabbros unit. 

Analyses of clinopyroxene for this investigation yielded similar compositions (table 5.4). Plagioclase of the 

gabbros, troctolites, and feldspathic wehrlites is extremely calcic (Ang2_97) and typically has bent and (or) 

dislocated twins. In plagioclase-bearing peridotites 

the feldspar is rounded and interstitial, and commonly is rimmed by brown pargasitic amphibole.

Di

En Fs

FIGURE 5.4. Compositions of clinopyroxenes from four-layer sequence in layered peridotites and gabbros unit, which forms 

transition zone from tectonite peridotite to overlying cumulate 

and isotropic gabbro. Sample 579A is olivine melagabbro, 579B 

is websterite, 579C is olivine and plagioclase-bearing websterite, and 579D is olivine-bearing websterite. Di, diopside; En, 

enstatite; Fs, ferrosilite; Hd, hedenbergite. Lines connecting 

data points indicate pyroxenes from same sample; curve across 

top of diagram is Skaergaard pyroxene trend.

GABBRO UNITS

Gabbro constitutes approximately one-third of the 

Canyon Mountain Complex and can be divided into 

four units. Hydrothermal alteration has affected all 

units and to some degree obscures textures, mineral 

compositions, and geochemistry of the original rocks. 

The informal but mineralogically distinct gabbroic 

rock units recognized are Gl, an olivine, two-pyroxene rock that occurs near the base of the harzburgite 

of the Canyon Mountain Complex and as dikes and 

sills within the harzburgite; G2, a two-pyroxene gabbro that overlies the layered peridotites and gabbros 

unit; G3, gabbronorite and norite that overly unit 

G2; and G4, hydrothermally altered gabbro (epidiorite) in which few original minerals remain.

Gabbro at the lowest stratigraphic level, unit Gl, 

has the least evolved mineralogy and geochemistry. 

Unit Gl is exposed only as a 25-m-wide pod along 

Gwyn Creek (a tributary of Pine Creek whose source 

is on the northwest side of Baldy Mountain, fig. 5.1) 

within the lower part of the tectonite peridotite unit. 

Float of layered gabbro and wehrlite occur near, and 

apparently envelop, the isotropic gabbro in Gwyn 

Creek. Clinopyroxenite occurs between the gabbro 

and adjacent harzburgite. Unit Gl is an olivine, twopyroxene gabbro with adcumulate to heteradcumulate textures. It displays subtle layering in its limited 

exposure at Gwyn Creek. Gabbro that is mineralogically and geochemically similar occurs as isotropic 

and layered gabbro dikes within the harzburgite (see 

section "Zone of Infiltration," below). In this gabbro 

unit, olivine (Fo85_89) may be rimmed by clinopyroxene. Orthopyroxene (En8o) is interstitial to anhedral, 

with fine exsolution lamellae. The clinopyroxene is 

calcic augite; it encloses both olivine and orthopyroxene. Plagioclase (An92_85) is subhedral, complexly 

twinned, and unzoned. Pargasitic hornblende replaces and rims mafic phases and some plagioclase.

Gabbro in which clinopyroxene is the dominant 

mafic phase occurs in the central and western parts 

of the Canyon Mountain Complex. This gabbro, unit 

G2, also called the gabbro of Bear Skull Rims, intrudes and overlies layered peridotites (layered peridotites and gabbros unit) near Gand Saddle (fig. 5.1).

Phase layering is common in unit G2 (fig. 5.5). In 

some locations, such as the northeast side of Bear 

Skull Rims, the layering is contorted. The absence of 

mineral deformation suggests that the contortion was 

caused by slumping or folding of layers prior to solidification. Pods and bands of melagabbro and clinopyroxene-rich peridotite are distributed throughout this 

unit and probably represent fragments of the layered 

peridotites and gabbros unit or local accumulations of 

mafic minerals resulting from the crystallization of 

the G2 gabbro unit.
228 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 5.4. Compositions of silicate minerals from the upper part of the layered peridotites andgabbros unit

of the Canyon Mountain Complex

[FeO*=FeO+Fe2O3; , not applicable]

Mineral 

Sample 

Lithic type 

Olivine

579A

Olivine 

websterite

579D

Werhlite

579A

Olivine 

websterite

Orthopyroxene

579B 579C

Olivine Metagabbro gabbro

Clinopyroxene

579D 579A

Werhlite Olivine 

websterite

579B

Olivine 

gabbro

579C

Metagabbro

579D

Werhlite

579A

Olivine 

websterite

Plaeioclase

579B

Olivine 

gabbro

579C

Metagabbro

Major-element oxides (weight percent)

Si02 ~

Ti02 --- 

A12Q3 

FeO* 

MnO -

MgO - 

CaO 

Na2O -

Cr2O3 

NiO - 

Total 

37.92

.00 

.00 

22.70

.35

38.80 

.04

.00

.02

.21

100.03

39.53

.00 

.00 

7.88

.26

1.84 

.04

.01

.00

.16

99.76

54.28

.16 

2.25 

14.32

.35

26.80 

.65

.00

.16

.12

99.08

55.39 54.80

.04 .00 

1.80 2.32 

12.55 13.17

.24 .23

28.56 28.14 

.80 .58

.00 .00

.11 .11

.00 .00

99.50 99.35

55.05 51.54

.04 .56 

2.26 3.81 

11.92 6.60

.24 .24

29.79 15.26 

.63 21.00

.00 .28

.22 .46

.04 .04

100.18 99.77

52.71

.17

2.54 

5.43

.08

16.03 

22.01

.16

.15

.06

99.33

53.31

.15 

2.92 

5.67

.12

15.42 

22.04

.23

.17

.00

100.03

51.98

.32 

2.97 

4.88

.09

16.07 

22.08

.19

.46

.03

99.07

44.80

.00 

36.11 

.30

.00

.02 

18.71

.85

.00

.00

100.79

45.17

.00 

35.52 

.19

.00

.02 

19.21

.66

.00

.00

100.75

44.71

.00 

36.32 

.20

.00

.02 

18.99

.63

.00

.00

100.67

Stoichiometry (number of ions per unit cell)1

Si 

Ti -____ 

Al _ _ 

Fe ________

Mn 

Mg 

Ca -

Na

Cr ~

Total 

1.07

.00

.00

.50

.01

1.51

.00

.00

.00

.01

3.09

1.00

.00

.00

.38

.33

1.57 

.00

.00

.00

.03

2.97

3.92

.01

.19

.87

.02

2.89 

.05

.00

.01

.01

7.97

3.95 3.92

.00 .00

.15 .20

.75 .78

.01 .01

3.04 3.00 

.06 .04

.00 .00

.01 .01

.00 .00

7.97 7.98

3.89 3.81

.00 .03

.19 .33

.70 .41

.01 .01

3.14 1.68 

.05 1.66

.00 .04

.00 .03

.00 .00

8.01 8.00

3.99

.01

.22

.33

.00

.00 

1.74

.02

.01

.00

8.00

3.90

.01

.25

.35

.01

.01 

1.73

.03

.01

.00

7.97

3.85

.02

.26

.30

.01

1.77 

1.75

.03

.03

.00

8.01

8.24

.00

7.80

.05

.00

.00 

3.67

.30

.00

.00

20.06

8.28

.00

7.67

.03

.00

.00 

3.77

.23

.00

.00

19.98

8.19

.00

7.83

.08

.00

.00 

3.70

.22

.00

.00

20.02

Relative proportions (percent)

An

Fo 

En2 

En3 

Wo 

_

75

___

_

80

_

77

___

80 79

_ ___

_ 

82

11 

45

44

___

8 

46

46

___

9 

45

46

___

8 

46

46

92

___

94 94

1Assuming 4 oxygen atoms per unit cell for olivine, 12 for orthopyroxene and clinopyroxene, and 32 for plagioclase. 

2Enstatite component of orthopyroxene. 

3Enstatite component of clinopyroxene.

FIGURE 5.5. Layering in unit G2 (gabbro of Bear Skull Rims). 

Coin is approximately 2 cm in diameter.
5. MAFIC AND ULTRAMAFIC ROCKS OF BAKER TERRANE, EASTERN OREGON 229

Two pyroxenes are commonly present and olivine is 

absent in the gabbro that forms unit G2. Clinopyroxene (calcic augite) is anhedral, and in deformed rocks 

may be pinched between grains of plagioclase. Orthopyroxene (En75_85) is most abundant in layered rocks, 

seldom exceeds 10 modal percent, is subhedral to euhedral, and usually is altered to serpentine+chlorite+amphibole. Plagioclase (An7o_75) is unzoned and nearly 

always partly altered to chlorite+hydrogrossular.

The two-pyroxene gabbro that forms unit G2 

grades upward into the more orthopyroxene-rich unit 

G3, also known informally as the gabbronorite of 

Pine Creek Mountain. Similar rocks extend east-west 

across most of the Canyon Mountain Complex. Most 

gabbros of unit G3 have heteradcumulate to adcumulate textures, and many display igneous lamination. 

The rocks are less deformed than those of the lower 

gabbro units; lattice defects in minerals are rare.

Olivine is absent in unit G3. Orthopyroxene is 

faintly pleochroic hypersthene (En^). Clinopyroxene 

is augite. Pyroxenes are rimmed and replaced by 

late-magmatic to subsolidus pargasitic hornblende. 

Plagioclase (Ang0-65) is unzoned. It is subhedral and 

lath-shaped in laminated gabbronorite, but anhedral 

in the uppermost, unlayered part of unit G3.

Gabbro that stratigraphically overlies unit G3 has 

been subjected to extensive hydrothermal alteration, 

and its original mineralogical and geochemical character is difficult to deduce. These rocks, which form 

unit G4, occupy the uppermost part of Canyon Mountain (fig. 5.1), which constitutes much of the western 

part of the Canyon Mountain Complex. Similar rocks 

are found on Little Canyon Mountain. This altered 

rock, strongly gabbroic in appearance, was termed 

"epidiorite" by Thayer (1972), meaning a hydrothermally altered rock in which chlorite and hornblende 

replace pyroxene, and plagioclase is largely altered to 

chlorite, hydrogrossular, and clays. Alteration of 

these rocks is so complete that igneous minerals cannot be identified, and geochemistry has been substantially changed from that of the original rock.

RELATIONS AMONG GABBRO UNITS

Lithologic, mineralogic, and stratigraphic relations 

in the gabbro units of the Canyon Mountain Complex 

support the origin of these rocks primarily by fractional crystallization of a single magma or as a rapidly intruded sequence of similar magmas. Although 

mineralogic units can be distinguished in the gabbro 

units of the Canyon Mountain Complex, their contacts are gradational and poorly defined; the lithic 

types progress upward from olivine gabbro to gabbro

to gabbronorite, suggesting a fractionation sequence. 

Geochemical data support this interpretation. Although complex processes were undoubtedly involved 

in the generation and emplacement of the magma, 

field evidence supports the model of a single, fractionating magma chamber for the origin of the Canyon Mountain Complex.

LATE INTRUSIVE GABBRO: THE ZONE OF INFILTRATION

A 3.5-km-wide, northeast-trending zone of gabbroic 

veins and dikes, as well as feldspar-bearing harzburgite (fig. 5.6), occurs in the central part of the tectonite 

peridotite unit. It extends from the serpentinized 

base of the Canyon Mountain Complex upward to the 

layered peridotites and gabbros unit, and is 4 km 

wide. Unit Gl, also called the gabbro Gwyn Creek, is 

included in this zone. The distribution and abundance 

of feldspar-bearing harzburgite and gabbroic dikes 

are remarkably symmetrical, changing from an outer 

subzone of feldspathic harzburgite and dikelets less 

than 1 cm wide (zone A), inward to a subzone of gabbroic dikelets 1 to 10 cm wide (zone B), inward to a 

subzone of gabbro dikes 10 cm or wider (zone C), and 

finally including some localities, such as the south 

face of Baldy Mountain, where gabbro dikes and pods 

occupy greater than 10 percent of the outcrop area 

(zone D). The textures, distribution, and compositions 

of the rocks suggest that basaltic magma intruded 

through the harzburgite to the overlying magma 

chamber along this zone. Textural evidence does not 

preclude the migration of gabbroic magma downward 

into the tectonite peridotite unit.

FIGURE 5.6. Harzburgite of the Canyon Mountain Complex 

with plagioclase and gabbro dikes. Plagioclase-bearing harzburgite is found only adjacent to small gabbro veins in the 

central part of the Canyon Mountain Complex. Pocketknife, 

for scale, is approximately 6 cm long.
230 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

The dikes in all zones are principally melagabbro 

and olivine, two-pyroxene gabbro. Their coarse textures suggest formation and crystallization at depth, 

probably early in the history of the Canyon Mountain 

Complex. Where gabbro dikes are small or absent 

near the margins of the zone of infiltration, diffusion 

of basaltic melt through the host tectonite harzburgite has locally transformed the peridotite to plagioclase-bearing harzburgite, plagioclase Iherzolite, 

troctolite, or even melagabbro.

Compositions of minerals in the intrusive gabbro 

are similar throughout the zone of infiltration. Feldspars are calcic plagioclase (Angg.gg), olivine is Mgrich (Fo85). Clinopyroxene is calcic augite, and orthopyroxene is bronzite (Engo). These compositions are 

all slightly more evolved than those of minerals of 

the host harzburgite, and they are similar to those of 

both the layered peridotites and gabbros unit and 

unit Gl, the gabbro of Gwyn Creek.

In the outermost subzone, zone A, dikelets of plagioclase, undeformed clinopyroxene, or olivine, two-pyroxene gabbro, are less than 1 cm wide. Harzburgite 

may contain as much as 50 percent undeformed plagioclase or as little as one plagioclase crystal in 

10 m of depleted harzburgite. This feldspar is undeformed and commonly surrounded by a magmatic reaction rim of pargasitic hornblende (fig. 5.7).

The transition from unaffected harzburgite of the 

tectonite peridotite unit to the feldspathic harzburgite and gabbroic dikelets of zone A is gradational. 

Blobs of feldspar and thin gabbroic dikelets increase 

in abundance inward. Early dikelets are contorted. 

Some are oblique to foliation in the host harzburgite,

FIGURE 5.7. Isolated plagioclase (dark, rounded form in middle 

of picture) in harzburgite is rimmed by pargasite reaction 

rim and is undeformed, indicating that plagioclase and gabbro infiltration postdated deformation of harzburgite. Field of 

view is about 1 mm across.

but many are parallel to host-harzburgite foliation or 

are offset parallel to it. Later veins are straight and 

crosscut early gabbro veins as well as structure in 

the host harzburgite. These relations suggest that 

early gabbro dikelets were emplaced before or during 

plastic deformation of the harzburgite, and later gabbro was emplaced after the last plastic deformation.

Field relations in zone A strongly suggest that gabbroic magma moved from veins and small dikes into 

the host harzburgite, rather than being a result of 

partial melting of a fertile peridotite, that is, a peridotite that will produce basaltic magma by partial 

melting. Gabbroic constituents decrease rapidly with 

distance from a vein, suggesting infiltration of fluid 

from the vein into depleted peridotite. There is no 

plagioclase, and only small (5 percent) amounts of 

clinopyroxene, in the harzburgite of the Canyon 

Mountain Complex outside the zone of infiltration.

A subzone containing broader gabbro dikes, 1 to 10 

cm wide, occurs inward from zone A and is denoted 

as zone B. The gabbro dikes in zone B have been altered to rodingite. Larger dikes have broad serpentinized contacts with the host tectonite peridotite unit. 

Feldspar-bearing harzburgite is less abundant than 

in zone A.

Tectonite peridotite that was intruded by gabbro 

dikes greater than 10 cm wide was designated zone 

C. Zone C is exposed in the north-central part of the 

tectonite peridotite unit. It constitutes about 25 percent of the area mapped in the gabbroic zone. Significantly, unit Gl (gabbro of Gwyn Creek) is enclosed by 

zone C. This gabbro is very similar to the dike gabbros in mineralogy, mineral composition, and wholerock geochemistry.

In zone D, gabbro forms at least 10 percent of the 

visible outcrops. The dikes are as much as 3 m wide. 

They crosscut foliation of the host tectonite peridotite 

unit and also crosscut one another. Some broader dikes, 

especially those exposed on the south face of Baldy 

Mountain, are banded or layered parallel to their borders. The banding may have resulted from flow segregation during magma movement or from pulses of 

magma flowing through an expanding fracture.

ORIGIN OF THE ZONE OF INFILTRATION

The zone of infiltration of the Canyon Mountain 

Complex is interpreted as a zone of depleted uppermantle tectonite peridotite that was infiltrated and 

intruded by the chemically and mineralogically 

evolved gabbroic magmas that fed an overlying 

magma chamber, or were forced downward from that 

chamber. Similar rocks and relations in the Oman
5. MAFIC AND ULTRAMAFIC ROCKS OF BAKER TERRANE, EASTERN OREGON 231

ophiolite, in the Trinity (Calif.) ophiolite, and in ophiolites elsewhere have been interpreted as various 

stages in the partial to complete segregation and expulsion of a melt fraction from a rising mantle diapir 

(Clifford Hopson, oral commun., 1986). However, the 

infiltration/intrusion model is favored for the Canyon 

Mountain Complex for the following reasons:

(1) Feldspathic harzburgite does not occur elsewhere in the Canyon Mountain Complex. Depleted 

harzburgite, with low modal clinopyroxene and no 

other fertile components, occupies 95 percent of the 

tectonite peridotite unit of the Canyon Mountain 

Complex. Lherzolite also has low modal abundances 

of clinopyroxene. Hence, it is unlikely that the tectonite peridotite unit of the Canyon Mountain Complex was the source of the comparatively voluminous 

gabbro in the zone of infiltration.

(2) Plagioclase in harzburgite within the zone of infiltration is undeformed. Hence plagioclase was not 

present during the principal deformation of the harzburgite in the upper mantle.

(3) Plagioclase has magmatic reaction rims of pargasitic amphibole at its contacts with enstatite. 

These rims suggest magmatic or fluid emplacement 

of the plagioclase rather than subsolidus crystallization by phase change from spinel.

(4) Plagioclase and augitic clinopyroxene are abundant adjacent to small veins in zone A and decrease

rapidly with distance from the veins. This strongly 

suggests that the veins are the source of the fertile 

components, rather than the plagioclase and clinopyroxene being contributed to the veins from partial 

melting in fertile peridotite. There is no plagioclase 

in the harzburgite of the Canyon Mountain Complex 

except in the zone of infiltration.

(5) The mineralogical and chemical compositions of 

the larger dikes (zone D) are somewhat evolved, as 

shown by their relatively high Fe and Na contents. 

Hence, it is unlikely that they represent a primary 

melt.

MIGMATITIC AND RECRYSTALLIZED GABBRO: 

ZONE OF PLAGIOGRANITE GENESIS

Elongate areas of unit G3 (the gabbronorite of Pine 

Creek Mountain) show evidence of generating the plagiogranites of the Canyon Mountain Complex. These 

zones, designated as migmatitic zones on figure 5.2, 

are more than 200 m wide and characteristically have 

a core of migmatitic textures (fig. 5.8), an intermediate band of granoblastic or granulitic fabrics, and an 

outer zone where lineated poikiloblastic hornblende is 

well developed in gabbronorite. The distinctive character of the migmatitic zones was first recognized by 

T.P. Thayer (written commun., 1980), who considered

Approximately 1 meter

FIGURE 5.8. Drawing of textures in the zone of infiltration, West Fork Pine Creek. Darkest areas are 

amphibolized diabase, intermediate are recrystallized gabbronorite, and lightest areas are plagiogranite. Contacts between plagiogranite and gabbronorite are gradational. Relict cumulate layering 

(vertical in this view) is shown at left.
232 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

the oddly textured rocks in the West Fork of Pine 

Creek (which flows between Gand Saddle and Bear 

Skull Rims, fig. 5.1) and the West Fork of Indian 

Creek (which flows along the southeast flank of Baldy 

Mountain, fig. 5.1) to be zones of penecontemporaneous diabase and plagiogranite intrusion and brecciation. Thayer indicated that these zones are important 

to the history of the Canyon Mountain Complex.

These zones of migmatitic and recrystallized rocks 

trend east-west and are parallel to diabase dikes and 

plagiogranite bodies. Gabbroic rocks in the narrow, 

migmatitic central zone are intimately mixed with 

plagiogranites (fig. 5.8). Gabbronorite grades into 

mafic diorite, which grades texturally and mineralogically into plagiogranite on scales ranging from meters to millimeters. Fabrics are migmatitic. Layered 

gabbro is plastically deformed into irregular, patchy 

mixtures of fine-grained and medium-grained gneissic rock. Angular to subrounded xenoliths of amphibolized diabase, up to 3 m in length, are contained 

within the migmatitic gabbros (figs. 5.9, 5.10). Diabase dikes, from less than 1 m to about 10 m in 

width, are crosscut and brecciated by plagiogranite. 

Mafic clots and fine-grained, amphibolized schlieren 

of gabbronorite also occur in the inner zones (figs. 

5.9, 5.10).

Although gabbro immediately outside this migmatitic zone appears unaffected, in thin section these

FIGURE 5.9. Gabbronorite schlieren in matrix that is mixture of quartz gabbro and plagiogranite. West Fork Indian Creek. 

Ruler is approximately 10 cm long. Photograph by T.P. Thayer, 1966.
5. MAFIC AND ULTRAMAFIC ROCKS OF BAKER TERRANE, EASTERN OREGON 233

rocks have a gramiloblastic to xenomorphic-granular 

texture (fig. 5.11). Triple-point (120°-angle) junctions 

are well developed. Relicts of euhedral plagioclase 

and subhedral augite are present in a matrix of granuloblastic albite. Clinopyroxene commonly is recrystallized into aggregates of small clinopyroxene 

crystals with triple-point junctions, or has a sieve 

texture. Orthopyroxene appears largely unaffected by 

recrystallization.

Poikiloblastic hornblende up to 2 cm in length appears in an outer zone, where granulitic textures are 

not apparent (fig. 5.12). In some outcrops the hornblende is lineated obliquely to the trend of the zone. 

Plagiogranite intrudes this outer zone, but no migmatitic textures are developed in the host gabbro. 

The hornblende commonly has clinopyroxene cores 

and appears to be a product of high-temperature hydration of clinopyroxene.

ORIGIN AND SIGNIFICANCE OF THE MIGMATITIC ZONE: 

GENERATION OF PLAGIOGRANITE

Zones of recrystallization and migmatitic textures 

similar to that of the Canyon Mountain Complex are 

rare. Annealed fabrics attributed to strain-enhanced 

recrystallization are present in the Gosse Pile stratiform intrusion, central Australia (Moore, 1973).

FIGURE 5.10. Migmatitic fabrics and brecciated diabase, migmatitic zone, West Fork Pine Creek. A, Light-colored plagiogranite is 

segregated into some poorly defined veins. Metal field stereoscope is 18 cm wide. Photograph by T.P. Thayer, 1966. B, See next 

page.
234 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Other occurrences of fine-grained, granular-textured 

rocks have been attributed to supercooling and release of volatile pressure (Fonger-Hyllingen complex, 

Norway; Thy and Esbensen, 1982), and to recrystallization along slip planes (Josephine peridotite, Oreg.; 

Jorgenson, 1979). Generation of plagiogranite by amphibolite anatexis has been recognized in the Fournier ophiolite, Prince Edward Island, New Brunswick 

(Flager and Spray, 1991).

The extensive migmatitic textures and recrystallization in unit G3 (gabbronorite) of the Canyon Mountain Complex represent a significant postmagmatic 

event. These zones of partial melting and recrystallization appear to be related to the production of plagiogranites of the Canyon Mountain Complex by 

partial melting of altered (hydrated, greenschist-facies) gabbro. Gerlach (1980) used trace-element data 

to calculate that the plagiogranites of the Canyon 

Mountain Complex were produced by partial melting 

of altered oceanic (tholeiitic) gabbro. The migmatitic 

zones appear to be the locations where this partial 

melting and plagiogranite generation occurred. They 

are spatially associated with plagiogranites near the 

top of the gabbro section of the Canyon Mountain 

Complex. Although plagiogranites of the migmatitic 

zone cannot be traced directly to the overlying plagiogranite bodies due to lack of appropriate bedrock exposure, the abundance of plagiogranite in the

FIGURE 5.10. Continued. B, Detail of migmatitic zone. Note complex mixture of diabase (very dark), recrystallized gabbro (dark), 

plagiogranite (coarse-grained and light), and mixed rock. Postmagmatic brittle deformation has offset some parts of this outcrop. 

Ruler is approximately 10 cm long. Photograph by T.P. Thayer, 1966.
5. MAFIC AND ULTRAMAFIC ROCKS OF BAKER TERRANE, EASTERN OREGON 235

migmatitic zone and the presence of plagiogranite 

dikes and veins leading upward indicate that this is 

a viable model.

The role of diabase dikes in this scenario is unclear. 

However, crosscutting field relations, the incorporation of diabase xenoliths within the plagiogranite, 

and the brecciation of diabase by plagiogranite dikes 

and stringers suggest that diabase dikes intruded the 

zone prior to plagiogranite generation.

Previously, the triple-point junctions and other 

granuloblastic to granulitic textures of the upper 

gabbros as well as their anomalously dry, or unaltered and fresh, character were attributed to adcumulate crystallization (Himmelberg and Loney, 

1980). There is a marked resemblance between adcumulate texture and granuloblastic texture (Vernon, 

1970). Both are characterized by straight grain 

boundaries and triple-point junctions. However, the 

reorganization of clinopyroxene into small, clustered 

crystals with triple-point boundaries and the absence 

of any intergranular phases throughout the recrystallized zone indicate that the texture of the Canyon 

Mountain Complex is due to recrystallization.

The presence of similar granuloblastic textures in 

the Fonger-Hyllingen complex, where they have been 

attributed to release of volatile pressure, and the association of similar textures in the Josephine peridotite with shear along slip planes suggest that 

shearing or some other tectonic event may have 

played a role in the creation of these zones during 

the late-magmatic or post-magmatic history of the 

Canyon Mountain Complex. Shear is recognized as 

an essential component of plagiogranite generation in 

the Fournier ophiolite.

FIGURE 5.11. Photomicrograph of granulite textures, migmatitic 

zone. Orthopyroxene, clinopyroxene, and albite in recrystallized gabbro, West Fork Pine Creek. Field of view, 2.5 mm.

The plagiogranites of the migmatitic zone of the 

Canyon Mountain Complex were produced by partial 

melting of altered Canyon Mountain gabbros; they 

are not the product of fractional crystallization nor 

the product of intrusion of a separate magma.

All the plagiogranite of the Canyon Mountain Complex is likely to have originated from this same 

mechanism and source. Field and geochemical evidence suggest that plagiogranites were generated 

during an event that released volatiles (^O) from 

altered gabbros and concentrated these volatiles in a 

central area under shear stress. Altered gabbros in 

the middle of the zone partly melted, producing plagiogranite magma, in accordance with Gerlach's 

(1980) geochemical modeling for plagiogranite genesis. If the partial melting occurred close to the surface, then small eruptions of keratophyres probably 

followed.

The production of the two other textural and mineralogical zones (granuloblastic fabrics and hornblende-gabbro) are also linked to this process. Altered 

gabbros outside this zone lost water and recrystallized with dry, granuloblastic textures. Volatiles (water) not directed inward or captured by partial 

melting would continue outward into gabbros cooled 

further by volatile release and expansion, producing 

amphiboles and amphibole reaction rims on clinopyroxenes. The alignment of hornblende in the zone 

strongly suggests crystallization under stress.

The operation of these processes is evident in the 

field and in thin section. Detailed work is needed to 

quantify the stress, water vapor pressures, and other 

parameters under which these processes occurred, 

and to elucidate the relation of diabase dikes. This

FIGURE 5.12. Poikiloblastic hornblende in altered gabbro. Rim 

of migmatitic zone. West Fork Pine Creek. Coin is approximately 2 cm in diameter.
236 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 5.5. Major- and trace-element analyses of rocks of the Canyon Mountain Complex 

[FeOJ^=FeO+Fe2% , not determined]

Lithic type Dunite Harzburgite

Sample 1 2

Lherzolite,

Lherzolite zone of

infiltration

3 4

Olivine

websterite,

transition

zone

5

Gabbro,

upper

transition

zone,

Celebration

Ridge

6

Gabbro,

lower

transition

zone,

Baldy

Mountain

7

Unit Gl

(gabbro

of Gwyn

Creek)

8

Unit G2

(gabbro

of Bear

Skull

Rims)

9

UnitGS

(gabbronorite of

Pine Creek

Mountain)

10

Gabbro dike,

zone of

infiltration,

Baldy

Mountain

11

Major-element oxides (weight percent)

Si02 -

A12Q3 -

TiO2 -

FeO* - 

MnO - 

MgO

CaO 

Na2O ----

K20 

P205 ---

Total36.1

.1

.0

9.1

.1

35.2

.2

.1

.0

.0

80.9

44.3

1.8

.0

8.3

.1

32.8

1.7

.6

.0

.0

89.6

42.5

1.1

.0

8.6

.1

35.8

1.5

.3

.1

.0

90.0

42.7

1.5

.0

8.6

.1

33.8

1.7

.1

.1

.0

88.6

48.3

4.6

.4

11.1

.2

17.5

13.4

1.4

.1

.0

97.0

49.8

24.9

.1

3.5

.1

5.1

13.6

2.7

.1

.0

99.9

50.3

17.9

.2

6.6

.1

10.7

12.2

2.5

.1

.0

100.6

45.7

11.6

.2

9.9

.2

17.3

10.3

1.2

.1

.0

96.5

48.6

19.6

.2

5.0

.1

8.8

14.4

2.3

.1

.0

99.1

50.6

16.6

.2

8.5

.1

8.9

12.8

1.7

.1

.0

99.5

_

__

...

 .

_.

Trace elements (parts per million)

Sr 

Co - 

Sc 

Ni 

cr _________

La ____ __

Ce - 

Nd -- 

gm ________

Eu - 

Yb _________

Lu - 

_

43

11

2800

3700

.06

___

.03

44

40

11

2800

3800

.01

.08

.09

.03

450

44

6

2750

5800

.06

__

.35

.19

.41

.12

84

26

36

50

3500

.12

.41

.53

.13

.18

.20

.06

190

42

54

100

8000

.20

.59

.54

.28

.29

.40

.09

366

30

24

190

2700

.08

.39

.67

.32

.58

1.30

.06

_

90

24

300

4850

.08

.34

.53

.34

.14

.36

.08

_

42

40

200

2700

.06

.28

.81

.39

.29

.59

.17

88

31

44

20

1050

.20

.71

1.35

.78

.26

1.09

.28

329

25

40

370

1300

.02

.13

.42

.25

.59

.40

.09

information could prove vital for understanding the 

late history of the Canyon Mountain Complex and 

the Baker terrane.

The generation of plagiogranite from shear stress 

applied to the Canyon Mountain Complex fits with 

several dates in Baker terrane history. The postmagmatic, metasomatic generation of plagiogranite in the 

gabbros of the Canyon Mountain Complex would explain the 10-m.y. hiatus between the age of the gabbros (278 Ma) and the age of the plagiogranites (268 

Ma)(Gerlach and others, 1988). Metagabbro amphibolites dated at 269 Ma, with a range of dynamothermal metamorphic ages between 255 and 269 Ma, are 

reported in the serpentinite-matrix melange in the 

Baker terrane just west of the Canyon Mountain 

Complex (Carpenter and Walker, 1992). These rocks 

are related to a "dynamothermal aureole" created by 

"emplacement of a young arc over mafic crust" (Carpenter and Walker, 1992).

The nature of the event that produced this significant increase in shear stress and "emplacement of a

young arc over oceanic crust" is enigmatic. Volcanic 

activity in the Wallowa terrane peaked at about 270 

Ma. Activity in the Olds Ferry terrane began at 

about 240 Ma. One possibility is that the high shear 

stresses and generation of amphibolites may be related to a change in arc polarity and (or) the inception 

of east-facing subduction at about 270 Ma that would 

eventually generate the Olds Ferry terrane (Vallier, 

chap. 3, this volume).

GEOCHEMISTRY

Geochemical studies of gabbro and peridotite of the 

Canyon Mountain Complex (table 5.5} suggest that 

the Canyon Mountain Complex is related to a calcalkaline trend. The AFM plot is not diagnostic for 

gabbro and peridotite, as these units plot at the base 

of both calc-alkaline and tholeiitic trends. However, 

cumulate gabbros, as well as unit G3 (gabbronorite), 

are more calcic than in most ophiolites and plot in a
5. MAFIC AND ULTRAMAFIC ROCKS OF BAKER TERRANE, EASTERN OREGON 237

field of alkalic to calc-alkaline cumulates rather than 

in the field for oceanic ophiolites on an Al2O3/MgO/ 

CaO diagram. These discriminants must be used with 

caution because all the units of the Canyon Mountain 

Complex are altered to some degree, and discriminants may not be as accurate for cumulate sections as 

for liquid compositions. The overall geochemistry of 

the complex is arc tholeiite to calc-alkaline.

The rare-earth-element (REE) compositions of 

rocks of the Canyon Mountain Complex are similar 

to those of analogous units of other arc-related ophiolites, including the Troodos (Cyprus) and Point Sal 

(Calif.) ophiolites. Light rare-earth elements 

(LREE's) increase and the positive europium anomaly is more pronounced upward through the cumulate 

units. Fractionation modeling of REE's from the 

lower gabbro units (Gl, G2, G3) suggests that fractionation, with only very limited replenishment by 

more primitive magma, can account for the composition of the lower gabbros. On the basis of this model, 

the Canyon Mountain Complex appears to have been 

a nearly closed magmatic system with limited influx 

of new magma, rather than the open magmatic system characteristic of midocean ridges. However, the 

extreme alteration of the uppermost, isotropic gabbro 

units, in concert with the cumulate nature of (lower) 

unit G2 (gabbro of Bear Skull Rims), precludes reliable fractionation modeling of the gabbros of the 

Canyon Mountain Complex. The general requisites 

for a fractionating magma chamber are present.

Sr and Nd isotopic work by Gerlach and others 

(1988) indicates that the lower parts of the Canyon 

Mountain Complex were derived from an isotopically 

homogeneous, depleted source. Isotropic and layered 

gabbros of the Canyon Mountain Complex have initial 87Sr/86Sr ratios of 0.7024 to 0.7031. Keratophyres 

and rocks from the alternating diabase-keratophyre 

sill sequence that overlies the gabbros of the Canyon 

Mountain Complex have 87Sr/86Sr ratios of 0.7031 to 

0.7037, within the average range of values for arcderived basalts and andesites.

ORIGIN AND TECTONIC SETTING

The mineralogy and geochemistry of the Canyon 

Mountain Complex are more characteristic of islandarc magmas than of oceanic tholeiites or oceanic alkalic 

rocks. The gabbros of the Canyon Mountain Complex 

are more enriched in CaO and AL^Os than are oceanic 

tholeiites; they are more like calc-alkaline to mildly 

alkalic continental intrusive bodies such as the Kiglapait (Morse, 1969) in this respect. The pyroxenes of the 

layered peridotites and gabbros unit and the gabbro

units Canyon Mountain Complex are also enriched in 

CaO and A12O3 , and are quite low in TiO2 and alkalis 

compared to pyroxenes from equivalent units in midocean-ridge ophiolites such as the Oman ophiolite. 

These traits are characteristic of arc-derived magmas. 

Other discriminants, including TiO2/Fe*O:MgO and 

MnO/TiO2/P2O5 , further suggest calc-alkaline to arctholeiite affinity for gabbro of the Canyon Mountain 

Complex. The initial 87Sr/86Sr ratios of cumulate gabbro from modern island arcs is not known. However, 

the ratios of 0.7024 to 0.7031 are within the range of 

ratios reported from arc basalts, as well as from normal midocean-ridge basalts (NMORB), and an arc origin cannot be ruled out on the basis of initial 87Sr/86Sr 

ratios. Thus, the geochemical signatures of the gabbros 

of the Canyon Mountain Complex support the complex's affiliation with an island arc.

The mineralogy and chemistry of the Canyon 

Mountain Complex also suggest that it is a remnant 

magmatic system produced immediately above uppermantle peridotite in much the same setting where we 

might expect to find midocean-ridge magma chambers. This setting is also appropriate for early islandarc magmatism. The scale of the Canyon Mountain 

Complex is smaller than that of ophiolites related to 

midocean-ridge basalt (MORE). Its transition zone of 

layered peridotites and gabbros averages less than 1 

km thick and is less than 100 m thick in the western 

part of the complex. Its gabbros and small transition 

zone suggest fractionation with a fairly limited rate 

of magma replenishment, which is consistent with 

early island-arc magmatism.

The dominantly arc-tholeiite geochemical signature 

of the gabbros of the Canyon Mountain Complex, its 

setting as a relatively small and quiet magma chamber perched directly above mantle harzburgite, and 

the presence of the Canyon Mountain Complex as the 

only coherent ophiolite in a sea of serpentinite-matrix melange suggest that this ophiolite may have 

originated as part of the fore-arc system or prior to 

the inception of significant arc-related magmatism. 

Limited arc-tholeiite magmatism occurs in the Marianas fore arc (in the western Pacific Ocean) more 

than 100 km from the active arc (Marlow and others, 

1992). However, fore-arc magmatism, in general, is 

rare. A source directly beneath Marianas fore-arc basalts would require partial melting at depths of 20 to 

30 km; partial melting at such shallow depths would 

not be possible without extraordinary amounts of hydration (Marlow and others, 1992). The source of 

Marianas fore-arc magmas is therefore assumed to 

be larger chambers beneath the active arc.

Pearcy (1991) has suggested that the Canyon 

Mountain Complex is a nascent island arc related to
238 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Olds Ferry terrane volcanism. The generally arc-tholeiite geochemistry supports this hypothesis. However, Olds Ferry volcanism is Triassic in age, 235 to 

216 Ma, and possibly as young as 181 Ma (Vallier, 

chap. 3, this volume), whereas the plagiogranites and 

gabbros of the Canyon Mountain Complex are Early 

Permian in age, 278 to 268 Ma (Gerlach and others, 

1988).

Age and chemistry suggest that the Canyon Mountain Complex may be closely related to the Wallowa 

terrane. It may represent the early roots of arc volcanism, or it may represent fore-arc volcanism on a 

larger scale than we currently know and understand.

Two periods of igneous activity may be represented 

in the Canyon Mountain Complex. The first period of 

activity produced the bulk of the transition zone and 

gabbroic rocks of the Canyon Mountain Complex, 

ending about 278 Ma. The second period of activity 

involved relatively minor volumes of intrusion (zone 

of infiltration) and partial melting (migmatitic 

zones). This episode, at about 270 to 265 Ma, generated the plagiogranites of the Canyon Mountain 

Complex (268 Ma; Gerlach and others, 1988). Walker 

(chap. 6, this volume) has found evidence for two 

ages of igneous activity in discordant U-Pb ages. The 

second episode of activity may have been related 

more to tectonics than to igneous processes.

OTHER MAFIC AND ULTRAMAFIC COMPONENTS 

OF THE BAKER TERRANE

The diverse volcanic and plutonic components of the 

Baker terrane include a variety of volcanic greenstones and metagabbros (Mullen, 1985). Many are 

thoroughly altered to greenschist-facies assemblages 

and pervasively sheared. However, some contain relict 

clinopyroxenes and (or) well-preserved textures. Such 

rocks provide valuable mineralogical and geochemical 

insights into their petrotectonic affinities, and reasonable clues to the nature of the Baker terrane.

On the basis of rock textures, trace-element data, 

and the compositions of relict clinopyroxenes, volcanic 

greenstones from several locations have island-arc affinities (fig. 5.13). Metabasalts with relict intergranular textures occur within the Burnt River Schist near 

Durkee, Oreg. (fig. 5.1). They contain low-Ti augite 

(TiO2<0.5 percent) (samples BR-52, BR-61, table 5.6). 

The pyroxene compositions are characteristic of lowK, early arc tholeiites (Nesbitt and Pearce, 1977). 

Whole-rock compositions (table 5.7) fall within arc 

fields on discriminant diagrams (fig. 5.14). Basaltic 

greenstones that occur in melange near the town of 

Greenhorn also have intergranular texture and isTiO2

Arc + Alkalic' within plate 

plate tholeiite

MnO Na2O

£8iszQOt= °0 

A1 

OLU

w§

^ 46

B

Ocean-floor basalt

-^ Within-plate 

 vy/alkalic basalt

N \. 

I '\

Island-arc.basalt

1

TiO2 CONTENT, 

IN WEIGHT PERCENT

Hf/3

Th Ta

FIGURE 5.13. Discriminant diagrams for greenstones of 

Baker terrane. Filled circles, pillowed greenstone 

samples V-27 (Greenhorn Mountains) and BRS 

(Sturgill); open circles, all other greenstones of Baker 

terrane. Compositions of relict clinopyroxenes are 

shown on A, MnO/Ti02/Na20 diagram, and B, graph 

of Si02 versus Ti02 content. Fields of basalt compositions modified from Pearce and others, 1977. C, 

Whole-rock trace-element analyses plotted on Th/Hf7 

Ta diagram. Fields of basalt compositions modified 

from Wood and others (1979). NMORB, normal midocean-ridge basalt; EMORB, enriched midocean-ridge 

basalt.
5. MAFIC AND ULTRAMAFIC ROCKS OF BAKER TERRANE, EASTERN OREGON

TABLE 5.6. Compositions of clinopyroxene and olivine from greenstones of the Baker terrane, northeastern Oregon

[FeO*=FeO+Fe2 03; n.a., not applicable; , not determined]

239

Sample -

Mineral 

grain 1 2 3 4

Clinopyroxenes from MV-48A

5678 9 10 11 12 13 14

Major-element oxides (weight percent)

Si02 ----- 

Ti02 -

A1203 - 

FeO* -----

MnO -

TVfcrO .......

CaO -

Na2 0 

Cr203 ----- 

NiO 

Total -

51.33 

.94 

3.61 

6.41 

.14 

17.60 

19.18 

.25 

.10 

.04

99.11

49.02 

1.21 

4.50 

7.93 

.16 

16.56 

18.09 

.30 

.00 

.06

97.85

53.58 

.46 

2.01 

7.82 

.23 

19.52 

16.86 

.22 

.43 

.00

100.74

51.17 

.84 

4.14 

6.40 

.15 

16.51 

19.41 

.28 

.39 

.00

99.28

52.08 52.33 52.27 53.78 

.86 .47 .81 .73 

3.94 3.15 4.34 3.97 

7.20 5.51 6.82 7.79 

.21 .12 .12 .15 

17.06 17.25 17.31 17.70 

19.15 19.95 19.16 17.60 

.27 .28 .25 .27 

.15 .55 .05 .15 

.00 .00 .13 .03

100.91 99.62 101.24 102.70

51.19 

.82 

4.05 

6.46 

.14 

17.10 

19.37 

.30 

.22 

.08

99.74

51.26 

.80 

3.84 

5.86 

.17 

17.00 

19.86 

.30 

.66 

.14

99.90

50.88 

1.13 

4.26 

7.99 

.22 

16.32 

18.31 

.27 

.00 

.00

99.37

49.91 

1.18 

3.77 

9.95 

.23 

15.98 

17.34 

.34 

.00 

.06

98.75

50.25 

.93 

3.78 

7.03 

.15 

16.91 

19.04 

.32 

.10 

.09

98.595

48.59 

1.46 

3.62 

12.08 

.19 

12.58 

17.84 

.38 

.00 

.09

96.85

Stoichiometry (number of ions per unit cell)1

Ti ...........

Al

Fe ... ...._

Mn-" -- 

Mg~~ - 

Ca - - 

Na ____

Cr 

Nj _____

Total

1.896 

.026 

.157 

.198 

.004 

.942 

.759 

.018 

.002 

.001

4.006

1.85 

.03 

.19 

.25 

.005 

.931 

.731 

.021 

.000 

.001

4.026

1.941 

.012 

.085 

.236 

.007 

1.054 

.654 

.015 

.001 

.000

4.010

1.888 

.023 

.179 

.197 

.004 

.908 

.767 

.020 

.011 

.000

4.002

1.893 1.918 1.889 1.920 

.023 .012 .002 .019 

.168 .136 .184 .166 

.218 .168 .206 .232 

.006 .003 .008 .004 

.924 .942 .933 .942 

.746 .783 .741 .673 

.019 .019 .017 .018 

.004 .015 .001 .004 

.000 .000 .003 .000

4.006 4.002 4.004 3.984

1.882 

.022 

.175 

.198 

.004 

.397 

.763 

.021 

.006 

.002

4.015

1.882 

.02 

.17 

.18 

.01 

.93 

.78 

.02 

.02 

.01

4.013

1.883 

.031 

.185 

.247 

.006 

.901 

.726 

.018 

.000 

.000

4.001

1.876 

.033 

.167 

.312 

.007 

.895 

.698 

.024 

.000 

.001

4.018

1.876 

.024 

.166 

.219 

.004 

.94 

.761 

.023 

.003 

.002

4.025

1.88 

.042 

.165 

.392 

.006 

.729 

.743 

.029 

.000 

.002

4.000

Relative proportions (percent)

Fs 

En -

Wo

Sample 

Mineral 

grain 

10 

50 

40

1

13 

49 

38

2

12

54 

34

3

11

49 

40

4

12 9 11 13 

49 50 50 51 

40 41 39 36

Clinopyroxenes from BR-52

5 6A 6B 6C

10 

49 

40

7

10 

49 

41

8A

13 

48 

39

8B

16

47 

37

Core

11

49 

40

9

Intermediate

21 

39 

40

Rim

Major-element oxides (weight percent)

SiQ2 -

Ti02 - 

A1203 ~- 

FeO* ----- 

MnO - 

MgO -----

poO

Na2 0 

Cr203 ----- 

NiO 

Total

53.99 

.24 

1.68 

6.07 

.15 

18.91 

18.09 

.14 

.04 

.06

99.38

53.46 

.26 

1.67 

5.34 

.12 

19.03 

18.73 

.14 

.24 

.16

99.00

52.23 

.31 

2.19 

5.69 

.13 

18.14 

19.83 

.19 

.09 

.03

98.82

53.28 

.26 

2.34 

6.08 

.20 

18.29 

17.95 

.17 

.10 

.11

98.69

53.04 51.65 52.62 51.19 

.29 .41 .42 .56 

2.73 3.89 3.47 3.33 

5.99 6.70 6.06 6.20 

.13 .10 .11 .12 

18.25 17.07 17.42 16.59 

18.46 19.21 19.54 19.54 

.19 .32 .18 .24 

.16 .03 .99 .00 

.08 .00 .27 .04

99.35 99.40 99.96 97.79

53.43 

.19 

1.90 

6.01 

.13 

19.12 

18.49 

.17 

.06 

.09

99.61

52.47 

.35 

2.93 

5.04 

.15 

18.12 

18.96 

.19 

.32 

.04

98.62

52.09 

.33 

3.13 

4.42 

.13 

17.46 

20.50 

.22 

.57 

.05

99.15

53.70 

.23 

1.87 

5.74 

.13 

19.17 

17.62 

.17 

.12 

.07

98.82

52.95 

.33 

3.39 

6.07 

.12 

17.03 

19.30 

.17 

.19 

.05

99.62

52.46 

.31 

2.91 

6.17 

.16 

18.55 

18.46 

.23 

.16 

.12

99.52

Stoichiometry (number of ions per unit cell)1

Si --- 

Ti - 

Al~

Fe 

Mn - 

Mg~~~ 

Ca - 

Na -----

Or

Ni -

Total

1.97 

.006 

.072 

.185 

.004 

1.03 

.707 

.009 

.001 

.001

3.990

1.959 

.007 

.071 

.163 

.003 

1.040 

.735 

.010 

.006 

.000

3.944

1.931 

.008 

.095 

.175 

.003 

.999 

.785 

.014 

.002 

.001

4.018

1.959 

.007 

.101 

.186 

.006 

1.002 

.707 

.011 

.003 

.000

3.987

1.941 1.902 1.920 1.915 

.007 .011 .011 .015 

.117 .169 .149 .146 

.183 .206 .185 .194 

.004 .003 .003 .003 

.996 .937 .947 .925 

.724 .758 .764 .783 

.013 .023 .013 .017 

.004 .000 .002 .000 

.002 .000 .000 .001

3.996 4.013 3.998 4.004

1.95 

.005 

.081 

.183 

.004 

1.040 

.723 

.012 

.001 

.002

4.007

1.931 

.009 

.127 

.155 

.004 

.994 

.747 

.013 

.009 

.001

3.997

1.914 

.009 

.135 

.145 

.003 

.957 

.807 

.015 

.016 

.001

4.008

1.967 

.006 

.08- 

.175 

.003 

1.046 

.691 

.011 

.003 

.002

3.990

1.936 

.009 

.146 

.185 

.003 

.928 

.756 

.012 

.005 

.001

3.985

1.922 

.003 

.125 

.189 

.004 

1.013 

.724 

.016 

.004 

.003

4.012

Relative proportions (percent)

Fs - 

En 

Wo -

10 

54 

37

8 

54 

38

9

51 

40

10 

53 

37

10 11 10 10 

52 49 50 49 

38 40 40 41

9 

53 

37

8 

52 

39

8 

50 

42

9 

55 

36

10 

50 

40

10 

53 

38

1Assuming 6 oxygen atoms per unit cell.

Table continued on next page.
240 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 5.6. Compositions of clinopyroxene and olivine from greenstones of the Baker terrane, northeastern Oregon Continued

Sample -

Mineral 

grain 

Clinopyroxenes from BR-61

1 2 3 4 567

Groundmass

V252-1

Clinopyroxenes from V-252

Phenocryst

Rim In*er; Core 

mediate

Intermediate Rim

Major-element oxides (weight percent)

SiO2 -

TiO2 - 

A1203 ~- 

FeO ----- 

MnO - 

MgO - 

CaO ----- 

NaO 

Cr203 ----- 

NiO 

Total -

52.84 

.31 

2.75 

4.12 

.11 

17.84 

20.54 

.25 

.56 

.00

99.31

53.89 

.23 

3.03 

4.24 

.12 

18.62 

20.37 

.21 

.47 

.03

101.21

52.50 52.71 

.31 .27 

2.90 2.62 

4.24 4.57 

.11 .14 

18.01 18.17 

20.55 19.95 

.19 .19 

.47 .32 

.38 .13

99.32 99.05

52.37 51.58 51.64 

.31 .26 .26 

2.64 3.00 2.07 

4.81 3.95 4.29 

.10 .10 .01 

18.69 17.35 17.55 

20.22 20.08 20.53 

.20 .22 .49 

.81 .85 .27 

.13 .04 .01

99.77 97.45 97.20

48.8 

.9 

2.8 

18.4 

.0 

.5 

9.0 

15.8 

1.6 

.0

97.9

50.7 

.9 

3.9 

12.7 

.0 

.3 

14.1 

17.6 

.3 

.1

97.5

50.0 

.6 

2.9 

12.9 

.0 

.5 

14.4 

16.5 

.2 

.0

97.9

50.6 

.9 

4.5 

13.0 

.0 

.4 

14.2 

17.6 

.2 

.0

101.5

50.1 

.7 

4.2 

13.3 

.0 

.4 

13.8 

17.5 

.3 

.0

100.2

50.2 

.9 

3.9 

12.7 

.0 

10.3 

14.1 

17.6 

.2 

.2

100.0

Stoichiometry (number of ions per unit cell)1

Si - 

Ti - 

AlFe -----

Mn

Mg 

Q a _______

Na -

Ni ____

Total

1.932 

.008 

.118 

.125 

.003 

.972 

.804 

.017 

.016 

.000

4.000

1.930 

.006 

.127 

.126 

.003 

.994 

.781 

.014 

.013 

.000

4.000

1.922 1.934 

.008 .007 

.125 .113 

.129 .140 

.003 .004 

.983 .993 

.806 .784 

.013 .013 

.013 .009 

.001 .003

4.007 4.004

1.914 1.923 1.936 

.008 .007 .007 

.113 .132 .091 

.146 .123 .134 

.002 .003 .003 

1.018 .964 .98 

.791 .802 .824 

.014 .015 .035 

.008 .025 .008 

.003 .001 .000

4.023 3.998 4.024

...

1.95 

.03 

.12

.40 

.02 

.73 

.71 

.02 

.00 

.00

3.98

1.92 

.02 

.13 

.42 

.01 

.83 

.67 

.02 

.00 

.00

4.02

1.88 

.02 

.20 

.40 

.01 

.79 

.70 

.01 

.00 

.00

4.01

1.88 

.02 

.18 

.42 

.01 

.76 

.71 

.02 

.00 

.00

4.00

1.88 

.03 

.17 

.40 

.01 

.77 

.71 

.04 

.00 

.00

4.01

Relative proportions (percent)

ps ____ 

En ____

Wo

Sample 

Mineral 

grain 

7 

51 

42

7 

52 

41

7 7 

51 52 

42 41

777 

52 51 51 

40 42 42

Clinopyroxenes from V-27A

1 2 345 123

22 

39 

39

22 

43 

35

21

42 

37

22 

41

37

21

41 

37

Clinopyroxenes from BUS

4 5 6 7 8 9 10 11

Major-element oxides (weight percent)

Si02 ----- 

TiO2 - 

A1203 ~- 

FeO* 

Cr203 ----- 

MnO - 

MgO 

CaO --- 

Na2O 

K2O

Total

45.3 

3.5 

5.4 

13.2 

.6 

.3 

9.9 

21.0 

.6 

.0

99.4

44.5 

4.6 

6.9 

13.6 

.2 

.3 

9.2 

21.2 

.7 

.0

46.1 44.6 46.6 

3.4 4.1 3.3 

6.2 5.9 6.3 

12.3 13.5 11.4 

.1 .0 .0 

.2 .3 .2 

10.0 9.5 8.8 

21.6 21.4 20.1 

.6 .7 .8 

.0 .0 .0

101.1 100.5 100.0 97.7

46.2 46.9 46.1 

3.7 2.8 3.6 

5.4 5.0 5.9 

12.5 12.1 11.7 

.0 .0 .1 

332 

10.9 11.3 11.5 

20.7 20.7 21.1 

.5 .5 .5 

.0 .0 .2

100.3 99.8 100.6

43.9 43.7 

4.5 4.4 

6.2 5.5 

11.6 12.3 

.0 .0 

2 2 

10.'8 10.4 

21.0 21.6 

.5 .5 

.01 .0

98.8 98.8

44.6 

4.4 

6.1 

11.6 

.0 

.3 

10.2 

21.6 

.5 

.0

99.5

44.8 

4.1 

5.9 

11.9 

.1 

.2 

10.3 

21.1 

.5 

.0

99.1

45.0 47.8 

3.9 2.4 

5.8 4.3 

12.3 12.9 

.0 .0 

.3 .3 

10.7 10.8 

21.4 20.3 

.5 .4 

.0 .0

100.1 99.4

46.7 

3.5 

5.1 

12.2 

.0 

.3 

10.9 

20.7 

.4 

.0

99.7

48.3 

2.4 

4.1 

12.5 

.0 

.3 

11.3 

20.0 

.4 

.0

99.4

Stoichiometry (number of ions per unit cell)1

gj __ ...

Ti - 

Al .-. 

Fe ___ 

Mn 

Mg 

Ca - -

Na - 

K _______ __

Total

1.75 

.10 

.25 

.43 

.01 

.58 

.87 

.05 

.02

4.06

1.70 

.13 

.31 

.44 

.01 

.52 

.87 

.05 

.00

4.03

1 T\ 1 79

.10 .12 

.28 .27 

.39 .57 

.01 .02 

.57 .87 

.88 .56 

.05 .68 

.00 .00

4.03 4.81

1.76 1.79 1.75 

.11 .08 .10 

.24 .23 .26 

.40 .38 .37 

.01 .01 .01 

.62 .65 .65 

.85 .85 .86 

.04 .4 .04 

.00 .01 .00

4.03 4.40 4.04

1.70 1.71 

.13 .13 

.28 .26 

.38 .40 

.01 .01 

.63 .61 

.87 .90 

.04 .04 

.00 .00

4.04 4.06

1.72 

.13 

.27 

.38 

.01 

.59 

.89 

.04 

.00

4.03

1.73 

.12 

.27 

.39 

.01 

.60 

.87 

.03 

.00

4.02

1.73 

.11 

.26 

.40 

.01 

.66 

.84 

.03 

.00

4.04

1.84 

.07 

.20 

.42 

.01 

.62 

.84 

.03 

.00

4.03

1.79 

.10 

.23 

.39 

.01 

.62 

.85 

.03 

.01

4.03

1.84 

.07 

.18 

.40 

.10 

.65 

.82 

.31 

.02

4.39

Relative proportions (percent)

Fs ____

En 

Wo -

23 

31 

46

24 

28 

48

21 29 

31 44 

48 28

21 20 20 

33 35 35 

45 45 46

20 21 

34 32 

46 47

20 

32 

48

21 

32

47

21

35 

44

22 

33

45

21 

33 

46

21 

35

45

1Assuming 6 oxygen atoms per unit cell.

Table continued on next page.
5. MAFIC AND ULTRAMAFIC ROCKS OF BAKER TERRANE, EASTERN OREGON 241

TABLE 5.6. Compositions of dinopyroxene and olivine from greenstones of the Baker terrane, northeastern Oregon Continued

Sample -

Mineral

grain 

SiO2 - 

Ti02 ----- 

A1203 ~- 

FeO* - 

Cr203 ----- 

MnO - 

MgO

CaO ~ 

Na2 0 --- 

K20--~

Total -

Clinopyroxenes from M-42 (cumulate)

Rim

50.6 

.7 

5.2 

9.1 

.0 

.3 

13.4 

21.0 

.4 

.0

100.7

Intermediate

49.6 

.6 

4.9 

7.5 

.0 

.3 

13.4 

22.0 

.5 

.0

98.9

1

Inter- c 

mediate Oore

48.9 

.3 

5.2 

8.3 

.0 

.3 

12.6 

22.1 

.4 

.0

98.2

47.3 

.6 

7.8 

12.5 

.0 

.3 

13.7 

16.5 

.4 

.0

99.1

Intermediate

47.9 

.7 

6.3 

8.6 

.3 

.3 

13.5 

22.3 

.4 

.0

100.0

2

mediate Rim Rim Core

51.8 49.3 48.7 50.0 

.3 .6 .5 .7 

3.9 6.2 4.6 5.2 

8.0 10.6 8.1 8.5 

.0 .0 .0 .0 

.3 .3 .3 .3 

13.5 13.1 13.2 12.9 

21.5 19.2 22.0 22.5 

.3 .4 .4 .4 

.0 .0 .0 .0

Rim

51.5 

.5 

4.0 

8.7 

.0 

.3 

13.6 

22.0 

.5 

.0

99.6 99.6 97.9 100.5 101.1

3

49.5 

.8 

1.8 

12.0 

.0 

.4 

12.9 

22.3 

.5 

.0

100.2

4 5

49.9 49.7 

.9 .8 

5.4 5.8 

8.4 8.3 

.0 .0 

.3 .3 

12.8 12.9 

21.6 22.3 

.5 .4 

.0 .0

99.9 100.6

6

48.8 

.8 

6.9 

9.4 

.0 

.3 

12.7 

20.5 

.5 

.0

100.1

Stoichiometry (number of ions per unit cell)1

Si -

Ti -----

Al - -

pe ____

Mn»-» - 

Mg 

Ca

Na

K - 

Total

1.87 

.02 

.23 

.28 

.00 

.74 

.83 

.04 

.01

4.02

1.87 

.02 

.22 

.24 

.01 

.75 

.89 

.04 

.00

4.04

1.86 

.01 

.23 

.26 

.01 

.71 

.90 

.03 

.00

4.01

1.79 

.02 

.35 

.90 

.01 

.77 

.67 

.03 

.00

4.54

1.80 

.02 

.28 

.27 

.01 

.75 

.90 

.03 

.00

4.06

1.93 1.85 1.86 

.01 .02 .01 

.17 .27 .21 

.25 .33 .26 

.01 .08 .01 

.75 .73 .75 

.86 .77 .90 

.02 .03 .03 

.00 .001 .00

1.86 

.02 

.23 

.26 

.01 

.72 

.89 

.03 

.00

4.00 4.081 4.03 4.02

1.90 

.01 

.17 

.27 

.01 

.75 

.87 

.04 

.00

4.02

1.88 

.02 

.08 

.38 

.01 

.74 

.91 

.04 

.00

4.06

1.86 

.02 

.24 

.26 

.01 

.71

o/;

.04 

.00

4.00

1.82 

.02 

.31 

.29 

.01 

.71 

.82 

.04 

.00

4.02

Relative proportions (percent)

ps ____ 

En -

Wo

Sample 

Mineral 

grain 

15

40 

45

13

40 

47

14 

38 

48

22 

42 

36

Clinopyroxenes from M-46 

(mafic gabbro)

1 2 3 4

14 

39

47

13 18 14 

40 40 39 

46 42 47

14 

39

47

14 

40 

46

Clinopyroxenes from 

M80b3 (layered Clinopyroxenes from M-91 

ultramafic rock)

1

23 1 Rim

2

Core Rim

19 

36

45

14 

39

47

17 

39

45

Olivines from M-91

1 2 3 4

Major-element oxides (weight percent)

SiO2 - 

Ti02 ----- 

A1203 »- 

PeO* - 

Cr203 ----- 

MnO 

MgO

CaO 

Na2O 

K2O

Total

47.9 

.8 

11.6 

12.2 

.03 

.2 

14.7 

11.6 

1.7 

.04

100.9

46.4 

.7 

10.8 

11.7 

.03 

.3 

13.7 

12.4 

2.1 

.06

98.1

47.5 

1.4 

8.6 

11.7 

.10 

.2 

14.6 

12.0 

1.9 

.20

98.1

54.8 

.1 

2.0 

8.3 

.02 

.20 

17.5 

14.5 

.44 

.00

98.0

53.4 

.1 

3.1 

2.9 

.33 

.1 

15.7 

23.6 

.07 

.01

99.2

53.0 52.8 54.9 53.8 

.01 .20 .22 .18 

4.7 4.3 1.48 1.7 

2.1 3.2 2.0 2.4 

.13 .48 .14 .17 

.11 .10 .04 .10 

18.2 15.7 17.0 16.7 

21.0 24.2 24.7 23.7 

.01 .09 .07 .08 

.00 .00 .00 .02

99.2 101.0 100.5 98.8

54.2 

.21 

1.2 

2.1 

.17 

.08 

16.7 

23.7 

.06 

.01

99.0

54.5 

.10 

1.5 

1.9 

.17 

.10 

17.0 

24.0 

.02 

.01

99.2

40.6 

.11 

.03 

12.2 

.0 

.8 

45.9 

.0 

.0 

.0

99.6

40.6 40.1 

.03 .08 

.00 .03 

14.8 12.5 

.0 .03 

.20 .6 

44.1 46.0 

.0 .0 

.0 .0 

.0 .0

99.7 100.3

40.5 

.04 

.04 

14.3 

.00 

.4 

44.72 

.0 

.0 

.0

100.0

Stoichiometry (number of ions per unit cell) 2

Si - 

Ti -

Al

pe ____

Mn - 

Mg----

K 

Cr 

Total

3.50 

.04 

1.00 

.75 

.014 

1.60 

.91 

.24 

.00 

.02

8.074

3.51 

.04 

.96 

.74 

.02 

1.55 

1.00 

.31 

.01 

.00

8.14

3.59 

.08 

.76 

.74 

.01 

1.64 

.97 

.29 

.02 

.01

8.11

4.00 

.01 

.98 

.52 

.01 

1.55 

.70 

.063 

.00 

.001

7.834

3.91 

.01 

.27 

.17 

.01 

1.71 

1.85 

.01 

.00 

.019

7.959

3.84 3.81 3.96 3.95 

.001 .01 .01 .01 

.40 .37 .13 .15 

.13 .19 .12 .15 

.01 .01 .00 .01 

1.96 1.69 1.83 1.83 

1.63 1.87 1.91 1.86 

.00 .01 .01 .01 

.00 .00 .00 .002 

.01 .03 .01 .01

7.981 7.99 7.98 7.982

3.96 

.01 

.14 

.13 

.01 

1.82 

1.86 

.01 

.00 

.01

7.95

3.97 

.01 

.13 

.11 

.01 

1.85 

1.87 

.01 

.01 

.01

7.98

...

_

...

...

...

-

Relative proportions (percent)

Po 

Ps .. 

En -

Wo - 

n.a.

23 

49 

28

n.a.

22 

47 

30

n.a.

22 

49 

29

n.a.

29 

31

40

n.a.

5 

46 

50

n.a. n.a. n.a. n.a.

35 34 

53 45 47 48 

44 50 49 48

n.a.

4 

48 

48

n.a.

3 

48 

49

89

n.a. 

n.a. 

n.a.

85 90

n.a. n.a. 

n.a. n.a. 

n.a. n.a.

84

n.a. 

n.a. 

n.a.

1 Assuming 6 oxygen atoms per unit cell of dinopyroxene. 

2Assuming 12 oxygen atoms per unit cell of dinopyroxene.
242 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 5.7. Major-element analyses of greenstones of the Baker terrane, northeastern Oregon 

[Values in weight percent; , not determined]

Sample- PH-76 MV-48 MV-48A DXB-17 DXB-19 DXB-20 V-252 NF-64 NF-65 BR-49

Si02 ----- 

A12Q3 

TiO2 - 

Fe2OQ 

£j tj

FpO __

MnO -

MgO ----- 

CaO 

Na2O -

K2O 

P205 ------

Total--

Sample 

SiO2 -

AloOq -

£t VO

Ti02 -----

Fe203 --- 

FeO 

MnO -

MgO -

faO

MaoO

K20-~- -

P2Oc -

& D

Total 

47.5

99 fl

1.1

4.5

c 1

1 a

3

/1

11 Q

q K

.01

iq

99.5

BR-52

Kq A

16.4

q

4

j^

5

C

IS

5

Q

8 7

3 7

09

09

99.3

^9 9

i a c\

1 ^

K 0

C Q

i a

6

0

7 9.

q q

1 ^7

1 a

99.7

BR-59

co q

17.7

4.2

A Q

iq

6

c

9

0

1 9

09

OQ

99.6

C1 K

-ten

1.4

5

C

6

0

.14

6

1

o q

q fl

1.07

1 C

99.8

BR-59A

56.4

16.4

q

4.2

4.8

.14

C 7

fi Q

4.1

nq

OS

99.7

CC A

91 ^

1.0

q 9

3

/1

.11

q fl

7 S

q q

.54

is

99.6

V-27

47 Q

i a i

9 0

£*.\J

4.5

7.1

9.5

«7.U

q 1

.04

q

99.9

cq A

17 c

1 ^

5.4

6

0

O/J

5.4

2

0

.54

oe

99.7

BR-5

4Q 8

1 Q C

q o

5 1

S 9

9

6

Q

A O

q c

flq

.4

99.6

cq c

IS 1

1 3

q o

4.4

IS

5.9

7 S

2

0

O£?

99.7

M-42

4Q 97

1 Q q9

9Q

1 3

6.5

fi ^7

15.62

C1

OQ

99.15

C7 i

1 R f\

1.1

c q

6

r\

1 S

q c

fi Q

o q

09

iq

99.5

M-46

C1 CC

IQ QO

91

I

n a

4 91

15.10

CQ

.11

99.82

CC fl

01 a

q

q q

4.5

.11

3 q

4 Q

4.6

f\C

.11

99.6

M-208

45.87

19 cq

qo

1 ^9

7 09

14.40

IE qc

91

rv/j

97.88

cq q

IS S

q

4.4

5 1

.17

4

rj

8 q

3 A

19

1^

99.5

M-213

48.19

9fl 1 q

99

3

1 a

7 QO

1 7 3^

/>i

flC

99.77

ci 9

i Q a

.7

4.1

4.7

oc

5

Q

U o

q 9

.01

.11

99.9

M-240

1 *7 £Q

.46

q 07

5 /jo

19 79

15.80

.41

flC

99.70

land-arc to island-arc-tholeiite geochemical signatures. Silicic metavolcanic rocks (53-55 percent SiO2) 

are present along the North Fork of the John Day 

River. These rocks are strongly deformed, but in some 

localities an intergranular texture is preserved. They 

plot in the destructive-margin field on a Hf-Ta-Th discriminant diagram (Wood and others, 1979) and in 

fields for island-arc tholeiite and MORE on a MnO/ 

TiO2/P2O5 discriminant diagram (Mullen, 1983b). 

These rocks have a slightly LREE-depleted REE pattern, with a substantially negative Eu anomaly. The 

fairly silicic composition of these greenstones, their 

intergranular character, and the geochemical discriminants suggest an island-arc origin for these rocks as 

well.

Metavolcanic rocks that may have MORE affinity 

occur as knockers of pillowed greenstone in serpentinite-matrix melange near Mount Vernon, Oreg. (fig. 

5.1). They fall within MORE fields on geochemical 

discriminant diagrams, and have flat lOxchondrite 

REE patterns. Relict clinopyroxenes have compositions typical of clinopyroxene from midocean-ridge 

basalts or arc tholeiites (sample MV-48A, table 5.6).

Similar compositions have been reported for the 

same area by Carpenter and Walker (1992).

Alkalic mineralogy and geochemistry characterize 

pillowed metabasalts at two widely separated localities. Well-preserved titanaugite has nearly identical 

compositions in the pillowed basalts from both the 

Greenhorn Mountains (sample V-27A, table 5.6) and 

Sturgill, Oreg., on the west bank of the Snake River 

(fig. 5.1; sample BRS, table 5.6). Pyroxene compositions indicate alkalic affinity for these rocks. The 

metabasalts are high in TiO2 (2-3.4 weight percent) 

and are mildly LREE-enriched. They are similar in 

petrography and geochemistry to oceanic transform 

basalts, as well as to many basalts recently reported 

from rifted-arc environments.

Fragments of plutonic ophiolitic rocks, principally 

gabbro, occur throughout the Baker terrane. They 

range in dimension from meter-sized knockers in serpentinite matrix to masses several kilometers square. 

They are usually deformed and metamorphosed to 

lower greenschist facies. Metagabbro exposed in the 

North Fork of the John Day River (fig. 5.1) is so 

strongly sheared that its phaneritic character is vir-
5. MAFIC AND ULTRAMAFIC ROCKS OF BAKER TERRANE, EASTERN OREGON 243

tually obliterated. Other gabbro, especially southsoutheast of Olive Lake in the Greenhorn Mountains, 

has been locally transformed into flaser gneiss. Most 

plutonic rocks are metamorphosed to greenschist facies. Amphibolite and garnet amphibolite facies rocks 

are also present, mostly in the western part of the 

terrane. There is little apparent stratigraphic order 

to mafic and ultramafic plutonic blocks within the 

Baker terrane. Recognizable lithic types include isotropic clinopyroxene-rich gabbro, cumulate clinopyroxene-rich gabbro, layered two-pyroxene gabbro, 

layered peridotite and gabbro, clinopyroxenite (wehrlite), dunite, and harzburgite.

Clinopyroxenes from plutonic ophiolitic rocks of the 

Baker terrane plot in a broad scatter on a pyroxene 

quadrilateral. Immediately west of the Canyon 

Mountain Complex, these rocks resemble the complex

MnOoxlO P2O5x10

EXPLANATION 

Greenstones from northeastern Oregon

 BR Burnt River Schist

O NF North Fork John Day River

 V252 Greenhorn Mountains 

D DXB Dixie Butte 

A PH Pleasant Hill 

A MV Mount Vernon 

O BRS-3 Snake River 

^ V-27 Greenhorn Mountains 

^^ Field of samples from the Canyon Mountain Complex

FIGURE 5.14. MnO/TiO2/P2O5 plot for mafic and ultramafic 

rocks of the Canyon Mountain Complex (shaded area) and 

greenstones of northeastern Oregon (data points). Abbreviations for fields of basalt compositions are as follows: CAB, 

calc-alkaline basalt; IAT, island-arc tholeiite; MOR, midocean 

ridge; OIA, ocean-island alkalic; OIT, ocean-island tholeiite. 

Modified from Mullen, 1983b.

in composition (Carpenter and Walker, 1992). Elsewhere, the chemistry and pyroxene compositions of 

the examined blocks do not overlap those of the Canyon Mountain Complex. Pyroxene samples from the 

Greenhorn Mountains, near Olive Lake (fig. 5.1; sample M42, table 5.6), are low-Cr augite, are enriched 

in TiO2 and M^O^, and plot in alkalic or island-arc 

fields on discriminant diagrams. Gabbro from Ben 

Harrison peak (sample M46, table 5.6), also in the 

Greenhorn Mountains, yields low-Ti augite similar to 

arc-tholeiite clinopyroxene. Clinopyroxenes from ultramafic rocks in the Baker terrane (samples M80, 

M91, table 5.6) are less calcic than clinopyroxene 

from lithically equivalent units in the Canyon Mountain Complex, and are subalkalic. On pyroxene discriminant diagrams, Clinopyroxenes from the plutonic 

rocks plot in island-arc fields (fig. 5.13).

Whole-rock geochemistry of the gabbro s and peridotites in the Baker terrane is of dubious value as a 

tectonic discriminant because the rocks are thoroughly altered and deformed. On a plot of SiO2 versus 

FeO*/MgO and on an AFM diagram, several compositions are apparently more tholeiitic than those of 

rocks of the Canyon Mountain Complex. However, 

this difference could be due simply to greater mobilization of MgO in the Baker terrane samples.

Trace-element data also show some differences between the Canyon Mountain Complex and lithically 

equivalent plutonic fragments from the Baker terrane. The gabbro that contains clinopyroxene with 

compositions characteristic of arc tholeiite has a relatively flat 4-12xchondrite REE pattern with slight 

LREE depletion and a small positive Eu anomaly, 

similar to the REE pattern of oceanic cumulate gabbro (Tiezzi and Scott, 1980). Peridotites of the Baker 

terrane analyzed in this study are far more REE-enriched and especially LREE-enriched than lithically 

equivalent units of the Canyon Mountain Complex. 

Although this enrichment may be due to alteration or 

to limited metasomatism from nearby granitic intrusions, it may also be related to original differences.

IMPLICATIONS FOR THE ORIGIN 

OF THE BAKER TERRANE

The origin of the Baker terrane can be read in the 

lithic types, geochemistry, and metamorphic facies of 

its components. The greenstones of the Baker terrane 

come from arc, midocean-ridge, and rift or seamount 

environments. The greenstones are incorporated into 

serpentinite-matrix melange with moderate- to highpressure schists, allochthonous limestones, and diverse clastic sedimentary rocks. The environment
244 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

that best fits this assemblage is a subduction zone or 

deep fore-arc setting.

The Canyon Mountain Complex clearly carries an 

island-arc signature. Its gabbros have arc-tholeiite affinities. Its overall stratigraphy, including the upper 

plagiogranites and keratophyres with calc-alkaline geochemistry, is more characteristic of arc than of oceanic crust. The timing and nature of later magmatic 

events (gabbro infiltration, migmatitic melting and recrystallization) need further clarification. However, 

high shear stress may have played a role in them.

The greenstones represent environments that fit 

well in a subduction zone or deep fore-arc environment. The most abundant greenstones in the Baker 

terrane have island-arc geochemical affinities and 

contain early-crystallizing clinopyroxenes with intergranular textures generally characteristic of islandarc basalts. These rocks probably represent the early, 

basal material of arc volcanoes, rather than volcanic 

cones or segments of stratovolcanoes. Alkalic greenstones may be related to intra-arc rifting, transform 

(enriched-MORB, or EMORB) basalts, or oceanic-island environments. MORE affinities are most abundant in the western part of the terrane, fitting well 

with a model of east-facing subduction, at least in 

the Triassic.

Many plutonic components of the Baker terrane, 

aside from the Canyon Mountain Complex, seem to 

be more tholeiitic than the volcanic greenstones. They 

are distinct from the arc-related Canyon Mountain 

Complex in mineralogy and geochemistry. They may 

represent oceanic crust or primitive arc magmas.

The data presented above indicate that the Baker 

terrane contains many petrotectonic units that formed 

in or could have been associated with a subductionzone or deep fore-arc environment. Detailed studies of 

fore arcs in the Marianas, Tonga, and Kermadec active 

arc systems in the western Pacific Ocean (Bloomer, 

1983; Bloomer and others, 1988) have revealed that a 

combination of calc-alkaline and arc-tholeiite affinities 

in volcanic and plutonic components is characteristic 

of the submarine parts of fore arcs. Additional components of the erosive, active Marianas fore arc reported 

by Bloomer and Hawkins (1983) are fragments of oceanic crust, serpentinite diapirs, and alkalic basalts. 

Plutonic rocks are deformed. Siliceous sedimentary 

rocks, cherts, and allochthonous shallow-water limestones are also recognized by Bloomer (1983). Coarse 

sediments recognized in Deep Sea Drilling Project 

(DSDP) cores from the Marianas fore arc include conglomerates and breccias. All these components are 

present in the Baker terrane, along with a seemingly 

coherent ophiolite that represents early island-arc or 

fore-arc magmatism.

Thus, a reasonable model for the Baker terrane, 

based upon the mineralogy and geochemistry of its 

mafic and ultramafic components, the variety of its 

incorporated igneous, sedimentary, and metamorphic 

lithic types, and the preponderance of sheared serpentinites, is an erosive fore arc and remnant subduction zone. The rarity of glaucophane blueschists 

in the Baker terrane (Bishop, chap. 4, this volume; 

Mullen, 1978) suggests that shear stresses in the 

subduction zone were high, exceeding 50 megapascals (about 5 kilobars), and that subduction rates 

were high (greater than about 10 cm/yr) according to 

amphibole stability work by Peacock (1992). Subduction rates and shear stress may have been lower in 

the Triassic, during eastward-directed subduction 

proposed by Vallier (1992), if the blueschists near 

Mitchell reflect lower shear stresses related to Triassic subduction and formation of the Olds Ferry terrane. The mostly Permian to Early Triassic age of 

rocks in the Baker terrane suggests that the Baker 

terrane began as a fore arc and subduction zone related to the Wallowa arc and that it underwent a 

complex history of rapid subduction, high shear 

stresses, internal mixing, and a possible subductionzone polarity reversal (Vallier, 1992). Generation of 

plagiogranites of the Canyon Mountain Complex at 

about 268 Ma, and the dynamothermal metamorphic 

ages of 255 to 269 Ma from melange just west of the 

Canyon Mountain Complex, may be related to a tectonic or petrotectonic event such as subduction-zone 

polarity reversal, or to the inception of a different 

subduction zone. This event may also be tied to slower subduction in the Triassic and eventual generation 

of the Olds Ferry terrane.

REFERENCES CITED

Ave Lallemant, H.G., 1976, Structure of the Canyon Mountain (Oregon) ophiolite complex and its implication for seafloor spreading: Geological Society of America Special Paper 173, 49 p.

Bishop, E.M., 1988, Igneous petrology of the Baker Melange Terrane and the Canyon Mountain complex, eastern Oregon: Geological Society of America Abstracts with Programs, v. 20, 

p. 407.

Blome, C., and Nestell, M., 1992, Field guide to the geology and 

paleontology of pre-Tertiary volcanic arc and melange rocks, 

Grindstone, Izee, and Baker terranes, east-central Oregon: 

Oregon Geology, v. 554, p. 123-141.

Bloomer, S., 1983, Distribution and origin of igneous rocks from 

the landward side of the Mariana trench Implications for its 

structure and evolution: Journal of Geophysical Research, 

v. 88, p. 7411-7428.

Bloomer, S.H., and Hawkins, J.W., 1983, Gabbroic and ultramafic 

rocks from the Mariana Trench An island arc ophiolite, in 

Hayes, D.E., ed., The tectonic and geologic evolution of Southeast Asia seas and islands, Part II: American Geophysical 

Union Monograph 27, p. 294-317.
5. MAFIC AND ULTRAMAFIC ROCKS OF BAKER TERRANE, EASTERN OREGON 245

Bloomer, S.H., Parsons, L., Huggett, Q., and Pearce, J., 1988, 

GLORIA Surveys of the Tonga and Kermadec trenches: Eos 

(Transactions, American Geophysical Union), v. 69, p. 1442.

Boudier, F, and Coleman, R.G., 1981, Cross section through the 

peridotite of the Samail ophiolite, southeastern Oman Mountains: Journal of Geophysical Research, v. 86, p. 2573-2592.

Brooks, H.G., 1979, Plate tectonics and the geologic history of the 

Blue Mountains: Oregon Geology, v. 41, p. 71-80.

Carpenter, PS., and Walker, N.W., 1992, Origin and tectonic significance of the Aldrich Mountain serpentinite matrix melange, 

northeastern Oregon: Tectonics, v. 11, p. 690-708.

Coleman, R.G., 1977, Ophiolites: New York, Springer Verlag, 229 p.

Dickinson, W.R., 1979, Mesozoic forearc basin in central Oregon: 

Geology, v. 7, p. 166-170.

Dungan, M.A., and Ave Lallemant, H.G., 1977, Formation of small 

dunite bodies by metasomatic transformation of harzburgite 

in the Canyon Mountain ophiolite, northeast Oregon, in Dick, 

H.J.B, ed., Magma genesis: Oregon Department of Geology 

and Mineral Industries Bulletin 96, p. 109-128.

Flager, PA., and Spray, J.G., 1991, Generation of plagiogranite by 

amphibolite anatexis in oceanic shear zones: Geology, v. 19, 

p. 70-73.

Gerlach, D., 1980, Petrology and geochemistry of plagiogranite and 

related basic rocks of the Canyon Mountain ophiolite, Oregon: 

Houston, Rice University, M.S. thesis, 198 p.

Gerlach, D., Ave Lallemant, H.G., and Leeman, W.P., 1981, An island arc origin for the Canyon Mountain complex, eastern Oregon, USA: Earth and Planetary Science Letters, v. 53, p. 255- 

265.

Gerlach, D., Leeman, W.P., Bishop, E.M., and Pearcy, L.A., 1988, 

Isotopic and trace element constraints on timing and the nature of magmatism in the Canyon Mt. Ophiolite, Oregon: Geological Society of America Abstracts with Programs, v. 20, 

p. A158.

Himmelberg, G.H., and Loney, R.A., 1980, Petrology of ultramafic 

and gabbroic rocks of the Canyon Mountain ophiolite, Oregon: 

American Journal of Science, v. 280A, p. 232-268.

Irvine, T.N., 1980, Magmatic infiltration metasomatism, doublediffusive fractional crystallization, and adcumulus growth in 

the Muskox intrusion and other layered intrusions, in Hargraves, R.B., ed., Physics of magmatic processes: Princeton, 

N.J., Princeton University Press, p. 325-383.

Jorgenson, D.B., 1979, Textural banding in igneous rocks An example from SW Oregon: American Mineralogist, v. 64, p. 527- 

530.

Marlow, M.S., Johnson, L.E., Pearce, J.A., Fryer, P.B., Pickthorn, 

L.B., and Murton, B.J., 1992, Upper Cenozoic volcanic rocks 

in the Mariana forearc recovered from drilling at Ocean Drilling Program site 781 Implications for forearc magmatism: 

Journal of Geophysical Research, B, Solid Earth and Planets, 

v. 97, no. 11, p. 15,085-15,097.

Moore, A.C., 1973, Studies of igneous and tectonic textures and 

layering in rocks of the Gosse Pile intrusion, central Australia: Journal of Petrology, v. 14, p. 49-79.

Morse, S.A., 1969, The Kiglapait layered intrusion, Labrador: Geological Society of America Memoir 112, 204 p.

Mullen, E.D., 1978, Geology of the Greenhorn Mountains, northeastern Oregon: Corvallis, Oregon State University, M.S. thesis, 372 p.

 1983a, Petrology and regional setting of peridotite and gabbro of the Canyon Mountain Complex, northeast Oregon: Corvallis, Oregon State University, Ph.D. dissertation, 277 p.

 1983b, MnO/TiO2/P2O5 A minor element discriminant for

basaltic rocks of oceanic environments and its implications for

petrogenesis: Earth and Planetary Science Letters, v. 62,

p. 53-62.

-1985, Petrologic character of Permian and Triassic greenstones from the Melange terrane of eastern Oregon, and their 

implications for terrane origin: Geology, v. 13, p. 131-134.

Nesbitt, E.G., and Pearce, J., 1977, Clinopyroxene composition in 

mafic lavas from different tectonic settings: Contributions to 

Mineralogy and Petrology, v. 63, p. 149-160.

Nicolas, A., and Jackson, M., 1982, High temperature dikes in 

peridotites Origin by hydraulic fracturing: Journal of Petrology, v. 23, p. 568-582.

Peacock, S.M., 1992, Blueschist-facies metamorphism, shear heating, and P-T-t paths in subduction zones: Journal of Geophysical Research, v. 97, p. 17,693-17,708.

Pearce, T.H., German, B.E., and Birkett, T.C., 1977, The relationship between major element geochemistry and tectonic environment of basic and intermediate volcanic rocks: Earth and 

Planetary Science Letters, v. 36, p. 121-132.

Pearcy, L.G., 1987, Clinopyroxene compositions and petrogenesis of 

the mantle-crust transition zone: Geological Society of America Abstracts with Programs, v. 19, p. 439.

 1991, Nascent island arc petrogenesis for the basal crustal 

cumulates of the Canyon Mountain Complex, northeastern 

Oregon: Geological Society of America Abstracts with Programs, v. 23, no. 5, p. A272.

Silberling, N.J., and Jones, D.L., 1984, Lithotectonic terrane maps 

of the North American Cordillera: U.S. Geological Survey 

Open File Report 84-523, 43 p.

Thayer, T.P., 1972, Gabbro and epidiorite versus granulite and amphibolite A problem in the ophiolite assemblage: Sixth Caribbean Geological Conference, Margarita, Venezuela, p. 315-320.

 1976, Metallogenic contrasts in the plutonic and volcanic 

rocks of the ophiolite assemblage: Geological Association of 

Canada Special Paper No. 14, p. 211-219. 

-1977, The Canyon Mountain complex, Oregon, and some

problems of ophiolites, in Coleman, R.G., and Irwin, W.P., eds., 

North American ophiolites: Oregon Department of Geology 

and Mineral Industries Bulletin 95, p. 93-105.

Thy, P., and Esbensen, K.H., 1982, Origin of fine-grained granular 

rocks in layered intrusions: Geological Magazine, v. 119, 

p. 405-412.

Tiezzi, L.J., and Scott, R.B., 1980, Crystal fractionation in a cumulate gabbro, mid-Atlantic ridge, 26 N: Journal of Geophysical 

Research, v. 85, p. 5454-5488.

Vallier, T.L., 1992, Late Triassic change in convergence direction, 

Blue Mountains island arc, Oregon and Washington: Geological Society of America Abstracts with Programs, v. 24, p. 88.

Vernon, R.H., 1970, Comparative grain boundary studies of some 

basic and ultrabasic granulites: Scottish Journal of Geology, 

v. 6, p. 337-351.

Wood, D.A., Jordon, J.L., and Treuil, M., 1979, A re-appraisal of 

the usage of trace elements to classify and discriminate between magma series erupted at different tectonic settings: 

Earth and Planetary Science Letters, v. 45, p. 326-336.

6. TECTONIC IMPLICATIONS OF U-Pb ZIRCON AGES OF

THE CANYON MOUNTAIN COMPLEX, SPARTA COMPLEX, AND

RELATED METAPLUTONIC ROCKS OF THE BAKER TERRANE,

NORTHEASTERN OREGON

By NICHOLAS W. WALKER!

CONTENTS

Page

Abstract 247

Introduction 247

Acknowledgments 248

Description of terranes 250

U-Pb geochronology 250

Analytical methods 250

Mass spectrometry 251

Procedural Pb blanks 251

Analytical uncertainty 251

Geochronologic results 251

Elkhorn Ridge argillite 251

Ultramafic-mafic-silicic ophiolitic igneous suites 256

Canyon Mountain Complex 256

Previous geochronologic investigations 259

U-Pb zircon geochronologic results 259

Sparta comolex_________________________ 260

Previous geochronologic investigations 261

U-Pb zircon geochronologic results 263

Serpentinite-matrix melange 263

Melange in the Aldrich Mountains 263

Timing of melange formation 265

Paleotectonic setting of the Baker terrane 265

Summary and conclusions 267

References cited 267

ABSTRACT

The Baker terrane consists of plutonic, metamorphic, and sedimentary rocks of oceanic origin that now are the components of 

kilometer-scale tracts of melange and broken formation. The structural character of the terrane resulted from the breaking and mixing 

of Permian to Triassic volcanic-plutonic arc crust and sediment cover. Plutonic and metaplutonic constituents of this terrane represent 

this volcanic-plutonic arc crust and include the ophiolitic Canyon

Present address: Department of Geological Sciences, Brown University, Providence, 

RI 02912

Mountain Complex, the Sparta complex, and other tectonic blocks in 

melange and broken formation. U-Pb zircon ages for separate intrusive phases in the Canyon Mountain Complex vary from 276 to 268 

Ma; this age range indicates an Early Permian multistage magmatic 

history for the complex. The Sparta complex exhibits a two-stage 

igneous history that involved magmatic activity at 253 Ma and 215 

Ma. Other plutonic bodies in the Baker terrane range in crystallization age from 279 Ma to 230 Ma; this range overlaps the age of 

Permian to Triassic sedimentation within the terrane. These ages 

are also synchronous with the age of plutonism within the Wallowa 

and Olds Ferry terranes (268-225 Ma), which have been interpreted 

as vestiges of ensimatic volcanic-plutonic complexes. Inferences 

drawn from geochronology, lithology, and structure of the Baker 

terrane suggest that it constituted the fore-arc region of a PermianTriassic arc-type ensimatic volcanic-plutonic complex now represented by the Wallowa and Olds Ferry terranes.

INTRODUCTION

Numerous erosional inliers in the extensive Cenozoic volcanic and sedimentary blanket of northeastern Oregon and westernmost central Idaho unveil a 

lithologically diverse and structurally complex assemblage of late Paleozoic to late Mesozoic rocks of oceanic affinity (Vallier and others, 1977; Brooks and 

Vallier, 1978; Dickinson and Thayer, 1978). These 

rocks are exposed in a generally northeasterly trending, discontinuous belt of inliers extending approximately 350 km between the towns of Mitchell, Oreg., 

and Grangeville, Idaho (fig. 6.1).

This heterolithologic assemblage includes variably 

deformed and metamorphosed igneous rocks and sedimentary rocks derived from volcanic-plutonic complexes, ultramafic-mafic-silicic igneous rocks, chertargillite broken formation, polymict melange tracts, 

and thick sequences of chiefly Upper Triassic to 

Upper Jurassic volcaniclastic sedimentary rocks all 

of which have been intruded by numerous gabbroic to 

granodioritic plutons of Late Jurassic to Early 

Cretaceous age (Taubeneck 1957, 1958, 1964, 1967;

247
248 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Armstrong and others, 1977; Walker, 1986, 1989). 

The entire assemblage is regarded as allochthonous 

(Vallier and others, 1977; Hillhouse and others, 1982) 

and is among the most inboard of paleo-oceanic terranes within the Cordillera.

Rocks within this assemblage are considered to 

represent the five terranes distinguished by Silberling and others (1984): the Wallowa, Olds Ferry, 

Grindstone, Izee, and Baker terranes. These terranes, whose distribution is shown in figure 6.2, are 

bounded by high- to low-angle reverse faults or highangle normal faults, commonly marked by zones of 

serpentinite. The terranes are crosscut by latest Jurassic to Early Cretaceous plutons (Armstrong and 

others, 1977, Walker, 1986, 1989). At some localities, 

Early Cretaceous plutons intrude interterrane bound120°

aries (Walker, 1986, 1989). Such a relationship indicates that structural juxtaposition of terranes dates 

from the Early Cretaceous. After the Early Cretaceous but prior to the Eocene, the entire terrane assemblage rotated clockwise 60°±29° (Wilson and Cox, 

1980).

This chapter presents U-Pb zircon geochronologic 

data for igneous and metaigneous components of the 

Baker terrane and describes the tectonic implications 

of these age data.

ACKNOWLEDGMENTS

James Mattinson, George Tilton, and Cliff Hopson 

of the University of California at Santa Barbara

118° 116°

44

EXPLANATION

Surficial deposits and sedimentary rocks (Cenozoic)

Idaho batholith and related plutons

Rocks of border zone of Idaho batholith

Rocks of Blue Mountains province (pre-Cenozoic)

Belt Supergroup (Proterozoic)

FIGURE 6.1. Regional geologic setting of pre-Cenozoic rocks of the Blue Mountains

province.
6. CANYON MOUNTAIN COMPLEX, SPARTA COMPLEX, AND RELATED ROCKS, NORTHEASTERN OREGON 249

supervised the dissertation research on which this 

report is based. Bob Fleck and Tracy Vallier of the 

U.S. Geological Survey and Howard Brooks of the 

Oregon Department of Geology and Mineral Industries (DOGAMI) provided helpful reviews of the 

manuscript. Tracy Vallier, Trudy Vallier, David 

Scholl, and Ellery Ingall assisted in fieldwork. Tracy 

Vallier, Howard Brooks, and Mark Ferns (DOGAMI)

120° 119° 118° 117° 116°

m

EXPLANATION

Cretaceous to Jurassic plutons

Wallowa terrane

Baker terrane

Izee terrane

Olds Ferry terrane

Grindstone terrane

Contact

Thrust fault Sawteeth on upper plate

FIGURE 6.2. Terrane map of northeastern Oregon and westernmost central Idaho (modified from Silberling and others, 1984).
250 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

led me on numerous field trips and shared their 

abundant knowledge of regional geologic relations in 

the Baker terrane and in northeastern Oregon and 

west-central Idaho. This research was supported by 

the U.S. Geological Survey, by grants from the Oregon Department of Geology and Mineral Industries, 

Sigma Xi, and the Penrose fund of the Geological Society of America, and by a National Science Foundation grant to James Mattinson.

DESCRIPTION OF TERRANES

The Wallowa terrane consists of a thick sequence of 

volcanic flows, volcaniclastic rocks, and intercalated 

sedimentary rocks of the Permian and Trias sic Seven 

Devils Group (Vallier, 1977; chap. 3, this volume) 

which is capped by Triassic platform-carbonate and 

fine-grained clastic rocks of Triassic and Jurassic age. 

Crystalline rocks that represent the basement and 

plutonic infrastructure of the Seven Devils Group 

have magmatic ages that range from 309 Ma to 225 

Ma (Walker, 1986). The Wallowa terrane is interpreted as part of an arc-type ensimatic volcanic-plutonic 

complex capped by intra-arc sedimentary rocks (Vallier, 1977; Vallier and others, 1977).

The Olds Ferry terrane consists of Upper Triassic 

volcanic and volcaniclastic rocks, intercalated sedimentary rocks, and volumetrically minor plutonic 

rocks; one representative pluton has yielded Middle 

Triassic U-Pb zircon ages (Walker, 1986). This assemblage, considered to be part of an arc-type ensimatic 

volcanic-plutonic complex (Vallier and others, 1977), 

is unconformably overlain by Lower and Middle Jurassic clastic strata of the Izee terrane, which are interpreted as fore-arc deposits by Dickinson (1979).

The Grindstone terrane is a melange tract consisting 

of structurally chaotic, generally lenticular blocks of 

silicic metavolcanic rocks, volcaniclastic greywacke, 

limestone, and sandstone in a matrix of deformed radiolarian chert and argillite. Although the major lithologic boundaries are faults, relatively coherent and 

stratigraphically distinct bodies are present. Faunal 

ages of these rocks range from Middle Devonian to 

Permian. Dickinson and Thayer (1978) suggested that 

the overall lithologic association and structural juxtaposition of silicic metavolcanic rocks, greywacke, and 

limestone in a matrix of deformed chert and argillite 

are evidence that this terrane represents a subductionrelated complex.

The Izee terrane is a thick (locally exceeding 15 

km) sequence of Upper Triassic to Upper Jurassic 

turbidites rich in volcanogenic components and lesser 

amounts of volcanic flows and tuffs. Dickinson (1979) 

interpreted this terrane to be of fore-arc origin.

The Baker terrane is an elongate tract with a general northeasterly trend and is exposed over a wide 

area (fig. 6.2). The terrane consists of a mixture of 

structural blocks comprising variably metamorphosed 

and deformed igneous suites (some of which are characteristically ophiolitic in stratigraphy), chert-argillite broken formation, polymict serpentinite-matrix 

melange, and limestone slabs. The southwestern part 

consists mainly of melange whereas more structurally coherent domains of broken formation distinguish 

the northeastern part. No basement to this terrane 

has been recognized. The overall dismembered to 

chaotic character of the terrane resulted from the 

breaking and mixing of the original plutonic and 

metamorphic crust upon which sedimentary and volcanic rocks had accumulated. The only likely presentday exposures of such crystalline rocks are represented by kilometer-scale slabs and blocks such as the 

Canyon Mountain Complex, the informally named 

Sparta complex of Phelps (1979), and other smaller 

plutonic and metamorphic fragments in melange. 

Considering that these plutonic rocks appear similar 

in the field, it is of critical importance to establish 

the crystallization ages of such components within 

the Baker terrane because their ages bear significantly on the tectonic evolution of this terrane as 

well as the entire pre-Tertiary terrane complex of 

northeastern Oregon.

U-Pb GEOCHRONOLOGY

ANALYTICAL METHODS

Isotopic analyses were conducted from 1980 

through 1984 in the laboratories of J.M. Mattinson 

and G.R. Tilton at the University of California at 

Santa Barbara.

Zircon concentrates were obtained from 20 to 100 

kg of sample by standard procedures of crushing and 

mineral separation using magnetic, heavy liquid, and 

Wilfley table techniques. Zircon concentrates were 

split into fractions on the basis of physical properties 

such as size and magnetic character and then handpicked to remove impurities. Purified zircon fractions 

weighing from 15 to 30 mg were weighed into Teflon 

dissolution capsules and given final acid washes in 

ultrapure 6N HC1 and 7N HNO3 . Approximately 1 ml 

of ultrapure concentrated HF and 25 ul of concentrated HNO3 were added to the dissolution capsule 

for digestion. A mixed U-Pb tracer was either delivered to a capsule that contained a representative dry 

split of the zircon fraction or delivered to an aliquot 

of the total solution following digestion. The dissolu-
6. CANYON MOUNTAIN COMPLEX, SPARTA COMPLEX, AND RELATED ROCKS, NORTHEASTERN OREGON 251

tion capsules were inserted in steel jackets, sealed, 

and placed in a 210°C oven for four to six days. Then 

the HF-HNC>3 mix was evaporated and fresh acid 

added before an additional four-day digestion. The 

HF-HNC>3 was evaporated in a clean-air environment 

and ultrapure 3N HC1 was added to prepare the zirconium salts for ion exchange chemistry. The method 

of chemical separation of U and Pb employed was 

similar to that of Krogh's (1973).

MASS SPECTROMETRY

Pb and U were loaded onto single rhenium filaments for mass spectrometry. Pb was loaded using a 

silica gel-phosphoric acid technique similar to that 

described by Cameron and others (1969); U was loaded following a phosphoric acid-graphite method similar to that of Arden and Gale's (1974).

Isotopic ratios of Pb and U were measured on a 35- 

cm-radius, 90°-sector, single-focusing, solid-source 

mass spectrometer. Within-run precision (at the twosigma level) of the measured isotopic ratios in every 

analysis was better than ±0.1 percent for 235U/238U, 

±0.2 percent for 208Pb/206Pb, ±0.15 percent for 

207Pb/206Pb; for 206Pb/204Pb, it ranged from ±0.1 percent for low 206Pb/204Pb ratios (<1,000) to as great as 

±11 percent for very high 206Pb/204Pb ratios (>17,000).

Measured U and Pb ratios are corrected for instrumental mass fractionation by replicate analyses of 

National Bureau of Standards (NBS) Pb standards 

SRM 981, SRM 982, and SRM 983, and NBS U 

standard SRM U500.

PROCEDURAL Pb BLANKS

Procedural Pb blanks were less than 0.3 ng whereas the total Pb concentration of individual analyzed 

fractions ranged from approximately 50 to 200 ng. 

Thus, blank Pb accounts for no more than 0.15 to 0.6 

percent of the total analyzed Pb. The procedural Pb 

blank is therefore negligible and no blank corrections 

have been applied to the data.

ANALYTICAL UNCERTAINTY

Uncertainty in the calculated ages stems mainly 

from the precision of the within-run isotopic measurements, uncertainty in the common Pb composition 

(that is, blank plus initial Pb), and long-term reproducibility of analyses of NBS Pb and U standards. 

The combined uncertainty for individual ages reported herein was calculated on the basis of these factors.

GEOCHRONOLOGIC RESULTS

Samples used in geochronologic studies are described in table 6.1, zircon isotopic data are presented in table 6.2, and sample localities are described in 

figure 6.1 and figures 6.3, 6.4, and 6.6 through 6.8.

In many cases, the data presented in this chapter 

are slightly discordant (that is, 206pb*_238u age ^

207pb*_235u age ^ 207pb*_206pb* age) Assigning pre.

cise geologic meaning to slightly discordant U-Pb 

ages is difficult. For geologically young zircons obtained from magmatic rocks in which there are no 

textural grounds for a metamorphic or thermal overprint and in which there is little cause to expect inheritance of zircons on the basis of regional geologic 

considerations, it is commonly observed that the 

207Pb*-206Pb* ages are typically slightly older than 

the Pb-U ages. The cause of this age disparity may 

stem from one or more of the following: (1) uncertainty in the isotopic composition of common Pb used in 

the age calculation, (2) Pb-loss, despite the lack of 

textural evidence for mechanical or thermal metamorphism of the rock that houses the zircon, (3) uncertainty in the decay constants of 235U and 238U, 

and (4) volumetrically small amounts of older, entrained ("inherited") zircon around which magmatic 

zircon grew.

Distinguishing among the possibilities listed above 

is sometimes possible but was beyond the scope of 

this study. In no analyzed zircon fraction does strong 

evidence of inheritance exist and there is no evidence 

in the sampled rocks (other than for metagabbro of 

the Canyon Mountain Complex) of moderate- to highgrade metamorphism following crystallization.

Because insufficient zircon fractions were analyzed 

to assess the cause of the discordance, the 206pb*-238U 

ages are cited as the best estimate of the crystallization 

ages of the sampled rocks.

ELKHORN RIDGE ARGILLITE

The most abundant supracrustal constituent of the 

Baker terrane is the Elkhorn Ridge Argillite (fig. 

6.3), named by Gilluly (1937). The Elkhorn Ridge Argillite is mostly metamorphosed and deformed broken formation composed chiefly of chert-argillite, 

which locally contains crudely defined interbeds or 

blocks of greenstone, tuffaceous volcanic rocks, 

greywacke, limestone, plutonic rocks, and rare conglomerate. The stratigraphic thickness of the 

Elkhorn Ridge Argillite is unknown due to discontinuous exposure, structural complexity, and absence of 

marker horizons.
252 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 6.1. Descriptions and field localities for rock samples used to obtain U-Pb zircon ages, Baker terrane, Oregon

Sample Locality Field setting Rock type, mineralogy, and texture

Plutonic and metaplutonic rocks of the Canyon Mountain Complex

CM79-1 Cany on Mountain, Oregon, 7.5- 

minute quadrangle. Lat 44°18' 

49"N.,longll8°55'39"W.

CM79-3 Canyon Mountain, Oregon, 7.5- 

minute quadrangle. Lat 44° 

20'02" N., long 118°53'06" W.

CM79-4 Canyon Mountain, Oregon, 7.5- 

minute quadrangle. Lat 44° 

19'06" N., long 118°53'27" W.

Small pluton that 

intrudes cumulus 

gabbro.

Structurally concordant 

layer in metacumulate 

sequence.

Sample from sill 

complex.

Hornblende tonalite. Plagioclase, quartz, 

hornblende; secondary sericite, actinolite. 

Hypidiomorphic-granular texture.

Quartz-bearing gabbro. Plagioclase, 

orthopyroxene, clinopyroxene, magnetite, 

quartz; secondary serpentine, saussurite. 

Decussate texture.

Metatrondhjemite. Plagioclase, quartz; 

secondary epidote, saussurite. Cataclastic 

texture.

Plutonic rocks of the Sparta complex

SP-AG-81 Sparta, Oregon, 15-minute

quadrangle. Lat 44°51'13" N., 

longll7°24'49"W.

SP79-1 Sparta, Oregon, 15-minute

quadrangle. Lat 44°49'05" N., 

longll7°24'49"W.

Hillside exposure of 

silicic intrusive rocks.

Roadcut exposure at 

J.N. Bishop Spring.

Trondhjemite. Embayed quartz, plagioclase, 

biotite, hornblende; secondary chlorite. 

Porphyritic texture.

Hornblende-biotite tonalite. Plagioclase, 

quartz, biotite, hornblende, titanite; 

secondary chlorite, saussurite. 

Hypidiomorphic-granular texture.

Other plutonic and metaplutonic rocks of the Baker terrane

ERA-81 -- Granite, Oregon, 7.5-minute

quadrangle. Lat 44°52'11" N., 

long 118°29'51" W.

OX79-1 Oxman, Oregon, 7.5-minute

quadrangle. Lat 43°30'42" N., 

long 117°31'10" W.

BRC79-1 Lost Basin, Oregon, 7.5-minute 

quadrangle. Lat 44°34'47" N., 

long 117°33'16" W.

AM80-1 Aldrich Mountains Oregon, 15- 

minute quadrangle. Lat 45° 

23'48" N., long 119°23'09" W.

Tectonic block within 

Elkhorn Ridge 

Argillite.

Tectonic block of gabbro 

cut by silicic dikes 

within Elkhorn Ridge 

Argillite.

Small pluton cutting 

Burnt River Schist.

Tectonic block within 

serpentinite-matrix 

melange.

Metatrondhjemite. Plagioclase, quartz; 

secondary chlorite, epidote. Cataclastic 

texture.

Metatrondhjemite. Plagioclase, quartz, 

hornblende; secondary sericite, actinolite. 

Allotriomorphic-granular texture.

Metatrondhjemite. Plagioclase, quartz, 

titanite; secondary chlorite, epidote, 

saussurite, carbonate. Cataclastic texture.

Trondhjemite. Quartz, plagioclase, 

hornblende, titanite; secondary sericite, 

epidote, chlorite. Protoclastic texture.

Chert in the Elkhorn Ridge Argillite is pervasively 

recrystallized and, hence, radiolarians are not well 

preserved. Chert samples that contain identifiable radiolarian assemblages yield Permian and Late Triassic fauna (Vallier and others, 1977; Blome and 

others, 1986).

Middle to Late Devonian, Middle Pennsylvanian, 

and Early Permian fossils are reported from limestone pods (olistoliths?) within the Elkhorn Ridge Argillite (Bostwick and Koch, 1962; Brooks and others,

1976; Blome and others, 1986; Mullen-Morris and 

Wardlaw, 1986). Structurally separate Permian limestone blocks contain either Tethyan or North American fusulinids (Nestell, 1983). Some of the limestone 

blocks are recognized as allochthonous and hence at 

least part of the Elkhorn Ridge Argillite may be melange rather than semicoherent broken formation.

At several localities metaplutonic rocks occur as 

blocks within the chert-argillite matrix of the Elkhorn 

Ridge Argillite. These fragments range in composition
6. CANYON MOUNTAIN COMPLEX, SPARTA COMPLEX, AND RELATED ROCKS, NORTHEASTERN OREGON 253

TABLE 6.2. Zircon isotopic data for plutonic rocks of the Baker terrane, including data for the Canyon Mountain Complex and the Sparta

complex, Oregon

Sample Zircon fraction data

Amount

Properties analyzed 

(mg)

Concentrations2

206pb*

Pb isotopic composition3

Ages and uncertainty 

(Ma)

208pb/206pb 207pb/206pb 206pb/204pb 206pb*/238u 207pb*/235u 207pb*/206pb*

Plutonic rocks of the Baker terrane

OX79-1

ERA-81

BRC79-1

AM80-1

c, nm, 0° 

f, nm, 0°

c, nm, 0° 

c, m, 0°

c, nm, 0° 

c, m, 0°

bulk

18.6 

14.1

18.5 

32.5

15.1 

29.1

23.7

13.7 

14.1

6.04 

6.65

2.89 

3.06

9.02

416

428

180 

199

91.7 

96.2

236

0.0849 

.08562

.0869 

.0872

.0943 

.0879

.1971

0.05596 

.05422

.05220 

.05231

.05423 

.05238

.05365

3,171 

4,638

17,752 

15,360

4,422 

10,558

8,170

240+1 

241±1

245±1 

244±1

230±1 

233±1

279±1

242±2 

241±2

246±2 

245±2

231±2 

234±2

279±2

256+9 

244±6

258±4 

257±5

237+7 

240±5

279±6

Canyon Mountain Complex

CM79-1

CM79-3

c, nm, 3C 

c, m, 0°

bulk

25.5

15.6

17.8

16.6

17.4

452

473

1.68 44.0

0.2095

.2455

.2255

0.05464

.05451

.09323

4,832

5,224

360

268±1 

268±1

278±1

268±1 

268±2

282±2

269±4 

272±5

314±29

CM79-4 c, nm, 0°

c, m, 1°

c, m, 0°

18.4

38.2

18.0

1.99

2.35

2.58

52.5

62.2

67.5

.1184

.1318

.1224

.05501

.05409

.05467

4,631

6,833

5,197

276±1

276+1

276±1

278±2

276±2

27812

278±6

283±5

279±5

Sparta complex

SP-AG-81

SP79-1

c, nm, 0°

f, nm, 0°

c, m, 1°

c, nm, 0°

c, m, 0°

30.7

15.3

31.6

16.0

24.9

9.57

10.3

16.8

6.90

6.96

277

299

484

236

237

0.1180

.1199

.1243

.0832

.0918

0.05298

.05295

.05288

.05318

.05439

10,361

9,644

9,552

5,872

4,028

253±1

252±1

253±1

215+1

215+1

254±2

252±2

254±2

216±2

216±2

266±6

260±6

257±5

226±11

230+8

Abbreviations: c, coarse (>100 |im); f, fine (<100 (am); nm, nonmagnetic; m, magnetic. Number in degrees (°) is side tilt of Frantz magnetic separator 

operated at a magnet current of 1.6 amperes.

2Chemical processing similar to that described by Krogh (1973). Procedural Pb blanks are less than 0.3 ng. Asterisk (*) denotes radiogenic Pb corrected for 

blank and initial common Pb.

3Ratios are corrected for mass fractionation on the basis of replicate analyses of National Bureau of Standards Pb standards. Uncertainty (at the 2o level) in 

the measured ratio is less than ±0.2 percent for 208Pb/206Pb, less than ±0.15 percent for 207Pb/206Pb, and ranges from ±0.1 to ±11 percent for 206Pb/204Pb.

4Decay constants employed: X = 1.5513 x e~10 for 238U; X - 9.8485 x e~10 for 235U. Uncertainty in ages (stated at the 2o level) is based on combined 

uncertainty in sample isotopic ratio measurements, analytical reproducibility of National Bureau of Standards Pb and U standards, uncertainty in 

fractionation corrections, and uncertainty in the isotopic composition of common Pb.

from serpentinized ultramafic rocks through gabbro 

to trondhjemite. Two such fragments were sampled 

(fig. 6.3) for geochronologic investigation.

In the Oxman quadrangle, east of Baker City, 

Oreg., a block engulfed by the Elkhorn Ridge Argillite consists of uralitized gabbro cut by dikes of 

trondhjemite. Two zircon fractions separated from 

one dike (sample OX79-1) have yielded Early Triassic 

206pb*_238u ageg of 241 and 240 Ma (table 6.2).

In the Granite quadrangle (Brooks and others, 

1982) metamorphosed intrusive rocks appear to be

rootless fragments surrounded by Elkhorn Ridge 

Argillite. These metaigneous rocks consist of metamorphosed gabbro, diorite, quartz diorite, and tonalite. Two zircon fractions from tonalite (sample ERA81) within this area have 206pb*_238u ageg of 245 

and 244 Ma, indicating an Early Triassic crystallization age.

The crystallization ages for the plutonic rocks indicate that their emplacement was, at least in part, 

synchronous with deposition of the Elkhorn Ridge 

Argillite. Whether these plutons represent part of the
254 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

plutonic foundation upon which the Elkhorn Ridge 

Argillite was deposited or whether they are simply 

fragments of plutons that intruded the basin into

which sedimentation was occurring remains unclear. 

The zircon ages, however, constrain major deformation and metamorphism of the Elkhorn Ridge Argil1200 119° 118° 117° 116°

46C WASHINGTON 

OREGON

45C

44°

D 

ERA-81

EXPLANATION

Cretaceous to Jurassic plutons 

Wallowa terrane

Baker terrane Showing Elkhorn Ridge 

Argillite (dark gray)

Izee terrane

Olds Ferry terrane

Grindstone terrane

Contact

Thrust fault Sawteeth on upper plate

Sample locality

FIGURE 6.3. Areal distribution of the Elkhorn Ridge Argillite in the Baker terrane, showing U-Pb zircon sample localities.
6. CANYON MOUNTAIN COMPLEX, SPARTA COMPLEX, AND RELATED ROCKS, NORTHEASTERN OREGON 255

lite to post-240 Ma time. The formation is cut by 

nonmetamorphosed, undeformed Early Cretaceous 

plutons, the oldest of which yields U-Pb zircon ages

of 144 Ma (Walker, 1989) that establish an upper 

limit on the time of deformation and metamorphism 

of the Elkhorn Ridge Argillite.

120° 119° 118° 117° 116°

EXPLANATION

Cretaceous to Jurassic plutons 

Wallowa terrane

Baker terrane Showing Burnt River Schist 

(dark gray)

Izee terrane

Olds Ferry terrane

Grindstone terrane

Contact

Thrust fault Sawteeth on upper plate

BRC 79-1 Sample locality

FIGURE 6.4. Areal distribution of the Burnt River Schist in the Baker terrane, showing U-Pb zircon sample locality.
256 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

BURNT RIVER SCHIST

Another major supracrustal component of the 

Baker terrane is the Burnt River Schist (fig. 6.4). The 

Burnt River Schist is dominantly quartzose phyllite 

with lesser amounts of greenstone, marble, metavolcanic and metavolcaniclastic rocks, and rare metaconglomerate; these rocks have undergone greenschistto epidote-amphibolite-facies metamorphism (Gilluly, 

1937; Ashley, 1967; chap. 12, this volume). Middle to 

Late Triassic conodonts were recovered from marble 

beds (Mullen-Morris and Wardlaw, 1986) within the 

formation.

Several plutonic bodies that range in composition 

from gabbro to trondhjemite cut the protolith and 

were metamorphosed with it (Ashley, 1967; chap. 12, 

this volume). These plutons may represent exotic tectonic fragments introduced into the protolith of the 

Burnt River as olistostromal blocks rather than magmas that intruded the protolith. It should be noted 

however, that these plutons are areally extensive in 

outcrop (up to several square kilometers), and are 

the only volumetrically significant rock type enclosed 

by, but not considered to be part of, the formation.

A trondhjemite within the Burnt River Schist was 

sampled for geochronologic investigation (sample 

BRC79-1, fig. 6.4). Two zircon fractions from this pluton have Late Triassic 206pb*_238TJ &geg of 233 and

230 Ma. These ages set a maximum age of metamorphism of the protolith of the Burnt River Schist. 

Along its southern margin the Burnt River Schist is 

intruded by two undeformed plutons that have 

206pb*_238u ageg of approximately 124 Ma and 120

Ma (Walker, 1986, 1989). Thus, metamorphism and 

deformation of the formation occurred between 230 

and 124 Ma.

The Elkhorn Ridge Argillite and Burnt River 

Schist are lithologically similar, although the Burnt 

River Schist is more siliceous overall, more deformed, 

and metamorphosed to higher grades than the 

Elkhorn Ridge Argillite. On the basis of such observations, Gilluly (1937) proposed that the Burnt River 

is the older of the two formations. Ashley (1967), 

however, tentatively concluded that part of the Burnt 

River is correlative with the Elkhorn Ridge. Exotic or 

olistostromal components, which are present within 

the Elkhorn Ridge, have not been recognized in the 

Burnt River. In addition, the age of the protolith of 

the chert-argillite within the Elkhorn Ridge is Permian to Triassic whereas the age of deposition of the 

protolith to the Burnt River Schist is Triassic. It is 

likely therefore that the two formations represent 

different depositional environments, although they 

could be components of the same tectonic system.

The original tectonic environment of the protoliths is 

not well established. Coward (1982) suggested that 

the Elkhorn Ridge Argillite represents a deformed accretionary prism, whereas Bishop (chaps. 4 and 5, 

this volume; Mullen, 1985) has proposed a fore-arc 

origin for the entire Baker terrane. This problem is 

considered in more detail later in this chapter.

ULTRAMAFIC-MAFIC^SILICIC OPHIOLITIC IGNEOUS SUITES

Several structural blocks of igneous and metaigneous rocks, two of which are ultramafic-mafic-silicic 

suites characterized by partially intact ophiolitic 

stratigraphy, are important constituents of the Baker 

terrane. Geochronologic investigation of the two most 

complete suites is described below.

CANYON MOUNTAIN COMPLEX

The Canyon Mountain Complex is the largest and 

the most structurally and lithologically coherent metaigneous suite of the Baker terrane. Structurally, the 

complex is a slab immersed in an extensive serpentinite-matrix melange tract in the southwestern part of 

the terrane (fig. 6.5). The complex is in contact with 

serpentinite-matrix melange on its west and northwest margins and with Quaternary sediments along 

its north margin, and it is faulted against Tertiary 

volcanic rocks on its east side. The southern extension of the complex is obscured by an unconformable 

cover of Tertiary volcanic rocks.

From south to north, the Canyon Mountain Complex 

has been divided into the following major lithologic 

units, each of which has a general east-west trend and 

steep dips (fig. 6.6; Ave Lallemant, 1976; Thayer, 1977; 

Himmelberg and Loney, 1980; Gerlach and others, 

198la; Leeman and others, chap. 1, this volume): tectonite harzburgite; ultramafic and gabbroic metacumulates; gabbroic cumulates with minor amounts of 

wehrlite and pyroxenite and isotropic hornblende gabbro; and an uppermost zone of irregular trondhjemite 

intrusive masses, diabase and trondhjemite sills, and 

keratophyre and quartz keratophyre flows.

The major part of the ultramafic unit consists of 

penetratively deformed tectonite composed of harzburgite, with subordinate amounts of dunite and pyroxenite, and volumetrically minor chromitite. This 

unit has the chemical characteristics of a refractory 

residuum resulting from extraction of a basaltic melt 

(Himmelberg and Loney, 1980).

The harzburgite tectonite is overlain by metacumulates (Himmelberg and Loney, 1980) ultramafic and
6. CANYON MOUNTAIN COMPLEX, SPARTA COMPLEX, AND RELATED ROCKS, NORTHEASTERN OREGON 257

gabbroic rocks of cumulus origin but now overprinted 

with metamorphic textures and a folded, penetrative 

foliation and (or) lineation that are shared by the underlying tectonite harzburgite. The metacumulates

consist of complexly interlayered ultramafic and gabbroic rocks. Ultramafic rock types of this unit range 

from dunite through wehrlite to clinopyroxenite, 

according to their relative proportions of olivine and

120° 119° 118° 117° 116°

WASHINGTON 

OREGON

-CANYON 

MOUNTAIN 

COMPLEX

EXPLANATION

Cretaceous to Jurassic plutons 

Wallowa terrane

Baker terrane Showing Canyon Mountain

Complex and Sparta complex (dark gray areas)

Izee terrane 

Olds Ferry terrane 

Grindstone terrane 

Contact

Thrust fault Sawteeth on upper plate 

FIGURE 6.5. Areal distribution of the Canyon Mountain Complex and the (informal) Sparta complex of Phelps (1979).
258 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

clinopyroxene. Volumetrically minor varieties of the 

now-metamorphosed ultramafic rocks include plagioclase-bearing varieties of those rocks named above 

and orthopyroxenite, Iherzolite, and websterite. The 

metagabbroic rocks within this ophiolitic sequence 

formed chiefly from gabbro and gabbronorite and 

from minor amounts of troctolite, anorthosite, and 

quartz norite. The texture of the majority of the 

metacumulate rocks is allotriomorphic-granular, however equigranular-mosaic recrystallization textures 

are common as are deformation textures such as 

kink bands in olivine and orthopyroxene.

The metacumulate sequence is separated from 

overlying cumulate gabbroic rocks by a complex, interdigitating contact zone. The cumulate rocks are 

predominantly gabbro and gabbronorite with lenses 

of wehrlite and feldspathic wehrlite. Centimeter- to 

meter-scale layering, defined by cyclic variation in 

the pyroxene/plagioclase ratio, is ubiquitous in the 

gabbro yet absent in the enclosed ultramafic lenses. 

The cumulate sequence is, in part, overlain by a thin, 

irregular zone (not shown in fig. 6.6) of isotropic, 

hornblende-bearing gabbro.

The stratigraphically uppermost unit of the Canyon Mountain Complex consists of a complex zone of 

keratophyre and quartz keratophyre flows, irregular 

trondhjemite intrusive masses, and dikes and sills of 

diabase and trondhjemite. Gerlach and others

118° 58' 118° 43'

JOHN DAY fi

44° 

20'

CANYON TV

CREEK CM79-4

FAULT

2 KILOMETERS

EXPLANATION

Surficial deposits (Quaternary) 

Volcanic rocks (Tertiary)

Fields Creek Formation of the 

Aldrich Mountains Group

Melange 

Sill complex

Gabbroic and ultramafic cumulates

Gabbroic and ultramafic metacumulates

Harzburgite tectonite

Contact

Fault Dashed where approximate

D Sample locality

CM79-1

FIGURE 6.6. Simplified geology of the Canyon Mountain Complex (modified from Himmelberg 

and Loney, 1980, and Gerlach and others, 198 la), showing U-Pb zircon sample localities. Area 

of map is shown in figure 6.5.
6. CANYON MOUNTAIN COMPLEX, SPARTA COMPLEX, AND RELATED ROCKS, NORTHEASTERN OREGON 259

(1981b) interpreted this uppermost zone as representing a sill complex on the basis of the fact that 

the zone and lithologic contacts within it share the 

general east-west trend and steep dip of the underlying units.

The metamorphic textures and folded penetrative 

lineation and foliation developed in the basal peridotite 

and metacumulates were interpreted by Ave Lallemant 

(1976) as indicating sub-solidus flow after solidification 

of this part of the Canyon Mountain Complex but prior 

to the formation of the cumulates and sill complex, 

which lack such structures. Alternatively, as George 

(1978) suggested for the Troodos complex, the absence 

of such features in the gabbroic-cumulate sequence 

may be the result of deformation processes operative 

on a partially crystallized magma chamber. In his model, the metacumulate and cumulate sequences are essentially coeval. The undeformed cumulates represent 

a stratigraphically higher level of the chamber that 

remained largely liquid during deformation. Thus, if 

George's (1978) model is applicable to the Canyon 

Mountain Complex, previously crystallized magmatic 

rocks and the basal peridotite were overprinted by this 

deformation whereas crystallization of the nonmetamorphosed cumulates took place shortly after and, in 

part, contemporaneous with deformation.

Although the complex has been interpreted as an 

"ophiolite," sheeted mafic dikes, although reported 

from the complex (Thayer, 1977) are rare. There are 

domains of multiple dike and sill injection but they 

lack one-way chilling characteristic of sheeted complexes in ophiolites of midocean-ridge origin. In addition, as previously noted, the volcanic subunit of this 

complex is dominated by keratophyre and quartz keratophyre flows rather than mafic pillow lavas. Unlike 

most ophiolites of midocean-ridge origin, the Canyon 

Mountain Complex contains abundant orthopyroxene 

as a cumulus phase, suggesting an overall more siliceous magma composition than that of typical midocean-ridge basalt (MORE). These observations 

combined with mineral-chemistry data (Himmelberg 

and Loney, 1980; Gerlach and others, 1981a, 1981b; 

Leeman and others, chap. 1, this volume) have led to 

the interpretation that the Canyon Mountain Complex 

represents an arc-related magmatic complex. U-Pb geochronologic data presented herein indicate a multistage magmatic history spanning at least 8 m.y for 

this complex.

PREVIOUS GEOCHRONOLOGIC INVESTIGATIONS

Previously determined isotopic ages for the Canyon 

Mountain Complex include 250- to 240-Ma K-Ar ages

determined on hornblende from hornblende-rich pegmatite dikes cutting gabbro and 210 Ma from quartz 

diorite (Vallier and others, 1977). Ave Lallemant and 

others (1980) reported a 40Ar/39Ar age of 262 Ma 

from a hornblende-rich pegmatite.

U-Pb ZIRCON GEOCHRONOLOGIC RESULTS

Five samples were collected from the Canyon 

Mountain Complex for geochronologic investigation. 

Three yielded sufficient zircon for isotopic analysis. 

Sample localities, field occurrence, and petrographic 

descriptions are given in table 6.1, sample localities 

are shown in figure 6.6 and isotopic data are presented in table 6.2.

The oldest ages obtained for the Canyon Mountain 

Complex are from the metacumulate sequence. Sample 

CM79-3 was collected from a quartz-bearing gabbro 

within this sequence; its granoblastic-polygonal texture is evidence of the post-magmatic metamorphism 

imposed on the rock. Zircons of this sample are a deep 

red color and translucent to opaque. Due to low zircon 

yield, the bulk zircon fraction was analyzed and gave 

slightly discordant ages, which are interpreted to indicate a minimum age of the cumulate protolith. The 

206pb*_238u age of thig fraction is 278 Ma whereas the 

206pb*_207pb* age ig 314±29 Ma. The large uncertainty 

in the 206pb*_207pb* age resuiis from the combined 

effect of the very low 206Pb/204Pb ratio and the uncertainty in the 207Pb/204Pb ratio of common Pb. Interpretation of such discordant ages is difficult from a single 

fraction, particularly considering the large analytical 

uncertainty in the 206pb*_207pb* age In view of this

large analytical uncertainty, I interpret the 278 Ma 

206pb#_238|j age ag fae mjnjmum age for the protolith

of the metacumulate sequence and suggest that the 

protolith age is no older than 314 Ma.

The zircon concentrate of this sample has low U 

and Pb concentrations (44 and 1.68 ppm, respectively) 

that may be accounted for in terms of both the composition and thermal history of the rock from which the 

sample was taken. The sample is from a quartz-bearing norite that is interlayered on a centimeter scale 

with norite and anorthosite. These rocks are interpreted to be metamorphosed fractional-crystallization 

products of a basaltic melt (Himmelberg and Loney, 

1980); if so, they are naturally low in concentrations 

of large-ion-lithophile and incompatible elements 

such as U. As reviewed earlier, Ave Lallemant (1976) 

views the formation of the Canyon Mountain Complex as a two-stage process. In his model, the basalperidotite tectonite and metacumulate sequence represent a distinctly older phase of magmatism and
260 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

deformation and metamorphism than do the overlying cumulates, sill complex, and volcanic rocks. The 

age discordance observed in zircons from sample 

CM79-3 therefore could be attributed to loss of Pb 

from the zircons due to metamorphism and deformation, which overprinted the original magmatic textures in the metacumulate sequence. Himmelberg 

and Loney (1980) have determined subsolidus reequilibration temperatures, based on pyroxene geothermometry, of approximately 880 to 975°C for the 

basal peridotite, metacumulates, and cumulates. Certainly, prolonged exposure to such temperatures during slow cooling or re-heating to such temperatures is 

sufficient to cause accelerated ionic diffusion resulting 

in the loss of radiogenic Pb from the zircon crystal lattice. Mattinson (1978) has empirically deduced that 

the blocking temperature of the U-Pb geochronometer 

for geologically young zircons with low U concentrations is approximately 650°C. The 206pb*-238U age reported here for the metacumulate may thus represent 

the time at which the metacumulates cooled below 

the blocking temperature for zircon. The low concentration of Pb in zircons from this sample is attributed 

to the original low concentration of U in the melt that 

produced the zircon and the effects of subsolidus reequilibration that likely enhanced Pb diffusion. It is 

irresolvable from this single zircon fraction whether 

the observed discordance is a result of a dynamic 

metamorphic overprint, of Pb-loss related to sustained high subsolidus temperatures, or is a combination of the two processes.

Sample CM79-4 was collected from a coarsegrained trondhjemite intrusive mass that is part of 

the sill complex described by Gerlach and others 

(1981a; 1981b). This leucocratic mass, locally characterized by cataclastic domains, cuts gabbro in the 

southwestern part of the complex. Three zircon fractions, also characterized by low U concentrations, 

from this sample have yielded Early Permian 206Pb*- 

238U ages of 276 Ma. On the basis of field relations, 

the trondhjemite is clearly younger than the cumulates and metacumulates. However, the ages of 276 

Ma for zircons from this sample are similar to the 

206pb*_238u age for that of sample CM79-3; this suggests either that the sill complex was affected by sustained high temperatures during sub-solidus reequilibration or that the sills intruded at this time.

The youngest age measured from the Canyon 

Mountain Complex is that of sample CM79-1. This 

sample was collected from a medium-grained, hypidiomorphic-granular hornblende tonalite that cuts gabbro near the west margin of the complex. The 

outcrop area of this pluton is small (<5 km ). Zircons 

from this sample are characterized by distinctly higher U concentrations than those separated from samples CM79-3 and CM79-4 (table 6.2). Two zircon 

fractions from sample CM79-1 have yielded concordant 206pb*_238u ageg of 268 Ma These ages indicate

that magmatic activity in the Canyon Mountain 

Complex continued after emplacement of the sill 

complex.

The age range of magmatism within the Canyon 

Mountain Complex is inconsistent with geologically 

instantaneous generation of the complex at a midocean-ridge spreading center. Rather, the geochronologic data indicate two stages of Early Permian 

magmatism: formation of the metacumulate and cumulate section and emplacement of trondhjemite 276 

Ma (or earlier?) followed by intrusion of quartz diorite at 268 Ma.

SPARTA COMPLEX

The informally named Sparta complex, which crops 

out over a 70 km2 area, is included in the Wallowa 

terrane by Silberling and others (1984) but had been 

considered by previous investigators to be part of the 

Baker terrane and is therefore discussed in this chapter. From north to south, the complex comprises serpentinite, cumulus clinopyroxenite, cumulus gabbroic 

rocks, diorite-tonalite, and trondhjemite (fig. 6.7). The 

south margin of the complex is in fault contact with 

the multiply deformed Elkhorn Ridge Argillite whereas the north margin of the complex has an uncertain 

contact relationship with Permian and Triassic rocks 

of the supracrustal Clover Creek Greenstone.

The serpentinite, whose protolith is indeterminate 

because of extensive serpentinization, composes a relatively small part of the complex and is locally gradational into partly serpentinized clinopyroxenite 

(Phelps, 1979). The clinopyroxenite is interlayered 

with and overlain by cumulus gabbroic rocks (chiefly 

gabbronorite) which become increasingly hornblenderich near the contact with the diorite-tonalite.

The contact between the gabbroic rocks and the diorite-tonalite is marked by a zone of shearing and alteration. Locally, gabbroic xenoliths occur in the 

diorite-tonalite near the contact with the underlying 

gabbro. The contact appears intrusive at most localities although locally the contact seems gradational. 

Rock types within the diorite-tonalite unit grade 

from hornblende diorite to biotite-hornblende tonalite 

(Almy, 1977; Phelps 1979).

The contact between the diorite-tonalite and the 

trondhjemite is an indistinct zone characterized by 

complex contact relationships and further obscured 

by shearing and alteration. Thus the age of the dio-
6. CANYON MOUNTAIN COMPLEX, SPARTA COMPLEX, AND RELATED ROCKS, NORTHEASTERN OREGON 261

rite-tonalite unit relative to that of the trondhjemite 

is not certain from field relations. Texturally, the 

trondhjemite is mostly massive but includes porphyritic, granophyric, hypidiomorphic-granular, and allotriomorphic granular varieties, most of which have 

been locally overprinted with brittle tectonite fabrics. 

The origin of the trondhjemite has been the subject 

of several studies (Gilluly, 1933; Almy, 1977; Phelps, 

1979; Phelps and Ave Lallemant, 1980). Gilluly believed it to be the product of metasomatic origin, derived by Na-metasomatism of the diorite-tonalite, 

whereas subsequent workers considered it to be of 

magmatic origin. Given the variety of textures within 

the trondhjemite and the generally poor outcrop, it is 

probable that several intrusive phases are present 

within the area that Prostka (1962) mapped as 

trondhjemite.

The trondhjemite, the uppermost unit of the Sparta complex in this area, is overlain by the Clover 

Creek Greenstone along a poorly exposed contact. 

Prostka (1962) interpreted the contact to be depositional with the Clover Creek resting on the eroded 

surface of the trondhjemite. He based this interpretation on the occurrence of silicic plutonic clasts, which 

he believed to be derived from the underlying Sparta 

complex, in the basal part of the Clover Creek Greenstone. The Clover Creek Greenstone thickens to the 

northeast of the Sparta quadrangle, and therefore 

these clasts could have been derived from exhumed 

Permian plutons such as those exposed to the northeast in Hells Canyon and the Seven Devils Mountains. Locally within Hells Canyon, Middle Triassic 

rocks of the Seven Devils Group rest unconformably 

on Permian silicic plutons (Vallier, 1977); therefore, 

by Middle Triassic time, uplift and erosion had exhumed tonalitic-trondhjemitic plutons of the Wallowa 

terrane as potential contributors of clasts to the 

lower Clover Creek Greenstone. In an area west of 

the Sparta quadrangle, Brooks and others (1976) 

mapped trondhjemitic plutons intruding the Clover 

Creek Greenstone and Brooks and Vallier (1978) reported thermal alteration of the Clover Creek near 

trondhjemite of the Sparta complex. Almy (1977) suggested that keratophyre and quartz keratophyre of 

the Clover Creek Greenstone and trondhjemite of the 

Sparta complex are cogenetic on the basis of majorelement chemical data from these rocks. Such a conclusion further implies that the contact of the trondhjemite of the Sparta complex and the Clover Creek 

Greenstone is intrusive.

The Clover Creek Greenstone consists of interbedded spilite, keratophyre and quartz keratophyre 

flows, pyroclastic rocks, volcanogenic sandstones, and 

volcanic breccia with minor amounts of intercalated

argillite and limestone. The various rock types within 

the Clover Creek thicken and thin markedly along 

strike, and bedforms tend to be lensoid. The type locality of the Clover Creek Greenstone, originally described by Gilluly (1937), lies to the west of the 

Sparta complex. In the area around Sparta, Oreg., 

Prostka (1962) subdivided rocks correlative with the 

Clover Creek Greenstone into Gold Creek Greenstone 

and "Lower Sedimentary Series" and estimated a 

combined stratigraphic thickness of approximately 

1,800 m for these units. Gilluly (1937) originally assigned the Clover Creek Greenstone a Permian age; 

however, it was demonstrated by subsequent workers 

(Wetherill, 1960; Bostwick and Koch, 1982; Prostka, 

1962; Nolf, 1966) that the Clover Creek includes 

rocks as young as Late Triassic. The age of the Clover Creek in the Sparta area is uncertain but herein 

is considered to include rocks of Late Permian as 

well as Triassic age on the basis of the conformable 

contact relationship (Prostka, 1962) of the Clover 

Creek Greenstone with the Upper Triassic Martin 

Bridge Limestone.

The Sparta complex has been termed an "ophiolite" 

(Vallier and others, 1977; Phelps and Ave Lallemant, 

1980) although it is lithologically unlike most welldocumented ophiolites. As shown in figure 6.7, the 

Sparta complex includes a large volume of silicic 

rocks. In fact, unlike most ophiolites, the bulk of the 

exposed part of the complex consists of rocks of tonalitic through trondhjemitic composition. In addition, 

the Sparta complex lacks tectonite harzburgite, has 

an extremely thin exposed ultramafic component, 

and although numerous mafic and silicic dikes cut all 

lithologic units within the complex, no sheeted dike 

or sill complex has been recognized. The absence of a 

dike or sill complex strongly suggests that this igneous suite was not generated in a long-lived tensional 

regime such as that of an oceanic spreading center. 

The unconformable cover of Cenozoic volcanic rocks 

may mask the actual proportional distribution of the 

various rock types. Nevertheless, the stratigraphy 

and bulk composition of the Sparta complex is unlike 

those ophiolites, such as the Oman ophiolite and the 

Bay of Islands ophiolite, Newfoundland, inferred to 

have originated at some type of oceanic spreading 

center.

PREVIOUS GEOCHRONOLOGIC INVESTIGATIONS

Previous geochronologic investigation of the Sparta 

complex, limited to K-Ar and 40Ar/39Ar age determinations of the diorite-tonalite unit, yielded Late Triassic 

ages: 213 Ma K-Ar ages (reported in Vallier and others, 

1977); a 223 Ma 40Ar/39Ar plateau age and a 212 Ma
262 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

total-gas age have been measured from this unit (Ave 

Lallemant and others, 1980). Assuming that the diorite-tonalite unit is coeval with the entire complex, the 

entire suite had been assigned to the Late Triassic.

117°30'

EXPLANATION

Sparta complex of Phelps 

(1979) Consists of:

Trondhjemite

Diorite-tonalite

Gabbro

Pyroxenite

Serpentinite

I____| Volcanic rocks (Tertiary)

[+_+J Tonalite (Cretaceous to Jurassic)

Mi || Martin Bridge Limestone

[- ^ i_\ Clover Creek Greenstone

[jiji|jiljj|| Elkhorn Ridge Argillite

 - - Contact Dashed where approximate

 Fault

SP79-1 Sample locality

FIGURE 6.7. Geology of the (informal) Sparta complex of Phelps (1979), showing UPb zircon sample localities. Geology is modified from Prostka (1962). Area of map 

is shown in figure 6.5.
6. CANYON MOUNTAIN COMPLEX, SPARTA COMPLEX, AND RELATED ROCKS, NORTHEASTERN OREGON 263

However, as discussed herein, the geochronologic history of this igneous suite is somewhat more complex.

U-Pb ZIRCON GEOCHRONOLOGIC RESULTS

For geochronologic investigation, five samples were 

collected from three lithologic units within the Sparta 

complex two from the cumulus gabbro, one from the 

diorite-tonalite unit, and two from the trondhjemite. Of 

these samples, only those of the diorite-tonalite unit 

and one of the samples of trondhjemite yielded sufficient zircon for isotopic analysis. Zircon isotopic data 

for the complex are presented in table 6.2.

Sample SP79-1, a hornblende-biotite tonalite, was 

collected at J.N. Bishop Spring in the diorite-tonalite 

unit. Two zircon fractions from this sample have 

identical Late Triassic 206pb*_238u ageg of 215 M^

in good agreement with K-Ar and total-gas 40Ar/39Ar 

ages previously reported for this unit (Ave Lallemant 

and others, 1980). Inasmuch as the underlying gabbro is more altered than the diorite-tonalite and xenoliths of gabbro are locally present within the dioritetonalite, the gabbro is therefore older than the diorite-tonalite. Gabbro of this complex is therefore at 

least 215 Ma in age and may be considerably older.

The oldest zircon ages measured from the complex 

are for the trondhjemite. The sampled trondhjemite 

is porphyritic and comprises phenocrysts of sodic plagioclase and rounded, embayed quartz in a finegrained matrix of quartz, plagioclase, and minor 

hornblende, biotite, and titanite. The three analyzed 

zircon fractions from the trondhjemite yield Late Permian 206pb*_238u ages of 253 Ma. Local intrusive relationships between the trondhjemite and the 

overlying Clover Creek Greenstone require that at 

least part of the Clover Creek Greenstone in the 

Sparta quadrangle is at least as old as Late Permian.

Phelps (1979) suggested that the trondhjemite and 

diorite-tonalite are cogenetic on the basis of petrologic modeling of major- and trace-element data. The UPb geochronologic data presented in table 6.2 and 

previously discussed demonstrate that the trondhjemite and the diorite-tonalite represent two distinct magmatic events separated by about 38 m.y. 

and preclude any likelihood that the diorite-tonalite 

and trondhjemite are cogenetic.

In summary, the Sparta complex comprises Late Permian and Late Triassic plutonic rocks. The trondhjemite and the diorite-tonalite unit are not coeval and 

represent temporally and compositionally separate intrusive events. The temporal relationship of the gabbro 

to the trondhjemite is not clear in the absence of field 

evidence such as observable contact relationships. An 

argument against a genetic relationship between the

trondhjemite and gabbro is the tremendous volume of 

magma that would have had to fractionate in order to 

produce the large volume of trondhjemite relative to 

the gabbro. The Sparta complex is not the product of 

magmatic differentiation of a single parental magma. 

Rather, it represents at least two distinct magmatic 

events separated by some 38 m.y. Because in the Sparta 

quadrangle the trondhjemite of the complex locally intrudes the overlying Clover Creek Greenstone, part of 

the Clover Creek Greenstone is at least Late Permian.

SERPENTINITE-MATRIX MELANGE

Polymict serpentinite-matrix melange is a major 

component of the Baker terrane (fig. 6.8). Included 

tectonic blocks range from less than a meter to several kilometers in length. The blocks are compositionally as well as texturally diverse and range in degree 

of metamorphic grade from low greenschist to garnet 

amphibolite grade. Blocks of chert, greenstone (some 

showing pillow structure), fragmented ultramaficmafic-silicic plutonic suites (ophiolites?), chert-argillite, foliated amphibolite, gabbro, silicic metaplutonic 

rocks, mafic and ultramafic plutonic rocks, mafic volcanic rocks, tuffaceous volcanic rocks, limestone, and 

rare blueschist are present.

The range of published isotopic ages of the tectonic 

blocks is Early Permian to Late Triassic. Blueschist 

blocks in melange near Mitchell, Oreg. (Swanson, 

1969) have yielded a 40Ar/39Ar age of 223±3 Ma (Hotz 

and others, 1977) whereas amphibole from foliated amphibolite in serpentinite-matrix melange south of the 

Canyon Mountain Complex has a 40Ar/39Ar age of 

255±3 Ma (Ave Lallemant and others, 1980). Carpenter 

and Walker (1992) reported a hornblende K-Ar age of 

267 Ma for amphibolite in the Aldrich Mountains.

Early Permian fusulinids and Late Triassic radiolarians are reportedly found in melange blocks east 

and southwest of John Day, Oreg. (Dickinson and 

Thayer, 1978; Blome and others, 1986).

MELANGE IN THE ALDRICH MOUNTAINS

Serpentinite-matrix melange in the Aldrich Mountains is exposed approximately 25 km southwest of 

the town of John Day, Oreg. (Dickinson and Thayer, 

1978). The melange matrix, composed of foliated serpentinite, engulfs blocks and slabs of heterogeneous 

lithology. Low-grade metavolcanic rocks, chert, amphibolite, argillite, gabbro, and peridotite are the 

dominant rock types that compose the blocks. Rare 

blocks consisting of metatrondhjemite are also 

present. A single zircon fraction from one such block
264 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

(sample AM80-1, fig. 6.8) yielded a concordant age of 

279 Ma, which is analytically identical to the age of 

trondhjemite in the sill complex of the Canyon Mountain Complex. The block is therefore interpreted as a 

tectonic fragment of the same crust from which the 

Canyon Mountain Complex was derived.

120° 119° 118° 117° 116°

46<

AREA OF MAP

45° -

> Mitchell

John

44° _

WASHINGTON 

OREGON

100 KILOMETERS v' 

_I t

EXPLANATION

Cretaceous to Jurassic plutons 

Wallowa terrane

Baker terrane Showing serpentinite-matrix melange 

(dark gray)

Izee terrane

Olds Ferry terrane

Grindstone terrane

Contact

Thrust fault Sawteeth on upper plate

D AM80-1 Sample locality

FIGURE 6.8. Areal distribution of serpentinite-matrix melange in the Baker terrane, showing U-Pb zircon sample locality (in the Aldrich Mountains).
6. CANYON MOUNTAIN COMPLEX, SPARTA COMPLEX, AND RELATED ROCKS, NORTHEASTERN OREGON 265

TIMING OF MELANGE FORMATION

In the John Day, Oreg., area, melange of the Baker 

terrane, which includes the Canyon Mountain Complex as a large slab, is unconformably overlain by a 

sequence of Upper Triassic to Upper Jurassic clastic 

rocks (Dickinson and Vigrass, 1965; Dickinson and 

Thayer, 1978). Therefore melange formation in this 

area is constrained to pre-Late Triassic time. However, these overlying clastic strata are deformed, numerous unconformities occur within the sequence, 

and the lowermost rocks of this sequence contain detritus that was most likely derived from the melange. 

Although intense deformation had ceased by Late Triassic time, deformation evident in this area continued 

until Early Jurassic time. The time of initiation of 

breaking and mixing that led to the development of 

the melange is difficult to determine. Given the abundance of Permian and Triassic faunal and isotopic 

ages of the blocks, it seems reasonable to infer that 

melange formation began no earlier than the Permian. Whether melange formation was geologically instantaneous, continuous, or episodic is not known.

Serpentinite-matrix melange in the Greenhorn 

area, west of Baker City, Oreg., is interpreted by 

Ferns and others (1983a) to be depositionally overlain by a sedimentary unit comprised of argillite and 

sandstone with lesser amounts of conglomerate, 

chert, and limestone. A similar unit mapped by 

Mullen (1978) has yielded both Early Permian conodonts and Early Permian fiisulinids from limestone 

pods. As Ferns and Brooks (1983) point out, it is possible that the limestone pods represent reworked 

blocks from the underlying melange and therefore 

the depositional age of the unit could be younger 

than Early Permian. The minimum age of melange 

formation within the Elkhorn Ridge inlier is thus not 

well constrained.

PALEOTECTONIC SETTING OF THE BAKER TERRANE

Any interpretation regarding the paleotectonic setting for the Baker terrane must account for the following characteristics: (1) The overall character of the 

Baker terrane is that of a melange. (2) Rocks of different 

metamorphic grade, including lawsonite blueschist, are 

commonly juxtaposed. (3) Both Tethyan and non-Tethyan Permian fusulinids are present in structurally separate blocks. (4) Devonian conodonts and Pennsylvanian 

and Permian fusulinids are present in limestone olistoliths. (5) Serpentinite-matrix melange, chert-argillite 

broken formation, and siliceous phyllite are intermixed 

on a broad scale. (6) The age of plutonic rocks within the 

terrane ranges from about 279 to 215 Ma.

The Early Permian to Late Triassic isotopic age 

range coincides with that of sedimentation within the 

terrane as deduced from faunal ages for the Elkhorn 

Ridge Argillite and the Burnt River Schist. This observation suggests that the deposition of sediments 

and the formation of plutonic crust were largely synchronous processes. This age range is also very similar to the 264- to 225-Ma age range of plutonism 

within the Wallowa and Olds Ferry terranes (Walker, 

1986). These terranes have been interpreted as vestiges of arc-type ensimatic volcanic-plutonic complexes. The ophiolitic ultramanc-mafic-silicic plutonic 

suites (such as the Canyon Mountain Complex) are 

characterized by a large component of silicic rocks, 

by the abundance of cumulus orthopyroxene in their 

gabbroic sections, by the rarity or complete absence 

of sheeted dike complexes, and by their associated 

keratophyric and (or) tuffaceous volcanic rocks. Features such as these however, are not characteristic of 

those ophiolites inferred to have originated at some 

type of oceanic spreading center. Rather, the Canyon 

Mountain Complex and the Sparta complex have features more characteristic of those ophiolites believed 

to have an arc-related origin (see Hawkins and 

Evans, 1983). Given the geologic characteristics reviewed above, only the following paleotectonic environments can reasonably account for the lithologic 

and structural complexity of the Baker terrane: an 

accretionary complex, a marginal basin, or a fore arc.

Dickinson (1979) cited the overall melange-like character of the terrane, the presence of blueschist blocks, 

and faunal ages as evidence that this terrane marks 

the existence of a former Permian to Triassic subduction zone. The juxtaposition of metamorphic rocks and 

tectonic blocks composed of metamorphic rocks of various grade within the terrane strongly implies that 

some type of mixing process has occurred. The requisite kinematic environment to accomplish such mixing 

may be an accretionary prism associated with an active 

arc-type volcanic-plutonic complex. Karig (1979), however, suggested that in active ensimatic volcanic-plutonic arc systems such admixtures of rock types may 

also result from strike-slip faulting along trends that 

are subparallel to the active magmatic-arc crest. The 

presence of rocks of various metamorphic grade in a 

melange-like setting thus is not unique to a subduction 

complex. The presence of both Tethyan and non-Tethyan fusulinids implies juxtaposition of different paleobiogeographic settings, but is not in and of itself evidence 

of a subduction complex. The ophiolitic rocks of this 

terrane have structural and lithologic characteristics 

unlike those of midocean-ridge origin (see earlier discussion) and yield U-Pb zircon ages that are synchronous with the inferred age of subduction (Permian to
266 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Triassic). It is therefore highly unlikely that these represent obducted fragments of subducting oceanic crust. 

Furthermore, extensive tracts of serpentinite matrix 

melange, abundant in the Baker terrane, have not 

been reported from present-day ensimatic accretionary 

complexes.

Another possible paleotectonic environment for the 

Baker terrane is a back-arc setting between an active 

arc and remnant arc or a back-arc basin behind an 

active arc. The structural character, type of sediment, 

and rock types associated with such basins may be 

similar to those characteristic of a fore-arc setting, 

thus these environments can be difficult to distinguish 

in the rock record. Disorganized spreading in a backarc basin (Weissel, 1981) behind an active ensimatic 

arc may result in the formation of oceanic crust bearing 

numerous fracture zones within which serpentinite diapirs could rise. These serpentinite diapirs could then 

mix with the overlying sediments and thereby produce 

such rock admixtures as are found in the Baker terrane. The main objections, however, to a marginalbasin or back-arc basin setting for the Baker terrane 

are the occurrence of blocks of Triassic blueschist, juxtaposed Tethyan and non-Tethyan fauna, and blocks of 

Devonian and Pennsylvanian limestone in a Permian 

to Triassic back-arc basin.

There is little supporting evidence that the Baker 

terrane represents either an accretionary-prism complex or a back-arc basin. Rather, the constraints just 

outlined are met more reasonably by a fore-arc environment of an ensimatic volcanic-plutonic arc that was 

tectonically and magmatically active in Permian to 

Triassic time an idea first proposed by Mullen (1985).

Investigations of the fore-arc region of the active 

ensimatic Mariana arc in the southwestern Pacific 

demonstrate that it is composed of a great variety of 

volcanic, plutonic, sedimentary, and metamorphic 

rocks (Hussong and Uyeda, 1982; Bloomer, 1983). 

Structurally and morphologically, the Mariana fore-arc 

is characterized by numerous seamounts (of both volcanic and diapiric origin), horst-and-graben structures, 

and hummocky terrane as well as domains of relatively 

undisturbed volcanogenic sedimentary rocks. Lithologic components of the Mariana fore-arc reported by 

Bloomer and Hawkins (1982), Bloomer (1983), and 

Fryer and others (1985) include fragments of oceanic 

crust, serpentinite diapirs that have flowed laterally 

(some more than 10 km) upon subaqueous extrusion, 

alkalic basalts, siliceous tuffaceous sediments, chert, 

allochthonous shallow-water limestone of Cretaceous 

age, conglomerate, and breccia.

Some present-day fore arcs are characterized by extension-related normal faults, not uncommonly with

significant vertical displacement. In the Mariana fore 

arc, such faults apparently have acted as conduits for 

protrusion of serpentinite diapirs into overlying forearc sediments (Fryer and others, 1985). Depending 

upon the depth of generation, serpentinite diapirs 

could entrain blocks of differing metamorphic grade 

during their ascent towards the surface. Such a 

mechanism could account for the intimate mixture of 

chert-argillite and serpentinite-matrix melange with 

blocks of varying metamorphic grade observed in the 

Baker terrane. The presence of limestone bearing 

Middle Devonian and Pennsylvanian fossils and both 

Tethyan and non-Tethyan fusulinids may have resulted from subduction processes welding such material 

to an inner trench wall. Thus, mechanisms exist by 

which to incorporate much older material into the 

fore-arc region of active arc systems.

Magmatism in fore-arc regions has been recognized in a number of volcanic-plutonic arc complexes 

(Gill, 1981), and ophiolitic assemblages can be generated in fore-arc settings (Leitch, 1984). In such a setting the timing of ophiolitic generation is similar to 

the age of magmatism in the arc with which the 

ophiolite is associated. Most igneous crystallization 

ages of ophiolite sequences and other plutonic components of the Baker terrane are synchronous with 

magmatic crystallization ages in the Wallowa and 

Olds Ferry terranes (Walker, 1986). The Wallowa 

and Olds Ferry terranes have been interpreted as 

arc-type ensimatic volcanic-plutonic arc-type complexes formed during Permian to Triassic subduction. Thus, it is possible that plutonic components of 

the Baker terrane are vestiges of Permian to Triassic 

fore-arc plutonism.

The Baker terrane is herein considered to represent 

the fore-arc region of an ensimatic arc-type volcanicplutonic complex whose other components are represented by the Permian to Triassic Wallowa terrane 

and Triassic volcanic and plutonic rocks of the Olds 

Ferry terrane. The overall melange-like character of 

the Baker terrane may be the normal expression of 

tectonically active forearc environments (see fig. 1, 

Fryer and others, 1985). The Baker terrane, however, 

has been affected by post-Triassic structural modification. In the John Day, Oreg., area, rocks of the Baker 

terrane have been folded and faulted with overlying 

Triassic to Cretaceous volcanic and volcaniclastic 

rocks of the Izee Terrane. The Baker terrane has also 

been structurally telescoped against the Olds Ferry 

terrane along the Connor Creek fault whose age is 

bracketed as Late Jurassic to Early Cretaceous 

(Walker, 1986). Numerous Tertiary normal faults cut 

the terrane as well.
6. CANYON MOUNTAIN COMPLEX, SPARTA COMPLEX, AND RELATED ROCKS, NORTHEASTERN OREGON 267

SUMMARY AND CONCLUSIONS

(1) Two metaplutonic fragments within the Elkhorn 

Ridge Argillite have crystallization ages of approximately 244 Ma and 240 Ma. These ages place lower 

limits on the age of deformation and metamorphism 

of the Elkhorn Ridge.

(2) A metamorphosed intrusive body in the Burnt 

River Schist has an igneous-crystallization age of approximately 233 to 230 Ma. This age represents a 

minimum age of deposition and maximum age of 

metamorphism for the Burnt River Schist.

(3) The Canyon Mountain Complex, of Early Permian age, had a protracted magmatic history spanning at 

least 8 m.y. The age of formation of the metacumulate 

sequence and cumulate sequence is at least 278 Ma, 

whereas the age of the sill complex is 276 Ma. A younger pluton intruded the complex at 268 Ma. The temporally protracted igneous history and lithologic 

characteristics are consistent with a volcanic-plutonicarc origin for the Canyon Mountain Complex.

(4) The Sparta complex comprises rocks of Late Permian as well as Late Triassic age. The main trondhjemite unit yields U-Pb zircon ages of 253 Ma, whereas 

zircons from the diorite-tonalite unit are 215 Ma. The 

Sparta complex is the result of at least two temporally 

unrelated magmatic events. The complex lacks sheeted 

dikes or sheeted sills and is characterized by a large 

volume of silicic plutonic rocks, and these characteristics seem to preclude its origin in a midocean-ridge 

environment. Rather, it likely represents the plutonic 

core of a volcanic-plutonic ensimatic arc-type complex.

(5) The age of the Clover Creek Greenstone in the 

Sparta, Oreg., area is no younger than Late Permian 

on the basis of the Late Permian age for the crosscutting trondhjemite of the Sparta complex.

(6) A metatrondhjemite block in serpentinite-matrix melange in the Aldrich Mountains yielded a concordant U-Pb age of 279 Ma. Thus it appears to 

represent a tectonically broken fragment of crust genetically related to the Canyon Mountain Complex.

(7) The age range of magmatism recorded by U-Pb 

zircon ages of plutonic rocks in the Baker terrane 

overlap and are nearly synchronous with U-Pb ages 

of plutons in the Wallowa and Olds Ferry terranes.

(8) The Baker terrane probably represents the 

structurally telescoped fore arc of a Permian to Triassic ensimatic volcanic-plutonic arc complex.

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v. 53, p. 255-265.

 1981b, Petrology and geochemistry of plagiogranite in the 

Canyon Mountain ophiolite, Oregon: Contributions to Mineralogy and Petrology, v. 77, p. 88-92.

Gill, J.B., 1981, Orogenic andesites and plate tectonics: New York, 

Springer-Verlag, 390 p.

Gilluly, J., 1933, Replacement origin of the albite granite near 

Sparta, Oregon: U.S. Geological Survey Professional Paper, 

175-C, p. 65-81.

 1937, Geology and mineral resources of the Baker quadrangle, Oregon: U.S. Geological Survey Bulletin 879, 119 p.

Hawkins, J.W., and Evans, C.A., 1983, Geology of the Zimbales 

Range, Luzon, Philippine Islands Ophiolite derived from an 

island arc-back arc basin pair, in Pt. 2 of Hayes, D.E., ed., The 

tectonic and geologic evolution of Southeast Asian seas and 

islands: Geophysical Monograph, v. 27, p. 95-123.

Hillhouse, J.W., Gromme, C.S., and Vallier, T.L., 1982, Paleomagnetism and Mesozoic tectonics of the Seven Devils arc in 

northeastern Oregon: Journal of Geophysical Research, v. 87, 

p. 3777-3794.

Himmelberg, G.R., and Loney, R.A., 1980, Petrology of ultramafic 

and gabbroic rocks of the Canyon Mountain ophiolite, Oregon: 

American Journal of Science, v. 280, p. 232-268.

Hotz, P.E., Lanphere, M.A., and Swanson, D.A., 1977, Triassic 

blueschists from northern California and north-central Oregon: Geology, v. 5, p. 659-663.

Hussong, D.M., and Uyeda, S., 1982, Tectonic processes and the 

history of the Mariana arc A synthesis of the results of the 

Deep-Sea Drilling Project, Leg 60, in Hussong, D.M., and Uyeda, S., eds., Initial reports of the Deep-Sea Drilling Project, 

Leg 60.

Karig, D.E., 1979, Material transport within accretionary prisms 

and the "knocker" problem: Journal of Geology, v. 88, p. 27-39.

Krogh, T.E., 1973, A low contamination method for hydrothermal 

decomposition of zircon and extraction of U and Pb for isotopic 

age determination: Geochimica et Cosmochimica Acta, v. 37, 

p. 485-494.

Leitch, E.G., 1984, Island arc elements and arc-related ophiolites: 

Tectonophysics, v. 106, p. 177-203.

Mattinson, J.M., 1978, Age, origin, and thermal histories of some 

plutonic rocks from the Salinian block of California: Contributions to Mineralogy and Petrology, v. 67, p. 233-245.

Mullen, E.D., 1978, Geology of the Greenhorn Mountains, northeastern Oregon: Corvallis, Oregon State University, M.S. thesis, 372 p.

Mullen, E.D., 1985, Petrologic character of Permian and Triassic 

greenstones from the melange terrane of northeastern Oregon 

and their implications for terrane origin: Geology, v. 13, 

p. 131-134.

Mullen-Morris, E., and Wardlaw, B.R., 1986, Conodont ages for 

limestones of eastern Oregon and their implication for preTertiary melange terranes, in Vallier, T.L., and Brooks, H.C., 

eds., Geology of the Blue Mountains Regions of Oregon, Idaho, 

and Washington Geologic implications of Paleozoic and Mesozoic paleontology and biostratigraphy, Blue Mountains province, Oregon and Idaho: U.S. Geological Survey Professional 

Paper 1435, p. 59-63.

Nestell, M.K, 1983, Permian foraminiferal faunas of central and 

eastern Oregon [abs.]: Geological Society of America Abstracts 

with Programs, v. 45, p. 65-68.

Nolf, B., 1966, Structure and stratigraphy of part of the northern 

Wallowa mountains, Oregon: Princeton, Princeton University, 

Ph.D. dissertation, 138 p.

Phelps, D., 1979, Petrology, geochemistry, and origin of the Sparta 

quartz diorite-trondhjemite complex, northeastern Oregon, in 

Barker, F. ed., Trondhjemites, dacites and related rocks: Amsterdam, Elsevier, p. 547-580.

Phelps, D., and Ave Lallemant, H.G., 1980, The Sparta ophiolite 

complex, northeast Oregon A plutonic equivalent to low-^O 

island-arc volcanism: American Journal of Science, v. 280-A, 

p. 345-358.

Prostka, H.J., 1962, Geology of the Sparta quadrangle: Oregon Department of Geology and Mineral Industries Geologic Map Series Map GMS-1, scale 1:62,500.

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 84-523, p. C1-C43.

Swanson, D.A., 1969, Lawsonite blueschists from north-central Oregon, in Geological Survey Research 1969: U.S. Geological 

Survey Professional Paper 650-B, p. 8-11.

Taubeneck, W.H., 1957, Geology of the Elkhorn Mountains, northeastern Oregon Bald Mountain batholith: Geological Society 

of America Bulletin, v. 68, p. 181-238.

 1958, Wallowa batholith, Wallowa Mountains, northeastern 

Oregon [abs.]: Geological Society of America Bulletin, v. 69, 

p. 1650.

 1964, Cornucopia stock, Wallowa Mountains, northeastern 

Oregon Field relationships: Geological Society of America 

Bulletin, v. 75, p. 1093-1116.

 1967, Petrology of the Cornucopia tonalite unit, Cornucopia 

stock, Wallowa Mountains, northeastern Oregon: Geological 

Society of America Special Paper 91, 56 p. 

-1969, Granitic clasts in Pennsylvanian conglomerates of

central Oregon [abs.]: Geological Society of America Abstracts 

with Programs, Cordilleran Section, v. 1, p. 67-68.

Thayer, T.P., 1977, The Canyon Mountain Complex, Oregon, and 

some problems of ophiolites, in Coleman, R.G., and Erwin, 

W.P., eds., North American ophiolites: Oregon Department of 

Geology and Mineral Industries Bulletin 95, p. 93-106.

Thayer, T.P, and Brown, C.E., 1966, Geologic map of the Aldrich 

Mountain quadrangle, Grant County, Oregon: U.S. Geological 

Survey Geologic Quadrangle Map GQ-438.
6. CANYON MOUNTAIN COMPLEX, SPARTA COMPLEX, AND RELATED ROCKS, NORTHEASTERN OREGON 269

Vallier, T.L., 1977, The Permian and Triassic Seven Devils Group, 

western Idaho and northeastern Oregon: U.S. Geological Survey Bulletin 1437, 58 p.

Vallier, T.L., Brooks, H.C., and Thayer, T.P., 1977, Paleozoic rocks 

of eastern Oregon and western Idaho, in Stewart, J.H., 

Stevens, C.H., and Fritsche, A.E., eds., Paleozoic paleogeography of the Western United States (Pacific Coast Paleogeography Symposium 1): Los Angeles, Society of Economic 

Paleontologists and Mineralogists, Pacific Section, p. 455-465.

Walker, N.W, 1986, U-Pb geochronologic and petrologic studies of 

the Blue Mountains terrane, northeastern Oregon and westcentral Idaho Implications for pre-Tertiary tectonic evolution: 

Santa Barbara, University of California, Ph.D. dissertation, 

224 p.

 1989, Early Cretaceous initiation of post-tectonic plutonism 

and the age of the Connor Creek fault, northeastern Oregon 

[abs.]: Geological Society of America Abstracts with Programs, 

v. 21, no. 5, p. 155.

Weissel, J.K., 1981, Magnetic lineations in marginal basins of the 

western Pacific: Philosophical Transactions of the Royal Society of London, ser. A, v. 300, p. 223-247.

Wetherill, C.E., 1960, Geology of part of the southern Wallowa 

Mountains, northeastern Oregon: Eugene, University of Oregon, M.S. thesis, 208 p.

Wilson, D., and Cox, A., 1980, Paleomagnetic evidence for tectonic 

rotation of Jurassic plutons in the Blue Mountains, eastern 

Oregon: Journal of Geophysical Research, v. 85, no. B7, 

p. 3681-3689.

7. PRE-CRETACEOUS TECTONIC EVOLUTION OF 

THE BLUE MOUNTAINS PROVINCE, NORTHEASTERN OREGON

By HANS G. AVE LALLEMANT*

CONTENTS

Page

Abstract- - . . . ._...... ...._.-..... .-.. 271

Introduction 272

Terranes of the Blue Mountains province 274

Wallowa terrane 274

Baker terrane __________________________ 275

Olds Ferry terrane 275

Izee terrane 276

Early Cretaceous plutons 276

Mid-Cretaceous to Tertiary 276

Structure 277

Megascopic analysis 277

Mesoscopic analysis 277

Microscopic analysis 277

terrane ________________________ 283

deformation - 283

Mesoscopic analysis 283

Microscopic analysis 283

Interpretation 285

Timing _________________________ 285

deformation 285

D

D

Mesoscopic analysis 

Microscopic analysis

Timing

287

287

292

D3 deformation 292

D4 deformation 292

D, deformation - 293

Mesoscopic analysis 293

Microscopic analysis 293

Interpretation 294

Timing 294

D2 deformation 295

Mesoscopic analysis 295

Interpretation 295

Timintr _________________________ 295

D3 deformation 295

Izee terrane - - 295

Department of Geology and Geophysics, Rice University, Houston, Texas 77251

Structure Continued

Izee terrane Continued

D! deformation Continued 

Mesoscopic analysis 

Microscopic analysis 

Interpretation 

Timing

.......... 295

.......... 295

. ... 296

.... -.. 296

D2 deformation 296 

Mesoscopic analysis 296 

Microscopic analysis 296 

Interpretation 296 

Timing 297

D3 deformation 297 

Olds Ferry terrane 297

Dj deformation 297 

D2 deformation 299

J_/3 Q.6101*1X1 clvlOIl ~ __---___ . -__ _ _____ _____________ «_. .___ _Sl/t7

Left-oblique plate convergence 299 

Post-D3 deformation 300

Conclusions 300

rifarl l*ll*Cvl -- .... ... . _ .. _ .. __ ... ___ _____ 301 uvx

ABSTRACT

The Blue Mountains province of northeastern Oregon has a 

complex tectonic history. The province consists of four major eastwest-trending belts or terranes, which from north to south are the 

Wallowa, Baker, Izee, and Olds Ferry terranes; the small Grindstone terrane in the southwestern part of the province is not discussed here. On the basis of mesoscopic and microscopic structural 

analysis and previously published geologic, geochronologic, geochemical, and paleomagnetic data, a coherent although not yet 

completely constrained story emerges about the tectonic evolution 

of the province. The Olds Ferry terrane was a north-south-trending, west-facing intraoceanic volcanic island arc, active possibly 

from Permian but certainly from Middle Triassic to Middle Jurassic time. The Izee terrane consists of intra-arc (including fore-arc) 

basin deposits related to the Olds Ferry arc. The Baker terrane 

consists in part of melange; structures in the melange indicate 

that the terrane was accreted by eastward subduction underneath 

the Olds Ferry volcanic arc. There are indications that the convergence between the Baker oceanic terrane and the Olds Ferry arc 

was left-oblique. The Wallowa terrane constitutes a volcanic arc 

that was active during Pennsylvanian to Late Triassic time, with 

minor volcanism occurring until Jurassic time. The polarity of the 

Wallowa arc is not constrained, but it may have been west-facing. 

Major mylonite zones indicate that the plate convergence related

271
272 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

to the Wallowa arc was also left-oblique. The Wallowa arc, mostly 

inactive since the Late Triassic, collided with the Baker, Izee, and 

Olds Ferry terranes in the Late Jurassic; again, the convergence 

may have been left-oblique. During the collision, rocks of the Baker and Izee terranes were thrust eastward across the Olds Ferry 

terrane, whereas in the Wallowa terrane the Jurassic assemblage 

and fragments of the Permian and Triassic arc were thrust westward. After the amalgamation of all terranes, the Blue Mountains 

were intruded by granitic plutons during the Early Cretaceous. 

The collision of this amalgamated terrane with the North American craton in late Early Cretaceous time cannot be documented in 

Oregon, but it has been recognized in Idaho.

INTRODUCTION

The Blue Mountains province of northeastern Oregon consists of deformed pre-Cretaceous rocks, undeformed Cretaceous plutons, and Upper Cretaceous 

and Cenozoic sedimentary and volcanic rocks. In this 

chapter only the pre-Cretaceous tectonic history is 

described. Other names for the Blue Mountains province, as used here, are the Blue Mountains region or, 

for short, Blue Mountains.

All igneous rocks of the Blue Mountains province 

have initial Sr-isotope ratios lower than 0.704 (Armstrong and others, 1977; Fleck and Criss, 1985; Criss 

and Fleck, 1987). Thus, they are unrelated to the 

North American craton. Consequently, the entire 

Blue Mountains region can be called a "suspect" terrane (Coney and others, 1980) and may have traveled far before juxtaposition to the North American 

craton during Cretaceous time (Sutter and others, 

1984; Lund and Snee, 1988; Snee and others, chap. 

15, this volume). Paleomagnetic evidence suggests 

that, relative to the site of this juxtaposition, the 

Blue Mountains province was farther to the north in 

Permian time (Harbert and others, 1988) and farther 

to the south in Triassic time (Hillhouse and others, 

1982).

The Blue Mountains region was subdivided into 

several terranes by Dickinson (1979), in part on the 

basis of previous work (Vallier and others, 1977; 

Brooks and Vallier, 1978; Dickinson and Thayer, 

1978), each terrane having a distinct stratigraphy 

and tectonic evolution. The terranes are elongate, approximately east-west-trending belts. Boundaries between terranes are always tectonic; they may be 

compressional boundaries related to subduction processes or strike-slip faulting, or they may have characteristics of both. Silberling and others (1984) 

somewhat changed the boundaries of the terranes in 

the Blue Mountains and renamed them (from north 

to south) the "Wallowa," "Baker," "Izee," and "Olds 

Ferry" terranes (fig. 7.1). They also named a new terrane, the "Grindstone" terrane, which is outside the

map area of figure 7.1 and about 75 km southwest of 

the town of John Day, Oreg.

The goal of the present study was to establish the 

tectonic evolution of each terrane by carrying out 

mesoscopic structural and strain analyses and microscopic strain and kinematic analyses. Structures 

common to two adjacent terranes may indicate the 

mechanism of juxtaposition of these terranes 

whether they collided by normal convergence, oblique 

convergence, or strike-slip movement. Studies of recent deformation related to oblique plate convergence 

in New Zealand (Walcott, 1978) and Sumatra (Beck, 

1983) indicate that the component of convergence 

normal to the plate boundary causes thrust faults 

and folds parallel to the boundary to form in the accretionary wedge; the component of convergence parallel to the plate boundary causes strike-slip 

displacements along fault zones parallel to the 

boundary in the accretionary wedge or the volcanic 

arc.

Because of the large areal extent of the Blue 

Mountains region, the results of the present study 

are clearly preliminary and the hypotheses tentative. 

The hypothesis proposed by Dickinson (1979) that 

the Baker terrane is the accretionary wedge or fore 

arc of the Olds Ferry volcanic-arc terrane is consistent with the structural data. The convergence may 

have been left-oblique; it may have started in the 

Permian and continued in the Late Triassic to Early 

Jurassic. The amalgamation of all five terranes of the 

Blue Mountains region took place when the Wallowa 

terrane collided with the amalgamated Baker-IzeeOlds Ferry terranes during Late Jurassic time. During this collision west-directed thrusting occurred in 

the Wallowa terrane (north-northwest-directed in 

present-day coordinates) and east-directed thrusting 

occurred in the Baker and Izee terranes (southsoutheast-directed in present-day coordinates). It is 

possible that this convergence also had a left-lateral, 

strike-slip component consistent with left-oblique 

plate convergence elsewhere in the North American 

Cordilleran orogenic belt (Oldow and others, 1984, 

1989; Ave Lallemant and Oldow, 1988). Collision of 

the Blue Mountains province with the North American craton occurred during the Cretaceous (Sutter 

and others, 1984; Lund and Snee, 1988; Snee and 

others, chap. 15, this volume), but no pervasive 

structures related to this collision have been identified in Oregon. Paleomagnetic evidence (Wilson and 

Cox, 1980; Hillhouse and others, 1982) indicates that 

the entire Blue Mountains province has rotated 

clockwise by about 65° since the intrusion of the Early Cretaceous granitic plutons. Part of the rotation 

may have occurred prior to or during the collision;
46'

120 C 119 C 118° 

45'

WALLOWA TERRANE

WEAKER TERRANE

IZEE TERRANE.

FIGURE 7.1. Index map of Blue Mountains province, showing localities and areas mentioned in text. See figure 7.2 for explanation of m
274 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

part of it has been attributed to the post-Eocene extension of the Basin and Range province (Magill and 

Cox, 1981).

ACKNOWLEDGMENTS

I am grateful to Tom Thayer and Howard Brooks 

who introduced me to the geology of the Blue Mountains. Reviews of earlier versions of this paper by H.C. 

Brooks, J.G. Evans, R.L. Thiessen, T.L. Vallier, and 

B.P. Wernicke were very helpful. I thank Gordon and 

Lura Glass of John Day, who have been so hospitable 

and helpful through the years that I worked in Oregon. The research was primarily supported by the National Science Foundation through grants GA-25894, 

EAR-76-13372, EAR-79-04829, and EAR-82-06531.

TERRANES OF THE BLUE MOUNTAINS PROVINCE

WALLOWA TERRANE

The Wallowa terrane, as named by Silberling and 

others (1984), is identical to the Seven Devils terrane 

of Dickinson (1979). It underlies most of the Wallowa 

Mountains in Oregon and the Seven Devils Mountains in Idaho (fig. 7.1). The Wallowa terrane consists 

of Permian and Triassic intrusive, volcanic, and volcaniclastic rocks; Triassic and Jurassic sedimentary 

rocks; and Early Cretaceous plutonic rocks.

The oldest rocks of the volcanic sequence are 

undated quartz keratophyre flows and tuffs overlain 

by Lower Permian keratophyre and spilite flows interbedded with volcaniclastic rocks, and Middle to 

Upper Triassic (Ladinian, Carnian) andesite, keratophyre, and spilite flows interbedded with clastic 

rocks (Vallier, 1977). The volcanic rocks follow a lowK2O basalt-andesite-dacite-rhyolite differentiation 

trend, typical of volcanic arcs (Vallier and Batiza, 

1978). The main volcanic activity ceased in Carnian 

time (Vallier and others, 1977); sporadic volcanism 

continued into the Callovian (Folio, 1994; White and 

Vallier, 1994).

The Permian and Triassic volcanic sequence is 

overlain by a thick platform limestone of late Carnian to Norian age (Vallier, 1977; Newton, 1983, 1986; 

Folio, 1994). It is overlain by Upper Triassic (Norian) and Lower Jurassic (Toarcian) shale, thin limestone, argillite, and siltstone of the Hurwal 

Formation (Nolf, 1966). There is no sedimentary 

record between the Toarcian and Callovian. Callovian 

to lower Oxfordian tuff and clastic beds of the Coon 

Hollow Formation occur in two areas of the Snake 

River canyon, directly overlying Triassic and older

rocks (Morrison, 1964; Vallier, 1977; Goldstrand, 

1994; White and Vallier, 1994).

In the Snake River canyon are three igneous-metamorphic complexes (Vallier, 1974) consisting mostly 

of gabbro, amphibolite, plagiogranite, and schistose 

metavolcanic rocks, which Vallier and others (1977) 

considered to be the basement upon which the Permian and Triassic volcanic sequences were deposited. 

Walker (1981, 1982, 1986) obtained U-Pb zircon ages 

of 265 to 233 Ma for plagiogranites from all three 

complexes, indicating that the plagiogranites are 

mostly synchronous with the volcanic activity. A discordant age of 309 Ma was obtained for zircons from 

a gneiss (Walker, 1986). Geochemical data are consistent with a volcanic-arc environment for the formation of these complexes (Balcer, 1980; Walker, 1982; 

Vallier, chap. 3, this volume). Major mylonite zones 

that occur in all three basement complexes of the 

Snake River canyon formed between 262 and 218 Ma 

(Ave Lallemant and others, 1980; Balcer, 1980; Walker, 1986).

The (informal) Sparta complex of Phelps (1979), 

about 40 km east of Baker City, consists of pyroxenite, gabbro, and plagiogranite. Phelps (1979) and 

Phelps and Ave Lallemant (1980) have shown that 

this complex has volcanic-arc affinities and probably 

represents the root of the Wallowa volcanic arc, because there is no obvious tectonic boundary between 

the Sparta complex and the Triassic volcanic rocks of 

the Seven Devils Group (see also Brooks and Vallier, 

1978). The plagiogranites were dated at 223 Ma 

(40Ar/39Ar: Ave Lallemant and others, 1980), and 

262, 223, and 219 Ma (U-Pb: Walker, 1981, 1983, 

1986); a gabbro sample yielded an age of 227 Ma 

(40Ar/39Ar: Ave Lallemant, 1984).

Many plate tectonic settings have been proposed 

for the Wallowa volcanic-arc terrane. Vallier and others (1977) and Brooks and Vallier (1978) correlated 

the Wallowa terrane with the Olds Ferry terrane. Silberling (1983) disagreed on the basis of stratigraphic 

and lithologic differences. Jones and others (1977) 

correlated the Wallowa terrane with the Wrangellia 

terrane, an exotic terrane in southeastern Alaska, 

Queen Charlotte Islands, and Vancouver Island, on 

the basis of paleomagnetic and faunal evidence. Newton (1983) agreed with the faunal correlation, but 

Stanley (1986) saw too many faunal differences. 

Sarewitz (1983) and Scheffler (1983) did not accept 

this correlation either, because of the large geochemical differences between the two terrane s: the Upper 

Triassic volcanic rocks of the Wrangellia terrane are 

oceanic-plateau-type basalts, and those of the Wallowa terrane are calc-alkaline (Vallier and Batiza, 

1978). Mortimer (1986) correlated the Wallowa ter-
7. PRE-CRETACEOUS TECTONIC EVOLUTION OF THE BLUE MOUNTAINS PROVINCE, NORTHEASTERN OREGON 275

rane with the Stikinia terrane in British Columbia 

(Monger and others, 1982; Monger and Berg, 1987). 

The Stikinia terrane consists of Mississippian to Jurassic arc-related volcanic and sedimentary rocks. It 

lies west of the Cache Creek assemblage, which is 

very similar in lithology and age to the Baker terrane of the Blue Mountains.

Paleomagnetic data (Harbert and others, 1988) 

suggest that during the Permian the Wallowa terrane resided at higher latitudes than the site of its 

ultimate accretion to the North American craton, and 

that in the Late Triassic (Hillhouse and others, 1982) 

it resided at lower latitudes. Thus, the Wallowa terrane may have migrated southward in post-Permian 

time, whereas it moved northward in post-Triassic 

time. This displacement history is in agreement with 

a model proposed by Ave Lallemant and Oldow 

(1988) that was based on field relationships in several Cordilleran terranes and on plate-motion reconstructions by Engebretson and others (1985). Ave 

Lallemant and Oldow (1988) suggested that several 

Cordilleran terranes, including the Wallowa terrane, 

moved southward along the continental margin until 

about 120 Ma, and northward from about 100 Ma to 

the present.

BAKER TERRANE

The Baker terrane lies south of the Wallowa terrane. It was previously called the Oceanic terrane 

(Vallier and others, 1977) and the Central Melange 

terrane (Dickinson and Thayer, 1978). The Central 

Melange terrane included Silberling and others' 

(1984) Grindstone terrane, an allochthonous sequence of Devonian to Lower Triassic sedimentary 

rocks about 75 km southwest of John Day (Blome 

and Nestell, 1991). The Grindstone terrane is not included in the present study.

The Baker terrane consists of large blocks that are 

internally coherent, separated by chaotic zones and 

shear zones. The main rock types are argillite and 

chert (of the Elkhorn Ridge Argillite), phyllite, limestone, serpentinite, gabbro, and subordinate volcanic 

flows and tuffs, graywacke, and conglomerate, and 

rare blueschists.

The ages of formation of rocks in the Baker terrane 

range from Devonian to Late Triassic or possibly 

Early Jurassic (Evans, 1986). Radiolarians in the 

cherts have been dated as Permian and Triassic and 

possibly Early Jurassic (Blome and others, 1986). 

Limestone blocks in the Baker terrane have Permian 

fusulinid faunas in some places only of Tethyan affinities, in others only of North American affinities

(Nestell, 1983). Devonian through Late Triassic (Morris and Wardlaw, 1986) conodonts are present in 

limestone blocks throughout the Baker terrane.

The volcanic rocks of the Baker terrane have been 

studied in detail (Mullen, 1985; Bishop, chap. 5, this 

volume). Some have MORE characteristics; others 

are intraplate basalts. Some basalts of volcanic island arc affinities are also present.

Blueschists near Mitchell (fig. 7.1; see also fig. 7.2, 

area J) were dated by Hotz and others (1977). They 

yielded an age of 223 Ma (40Ar/39Ar). Amphibolites 

with blue amphiboles have also been described in the 

Vinegar Hill area (K, fig. 7.2; Mullen, 1978; Bishop, 

chap. 4, this volume).

Dickinson (1979) suggested that this terrane is a 

subduction complex related to the Triassic volcanic 

arc of the Olds Ferry terrane in the southern part of 

the Blue Mountains region (fig. 7.1). Structural data 

presented here support this hypothesis.

The Baker terrane has been correlated invariably 

with the Cache Creek assemblage in British Columbia (Klepacki and Wernicke, 1985; Mortimer, 1986; 

Monger and Berg, 1987), with the Hayfork, North 

Fork, and Fort Jones terranes in the Klamath Mountains, and with the Calaveras Formation in the foothills of the Sierra Nevada (Burchfiel and Davis, 

1981; Mortimer, 1986). All of these rocks are lithologically very similar and are approximately of the same 

age as those of the Baker terrane.

OLDS FERRY TERRANE

The most southerly terrane in the study area is the 

Olds Ferry terrane, previously called the Juniper 

Mountain-Cuddy Mountain terrane by Brooks and 

Vallier (1978) and the Huntington terrane by Dickinson (1979). Silberling and others (1984) included the 

Weatherby Formation of Brooks (1979a) in the Olds 

Ferry terrane; the Weatherby consists of Lower Jurassic sedimentary rocks overlying the Triassic volcanic and volcaniclastic rocks. Dickinson (1979), 

however, put these Jurassic rocks in the Mesozoic 

clastic terrane, and in the present study, as well as 

in Vallier (chap. 3, this volume), his boundaries are 

used. The name Olds Ferry terrane is used here only 

for the Triassic volcanic and volcaniclastic rocks.

The Olds Ferry terrane is relatively small and 

somewhat poorly exposed. Only Middle and Upper 

Triassic (Carnian to Norian) rocks have been recognized. They consist of basalt, andesite, rhyolite, volcaniclastic rocks, and limestone (Brooks and Vallier, 

1978; Vallier, chap. 3, this volume). It is generally 

accepted that these rocks formed in a volcanic-arc
276 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

environment. Quartz diorite and gabbro stocks occur 

throughout the terrane (Brooks and others, 1976). 

One pluton in Oregon has been dated at 235 Ma 

(U-Pb: Walker, 1986). Several plutons in the Cuddy 

Mountains, Idaho, were dated at 217 to 187 Ma 

(K-Ar: Henricksen, 1975). They are believed to have 

formed in an intraoceanic volcanic arc (Brooks and 

Vallier, 1978). Although no volcanic rocks younger 

than Norian are exposed, Dickinson (1979) proposed 

that the arc was active through the Middle Jurassic, 

assuming that the Callovian lavas and tuffs in the 

western part of the Izee terrane were derived from 

the Olds Ferry terrane.

In most studies, the Canyon Mountain (ophiolite) 

Complex near John Day is included in the Baker or 

Central Melange terrane. Thayer (1969) believed it to 

be a large fragment of oceanic crust. Ave Lallemant 

(1976) suggested that it had formed in a marginal 

basin adjacent to a volcanic island arc. On the basis 

of geochemical studies, Himmelberg and Loney 

(1980), Gerlach, Ave Lallemant, and Leeman (1981), 

Gerlach, Leeman, and Ave Lallemant (1981), and 

Mullen (1983) proposed that the complex formed in a 

volcanic island arc. The age of the Canyon Mountain 

Complex is Permian (U-Pb data: Walker and Mattinson, 1980; Walker, 1981, 1986; 40Ar/39Ar data: Ave 

Lallemant and others, 1980; Ave Lallemant, 1984); it 

is plausible that the complex represents the Permian 

part of the Olds Ferry volcanic island arc terrane.

Hillhouse and others (1982) have obtained paleolatitudes for the Olds Ferry terrane that are the 

same as those for the Wallowa. Brooks and Vallier 

(1978) suggested that the two terranes were fragments of the same arc. However, because the Norian 

limestone typical of the Wallowa terrane is absent 

from the Olds Ferry terrane and because the abundance of Jurassic volcanic rocks in the Izee terrane 

suggests that the Olds Ferry arc was active through 

the Middle Jurassic, Silberling (1983) concluded that 

the Olds Ferry terrane should not be correlated with 

the Wallowa terrane. Mortimer (1986) correlated the 

Olds Ferry terrane with the Quesnellia terrane, a 

Trias sic to Jurassic volcanic arc terrane in British 

Columbia (Monger and Berg, 1987), on the basis of 

lithologic and age relations.

IZEE TERRANE

The Izee terrane of Silberling and others (1984) 

does not include the Jurassic sequence near the 

Snake River canyon because they assumed that the 

Weatherby Formation was deposited directly on the 

Olds Ferry terrane. Although these Jurassic sedimentary rocks most probably are related to the Olds Ferry volcanic-arc terrane (Brooks and Vallier, 1978), the 

contact between the two sequences is tectonic (Ave 

Lallemant, 1983). Furthermore, the lithologies and 

age of the Weatherby Formation are very similar to 

those of the Mesozoic clastic sedimentary rocks near 

John Day. Thus, these rocks are herein included in 

the Izee terrane.

In the John Day area the sedimentary rocks are of 

Late Triassic (Carnian) to Middle Jurassic (Callovian) age (Dickinson and Thayer, 1978); they are 

flysch-like turbidites consisting of shales and 

graywackes with chert, serpentinite, and volcanicrock particles and crystal fragments of quartz, calcite, plagioclase, and spinel. Volcanic flows are ubiquitous in the lower part of the sedimentary section 

and become subordinate near the top. In the Snake 

River canyon, sedimentary rocks of similar lithology 

occur, but they are restricted to the Early and Middle 

Jurassic (Sinemurian to Bajocian), although there 

are Upper Jurassic (Callovian) sedimentary rocks 

across the Snake River in Idaho (Brooks, 1979a).

Dickinson and Thayer (1978) and Dickinson (1979) 

suggested that most sedimentary rocks of the Izee 

terrane were deposited in intra-arc basins and that 

the youngest Callovian sediments were deposited in 

a deep fore-arc basin of the Huntington or Olds Ferry 

volcanic arc.

EARLY CRETACEOUS PLUTONS

All four terranes of the Blue Mountains region that 

are discussed here were intruded post-tectonically by 

mostly granitic stocks and plutons. They were dated 

by Armstrong and others (1977) using the K-Ar and 

Rb-Sr methods as Late Jurassic to Early Cretaceous 

(160 to 95 Ma). Most of these plutonic rocks, however, have been thermally disturbed. Recent U-Pb studies (Walker, 1989) suggest that the oldest intrusions 

are of earliest Cretaceous age (-143 Ma); the main 

plutonic activity appears to have ceased at about 120 

Ma (Walker, 1986, 1989). Some of the intrusions 

straddle the contact between terranes (Brooks and 

others, 1976). Paleomagnetic data (Wilson and Cox, 

1980; Hillhouse and others, 1982) indicate that the 

granites were emplaced at the latitude of present-day 

Idaho but before the 65° clockwise rotation occurred.

MID-CRETACEOUS TO TERTIARY

In the western part of the Blue Mountains, Cretaceous (Albian, Cenomanian) sedimentary rocks have
7. PRE-CRETACEOUS TECTONIC EVOLUTION OF THE BLUE MOUNTAINS PROVINCE, NORTHEASTERN OREGON 277

been recognized; they consist of shallow-marine 

coarse-grained clastic rocks (Dickinson and Thayer, 

1978). The remainder of the region is covered by volcanic and clastic rocks of Eocene to Pleistocene age, 

including the Miocene Columbia River Basalt Group.

STRUCTURE

In this chapter, structures are classified as megascopic, mesoscopic, and microscopic. Megascopic 

structures relate to the map scale; mesoscopic structures have a scale ranging from hand specimen to 

outcrop; and microscopic structures are sometimes 

recognized with the hand lens but generally only 

with the petrographic microscope.

MEGASCOPIC ANALYSIS

Megascopic folds were recognized in the Wallowa 

terrane by Smedes (1959), Wetherell (1960), and Vallier (1974) and in the Izee terrane by Dickinson and 

Vigrass (1965), Brown and Thayer (1966a, b), Thayer 

and Brown (1966), and W.D. Lowry and C.F. Wray 

(unpub. data, 1968). Because of the apparent chaotic 

character of the rocks, only a few megascopic folds 

were recognized in the Baker terrane (Gilluly, 1937; 

Ashley, 1966); but recently, many more megascopic 

folds were mapped on the basis of structural analysis 

of asymmetric mesoscopic folds in the Elkhorn Ridge 

Argillite (Coward, 1983). On the same basis, megascopic folds were recognized in the Weatherby area of 

the Izee terrane (Ave Lallemant, 1983). Map-scale 

thrust faults were observed in the Wallowa terrane 

by Smedes (1959) and Vallier (1974) and in the Izee 

terrane by Brooks and others (1976) and Ave Lallemant (1983).

MESOSCOPIC ANALYSIS

Observed mesoscopic structures include bedding 

planes (S0), fold axial planes and axial-plane cleavages 

or foliations (Si to S4), fold axes (Bi to B4), and lineations (Li to L4) (fig. 7.2). The subscripts 1 to 4 refer to 

deformation phases DI to D4, from the oldest to the 

youngest. Fold geometries are described according to 

the classification of Ramsay (1967). Lineations are intersection lineations (generally intersections of bedding and cleavage planes, the subscript of L 

corresponding to that of the cleavage involved), mineral lineations, and stretching lineations. In stretchedpebble conglomerates, the mean orientations of the 

three principal axes were determined. The mean short,

intermediate, and long dimensions are approximately 

parallel, respectively, to the major principal compressive strain axis (Z), the intermediate principal strain 

axis (Y), and the major principal extensile strain axis 

(X), but exceptions are observed.

Mylonites occur in several places in the Blue 

Mountains province. Along fine-grained mylonitic 

shear zones (Sm, fig. 7.2) the cleavage or X-Y plane is 

subparallel to Sm, but with increasing distance from 

Sm the strain magnitude decreases and the angle between Sm and the X-Y plane increases. Such shear 

zones are called "S-C mylonites" by Berthe and others (1979). The sense of shear that caused the mylonite can be determined: The displacement is parallel 

to the projection of the X axis onto the Sm plane and 

toward the acute angle between Sm and X-Y planes.

Extension veins were recorded in several areas. 

The major extension axis is subperpendicular to 

these veins.

In each outcrop, the relative age of different structures was determined by crosscutting relationships. 

Locally, such as in the Canyon Mountain Complex 

(Ave Lallemant, 1976; Ave Lallemant and others, 

1980) and in one of the basement complexes in the 

Snake River canyon (Walker, 1986), the oldest penetrative deformation occurred in the Permian. (These 

penetrative structures are not discussed in this chapter). In all other places, the first deformation (Dj) 

took place during Late Triassic to Early Jurassic 

time; the other three deformations (D2 to D4) appear 

to have occurred during the Late Jurassic.

All mesoscopic data are plotted on lower-hemisphere, equal-area stereographic projections (fig. 7.3), 

simplified versions of which are shown on the general 

map of the Blue Mountains region in figure 7.2. Axial 

trends of the four deformational events are also plotted in figure 7.2; these trends are based mainly on 

mesoscopic analysis but also on published maps by 

Gilluly (1937), Smedes (1959), Dickinson and Vigrass 

(1965), Brown and Thayer (1966a, b), Thayer and 

Brown (1966), Walker (1977), and Brooks (1979a).

MICROSCOPIC ANALYSIS

Several rock samples, oriented in the field, were 

thin-sectioned for microfabric and microtextural analysis. The word "microfabric" is used here to describe 

the preferred orientation of crystal lattice or opticalindicatrix axes, whereas the term "microtexture" is 

used to describe the spatial relationships of grains or 

the network of grain boundaries.

For the microfabric analysis, a four-axis universal 

stage was used to measure the c axes of quartz and
46C

45 c

120°

25

119C

50 KILOMETERS

118C 

WASHINGTON

AREA OR 

MAP

IDAHO

BAKER TERRANE

WALLOWA TERRANE

IZEE TERRANE

FIGURE 7.2. Simplified geologic map of Blue Mountains province (from Walker, 1977). Lower-hemisphere, equal-area projections of averag

elements characteristic of 22 areas on the map. Mesoscopic-structure data that correspond to these 22 stereographic diagrams are shown 

on figure 7.3.
EXPLANATION

MAP SYMBOLS LOWER-HEMISPHERE EQUAL-AREA PROJECTION 

Granite ijj Sm Mylonitic shear plane Arrow

/x~"^V~Sm relative movement 

Metasedimentary and metavolcanic rocks f ^^

^g Ophiolitic rocks \ -/ S0 Bedding plane

^>- Shear zones Arrows show relative N SQ 

"^ movement Sg^x^T^x . D

/> ~^ 2 ^'~4 ^ ax'a' P'0116 anc'

..-- Terrane boundary \A l~-l\.~ cleavage for deformatio

x^^vy ^i n

V^/J^L B 4

^ Trace of S, n r- ,^ , . *

B|_4 Fold axis for deformation

^ " ' Trace of S2 , n /^W_j - t0 °4 L4 ,b4-^--(^

Trace of S3 f / 'y"S4 . ^ -' i L|.4 Lmeation for deformation

to D4

ax

t!!f£

^B.S=^

3

1 * << >^- 

HI- i-3. JT

B-8 §-- ? > 

s- E § p ff

S^^cl

.t^iR-s 

l??|l

-agLB-S-'S 

o | ^3 BI P S:B.§'

i CD M J5 M ' '

 *I*|I3 

ffil 3 ff

» oq* » ^ P" 

rt- crq Q- H< a

^ M CD t3 3

P ~ S B 5"

!» ^ S 5 CD

" Ohrt & ^

fpll£

3 HJ &^ 5* 

?' CD' S- ct ?
280 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

WALLOWA TERRANE

N N N

A

FIGURE 7.3. Lower-hemisphere, equal-area projections of 

mesoscopic structural elements. Wallowa terrane: A, 

Pendleton inliers; B, Northern Wallowa Mountains; C, 

Confluence of Imnaha and Snake Rivers; D, Pittsburg 

Landing area; E, Cougar Creek area; F, Central Wallowa 

Mountains; G, Hells Canyon; H, Oxbow Dam area; I, 

Southern Wallowa Mountains. Baker terrane: J, Mitchell 

area; K, Vinegar Hill area; L, North Fork John Day River 

area; M, Elkhorn Ridge; N, Burnt River canyon; O, 

Brownlee Reservoir area. Izee terrane: P, Izee area; Q, Aldrich Mountains; R, Wickiup area; S, John Day area; T, 

Mine Ridge; U, Weatherby area. Olds Ferry terrane: V,

B,, L3

Huntington area. See figure 7.2 for localities corresponding 

to A through V. SQ, bedding plane; S, fold axial plane and 

axial-plane cleavage and foliation; B, fold axis; L, lineation. 

Subscripts 1 to 4 refer to phase of deformation Dj to D4. Sm, 

mylonitic shear plane (arrows show relative movement; 

dashed great circle indicates rarer orientation); a, tectonictransport direction; circle, pole to bedding plane; dot, pole to 

fold axial plane and axial-plane cleavage and foliation; solid 

triangle, fold axis; open triangle, lineation; diamond, tectonic-transport direction constructed from S and Sm; square, 

pole to mylonitic shear plane. Contour (only in /) drawn at 

1.7 percent per 1-percent area.
7. PRE-CRETACEOUS TECTONIC EVOLUTION OF THE BLUE MOUNTAINS PROVINCE, NORTHEASTERN OREGON 281

Poles to {0001} of biotite form great-circle girdles 

perpendicular to a strong mineral lineation (A, B, 

and G, fig. 7.5) and a point maximum perpendicular 

to the foliation. Because of basal slip in biotite, the 

{0001} planes are expected to rotate during deformation into parallelism with the foliation or X-Y plane.

In syntectonic-recrystallization experiments Tullis 

(1976) demonstrated that micas form [0001] girdles 

perpendicular to the X axis with a point maximum 

parallel to the Z axis.

Preferred orientations of amphibole (both glaucophane and green amphibole) are shown in projections

BAKER TERRANE

J

L

So, S, S3 , B3 , L3 

FIGURE 7.3. Continued.

, BZ, La
282 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

IZEE TERRANE

T

P

OLDS 

PER RY 

TERRANE

V

FIGURE 7.3. Continued.
7. PRE-CRETACEOUS TECTONIC EVOLUTION OF THE BLUE MOUNTAINS PROVINCE, NORTHEASTERN OREGON 283

D and F in figure 7.4 and C in figure 7.5. The a axes 

form a point maximum perpendicular to the foliation 

and a y point maximum parallel to a well-developed 

mineral lineation. Experimental work (Dollinger and 

Blacic, 1975; Nielsen and Ross, 1979) indicates that 

the favored slip system of amphiboles is (100)[001]. 

Accordingly, it is expected that during deformation 

the (100) planes rotate toward parallelism with the 

X-Y plane and the [001] axes rotate toward the X axis. The amphibole fabrics in the present study are 

consistent with these experimental conclusions.

Two plagioclase fabrics were determined here (A 

and C, fig. 7.5). Few experimental studies on plagioclase have been performed, but none of the experimentally determined slip systems are consistent with 

the present results. These two plagioclase fabrics can 

be related to the strain axes derived from the fabrics 

of biotite (A, fig. 7.5) and of quartz and amphibole (C, 

fig. 7.5). If indeed translation glide planes rotate toward the X-Y plane and the slip direction rotates toward the X axis, the required slip system for the 

plagioclase is (001)[100].

Only two of the microfabrics have monoclinic symmetry (B, M, fig. 7.4), and both are quartz c axis fabrics. These fabrics can be interpreted in terms of 

strain axes (see above) and in terms of kinematics. 

Studies have shown (Bouchez and Pecher, 1981) that 

a hypothetical shear plane can be constructed subperpendicular to the strongest of the two c axis girdles; 

the sense of shear or the displacement is then toward 

the acute angle between the shear plane and the X 

axis. Passchier (1983) warned, though, that this construction sometimes yields results opposite to the 

shear sense derived from other structural features.

All fabric diagrams in figures 7.4 and 7.5, except B 

and M in figure 7.4, can be considered orthorhombic. 

That does not necessarily mean that they formed by 

coaxial deformation. Orthorhombic fabrics have been 

produced in noncoaxial deformation experiments (see 

Tullis, 1977), but the sense of shear cannot be determined from the microfabric data. However, microtextural studies (figs. 7.6, 7.7) of thin sections cut parallel 

to the X-Z plane of samples shown in figures 7.4 and 

7.5 and of several additional samples revealed the 

sense of displacement. The sense of shear can be determined from asymmetric pressure shadows and rotated 

porphyroblasts (Simpson and Schmid, 1983), curved 

quartz and calcite fibers in veins and pressure shadows (Ramsay and Huber, 1983), rotated chlorite flakes 

and quartz fibers in pressure shadows of magnetite 

(Spry, 1969), microscopic S-C structures (Berthe and 

others, 1979), and displacements along microfaults.

Directions of tectonic transport derived from these 

microfabric and microtextural studies and mesoscopic

data are shown in figures 7.8 and 7.9. All geometric 

data are plotted on lower-hemisphere, equal-area 

stereographic projections.

WALLOWA TERRANE

D] DEFORMATION

The oldest penetrative deformation (Dj) structures 

in the Wallowa terrane occur in three so-called basement complexes exposed in the Snake River canyon 

(fig. 7.1; C, E, and H, fig. 7.2). These complexes consist of gabbro, quartz diorite, and volcanic rocks and 

their metamorphic and deformed equivalents. The 

grade of metamorphism is variable, from lowest 

greenschist to epidote amphibolite facies.

Both the plutonic and metavolcanic rocks are penetratively deformed in mylonitic shear zones. The Imnaha shear zone (C, fig. 7.2) at the confluence of the 

Imnaha and Snake Rivers, trends east-west and dips 

steeply to the north. The Cougar Creek shear zone 

(E, fig. 7.2) near Pittsburg Landing, Idaho, and the 

Oxbow shear zone (H, fig. 7.2) near the Oxbow Dam, 

Oregon and Idaho, strike northeast-southwest and 

dip steeply to the southeast.

MESOSCOPIC ANALYSIS

Each of three complexes is characterized by a multitude of mesoscopic shears subparallel to the megascopic shear zone. Fine-grained ultramylonites occur 

along these shears. With increasing distance to the 

shears, the grain size of the mylonites increases, the 

aspect ratio of porphyroclasts decreases, and the angle between the foliation planes and the shears increases. Thus, these structures are S-C mylonites 

(Berthe and others, 1979). The aspect ratio of porphyroclasts is the same as the finite strain ratio X/Z 

(Ave Lallemant and others, 1985). In general, mineral lineations are subhorizontal. The general trend of 

foliations is more easterly to southeasterly than the 

strikes of the shears (C, E, H, fig. 7.3). This relationship indicates that the shears are left-lateral. Rarely, 

northwest-southeast- and east-west-striking shears 

have been recognized in the Imnaha and Cougar 

Creek shear zones, respectively (C, E, fig. 7.3). The 

foliation/shear zone relationship in these two cases is 

opposite to the general relationship just described 

and indicates right-lateral displacement.

MICROSCOPIC ANALYSIS

Ave Lallemant and others (1985) obtained five 

quartz c axis microfabrics from quartz diorite mylonite
284 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

/?= 200 c, quartz

/?= 200 c, quartz n - 50 a, glaucophane /7=50 0, glaucophane n=50 7 , glaucophane

F

" = 50 a,amphibole /7=50 /3, amphibole /?= 50 7, amphibole

c, quartz /?= 200 c, quartz
7. PRE-CRETACEOUS TECTONIC EVOLUTION OF THE BLUE MOUNTAINS PROVINCE, NORTHEASTERN OREGON 285

samples from the Oxbow shear zone, and Chen (1985) 

obtained a large number of microfabric data from 

quartz diorite and granodiorite mylonites from the 

Imnaha shear zone: 12 quartz, 2 amphibole, and 3 

plagioclase fabrics. In the present study, only one 

sample of quartz diorite mylonite from the Cougar 

Creek shear zone was analyzed; it has a very strong 

quartz c axis preferred orientation (B, fig. 7.4). This 

fabric and those from the other shear zones (Ave Lallemant and others, 1985; Chen, 1985) display crossgirdle patterns. Many of the fabrics have monoclinic 

symmetry. Microscopic S-C mylonite structures are 

common (B, fig. 7.6).

The amphibole fabrics of Chen (1985) are very well 

defined. They consist of an a point maximum perpendicular to the foliation, a y point maximum subparallel to the lineation, and a P maximum perpendicular 

to the a and y maxima. Some of the fabrics have 

monoclinic symmetry; the y point maxima are not exactly parallel to the lineation.

The oligoclase microfabrics of Chen (1985) are also 

strongly defined. They consist of a y point maximum 

normal to the foliation, an a maximum parallel to the 

lineation, and a P maximum perpendicular to a and y.

INTERPRETATION

The orientation of the principal finite strain axes 

can be derived from the quartz and amphibole microfabrics. Comparison of these strain axes and the plagioclase microfabric indicates that the plagioclase in 

the study area has deformed by translation glide on 

the (010)[001] slip system.

The orientations of the strain axes X, Y, and Z for 

all three shear zones are shown in projections A

FIGURE 7.4. Lower-hemisphere, equal-area projections of Djdeformation microfabric elements of samples from selected areas shown in figure 7.8: B, Quartz diorite mylonite; 

D, blueschist; F, epidote amphibolite; G, marble; /, L, 

phyllites; M, metachert. Microfabric elements are c axes 

of quartz and calcite, and a, (3, and y (optical-indicatrix) 

axes of glaucophane and green amphibole. n, number of 

measurements. Contours drawn at intervals of 2 percent 

per 1-percent area where rc=50, and at intervals of 1 percent per 1-percent area where n >200. Dashed and heavy 

(solid) contour lines represent 2- and 4-percent contours, 

respectively, where n=50; heavy contour is 1-percent contour where n>200. Great circles with dots show geographic horizontal, where N, S, E, and W represent 

north, south, east, and west, respectively. Sj^ represents 

Si cleavage and foliation; L± (solid diamond) represents 

first (Dj-phase) lineation; Sm represents mylonitic shear 

plane (arrows show relative movement).

through C on figure 7.8. Projections A and C represent average orientations, and B was derived from 

projection B on figure 7.4.

The monoclinicity of some of the microfabrics in the 

Wallowa terrane samples and the mesoscopic relationship between foliations and shear planes indicate that 

all three shear zones were formed by left-lateral 

strike-slip motion. The rare shears oriented northwest-southeast and east-west in the Imnaha and Cougar Creek shear zones are right-lateral (C, E, fig. 7.3) 

and are conjugate to the major shear zones.

Ave Lallemant and others (1985) estimated that 

the minimum displacement along the Oxbow shear 

zone is 65 km. Because the other complexes are similarly deformed, the total left-lateral displacement 

along all three shear zones combined could easily 

have been 200 km or more.

TIMING

The timing of the shear displacement in the Imnaha, Cougar Creek, and Oxbow areas is rather well 

constrained. Several 40Ar/39Ar analyses date the timing of the metamorphism that was coeval with the 

mylonitization at 263 to 213 Ma (Ave Lallemant and 

others, 1980; Balcer, 1980; Walker, 1986). An undeformed lamprophyre crosscutting the mylonites of the 

Oxbow shear zone has a total 40Ar/39Ar gas age of 

197.4 Ma (Ave Lallemant and others, 1985).

D2 DEFORMATION

After the Late Triassic Dj deformation in the three 

basement complexes, the next major penetrative deformation in the Wallowa terrane occurred in the 

Late Jurassic. The associated mesoscopic deformation 

structures (fold axial planes, foliations, cleavages, 

and lineations) are plotted in figure 7.3 (projections 

A, B, D, F, G, I); D2 microfabrics are shown in figure 

7.5; and megascopic structural data are shown in figure 7.2 (projections A-/).

In some areas (such as Hells Canyon, G, fig. 7.2) 

the deformation is rather mild with only the development of crude cleavages subparallel to fold axial 

planes; in these areas, rocks have been metamorphosed only to lowest greenschist facies. In other areas (northern Wallowa Mountains, B, fig. 7.2, and the 

Pendleton inliers, A, fig. 7.2) the deformation is severe and the metamorphic grade is much higher. In 

one of the Pendleton inliers the mica schists contain 

sillimanite, and they are partly migmatized.

Northwest-vergent thrust faults (fig. 7.9), recognized at Pittsburg Landing (Vallier, 1974) and in the
286 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

= IOO {0001}, biotite n = 100 a, plagioclase n- 100 /3, plagioclase /? = IOO *V, plogioclase

= IOO {0001}, biotite /?= 200 c, quartz

/MOO

/?= 100 a, plagioclase n- 100 /3, plagioclase /?=IOO "Y , plagioclase

/;= 200 c, quartz {OOOI}, biotite /; = 200 c, quartz c, calcite
7. PRE-CRETACEOUS TECTONIC EVOLUTION OF THE BLUE MOUNTAINS PROVINCE, NORTHEASTERN OREGON 287

central Wallowas (Smedes, 1959), are contemporaneous with the D2 folding (D, F, fig. 7.2). Megascopic 

folds in the central and southern Wallowas are overturned to the northwest (Smedes, 1959; Mirkin, 1986).

MESOSCOPIC ANALYSIS

The orientations of the mesoscopic D2 structures 

are shown in figure 7.3 (projections A, B, D, F, G, and 

/), and average orientations are presented on figure 

7.2 (projections A, B, D, F, G, and /). The style of 

folding strongly correlates with the metamorphic 

grade and is a function of the amount of strain. In 

Hells Canyon (G, fig. 7.2) and the southern Wallowas 

(/, fig. 7.2), folds are open and have a class Ib (parallel) to Ic (weakly converging dip isogons) geometry 

(Ramsay, 1967). In the central Wallowas (F, fig. 7.2), 

the northern Wallowas (B, fig. 7.2), and the Pendleton inliers (A, fig. 7.2), folds approach class 2 (similar-fold) geometry.

The orientation of bedding planes (S0) is highly 

variable owing to the later (D3) refolding. Two very 

indistinct girdles of bedding-plane poles trend northwest and northeast; the first girdle is subnormal to 

the B2 fold axes (D, G, fig. 7.3); the second is related 

to the D3 deformation (B, fig. 7.3).

Fold axial planes and axial-plane cleavages (S2) 

strike generally northeast-southwest (A, D, F, G, and 

/, fig. 7.3). In the northern Wallowas (B, fig. 7.3) the 

S2 cleavages are variable. The fold axes (B2) and lineations (L2) trend roughly northeast-southwest in 

most areas (B, D, F, G, I, fig. 7.3). However, in the

FIGURE 7.5. Lower-hemisphere, equal-area projections of D2- 

deformation microfabric elements of samples from selected areas shown in figure 7.9: A, Metaconglomeratic biotite-hornblende gneiss; B, garnet-biotite-sillimanite 

gneiss; C, garnet-biotite-hornblende gneiss; D, chloritemuscovite schist; G, biotite gneiss; I, marble. Microfabric 

elements are c axes of quartz and calcite, poles to {0001} 

in biotite, and a, (3, and y (optical-indicatrix) axes of plagioclase and green amphibole. n, number of measurements. Contours drawn at intervals of 2 percent per 1- 

percent area where n=50, and at intervals of 1 percent 

per 1-percent area where n>100. Biotite contours at 1, 5, 

10, and 15 percent per 1-percent area in B and 1, 2, 5, 

10, and 15 percent per 1-percent area in G. Dashed and 

heavy (solid) contour lines represent 2- and 4-percent 

contours, respectively, where n=50, and 1- and 2-percent 

contours, respectively, where n=100 (except biotite in B); 

heavy contour is 1-percent contour where n=100 (except 

biotite in G). S2 (heavy great circle), foliation plane; S2' 

(dashed great circle), rare second foliation plane; L2 (solid 

diamond), mineral lineation; L2' (solid diamond), rare second mineral lineation. Light great circle represents geographic horizontal; N, S, E, and W stand for north, south, 

east, and west, respectively.

high-grade metamorphic rocks of the Pendleton inliers (A, fig. 7.2; A, fig. 7.3) fold axes and lineations 

trend generally north-south; many of these lineations 

are mineral lineations, in particular due to parallelism of amphibole crystals.

MICROSCOPIC ANALYSIS

Six samples were chosen for petrofabric analysis. 

The preferred orientations were determined for 

quartz (in four samples), biotite (three), plagioclase 

(two), amphibole (one), and calcite (one) and are presented in figure 7.5. All fabrics have orthorhombic 

symmetry. Quartz c axis fabrics display moderate to 

strong preferred orientations; each consists of two 

cross-girdles, the intersection of which lies in the foliation plane and is approximately 90° from a well-developed mineral lineation (B, C, G, fig. 7.5). Biotite 

[0001] axes form girdles normal to the lineation with 

point maxima perpendicular to the foliation (A, B, G, 

fig. 7.5). Plagioclase (oligoclase) axes show preferred 

orientations with an a axis small-circle girdle around 

the lineation; the (3 axes form a great circle around 

the lineation and have a weak point maximum normal to the foliation (A, C, fig. 7.5). Amphiboles have 

a strong y point maximum parallel to the lineation 

and an oc maximum perpendicular to the foliation (C, 

fig. 7.5). Calcite c axes are preferentially oriented in 

a great circle perpendicular to the lineation (/, fig. 

7.5). Mirkin (1986) carried out microfabric analysis 

on ten samples from the southern Wallowas (/, fig. 

7.2) resulting in seven quartz and three calcite fabrics. His fabric patterns are nearly identical to the 

quartz and calcite fabrics described above.

All these Wallowa terrane D2 microfabrics have orthorhombic symmetry. They can be interpreted in 

terms of the finite-strain axes, but they do not show 

whether or not the deformation was coaxial. Therefore, all six samples and three additional ones were 

studied to find textural criteria indicating noncoaxial 

deformation. Criteria that were used to determine the 

shear sense are asymmetric pressure shadows (A, B, 

J, fig. 7.7), grain elongations at an angle to finegrained shear zones subparallel to the foliation (A, B, 

fig. 7.7), rotated chlorite and quartz fibers in pressure 

shadows of magnetite crystals (D, E, F, fig. 7.7), and 

rotated synkinematic porphyroblasts (snowball garnet, B, fig. 7.7). In six of the nine samples, definite 

conclusions as to the sense of shear could be made.

INTERPRETATION

The quartz, biotite, amphibole, and calcite microfabrics can be interpreted in terms of the orientation
288 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

r

F iJ^k ch-7^^jl.<'

am,

G

CC

0 2.5

 CC

0 .5

CC

K

ph

FIGURE 7.6. Selected microtextures indicating Dpdeformation sense of shear (arrows) in thin sections cut parallel 

to X-Z plane from samples collected in Blue Mountains 

province. Letter designations B, F, G, J, K, and L refer to 

corresponding projections of D± fabrics in figure 7.8. 

Sample B is from Wallowa terrane; other samples from

Baker terrane. B, Quartz diorite mylonite; F, epidote amphibolite; G, marble mylonite; J, calcareous phyllite; K, 

L, phyllite; am, amphibole; cc, calcite; ch, chlorite; ep, 

epidote; M, muscovite; mt, magnetite; pg, plagioclase; ph, 

phyllite; Q, quartz; serp, serpentine. Heavy lines (in F, J, 

and L), microfaults. All scales in millimeters.
7. PRE-CRETACEOUS TECTONIC EVOLUTION OF THE BLUE MOUNTAINS PROVINCE, NORTHEASTERN OREGON 289

B

Q,bl, ,!» Q,bl,am ] 

am ^Vjfffil i

/"/TOP *% o«;r

0 .4

.15

Q, M

Q, M

FIGURE 7.7. Selected microtextures indicating D2-deformation sense of shear (arrows) in thin sections cut parallel 

to X-Z plane from samples collected in Blue Mountains 

province. Letter designations A, B, D, E, F, J, L, M, and 

P refer to corresponding projections of D2 fabrics in figure 7.9. Samples A, B, and D through F, from Wallowa 

terrane; samples J, L, M, and P, from Izee terrane. A, 

Metaconglomeratic biotite-hornblende gneiss; B, garnetbiotite-sillimanite gneiss; D, E, F, chlorite-muscovite 

schist; J, graphitic marble; L, calcareous metaconglomerate; M, interbedded graywacke and phyllite; P, metaconglomerate; am, amphibole; bi, biotite; cc, calcite; ch, 

chlorite; gt, garnet; gw, graywacke; M, muscovite; mt, 

magnetite; ph, phyllite; Q, quartz; serp, serpentine; sp, 

spinel. Heavy lines (in F, L, M, and P) are microfaults. 

All scales in millimeters.
46'

120 C

45 C

25 

I

119 C

50 KILOMETERS

118 C 

WASHINGTON

WALLOWA TERRANE

BAKER TERRANE

D N

IZEE TERRANE.

FIGURE 7.8. Geologic map showing Late Triassic to Early Jurassic deformation pal compressive-strain axis. Solid great circle, X-Y pl

(D]) structures in Blue Mountains province. Inset diagrams A through O are great circle, X-Z plane or movement plane. Heavy a

lower-hemisphere, equal-area projections of principal strain axes. X, major prin- tion; double-headed arrows, left-lateral strike-slip d

cipal extensile-strain axis; Y, intermediate principal strain axis; Z, major princi- of X-Z plane. See figure 7.2 for explanation of geologi
46'

120 C 119 C 118 C 

45 <

WALLOWA TERRANE

F

BAKER TERRANE

IZEE TERRANE.

FIGURE 7.9. Geologic map showing Late Jurassic (deformation D2) structures in X-Y plane or cleavage plane; dashed great circle, X

Blue Mountains province. A-R, lower-hemisphere, equal-area projections of Heavy arrow, tectonic-transport direction; double-he

principal strain axes. X, major principal extensile strain axis; Y, intermediate strike-slip displacement; heavy bar, strike of X-Z pla

principal strain axis; Z, major principal compressive strain axis; X', Y, and Z', fault (sawteeth on upper plate; dotted where conce

principal strain axes of rare second penetrative deformation. Solid great circle, nation of other geologic map symbols.
292 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

of the principal finite strain axes (see section "Structure, Microscopic Analysis"). On the basis of the deduced strain axes, the plagioclase fabrics appear to 

have formed by translation glide on (001)[100]. The 

orientation of the X, Y, and Z strain axes and the 

sense of shear are plotted on figure 7.9. The mean 

orientation of the strain axes and sense of tectonic 

transport as determined by Mirkin (1986) are shown 

in projection H on figure 7.9.

These results are consistent with the mesoscopic 

structures (A, F, fig. 7.2). Generally, the X axes are 

parallel to south- to southeast-plunging mineral lineations (L2 ) and fold axes (B2 ) (A and F, fig. 7.3), and 

the Z axes are perpendicular to the S2 cleavages. The 

parallelism of fold axes and X axes indicates large 

rotations (Escher and Watterson, 1974) and thus very 

large strains; this is also supported by the class 2 

(similar-fold) geometry (Ramsay, 1967) of the folds in 

the Pendleton inliers. In the southern Wallowa Mountains and in Hells Canyon, where the rocks are hardly recrystallized, fold axes trend northeast-southwest 

and the folds have a class Ib (parallel) to Ic (weakly 

converging dip isogons) geometry, indicating much 

lower strains and no or little strain rotations.

The strike of the X-Z planes (main planes of distortion or so-called movement planes) are accentuated in 

figure 7.9 with heavy bars, and where the sense of 

shear could be established (on the basis of microtextures or of the vergence of megascopic folds), the 

sense of tectonic transport is indicated. It is clear that 

shearing and thrusting in the Wallowa Mountains 

were toward the north and northwest during the D2 

deformation. The shortening direction in the northern 

Wallowas (G, fig. 7.9) is northeast-southwest; the 

sample from which this conclusion was drawn must 

have been rotated clockwise, as similarly indicated by 

the mesoscopic and megascopic structures. This rotation may have been produced by either the intrusion 

of the Wallowa batholith or younger (Tertiary) faulting (Smedes, 1959). One sample from the Pendleton 

inliers (A, fig. 7.9) shows an opposite, southerly vergence, perhaps the result of back thrusting. Another 

sample (B, fig. 7.9) shows two noncoaxial superposed 

strains, both probably related to the D2 deformation.

TIMING

There are not many constraints on the timing of 

the D2 deformation. The youngest rocks in the Wallowa Mountains deformed during D2 are of late Early 

Jurassic (Toarcian) age (Nolf, 1966). The oldest intrusives in the Wallowa terrane were dated at about 160 

Ma (Armstrong and others, 1977), but, as previously 

mentioned, more recent work suggests that the main

magmatism did not start until about 143 Ma (Walker, 

1989). The plutonic rocks are all post-tectonic and cut 

across the D2 structures. Thus, a Middle to Late Jurassic age for D2 is indicated. In the Snake River canyon, Middle to Upper Jurassic (Callovian, Oxfordian) 

sedimentary rocks have been recognized unconformably overlying Trias sic and older rocks (Morrison, 

1964; Vallier, 1977; Goldstrand, 1994; White and 

Vallier, 1994). These Middle and Upper Jurassic rocks 

are folded (D, fig. 7.2; D, fig. 7.3). Fold axial planes 

and cleavages are steep and strike northeast-southwest; fold axes are subhorizontal. It can be argued 

that the deformation responsible for these folds occurred during the Cretaceous, at the same time as 

the major deformation in the Riggins, Idaho, area 

(Sutter and others, 1984), and that the unconformity 

is related to the post-Toarcian prebatholith deformation. However, no penetrative structures related to a 

Cretaceous event have been recognized elsewhere in 

Oregon. Thus it is suggested that the major Jurassic 

deformation of the Wallowa terrane occurred in the 

Late Jurassic, between the early Oxfordian and the 

time of the main batholith formation (-143 Ma).

D3 DEFORMATION

The third penetrative deformation D3 resulted only 

in small folds, kink folds, and crenulations (A, B, F, 

G, and /, fig. 7.3). It did not involve large amounts of 

strain. In general, folds have a class Ib (parallel) geometry (Ramsay, 1967). The axial planes are steep 

and strike northwest-southeast to north-south (A, B, 

F, G, and /, fig. 7.2); they are almost perpendicular to 

the D2 structures.

The timing of the D3 deformation is rather 

straightforward. The D3 structures deform the older 

D2 structures and are crosscut by the Wallowa batholith. Thus they may be a late expression of the same 

event that caused the D2 structures. Both the D2 

folds and the D3 cross folds may be the expression of 

only one displacement field. If the convergence-rate 

vector between two terranes is oblique to their mutual boundary, strain partitioning will occur in general: 

the vector component perpendicular to the boundary 

causes boundary-parallel thrust faults and related 

folds, whereas the boundary-parallel component results in boundary-parallel strike-slip faults (Oldow 

and others, 1990). The D3 cross folds may be the result of restraining bends in such strike-slip faults.

D4 DEFORMATION

A few post-D2 deformation subhorizontal kinks 

with east-west-trending fold axes were identified in
7. PRE-CRETACEOUS TECTONIC EVOLUTION OF THE BLUE MOUNTAINS PROVINCE, NORTHEASTERN OREGON 293

the Pendleton inliers (A, fig. 7.3); they have been attributed to a fourth deformation (D4). Because they 

have not been found elsewhere, they are considered 

unimportant and are not further discussed in this 

chapter.

BAKER TERRANE

D! DEFORMATION

The Permian and Triassic rocks of the Baker terrane are thought to represent a fore arc. Locally, the 

rocks are severely disrupted and chaotic, but in general the terrane consists of large coherent blocks in 

which the structures, although complex, have a certain regularity. These blocks are separated from each 

other by chaotic zones and shear zones. All rocks are 

metamorphosed, but the grade of metamorphism is 

extremely variable. Most of the rocks have undergone 

only the lowest greenschist-facies metamorphism, but 

amphibolites and blueschists do occur. Megascopic D± 

folds have been recognized only in the Elkhorn Ridge 

area (M, fig. 7.2); they were delineated there by mapping asymmetric mesoscopic folds and cleavage-bedding relationships (Coward, 1983).

MESOSCOPIC ANALYSIS

The metasedimentary rocks are deformed into tight 

to isoclinal Dj folds, which have a class Ic (weakly 

converging dip isogons) to class 2 (similar) geometry 

(Ramsay, 1967). Several generations of cleavages 

commonly developed (Coward, 1983) cozonally, their 

mutual intersections paralleling the fold axes. Apparently cleavage formation was accompanied by strong 

dissolution. The major trend of the axial-plane foliation and cleavage (S^ is shown in projections J 

through O on figure 7.2. In the eastern part of the 

terrane the foliation strikes northeast-southwest; in 

the central and western parts it strikes approximately east-west. Fold axes (Bj) and intersection lineations (Li) in the Baker terrane are generally steeply 

plunging and trend northwest-southeast. Mineral lineations in the blueschists (J, fig. 7.2; J, fig.7.3) and 

amphibolites (K, fig. 7.2; K, fig.7.3) also plunge steeply to the southeast. Fragments in sedimentary breccias and pebbles in conglomerates occurring in the 

Vinegar Hill area are flattened parallel to the foliation and have a slight elongation (parallel to the 

strain axis X) trending northwest-southeast (E, fig. 

7.8). In the Elkhorn Ridge area, bedding (S0 ) and foliations (Si) are moderately to steeply dipping to the 

south, and fold axes (B^ and intersection lineations

form a girdle with a south-plunging maximum 

(M, fig. 7.2; M, fig. 7.3). It is here that Coward (1983) 

discovered several generations of cleavages, all cozonal with an east-west-trending subhorizontal axis.

Also in the Elkhorn Ridge area, shear zones were 

identified in particular where they crosscut marble 

bodies. The sheared marble is strongly mylonitized 

and has subhorizontal, east-west-trending mineral 

lineations.

In the Burnt River canyon area, argillite fragments 

in a mud breccia are flattened and elongated. The 

elongations (L^) have variable orientations: in some 

places they are subhorizontal and trend east-west, 

and in other places they plunge steeply to the south 

or north (N, fig. 7.2; N, fig. 7.3).

Near the Snake River canyon (O, fig. 7.2), chert, 

argillite, and limestone are exposed with steep northdipping cleavages (S^). Intersection lineations (L^) 

plunge moderately to the west (O, fig. 7.3).

MICROSCOPIC ANALYSIS

Microfabric analysis was performed on six samples 

from the Baker terrane. Microtextures of these and 

some additional samples were examined in a search 

for kinematic indicators.

The preferred orientation of quartz c axes and 

glaucophane optical-indicatrix axes in a blueschist 

from the Mitchell area (D, fig. 7.8) are shown in D on 

figure 7.4. The quartz grains are recrystallized and 

show poorly developed c axis cross-girdles, which intersect in the foliation plane (8j) approximately at a 

right angle to the mesoscopic mineral lineation (L^). 

The glaucophane crystals have an a point maximum 

perpendicular to Si and a y point maximum parallel 

to Lj_. No textural indicators were found to establish 

the shear direction.

The orientation of amphibole in an amphibolite 

from the Vinegar Hill area (F, fig. 7.8) has fabric patterns (F, fig. 7.4) that are nearly identical to those of 

the glaucophane in the blueschist. According to E.M. 

Bishop (Mullen, 1978; Bishop, chap. 4, this volume) 

these amphiboles are also sodic and thus may have 

formed at relatively high pressures in a subduction 

complex. Several kinematic indicators were found in 

thin sections cut parallel to the X-Z plane: rotated 

porphyroblasts, extension cracks, and displacements 

across microfaults that parallel cleavage planes (F, 

fig. 7.6).

In a marble mylonite from the Elkhorn Ridge area 

(G, fig. 7.8), the c axes of calcite lie in a girdle (G, fig. 

7.4) perpendicular to a strong mesoscopic subhorizontal lineation trending east-west. The mylonite has a 

typical S-C structure indicative of the sense of
294 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

displacement: the longest dimensions of large calcite 

grains (subparallel to the X strain axis) are oriented 

at a small angle to shear zones containing finegrained calcite (G, fig. 7.6).

Two phyllite samples from the Burnt River canyon 

yielded surprising results. These phyllites contain 

quartz augen and very flattened and elongate argillite pebbles that show no significant recrystallization. 

In one of these samples (L, fig. 7.8) the quartz c axes 

are oriented in two cross-girdles (L, fig. 7.4) which 

intersect in the foliation plane (Sj), at 90° to the lineation (L1) that is caused by parallelism of the longest axes of the stretched argillite pebbles. Quartz 

fibers in pressure shadows of magnetite grains and 

displacements along microfaults yielded consistent 

shear senses (L, fig. 7.6). The other sample (7, fig. 

7.8) also displays a cross-girdle pattern of quartz c 

axes, but the intersection of the two girdles is parallel to the lineation Llt To verify these results, another thin section perpendicular to the first was studied, 

and it gave the same result (7, fig. 7.4).

In a sample of slightly recrystallized metachert collected in the Snake River canyon (M, fig. 7.8) the 

quartz c axes form two cross-girdles, intersecting in 

the Si plane at about 90° from a mesoscopic intersection lineation (M, fig. 7.4). One girdle is more strongly 

populated than the other, giving rise to a monoclinic 

symmetry which is an indicator of the shear sense.

Two other phyllite samples from the Burnt River 

canyon (J, K, fig. 7.8) have consistent kinematic indicators, such as rotated calcite and quartz fibers in 

veins (J, fig. 7.6), and rotated quartz fibers in pressure shadows of magnetite, asymmetric pressure 

shadows, and rotated porphyroclasts (K, fig. 7.6).

INTERPRETATION

The microfabrics of five samples have orthorhombic 

symmetry, from which the orientation of the principal 

strains can be identified: in all instances the foliation 

plane Si coincides with the X-Y plane of the finitestrain ellipsoid. In all samples, the major principal 

extension axis (X) can also be identified. In four samples X is parallel to a pervasive lineation, but in one 

sample (7, fig. 7.4) X is perpendicular to the lineation. 

The latter situation may be the result of flattening 

during deposition and diagenesis, followed by tectonic 

deformation (see Ramsay and Wood, 1973).

The orientations of X, Y, and Z axes in the Baker 

terrane, as derived from microfabrics and microtextures, are shown in projections D through M on figure 

7.8. These results are consistent with the mesoscopic 

data presented in projections J through O on figure 

7.3. In all comparable instances the Z axes correspond to the maximum of poles to Si cleavage planes 

and the X axes are parallel to the maximum of BI 

fold axes (for example, compare Z and X in projection 

M, fig. 7.8, with Si and BI, respectively, in projection 

O, fig. 7.3). The style of deformation (class 2, or similar-fold, geometry; Ramsay, 1967) and the parallelism 

of X axes and BI fold axes indicate large strains and 

large strain rotations (Escher and Watterson, 1974). 

Left-lateral displacement in the Elkhorn Ridge area 

(G, fig. 7.8) along a steep south-southeast-dipping 

shear plane is clearly different from the displacements in most other areas of the Baker terrane (D F 

and 77-M, fig. 7.8). The sense of shear, derived from 

the monoclinicity of one quartz fabric (M, fig. 7.4) and 

from several kinematic indicators (F, J, K, L, fig. 7.6), 

shows that most displacements in the Baker terrane 

are dip-slip motions related to northwest-directed 

thrusting or southeast-directed underthrusting. (The 

curved quartz fibers in the pressure shadows of drawings K and L, figure 7.6, may have formed by a small 

amount of simple shear followed by a large amount of 

pure shear; this interpretation is different from that 

of Etchecopar and Malavieille (1987) but is consistent 

with other kinematic indicators and the orthorhombic 

symmetry of the quartz fabrics.)

TIMING

It is difficult to constrain the timing of the DI deformation over such a large area. It is very well possible that the deformation was diachronous, that it 

occurred earlier in one area and later in another, 

within the Baker terrane. The youngest sedimentary 

rocks deformed by DI are of Late Triassic and possibly Early Jurassic age (Blome and others, 1983, 

1986). An 40Ar/39Ar age of 223 Ma (Hotz and others, 

1977) for the blueschists at Mitchell indicates that 

the DI deformation at Mitchell was at least constrained to the Late Triassic. The deformation was 

certainly pre-Cretaceous in parts of the Baker terrane where the rocks are crosscut by Early Cretaceous granites (Armstrong and others, 1977; Walker, 

1989). Although the contact between the Elkhorn 

Ridge Argillite and the Jurassic sedimentary rocks of 

the Izee terrane in the Snake River canyon along 

the Brownlee Reservoir is a tectonic contact, it is 

plausible that the D2 structures in the Baker terrane 

(see below) were formed at the same time as the first 

structures (here also called D2) in the Lower Jurassic 

sedimentary rocks of the Weatherby Formation; they 

are similar in style and orientation. The DI structures in the Baker terrane do not occur in the Jurassic sedimentary rocks, and thus the DI structures 

probably formed before the Early Jurassic.
7. PRE-CRETACEOUS TECTONIC EVOLUTION OF THE BLUE MOUNTAINS PROVINCE, NORTHEASTERN OREGON 295

D2 DEFORMATION

Megascopic D2 folds in the Baker terrane were recognized only in the Burnt River canyon (fig. 7.1; N, 

fig. 7.2; Ashley, 1966). Large north-dipping reverse 

faults were recognized along the contact of the Baker 

and Izee terranes (fig. 7.2 between areas O and U; 

Brooks, 1979a). The axial trends of the D2 folds 

shown in figure 7.2 are based mainly on the mesoscopic data; these structures (B2 and S2) are generally subparallel to the D! trends. Although data on 

mesoscopic structure were collected and are discussed below, microscopic studies were inconclusive 

and are not reported in this chapter.

MESOSCOPIC ANALYSIS

D2 folds generally have a class Ib (parallel) to Ic 

(weakly converging dip isogons) geometry (Ramsay, 

1967). They are often crenulations and kink folds involving little strain. Generally, the axial planes and 

cleavages dip steeply to the north and northwest (JO, fig. 7.2; J-O, fig. 7.3). The orientations of the B2 

fold axes and L2 lineations are variable but are often 

subhorizontal, trending roughly east-west.

INTERPRETATION

Although D2 structures are locally penetrative, no 

kinematic-indicator data have been gathered because 

microtextural features are generally reliable only in 

rocks deformed just once. Nonetheless, the consistent 

northward dip of the axial planes and cleavages in the 

rocks of the Baker terrane suggests a south vergence, 

opposite to the D2 vergence in the Wallowa terrane but 

similar to the vergence of D2 folds in the Izee terrane.

TIMING

The timing of the D2 deformation in the Baker terrane is not well constrained. This deformation postdates the D! deformation and certainly predates the 

intrusion of the granitic plutons, such as the Bald 

Mountain batholith, into the Elkhorn Ridge Argillite, 

which has been dated at about 145 Ma (Armstrong 

and others, 1977; Walker, 1989).

D3 DEFORMATION

A third set of folds crosscutting the DI and D2 

folds is recognized almost everywhere in the Baker 

terrane. These D3 folds are of only mesoscopic and

microscopic size. Their geometry belongs to class Ib 

(parallel) to Ic (weakly converging dip isogons; Ramsay, 1967); they are generally crenulations and kink 

folds. Their axial planes generally are steep and 

strike north-south. 63 fold axes and L3 lineations 

have variable orientations (J, L-N, fig. 7.3). The general trends of D3 structural elements are shown in 

projections J and L through N on figure 7.2.

The timing of the D3 folding is post-D2 deformation 

and prebatholith intrusion. As discussed above for 

the Wallowa terrane, the D3 cross folds may have 

formed in response to oblique convergence.

IZEE TERRANE

D! DEFORMATION

The oldest structures in the Izee terrane that are 

ascribed to the DI deformation occur in the area 

south and southwest of the town of John Day (P S, 

fig. 7.2). These structures are found in the oldest sedimentary rocks, which are Late Triassic (Carnian) in 

age, and seem to die out in rocks of Early Jurassic 

(Sinemurian) age. Dickinson and Vigrass (1965), 

Brown and Thayer (1966a, b), and Thayer and Brown 

(1966) recognized that the deformation was syndepositional; they mapped several north-south-trending 

megascopic DI folds.

MESOSCOPIC ANALYSIS

In the Aldrich Mountains (Q, fig. 7.2) mesoscopic 

folds are open and generally of class Ib (parallel) to 

Ic (weakly converging dip isogons) geometry (Ramsay, 1967). Fold axes plunge to the south, and axial 

planes dip steeply to the east (Q, fig. 7.3). Elsewhere, 

the folds are tight to isoclinal and have class Ic to 2 

(similar) geometries. Axial planes and axial-plane 

cleavages dip steeply eastward in the Izee area (P, 

fig. 7.2; P, fig. 7.3); east of Izee, in the John Day area 

(S, fig. 7.2) and the Wickiup area (R, fig. 7.2), the 

cleavages are rotated and dip moderately to the 

northeast (R, S, fig. 7.3).

MICROSCOPIC ANALYSIS

None of the samples collected were suitable for petrofabric analysis. Microscopic strain analysis was 

performed on one graywacke sample that contained 

numerous fragments of chert and volcanic rocks and 

grains of serpentine, spinel, plagioclase, calcite, and 

quartz. The aspect ratio of the chert grains is about 

1.5:1.0:0.5 for X:Y:Z; the orientation of these axes is 

shown in projection O on figure 7.8.
296 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

INTERPRETATION

The asymmetry of mesoscopic folds and the general 

east to northeast dip of the axial planes and axialplane cleavages suggest that the folds formed in a 

regime of west- to southwest-directed thrusting (P-R, 

fig. 7.2; P-R, fig. 7.3), although no kinematic indicators were found. The north-south to northwest-southeast trend of the Dj structures in the Izee terrane is 

markedly different from the east-west to northeastsouthwest trend of Dj structures in the Baker terrane (J-O, fig. 7.2; J-O, fig. 7.3). Two factors may 

have been involved: (1) the ramping effect due to the 

rigid Canyon Mountain Complex may have caused 

the deviation, although this complex was itself deformed somewhat during the DI event (Ave Lallemant, 1976); (2) the subsequent D2 deformation may 

have caused rotation of the Dj structures.

TIMING

The timing of the Dx deformation in the Izee terrane was relatively well established by Dickinson 

and Vigrass (1965), Brown and Thayer (1966a, b), 

and Thayer and Brown (1966). The youngest sedimentary rocks deformed by Dx are of Early Jurassic 

(Sinemurian) age; upper Lower Jurassic (Pliensbachian) rocks are not affected. Thayer and Brown (1966) 

have shown that the deformation began during the 

Late Triassic, because Upper Triassic rocks are more 

heavily strained than Lower Jurassic rocks.

D2 DEFORMATION

The results of the D2 deformation are recognized in 

all parts of the Izee terrane. In the John Day area, D2 

structures crosscut DI structures; elsewhere, D2 

structures are the first penetrative structures. Megascopic folds have been recognized throughout the terrane (Dickinson and Vigrass, 1965; Brown and 

Thayer, 1966a, b; Thayer and Brown, 1966; Ave Lallemant, 1983). They trend generally from east-west (in 

the western part of the terrane) to northeast-southwest (in the eastern and central parts) (fig. 7.2). Jurassic conglomerates have been found in contact with 

pre-Jurassic rocks in the Mine Ridge area (T, fig. 7.2) 

and Weatherby area (U, fig. 7.2). These conglomerates 

are much more strongly deformed than the older 

rocks, indicating that the contact is structural. In the 

Weatherby area, Sinemurian sedimentary rocks overlie younger (Pliensbachian) conglomerates, suggesting 

a thrust relationship (Brooks and Vallier, 1978). Also 

in the Weatherby area (near M, N, R, fig. 7.9) the

contact between the Baker and Izee terranes is a 

north-dipping reverse fault (Brooks and others, 1976) 

intruded by a post-tectonic stock dated at 124 Ma 

(Walker, 1989).

MESOSCOPIC ANALYSIS

Small-scale folds in the graywackes have a class Ib 

(parallel) geometry and in pelitic rocks a class Ic 

(weakly converging dip isogons) to 2 (similar) geometry (Ramsay, 1967). In the west, fold axes (B2) and 

bedding-cleavage intersection lineations (L2) are subhorizontal and trend east-west (P-T, fig. 7.2; P-T, fig. 

7.3). Near the Snake River canyon in the east, fold 

axes and lineations form great-circle girdles with a 

steep point maximum (U, fig. 7.2; U, fig. 7.3). Axialplane cleavages (S2) are well developed locally, in 

particular near the Snake River (U, fig. 7.3); they 

generally dip steeply to the north. In stretched pebble conglomerates of the Mine Ridge area (T, fig. 7.2), 

the longest pebble axes (subparallel to the strain axis 

X) plunge at low angles to the east-southeast. The X 

axes of stretched pebbles in the Weatherby area (U, 

fig. 7.2) plunge to the northwest, parallel to the dip 

direction of the S2 cleavage; axial ratios X/Z are locally as large as 10.

MICROSCOPIC ANALYSIS

Previously (Ave Lallemant, 1983), four quartz c axis 

fabrics were obtained from the Jurassic sedimentary 

sequence near Weatherby. They all display the crossgirdle pattern, and the fabrics have orthorhombic 

symmetry. Microtextural studies of these rocks yielded kinematic information in one sample only, in 

which displacements were noted along microfaults 

parallel to the cleavage (M, fig. 7.7). The Jurassic 

stretched-pebble conglomerates from the Mine Ridge 

area (T, fig. 7.2) and the Weatherby area (U, fig. 7.2) 

have several consistent kinematic indicators such as 

quartz and calcite fibers in extension veins and in 

pressure shadows and displacements along microfaults parallel to the cleavage plane (L, P, fig. 7.7).

INTERPRETATION

The geometries and orientations of the D2 folds in 

the Izee terrane (P-U, fig. 7.3) can be explained in 

terms of strain (see section "Structure, Microscopic 

Analysis"). Principal strain axes were derived from 

the quartz fabrics (M, N, Q, R, fig. 7.9; Ave Lallemant, 1983), and the sense of shear from microtex-
7. PRE-CRETACEOUS TECTONIC EVOLUTION OF THE BLUE MOUNTAINS PROVINCE, NORTHEASTERN OREGON 297

tural kinematic indicators (L, M, P, fig. 7.7). The 

shear senses in the rocks near Wickiup and Weatherby (K, O, fig. 7.9) are based on the asymmetry of 

the folds on all scales. All D2 data for the Izee terrane are consistent with south- to southeast-directed 

thrusting during the Late Jurassic.

The kinematic results from the stretched-pebble 

conglomerate in the Mine Ridge area (L, fig. 7.9) can 

be interpreted in two ways. The first interpretation is 

that the rocks were thrust to the south and later, 

during the D3 deformation, rotated into their present 

position. Although this interpretation is feasible, D3 

structures were not recognized in the area, and D3 

structures generally show only very minor strain. 

The second and preferred alternative is that these 

conglomerates were deformed by left-lateral strikeslip displacement along a west-northwest-east-southeast shear zone (L, fig. 7.9).

TIMING

In the John Day area, the youngest sedimentary 

rocks affected by the D2 deformation are of late Middle Jurassic (Callovian) age, whereas the oldest undeformed sedimentary rocks are of Late Cretaceous 

(Cenomanian) age (Dickinson and Vigrass, 1965). 

Small granitic plutons that intruded the Jurassic 

sedimentary rocks post-tectonically have not been 

dated but are believed to be of Early Cretaceous age 

(Thayer and Brown, 1966).

The youngest sedimentary rocks deformed by the 

D2 deformation in the Weatherby area are of Middle 

Jurassic (Bajocian) age (Brooks, 1979a), but across 

the Snake River in Idaho lower Callovian sedimentary rocks show D2 deformational effects (Henricksen, 

1975). These rocks were intruded post-tectonically by 

a granitic stock, which has been dated at 124 Ma 

(U-Pb: Walker, 1989).

D3 DEFORMATION

In most areas the Late Jurassic (D2) structures 

have been refolded. D3 folds are generally open and 

of class Ib (parallel) to Ic (weakly converging dip isogons) geometry (Ramsay, 1967). In slate and phyllite 

this deformation is characterized by crenulations and 

kink bands. Generally, the axial planes (S3) are steep 

and strike north-south, whereas the fold axes (B3) 

have variable orientations (P, R, S, U, fig. 7.2; P, R, 

S, U, fig. 7.3). The timing of the D3 deformation is 

rather unconstrained. It postdates the D2 event, but 

an upper age limit has not been established.

OLDS FERRY TERRANE

Although there exists a hiatus between the Upper 

Trias sic volcanic and volcaniclastic rocks of the 

Huntington Formation of Brooks (1979a) and the 

Lower Jurassic sedimentary rocks of the Weatherby 

Formation, and although the lowermost rocks of the 

Weatherby Formation are sheared and thrust southeastward, no important deformation structures have 

been recognized in the Huntington Formation. No 

mesoscopic folds have been found, and no distinct 

cleavages have been observed. Poles to bedding 

planes have been plotted in projection V on figure 

7.3; they form a crude girdle the normal of which 

plunges moderately to the northeast. On this basis it 

is assumed that the rocks in the Olds Ferry terrane 

are folded in major open folds and that this deformation took place during Late Jurassic time, although 

there exist no constraints on the timing.

TECTONIC MODEL

It should be reemphasized that the hypotheses proposed here are tentative and that the tectonic model 

is preliminary. Detailed studies should be carried out 

in many areas to support or reject the overall picture 

that emerges from the data presented here. Paleomagnetic evidence (Wilson and Cox, 1980; Hillhouse 

and others, 1982) indicates that the entire Blue 

Mountains region has to be rotated counterclockwise 

about 65° to place the structures within the context 

of regional tectonic models and pre-Cretaceous geography (fig. 7.10). Unless stated otherwise, in the following discussion of the proposed tectonic model, 

geographic orientations are given in terms of preCretaceous coordinates.

D! DEFORMATION

The oldest (D^) structures in most areas of the Baker terrane formed during the Late Triassic to Early 

Jurassic, probably in a subduction environment as 

suggested by the occurrence of 223-Ma blueschists 

(Hotz and others, 1977). However, a 255-Ma (40Ar/ 

39Ar: Ave Lallemant and others, 1980) amphibolite 

knocker in the melange near the Canyon Mountain 

ophiolite suggests that the DI deformation may have 

started in the Permian. The Izee terrane was first 

deformed during Late Triassic to Early Jurassic time. 

Kinematic evidence, presented above, suggests that 

the DI deformation was caused by eastward subduction underneath the Olds Ferry volcanic-arc terrane 

(fig. 7.1QA). Although the deformation appears to
298 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

have ceased in the Early Jurassic, subduction continued into Late Jurassic time as indicated by the occurrence of volcanic strata in the Izee terrane, which was 

interpreted as a fore-arc basin of the Olds Ferry volcanic arc (Dickinson and Vigrass, 1965). Although melange and structures related to Middle Jurassic 

subduction have not been recognized, they may be 

deeply buried or may have been removed by strikeslip displacement. A mylonitic shear zone in the Baker terrane, trending approximately north-south in the 

pre-Cretaceous, was formed by left-lateral slip. Thus, 

the convergence between an oceanic plate (part of 

which is preserved in the Baker terrane) and the volcanic arc (Olds Ferry) was most probably left-oblique 

(fig. 7.10A).

The most important components of the Baker terrane are the chert-argillite sequence, Permian limestone of diverse affinities (Tethyan and American), 

and blueschists. Very similar and coeval assemblages 

have been found to the south in the Klamath Mountains and the Sierra Nevada (see Schweickert and 

Cowan, 1975; Burchfiel and Davis, 1981; Mortimer, 

1986). Thus, the Olds Ferry arc may be related to the

west-facing Triassic to Early Jurassic volcanoplutonic-arc terranes in the Klamaths and Sierra Nevada. 

To the north, in British Columbia, the Cache Creek 

assemblage is virtually identical to the Baker terrane 

(see Monger and others, 1982; Monger and Berg, 

1984). The Cache Creek assemblage lies west of the 

Quesnellia terrane, a volcanic-arc terrane of Triassic 

and Jurassic age. Thus, the Olds Ferry arc may be 

correlated also with the Quesnellia terrane. Mortimer 

(1986) made a similar correlation and, on the basis of 

geochemical data, suggested that the Quesnellia terrane represents a west-facing arc, which is consistent 

with the structural data presented herein.

The Wallowa terrane represents a volcanic arc that 

was active during Permian and Triassic time; volcanic activity ceased during Late Triassic (Norian) time, 

except for some minor enigmatic tuffs in the Middle 

and Upper Jurassic (Callovian to Oxfordian) Coon 

Hollow Formation (Goldstrand, 1994; White and 

Vallier, 1994). The oldest structures in the Wallowa terrane described herein are the three mylonitic 

shear zones (Oxbow, Cougar Creek, and Imnaha) 

along the Snake River. They were active during MidWASH_

'OREG

4 Olds Ferry 

Terrane

'Izee Terrane

_._OREG____ 

CALIF~I NEV

Wallowa 

Terrane

Baker 

Terrane

$ _WASH_

^* 6 REG

Olds Ferry 

Terrane

Izee Terrane

_OREG_ _ 

CALIF F NEV

i

I

B

EXPLANATION

Volcanic arc 

Accretionary complex 

Fore-arc basin

Subduction zone 

Sawteeth on upper plate

N- Thrust fault 

Sawteeth on upper plate

^^Strike-slip fault 

Arrows show relative movement

300 KILOMETERS

FIGURE 7.10. Tectonic model of Blue Mountains province. A, 

Late Triassic to Early Jurassic. B, Late Jurassic. In A, Olds 

Ferry terrane represents a Late Triassic to Early Jurassic 

west-facing volcanic island arc. Concurrently, Baker terrane 

melange and intra- and fore-arc basin sediments of Izee 

terrane were underthrust eastward. Plate convergence was 

left-oblique; left-lateral shear zones developed in Baker terrane. Main volcanic activity in Wallowa arc ceased during 

Late Triassic time, when major left-lateral strike-slip zones 

developed. Olds Ferry volcanic arc was active until Middle

Jurassic time; it may be correlated with Jurassic arc in California (Oldow and others, 1984). Collision of Wallowa terrane 

and the other three terranes occurred during Late Jurassic 

(B), at same time that collision apparently occurred in Sierra 

Nevada (Schweickert and Cowan, 1975; Oldow and others, 

1984, 1989). This collision is expressed by westward thrusting 

in Wallowa terrane and eastward thrusting in Baker, Izee, and 

Olds Ferry terranes. Some left-lateral slip along north-northeast-south-southwest shear zones may have occurred. (All orientations refer to pre-Cretaceous geographic coordinates.)
7. PRE-CRETACEOUS TECTONIC EVOLUTION OF THE BLUE MOUNTAINS PROVINCE, NORTHEASTERN OREGON 299

die to Late Triassic time. After applying the 65° counterclockwise rotation, their original orientation is 

found to be approximately north-south. All three mylonite zones were formed by left-lateral strike-slip 

displacement (fig. 7.10A). The minimum amount of 

displacement along the Oxbow shear zone has been 

estimated as 65 km by Ave Lallemant and others

(1985); the aggregate displacement along all three 

shear zones may have been hundreds of kilometers. 

The displacements along the shear zones may have 

coincided with the end of the volcanic activity in the 

Wallowa volcanic arc.

The data presented herein are not sufficient to indicate the polarity of the Wallowa arc. All kinematic 

indicators in the Baker terrane indicate that this melange is related to the west-facing Olds Ferry arc. If 

the Wallowa arc was east-facing, either its accretionary wedge has not been recognized or it has been 

buried or completely displaced by strike-slip motion. 

Without evidence to the contrary, it is proposed here 

that the Wallowa arc was also west-facing (fig. 

7.10A).

The Wallowa terrane has been correlated with the 

Wrangellia terrane (Jones and others, 1977; Newton, 

1983; KLepacki and Wernicke, 1985; Wernicke and 

Klepacki, 1988), but this correlation has been refuted 

on the basis of the geochemical composition of Triassic volcanic rocks (Sarewitz, 1983; Scheffler, 1983) 

and of Triassic faunas (Stanley, 1986). Mortimer

(1986) correlated the Wallowa terrane with the Stikinia terrane (Monger and others, 1982; Monger and 

Berg, 1984) in British Columbia. Monger and Berg

(1984) suggested that the Cache Creek assemblage is 

related both to the Quesnellia terrane to the east and 

the Stikinia terrane to the west. A very different interpretation was favored by Klepacki and Wernicke

(1985) and Wernicke and Klepacki (1988). They proposed that the Stikinia, Quesnellia, and Olds Ferry 

terranes are fragments of one volcanic arc terrane 

that was disrupted and displaced northward along 

right-lateral strike-slip faults resulting from the Cretaceous collision of the Wrangellia terrane (including 

the Wallowa terrane) with North America. Their 

model, while interesting and thought-provoking, is 

not compatible with stratigraphic and faunal data 

(see Oldow and others, 1989).

D2 DEFORMATION

The D2 deformation may have resulted from a relatively short-lived event, related to the collision of the 

Wallowa terrane with the amalgamated Baker-IzeeOlds Ferry terranes (fig. 7.10A). The youngest sedimentary rocks deformed during the D2 event are of

Oxfordian age, whereas post-tectonic plutons in all 

four terranes were emplaced between 143 and 120 

Ma (Armstrong and others, 1977; Walker, 1989). Kinematic analysis indicates that during collision rocks 

in the Wallowa terrane were thrust westward, whereas the orientation of D2 structures in the Baker and 

Izee terranes indicate eastward thrusting. In one 

area mylonites indicate left-lateral displacements 

along originally northeast-southwest-trending shear 

zones. This suggests that plate convergence during 

the Late Jurassic was left-oblique.

The Late Jurassic amalgamation of the Blue Mountains province is approximately coeval with probable 

collision in the Sierra Nevada (Schweickert and Cowan, 1975) and a back-arc basin collapse in the KLamath Mountains (Harper and Wright, 1984).

D3 DEFORMATION

A persistent but, in terms of strain, insignificant 

cross-folding event occurred subsequent to the D2 deformation. This D3 event took place also before the 

emplacement of the Early Cretaceous plutons. Its origin may be related to the strike-slip component of 

the oblique convergence vector between the Wallowa 

terrane and the amalgamated Baker-Izee-Olds Ferry 

terranes.

LEFT-OBLIQUE PLATE CONVERGENCE

Major north-south-striking left-lateral strike-slip 

zones in the Blue Mountains province were active 

first during the Middle to Late Triassic and second 

during the Late Jurassic. This implies that the oceanic plates being subducted beneath the Wallowa and 

Olds Ferry arcs had a southward component of motion relative to the North American continent during 

these times. On the basis of paleomagnetic evidence, 

Stone and others (1982) proposed two migration models for terranes of southeastern Alaska (including the 

Wrangellia terrane). During the Permian, these terranes resided at high, northerly latitudes (see also 

Harbert and others, 1988). According to their second 

model, these terranes moved southward with respect 

to continental North America until Late Triassic 

time; they moved northward during the Early and 

Middle Jurassic, southward during the Late Jurassic 

and Early Cretaceous, and finally northward again 

from the mid-Cretaceous onward. The timing of the 

shearing in the Wallowas is consistent with the first 

southward motion of the Wrangellia terrane, and the 

timing of the shearing in the Izee terrane fits the 

second southward translation proposed by Stone and
300 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

others (1982). Left-oblique convergence during the 

Late Jurassic to Early Cretaceous has also been invoked to explain the structural evolution of terranes 

elsewhere in the North American Cordillera (Oldow 

and others, 1984; Ave Lallemant and Oldow, 1988).

Paleomagnetic data from the Blue Mountains province (Hillhouse and others, 1982; Harbert and others, 

1988) suggest that in the Late Triassic the province 

was located farther south with respect to cratonic 

North America than at present and that it had to 

have moved northward again since the Late Triassic. 

This hypothesis is compatible with paleomagnetic 

data from several Cordilleran terranes which indicate that they were displaced northward since the 

mid-Cretaceous (see Beck, 1983; Irving and others, 

1985; Hagstrum and others, 1985; Umhoefer, 1987). 

These data, however, were derived from igneous intrusions in which paleohorizontal is not constrained. 

May and Butler (1986) rightfully pointed out that the 

paleomagnetic data were also consistent with a systematic tilt of the terranes not necessitating any latitudinal displacements. Recently, however, Irving and 

Thorkelson (1990) obtained new paleomagnetic data 

from layered sedimentary and volcanic rocks compatible with large northward displacements. These displacements are also in accord with models of absolute 

plate motions based on the hot-spot model as developed by Engebretson and others (1985). Furthermore, 

major northward displacements of the Stikinia and 

Quesnellia terranes in Late Mesozoic time are required to explain the occurrence at high latitudes of 

Jurassic tropical fauna (Tipper, 1981; Tozer, 1982; 

Taylor and others, 1984).

POST-D3 DEFORMATION

Sutter and others (1984), Lund and Snee (1988), 

and Snee and others (chap. 15, this volume) identified a major mid-Cretaceous to Late Cretaceous regional metamorphic event in west-central Idaho that 

may be related to the collision of the amalgamated 

Blue Mountains province with the North American 

craton. However, no pervasive structures related to 

the collision were identified in Oregon.

In eastern Oregon and western Idaho, the contours 

of initial Sr-isotope ratios typical of continental crust 

(87Sr/86Sr>0.706) and typical of oceanic crust (87Sr/ 

86Sr<0.704) are very close together (Armstrong and 

others, 1977; Fleck and Criss, 1985; Criss and Fleck, 

1987). In the same area, the Paleozoic miogeoclinal 

sequence, which elsewhere in the Cordillera drapes 

the North American continental margin, is missing. 

These two pieces of evidence strongly suggest that

the contact between the Blue Mountains province 

and the craton is a strike-slip boundary. Such a 

boundary may have been the result of escape tectonics (Wernicke, 1984; Klepacki and Wernicke, 1985; 

Wernicke and Klepacki, 1988) or oblique plate convergence (Ave Lallemant and Oldow, 1988; Oldow 

and others, 1989). However, no evidence for mid-Cretaceous transcurrent motion along this boundary has 

been obtained; such structures may have been destroyed by the pervasive dip slip that constitutes the 

latest movement on this boundary (Strayer and others, 1989), or they may have been buried or eroded.

CONCLUSIONS

The Blue Mountains province of northeastern Oregon consists of five terranes, four of which were discussed in this chapter. These are, from north to 

south, the Wallowa, Baker, Izee, and Olds Ferry terranes. The pre-Cretaceous tectonic evolution of the 

terranes (fig. 7.10) is not easy to decipher because so 

much of the area is covered by Cretaceous and 

younger rocks and because many more detailed studies are necessary to better constrain the geologic history. Nevertheless, the following tentative conclusions 

are drawn (note that all geographic directions mentioned refer to pre-Cretaceous coordinates):

(1) The Baker terrane consists of a Permian and 

Triassic fore-arc terrane/subduction complex related 

to eastward subduction underneath the west-facing 

Olds Ferry volcanic island arc; both intra- and forearc-basin deposits are included in the Izee terrane 

(Dickinson, 1979). The Olds Ferry terrane contains 

only Upper Triassic island-arc deposits, but may 

have been active from Permian to Middle Jurassic 

time if the Permian Canyon Mountain Complex is 

indeed part of the arc (Ave Lallemant, 1976) and if 

the Middle Jurassic volcanic rocks in the Izee terrane 

were derived from the Olds Ferry arc (Dickinson and 

Thayer, 1978). No melange of Jurassic age and no obvious Middle Jurassic structures related to subduction have been recognized.

(2) The simultaneous formation of north-southtrending folds and left-lateral strike-slip shear zones 

in the Baker terrane indicates that during the Late 

Triassic to Middle Jurassic, plate convergence was 

left-oblique, with the convergence vector partitioned 

into components perpendicular and parallel to the 

subduction zone (fig. 7.10A).

(3) The Wallowa terrane is a Permian and Triassic 

volcanic island arc. Magmatic activity virtually 

ceased in Late Triassic time. The polarity of the arc 

is unknown, but because the melanges of the Baker
7. PRE-CRETACEOUS TECTONIC EVOLUTION OF THE BLUE MOUNTAINS PROVINCE, NORTHEASTERN OREGON 301

terrane are probably all related to the Olds Ferry 

arc, it is proposed that the Wallowa arc was also 

west-facing (fig. 7.10A).

(4) North-south-striking left-lateral strike-slip shear 

zones in the Wallowa terrane, active in the Middle to 

Late Triassic, indicate left-oblique plate convergence. 

No coeval compressional structures were recognized 

however (fig. 7.10A).

(5) The Wallowa terrane collided with the Olds 

Ferry, Baker, and Izee terranes in the Late Jurassic. 

This collision caused east-directed thrusting and 

east-vergent folding in the Izee and Baker terranes 

and west-directed thrusting and west-vergent folding 

in the Wallowa terrane (fig. 7.10B).

(6) The mid-Cretaceous collision of the Blue Mountains province with continental North America 

caused no penetrative deformation of the rocks in Oregon, but intense deformation took place in Idaho 

(Sutter and others, 1984; Lund and Snee, 1988; 

Strayer and others, 1989; Snee and others, chap. 15, 

this volume).

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8. PRE-TERTIARY DEFORMATION IN THE 

DESOLATION BUTTE QUADRANGLE, NORTHEASTERN OREGON

By JAMES G. EVANS

CONTENTS

Page

Abstract 305

Introduction 306

Acknowledgments 306

Descriptions of the rock units 306

Elkhorn Ridge Argillite 306

Pyroclastic and volcanic rock assemblage 309 

Early Triassic diorite - 310 

Brecciated argillite and chert unit 311 

Peridotite and associated rocks assemblage 311 

Andesite - . 313

Late Jurassic diorite 313 

Quartz monzonite 313 

Tertiary and Quaternary rocks and deposits 313 

Genetic relations between the Elkhorn Ridge Argillite, the 

pyroclastic and volcanic rock assemblage, and the 

Early Triassic diorite - 313 

Metamorphism 314 

Structure 317

Macroscopic structure 317

Mesoscopic structure 317

Methods -~-~~~ ~--~~~~_~~ ~- ------------------- 317

Elkhorn Ridge Argillite 317

Pyroclastic and volcanic rock assemblage 319 

Relations among structures in the Elkhorn Ridge 

Argillite and the pyroclastic and volcanic rock 

assemblage 320 

Peridotite and associated rocks assemblage 320 

Relations among fabric elements in the Elkhorn Ridge 

Argillite and the peridotite and associated rocks 

assemblage 321 

Deformations 323 

Comparison of fabric in the Elkhorn Ridge Argillite,

Desolation Butte quadrangle, with fabric in other areas 324 

Discussion 326 

Conclusions 328

ABSTRACT

Pre-Jurassic rocks of the Desolation Butte quadrangle in northeastern Oregon comprise four main assemblages: the Elkhorn 

Ridge Argillite (Devonian to Triassic?), a pyroclastic and volcanic 

rock assemblage (Lower Triassic), a brecciated argillite and chert

unit (Upper Jurassic), and a peridotite and associated rocks assemblage (Late Triassic to Middle? Jurassic). The Elkhorn Ridge 

is most likely a tectonostratigraphic assemblage of sedimentary 

rocks, which, in the study area, is Paleozoic and Triassic(?) in 

age. The pyroclastic and volcanic rock assemblage may be partly 

contemporaneous with the Elkhorn Ridge Argillite or may even 

have been deposited on it as a developing andesitic arc. The brecciated argillite and chert may be largely equivalent to the 

Elkhorn Ridge Argillite but could contain additional sedimentary 

rocks emplaced by faulting during Late Jurassic time. The peridotite and associated rocks assemblage may represent a dismembered ophiolite-like assemblage that includes peridotite as well as 

gabbro, hornblende diorite, and metasedimentary and metavolcanic rocks.

Metamorphism of the Elkhorn Ridge Argillite and the pyroclastic and volcanic rock assemblage includes contact metamorphism adjacent to Triassic diorite intrusions and regional 

greenschist-facies metamorphism, which also affected the diorite. The peridotite and associated rocks assemblage underwent 

complex metamorphism in an ocean floor environment. Broad 

contact-metamorphic aureoles occur in all these rocks surrounding Late Jurassic diorite plutons. Narrow contact-metamorphic 

aureoles occur adjacent to Jurassic or Cretaceous quartz monzonite intrusions.

Deformations recorded in the rocks include pre-Mesozoic development of broken formation in the Elkhorn Ridge Argillite, Triassic imbrication of the Elkhorn Ridge Argillite, pyroclastic and 

volcanic rock assemblage, and Triassic diorite, and three later 

episodes of folding and cleavage development. The Late Triassic 

to Middle(?) Jurassic folding is the most widespread deformation. 

Folds formed during this event are subhorizontal and trend eastwest. Axial planes of the folds are parallel to a generally steep 

cleavage. This folding event may account for the principal deformations in the Elkhorn Ridge near Baker City and in the Burnt 

River Schist, a possible partial lithic correlative of the Elkhorn 

Ridge, and may be related to the Early Triassic to Middle(?) Jurassic accretion of the Baker terrane of the Blue Mountains. Emplacement of the peridotite and associated rocks assemblage in 

an extensional tear in the Elkhorn Ridge may have occurred 

during this event. Poles of cleavages in the Elkhorn Ridge and 

the strike of the tear may indicate the direction of principal 

shortening approximately east-west in pre-Cretaceous time. 

Later deformation of the peridotite and associated rocks assemblage moderately affected the Elkhorn Ridge nearby and could 

be related to Late Jurassic diorite emplacement that resulted in 

compression at right angles to the zone of peridotite and associated rocks and in widespread metamorphism that locally obliterated minor structures in the ultramafic matrix of the unit. Most 

deformation in the Baker terrane ceased after the Late Jurassic.

305
306 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Subsequent tectonism involved rotations of large blocks that may 

have included the Blue Mountains.

INTRODUCTION

The Desolation Butte 15-minute quadrangle, located 55 km west of the town of Baker City, Oreg. (fig. 

8.1), includes part of the melange terrane of the 

north-central Blue Mountains (Baker terrane of Silberling and others, 1984). This part of the Blue 

Mountains consists of pieces of oceanic and island arc 

crust and overlying sedimentary and volcanic rocks 

that have been greatly deformed before and during 

the Mesozoic, when the terrane was accreted to western North America (Brooks and others, 1976; Vallier 

and others, 1977; Brooks and Vallier, 1978; Dickinson 

and Thayer, 1978; Brooks, 1979). Stratigraphic and 

(or) tectonic relations between pieces of the oceanic 

crust and volcanic and sedimentary rock assemblages 

in the Blue Mountains are yet to be resolved. Stratigraphic relations between some geologic units may 

never be clarified because of the destruction of the 

evidence, presumably during subduction and pre-accretionary tectonic transport. The varied rocks in the 

Desolation Butte quadrangle, however, provide some 

clues to the geologic history of this belt of fragmented 

rocks.

The purpose of this chapter is to illustrate, using 

two geological vignettes, some of the details of the 

pre-Tertiary structure encountered in the Desolation 

Butte quadrangle (Evans, 1989) and to infer relations 

between the rock units. Some fabric features are revealed, although the data obtained during geologic 

mapping of the area are insufficient to characterize 

all aspects of the complex deformation these rocks 

have undergone.

The pre-Tertiary rocks of the Desolation Butte 

quadrangle, which are the chief geologic interest of 

this chapter, comprise four main assemblages that 

have had complex strain and thermal histories. Deformed and metamorphosed dioritic rocks, undeformed and metamorphosed andesite porphyry, and 

unaltered and unstrained diorite and granitic rocks 

intrude these assemblages (fig. 8.2). The main assemblages consist of the Paleozoic and Triassic(?) 

Elkhorn Ridge Argillite C&Rze) that is predominantly 

argillite and chert, a Lower Triassic pyroclastic and 

volcanic rock assemblage (~fcp), an Upper Jurassic 

brecciated argillite and chert (Jb), and a peridotite 

and associated rocks assemblage, including gabbro 

(J~fcg), hornblende metadiorite (J~fcm), and metamorphosed sedimentary and volcanic rocks (J~fcs). Some of 

the components of the peridotite and associated rocks

assemblage (fig. 8.2) may be as old as Paleozoic and 

as young as Triassic. However, in its present melange 

form, the unit is herein regarded as Late Triassic to 

Middle(?) Jurassic in age. Original relations between 

the Elkhorn Ridge, the pyroclastic and volcanic rock 

assemblage, and the peridotite and associated rocks 

assemblage are not known because they are in fault 

contact with one another (fig. 8.2). Several intrusions 

occur in the study area: metamorphosed and brecciated diorite ("fcd) of Early Triassic age intruding the 

Elkhorn Ridge and the pyroclastic and volcanic rock 

assemblage; metamorphosed but undeformed andesite (Ja) of Jurassic age intruding the Elkhorn Ridge 

and the Triassic diorite; fresh Late Jurassic diorite 

(Jd); and fresh Jurassic or Cretaceous quartz monzonite (KJqm) (figs. 8.2, 8.3, 8.4).

ACKNOWLEDGMENTS

I wish to acknowledge the ideas, comments, and 

suggestions contributed during the course of this 

study by H.C. Brooks and M.L. Ferns of the Oregon 

Department of Geology and Mineral Industries and 

by T.L. Vallier of the U.S. Geological Survey. Brooks 

and Ferns visited me in the field, showed me some of 

the relevant geology that they had mapped in adjoining areas, and in general shared with me their expertise in the geology of northeastern Oregon. Brooks, 

Ferns, and Vallier suggested many improvements to 

the manuscript.

DESCRIPTIONS OF THE ROCK UNITS

ELKHORN RIDGE ARGILLITE

Large parts of the Desolation Butte quadrangle 

(fig. 8.2) and of the area shown in figure 8.3 are underlain by the black to gray argillite and chert of 

the Elkhorn Ridge Argillite. This formation was 

named by Gilluly (1937, p. 14) for exposures on 

Elkhorn Ridge near Sumpter, 25 km east of the 

Desolation Butte area. The formation contains minor 

amounts of other rock types: metabasalt (flows in 

southern part of the adjacent Trout Meadows 7.5- 

minute quadrangle), metatuff, limestone, and conglomerate. The conglomerate contains angular to 

well-rounded clasts of chert, siliceous siltstone, 

quartzite, argillite, granitoid rocks, and felsic volcanic rocks. Rare pods of talc-carbonate rock in the formation may be metamorphosed dolomite or could be 

altered ultramafic rocks that were emplaced within 

the argillite by faulting.
8. PRE-TERTIARY DEFORMATION IN THE DESOLATION BUTTE QUADRANGLE, NORTHEASTERN OREGON 307

It is not clear whether the Elkhorn Ridge Argillite 

in the study area originally consisted of a section 

several thousand meters thick or whether the unit is 

composed of several unrelated sections assembled by 

accretionary processes. An unfaulted stratigraphic 

section is not preserved in the quadrangle, nor has 

the base of the formation been identified anywhere. 

Much of the Elkhorn Ridge has been pervasively deformed so that it is a broken formation (Hsii, 1968).

Ave Lallemant and others (1980) concluded that 

the Elkhorn Ridge Argillite was deposited in a volcanic-arc-back-arc-basin environment, which is at 

least superficially a reasonable conclusion considering the volcanic clasts in parts of the Elkhorn Ridge 

and the possibly large original thickness of the unit. 

Mullen (1985), however, concluded that a fore-arc setting is most consistent with the chemistry and petrography of the mafic metavolcanic rocks from the 

Baker terrane. The origin of the Elkhorn Ridge in a 

fore-arc setting would also be consistent with its 

complex internal stratigraphy. Bostwick and Nestell 

(1967) point out that Tethyan and American fusulinid 

faunas occur in separate pods of limestone in the 

Elkhorn Ridge. Dickinson (1979, p. 166) has interpreted these apparent faunal juxtapositions to imply 

tectonic juxtapositions of strata that were deposited 

far apart in the ancestral Pacific Ocean.

East of the Desolation Butte quadrangle, the 

Elkhorn Ridge Argillite has yielded Pennsylvanian 

and Permian fiisulinids and Early to Middle Triassic 

pentacrinids, hydrozoans, and hexacorals (Brooks 

and others, 1976; Evans, 1986, table 1; Evans, 1989). 

A conodont fragment from a 1-m-thick limestone bed 

within the quadrangle (locality J299, fig. 8.3) was 

identified as Polygnathus sp., making parts of the 

Elkhorn Ridge in this area as old as Middle to Late 

Devonian in age (Morris and Wardlaw, 1986). Although regionally the Elkhorn Ridge has a Paleozoic 

and Triassic overall age, in the Desolation Butte 

quadrangle and vicinity it may be no younger than 

Permian. The large age range of the Elkhorn Ridge 

(Middle Devonian to Late Triassic, or about 150 m.y.) 

and its chaotic stratigraphy are consistent with the 

concept that the Elkhorn Ridge may have been assembled from unrelated but similar sedimentary 

rocks during several plate collisions, a model that 

Brooks (1979, p. 74) proposed for the entire oceaniccrust terrane of northeastern Oregon.

119°00' 118°30' 118°00' 117°30' 117°00'

45°00' __ UMATJLLA^COUNTY | 

GRANT

44°30'

UNIONCOUNTYj WALLQWACpUNTY__ 

J~ BAKER COUNTY

B/y<ERCqUNTY_ 

_r-^ALHEUR COUNTY 'T-^ AREA OF MAP

15 30 45 KILOMETERS

15 30 MILES

FIGURE 8.1. Index map of northeastern Oregon, showing location of Desolation Butte (shaded) and other quadrangles mentioned in

text.
308 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

118°45' 

45°00'

118°30'

44°45'

-KJqm

5 KILOMETERS

FIGURE 8.2. Geologic sketch map of Desolation Butte quadrangle. Quaternary units not shown in this figure. Geology

modified from Evans (1989).
8. PRE-TERTIARY DEFORMATION IN THE DESOLATION BUTTE QUADRANGLE, NORTHEASTERN OREGON 309

PYROCLASTIC AND VOLCANIC ROCK ASSEMBLAGE

The pyroclastic and volcanic rock assemblage consists chiefly of tuff, some of which is clearly waterlaid, lapillistone (containing fragments of andesite, 

tuff, and chert), and andesite flows and flow breccia. 

The andesite appears to have been extruded in a submarine environment because of its proximity to 

water-laid tuffs, but no pillows are preserved. The assemblage also includes minor siltstone, limestone, 

conglomerate (clasts of chert, siltstone, and tuff), and 

argillite. In places limestone is interbedded with tuff.

The mostly angular fragments in pyroclastic and sedimentary rocks include andesite, tuff, siltstone, and 

chert. Estimated minimum thickness of this unit, 

taken from the south limb of the fold shown in cross 

section B-B' (fig. 8.5) is 850 m.

No fossils were found in the pyroclastic and volcanic rock assemblage. The unit is intruded by a diorite 

body dated as earliest Triassic (see section "Early 

Triassic Diorite"). The andesitic volcanic rocks of the 

pyroclastic and volcanic rock assemblage may be considered contemporaneous with the Early Triassic diorite if the diorite was the magmatic source of the

EXPLANATION FOR FIGURES 8.2 THROUGH 8.5

Alluvium and glacial deposits (Quaternary)

Qls Landslide deposits (Quaternary)

QTm Mudflow deposits and colluvium (Quaternary and

Volcanic and sedimentary rocks (Tertiary)

Th Basalt (Tertiary)

Ta Andesite (Tertiary)

KJqm Quartz monzonite (Cretaceous or Jurassic)

Jd Diorite (Late Jurassic)

Jb Brecciated argillite and chert unit (Late Jurassic) 

Locally includes serpentinite (sp)

Ja Andesite (Jurassic)

Peridotite and associated rocks assemblage (Middle? 

Jurassic to Late Triassic) Divided into:

80

85

Pyroclastic and volcanic rock assemblage 

(Early Triassic)

Diorite (Early Triassic)

Elkhom Ridge Argillite (Triassic? and Paleozoic)

Contact Dotted where concealed

Faults Dotted where concealed. Arrows in cross sections 

(fig. 8.5) show displacement

Steeply dipping

Thrust Sawteeth on upper plate

Strike and dip of beds

Inclined 

Vertical

Strike and dip of foliation

Strike and dip of joints 

Inclined 

Vertical

Hornblende metadiorite - 45 Bearing and plunge of lineation

Metamorphosed sedimentary and volcanic rocks A Locality J299

FIGURE 8.2. Continued.
310 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

volcanic rocks. In this chapter the pyroclastic and 

volcanic rock assemblage is tentatively assigned an 

Early Triassic age.

The pyroclastic and volcanic rock assemblage may 

be a facies of the volcanic assemblage that includes 

the Dixie Butte Meta-andesite of Brooks and others 

(1984) located 15 km south of the quadrangle. The 

pyroclastic and volcanic rock assemblage differs from 

the Dixie Butte rocks in that it contains predominantly pyroclastic rocks and subordinate andesite 

flows. Nevertheless, the pyroclastic and volcanic rock 

assemblage could be a distal facies of a volcanic center located closer to or underlying the meta-andesite.

EARLY TRIASSIC DIORITE

Early Triassic dioritic rocks intrude the Elkhorn 

Ridge Argillite and the pyroclastic and volcanic rock

assemblage; these dioritic rocks are metamorphosed 

to greenschist facies and brecciated. Preserved primary textures and minerals indicate that original 

compositions included diorite, quartz diorite, hornblende diorite, granodiorite, and quartz monzonite, 

but most samples are diorite. More than one episode 

of intrusion may have been involved, as suggested by 

quartz monzonite dikes cutting the diorite.

A metamorphosed and brecciated diorite in the 

northwest corner of the adjacent Granite 7.5-minute 

quadrangle (fig. 8.1) was dated as earliest Triassic (243 

Ma, N.W. Walker, in Brooks and others, 1982; Palmer, 

1983). This intrusion is continuous with a diorite that 

is shown in figure 8.2 and near the northeast corner of 

figure 8.3. Because of general lithologic and textural 

similarities, all of the metamorphosed and brecciated 

dioritic plutons are tentatively assigned an Early Triassic age. Such an age for this pluton would be consistent with the Devonian age for part of the Elkhorn

o 1

B

2 KILOMETERS

FIGURE 8.3. Geologic map of Rabbit Butte area. See figure 8.2 for location and explanation. Base from U.S. 

Geological Survey l:62,500-scale Desolation Butte quadrangle, 1950. Contour interval 80 ft. Geology from 

Evans (1989). See figure 8.5 for cross sections.
8. PRE-TERTIARY DEFORMATION IN THE DESOLATION BUTTE QUADRANGLE, NORTHEASTERN OREGON 311

Ridge Argillite in the study area and with other radiometric ages in the region (Walker, 1983).

BRECCIATED ARGILLITE AND CHERT UNIT

A zone of brecciated argillite and chert as much 

as a few hundred meters thick is technically interlayered between part of the Elkhorn Ridge Argillite 

and the pyroclastic and volcanic rock assemblage. 

The breccia contains blocks of relatively undeformed 

chert, in which bedding is well preserved, surrounded by intensely fractured argillite and chert, in 

which no mesoscopic sedimentary structures are preserved due to recrystallization and later brecciation. 

Most of these rocks are superficially like the argillite and chert of the Elkhorn Ridge and may well 

have been derived from it. It is possible that the 

brecciated argillite and chert unit was originally 

part of a slope basin and that the strata were subsequently thrust over the Elkhorn Ridge rocks. Such 

an origin may account for some of the rock types 

that are found in the brecciated argillite and chert 

unit but are absent in the Elkhorn Ridge such as a 

thin-bedded limestone lens, having outcrop dimensions of approximately 15 by 60 m, on the north 

slope of Rabbit Butte, and, near the mouth of Lake 

Creek, conglomerate containing rounded granitoid 

clasts (as much as 30 cm across) and sandstone. The 

brecciated argillite and chert unit includes serpentinite intrusions (protrusions?) (fig. 8.3, 8.4), which 

are common in some fore-arc environments 

(Bloomer, 1983; Taylor and Smoot, 1984; Fryer and 

others, 1985) and could have been emplaced in slope 

basin deposits. The brecciated argillite and chert 

unit also contains pods of Triassic(?) diorite. Part of 

the brecciated argillite and chert unit is white, suggesting bleaching of organic matter that is ordinarily abundant in the Elkhorn Ridge. Hematite and 

quartz veins are locally abundant in the brecciated 

argillite and chert unit.

The brecciated argillite and chert unit must be at 

least as old as Early Triassic because it contains 

blocks of Early Triassic diorite. A minimum age of 

Jurassic can be inferred from the occurrences of undeformed Jurassic andesite (see section "Andesite") 

in the area. This possible age range of Early Triassic to Jurassic brackets the Middle Triassic age of 

melange in the John Day area of central Oregon, 

135 km southwest of the Desolation Butte quadrangle (Dickinson and Thayer, 1978, p. 150; Dickinson, 

1979, p. 166). The brecciated argillite and chert 

unit, however, is tentatively assigned a Late Jurassic age because the brecciation appears to have

been caused by Jurassic thrusting (see section 

"Structure").

PERIDOTITE AND ASSOCIATED ROCKS ASSEMBLAGE

The assemblage consisting of peridotite and associated rocks is thrust over the Elkhorn Ridge Argillite 

and the brecciated argillite and chert unit, and it is 

truncated on the southwest and south by Jurassic intrusions (figs. 8.2 and 8.4). The peridotite of this assemblage has a complex history of cataclasis, 

serpentinization, alteration to talc, and recrystallization of serpentine to olivine. Xenoliths in the peridotite include gabbro, hornblende diorite, and metamorphosed igneous and sedimentary rocks. The 

gabbro of this unit, recognized by relict textures and 

minerals, is mostly altered to amphibolite. The associated hornblende diorite, which may have intruded 

the peridotite and the metasedimentary rocks, has 

been recrystallized to amphibolite and, in places, is 

mylonitized. Some of the diorite of this unit is compositionally and texturally like the Early Triassic diorite; most of it, however, has undergone amphibolitegrade metamorphism and mylonitization in contrast 

to the greenschist-facies metamorphism and brecciation of the Early Triassic diorite (see section "Early 

Triassic Diorite"). The metasedimentary rocks of this 

unit consist of argillite, chert, and conglomerate containing angular to rounded clasts of chert, siltstone, 

quartz, sandstone, and felsic volcanic rocks. These 

clasts could have come from a terrane like the one 

that supplied clasts to the Elkhorn Ridge, suggesting 

the possibility that these metasedimentary rocks are 

related to the Elkhorn Ridge. Quartzite clasts containing recrystallized quartz grains that are sharply 

truncated by clast boundaries indicate that the recrystallization occurred in the source terrane. Some 

of the metasedimentary rocks recrystallized under 

nearly lithostatic conditions, as indicated by the lack 

of metamorphic fabric, during greenschist- to amphibolite-grade metamorphism. Some of the metasedimentary rocks have well-developed cleavage. In some pods 

of metasedimentary rocks that are as much as 0.5 

km2 , strain intensity ranges from slightly warped bedding to isoclinally folded rocks in which bedding is 

transposed parallel to cleavage. Metaigneous rocks of 

this unit consist of amphibolite, derived from gabbro 

and hornblende diorite, and probable metavolcanic 

rocks (inferred from a few relict primary textures, 

high calcic plagioclase and amphibole contents, and 

low quartz and mica contents). A small amount of 

fault breccia along the contact between the peridotite 

and associated rocks assemblage and the brecciated
312 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

argillite and chert unit appears to be a tectonic mix 

of hornblende diorite and intensely brecciated chert 

and argillite. The clustering of small pods of metamorphosed igneous and sedimentary rocks in the peridotite suggests either that the pods are xenoliths 

in a peridotite intrusion (magmatic interpretation), or 

that they are pieces of a metamorphic, sedimentary, 

and igneous terrane dismembered during emplacement of a very ductile serpentinite mass (protrusion?).

One model for the peridotite and associated rocks 

assemblage would be a protrusion along a tear fault in 

an obducted plate, similar to the tectonic situation 

described for the Batinah melange in Oman (Robertson 

and Woodcock, 1983), with later structural complications (see section "Structure")- This geometry would 

explain why the peridotite and associated rocks assemblage is not regionally extensive. No evidence of sedimentary peridotite was found in the Desolation Butte

118°42' 118°38'

44° 

50'

44° 

46'

1 2 KILOMETERS

FIGURE 8.4. Geologic map of Lost Camp area. See figure 8.2 for location and explanation. 

Base from U.S. Geological Survey l:62,500-scale Desolation Butte quadrangle, 1950. Contour interval 80 ft. Geology from Evans (1989). See figure 8.5 for cross section.
8. PRE-TERTIARY DEFORMATION IN THE DESOLATION BUTTE QUADRANGLE, NORTHEASTERN OREGON 313

quadrangle to indicate that the protrusion might have 

reached the sea floor, but serpentine matrix melange 

in the Bates quadrangle (fig. 8.1) is overlain by conglomerate containing clasts from the melange (Brooks 

and others, 1983). According to this model, the peridotite and associated rocks assemblage could represent 

basement upon which the Elkhorn Ridge Argillite was 

deposited or assembled, and some of the metasedimentary rocks in the peridotite and associated rocks assemblage could be fragments of the Elkhorn Ridge 

engulfed by the rising serpentinite protrusion.

Age of the components of the peridotite and associated rocks assemblage could be older than Devonian if 

the unit represents an ophiolite upon which the 

Elkhorn Ridge Argillite was deposited. In the dismembered-ophiolite model, the peridotite and associated 

rocks assemblage is presumably Late Triassic to Middiet?) Jurassic in age, on the basis of the estimated 

time of emplacement of the ophiolite (see section 

"Structure"). The emplacement age (Late Triassic to 

Middle? Jurassic) of the peridotite and associated rocks 

assemblage is the age used for this unit in this chapter.

ANDESITE

Andesite intruded the Early Triassic diorite and 

the Elkhorn Ridge Argillite (fig. 8.2) and was subsequently metamorphosed. The andesite in the northern part of the Desolation Butte quadrangle contains 

less than 5 percent metamorphic amphiboles, which 

are tentatively identified as actinolite and richterite. 

Richterite replaces actinolite locally, but calcic plagioclase has not undergone intensive alteration to epidote. Crystallization of richterite could be due to 

sodium metasomatism (from seawater?) or due to redistribution of sodium in the rocks during recrystallization. An andesite dike in SE1/4 sec. 13, north of 

North Fork John Day River, has been much altered 

to white mica and contains less than 1 percent brown 

biotite. The unstrained appearance of these intrusions suggests that they are younger than the deformations associated with the tectonic emplacement of 

the Triassic host rocks. The andesite, however, must 

have been emplaced before the unaltered Jurassic or 

Cretaceous plutons (see sections "Late Jurassic Diorite" and "Quartz Monzonite"). The andesite is here 

tentatively given a Jurassic age.

LATE JURASSIC DIORITE

Fresh or slightly altered, undeformed, mostly dioritic rocks quartz diorite, tonalite, gabbro (Streckeisen, 1976) compose several plutons within the 

Desolation Butte quadrangle (figs. 8.2-8.5). These 

plutons may be related to the Bald Mountain batholith exposed to the east of the quadrangle. That batholith was dated at 144±17 Ma (by the whole-rock 

Rb/Sr isochron method) by Armstrong and others 

(1977, p. 400). Although this age spans the range 

from Late Jurassic to Early Cretaceous (161 to 127 

Ma), they postulated that the batholith was emplaced 

in the Late Jurassic but remained hot or was reheated during the Early Cretaceous. For this reason the 

diorite intrusions are here assigned a Late Jurassic 

emplacement age (161 Ma).

QUARTZ MONZONITE

Generally fresh quartz monzonite plutons and 

granite (Streckeisen, 1976) dikes intrude the Late 

Jurassic diorite and the Elkhorn Ridge Argillite (figs. 

8.1, 8.2). Zones of greisen as much as 15 m wide 

occur in some of the intrusions. The spatial association of most of the quartz monzonite bodies with the 

Late Jurassic diorite suggests that there may be a 

genetic relation between the two types of intrusions. 

The quartz monzonite bodies may be a later product 

of differentiation of the magma from which the diorite plutons were derived. The age of the quartz 

monzonite intrusions is presumably Jurassic or Cretaceous, but a Tertiary age cannot be ruled out. In 

this chapter, the quartz monzonite intrusions are assigned a Jurassic or Cretaceous age.

TERTIARY AND QUATERNARY ROCKS AND DEPOSITS

Tertiary volcanic and sedimentary rocks (units Ta, 

Tb, and TVS) and Tertiary and Quaternary colluvium, 

mudflow and landslide deposits, glacial moraine, fanglomerate, and alluvium (units QTm, Qls, and Qag) 

are shown on the geologic maps and cross sections 

(figs. 8.2-8.5) but will not be discussed here.

GENETIC RELATIONS BETWEEN

THE ELKHORN RIDGE ARGILLITE,

THE PYROCLASTIC AND

VOLCANIC ROCK ASSEMBLAGE, 

AND THE EARLY TRIASSIC DIORITE

Several lines of evidence suggest but do not prove 

that the Elkhorn Ridge Argillite and the pyroclastic 

and volcanic rock assemblage were deposited at sites 

that may not have been very far apart: (1) The units
314 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

are imbricated; (2) Early Triassic diorite intruded 

both units, possibly before the imbrication (see northern part of fig. 8.2); and (3) clasts in tuff and conglomerate of the pyroclastic and volcanic rock 

assemblage include chert that could have been derived from the Elkhorn Ridge. The similarity in composition between the Early Triassic diorite intrusions 

and the andesitic volcanic rocks in the pyroclastic 

and volcanic rock assemblage suggests that the intrusions represent the magmatic feeders for the andesitic rocks, making at least part of the pyroclastic 

and volcanic rock assemblage roughly the same age 

as the diorite. In this case the pyroclastic and volcanic rock assemblage would appear to be younger than 

the Elkhorn Ridge in the study area and could have 

been deposited on it.

METAMORPHISM

The rocks in the Desolation Butte quadrangle 

record several metamorphic events, which are listed 

in table 8.1. Some of these events may have occurred 

before the Elkhorn Ridge Argillite, the pyroclastic 

and volcanic rock assemblage, and the peridotite and 

associated rocks assemblage were juxtaposed by 

faulting, and many of the events may have occurred 

before these three units were accreted to the North 

American continent.

The earliest metamorphic event (Ml, table 8.1) for 

which there is evidence in the quadrangle is the recrystallization of quartzite in the Paleozoic(?) terrane 

that supplied clasts to the Elkhorn Ridge Argillite 

and possibly to the metasedimentary rocks in the peridotite and associated rocks assemblage.

The metamorphic history of the peridotite and associated rocks assemblage was already complex before the assemblage came into contact with the block 

composed of the Elkhorn Ridge Argillite and the pyroclastic and volcanic rock assemblage. Early contact 

metamorphism, (M2, table 8.1) probably to hornblende hornfels facies, affected sedimentary and volcanic rocks intruded by hornblende diorite. Evidence 

of this metamorphism consists of xenoliths of conglomerate (JTis) in the diorite, and dikes of the diorite in metasedimentary rocks (JTis) that have 

hornfels texture. Contact relations between the hornblende diorite and the peridotite are ambiguous. It is 

not clear from this evidence whether the peridotite 

underwent contact metamorphism from intrusion of 

the hornblende diorite or from another source, or 

whether the hornblende diorite was later engulfed by 

a peridotitic magma or engulfed tectonically by a serpentinized peridotite protrusion.

Gabbro and hornblende diorite in the peridotite and 

associated rocks assemblage have complex textural 

relations that could be the result of a single multistage 

event (M3, table 8.1) in which conditions of metamorphism temperature, pressure, and PH2O varied. 

The diorite in the southern part of the peridotite and 

associated rocks assemblage (fig. 8.4) is less deformed 

than are the diorite and gabbro close to the contact 

between the peridotite and associated rocks assemblage and the composite block of the Elkhorn Ridge 

Argillite and the pyroclastic and volcanic rock assemblage; the relatively undeformed diorite shows 

abundant primary igneous textures. This diorite and 

the amphibolites that developed from it show equilibrium relations between calcic plagioclase and actinolite and between actinolite and hornblende, and 

show alteration of primary ilmenite to sphene. In places, the plagioclase was altered to epidote and the 

hornblende to chlorite. Subsequently, hornblende crystallized from chlorite, and calcic plagioclase veins cut 

across the epidote. Annealing textures are developed 

in amphibolites that formed from the gabbro and diorite; recrystallized amphibole grains obscure a preexisting foliation. The annealing recrystallization 

could be due to contact metamorphism. The source 

of the heat could be the Late Jurassic diorite, although 

no firm conclusions regarding the heat source can 

be drawn. Late hydrothermal activity, indicated by 

quartz-chlorite veins in the gabbro and diorite, suggests that the Late Jurassic diorite was the heat source 

for the veins. However, the veins are restricted to 

the pods in the peridotite and thus are likely to be 

older than Jurassic.

A model of the kind of complex metamorphism described for the gabbro and diorite has been proposed 

for oceanic lithosphere at spreading centers (Bonatti 

and others, 1975; Elthon and Stern, 1978; Elthon, 

1981), in which contact and dynamic metamorphic 

processes and hydrothermal alteration operate. The 

metamorphosed gabbro and diorite may fit into the 

facies scheme proposed by Elthon and Stern (1978, 

p. 464) and Elthon (1981, p. 286). According to their 

scheme, the highest metamorphic grade attained by 

the gabbro and diorite in the study area was the 

lower actinolite facies. This facies assignment is suggested by the widespread presence of calcic plagioclase in equilibrium with actinolite and by the 

alteration of primary ilmenite to sphene. The lower 

actinolite facies is approximately equivalent in temperature range (525-600°C) to the lower part of the 

almandine amphibolite facies of regional metamorphism and the lower hornblende hornfels facies of 

contact metamorphism (Turner and Verhoogen, 1960, 

p. 550, 552) but probably was achieved under lower
8. PRE-TERTIARY DEFORMATION IN THE DESOLATION BUTTE QUADRANGLE, NORTHEASTERN OREGON 315

TABLE 8.1. Metamorphic and tectonic events in the Desolation Butte quadrangle

[M1-M7 and D1-D7, metamorphic and tectonic events, respectively, the numbers indicating chronological order from earliest to latest event; 

events, the subscript numbers indicating relatively decreasing concentrations of fold axes]

^E^, folding

Symbol

Metamorphic event

Description

Age

Symbol

Tectonic event

Description

M7-

M6-

Contact metamorphism (due to 

quartz monzonite intrusion) of 

Elkhorn Ridge Argillite and 

pyroclastic and volcanic rock 

assemblage.

Contact metamorphism (due to 

diorite intrusion) of Elkhorn Ridge 

Argillite, pyroclastic and volcanic 

rock assemblage, and possibly also 

the peridotite and associated rocks 

assemblage and andesite.

M5- Greenschist-facies metamorphism 

of Elkhorn Ridge Argillite, 

pyroclastic and volcanic rock 

assemblage, and diorite.

Jurassic or 

Cretaceous.

Late Jurassic 

(161 Ma).

Late Triassic to 

Middle(?) Jurassic.

Early Triassic to 

Middle(?) Triassic.

D7 Rotation, translation, and uplift of 

large blocks that include Blue 

Mountains.

D6 Steep faulting and minor thrusting. 

Thrusting of the peridotite and 

associated rocks assemblage over 

composite block of Elkhorn Ridge 

Argillite and pyroclastic and volcanic 

rock assemblage. Development of 

brecciated argillite and chert unit. 

Formation of cleavages and folds in 

the peridotite and associated rocks 

assemblage, and 62- and Bs-type 

folding in Elkhorn Ridge Argillite.

D5 B4 folding (minor), cleavage in

Elkhorn Ridge Argillite, rotation of 

Brgroup folds, emplacement of the 

peridotite and associated rocks 

assemblage in extensional tear.

D4 B! folding (minor).

D3 Imbrication of Elkhorn Ridge

Argillite, pyroclastic and volcanic 

rock assemblage, and diorite. 

Pervasive brecciation.

M4 Contact metamorphism (due to 

diorite intrusion).

M3 Actinolite-facies metamorphism 

within protolith of the peridotite 

and associated rocks assemblage.

M2 Contact metamorphism (due to 

hornblende diorite intrusion) 

within protolith of the peridotite 

and associated rocks assemblage.

Ml Recrystallization within source

terrane of Elkhorn Ridge Argillite 

and metasedimentary rocks of the 

peridotite and associated rocks 

assemblage.

Early Triassic.

Devonian to Triassic D2 Development of broken formation in 

Elkhorn Ridge Argillite.

Pre-Devonian or Dl 

Devonian.

Uplift and other tectonism affecting 

source terrane of Elkhorn Ridge 

Argillite and metasedimentary rocks 

of the peridotite and associated rocks 

assemblage.
316 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

pressures than would be typical for regional metamorphism (Bonatti and others, 1975, p. 67). Elthon 

(1981, p. 288) characterized the actinolite facies as 

generally undeformed, while amphibolites from oceanic fracture zones have complex strain histories. 

However, Bonatti and others (1975, p. 74) pointed out 

that dynamic metamorphism is likely during seafloor spreading and that evidence of recrystallization 

under directed stress was observed in metagabbros 

from the mid-Atlantic Ridge. Most likely the actinolite-facies metamorphism (M3, table 8.1) occurred before emplacement of the peridotite and associated 

rocks assemblage that is in contact with the composite block of the Elkhorn Ridge Argillite and pyroclastic and volcanic rock assemblage, and it could have 

occurred in the Paleozoic.

Early Triassic dioritic intrusions into this composite block were probably accompanied by contact metamorphism (M4, table 8.1); if present, the effects of 

such contact metamorphism would not be easily distinguished from later greenschist-facies metamorphism (M5, table 8.1) if the high-temperature parts 

of the aureoles had been faulted away. Greenschistfacies metamorphism (M5, table 8.1) affected the 

Elkhorn Ridge Argillite, the pyroclastic and volcanic 

rock assemblage, and the Early Triassic diorite in the 

Triassic, assuming that the peridotite and associated 

rocks assemblage and the composite block of the 

Elkhorn Ridge Argillite and pyroclastic and volcanic 

rock assemblage were already juxtaposed by the Middle Triassic. If pre-Triassic regional metamorphism 

occurred, it has not been distinguished from the later 

metamorphic events.

The peridotite component of the peridotite and associated rocks assemblage underwent a complex history of metamorphism. Stages in the development of 

the peridotite include recrystallization of intensely 

strained dunite, serpentinization, alteration to talccarbonate rock, and recrystallization of talc and serpentine to olivine. Some of these stages may correlate with the Devonian(?) contact metamorphism 

and the Devonian to Triassic regional metamorphism (M2 and M3, table 8.1). Other stages may 

postdate these metamorphic episodes (see below, this 

section), and still others may have been obliterated 

altogether.

Recrystallization of the peridotite obliterated most 

evidence of the processes of emplacement. The extreme fragmentation of the metasedimentary and 

metavolcanic rocks in the peridotite and associated 

rocks assemblage, however, suggests that the peridotite engulfed these rocks either as a peridotitic 

magma or as a cooler ductile serpentinite possibly 

during diapiric emplacement in the Late Triassic to

Middle(?) Jurassic. Evidence for hot emplacement of 

peridotite should include high-temperature metamorphic rinds on the inclusions in the peridotite. 

Pyroxene hornfels-facies mineral assemblages would 

likely characterize the rinds because the peridotite 

magma would have a temperature in excess of 

1,000°C (Wyllie, 1971) whereas the lower temperature limit of the pyroxene hornfels facies is less 

than 700°C (Turner and Verhoogen, 1960, p. 520). 

The abundance of hornblende and muscovite 

throughout the metasedimentary and metavolcanic 

rocks that are surrounded by peridotite strongly 

suggests that the peridotite was not emplaced at 

high temperature. High-grade metamorphic rinds 

could conceivably have been tectonically eroded off 

the inclusions during postemplacement deformations. However no vestige of any high-temperature 

metamorphic rind was found during this study. Superficially, at least, the evidence suggests that the 

peridotite was emplaced as a cooler plastic serpentinite mass.

Broad contact-metamorphic aureoles as much as 

1,200 m wide developed in the Elkhorn Ridge Argillite and the pyroclastic and volcanic rock assemblage 

in response to the Late Jurassic diorite intrusions 

(M6, table 8.1). The rocks in the aureoles in the 

northwestern and southeastern parts of the quadrangle attained hornblende hornfels-facies metamorphic 

grade. Intrusion of the Late Jurassic diorite also may 

have been responsible for some of the jackstraw-textured olivine in contact-metamorphosed peridotite of 

the peridotite and associated rocks assemblage. 

Greenschist- to amphibolite-grade metamorphism, 

some of it at least a thermal equivalent of the actinolite-facies metamorphism (M3, table 8.1) already discussed, accompanied development of cleavage in 

metasedimentary and metaigneous rocks of this assemblage (D6, table 8.1). The same metamorphic 

event (M6, table 8.1) may also have affected andesite 

in the north-central part of the quadrangle.

Contact-metamorphic aureoles up to 200 m wide 

occur in the composite block of the Elkhorn Ridge Argillite and the pyroclastic and volcanic rock assemblage around intrusions of Jurassic or Cretaceous 

quartz monzonite (M7, table 8.1). The rocks attain 

hornblende hornfels facies very close to the contacts 

with these intrusions. The outer parts of the aureoles 

have low-grade mineral assemblages indistinguishable from the low-grade regional-metamorphic assemblages in the host rocks but exhibit partial 

obliteration of sedimentary textures and bedding by 

recrystallization. No contact-metamorphic effects of 

the quartz monzonite intrusion on the Late Jurassic 

diorite were found.
8. PRE-TERTIARY DEFORMATION IN THE DESOLATION BUTTE QUADRANGLE, NORTHEASTERN OREGON 317

STRUCTURE

MACROSCOPIC STRUCTURE

The cross sections (fig. 8.5) illustrate the imbricate 

structure of the pre-Jurassic rocks in the Desolation 

Butte quadrangle. Arrows on southwest-dipping faults 

are shown to indicate presumed relative northward 

displacement of the hanging walls. Folds in cross sections A-A' and B-B' were drawn assuming concentric 

to subconcentric form and no overturning of beds (Busk, 

1929). The folds are not directly mappable because of 

the absence of marker beds in the Elkhorn Ridge Argillite and the pyroclastic and volcanic rock assemblage 

but are consistent in having their axial planes vertical 

or steeply dipping to the south and their subhorizontal 

axes trending generally west-southwest. The simplicity 

of the fold profiles is consistent with the absence of 

evidence of intricate folding on the mesoscopic scale 

and the uniform attitudes of bedding. The attitudes of 

the axial planes are consistent with assumed northward dip-slip of the upper plates of the thrusts. These 

relations, although resting on several assumptions, 

suggest that the latest movements on the thrusts between the Elkhorn Ridge Argillite and the pyroclastic 

and volcanic rock assemblage were synchronized with 

or postdated the major folding events.

The structure within the peridotite and associated 

rocks assemblage, shown in cross section C-C" 

(fig. 8.5), is highly conjectural. The outcrops of the 

metasedimentary and metavolcanic rocks in the assemblage suggest the outline of an isoclinal fold (fig. 

8.5) having a subhorizontal northwest-trending axis. 

Continuity of the folded rocks is interrupted by peridotite. The principal fold is drawn as an antiform; 

to the south a less clearly defined synform is drawn, 

its hinge zone shown dismembered by peridotite. 

Exact orientation of the axis of the isoclinal fold is 

not clear, but it may be plunging at a low angle to 

the northwest. Evidence for this fold geometry is very 

weak; this interpretation is presented as a possible 

example of the kind of structural complexity one 

might expect within the assemblage. An alternative 

interpretation suggests a similar style with local large 

noncylindroidal folds of lenses of metamorphic rocks 

that are disrupted by peridotite.

MESOSCOPIC STRUCTURE

METHODS

The orientations of bedding, cleavage, fold axes, 

crenulations, and axes of elongate chert nodules were

collected during geologic mapping in the Desolation 

Butte quadrangle. Although the data, especially from 

the peridotite and associated rocks assemblage, is 

scanty, it is enough to outline part of the strain history of the rocks within the quadrangle and to describe 

their mesoscopic fabric in a general way.

The structural data were plotted as points on the 

lower hemispheres of equal-area projections. Some 

structural data (w>25) were contoured on equal-area 

lower-hemisphere projections by a method originally 

described by Kamb (1959) whereby a variable counting area is used. The counting area is a function of 

the number of data points and the frequency of significant deviations from a uniform spatial distribution. Stereograms are contoured in intervals of 2o 

and 4o, where 2o is the expected number of data 

points within a counting area for a uniform distribution across the entire stereogram. A point density 

greater than 3o is considered to be a significant deviation from a random distribution, while pole-free 

areas indicate a possibly significant absence of data 

points.

ELKHORN RIDGE ARGILLITE

In much of the Elkhorn Ridge Argillite, relatively 

well developed shear zones with a nearly penetrative 

shearing foliation subparallel to discontinuous bedding 

in broken formation surround blocks as much as 100 

m long of little-sheared chert and argillite in which 

bedding is well preserved. Many of the unsheared 

blocks are elongate parallel to the shear foliation in the 

broken formation. In many places this shear foliation 

is difficult to distinguish from the bedding. Little quantitative information was gathered about this shearing 

deformation, but it may have caused most of the strain 

in the Elkhorn Ridge Argillite.

Bedding in the Elkhorn Ridge Argillite has a mean 

strike of N. 70° W. and dip of 52° SW. (rc=303; fig. 

8.6A). Fracture cleavage is developed locally; in places it obliterates bedding. Cleavage has an average 

strike of N. 75° E. and dip of 63° SE (n=56; fig. 8.6B). 

Poles to cleavage are spread along a great circle, the 

pole of which defines a B-axis plunging 62°, S. 50° E.

Nearly all mesoscopic folds have a concentric form, 

amplitudes as much as 60 cm, and wavelengths as 

much as 1 m; few, seen only in float, are tight, nearly 

isoclinal folds. In most outcrops the folds are not 

penetrative.

Axes of folds and bedding crenulations plunge in 

many directions, mostly in the southeast quadrant 

(fig. 8.6C). The folds cluster into four populations 

based on relative density of fold axes; in table 8.1
318 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

1000 1000

1000 feet = 305 meters

1000 feet = 305 meters

LEVEL 10QO feet = 3Q5 meters

FIGURE 8.5. Cross sections. Locations of A-A' and B-B' shown in figure 8.3; location of C-C' shown in figure 8.4. See figure 8.2 for

explanation.
8. PRE-TERTIARY DEFORMATION IN THE DESOLATION BUTTE QUADRANGLE, NORTHEASTERN OREGON 319

and in figures 8.6 and 8.11, these groups are represented by "B/ "B2 ," "B3 ," and "B4 ," ordered by decreasing concentrations of fold axes. The center of 

the group with the largest concentration of axes (B^ 

plunges 44°, S. 22° E. These folds appear to lie along 

a great circle, the pole of which defines a rotation 

axis, R, that plunges 47°, N. 10° W., 17° from the 

mean of poles to cleavage. No fold axes parallel to R 

were measured. Also, the B-axis of cleavage (fig. 8.6B) 

is 26° from B: and is on a part of the contour diagram for which concentrations of fold axes and crenulations are low (fig. 8.6C). Consequently, it seems 

unlikely that cleavage fanning was directly related to 

B^group folding. These relations suggest that the 

folds of the B: group were rotated during deformation 

that occurred at the same time as or later than the 

cleavage development, and it is possible that the

rotation involved differential slip along cleavage 

planes. 

The three other fold groups (represented by B2 , B3'

and B4 in fig. 8.6C) do not appear rotated and, for that 

reason, are considered to be younger than Br Macroscopic fold axes trend northwest, subparallel to B2 .

Three of the four localities containing elongate 

chert nodules have nodule axes subparallel to minor 

B^group folds. At the fourth locality the nodule axes 

plunge steeply, closely parallel to the trend of B3- 

group fold axes.

PYROCLASTIC AND VOLCANIC ROCK ASSEMBLAGE

Bedding in the pyroclastic and volcanic rock assemblage has a mean strike of N. 85° W. and a mean dip

B

FIGURE 8.6. Lower hemisphere, equal-area projections of fabrics for Elkhorn Ridge Argillite in eastern and central parts of Desolation Butte 

quadrangle. Contours at 2a, 4a, 6a, and 8a; dotted 

line surrounds pole-free area. See text for explanation of ^pntouring method. A, Bedding. 303 

points. JLb, mean pole to bedding, which strikes N. 

70° W. and dips 52° SW. B, Cleavage. 56 

points. JLC1, mean pole to cleavage, which strikes 

N. 75° E. and dips 63° SE., 6, B-axis, plunging 62°, 

S. 50° E., defined by great circle (dashed-line) distribution of poles to cleavage. C, Fold axes and axes 

of elongate chert nodules. 86 points. Bj, B^, B^, 

and B4, mean poles to groups of folds CBi~E4 

numbered according to relatively decreasing concentrations of fold axes). Bj^ plunges 44°, S. 22° 

E. Bg plunges 5°, S. 63° E. B3 plunges 72°, S. 25° 

W. B4 plunges 2°, due E. R, rotation axis of great 

circle (dashed line) distribution of Bj-group folds.
320 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

of 61° SW. (fig. 8.7A), and poles of beds appear to lie 

along a great circle, defining a fi-axis plunging 61°, S. 

6° W. The axes of the few minor folds measured in 

the assemblage plunge southeast (fig. 8.7B).

RELATIONS AMONG STRUCTURES IN

THE ELKHORN RIDGE ARGILLITE AND

THE PYROCLASTIC AND VOLCANIC ROCK ASSEMBLAGE

The mean attitudes of bedding in the Elkhorn 

Ridge Argillite and the pyroclastic and volcanic rock 

assemblage are 13° apart. This angular difference 

could be a measure of an original angular discordance between these units. The fi-axis of bedding in 

the pyroclastic and volcanic rock assemblage (fig. 

8.7A) is 22° from Brgroup folds in the Elkhorn Ridge 

Argillite (fig. 8.6C) and falls between the 4a and 60 

contours on figure 8.6C close enough to each other 

to suggest that the fi-axis of the pyroclastic and volcanic rock assemblage and the Brgroup folds of the 

Elkhorn Ridge Argillite are related. With allowances 

for some rotation of thrust slivers after juxtaposition, 

these combined similarities in orientations of structural elements in the two assemblages suggest that 

these units were imbricated before the earliest folding (Bx group). The greater development of mesoscopic folds in the Elkhorn Ridge may be due to the 

lesser competence of the unit. The greater disruption 

of bedding (broken formation) in the Elkhorn Ridge, 

however, may be due to deformation that preceded 

imbrication.

PERIDOTITE AND ASSOCIATED ROCKS ASSEMBLAGE

Foliation and cleavage are present but not common 

in the peridotite and associated rocks assemblage. 

Annealing recrystallization during Late Jurassic contact metamorphism (M6, table 8.1) obliterated much 

of the early structure that had developed in the assemblage, especially in the peridotite. Planar structures are best preserved in some of the 

metasedimentary and metavolcanic rocks of the assemblage. In these rocks cleavage coincides with 

banding and a penetrative foliation defined by concentrations and arrays of phyllosilicate minerals and 

amphibole. These structural elements are steeply 

dipping or vertical and strike northeast to north (fig. 

8.8A). Rotation of pods of these metasedimentary and 

metavolcanic rocks in the peridotitic matrix after 

cleavage had formed may account for some of the 

variation in attitude.

Planar structures are much more poorly preserved 

in the hornblende metadiorite and peridotite of the 

assemblage. In the metadiorite, cleavage mostly coincides with a foliation that is defined by arrays of long 

axes of amphiboles. Some of the foliation in finegrained amphibolites that developed from the metadiorite may be mimetic after a mylonitic foliation.

Cleavage is well preserved locally in serpentinite, 

especially near pods of metamorphic and igneous 

rocks. The cleavages are developed in narrow domains (0.5 mm) of serpentine minerals (chrysotile?) 

showing subparallel elongation; these narrow doB

FIGURE 8.7. Lower hemisphere, equal-area projections of fabrics for pyroclastic and volcanic rock assemblage. A, Bedding. 86 points. Contours at 2a, 4a, jto, and 12a; dotted line surrounds pole-free 

area. See text for explanation of contouring method. _L b, mean pole to bedding, which strikes N. 85° 

W. and dips 61° SW. B, B-axis, plunging 61°, S. 6° W., denned by great-circle (dashed-line) distribution 

of poles to bedding. B, Fold axes. 4 points.
8. PRE-TERTIARY DEFORMATION IN THE DESOLATION BUTTE QUADRANGLE, NORTHEASTERN OREGON 321

mains are separated by broader domains in which 

serpentine minerals, talc, and carbonate are randomly oriented. Before the annealing recrystallization to 

form massive olivinite, most of the peridotite may 

have had this kind of fabric.

Attitudes of cleavages in the peridotite and hornblende metadiorite are similar and are combined in 

figure 8.85. The cleavages are vertical to steeply dipping and strike north-northeast, similar to the cleavages in the metasedimentary and metavolcanic rocks 

(fig. 8.8A), and show less scatter than the cleavages 

in the metamorphic rocks.

A few crenulations and minor folds were found on 

cleavages and foliations in the peridotite and associated rocks assemblage (figs. 8.8C, 8.8.D). The axes of 

minor folds and crenulations in the metasedimentary and metavolcanic rocks of the assemblage 

plunge at various angles in a northwest-southeast 

zone across the orientation diagram in figure 8.8C. 

The few linear structures in serpentinite plunge in 

widely varying directions at moderate to steep angles (fig. 8.8.D). Although the linear elements in the 

peridotite and associated rocks assemblage do not 

clearly define a subfabric, they suggest the presence 

of sets of steep and gently plunging minor folds. 

These fold sets may correspond to the fold sets 

found in the Elkhorn Ridge Argillite (see subsection 

"Elkhorn Ridge Argillite" under "Mesoscopic Structure"). One group of minor folds in the peridotite 

and associated rocks assemblage appears to be subparallel to the northwest-trending axis of the major 

antiform outlined by metamorphic rocks (figs. 8.2, 

8.4, 8.5; cross section C-C").

The peridotite and associated rocks assemblage 

would probably have had to undergo a two-stage development to explain the extensional and compressional features and the annealing textures. If the 

assemblage developed as a serpentine-matrix melange protrusion into a plate composed of Elkhorn 

Ridge Argillite, it would have been emplaced in an 

environment of extension at right angles to the 

N. 40° W. trend of the zone occupied by the assemblage. Structures developed in the assemblage at the 

time of protrusion should reflect the probable upward movement of the assemblage into this large 

extensional fracture. The thrusting and mylonitization along the northeast margin of the assemblage 

and the cleavage developed in it, however, reflect 

compression, at a high angle to the trend of the 

zone, rather than extension. These structures most 

likely resulted from a later event, possibly associated with the emplacement of Late Jurassic diorite 

along the southwest margin of the zone. As the dioritic magma rose preferentially along this contact,

possibly to feed andesitic volcanism on the Late Jurassic surface, the peridotite and associated rocks 

assemblage was pushed aside, with the resulting development of the compressional structures. Heat 

from the diorite caused widespread recrystallization 

of deformed serpentinite to a massive olivinite and 

also affected the metamorphic rocks of the assemblage, including the mylonite. Most of the minor 

structures in the serpentinite were obliterated during recrystallization.

RELATIONS AMONG FABRIC ELEMENTS

IN THE ELKHORN RIDGE ARGILLITE AND

THE PERIDOTITE AND ASSOCIATED ROCKS ASSEMBLAGE

Part of the Elkhorn Ridge Argillite in the western 

part of the Desolation Butte quadrangle (fig. 8.2) 

has a subfabric that may bear on the relative timing 

of the development of fabric elements in the 

Elkhorn Ridge and in the peridotite and associated 

rocks assemblage. Cleavage there cuts intercalated 

chert and argillite. Beds are disrupted along cleavages, and cleavage-bound chert fragments are flattened perpendicular to cleavage. Long axes of chert 

fragments parallel minor folds and crenulations. Attitudes of cleavages vary widely (fig. 8.8.B), but 

many resemble those in the peridotite and associated rocks assemblage (figs. 8.8A, 8.8B) more than 

those in the Elkhorn Ridge Argillite in the eastern 

and central parts of the Desolation Butte quadrangle (fig. 8.6B). In the western part of the quadrangle 

minor folds and elongate chert nodules in the 

Elkhorn Ridge plunge steeply or gently (fig. 8.8F). 

Many of these linear elements have attitudes like 

the ones in the peridotite and associated rocks assemblage (figs. 8.8C, 8.8D). These similarities in fabric between the part of the Elkhorn Ridge in the 

western part of the quadrangle and the peridotite 

and associated rocks assemblage suggest that the 

fabric in the assemblage developed during or after 

juxtaposition of the assemblage and the composite 

block of the Elkhorn Ridge and the pyroclastic and 

volcanic rock assemblage. The steep and gently 

plunging folds in the peridotite and associated rocks 

assemblage (figs. 8.6C, 8.6.D) may correspond to the 

B2- and B3-group folds in the Elkhorn Ridge. If 

these correlations are correct, then B^ and B4-group 

folds in the Elkhorn Ridge most likely formed before 

the peridotite and associated rocks assemblage was 

juxtaposed with the composite block of the Elkhorn 

Ridge and the pyroclastic and volcanic rock assemblage, because folds with the B^ and B4-type orientations appear to be lacking in the fabric of the 

peridotite and associated rocks assemblage.
322 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

A B

C
8. PRE-TERTIARY DEFORMATION IN THE DESOLATION BUTTE QUADRANGLE, NORTHEASTERN OREGON 323

DEFORMATIONS

Rocks of the Desolation Butte quadrangle have undergone several deformations (Dl to D7, table 8.1). 

Like the metamorphic events, some of the deformations occurred before the peridotite and associated 

rocks assemblage and the composite block of the 

Elkhorn Ridge Argillite and the pyroclastic and volcanic rock assemblage were juxtaposed by faulting. Other 

deformations may have occurred jointly in both blocks 

before the rocks arrived in their present location.

The earliest deformation (Dl, table 8.1) recorded is 

uplift of the terrane that supplied sediment to the 

Elkhorn Ridge Argillite, to the pyroclastic and volcanic rock assemblage, and possibly to the metasedimentary rocks in the peridotite and associated rocks 

assemblage. Some of this uplift may date from preDevonian time.

The Elkhorn Ridge Argillite developed broken 

formation during Devonian to Triassic time (D2, 

table 1). This event probably occurred before the 

Elkhorn Ridge and the pyroclastic and volcanic rock 

assemblage were imbricated, because the assemblage 

does not have the appearance of broken formation. 

Alternatively, the absence of broken formation in the 

assemblage could be due to the greater competence of 

the assemblage, but this explanation would be more 

convincing if the pyroclastic and volcanic rock assemblage consisted entirely of volcanic flows or of flows 

and thick-bedded well-cemented volcaniclastic rocks. 

The thick zones of thin-bedded or laminated tuff and 

argillite in the assemblage most likely would not 

have remained intact if the development of a broken 

formation had occurred after the imbrication (D3, 

table 8.1). Possibly some of the Elkhorn Ridge was 

not completely lithified before the development of 

broken formation, whereas the pyroclastic and volcanic rocks assemblage had a greater bulk competence than the Elkhorn Ridge had during that event

FIGURE 8.8. Lower hemisphere, equal-area projections of fabrics 

for the peridotite and associated rocks assemblage and for 

Elkhorn Ridge Argillite in northwestern and western parts of 

Desolation Butte quadrangle. A, Poles to cleavage and foliation 

in metamorphic rocks of the peridotite and associated rocks assemblage (rc=14). B, Poles to cleavage in peridotite and associated hornblende metadiorite (rc=16). C, Fold axes in metamorphic 

rocks of the peridotite and associated rocks assemblage (n=l). 

Double-headed arrows show angles between intersecting fold 

axes. D, Fold axes in peridotite (n=3). E, Poles to cleavage in 

Elkhorn Ridge Argillite in northwestern part of quadrangle 

(rc=15). F, Fold axes and elongate chert nodules in Elkhorn 

Ridge Argillite in western part of quadrangle (rc=6). Dots, fold 

axes; squares, axes of elongate chert nodules.

(D2, table 8.1). If the Elkhorn Ridge (Paleozoic and 

Triassic?) within the study area is older than the pyroclastic and volcanic rock assemblage (Early Triassic), the Elkhorn Ridge was probably well lithified by 

Triassic time and had existed in the paleo-oceanic 

(Pacific) crust for about 125 m.y., long enough to have 

developed complex internal structure before imbricate thrusting with (or even before the deposition of) 

the pyroclastic and volcanic rock assemblage. The development of broken formation (D2, table 8.1) can be 

given a tentative age of Devonian to Triassic assuming that broken formation in the Elkhorn Ridge 

records a structurally distinct strain event that is not 

shown in other rocks of the study area.

Evidence of deformation of the peridotite and associated rocks assemblage, especially the peridotite, 

during the Triassic and perhaps even pre-Mesozoic 

time, was largely obliterated during subsequent recrystallization. If this assemblage was once an ophiolite, disruption of the original, lithologically varied 

stratigraphy has resulted in the present-day nearly 

chaotic outcrop pattern of lithologic components.

Early Triassic to Middle(?) Jurassic imbrication of 

the Elkhorn Ridge Argillite, the pyroclastic and volcanic rock assemblage, and the Triassic diorite (D3, table 

8.1) occurred after intrusion of the Triassic diorite.

Mesoscopic folds (B15 fig. 8.6C and table 8.1; D4, 

table 8.1) appear to have developed during or after 

imbrication of the Elkhorn Ridge Argillite and the 

pyroclastic and volcanic rock assemblage, on the 

basis of similarities between the fabric of the assemblage and a subfabric of the Elkhorn Ridge, especially the orientations of certain linear elements (fig. 

8.9A). The Enfolding event (D4, table 8.1) probably 

occurred during the interval from the Early Triassic 

to the Middle(?) Triassic.

Southeast-dipping cleavage developed in the 

Elkhorn Ridge during an event (D5, table 8.1) that 

may also have resulted in rotation of Bj-group folds. 

Minor B4-group folds as well as some macroscopic 

folds (such as are shown in cross sections A-A and 

B-B', fig. 8.5) may have formed during this same 

event. The peridotite and associated rocks assemblage is presumed to have been emplaced at about 

this time (Late Triassic to Middle? Jurassic).

During a later event (D6, table 8.1), the peridotite 

and associated rocks assemblage was thrust over the 

composite block of the Elkhorn Ridge Argillite and 

pyroclastic and volcanic rock assemblage, and the 

brecciated argillite and chert unit formed and was 

thrust into contact with part of the composite unit. 

At this time new cleavage formed in serpentinite and 

metamorphic rocks of the peridotite and associated 

rocks assemblage, and mylonite zones developed in
324 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

the hornblende diorite in response to emplacement of 

the Late Jurassic diorite.

B2- and B3-group folds in the Elkhorn Ridge Argillite, folds in the peridotite and associated rocks 

assemblage, and steeply dipping to vertical northwest-striking cleavage and foliation in the peridotite 

and associated rocks assemblage and in part of the 

Elkhorn Ridge may have formed when the peridotite 

and associated rocks assemblage was thrust over the 

composite block of the Elkhorn Ridge Argillite and 

pyroclastic and volcanic rock assemblage (D6, table 

8.1). It is not clear whether these structures formed 

during the same 161-Ma tectonic events, dated by 

Oldow and others (1984, p. 214) that resulted in penetrative deformation in the Blue Mountains. They 

suggested that three generalized episodes of Mesozoic 

deformation affected the Blue Mountains: Triassic to 

Early Jurassic, Late Jurassic, and Cretaceous. The 

two earlier deformations were penetrative and may 

correspond to deformations D5 and D6 (table 8.1). No 

evidence was found for Cretaceous deformation in the 

Desolation Butte quadrangle.

Steep northwest- and northeast-striking faults cut 

the imbricated assemblages the peridotite and associated rocks assemblage and the (composite) Elkhorn

Ridge Argillite-pyroclastic and volcanic rock assemblage-brecciated argillite and chert unit before emplacement of the Late Jurassic diorite. These faults 

could be tear faults in thrust sheets developed during 

imbrication (D3, table 8.1), but movement on some of 

them may have continued until long after D6 (table 8.1).

Ave Lallemant and others (1980, p. 374) concluded 

that the Blue Mountains may have been accreted to 

the continent in the Late Jurassic. If so, D6 may be 

related to accretionary processes.

Post-accretionary deformation in the Jurassic or 

Cretaceous (D7, table 8.1) included rotation, translation, and uplift of large blocks that include the Blue 

Mountains.

COMPARISON OF FABRIC IN

THE ELKHORN RIDGE ARGILLITE,

DESOLATION BUTTE QUADRANGLE,

WITH FABRIC IN OTHER AREAS

The fabric of the Elkhorn Ridge Argillite in the 

Desolation Butte quadrangle (hereinafter referred to 

as "Elkhorn Ridge Argillite-DB" for the sake of convenience) is similar to the fabric of the Elkhorn

A

iCI! pd

B

FIGURE 8.9. Lower hemisphere equal-area^ summary projections for assemblages in Desolation Butte 

quadrangle. A, Linear elements. Bj, B2, 63, and 84, mean orientations of fold groups from figure 

8.6C. R, rotation of axis of Bj-group folds, fi^, 6-axis of poles to bedding, pyroclastic and volcanic 

rock assemblage; BCIE> fi-axis of poles to cleavage, Elkhorn Ridge Argillite; B^E, mean orientation 

of subhorizontal folds jn Elkhorn Ridge Argillite in western part of Desolation Butte quadrangle. 

B, Planar elements. J-b^, mean pole to bedding, Elkhorn Ridge Argillite. -Lbp, mean pole to bedding, 

pyroclastic and volcanic rock assemblage. -LCl^, mean pole to cleavage, Elkhorn Ridge Argillite. J-Clpd, mean pole to cleavage, peridotite and diorite. lClm, mean pole to cleavage, metasedimentary and metavolcanic rocks of the peridotite and associated rocks assemblage.
8. PRE-TERTIARY DEFORMATION IN THE DESOLATION BUTTE QUADRANGLE, NORTHEASTERN OREGON 325

Ridge Argillite in the Elkhorn Peak quadrangle 

(hereinafter referred to as "Elkhorn Ridge ArgilliteEP") west of Baker City (fig. 8.1; Coward, 1983) and 

to that of the Burnt River Schist (Ashley, 1966), a 

possible correlative of part of the Elkhorn Ridge Argillite (Gilluly, 1937, p. 13). The Elkhorn Ridge Argillite-EP and the Burnt River Schist have 

undergone compound, nearly coaxial deformations in 

which at least two similar cleavages, and two sets of 

subparallel fold axes formed. Orientations of corresponding fabric elements (bedding, cleavages, and 

fold axes) are similar for the two areas (fig. 8.10). 

The deformation in the Burnt River Schist seems to 

be more intense than in the Elkhorn Ridge ArgilliteEP, and bedding and cleavages in the Burnt River 

tend to have great-circle pole distributions defining 

13-axes.

Ages of the deformations in the Elkhorn Ridge Argillite-EP and the Burnt River Schist may be similar. 

Ashley (chap. 12, this volume) considers the age of 

deformation in the Burnt River Schist to be no older 

than Middle Triassic and no younger than Early CreFIGURE 8.10. Lower hemisphere, equal-area summary 

projection for fabrics in Desolation Butte quadrangle, Elkhorn Peak quadrangle, and Burnt River 

area. Fabric elements from Desolation Butte_ quadrangle: -L b^, pole tojiverage bedding; J.C1, pole 

to average cleavage; B4, average of group of eastwest-trending folds. Elements from Elkhorn Peak 

quadrangle: _L bgp, pole to average bedding; _L S IEP 

and J.S2EP, poles to average cleavages; B, average 

of minor-fold axes. Elements from Burnt River 

area: -Lbg, pole to average bedding; -LSjB 

and -LS2B. poles to average cleavages; F^ and F2, 

average minor-fold axes; %, JJ-axis of bedding; JJgi> 

13-axis of Sj^ cleavage; J3g2> 8-axis of S2 cleavage.

taceous. Coward (1983, p. 109-110) concluded that 

the main deformation in the Elkhorn Ridge ArgilliteEP occurred in Late Triassic to earliest Jurassic 

time, apparently on the basis of Late Triassic to 

Early Jurassic radiolarians from a locality several kilometers south of Sumpter (fig. 8.1). Brooks and 

Ferns (1986) reported Late Triassic (early Norian) 

conodonts in the Elkhorn Ridge Argillite-EP, indicating post-early Norian deformation. If the deformations in the Elkhorn Ridge Argillite-EP and Burnt 

River Schist are correlative, as suggested by similarities in rock fabric, the main deformation must have 

occurred in the Late Triassic or Jurassic and possibly 

no later than the Middle Jurassic.

The average attitude of bedding in the Elkhorn 

Ridge Argillite-DB is not close to that of either the 

Elkhorn Ridge Argillite-EP or the Burnt River Schist, 

although data points overlap considerably. The orientations in the subfabric that includes B4-group folds 

and cleavage in the Elkhorn Ridge Argillite-DB resemble those of folds and early cleavage in the 

Elkhorn Ridge Argillite-EP and the Burnt River 

Schist (fig. 8.10); this similarity further supports a 

Late Triassic or Jurassic age for the D5 deformation 

in the Elkhorn Ridge Argillite-DB. However, these 

fabric elements do not seem as strongly developed in 

the Elkhorn Ridge Argillite-DB as they are in the 

Elkhorn Ridge Argillite-EP

The B^group of folds in the Elkhorn Ridge Argillite-DB does not have counterparts in the fabrics of 

the Elkhorn Ridge Argillite-EP and the Burnt River 

Schist. Because the Elkhorn Ridge Argillite-DB may 

be much older (Devonian, in part) than the Elkhorn 

Ridge Argillite-EP and the Burnt River Schist, it may 

have undergone several deformations before the Paleozoic sedimentary rocks were gathered into the 

Elkhorn Ridge tectonostratigraphic package. If B4- 

group folds and correlative fabric elements in the 

Elkhorn Ridge Argillite-EP and the Burnt River 

Schist correspond to Late Triassic or Jurassic deformations, then El folding may have occurred in the 

Early to Middle(?) Triassic.

Folds with attitudes like those of the B2 and Bg 

groups are not common in the Elkhorn Ridge Argillite-EP and the Burnt River Schist. In the Elkhorn 

Ridge Argillite-DB, B2- and Bg-type folding may have 

been associated with the Late Jurassic diorite intrusion. Structures correlative with this event would 

probably be uncommon in the Elkhorn Ridge Argillite-EP and absent in the Burnt River Schist. Present 

evidence suggests that B2- and Bg-type folds and 

cleavage in the western part of the Elkhorn Ridge 

Argillite-DB and in the peridotite and associated 

rocks assemblage represent local strain.
326 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

DISCUSSION

The history of the rocks in the Desolation Butte 

quadrangle reveals the long time involved in the 

structural development of the rock units, about 

200 m.y. from the end of the Devonian to the intrusion of the Late Jurassic diorite. The structural and 

metamorphic complexity of these rocks is consistent 

with a history spanning several geologic periods and 

with assembly from a variety of rock types that may 

have originated in widely separated and different environments. More than one subduction-obduction 

event can be postulated to explain the heterogeneous 

nature of the rocks. It is not clear how much of the 

rock fabric is a result of accretionary processes or is 

due to later tectonism. Regional foliation, mesoscopic 

folds, and coaxially refolded folds, such as in the 

Elkhorn Ridge Argillite-EP and the Burnt River 

Schist, can all be produced in accretionary wedges 

(Moore and Wheeler, 1978; Moore and Karig, 1980; 

Speed, 1983; Cloos, 1984) as well as in nonobduction 

environments; the rock fabric by itself is unlikely to 

provide enough clues to the tectonic environment.

Geologic relations point to nonpenetrative deformation occurring before the earliest (Triassic) folding. 

With regard to the total strain in the Elkhorn Ridge 

Argillite, the broken formation that developed during 

the Devonian to Triassic in the Elkhorn Ridge (D2, 

table 8.1) may have been the most significant deformation. Unfortunately, it is also the least studied or 

understood of the deformations. During the Early to 

Middle(?) Triassic, the Elkhorn Ridge Argillite may 

have been uplifted and eroded to supply chert clasts 

to the pyroclastic and volcanic rock assemblage. Later, the Elkhorn Ridge, the pyroclastic and volcanic 

rock assemblage, and the Early Triassic diorite were 

imbricated (D3, table 8.1).

The nearly penetrative deformations (D4 through 

D6, table 8.1) occurred over the time interval Early 

Triassic to Late Jurassic in a series of intermittent 

events, the last of which may have been regionally 

unimportant.

Before drawing conclusions about structural history from the fabric data, the 60° clockwise post-Jurassic rotation postulated for the Blue Mountains 

(Wilson and Cox, 1980) should be removed from the 

fabric orientation. The pre-Cretaceous orientations of 

the main fabric features of the Desolation Butte 

quadrangle are shown in figure 8.11, after 60° counterclockwise rotation about a vertical axis.

The oldest and best represented group of folds (Bxtype; D4, table 8.1) plunged on the average approximately 45°, S. 78° E. during the Early to Middle(?) 

Triassic (fig. 8.11A); later (Late Triassic to Middle?

Jurassic) they were rotated during B4-type folding 

(D5, table 8.1). Orientations of B1-type axial surfaces 

as well as later linear elements within the study 

area are difficult to measure because the structures 

are mostly crenulations and open concentric folds in 

which the axial surfaces are not precisely determinable. The general impression is, however, that in preCretaceous time axial planes were at large angles to 

bedding and were vertical or steeply dipping and 

west-northwest striking. If axial planes of concentric 

folds are expected to be approximately perpendicular 

to the greatest total compressive strain in the rocks 

(Dieterich and Carter, 1969; Dieterich, 1969, 1970), 

then El folding may be due to shortening along an 

Early to Middle(?) Triassic axis oriented north-northeast-south-southwest.

If cleavage is a shortening phenomenon (Ramsay, 

1967, p. 180; Siddans, 1972, p. 219; Wood, 1974, 

p. 398-399; Hobbs and others, 1976, p. 252), then 

cleavages in the Elkhorn Ridge Argillite in the Desolation Butte and Elkhorn Peak quadrangles and in 

the Burnt River Schist give clues to the directions of 

principal shortening during Late Triassic to Late Jurassic deformations. During the Late Triassic to Middle(?) Jurassic tectonism (D5, table 8.1) in the 

Desolation Butte quadrangle, the direction of shortening was oriented approximately east-west (fig. 

8.11A). Equivalent shortening in the Elkhorn Peak 

quadrangle and the Burnt River area (fig. 8.1) was 

oriented about west-northwest (fig. 8.1LB). This 

cleavage-forming event is correlative with tectonic 

shortening described by Oldow and others (1984, Dl 

in fig. 2), whose data suggest that this deformation 

in northeastern Oregon was part of a regional tectonic episode that also affected the northern Sierra Nevada. Associated fold axes in the northern Sierra 

Nevada, however, range in orientation from gently to 

steeply plunging, indicating that the total movement 

picture there is complex and probably differs appreciably in detail from the movement picture in the 

Desolation Butte quadrangle-Elkhorn Peak quadrangle-Burnt River area.

If the peridotite and associated rocks assemblage 

was emplaced along a tear in an obducted plate, as 

occurred in the Batinah melange (Robertson and 

Woodcock, 1983), the long axis of the tear could be an 

extension fracture oriented at a low angle to the local 

movement direction of the plate. Corrected for postJurassic rotation, this direction would have been approximately N. 80° E. (fig. 8.11), which is close to the 

orientation of the Late Triassic to Middle(?) Jurassic 

shortening in the Desolation Butte quadrangleElkhorn Peak quadrangle-Burnt River area. During 

this event the ultramafic component of the peridotite
8. PRE-TERTIARY DEFORMATION IN THE DESOLATION BUTTE QUADRANGLE, NORTHEASTERN OREGON 327

and associated rocks assemblage was most likely 

largely serpentinite.

After its development, the tear occupied by the peridotite and associated rocks assemblage remained a 

zone of weakness that was used during intrusion of 

the Late Jurassic diorite. The greatest principal 

strain, oriented north-northwest in the Late Jurassic 

(fig. 8.7) and arising from the diorite's pushing aside 

the peridotite and associated rocks assemblage, resulted in a cleavage in the serpentinite and metasedimentary rocks of the assemblage and in mylonitization of the unit. In addition, the peridotite and associated rocks assemblage was thrust over the block 

composed of the Elkhorn Ridge Argillite and the pyroclastic and volcanic rock assemblage along the 

northern (Late Jurassic) border of the extensional 

tear zone (D6, table 8.1). Thrust faulting brought 

rocks (brecciated argillite and chert unit) that are 

somewhat different from the Elkhorn Ridge Argillite 

into the area and caused moderate nearly penetrative disruption of the earlier fabric of the Elkhorn 

Ridge Argillite in the western Desolation Butte quadrangle. The fabric resulting from the 161-Ma tectonism (D6, table 8.1) does not appear to be of regional 

extent like the fabric that was produced by Late Triassic to Middle(?) Jurassic events (D5, table 8.1), because the younger (D6-type) fabric is not developed 

in either the Elkhorn Peak quadrangle (Coward,

1983) or the Burnt River Schist (Ashley, 1966). Also, 

the orientations of axial planes of possibly geometrically and temporally equivalent folds in the northern 

Sierra Nevada (Oldow and others, 1984, fig. 3) are 

different from D6-type cleavages.

As the diorite was emplaced, possibly feeding andesitic extrusions on the Late Jurassic surface, heat 

and possibly fluids penetrated and generally recrystallized the peridotite and associated rocks; this 

widespread recrystallization nearly obliterated cleavage and other D6-type structural elements in the ultramafic matrix of the unit.

The presence of the peridotite and associated rocks 

assemblage and similar serpentine-matrix melanges 

in the Desolation Butte quadrangle and vicinity implies the existence, at least in the past, of an ophiolitic layer below the Elkhorn Ridge Argillite. One can 

only speculate whether the ophiolitic layer was basement rock on which the Elkhorn Ridge was deposited 

or whether the peridotite and associated rocks assemblage was part of a regional ultramafic plate over 

which the Elkhorn Ridge was tectonically emplaced.

Dickinson (1981) showed northeastern Oregon as 

an extension of the Late Jurassic to middle Tertiary 

zone of the Franciscan Complex. This view is partly 

consistent with the Late Jurassic deformational ages 

inferred by Ave Lallemant and others (1980) and 

Coward (1983) for the Blue Mountains. However,

-LCI

A

FIGURE 8.11. Lower hemisphere, equal-area summary projections for fabrics of rocks in Desolation Butte 

quadrangle, Elkhorn Peak quadrangle, and Burnt River area. Fabric rotated 60° counterclockwise about 

a vertical axis to remove post-Jurassic rotation of Blue Mountains. PAR, peridotite and associated rocks 

assemblage, possibly emplaced along tear in an obducted plate (see section "Discussion"). See figures 8,9 

and 8.10 for explanation of symbols. A, Desolation Butte quadrangle. B, Fabric resulting from 161-Ma 

tectonism (D6, table 8.1) for area comprising the Desolation Butte quadrangle-Elkhorn Peak quadrangleBurnt River area.
328 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

penetrative deformation in the Desolation Butte 

quadrangle and possibly in other parts of the oceanic terrane of northeastern Oregon must have largely 

ceased after the intrusion of the Bald Mountain 

batholith and its satellitic plutons in the early Late 

Jurassic, about 161 Ma (Armstrong and others, 

1977). By this time the accreting oceanic terrane 

had ceased deforming internally, but may not yet 

have firmly docked at the continental margin. PostJurassic deformations involved rotation, translation, 

and uplift of large blocks that include the Blue 

Mountains.

CONCLUSIONS

1. Pre-Jurassic rocks of the Desolation Butte quadrangle are divided into four main assemblages: the Elkhorn Ridge Argillite, a pyroclastic 

and volcanic rock assemblage, a brecciated argillite and chert unit, and a peridotite and associated rocks assemblage.

2. The Elkhorn Ridge Argillite contains Devonian fossils in the Desolation Butte quadrangle but 

ranges to Late Trias sic (and may possibly 

range to Early Jurassic) age in areas to the 

south and east. It is apparently a tectonostratigraphic unit with a 200-m.y history and 

was assembled from unrelated packets of oceanic sedimentary and volcanic rocks.

3. The pyroclastic and volcanic rock assemblage, presumably Early Triassic in age, may represent 

an andesitic-arc environment that was not far 

from or was on top of the Elkhorn Ridge Argillite. The Triassic dioritic rocks in the area 

may have fed the vents from which the extrusive rocks originated.

4. The brecciated argillite and chert unit may consist 

mostly of brecciated fragments derived from 

the Elkhorn Ridge Argillite, although sedimentary rocks derived from other sources, including fore-arc basins, may have been emplaced 

by thrust faulting during Late Jurassic time.

5. The peridotite and associated rocks assemblage, 

possibly in part a dismembered ophiolite, contains rocks that may have undergone complex 

metamorphism in an oceanic crustal environment before juxtaposition with the composite 

block of Elkhorn Ridge Argillite, pyroclastic 

and volcanic rock assemblage, and brecciated 

argillite and chert unit. Part of the metasedimentary rock component of the peridotite and 

associated rocks assemblage may be related to 

the Elkhorn Ridge Argillite. The emplacement

of the peridotite and associated rocks assemblage may have been by local protrusion from 

a peridotite-rich basement during the Late 

Triassic to Middle(?) Jurassic.

6. Pre-Jurassic metamorphism in the study area includes low-grade regional and contact metamorphism of the Elkhorn Ridge Argillite, the 

pyroclastic and volcanic rock assemblage, and 

the Triassic diorite, as well as complex metamorphism of the peridotite and associated 

rocks assemblage. All these rocks were affected by widespread contact metamorphism as a 

result of intrusion of the Late Jurassic diorite 

and minor contact metamorphism around intrusions of Jurassic or Cretaceous quartz 

monzonite.

7. Tectonism in the study area includes possibly Paleozoic development of broken formation, Triassic thrusting, Triassic and Jurassic folding 

and cleavage-formation, and Late Jurassic 

thrusting and folding caused by emplacement 

of a diorite pluton. Most deformation within 

the Baker terrane ceased by the end of the 

Jurassic. Subsequent tectonism involved 

movements of large crustal blocks that may 

have included the rest of the Blue Mountains.

8. The most widespread deformation in the oceanic 

terrane in the Desolation Butte quadrangleElkhorn Peak quadrangle-Burnt River area 

seems to be east-west folding with subhorizontal axes and a steep cleavage developed 

approximately parallel to axial planes of the 

folds. Structures related to this deformation 

are less developed in the Desolation Butte 

quadrangle than in areas to the east and are 

most abundant in the Burnt River Schist, a 

possible partial lithic correlative of the 

Elkhorn Ridge Argillite. This deformation 

may be a result of Early Triassic to Middle(?) 

Jurassic oceanic crust-island arc convergence 

with its resultant obduction-subduction. The 

peridotite and associated rocks assemblage 

may have been emplaced as a protrusion 

along an extensional tear in the obducted 

plate. The orientation of pre-Cretaceous shortening was approximately east-west.

9. Deformation in the Desolation Butte quadrangle 

occurred during the Late Jurassic intrusion of 

diorite alongside the peridotite and associated 

rocks assemblage. Emplacement of the diorite 

forced aside the peridotite and associated 

rocks assemblage, resulting in northward 

(pre-Cretaceous) thrusting of this unit, development of steep east-west cleavage in the ser-
8. PRE-TERTIARY DEFORMATION IN THE DESOLATION BUTTE QUADRANGLE, NORTHEASTERN OREGON 329

pentinite-matrix melange, mylonitization of 

mafic igneous rocks in this unit, and moderate 

disruption of the earlier fabric of part of the 

Elkhorn Ridge Argillite. Heat and possibly fluids from the intrusion later resulted in widespread recrystallization of the peridotite and 

associated rocks assemblage.

10. The rocks of the Desolation Butte quadrangle 

were rotated, translated, and uplifted in the 

Jurassic and (or) Cretaceous.

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9. THE BOURNE AND GREENHORN SUBTERRANES OF 

THE BAKER TERRANE, NORTHEASTERN OREGON:

IMPLICATIONS FOR THE EVOLUTION OF 

THE BLUE MOUNTAINS ISLAND-ARC SYSTEM

By MARK L. FERNS! and HOWARD C. BROOKS1 -2

CONTENTS

Abstract 331

Introduction 331

Bourne subterrane 334

Elkhorn Ridge Argillite 335

Goodrich Creek and North Fork units 337

Olive Creek unit 339

Mixed-rock zones 340

Greenhorn subterrane 342

Badger Creek unit 342

Melange of Vinegar Hill 343

Dixie Butte Meta-andesite 346

Structure 347

Planar elements 347

Folds 349

Structure summary 350

Summary 350

Discussion 353

ABSTRACT

The central part of the Baker terrane in northeastern Oregon is 

made up of two lithologically distinctive belts of late Paleozoic and 

early Mesozoic age. The northernmost belt, herein informally referred to as the Bourne subterrane, consists of slabs and blocks of 

argillaceous broken formation (Elkhorn Ridge Argillite) that are 

tectonically intercalated with stratally disrupted slabs and blocks 

of island-arc rocks and oceanic rocks and chaotic mixed-rock zones. 

The Elkhorn Ridge Argillite contains limestone olistoliths of Devonian, Pennsylvanian, and Permian age, and ribbon cherts of Permian and Late Triassic (Karnian) age. Permian fusulinids in 

limestone blocks are mostly of Tethyan affinity. The association of 

presumably deepwater chert and siliceous argillite with alkalic pillow basalt and ophiolitic mixed-rock zones suggests ocean-floor 

(nonvolcanic arc) origins for parts of the Bourne subterrane. Intercalated low-K diorite and quartz diorite intrusions, basalt and andesite flows and breccias, tuffs, epiclastic sandstones, and Late

1Oregon Dept. of Geology and Mineral Industries, Baker City Field Office, 1831 First 

Street, Baker City, OR 97814

2Present address: 950 llth Street, Baker City, OR 97814

Triassic (Norian) bedded limestones are interpreted to be volcanicarc fragments. The Bourne subterrane includes severely deformed 

zones similar in appearance to type II and type IV melanges described by Cowan (1985). Deformational style and rock types are 

consistent with an accretionary wedge setting, as suggested by 

Coward (1983).

The southern belt, herein informally named the Greenhorn subterrane, is composed mainly of serpentinite-matrix melange that 

resembles the type III melange described by Cowan (1985). Other 

components include late Paleozoic clastic sedimentary rocks and 

basaltic andesite and andesite flows and tuffs. Blocks incorporated 

within the melanges include both oceanic and island-arc volcanic 

rocks as well as epidote amphibolites that contain blue amphibole. 

Permian fusulinids are similar to fusulinids from the Grindstone 

terrane to the south and the McCloud area in the Klamath Mountains of northern California. The close association of serpentinitematrix melanges with coarse clastic sedimentary rocks suggests 

that the Greenhorn subterrane may be part of the outer high of a 

fore-arc basin as suggested by Mullen (1985).

Lithodemic units in both assemblages are cut by a prominent 

south-dipping fracture cleavage that apparently developed in conjunction with high-angle reverse and thrust faults. Both the fracture cleavage and the faults appear to have formed during an 

intense, short-lived period of penetrative deformation that was followed by post-orogenic emplacement of Late Jurassic batholiths 

and stocks.

INTRODUCTION

The Baker terrane is an eastward-tapering tectonostratigraphic terrane (fig. 9.1) that separates the 

Grindstone, Izee, and Olds Ferry terranes on the south 

from the Wallowa terrane on the northeast (Silberling 

and others, 1984). The Baker terrane includes extensively disrupted late Paleozoic and early Mesozoic 

ocean-floor (non-arc) and island-arc material that separates the Permian and Late Triassic island-arc volcanic rocks (Brooks and others, 1976; Brooks and 

Vallier, 1978) of the Wallowa terrane on the northeast 

from the Late Triassic volcanic rocks and Late Triassic 

to Late Jurassic clastic sedimentary rocks of the Olds 

Ferry and Izee terranes on the south (Silberling and

331
332 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

others, 1984). Previous workers suggested that the 

Baker terrane is part of a paleosuture zone between 

two separate and distinct volcanic arcs now represented by the Wallowa and Olds Ferry terranes (Dickinson, 1979; Hietanen, 1981; Mortimer, 1986; Blome and 

Nestell, 1991). These workers generally agree that the 

Grindstone, Olds Ferry, and Izee terranes represent 

part of an island-arc system that bordered the continental margin, and that the Baker and Wallowa terranes are more "exotic" ocean-floor and island-arc

material initially formed at some distance from the 

continent. Other workers (including Vallier, chap. 3, 

this volume) suggest that the Baker terrane is part of 

a marginal basin developed within a single complex 

island-arc system.

Detailed mapping (fig. 9.2; see Mullen, 1978; Brooks, 

Ferns, Coward, and others, 1982; Brooks, Ferns, and 

Mullen, 1982; Ferns and others, 1987; and Evans, 

1988), coupled with geochemical analyses of volcanic 

rocks (samples and locations listed in table 9.1), faunal

119° 118° 117°

44

FIGURE 9.1. Pre-Tertiary terranes in the Blue Mountains province. Terrane names are modified from Silberling and others 

(1984). Dotted lines, queried where uncertain, are terrane 

boundaries modified from Silberling and others (1984). Line

of x's, queried where uncertain, marks boundary between the 

Bourne and Greenhorn subterranes (of the Baker terrane) discussed in text. Solid lines outline areas of exposure of pre-Tertiary rocks.
9. BOURNE AND GREENHORN SUBTERRANES OF THE BAKER TERRANE, NORTHEASTERN OREGON 333

TABLE 9.1. Location of samples used for geochemical analyses

Sample

EP-105

B-125

B-57

EP-89

EP-102

EP-38

B-//O'7 

EP-52

B-90

EP-48

B-261

B-259A

B-259B

153

B-289

77

MB-45

BB-133

B-133A

B-135B

74

B-364

55

V-l

V-2

V-3

60A

60B

63A

B-372

185

DBV

VDB-1

VDB-2

Subterrane

Bourne 

Bourne -

Bourne -

Bourne ~

Bourne -

Bourne 

Bourne 

Bourne -

Bourne -

Bourne 

Bourne 

Bourne 

Bourne 

Bourne 

Bourne 

Bourne 

Bourne 

Bourne -

Bourne 

Bourne -

Bourne 

Greenhorn -

Greenhorn -

Greenhorn -

Greenhorn -

Greenhorn -

Greenhorn -

Greenhorn -

Greenhorn -

Greenhorn -

Greenhorn -

Greenhorn -

Greenhorn -

Greenhorn -

Unit

Goodrich Creek unit 

Goodrich Creek unit 

Goodrich Creek unit 

Goodrich Creek unit 

Goodrich Creek unit 

Goodrich Creek unit 

Goodrich Creek unit 

Goodrich Creek unit 

Goodrich Creek unit 

Goodrich Creek unit 

North Fork unit 

North Fork unit 

North Fork unit 

North Fork unit 

North Fork unit 

McCully Fork mixed-rock zone

McCully Fork mixed-rock zone

McCully Fork mixed-rock zone

McCully Fork mixed-rock zone

McCully Fork mixed-rock zone

Olive Creek unit 

Melange of Vinegar Hill 

Melange of Vinegar Hill 

Melange of Vinegar Hill 

Melange of Vinegar Hill 

Melange of Vinegar Hill 

Melange of Vinegar Hill 

Melange of Vinegar Hill 

Melange of Vinegar Hill 

Melange of Vinegar Hill 

Dixie Butte Meta-andesite -

Dixie Butte Meta-andesite 

Dixie Butte Meta-andesite 

Dixie Butte Meta-andesite 

174

Tvn? JNJi/

NE

NE

NE

NW

NE

O1I7 ow

SE

NW

sw

sw

SE

SE

NE

NE

SW

NE

SE

SE

SE

SW

NW

SW

sw

sw

sw

sw

NW

SW

sw

NW

NW

NE

NE

174

NW

NW

SE

NE

NW

SW

XTT? iNEj

SE

NW

NE

SW

SW

sw

NW

SE

SW

NE

NW

NW

SE

SE

NE

SW

NE

NE

NE

NE

NE

SW

SW

NW

NW

SW

sw

Section

90Zo

33

33

10

19

22

&o 90

30

1

28

15

15

15

22

16

5

28

29

29

24

34

5

5

12

12

12

16

16

9

12

35

33

28

28

Township

o q o O.

8S.

8S.

9S.

9S.

9S.

Q S5 y o.

8S.

9S.

9S.

8S.

8S.

8S.

8S.

8S.

9S.

9S.

10 S.

10 S.

10 S.

9S.

10 S.

10 S.

10 S.

10 S.

10 S.

10 S.

10 S.

10 S.

10 S.

11,3.

us.

us.

us.

Range

QQ IT1 OO EJ.

38 E.

38 E.

38 E.

38 E.

38 E.

00 TT> OO EJ.

38 E.

38 E.

38 E.

35 E.

35 E.

35 E.

35 E.

35 E.

36 E.

37 E.

39 E.

39 E.

38 E.

35 E.

35 E.

35 E.

34 E.

34 E.

34 E.

35 E.

35 E.

35 E.

35 E.

33 E.

34 E.

34 E.

34 E.

Elevation 

(ft)

C OCA o,ot>U

6,600

7,000

6,500

5,500

7,700

7,400

8,700

4,400

7,000

6,000

5,760

5,760

5,440

5,900

5,800

5,000

3,840

3,840

4,160

5,600

6,320

7,320

8,000

8,000

8,000

6,200

6,400

6,800

5,600

4,900

6,700

7,400

7,400

Quadrangle

Elkhorn Peak

Elkhorn Peak

Elkhorn Peak

Elkhorn Peak

Elkhorn Peak

Elkhorn Peak

T?llrVir»vr» Poalr .CjlKnUiil t cclK

Elkhorn Peak

Elkhorn Peak

Elkhorn Peak

Granite

Granite

Granite

Granite

Granite

Mt. Ireland

Bourne

Blue Canyon

Blue Canyon

Blue Canyon

Greenhorn

Bates NE

Bates NE

Bates NE

Bates NE

Bates NE

Bates NE

Bates NE

Bates NE

Greenhorn

Bates SW

Bates SW

Bates SW

Bates SW

118°45' 118°30' |

45°

44°

45'

44°

KELSAY

BUTTE

SILVER

BUTTE

DESOLATION

 BUTTE 

GQ-1654

DESOLATION

BUTTE

BOULDER

BUTTE

GMS-31 BA"

DIXIE 

/lEADOWS

G MS-35

OLIVE 

LAKE

VINEGAR

HILL

GMS-29 

rcc

BATES

\

\ UHtbUIN

AREA OF MAP

GRANITE

GMS-25

GREENHORN

GMS-28

MT 

IRELAND

GMS-22

118°

BOURNE

GMS-19

ELKHORN 

PEAK

GMS-41

0

I 

0

BLUE

CANYON

5 10 MILES

i i

i i

5 10 KILOMETERS

30'

FIGURE 9.2. Quadrangles mentioned in text, and recently published geologic maps covering central part of Baker terrane. 

GMS-19, Brooks, Ferns, Coward, and others (1982); GMS-22, 

Ferns and others (1982); GMS-25, Brooks, Ferns, and Mullen 

(1982); GMS-28, Ferns and others (1983); GMS-29, Brooks 

and others (1983); GMS-31, Ferns and others (1984); GMS35, Brooks and others (1984); GMS-41, Ferns and others 

(1987); GQ-1654, Evans (1988). Generalized geology of all 

but Blue Canyon and Bates 7J4-minute quadrangles is shown 

in figure 9.3.
334 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

and radiometric ages, and sparse structural data, indicates to us that the Baker terrane consists of two 

distinct subterranes, herein referred to as the Bourne 

and Greenhorn subterranes (fig. 9.3). The configuration and composition of the subterranes indicate that 

the Baker terrane may be a paleosuture zone that 

contains part of an accretionary wedge (represented by

the Bourne subterrane) and a fore-arc high (represented by the Greenhorn subterrane).

BOURNE SUBTERRANE

The main lithologic unit in the Bourne subterrane is 

the Elkhorn Ridge Argillite (Gilluly, 1937), which is

118°30

Area of 

figure 9.4

1186

10 KILOMETERS

EXPLANATION

Cover deposits (Cenozoic)

Intrusive rocks (Early Cretaceous to Late Jurassic)

Bourne subterrane Consists of:

Elkhorn Ridge Argillite (Triassic to 

Pennsylvanian) Mainly Permian and Late 

Triassic chert and argillite; also includes 

Pennsylvanian, Permian, and Late Triassic 

limestone

Volcanic rocks (early Mesozoic and 

Paleozoic) Island-arc-related rocks of the 

North Fork (nf) and Goodrich Creek (gc) units, 

and oceanic rocks of the Olive Creek (oc) unit

McCully Fork mixed-rock zone

Greenhorn subterrane Consists of: 

Melange of Vinegar Hill As mapped, includes

some unnamed serpentinite and serpentinitematrix melanges 

Badger Creek unit (Mesozoic? and Paleozoic) As

mapped, includes some unnamed sedimentary

rocks 

Dixie Butte Meta-andesite of Brooks and others

(1984) (Mesozoic?) Island-arc volcanic rocks

FIGURE 9.3. Simplified geologic map of central part of Baker terrane. Line of x's, queried where uncertain, marks boundary between the Bourne and Greenhorn subterranes.
9. BOURNE AND GREENHORN SUBTERRANES OF THE BAKER TERRANE, NORTHEASTERN OREGON 335

TABLE 9.2. Melange types and characteristics in the central part of the Baker terrane

[Modified from Cowan, 1985]

Melange type

Type I Stratified sandstone 

and mudstone 

sequences.

Description

Broken zones with elongate and 

lenticular sedimentary blocks 

formed by layer-parallel extension.

Process of formation

In situ disruption of layered 

sediments by gravity slides or 

thrust faults in accretionary 

prisms.

Baker terrane 

examples

No examples 

known.

Type II Progressively

disrupted sequences 

of mudstone, tuff, 

chert, and 

sandstone.

Type III Block-in-matrix 

mudstone chaos.

Type IV Mudstone-dominated 

brittle fault zones.

Grossly layered, stratally disrupted 

zones subparallel to originally 

stratigraphic layering formed by 

layer-parallel extension with local 

zones of folding and transposition.

Polymict, chaotically disposed 

blocks in massive, weakly fissile, 

or scaly matrix.

Semipenetrative foliation defined 

by myriad anastomosing 

subparallel slip surfaces that 

envelop lenticular tectonic 

inclusions.

Progressive disruption of Elkhorn 

layered sediments by gravity Ridge(?) 

slides or thrust faults in Argillite. 

accretionary prisms.

Incorporation of exotic blocks in Melange of 

a remobilized matrix as Vinegar Hill, 

submarine debris flows or as 

diapirs.

Progressive dismemberment in McCully Fork 

large fault zones. mixed-rock

zone.

best described as argillaceous broken formation. Discrete, mappable zones of oceanic volcanic rocks (mainly alkalic pillow basalts) and island-arc and intrusive 

rocks, including low-K tholeiites, calc-alkaline basalts, 

andesites, dacites, diorites and albite granites, are locally interleaved as tectonic slabs within the Elkhorn 

Ridge Argillite (see Brooks, Ferns, and Mullen, 1982; 

Ferns and others, 1987). In this chapter, we use the 

term "slab" for a coherent, tabular-shaped body of rock 

and the term "block" for a coherent, equidimensional 

body of rock.

ELKHORN RIDGE ARGILLITE

The Elkhorn Ridge Argillite (Gilluly, 1937) consists 

mainly of stratally disrupted, fine-grained siliceous 

sedimentary rocks (argillite, massive chert, and ribbon chert). On Elkhorn Ridge (fig. 9.3), the unit occurs in thick, southward-dipping fault-bounded 

packets of pervasively cleaved, dark-colored argillite 

and chert that contain only minor amounts of intercalated limestone, tuff, sandstone, and conglomerate. 

Individual packets of the Elkhorn Ridge Argillite are 

generally coherent, lithologically homogeneous rock 

masses that are cut by a pervasive fracture cleavage 

(Coward, 1983). The northernmost exposures form 

generally coherent rock masses that are considerably

less disrupted than exposures to the south and west, 

where Evans (1986, 1988) recognized a wide crushed 

area interpreted as a megabreccia zone along the 

southernmost exposures of the Elkhorn Ridge Argillite. Internal chert and argillite fabrics are nearly 

obliterated within this and a similar zone mapped in 

the Granite quadrangle to the east (Brooks, Ferns, 

and Mullen, 1982).

In gross aspect, the Elkhorn Ridge Argillite is best 

described as broken formation (after Hsii, 1968). 

Much of the Elkhorn Ridge Argillite also fits the description of a type II melange (table 9.2; Cowan, 

1985), consisting of stratigraphically continuous zones 

of isoclinally folded chert and argillite that are bounded by either faults or cataclasite zones (Coward, 

1983).

ROCK TYPES

The Elkhorn Ridge Argillite consists mainly of darkcolored siliceous argillite, argillite, and chert. Carbonaceous argillite, feldspar-bearing tuffaceous argillite, 

and ribbon chert are locally abundant, whereas sandstone and conglomerate are rare. Most of the Elkhorn 

Ridge Argillite consists of massive, homogeneous outcrops of blocky argillite in which bedding features are 

generally absent. Chert and metachert are present as
336 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

both coherent and dismembered strata, with individual 

chert beds 5 to 10 cm thick. Chert boudins and 

pseudoconglomerates, that is, broken elongate lenses 

of chert in sheared dark argillite (Pardee and Hewett, 

1914), are present where chert and argillite are interbedded.

Limestone is generally not common in the central 

part of the Baker terrane; however, large elongate 

limestone masses up to 1,000 m in length are present 

in several limestone-rich zones in the Elkhorn Peak 

quadrangle (Ferns and others, 1987). Both individual 

limestone pods and limestone-rich zones parallel the 

regional strike of the enclosing fault-bounded chert 

and argillite packages (fig. 9.4). Very little limestone 

is associated with the Elkhorn Ridge Argillite west of 

Elkhorn Ridge. Margins of some limestone lenses on 

Elkhorn Ridge are marked by limestone-clast breccias. Oolites, peloids, and fossil detritus in some of 

the limestone masses indicate shallow-water deposition (Coward, 1983). Because the limestones are now 

enclosed in presumably deepwater siliceous sediments, the limestones are interpreted as olistoliths 

that formed as shallow-water bioherms and were 

subsequently transported with the limestone breccias 

as carbonate slide blocks into the enclosing deepwater sediments (Coward, 1983).

AGE

Devonian, Pennsylvanian, and Permian limestones 

are enclosed within the Elkhorn Ridge Argillite. 

Evans (1988) and Morris and Wardlaw (1986) reported a Devonian conodont from a small limestone block 

in the Desolation Butte area. According to Coward 

(1983), the limestone masses on the east end of 

Elkhorn Ridge are Pennsylvanian and Permian in 

age. Permian fusulinids include Yabeina sp., Cusenealla sp., Neoschwagerina sp., and Pseudoliolina sp. 

and are considered to be a Tethyan fauna (Coward, 

1983; Nestell, 1983). Similar Tethyan fusulinids are 

reported from the Virtue Hills area east of Baker 

City, Oreg. (fig. 9.1), where limestone of Pennsylvanian, Permian, and Late Triassic ages have been 

identified (C.D. Blome and M.K. Nestell, written 

commun., 1987), whereas a fusulinid assemblage 

similar to those in the Grindstone terrane and the 

McCloud area in northern California is reported 

south of Sumpter, Oreg. (Bostwick and Koch, 1962).

Recent collections by C.D. Blome (Ferns and others, 1987) indicate that cherts in the Bourne subterrane are Permian and Late Triassic in age, although 

Coward (1983) reported radiolarians (Canoptum sp.) 

possibly as young as Early Jurassic from cherts 

south of Elkhorn Ridge. Blome has identified six localities in the Elkhorn Peak quadrangle (Ferns and

118°021 30"

BALD MOUNTAIN^BATHQUTH

44° 

47' 

30"

1 MILE

0 1 KILOMETER

EXPLANATION

Intrusive rocks (Early Cretaceous to Late Jurassic)

Elkhorn Ridge Argillite (Late Triassic to

Devonian) Consists of: 

Chert and argillite

Limestone

Goodrich Creek unit (Late Triassic to

Permian?) Consists of: 

Metamorphosed intrusive rocks

Volcanic and sedimentary rocks

Thrust fault Sawteeth on upper plate 

Strike and dip of bedding (So) 

Strike and dip of foliation (Sj) 

Strike of vertical foliation (Sj)

FIGURE 9.4. Geologic map of the Elkhorn Ridge area. Shows 

contact relations between the Elkhorn Ridge Argillite and intrusive and volcanic and sedimentary rocks on Goodrich 

Creek. Note imbricate thrust faults separating rock types on 

Goodrich Creek (modified from Ferns and others, 1987).
9. BOURNE AND GREENHORN SUBTERRANES OF THE BAKER TERRANE, NORTHEASTERN OREGON 337

others, 1987) that contain Late Triassic (Karnian) radiolarians, including one locality adjacent to a Permian fusulinid-bearing limestone lens; he has also 

identified one locality that contains late Paleozoic radiolarians. Similar chert outcrops east of Baker City 

also are Permian and Late Triassic in age (Blome 

and others, 1986).

METAMORPHISM, STRUCTURE, AND CONTACT RELATIONSHIPS

Metamorphic mineral assemblages are not easily 

recognized in the fine-grained argillite and chert. 

Coarser grained clastic rocks contain low-grade mineral assemblages typical of greenschist-facies conditions, including biotite+chlorite+white mica (Evans, 

1986) and white mica+chlorite+quartz (Kays and others, 1987). Argillite and chert have locally been converted to coarser grained biotite and biotite-garnet 

schists within thermal aureoles of Late Jurassic and 

Early Cretaceous intrusions.

GOODRICH CREEK AND NORTH FORK UNITS

Disrupted, southward-dipping masses of metamorphosed intrusive, extrusive, and sedimentary rocks of 

probable island-arc origin are tectonically intercalated 

with the Elkhorn Ridge Argillite in the northernmost 

exposures of the Bourne subterrane (Stimson, 1980; 

Brooks, Ferns, Coward, and others, 1982; Brooks, 

Ferns, and Mullen, 1982; Evans, 1986, 1988; Ferns 

and others, 1987). The two main areas of island-arc 

volcanic rocks are herein referred to as the North Fork 

and Goodrich Creek units. The North Fork unit consists of andesitic flows and volcaniclastic rocks and 

intrusive rocks that crop out along the North Fork of 

the John Day River west of Granite, Oreg. (fig. 9.3) 

(Brooks, Ferns, and Mullen, 1982; Evans, 1986). Similar volcaniclastic rocks, siliceous tuffs, lava flows, and 

intrusive rocks are tectonically intercalated with slabs 

of the Elkhorn Ridge Argillite on Elkhorn Ridge and 

are well exposed along Goodrich Creek (fig. 9.4) (Stimson, 1980; Brooks, Ferns, Coward, and others, 1982; 

Coward, 1983; Ferns and others, 1987). Both the 

North Fork and the Goodrich Creek units are separated from adjacent overlying and underlying slabs of 

Elkhorn Ridge Argillite by generally low-angle, southward-dipping reverse or thrust faults.

ROCK TYPES

Gray calcareous argillite, sandstone, dark-gray impure argillaceous limestone, and white and blue keratophyre tuff and tuff breccia are complexly interlayered

with limestone, argillite, basalt, and andesite flows, all 

of which are in thrust contact with intrusive rocks. The 

volcaniclastic rocks include siliceous tuff, quartz keratophyre, tuff, and calcareous wacke. Intrusive rocks include mainly equigranular diorite and subordinate 

quartz diorite, albite granite, and foliated hornblende 

gabbro, all of which are similar to the intrusions described by Gilluly (1937). Coarse-grained pyroxenite 

and hornblende pegmatite are rare. Serpentinite and 

its metamorphic derivatives are notably scarce, 

present only as widely separated, narrow serpentinitic 

or talcose selvages along faults and within narrow, 

discontinuous mixed-rock or brittle fault zones similar 

to type IV melanges (table 9.2; Cowan, 1985). Individual beds cannot be traced for more than 100 m. Unlike 

the argillites in the Elkhorn Ridge Argillite, the sediments in the Goodrich Creek unit locally preserve sedimentary structures, including graded bedding and 

flame structures. Much of the Goodrich Creek unit 

consists of widespread, apparently randomly oriented 

cataclasite zones.

Cataclasite zones generally obscure contacts between intrusive and volcanic or volcaniclastic rocks; 

however, small exposures of amphibolite along contacts in the Goodrich Creek area may be relicts of 

contact aureoles between greenstones and diorite intrusions. Strongly deformed hornblende gabbros are 

cut by equigranular hornblende diorites that are in 

turn cut by quartz diorite and albite granite masses. 

The quartz diorite and albite granite masses locally 

intrude volcanic and sedimentary rocks.

Intrusive rocks are separated by faults from the 

volcanic and volcaniclastic rocks. The imbricate nature of these faults is well demonstrated on the high 

ridge northwest of Pine Creek (fig. 9.4) where a 400- 

m-thick sequence of tuff breccia, argillite, and argillaceous limestone is tectonically overlain and underlain by southward-dipping fault slices of diorite 

(Ferns and others, 1987).

GEOCHEMISTRY

Andesite is the most abundant volcanic rock type 

on the North Fork of the John Day River. Major-element geochemistry (table 9.3) suggests that most if 

not all of the Goodrich Creek and North Fork volcanic 

flow rocks formed in an island-arc environment. Generally, the basaltic rocks have low TiO2 contents and 

low K2O contents (table 9.3) typical of low-K tholeiites. Mullen (1985) concluded that North Fork rocks 

have an island-arc origin on the basis of their relatively high SiO2 contents (53-55 percent), low TiO2 

contents, depleted light rare-earth elements, and position within the destructive margin field on the Hf/3- 

Th-Ta plot of Wood and others (1979). Calder (1986)
338 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 9.3. Major-element data for greenstones from the Bourne subterrane

[Values in weight percent. All analyses recalculated to 100 percent on a volatile-free basis, n.d., not determined]

Volcanic rocks

Unit

Sample -

SiO2 -

Al-A 

FeO*

MnO -

PaO

MgOIT n

Ma O ~

P2Oo5 -

H2(-)+ u - 

H

/-VC02

Olive 

Creek 

unit

174

46.27 

16.00 

1.75 

10.22 

0.13 

15.30 

4.88 

0.38 

4.80 

0.30

n.d. 

n.d. 

n.d.

McCully Fork mixed-rock zone

2B-135B 3BB-133

51.39 

17.55 

1.43 

7.12 

0.16 

11.95 

6.80 

0.25 

3.27 

0.18

2.35 

0.38 

0.17

51.63 

16.52 

1.30 

5.45 

0.15 

14.69 

6.69 

0.17 

3.26 

0.13

n.d. 

n.d. 

n.d.

3MB-45 177 2"R 133A 1"R *S7 1] 1 I O-J.OO.rV JJ-O / J

52.14 53.18 

17.18 16.45 

0.84 1.00 

8.95 9.20 

0.15 0.16 

10.37 8.54 

6.78 7.39 

0.10 0.02 

3.36 3.95 

0.11 0.10

n.d. 

n.d. 

n.d.

n.d. 

n.d. 

n.d.

53.50 

14.14 

1.11 

6.93 

0.13 

13.95 

6.19 

0.29 

3.64 

0.12

2.13 

0.12 

1.51

Goodrich Creek unit

i?T3 inK IT?D CQ ID 07 IT?D QC diJr-lUO Hii -oy D-Zi 1 Eir-oo

49.66 49.76 51.92 50.32 

18.24 16.63 14.18 18.22 

1.19 1.70 1.67 1.22 

10.23 10.36 11.21 10.21 

0.17 0.18 0.18 0.15 

11.37 12.31 10.49 9.97 

6.51 6.52 7.46 4.26 

0.14 0.06 0.02 1.57 

2.30 2.31 2.69 3.38 

0.19 0.17 0.18 0.70

n.d. 

n.d. 

n.d.

Volcanic rocks Continued

Unit 

Sample 

TO

FeO* -----

MnO

CaO 

MgO -

V (\

Na20 

P2nU5 - 

H n+

H

/->- 2cr 

Goodrich Creek 

unit Continued

1EP-102

55.19 

14.20 

0.96 

9.49 

0.17 

10.64 

6.59 

0.04 

2.61 

0.10

n.d. 

n.d. 

n.d.

1B-125

57.30 

18.64 

1.01 

7.54 

0.13 

6.27 

3.97 

1.75 

3.23 

0.18

n.d. 

n.d. 

n.d.

'B-259A

58.29 

18.33 

1.27 

10.74 

0.06 

2.23 

5.73 

0.27 

2.80 

0.17

n.d. 

n.d. 

n.d.

North Fork unit

:B-259B

62.28 

15.66 

0.93 

9.02 

0.06 

4.72 

4.20 

0.15 

2.70 

0.26

n.d. 

n.d. 

n.d.

^53

74.07 

14.75 

0.36 

2.46 

0.04 

2.22 

1.35 

1.19 

3.47 

0.08

n.d. 

n.d. 

n.d.

'M61

75.63 

14.15 

0.43 

2.25 

.05 

0.15 

1.53 

3.52 

2.23 

0.03

n.d. 

n.d. 

n.d.

n.d. n.d. 

n.d. n.d. 

n.d. n.d.

n.d. 

n.d. 

n.d.

53.40 

17.54 

1.16 

9.35 

0.19 

9.37 

4.45 

0.00 

4.33 

0.21

n.d. 

n.d. 

n.d.

Plutonic rocks

Goodrich Creek unit

^-90

55.87 

15.16 

0.70 

9.07 

0.18 

7.97 

7.32 

0.49 

3.17 

0.07

n.d. 

n.d. 

n.d.

1EP-52

58.80 

20.05 

0.93 

6.57 

0.10 

4.90 

2.88 

0.79 

4.73 

0.24

n.d. 

n.d. 

n.d.

1EP-48

74.07 

13.81 

0.61 

3.52 

0.07 

2.56 

0.65 

0.10 

4.49 

0.11

n.d. 

n.d. 

n.d.

North 

Fork unit

1B-289

64.48 

16.62 

0.67 

5.99 

0.13 

6.52 

2.24 

0.03 

3.22 

0.10

n.d. 

n.d. 

n.d.

1Sample analyzed by X-ray fluorescence at Washington State University, Pullman, Wash. 

2Sample analyzed by X-ray fluorescence at U.S. Geological Survey, Menlo Park, Calif. 

3Sample analyzed by atomic absorption at University of Oregon, Eugene, Oreg.

presented trace-element data that suggest that Goodrich Creek greenstones also include low-K island-arc 

tholeiites.

Coarse-grained plutonic rocks have major-element 

contents (table 9.3) similar to those of rocks in the 

(informal) Sparta complex of Phelps (1979) (fig. 9.5),

which are considered to be genetically related to the 

overlying volcanic rocks of the Wallowa arc (Almy, 

1977; Phelps, 1979; Ave Lallemant and others, 1980). 

Analyzed flows from both the Goodrich Creek and the 

North Fork units are similar to Wallowa arc rocks in 

that they have low amounts of K2O and TiO2. High-
9. BOURNE AND GREENHORN SUBTERRANES OF THE BAKER TERRANE, NORTHEASTERN OREGON 339

silica intrusive rocks in the North Fork and Goodrich 

Creek units are high in Na2O and low in K2O content, 

which in part may reflect mobilization of these oxides 

during metamorphism.

AGE AND CORRELATIONS

Available radiometric data indicate that the volcanic-arc rocks are about the same age, Permian and Late 

Triassic, as the adjoining fine-grained siliceous sedimentary rocks of the Elkhorn Ridge Argillite. An Early 

Triassic low-K quartz diorite intrusion from the North

FeO'

CaO MgO

EXPLANATION

Baker terrane 

Bourne subterrane 

Goodrich Creek unit 

A 3 ODGMI analyses

 5 analyses by Stimson (1980) 

North Fork unit

* 1 ODGMI analysis

Greenhorn subterrane 

Melange of Vinegar Hill 

n 1 ODGMI analysis

x 6 analyses by Hunt (1985) 

+ 7 analyses by Mullen (1978)

FIGURE 9.5. FeO*-CaO-MgO diagram of analyzed metamorphosed intrusive rocks from central part of Baker terrane. 

Shaded fields are from 24 analyses by Almy (1977) and 

Phelps (1978, 1979) of the (informal) Sparta complex of 

Phelps (1979). Baker terrane data are from table 9.4 and 

from master's theses by Mullen (1978), Hunt (1985), and 

Stimson (1980). ODGMI, Oregon Department of Geology and 

Mineral Industries.

Fork of the John Day River has yielded a U-Pb zircon 

age of 243 Ma (Brooks, Ferns, and Mullen, 1982; Walker, 1983). Intruded greenstones are therefore earliest 

Triassic or older in age. Limestones interbedded with 

pyroclastic rocks on Goodrich Creek contain Late Triassic (Norian) conodonts (Wardlaw, written commun., 

1986; Ferns and others, 1987). These data, although 

sparse, suggest earliest Triassic or older volcanism, 

followed by emplacement of earliest Triassic plutons, 

in turn followed by renewed pyroclastic volcanism and 

limestone deposition in the Norian.

METAMORPHISM

Mineral assemblages outside the thermal aureoles of 

the Late Jurassic intrusions are generally typical of 

greenschist-facies conditions. In the greenstones, metamorphic assemblages include actinolite+albite+epidote+ 

chlorite+calcite, epidote+albite+chlorite+calcite, and 

chlorite+actinolite+albite+calcite. One foliated sample 

from a fault zone on the east end of Elkhorn Ridge 

contains stilpnomelane.

OLIVE CREEK UNIT

Alkalic pillow basalt (Mullen, 1982, 1985) and intercalated felsic tuff, volcanic breccia, limestone lenses, and contorted ribbon chert are exposed along 

Olive Creek about 2 km north of Greenhorn, Oreg. 

(figs. 9.3, 9.9). The volcanic rocks are tectonically interleaved with slices and slivers of metamorphosed 

intrusive rocks and chert and argillite of the Elkhorn 

Ridge Argillite (Ferns and others, 1983) near what 

we interpret as the southern margin of the Bourne 

subterrane. The Olive Creek volcanic rocks and the 

cherts and intrusive rocks appear to be components 

of a slab melange in which little or no matrix is 

present. Individual slabs are large elongate masses, 

as much as 2,400 m in length.

PILLOW LAVA GEOCHEMISTRY

Mullen (1978, 1982, 1985) identified relict titaniferous clinopyroxenes in the Olive Creek pillow lavas; 

such clinopyroxenes are typical of alkalic basalts. 

The pillow lavas' major-element contents, including 

high TiO2 contents (table 9.3), are also indicative of 

alkalic basalts (Mullen, 1978, 1982; Ferns and others, 1983). Rare-earth-element data presented by 

Mullen (1985) are consistent with the pillow lavas 

originating in an intra-plate or seamount setting.
340 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

AGE

Rocks of the Olive Creek unit are apparently Permian and Late Triassic in age. Dickinson and Thayer 

(1978) reported Permian conodonts from ribbon chert 

in the Olive Creek area. C.D. Blome (written commun., 1987) identified Late Triassic radiolarians from 

contorted ribbon chert exposed along Olive Creek. 

Early Permian (Leonardian) conodonts were reported 

by Mullen (1978) from a limestone lens encased in 

mafic tuff and pillow lava exposed in a small pre-Tertiary inlier east of Olive Creek.

METAMORPHISM

Pillow lava samples usually retain relict igneous 

textures. Massive greenstones are nonschistose and 

contain typical greenschist-facies mineral assemblages. Mafic phenocrysts are partially replaced by felted 

mats of chlorite, epidote, and (or) actinolite (Mullen, 

1978). Regional metamorphic mineral assemblages 

include chlorite+albite+epidote and chlorite+albite+ 

actinolite.

MIXED-ROCK ZONES

The term "mixed-rock zone" is used (Brooks, Ferns, 

Coward, and others, 1982; Ferns and others, 1982, 

1983; Ferns and Ramp, 1988) to describe mappable 

lithodemic units in which individually unmappable (at 

a reasonable scale) slabs and slices of diverse rock 

types are juxtaposed. These zones often contain relatively small slices and anastomosing lenses of intervening serpentinite-matrix melange and are similar 

in overall appearance to the slab melange described 

by Robertson and Woodcock (1983) in the Batinah melange in Oman. The mixed-rock zones are analogous 

to the brittle fault zones or type IV melanges of 

Cowan (1985; table 9.2). The zones are composed of 

elongate, lensoid slices of diverse rock types such as 

argillite, greenstone, serpentinite, and metaserpentinite, all of which are bounded by faults generally 

parallel to the dominant foliation or cleavage.

The McCully Fork mixed-rock zone (figs. 9.3, 9.6) is 

the largest mixed-rock zone in the Bourne subterrane. This northwest-trending zone is traceable for 

about 18 km along strike and ranges from a few hundred meters to 3,200 m in width. It is composed of 

randomly juxtaposed slices and blocks of argillite, 

serpentinite, pyroxenite, gabbro, diorite, quartz diorite, basalt, volcaniclastic breccia and conglomerate, 

limestone, and chert that are locally separated by ultramafic melange zones containing from about 10 to

70 percent matrix (Ferns and Ramp, 1988). Individual slices outside the melange zones are elongate parallel to the northwesterly strike of the McCully Fork 

mixed-rock zone and seldom reach more than 150 m 

in width and 1,000 m in length. Blocks within the 

melange zones are equidimensional, ranging from 1

118°22'30"

1 MILE

0 1 KILOMETER

EXPLANATION

Cover rocks (Tertiary)

Intrusive rocks (Early Cretaceous to Late Jurassic)

Elkhorn Ridge Argillite (Late Triassic to Devonian)

McCully Fork mixed-rock zone Consists of: 

Serpentinite and meta-serpentinite

Slab melange with little or no intervening matrix 

Block melange with intervening matrix

Thrust fault Sawteeth on upper plate 

Strike and dip of bedding (S0 ) 

Strike and dip of foliation (Sj) 

Strike of vertical foliation (S)

FIGURE 9.6. Geologic map of McCully Fork mixed-rock zone east 

of Sumpter, Oreg. (modified from Ferns and others, 1982, and 

Ferns and Ramp, 1988).
9. BOURNE AND GREENHORN SUBTERRANES OF THE BAKER TERRANE, NORTHEASTERN OREGON 341

TABLE 9.4. Trace-element data for greenstones from the central part of the Baker terrane

[Values in parts per million. All samples analyzed at U.S. Geological Survey, Menlo Park, Calif. Samples VDB-1 and VDB-2 analyzed by X-ray fluorescence, all 

others by instrumental neutron activation analysis. , not determined or below instrumental detection limit]

Bourne subterrane

Unit McCully Fork mixed-rock 

zone

Sample- B-133A B-135B

Rb . 2 4

Sr - 223 162

Ba - 32 56

Th - 0.5 0.9 

U .. 

La - 7 11

Ce - 15 20 

Sm -

Eu - 1.04 1.17

Gd - 3.2 3.9

Tb - 0.71 0.65 

TinYb - 2.5 2.1

LU - 0.41 0.32

Y - - 23 22

Zr - 95 97

Hf - 2.3 2.2

Ta - 0.48 0.84

Nb - 7 13

Ni ~ 96 73

Co - 37 36

Cr - 310 121

Sc - 34.8 33.9

Zn - 72 75

Greenhorn subterrane

Melange of Vinegar Hill

V-l V-2 V-3

Dixie Butte Meta-andesite

VDB-1 VDB-2

13

373

677

2.61

0.73

22.5

44.7

6.3

1.70

5.5

0.85

0.42

2.26

0.33

27

212

4.74

1.75

25

47

295

20.8

100

10

172

189

0.50

0.45

3.54

.0

0.89

0.70

3.42

0.52

30

60

1.54

57

320

61.6

134

21

188

99

0.96

0.59

10.4

1.6

4.0

1.31

0.70

2.41

0.37

26

133

2.89

0.72

16

66

679

32.6

104

20

230

320

20

5

20

230

60

20

0

15

110

25

170

10

to 40 m in diameter, and consist mainly of greenstone, but also include basalt, pyroclastic breccias, 

metagabbro, and metadiorite.

Serpentinite- and metaserpentinite-matrix melanges are best developed in the center of the McCully 

Fork mixed-rock zone where it widens east of Granite, Oreg. A high-angle reverse fault separates the 

mixed-rock zone from the structurally underlying 

Elkhorn Ridge Argillite to the north (Brooks, Ferns, 

Coward, and others, 1982; Ferns and others, 1982). 

Thin mylonite zones in the underlying Elkhorn Ridge 

Argillite strike parallel to the faults. The contact between the McCully Fork mixed-rock zone and the 

overlying Elkhorn Ridge Argillite to the south is a 

diffuse zone of intercalated slices of argillite and serpentinite. Unusually coarse grained sedimentary 

rocks, including chert-pebble conglomerate and 

graywacke interbedded with small limestone lenses, 

locally crop out along the southern contact zone. The 

Late Jurassic Bald Mountain batholith intrudes the

contact and has converted the serpentinite matrix of 

the melange zones to a variety of metaserpentinites, 

including talc- and enstatite-olivinites (Ferns and 

Ramp, 1988).

GEOCHEMISTRY

Ophiolitic components of the mixed-rock zones include serpentinite, gabbro, basalt, and chert. Basalts 

(table 9.3) are low-K tholeiites that are slightly enriched in light rare-earth elements (table 9.4). Some 

basalts are similar to the low-K tholeiites of the Goodrich Creek unit (low K^O and moderate Na2O contents). The mixed-rock-zone basalts have generally 

lower FeO* (total iron as FeO) contents than the 

Goodrich Creek and North Fork rocks. Oceanic affinities are suggested by various trace-element plots, including the Ti-Zr-Y plot (fig. 9.7) and the Hf-Ta-Th 

plot (fig. 9.8).
342 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

METAMORPHISM

Most of the rocks within the mixed-rock zones have 

been recrystallized by Late Jurassic intrusions. Metamorphic zonation around the Bald Mountain batholith 

is well evidenced in the metaserpentinites that grade 

from enstatite-olivine rock outward through talc-amphibole and talc-carbonate rock into antigorite-carbonate rock (Ferns and Ramp, 1988). Metamorphic 

mineral assemblages in greenstones outside the thermal aureoles are typical albite-epidote facies assemblages (epidote + chlorite + albite + calcite and epidote 

+ chlorite + albite + actinolite + calcite).

GREENHORN SUBTERRANE

Serpentinite-matrix melange, best exposed at Vinegar Hill (see fig. 9.9), is the main unit in the Greenhorn subterrane. Other major Greenhorn subterrane 

units include the clastic-rich Badger Creek unit and 

volcanic rocks of the Dixie Butte Meta-andesite of 

Brooks and others (1984) (fig. 9.3).

Ti/100

Y + 3

FIGURE 9.7. Triangular diagram of Ti-Zr-Y contents in greenstones from central part of Baker terrane. Bourne subterrane samples include samples of Goodrich Creek unit 

(crosses) from Calder (1986) and samples from McCully Fork 

mixed-rock zone (triangles). Greenhorn subterrane samples 

are from the melange of Vinegar Hill (open squares) and 

from the Dixie Butte Meta-andesite of Brooks and others 

(1984) (solid squares). Fields from Pearce and Cann (1973): 

A, ocean-island and continental basalts; B, calc-alkaline basalts; C, ocean-floor basalts, low-K tholeiites, and calc-alkaline basalts; D, low-K tholeiites.

BADGER CREEK UNIT

Greenhorn subterrane supracrustal rocks include 

Paleozoic and possibly Mesozoic clastic sedimentary 

rocks previously informally referred to as undivided 

sedimentary rocks (J~fcu) by Brown and Thayer (1966), 

the Badger Creek unit by Wheeler (1976) and Mullen 

(1978), and clastic sedimentary rocks (~EPa) by Ferns 

and others, (1983). In this chapter, we use the name 

"Badger Creek unit" (Wheeler, 1976; Mullen, 1978).

ROCK TYPES

The Badger Creek unit includes sedimentary rocks 

made up chiefly of fine- to medium-grained calcareous 

sandstone, chert-pebble conglomerate, tuff, limestone 

lenses, argillite, and polymict breccia that have been 

folded and pervasively sheared. Although the Badger 

Creek unit is discontinuously exposed over an area of 

about 155 km2, extensive Tertiary cover has prevented 

the determination of its stratigraphic relationships.

Sedimentary features such as graded bedding and 

flame structures are locally preserved in fine-grained 

calcareous sandstone and argillite. In some places inverted graded bedding suggests that the sandstones 

are locally overturned. Conglomerate units contain 

well-rounded clasts of sandstone, chert, argillite, and 

andesite, as well as detrital quartz and feldspar grains. 

Chert-pebble conglomerates locally fill channels cut 

down into argillite and sandstone. Chert pebbles are

Hf/3

FIGURE 9.8. Triangular diagram of Hf73-Ta-Th contents in 

greenstones from central part of Baker terrane. Bourne subterrane (triangles) and Greenhorn subterrane (hexagons) 

samples and localities listed in tables 9.1 and 9.4. Fields 

from Wood and others (1979). NMORB, normal midoceanridge basalts; EMORB, evolved midocean-ridge basalts.
9. BOURNE AND GREENHORN SUBTERRANES OF THE BAKER TERRANE, NORTHEASTERN OREGON 343

well rounded and typically range from 1 to 5 cm in 

diameter. In some areas, such as east of Vinegar Hill 

(Brooks and others, 1983), conglomerate beds are carbonate cemented and contain crinoidal debris. These 

chert-pebble conglomerate beds are interbedded with 

1- to 25-m-thick fossiliferous limestone lenses, which, 

at Vinegar Hill, contain Pennsylvanian conodonts 

(Morris and Wardlaw, 1986).

Polymict breccias occur in the Badger Creek unit 

near its contact with the underlying melange of Vinegar Hill. Clasts are subangular to subrounded and 

include metamorphosed intrusive rocks, greenstones, 

and chert. Clasts are up to 12 cm in diameter and 

occur in both argillaceous matrix-supported and 

clast-supported breccias. Distinctive lineated gabbro 

clasts appear to be derived from immediately adjacent gabbro blocks within the underlying melange. 

Serpentinite-matrix and talc-matrix sedimentary 

breccias (Ferns and Ramp, 1988), although rare, are 

important because they indicate sporadic introduction of ultramafic olistostromes as the Badger Creek 

unit was being deposited.

AGE AND FAUNAL CORRELATIONS

Pennsylvanian, Permian, and Late Triassic(?) faunas have been reported from the Badger Creek unit. 

Pennsylvanian (Desmoinesian) conodonts were identified from a limestone pod interbedded with chertpebble conglomerate on the east flank of Vinegar Hill 

(Morris and Wardlaw, 1986). Both Permian conodonts 

and fusulinids have been identified in a limestone 

pod west of Granite Boulder Creek (see fig. 9.13) 

(Mullen, 1978). Permian (Guadalupian) radiolarians 

have been found in siliceous interbeds in sandstone 

and siltstone on Granite Boulder Creek near the fusulinid locality (Ferns and others, 1984; Blome and 

others, 1986), indicating that the limestones here are 

not olistoliths of older Paleozoic material incorporated into a younger Mesozoic strata.

Fusulinids include Pseudofusulinella and Schwagerina (Mullen, 1978). Nestell (oral commun., 1986) 

suggested that the fusulinid assemblage of the Badger Creek unit is similar to those from the McCloud 

area in the Klamath Mountains of northern California and is different from most of the Tethyan fusulinid assemblages found in similar-aged limestones in 

the Elkhorn Ridge Argillite in the Bourne subterrane.

The Badger Creek unit may include Mesozoic (Late 

Triassic?) sediments as evidenced by poorly preserved 

cone-shaped radiolarians (Blome, oral commun., 

1983) collected from green chert beds along a depositional contact with underlying melange south of the 

town of Greenhorn (Ferns and others, 1982). This 

chert separates the melange of Vinegar Hill from a 

coarse polymict breccia that contains clasts derived 

from blocks within the melange.

METAMORPHISM

Most of the outcrop area of the Badger Creek unit 

lies within thermal aureoles of Late Jurassic to Early 

Cretaceous intrusions. Sedimentary rocks outside of 

the aureoles are poorly recrystallized and lack recognizable metamorphic fabrics. The typically observed 

metamorphic mineral assemblage of epidote+actinolite+chlorite+ albite would be characteristic of either 

low-grade regional or low-grade contact metamorphism.

MELANGE OF VINEGAR HILL

The northernmost unit of the Greenhorn subterrane is the serpentinite-matrix melange of Vinegar 

Hill (fig. 9.9). This melange is exposed over an area 

of about 180 km2 in the Greenhorn Mountains and is 

separated from units of the Bourne subterrane by a 

major, southward-dipping high-angle reverse or 

thrust fault (Ferns and others, 1983; Evans, 1988). 

The melange of Vinegar Hill is composed of mafic 

plutonic and volcanic rocks, silicic volcanic rocks, 

chert, and argillite in irregularly shaped blocks ranging from 2 m to more than 3 km in length, which are 

encased in a serpentinite matrix that makes up over 

50 percent of the unit. Blocks include monolithologic 

greenstone, gabbro, and argillite and heterolithologic 

blocks made up of intrusive, extrusive, and sedimentary rocks that occasionally retain primary depositional and (or) intrusive contacts. The melange of 

Vinegar Hill is analogous to the type III melange of 

Cowan (1985) in that diverse rock types are apparently randomly arrayed as blocks in a chaotic, finegrained matrix. Some of the blocks are cut by narrow 

schistose serpentinite veins that appear similar to 

those described in the Kaweah belt in California by 

Saleeby (1979), which document the rafting apart of 

the greenstones as serpentinite was injected. Metasomatic reaction rinds (rodingite) occur only along the 

margins of the gabbro blocks. Since rodingites form 

during serpentinization (Coleman, 1963), the lack of 

metasomatic reaction rinds in the greenstones suggests that the greenstones were incorporated into the 

melange after initial serpentinization, presumably as 

the serpentinite was injected along fractures.
344 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

118°32'30"

Hbgvspi

"bgyny

EXPLANATION

Cover deposits (Cenozoic)

Intrusive rocks (Early Cretaceous and Late Jurassic)

Baker terrane Divided into: 

Bourne subterrane Consists of: 

Olive Creek unit Divided into: 

Ribbon chert and argillite

Alkalic pillow basalts and volcaniclastic rocks

Greenhorn subterrane Consists of: 

Badger Creek unit

Melange of Vinegar Hill Divided into: 

Serpentinite

Thrust fault and subterrane boundary 

Sawteeth on upper plate 

Strike and dip of bedding (S0)

Strike and dip of foliation (Sj) 

Strike of vertical foliation (Sj)

Metavolcanic rocks 

Meta-intrusive rocks 

Foliated amphibolite

FIGURE 9.9. Geologic map of melange of Vinegar 

Hill showing boundary thrust fault between 

Greenhorn and Bourne subterranes (modified 

from Ferns and others, 1983, and Brooks and 

others, 1983).
9. BOURNE AND GREENHORN SUBTERRANES OF THE BAKER TERRANE, NORTHEASTERN OREGON 345

TABLE 9.5. Major-element data for greenstones from the Greenhorn subterrane

[All samples normalized to 100 percent on a volatile-free basis, n.d., not determined]

Unit

Sample 

SiO2 

FeO* - 

MnO 

faO- - -

MgO -

V O

Na2O 

P2Ur\5 -

H

rvH2O- - 

r1 n

Melange of Vinegar Hill

2B-372

44.14 

16.62 

1.49 

15.67 

0.29 

11.29 

7.73 

0.26 

2.37 

0.14

n.d. 

n.d. 

n.d.

 V-8

45.75 

15.62 

1.77 

11.88 

0.20 

10.44 

9.02 

1.58 

3.47 

0.28

4.10 

0.27 

4.79

263A

46.12 

19.40 

0.63 

9.64 

0.17 

12.22 

9.24 

0.20 

2.34 

0.03

n.d. 

n.d. 

n.d.

2B-364

47.07 

16.60 

1.84 

10.34 

0.12 

11.85 

4.95 

1.29 

5.30 

0.64

n.d. 

n.d. 

n.d.

'V-2

47.18 

15.22 

1.01 

13.80 

0.20 

11.61 

8.54 

0.36 

2.36 

0.16

2.23 

0.17 

0.08

 V-l

49.75 

15.05 

2.43 

11.98 

0.20 

9.02 

8.29 

0.35 

2.54 

0.39

2.25 

0.38 

0.01

260A

50.51 

16.43 

0.79 

9.18 

0.15 

9.74 

9.18 

0.16 

3.69 

0.15

n.d. 

n.d. 

n.d.

260B

70.85 

14.56 

0.35 

4.22 

0.09 

1.57 

2.56 

0.60 

5.13 

0.07

n.d. 

n.d. 

n.d.

255

49.17 

19.82 

0.44 

7.19 

0.14 

11.19 

8.57 

0.17 

3.27 

0.04

n.d. 

n.d. 

n.d.

Dixie Butte Meta-andesite

2 185

53.37 

15.68 

0.85 

7.72 

0.15 

10.27 

8.93 

0.74 

2.22 

0.14

n.d. 

n.d. 

n.d.

2VDB-1

53.61 

16.16 

1.00 

8.50 

0.16 

8.56 

7.96 

1.00 

2.85 

0.18

n.d. 

n.d. 

n.d.

2VDB-2

55.04 

16.45 

1.36 

8.73 

0.14 

8.48 

5.61 

1.01 

2.93 

0.25

n.d. 

n.d. 

n.d.

2DBV

55.46 

18.94 

1.52 

7.56 

0.14 

6.82 

3.31 

2.22 

3.74 

0.30

n.d. 

n.d. 

n.d.

1Analyzed at U.S. Geological Survey, Menlo Park, Calif. 

2Analyzed at Washington State University, Pullman, Wash.

GEOCHEMISTRY

Major- and trace-element contents (tables 9.4, 9.5) 

indicate that greenstone blocks within the melange 

include both oceanic and island-arc-related material. 

The large heterolithologic block south of Greenhorn 

includes island-arc basalts (Mullen, 1978, 1982) and 

dacite (Ferns and others, 1983). Both alkalic and 

low-K tholeiitic basalts (tables 9.4, 9.5) occur in the 

large heterolithologic block that culminates in the 

summit of Vinegar Hill. This block, which is made up 

of interlayered pillow lava and banded chert intruded 

by narrow diorite dikes, is apparently a fragment of 

oceanic crust. It contains a low-K tholeiite geochemically similar to the Bourne subterrane mixed-rockzone rocks that plots within the evolved-midoceanridge-basalt (E-MORB) field on the Hf/3-Th-Ta diagram (fig. 9.8). The one Vinegar Hill low-K tholeiite 

has a flatter rare-earth-element pattern than the 

Bourne subterrane mixed-rock-zone low-K tholeiites 

(fig. 9.10).

Work by E.M. Bishop (Mullen, 1985) indicates that 

Vinegar Hill gabbros are tholeiitic in character and 

may have more oceanic affinities than the gabbros in 

the Canyon Mountain Complex. Major-element abundances (table 9.5) show that the Vinegar Hill intrusive blocks are chemically distinct from the Bourne 

subterrane island-arc intrusive rocks. Vinegar Hill 

blocks are enriched in CaO and MgO and depleted in 

TiO2 and FeO* as compared to Bourne subterrane intrusive rocks of similar silica content (fig. 9.5).

METAMORPHISM

Blocks in the melange of Vinegar Hill range from 

prehnite-pumpellyite to epidote-amphibolite in metamorphic grade. Igneous textures in the lower grade 

rocks are obscured by felted masses of chlorite, albite, and epidote. Epidote amphibolites are typically 

tightly crenulated foliated rocks either associated 

with sheared serpentinite along major fault zones or

100

80

2 60

£ 40

20

La Ce Nd Sm Eu Gd Tb Ho Tm Yb Lu 

RARE-EARTH ELEMENTS

FIGURE 9.10. Rare-earth-element (REE) plot for greenstones 

from Baker terrane. Dots, basalts from mixed-rock zones in 

Bourne subterrane; crosses, basalts of the melange of Vinegar Hill (Greenhorn subterrane). Data are listed in table 9.4 

and Vallier (1985 and chap. 3, this volume).
346 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

along the outer margins of metagabbro blocks. Metamorphic minerals include greenish-brown hornblende, epidote, sodic plagioclase, and quartz.

Blue amphiboles occur in foliated metamorphic 

rocks encased in melange in small inliers southeast 

of the main exposures of the melange of Vinegar Hill 

(Mullen, 1978, 1980; Ferns and others, 1983). Associated minerals include chlorite, epidote, albite, quartz, 

white mica, and garnet. Mullen (1978, 1980) reported 

that the amphiboles recorded two periods of metamorphism. An early low-temperature/high-pressure 

event (about 350°C and 6.5 kb) is indicated by barroisitic amphibole cores. A later, moderate-temperature/moderate-pressure event (400-450°C and 3-5 

kb) is indicated by more calcic amphibole rims.

ORIGIN AND AGE

Fossil data pertaining to the age of the melange of 

Vinegar Hill are fragmentary and contradictory. Ages 

are based mainly on the overlying Badger Creek unit, 

which contains Pennsylvanian, Permian, and Late 

Triassic fossils. Morris and Wardlaw (1986) reported 

Pennsylvanian conodonts from a limestone lens interbedded with chert-pebble conglomerate and calcareous sandstone and siltstone in apparent depositional 

contact with the Vinegar Hill block. The sediments 

were considered by Brooks and others (1983) to be 

part of the depositionally overlying Badger Creek 

unit. Provided that the limestone is not an exotic clast 

in the overlying sediments, the underlying block in 

the melange of Vinegar Hill must be Pennsylvanian or 

older in age. However Blome (oral commun., 1983) 

reported that radiolarians from green cherts near the 

base of the Badger Creek unit near Greenhorn and 

from a bedded chert block encased within the melange 

south of Greenhorn, although poorly preserved, were 

cone-shaped and likely Mesozoic in age.

The apparent wide range of faunal ages in the overlying Badger Creek unit sediments suggests that the 

melange of Vinegar Hill may be polygenetic, that is, it 

may have developed over a period of time under 

changing tectonic conditions. Scattered exposures of 

overlying polymict sedimentary breccia and conglomerate that contain clasts locally derived from blocks 

within the melange, along with rare exposures of 

overlying serpentinite-matrix and talc-matrix sedimentary breccias, indicate that the melange of Vinegar Hill locally shed olistostromal debris into flanking 

sedimentary basins. Many of the smaller sedimentary 

blocks are completely engulfed by serpentinite and 

now compose part of the melange, indicating that melange continued to develop after deposition of the sediments began. The elongate slices of foliated epidote

amphibolite and sheared serpentinite that are tectonically interleaved with sheared sedimentary rock in 

fault zones along the contact between the melange 

and the Badger Creek unit (Brooks and others, 1983) 

indicate that tectonic mixing continued after the sediments were initially deposited.

Small patches of serpentinite-matrix sedimentary 

breccias south of Greenhorn that locally overlie the 

melange of Vinegar Hill do not appear to be as intensely deformed as the rest of the Badger Creek 

unit, indicating that serpentinite did not breach the 

oceanic floor until later stages of basin development. 

Late-stage breaching by serpentinite is supported 

also by Hunt (1985), who reported deformation of melange blocks prior to serpentinite injection. Widespread introduction of ultramafic rocks may not have 

occurred until the late stages of melange evolution, 

as suggested by the general absence of ultramafic detritus in Mesozoic polymict breccias that contain 

abundant block-derived detritus in the form of gabbro and greenstone clasts.

Earlier periods of serpentinite emplacement may 

be evidenced by local ophicarbonate (talc-carbonate) 

beds within the central part of the Badger Creek unit 

in the Granite Boulder Creek area (see fig. 9.13; 

Ferns and Ramp, 1988). Permian fusulinid and radiolarian localities occur in the immediate vicinity of 

the ophicarbonate bed.

The melange of Vinegar Hill may have begun to 

form in a fore-arc setting as a rising serpentinite diapir similar to those reported in the Mariana and 

Bonin arcs in the western Pacific Ocean (Bloomer, 

1983; Taylor and Smoot, 1984). Serpentinites apparently form from fluxing of ultramafic rocks at the base 

of the outer wall of the fore arc by waters derived from 

the underlying subducted slab. Presumably, the decrease in specific gravity during conversion of ultramafic rock to serpentinite would allow for diapiric 

injection of serpentinite into a fracturing cover of 

overlying, foundering oceanic and island-arc crust in 

the fore-arc wall. Olistostromal debris derived from 

the rising diapir and fractured cover could then be 

shed as gravity slides into flanking basins. Blue-amphibole-bearing epidote amphibolites and disrupted 

intrusive rocks may be fragments of the underlying 

crust which were later incorporated into the diapir.

DIXIE BUTTE META-ANDESITE

The southernmost major lithologic unit mapped in 

the central part of the Baker terrane is the volcanicarc-related Dixie Butte Meta-andesite of Brooks and 

others (1984). These volcanic rocks crop out over an
9. BOURNE AND GREENHORN SUBTERRANES OF THE BAKER TERRANE, NORTHEASTERN OREGON 347

area of about 80 km2 and are mainly massive plagiophyric andesite and basaltic andesite flows that are 

interbedded with island-arc basalts (Mullen, 1982).

ROCK TYPES

Major rock types are pyroxene basalts and plagiophyric basaltic andesites that contain feldspar phenocrysts up to 5 cm in length. The Dixie Butte Metaandesite is estimated to be about 2,100 m thick, with 

the lowermost 600 m being mainly thick-bedded andesitic lithic tuff, tuff breccia, and coarse tuffaceous 

sedimentary rocks (Brooks and others, 1984). Scattered round chert clasts in the basal sedimentary 

rocks indicate that the Dixie Butte Meta-andesite is 

depositional onto older ocean-floor material made up 

of chert and argillite. Melange with chert and argillite is exposed along the eastern and southern flanks 

of Dixie Butte (see fig. 9.3; Brooks and others, 1984).

GEOCHEMISTRY

Dixie Butte basalts and basaltic andesites are calcalkaline (table 9.5) and have volcanic-arc affinities 

(Mullen, 1982). They differ from the Bourne subterrane volcanic-arc rocks in that the Dixie Butte rocks 

have considerably higher abundances of K2O and Zr 

(tables 9.3-9.5).

AGE AND REGIONAL CORRELATIONS

The eastern exposures of the Dixie Butte Meta-andesite are in high-angle contact with a serpentinitematrix melange similar in appearance to the melange 

of Vinegar Hill. This melange contains diverse lithologic blocks whose compositions range from greenschist to foliated epidote amphibolite. The melange 

also contains porphyritic andesite blocks and Permian 

(Guadalupian) ribbon cherts (Brooks and others, 

1984; Blome and others, 1986).

Chert clasts presumably derived from the underlying Permian ribbon cherts are present in tuffaceous 

sedimentary rocks exposed near the base of Dixie 

Butte. The rounded chert clasts indicate that the 

Dixie Butte flows and tuffs were emplaced after the 

cherts were eroded, rounded, and redeposited, and 

are evidence that the Dixie Butte is younger than 

Guadalupian in age.

The low metamorphic rank and absence of S^ 

cleavage further suggest that the Dixie Butte Metaandesite may be younger than all the other Paleozoic

and Triassic components of the Baker terrane and 

may be one of the postulated northern volcanic sources for andesitic tuffs and lava flows in the Late Triassic Fields Creek Formation in the Izee terrane to the 

south (Brown and Thayer, 1966). High K^O contents 

suggest that the Dixie Butte rocks are not related to 

the Wallowa terrane (Mortimer, 1986).

METAMORPHISM

Metamorphic recrystallization and hydrothermal 

alteration due to emplacement of numerous small Late 

Jurassic and early Tertiary intrusions (Brooks and 

others, 1984) have generally obscured early metamorphic mineral assemblages at Dixie Butte. Regional metamorphic assemblages include chlorite+albite+calcite+ 

zeolite and chlorite+albite+actinolite.

STRUCTURE

Major structural characteristics of the Bourne and 

Greenhorn subterranes are summarized below. For a 

more detailed treatment of structural development in 

the Baker terrane, the reader is referred to Coward 

(1983), Evans (chap. 8, this volume), and Hunt 

(1985).

Four major structural events can be recognized: (1) 

early disruption with development of type II melanges and cataclasite zones in the Bourne subterrane 

and development of serpentinite-matrix melange and 

ophicarbonates in the Greenhorn subterrane, (2) folding of Bourne and Greenhorn subterrane units about 

north- and northeast-trending fold axes without development of penetrative cleavages, (3) later development of penetrative and subpenetrative fracture 

cleavages (S^ axial planar to east-west-trending 

folds in both subterranes, and (4) local development 

of penetrative and spaced cleavages (82), mainly in 

the matrixes of melange zones in both subterranes.

PLANAR ELEMENTS

Bedding is preserved only locally in large Bourne 

subterrane blocks where bedded sequences are separated by cataclasite or shear zones. Both Coward 

(1983) and Evans (chap. 8, this volume) describe the 

Elkhorn Ridge Argillite as a broken formation whose 

initial fragmentation predated development of penetrative and subpenetrative cleavages. Bedding (S0) in 

the Bourne subterrane is most common in the layered cherts of the Elkhorn Ridge Argillite and lay-
348 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

ered cherts in the Olive Creek area; bedding is only 

rarely preserved in argillites of the Elkhorn Ridge 

Argillite and finer grained sedimentary fades of the 

Goodrich Creek and North Fork units. Inverted graded beds on the east end of Elkhorn Ridge indicate at 

least local overturning of units in the Goodrich Creek 

area (Ferns and others, 1987). Attitudes of bedding 

planes in the greenstones are more varied than those 

of bedding planes in the adjacent Elkhorn Ridge Argillite, which generally strike east-west and dip 

steeply to the south.

Bedding is only locally apparent in Greenhorn 

subterrane units. Inverted graded beds and flame 

structures in the Badger Creek unit suggest some 

overturning of units. Bedding-plane relationships 

are difficult to determine in the Dixie Butte Metaandesite, owing to the massive nature of flow units. 

Vhay (1960) notes that steeply dipping and locally 

overturned tuff beds indicate extensive folding of the 

Dixie Butte. Basal tuffs at Dixie Butte generally 

strike to the north and dip to the east and west.

FAULTS, MICROBRECCIAS, AND CATACLASITE ZONES

Pervasive, apparently randomly oriented, anastomosing microbreccia zones as much as 1 cm in width 

cut all rock types in the Goodrich Creek unit. Microbreccia zones are made up of angular and flattened 

porphyroclasts arranged in an isotropic groundmass. 

Extensive areas of protocataclasites mark contacts 

within the Goodrich Creek between metamorphosed 

intrusive rocks and all other rock types. The protocataclasites are zones as much as a hundred meters 

wide in which protoliths are difficult to distinguish. 

The protocataclasite zones contain granular pophyroclasts in an isotropic matrix.

Narrow cataclasite zones, with partially recrystallized matrix, and microbreccias also cut across the 

Elkhorn Ridge Argillite, which bears a strong resemblance to Cowan's (1985) type II melange. Coward 

(1983) noted that bedding and folds were transposed 

by movement along shear planes that are subparallel 

or at low angles (<45°) to bedding or axial-plane surfaces. Evans (chap 8, this volume) notes that the 

shear zones surround unsheared blocks that are generally elongate parallel to both bedding (S0) and 

cleavage (Si).

The shear planes and cataclasite zones apparently 

parallel SQ in the Elkhorn Ridge Argillite. Orientation and distribution of the Goodrich Creek cataclasite zones have not been determined. In outcrop, 

multiple generations of apparently randomly oriented 

shears may indicate progressive rotation or grinding

of blocks within the Goodrich Creek cataclasite 

zones.

Cataclasites comparable to those in the Bourne 

subterrane have not been recognized in the Greenhorn subterrane. Mullen (1978) noted cataclasites in 

metagabbros along the boundary fault separating the 

subterranes, whereas Hunt (1985) noted a lack of 

cataclasites in blocks of the melange of Vinegar Hill, 

about 5 km to the south.

CLEAVAGES

The most prominent structural element in nearly 

all the units of the Baker terrane is a generally eastwest-striking, south-dipping spaced fracture cleavage 

(S{). Si is best developed in the finer grained argillites of the Elkhorn Ridge Argillite and the Goodrich 

Creek unit in the Bourne subterrane and the Badger 

Creek unit in the Greenhorn subterrane.

Si in stratigraphically coherent ribbon-chert blocks 

of the Elkhorn Ridge Argillite generally parallel the 

narrow microbreccia and cataclasite zones that 

bound the blocks. Similarities in orientation of SQ 

and Si in the Elkhorn Ridge Argillite suggest that 

bedding may be transposed or that S^ is axial planar 

to east-west-trending folds in the Elkhorn Ridge (F2 

of Ave Lallemant and others, 1980). S^ cuts across 

cataclasite zones in the Goodrich Creek unit and 

tends to parallel the trend of the faults that separate 

the Goodrich Creek unit from overlying and underlying slabs of Elkhorn Ridge Argillite. Si attitudes in 

the Goodrich Creek unit, although varied, generally 

parallel the east-west striking and south-dipping Si 

of the Elkhorn Ridge. Superposition of Si on the cataclasite zones in the Goodrich Creek unit indicates 

that the greenstone cleavage developed after formation of the cataclasite zones. Si fracture cleavage in 

the Olive Creek rocks is poorly developed in comparison to Si in the Elkhorn Ridge Argillite. Si in the 

Olive Creek unit appears to be axial planar to eastwest-trending fold axes and parallels the strike of 

the thrust faults that separate the Elkhorn Ridge 

slices from enclosed Goodrich Creek and North Fork 

slices (see Ferns and others, 1987). Steep Si cleavages in greenstone units also generally parallel the predominantly east-west regional strike of the faults.

Si in the Greenhorn subterrane is best developed 

in argillites of the Badger Creek unit as a penetrative slaty cleavage. Si in coarser grained Badger 

Creek conglomerates is defined by subparallel alignment of anastomosing stringers of sheet silicates 

(Hunt, 1985). Si generally cuts bedding in the Badger Creek rocks and strikes east-west. The general
9. BOURNE AND GREENHORN SUBTERRANES OF THE BAKER TERRANE, NORTHEASTERN OREGON 349

east-west strike of Si in the Greenhorn subterrane is 

similar to that of Si fracture cleavage in the Bourne 

subterrane, and the two sets of cleavage are presumed to be of the same age. Steep northerly dips 

are more common in the Greenhorn subterrane units, 

and the average dip of Si is closer to vertical in the 

Greenhorn than in the Bourne subterrane units (fig. 

9.11). Si is poorly developed in greenstones, is defined by widely spaced fracture cleavage in blocks of 

the melange of Vinegar Hill (Hunt, 1985), and has 

not been recognized in the Dixie Butte Meta-andesite. The intervening serpentinite matrix of the melange of Vinegar Hill apparently lacks the Si 

cleavage and is cut by a younger, phacoidal cleavage 

(S2). Hunt (1985) suggested that the similarity in Si 

orientation in different Vinegar Hill blocks indicates 

that serpentinite was injected after development of 

Si with little or no rotation of individual blocks. Alternatively, early Si fabrics in the serpentinite maFlGURE 9.11. Lower-hemisphere, equal-area plot comparing visually determined mean poles to bedding (S0) and cleavage 

(Si). Ens1( 8} cleavages from outcrops of the Elkhorn Ridge 

Argillite north of the Goodrich Creek unit (number of data 

points (n)=36); Est, S^ cleavages from outcrops of Elkhorn 

Ridge Argillite in area between the Goodrich Creek unit and 

McCully Fork mixed-rock zone (n=108); EoSj, Sx cleavages 

from outcrops in the Elkhorn Ridge Argillite and Olive Creek 

unit south of the McCully Fork mixed-rock zone (n=40); ESO , 

bedding-plane orientations in Elkhorn Ridge Argillite 

(n=157); GSI( Sx cleavages from the Goodrich Creek unit 

(n=31); BSI; S} cleavages from the Badger Creek unit (n=48). 

Note the progressive steepening of cleavage from northern to 

southern parts of the study area.

trix may have been destroyed by serpentinite flowage 

during subsequent deformation.

Evans (chap. 8, this volume) notes that the shear 

zones surround unsheared blocks that are generally 

elongate parallel to bedding (S0) and fracture cleavage (Si). Coward (written commun., 1993) suggested 

that these features indicate isoclinal folding and brittle shortening that has formed a type II melange 

(Cowan, 1985) in which stratigraphically continuous 

zones of isoclinally folded chert and argillite are 

bounded by either faults or cataclasite zones.

A younger cleavage (S2) affects some Baker terrane 

units. Coward (1983) noted a northerly dipping 

spaced cleavage that cuts Si of the Elkhorn Ridge 

Argillite; the spaced cleavage is defined by alignment 

of fine-grained phyllosilicates. S2 postdates the Si 

fracture cleavage and may have formed in response 

to emplacement of the Late Jurassic to Early Cretaceous intrusions.

A phacoidal cleavage defined by augen of unsheared serpentinite in intensely sheared serpentinite was recognized by Hunt (1985) in the melange of 

Vinegar Hill. The strike of the cleavage parallels that 

of contacts between matrix and blocks in the melange of Vinegar Hill and appears randomly oriented 

(Hunt, 1985). Serpentinite injected along fault zones 

south of Hunt's area exhibits a strongly developed 

shear foliation that parallels both the direction of 

elongation of metamorphic and sedimentary blocks 

and the Si fracture cleavage within the blocks 

(Brooks and others, 1983). A similar matrix cleavage 

(herein defined as S2) is developed in the McCully 

Fork mixed-rock zone and parallels the regional 

strike of the zone and the direction of elongation of 

included slabs. S2 occurs in talc-amphibole hornfels 

within the contact aureole of the Bald Mountain 

batholith and apparently developed in conjunction 

with recrystallization of serpentinite matrix in the 

McCully Fork mixed-rock zone during batholith emplacement. Late development of the McCully Fork 

mixed-rock zone is suggested by Coward (1983) and 

supported by narrow cataclasite zones in the underlying Elkhorn Ridge Argillite that parallel the fault 

contact between the Elkhorn Ridge Argillite and the 

McCully Fork mixed-rock zone, and cut across Si of 

the Elkhorn Ridge Argillite.

FOLDS

Mesoscopic folds are relatively uncommon in all 

Bourne subterrane units. Two sets of fold axes (fig. 

9.12) have been identified in the Elkhorn Ridge Argillite: a steeply plunging north-south-trending set
350 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

and a more gently plunging east-west-trending 

set (F2) (Ave Lallemant and others, 1980).

F! includes rare open folds in the Goodrich Creek 

unit and tight isoclinal folds in Elkhorn Ridge and 

Olive Creek cherts. Fj axes plunge 30° to 35° to the 

south and parallel the southeast-plunging outcrop 

pattern of metamorphosed intrusive rocks of the 

Goodrich Creek unit on the east end of Elkhorn 

Ridge (Ferns and others, 1987).

F2 fold axes are nearly horizontal and trend eastwest. Coward (1983) noted two varieties: prelithification folds that lack axial-planar cleavage (his Flnc), 

and folds with axial-planar cleavage (his Flwc). Fold 

axes of both varieties are oriented east-west with 

southerly dipping axial planes and are herein designated F2. Similarly oriented F2 folds that occur in 

the Goodrich Creek unit include small folds with 5- 

cm wavelengths in contorted limestones and folded 

contacts between Goodrich Creek rock types.

Mesoscopic Fj folds have not been identified in 

Greenhorn subterrane units. Possible north- to northeast-trending F! axes are suggested by overturning 

of SQ in the Badger Creek unit and the Dixie Butte 

Meta-andesite. A thrust fault separating the Badger 

Creek unit from ultramafic rocks on Badger Creek

FIGURE 9.12. Lower-hemisphere, equal-area plot of fold axes. 

Filled symbols represent Fj axes; open symbols, F2 axes. 

Squares represent data from Elkhorn Ridge Argillite; circles, 

data from Olive Creek unit; triangles, data from Goodrich 

Creek unit.

(fig. 9.13) may also be folded about a northeast-trending axis (Ferns and others, 1983).

F2 folds related to an axial-planar cleavage have been 

recognized by Hunt (1985) in the Badger Creek unit 

along Snow Creek (fig. 9.9). Axial planes of tight, upright symmetrical folds parallel the northeast strike 

and near-vertical dip of the cleavage here (Hunt, 1985).

Mullen (1978) recognized a later folding event 

(herein referred to as F3) affecting the amphibolitic 

blocks of the melange of Vinegar Hill. Axial planes to 

F3, which consists of small isoclinal folds and crenulations in foliated metamorphic rocks, apparently 

strike northwest (Mullen, 1978), parallel to the subterrane boundary.

STRUCTURE SUMMARY

While gross similarities in late structural styles 

and fracture cleavages in the Bourne and Greenhorn 

subterranes indicate contemporaneous late deformation, distinctly different styles of early deformation 

indicate separate early-deformation paths for the 

subterranes. Early deformation in the Bourne subterrane was characterized by severe internal disruption 

of units, including the development of type II melange in the Elkhorn Ridge Argillite and the development of cataclasite zones in the Goodrich Creek unit. 

Later structures include poorly preserved folds about 

north-trending fold axes (Fj), thrust faults, and 

axial-planar cleavage (S]\ about east-west-trending 

F2 fold axes. In the Greenhorn subterrane, development of serpentinite-matrix melanges was followed 

by clastic basinal sedimentation prior to thrust faulting and folding about north- to northeast-trending Fj 

axes. The folds are cut by the penetrative Sj cleavage 

in some but not all Greenhorn subterrane units. A 

later cleavage (S2) is locally developed in melange 

matrixes in both subterranes and is apparently related to remobilization and recrystallization of serpentinite accompanying Late Jurassic plutonism.

SUMMARY

The Bourne subterrane of the Baker terrane is composed mainly of southward-dipping, fault-bounded 

packages of the Elkhorn Ridge Argillite, a sedimentary sequence made up chiefly of deepwater chert and 

siliceous argillites of Permian and Late Triassic (Karnian) age. These deepwater cherts and argillites enclose olistoliths of shallow-water limestones that 

range from Devonian to Permian in age (fig. 9.14). 

Although most of the known Permian fusulinid locali-
9. BOURNE AND GREENHORN SUBTERRANES OF THE BAKER TERRANE, NORTHEASTERN OREGON 351

118°45'

Tc

.' .:T*Pbc:

EXPLANATION

Cover deposits (Tertiary)

Intrusive rocks (Early Cretaceous and Late Jurassic) 1~

Badger Creek unit (Late Triassic? to Pennsylvanian) f~

_i_

Serpentinized peridotite *

Metagabbro t

Thrust fault Sawteeth on upper plate

Anticline

Syncline

Strike and dip of bedding (S0)

Strike and dip of foliation (S^

Strike of vertical foliation (S^

FIGURE 9.13. Relations between bedding (S0), cleavage (Sj), and intercalated ultramafic sheets in the Badger Creek unit. A, Geologic map. B, Lower-hemisphere, equal-area projection of poles to bedding (S0) in the Badger Creek unit. C, Lower-hemisphere, 

equal-area projection of poles to cleavage (Si) in the Badger Creek unit.

ties yielded fusulinids of Tethyan affinities, one locality with fiisulinids similar to those of the McCloud 

Limestone in northern California has been described 

in the literature. Slices of disrupted volcanic rocks of 

island-arc affinities are intercalated with the Elkhorn 

Ridge Argillite along southward-dipping high-angle 

reverse and thrust faults along the northern margin 

of the Bourne subterrane. The island-arc packages

consist of Late Permian to Early Triassic intrusive 

and extrusive rocks and Late Triassic (Norian) sedimentary rocks. Alkalic pillow basalts are intercalated 

with Permian and Triassic cherts along the southern 

margin of the subterrane.

Bourne subterrane rocks are strongly faulted, locally brecciated, and cataclasized, and they generally 

correspond to the type II melanges of Cowan (1985)
352 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

(tables 9.2, 9.6). Strongly disrupted mixed-rock zones 

resembling the type IV melanges of Cowan (1985) cut 

across the central and southern parts of the Bourne 

subterrane. The Elkhorn Ridge Argillite along the 

southern margin of the Bourne subterrane has locally been converted to large megabreccia zones. A generally south-dipping fracture cleavage (Sj) affects all 

Bourne subterrane units.

The Greenhorn subterrane of the Baker terrane is 

composed mainly of serpentinite-matrix melange that

118°30

118°

10 MILES

10 KILOMETERS

EXPLANATION

Cover deposits (Cenozoic)

Intrusive rocks (Early Cretaceous to Late Jurassic)

Bourne subterrane Consists of:

Elkhorn Ridge Argillite (Triassic to 

Pennsylvanian) Mainly Permian and Late 

Triassic chert and argillite; also includes 

Pennsylvanian, Permian, and Late Triassic 

limestone

Volcanic rocks (early Mesozoic and 

Paleozoic) Island-arc-related rocks of the 

North Fork (nf) and Goodrich Creek (gc) units, 

and oceanic rocks of the Olive Creek (oc) unit

McCully Fork mixed-rock zone

Greenhorn subterrane Consists of:

Melange of Vinegar Hill As mapped, includes 

some unnamed serpentinite and serpentinitematrix melanges

Badger Creek unit (Mesozoic? and Paleozoic) As 

mapped, includes some unnamed sedimentary 

rocks

Dixie Butte Meta-andesite of Brooks and others 

(1984) (Mesozoic?) Island-arc volcanic rocks

FIGURE 9.14. Fossil and amphibolite localities in central part of Baker terrane. Fossil localities: D, Devonian; P, Pennsylvanian; P, 

Permian; ~fi, Triassic, queried where uncertain. Paleontological ages are from Blome and others (1986), Morris and Wardlaw 

(1986), Ferns and others (1987), C.D. Blome (written commun., 1987), and M.K. Nestell, (oral commun., 1986). a, foliated amphibolite localities.
9. BOURNE AND GREENHORN SUBTERRANES OF THE BAKER TERRANE, NORTHEASTERN OREGON 353

TABLE 9.6. Comparison of lithologic, structural, and paleontologic characteristics of the Greenhorn and Bourne subterranes

Greenhorn subterrane Bourne subterrane

Main lithologic unit Serpentinite-matrix melange Chert and argillite (Elkhorn Ridge Argillite).

(melange of Vinegar Hill).

Volcanic rock types High-K calc-alkaline volcanic rocks High-Na calc-alkaline volcanic rocks (Goodrich

(Dixie Butte Meta-andesite); calc- Creek and North Fork units); alkalic pillow 

alkaline, alkalic, and oceanic lavas (Olive Creek unit); ocean-floor(?) 

melange blocks (melange of tholeiites (McCully Fork mixed-rock zone). 

Vinegar Hill).

Intrusive rock types Mafic gabbros, minor pyroxenites Low-K diorites, quartz diorites, and gabbros

and harzburgites (melange of similar to the Sparta complex (Goodrich

Vinegar Hill). Creek and North Fork units).

Melange types (see table 9.2) Serpentinite-matrix block-in-matrix, Type IV, brittle fault zones (McCully Fork

analogous to type III. mixed-rock zone); type II, disrupted chert

and argillite sequences (Elkhorn Ridge 

Argillite).

Faunal ages Pennsylvanian, Permian, and early Devonian, Pennsylvanian, Permian, and Late

Mesozoic(?) (Badger Creek unit); Triassic limestone blocks and Permian and 

Permian and early Mesozoic(?) Late Triassic cherts (Elkhorn Ridge 

blocks (melange of Vinegar Hill). Agillite); Late Triassic limestones (Goodrich

Creek unit).

Faunal affinities McCloud, Calif., area (Badger Creek Separate blocks withTethyan and McCloud

unit). (Calif.) area affinities (Elkhorn Ridge

Argillite).

Metamorphism Epidote-amphibolite blocks with Greenschist facies in all units.

blue amphibole (melange of 

Vinegar Hill).

Sedimentary rocks Interbedded sandstone, siltstone, Chert and argillite with limestone olistoliths

and limestone (Badger Creek (Elkhorn Ridge Argillite), tuffaceous 

unit). siltstones and limestones interbedded with

tuffs and volcaniclastic rocks (Goodrich

Creek unit).

Proposed tectonic environment Fore arc- Accretionary wedge/marginal basin.

contains a wide variety of metamorphic and igneous 

rocks, including blue-amphibole-bearing epidote amphibolites (fig. 9.14), lower greenschist and greenschist facies greenstones, and a variety of island-arc, 

oceanic-island (seamount?), and tholeiitic pillow basalts. Associated Paleozoic and Mesozoic(?) clastic sedimentary rocks include coarse polymict breccias 

derived from blocks within the melange, and limestone lenses containing Permian fusulinids with affinities to those of the McCloud Limestone of northern 

California. Overall, the Greenhorn subterrane bears a 

strong resemblance to modern fore-arc zones in the 

Marianas and Bonin arcs, where various island-arc 

and oceanic volcanic rocks are intimately associated

with serpentinite diapirs, polymict breccias, and clastic basinal sediments.

DISCUSSION

Previous workers (Dickinson and Thayer, 1978; 

Brooks, 1979; Hietanen, 1981), citing the blueschist 

blocks at Mitchell, Oreg., and the mixed Tethyan and 

non-Tethyan fusulinid faunas, suggested that the 

Baker terrane represents a paleosuture zone. Dickinson and Thayer (1978) postulated that the suture is 

the product of subduction that lasted until at least 

Middle Triassic time. Hietanen (1981) noted that
354 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

long-lasting subduction during late Paleozoic to early 

Mesozoic time also occurred in the Sierra Nevada 

and Klamath Mountains in northern California and 

southwestern Oregon, and she suggested that rocks 

of the Baker terrane are an extension of the Sierra 

Nevada-Klamath suture system.

Our work, in conjunction with that of Mullen 

(1978), Evans (chap. 8, this volume), and Coward 

(1983) leads us to suggest that the Bourne and 

Greenhorn subterranes evolved along separate earlydeformation paths and represent different tectonic 

environments within the suture zone. The Greenhorn 

subterrane appears to have been a fore-arc high 

(Mullen, 1985), as evidenced by serpentinite-matrix 

melanges, clastic basinal sedimentary rocks, mixed 

island-arc and oceanic volcanic rocks, and moderately 

high pressure/low temperature foliated metamorphic 

rocks. Permian fusulinid assemblages are similar to 

those of the Coyote Butte area in the Grindstone terrane to the southwest (fig. 9.1) and the McCloud area 

in the eastern Klamath Mountains in northern California, and may indicate a link between the Grindstone terrane and the Greenhorn subterrane. The 

apparent age range of the basinal supracrustal rocks 

of the Badger Creek unit from Pennsylvanian to 

possibly Triassic time suggests that the Greenhorn 

subterrane was a site of active sedimentation over a 

long time period.

Intercalation of Permian to Late Triassic deepwater chert and siliceous argillite packages with Permian to Late Triassic island-arc packages in the Bourne 

subterrane suggests significant post-Late Triassic 

crustal shortening during tectonic mixing of coeval 

island-arc and oceanic material. The Permian to Late 

Triassic island-arc packages may correlate with the 

Wallowa terrane now exposed immediately to the 

north. Volcanic rocks exposed in the central and 

southern portions of the Bourne subterrane have oceanic characteristics and may represent oceanic-island 

or ocean-floor material peripheral to the island-arc 

rocks of the Wallowa terrane.

Cowan (1985) noted that type II and type IV melanges might develop in an accretionary wedge setting above a subduction zone. Assuming that 

boundary faults separating lithologic units within the 

Bourne subterrane dip toward the subduction zone, 

then the sense of subduction, discounting postsuture 

rotation, would have been roughly from the north 

down to the south. The present distribution of intercalated island-arc slabs in the north and oceanic 

slabs in the south may record the impingement of a 

Permian to Late Triassic island-arc system (Wallowa 

terrane?) on a southward-dipping subduction zone in 

the post-Late Triassic (fig. 9.15).

Evans (chap. 8, this volume) and Coward (1983) 

recognize early, nonpenetrative deformations in the 

Bourne subterrane that may record early development of argillaceous broken formation (type II melange of Cowan, 1985) in an accretionary wedge 

setting. Tethyan Permian fusulinids in the Bourne 

subterrane indicate an environmental or paleogeographic setting distinct from that of the Greenhorn 

subterrane. Evans (chap. 8, this volume) indicates 

that rocks which constitute the Bourne and Greenhorn subterranes in the Desolation Butte quadrangle 

underwent separate deformations prior to their juxtaposition in Late Triassic or Early Jurassic time. 

The presence of Late Triassic cherts in the Elkhorn 

Ridge Argillite and Late Triassic limestones in 

Bourne subterrane island-arc packages indicates that 

nonpenetrative deformation continued through the 

Late Triassic as the Bourne subterrane rocks were 

tectonically intercalated within the suture zone.

Other workers (Evans, chap. 8, this volume; Ave 

Lallemant and others, 1980) have recognized an early, locally penetrative deformation that affected some 

Bourne subterrane components in the Baker terrane 

and is apparently related to a regionally extensive 

Late Triassic to Early Jurassic erogenic event (Dickinson and Thayer, 1978; Ave Lallemant and others, 

1980). We suggest that this event may represent the 

cessation of subduction within the suture zone due to 

the incomplete consumption of the lighter (as compared to oceanic) island-arc material now preserved 

on the North Fork of the John Day River and on 

Elkhorn Ridge. This interpretation is suggested by 

extensive cataclasite zones in the Bourne subterrane. 

The Late Triassic to Early Jurassic event may be recorded in the Greenhorn subterrane by the poorly 

documented north- to northeast-trending fold axes in 

the Badger Creek unit and the Dixie Butte Meta-andesite of Brooks and others (1984). The northerly 

trending Greenhorn subterrane fold axes are similar 

in orientation to northerly trending fold axes identified by Dickinson and Thayer (1978) and Ave Lallemant and others (1980) in Late Triassic and Early 

Jurassic clastic sedimentary rocks in the Izee terrane 

to the south. Cessation of subduction at about this 

time may be marked by the 225-Ma blueschist block 

near Mitchell, Oreg. (Hotz and others, 1977).

The readily apparent, widespread, southerly dipping S^ cleavage records a regional, short-lived, penetrative deformation event in the Late Jurassic. Sj 

cleavage is apparent in most of the Baker terrane 

components as well as the overlying Jurassic clastic 

basinal sediments of the Olds Ferry terrane to the 

south. BI apparently developed in conjunction with 

the thrust and high-angle reverse faults that now
9. BOURNE AND GREENHORN SUBTERRANES OF THE BAKER TERRANE, NORTHEASTERN OREGON 355

GRINDSTONES) TERRANE 

GREENHORN SUBTERRANE

"McCLOUD" FAUNAL REALM

-BOURNE SUBTERRANE 

WALLOWA(?) TERRANE

TETHYAN FAUNAL REALM

SW

BAKER TERRANE

IZEE AND

OLDS FERRY

TERRANES

-BOURNE SUBTERRANEGREENHORN SUBTERRANEWALLOWA 

TERRANE

NE

SW

EXPLANATION

Greenhorn subterrane Includes: Bourne subterrane Includes: 

Badger Creek unit ||||||||| Elkhom Ridge Argillite

Melange of Vinegar Hill Metamorphosed 

to amphibolite facies in basal part

Dixie Butte Meta-andesite of Brooks and 

others (1984)

Olive Creek unit

Goodrich Creek and North Fork units, 

undivided

Contact Approximately located 

Reverse fault

Surface of active subduction Arrows 

show relative movement

FIGURE 9.15. Schematic diagram showing possible evolutionary 

path for development of Baker terrane. Compass directions, 

on diagrams and in caption, are approximate and are in 

present-day coordinates. A, The Pennsylvanian to the Late 

Triassic was a time of active subduction of open-ocean and 

(or) Wallowa terrane back-arc basin material beneath a volcanic arc southwest of the Greenhorn subterrane. Significant 

events in formation of Greenhorn subterrane units include 

deposition of Badger Creek unit as basinal sediments in a 

faunal realm similar to that of the McCloud Limestone, initial protrusion of serpentinite diapirs in the melange of Vinegar Hill, initial formation of high-pressure metamorphic rocks 

beneath the melange of Vinegar Hill, and (not shown on this 

diagram) eruption of the Dixie Butte Meta-andesite of Brooks 

and others (1984). Significant events in formation of Bourne 

subterrane units include deposition and soft-sediment deformation of Elkhorn Ridge Argillite, incorporation of Tethyan 

limestone olistoliths into Elkhorn Ridge Argillite, formation of 

alkalic seamounts of the Olive Creek unit in the open ocean, 

and formation of volcanic and sedimentary rocks of the Goodrich Creek and North Fork units as parts of the Wallowa arc. 

B, In Late Triassic to Early Jurassic time, units of the 

Bourne and Greenhorn subterranes were amalgamated into 

the Baker terrane owing to clogging of the subduction zone by 

the Wallowa terrane. Important events include incorporation 

of fragments of Wallowa terrane (Goodrich Creek and North 

Fork units) into the Bourne subterrane, initial development 

of cataclasite zones within Goodrich Creek unit, folding of all 

Baker terrane units about north-south and northeast-southwest fold axes, cessation of deposition of basinal sediments of 

Badger Creek unit, and shifting of depocenters to Izee and 

Olds Ferry terranes to the southwest.
356 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

separate lithic packages in the central part of the 

Baker terrane (Ave Lallemant and others, 1980; 

Engh, 1984).

In conclusion, there is evidence for four separate 

periods and (or) styles of deformation within the suture zone. (1) Early, long-lasting subduction-related 

deformations were marked by soft-sediment deformation, incorporation of limestone olistoliths, and development of type II melange in the Elkhorn Ridge 

Argillite of the Bourne subterrane and by development of polymict breccias in the Badger Creek unit of 

the Greenhorn subterrane. Subduction-related deformation may have started as early as the Pennsylvanian. (2) A later, locally penetrative deformation in 

the Late Triassic to Early Jurassic was related to incomplete subduction of the Bourne subterrane islandarc rocks, as evidenced by development and folding of 

cataclasite zones in the island-arc packages, development of mixed-rock zones (type IV melanges) in the 

Bourne subterrane, and development of north-trending fold axes in the Greenhorn subterrane. (3) Imposition of regionally extensive fracture cleavage took 

place in conjunction with development of major thrust 

and high-angle reverse faults in the Late Jurassic. (4) 

Melange matrixes were recrystallized and remobilized 

during emplacement of Late Jurassic intrusions.

The above discussion does not take into account 

the 60°+ postaccretion clockwise rotation of the Blue 

Mountains region identified by Wilson and Cox 

(1980). Assuming no significant rotation of the Baker 

terrane components during amalgamation from Late 

Triassic to Jurassic time, the original tectonic configuration would have placed an island arc (Wallowa 

terrane) to the northwest of the Bourne subterrane. 

The sense of subduction would have been from west 

to east, with the Wallowa terrane and intervening 

Bourne subterrane collapsing against the Olds Ferry 

terrane and flooring the Greenhorn subterrane to the 

east in the Late Triassic (see Dickinson, 1979; Hietanen, 1981).

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358 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

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10. AN 40Ar/39Ar CHRONICLE OF THE TECTONIC DEVELOPMENT OF 

THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO

By LAWRENCE W. SNEE, KAREN LUND, JOHN F. SUTTER, DAVID E. BALCER/ and KARL V. EVANS

CONTENTS

Page

Abstract 359

Introduction 359 

Acknowledgments 360 

Geologic setting - - 360 

Analytical methods - - 366 

40Ar/39Ar techniques 366

U-Pb techniques - - -- 368

Sample distribution and isotopic data 369 

Interpretation 369 

Basement rocks of the Snake River canyon area and

associated plutons of island-arc affinity 369 

Metamorphic rocks of the Riggins Group and Seven Devils

island arc 385

Deformed plutons of the Salmon River suture zone 396

Undeformed plutons of the Idaho batholith 398

Discussion 402

Metamorphism and plutonism

in eastern Oregon and western Idaho 403 

Cooling and uplift history of the area east of the

Salmon River suture zone 406 

Depth of emplacement of muscovite-biotite granite 407 

Erosion and Cretaceous sedimentation 408 

Summary of Cretaceous tectonic history central Idaho and

southwestern Montana 409 

Conclusion 409

ABSTRACT

In west-central Idaho, rocks of island-arc origin are in sharp 

contact with North American continental rocks. The contact, 

called the Salmon River suture zone, is marked by lithologic and 

chemical differences, and rocks directly adjacent to the suture 

zone are characterized by high-grade metamorphism and structural complexity. The origin and age of the suture zone have been 

the subject of debate. 40Ar/39Ar age-spectrum data for hornblende, 

biotite, muscovite, and microcline from 48 samples of metamorphic and plutonic rocks in both island-arc and continental settings, and U-Pb data for one sample of continental plutonic rock, 

record the geologic development of this area between 244 and 55 

Ma. Rocks of island-arc origin, which include the Riggins and 

Seven Devils Groups, contain evidence for a period of metamoriChevron, USA, 935 Gravier St., Rin. 910, New Orleans, LA 70112

phism at about 244 Ma and two periods of plutonism at about 227 

and 145 Ma, all of which occurred before the island-arc rocks 

were accreted to the North American continent. Accretion of the 

island-arc terrane took place along the Salmon River suture zone 

in western Idaho and is recorded by at least four episodes of metamorphism at 130, 118, 109, and 101 Ma. Minor plutonism took 

place within this interval; the most extensive of this plutonism 

has been dated at about 115 Ma. The island-arc terrane was 

stitched to the North American continent at 93 Ma by the emplacement of tonalitic plutons, some of which were emplaced as 

deep as 20 km or more. These plutons are the first intrusions of 

the Idaho batholith, and chemically similar rocks were emplaced 

up to 139 km east of the suture zone until about 86 Ma. Porphyry 

mineralization is associated with this stage of plutonic emplacement. At or before 88 Ma, rapid uplift of continental rocks at 

rates greater than 3 mm/yr began along high-angle structures 

within and adjacent to the suture zone; this rapid uplift moved 

eastward about 3 m.y. later. Postdeformational cooling or uplift 

continued in some places along the suture zone until 79 Ma. More 

than 15 km of uplift affected a region extending at least 50 km 

eastward from the suture zone during a 10-m.y. period from about 

88 to 78 Ma; no plutonic activity during this uplift is recorded by 

the argon data. Beginning at about 78 Ma, muscovite-biotite granites of the Idaho batholith were emplaced at depths less than 9 

km; the granites cooled below 300°C by 70 Ma. From 78 to 68 Ma, 

mineralized quartz veins filled fractures and faults in roof pendants of the granites and are apparently genetically related to the 

granites. By 55 Ma, a large part of the presently exposed Cretaceous plutons of the Idaho batholith was at a crustal depth of 

about 5 km or less and had cooled to less than 130°C. The sedimentary record to the east and west of the Salmon River suture 

zone reflects this complex history of metamorphism, deformation, 

plutonism, and uplift that took place between 130 and 55 Ma.

INTRODUCTION

In western Idaho, oceanic rocks of island-arc origin 

are in sharp contact with continental units (Hamilton, 1963; Lund, 1984); the contact is north-trending 

and is thought to be the western extent of North 

American sialic crust (Criss and Fleck, 1987; Lund 

and Snee, 1988). Nonplutonic rocks west of this 

boundary, which was named the Salmon River suture 

zone by Lund and Snee (1988), consist mainly of volcanic and metavolcanic rocks of the Seven Devils and

359
360 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Riggins Groups (Hamilton, 1963; Vallier, 1977). To 

the east, the metasedimentary rocks are correlative 

with Belt Supergroup and pre-Belt units of Proterozoic age and with continental sedimentary units of uncertain, but probable Late Proterozoic or Paleozoic, 

age (fig. 10.1).

Although there is a clear difference in lithology 

across the boundary, the exact location of the lithologic change is generally obscured by high-grade metamorphism and by younger deformed and undeformed 

plutons that clearly crosscut the regional metamorphic fabric. However, within the plutonic rocks, the 

boundary is sharply defined by a major discontinuity 

in initial 87Sr/86Sr ratios. To the west, initial 

8?Sr/86Sr ratios are 0.704 or less; to the east, initial 

8?Sr/86Sr ratios are 0.706 or greater (Armstrong and 

others, 1977; Fleck and Criss, 1985; Davidson, 1989). 

In some ways, the discontinuity in initial 87Sr/86Sr 

ratios is the easiest way to trace the boundary because the discontinuity is sharply defined within the 

younger, crosscutting plutons. The discontinuity 

trends northward in Idaho from near McCall to Orofino where it abruptly swings westward (Armstrong 

and others, 1977; Fleck and Criss, 1985). To the 

south, the discontinuity is covered by Cenozoic rocks.

Even though a major suture zone has been suspected in this area since 1976 (Hamilton, 1976; Onasch, 

1976; Hyndman and Talbot, 1976), its tectonic origin 

remains obscure. In an attempt to provide some 

regional time and temperature constraints on the formation of this unusual crustal boundary in west-central Idaho, we have studied 40Ar/39Ar age spectra of 

hornblende, biotite, muscovite, and microcline from 

metamorphic and plutonic rocks within a large area 

extending from eastern Oregon to western Montana 

and straddling the suture zone. U-Pb data for zircon 

and sphene from one sample of tonalite from the 

Idaho batholith are also included. The purpose of our 

work is to determine the age of metamorphism, deformation, plutonism, cooling, and uplift of this large 

area and to use these data to unravel its tectonic evolution during the Cretaceous. In addition, we investigate whether a temporal relationship exists between 

the Cretaceous tectonic development of this area and 

the structural activity along the suture zone between 

the island-arc terranes and the North American continent. This project has been done in conjunction with 

a mapping study of the suture zone; the results of the 

mapping study are presented in Lund (1984; chap. 

14, this volume), and an interpretation of the tectonic 

development of the suture zone is discussed within 

the framework of the mapping and 40Ar/39Ar thermochronologic study in Lund and Snee (1988). In the 

following discussion, some of the details from previously published works are summarized and placed 

within the context of this study.

ACKNOWLEDGMENTS

The authors have benefited from the insight and 

ideas of E-an Zen, Tracy Vallier, Jane Selverstone, 

Robert Scholten, Cathy Manduca, Mel Kuntz, Anna 

Hietanen, Jane Hammarstrom, Warren Hamilton, 

Jad D'Allura, Gary Davidson, and Howard Brooks. 

Earl Bennett, Wayne Hall, Chris Gammons, and 

Thor Kiilsgaard kindly provided some samples from 

the Idaho batholith. Special thanks go to George 

Simmons for spectacular helicopter rides to remote 

sample sites in the Snake River canyon. Careful reviews of the manuscript were provided by Tracy Vallier and Howard Brooks.

GEOLOGIC SETTING

In the Riggins-Slate Creek area of west-central 

Idaho (fig. 10.2), pre-Cretaceous rocks east of the 

Salmon River suture zone are metasedimentary rocks 

whose protoliths were derived from terrigenous 

sources. These rocks have been divided into five tectonostratigraphic packages by Lund (1984). The metamorphic rocks are preserved as roof pendants and 

xenoliths in the Idaho batholith. The structurally 

lowest and probably oldest tectonostratigraphic unit 

is composed of mica schist and feldspathic gneiss 

with interleaved calcsilicate and quartzitic gneiss; it 

was intruded by igneous rocks that predate the Idaho 

batholith and are now amphibolite and augen gneiss 

(Lund, 1984). This lowermost unit is probably Middle 

Proterozoic in age on the basis of an inferred U-Pb 

zircon crystallization age of the augen gneiss of about 

1,370 Ma (Evans and Fischer, 1986). Above this unit 

lie two Middle Proterozoic(?) tectonostratigraphic 

units composed of calcsilicate gneiss and quartzite 

with a combined thickness of about 3,000 m. The uppermost two tectonostratigraphic units are probably 

either Late Proterozoic or Paleozoic in age and form

FIGURE 10.1. Index maps of region discussed in text. A, General 

geology and location map. Numbered areas are 1, Bald Mountain batholith; 2, Wallowa batholith; 3, The Oxbow on the Snake 

River; 4, Pittsburg Landing; 5, confluence of Imnaha and Snake 

Rivers; 6, Round Valley; 7, Riggins; 8, Slate Creek; 9, Orofino; 

10, McCall; 11, Buffalo Hump area; 12, Yellow Pine; 13, Stanley; 

14, Thompson Creek; 15, Warm Spring Forest Service Station; 

16, Warm Lake; 17, Pioneer Mountains; and 18, Hailey. B, Areas of figures 10.2 through 10.6, 10.21, and 10.23.
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 361

118° 116° 114° 112°

46°

44°

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REG

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(

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MONTANA

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100

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200 KILOMETERS

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EXPLANATION 

Cenozoic volcanic and sedimentary rocks

Tertiary and Mesozoic igneous rocks

Mesozoic and Paleozoic oceanic and island-arc volcanic, plutonic, and 

sedimentary rocks

Mesozoic and Paleozoic sedimentary rocks

Precambrian sedimentary, plutonic, and sedimentary rocks

Salmon River suture zone Queried where covered by younger rocks

118° 116° 114° 112°

1 !g

Snat^-Ol

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WASH \ |

OREG

Area of

figure 10.3 "

Area of figure 10.21 

i

PVi. 

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' &\%> ' '

"*^~»

Area of (

figure 10.2 /

- /-. .i

K VN

r/T"""V Area of \ I/J \ figure 10.23 

/ l x Area of ^*% / \ figure 10.6 >

1 Area of

v-» figure 10.4

' fe l''^

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B

46'

44'
362 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

116°30' 116°00'

45° 

45'

45° 

30'

45° 

15'

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10 KILOMETERS

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EXPLANATION

| Tb | Basalt (Tertiary) 

11 | |[ Undeformed plutons (Cretaceous) 

y^//fy Deformed plutons (Cretaceous) 

I- :: :| Seven Devils Group (Triassic and Permian) 

Riggins Group (Mesozoic or Paleozoic) 

Basement (?) rocks (Paleozoic) 

IIIIIIIIIIH Umbrella Butte unit (Paleozoic ?) 

Moores unit (Paleozoic ?) 

Quartzite Butte unit (Middle Proterozoic ?) 

Concord Butte unit (Middle Proterozoic ?)

Basal metasedimentary rocks unit (Middle 

Proterozoic ?)

FIGURE 10.2. Geologic map of study area along Salmon 

River suture zone in west-central Idaho. Geology modified from Lund and Snee (1988); stratigraphic nomenclature directly from Lund and Snee (1988, fig. 11-2).

Contact

Fault

Thrust fault Sawteeth on upper plate

Antiform Showing Crestline and 

direction of plunge

Synform Showing troughline and 

direction of plunge

 Initial 87Sr/86Sr discontinuity

Salmon River suture zone marked by discontinuity in initial 87Sr/86Sr ratio is shown as dot-and-dashed line. Geology 

of unpatterned areas is not compiled.
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 363

a distinctive multilithologic sequence composed of 

quartzite, quartzite-pebble quartzose conglomerate, 

calcsilicate gneiss, mica schist, and marble.

West of the suture zone, the metamorphic country 

rocks consist of felsic to mafic volcanic, volcaniclastic, 

intrusive, volcanogenic-sedimentary, and calcareous 

rocks (Hietanen, 1962: Hamilton, 1963; Onasch, 1977; 

Vallier, 1977; Myers, 1982; Lund, 1984). Nearest the 

suture zone, units of pre-Tertiary age are of low, middle, and high metamorphic grade and were named 

the Riggins Group by Hamilton (1963). The Riggins 

Group consists of two lower units composed of volcanic rocks and a structurally higher unit composed of 

carbonaceous and calcareous phyllite to schist and 

minor carbonate rocks. Some sheared serpentinite 

bodies are present along faults (Hamilton, 1963; 

Onasch, 1977).

The Riggins Group structurally overlies more extensive Permian and Middle and Upper Triassic 

rocks of the Seven Devils island-arc sequence of 

Dickinson (1979) (also known as the Wallowa terrane 

of Silberling and others, 1984). In the vicinity of 

Snake River canyon, rocks of the Seven Devils Group 

(Vallier, 1977) are predominantly very low grade to 

medium-grade metamorphosed volcanic rocks with 

minor interbedded volcaniclastic conglomerate and 

sedimentary rocks. Stratigraphically overlying the 

Seven Devils Group is a sequence of interfingered 

volcaniclastic, carbonate, and minor volcanic rocks 

(Lund and others, 1983a; Lund, 1984 and chap. 14, 

this volume). A similar sequence overlies the Seven 

Devils Group in northeastern Oregon (Folio and Siever, 1984; Folio, 1992, 1994). This similarity suggests 

that the change upward in the stratigraphic section 

from volcanic units of the Seven Devils Group to carbonaceous and calcareous shale and carbonate rocks 

(and in eastern Oregon, younger noncalcareous shale 

and conglomerate) is of regional extent.

Metamorphic grade both east and west of the 

Salmon River suture zone decreases away from the 

suture zone. East of the suture zone in the Riggins 

and Slate Creek areas the highest grade rocks are of 

upper amphibolite facies and occur in a band about 6 

km wide adjacent to the suture zone. The rocks of 

this band include sillimanite-bearing metapsammitic 

rocks, garnet- and sillimanite-bearing metapelitic 

rocks, and diopside-bearing calcsilicate gneiss. Because of stacking by thrust faulting, metamorphic facies are inverted to the east, where higher grade 

rocks structurally overlie lower grade rocks (Lund 

and Snee, 1988).

Metamorphic grade in the Riggins Group and 

Seven Devils island-arc sequence decreases westward 

from the suture zone. The three formations of the

Riggins Group are at lowest metamorphic grade near 

Riggins and increase in grade toward the northeast, 

east, and south (Hamilton, 1963; Onasch, 1977, 

1987). At lowest grades, greenschist-facies rocks are 

muscovite and (or) chlorite schist, with plagioclase, 

quartz, and epidote in the groundmass and, at slightly higher grade, with porphyroblasts of biotite flakes, 

unoriented hornblende needles, and carbonate 

rhombs (Hamilton, 1963). Closer to the Salmon River 

suture zone, the rocks are garnet-muscovite schist, 

biotite schist, garnet-hornblende-biotite schist, and 

garnet-amphibolite gneiss, depending on original 

composition. Along a narrow belt adjacent to the suture zone in this area, the rocks of the Riggins Group 

contain clear evidence of more than one dynamothermal event. This high-grade overprinted zone is about 

10 km wide, but it widens south of the Salmon River. 

In these rocks, garnet, hornblende, and other porphyroblasts were rolled during formation of a secondary 

schistosity (Lund and Snee, 1988; see also figs. 10.15 

and 10.16); growth of new minerals is indicated in 

many places by the presence of two generations of 

the same phase that have different orientations and 

(or) fabric characteristics. In some areas, fabric and 

complex crosscutting relations are consistent with 

emplacement of some dikes and small plutonic bodies 

between metamorphic pulses.

Along the Snake River (Vallier, 1977) and near Riggins (Hamilton, 1963), the Seven Devils Group is 

mostly metamorphosed volcanic and volcaniclastic 

rocks. Metamorphic grade in these rocks increases to 

the northeast from the Riggins area. In the Slate 

Creek area, the Seven Devils Group is at greenschist 

facies and includes rocks containing metavolcanic 

clasts in a metatuffaceous matrix (Lund, 1984). Recrystallized quartz, untwinned plagioclase, chlorite, 

tremolite, epidote, biotite, and muscovite constitute 

the mineralogically variable groundmass. Plagioclase, 

garnet, and chlorite porphyroblasts are common.

The stratigraphic section of the Seven Devils island-arc sequence reflects a complex history of 

volcanic-island-arc formation through late-stage quiescence (Lund and others, 1983a; Folio and Siever, 

1984; Lund, chap. 14, this volume; Folio, 1992, 1994). 

Chemistry of the Riggins Group was described by 

Hamilton (1963) as "similar in many respects to the 

Seven Devils volcanics"; in addition, he noted that 

"the distribution of minor elements [in the Seven 

Devils Group] is *** identical with that in the metavolcanic rocks of the Riggins Group." Furthermore, 

detailed mapping by Onasch (1977, 1987), Lund 

(1984 and chap. 14, this volume), and McCollough 

(1984) has shown that units of the Riggins Group are 

traceable northward along strike and across meta-
364 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

morphic isograds into rocks earlier mapped as part of 

the Seven Devils island-arc sequence by Hamilton 

(1963).

Hamilton (1963) named the Riggins Group as a 

separate stratigraphic unit from the Seven Devils 

Group on the basis of differences in metamorphic 

grade and the presence of a separating structure. 

However, the metamorphic grade of the Seven Devils 

Group clearly increases to upper greenschist or amphibolite facies near its eastern extent where it is in 

thrust contact with the Riggins Group of similar 

grade (Lund, 1984; Onasch, 1987). The structure that 

separates these two sequences was named the Rapid 

River thrust fault by Hamilton (1963), who later stated that it is the suture that marks the collision of the 

Seven Devils ensimatic island arc with the continent 

during Late Permian or Triassic time (Hamilton,

1976). In contrast, Lund and Snee (1988) have suggested that there is no direct evidence that the Seven 

Devils and Riggins Groups represent different units 

or terranes; instead, the Riggins Group is simply the 

more highly metamorphosed equivalent of the Seven 

Devils Group. The Rapid River thrust fault, then, is 

not a suture between two different terranes but is 

one of a group of structures of similar age within a 

single package of rocks. In fact, the Rapid River 

thrust fault and other associated structures commonly place higher grade metamorphic rocks over their 

lower grade metamorphic equivalents; this type of 

structural stacking is probably related to development of the Salmon River suture zone (Lund, 1984 

and chap. 14, this volume; Lund and Snee, 1988). 

Within the Snake River canyon, Vallier (1974,

1977) mapped and described three areas of basement 

rocks (Cougar Creek Complex of Vallier, 1968) underlying units of the Seven Devils Group (see figs. 10.3, 

10.4, and 10.5). From north to south, these areas are 

the confluence of the Imnaha and Snake Rivers, the 

Pittsburg Landing area, and the area of The Oxbow. 

The exposures of basement rocks were mapped in 

part as shear zones that include crushed and recrystallized breccia, schist, mylonite, gneissic mylonite, 

amphibolite, metadiorite, metagabbro, and metatrondhjemite (Vallier, 1974). The rocks are generally 

metamorphosed to amphibolite facies, but retrograde 

chlorite, especially after biotite, is common. In the 

Imnaha River and Pittsburg Landing areas, younger 

plutons and dikes of tonalite and quartz diorite intruded the metamorphic units; these relatively unmetamorphosed intrusive rocks commonly also 

contain chloritized biotite. Ages of metamorphosed 

rocks and unmetamorphosed plutonic rocks in the 

basement complexes, based on U-Pb isotopic ages on 

zircons, range from 266 to 145 Ma (Walker, 1986).

The relation between the basement complexes and 

the Seven Devils Group has been the subject of speculation but is becoming clarified (White and Vallier, in 

press). The basement rocks are generally amphibolite 

facies gneisses and highly deformed. In contrast, 

nearby rocks of the Seven Devils Group are low grade 

to unmetamorphosed and exhibit primary volcanic 

and sedimentary textures. Plutons are common in 

both the basement rocks and the Seven Devils Group. 

Some of the plutons were emplaced before metamorphism of the basement complex; others were emplaced after metamorphism of the basement rocks; 

still others were emplaced after metamorphism of the 

Seven Devils Group. Wherever in contact, the basement rocks and the Seven Devils Group are separated 

by faults. Generally these are high-angle structures, 

but at Pittsburg Landing, Vallier (1968, 1977; White 

and Vallier, in press) mapped a thrust fault separating lower-plate Middle Jurassic (Bajocian to Callovian) strata from the overriding metamorphic and 

plutonic rocks of the Permian Cougar Creek (basement) Complex and an unnamed member of the Wild 

Sheep Creek Formation, a unit of the Seven Devils 

Group and Middle Triassic (Ladinian; between about 

235 and 230 Ma) in age. This thrust fault, therefore, 

must be younger than Middle Jurassic. In contrast, 

other structures, such as the ductile shear zones developed in rocks at the confluence of the Imnaha and 

Snake Rivers (fig. 10.5) and the area of The Oxbow 

(fig. 10.4), are likely older. Therefore isotopic age constraints are needed for several important questions, 

such as, what is the age of metamorphism of the Cougar Creek (basement) Complex, what is the age of the 

various plutonic rocks, and when did the thrusts, 

high-angle faults, and shear zones form?

Postmetamorphic plutonic rocks intrude the Seven 

Devils and Riggins Groups or correlative units in numerous places in the region in addition to Snake 

River canyon. Two of the largest of these plutons are 

the Wallowa and Bald Mountain batholiths (Taubeneck, chap. 2, this volume), which intrude units correlative with the Seven Devils Group in 

northeastern Oregon (fig. 10.1). Previously published 

Rb-Sr and K-Ar dates for these intrusive bodies 

(Armstrong and others, 1977) are widely discordant 

but generally bracket emplacement within the Jurassic to the Cretaceous. In addition, stocks that were 

emplaced after metamorphism of the basement rocks 

are exposed east of Snake River canyon near the 

Rapid River thrust fault. Several small plutons along 

the Rapid River southwest of Riggins are crosscut by 

structures like the Rapid River thrust fault but appear to intrude mylonites and faults (Sarewitz, 1982) 

that are apparently older than the Rapid River
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 365

thrust fault. These plutons along the Rapid River exhibit metamorphic effects similar in grade to that 

displayed by the surrounding Seven Devils Group. 

This similarity in metamorphic grade indicates that 

the plutons were emplaced after regional metamorphism of the Cougar Creek (basement) Complex but 

prior to metamorphism of the Seven Devils Group, 

which has been attributed to activity on the Salmon 

River suture zone (Lund and Snee, 1988).

Within the highest grade and most deformed rocks 

of some parts of the Salmon River suture zone are 

large areas of strongly deformed plutons of tonalitic 

and trondhjemitic compositions. The most extensive 

exposures of these plutons are near McCall and near 

Orofino, Idaho. Although some of these plutons are 

about 90 Ma (see next paragraph), an increasing 

amount of isotopic data indicate that some are about 

115 Ma (Davidson, 1990; C.A. Manduca, California 

Institute of Technology, unpub. data, 1990; Snee and 

others, in press). The 115-Ma plutons are of islandarc chemical affinity (Davidson, 1990) and generally 

have initial 87Sr/86Sr ratios of less than 0.704 (Armstrong and others, 1977; Criss and Fleck, 1987; Snee 

and others, in press). Temporal and spatial relations 

indicate that these plutons were emplaced during the 

period of metamorphism that affected Riggins Group 

rocks (Onasch, 1987).

A younger group of hornblende-biotite tonalite and 

granodiorite plutons (about 95 to 90 Ma; Lund and 

Snee, 1988), some of which contain magmatic epidote, 

intruded along the Salmon River suture zone after the 

youngest metamorphism and deformation affecting 

the Riggins Group and Seven Devils island arc had 

taken place (fig. 10.2). These plutons cut rocks of both 

island-arc and North American continental affinities. 

Where primary epidote is present, petrologic evidence 

indicates depths of emplacement greater than 25 km 

(Zen and Hammarstrom, 1984); however, not all tonalites contain epidote and, as is argued below, this 

structurally complex zone may juxtapose various levels of the same pluton. These plutons are the firstphase intrusions of the Idaho batholith, and their emplacement occurred after the island-arc terrane(s) finally accreted to the North American continent (Lund, 

1984; Lund and Snee, 1988). The plutons and roof 

pendants within them and nearby metamorphic rocks 

of the Riggins Group were deformed together by one 

or more postemplacement events that were confined 

to a zone about 5 km wide along the suture zone. The 

fabrics that formed during this deformation are manifested by steep schistosity that parallels the suture 

zone and by steep southeast-plunging mineral lineations that dominate the fabric in plutonic rocks on the 

west side of the suture zone (Lund, 1984).

Away from the suture zone, hornblende-biotite tonalite and biotite granodiorite plutons intruded 

both terranes (fig. 10.2). Epidote is present as a primary magmatic mineral in some of the undeformed 

tonalites (Hammarstrom and Zen, 1988). The rocks 

are mostly homogeneous in texture, although a foliation is present in some places. The magmas intruded by magmatic stoping (Lund, 1984). These 

plutonic rocks may be undeformed equivalents of, or 

cogenetic with, tonalitic rocks that were emplaced 

into and deformed along the suture zone. They may 

also be laterally equivalent to similar plutons along 

the east side of the Idaho batholith near Thompson 

Creek, Hailey, and Stanley (fig. 10.1; see also 

fig. 10.21).

In the eastern, continental terrane, the tonalite, 

granodiorite, and metamorphic country rocks were 

intruded by biotite granite to granodiorite and muscovite-biotite granite plutons with peraluminous to 

near-peraluminous chemical compositions. Muscovite 

is primary in the muscovite-biotite granites and is a 

common deuteric or hydrothermal alteration phase in 

the biotite granite to granodiorite. These plutons 

have no directional fabric (Lund and others, 1986) 

and, like the tonalite, intruded passively by magmatic stoping. Country-rock xenoliths, which are prevalent in the upper parts of the plutons, were not 

rotated during stoping (Lund, 1984). Some workers 

(for example, Hyndman, 1981; WB. Hamilton, U.S. 

Geological Survey, oral commun., 1988) have argued 

that because the muscovite-biotite granite plutons 

contain magmatic muscovite, they must have been 

emplaced at depths greater than 17 km. Lund and 

Snee (1988) and Lund and others (1986) presented 

evidence, based on cooling calculations, that the 

granites must have been emplaced at depths less 

than 9 km (see below).

The ages of emplacement of plutons within the 

Idaho batholith have been difficult to define despite 

the numerous attempts at isotopic dating that have 

been made. The earliest numerical ages on the plutons were Pb-alpha determinations by Larsen and 

Schmidt (1958) that ranged between 90 and 135 Ma. 

This method has since been shown to provide unreliable results, and in the case of the Idaho batholith, 

the Pb-alpha dates are at least 20 percent too old. 

The vast majority of isotopic dates for Idaho batholith plutons have been determined by the conventional K-Ar method, but the complex subsolidus thermal 

and tectonic history of the batholith has caused inconsistent results which range from 117 to about 

50 Ma (for example, Armstrong, 1975; Armstrong and 

others, 1977; Criss and others, 1982). Many of these 

dates are geologically meaningless when they are
366 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

carefully evaluated. Because more accurate and precise results are now available, these earlier dates are 

not used in this paper for evaluating ages of intrusion of plutons of the Idaho batholith.

ANALYTICAL METHODS

40Ar/39Ar TECHNIQUES

The 4°Ar/39Ar dating technique is a variant of the 

conventional K-Ar method. To obtain a date by this 

technique, the sample of unknown age and a standard of known age are irradiated together in a nuclear 

reactor to produce 39Ar from 39K by fast-neutron 

bombardment. After irradiation, the 4°Ar/39Ar ratios 

of sample and standard are measured. The date of a 

sample can be calculated from its 4°Ar/39Ar ratio in 

comparison with the 4OAr/39Ar ratio of the standard. 

Only the isotopic composition of the argon of the 

sample and standard need be measured and this is 

done by gas-source mass spectrometry, potentially a 

very precise analytical technique. In contrast, for a 

conventional K-Ar date, both 40K and 4°Ar must be 

measured quantitatively To do this, argon (a gas) in 

one aliquot of sample is measured by isotope-dilution, gas-source mass spectrometry. Potassium (a solid) in a different aliquot of sample is determined by 

some other analytical method such as flame photometry, X-ray fluorescence, or isotope-dilution, solidsource mass spectrometry. Thus, one inherent problem of the conventional K-Ar technique is the necessity of measuring isotopic abundances for separate 

aliquots of the same sample. This poses the danger 

that, because of sample inhomogeneity, different potassium and (or) argon contents may exist in each 

aliquot. Two major advantages of the 4OAr/39Ar dating method are (1) only isotopic ratios of argon need 

be determined, and (2) all measurements are made 

on the same sample aliquot, thus avoiding the question of inhomogeneity. In addition, by the 4OAr/39Ar 

method, it is possible to obtain a series of dates from 

a single sample when argon is extracted by stepheating. The combination of these advantages potentially increases both the accuracy and precision of 

the 4°Ar/39Ar method over the conventional K-Ar 

technique. However, the 4OAr/39Ar technique will suffer if proper corrections are not made for interfering 

radiation-induced isotopes; these corrections are well 

known and routinely made.

Details of the 4°Ar/39Ar dating technique have been 

discussed by Dalrymple and others (1981). The procedures specific to our study are described in Snee and 

others (1985, 1987b, 1988). To place the data of this

paper in the proper perspective, however, a summary 

of the formulation of the age equation, a discussion of 

mineral argon-retention temperatures, and an evaluation of various types of age spectra are presented.

The 4OAr/39Ar date of a sample is calculated according to the age equation,

where tu is the calculated date of the "unknown" 

sample, X is the decay constant for decay of 40K to 

4°Ar and 40Ca, J is related to neutron flux during 

irradiation, and F is the ratio of 40ArR (radiogenic 

4°Ar) to 39ArK (potassium-derived 39Ar) of the sample. The decay constants used in this study are those 

recommended by Steiger and Jager (1977), that is, 

^=0.581x10-10/3^, XB_=4.962xlO-iO/yr, and X=X£+XB_ 

=5.543xlO-10/yr. The flux parameter, J, is calculated 

according to the relationship

where e is a constant, tm is the age of the primary flux 

"monitor" (that is, standard), and (40ArR/39ArK)m is 

the measured ratio of the standard. The standard for 

this experiment is hornblende MMhb-1 with percent 

K=1.555, 4OArR=1.624xlO-9 mol/g, and K-Ar age= 

519.4 Ma (Alexander and others, 1978; Dalrymple and 

others, 1981). F is calculated from the measured 

argon isotopic abundances by mass spectrometry. In 

order to determine the actual F, that is, 4°ArR/39ArK 

ratio, of a sample or standard, it is necessary to correct for the presence of atmospheric 4°Ar and irradiation-produced interfering isotopes such as 40Ar (from 

4<>K), 39Ar (from 42Qa), and 36Ar (from 4<>Ca and 35C1). 

These corrections are well defined and were all made 

on the data used in this study. Because the mathematical formulation for F is complex, it is not presented 

here.

40Ar/39Ar age-spectrum dating of hornblende, muscovite, biotite, and microcline, which formed during 

metamorphism or crystallization from a magma, is 

currently the best method for determining thermal 

history and age of complex metamorphic and plutonic 

terranes. Ideally, a mineral date determined by this 

method marks the time when that mineral became 

closed to the diffusion of argon. Closure of a particular mineral to diffusion is controlled chiefly by temperature, to a smaller extent by cooling rate (Dodson, 

1973), and possibly by chemical compositional variation, structural-state variation, or grain-size variation 

within a particular kind of mineral. Each mineral 

that is appropriate for argon dating has a characteristic closure temperature (or argon-retention temperature) that is known with a precision of about ±20°C. 

The closure temperature is higher for minerals that
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 367

cooled rapidly and lower for minerals that cooled 

slowly. Commonly accepted closure temperatures that 

span a range from rapid cooling (l,000°C/m.y.) to slow 

cooling (5°C/m.y.) are 580-480°C for hornblende (Harrison, 1981), 325-270°C for muscovite (Snee and 

others, 1988), 300-260°C for biotite (Harrison and 

McDougall, 1980; Snee, 1982), and 160-100°C for microcline (Harrison and McDougall, 1982). For simplicity, intermediate closure temperatures, that is, 

closure temperatures for intermediate cooling rates of 

500-100°C/m.y, are assumed for most of this study: 

530°C, 300°C, 280°C, and 130°C for hornblende, muscovite, biotite, and microcline, respectively.

The 4OAr/39Ar method was first used in "total-fusion" experiments in which an irradiated sample was 

completely melted and all isotopes of argon measured 

in a single analysis to calculate a date for the sample. This total-fusion date is roughly analogous to a 

conventional K-Ar date for the sample except that no 

isotopic concentration measurements are required. 

Very soon after the first uses of the 40Ar/39Ar method, it was realized that a sample could be progressively degassed in temperature increments (Merrihue 

and Turner, 1966). A date can be calculated for each 

increment of gas released, and the dates can be plotted against percent of released argon to form an age 

spectrum. The character of the spectrum can be evaluated within a theoretical framework to interpret the 

apparent distribution of potassium and argon within 

the sample.

Turner (1968) showed that the dates for the temperature increments of some meteorites were identical within analytical precision and that when plotted 

on an age-spectrum diagram, which shows date of 

each temperature increment as a function of percent 

39Ar released during the experiment, the spectra 

formed "plateaus." In contrast, some age spectra exhibit a distinct increase in dates for the temperature 

increments from low-temperature to high-temperature extraction-steps in the experiment. Turner 

showed theoretically that an age spectrum will exhibit a step-up in dates if argon was lost from the sample 

in the geologic environment by thermally activated 

volume diffusion. The age spectrum will exhibit a plateau if the sample had never been disturbed after formation or if the sample had been completely reset by 

a younger event. Depending on the amount of thermal disturbance, that is, the percentage of argon lost, 

the date of the younger, low-temperature fractions 

will be equal to, or older than, the age of the thermal 

disturbance that affected the sample, whereas the 

date of the older, high-temperature fractions will be 

equal to, or younger than, the original age of closure. 

This step-up in dates exhibited by an age spectrum

and apparently resulting from argon loss due to geologic activity has been experimentally reproduced by 

Harrison (1981) from hornblende and has been observed by numerous investigators in data from hornblende, muscovite, and K-feldspar (see below and, for 

example, Snee and others, 1988, and Harrison, 1982).

Recently, the reliability of hornblende and other 

minerals, including muscovite and K-feldspar, as recorders of geologically induced, thermally activated 

argon loss in an age spectrum produced during experimental degassing in the laboratory has been 

challenged by Lee and others (1991). Their contention is that hydrous minerals in particular, but also 

K-feldspar, undergo changes including dehydration, 

melting, and phase conversions during heating 

under vacuum that prevent release of argon by thermally-activated volume diffusion from the mineral. 

Thus, according to Lee and others (1991), experimental reproduction of the geologic phenomenon of argon 

loss is impossible.

Although the mechanism responsible for the development of apparent argon-loss age spectra from some 

hornblendes, K-feldspars, and muscovites may not be 

completely understood, there is no question that this 

pattern is exhibited by many, if not most, of these 

minerals after they undergo partial argon loss due to 

reheating by later thermal activity. Besides thermally activated volume diffusion, several other possible 

mechanisms can explain apparent argon-loss spectra. 

For example, two or more phases with different closure ages may coexist in one rock. If these phases of 

different age are similar enough in physical character to coexist in a mineral separate (for example, if 

one is phengite, a white mica formed under highpressure metamorphic conditions, and the other is 

muscovite, a white mica formed under high-temperature conditions), and if each phase degasses under 

vacuum over different temperature ranges, then the 

resultant age spectrum will represent a mixture of 

the age spectra for both samples and will resemble 

an argon-loss spectrum (Roeske and others, 1995). 

Similarly, a single-phase mineral separate that has 

several grain sizes or structural types that degas at 

different temperatures under vacuum can produce 

the same result (Cosca and others, 1992). In both of 

these examples, the apparent argon-loss spectrum 

can be interpreted as if it had resulted from thermally activated volume diffusion. That is, the apparent 

age of the younger, lower temperature fractions is approximately the age of, or older than, the thermal 

disturbance, and the apparent age of the older, higher temperature fractions is approximately the age 

of, or younger than, the original closure. Therefore, 

apparent argon-loss spectra can be potentially very
368 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

useful for understanding the thermal histories of 

complex geologic environments. However, caution 

must be taken when interpreting apparent argon-loss 

spectra because other, nongeologic factors, such as 

extraction-system blanks and sample impurity, can 

affect the age spectra.

Besides argon loss, excess 40Ar can be detected in a 

sample. Excess 40Ar is incorporated into rocks and 

minerals by processes other than in situ decay of 

40K. The mechanism for the incorporation of excess 

40Ar into a rock or mineral is unknown, but several 

possible explanations exist. If a rock or mineral crystallizes in an environment that contains argon with 

an 40Ar/36Ar ratio greater than 295.5 (present-day 

ratio for atmospheric argon), the extra 40Ar is "excess." Similarly, hydrothermal fluids are known to 

carry argon that commonly has an 40Ar/36Ar ratio 

composition greater than 295.5, perhaps due to their 

interaction with rocks containing radiogenic argon. If 

these fluids alter or become incorporated in a rock or 

mineral, excess 40Ar can be added to that rock or 

mineral. Several studies (for example, Lanphere and 

Dalrymple, 1971, 1976; Kaneoka, 1974) have documented that large quantities of excess 40 Ar in a sample will produce a saddle-shaped 40Ar/39Ar age 

spectrum with anomalously old dates for the lowtemperature and high-temperature extraction steps. 

Small quantities of excess 40Ar commonly only affect 

the low-temperature steps and form an L-shaped 

spectrum.

Many studies have shown that meaningful ages of 

samples could be determined by the age-spectrum 

technique even though some loss or gain of 40Ar had 

occurred during the geologic history of the sample. 

However, loss or gain of 40Ar by a sample will result 

in an erroneous conventional K-Ar date, and because 

loss or gain of 40Ar by a sample is primarily thermally controlled, samples from thermally complex areas 

such as metamorphic and plutonic complexes should 

by analyzed by the 40Ar/39Ar age-spectrum technique. Similarly, conventional K-Ar data from these 

geologic settings should be viewed with caution.

Ideally a sample dated by the 40Ar/39Ar age-spectrum 

technique will yield concordant dates for all temperature steps. If a sample's age spectrum is 100 percent 

concordant, that is, if all temperature steps yield identical dates within 2 standard deviations (2a) of the 

weighted mean date for all temperature steps, then 

clearly the best date for the sample is a weighted mean 

of the dates of all temperature steps. The geologic 

significance of the resultant date must then be evaluated using geologic or other independent constraints. 

(In this paper, "date," "apparent age," or "numerical 

age" is used to refer to the analytical number determined by solution of the age equation using isotopic 

data produced for a sample in the laboratory. In contrast, "age" or a specific type of "age," such as "emplacement age," "cooling age," or "age of metamorphism," 

refers to a date that has been constrained by geologic 

data and is interpreted to have geologic meaning.)

In reality most dated samples display some discordancy in their age spectra, either because of a physical or chemical disturbance, because of 40Ar not 

derived from in situ decay of 40K, or because of loss 

of some 40Ar after the original closure of the samples 

to 40Ar diffusion. Even if discordancy exists in an age 

spectrum, the interpreted date may still hold geologic 

significance. Some terms for interpreted dates that 

are used in this paper include "total-gas date," "plateau date," and "preferred date." The "total-gas date" 

of an age spectrum results from weight-averaging 

dates of all temperature steps for a sample; the totalgas date is comparable to a conventional K-Ar date 

but is more precise and more accurate. The "plateau 

date" is that part of an age-spectrum diagram composed of contiguous gas fractions that together represent more than 50 percent of the total 39Ar released 

from the sample and for which no difference in date 

can be detected between any two temperature fractions at the 95-percent confidence level (Fleck and 

others, 1977). The term "plateau date" has been used 

by others (for example, Lanphere and Dalrymple, 

1971; Dalrymple and Lanphere, 1974) to refer to a 

"high-temperature segment" of an age spectrum over 

which the dates of the plateau-defining steps are concordant at the 95-percent confidence level. In some 

cases, by this usage, a single temperature step may 

define a plateau. In this paper, we use the definition 

for plateau of Fleck and others (1977). Finally, some 

age spectra show near-concordancy and yield weightaverage dates of the near-concordant parts of the 

spectra that seem to have geologic meaning. We refer 

to these dates as "preferred dates," which either consist of an apparent plateau with less than 50 percent 

of the released 39Ar or consist of fractions of gas 

whose dates overlap within 3 standard deviations 

(3a) of the weighted mean.

U-Pb TECHNIQUES

Dissolution and chemical preparation of zircon and 

sphene samples were accomplished using a procedure 

slightly modified from Krogh (1973). Samples of approximately 10 to 20 mg were dissolved in hydrofluoric 

acid in Teflon bombs at 200°C, followed by Pb purification on a bromide-form anion-exchange column. Lead 

content was determined by isotope dilution on an ali-
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 369

quot of the dissolved sample. A combined 235U and 

230Th spike was added to the sample before dissolution, 

and eluate off the bromide-form column was passed 

through a nitrate-form anion-exchange column to isolate uranium and thorium. Isotopic ratios were determined using a 15-cm (U, Th) or 30-cm (Pb) automated 

solid-source mass spectrometer.

Uranium, thorium, and lead contents are considered accurate to ±1 percent (2a). Lead concentration 

ratios have uncertainties of ±0.1 percent (2o), except 

206pb/204pb ratios greater than 1,000, which probably 

have uncertainties of 1-3 percent. A common lead 

correction was made by first subtracting lead equal 

in amount and composition to the analytical blank, 

which over the course of this study was about 0.7 ng. 

The remaining 204Pb was assumed to be 100-Mamodel lead with the composition 204pb:206pb:207pb: 

208pb=l:18.55:15.62:38.45. Model lead compositions 

are from Stacey and Kramers (1975); all constants 

are from Steiger and Jager (1977). Uncertainty of the 

lower-intercept date is the 95-percent confidence 

limit calculation according to Ludwig (1980, 1982).

SAMPLE DISTRIBUTION AND ISOTOPIC DATA

To evaluate the age of metamorphism, deformation, plutonism, cooling, and uplift of rocks in the 

vicinity of the suture zone in western Idaho, this 

paper synthesizes the argon data for 48 samples 

from which we analyzed 34 hornblendes, 12 biotites, 

11 muscovites, and 5 microclines that were collected in an area extending from the Snake River canyon in Oregon and Idaho eastward to Thompson 

Creek near Stanley, Idaho. U-Pb dates on one sample of plutonic rock from the west side of the Idaho 

batholith are also included. Preliminary argon data 

for five additional samples of the Bald Mountain 

batholith are discussed. The data for this paper are 

synthesized from studies of several areas which include (1) the Bald Mountain batholith in the Blue 

Mountains, Oregon (L.W. Snee and W.H. Taubeneck, 

unpub. data, 1987), (2) the basement metamorphic 

and plutonic rocks of the Cougar Creek Complex in 

the Wallowa terrane in the Snake River canyon area 

between the confluence of the Imnaha and Snake 

Rivers and The Oxbow, Oregon (Baker, 1980), (3) 

metamorphic rocks of the Riggins Group and Seven 

Devils island arc, and deformed and undeformed 

plutons and dikes from the Slate Creek-RigginsRound Valley area, Idaho (Lund and Snee, 1988), (4) 

plutonic rocks of the Idaho batholith and metasedimentary roof pendants in the Buffalo Hump area, 

Idaho (Lund and others, 1986), and (5) plutonic

rocks of the Idaho batholith near Stanley, Warm 

Lake, and Warm Springs Forest Service Station, 

Idaho, and alteration associated with the Thompson 

Creek porphyry molybdenum deposit (L.W. Snee, unpub. data, 1987). Distribution of these samples is 

shown on figures 10.3 through 10.6, 10.21, and 

10.23, and locations are listed in table 10.1. Argon 

data are summarized in table 10.2, and detailed 

data are listed in table 10.3. U-Pb data are listed in 

table 10.4, and a Concordia diagram of the U-Pb 

data is plotted on figure 10.20.

INTERPRETATION

Because of the diversity of these data, the following 

discussion is divided into four sections: (1) basement 

rocks of the Snake River canyon area and associated 

plutons of island-arc affinity, (2) metamorphic rocks 

of the Riggins Group and Seven Devils island arc, (3) 

deformed plutons of the Salmon River suture zone, 

and (4) undeformed plutons elsewhere in the Idaho 

batholith.

BASEMENT ROCKS OF THE SNAKE RIVER CANYON AREA AND 

ASSOCIATED PLUTONS OF ISLAND-ARC AFFINITY

In an attempt to determine age of the metamorphism, shearing, and later plutonism that affected 

the basement rocks in the Snake River canyon area, 

we used the 40Ar/39Ar age-spectrum technique to analyze samples from the three areas of exposure: Pittsburg Landing, The Oxbow, and the confluence of the 

Imnaha and Snake Rivers (fig. 10.1). Balcer (1980) 

analyzed five hornblende and two biotite samples 

from amphibolite and gneiss near Pittsburg Landing, 

and two hornblende samples from two sheared amphibolites formed from metamorphism of mafic dikes, 

one intrusive into the basement rocks and the other 

intrusive into the Windy Ridge Formation near The 

Oxbow. Balcer also analyzed four hornblende samples 

from unmetamorphosed plutons that intruded the 

basement rocks near the confluence of the Imnaha 

and Snake Rivers, and Snee and others (1987b) dated 

hornblende from unmetamorphosed tonalitic dikes 

that intruded the basement complex about 8 km east 

of Pittsburg Landing. In addition, a hornblende sample from a pluton that intruded the Seven Devils 

Group along the Rapid River and later underwent 

greenschist-facies metamorphism along with the 

Seven Devils Group was also dated. The sample locations for each area are shown on figures 10.3 through 

10.6, and a summary of the dates is given in table
370 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 10.1. Sample locations

[Mineral abbreviations: bi, biotite; hb, hornblende; mi, microcline; mu, muscovite. All quadrangles are 7.5-minute quadrangles, except Burgdorf, Copperfield,

and Warm Lake, which are 15-minute quadrangles]

Sample

R3 

R7 

R8a 

RIO 

Rll 

R12 

R14 

R16 

R17 

R18 

R20 

R21 

R22 

R23 

R24 

R25 

R26 

R27 

R28 

R29 

R30 

R34 

R35 

R36

Mineral

mu 

bi, hb 

bi, mi, mu 

mu 

hb 

bi, hb 

hb 

bi, hb 

hb 

hb 

hb 

bi, hb, mi 

hb 

hb 

hb, mi, mu 

bi 

bi, hb 

hb 

hb 

hb 

hb 

hb 

mu 

bi

Quadrangle

Riggins Hot Springs 

Riggins 

Riggins 

Dairy Mountain ~ 

McKinzie Creek 

McKinzie Creek 

Bally Mountain -

Pollock __ _ __

Riggins 

Riggins Hot Springs ~

Location

45°23'26" N., 116°25'14" W., SE1/4 sec. 26, T.24N., 

R.1E., Idaho County, Idaho.

45°24'30" N., 116°20'13" W., NE1/4 sec. 21, T.24N., 

R.1E., Idaho County, Idaho.

45°27'30" N., 115°54T 29" W., not surveyed, Idaho County, 

Idaho.

45°25'23" N., 116°01'55" W, NE1/4 sec. 13, T.24N., R.3E., 

Idaho County, Idaho.

45°24'35" N., 116°07'34" W., NE1/4 sec. 19, T.24N., 

R.3E., Idaho County, Idaho.

45°25'43" N, 116°08'34" W, SW1/4 sec. 7, T.24N., R.3E., 

Idaho County, Idaho.

45°24'02" N., 116°12'51" W., SE1/4 sec. 21, T.24N., 

R.2E., Idaho County, Idaho

45°24'46" N., 116°15'18" W., SEl/4 sec. 18, T.24N., 

R.2E., Idaho County, Idaho.

45°25'15" N., 116°16'48" W., NW1/4 sec. 13, T.24N., 

R.1E., Idaho County, Idaho.

45°24'38" N., 116°18'24" W., NW1/4 sec. 23, T.24N., 

R.1E., Idaho County, Idaho.

45°39'25" N., 116°23'54" W., SEl/4 sec. 24, T.24N., 

R.1W., Idaho County, Idaho.

45°37'58" N., 116°00'45" W., not surveyed, Idaho County, 

Idaho.

45°37'22" N., 116°03'43" W., NE1/4 sec. 3, T.26N., R.3E., 

Idaho County, Idaho.

45°38'14" N., 116°06'14" W., NE1/4 sec. 32, T.27N., 

R.3E., Idaho County, Idaho.

45°38'25" N., 116°07'18" W., NE1/4 sec. 31, T.27N., 

R.3E., Idaho County, Idaho.

45°38T 09" N., 116°08'31" W., NE1/4 sec. 36, T.27N., 

R.2E., Idaho County, Idaho.

45°38'08" N., 116°11T 36" W., NW1/4 sec. 34, T.27N., 

R.2E., Idaho County, Idaho.

45°38'07" N., 116°14'12" W., NW1/4 sec. 32, T.27N., 

R.2E., Idaho County, Idaho.

45°07'28" N., 116°17'40" W., SE1/4 sec. 26, T.21N., 

R.1E., Adams County, Idaho.

45°09'43" N., 116°17'54" W., NE1/4 sec. 14, T.21N., R.1E. 

Idaho County, Idaho.

45°15'10" N., 116°20'5" W., NE1/4 sec. 16, T.22N., R.1E., 

Adams County, Idaho.

45°27'41" N., 116°18'26" W., SW1/4 sec. 35, T.25N., 

R.1E., Idaho County, Idaho.

45°29'17" N., 116°18'05" W., NW1/4 sec. 23, T.25N., 

R.1E., Idaho County, Idaho.

45°24'35" N., 116°07'34" W., NE1/4 sec. 19, T.24N., 

R.3E., Idaho County, Idaho.
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 371

TABLE 10.1. Sample locations Continued

Sample Mineral Quadrangle Location

7-30-5 hb

Hem-23 hb

Hdh-24 hb

Hdh-25 hb

Hdh-27 hb

Hkc-29 bi

Hkc-32 hb

Hkc-34 hb

Hkc-35 hb

Hkc-36 bi, hb

Hkc-37 hb

Hcop-39 hb

Hcop-40 hb

D81-1 hb

1KE052 mi, mu

2KE080 hb, bi

2KE086 hb

2KE087 mu

MP201 hb

TC-S-41-16 mu

TC-S-124-23 mu

D3113M mu

B12178201 mu

B12198202 bi, mi

Indian Mountain ----

Cactus Mountain - 

Deadhouse Ridge - 

Deadhouse Ridge -- -- 

Deadhouse Ridge - 

Kirkwood Creek - -- 

Kirkwood Creek - - 

Kirkwood Creek 

Kirkwood Creek 

Kirkwood Creek 

Kirkwood Creek 

Copperfield 

Copperfield 

Heavens Gate 

Buffalo Hump 

Buffalo Hump 

Buffalo Hump 

Buffalo Hump 

East Basin Creek 

Thompson Creek 

Thompson Creek 

Warm Lake 

Teapot Mountain - 

Eightmile Mountain 

45°08'05" N., 116°16'56" W, NW1/4 sec. 25, T.21N., 

R.1E., Adams County, Idaho.

45°49'26" N., 116°41'46" W., NE1/4 sec. 28, T.29N., 

R.3W., Idaho County, Idaho.

45°49'00" N., 116°45'45" W., not surveyed, Wallows 

County, Oregon.

45°49'25" N., 116°46'31" W., not surveyed, Wallows 

County, Oregon.

45°49'00" N., 116°45'48" W., not surveyed, Wallows 

County, Oregon.

45°36'48" N., 116°27'35" W., Sl/2 sec. 4, T.26N., R.1W., 

Idaho County, Idaho.

45°36'46" N., 116°27"35" W., Sl/2 sec. 4, T.26N., R.1W., 

Idaho County, Idaho.

45°36'25" N., 116°27'35" W., Nl/2 sec. 9, T.26N., R.1W., 

Idaho County, Idaho.

45°36'25" N., 116°27'35" W., Nl/2 sec. 9, T.26N., R.1W., 

Idaho County, Idaho.

45°36'36" N., 116°27'35" W., Nl/2 sec. 9, T.26N., R.1W., 

Idaho County, Idaho.

45°36'36" N., 116°27'35" W., Nl/2 sec. 9, T.26N., R.1W., 

Idaho County, Idaho.

44°58'45" N., 116°51'30" W., SE1/4 sec. 17, T.19N., 

R.4W., Adams County, Idaho.

44°56'55" N., 116°50'45" W., NE1/4 sec. 21, T.5S., 

R.48E., Baker County, Oregon.

45°19' N., 116°25' W., NE1/4 sec. 23.T.23N., R.1W., 

Idaho County, Idaho.

5°32'24" N., 115°40'10" W., not surveyed, Idaho County, 

Idaho.

45°31'50" N., 115°43'11" W., not surveyed, Idaho County, 

Idaho.

45°33'10" N., 115°41'38" W., not surveyed, Idaho County, 

Idaho.

45°35'14" N., 115°40'53" W., not surveyed, Idaho County, 

Idaho.

44°20' N., 114°50' W., sec. 24, T.11N., R.14E., Custer 

County, Idaho.

44°19'00" N., 114°33'45" W., SE1/4 sec. 34, T.12N., 

R.16E., drill hole S41, 495.4-ft depth, Custer County, 

Idaho.

44°19'00" N., 114°33'45" W., SE1/4 sec.34, T.12N., 

R.16E., drill hole S124, 1606.5-ft depth, Custer 

County, Idaho.

44°38'44" N., 115°36'33" W., SW1/4, sec. 9, T.15N., 

R.7E., Valley County, Idaho.

44°57'25" N., 115°38'00" W., NE1/4, sec. 29, T.19N., 

R.7W., Valley County, Idaho.

44°08'09" N., 115°18'53" W., sec. 7, T.9N., R.10E., Boise 

County, Idaho.
372 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 10.2. Summary of 40Ar/39Ar age-spectrum data

[Mineral abbreviations: bi, biotite; hb, hornblende; mi, microcline; mu, muscovite. Metamorphic-zone abbreviations: An, andesine; Bi, biotite; Ch, 

chlorite; Gt, garnet; (Hi), upper amphibolite facies or higher; 01, oligoclase; Si, sillimanite; , unmetamorphosed. Other abbreviations: Ar 

loss, apparent argon loss; do, ditto; Tmp, preferred age of sample; Tp, plateau age. All errors are 1 standard deviation from mean. Decay 

constants and isotopic abundances are those recommended by Steiger and Jager (1977)]

Sample Mineral Metamorphic 

zone

Unit or rock type Age, and characteristics of 

40Ar/39Ar spectrum

Basement rocks and associated plutons

Hdh-24

Hdh-25 

Hdh-27 

Hem-23

Hkc-29 

Hkc-32 

Hkc-34 

Hkc-35 

Hkc-36

 do 

Hkc-37

Hcop-39 

Hcop-40

R20 

D81-1

R7

 do - 

R30

hb

hb 

hb 

hb

bi 

hb 

hb 

hb 

hb

bi 

hb

hb 

hb

hb 

hb

hb

bi 

hb

Tonalite - T =215.6±3.6 Ma;

Ar loss from 220 Ma.

.....do T =213.0±2.0 Ma.

Tonalite inclusion in mafic T =223.3±5.9 Ma. 

dike. P

(Hi) Deformed plagiogranite T =233.5±2.5 Ma;

disturbed.

(Hi) Amphibolite T =229.3±2.6 Ma;

Ar loss from 232 Ma.

(Hi) - do T =225.8±6.7 Ma;

Ar loss? from 229 Ma.

(Hi) do Tp=239.2±7.6 Ma;

Ar loss from 243 Ma.

(Hi) do T =238.6±3.4 Ma;

Ar loss from 244 Ma.

(Hi) do Disturbed.

(Hi) - do T =235.0±3.5 Ma;

disturbed.

(Hi) ~ do T =227.8±5.3 Ma.

(Hi) .....do T =219.2±7.6 Ma;

Ar loss from 224 Ma.

Tonalite T =227.1±1.2 Ma.

Ch .....do T =145.1+1.5 Ma.

Rocks near the suture zone

Gt Riggins Group Tmp=117' 0±0- 6 Ma;

minor excess argon.

..do - -do T =116.1±0.6Ma.

(Hi) Riggins(?) Group T =118.1±0.6 Ma;

Ar loss to 114 Ma.

R3 mu Ch Seven Devils island arc Ar loss: 105-100 Ma.
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 373

TABLE 10.2. Summary of 40Ar/ 39Ar age-spectrum data Continued

Sample Mineral Metamorphic 

zone

Unit or rock type Age, and characteristics of 

40Ar/39Ar spectrum

Rocks near the suture zone Continued

R28 

R29 

R18 

R17

R34 

R27

R26

 do -

R35

R16

 do -

R14 

R25

7-30-5

R24

 do

 do 

R23 

R22

R12

 do

hb 

hb 

hb 

hb

hb 

hb

hb 

bi

hb 

bi

hb 

bi

hb

hb 

mu

mi 

hb 

hb

hb 

bi

(Hi) Riggins(?) Group T =107.7+0.5 Ma ;

Ar loss: 108-100 Ma.

(Hi) Riggins(?) Group T =108.1±0.5 Ma;

Ar loss to 107 Ma.

Ol Riggins Group T =106.5±1.4 Ma;

disturbed.

Ol do T =106.8±0.5 Ma;

Ar loss: 107-105 Ma.

Gt do T =109.1±0.5Ma.

Ol do T =108.9±0.6 Ma;

minor excess Ar.

An ----- do Ar loss: 109-94 Ma.

 do - do T =89.9±0.6 Ma.

Bi do T =98.3±0.2 Ma;

Arloss: 101-97 Ma.

An ----- do Ar loss: 100.5-93 Ma.

.. do - - do T =87.9±0.4 Ma.

Si . do Arloss: 101-89 Ma.

An Seven Devils island arc T =84.1+0.4 Ma.

Tonalite T =112.0±0.7 Ma.

Suture-zone rocks

(Hi) Riggins(?) Group Ar loss: 92.5-88 Ma.

.. do - Deformed pegmatite in R24 T =83.4±0.4 Ma. 

gneiss.

.. do do - Ar loss: 79-69.5 Ma.

(Hi) Riggins(?) Group inclusion Ar loss: 99.5-93 Ma.

(Hi) do T =89.3±0.6 Ma.

Riggins(?) Group Ar loss: 93.5-88 Ma. 

 do T =82.6±0.4 Ma.

10.2 under the heading "Basement Rocks and Associated Plutons." Samples from plutons of the Bald 

Mountain batholith (L.W. Snee and W.H. Taubeneck, 

unpub. data, 1987) have also been analyzed.

The 40Ar/39Ar age spectra of the five hornblende 

samples from metamorphic rocks near Pittsburg 

Landing (Hkc-32, Hkc-34, Hkc-35, Hkc-36, and Hkc37; fig. 10.3) show variable amounts of discordance.
374 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 10.2. Summary of 40Ar/ 39Ar age-spectrum data Continued

Sample Mineral Metamorphic 

zone

Unit or rock type Age, and characteristics of 

40Ar/39Ar spectrum

Suture-zone rocks Continued

Rll hb Deformed tonalite - T =84.9+0.3 Ma.

R36 bi Deformed granodiorite-- T =80.5±0.4 Ma.

Undeformed plutons and associated rocks

R21 hb Tonalite T =83.9±0.4 Ma.

...do- bi .....do T =81.1±0.4Ma.

 do mi do - Ar loss: 78-75 Ma.

RIO mu Pegmatite in tonalite- T =76.7±0.4 Ma.

R8a mu Muscovite-biotite granite T =75.3±0.4 Ma.

... do - bi do T =74.7±0.4 Ma.

 do mi do Ar loss: 69-47 Ma.

2KE080 hb (Hi) Country rock (amphibolite)-- T =81.2+0.4 Ma.

... do - bi - do - - do -- T =78.2±0.4 Ma.

2KE086 hb (Hi) - do T =80.5±0.6 Ma.

1KE052 mu Muscovite-biotite granite- T =73.8±0.4 Ma.

 do mi do Ar loss: 67-55 Ma.

2KE087 mu Quartz vein T =71.0±0.4Ma.

MP201 hb Hornblende-biotite T =86.3±0.4 Ma.

	granodiorite.

TC-S-41-16 mu Molybdenum porphyry - T =87.4+0.4 Ma.

TC-S-124-23 mu - do T =87.6±0.4 Ma.

D3113M mu Muscovite-biotite granite T =72.4±0.4 Ma.

B12178201 mu .....do T =70.1+0.4 Ma.

B12198202 bi Leucogranite Disturbed.

 do mi do Ar loss: 66-55 Ma.

Even though the analytical precision for these analyses is relatively large (ranging from 3 to 8 m.y) compared to that of our other data, the data are precise 

enough to show that each of the samples has lost some 

argon over geologic time. Because the hornblendes

define the metamorphic foliation of the amphibolites, 

their plateau dates should be minimum estimates for 

the age of metamorphism. However, the superposed 

argon loss probably occurred during a thermal event, 

such as heat flow associated with later plutonism,
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 375

TABLE 10.3. 40Ar/ 39Ar age-spectrum data for samples from west-central Idaho and eastern Oregon

[All errors are 1 standard deviation from mean. Italic entries indicate data used to calculate plateau dates. J, flux parameter (see text); mol, mole; pet, percent; wt., weight; , not

determined]

Temperature

( °C)

40Ar/ 39Ar 37Ar/ 39Ar 36Ar/ 39Ar

39Ar

(pet of

total)

pet 40ArR

39Ar

(mol

xlO'13 )

Apparent

K/Ca

(mol/mol)

Apparent

age and error

(Ma)

R3, MUSCOVITE 

Total-gas date: 102.0 Ma; no plateau; J=0.007387; wt., 0.0226 g

750

875

925

975

1,000

1,050

Fuse

9.403

8.512

8.415

8.569

10.75

10.61

9.82

0.5345

.0266

.0070

.0079

.2273

.8400

.3253

0.0110

.0027

.0015

.0017

.0089

.0109

.0064

3.9

13.3

38.1

21.8

4.4

3.6

15.2

65.7

90.7

94.7

94.2

75.7

70.3

81.1

0.129

.449

1.29

.737

.148

.121

.514

0.97

19.5

74.8

66.0

22.9

.62

1.60

80.6+0.5

100.1+0.7

103.2+0.5

104.5+0.5

105.3+0.6

96.7+0.6

103.1+0.5

R7, HORNBLENDE 

Total-gas date: 133.7 Ma; no plateau; J=0.007083; wt.,0.7261 g

850

950

1,000

1,050

1,100

1,125

1,150

1,175

1,200

1,250

Fuse

112.45

51.49

29.67

24.76

19.85

15.69

15.43

16.20

12.11

10.95

15.98

2.

7.

9.

10.

12.

13.

14.

14.

14.

14.

14.

826

399

483

85

84

80

04

25

45

43

29

0.1420

.0902

.0639

.0443

.0301

.0217

.0225

.0255

.0133

.0088

.0261

1.5

1.0

1.1

2.7

5.2

8.4

9.0

7.2

16.0

39.5

8.5

63

49

38

50

60

66

64

60

76

86

58

.0

.4

.8

.6

.2

.0

.0

.3

.8

.5

.8

0.067

.045

.049

.122

.235

.375

.402

.320

.714

1.76

.380

0.18

.07

.06

.05

.04

.04

.04

.04

.04

.04

.04

733

298

141

153

146

127

121

120

115

117

116

.7+4.1

.6+2.7

.3+0.9

.2+0.8

.5+0.7

.6+0.7

.8+0.6

.5+0.6

.0+0.6

.0+0.6

.0+0.6

R7, BIOTITE 

Total-gas date: 115.7 Ma; plateau date: 116.1+0.6 Ma; J=0.007272; wt., 0.0712 g

850

1,000

1,050

1,100

1,200

Fuse

10.03

10.01

10.08

10.43

12.19

53.29

0.0127

.0118

.0217

.1098

.0792

.2905

0.0038

.0027

.0032

.0045

.0104

.1510

11.6

30.5

21.8

20.7

13.5

1.8

88.9

91.9

90.6

87.2

74.8

16.3

1.35

3.55

2.54

2.41

6.56

.215

40.9

44.2

24.0

4.74

6.56

1.79

113.4+0.6

116.9+0.6

116.0+0.6

115.5+0.6

115.9+0.6

110.5+0.9

R8a, MUSCOVITE 

Total-gas date: 75.1 Ma; plateau date: 75.3+0.4 Ma; J=0.007296; wt. 0.0642 g

750

875

925

975

1,015

1,050

1,200

Fuse

32

10

8

6

6

6

7

8

.87

.45

.386

.622

.191

.471

.285

.363

0.9163

.0143

.0086

.0003

.0006

.0007

.0004

.0005

0.094

.0167

.0089

.0026

.0011

.0022

.0050

.0086

40.3

.8

2.0

13.6

32.1

16.1

16.8

18.4

15.3

52.8

68.6

88.5

94.5

90.1

79.5

69.5

0.035

.111

.256

1.78

4.20

2.11

2.20

2.40

0.

36.

60.

1,750.

845.

73.

116.

95.

57

4

7

0

0

0

0

1

65

71

74

75

75

75

74

74

.0+2.2

.1+0.4

.1+0.4

.5+0.4

.4+0.4

.2+0.4

.7+0.4

.9+0.4

younger than the metamorphism. Therefore, the minimum cooling age of the metamorphic event is best 

represented by the highest temperature, oldest step of 

the oldest spectrum. The age spectra for samples Hkc35 and Hkc-36 show relatively little disturbance and 

step up to 242.4+4.2 and 243.8+3.7 Ma, respectively.

Therefore, we interpret the cooling age of the metamorphism that affected the basement rocks of the 

Pittsburg Landing area to be about 244 Ma. Because 

this is a cooling age that represents the time when the 

rocks cooled below 530°C (hornblende argon-closure 

temperature) and because these rocks were affected
376 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 10.3. 40Arl39Ar age-spectrum data for samples from west-central Idaho and eastern Oregon Continued

Temperature 40 Ar/ 39Ar 37Ar/ 39Ar 36Ar/ 39Ar 

<°C)

650 

850

1,000

1,050

1,125

1,200

Fuse

Total -gas

21.24 

7.164

6.174

7.083

6.613

7.156

13.41

date:

1.

.

.

.

.

74.3 Ma

036 0. 

0168

0017

0420

0178

0134

0713

Total-gas date:

600 

700

800

900

1,000

1,050

1,100

1,150

1,250

Fuse

750 

875

925

975

1,000

1,050

Fuse

850 

1,000

1,050

1,100

1,125

1,150

1,175

1,200

1,250

Fuse

22.90 

9.724

5.203

5.538

5.299

5.361

5.476

5.796

6.075

5.928

Total-gas

10.85 

7.093

6.613

6.477

6.367

6.343

6.303

Total-gas

28.22 

17.17

21.92

9.478

7.926

7.632

7.420

7.448

7.805

8.841

0.

.

.

.

.

.

.

.

date:

0.

.

.

.

.

date:

5. 

2.

2.

3.

3.

3.

3.

3.

3.

3.

3228 0. 

0465

0147

0168

0118

0057

0042

0098

0182

0057

76.8 Ma

2755 0. 

2519

0273

0192

0134

0144

0104

85.3 Ma

812 0. 

094

490

533

656

656

635

622

612

918

39Ar pet 

(pet of 40 ArR 

total)

R8a , BIOTITE

; plateau date: 74.7+0.4

0615 

0056

0013

0044

0028

0047

0262

67.4

0495 

0206

0031

0018

0005

0005

0007

0017

0024

0017

0.4 

6.9

38.8

7.8

16.4

25.0

4.7

R8a, 

Ma; no

0.2

.5

1.7

4.3

16.4

16.9

14.1

7.9

10.7

27.3

RIO,

14.8 

77.0

93.9

81.5

87.3

80.6

42.3

MICROCLINE 

plateau ; J=

36.2 

37.4

82.2

90.2

96.9

97.0

95.9

91.3

88.4

91.4

MUSCOVITE

; plateau date: 76.7+0.4

0179 

0037

0021

0019

0015

0015

0012

1.0 

4.2

6.5

11.8

23.2

20.0

33.4

51.4 

84.9

90.4

91.4

93.0

93.2

94.2

Rll, HORNBLENDE 

; plateau date: 84.9+0.3

0613 

0347

0497

0101

0051

0041

0032

0033

0045

0076

0.2 

.6

.5

5.1

9.9

12.9

22.4

21.4

14.5

12.4

37.4 

41.3

33.8

71.4

84.6

87.8

90.9

90.0

86.6

77.9

39Ar

(mol 

xlO'13 )

Ma; J=0

0.047 

.849

4.78

.965

2.03

3.08

.574

0.007244

0.049 

.126

.427

1.09

4.11

4.23

3.51

1.98

2.67

6.82

Ma; J=0

0.128 

.539

.833

1.51

2.67

2.56

4.28

Ma; J=0

0.051 

.126

.110

1.07

2.10

2.74

4.75

4.54

3.06

2.63

Apparent Apparent 

K/Ca age and error 

(mol /mol) (Ma)

.007309; wt.

0 

30

304

12

29

38

7

; wt.

1 

11

35

31

43

91

124

52

28

90

.50 

.9

.0

.4

.2

.8

.29

0.0938

.61 

.2

.5

.0

.9

.3

.0

.9

.6

.7

.007333; wt.

1 

2

19

27

38

36

50

.89 

.06

.0

.0

.7

.2

.2

.007156; wt.

0 .09 

.25

.21

.15

.14

.14

.14

.14

.14

.13

, 0.0630 g

41

71

74

74

74

74

73

g

105 

46

55

64

65

66

67

67

68

69

.1+1.0 

.3+0.4

.9+0.4

.6+0.4

.6+0.4

.6+0.4

.3+0.4

.2+1.2 

.9+0.7

.0+0.3

.1+0.3

.9+0.3

.7+0.3

.4+0.4

.8+0.3

.9+0.4

.4+0.4

, 0.0630 g

72 

78

77

76

76

76

76

.3+0.6 

.0+0.4

.4+0.4

.7+0.4

.7+0.4

.6+0.4

.9+0.4

, 0.7533 g

131 

89

93

85

84

84

85

85

85

86

.2+0.9 

.2+0.9

.2+0.5

.3+0.4

.5+0.4

.5+0.4

.0+0.4

.1+0.4

.1+0.4

.8+0.5

by later thermal activity, 244 Ma is a minimum estimate of the age of the oldest metamorphism that is 

preserved in argon systematics of these rocks. The age 

spectra of the two biotite samples from the Pittsburg 

Landing area (Hkc-29 and Hkc-36) are disturbed and 

corroborate the interpretation that these rocks were 

thermally or chemically affected after metamorphism.

The age spectra for two hornblende samples 

from sheared amphibolites near The Oxbow (Hcop-39 

and Hcop-40; fig. 10.4) also show variable discordance, but sample Hcop-39 yields a plateau date of 

227.8±5.3 Ma with some apparent argon loss. This 

date may represent cooling after the 244-Ma metamorphism, a later metamorphic event (that is, after
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 377

TABLE 10.3. 40Ar/ 39Ar age-spectrum data for samples from west-central Idaho and eastern Oregon Continued

Temperature 40Ar/ 39Ar 37Ar/ 39Ar 36Ar/ 39Ar

39Ar pet 

(pet of 40ArR 

total)

39Ar Apparent

(mol K/Ca

xlO'13 ) (mol/mol)

Apparent

age and error

(Ma)

R12, HORNBLENDE 

Total-gas date: 92.0 Ma; no plateau; J=0.007363; wt., 0.8408 g

850

950

1,000

1,050

1,100

1,125

1,150

1,175

1,200

1,250

Fuse

33

42

72

20

9

8

8

10

9

12

34

.43

.14

.72

.58

.647

.616

.471

.80

.99

.84

.45

1.789

1.635

2.636

4.333

4.545

4.520

4.573

4.577

4.462

4.432

4.403

0.0855

.1170

.2220

.0477

.0102

.0064

.0056

.0133

.0110

.0207

.0944

0.2

.3

.2

1.5

11.7

22.6

30.0

9.1

13.6

8.2

2.7

24.8

18.2

10.0

33.1

72.4

82.0

84.8

66.8

71.0

55.1

20.0

0

2

4

5

1

2

1

4

.038

.058

.034

.265

.10

.05

.38

.64

.45

.47

.76

0.29

.32

.20

.12

.11

.11

.11

.11

.11

.11

.11

106.

99.

94.

88.

90.

91.

92.

93.

91.

91.

89.

9+1.8

2+1.4

1+2.2

3+0.5

4+0.4

5+0.5

9+0.5

4+0.5

8+0.5

5+0.5

1+0.6

R12, BIOTITE 

Total-gas date: 80.3 Ma; plateau date: 82.6+0.4 Ma; J=0.007332; wt., 0.0680 g

650

850

1,000

1,050

1,125

1,200

Fuse

13

6

6

7

7

7

10

.73

.876

.887

.155

.469

.914

.50

0.0164

.0042

.0045

.0065

.0080

.0173

.0194

0.0366

.0030

.0016

.0026

.0037

.0054

.0143

1.52

26.8

27.3

13.3

11.7

12.2

7.2

1.2

86.9

92.9

89.2

85.2

79.8

59.7

0.187

3.25

3.31

1.61

1.43

1.48

.875

31

123

117

80

65

30

26

.7

.0

.0

.4

.4

.0

.8

38

77

82

82

82

81

81

.1+0.4

.3+0.4

.7+0.4

.5+0.4

.3+0.4

.7+0.4

.1+0.4

R14, HORNBLENDE 

Total-gas date: 97.9 Ma; no plateau; J=0.007400; wt., 0.7880 g

850

1,000

1,050

1,100

1,125

1,150

1,175

1,200

1,250

Fuse

47.81

30.33

22.03

14.98

16.49

15.07

14.80

14.50

14.00

22.27

3

7

13

16

16

16

17

17

17

17

.512

.545

.65

.36

.77

.96

.26

.54

.47

.16

0.129

.0804

.0551

.0303

.0347

.0301

.0291

.0286

.0265

.0536

2.3

3.2

3.1

10.9

7.0

11.8

13.6

14.0

16.3

17.7

20

23

30

48

45

49

51

51

58

34

.6

.6

.9

.7

.7

.7

.0

.1

.7

.9

0.098

.135

.130

.451

.290

.488

.564

.578

.674

.732

0.14

.07

.04

.03

.03

.03

.03

.03

.03

.03

126.

93.

88.

94.

97.

97.

97.

96.

97.

100.

7+2.0

1+0.7

5+0.6

8+0.5

8+0.6

1+0.6

8+0.5

2 + 0.5

6+0.5

7+0.6

R16, HORNBLENDE 

Total-gas date: 99.7 Ma; no plateau; J=0.007254; wt., 0.7839 g

850

1,000

1,050

1,100

1,125

1,150

1,175

1,200

1,250

Fuse

46

31

20

11

15

14

16

14

10

10

.42

.54

.71

.84

.37

.06

.86

.77

.60

.01

2.

3.

8.

9.

9.

9.

9.

9.

9.

9.

390

964

198

259

373

186

326

683

538

266

0.116

.0811

.0476

.0164

.0280

.0235

.0335

.0266

.0121

.0096

0.8

1.7

2.4

10.5

12.1

6.1

3.4

7.3

16.9

39.0

26.2

24.9

35.2

65.1

50.8

55.7

45.6

51.9

73.2

79.0

0.061

.132

.186

.822

.949

.475

.265

.569

1.33

3.06

0.22

.13

.06

.06

.06

.06

.06

.05

.05

.06

152

100

92

98

99

99

97

97

98

100

.5+1.0

.1+0.7

.8+0.6

.1+0.5

.4+0.5

.7+0.5

.7+0.5

.5+0.5

.6+0.5

.6+0.5

244 Ma), complete resetting by postmetamorphic 

thermal activity or shearing, or time of emplacement 

of the dikes. Considering the data from postmetamorphic igneous rocks, below, we suggest that 227 Ma 

may be the age of emplacement and (or) shearing of 

these rocks.
378 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 10.3. 40Ar/39Ar age-spectrum data for samples from west-central Idaho and eastern Oregon Continued

Temper- 39Ar pet 

ature 40Ar/ 39Ar 37Ar/39Ar 36Ar/ 39Ar (pet of 40ArR 

(°C) total)

Total-gas date: 83.8 Ma

650 

850

1,000

1,050

1,125

1,200

Fuse

850 

1,000

1,050

1,100

1,125

1,150

1,175

1,200

1,250

Fuse

5.955 

7.926

8.151

8.224

8.300

11.19

58.42

Total-gas

73.67 

48.67

26.44

13.25

12.01

12.17

11.12

10.09

10.14

12.14

0.0262 0. 

.0067

.0077

.0160

.0375

.0663

.5734

date: 121.1 Ma

9.917 0. 

5.095

8.583

9.668

9.918

9.904

10.16

10.28

10.31

10.24

R16 , BIOTITE

39Ar 

(mol 

xlO- 13 )

; plateau date: 87.9+0.4 Ma; J=0 .

0103 3.8 

0045 21.9

0024 31.7

0027 23.2

0032 14.3

0131 4.6

1770 .5

48.9 

83.3

91.1

90.1

88.7

65.3

10.4

R17, HORNBLENDE 

; plateau date: 106.8+0.5

1050 0.8 

0580 1.6

0342 3.6

0147 10.4

0142 11.8

0156 8.0

0127 8.6

0088 15.1

0088 23.5

0157 16.7

R18, 

Total-gas date: 110.5 Ma; no

900 

1,000

1,050

1,100

1,125

1,150

1,175

1,200

1,250

Fuse

850 

950

1,000

1,050

1,100

1,125

1,150

1,175

1,200

1,250

Fuse

38.09 

25.57

20.11

12.80

13.32

11.22

11.12

11.11

11.83

18.83

Total-gas

6$. 45 

59.35

56.01

42.17

34.14

31.37

23.89

22.07

21.36

21.23

22.54

4.151 0. 

8.557

10.54

13.12

13.31

13.22

12.96

13.48

13.31

13.31

date: 227.5 Ma

6.893 0. 

9.580

10.42

10.36

9.420

8.953

8.612

8.523

8.478

8.475

8.834

0680 0.7 

0536 3.2

0398 4.2

0164 21.5

0187 13.0

0130 17.8

0121 6.9

0128 11.3

0145 16.8

0389 4.6

58.9 

65.6

64.3

72.9

71.4

68.3

73.3

82.0

82.3

68.3

HORNBLENDE 

plateau ; J=0

48.1 

40.7

45.6

70.0

66.2

74.9

77.0

75.3

72.6

44.5

R20, HORNBLENDE 

; plateau date: 227.1+1.2

1620 0.4 

1430 .6

1330 .4

0790 1.4

0532 2.1

0447 2.9

0195 7.7

0136 14.7

0118 20.6

0115 25.8

1540 23.4

31.8 

30.1

31.4

46.6

56.1

60.1

78.7

84.8

86.8

87.1

82.9

0.347 

1.99

2.88

2.11

1.30

.420

.048

Ma; J=0

0.056 

.112

.243

.708

.808

.546

.584

1.03

1.60

1.14

.007200;

0.033 

.146

.192

.981

.595

.811

.315

.517

.765

.210

Ma; J=0

0.032 

.048

.033

.106

.165

.223

.602

1.15

1.61

2.02

1.83

Apparent Apparent 

K/Ca age and error 

(mol/mol) (Ma)

006750;

19.8 

77.4

67.3

32.5

13.5

7.85

.91

.007350;

0.05

.10

.06

.05

.05

.05

.05

.05

.05

.05

wt., 0.

0.13 

.06

.05

.04

.04

.04

.04

.04

.04

.04

.007210;

0.08 

.05

.05

.05

.06

.06

.06

.06

.06

.06

.06

wt. , 0.0609 g

35.1+0.4 

78.6+0.4

88.2+0.4

88.1+0.4

87.4+0.4

86.9+0.4

72.2+2.7

wt. , 0.7460 g

499.7+3.2 

380.5+1.8

212.4+1.0

123.6+0.6

110.1+0.6

107.0+0.5

104.9+0.6

106.4+0.5

107.3+0.5

106.6+0.5

7413 g

223.5+2.2 

130.2+0.7

115.1+1.7

112.7+0.6

111.1+0.6

105.9+0.5

107.9+0.6

105.4+0,6

108.2+0.6

105.2+0.6

wt. , 0.8099 g

266.7+4.1 

218.6+3.8

215.1+1.7

238.8+1.5

233.2+1.3

229.8+1.2

229.1+1.1

228.3+1.1

226.3+1.3

225.8+1.1

227.9+1.1

Age spectra for hornblendes from postmetamorphic 

plutons at the confluence of the Imnaha and Snake 

Rivers (Hem-23, Hdh-24, Hdh-25, and Hdh-27; fig. 10.5) 

and near Pittsburg Landing (R20; fig. 10.6) are less

complicated and generally display plateaus that represent the time when these rocks cooled below 530°C. At 

the Imnaha and Snake River confluence, the dates 

range from 213.0±2.0 to 225.0±5.3 Ma. This dispersion
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 379

TABLE 10.3. 40Ar/ 39Ar age-spectrum data for samples from west-central Idaho and eastern Oregon Continued

Temper- 39Ar pet 

ature 40 Ar/ 39Ar 37Ar/ 39Ar 36Ar/ 39Ar (pet of 40 ArR 

(°C) total)

850 

950

1,000

1,050

1,100

1,125

1,150

1,175

1,200

1,250

Fuse

850 

1,000

1,050

1,100

1,250

Fuse

Total-gas

41.19 

16.72

22.87

15.85

9.506

8.918

8.725

8.695

8.184

8.412

39.13

Total-gas

7.113 

6.895

6.998

7.292

10.22

55.45

date:

2. 

1.

2.

3.

3.

3.

3.

3.

3.

3.

3.

date:

0.

.

.

.

83.9 Ma

472 0. 

181

449

403

694

695

700

702

700

767

855

80.5 Ma

0040 0. 

0045

0081

0128

0448

2324

Total-gas date:

600 

700

800

900

1,000

1,050

1,100

1,150

Fuse

850 

950

1,000

1,050

1,100

1,125

1,150

1,175

1,200

1,250

Fuse

33.85 

7.964

6.153

5.943

6.031

6.261

6.705

7.469

8.209

Total-gas

33.35 

20.98

14.87

8.267

8.317

13.30

10.59

9.021

8.662

7.909

8.055

0.

a

.

.

.

,

.

date:

2. 

3.

3.

4.

4.

4.

4.

4.

4.

4.

4.

0047 0. 

0014

0014

0008

0007

0004

0004

0004

0002

89.2 Ma;

632 0. 

023

844

180

270

225

285

335

338

585

495

R21, HORNBLENDE

; plateau date : 83 . 9+0 . 4

1120 

0338

0533

0304

0090

0070

0064

0063

0046

0053

1110

0.9 

1.1

.5

1.1

7.0

8.5

13.7

16.6

24.8

23.9

1.9

19.8 

40.8

31.9

44.9

74.9

79.9

81.6

81.8

86.8

84.9

17.2

R21, BIOTITB 

; plateau date: 81.1+0.4

0024 

0011

0014

0025

0126

1680

76.8

0882 

0073

0013

0004

0006

0015

0025

0047

0073

21.2 

32.6

20.1

16.6

8.4

1.1

R21, 

Ma; no

0.1 

.5

3.8

11.2

24.2

8.0

8.6

19.9

16.8

89.9 

95.4

93.9

89.9

63.6

10.4

MICROCLINE 

plateau; J=0

23.0 

72.9

93.9

97.7

97.2

93.0

88.9

81.5

73.6

39Ar 

(mol 

xlO'13 )

Ma; J=0.

0.165 

.187

.086

.96

1.23

1.51

2.42

2.93

4.39

4.23

3.40

Ma; J=0.

2.60 

4.01

2.47

2.03

1.03

.140

.007313;

0.025 

.109

.809

2.41

5.21

1.72

1.85

4.29

3.62

R22, HORNBLENDE 

plateau date: 89.3+0.6 Ma ; J=0

0848 

0481

0277

0058

0061

0233

0138

0087

0074

0048

0050

0.5 

.7

.9

7.5

5.3

1.6

3.8

5.3

12.8

24.6

37.0

25.5 

33.4

47.0

83.3

82.2

50.8

64.7

75.3

78.6

86.6

85.9

0.089 

.131

.170

1.42

.994

.297

.720

.989

2.41

4.63

6.97

Apparent Apparent 

K/Ca age and error 

(mol/mol) (Ma)

006690; wt . ,

0 .21 

.44

.21

.15

.14

.14

.14

.14

.14

.14

.14

007000; wt. ,

131 

115

63

40

11

2

wt.,

Ill 

365

380

651

790

1360

1210

1430

2220

.0 

.0

.9

.5

.6

.24

0.0823

.007373; wt. ,

0 .20

.17

.14

.12

.12

.12

.12

.12

.12

.11

.12

0.7187

96.9+1 

80.5+0

86.0+0

83.8+0

83.9+0

84.0+0

83.9+0

83.8+0

83.7+0

84.2+0

79.4+0

0.0698

79.1+0 

81.3+0

81.1+0

81.1+0

80.5+0

71.6+0

g

99.7+2 

75.0+0

74.7+0

75.0+0

75.7+0

75.2+0

77.0+0

78.6+0

78.0+0

0.8212

109.7+1 

90.7+0

90.6+0

89.3+0

88.7+0

87.6+0

88.8+0

88.1+0

88.3+0

88.9+0

89.7+0

g

.0

.5

.9

.5

.4

.4

.4

.4

.4

.4

.5

g

.4

.4

.4

.4

.4

.3

.1 

.6

.7

.4

.4

.4

.4

.4

.4

g

.0

.7

.5

.5

.5

.5

.4

.4

.4

.4

.4

in dates may be due to analytical error, differential 

cooling of the pluton, or later thermal effects. Sample 

R20 was analyzed by higher precision methods and 

yielded a date of 226.7±1.1 Ma. We interpret this date 

to record a period of postmetamorphic plutonism that

occurred at about 227 Ma and after the metamorphic 

rocks had cooled below 280°C (argon-retention temperature of biotite). The thermal effect of this plutonism 

was large enough to cause partial to complete argon 

loss from some of the hornblendes in the metamorphic
380 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 10.3. 40Ar/39Ar age-spectrum data for samples from west-central Idaho and eastern Oregon Continued

Temperature 40Ar/ 39Ar 

(°C)

39Ar pet 

37 Ar/ 39Ar 36Ar/ 39Ar (pet of 40ArR 

total)

Total-gas date: 96.9 Ma

850 

950

1,000

1,050

1,100

1,125

1,150

1,175

1,200

1,250

Fuse

4T 

30

40

36

15

10

9

9

9

9

10

.74 

.64

.32

.57

.02

.65

.471

.209

.261

.505

.09

3.891 

2.846

3.806

5.822

6.494

5.547

5.491

5.304

5.105

5.045

5.087

Total -gas date

850

950

1,000

1,050

1,100

1,125

1,150

1,175

1,200

1,250

Fuse

53

31

41

49

31

20

14

12

19

13

15

.59

.43

.07

.52

.95

.89

.74

.42

.19

.97

.61

Total-gas date;

750 

875

925

975

1,050

1,150

1,200

Fuse

15 

8

7

7

7

7

7

11

.95 

.797

.856

.749

.711

.235

.374

.76

2.530

1.839

4.490

8.973

12.35

13.26

13.56

13.63

13.67

13.65

13.56

83.4 Ma:

0.0350 

.0117

.0016

.0007

.0

.0

.0004

.0026

0.1200 

.0827

.1120

.0997

.0265

.0114

.0072

.0059

.0052

.0058

.0076

R23,

; no

0.3 

.3

.2

.4

2.3

4.6

10.6

18.0

21.0

19.7

22.6

R24, 

: 91.8 Ma; no

0.140

.0797

.118

.148

.0866

.0486

.0275

.0197

.0432

.0248

.0308

1,1

1.4

1.1

1.3

3.2

4.3

11.1

26.6

8.8

19.2

21.9

HORNBLENDE

plateau ; J=0

15.6 

20.9

18.3

20.7

51.1

72.5

82.2

85.6

87.5

86.0

81.7

HORNBLENDE 

plateau ; J=0 .

23.1

25.5

16.1

13.2

22.9

36.2

52.0

61.6

39.1

55.2

48.5

R24, MUSCOVITE 

plateau date: 83.4+0.4 Ma;

0.0342 

.0068

.0037

.0037

.0035

.0019

.0024

.0174

Total-gas date: 77.8 Ma

600 

700

800

900

1,000

1,050

1,100

1,100

1,250

Fuse

21 

6

6

6

6

6

6

7

7

8

.35 

.470

.118

.141

.245

.611

.967

.426

.669

.346

0.0344 

.0273

.0207

.0073

.0097

.0091

.0094

.0127

.0118

.0062

0.0343 

.0033

.0011

.0005

.0006

.0017

.0029

.0043

.0071

.0071

0.7 

2.0

3.6

5.7

12.6

23.3

42.4

9.6

R24, 

; no

0.6 

1.3

4.7

12.6

13.4

10.6

10.0

12.9

15.8

18.0

36.6 

77.2

86.1

85.9

86.4

92.0

90.4

56.2

MICROCLINE 

plateau ; J=0

54.7 

84.7

94.5

97.4

96.9

92.2

87.7

82.9

81.2

74.8

39Ar 

(mol 

xlO-13 )

.006873;

0.038 

.046

.031

.058

.312

.643

1.47

2.50

2.91

2.73

3.13

006832;

0.057

.071

.057

.068

.160

.218

.556

1.33

.442

.963

1.10

Apparent Apparent 

K/Ca age and error 

(mol/mol) (Ma)

wt.,

0

wt.,

0

J*0. 007107;

0.095 

.252

.465

.731

1.61

2.98

5.42

1.23

.007168;

0.097 

.225

.793

2.15

2.28

1.81

1.71

2.19

2.69

3.06

14 

44

333

743

1160

196

wt.,

15 

19

25

71

53

57

55

41

44

83

0.7618

,13 

.18

.14

.09

.08

.09

.09

.10

.10

.10

.10

0.8055

.21

.28

.12

.06

.05

.04

.04

.04

.04

.04

.04

wt., 0.

.9 

.4

.0

.0

.0

.0

.0

.0

0.0713

.1 

.1

.1

.6

.8

.2

.3

.0

.1

.9

g

79 

77

89

91

92

93

94

95

97

98

99

g

146

96

79

78

88

90

91

91

90

92

90

580

73 

85

84

83

83

83

83

82

g

151 

69

73

75

76

77

77

77

78

79

.0+2.1 

.8+1.2

.4+2.2

.5+1.2

.8+0.5

. 2 TO . 5

.0+0.5

.1+0.5

.8+0.5

.6+0.5

.4+0.5

.3+4.3

.2+0.7

.7+1.7

.5+0.9

.0+0.8

. 8+0 . 5

.9+0.5

.9+0.5

.1+0.5

.6+0.5

.8+0.5

g

.3+1.2 

.0+0.5

.7+0.4

.8+0.4

.5+0.4

.4+0.4

.5+0.4

.7+0.4

.5+0. 8 

.5+0.4

.3+0.4

.8+0.4

.6+0.4

.2+0.4

.3+0.4

.9+0.4

.8+0.4

.0+0.4
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 381

TABLE 10.3. 40Ar/ 39Ar age-spectrum data for samples from west-central Idaho and eastern Oregon Continued

Temperature 40Ar/ 39Ar 

(°C)

37Ar/ 39Ar 36Ar/ 39Ar

39Ar pet 39Ar Apparent Apparent 

(pet of 40ArR (mol K/Ca age and error 

total) xlO'13 ) (mol /mol) (Ma)

R25 , BIOTITE

Total-gas date: 83.7 Ma; plateau date: 84.1+0.4 Ma; J=0. 007340; wt.,

650

850

1,000

1,050

1,125

1,200

Fuse

850 

950

1,000

1,050

1,100

1,125

1,150

1,175

1,200

1,250

Fuse

23 

7

7

7

7

8

29

.95 

.834

.024

.110

.052

.109

.95

0.2735 

.1717

.0105

.0080

.0105

.0975

.0975

Total-gas date;

36 

17

17

11

10

11

11

18

17

13

11

.21 

.82

.45

.75

.57

.06

.73

.81

.77

.53

.75

Total-gas

650

850

1,000

1,050

1,125

1,200

Fuse

9

7

7

7

7

10

49

.944

.476

.423

.440

.549

.84

.72

3.767 

3.902

5.923

6.486

6.172

6.127

6.201

6.099

6.254

6.337

6.281

date: 89.0

0.0634

.0133

.0099

.0211

.0257

.0775

.1644

Total-gas date

850 

950

1,000

1,050

1,100

1,125

1,150

1,175

1,200

1,250

Fuse

47 

31

26

19

18

17

18

14

18

14

15

.94 

.97

.49

.93

.00

.27

.86

.09

.51

.13

.17

5.651 

6.827

8.640

14.02

15.19

14.51

15.44

15.09

15.19

15.20

15.30

0.0674 

.0054

.0018

.0021

.0019

.0054

.0809

: 107.5

0.0903 

.0343

.0329

.0118

.0065

.0076

.0100

.0341

.0310

.0162

.0099

0.3 

5.8

27.0

16.0

35.0

14.3

1.7

R26,

Ma; no

2.3 

2.1

1.6

6.3

22.3

14.7

7.1

2.0

2.8

8.1

30.9

16.9 

79.8

92.5

91.4

92.2

80.3

20.2

HORNBLENDE

plateau ; J=0 .

27.1 

44.7

46.8

74.6

86.3

84.8

79.0

49.0

51.2

68.2

79.3

R26, BIOTITE 

Ma, ; plateau date : 89. 9+0 . 6

0.0217

.0028

.0015

.0015

.0020

.0132

.1480

: 111.8

0.1280 

.0723

.0634

.0400

.0344

.032

.0383

.0226

.0371

.0220

.0251

0.8

13.8

35.4

21.6

23.1

4.6

.6

R27, 

Ma; no

2.0 

1.6

1.5

4.1

8.5

9.0

7.3

14.3

9.8

16.7

25.2

35.5

89.1

93.9

94.1

92.3

64.1

11.9

HORNBLENDE 

plateau ; J=0 .

21.9 

34.8

31.8

46.2

50.0

51.6

46.4

61.6

47.2

62.4

59.0

0.034 

.646

2.99

1.77

3.87

.58

.183

006691;

0.223 

.199

.151

.605

2.15

1.42

.684

.192

.783

.783

2.98

1. 

3.

49.

65.

49.

36.

5.

wt.,

0.

.

.

.

.

.

.

.

.

; J=0.007327;

0.093

1.59

4.07

2.49

2.67

.529

.074

007171;

0.099 

.078

.074

.199

.409

.436

.350

.689

.472

.805

1.21

8.

39.

52.

24.

20.

6.

3.

wt.,

0.

.

.

.

.

.

.

.

.

90

03

5

3

5

1

34

0.7132

14 

13

09

08

08

09

08

09

08

08

08

wt. , 0

20

0

7

7

2

71

16

0.7382

09 

08

06

04

03

04

03

03

03

03

03

0.0628 g

52.8+0. 

81.0+0.

84.0+0.

84.0+0.

84.1+0.

84.2+0.

78.3+0.

g

114.6+0. 

93.7+0.

96.0+0.

102.8+0.

106.8+0.

108.8+0.

108.5+0.

107.8+0.

106.7+0.

108.0+0.

109.0+0.

.0697 g

46.1+0.

85.9+0.

89.9+0.

90.8+0.

89.8+0.

89.7+0.

76.8+0.

g

130.8+1. 

138.4+1.

105.7+0.

115.2+0.

112.8+0.

111.6+0.

109.7+0.

108.9+0.

109.5+0.

110.6+0.

112.1+0.

7 

4

4

4

4

47 '

8 

5

5

5

5

5

5

6

6

6

6

7

4

4

5

5

5

8

6 

0

7

6

6

6

6

6

6

6

6

rocks and to disturb the argon age spectra of metamorphic biotite.

Preliminary age-spectrum data for seven samples 

from plutons in the Bald Mountain batholith (fig. 10.1; 

five biotites and two hornblendes; L.W. Snee and W.H. 

Taubeneck, unpub. data, 1987) range from 142.3+0.8

to 144.6±1.0 Ma. This narrow range in dates recorded 

by both hornblende and biotite within five distinct 

plutons of the batholith suggests rapid cooling, which 

in turn indicates emplacement over a very short period 

of time and at high crustal levels. This period of intrusive activity was apparently regional in scope as
382 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 10.3. 40Ar/ 39Ar age-spectrum data for samples from west-central Idaho and eastern Oregon Continued

Temperature 40Ar/ 39Ar 37Ar/ 39Ar 36Ar/ 39Ar

39Ar pet 

(pet of 40ArR 

total)

39Ar Apparent

(mol K/Ca

xlO' 13 ) (mol/mol)

Apparent

age and error

(Ma)

R28, HORNBLENDE

Total-gas

850 13.40 

1,000 9.694

1,050 12.61

1,100 10.57

1,125 9.608

1,150 9.457

1,175 9.301

1,200 9.451

1,250 8.882

Fuse 9.043

Total-gas

850 37.06 

1,000 26.75

1,050 25.30

1,100 11.08

1,125 9.86

1,150 9.86

1,175 9.68

1,200 9.17

1,250 11.76

Fuse 11.00

Total-gas

850 32.43 

1,000 34.83

1,050 21.03

1,100 12.58

1,125 11.19

1,150 16.73

1,200 10.65

1,250 10.80

Fuse 10.31

Total-gas

850 42.59 

950 20.31

1,000 18.02

1,050 13.97

1,100 12.28

1,125 11.69

1,150 11.87

1,175 12.55

1,200 13.44

1,250 14.54

Fuse 18.87

date: 107.0

1.383 

.2511

1.384

3.429

4.211

4.454

4.538

4.582

4.737

4.755

date: 107.8

1.997 

2.067

4.713

5.556

5.528

5.524

5.469

5.486

5.495

5.439

date: 118.8

4.138 

2.530

4.717

5.424

5.471

5.416

5.376

5.386

5.353

date: 109.2

2.597 

12.60

13.72

13.57

13.39

13.42

13.55

13.61

13.65

13.56

13.52

Ma; plateau date:

0.0187 

.0065

.0164

.0089

.0055

.0050

.0045

.0049

.0027

.0034

1.3 

3.8

1.9

6.2

10.4

9.5

13.2

8.8

20.6

24.2

107.7+0.5

59.5 

80.3

62.4

77.7

86.5

88.1

89.6

88.5

95.1

93.1

R29, HORNBLENDE 

Ma; plateau date: 108.1+0.5

0.0974 

.0671

.0622

.0107

.0064

.0064

.0057

.0039

.0126

.0101

0.5 

.6

.7

6.7

12.0

12.2

16.9

29.0

6.8

14.6

22.8 

26.4

28.8

75.4

85.2

85.3

87.1

92.0

71.8

76.8

R30, HORNBLENDE 

Ma; plateau date: 118.1+0.6

0.0394 

.0802

.0412

.0124

.0070

.0054

.0051

.0059

.0041

0.6 

.0

1.1

4.6

6.7

11.2

28.7

12.2

34.8

65.1 

32.5

43.8

74.5

85.2

89.1

87.7

87.7

92.4

R34, HORNBLENDE 

Ma; plateau date: 109.1+0.5

0.116 

.0428

.0354

.0214

.0159

.0142

.0146

.0170

.0202

.0238

.0382

1.9 

3.5

3.8

8.1

16.6

16.2

14.1

8.7

9.3

9.6

8.3

19.7 

42.6

47.8

62.3

70.1

73.1

72.5

68.5

63.4

58.9

45.7

Ma; J=0

0.232 

.664

.332

1.10

1.83

1.67

2.31

1.55

3.62

4.24

Ma; J=0

0.0634 

.0753

.0902

.847

1.52

1.55

2.14

3.67

.860

1.85

Ma; J=0

0.0737 

.0051

.135

.545

.797

1.34

1.44

1.44

4.13

Ma; J=0

0.105 

.188

.208

.438

.902

.881

.765

.470

.503

.519

.453

.007334;

0.38 

2.07

.38

.15

.12

.12

.11

.11

.11

.11

.007326;

0.26 

.25

.11

.09

.09

.09

.09

.09

.09

.09

.007093;

0.13 

.21

.11

.10

.10

.10

.10

.10

.10

.007270;

0.20 

.04

.04

.04

.04

.04

.04

.04

.04

.04

.04

wt. , 0.7885 g

102. 4+0. 6 

100.1+0.5

101.3+0.5

105.5+0.5

106.7+0.5

107.0+0.5

107.9+0.5

107.4+0.5

108.4+0.5

108.0+0.5

wt. , 0.7110 g

108.3+1.0 

91.0+0.7

93.9+0.7

107.1+0.6

107.7+0.5

108.0+0.5

108.1+0.5

108.2+0.6

108.3+0.6

108.3+0.6

wt., 0.7041 g

251.7+1.3 

139.3+2.6

114.2+0.6

116.2+0.6

118.0+0.6

118.4+0.6

118.5+0.6

117.3+0.6

117.9+0.6

wt. , 0.8333 g

106.6+1.7 

109.9+0.6

109.5+0.6

110.5+0.6

109.5+0.6

108.7+0.6

109.3+0.6

109.3+0.6

108. 3tf. 6

108.8+0.6

109.5+0.6
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 

TABLE 10.3. 40Ar/ 39Ar age-spectrum data for samples from west-central Idaho and eastern Oregon Continued

383

Temperature 40Ar/ 39Ar 37Ar/ 39Ar 

<°C)

39Ar

36Ar/39Ar (pct Qf

total)

pet

40Ar Ar R

39Ar 

(mol 

xlO'13 )

Apparent Apparent 

K/Ca age and error 

(mol/mol) (Ma)

R35, MUSCOVITE

Total-gas date: 98.6

750 12.49 

875 9.735

925 8.667

975 8.195

1,015 7.959

1,200 8.062

Fuse 10.36

Total-gas

500 8.568 

650 6.861

850 6.570

1,050 6.615

1,250 6.549

Fuse 6.882

Total-gas

650 204.1 

850 44.28

950 15.08

1,000 15.26

1,050 13.78

1,100 12.72

1,150 12.54

1,200 12.52

1,250 16.17

1,350 17.16

1,450 49.32

Fuse 81.01

0.8865 

.6722

.0384

.0112

.0010

.0023

.2632

date: 80.1

___

 

 

 

_ .

date: 115.4

5.543 

3.574

3.792

3.804

3.789

3.774

3.784

3.823

3.791

3.782

3.781

3.790

Total-gas date: 224.5 Ma

900 789.8

1,000 198.1

1,050 61.22

1,120 38.29

Fuse 51.16

Total-gas

800 99.97 

1,000 66.40

1,050 51.58

1,150 38.58

Fuse 35.38

39.78

17.02

19.60

16.40

16.78

date: 214.7

13.86 

14.48

12.02

7.96

11.74

Ma; plateau date

0.0205 0.6 

.0077 2.0

.0041 4.6

.0021 13.6

.0012 25.4

.0011 36.2

.0088 4.3

R36, 

Ma; plateau date

0.0094 6.0 

.0017 14.7

.0007 25.5

.0008 29.3

.0006 20.9

.0014 3.6

7-30-5,

Ma; plateau date

0.3470 0.4 

.1020 .8

.0124 6.0

.0130 4.6

.0084 8.4

.0046 26.4

.0039 25.3

.0035 21.9

.0156 2.8

.0193 2.7

.1270 .4

.2330 .2

: 98.3+0.2 Ma; J«0.

51.9 

77.3

86.0

92.4

95.3

95.8

75.1

0.069 

.216

.506

1.49

2.79

3.98

.475

BZOTZTE 

: 80.5+0.4 Ma; J=0 .

67.6 

92.5

96.7

96.4

97.0

94.0

HORNBLENDE

: 112.0+0.7

50.0 

32.7

77.7

76.7

84.0

91.5

93.0

94.1

73.2

68.5

24.4

15.3

Hem-23, HORNBLENDE 

; plateau date: 223.3+5.9 Ma

2.579 3.53

.5702 2.44

.1165 16.44

.0380 24.35

.0816 53.24

Hdh-24, 

Ma; plateau date

0.2784 2.48 

.1543 2.76

.1007 2.60

.5670 69.05

.0448 23.11

3.9

15.7

46.4

74.2

55.6

HORNBLENDE 

: 215.6+3.6

18.9 

33.1

44.3

3.72 

9.06

15.7

18.0

12.8

2.18

Ma; J*0

0.821 

1.84

13.2

10.1

18.5

57.9

55.6

48.0

6.15

6.02

.862

.419

007340; wt., 0.0628 g

0.59 

.77

13.6

46.5

53.4

23.0

1.98

84.4+0.7 

97.5+0.5

96.7+0.5

98.2+0.5

98.4+0.5

99.4+0.5

100.9+0.5

007174; wt., 0.1243 g

_ _ _

 

 

 

 

.005480;

0.09 

.15

.14

.14

.14

.14

.14

.14

.14

.14

.14

.14

; J=0. 004577; wt

1.5

1.0

6.9

10.2

22.3

Ma; J=0

4.1 

4.6

4.3

58.3 114.1

65.3 38.2

0.01

.03

.03

.03

.03

.005570;

0.04 

.04

.04

.07

.04

73.5+0.4 

80.3+0.4

80.4+0.4

80.7+0.4

80.4+0.4

81.8+0.8

wt., 1.1493g

801.8+4.2 

137.8+0.8

112.2+0.6

112.1+0.7

111.0+0.6

111.5+0.5

111.8+0.6

112.8+0.6

113.4+0.6

112.5+0.6

115.4+1.2

118.3+3.6

. , approx . 1 . 5 g

246.6+214.7

242.7+96.9

223.5+11.4

223.1+6.1

223.2+5.8

approx . 1 . 5 g

181.8+22.4 

210.7+18.1

217.7+14.6

214.1+3.2

220.1+3.8
384 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 10.3. 40Ar/39Ar age-spectrum data for samples from west-central Idaho and eastern Oregon Continued

Temperature 40Ar/ 39Ar 

(°C)

Total-gas date:

1,025 

1,075 

1,150 

Fuse

33.34

24.39 

24.88 

24.92

Total-gas date:

900 

1,025 

1,075 

1,100 

Fuse

171.3 

27.54 

25.35 

28.20 

26.95

Total-gas date:

600 

800 

900 

1,025 

Fuse

16.68 

38.32 

40.73 

40.28 

41.25

Total-gas date :

900 

1,025 

2,075 

1,100 

Fuse

32.33 

27.05 

29.12 

25.78 

26.76

37Ar/ 39Ar 36Ar/ 39Ar

213.0

7.68 

7.64 

7.79 

7.96

225.5

13.19 

9.94 

9.80 

10.65 

11.25

218.3

 

225.8

5.45 

6.61 

6.51 

6.49 

6.69

Total-gas date: 224.6 Ma

900 

1,025 

1,075 

1,100 

Fuse

211.8 

298.4 

59.33 

36.56 

63.50

Total-gas date:

900 

1,025 

1,075 

Fuse

154.8 

75.12 

36.01 

32.65

Total-gas date:

800 

1,025 

1,070 

1,100 

Fuse

53.83 

44.59 

42.50 

43.68 

45.07

29.78 

28.81 

30.56 

29.54 

31.20

241.1

23.87 

32.31 

31.56 

33.51

231.0

4.64 

24.60 

26.51 

27.47 

27.11

39Ar 

(pet of 

total)

Hdh-25 

Ma; plateau date:

0.0386 

.0089 

.0112 

.0115

18.28 

54.25 

11.91 

15.57

Hdh-27 

Ma; plateau date:

0.4985 

.0153 

.0089 

.0181 

.0142

Ma ; plateau

0.0202 

.0057 

.0078 

.0036 

.0114

2.26

40.60 

35.14 

7.35 

14.65

pet 

4 °ArR

39Ar Apparent Apparent 

(mol K/Ca age and error 

xlO'13 ) (mol/mol) (Ma)

, HORNBLENDE 

213.0+2.0 Ma; J=0. 005570

67.7 39.4 0 

91.8 116.8 

89.2 25.6 

89.0 33.5

, HORNBLENDE 

225.0+2.3 Ma

14.6 

86.6 

92.8 

84.2 

87.9

r

3 

64 

55 

11 

23

HRc-29, BZOTZTE 

date: 233.5+2.5 Ma;

7.22 

32.41 

13.39 

27.89 

19.09

Hkc-32 

Ma; plateau date:

0.0345 

.0137 

.0188 

.0078 

.0096

9.33 

36.40 

10.14 

21.45 

22.68

64.2 

95.6 

94.3 

97.3 

91.8

11 

51 

21 

44 

30

, HORNBLENDE 

229.3+2.6 Ma;

69.9 

87.0 

82.7 

93.1 

91.5

21 

84 

23 

49 

52

Hkc-34, HORNBLENDE 

; plateau date 225.8+6.7 Ma; J=0 .

0.6497 

.9512 

.1318 

.0534 

.1433 -

2.02 

2.97 

12.02 

40.63 

42.36

Hkc-35 

Ma; plateau date:

0.4253 

.1802 

.0474 

.0346

2.84 

17.82 

31.31 

48.04

Hkc-36 

Ma; plateau date:

0.0935 

.0275 

.0217 

.0243 

.0262

10.31 

35.09 

24.28 

11.17 

19.15

10.5 

6.6 

38.6 

63.5 

37.4

, HORNBLENDE 

239.2+7.6 Ma

20.1 

32.7 

68.4 

77.2

, HORNBLENDE 

238.6+3.4 Ma

49.4 

86.3 

90.1 

88.8 

87.8

0 

1 

5 

17 

18

r

1

8 

14 

21

r

8 

29 

20 

9 

16

; wt . ,

.07 

.07 

.07 

.07

JsO. 005570; wt. ,

.6 0 

.3 

.6 

.6 

.2

JsO. 003578

.5 

.5 

.3 

.3 

.3

J«0. 005570

.7 0 

.6 

.6 

.9 

.7

005570; wt

.9 0 

.3 

.1 

.3 

.1

J=0. 005570

.3 0

.1 

.3 

.9

J*0. 003585

.6 0 

.4 

.3 

.4 

.0

.04 

.05 

.05 

.05 

.OS

: wt.,

___

; wt . ,

.10 

.08 

.08 

.08 

.08

approx . 1.5

214.8+3.5 

213.2+2.5 

211.4+3.1 

211.3+2.9

g

approx . 1 . 5 g

238.0+33.1 

226.5+2.9 

223.7+2.8 

225.6+4.1 

225.3+3.6

approx. .25

67.8+2.5 

222.2+2.6 

232.4+3.1 

236.8+2.8 

229.3+2.8

approx . 1.5

214.6+4.1 

223.3+2.8 

228.23.7 

227.5+2.8 

231.6+2.9

9

g

. , approx . 1 . 5g

.02 

.02 

.02 

.02 

.02

; wt . ,

.02 

.02 

.02 

.02

; wt.,

.11 

.02 

.02 

.02 

.02

215.4+103. 

191.9+112. 

221.4+15.3 

224.0+5.4 

228.9+7.8

approx . 1.5

293.3+60.9 

236.4+11.8 

236.9+6.0 

242.4+4.2

approx . 1.5

164.8+5.8 

237.2+3.1 

236.4+3.2 

239.2+3.9 

243.8+3.7

i

8

g

g
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 385

TABLE 10.3. 40Ar/ 39Ar age-spectrum data for samples from west-central Idaho and eastern Oregon Continued

Temperature 40Ar/ 39Ar 

(°C)

39Ar pet 

37 Ar/ 39Ar 36Ar/ 39Ar (pet of 40ArR 

total)

39Ar Apparent Apparent 

(mol K/Ca age and error 

xlO'13 ) (mol/mol) (Ma)

Hkc-36, BIOTITE

Total-gas date: 216.8 Ma; no plateau; J»0. 003578; wt., approx. .25

600

800

900

1,025

Fuse

657 

50.

73.

59.

73.

10

52

23

43

24

Total-gas date

950

1,025

1,070

1,125

Fuse

557 

43.

37.

33.

34.

11 

41

68

17

07

Total-gas date:

900 

1,000

1,050

1,100

Fuse

307. 

274.

129.

43.

43.

1

7

9

43

92

Total-gas date:

900 

1,000

1,050

1,100

Fuse

850 

975

1,075

1,175

1,275

fuse

29. 

23.

34.

32.

28.

29 

20

28

66

75

~ ""* ~~

 

 

 _

: 239.1

9.60 

16.08

14.23

11.93

13.44

224.9

20.32 

11.73

19.28

31.73

20.92

215.4

14.01 

14.75

14.55

14.23

15.29

Total-gas date: 148

87. 

67.

57.

37.

32.

55.

20 

87

72

43

99

65

5.806 

3.541

23.22

23.81

23.75

24.68

0.1944 

.0459

.1140

.0649

.1209

; plateau

0.0859 

.0450

.0253

.0143

.0164

8.64 

34.66

15.06

24.10

17.53

11.8 

73.1

54.0

67.7

51.2

Hkc-37, HORNBLENDE 

date: 235.0+3.5 Ma;

7.29 

11.85

28.25

26.85

25.77

Hcop-39 

Ma; plateau date:

0.9538 

.8461

.3603

.0554

.0562

7.35 

3.02

3.23

44.42

41.97

Hcop-40

Ma; plateau date;

0.0302 

.0090

.0414

.0396

.0223

29.72 

41.14

3.66

4.34

21.15

55.4 

72.4

83.3

90.2

89.0

8 

35

15

24

18

Ja

6

10

25

24

23

, HORNBLENDE 

227.8+5.3 Ma;

8.8 

9.3

19.3

66.4

66.1

, HORNBLENDE

3

1

1

18

17

219.2+7.6 Ma;

73.5 

94.1

67.9

67.8

81.4

D81-1, HORNBLENDE 

.5 Ma; plateau date: 145.1+1.5

0.2214 

.1583

.1280

.0645

.0522

.1276

2.3

2.5

8.2

27.8

43.0

16.2

25.5 

31.5

37.6

54.0

58.8

35.7

28 

40

3

4

20

.9

.7

.5

.8

.1

0.004577;

.7 0

.9

.9

.6

.6

J-0. 004577

.1 0 

.3

.4

.9

.9

J*0. 005570

.9 0 

.0

.6

.2

.6

-_ 

 

 

___

wt.,

.05 

.03

.04

.04

.04

; wt

.03 

.04

.03

.02

.02

; wt

.04 

.04

.04

.04

.03

Ma; J«0.004290;

0

1

6

10

3

.544 0 

.598

.96

.64

.3

.87

.09 

.15

.02

.02

.02

.02

48 

223

238

242

227

approx

237 

245

244

233

236

g

.8+8.3

.9+3.3

.7+5.6

.6+3.9

.3+5.1

. 1.5 g

.4+10.6 

.3+6.8

.4+4.1

.5+3.4

.6+3.6

. , approx . 1 . 5 g

212 

201

198

227

228

.8+74.2 

.7+99.7

.6+56.8

.1+4.9

.4+5.9

. , approx . 1 . 5 g

206 

217

221

211

223

wt. 1.

164 

158

160

150

144

147

.4+3.6 

.5+2.9

.9+13.7

.8+13.2

.5+4.1

1024 g

.2+6.5 

.1+5.2

.4+2.0

.0+1.3

.2+1.2

.5+1.8

indicated by the hornblende plateau date of 145.1±1.5 

Ma for slightly metamorphosed tonalite that intruded 

the Seven Devils Group along Rapid River southwest 

of Riggins (sample D81-1; fig. 10.6) and by an 40Ar/39Ar 

date of about 145 Ma on hornblende from pre-accretionary quartz diorite near Orofino (fig. 10.1; Fleck 

and Criss, 1987; Davidson and Snee, 1989; Davidson, 

1990).

METAMORPHIC ROCKS OF THE RIGGINS GROUP AND 

SEVEN DEVILS ISLAND ARC

In an attempt to determine the age of metamorphism and deformation that affected the island-arc 

rocks near the Salmon River suture zone, 15 samples 

were analyzed by the 4°Ar/39Ar method. Rocks from 

the Riggins Group and Seven Devils island arc were
386 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 10.3. 40Arl39Ar age-spectrum data for samples from west-central Idaho and eastern Oregon Continued

Temperature 40Ar/ 39Ar 37Ar/ 39Ar 

(°C)

550 

700

800

900

950

1,000

1,025

1,050

1,075

1,100

1,125

1,150

1,200

Fuse

750 

1,050

1,100

Fuse

550 

700

800

900

950

975

1,000

1,025

1,050

1,075

1,100

1,125

1,150

1,175

1,225

Fuse

750 

875

925

975

1,015

1,050

1,100

Fuse

Total -gas

66.33 

50.94

115.9

63.77

45.93

7.256

6.650

7.131

7.855

7.478

6.881

6.479

6.503

21.21

Total-gas

23.175 

6.491

7.718

10.272

Total -gas

142.0 

100.2

114.6

71.34

60.74

91.15

32.49

15.30

9.167

7.388

7.936

11.99

12.37

8.667

8.725

8.557

Total-gas

6.193 

5.597

5.563

5.738

5.600

5.582

6.189

113.488

date:

5. 

3.

4.

6.

7.

5.

5.

4.

4.

4.

4.

4.

4.

4.

date:

8.

.

date:

9. 

5.

4.

4.

8.

10.

11.

9.

8.

8.

8.

8.

8.

8.

8.

8.

date:

0.

m

.

^

m

80.9

629 

856

291

187

441

468

102

914

850

794

849

849

835

803

78.2

719 

02527

2476

8158

80.4

144 

191

785

220

194

97

29

169

443

190

253

235

159

061

208

387

73.7

0035 

0016

0012

0042

0015

0009

0051

1614

36Ar/ 39Ar

39Ar 

(pet of 

total)

2KE080,

Ma; plateau date

0.2184 

.1575

.3708

.1964

.1386

.0073

.0049

.0066

.0088

.0074

.0055

.0041

.0041

.0544

0.0

.1

.1

.2

.5

10.5

15.7

9.3

6.0

6.5

9.8

15.1

23.2

3.0

2KE080 

Ma; plateau date

0.0606 

.0030

.0074

.0163

1.8 

65.2

21.9

11.1

2KE086, 

Ma; plateau date

0.4356 

.3118

.3800

.2256

.1897

.2968

.0939

.0354

.0146

.0083

.0103

.0238

.0249

.0127

.0128

.0121

0.0

.1

.1

.5

.2

.2

.6

2.4

5.7

16.1

17.7

2.9

3.1

8.2

13.9

28.2

1KE052, 

Ma; plateau date

0.0040

.0018

.0017

.0024

.0018

.0017

.0037

.3750

5.4

16.9

27.2

10.3

13.9

17.2

9.0

.1

pet

40 ArArR

HORNBLENDE

: 81.2+0.4

3.4 

9.2

5.7

9.7

12.1

76.2

84.2

78.1

71.6

75.5

81.8

87.3

87.2

26.0

, BIOTITE 

: 78.2+0.4

25.7 

86.4

71.7

53.6

HORNBLENDE 

: 80.5+0.6

9.9 

8.5

2.3

7.0

8.8

4.7

17.3

36.2

60.1

75.2

69.8

46.5

45.7

63.8

63.8

65.7

MUSCOVITE 

: 73.8+0.4

80.7 

90.4

90.9

87.7

90.2

91.0

82.4

2.4

39Ar

(mol 

xlO-13 )

Ma;

0.

.

.

.

14.

21.

12.

8.

8.

13.

20.

31.

4.

Ma;

0. 

18.

6.

3.

Ma;

0.

.

.

.

.

.

1.

4.

12.

14.

2.

2.

6.

11.

22.

Ma;

5. 

18.

29.

11.

14.

18.

9.

.

J=0.

0418 

0867

123

216

667

3

3

6

19

89

4

6

6

01

J=0.

523 

5

21

16

J=0.

0381 

0607

0910

384

199

137

437

95

55

9

2

34

48

58

2

6

J=0.

70 

0

0

0

8

3

62

133

Apparent Apparent 

K/Ca age and error 

(mol/mol) (Ma)

008163; wt. ,

0.

.

.

.

.

.

.

.

.

.

.

.

09

13

12

08

07

09

10

10

11

11

11

11

11

11

007930; wt. ,

0. 

20.

2.

0593 

7

10

637

008190; wt. ,

0.

.

.

.

.

.

.

.

.

.

.

.

.

.

06 

10

11

12

06

05

05

06

06

06

06

06

06

06

06

06

008244; wt. ,

149 

328

434

122

351

554

102

3. 22

1.0388 g

32.5+29.8 

68.0+7.4

95.4+19.9

89.0+3.2

79.9+1.8

79.6+0.4

80.6+0.4

80.2+0.4

80.9+0.4

81.2+0.4

81.0+0.4

81.3+0.4

81.6+0.5

79.4+0.7

0.0446 g

83.2+1.0 

78.5+0.4

77.5+0.4

77.2+0.4

0.9842 g

195.9+62.2 

121.0+30.9

39.1+12.8

72.4+5.3

77.0+6.8

62.4+17.3

81.0+0.8

79.9+0.5

79.5+0.4

80.2+0.5

80.0+1.0

80.5+0.5

81.5+0.5

79.8+0.4

80.4+0.5

81.1+0.4

0.1812 g

72.9+0.4 

73.8+0.4

73.7+0.4

73.3+0.4

73.6+0.4

74.0+0.4

74.3+0.4

39.6+13.2
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 387

TABLE 10.3. 40Ar/ 39Ar age-spectrum data for samples from west-central Idaho and eastern Oregon Continued

Temperature 40 Ar/ 39Ar 37Ar/ 39Ar

39Ar 

36Ar/ 39Ar (pet of 

total)

pet 

4 °ArR

39Ar 

(mol 

xlO'13 )

Apparent Apparent 

K/Ca age and error 

(mol/mol) (Ma)

1KE052, MICROCLINE

Total-gas date: 63.5 Ma; no plateau; J=0.007990;

400 

500

600

700

800

900

1,000

1,050

1,100

1,150

1,200

1,260

1,300

Fuse

750 

875

925

975

1,050

1,100

Fuse

850 

950

1,000

1,050

1,100

1,150

1,200

1,250

1,450

Fuse

550 

650

800

900

1,000

1,125

Fuse

23.500 

34.384

11.282

4.656

4.661

4.422

4.506

4.938

5.632

6.253

7.336

7.425

6.785

6.938

Total-gas

5.405 

5.208

5.149

5.184

5.404

7.066

38.81

Total -gas

0.75 

10.52

9.281

9.058

8.718

8.442

8.286

8.465

8.434

10.93

Total-gas

22.72 

10.65

8.237

8.003

8.010

7.898

8.102

0.0025 

.0068

.0025

.0028

.0030

.0037

.0049

.0042

.0024

.0018

.0019

.0028

.0017

.0010

date: 70.7

0.0024

 

 

.0001

.0017

.0197

date: 86.6

2.405 

4.575

4.829

4.881

4.916

4.987

5.003

5.019

5.166

5.240

date: 87.4

0.0880 

.0057

.0013

.0006

.0021

.0014

.0079

0.0627 0.1 

.0794 .1

.0113 .6

.0027 1.4

.0026 3.8

.0013 5.3

.0011 13.4

.0021 9.7

.0042 7.8

.0050 7.3

.0097 6.9

.0096 9.1

.0069 19.6

.0074 15.0

2KE087, 

Ma; plateau date:

0.0024 11.2 

.0012 25.5

.0009 25.7

.0011 18.3

.0016 14.4

.0074 4.5

.0117 0.3

21.2 

31.8

70.4

82.5

83.4

91.1

93.0

87.5

77.8

71.6

61.0

61.9

69.8

68.2

MUSCOVITE 

71.0+0.4

86.8 

93.0

94.5

93.8

91.0

69.0

10.9

MP201, HORNBLENDE 

Ma; plateau date: 86.3+0.4

0.0740 2.2 

.0107 10.0

.0067 8.3

.0060 7.2

.0047 9.3

.0040 14.2

.0035 17.4

.0039 13.8

.0040 12.8

.0125 4.8

TC-S-41-16 

Ma; plateau date:

0.0509 2.2 

.0093 3.4

.0014 7.8

.0008 15.7

.0010 11.2

.0005 33.4

.0011 26.4

29.5 

73.2

82.7

84.7

88.3

90.7

92.1

91.0

90.7

69.9

0.

.

.

2.

2.

7.

5.

4.

3.

3.

4.

10.

7.

Ma;

9. 

22.

22.

16.

12.

3.

Ma;

2. 

9.

7.

6.

8.

13.

16.

12.

12.

4.

, MUSCOVITE 

87.4+0.4 Ma;

33.8 

74.2

94.8

96.9

96.2

97.9

95.9

4.

6.

15.

30.

21.

65.

51.

0414 

0485

295

763

00

82

09

17

13

85

67

82

40

94

wt.,

20 

7

21

18

17

14

10

12

21

28

27

18

30

51

0.0785

.9 

.63

.0

.3

.5

.2

.6

.5

.7

.1

.3

.7

.1

.6

J=0. 008244; wt. ,

87 

3

5

1

7

99

303

J=0.

05 

33

79

75

73

3

3

9

0

49

J=0.

2V 

53

1

6

9

0

5

219

 

5400

312

26 .4

006390; wt . ,

0 .22

.11

.11

.11

.11

.10

.10

.10

.10

.10

006400; wt. ,

5

91

396

817

252

360

66

.91 

.5

.1

g

70

151

111

54

55

57

59

61

62

63

63

65

67

67

.4+3

.1 + 5

.0+0

.5+0

.2 + 0

.1+0

.4 + 0

.2 + 0

.0+0

.4 + 0

.3 + 0

.0+0

.0+0

.0 + 0

0.1946

68 

70

71

70

71

71

61

.4+0 

.7 + 0

.0+0

.9+0

.7 + 0

.1 + 0

.7 + 0

0.5054

101 

86

86

86

86

86

85

86

86

86

.5+0 

.6+0

.4 + 0

.3+0

.6+0

.1 + 0

.9+0

.7+0

.1+0

.0+0

0.1094

86 

89

88

87

86

87

87

.5+0 

.0 + 0

.0+0

,4 + 0

.8 + 0

.2 + 0

.6+0

.8 

.5

.7

.4

.3

.3

.3

.3

.3

.4

.3

.4

.4

.3

g

.3 

.4

.4

.4

.4

.4

.9

g

.6

.5

.4

.5

.4

.5

.5

.4

.4

.5

g

.5 

.4

.4

.4

.5

.4

.4

collected from an area extending from Slate Creek in 

the north to Round Valley in the south, and from 7 

mi west of Riggins on Papoose Creek to the easternmost exposures of the Riggins Group or correlative

rocks on the Salmon River (fig. 10.6). Samples of 

hornblende, muscovite, and biotite from all metamorphic grades were dated. The argon age spectra are 

simple for some of these samples and complex for
388 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 10.3. 40Ar/ 39Ar age-spectrum data for samples from west-central Idaho and eastern Oregon Continued

Temper- 39Ar 

ature 40Ar/ 39Ar 37 Ar/ 39Ar 36Ar/ 39Ar (pet of 

(°C) total)

Total-gas

550 

650

800

900

1,000

1,125

Fuse

10. 

8.

8.

8.

7.

7.

7.

06 

736

207

100

989

879

930

Total-gas

550 

650

800

900

1,000

Fuse

16. 

12.

9.

9.

9.

9.

33

941

970

392

121

Total-gas

550 

650

900

950

1,000

1,125

Fuse

23. 

9.

8.

8.

8.

8.

8.

57 

757

886

618

566

414

567

date: 87.6 Ma

0.1233 0. 

.0079

.0037

.0014

.0010

.0004

.0012

date: 72.5 Ma

0.2556 0. 

.0968

.0233

.0107

.0139

.0115

date: 70.5 Ma

0.0338 0. 

.0032

.0009

.0007

.0018

.0020

.0069

Total-gas date :

550 

750

850

950

1,050

1,150

Fuse

10. 

10.

12.

12.

10.

8.

8.

24 

03

34

04

92

662

810

0.0503 0. 

.0159

.0209

.0240

.0271

.0256

.0561

Total-gas date:

600 

700

800

900

1,000

1,050

1,100

1,200

1,300

Fuse

9. 

7.

7.

7.

7.

7.

7.

8.

8.

8.

132 

801

469

304

374

578

791

005

511

741

0.0038 0. 

.0030

.0023

.0026

.0025

.0028

.0043

.0054

.0049

.0017

TC-S-124-23 

; plateau date :

0083 

0026

0012

0011

0008

0003

0004

3.6

3.1

4.8

13.9

24.3

25.0

25.3

pet 

4 °ArR

39Ar 

(mol 

xlO'13 )

, MUSCOVITE 

87.6+0.4 Ma; J=0 .

75.8

91.1

95.6

96.1

97.0

98.6

98.3

D3113M, MUSCOVITE 

; plateau date: 72.4+0.4

0237 

0121

0044

0045

0026

0017

1.5 

1.6

3.9

12.2

20.6

60.2

B12178201, 

; plateau date:

0539 

0055

0023

0018

0017

0010

0011

68.7

0098 

0042

0060

0042

0059

0030

0036

60.5

0085 

0031

0019

0012

0009

0012

0017

0022

0033

0028

1.3 

7.7

29.2

19.8

12.5

19.7

10.0

56.4 

70.9

86.9

86.6

91.7

94.6

7.64 

6.75

10.4

29.9

52.2

53.6

54.2

Ma; J=0.

2.07 

2.21

5.38

16.6

28.1

82.1

MUSCOVITE 

70.1+0.4 Ma; J=0 .

32.4 

83.1

92.3

93.6

94.0

96.6

96.3

B12198202, BIOTITE 

Ma; no plateau; J=0

5.9 

27.4

10.1

11.2

6.7

11.4

27.3

71.7 

87.7

85.6

89.8

83.9

89.6

88.0

1.66 

10.1

38.5

26.1

16.5

26.0

13.2

.004445;

4.44 

20.6

7.58

8.41

5.03

8.57

20.5

B12198202, MI CROC LINE 

Ma; no plateau; J=0. 004695;

6.0 

7.3

5.1

8.5

12.3

10.3

10.5

12.7

13.0

14.4

72.5 

88.1

92.3

94.9

96.4

95.3

93.5

91.9

88.4

90.3

5.89 

7.14

5.02

8.35

12.1

10.1

10.4

12.5

12.8

14.2

Apparent Apparent 

K/Ca age and error 

(mol /mol) (Ma)

006400; wt. ,

4 

65

139

367

519

1480

436

.22

.6

004750; wt . ,

2 

5

22

48

37

45

.04 

.37

.4

.5

.4

.3

004897; wt. ,

15 

160

592

695

294

264

75

wt.,

10 

32

24

21

19

20

9

wt.,

138 

175

225

197

208

186

121

95

106

300

.4

.4

0.1012

.3 

.7

.9

.7

.2

.3

.27

0.1042

.4

0.1194

86.0+0 

89.6+0

88.4+0

87.7+0

87.3+0

87.6+0

87.8+0

0.1019

75.9+0 

73.4+0

72.5+0

72.5+0

72.3+0

72.4+0

0.1012

66.3+0 

70.3+0

71.0+0

69.9+0

69.8+0

70.4+0

71.4+0

g

57.9+0 

69.2+0

82.8+0

84.6+0

72.0+0

61.2+0

61.1+0

g

55.2+0 

57.3+0

57.5+0

57.8+0

5-9.2 + 0

60.2+0

60.7+0

61.3+0

62.7+0

65.7+0

g

.4 

.4

.4

.4

.4

.4

.5

g

.5 

.4

.4

.4

.4

.4

g

.4 

.4

.4

.3

.3

.4

.4

.3 

.4

.4

.4

.4

.3

.3

.3 

.3

.3

.3

.3

.3

.3

.3

.3

.4
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 389

others; many exhibit argon loss. A brief summary is 

presented in table 10.2 under the heading "Rocks 

Near the Suture Zone."

Even though 4OAr/39Ar age spectra of metamorphic 

minerals from the Riggins Group and Seven Devils 

island arc are complex (figs. 10.7, 10.10, 10.14, 10.19, 

and 10.25), some general observations are evident. Age 

spectra of 7 of the 12 dated hornblendes exhibit plateaus 

or near-plateaus (figs. 10.7, 10.10, and 10.17); five of 

these display minor argon loss, indicated by progressively increasing dates from left to right on the agespectrum diagrams (figs. 10.7, 10.10, and 10.17). Two 

of the five age spectra that do not exhibit plateaus

116°35' 116°25'

EXPLANATION

Surficial deposits 

(Quaternary)

|ib | Basalt (Tertiary)

Coon Hollow Formation 

(Jurassic)

"> Contact

- Fault Dashed where 

inferred

Doyle'Creek Formation (Triassic) 

Kurry unit

Wild Sheep Creek Formation 

(Triassic)

Cougar Creek Complex of 

Vallier (1968) (Permian)

Hkc-35

Thrust fault Sawteeth 

on upper plate

Sample locality

FIGURE 10.3. Sample locality map for Pittsburg Landing area 

on Snake River. Geologic base from Vallier (1974 and written 

commun., 1987) and White and Vallier (1994).

display argon loss (figs. 10.10 and 10.17). Hornblende 

dates form at least three statistically distinct groups 

at about 118, 109, and 101 Ma (fig. 10.25); a younger 

93-Ma thermal event is documented by the 40Ar/39Ar 

dates of hornblende from suture-zone plutons (see 

below). These groups of dates show a clear geographical

116°52' 116°48'

EXPLANATION 

Surficial deposits (Quaternary)

Tb Basalt (Tertiary)

^^ Oxbow Amphibolite of Vallier (1974) (Triassic) 

Oxbow Quartz Diorite of Vallier (1974) (Triassic) 

Hunsaker Creek Formation (Permian) 

Windy Ridge Formation (Permian?) 

Oxbow Shear Zone of Vallier (1974)

'\ Contact

 -* Fault 

Hcop-39 Sample locality

FIGURE 10.4. Sample locality map for area of The Oxbow on 

Snake River. Geologic base modified from Vallier (1974).
390 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

distribution: the 118-Ma group is located south and 

west of Riggins; the 109-Ma group forms an apparent 

belt to the north, east, and south of the 118-Ma group; 

and the 101-Ma group is located farther east along the 

west side of the suture zone. This geographic distribution is schematically shown on figure 10.6. The groups 

of dates also exhibit a relation, albeit less clearly defined, to the metamorphic grade described by Hamilton 

(1963), Onasch (1977,1987), and Lund (1984). The 118- 

Ma group occurs not only in greenschist-facies, garnetzone rocks but also in upper-amphibolite-facies rocks 

in this region; the 109-Ma group occurs in amphibolitefacies, oligoclase- and andesine-zone rocks; and the 101- 

Ma group occurs in upper-amphibolite-facies, sillimanite-zone rocks. Although argon loss from hornblende is 

present in all age groups, it is found only in rocks of 

oligoclase-zone or higher grade metamorphism. Argon 

loss is most severe where textural evidence shows that 

the rock contains hornblende of two different ages or 

has undergone multiple deformations. The dates of 

muscovites and biotites are generally younger than the

116°50' 116°40'

EXPLANATION 

Basalt (Tertiary)

Intrusive rocks (Jurassic and Late Triassic) 

mill Wild Sheep Creek Formation (Triassic) 

Shear zone 

Contact 

Fault

Hem-23

Sample of plutonic rock 

Sample of shear-zone rock

FIGURE 10.5. Sample locality map for area at confluence of Imnaha and Snake Rivers. Geologic base modified from Vallier 

(1974).

dates of hornblendes from any particular group. One 

biotite preserves a cooling age that is statistically identical to the date of a coexisting hornblende that records 

the oldest metamorphic event. Two dated muscovites 

from biotite-zone rocks exhibit argon loss. Three biotites from andesine-zone rocks display significantly 

younger dates than coexisting hornblendes.

An important conclusion from these data is that the 

oldest metamorphism associated with the Salmon 

River suture zone in the Riggins-Slate Creek area occurred at 116 to 118 Ma. One sample (R7), which yielded this date (116.5±0.6 Ma), is a garnet-zone rock 

containing hornblende crystals that define a single foliation but no lineation (figs. 10.7, 10.8). The mineral 

assemblage indicates that the hornblende formed 

within the lower limits of the stability field of hornblende and probably at a temperature of 500°C or

EXPLANATION 

Volcanic rocks and sedimentary deposits (Cenozoic)

Undeformed plutons (Cretaceous) 

Deformed plutons (Cretaceous) 

Seven Devils Group (Triassic and Permian) 

Riggins Group (Mesozoic or Paleozoic) 

Basement (?) rocks (Paleozoic) 

Umbrella Butte unit (Paleozoic?)

OzV

Kup

Kdp

IPsd

Mzftr

Rzb

ftub

Postmetamorphic or postplutonic hornblende cooling age: 

Less than 88 Ma

88-93 Ma 

101 Ma 

109 Ma

118 Ma

244 or 226 Ma (Pre-suture 

zone-related events)

No metamorphic overprint 

related to suture zone

Cenozoic rocks

Contact

Fault U, upthrown side; D, downthrown side

Thrust fault Sawteeth on upper plate

- Initial ^Sr/^Sr discontinuity

 Sample of metamorphic rock 

A Sample of plutonic rock or dike

FIGURE 10.6. Explanation.
10 TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 391

FIGURE 10.6. Sample locality map for Riggins, Idaho, area. Age-distribution pattern is based on hornblende 40Ar/39Ar dates and is highly generalized. Geology is modified from Lund and Snee (1988). See 

also figure 10.2, which covers part of the same area. Geology of unpatterned area is not compiled.
392 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

lower (upper greenschist fades; Winkler, 1979). Because the argon closure-temperature of hornblende is 

over 500°C even at slow cooling rates, this hornblende 

probably closed to diffusion of argon upon formation, 

that is, this age does not reflect any significant period 

of cooling. The coexisting biotite of this sample (R7) 

yielded a plateau date of 115.8±0.3 Ma, indicating that 

this rock cooled very quickly after metamorphism to a 

temperature lower than 280°C (biotite argon-retention 

temperature). Hornblende (R30) from amphibolite-facies rocks also has a single hornblende-defined foliation (figs. 10.7, 10.9); this sample, which probably 

formed at about 600°C (above the hornblende argonclosure temperature), also gives an age of 118 Ma. 

Therefore, it is unlikely that an extended period of 

cooling could have occurred after metamorphism in 

either upper-amphibolite-facies or greenschist-facies 

rocks of this age. In fact, sample R30 preserves evidence for a second period of metamorphism that 

caused the formation of symplectic epidote and hornblende at the expense of grossular garnet. Thermodynamically, this reaction indicates an increase in 

pressure from the conditions of the first metamorphism when the garnet formed (A.B. Till, U.S. Geological Survey, written commun., 1989); however, the 

temperature of this reaction is not well constrained. 

The argon spectrum for this sample displays argon 

loss, which we interpret to have occurred during the 

second, higher pressure metamorphic episode. If our

co 130 

IT

55

LL

°120

C/D

110

100

90

50 100

39ArK RELEASED, IN PERCENT

FIGURE 10.7. Composite 40Ar/39Ar age-spectrum diagram for 

118-Ma metamorphic rocks from Riggins Group. 1, Hornblende from sample R7; Tmp (preferred date)=116.5±0.6 Ma 

(preferred date defined by 48 percent of total 39ArK released).

2. Biotite sample from R7; Tp (plateau date)=116.1+0.6 Ma 

(plateau date denned by 87 percent of total 39ArK released).

3. Hornblende sample from R30; Tp=118.0±0.5 Ma (plateau 

date defined by 94 percent of total 39ArK released; minor argon loss to 114 Ma). See text for discussion of plateau date, 

preferred date, and argon loss.

interpretation is correct, the temperature of this second event must have been close to but lower than 

about 530°C, the argon closure-temperature of hornblende, because the argon loss was minor. This conclusion is consistent with geobarometry data (Selverstone 

and others, 1992) that indicate that the metamorphic 

rocks of the area of sample R30 and southward underwent an increase in pressure after the higher temperature metamorphic event. We suggest that this 

increase in pressure occurred after 118 Ma and possibly is our 109-Ma metamorphic event (see below).

2 MILLIMETERS

J

FIGURE 10.8. Photomicrograph of sample R7: non-lineated, garnet-zone amphibolite from the Riggins Group. Hornblende 

(darker prismatic mineral among lighter colored feldspar and 

quartz) is randomly to weakly oriented, but rock has moderately 

defined foliation in hand specimen. Garnet (dark, poorly defined 

grain in left center of photomicrograph) is uncommon; biotite is 

common in other layers. Plane light.
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 393

Formerly, the oldest documented metamorphic 

event associated with tectonic activity along the 

Salmon River suture zone was the 118-Ma metamorphism described above. But recently Davidson and 

Snee (1989, 1990), Davidson (1990), and Snee and 

others (in press) have demonstrated that regional dynamothermal metamorphism along the Salmon River 

suture zone in the Orofino, Idaho, area began at 130 

Ma. This period of metamorphism is preserved in

2 MILLIMETERS

I

FIGURE 10.9. Photomicrograph of sample R30: foliated, banded, 

non-lineated amphibolite from the Riggins Group. Band in upper half of photomicrograph is unoriented hornblende (darker 

mineral) and quartz and feldspar (lighter minerals). Band in 

lower half consists of unoriented symplectic epidote, symplectic 

hornblende, and altered grossular garnet (darker minerals) and 

quartz, feldspar, and calcite (lighter minerals). Mineral assemblage in lower band formed during a second, higher pressure 

metamorphic event than that forming the assemblage in the upper band. Plane light.

40Ar/39Ar age spectra for hornblendes from uppergreenschist-facies metamorphic rocks adjacent to the 

suture zone. If metamorphism as old as 130 Ma affected rocks in the Riggins-Slate Creek area, no 

argon isotopic record is preserved. It is possible that 

metamorphism associated with the suture zone was 

diachronous. It is equally possible, however, that 130- 

Ma metamorphic rocks in the Riggins-Slate Creek 

area were tectonically removed by right-slip movement along the suture or were structurally buried 

and(or) thermally overprinted during overthrusting 

of higher grade, younger metamorphic rocks.

The 40Ar/39Ar age-spectrum data for metamorphic 

rocks from the Riggins-Slate Creek area reveal a second period of hornblende growth and argon closure 

that occurred at about 109 Ma. Samples that best 

show this are two postkinematic hornblendes (samples 

R27 and R34; figs. 10.10-10.12); these hornblendes 

formed after deformation that produced strong foliation in garnet-zone rocks. No argon loss is exhibited by 

either spectrum, and sample R34 has a well-defined 

plateau. Like the 118-Ma garnet-zone sample, these 

younger garnet-zone hornblendes probably formed at a 

temperature below that for hornblende argon closure. 

Thus, the dates closely approximate the age of static 

hornblende formation. That this thermal activity was 

accompanied in places by deformation is supported by 

data from five other synkinematic hornblendes from 

middle- to upper-amphibolite-facies rocks (samples 

R17, R18, R26, R28, and R29; figs. 10.10 and 10.13),

co 130

DC

0 120

110

100

90

50 100

39ArK RELEASED, IN PERCENT

FIGURE 10.10. Composite 40Ar/39Ar age-spectrum diagram for 

109-Ma metamorphic rocks from Riggins Group. 1, Hornblende from sample R34; Tp (plateau age)=109.1±0.5 Ma (plateau defined by 83 percent of 39ArK released). 2, Hornblende 

from sample R29; Tp=108.1±0.7 Ma (plateau defined by 98 

percent of 39ArK released). 3, Hornblende from sample R26; 

argon loss; Tmax (maximum date)=109.0±0.6 Ma; Tmin (minimum date)=93.7±0.5 Ma. Omitted from this diagram for clarity but of the same group are hornblende samples R17, R18, 

R27, and R28.
394 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

which show variable amounts of argon loss superposed 

on a maximum age of about 109 Ma. Argon loss recorded by these five hornblende samples probably resulted 

from one or more overprinting events.

Sillimanite-zone samples (R14, R16, and R23) 

which clearly show two foliations, a lineation-dominated fabric, or two different episodes of hornblende

o 2 MILLIMETERS

J

FIGURE 10.11. Photomicrograph of sample R34: garnet-zone 

gneiss from Riggins Group. Foliation is defined by oriented biotite (uncommon planar minerals) and layers rich in epidote 

(small, clear, high-relief minerals) alternating with layers rich 

in quartz and feldspar. Hornblende (large, dark crystals) is unoriented and postkinematic. Plane light.

2 MILLIMETERS

FIGURE 10.12. Photomicrograph of sample R27: garnet-zone 

gneiss from Riggins Group. Unoriented, postkinematic amphibole (dark, prismatic crystals) formed over compositional banding that defines an earlier foliation. Plane light.

growth (figs. 10.14-10.16) have age spectra with 

minor argon loss from high-temperature ages of 

99.5-101 Ma to low-temperature ages of 89-93 Ma 

(fig. 10.17). Although it is difficult to confidently interpret the high-temperature ages of these three 

samples, these data indicate another pulse of thermal activity at about 101 Ma. Argon loss exhibited by 

the low-temperature ages of these three samples 

probably resulted during emplacement and deformation of the suture-zone plutons.

Argon age-spectrum data also constrain the timing 

of deformation. Primary foliation, which in places 

partially transposes bedding (fig. 10.18), is defined 

(in the appropriate rock types) by 118-Ma hornblende

2 MILLIMETERS

FIGURE 10.13. Photomicrograph of sample R29: amphibolite from 

Riggins Group. Unoriented amphibole (dark minerals) and 

quartz and feldspar (light minerals) in a strongly banded gneiss. 

Earlier, deformed garnet (large grain in lower right) is partly 

altered to chlorite. Plane light.
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 395

that is associated with the onset of metamorphism. 

Thus, structures that deform metamorphic fabric in 

these rocks are all 118 Ma or younger. Although we 

feel that the Rapid River thrust fault is only one of 

several important postmetamorphic fault systems 

within the Riggins Group and Seven Devils island 

arc, it is a well-known feature for which an age can 

be derived from the argon data. Because this fault 

system cuts 118- and 109-Ma metamorphic rocks that 

have short cooling histories, it must have formed 

after 109 Ma. (As discussed below, it must be older 

than 93 Ma because it is intruded by 93-Ma plutons.) 

This conclusion is corroborated by the age spectrum 

for muscovite sample R3, which displays argon loss 

from 105 to 100 Ma. Sample R3 was collected from a 

pegmatite that intruded Seven Devils island-arc 

rocks near the Rapid River thrust fault on Papoose 

Creek (fig. 10.6) and was later deformed by boudinage. Both the pegmatite's emplacement and its later 

deformation may have been related to activity on the 

Rapid River thrust fault.

Within the island-arc metamorphic rocks of the Riggins Group, small bodies of deformed to undeformed 

tonalite and leucotonalite are exposed in places, and 

pegmatites are also present in both the Riggins 

Group and Seven Devils island-arc rocks. Whether 

these bodies are pre- or synmetamorphic has been an 

important question. Hornblende from a weakly foliated tonalite that intruded the Riggins(?) Group near 

Round Valley (sample 7-30-5) was dated (fig. 10.6)

2 MILLIMETERS

J

FIGURE 10.14. Photomicrograph of sample R16: anorthite-zone 

amphibolite from Riggins Group. Foliation is well defined by 

younger generation of hornblende, which can be seen here wrapping around older hornblende microaugen. Fractured, postkinematic garnets (gray mineral in upper center) overgrew 

hornblende foliation. Plane light.

and yielded a plateau date of 112.0±0.7 Ma. This 

date, as well as the sample R3 muscovite date from 

pegmatite on Papoose Creek, provides evidence that 

pluton emplacement, albeit apparently minor in the 

Riggins area, occurred during the period of metamorphism and deformation, that is, 101-118 Ma. Additional data from other locations within the suture 

zone, such as Orofino, Idaho (fig. 10.1; Snee and others, 1987a; Davidson and Snee, 1989; Davidson, 1990; 

Snee and others, in press), suggest that magmatic activity occurred after both the 118-Ma and the 109-Ma 

metamorphic events and that later deformation commonly did not entirely reset the mineral argon systems. Preliminary results of U-Pb zircon dating (C.A. 

Manduca, California Institute of Technology, unpub.

2 MILLIMETERS

FIGURE 10.15. Photomicrograph of sample R14: anorthite-zone 

amphibolite from Riggins Group. Foliation is well defined by 

bands rich in hornblende (dark minerals) and bands rich in 

quartz and feldspar (light minerals). Older hornblende microaugen (for example, large grain in center) are wrapped by 

the compositional bands. Plane light.
396 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

data, 1990; Snee and others, in press) indicate that 

some deformed plutons with island-arc chemical affinities located west of the suture zone were emplaced at 

115 Ma. Current work is attempting to determine the 

extent of these plutons.

DEFORMED PLUTONS OF THE SALMON RIVER SUTURE ZONE

Determination of the age of tonalitic plutons within 

the suture zone by the 4OAr/39Ar age-spectrum technique is difficult because minerals may remain open 

to argon diffusion during postcrystallization cooling

2 MILLIMETERS

J

FIGURE 10.16. Photomicrograph of sample R23: amphibolite roof 

pendant in foliated tonalite of suture zone. Older, coarse-grained 

hornblende (see right half of photo) has been partly replaced by 

finer grained hornblende (see left half of photo) that forms lineation perpendicular to plane of photo. Plane light.

and deformation. (Age spectra for suture-zone plutons are summarized in table 10.2 under "Suturezone Rocks.") Because the emplacement temperature 

of tonalite is about 850°C (Naney, 1983) and because 

the argon-closure temperature for hornblende is 

about 530°C, hornblende dates should provide a minimum estimate that, unless the plutons had an extended early cooling history, should be close to the 

age of crystallization.

The 4OAr/39Ar age spectrum for a sample of hornblende taken from a metasedimentary xenolith (sample R23; figs. 10.16, 10.19) in a suture-zone pluton 

provides important constraints on the emplacement 

history of the enclosing pluton. This age spectrum 

shows argon loss superposed on an earlier cooling 

age. The cooling age indicates that the metamorphic 

rock from which the hornblende was taken had 

cooled below the hornblende argon-closure temperature at or before 99.5 Ma; thus, this hornblende 

formed during a metamorphic event that preceded 

emplacement of the enveloping tonalite. Argon loss at 

about 91.5 Ma, which is preserved in the spectrum, 

probably resulted from emplacement of the tonalite. 

Because this hornblende was not completely reset 

during pluton emplacement, it seems likely that the 

heat of the pluton was lost quickly. Quick loss of heat 

suggests a short cooling history and implies a high 

crustal emplacement level for this tonalite at some 

time after 99.5 Ma. Therefore, the date of 91.5 Ma,

en 130

DC

LL

0 120 

CO

110

ArK RELEASED, IN PERCENT

FIGURE 10.17. Composite 40Ar/39Ar age-spectrum diagram for 

101-Ma metamorphic rocks from Riggins Group. 1, Hornblende from sample R14; complex age spectrum showing 

argon loss from Tmax (maximum date)=100.7±0.6 Ma to Tmin 

(minimum date)=88.5±0.6 Ma. 2, Hornblende from sample 

R16; complex age spectrum showing argon loss from 

Tmax=100.6±0.5 Ma to Tmin=92.8±0.6 Ma. 3, Hornblende from 

sample R23; argon-loss spectrum showing consecutively 

increasing date, with Tmax=99.4±0.5 Ma and Tmin=91.5±1.2 

Ma. Tmjn is chosen according to apparent K/Ca ratio that 

is consistent with higher temperature steps.
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 397

an estimate of the time of the argon loss from the 

metamorphic hornblende, may be close to the age of 

emplacement of the suture-zone tonalite.

The age spectrum of hornblende sample R12 (fig. 

10.19) from suture-zone tonalite supports the interpretation that plutons were emplaced within the suture zone at or before about 91.5 Ma. Hornblende in 

this sample, which contains no epidote, was deformed 

by a strong postemplacement event that produced a 

steep southeast-plunging lineation defined by the deformed hornblende crystals. Argon loss exhibited by 

the age spectrum apparently occurred at 88 Ma, probably during deformation; the date of 93.5 Ma for the

FIGURE 10.18. Photomicrograph of sample R6: garnet-zone 

gneiss from Riggins Group, collected near sample R7. Contact 

between compositional layers extends from upper right to lower left corner of photo. Incipient transposition of compositional 

layering to an orientation extending from top to bottom is evident in hornblende and biotite crystals (dark minerals) and 

iron oxide minerals (opaque grains) along the contact. Plane 

light.

highest temperature extraction step is a minimum estimate of the age of emplacement of the tonalite. The 

virtually identical data of hornblende samples R24 

and R22 (fig. 10.19), which were collected from the 

suture zone more than 25 km north of the location of 

sample R12, lend additional support to the interpretation that at least some tonalites within the suture 

zone were emplaced at a relatively shallow level at or 

before 93.5 Ma and that a major episode of deformation caused argon loss at about 88 Ma.

The ages of emplacement and deformation of the 

suture-zone plutons provide additional constraints on 

the age of the Rapid River thrust fault. In the Slate 

Creek area, as mentioned above, the Rapid River 

thrust fault cuts 118-Ma and 109-Ma metamorphic 

rocks that had short cooling histories; therefore the 

thrust fault is younger than 109 Ma. The Rapid 

River thrust fault in the Slate Creek area is folded 

and cut by later thrust faults (Lund, 1984), and all of 

the thrust faults are intruded by the 93.5-Ma or 

older suture-zone plutons. Thus the original Rapid 

River thrust-faulting event, plus the later folding and 

rethrusting, all occurred between 109 and 93.5 Ma. 

Finally, after the emplacement of the suture-zone 

plutons, high-angle deformation occurred at about 88

OF YEARS

1 i

ILLIONS

i

2 90

z

ul 3

I- 80

HI 

DC

<

a. 

a.

< 70

i i i i i i i i i

1h

, ,___.__.. ____..._: , yh , ,t__^ 

-if--, ,-,.__ ^ r~ --

s 4h

! 5b > ^ ^___ j < r~r~~-"^^-~7^-^«_^r3

 i i i i i i ii

50 100

39ArK RELEASED, IN PERCENT

FIGURE 10.19. Composite 40Ar/39Ar age-spectrum diagram for 

93-Ma plutonic rocks from suture zone in Riggins area. Ih, 

Hornblende from sample R23; same spectrum as shown in 

figure 10.17. 2h, Hornblende from sample R24; disturbed 

spectrum; Tmp (preferred date)=92.6±0.5 Ma. 3h, Hornblende 

from sample R22; Tp (plateau date)=89.3±0.6 Ma (plateau 

date defined by 62 percent of 39ArK released). 4h, Hornblende 

from sample R21; Tp=83.9±0.4 Ma (plateau date defined by 

96 percent of 39ArK released). 5b, Biotite from sample R21; 

Tp=81.110.4 Ma (plateau date defined by 69 percent of 39ArK 

released). 6m, Microcline from sample R21; disturbed spectrum; Tmax (maximum date)=78.6±0.4 Ma. 7m, Microcline 

from sample R24; argon loss from Tmax=79.0±0.4 Ma to Tmin 

(minimum date)=69.5±0.4 Ma. 8h, Hornblende from sample 

R12; argon loss from Tmax= 93.4±0.5 Ma to Tmin=88.3±0.5 Ma. 

9b, Biotite from sample R12; T =82.6±0.4 Ma.
398 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Ma. This 88-Ma fabric is most strongly developed in 

the suture-zone plutons, but this high-angle fabric 

also cuts off the 118- and 109-Ma metamorphic fabric 

near the suture zone.

Regionally, the younger, steeply plunging, lineationdominated fabric is well-developed in a narrow belt of 

highly deformed rocks straddling the suture zone and 

in many places completely transposes any older 

fabrics. For example, sample Rll, collected near the 

location of sample R12, is, like sample R12, from deformed plutonic rock on the Salmon River (fig. 10.6) 

and on structural trend with samples R24 and R22 

from the Slate Creek area to the north. A plateau date 

of 85.1±0.4 Ma for deformed hornblende from this 

sample is evidence that deformation or cooling affected this area after 88 Ma. Furthermore, hornblende 

plateau dates from similarly deformed tonalites on 

structural trend about 50 km south of the location of 

sample Rll range from 81 to 79 Ma (L.W. Snee, unpub. data, 1990); for similarly deformed rocks in the 

McCall area, hornblende plateau dates range from 79 

to 88 Ma (L.W. Snee and M.A. Kuntz, U.S. Geological 

Survey, unpub. data, 1990). For samples of similarly 

deformed plutonic rocks on structural trend in the suture zone to the north near Orofino (fig. 10.1), plateau 

dates for hornblendes range from 83 to 81 Ma (Snee 

and others, 1987a; Davidson and Snee, 1989; Davidson, 1990); U-Pb zircon dates indicate that some of 

these plutons were emplaced at 115 Ma (Snee and 

others, in press). Furthermore, long periods of cooling 

after deformation along parts of the suture zone, such 

as at Orofino, are supported by the coexistence of 74- 

Ma biotite with 83- to 81-Ma hornblende (Snee and 

others, 198 7a; Davidson and Snee, 1989; Davidson, 

1990; Snee and others, in press) and 115-Ma zircon. In 

summary, cooling ages for hornblendes from this fabric throughout the region range from 88 to 79 Ma and 

indicate the regional extent of the vertical movement 

that affected a narrow belt of rocks near the suture 

zone. The complex and large-magnitude character of 

deformation in this narrow belt is confirmed by the 

juxtaposition of epidote-free (that is, shallow) and epidote-bearing (that is, deep-seated) plutons as well as 

plutons of different ages within this belt.

UNDEFORMED PLUTONS OF THE IDAHO BATHOLITH

40Ar/39Ar age-spectrum data for undeformed plutons east of the suture zone are summarized in table 

10.2 under the heading "Undeformed Plutons and Associated Rocks." Sample R21, tonalite that contains 

magmatic epidote (fig. 10.6), is from the oldest undeformed phase of the Idaho batholith. Hornblende

geobarometry (Hammarstrom and Zen, 1988), as well 

as the presence of magmatic epidote, indicates that 

this tonalite was emplaced at a depth of about 20 

km. The 40Ar/39Ar age spectra for hornblende, biotite, and microcline define the cooling history of this 

sample from about 530°C at 84 Ma, to 280°C at 81 

Ma, to 130°C at 78 Ma (fig. 10.19). The hornblende 

plateau date of 83.9±0.4 Ma is a minimum estimate 

for the age of emplacement. U-Th-Pb isotopic analyses of five zircon fractions and one sphene fraction 

from sample R21 yield a lower-intercept date of 90±3 

Ma (mean square of weighted deviates=7.4; fig. 

10.20; table 10.4). The lower-intercept date is interpreted to be the best estimate for age of emplacement of the tonalite pluton. Using U-Pb zircon data, 

C.A. Manduca (California Institute of Technology, unpub. data, 1992) has determined an identical emplacement age for similar undeformed tonalite in the 

McCall area, east of the Salmon River suture zone.

On the east side of the Idaho batholith, metaluminous granodiorite and tonalite were emplaced earlier 

than the extensive muscovite-biotite granites that 

are typical of the central part of the batholith (Hall, 

1985; Kiilsgaard and Lewis, 1985; Lewis and others, 

1987; fig. 10.21). Sample MP201 of hornblende-biotite 

porphyritic granodiorite from 15 km east of Stanley, 

Idaho, on the Salmon River yielded a hornblende plateau date of 86.3±0.5 Ma (fig. 10.22; table 10.3). (An 

unpublished U-Pb zircon date from the same outcrop 

of 88±6 Ma by Lynn Fischer, U.S. Geological Survey,

0.034

0.026

0.018

0.010

' 100-150nm

^ -3^00-250 nm 

 "" 250-325 nm 

. *325-400nm 

>400nm

'sphene

0.08 0.12 0.16 0.20 

207Pb/ 235U RATIO

0.28

FIGURE 10.20. Concordia diagram for sample R21, from undeformed pluton east of Salmon River suture zone. Solid line 

(concordia) is time curve that is the locus of all concordant 

U-Pb dates. Dashed line (discordia) is defined by U-Pb dates 

on minerals from sample R21. Five discordant dates are 

from five size fractions of zircon; one concordant date is from 

sphene. Error envelopes for each date are defined by symbols. Discordia intercepts concordia at 90.4±3 Ma (see diagram) and at 1,932+150 Ma (not shown). Best estimate of 

crystallization age is lower intercept date, 90.4+3 Ma. Mean 

square of weighted deviates (MSWD)=7.4.
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 399

TABLE 10.4. U, Th, and Pb analytical data and dates for sample R21

[Size-fraction ranges listed for zircon; sphene separate was not sized]

Size

fraction

(nm)

100-150

200-250

250-325

325-400

>400

Sphene

Concentration

U

413

412

408

402

416

120

(ppm)

Th

145

154

165

179

195

166

Pb

10.4

8.7

8.3

7.6

10.7

3.8

Measured

ratio

206pb

204pb

2,146

1,932

4,462

1,786

155

79.5

Isotopic composition of lead

(atom percent)

204pb

0.0312

.0275

.0110

.0359

.4164

.5688

206pb

82.17

82.79

83.32

81.79

65.32

49.29

207pb

6.740

6.295

5.793

5.716

10.020

10.744

208pb

11.06

10.89

10.87

12.46

24.24

39.40

2oepb

238U

152

130

125

114

110

91

Dates

(Ma)

207pb

235u

228

185

170

147

138

91

207pb

206pb

1,111

965

857

726

641

99

208pb

232Th

160

126

118

107

102

91

1985, and cited in Lewis and others, 1987, is statistically identical to the hornblende date.) Approximately 25 km farther east, the Thompson Creek porphyry 

molybdenum deposit formed during emplacement of a 

stock composed of biotite granodiorite and biotite 

granite of Cretaceous age (Hall and others, 1984). 

Two muscovite samples (TC-S-41-16 and TC-S-124- 

23) from a drill core of alteration assemblages associated with the molybdenite have plateau dates of 

87.4±0.4 and 87.6±0.4 Ma, respectively (fig. 10.22; 

table 10.3). These muscovite dates indicate that the 

minimum age of the associated plutons is about 87.5 

Ma. Numerous additional hornblende 4OAr/39Ar cooling ages for metaluminous plutons in the southern 

part of the Idaho batholith range from 93 to 88 Ma 

(L.W. Snee, unpub. data, 1991). Thus, on the basis of 

all the ages for metaluminous plutons of the Idaho 

batholith, an area extending from the Salmon River 

suture zone up to 139 km to the east was affected by 

this period of metaluminous magmatism.

Sample R8 is a representative sample of the voluminous muscovite-biotite granites of the Idaho batholith from a few miles east of the suture zone (fig. 

10.6). Muscovite and biotite plateau dates of 75.3±0.4 

and 74.7±0.4 Ma, respectively (tables 10.2, 10.3), are 

statistically indistinguishable and mark the time 

when this pluton cooled through 325°C and 280°C, 

the argon-closure temperatures of muscovite and biotite, respectively. Whether approximately 75 Ma is 

the age of emplacement of the muscovite-biotite granite depends upon how rapidly the granite cooled from 

its emplacement temperature of about 650-750°C to 

below 300°C. This cooling rate is controlled in part 

by the temperature of the country rocks at the time 

of emplacement of the granite and the depth of emplacement. The argon age spectrum of the microcline 

from sample R21, a sample from one of the metaluminous plutons discussed above, indicates that the 

country rocks, including the metaluminous plutons,

of the muscovite-biotite granite had cooled below 

about 130°C by 78 Ma. Using this temperature as a 

maximum for the country rocks of the muscovite-biotite granite at the time of its emplacement and assuming a minimum geothermal gradient of 25°C/km 

at the time of emplacement (see below), the maximum depth of emplacement of muscovite-biotite 

granite near the suture zone was about 5 km (that is, 

130°C, microcline argon-closure temperature, divided 

by 25°C/km). Both the low temperature of the country rocks and the relatively shallow depth support 

the contention that the muscovite-biotite granite 

cooled rapidly upon emplacement. The contention 

that the granites were emplaced about 75 Ma also is 

supported by the date of 76.7±0.4 Ma from muscovite 

sample RIO (fig. 10.6; tables 10.2, 10.3), from a pegmatite dike that intruded undeformed tonalite and 

may be part of the same granitic intrusive episode.

Additional 40Ar/39Ar muscovite dates from muscovite-biotite granite throughout the southern part of 

the Idaho batholith range from 70.1±0.3 to 76.7±0.4 

Ma (table 10.3). Muscovite sample 1KE052 from 

muscovite-biotite granite of the Buffalo Hump area 

(fig. 10.23) yields the best constrained of these cooling ages. Lund and others (1986) dated six mineral 

separates from this area (fig. 10.24): two hornblendes 

and one biotite from two amphibolites within a metasedimentary roof pendant (samples 2KE080 and 

2KE086), muscovite and microcline from muscovitebiotite granite (sample 1KE052), and muscovite from 

a younger gold-bearing quartz vein (sample 2KE087). 

These data indicate that the amphibolite and surrounding metasedimentary rocks cooled below 280°C 

by 78 Ma, that the granite intruded the metasedimentary roof at about 74 Ma, and that all rocks 

cooled below about 100°C by 55 Ma. (Lund and others, 1986, used a multidomainal diffusion model for 

interpretation of the microcline data in this study. 

This model recognizes the probable existence in any
400 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

46°

117° 115° 113°

43°

 \

25 MILES 

_____I

EXPLANATION

[ [ Pre- and post-batholithic rocks 

{^§j^| Eocene plutonic rocks 

IJfM-ffi Muscovite-biotite and biotite granite and granodiorite

Biotite granodiorite

Porphyritic granodiorite, hornblende-biotite granodiorite, and tonalite

(^^| Deformed plutons Primarily tonalite and hornblende-biotite granodiorite 

along western side of Idaho batholith

FIGURE 10.21. Sample locality map for central part of Idaho batholith. Geologic base modified from Kiilsgaard and 

Lewis (1985), Lund and others (1989), and Worl and others (1991). 1, Molybdenite porphyry (samples TC-S-41-16

and TC-S-124-23); 2, Hornblende-biotite granodiorite (sample MP201); 3, Muscovite-biotite granite (sample D3113M); 

4, Muscovite-biotite granite (sample B12178201); 5, Leucogranite (sample B12198202).
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 401

potassium feldspar of at least two argon diffusion domains with different closure temperatures, in this 

case, 150° and 100°C.) Interpolating the cooling history of these rocks from the biotite cooling age 

through the microcline cooling age and using a geothermal gradient of 25°C/km, the maximum depth of 

emplacement of the granite was less than 9 km (fig. 

10.27), which is in agreement with the conclusion 

above about depth of emplacement of muscovite-biotite granite near the suture zone. In addition, goldbearing quartz veins were formed at 71 Ma within 

fractures that had cut the already cool (less than 

280°C) granite and roof pendants.

The low-temperature cooling history preserved in 

the microcline sample from Buffalo Hump is nearly 

identical to that of a microcline sample (B12198202; 

fig. 10.21; table 10.3) from leucogranite near the 

Warm Springs Forest Service Station on the South 

Fork of the Payette River, which is almost 160 km 

south of Buffalo Hump (fig. 10.1). Although more 

than two samples are necessary to draw detailed conclusions about the cooling history of the central part 

of the Idaho batholith, these nearly identical age 

spectra indicate that a large part of the central Idaho 

batholith had cooled below 130°C by 67-55 Ma and 

was not affected by any later thermal activity

Sample B12198202 (fig. 10.21) is also important because the biotite separate yielded a highly discordant,

W 100 

cc

£ 

u.

O 90 

w

O 

5 80

70

60

39

50 

'ArK RELEASED, IN PERCENT

100

FIGURE 10.22. Composite age-spectrum diagram for Thompson 

Creek porphyry molybdenum deposit and for hornblendebiotite granodiorite near Stanley, Idaho. Numbers refer to 

specific samples: 1, Muscovite sample TC-S-124-23 from 

Thompson Creek porphyry molybdenum deposit; Tp (plateau 

date)=87.6+0.4 Ma (plateau date denned by 88 percent of 

39ArK released). 2, muscovite sample TC-S-41-16 from Thompson Creek porphyry molybdenum deposit; Tp=87.4±0.4 Ma 

(plateau date denned by 94 percent of 39ArK released). 3, 

Hornblende sample MP201 from hornblende-biotite granodiorite near Stanley, Idaho; Tp=86.3±0.4 Ma (plateau date defined by 98 percent of 39ArK released).

"hump-shaped" pattern (table 10.3) that is typical for 

slightly altered biotite (Kunk, 1982; Hess and Lippolt, 

1986). No single temperature step from this sample 

yielded a geologically meaningful age, and the totalgas date, which is a weight-averaged date including 

all data and is comparable to a conventional K-Ar 

date, also has no geologic meaning. Because the totalgas date is within the realm of reasonable geologic 

ages, it exemplifies the danger of dating biotites from 

complex areas by the conventional K-Ar method. As 

pointed out by Lewis and others (1987), most cited

115°45' 115037'30"

45° 

37' 

30"

45° 

32' 

30" X2 

0 1 KILOMETER

OREGON BUTTE

EXPLANATION 

Igneous rocks (Cretaceous)

Igneous and metasedimentary rocks, undivided 

(Cretaceous and Middle Proterozoic)

[\X\3 Metasedimentary rocks (Middle Proterozoic)

s ^ Contact

^ Vein

 Mine or prospect

X2 Sample locality

FIGURE 10.23. Geology and sample locality map of Buffalo 

Hump area modified from Lund and others (1986). 1, Amphibolite (sample 2KE080); 2, Muscovite-biotite granite 

(sample 1KE052); 3, Quartz vein from Mother Lode mine 

(sample 2KE087); 4, Amphibolite (sample 2KE086).
402 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

geochronologic data for the Idaho batholith are conventional K-Ar dates on biotite (or even less reliable 

Pb-alpha numerical ages on zircon). Scatter in the 

conventional K-Ar dates has led to much confusion 

about the actual age of the Idaho batholith and even 

has been used to construct elaborate uplift histories 

that have no basis in geologic reality (see, for example, Criss and others, 1982). However, Snee (1982), 

Lund and others (1986), Criss and Fleck (1987), and 

Meen and others (1988), among others, have demonstrated that 4°Ar/39Ar age-spectrum analyses of hornblende, biotite, muscovite, and potassium feldspar 

yield useful data that are providing an understanding 

of the age and extent of Cretaceous magmatism in the 

northwestern United States.

Other muscovite dates for granite, which are as 

young as 72.4±0.4 and 70.1±0.3 Ma (samples D3113M 

from near Warm Lake and B12178201 from north of 

Warm Lake and east of McCall, respectively; table 

10.3), were obtained from samples collected near the 

center of a large area of muscovite-biotite granite 

(fig. 10.21); muscovite dates from the center of the 

granite area tend to be younger than muscovite dates 

from the outer parts of the granite, as described

80

LU 60 

CJ

LU 

EC

£ 40 

<

2h 1h

3b

4mu

5mu

6mi

50 100

39ArK RELEASED, IN PERCENT

FIGURE 10.24. Composite age-spectrum diagram for Buffalo 

Hump district and Mother Lode mine. Ih, Hornblende sample 2KE086 from amphibolite; Tp (plateau date)=80.5±0.6 Ma 

(plateau date defined by 90 percent of 39ArK released). 2h, 

Hornblende sample 2KE080 from amphibolite; Tp=81.2±0.4 

Ma (plateau date defined by 61 percent of 39ArK released). 3b, 

Biotite sample 2KE080 from amphibolite; Tp=78.2±0.4 Ma 

(plateau date defined by 88 percent of 39ArK released). 4mu, 

Muscovite sample 1KE052 from muscovite-biotite granite; 

Tp=73.8±0.4 Ma (plateau date defined by 95 percent of 39ArK 

released). 5mu, Muscovite sample 2KE087 from quartz vein 

at Mother Lode mine; Tp=71.0±0.4 Ma (plateau date defined 

by 89 percent of 39ArK released). 6mi, Microcline sample 

1KE052 from muscovite-biotite granite; apparent argon loss 

from Tmax (maximum date)=67.0±0.3 Ma to Tmin (minimum 

date)=54.5±0.4 Ma.

above. This trend may be a result of multiple intrusive pulses, differential cooling, or exposure of different emplacement levels.

In addition to dates obtained directly from samples 

of the muscovite-biotite granites, muscovites from 

mineral deposits throughout the central part of the 

batholith and near roof zones of the muscovite-biotite 

granites yield dates ranging from about 78 to 68 Ma 

(Snee and others, 1985; Gammons and others, 1985; 

Gammons, 1986; Lund and others, 1986; Snee, 1987; 

and Snee and Kunk, 1989). These muscovites are 

generally from quartz veins that were emplaced in 

fractures and fault zones near or within roof pendants in the biotite and muscovite-biotite granite. This 

range in dates is consistent with the data for the 

granites and supports the conclusions that the deposits are genetically related to the granites and that 

granite emplacement and quartz vein formation 

spanned this time period.

DISCUSSION

The 4°Ar/39Ar age-spectrum data of this study provide time constraints on several metamorphic and 

plutonic events in western Idaho and eastern Oregon. 

These data can be used to reconstruct the geologic 

development of a large area straddling the Salmon 

River suture zone between accreted island-arc terranes and the pre-Cretaceous North American continent. The available argon data include (1) the age of 

one of the latest pre-accretion metamorphic events 

and two periods of subsequent pre-accretion plutonism within the accreted island-arc terranes; (2) the 

age of metamorphism that occurred during accretion 

(Lund and Snee, 1988) of previously amalgamated allochthonous island-arc terranes (Vallier and Engebretson, 1983; Hillhouse and others, 1982) to the 

North American continent along the Salmon River suture zone; (3) the time of final stitching of the islandarc terranes to the North American continent marked 

by emplacement of plutons along the Salmon River 

suture zone; (4) the age of postaccretion deformation 

caused by vertical movement along the suture zone; 

(5) the age of emplacement of metaluminous plutons 

of the Idaho batholith and the age of apparently associated porphyry deposits; and (6) the age of emplacement of the large bodies of muscovite-biotite granite of 

the Idaho batholith and associated mineral deposits. 

In addition to direct ages for these events, the 

40Ar/39Ar age-spectrum mineral dates used within a 

time-temperature framework provide information 

about depth of emplacement of plutons, cooling, and 

uplift.
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 403

METAMORPHISM AND PLUTONISM 

IN EASTERN OREGON AND WESTERN IDAHO

Our 40Ar/39Ar age-spectrum data record several periods of metamorphism and plutonism that affected 

island-arc rocks in west-central Idaho and eastern 

Oregon. The oldest metamorphism documented by 

this study took place at or before 244 Ma and affected basement rocks now exposed in Snake River canyon. This age of metamorphism is consistent with 

ages reported by numerous studies (for example, 

Walker, 1981, 1983, 1986; Ave Lallemant and others, 

1980, 1985) for plutonism and shearing within the allochthonous island-arc terranes and, on the basis of 

current data, is likely the youngest dynamothermal 

metamorphic event that took place during proposed 

amalgamation of separate island-arc terranes and associated shearing (Ave Lallemant and others, 1980; 

Lund and Snee, 1988) before the terranes were accreted to the North American continent.

At about 227 Ma, tonalite and quartz diorite plutons and dikes intruded the 244-Ma metamorphic 

rocks after the metamorphic rocks had cooled below 

280°C, the argon-closure temperature of biotite. T.L. 

Vallier (oral commun., 1988) regards these igneous 

rocks as part of magmatic roots of volcanoes that were 

active during the deposition of the contemporaneous 

volcanic rocks of the Wild Sheep Creek Formation. In 

turn, this age is very important for understanding the 

timing of pre-accretionary structural events within 

the island-arc terrane because, on the basis of the 

mapping of Vallier (1974) and White and Vallier 

(1994), the Triassic Wild Sheep Creek Formation of 

the Seven Devils Group was deposited on basement 

rocks metamorphosed during the 244-Ma event and 

subsequently intruded by these 227-Ma plutons. 

Thus, the cooling after metamorphism of the basement rocks and the deposition of the Wild Sheep 

Creek Formation, which is Middle Triassic (Ladinian; 

between 235 and 230 Ma), is bracketed between 244 

and 227 Ma. Therefore, the age of the plutons is consistent with the interpretation that their emplacement is temporally close to the beginning of Seven 

Devils Group deposition. Other studies (Hotz and others, 1977; Ave Lallemant and others, 1980, 1985; 

Walker, 1981, 1983, 1986) have recorded additional 

events, including structural emplacement of blueschist, ductile shearing, and plutonism, within the island-arc terranes at about 227 Ma.

On the basis of zircon U-Pb dates, Walker (1981, 

1983, 1986) has shown that several periods of plutonism occurred in the island-arc terranes between 266 

and 145 Ma. His data were derived from samples of 

deformed and metamorphosed as well as unmetamorphosed plutons. The argon data for rocks of Snake 

River canyon clarify and provide some constraints 

with respect to tectonic interpretation of the zircon 

U-Pb data. Metamorphism at 244 Ma was a major 

event that affected preexisting plutons and volcanic 

rocks, and plutonism and volcanism seem to have resumed almost immediately after metamorphism. 

High-grade metamorphism, like that of the 244-Ma 

event, is not generally associated with the numerous 

plutonic episodes documented by the U-Pb data. 

Thus, the 244-Ma metamorphism represents a singular dynamothermal event that dates the construction 

and (or) amalgamation of the island-arc terrane. Subsequent to this major amalgamation event, plutonism, blueschist emplacement, and ductile shearing 

affected the amalgamated rocks at 227 Ma and mark 

the reorganization of the terranes and resumption of 

island-arc volcanism.

At about 145 Ma, plutons with island-arc chemical 

affinity and compositions ranging from tonalite to 

granite (Armstrong and others, 1977) were emplaced 

into the Seven Devils island-arc terrane in several 

areas in the Blue Mountains and along the Rapid 

River, west of Riggins. Plutons of the Wallowa batholith of northeastern Oregon were emplaced into the 

Seven Devils island arc at about the same time 

(Armstrong and others, 1977; L.W Snee and WH. 

Taubeneck, unpub. data, 1987) and were possibly 

part of the same magmatic event. The 145-Ma tonalite pluton along the Rapid River, southwest of Riggins, which we dated, was later metamorphosed to 

greenschist facies by the events that affected rocks 

along the Salmon River suture zone to the east. Because some of these plutons were affected by metamorphism associated with Salmon River suture zone 

activity, their 145-Ma age provides the maximum age 

for accretion of the island-arc terranes to the North 

American continent.

Low- to high-grade metamorphism affected a belt 

of rocks along the Salmon River suture zone beginning about 130 Ma. At least four groups of metamorphic and deformational ages (130, 118, 109, and 101 

Ma; fig. 10.25) are documented in the Riggins Group 

and Seven Devils island-arc rocks near the suture 

zone. A summary of the thermal history is displayed 

on a time-temperature diagram (fig. 10.26). The oldest metamorphic event associated with accretion 

along the Salmon River suture zone is documented in 

rocks of the Orofino area and occurred at 130 Ma. 

The oldest documented event in the Riggins-Slate 

Creek area occurred at about 118 Ma. Both the 130- 

and 118-Ma events were dynamothermal and were 

marked by a rapid rise to maximum metamorphic 

conditions followed by rapid cooling. The next two
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10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 405

though some tonalitic magma with island-arc affinity 

was emplaced into the metamorphic rocks during the 

118- to 101-Ma interval, plutonism did not cause the 

dynamothermal metamorphism exhibited within the 

Salmon River suture zone. Instead, plutonism trailed 

metamorphism by a few million years and was more 

likely an effect of the tectonic activity that caused 

the metamorphism.

At about 93 to 90 Ma, the Salmon River suture zone 

was the site of a fifth period of metamorphism and 

deformation that was accompanied by emplacement of 

plutons with compositions ranging from tonalite to 

granodiorite (figs. 10.25, 10.26); these plutons reheated adjacent island-arc metamorphic rocks, which had 

cooled below 280°C before the 93- to 90-Ma event. Emplacement of the tonalitic plutons stitched the islandarc terranes and the North American continent 

together, and the island-arc terranes became part of 

the stable North American continent. Beginning at or

before 88 Ma and extending to at least 79 Ma, deformation resulting from vertical movement affected the 

Salmon River suture zone.

The emplacement of these 93-Ma tonalitic plutons 

marked the initiation of magmatism that formed the 

Idaho batholith. These plutons were emplaced into island-arc and North American continental rocks that 

had been juxtaposed along the Salmon River suture 

zone. The preservation of the initial 87Sr/86Sr ratio 

discontinuity in these plutons is conclusive evidence 

that both the island-arc and continental terranes influenced the composition of the magmas. Undeformed 

metaluminous plutons in the vicinity of the suture 

zone were emplaced at the same time as some of the 

deformed tonalitic plutons within the suture zone. 

Until about 86 Ma, metaluminous magmatism affected 

an area extending up to 139 km east of the suture zone, 

where plutons of the Idaho batholith with similar 

compositions were emplaced. The 87.5-Ma porphyry

y\j\j

700

03 

03

UJ

O 

03

Hi 500

DC 

0 

UJ 

Q

Z

UJ 

DC

^

gj 300 

Q. 2

UJ

100

n

I I 1 I 1 1 1 1 1

^-Tacking Preaccretion - ^

^--»^^

/

Uplift ^^ "*7

/

Cooling of ' \_^7\ 

eastern terrane I

/

H ;

<^ Tonalite emplacement ^ U

Episodes of metamorphism

and (or) deformation _

999

I 1 /

1 / 1 /

H H H H H H H

1 * /i ' 1 .1 / 1 Cooling of / ' /| j | / 

suture zone ^ VL^V ii li II / 

1 / 1 1 1 1 il ' / 

/ / ' ' 1 1 "l ll / 

Granite ? / / / 1 ' M ,1 / 

emplacement / // / i I I V / 

' / / 1 . 1 \ 1 /

M f M 

B BB

/ / / f

1 /

1 II / ^^^ / Is

rrl m* ,..,

/ \y \ ' \ \ /

B Y \ B \ / \ B

o^ / \

9 ?

^-^^^ Cooling of

western terrane

Suture-zone activity ^

^^*/ Vertical Translational

Emplacement of Idaho batholith

1 1 1 1 1 1 1 1 1

EXPLANATION 

Metamorphism and cooling of 

metamorphic rocks

Cooling of igneous rocks

|~| Temperature from Naney (1983)

	Temperature from this study: 

B Biotile

M Muscovite

H Hornblende

m Mlcrocline

? Temperature unknown

50 70 90 110

AGE, IN MILLIONS OF YEARS

130 150

FIGURE 10.26. Thermal history of Salmon River suture 

zone, Idaho, based on time and temperature relations derived from 40Ar/39Ar age-spectrum data (from Lund and 

Snee, 1988). Cooling of 115-Ma suture-zone plutons omitted for clarity. The metamorphic episode at 130 Ma is recorded in rocks near Orofino, Idaho, but not in rocks of the 

Riggins, Idaho, area.
406 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

molybdenum deposit at Thompson Creek is a vestige 

of the mineralization that accompanied this period of 

igneous activity.

Beginning about 78 Ma, after an apparent 10-m.y 

lull in plutonic activity, the voluminous muscovite-biotite granites, which are characteristic of the central 

part of the Idaho batholith, were emplaced. Muscovite dates indicate that emplacement of the granites 

and that cooling of the rocks to below 300°C were 

completed by 70 Ma. Many gold-bearing quartz veins 

that are within fractures and fault zones primarily in 

roof pendants of the muscovite-biotite granites were 

emplaced from 78 to 68 Ma during several stages of 

mineralization.

An important result of our study is the discovery 

that the Idaho batholith is younger than the metamorphism of the rocks within the Salmon River suture 

zone. Thus, even though the batholith is adjacent to 

and in some places cross-cuts the suture, its emplacement did not cause the metamorphism. Instead, it is 

reasonable to conclude that metamorphism and possibly plutonism along the suture zone were in response 

to events that occurred during accretion of the islandarc terrane to North America.

COOLING AND UPLIFT HISTORY OF THE AREA EAST OF THE 

SALMON RIVER SUTURE ZONE

The last structural event in the region before the 

onset of extension and basaltic magmatism in the 

Miocene is manifested by sets of north-northeasttrending, high-angle normal faults that are parallel 

to the Salmon River suture zone and along which uplift occurred. In some places (see fig. 10.6) the highangle faults coincide with the initial 8?Sr/86Sr ratio 

discontinuity that marks the Salmon River suture 

zone and apparently truncate the suture zone. The 

uplift event began during or immediately before emplacement of the unfoliated tonalitic plutons of the 

Cretaceous Idaho batholith (Lund and others, 1986) 

and continued at least until after emplacement of Eocene epizonal granitic magmas in the same region 

(Lund and others, 1983b).

The rate of this uplift for the suture zone area can 

be calculated from cooling rates derived directly from 

the argon data plotted on figure 10.26. The cooling 

rate for the undeformed tonalite during the interval 

from 84 to 81 Ma is 95°C/m.y, and the cooling rate 

from 81 to 78 Ma is about 45°C/m.y. This rapid cooling 

rate is consistent either with high-level emplacement 

or with rapid uplift following an unknown period of 

slow cooling after deep emplacement. The presence of 

magmatic epidote suggests emplacement at a depth of

greater than 25 km (Zen and Hammarstrom, 1984) if 

the epidote crystallized at the level of pluton emplacement (Evans and Vance, 1985). Additional study of the 

Rocky Bluff pluton by Hammarstrom and Zen (1988) 

indicates a depth of emplacement of about 20 km 

based on hornblende geobarometry Assuming that the 

tonalite was emplaced at a depth of 20 km or deeper 

and assuming a closure temperature of 530°C for diffusion of argon from hornblende, a geothermal gradient of about 27°C/km is derived. This value is similar 

to, but slightly larger than, that resulting from a synthesis of studies by Carmichael (1978), Hyndman and 

others (1979), Hollister (1982), and Zen (1985). Using 

a geothermal gradient of 25°C/km, the average uplift 

rate from 84 to 81 Ma was 3 mm/yr; the average rate 

from 81 to 78 Ma was about 2 mm/yr. If a higher 

geothermal gradient or greater depth of emplacement 

of the tonalite were assumed, the uplift would have 

been faster. These rates are less than, but on the same 

order as, the rapid recent uplift of 5 mm/yr for the 

Nanga Parbat massif in the Himalaya (Zeitler, 1985) 

and the 2.5 to 14 mm/yr uplift along the Alpine fault 

in New Zealand (Adams, 1981).

In the Buffalo Hump area, located about 50 km 

east of the suture zone (fig. 10.1), cooling rates can 

be determined from figure 10.27 (Lund and others,

600

500

400

300

200

100

1 (HORNBLENDE)

2 (MUSCOVITE) 

3 (MUSCOVITE) -L- i

1 (BIOTITE)

50 60 70 80 

AGE, IN MILLIONS OF YEARS

90

FIGURE 10.27. Cooling curve for Buffalo Hump area from 

Lund and others (1986). Numbers indicate the following 

rock types: 1, amphibolite; 2, muscovite-biotite granite; 3, 

quartz vein. Name of dated mineral is shown in parentheses. Two-domain diffusion model was used for microcline 

date, with closure temperature of most retentive domain 

equal to 130°C and closure temperature for least retentive domain equal to 100°C. Vertical bar in each cross 

shows approximate error estimate in selected closure 

temperatures; horizontal bar shows one-sigma analytical 

error of each date. Dotted lines indicate hypothesized local temperature increases during emplacement of muscovite-biotite granite and quartz veins. Queries indicate 

that relation between local temperature increases and regional cooling curve is unknown.
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 407

1986). The cooling rate for the metasedimentary roof 

pendant at Buffalo Hump from about 81 to 78 Ma, 

before emplacement of muscovite-biotite granite, was 

about 85°C/m.y. This rapid cooling was taking place 

during the same time that the cooling to the west 

along the suture zone had slowed to about 45°C/m.y. 

In the Buffalo Hump area, cooling from 78 Ma to the 

time of emplacement of muscovite-biotite granite at 

74 Ma cannot be directly calculated, but it must have 

been slower than 85°C/m.y and probably closer to 

45°C/m.y. on the basis of interpolation between biotite and microcline cooling ages. Assuming a geothermal gradient of 25°C/km, the rate of uplift of rocks in 

the Buffalo Hump area was probably similar to that 

near the suture zone, but the time of uplift of rocks 

of the Buffalo Hump area lagged behind that of rocks 

of the suture zone by about 3 m.y. The nearly identical age spectra for microcline from muscovite-biotite 

granite at Buffalo Hump and microcline from leucogranite near the Warm Springs Forest Service Station provide constraints on cooling of the central part 

of the Idaho batholith: By 55 Ma, this part of the 

batholith had cooled below 130°C, the argon-closure 

temperature for microcline, and therefore must have 

been no deeper than 5 km. Fluid-inclusion data from 

quartz in gold-bearing quartz veins emplaced at 71 

Ma in the Buffalo Hump area (Muniz and Brown, 

1984) and between 74 and 71 Ma in the Big Creek 

area (fig. 10.21; Gammons and others, 1985; Gammons, 1986) suggest that the depth to the granite 

was less than 4 km.

DEPTH OF EMPLACEMENT OF MUSCOVrTE-BIOTITE GRANITE

Much petrologic debate has focused recently on the 

depth of emplacement of plutons containing magmatic muscovite. The 40Ar/39Ar age-spectrum data of this 

study provide some independent control on the depth 

of emplacement of muscovite-biotite granites in the 

Idaho batholith. Muscovite-biotite granite near the 

Salmon River suture zone was emplaced after 76 Ma 

into country rock that had already cooled below 

130°C (microcline argon-closure temperature; fig. 

10.26). The emplacement of muscovite-biotite granite 

had little or no effect on the age-spectrum of a sample of microcline (sample R21, which cooled below 

about 130°C at 78 Ma) collected from nonfoliated tonalitic plutons only a few hundred yards from the 

granite contact. Assuming a reasonable minimum 

geothermal gradient of 25°C/km, the depth of emplacement of the granite must have been less than 6 

km. Lund and others (1986) used the cooling history 

shown on figure 10.27 to derive a similar estimate

(less than 9 km) for depth of emplacement of muscovite-biotite granite in the Buffalo Hump area. Thus, 

in general, the argon data support contentions of 

Miller and others (1981) and Anderson and Rowley 

(1981) that primary muscovite of non-endmember 

composition can form at relatively low pressure.

An apparent discrepancy exists when the 20-km 

minimum depth of emplacement for the epidote-bearing tonalite of the suture zone is compared to our 

less-than-9-km depth-of-emplacement estimate for 

muscovite-biotite granite of the Idaho batholith. 

However, the existing 4°Ar/39Ar age-spectrum data 

reveal a gap of about 8 to 10 m.y between the end of 

emplacement of metaluminous tonalites and granodiorites and the beginning of emplacement of biotite 

and muscovite-biotite granites. During this time gap, 

our cooling data indicate that more than 10 km of 

uplift affected the region of the Idaho batholith. We 

contend that the time gap and the large amount of 

uplift that occurred during this gap not only explain 

the difference in depth of emplacement but also are 

important clues that directly bear on the abrupt 

change during the Cretaceous in the Idaho batholith 

from metaluminous tonalitic magmatism to the voluminous muscovite-biotite (peraluminous) granitic 

magmatism.

This change in composition of magma and the intervening time gap were a direct result of progressive 

tectonic events along the Salmon River suture zone. 

These progressive events include accretion, tectonically driven crustal thickening, increasing crustal 

temperatures, buoyancy contrast between thickened 

continental crust and underlying mantle, uplift, depressurization melting, and granitic magmatism. During accretion, thrust plates moved both to the west 

and to the east away from the zone of convergence 

that is now the Salmon River suture zone. The degree 

of tectonic thickening to the east of the suture zone 

was 20 km or more on the basis of the hornblende 

geobarometry of Zen and Hammarstrom (1984) for tonalite along the eastern side of the suture zone. Roof 

pendants exposed in these plutons preserve some of 

the thrust plates that are evidence of this crustal thickening (see fig. 10.2). The greatly thickened crust east 

of the suture zone plus the emplacement of tonalitic 

magmas raised the geothermal gradient. This hot, 

thickened crust was of pelitic and psammitic composition and more buoyant than less thickened crust farther east within the craton, creating a buoyancy 

contrast in the underlying mantle. Uplift of the buoyant crust would have resulted. Our argon data indicate 

that this uplift began at about 88 Ma. With the onset 

of uplift, pressure in the lower part of the structurally 

thickened pelitic rocks would have quickly been
408 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

reduced and the trapped heat would have been sufficient to cause melting and the production of peraluminous magma. These magmas in turn were lighter 

than their surrounding source region and rose to less 

than 9 km before freezing. These peraluminous magmas derived abundant volatiles from their hydrated 

source rocks volatiles that were effective in transport 

of metals that now form important mineral deposits 

associated with the muscovite-biotite granites. Our 

cooling age constraints indicate that about 10 m.y. 

elapsed between last emplacement of metaluminous 

magma and first emplacement of peraluminous magma. Peraluminous magma emplacement and associated mineral deposit formation took place over an 

additional 10-m.y. period.

EROSION AND CRETACEOUS SEDIMENTATION

The onset of orogeny along the Salmon River suture zone caused by accretion of island-arc terranes 

to the North American continent occurred approximately 130 Ma; this is the earliest time that significant amounts of coarse continental sediments would 

have been deposited as an onlapping sequence on the 

island-arc rocks. Furthermore, on the basis of the 

cooling histories derived from the argon data, rapid 

erosional unroofing in the area would have occurred 

when 15 km of rapid uplift and erosion of parts of 

the Idaho batholith began about 90 Ma; the rate of 

uplift slowed somewhat by 76 Ma, but uplift certainly continued, albeit more slowly, until the Eocene. By 

55 Ma, much of the Idaho batholith was within a few 

kilometers of the present-day erosion level. The sedimentary record both to the east and to the west reflects these events.

In central Oregon, conglomeratic rocks of the Albian to Cenomanian (113 to 91 Ma) Gable Creek Formation of Wilkinson and Oles (1968) crop out in the 

Mitchell inlier. Partly correlative conglomeratic rocks 

of the Cenomanian Bernard Formation of Dickinson 

and others (1979) crop out in the John Day inlier. 

Detrital grains in sandstone layers include quartz, 

mica, and potassium feldspar. Conglomeratic layers 

include granitic and metaquartzite clasts as well as 

more common greenstone, volcanic, and phyllitic 

clasts. The clasts in these sedimentary rocks reflect a 

mixed provenance of metamorphic, plutonic, and volcanic rocks that is thought to have been a "mature 

and dissected arc terrane" (Dickinson and others, 

1979).

To the east, a major angular unconformity in 

southwestern Montana is located at the base of the 

Cretaceous Kootenai Formation and exhibits a discordance of up to 40° (for example, Myers, 1952; 

Snee, 1982; Zen, 1988). The erosion surface defining 

this unconformity is developed on rocks as young as 

the Jurassic Morrison Formation; however, Jurassic 

rocks are generally thin (a few tens of meters; Perry, 

1986) or absent, and commonly the youngest rocks 

below the unconformity are Triassic in age. Basal 

conglomerate of the Kootenai marks a resumption of 

deposition after a major period of erosion. The age of 

the oldest Kootenai sedimentary rocks in Montana is 

poorly constrained, but is thought to be earliest Aptian (118 Ma; Schwartz and DeCelles, 1988). In southwestern Montana the oldest Cretaceous sediments 

represented by the Kootenai Formation are generally 

only 200 m thick (Zen, 1988).

Major sedimentation began in early Albian time 

(about 113 Ma) and extended to the end of the Cretaceous (66 Ma). It produced a very thick sequence 

consisting of graywacke, sandstone, volcanogenic 

sandstone, ash, and conglomerate that was deposited 

conformably on the Kootenai Formation (Schwartz 

and DeCelles, 1988). In the Pioneer Mountains, Montana (fig. 10.1), more than 2,000 m of this sequence, 

the Colorado Group, was derived from the west and 

deposited in a fluvial environment (Snee, 1982; Zen, 

1988). Recent U-Pb dating of zircon from porcellanite 

beds in the upper, volcaniclastic Vaughn Member of 

the Blackleaf Formation within the Colorado Group 

from the eastern Pioneer Mountains, Montana (fig. 

10.1), yielded dates ranging from 97 to 93 Ma (Zartman and others, 1995). Although slightly older than 

our oldest dates for Idaho batholith magmatism, 

these zircon dates are reasonably consistent with an 

interpretation that the volcanic beds of the Vaughn 

Member were derived during earliest Idaho batholith 

magmatism.

The origin of Aptian through Upper Cretaceous 

sedimentary rocks both to the east and to the west 

of the Salmon River suture zone is thus directly 

linked in time to metamorphic, deformational, plutonic, and uplift events along the suture zone and 

within the area of the Idaho batholith. In fact, 

Schwartz and DeCelles (1988) have demonstrated a 

genetic relation between sedimentation in the Cretaceous and the timing of tectonism established by 

Lund and Snee (1988) along the Salmon River suture 

zone. According to Schwartz and DeCelles (1988), 

coarse sediments were deposited in the Cretaceous in 

three intervals at 118, 109, and 101-99 Ma, corresponding to three periods of metamorphism and deformation along the Salmon River suture zone. On a 

more regional scale, Heller and others (1986) concluded that thrusting during the Sevier orogeny 

began in the early Aptian; this conclusion was
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 409

reached by evaluating synorogenic conglomerates in 

Wyoming, Utah, and Colorado. In addition to the link 

between metamorphic and deformational events and 

the deposition of coarse sediments from 118 to 99 

Ma, we suggest that the thick sequence of Cretaceous 

sedimentary rocks now exposed in western Montana 

reflects the more than 15 km of uplift that occurred 

between about 90 and 75 Ma. The lack of significant 

amounts of Paleocene and Eocene sedimentary rocks 

in southwestern Montana probably reflects the decrease of uplift and erosion in the area of the Idaho 

batholith.

SUMMARY OF CRETACEOUS TECTONIC

HISTORY CENTRAL IDAHO AND

SOUTHWESTERN MONTANA

The Cretaceous tectonic history of central Idaho 

and southwestern Montana from 130 to 75 Ma is 

summarized in figure 10.28 a series of hypothetical 

cross sections from the Snake River canyon to the Pioneer Mountains along a line at latitude 45°30' N. 

These cross sections are intended to generally represent (and in some cases exaggerate) the major activity along the line of section during each of five time 

periods: 130-100 Ma, 95-90 Ma, 85 Ma, 80 Ma, and 

75 Ma. Important events portrayed on the figure are 

summarized below.

From 130 to 100 Ma, the Salmon River suture zone 

was the locus of right-slip transpressive movement 

between island-arc terranes and North America. Associated with this transpression were dynamothermal metamorphism, regional thrust faulting that 

verged both to the east and west away from the suture zone, and local pluton emplacement. Transpression caused thickening of the crust at the suture 

zone. Thrusting was assymetrical with the largest 

area affected lying to the east of the suture zone. The 

thickened crust west of the suture zone underwent 

partial melting at the crust-mantle boundary to form 

tonalitic magmas. Erosion driven by orogenesis resulted in sediments that were deposited in southwestern Montana.

From 95 to 90 Ma, right-slip transpression was virtually complete, and the sense of movement within 

the suture zone changed to predominantly vertical. 

This was a period of major metaluminous magmatism marking the beginning of Idaho batholith plutonism. Plutons were emplaced as far as 139 km to the 

east. These magmas were likely derived from the 

base of the thickened crust as a result of partial 

melting of material of intermediate composition at 

the crust-mantle boundary. The resultant magmas

were buoyant, and their rise drove uplift of blocks 

within the suture zone and to the east. In southwestern Montana, volcaniclastic rocks derived from the 

region of the Idaho batholith were deposited during 

continued sedimentation.

At about 85 Ma, uplift along high-angle faults within 

the Salmon River suture zone was well underway. Metaluminous magrnatism of the Idaho batholith was nearing completion. The area influenced by thrust faulting 

extended into southwestern Montana where some of 

the thrust sheets overrode the recently deposited Cretaceous sediments (Snee, 1982; Zen, 1988).

At about 80 Ma, uplift along high-angle faults in 

the Salmon River suture zone had decreased. Rapid 

uplift was underway in the central part of the Idaho 

batholith. Magrnatism within the area of the Idaho 

batholith was in a lull, but the earliest mafic magmas were emplaced in southwestern Montana (Snee, 

1982). Andesitic volcanism had begun in southwestern Montana (Snee, 1982).

At about 75 Ma, peraluminous magmatism affected 

a large area within the Idaho batholith. These magmas were likely derived from the partial melting of 

pelitic rocks at midcrustal levels within the thrustthickened crust. Gold-bearing quartz veins were emplaced in fractures in roof pendants of these granites. 

In southwestern Montana, this was the predominant 

period of metaluminous magmatism (Snee, 1982; 

Zen, 1988). Some dacitic volcanism was still active 

(Snee, 1982).

After 70 Ma, igneous and hydrothermal activity 

in the region of the Idaho batholith was nearly 

complete. Uplift also was nearly complete. In southwestern Montana, magmatism, faulting, uplift, mineral-deposit emplacement, and erosion continued 

until about 60 Ma (Snee, 1982).

CONCLUSION

The Salmon River suture zone marks an area of 

multiple metamorphic events, complex structure, and 

several periods of intrusion. It also marks a sharp 

lithologic and chemical break between younger island-arc rocks to the west and older continental 

rocks to the east. Lund (1984) and Lund and Snee 

(1988) have proposed that the Salmon River suture 

zone is a right-lateral transpressive fault along 

which exotic island-arc terranes were accreted to the 

North American continent. If this proposal is correct, 

transpression along the suture zone caused some, if 

not all, of the prograde metamorphism that affected 

the Riggins Group and Seven Devils island arc, as 

well as rocks of the North American continent in cen-
410 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

EXPLANATION

Plutons with peraluminous 

affinity

Volcanic rocks east of the 

suture zone

Cretaceous sedimentary 

rocks (sandstone and 

conglomerate) and 

volcaniclastic rocks

Plutons with metaluminous 

affinity

115-Ma deformed suturezone plutons

145-Ma island-arc plutons

Island-arc volcanic, sedimentary, metamorphic, 

and plutonic rocks 

Unpatterned west of 

suture zone

Continental sedimentary, 

plutonic, and metamorphic rocks 

Unpatterned east of 

suture zone

Mantle lithosphere 

SRSZ Salmon River suture zone 

 Contact

Fault

-=i Thrust fault Arrow shows 

relative movement of 

upper plate

I Strike-slip fault A,

movement away from 

observer; T, movement 

toward observer

Trend lines of foliation or 

lineation

FIGURE 10.28. Hypothetical cross sections from Snake Riv- ment of central Idaho and southwestern Montana from

er canyon, at the boundary between Oregon and Idaho, 130 to 75 Ma. (See text for explanation.) No vertical exto Pioneer Mountains, Montana, at latitude 45°30' N. aggeration. Some features are exaggerated to emphasize

From top to bottom, cross sections show tectonic develop- process.
10. TECTONIC DEVELOPMENT OF THE SALMON RIVER SUTURE ZONE, WESTERN IDAHO 411

tral Idaho. This metamorphism began at about 130 

Ma and continued until 93 Ma, when the terranes 

were stitched by the emplacement of plutons of the 

suture zone. If the Salmon River suture zone is the 

structure along which the island-arc terranes were 

accreted to North America, then events prior to 93 

Ma, which are preserved in rocks west of the suture 

zone, occurred when the island-arc terranes were not 

attached to this part of North America. These events 

include the 244-Ma metamorphism and 227- and 

145-Ma plutonism documented by argon dating in 

this report, and the 266- to 145-Ma plutonism demonstrated by the U-Pb dating of Walker (1981, 1983, 

1986).

The emplacement of the Idaho batholith did not 

cause the regional metamorphism exhibited by the 

rocks within the Salmon River suture zone. Indeed, 

emplacement of the Idaho batholith began with the 

intrusion of tonalitic plutons at about 93 to 90 Ma 

within the suture zone and as far away as 139 km 

to the east. The oldest plutons of the eastern part 

of the Idaho batholith thus far identified were emplaced 

at 91 Ma, virtually simultaneously with tonalites of 

the suture zone. Some of the suture-zone plutons were 

emplaced at depths of 20 km or more; others were 

emplaced at relatively high crustal levels and were 

associated with porphyry mineralization. At or before 

88 Ma, uplift at the rate of 3 mm/yr or more began 

along the suture zone; it spread eastward about 3 

m.y. later. An apparent interval of 10 m.y. elapsed 

between metaluminous and peraluminous plutonism. 

Beginning at 78 Ma, muscovite-biotite granites were 

emplaced at depths of less than 9 km after some of 

the terrain east of the suture zone had experienced 

more than 10 km of uplift. Quartz-vein mineralization 

accompanied granite emplacement at relatively shallow depths. By 55 Ma, much of the central part of 

the Idaho batholith had been uplifted to upper crustal 

levels where the temperature was less than 130°C. 

As much as 20 km of uplift (and overburden removal) 

are recorded in the cooling histories of some rocks 

east of the suture zone and in the sedimentary record 

observed both to the east and to the west of the suture 

zone.

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11. GEOLOGY OF THE NORTHERN PART OF 

THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON

By PETER R. HOOPER, MICHEL D. HOUSEMAN, JOHN E. BEANE, GREGORY M. CAFFREY, 

KENNETH R. ENGH, JAMES V. SCRIVNER, and A. JOHN WATKiNSON1

CONTENTS

Abstract 415

Introduction 416

Acknowledgments 416

Analytical methods 417

Rocks of the Baker terrane 418

Peridotite of Bullrun Mountain 418

Serpentinite 418

Chemical composition of peridotite and serpentinite 422

Melange of Mine Ridge 423

Metamorphosed gabbro and basalt 424

Plagiogranite 426

Schistose rocks of Mine Ridge 427

Ribbon chert 429

Pre-Jurassic volcanic rocks 431 

Origin of the rocks of the Baker terrane, Ironside

Mountain inlier 432 

Correlation and timing of the deformation of the

schistose rocks of Mine Ridge 433

Fore-arc basin deposits: the Weatherby Formation 433

Provenance 434

Deformation of the Weatherby Formation 437

Post-Jurassic igneous rocks 437

Granodiorite of Bullrun Creek 440

Granodiorite of Grouse Creek and related minor intrusions 441

Late-stage porphyritic dikes 442

Chemical composition of the Oligocene intrusions 443

Alteration and mineralization 444

Granodiorite of Bullrun Creek Record Mine 444

Granodiorite of Grouse Creek 445

Discussion and conclusions 449

References cited 453

ABSTRACT

The Ironside Mountain inlier exposes Permian and Triassic 

ophiolitic rocks along the southern margin of the Baker terrane at 

the center of the Blue Mountains province. Lying between the 

Canyon Mountain Complex and the (informal) Sparta complex, 

rocks of the Ironside Mountain inlier consist of dunite, harzburgiM\ at Department of Geology, Washington State University, Pullman, 

WA 99164-2812

ite, and a serpentinized melange containing knockers of peridotite, 

metagabbro, plagiogranite, schist, ribbon chert, and pre-Jurassic 

andesite. The peridotite has a high-temperature mineralogy and 

an equivalent bulk chemical composition depleted in A^Os and 

CaO. Its coarse metamorphic fabric predates two stages of serpentinization, the second of which may have been associated with the 

intrusion of the peridotite into the crust. These properties are consistent with an origin as the residuum of partial melting in the 

mantle.

The knockers of the schistose rocks of Mine Ridge include three 

fold phases. FI folds are isoclinal folds whose axes plunge down 

the dip of a steep east-northeast-striking penetrative foliation (S^ 

parallel to S0, henceforth referred to as SQ-SJ) associated with epidote amphibolite-facies metamorphism. The hornblende lies parallel to the FI fold axes and has been dated at 260 Ma. F2 fold axes 

also plunge down the dip of the SQ-SI plane, but the F2 fold axes 

have varied axial surfaces that are at high angles to the foliation 

and are not associated with a new metamorphic fabric. F% folds 

are a younger set of crenulations. All three fold phases are correlated with the regional DI deformation event, which is interpreted 

as a result of the subduction process. In this model the original 

fold axes were formed horizontally in the subduction plane perpendicular to the direction of subduction, but were then steepened and 

reoriented into parallelism with the direction of subduction by extensive transport down the subduction plane.

Rare-earth-element patterns, the presence of large andesite 

blocks in the melange, and the near coincidence between the age of 

crystallization of the ophiolitic rocks and the age of their first major metamorphism all support the hypotheses that (1) the rocks of 

the Ironside Mountain inlier are similar in composition and age to 

the rocks of the Canyon Mountain Complex, and (2) each of these 

ophiolitic sequences represents the root zones of neighboring island 

arcs, rather than the remains of midocean ridge-rift systems that 

were transported large distances to their present positions.

The Permian and Triassic rocks are overlain by calcareous 

wackes of the Jurassic Weatherby Formation, a part of the Izee 

terrane. The wackes were derived in large part from the Permian 

and Triassic rocks, on which they appear to lie unconformably, although observed contacts are now faulted. The Weatherby is deformed into tight chevron folds with horizontal axes and steep 

east-northeast-striking axial surfaces that parallel the So-Sj plane 

of the older rocks. These folds represent approximately horizontal 

north-northwest/south-southeast compression and are correlated 

with the regional D2 event of Jurassic age that others have associated with the accretion of these rocks to the North American plate 

in the Late Jurassic. Failure to recognize the D2 event in the fabric of the knockers of the schistose rocks of Mine Ridge is attributed to the incompetency of the serpentinite matrix of the melange;

415
416 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

the matrix apparently absorbed the strain so that the D2 compression resulted only in the reorientation of the long axes of the 

knockers into the steep east-northeast-striking D2 axial surface.

Undeformed granitic bodies of Late Jurassic, Cretaceous, and 

Oligocene age intrude the Blue Mountains province. Two large igneous bodies and numerous dikes and sills (36 to 34 Ma) intrude the 

Weatherby Formation at the north end of the Ironside Mountain 

inlier. All have a very similar tonalitic composition but are distinguished by texture and degree of alteration. The well-developed 

hydro thermal alteration associated with the younger intrusion 

caused Cu-Au-Ag and associated base-metal enrichment and is the 

first Tertiary porphyry to be documented in northeastern Oregon.

Interpretation of the Permian and Trias sic rocks as the roots of 

the primitive island-arc systems and the derivation of the Jurassic 

Weatherby Formation from these older oceanic rocks imply that 

the various "terranes" of the Blue Mountains province were in 

close proximity to each other at the time of their formation. To this 

extent the term "terrane" may not be entirely appropriate. The 

varied "terranes" appear to represent a complex system of primitive oceanic island arcs formed by subduction off the western edge 

of the North American plate at a latitude relative to the North 

American plate equivalent to that observed today. These terranes 

suffered compression during both the subduction (Permian, Dj) 

and the accretion (Jurassic to Cretaceous, D2) processes.

INTRODUCTION

Ironside Mountain is one of many inliers of preTertiary rocks exposed through the cover of Tertiary 

volcanic and sedimentary rocks of northeastern Oregon (Thayer and Brown, 1973). The inlier forms a 

north-northwest-trending faulted anticlinal ridge of 

Tertiary age immediately west and south of Unity on 

Highway 26, approximately 60 km southwest of Baker 

City (fig. 11.1, inset). The anticlinal ridge has a core of 

Tertiary granitic intrusions.

An important boundary between the Baker (oceanic) terrane to the north and the Izee (fore-arc basin) 

terrane to the south runs east-west across the northern part of the inlier. The Baker and Izee terranes 

are two of the four allochthonous Mesozoic terranes 

recognized in northeastern Oregon (Brooks and Vallier, 1978; Dickinson and Thayer, 1978; Brooks, 1979a; 

Silberling and others, 1984). The other two are the 

Wallowa (Seven Devils arc) terrane and the Olds Ferry (Huntington arc) terrane (fig. 11.1). The four terranes were welded together by Middle Jurassic time 

and subsequently accreted to the western edge of the 

North American craton, a process completed in the 

Late Cretaceous (75-65 Ma) (Fleck and Criss, 1985).

The two northern spurs of the inlier, Mine Ridge 

and Bullrun Mountain (fig. 11.1), are formed by rocks 

of the Baker terrane. The north end of Bullrun 

Mountain is a large fault-bounded peridotite body. 

Mine Ridge is composed of a serpentinite melange 

with blocks (knockers) of peridotite, gabbro, plagiogranite, and basalt of ophiolite affinity. The knockers 

are intermixed with ribbon cherts, relatively fresh

massive blocks of pre-Jurassic andesite, and various 

schistose units (the schistose rocks of Mine Ridge), 

which include hornblende schist, mica schist, and 

quartzite. The units of the melange are of mid-Permian to Triassic age (Lowry, 1968; Brooks and Ferns, 

1979; Brooks and others, 1979), are metamorphosed 

to the epidote amphibolite facies, and are complexly 

folded. The peridotite of Bullrun Mountain and the 

melange of Mine Ridge are separated by the Oligocene granodiorite of Bullrun Creek (fig. 11.1).

The Izee terrane is represented by the Jurassic 

Weatherby Formation of Brooks (1979b), which consists of calcareous graywacke and siltstone with minor horizons of conglomerate, tuff, limestone, arkosic 

sandstone, and shale. All exposed contacts between 

the Jurassic Weatherby Formation and the Permian 

and Triassic melange are faulted, but an originally 

unconformable contact is suggested both by the presence of the graywacke north and south of the older 

units and by conglomeratic horizons close to the contact containing clasts apparently derived from the 

melange (Lowry, 1968; Brooks and Ferns, 1979; 

Brooks and others, 1979). The Weatherby has been 

folded in a single episode of deformation.

Many dikes, sills, and stocks of fine-grained granodiorite intrude the Weatherby along the axis of the 

anticlinal ridge and are now known to be of Oligocene 

age (34-36 Ma) (N.W Walker, oral commun., 1986; 

Manville Products Corporation, oral commun., 1986). 

To the south of the mapped area, however, the diorite 

of Tureman Ranch intruded the Weatherby in the 

Early Cretaceous (Brown and Thayer, 1966; Lowry, 

1968; Thayer and Brown, 1973). Mild deformation of 

the inlier continued into the Holocene, as fault-controlled basins were filled with fluvial and lacustrine 

Pliocene to Holocene sediments that have steep faulted contacts against the older rocks (Brooks and Ferns, 

1979; Brooks and others, 1979).

The northern part of the Ironside Mountain inlier 

is a microcosm of the complex geology of the accreted 

terranes of the Blue Mountains province in northeastern Oregon. Over several years the authors have 

mapped critical parts of the inlier with the purpose 

of clarifying the structural relations between the 

units, correlating them with similar units exposed 

elsewhere in northeastern Oregon, and so gaining a 

better understanding of the origin of the various 

rocks represented.

ACKNOWLEDGMENTS

Study of the Ironside Mountain inlier was originally suggested, encouraged, and subsequently supported by Howard Brooks and Mark Ferns of the Oregon
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 417

Department of Mines and Geology, Baker City office. 

We greatly appreciate their continuing support. Particular aspects of the study were also supported by 

the Washington Mining and Mineral Resources Research Institute, Manville Corporation, and the Ladies Auxiliary to the American Institute of Mining, 

Metallurgical, and Petroleum Engineers. Peter Hooper acknowledges the kind cooperation of J.N. Walsh, 

Kings College, London, in making his inductively 

coupled plasma laboratory available for trace- and

rare-earth-element analyses. Detailed reviews by 

Mark Ferns, James Evans, and Tracy Vallier added 

significantly to the clarity of the paper and are gratefully acknowledged.

ANALYTICAL METHODS

Samples were analyzed for major elements on a 

manual Philips 1410 X-ray spectrometer using 2:1

118° 16' 118° 14'

EXPLANATION

Fluvial and lacustrine deposits 

(Holocene to Pliocene)

Ashflow tuff (Miocene)

Strawberry Volcanics (Miocene)

Granodiorite of Grouse Creek 

(early Oligocene)

Granodiorite of Bullrun Creek 

(early Oligocene)

Granodiorite intrusions (Tertiary)

Weatherby Formation of Brooks 

(1979b) (Middle to Early Jurassic)

Andesite (pre-Jurassic)

Peridotite and serpentinite 

(Triassic and Permian) Ribbon chert (Triassic and Permian)

Serpentinite melange (Triassic 

and Permian)

Contact Dashed where 

approximately located Plagiogranite (Triassic and Permian)

Metamorphosed gabbro and basalt 

(Triassic and Permian) Fault Bar and ball on 

downthrown side

FIGURE 11.1. Simplified geologic map of north end of Ironside Mountain inlier. Inset shows position of inlier in relation to the four major 

allochthonous terranes recognized in northeastern Oregon. Patterns in inset map show outcrop areas in the four terranes.
418 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Li2B4O7 :rock powder fused disks (Hooper and Atkins, 

1969; Hooper and others, 1976). Trace elements, including rare-earth elements, were analyzed on an 

automated Philips inductively coupled plasma spectrometer by the senior author at Kings College, London (Thompson and Walsh, 1983). The compositions of 

olivines and orthopyroxenes in the peridotite were determined on an automatic Cameca electron microprobe. X-ray diffraction analysis of four samples was 

used to verify the optical identification of serpentine 

polymorphs.

ROCKS OF THE BAKER TERRANE

PERIDOTITE OF BULLRUN MOUNTAIN

A large body of peridotite which is here called the 

peridotite of Bullrun Mountain covers an area of about 

4 km2 at the northeast end of the Ironside Mountain 

inlier (figs. 11.1, 11.2). Smaller bodies of ultramafic 

rock, most totally serpentinized, form pods and the 

matrix of the melange on Mine Ridge to the west. The 

main peridotite body forms bold rounded cliffs on 

Bullrun Mountain, broken by more easily eroded serpentinized zones. It is separated from the younger 

Weatherby Formation to the south by serpentinized 

fault zones intruded by younger granodiorite dikes. In 

the valley of Bullrun Creek the peridotite is intruded by, 

and occurs as inclusions in, the granodiorite of Bullrun 

Creek (fig. 11.2). To the north and northeast the ultramafic rocks are onlapped by the Pliocene to Holocene 

fluvial and lacustrine sediments of Unity basin. The 

covered contact between the inlier and the basin is 

assumed to be faulted (Thayer and Brown, 1973).

The peridotite is everywhere partially serpentinized 

(30-90 percent serpentine with an average of 60 percent; Caffrey, 1982), but in the main peridotite body 

the original mineralogy is generally apparent. It is 

restricted to the three phases olivine, orthopyroxene, 

and chromite. The rocks range from dunite (0-10 percent orthopyroxene) to harzburgite (10-25 percent orthopyroxene) (Streckeisen and others, 1973). Vague 

and irregular compositional banding is present locally 

(fig. 11.3). Only one crystal of clinopyroxene has been 

observed, and the orthopyroxene lacks exsolution lamellae. Plagioclase is absent.

The olivine forms small 0.5-1.0 mm grains separated by a mesh of serpentine. The small grains occur in 

optically continuous groups as much as 10.0 mm in 

diameter. The groups are elongate parallel to each 

other and represent primary olivine crystals. Deformation lamellae and, less commonly, kink bands are 

visible in approximately 60 percent of the primary 

crystals; irregular shadowy extinction is present in 

most others. Kink bands may be traced across many

smaller grains of a single primary crystal and demonstrate that serpentinization postdates the deformation 

that produced the kink bands (fig. 11.4). The olivine 

grains have optic angles close to 90° (2Vy=89° to 93°, 

estimated), and electron microprobe analyses indicate 

compositions ranging from Fo90.7 to Fo92.4 (table 11.1).

Orthopyroxene crystals are as much as 8.0 mm in 

length, enclose small grains of chromite, and tend to 

poikilitically enclose olivine. Crystals are commonly 

bent, and kink bands are developed in many. Orthopyroxene has been altered to anthophyllite or Mg-rich 

cummingtonite along its cleavage and rims, and may 

have been completely pseudomorphed in the more 

highly serpentinized rocks. Most crystals are optically 

positive (2Vy=87° to 93°±5°), and electron microprobe 

analyses give a compositional range of En910 to 

En95 3 , well within the enstatite range (table 11.1).

Chromite is ubiquitous in both the dunite and 

harzburgite in modal proportions varying from 1 percent to 4 percent. It appears as scattered brown or 

red-brown isotropic grains with both embayed and 

euhedral shapes. It alters to magnetite on rims and 

along fractures.

The primary-phase assemblage of the ultramafic 

rocks (olivine+orthopyroxene+spinel) is typical of 

mantle, not crustal, conditions. The coarse grain size, 

the lack of zoning in both olivine and orthopyroxene, 

the lack of significant rhythmic layering, and in particular the lack of either plagioclase- or clinopyroxenebearing layers suggest that the peridotite of Bullrun 

Mountain is a residue of partial melting in the mantle. However, the nonmetamorphic poikilitic relations 

between olivine and orthopyroxene are a common feature of cumulate rocks.

Similar cumulate-like textures are common in alpine 

peridotites and other ultramafic assemblages that are 

associated with ophiolite complexes and have generally 

been interpreted as residual mantle (Moores, 1969; 

England and Davies, 1973; Jackson and others, 1975; 

Coleman, 1977; Evarts, 1977; Malpas, 1977). The olivine alignment seen in the peridotite of Bullrun Mountain is typical of mantle metamorphic fabrics (Coleman, 

1977), and the embayed shape of many chromite grains 

in the peridotite mirrors those reported by Augustithis 

(1979) from mantle xenoliths in Ethiopian basalt and 

from mantle diapirs in Greece. The experimental work 

by Ave Lallemant and Carter (1968, 1970) indicates 

that recrystallization in the upper mantle can yield all 

of these typical alpine peridotite fabrics.

SERPENTINITE

Samples of peridotite described above have retained their primary structure and enough of their
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 419

primary mineralogy to permit unequivocal identification as peridotite. Elsewhere on Bullrun Mountain

the olivine has been completely serpentinized, and 

for these rocks the term serpentinite is used.

118° 15'00"

'« '<e VJ^

1 MILE

1 KILOMETER

EXPLANATION

Fluvial and lacustrine deposits 

(Holocene to Pliocene)

Granodiorite of Grouse Creek 

(early Oligocene)

Granodiorite of Bullrun Creek 

(early Oligocene)

Granodiorite intrusions (Tertiary) 

Most bodies belong to the granodiorite 

of Grouse Creek

Weatherby Formation of Brooks (1979b) 

(Middle to Early Jurassic)

Plagiogranite (Triassic and Permian)

Metamorphosed gabbro and basalt 

(Triassic and Permian)

Serpentinite (Triassic and Permian)

Peridotite (Triassic and Permian) 

Dunite and harzburgite

Unmapped

 1 Fault Bar and ball on downthrown side

FIGURE 11.2. Geologic map of Bullrun Creek and north end of Bullrun Mountain. Contour interval, 200 ft. Base from U.S. Geological Survey 

l:24,000-scale quadrangles: Bullrun Rock, 1972 (photorevised 1983); Rail Gulch, Rastus Mountain, and Unity, 1972 (photorevised 1984).
420 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 11.1. Compositions ofolivine and orthopyroxene in peridotite from Bullrun Mountain 

[ppm, parts per million; tr, trace; , not determined]

Olivine Orthopyroxene

A

Sample

GC-81-126G

GC-81-139B

GC-81-6

GC-81-107

Rock type

Dunite

Dunite

Harzburgite

Harzburgite

Forsterite

(percent)

91.9-92.4

90.8-91.8

90.7-91.2

90.8-91.4

Ni

(ppm)

1,070-1,100

1,070-1,130

1,040-1,150

900-1,050

Cr

(ppm)

tr

tr

tr

tr

Enstatite

(percent)

91.0-95.3

91.2-91.6

Ni

(ppm)

220-260

130-150

Cr

(ppm)

___

1,840-1,960

1,350-1,460

Area of figure 11.3 B

1.8 m

25.4 cm

B

FIGURE 11.3. Compositional banding in peridotite of Bullrun 

Mountain. A, Sketch of peridotite outcrop showing banding 

and shearing. B, Detail from figure 11.3A. 1, dunite; 2, 3, harzburgite. Large crystals shown are orthopyroxene; they lie in 

matrix of partially serpentinized olivine grains.

FIGURE 11.4. Photomicrograph of kink bands in olivine. Note continuity from grain to grain across serpentine veinlets. Rock is 

dunite, viewed in crossed nicols.

Serpentinite is less resistant to erosion than the peridotite and forms a more subdued outcrop pattern; it 

often is seen only as loose debris on the ground surface. Various textural types of serpentinite may be 

correlated in a general way with the serpentinite's 

position relative to the main peridotite body (fig. 11.2). 

Most common is a distinctive banded serpentinite (fig. 

11.5A) that forms broad zones striking east-west and 

dipping steeply northward in the peridotite of Bullrun 

Mountain. The banded serpentinite is also exposed 

over large areas to the east and west of the peridotite 

body (fig. 11.2). Brecciation, slickensides, and other 

evidence of shearing are generally absent in the banded serpentinite, which is composed of black bands of 

magnetite with a roughly planar orientation in a matrix of olive-green lizardite associated with some brucite (fig. 11.5A). Pseudomorphs of anthophyllite after 

orthopyroxene are present, and primary chromite remains, altered to magnetite around its margins. Locally, the banded serpentinite grades into a variety
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 421

that has a radiating acicular (felted) texture crosscut 

by trains of tiny magnetite grains.

Near the faulted southeastern margin of the peridotite body (fig. 11.2), the banded serpentinite grades 

into a mottled variety whose appearance is due to the 

accumulation of magnetite into irregular, slightly 

elongate clots (fig. 11.5B). Near the Thomason Mine 

(fig. 11.2), the mottled serpentinite is severely sheared 

and has a matrix of antigorite (Caffrey, 1982). Primary chromite is present, and some talc has formed; a 

small volume of a chlorite phase appears to be pseudomorphing orthopyroxene.

Talc-bearing serpentinite breccia is restricted to the 

faulted and hydrothermally altered southeastern 

mar-gin of the peridotite body, near the Thomason Mine 

(fig. 11.2). It contains cryptocrystalline silica derived

Weathered rim (white)

from hydrothermal activity and talc derived from the 

silicic alteration of serpentine (Caffrey, 1982). An iron 

oxide stain and slickensides are common. An oval outcrop of serpentinite breccia about 50 ft long within the 

main body of peridotite, unrelated to faulting and containing fragments coated with cryptocrystalline magnesite and (or) dolomite, is suggestive of a breccia pipe.

A narrow zone of intensely sheared and fractured 

serpentinite is present along the poorly exposed intrusive contact with the granodiorite of Bullrun 

Creek and is believed to have resulted from contact 

metamorphism and faulting (Caffrey, 1982).

Timing of serpentinization relative to the peridotite 

emplacement remains elusive. The mottled and 

sheared serpentinite along the southeastern margin 

of the peridotite body is associated, at least in part,

Serpentine matrix 

(olive gray, light olive gray)

Magnetite band (black) Anthophyllite after orthopyroxene

Magnetite (black)

Serpentine matrix Irregular coloration 

(dark greenish gray, pale yellowish green)

3cm

Serpentine matrix Irregular coloration 

(light olive gray, olive gray, pale olive)

Serpentine fragments (brownish black or dusky brown). 

Thin white rind (dolomite?) on most fragments

FIGURE 11.5. Three main varieties of serpentinite. A, Banded serpentinite; B, mottled serpentinite; C, serpentinite breccia.
422 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 11.2. Major- and trace-element analyses ofperidotite from Bullrun Mountain and Mine Ridge

[Analyses by X-ray-fluorescence spectroscopy, Washington State University (Hooper and Johnson, 1989). Major elements normalized on a volatile-free basis 

with 0.9 Fe expressed as FeO, 0.1 Fe expressed as Fe^s, and Fe* equal to FeO+Fe203. , not determined]

Dunite, Bullrun Mountain Harzburgite, Bullrun Mountain Mine Ridge

Sample GC-1 TC-5 GC-7 BRUM- JB-81-

A 34

GC-2 GC-3 GC-4 GC-6 GC-10 GC-12 BRUM-BRUM-BRUM- A15B R-68 R-84

B E F

Major elements (weight percent)

Si02 ~ 

A1203 -

TiOa JLIV/^ __

Fe*

MnO 

CaO 

MgO - -

K20 -----

Na20 - 

P205 

44.88

.00

.02

9.78

.17

.08

45.03

.04

.00

..00

44.29

.01

.02

9.53

.16

.08

45.89

.02

.00

.00

46.48

.00

.02

9.25

.17

.64

43.38

.04

.03

.00

43.01

.00

.02

9.39

.15

.10

47.31

.03

.00

.00

45.69

.00

.02

10.10

.17

.67

43.30

.03

.00

.00

47.56

1.18

.02

9.01

.15

1.03

41.01

.03

.01

.00

46.61

.63

.02

10.04

.15

.64

41.89

.02

.00

.00

45.94

.72

.02

9.78

.16

.36

43.01

.02

.00

.00

46.45

.23

.01

9.46

.15

.45

43.23

.01

.00

.00

46.51

.70

.02

9.53

.15

.55

42.49

.05

.00

.00

46.68

.51

.02

10.02

.19

.61

41.92

.05

.00

.00

47.92

.40

.03

9.54

.15

.59

41.34

.02

.00

.00

46.72

.50

.03

9.70

.16

1.10

41.77

.02

.00

.00

48.93

.35

.03

9.76

.15

.22

40.54

.02

.00

.00

48.52

.19

.03

11.04

.13

.07

40.01

.01

.00

.00

46.44

.22

.02

9.99

.17

.17

42.94

.05

.00

.00

50.58

.36

.03

10.21

.14

.10

38.56

.02

.00

.00

Trace elements (parts per million)

Ba 

Qr ____

cu _

Nb - - 

NJ _____

Sc -

gr _____

y __________ 

Y ___ __ 

Zn ________ 

Zr __________

10 

300

6 

- 2,120 

4 

3 

21 

\ 

70 

4

6

774

5 

- 1,938 

10 

2 

50 

1

70 

4

26 

1,370 

12

6 

1,875 

9 

4 

51 

1 

68 

4

2 

1,061 

15

6 

1,900 

9 

5 

48 

1 

70 

4

4 

1,008

4 

2,196 

9 

2 

47 

1 

69 

4

7 

1,442 

40

4 

1,962 

8 

2 

51 

1 

101 

4

8 

1,388 

17

4 

2,069 

8 

2 

47 

1 

93 

4

with marginal faults and was probably formed during the tectonic emplacement of the peridotite body 

into the upper crust, or during the reactivation of 

those faults at a later period, or both. The presence 

of antigorite rather than lizardite in the mottled and 

sheared serpentinite also implies shear stress during 

formation of the mottled and sheared serpentinite 

and probably a higher metamorphic grade than that 

of the lizardite-bearing banded serpentinite (Coleman, 1977; Evans, 1977). We conclude that the mottled and sheared serpentinite was formed during 

peridotite emplacement at considerable depth by the 

recrystallization of the banded lizardite serpentinite. 

The banded serpentinite predates emplacement and 

was formed at considerable depths in an environment of low directed stress. This early serpentinization affected rocks that already possessed a 

metamorphic fabric and well-developed kink bands.

CHEMICAL COMPOSITION OF 

PERIDOTITE AND SERPENTINITE

Seventeen peridotite samples (dunites and harzburgites) and 20 serpentinite samples have been analyzed for major elements (tables 11.2 and 11.3).

Seven peridotite samples were also analyzed for trace 

elements (table 11.2). The very low concentrations of 

Al2O3 and CaO in the peridotite samples reflect the 

almost total absence of clinopyroxene and plagioclase 

and place these rocks firmly in the MgO corner of the 

MgO-CaO-Al2O3 diagram, in the "metamorphic ultramafic rocks" field of Coleman (1977, p. 45). On an 

AFM diagram (fig. 11.6) the peridotite samples plot 

on the edge of the "metamorphic ultramafic rocks" 

field of Coleman (1977). The analyses support the 

mineralogical and textural evidence that the peridotite of Bullrun Mountain is residual mantle material 

formed by partial melting. The chemical composition 

of the peridotite body is similar to that of peridotites 

from Vourinos, Greece (Moores, 1969), Burro Mountain, Calif. (Coleman and Keith, 1971; Loney and 

others, 1971), Del Puerto, Calif. (Himmelberg and 

Coleman, 1968), and Point Sal, Calif. (Hopson and 

Frano, 1977). The peridotite of Bullrun Mountain has 

significantly less CaO than the clinopyroxene-bearing 

harzburgite of the Canyon Mountain Complex in the 

John Day basin about 70 km west of the Ironside 

Mountain inlier (Thayer, 1977).

Banded and mottled serpentinites (table 11.3) have 

Al2O3 values that imply they were derived from both
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 423

TABLE 11.3. Major-element analyses of serpentinite from Bullrun Mountain and Mine Ridge

[All values in weight percent. Analyses by X-ray fluorescence spectroscopy, Washington State University (Hooper and Johnson, 1989). Major elements 

normalized on a volatile-free basis with 0.9 Fe expressed as FeO, 0.1 Fe expressed as Fe203, and Fe* equal to FeO+Fe2O3]

Banded and mottled serpentinite Contact-metamorphic serpentinite

Sample GC-11 JB-81-74 GC-8 GC-9 JB-81-32 JB-81-45 GC-150 GC-14 GC-15

Sample- GC-17 GC-18 JB-81-33 GC-31 GC-28 GC-29 A-24 A-15a R-03

GC-16

SiO2 -

A1203 --

rr\'f\ liO2 ----- -- 

Fe* 

MnO

CaO- -

MgO - 

K20 -

Na2O 

P2O5

49.15 

.49 

.02 

9.25 

.13 

.09 

40.84 

.03 

.00 

.00

48.22 

.01 

.02 

11.01 

.15 

.08 

40.44 

.02 

.05 

.00

49.74 

.86 

.02 

9.74 

.10 

.03 

39.49 

.02 

.00 

.00

48.60 

.81 

.02 

11.11 

.07 

.06 

39.31 

.02 

.00 

.00

46.66

.74 

.02 

9.76 

.24 

.59 

41.98 

.01 

.00 

.00

Contactmetamorphic 

serpentinite 

Continued

52.18 

.00 

.02 

10.88 

.19 

.33 

36.28 

.03 

.08 

.00

46.15 

.00 

.02 

8.42 

.14 

4.65 

40.46 

.01 

.14 

.00

50.19 

.27 

.02 

8.74 

.12 

1.39 

39.28 

.02 

.00 

.00

48.72 

.00 

.02 

10.22 

.17 

.38 

40.47 

.02 

.00 

.00

49.35 

.38 

.02 

10.21 

.14 

.85 

38.91 

.04 

.10 

.00

Sheared talc-bearing serpentinite

R-61

Si02 

A1203 --

rrv/"\ TiO2 

pe* _____

MnO 

PoO -

MgO 

K20 -------

Na20 

P205 -~ 

48.42

.00

.01

10.73 

OQ

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dunite and harzburgite. In the less serpentinized peridotites, CaO abundances are generally lower and FeO/ 

MgO and SiO2/MgO values are significantly higher, 

suggesting loss of Ca and gain of Fe and Si relative to 

Mg in the serpentinization process (Barnes and O'Neill, 

1969; Coleman, 1977). Most serpentinite from Mine 

Ridge is similar to the banded and mottled serpentinite 

from Bullrun Mountain. Sheared, talc-bearing varieties occur in fault zones, and the more brecciated varieties tend to have higher SiO2 and lower MgO 

concentrations than do the unbrecciated varieties.

The narrow zone of serpentinized peridotite adjacent to the granodiorite intrusion is not noticeably 

different in chemical composition from the banded 

and mottled serpentinite varieties.

The serpentinization process that formed the banded 

and mottled serpentinite appears to have occurred at 

constant volume with loss of MgO, because SiO2/MgO 

is generally higher in the serpentinite than in the peridotite whereas SiO2/FeO is about the same in both 

rock types (fig. 11.7). These relations imply that MgO 

has decreased and SiO2 and FeO have remained unchanged during serpentinization. We conclude that water was added and MgO removed according to the 

following reaction:

2Mg2SiO4 + 2H2O -> Mg3Si2O5(OH)4 + MgO 

(olivine) (added) (serpentine) (removed) 

The scarcity of brucite in the serpentinite supports 

this conclusion, implying that the following reaction, 

which requires an increase in volume (Hostetler and 

others, 1966; Thayer, 1966), is not pertinent: 

2Mg2SiO2 + 3H2O -> Mg3Si2O5(OH)4 + Mg(OH)2 

(olivine) (added) (serpentine) (brucite)

MELANGE OF MINE RIDGE

West of the granodiorite of Bullrun Creek (fig. 11.2) 

the ultramafic rocks form the eastern slopes of Mine
424 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Ridge. From the crest of the ridge westward they are 

increasingly serpentinized, dismembered into smaller 

pods, and mixed with numerous pods of the more resistant ribbon chert, schists, and andesitic "knockers." We here refer to this assemblage as the melange 

of Mine Ridge (fig. 11.8). The ultramafic rocks of the 

melange (dunite, harzburgite, and serpentinite) are 

similar to, and an obvious continuation of, the main 

peridotite body to the east. But they are extensively 

sheared, a process accompanied by the formation of 

talc and substantial amounts of cryptocrystalline silica and magnetite. Serpentinite veins are locally tightly folded.

METAMORPHOSED GABBRO AND BASALT

Metamorphosed gabbro, basalt, and plagiogranite 

have been collected from poor exposures on the northeastern slopes of Mine Ridge (fig. 11.2; Beane, 1984). 

Metamorphosed basalt was found only as loose samples. The metamorphosed gabbro forms rounded boulder ledges that have a weak foliation and are elongate 

in an east-northeast direction. This direction is approximately parallel to the long axes of the knockers 

in the melange and many of the more obvious joints 

and fracture planes. Compositional layering was not 

observed in the gabbro, but wide variations in modal 

mineral proportions in samples from the same outcrop 

imply its presence. Many samples of both gabbro and

FeO

Field of metamorphic 

ultramafic rooks

K,O + Na,O MgO

FIGURE 11.6. Alkali (Na2O+K2O)-FeO-MgO diagram of peridotite at Bullrun Mountain and Mine Ridge. Cumulate and metamorphic fields from Coleman (1977).

basalt retain their igneous texture; others have been 

partly converted to a metamorphic fabric.

Primary mineralogy still recognizable in the metamorphosed gabbro includes cloudy plagioclase feldspar (40-60 modal percent) and augite (20-30 modal 

percent). Augite is partially replaced by green hornblende along margins and cleavages. Primary hornblende, subordinate to augite, occurs in some samples. 

Chlorite is rare. Serpentine minerals are commonly 

intergrown with tremolite and clinozoisite, the last replacing plagioclase. Opaque minerals or their alteration products are rare to absent. Poikilitic textures 

were not observed. Both mineralogy and texture suggest that the metamorphosed gabbro originated as an 

adcumulate.

Basaltic rocks are scarcer than gabbro, and pillow 

structures have not been observed. One sample has 

an aphyric plagioclase-hornblende composition; its 

original intergranular to subophitic igneous texture 

is still apparent, although the pyroxene has been 

pseudomorphed by blue-green hornblende. The metamorphic mineral assemblage is dominated by horn60

CO

O

+ 50

CJ

O

40

A O

70 80 

SiO, x 100

90

+ FeO + Fe 2 O,

EXPLANATION

A Dunite

A Harzburgite

 Serpentinized peridotite

o Banded and mottled serpentinite

FIGURE 11.7. Silicon-magnesium and silicon-iron variation in peridotite and serpentinite of Ironside Mountain inlier.
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 425

blende, sodic andesine, and opaque minerals, and 

also contains epidote, chlorite, and calcite.

Major- and trace-element analyses of gabbro and basalt samples are listed in table 11.4. The adcumulate

nature of the gabbro samples is reflected in the exceptionally low and variable concentrations of incompatible elements (TiO2, K2O, P2O5, Rb, Y, and Zr) in 

contrast to the high but equally variable abundances

R36E 118° 16'00"

44° 21'30"

T14S

500 1000 1500 FEET

150 300 METERS

EXPLANATION

Strawberry volcanics (Miocene)

Intrusions (Oligocene) Most bodies belong to the 

granodiorite of Grouse Creek

Weatherby Formation of Brooks (1979b) (Jurassic)

Andesite (pre-Jurassic)

Ribbon chert (Triassic and Permian)

FIGURE 11.8. Geologic map of melange of Mine Ridge and adjacent 

units. Melange of Mine Ridge includes the andesite, ribbon chert, 

quartzite and mica schist, amphibole schist, and hornblende-albite 

schist units, and the part of the peridotite-and-serpentinite unit

Peridotite and serpentinite (Triassic and Permian) 

Quartzite and mica schist (Triassic and Permian) 

Amphibole schist (Triassic and Permian) 

Hornblende-albite schist (Triassic and Permian) 

Contact 

Fault

that lies west of Mine Ridge. Contour interval, 200 ft. Section numbers for township and range are shown within map area. Base from 

U.S. Geological Survey l:24,000-scale quadrangles: Bullrun Rock, 

1972 (photorevised 1983), and Rail Gulch, 1972 (photorevised 1984).
426 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 11.4. Major- and trace-element analyses ofgabbro, basalt, plagiogranite, and related schists of the melange of Mine Ridge

[Analyses by X-ray-fluorescence spectroscopy, Washington State University (Hooper and Johnson, 1989). Major elements normalized on a volatile-free basis 

with 0.9 Fe expressed as FeO, 0.1 Fe expressed as Fe2C>3 and Fe* equal to FeO+Fe^s , not determined]

Metamorphosed gabbro Metamorphosed basalt and 

andesite

Hornblende schist

Major elements (weight percent)

Trace elements (parts per million)

Plagiogranite

Sample JB-81- 

03

JB-81-

42A

JB-81- 

42B

JB-81- 

82

JB-81- 

BR1

JB-82- 

120

JB-81- A-31 A-52 

86B

JB-81- 

83B

JB-81-

84

Si02 

A1203 -----

Ti02 ---

Fe* ----- -

MnO

CaO - -

MgO -

K20 -

Na2O 

P205~~-

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of such oxides as A12O3, CaO, and MgO, which are 

concentrated in the cumulate phases plagioclase and 

augite. The basalt samples have higher concentrations 

of incompatible elements than the gabbros and presumably have compositions closer to the original liquid 

from which the gabbro cumulates were derived. Both 

basaltic rocks analyzed are quartz normative. It may 

be noted in passing that the two hornblende schist 

samples of the schistose rocks of Mine Ridge (see 

below) have major- and trace-element compositions 

similar to those of the basalt (table 11.4).

PLAGIOGRANITE

A minor summit on the north spur of Mine Ridge is 

underlain by medium- to coarse-grained, equigranular, quartz-rich intrusive rock (fig. 11.2). Most outcrops are deeply weathered. Samples contain equal

amounts of quartz and albite in which opaque minerals, biotite, chlorite, muscovite, and zircon are present 

as accessories. Primary ferromagnesian minerals have 

been largely replaced, but a single crystal of augite, 

rimmed by biotite, was noted. Zones where this plagiogranite is composed of finer grained quartz and albite surrounded by mafic alteration minerals are 

interpreted as due to cataclastic shear. Bent albite 

twin lamellae and strain-induced extinction in quartz 

grains attest to the deformation of these rocks.

Two analyses of plagiogranite are listed in table 

11.4. The analyses reflect the samples' low K2O and 

high Na2O concentrations and so emphasize their 

tonalitic (trondhjemitic) affinities. Their rare-earth 

elements (REEs) (table 11.5; fig. 11.9A) are depleted 

in light rare-earth elements (LREEs) and have chondrite-normalized patterns similar to those of the 

basic rocks of Mine Ridge and the basic and silicic
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 427

TABLE 11.5. Abundances of rare-earth elements in rocks of the Ironside Mountain inlier 

[All values in parts per million. Analyses by inductively coupled plasma spectrometry, Kings College, London. Analyst, P.R. Hooper]

Rock type Peridotite

Sample Brum-A

Hornblende 

schist

A-31

Quartzmica

schist

A-02

Plagiogranite

JB-8183B JB-8184

PreJurassic 

andesite

A-33a

Strawberry 

Volcanics

A-28

Granodiorite of 

Grouse Creek

MH-51

La <0.3

Ce <.5

Pr __ <, K.

Nd <.2

Sm <.l 

Eu <.05

/~1 J ^ O

Dy~ <.l

Uf> ___ ^ nc

T?^ ^ c

Yh ________ < 1

Lu <.Q2

4.45

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9 19

11.68 

4.16 

1.44 

5.74

7

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4 ^9

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4 Q1

21.95 

4.86 

1.17

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rocks of Canyon Mountain (lat 44°20' N., long 118°53' 

W.; Gerlach and others, 1981).

SCHISTOSE ROCKS OF MINE RIDGE

The schistose rocks of Mine Ridge include four 

dominant rock types: hornblende-garnet schist, hornblende-albite schist, quartz-mica-garnet schist, and 

quartzite. Of these the hornblende schist is by far 

the most abundant. All these rock types are poorly 

exposed and form isolated pods in the serpentinite 

matrix of the melange (fig. 11.8). Some pods include 

more than one rock type; consistent lithologic relations in the pods, together with the direction in 

which beds young as defined by graded beds in the 

quartzite, suggest an original stratigraphic sequence 

(Scrivner, 1983). At the base, overlying the ultramafic 

materials, is the hornblende-garnet schist; it is overlain by hornblende-albite schist, quartz-mica-garnet 

schist, and quartzite on top. Contacts, where observed, are typically steep and strike east-northeast; 

the whole sequence apparently is overturned to the 

north.

The modal mineralogy of the schists is listed in 

table 11.6. The hornblende-garnet schist is a dense 

dark-green rock containing aligned elongate hornblende crystals and conspicuous equant garnet porphyroblasts. Sphene is present as inclusions in both 

the major phases. Primary sphene forms fractured 

euhedral prismatic grains with metamorphic euhedral overgrowths. Quartz and epidote fill fractures in 

the garnet.

Hornblende-albite schist contains green layers of 

fine- to medium-grained hornblende alternating with 

white albite layers 2+1 mm thick. The layers are clearly 

metamorphic, paralleling a penetrative cleavage that 

includes the orientation of white mica in other lithologies. Where the boundaries between rock units (S0) are 

observed, the layering and cleavage (S]_) are essentially 

parallel to them. Hornblende grains are also aligned 

in this plane (So-Sj). Garnet is absent, and the albite 

has a dusting of sericite and inclusions of epidote.

The quartz-mica-garnet schist and the quartzite 

tend to occur together with gradational contacts. The 

schist is a well-foliated, buff-colored rock in which 

orientated muscovite grains cause the penetrative 

cleavage (S]). Quartz forms layers (So-Sj) that pinch 

and swell between muscovite layers, the latter including pale pink garnet porphyroblasts (table 11.6). 

Epidote porphyroblasts and elongate grains of euhedral zircon are associated with mica-garnet-hornblende layers.

The quartzite is 95 percent quartz with wavy extinction and ragged crystal margins. Muscovite, 

sphene, and garnet occur as accessories closely associated with each other; the mica defines a cleavage 

cut at a high angle by quartz veins.

The quartzite contains graded beds and clearly had 

a sedimentary origin. The hornblende schists, in contrast, have basaltic compositions (table 11.4), and these 

rocks probably represent highly metamorphosed and 

deformed igneous rocks, rather than calcareous sediments as suggested by Lowry (1968). The quartz-micagarnet schist has an intermediate composition that
428 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

may reflect an origin from more evolved igneous rocks 

or, conceivably, from physical mixing between the basaltic units and the overlying psammitic material during intense metamorphism and deformation.

The various schistose units are tightly folded, and 

three fold phases are apparent in some outcrops 

(fig. 11.10). The metamorphic foliation and associated 

mineral alignment represent the earliest penetrative

100 100

Pr Nd Sm Eu Gd Tb Dy Ho Er Yb Lu La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Yb Lu

200

100 -

La Ce Pr Sm Eu Gd Tb Dy Ho Er Yb Lu La Ce Pr Nd 

RARE-EARTH ELEMENTS

Sm Eu Gd Tb Dy Ho Er Yb Lu

FIGURE 11.9. Chondrite-normalized rare-earth-element (REE) 

concentrations in igneous rocks of Blue Mountains province. A, 

Basic (open circles) and granitic (crosses) rocks of melange of 

Mine Ridge. Basic rocks (shaded area) and plagiogranite and 

keratophyre (diagonal lines) from Canyon Mountain Complex 

(Gerlach and others, 1981). B, Sparta complex of Phelps (1979). 

Solid lines, this paper; dashed lines, Phelps (1978). Dots, plagiogranite; x's, diorite; open circles, gabbro. C, Representative 

samples from Seven Devils Group (Hooper and Scheffler, unpub. data, 1983). Dots, Permian; x's, Triassic. D, REE concentrations in pre-Jurassic andesite of melange of Mine Ridge (triangles) 

and other calc-alkaline volcanic rocks of eastern Oregon. Open 

circles, andesite of Strawberry Volcanics, Mine Ridge; x's, Permian 

Hunsaker Creek Formation, Seven Devils Group; crosses, andesite of Strawberry Volcanics near Baker City, Oreg.; dots, hornblende schist of melange of Mine Ridge; diagonal lines, Triassic 

rocks of Seven Devils Group; vertical lines, Permian rocks of Seven Devils Group (fig. 11.9C).
11 GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 429

TABLE 11.6. Modal mineralogy of the schistose rocks of Mine Ridge 

[All values in volume percent, tr, trace; , not observed]

Hornblendegarnet schist

Sample A-31

Quartz 

Tj| Q/YlfkfM QOO

Hornblende 

Orthoclase 

Muscovite 

Almandine 

Zircon 

Ql 1

Pyrite 

Epidote 

Opaque minerals 

tr 

tr 

80

5 

tr 

tr

tr 

tr 

15 

tr

Hornblende-albite schist

Sample A-52 Average 

(9 samples)

tr 

50 

tr

tr 

tr

tr 

20 

tr

0^35

35-70 

0-5 

0-5

0-1 

tr 

0-10 

0-1 

0-30 

0-20 

tr 

0-5

Quartz-micagarnet schist

Average 

(2 samples)

5-40 

0-10 

0-25

20-30 

5-20 

0-10 

tr

0-15 

tr 

tr

All plagioclase in these samples is albite.

fabric (S]), with which the original premetamorphic 

lithologic boundaries (S0) have been brought into parallelism (SQ-SX). The earliest folds recognized in outcrop 

(F]) have a steeply dipping, predominantly east-northeast-striking penetrative axial-planar cleavage (S]). 

The FI folds are isoclinal with axes (L^) approximately 

parallel to the dip of the axial planes (figs. 11.11, 

11.12A). The steep SQ-SI plane is ubiquitous throughout the schistose rocks of Mine Ridge.

Folds of the second phase (F2) vary in style from 

tight folds with thickening of layers around the hinge 

zones (fig. 11.10) to open box folds with steep but 

variable S2 (fig. 11.11). S2 normally cuts across SQ-S! 

at moderate to large angles, but L! and L2 are approximately parallel (fig. 11.12C). The typical interference pattern of F2 and FI is shown in fig. 11.11C. 

The main fabric (S0-Si) is also crenulated by F3 folds 

(fig. 11.12D) whose axial planes lie at a high angle to 

SQ-S-L. In thin section, most minerals are seen to be 

bent by F3. Chlorite fills fractures in epidote perpendicular to the SQ-S! surface and is orientated parallel 

to S3. Evidence of slip along some 83 surfaces includes fracturing of pyrite developed in the slip 

planes where the pyrite crystals are also flattened 

and partly altered to goethite.

FIGURE 11.10. Three fold phases in hornblende-albite schist from 

melange of Mine Ridge. Sj, S2, and S3 are axial-planar surfaces 

produced, respectively, by F1; F2, and F3. Sx (to which S0 is now 

parallel) is tightly folded by F2 and crenulated by an irregularly 

oriented F3. Scale in inches.

RIBBON CHERT

If the apparent stratigraphic sequence of the 

schistose rocks of Mine Ridge, discussed above, is real, then the ribbon chert lies above the schists, but
430 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

the nature of the melange prevents this relationship 

from being demonstrated unequivocally. The chert is

A fold

B F2 fold

F2 axial planes 

(variable trends)

F fold axis

F2 fold axis(L 2)

FIGURE 11.11. Diagram to illustrate relations between Fj and F2 

folds in melange of Mine Ridge. A, Bedding (So) folded by Fj. B, 

Si axial-planar surfaces folded by F2. S2 axial-planar surfaces 

have variable orientations. C, Although Fj and F2 axial planes 

generally intersect at moderate to large angles, F± and F2 fold 

axes are approximately parallel.

composed of alternating layers of fine-grained chocolate-brown chert and argillite, 1-3 cm thick. Finegrained quartz forms about 95 percent of the chert, 

and carbonate forms up to 5 percent. Despite careful 

search no radiolarians have been found. The cherts 

are similar to those described from other parts of the 

central melange terrane of northeastern Oregon 

(Elkhorn Ridge Argillite, cherts of Vinegar Hill, and 

so on; Gilluly, 1937; Ashley, 1966; Mullen, 1979).

The layers are folded by a single set of overturned, 

moderately open to close, asymmetric chevron folds. 

Axial planes strike east-northeast and dip south at 

moderate angles; the fold axes plunge southeast at 

moderate angles.

The single set of folds lacks an axial-planar cleavage. The folds in the chert differ in style from both

 Poles to S0 

O L,

X Poles to S-|

A Poles to axial 

planes of 

specific F1 folds

Poles to S0 O Fold axes

Poles to X Pole to a single

axial plane foliation plane

Poles to S2 Poles to Fg 

crenulations

O Fold axes of 

Fg crenulations

FIGURE 11.12. Lower-hemisphere, equal-area projections of structural data for the schistose rocks of Mine Ridge. SQ, bedding 

planes; Sj, foliations axial planar to first-generation folds (F^); 

Lj, FI fold axes; Bl, poles to axial planes of specific F± folds; 

S2, axial planes to second-generation folds (F2); L2, F2 fold axes; F3, third-generation folds. A, Bedding planes and Fj folds in 

schistose rocks of Mine Ridge. B, Single fold set in ribbon 

chert. C, F2 folds in schistose rocks of Mine Ridge. D, Crenulations (F^) in schistose rocks of Mine Ridge.
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 431

TABLE 11.7. Major- and trace-element analyses of pre-Jurassic volcanic rocks, quartz-mica-garnet schist, chlorite schist, and ribbon

chert from the melange of Mine Ridge

[Analyses by X-ray fluorescence, Washington State University (Hooper and Johnson, 1989). Major elements normalized on a volatile-free basis with 

0.9 Fe expressed as FeO, 0.1 Fe expressed as Fe2O3; and Fe* equal to FeO+Fe2C>3]

Pre-Jurassic volcanic rocks

Quartzmicagarnet 

schist

Chlorite

schist

Ribbon

chert

Sample A-15 AD-01 AD-04 A-33a A-02 A-08a RC

Major elements (weight percent)

SiO2 - 

A1203 ------ 

Ti02 -----

Fe* -

MnO -

PnO

MgO -

K20 ------

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poO^ " -£ VX O 

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9 48

1.71

1 no

IQ

52.43

6

(\C

Kfl

8

I7C

OQ

9

qo

01 SQ

.04

7K

.04

96.01

2 1 Q

nq

.41

.07

.00

44

en

.17

.01

Trace elements (parts per million)

Do.. _____

Pr

Pn

MV, __________

Ni 

Q-.

V 

Y 

Zr 

A QQ

4QQ

1 ^

q7

591

204

17

70

50

QKO

CQ

74

14

q/5

qn

240

OQ

100

51

480

qi

10

54

OQ

Qfi

ono

24

7Q

34

4QQ

104

C-l

Q7

Oq

204

on

84

118

64/?/?

10

Q7

00

144

QO

Q5U*J

26

10

£on

q

QOC

on

q

1 7C

12

70

11

134

34

OQ

0

iq

q

11

iq

4

iq

16

F! and F2 folds in the schists; these differences probably reflect the chert's thin alternating layers of contrasting competence rather than a different stress 

regime.

PRE-JURASSIC VOLCANIC ROCKS

The melange of Mine Ridge also includes large, 

relatively coherent blocks of massive, dark-green, 

fine-grained volcanic rocks. The blocks are elongated 

east-west to northeast-southwest, parallel to the 

gross structural trend of the melange. Most are 

aphyric flows, although some phenocrysts of olivine 

and augite, and more rarely plagioclase, are present. 

In the stratigraphic framework described above, these

rocks appear to lie stratigraphically above the schists, 

but their age relative to the ribbon cherts is obscure.

These predominantly andesitic volcanic rocks are 

surprisingly fresh. They lack the penetrative deformation of the schistose rocks of Mine Ridge, but contain many discrete shear planes. Original textures 

are generally well preserved but are modified locally 

by a metamorphic fabric that includes mortar textures and triple junctions between grains. Plagioclase 

(Anso+io) is bent and altered to sericite and epidote.

Chemical analyses (table 11.7) show that the volcanic 

rocks range in composition from olivine basalt, through 

andesite, to (rare) rhyolite. Up to 30 modal percent 

hypersthene may be present. Olivine and augite dominate in some samples, hypersthene and pigeonite in
432 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

others. Opaque minerals are relatively scarce. Secondary minerals include quartz, epidote, and hornblende, 

and traces of chlorite, sericite, sphene, calcite, and 

goethite. The rhyolite has a cataclastic texture and may 

have formed from a silicic protolith or may be derived 

from a more mafic parent silicified in the metamorphic 

process.

It is not clear whether these volcanic rocks have 

suffered less deformation than the schistose rocks of 

Mine Ridge, or whether they have been shielded from 

the same deformation by their more massive form. 

The number of analyses available is not adequate to 

fully characterize the group, but the wide range in 

major- and trace-element compositions, their LREE 

enrichment (fig. 11.9Z)), and their original mineralogy 

suggest a calc-alkaline affinity.

ORIGIN OF THE ROCKS OF THE BAKER TERRANE, 

IRONSIDE MOUNTAIN INLIER

The chemical composition, the mineralogy, and the 

texture of the dunite and harzburgite that constitute 

the peridotite body forming the north end of Bullrun 

Mountain all support the thesis that these rocks represent the residuum of partial melting in the mantle. 

Serpentinization of the peridotite to form banded and 

mottled serpentinite occurred without production of 

significant brucite, but with loss of magnesium and 

consequently little change in volume. This process occurred at depth in peridotite that had already been 

deformed. Subsequent injection of the partially serpentinized body into the upper crust along major 

fault planes and the formation of the melange were 

accompanied by additional serpentinization involving 

severe shear and an increase in SiC>2.

The much smaller volumes of cumulate gabbro, basalt, and plagiogranite are best explained as crystallization products derived from the melt produced in 

the same partial melting process. Pillow basalts and 

a basic dike swarm are missing from the complex, as 

they are from both the Canyon Mountain Complex 

(Thayer, 1977) and the (informal) Sparta complex of 

Phelps (1979) and Almy (1977). Their absence from 

all three complexes probably means that they were 

never formed. Earlier workers (for example, Gerlach 

and others, 1981) use the absence of pillow basalts 

and basic dike swarms as evidence that the ophiolitic 

rocks of northeastern Oregon originated at the base 

of an island-arc system rather than at a midocean 

ridge. This argument is equally applicable to the 

rocks of the Ironside Mountain inlier.

Chondrite-normalized patterns of the rare-earth elements (REEs) of the Ironside Mountain ophiolitic 

rocks are particularly revealing when compared to

those of the Canyon Mountain Complex and the Sparta 

complex (fig. 11.9). The REE patterns for these two 

complexes are fundamentally different from each other. 

All samples from the Canyon Mountain Complex, gabbro through basalt to plagiogranite and keratophyre, 

have essentially similar patterns of light-rare-earthelement (LREE) depletion, usually accompanied by a 

well-marked negative europium anomaly (fig. 11.9A).

The REE data from the Sparta complex (Phelps, 

1978; PR. Hooper, unpub. data, 1984; fig. 11.95) appear to divide these rocks into two groups: (1) gabbro 

with LREE depletion and a positive europium anomaly 

and (2) diorite and plagiogranite with mild LREE enrichment and a strong negative europium anomaly. It 

is possible that the two patterns represent cumulates 

and residual magmas of a parent with an almost flat 

REE pattern at a value ten times the chondrite value. 

Alternatively, the two patterns may represent the 

products of two mantle sources with different degrees 

of LREE depletion. Whichever interpretation is favored, both patterns are distinct from the single pattern displayed by the rocks of the Canyon Mountain 

Complex. Three chondrite-normalized REE profiles 

from Mine Ridge, one for a basic rock and two for 

plagiogranite, clearly resemble the pattern from the 

Canyon Mountain Complex (fig. 11.9A). We conclude 

that the ophiolitic rocks of the Ironside Mountain inlier 

are most likely the lateral equivalents of the Canyon 

Mountain Complex and may, therefore, be assumed to 

have the same Late Permian age.

The schistose rocks of Mine Ridge include rocks of 

basic igneous origin, comparable to nonschistose metamorphic basalts of the ophiolite complex, and rocks 

of sedimentary origin such as the quartzites and 

probably the muscovite schists. The relationship between these metamorphic rocks and the less metamorphosed and deformed chert and andesite of the 

melange of Mine Ridge is not clear. The ribbon chert 

requires a relatively deep oceanic environment remote from a clastic source for its formation. The andesite has a LREE-enriched, typically calc-alkaline 

REE pattern (fig. 11.9Z)) similar to patterns observed 

in the Strawberry Volcanics and in a few samples of 

the Seven Devils Group. This pattern is distinct from 

the relatively flat REE pattern of the majority of both 

Permian and Triassic Seven Devils rocks, which closely resembles the pattern of the ophiolitic rocks of 

Mine Ridge (figs. 11.9C, 11.9Z)). Rare-earth-element 

patterns for island-arc suites vary from LREE-depleted in the earlier, more primitive island-arc tholeiites 

to LREE-enriched in the more typical calc-alkaline 

suites, so the striking differences in pattern observed 

in figure 11.9D need not imply two distinct tectonic 

settings. Indeed all these rocks have characteristics of
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 433

rocks found in primitive volcanic island arcs. But the 

REE differences, when added to the different degrees 

of deformation observed, reinforce the possibility that 

the coherent blocks of andesite and the ophiolites represent two suites of rocks of different age.

CORRELATION AND TIMING OF THE DEFORMATION OF 

THE SCHISTOSE ROCKS OF MINE RIDGE

All mesoscopic structures observed in the blocks of 

schist within the melange on Mine Ridge are correlated with the earlier (Dj) of two major deformations 

recognized throughout the Baker and Izee terranes 

by Ave Lallemant and others (1980). Dj on Mine 

Ridge, as elsewhere, is a relatively high grade metamorphic event (epidote amphibolite facies) that developed a strong axial-planar foliation (Sj) with isoclinal 

folds (FI) whose axes plunge steeply downdip. The 

isoclinal folds bring bedding planes (So) into parallelism with B! (So-Sj) except at fold hinges. SQ-SJ is 

characterized by oriented mica plates and elongate 

hornblende crystals aligned within SQ-SJ and parallel 

to the FI fold axes. On Mine Ridge SQ-SJ strikes east 

to northeast, and the fold axes (L^) plunge steeply 

south (fig. 11.12).

The second phase of more open, commonly boxshaped folds is approximately coaxial with the first 

phase, but the second-phase folds have steep axial 

planes at a high angle to So-S^ A third fold phase 

within the Dj event formed crenulation cleavage. The 

sequence of fold phases can be interpreted as progressive, noncoaxial shear deformation that occurred 

during, and subsequent to, the development of the 

dominant metamorphic fabric (Si).

The FI folds and associated mineral fabrics must 

have developed prior to the formation of the serpentinite melange matrix. Once the matrix formed, subsequent strain would be absorbed by the incompetent 

envelope the matrix leaving the more competent 

blocks of the schistose rocks of Mine Ridge to react 

only passively. If deformation developed on shear 

planes parallel to the plane of subduction (Dickinson 

and Thayer, 1978), then horizontal fold axes would 

develop about axial surfaces parallel to the subduction plane, at lower dips than are now seen. Thus, 

reorientation of the F^ folds is required to achieve 

the present uniformly steep plunge of the FI fold 

axes down the dip of the FI axial surfaces. In a subduction zone, in which extension parallel to the direction of subduction is likely to be extreme on the 

upper surface of the subducting slab, rotation of early folds is not only feasible but almost unavoidable. 

All early fold axes would be transposed into the direction of maximum extension, down the dip of their

axial surfaces. Transposition of this type would be 

greatly facilitated as serpentinites were incorporated 

into the shear system to form a melange. This is the 

model outlined by Ave Lallemant and others (1980).

The east-to-northeast strike of SQ-&I on Mine Ridge 

and elsewhere in the Baker terrane parallels the terrane boundaries, and the associated FI fold axes, 

which plunge down the dip of these nearly vertical 

axial surfaces, are also similar to the regional pattern. The pattern implies a subduction zone striking 

approximately east-west with subduction either toward the north, under the arc rocks of the Wallowa 

terrane, or toward the south, under the arc rocks of 

the Olds Ferry terrane. Wilson and Cox (1980) postulated 60° clockwise rotation of the Wallowa terrane in 

post-Jurassic time; if this rotation extended as far 

south as Ironside Mountain, then the direction of 

subduction would have been N. 60° W. or S. 60° E. 

These subduction directions, however, assume that 

the Dj structures were not reoriented during the D2 

event. As discussed later in the section "Deformation 

of the Weatherby Formation," reorientation during 

D2 is probable, so these apparent directions of subduction may have little meaning.

The present attitude of S0-Si and the FI fold axes 

is steeper than is required in a subduction model, 

but is readily explained by the reorientation that appears to have accompanied the second regional deformation (D2).

A recent K-Ar age on hornblende from the schistose rocks of Mine Ridge is 260 Ma (N.W. Walker, 

oral, commun., 1987). The Dj event is, therefore, 

Late Permian rather than Late Triassic in age (Ave 

Lallemant and others, 1980). This age is similar to 

the 262-Ma age determined for the Canyon Mountain 

Complex by Ave Lallemant and others (1980) and 

precludes the possibility of that ophiolite complex 

forming in an oceanic ridge-rift system far from the 

subduction zone.

FORE-ARC BASIN DEPOSITS: 

THE WEATHERBY FORMATION

A broad east-west zone of Paleozoic to Mesozoic 

clastic sedimentary deposits, the Izee terrane, lies 

across northeastern Oregon, south of the Baker terrane and northwest of the Olds Ferry terrane (Huntington arc; fig. 11.1; Brooks and Vallier, 1978; 

Dickinson and Thayer, 1978). Dickinson (1979) suggests that these sedimentary materials were deposited in a fore-arc basin between a ridge of ocean-floor 

melange (Baker terrane) to the northwest and a rising volcanic arc (the Huntington arc of the Olds 

Ferry terrane) to the southeast.
434 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

In the John Day basin on the west end of the zone 

(fig. 11.1), Dickinson (1979) and Dickinson and others 

(1979) have described a series of clastic units that 

range in age from Karnian (Late Triassic) to Cretaceous. Farther east, from Strawberry Mountain (about 

20 km southeast of the town of John Day; Robyn, 1977) 

to the Snake and Salmon River canyons of western 

Idaho, a less well known but apparently similar sequence of metasedimentary turbidites is exposed. 

These constitute the Weatherby Formation of Brooks 

(1979b). Crinoids, brachiopods, molluscs and ammonites distributed sparsely throughout these rocks date 

them as Early (Sinemurian) to Middle (late Bajocian) 

Jurassic (200-170 Ma) in age (Brooks and Vallier, 

1978). In the Snake River area the Weatherby includes 

clasts apparently derived from the Huntington arc 

(Olds Ferry terrane, fig. 11.1, inset), which the Weatherby appears to overlie unconformably (Brooks, 1979b). 

Farther north near the Snake River the Weatherby is 

in fault contact with rocks of the Baker terrane (Connor Creek fault; see Brooks and Vallier, 1978).

The Weatherby Formation is the most extensive 

unit of the Ironside Mountain inlier (Lowry, 1968; 

Brooks and Ferns, 1979; Brooks and others, 1979). It 

is strongly folded and is intruded by a series of Cretaceous to Oligocene diorite to granodiorite stocks 

along the anticlinal axis of the inlier.

The Weatherby Formation in the inlier consists of 

olive-gray to gray-brown lithic wacke, lithic arenite, 

brown-gray slate, light-brown-gray silty limestone, 

and limey siltstone (Engh, 1984). Typically, its beds 

range in thickness from 25 cm in the coarser units to 

5 cm in the finest pelitic units. Close to the Permian 

and Triassic rocks of the Baker terrane at the north 

end of the inlier, wacke, siltstone, and conglomerate 

predominate; minor horizons of arkose and breccia 

are also present. Large angular chert fragments up 

to 4.0 cm across occur in the breccias in a chert matrix. The conglomerate contains smaller angular fragments of chert, volcanic rock, and serpentinite. The 

Weatherby is in fault contact with the rocks of the 

melange of Mine Ridge, but the conglomeratic horizons and clasts derived from the melange and the 

presence of the Weatherby on both the north and 

south sides of the melange (fig. 11.1) suggest that its 

sediments were originally laid down unconformably 

on, and were largely derived from, rocks of the melange. No useful fossils have been found in the 

Weatherby Formation during the present study, but 

H.C. Brooks (oral, commun., 1983) has observed 

limestone pods with abundant but poorly preserved 

fossils that are tentatively assigned a Jurassic age.

Sedimentary structures are well developed in the 

turbidites. Among these structures are graded bedding (including some reverse grading), small ripples, 

load casts and flame structures, and wavy or convolute laminae. Rip-up clasts are common. Together the 

structures provide abundant evidence for identifying 

facing directions, that is, the directions in which beds 

young. In general, the beds fit the Bouma sequence 

model (Middleton and Hampton, 1976), although a 

complete Bouma sequence is rare.

The Weatherby Formation is strongly folded about 

a single axial-planar surface represented by a penetrative cleavage striking east-northeast and dipping 

steeply south. The consistent relations between facing directions and fold limbs, as defined by the cleavage-bedding relationship, and the lack of interference 

structures imply that the folds belong to a single episode of deformation. The effects of low-grade regional 

metamorphism are apparent throughout the formation. Chlorite, stilpnomelane, and quartz form pressure shadows. Calcite and albite are also present. 

Absent are zeolites, prehnite, pumpellyite, lawsonite, 

and any amphibole or pyroxene.

Contact metamorphism and mineralization occurred around the granodiorite of Grouse Creek. 

These processes obscured the cleavage in some areas 

and obliterated it in others, and they were accompanied by the formation of calc-silicate minerals. Close 

to the granodiorite body, garnet and epidote are 

present, disappearing approximately 450 m from the 

contact. Outside the garnet-epidote zone a spotted 

cordierite hornfels is found locally; the hornfels contains up to 30 percent radiating cordierite crystals in 

a matrix that appears to contain both andalusite and 

anthophyllite.

PROVENANCE

Lithic arenite and lithic wacke make up approximately 75 percent of the Weatherby section in the 

Rastus Mountain area. They contain 10-60 percent 

volcanic rock fragments (average 45 percent). Feldspar, quartz, chert, and detrital quartzite grains are 

the other major components, forming about 10 percent 

each. The lithic wacke contains greater than 15 percent of very fine grained matrix, but matrix estimates 

are complicated by the low-grade metamorphism.

Pelitic layers are composed of 45 percent sheet silicates, commonly stilpnomelane, and 30 percent or 

more very fine grained microcrystalline quartz. The 

remaining 25 percent consists primarily of lithic fragments and feldspar. The silty limestone has an average 

composition of 75 percent calcite, 7 percent volcanic 

rock fragments, 3 percent mica and chlorite, and about 

15 percent organic material and fine clay minerals. The
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 435

TABLE 11.8. Petrographic analyses of the Weatherby Formation 

[All values except "Average grain size" and "Total framework grains counted" are in volume percent. , not observed; tr, trace; n.d., not determined]

Volcanic fragments 

Chert and quartzite 

Sedimentary rock

ITT?

29

32 

14

qc

1

1

9

290

22 

5

CO

K

1

7

30

25 12 

2 

on en

2

0

10

1

30 tr

W q

52 

12 

17

19

7

14

30 

10 

40

10

4

q

4

Vf? 7

58 

5

4

9

10

vu* a

10 

70

10

q

7

PP

10 

10

en

7

4

1

7

PP

100

55 

3

14

1 q

T P

207

30

9fi

9

48

tff?

207B

20

40

40

Average grain size (urn) <1 <1 <1 0.5 1.5 >4 >4

Proportions used in figure 11.13

Quartz (Q)»- - 24

Feldspar (F) 9

Monocrystalline quartz 

(OtTO _____________ <i

Feldspar (F) - - 9

Total lithic fragments 

(Lt-Qp+Lv+Ls) 86

Polycrystalline quartz 

(Qp) --- - 22

Lithic volcanic fragments 

(Lv) 56

Lithic sedimentary 

fragments (Ls) 22

20 21 n.d. 5 12 8 n.d. 31 n.d. n.d. n.d. 

29 5 n.d. 4 19 7 n.d. 10 n.d. n.d. n.d. 

51 74 n.d. 91 69 84 n.d. 58 n.d. n.d. n.d.

7 1 n.d. 131 n.d. 4 n.d. n.d. n.d. 

29 5 n.d. 19 19 7 n.d. 10 n.d. n.d. n.d.

64 95 n.d. 80 78 92 n.d. 86 n.d. n.d. n.d.

19 22 n.d. 4 11 8 n.d. 32 n.d. n.d. n.d. 

75 59 n.d. 91 80 88 n.d. 52 n.d. n.d. n.d. 

6 20 n.d. 594 n.d. 16 n.d. n.d. n.d.

Total framework grains

counted 308 232 253 n.d. 307 233 327 n.d. 178 n.d. n.d. n.d.

lithology is summarized in table 11.8 and plotted in 

figure 11.13.

On the quartz-feldspar-lithic (QFL) diagram (fig. 

11. ISA) all the Weatherby rocks plot near the lithic 

pole; they are fairly similar to the clastic sediments 

of the John Day inlier (Dickinson and others, 1979), 

but have less feldspar. Plots of the Weatherby rocks 

lie within the transitional arc and undissected arc 

provenance fields as defined by Dickinson and others 

(1983). The inferred main source for these sediments 

is a volcanic arc; subordinate debris was derived 

from flanking metamorphic rocks and their sedimentary cover.

The monocrystalline quartz-feldspar-total lithic (QmFLt) diagram (fig. 11.13C) emphasizes provenance. Samples of the Weatherby Formation plot

near the lithic pole (Lt) within the field of a melange 

source, similar to samples from the Vester Formation 

of the John Day inlier (Dickinson and others, 1979). 

A single exception falls within the volcanic source 

field. On the polycrystalline quartz-lithic volcaniclithic sedimentary (Qp-Lv-Ls) diagram (fig. 11.135), 

which includes only polycrystalline grains, most samples plot near the volcanic pole in the magmatic-arc 

source field (Dickinson and Suczek, 1979).

Some limitations in the application of these diagrams to the determination of the provenance of the 

Weatherby Formation need to be stressed. Both chert 

and quartzite fragments are recorded together as a 

polycrystalline quartz. Quartzite makes up 50 percent 

or more of these fragments in some samples. Quartzite is one of the components of the older melange of
436 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Cl

Om

Ls F 

EXPLANATION

Samples from the Ironside Mountain inlier

Modified from Dickinson and others (1979):

V Vester Formation field

K Cretaceous sandstone field

L Lonesome Formation field

J Jurassic volcaniclastic field

m Field of detritus derived from melange

vol Field of detritus derived from volcanic rock

FIGURE 11.13. Composition and provenance of Weatherby Formation of Brooks (1979b), Ironside Mountain inlier. A, Ternary plot of quartz (Q), feldspar (F), and lithic (L) components. 

B, Ternary plot of polycrystalline quartz (Qp), volcanic lithic 

(Lv), and sedimentary lithic (Ls) components. Dashed line

Fields of tectonic setting from

Dickinson and others (1983):

Cl Craton interior

TC Transitional continental

RO Recycled erogenic

BU Basement uplift

DA Dissected arc

TA Transitional arc

DA Undissected arc

M Mixed

OR Ouartzose recycled

TR Transitional recycled

LR Lithic recycled

surrounds field of arc-orogen sources of Dickinson and Suczek 

(1979). C, Ternary plot of monocrystalline quartz (Qm), feldspar 

(F), and total lithic (Lt) components. Ten samples plotted in 

each diagram; some samples overlap in figures 11. ISA and 

11.13C.
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 437

Mine Ridge, but the relative abundance of quartzite 

fragments in the Weatherby could be construed as evidence of another, more continental source component. 

The low-grade metamorphism may also bias the interpretation of these diagrams by altering grain size 

and mineralogy. Such alteration may explain in part 

the low feldspar count. Small feldspar grains are commonly corroded and replaced by calcite and clay minerals, resulting in original feldspar grains being 

counted as matrix. Despite these limitations, the 

rocks of the Weatherby Formation appear to have 

been derived from sources whose origin and tectonic 

setting are similar to those of the John Day inlier. 

The Weatherby was probably deposited in an arctrench gap or fore-arc basin; the sediments were most 

likely derived from a volcanic arc underlain by metamorphic terrane, although some material (quartzite) 

may have been derived from a continental source.

DEFORMATION OF THE WEATHERBY FORMATION

The Weatherby Formation of the Ironside Mountain inlier is dominated by a well-developed penetrative cleavage striking east-northeast. Bedding is 

preserved in many places but is less evident than 

cleavage. Use of cleavage/bedding relations permits 

an accurate reconstruction of minor and major folds 

(figs. 11.14-11.16), although actual folds are rarely 

seen in outcrop.

Critical structural data are plotted in figure 11.15. 

The poles to the cleavage surface (fig. 11.15C) are 

more scattered than might be expected from a single 

strong episode of folding in a relatively homogeneous 

lithology. This scattering is due to refraction of the 

cleavage across beds of varying competence; dips of 

cleavage planes around a single fold vary by as much 

as 20°. Only the cleavage in the pelitic horizons is 

truly axial planar.

Fold style is illustrated in figure 11.16. One of the 

rare folds observed in the field is shown in figure 

11.16A. A reconstruction of the fold style based on 

cleavage/bedding relations and facing-direction data 

is shown in figure 11.16J3. The folds are strongly 

asymmetric and have a similar or chevron style. The 

axial plane lies close to the more southerly limb, and 

beds thicken over the fold hinges.

The axial-planar cleavage and the two fold limbs 

maintain relatively constant orientations across the 

mapped area, but the sense of asymmetry, or the relative length of one limb compared to the other, varies 

systematically and permits the identification of major 

fold hinges (fig. 11.14).

An analysis of strain in the Weatherby Formation 

using elliptical microlitic glass grains (Dunnet, 1969; 

Dunnet and Siddans, 1971) is shown in figure 11.17.

The strain occupies the oblate field of a Flinn diagram (fig. 11.17J5). Some volume loss could well have 

occurred, as evidenced by pressure-solution features 

in the cleavage planes, and the strain shape may, 

therefore, reflect components of both tectonic strain 

and volume loss. The flattening occurs with maximum shortening perpendicular to the axial surface 

and maximum elongation down the dip of the axial 

surface; the intermediate axis lies approximately parallel to the fold axis. Boudinage structures and 

crinoid fragments in suitably oriented two-dimensional sections (Engh, 1984) suggest a maximum extension of 35 percent down the dip of the axial plane 

and maximum shortening of 50 percent perpendicular to that plane.

Joints, dikes, and veins show little evidence of preferred orientations, but slickensides on a number of 

small faults, which dip at a slightly lower angle to 

the south-southeast than the fold axial planes, indicate oblique slip toward the southeast.

We consider the post-Jurassic folds in the Weatherby Formation to be the products of the second (02) 

episode of regional deformation recognized by Ave Lallemant and others (1980). The lack of evidence of 

the D2 deformation within the much older competent 

blocks of the schistose rocks of Mine Ridge is surprising, although it is possible that the F3 crenulation 

cleavage reflects the Late Jurassic (D2) event. The 

explanation, we assume, lies in the concept that once 

the melange was formed, all subsequent strain suffered by the melange, including that undergone during D2, was taken up by the incompetent matrix of 

serpentinite. The competent blocks only rotated passively, undergoing little internal deformation, but the 

D2 rotation brought the SQ-SX surfaces into near parallelism with the D2 axial surface and steepened the 

axes of the Fj folds.

The predominantly upright orientation of the major 

folds in the Weatherby Formation and their nearsymmetrical form suggest that the older Permian 

and the younger Jurassic rocks were flattened 

against each other so that the direction of maximum 

extension was close to vertical. Some small component of thrusting toward the northwest may have 

been the cause of the asymmetric distribution of the 

cleavage surfaces in the Weatherby folds of the D2 

deformation (fig. 11.165).

POST-JURASSIC IGNEOUS ROCKS

South of Ironside Mountain the diorite of Tureman 

Ranch intruded the Weatherby Formation in the Cretaceous (about 120 Ma) (Thayer and Brown, 1964). 

Within the mapped area of the inlier (fig. 11.1), early 

Oligocene intrusions were emplaced over a period of
438 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

0.5 1 MILE

1 KILOMETER

EXPLANATION

Sedimentary deposits 

(post-Tertiary)

Granodiorite of Bullrun Creek 

(Oligocene)

Granodiorite of Grouse Creek 

(Oligocene)

Weatherby Formation of Brooks 

(1979b) (Jurassic)

x' \ :-'.-j Hydrothermal alteration zone 

(pyrite halo) North End after 

Brooks and others (1979)

Phyllitic alteration zone 

Potassic alteration zone

Argillic Single area of argillic alteration

Post-alteration intrusion 

Contact

Major fold axes in Weatherby 

Formation

Anticline

Syncline

Minor fold axis, showing plunge 

Strike and dip of beds 

Strike and dip of cleavage

44° 18' 00"

FIGURE 11.14. Map of Ironside Mountain inlier, showing bedding/cleavage relations and major fold hinges in Weatherby Formation, 

and zones of hydrothermal alteration in and around granodiorite of Grouse Creek. Contour interval, 200 ft. Base from U.S. 

Geological Survey 1:24,000, Rastus Mountain quadrangle, 1972 (photorevised 1984).
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 439

Density of data points:

15 percent g 2 percent

10 percent Q 1 percent

Density of data points20 percent g 5 percent

10 percent [~] 1 percent

Density of data points 

| 8 percent g 2 percent 

[HI 5 percent

FIGURE 11.15. Lower-hemisphere, equal-area projections of structural data for Weatherby Formation of Brooks (1979b). A, Poles 

to bedding. Number of data points (n)=115. B, Contour diagram 

of poles to bedding. Density reported in percentage of total data 

points per 1-percent area of diagram. Open circle, minor fold 

axis (measured); x, fold axis (estimated). C, Poles to cleavage 

(n=117). D, Contour diagram of poles to cleavage. Density and 

symbols as in figure 11.15.5. E, Contour diagram of bedding/ 

cleavage intersections (n=60). Density as in figure 11.15.5. x, 

fold axis (measured); open circle, bedding/cleavage intersection 

(measured).

20 cm

Strike: 

N70°E

35-50 percent 

shortening

\

\

\

\

B

35 percent extension

FIGURE 11.16. Fold style in Weatherby Formation of Brooks 

(1979b). A, Minor fold traced from field photograph. B, Diagram 

of fold style as determined from observed folds and cleavage/ 

bedding relations. Large arrows indicate directions of principal 

strain.
440 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

B

x

ci) 

0) 

rt '"8 3

3 4 

Y _ Intermediate axis 

Z Minor axis

Strain ratios

Sample 

number

2g-----

290 ----

X

Z

2.5

2.9

Y

Z

1.9

2.3

X 

Y

1.2

1.2

Stretching 

(percent)

X

40

45

Y

7

15

Z

-33

-40

FIGURE 11.17. Orientation and degree of strain in Weatherby Formation of Brooks (1979b). A, Strain ellipsoid. X, major axis; Y, 

intermediate axis; Z, minor axis. Strain ellipse ratio X/Z on face 

a is greater than ratio Y/Z on face b, which is greater than ratio 

X/Y on face c. B, Flinn diagram. K=(X/Y-1)/(Y/Z-1). Line K=l 

separates field of apparent flattening (K<1) from field of apparent 

constriction (K>1). 29 and 290 are sample numbers. C, Strain 

ratios and stretching values for samples 29 and 290. X, Y, and Z 

are axes shown in figure 11.17A. Stretching values assume there 

was no change in volume during or after deformation.

at least 2 m.y. (36 to 34 Ma) (N.W. Walker, oral commun., 1986; Manville Products Corp., oral commun., 

1986) at the beginning of the Oligocene. The Tertiary 

intrusive activity can be divided into three periods: 

(1) emplacement of the granodiorite of Bullrun Creek 

between Mine Ridge and Bullrun Mountain (fig. 

11.1), associated with a few late-stage biotite-bearing 

dikes, at 36 Ma; (2) emplacement of the porphyritic 

granodiorite of Grouse Creek along the axis of the 

inlier to the south of Bullrun Mountain (fig. 11.1), accompanied by many minor intrusions and associated 

with extensive hydrothermal alteration and mineralization, at 34 Ma; (3) emplacement of several types of 

post-hydrothermal minor intrusions concentrated immediately southeast of Bullrun Mountain (fig. 11.14). 

Volcanic rocks of Tertiary age dip off the inlier to 

east and west and are described by Walker (1990).

GRANODIORITE OF BULLRUN CREEK

The intrusion that is here called the granodiorite of 

Bullrun Creek (hereafter also referred to as the Bullrun Creek stock) is a kidney-shaped stock 2.0x0.7 km 

that occupies the east side of Mine Ridge, crosses 

Bullrun Creek, and extends south into Koski Basin 

(figs. 11.1, 11.2).

The main phase of the stock Li a massive, mediumgrained, gray granodiorite that contains plagioclase, 

K-feldspar, quartz, biotite, and hornblende in modal 

proportions that span the tonalite-g/ant diorite boundary (table 11.9; fig. 11.18; Houseman, 1983). Inclusions are rare but are occasionally found near 

metasedimentary contacts. A seriate ptrpryritic border phase occurs in Payton Gulch, and a clarr, slightly 

mafic-rich tonalitic border phase is found OL th.3 lower, 

eastern slope of Mine Ridge. Both variants g/aue into 

the main body of the stock. Late aplite veins anu a; kes 

from 1 to 50 cm wide, with minor green amphiboie 'n 

a feldspar-rich matrix, are common within the stock.

The border zone is poorly exposed, but the contact 

against hornfelsed metasedimentary rocks can be observed in both Koski Basin and Payton Gulch where 

small-scale forked and braided dikes penetrate a few 

meters into the hornfels. The contact between the 

granodiorite and serpentinite is exposed underground 

in the Record Mine, where it dips 60° NE but is modified by faulting. Discontinuous shear zones containing lensoid slivers of serpentinite within argillized 

granodiorite parallel the contact and are cut by 

northeast-striking en echelon fractures and faults 

that dip steeply northwest. This latter set, although 

displaying only limited offsets, has strongly influenced later vein and dike orientations. In undisturbed 

granodiorite/serpentinite contacts, chill zones in the 

granodiorite are limited to 0.5 to 1.0 m in width.
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 441

TABLE 11.9. Estimated modal mineralogy of the granodiorite of Bullrun Creek

[All values in volume percent. Traces of apatite and zircon were detected in all samples. Epidote and muscovite are present in JB-81-18E, and

significant calcite occurs in JB-81-59A. tr, trace; , not observed]

Sample JB-81-18E JB-81-35 JB-81-59A JB-81-59C JB-81-61 JB-81-63 JB-81-71

K-feldspar 

fill QT*^" f7

Hornblende « »--

"Rintito

Chlorite 

Opaque minerals 

Sphene - - 

1

1 XL

1 K.

5 

3

tr

5

q

1 

1 

tr

6

Ofl

1

q

1 

3 

tr

CO

5 

oe

c

c

3

C1

5

Ofl

1

K

1

2 

tr

10

Ofl

1 F;

1

4

tr

5

Ofl

1

q

2 

3

AT7 XTTX

Fine-grained biotite tonalite dikes cut the granodiorite of Bullrun Creek at the Record Mine and to the 

south in Koski Basin. These dikes are cut by younger 

porphyritic dikes at both localities. The biotite tonalite dikes are quartzofeldspathic and equigranular, 

and are believed to represent a late phase of the 

Bullrun Creek stock.

Hydrothermal alteration associated with the emplacement of the granodiorite of Bullrun Creek is 

minimal. Amphibole-apatite-sulfide veins and quartzalbite-carbonate alteration is present along the eastern margins of the stock at the Record Mine and at 

least intermittently along the serpentinite contact to 

the south. The veins represent later infilling of 

northeast-striking en echelon fractures.

GRANODIORITE OF GROUSE CREEK 

AND RELATED MINOR INTRUSIONS

The main stock that is here called the granodiorite 

of Grouse Creek (hereafter also referred to as the 

Grouse Creek stock) is elongated northwest and is 

4.25 km long and 0.75 km wide (figs. 11.1, 11.14). It 

is a pale, medium- to fine-grained rock with phenocrysts (1-4 mm) of plagioclase, hornblende, and, in 

some areas, quartz. The groundmass of plagioclase,

FIGURE 11.18. Streckeisen diagram of samples of the equigranular granodiorite of Bullrun Creek. Q, quartz; A, alkali feldspar; P, plagioclase. Fields 1 through 6 (from 

Streckeisen and others, 1973): 1, granite; 2, granodiorite; 

3, tonalite; 4, quartz monzonite; 5, quartz monzodiorite 

and quartz monzogabbro; 6, quartz diorite, quartz gabbro, 

and quartz anorthosite. A, Normative values with normative albite and anorthite summed to create total normative plagioclase. B, Modal values. C, Normative values 

with plagioclase calculated to An40 and remaining albite 

placed in alkali-feldspar component.
442 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

hornblende, quartz, and opaque minerals, with traces 

of K-feldspar and biotite and secondary chlorite and 

sericite, is too fine grained for modal analysis, but a 

normative plot (fig. 11.19) indicates a granodiorite 

composition. The main stock and the many associated dikes, sills, and breccias are pervasively altered 

and lie in the central part of a larger alteration halo 

that overprints all pre-Tertiary rocks. Alteration assemblages and copper mineralization (Houseman, 

1983) imply that a copper porphyry system is associated with the intrusion.

The complex border zone is composed of alternating slivers of calc-silicate hornfels and porphyritic 

dikes and sills (Houseman, 1983). On a large scale, 

porphyritic dike orientation corresponds to the prevailing east-northeast strike of the axial-planar 

cleavage in the country rock (Weatherby Formation). 

In detail, dike margins are irregular and accompanied by conspicuous brecciation. Breccia zones range

EXPLANATION

 Granodiorite of Bullrun Creek

o Granodiorite of Grouse Creek 

and associated dikes

A Dikes from Rough Ridge younger 

than granodiorite of Grouse Creek

A Dikes from Rastus Mountain younger 

than granodiorite of Grouse Creek

FIGURE 11.19. Normative values of all analyses of Oligocene intrusive rocks of Ironside Mountain inlier plotted on Streckeisen 

diagram (Streckeisen and others, 1973) with normative albite 

and anorthite summed to create total normative plagioclase. 

Endpoints and fields as in figure 11.18.

from a few meters to several hundred meters across 

and consist of metasedimentary clasts from 1 cm to 

1 m in size in an igneous matrix. Large metasedimentary roof pendants are common and typically haloed by breccia. Smaller sedimentary inclusions occur 

throughout the stock. Breccias also occur within the 

main stock as dikes and pipes a few centimeters to a 

few hundred meters across. Simple breccia zones contain monolithologic assemblages of porphyritic granodiorite fragments, whereas the more complex zones 

contain fragments of both granodiorite and metasedimentary rock. In the latter the clasts are more obviously abraided and show evidence of multiple 

brecciation. Many breccia zones have a northwest to 

north-northwest orientation. The breccia matrix is 

typically aplitic, altered to a felty aggregate of chlorite, epidote, and clinozoisite.

Several northwest-striking fracture zones cross the 

stock northwest of Rough Ridge (figs. 11.1, 11.14). 

These appear to represent major zones of weakness 

and host one or more late-stage intrusive phases. 

Movement along the fractures does not appear to 

have been significant.

Numerous, apparently co-magmatic, porphyritic 

dikes associated with propylitic alteration crop out on 

Mine Ridge and on Bullrun Mountain. Chilling of the 

porphyritic dikes against the granodiorite of Bullrun 

Creek indicates that at least a brief period of cooling 

separated the two intrusive events. Dikes intruding 

along the eastern margins of the Bullrun Creek stock 

are commonly quartz-albite-carbonate altered. Along 

Bullrun Creek, the dikes tend to strike north-northwest. To the east, within the ultramafic rocks, the 

dikes tend to strike more northeasterly. These latter 

dikes intrude along faults that are parallel to, or 

form the contact between, the Weatherby Formation 

and ultramafic rocks (fig. 11.2).

LATESTAGE PORPHYRITIC DIKES

Several phases of porphyritic dikes postdate the 

granodiorite of Grouse Creek. Some are texturally 

and mineralogically similar to the stock and are distinguished only by their lack of hydrothermal alteration. Relative ages of the younger dikes are 

determined by their mutual crosscutting relations, 

and the increasingly younger dikes are progressively 

more distinct in appearance and texture from the 

granodiorite.

Plagioclase and hornblende phyric granodiorite 

dikes cut the granodiorite of Grouse Creek on the 

southeast side of Rough Ridge (fig. 11.1) and cut the 

granodiorite of Bullrun Creek and adjacent peridotite 

near the Record Mine (fig. 11.2). The dikes are elon-
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 443

gate, irregular bodies as much as 150 m long. One 

large body near the Record Mine is 100x30 m. These 

dikes resemble the rocks of the granodiorite of 

Grouse Creek but lack the alteration and are chilled 

against altered granodiorite.

Chemically similar dikes that postdate the hydrothermal alteration intrude both the granodiorite of 

Grouse Creek and its country rock near Rastus 

Mountain. These are aphyric to weakly phyric and 

contain biotite as the dominant phase. They have 

variable widths and are as much as 700 m long in a 

north-northwest direction. Some are circular-shaped 

bodies chilled against the granodiorite.

Another, younger group of post-hydrothermal plagioclase-hornblende-biotite-phyric dikes form a northnorthwest-oriented group that dips 30°-85° northeast 

between Rough Ridge and Bullrun Mountain. The 

dikes are 3-15 m wide and 50-1,500 m long. Some 

pinch and swell and appear to have intruded along 

prominent north-northeast-striking fracture zones 

that cut the Grouse Creek stock and the adjacent 

Weatherby Formation.

CHEMICAL COMPOSITION OF THE OLIGOCENE INTRUSIONS

The Bullrun Creek and the Grouse Creek stocks 

are readily distinguished in the field by texture and 

degree of hydrothermal alteration. Differences in 

bulk chemical composition are small.

Classification of coarse-grained calc-alkaline rocks 

is normally based on the relative proportions of modal 

quartz, alkali feldspar, and plagioclase (Streckeisen 

and others, 1973). Modal analysis of the fine-grained 

and altered porphyritic rocks of the granodiorite of 

Grouse Creek is impractical, and the bulk chemical 

composition provides a better means of recording and 

comparing their compositions. Superimposing normative values derived from chemical analyses on a 

Streckeisen modal diagram results in some distortion. 

Various methods for converting from modal to normative data have been attempted (Le Maitre, 1976).

Figure 11.18 shows the modal compositions of the 

equigranular Bullrun Creek stock plotted directly on a 

Streckeisen diagram (fig. 11.18.B). It also shows the 

differences produced by plotting the normative values 

of the same samples on the same diagram, first, with 

normative albite and anorthite summed to create a 

total normative plagioclase (fig. 11.18A), and second, 

with normative plagioclase calculated assuming a plagioclase composition of An40 (judged the average plagioclase composition), the remaining normative albite 

being added to the alkali feldspar component (fig. 

11.18C). Le Maitre's calculations were tested but not 

used because the results were at greater variance with

the modal plots than were the results of the two methods described above. We conclude from figure 11.18 

that a straightforward normative plot (fig. 11. ISA) 

shows the least divergence from the true modal plot, 

although it results in a slight bias toward the alkali 

feldspar side of the diagram. Plots of the normative 

values of all the analyzed samples of the various Oligocene intrusions (tables 11.10-11.12) are shown in 

figure 11.19.

Figures 11.18 and 11.19 demonstrate that all the 

Oligocene intrusive rocks from the north end of the 

Ironside Mountain inlier have compositions close to 

the tonalite-granodiorite boundary and that they are 

not obviously distinguishable one from another on the 

basis of major-element composition. Nor do the relatively few trace-element analyses provide an obvious 

means of distinguishing the rocks (table 11.13). The 

low-K tonalitic aspect of these rocks is further emphasized in Barker's (1979) diagram (fig. 11.20) in which 

the great majority of samples fall within Barker's tonalite field, some in>the trondhjemite field, and only a 

few in the granodiorite field. While clearly calc-alkaline in character (fig. 11.21), all these rocks belong to 

a low-K calc-alkaline suite (Barker and Arth, 1976).

The analyses are displayed in a Harker variation 

diagram in figure 11.22. Although some scatter is apparent, most obviously in the CaO, K2O, and Na2O 

concentrations of the hydrothermally altered Grouse 

Creek rocks, the overall trends are typical of calc-alkaline suites, with the exception that there is little 

evidence of the expected positive correlation between 

K2O and SiO2 concentrations.

Potassium-poor salic magmas similar to those of the 

Ironside Mountain inlier have been interpreted by 

Ewart (1979) as indicative of subcontinental or young 

continental crust, whereas Barker (1979) suggested 

that such tonalite-trondhjemite suites are frequently 

found within Mesozoic to Cenozoic continental margins 

and typically on the ocean side of large calc-alkaline 

batholiths. The Oligocene intrusions of the Ironside 

Mountain inlier were emplaced into recently accreted 

Permian and Triassic ophiolites and Mesozoic volcaniclastic rocks. During the Oligocene these terranes lay 

marginal to the continent, outboard (west) of the broad 

calc-alkaline Challis-Absaroka arc (which extends 

from northern Washington through central Idaho and 

into northwestern Wyoming) and the Cretaceous to 

Tertiary intrusions associated with the Idaho batholith. The age of the Ironside Mountain intrusions suggests that they were emplaced during the abrupt 

westward migration of calc-alkaline volcanism from 

the Challis-Absaroka arc into central and western Oregon and are best regarded as intrusive equivalents of 

the Clarno Formation of northeastern Oregon.
444 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 11.10. Chemical analyses of the granodiorite ofBullrun Creek and related dike and aplite

[All values in weight percent. Major-element analyses and Cross, Iddings, Pirsson, and Washington (CIPW) norms by X-ray-fluorescence spectroscopy, 

Washington State University (Hooper and Johnson, 1989). Major elements normalized on a volatile-free basis with 0.9 Fe expressed as FeO, 0.1 Fe 

expressed as Fe2C>3 and Fe* equal to FeO+FejjCg. , not present]

Sample ------ TT2 Ql

62

JB-82- 

133

JB-81- 

36

JB-81- 

55

JB-81- 

59

JB-82-

157

JB-82- GCA 

162

JB-82- 

295

JB-82- 

332

Major elements

SiO2 

A1203 -----

m: r\ 1 1U2 

Fe* ______

MnO ________

PaO 

MaO 

Tf r\ J12U 

Na2O ----- 

Pr\ 2U5 - 

66.84 

16.71 

fi^

Q Q^

n^

4.27

2 op.

1 ^7

4.10 

.17

63.34 

16.69

OK

5.04

.10

5 ifi

2

^°.

1 Q7

4.64

OQ

68.60 

16.54

£JQ

Q f^A

OK

3

d

9 91

1 QQ

3.33 

1 s.

65.28 

17.07 

.84

4.15

.07

A QQ rt.OG 

9 41

1

A O .43

3.72 

,£i£t 99

63.46 

18.01 

.84

4.34

no

K. ^fi

2 en

1 1 Q

3.71 

22

68.10 

15.98 

6Q

3

KQ

f\C

4 09 rr. \JU

2 07

1

KQ

3.68

1 Q

66.11 

16.12

CO

4

QQ

no

*=» 1 Q

2

AQ

1 9*=;

3.60

1Q

67.14 

15.83

ca

3 07

nc

4 90 rr.^v/

2 04

1

"7C

4.43

1Q

69.55 

16.48 

.43

2

CQ

no

3

CQ

1 19

.95

4.91 

1^

79.30 

12.16

1 Q

.00

.00

QQ

HP.

2 at

4.75 

.04

CIPW norms

Quartz 

Corundum 

ilfi" ri f\f* 1 Q GO

Albite 

A f-i j-vf* 4- f\ 1 4- f\

Wollastonite 

Enstatite 

Ferrosillite 

Ilmenite 

Rutile 

94 4*5

91

Q 9P.

34.70 

90 07

5.68 

1.02

2

OC

1 9Q

.40

1 a co

8 -in

39.26

on an

1.43 

6.43 

1.78 

3.41

1

C1

.54

QQ m

Q 49

7

QC

28.18

1C QQ

5.50 

1.16 

9 QQ

1 90

QK

24.30

11 s.

8.45

31.47 

99 P59

6.00 

1.21

2 on

I

cn

^9

21.59

1 n^

6 Q7

31.40 

26.15

6.48 

1.33

2 QQ

1 5Q

K.9

OQ on

1 91

Q Q4

31.14

1911

5.15 

1.10 

2 49

1 31

.43

9^ K.1

7 QQ

30.46 

24.13

.03 

6.18 

1.62 

2 96

1 29

.45

22 90

10.60

37.20 

18.10

.80 

5.10

1.20 

2 en

1 90

QO

97 P.H

113

5

f*f\ 

.60

41.55

n OQ

2.96 

.94

-I Q-l

00

QK,

qq 77

15.42

40.19 

4.15

.20

.18 

09

ALTERATION AND MINERALIZATION

The emplacement of the granodiorite of Grouse 

Creek was accompanied by extensive hydrothermal 

alteration that caused Cu-Au-Ag and associated basemetal enrichment. Within the vicinity of the Grouse 

Creek stock, this alteration was responsible for silver 

and base-metal enrichment. To the north, within the 

granodiorite of Bullrun Creek and the ultramafic 

rocks, porphyritic dike emplacement was associated 

with widespread gold and base-metal mineralization.

GRANODIORITE OF BULLRUN CREEK RECORD MINE

Gold mineralization occurred within amphibole 

stockwork zones in the granodiorite of Bullrun Creek 

(Record Mine, fig. 11.2) and within sheared, talc-carbonate-altered ultramafic rocks on Bullrun Mountain 

(Thomason Mine, fig. 11.2) and Mine Ridge (Bullrun 

Mine, fig. 11.2; Orion Mine, lat 44°20'40" N., long 

118°16'47" W.). These mines form the old Bullrun Mining District. The Record Mine, the largest mine in the 

district, is thought to have produced about $103,000, 

mainly during the 1930's (Brooks and Ramp, 1968).

Gold mineralization at the Record Mine is in northeast-striking amphibole veins and stockwork zones 

along the northeast margin of the Bullrun Creek 

stock. The veins extend 100-200 m into the stock away 

from the ultramafic contact, but few veins extend beyond the contact into the ultramafic rocks. Within the 

stock, vein densities and gold mineralization appear to 

increase outward toward the ultramafic contact. The 

veins consist of a dark-green coarse-bladed amphibole, 

coarse-grained apatite, magnetite, sulfides, and occa-
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 445

sional pyroxene or albite intergrowths. Sulfide minerals include molybdenite, pyrrhotite and lesser pyrite, 

chalcopyrite, and cobaltite. Molybdenite-cobaltite concentrations may locally reach several percent. Very 

fine to coarse free gold is present as intergrowths with 

amphibole, apatite, magnetite, or molybdenite, and as 

very fine inclusions in most sulfides. Actinolite, biotite, 

and chlorite replace amphibole and pyroxene, and, 

along with rare quartz, form late, crosscutting veins. 

These late veins also contain pyrrhotite, pyrite, and 

chalcopyrite and some free gold. Albite encloses most 

amphibole veins and locally replaces the equigranular 

wall rock. Within albitic zones, primary hornblende is 

replaced by a secondary pale or light-green amphibole. 

Alteration and vein formation must have occurred 

when the Bullrun Creek stock was cool enough to 

sustain brittle fracture. Mineralized veins cut both 

the stock and later biotite-tonalite dikes, but few 

veins are found crosscutting porphyritic dikes. Instead, many of the porphyritic dikes (associated with 

the Grouse Creek stock) in the vicinity of the Record

EXPLANATION

 Granodiorite of Bullrun Creek

O Granodiorite of Grouse Creek 

and associated dikes

A Dikes from Rough Ridge younger 

than granodiorite of Grouse Creek

A Dikes from Rastus Mountain younger 

than granodiorite of Grouse Creek

FIGURE 11.20. Normative anorthite(An)-albite(Ab)-orthoclase(Or) 

diagram (Barker, 1979) of all analyses of Oligocene intrusive 

rocks of north end of Ironside Mountain inlier.

Mine are pervasively altered to secondary amphibole 

and albite, suggesting alteration may have accompanied, or closely followed, the emplacement of the porphyritic dikes. These porphyritic dikes may contain 

several percent disseminated pyrrhotite, chalcopyrite, 

and pyrite, and many are gold-bearing.

GRANODIORITE OF GROUSE CREEK

The granodiorite of Grouse Creek forms the core of 

a large, elongate hydrothermal alteration zone (fig. 

11.14), or pyrite halo, that extends from the south 

end of Mine Ridge 13-14 km southeast toward Ironside Mountain (figs. 11.1, 11.14). The zone is formed 

mainly within the Weatherby Formation, but in Koski Basin it cuts the south end of the Bullrun Creek 

stock. Altered metasedimentary rock within the zone 

typically contains secondary pyrite in amounts ranging from trace to several percent; in addition the rock 

may be weakly silicified. Calcareous beds near the 

margins of the Grouse Creek stock contain minor 

skarn. Porphyritic dikes within the alteration zone 

are propylitically altered.

FeO

Na2O+K2O MgO

EXPLANATION

 Granodiorite of Bullrun Creek

O Granodiorite of Grouse Creek 

and associated dikes

A Dikes from Rastus Mountain younger 

than granodiorite of Grouse Creek

FIGURE 11.21. Alkali (Na2O+K2O)-FeO-MgO diagram of all Oligocene intrusive rocks of north end of Ironside Mountain inlier. Tholeiitic and calc-alkalic fields from Kuno (1968).
446 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 11.11. Chemical analyses of

[All values in weight percent. Major-element analyses and Cross, Iddings, Pirsson, and Washington (CIPW) norms by X-ray-fluorescence spectroscopy, Washington State University

MH2 MH11 MH36 MH49A MH51 MH65 MH101A MH140 MH165 MH174

Major

gi()2 ______________

AloOo

pe* __________

MfrO ______ ____ 

paO _ -

Na2O 

KoO ___

TiO2 _

PoOe £ £\J£ _

MnO

69.65 

16.26 

2.60 

1.02 

4.95 

3.32 

1.61 

.34 

.10 

.06

66.45 

17.13 

3.20 

1.99 

4.51 

5.07 

.90 

.52 

.15 

.08

66.52 

17.04 

3.39 

1.86 

3.88 

5.17 

1.26 

.62 

.19 

.07

68.03 

17.14 

2.98 

2.01 

3.86 

3.43 

1.74 

.53 

.16 

.13

67.99 

16.27 

3.27 

1.61 

3.56 

5.24 

1.31 

.46 

.16 

.13

66.89 

17.42 

2.55 

1.69 

4.48 

4.41 

1.78 

.57 

.15 

.05

68.04 

16.73 

.75 

1.84 

3.18 

5.15 

3.57 

.54 

.15 

.04

67.27 

17.03 

2.77 

1.85 

4.30 

4.52 

1.51 

.60 

.19 

.03

66.35 

16.43 

3.74 

2.07 

4.27 

4.87 

1.44 

.59 

.19 

.07

66.90 

17.31 

3.61 

1.98 

4.14 

3.88 

1.38 

.56 

.17 

.07

CIPW

^)ii 3. rtz _-- _- « --

f^iH~Vi r*/^l o c A

Alhifa

Enstatite 

Ferrosilite 

TW a cm oi"i f"o

Hematite 

30.27

9.99 

29.79 

23.52 

.24 

2.39 

1.13 

1.80

.65

.24

JB-81- 

25

23.64 

.60 

5.56 

38.42 

21.84

6.42 

5.23 

2.18

1.01 

.36

JB-81- 

26

23.52 

1.18 

7.86 

39.27 

18.46

4.66 

1.12 

2.31

1.20 

.45

JB-81- 

40

28.71 

2.43 

10.82 

30.13 

18.45

4.96 

1.13 

2.00

1.01 

.38

JB-81- 

48

25.56 

.93 

8.21 

39.26 

17.09

4.01 

1.43 

2.23

.89 

.40

JB-82- 

154

23.53 

.47 

11.17 

35.54 

21.63

3.81 

.65 

1.73

1.10 

.38

JB-82- 

156

18.26

21.80 

39.26 

14.00 

.43 

4.49

.33 

.91 

.17 

.36

JB-82- 

163

23.72 

.67 

9.40 

37.23 

20.04

4.78 

.71 

1.86

1.16 

.45

JB-81- 

JM1

21.06

8.92 

39.69 

19.21 

.43 

5.23 

2.60 

2.52

1.14 

.45

GCB

26.76 

2.08 

8.57 

32.75 

19.66

4.91 

1.35 

2.44

1.08 

.43

GCC

Major

SiO2 

AloOs

T?p* _______ ______

MpO 

PnO

MooO 

KoOTiO2 

PoOc JT^V-'O

MnO ____ ____

66.30 

16.02 

3.65 

2.34 

4.10 

4.88 

1.71 

.78 

.19 

.04

66.48 

16.08 

3.74 

2.04 

5.33 

4.49 

1.00 

.63 

.17 

.05

66.59 

17.35 

3.69 

2.54 

3.88 

3.88 

1.39 

.48 

.14 

.06

65.36 

17.76 

4.39 

2.33 

4.63 

3.33 

1.28 

.71 

.18 

.02

67.88 

16.22 

3.52 

1.87 

3.78 

4.77 

1.24 

.52 

.14 

.06

67.21 

16.12 

4.23 

1.95 

3.86 

3.91 

1.98 

.55 

.16 

.03

66.52 

16.31 

4.39 

2.22 

3.58 

4.80 

1.26 

.64 

.17 

.11

67.89 

17.30 

2.89 

.89 

5.18 

3.12 

1.49 

.43 

.13 

.07

65.17 

16.54 

4.38 

2.61 

3.84 

4.50 

1.98 

.88 

.26 

.05

69.00 

15.98 

3.07 

2.59 

1.70 

3.80 

3.35 

.47 

.13 

.05

CIPW

Quartz 

Orthoclase 

Alhif-p ___ ___ _

Anorthite 

Wollastoite 

Enstatite 

Ferrosilite 

T\TQ crT"ioi"ii"o

19.70

10.19 

41.30 

16.75 

0.98 

5.83 

0.96 

2.46

1.48

0.45

22.96

5.91 

37.99 

20.77 

1.91 

5.08 

1.27 

2.52

1.20 

0.40

26.31 

2.74 

8.21 

32.83 

18.33

6.33 

1.52 

2.49

0.91 

0.33

27.54 

2.91 

7.56 

28.18 

21.80

5.80 

1.47 

2.97

1.35 

0.43

24.29 

0.49 

7.33 

40.37 

17.84

4.66 

1.33 

2.38

0.99 

0.33

25.40 

0.91 

11.70 

33.08 

18.10

4.86 

1.67 

2.86

1.04 

0.38

22.48 

0.95 

7.45 

40.62 

16.65

5.53 

1.76 

2.96

1.22 

0.40

30.11 

0.30 

8.80 

26.40 

27.97

2.22 

1.14 

1.94

0.82 

0.31

19.30 

0.50 

11.70 

38.30 

17.50

6.50 

1.20 

2.80

1.70 

0.60

26.90 

3.40 

20.00 

32.00 

7.50

6.50 

0.10 

1.80

2.00 

0.30
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 447 

the granodiorite of Grouse Creek

(Hooper and Johnson, 1989). Major elements normalized on a volatile-free basis with 0.9 Fe expressed as FeO, 0.1 Fe expressed as Fe2C>3, and Fe* equal to FeO+Fe2C>3. , not present]

MH178C MH179 MH180 MH189 MH191 MH196A MH198B MH204 MH208 MH219 MH221 MH226

elements

67.64

17.02

3.70

2.15

3.24

3.95

1.50

.55

.17

.07

69.21

17.28

3.26

2.27

2.17

3.95

1.07

.56

.18

.05

66.97

16.71

5.07

2.23

2.37

4.08

1.68

.63

.19

.08

66.09

17.31

3.92

2.41

3.55

4.33

1.52

.59

.21

.07

66.51

17.21

3.52

1.81

3.82

4.79

1.55

.54

.19

.07

68.35

16.84

1.23

1.98

3.69

4.72

2.45

.52

.14

.07

68.16

16.48

3.56

2.14

4.88

3.48

.46

.61

.15

.07

66.40

17.13

3.88

2.03

3.93

4.50

1.31

.60

.16

.07

68.91

16.64

3.13

1.61

3.14

4.39

1.50

.46

.13

.07

69.30

17.31

1.71

1.71

3.12

4.81

1.40

.50

.13

.03

69.83

17.08

3.29

1.66

1.75

4.23

1.50

.48

.13

.05

64.43

17.62

4.82

1.84

4.16

4.62

1.37

.77

.27

.10

norms

28.15

3.21

9.34

33.25

15.19

5.48

1.42

2.49

...

1.06

.43

GCD

33.87

6.01

6.56

33.26

9.74

5.78

1.09

2.20

...

1.08

.38

GCE

27.66

4.21

10.46

34.02

10.65

5.80

2.09

3.44

...

1.22

.47

GCG

23.79

2.43

9.52

25.88

16.54

6.08

1.48

2.65

1.14

.50

GCH

22.18

1.18

9.63

39.09

17.94

4.81

1.28

2.38

...

1.05

.47

KEBR7

21.28

.01

15.19

38.84

17.67

4.88

.72

.05

1.01

.36

KEX9

32.00

1.65

2.54

29.79

23.31

5.53

1.22

2.41

...

1.18

.38

KE9

23.61

1.49

8.16

36.98

18.70

5.48

1.43

2.62

1.16

.38

KE17

28.14

2.33

9.34

36.39

15.03

4.26

1.22

2.12

...

.89

.31

KE19

27.48

2.58

8.75

39.35

14.91

4.28

.24

1.13

...

.95

.33

KE95

32.95

5.59

9.34

35.37

7.87

4.16

1.24

2.23

...

.95

.33

KE113

20.61

1.62

8.10

39.09

18.87

4.58

1.78

3.26

1.46

.64

KE214

elements

68.49

16.62

3.83

1.84

1.76

4.90

2.03

.53

.15

.04

65.76

17.58

3.14

2.67

4.03

4.98

.81

.92

.22

.03

68.84

16.98

3.34

1.88

3.12

4.52

.76

.54

.16

.02

68.53

17.37

3.91

3.02

1.99

4.08

.47

.59

.16

.06

64.17

18.47

4.74

2.07

4.60

3.50

1.50

.69

.17

.07

70.48

16.46

1.98

1.08

4.39

3.48

1.74

.27

.08

.04

68.40

16.75

3.28

1.64

2.93

5.12

1.19

.49

.15

.07

67.48

16.56

4.18

2.05

3.71

4.28

.94

.57

.16

.06

69.53

16.62

3.11

1.29

1.96

5.57

1.27

.47

.14

.06

68.21

15.98

.30

1.77

6.19

4.85

1.89

.49

.13

.05

66.41

17.63

4.89

2.57

1.81

4.21

1.41

.74

.24

.08

67.98

16.36

3.41

1.63

4.14

4.36

1.47

.51

.15

.06

norms

25.40

3.50

12.20

41.40

7.80

...

4.60

1.30

2.60

...

1.10

...

0.30

21.20

1.40

5.00

41.90

19.20

.40

6.70

2.10

1.70

...

.60

30.00

3.30

4.40

38.30

14.70

4.70

1.10

2.10

1.10

...

.30

34.00

6.80

2.80

34.60

8.90

...

7.60

1.30

2.60

1.10

...

.30

24.90

3.10

8.90

29.30

22.00

...

5.20

1.70

3.20

...

1.40

.30

33.20

1.20

10.00

29.30

20.80

...

0.40

2.60

2.10

...

.50

.30

25.00

1.90

7.20

43.50

13.60

...

4.00

1.20

2.30

...

.90

.30

22.70

2.00

5.60

36.20

17.50

...

5.10

1.70

2.80

1.10

.30

26.20

2.90

7.20

47.20

8.90

...

3.10

1.20

2.10

.90

.30

20.00

11.10

41.40

16.10

5.10

4.40

...

.20

.50

.60

.30

28.20

6.20

8.30

35.60

8.10

6.60

1.80

3.20

...

1.40

.60

24.90

8.90

36.70

20.60

...

3.90

1.50

2.30

...

.90

.30
448 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 11.12. Chemical analyses of dikes that postdate hydrothermal alteration

[All values in weight percent. Major-element analyses and Cross, Iddings, Pirsson, and Washington (CIPW) norms by X-ray fluorescence spectroscopy, 

Washington State University (Hooper and Johnson, 1989). Major elements normalized on a volatile-free basis with 0.9 Fe expressed as FeO, 0.1 Fe 

expressed as Fe2C>3 and Fe* equal to FeO+Fe203 , not present]

Sample MH891 MH2161 JB81-61 GCF1 KEX101 KE261 MH52 MH1882 MH533 MH1993 MH324 MH1844 MH2114

Major elements

SiO2 

A1203 ------- 

TiOo 

£1 FP* 

MnO _____ 

CaO 

M?O ______ IVAgW

KoO __ _

NnoO - _

P205 ------ ----- -

61.77 

16.27

QO

C CQ

.10

4.78 

3.53 

1.71

4 Q8

o/>

62.73 

16.86

QQ

C 10

AQ

4.62 

2.82

1 &*(

A crj

00

60.69 

17.84

1 OQ

K 71

.07

4.78 

3.50

1 CQ

A on

9C

63.75 

17.61 

2.40

O CO

A F7F7

5.71 

.82

-| f\£*

91

.04

59.73 

18.07

QQ

f* crj

.10

5.92 

3.61

Q9

0 Qfi

i a

63.00 

16.87 

7Q

K 10

.07

5.23 

2.57

1 09

A CO

on

66.81 

17.63

CQ

O CC

.07

2.64 

2.06

2 on

4.08

ift

67.11 

17.18 

.54

3.00

.07

4.25 

2.08

1 qA

A qq

1 c

68.43 

17.08 

.40

2

09

nc

3.19 

1.78

A qo

19

68.68 

16.82

3 9O

AC

3.36 

1.72

4.14

1 q

71.66 

16.08

97

1 oq

AO

3.22 

.95 

9 1Q

q C7

09

73.60 

15.73 

97

1 70

AO

2.45 

.45

9 qc

q 9K

OQ

71.45 

16.56 

97

1 Q7

.04

2.85 

.80

2

1 Q

q 7Q

OQ

CIPW norms

Corundum 

Orthoclase 

AlKifp ______

Wollastonite 

Enstatite 

Ferrosilite 

Magnetite 

Haematite 

Apatite 

1 q oq

10.64

qo AA

1 Q Id

1.18 

9.12 

1.95 

3.79

1 7Q

.88

1 C OQ

11.47

OQ r\(\

1Q 77

.47 

7.55 

1.70 

3.47

1.71

.81

14 94

.96 

9.34

O/> OQ

99 AQ

8.72 

1.40 

3.86

2.40

.59

19 70

5.00 

48.20

20.00

1.20 

9.10

1.20 

.30 

2.00

.30

1 q QA

.10 

5.60

oo en

28.40

2.80 

9.40 

4.40

1.70

.30

15.10

10.60

qo on

1 Q QA

2.20 

6.70 

2.20 

3.50

1 ^O

.60

OK CO

3.87 

14.30

oo na

19 ic

4.93 

1.32 

2.48

1 19

.45

24.81

1.18 

8.04

qc 07

5.18 

1.03 

2.02

1 no

.38

97 84

2.66 

10.28

qc 90

ic 07

4.33 

1.18 

1.97

70

.31

90 09

2.46 

15.19 

34.69

"1 C t F7

4.18 

1.29 

2.16

09

.31

00 Q .4

.60 

5.56

OQ AC)

10 IK

2.07 

.75 

1.29

1.01

.21

07 07

3.36 

14.59

OQ AQ

U 7C

1.67 

.65 

1.17

K1

.24

00 Qfi

2.77 

13.59

09 qo

1 Q QQ

1.69 

.80 

1.31

C1

.24

Porphyritic quartz monzodiorite. 

Porphyritic granodiorite, Rough Ridge. 

Bullrun Mountain granodiorite dikes. 

Rastus Mountain granodiorite dikes.

At the core of the pyrite halo, within the Grouse 

Creek stock, porphyritic granodiorite is altered to 

propylitic, phyllitic, argillic, and potassic alteration 

assemblages. These assemblages form complex, overlapping alteration zones that pervasively alter the 

stock (fig. 11.14; propylitic zone not shown). Typically, 

zones of early propylitic and potassic alteration are 

partially or completely overprinted by later phyllitic 

or argillic alteration.

The zone of potassic alteration in the granodiorite, 

characterized by the replacement of hornblende by 

secondary biotite, crops out in topographically low-lying areas along the north and south forks of West 

Camp Creek. Exploratory drilling in these areas encountered well-developed potassic alteration at 

depth. Throughout the Grouse Creek stock, the potassic alteration zone apparently forms a deeper alteration assemblage exposed only in eroded, low-lying

valleys. Geochemical sampling and drilling in the potassic zone encountered significant enrichment in disseminated chalcopyrite and molybdenite.

The most widespread alteration within the Grouse 

Creek stock is propylitic alteration (not shown in fig. 

11.14) formed at the same time as the potassic alteration. The propylitic alteration is characterized by the 

formation of secondary epidote, chlorite, calcite, and 

pyrite, as well as minor clinozoisite, sphalerite, and 

galena. Both the potassic and propylitic zones are partially overprinted by the younger quartz-sericite-pyrite 

assemblage of the phyllitic alteration. This younger 

phyllitic-alteration assemblage forms widespread, geometrically complex zones around the margins of the 

Grouse Creek stock, as well as blanket-like zones on 

Rough Ridge and Rastus Mountain. Porphyritic granodiorite within the phyllitic alteration zone is strongly 

silicified and pyritized. Plagioclase is typically altered
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 449

TABLE 11.13. Trace-element analyses ofOligocene intrusions, Ironside Mountain inlier

[All values in parts per million. Analyses by inductively coupled plasma spectrometry, Kings College, London; analyst, P.R. Hooper.

 , not determined]

Granodiorite of 

Bullrun Creek

Sample JB-81-55 JB-82-295

Granodiorite of Grouse Creek Postalteration dikes

MH-51 1 MH-55 MH-89 JB-81-48 MH-5 MH-53

Ro _____ ___

fr

Cu 

Nb ----- 

Mi

gc _____

Sr 

v ___

Y _____

Zn

Rb -~---

97

20

94

o-i

10

655

97

18

63

34

700

i °.

11

19

IS

6

456

48

13

28

26

7QQ

on

13

1°.

09

8

632

61

15

83

rr/jo

97

42

1 °.

0.7

9

675

73

15

125

7Q7

43

9n

14

755

123

21

125

91

29

1 C.

o.n

10

632

87

16

33

39

rr/n

OK

21

10

O A

10

643

82

13

67

OAA

oc

16

11

9Q

8

686

62

11

50

1Rare-earth-element data (in parts per million) for sample MH-51: La, 3.10; Ce, 23.97; Pr, 3.27; Nd, 13.07; Sin, 2.45; Eu, 0.79; 

Gd, 2.30; Dy, 2.19; Ho, 0.56; E, 1.36; Yb, 1.26; Lu, 0.18.

to sericite, and hornblende is altered to a sericite-chlorite aggregate.

Late-stage argillic alteration overprints propylitic, 

potassic, and phyllitic alteration zones on the lower 

western slopes of Rastus Mountain. The irregularly 

shaped argillic zone appears to represent the youngest alteration event in the Rastus Mountain area, 

and is characterized by intense acid-leaching of the 

porphyritic granodiorite. The rock's primary texture 

has been destroyed and altered to an aggregate of 

quartz, sericite, clay, and limonite. The late-stage argillic alteration zone contains anomalously high 

amounts of mercury and antimony. Coarse bladed 

stibnite is exposed locally in a roadcut in intensely 

acid-leached porphyritic granodiorite.

On the ridge west of Grouse Spring, younger hydrothermal breccias crosscut the Grouse Creek stock. 

Though exposures are poor, results of exploratory 

drilling in the area suggest the presence of an irregularly shaped, shallow, northeast-plunging hydrothermal breccia zone or pipe. Within the pipe, porphyritic 

granodiorite and metasedimentary fragments are cemented by chlorite, quartz, calcite, epidote, clinozoisite, pyrite, and chalcopyrite. Intergrowths of fine- to 

coarse-grained chalcopyrite with chlorite, quartz, and 

calcite have partially replaced earlier epidote and clinozoisite. Drilling within the breccia reveals copper 

mineralization averaging as much as 1 percent Cu, 

with locally higher grades with as much as 4 percent 

Cu. Mineralized zones are silver enriched, and their 

Cu:Ag ratios average around 1:1 (Manville Products 

Corporation, oral commun., 1986).

Hydrothermal breccias in the Grouse Spring area 

display a close spatial and temporal relation to aplite-cemented intrusion breccias cropping out on the 

same ridge. In drill core, hydrothermal breccias are 

both superimposed upon and crosscut by aplitic intrusion breccias. Both breccia types contain a variety 

of porphyritic fragments with different alteration assemblages. Both of these breccias may be related to 

deeper, late-stage magmatic activity. The alteration 

zones associated with the Grouse Creek stock constitute the first Tertiary porphyry system documented 

in northeastern Oregon.

DISCUSSION AND CONCLUSIONS

The pre-Jurassic rocks of the Ironside Mountain inlier include serpentinized peridotite, gabbro, and plagiogranite as well as some basalt, ribbon chert, and 

various schistose units that vary from crossbedded 

quartzite through mica- to hornblende-schist. All but 

the massive peridotite of Bullrun Mountain occur as 

oriented blocks in a melange with a highly sheared, 

serpentinized matrix. The melange also contains 

large blocks of calc-alkaline andesite. The association 

of rock types is similar to that of an ophiolite assemblage (Coleman, 1977), although the andesite is 

anomalous, and neither pillow basalt nor a basic dike 

complex is present.

Ironside Mountain lies between two larger and better exposed inliers of ophiolitic rocks within the Baker 

terrane: the Canyon Mountain Complex 50 km to the
450 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

18 -

16 -

14

12

6

A

- AI 2 0 3

A

FeO 0

°

A

A -

1.2

0.8

0.4

0.12

0.08

0.04

0

6

Ti0

<& 0

n

A A

. MnO

o o

A 0

0 A 

° °

0

0

A

A -

CaO

O

A

A -

60 62 64 66 68 70 72

SiO2 CONTENT, IN WEIGHT PERCENT

74 76

EXPLANATION

Granodiorite of Grouse Creek O Dikes associated with the

Monzodiorite of Bullrun Creek granodiorite of Bullrun Creek

Granodiorite of Bullrun Creek Dikes on Rastus Mountain

Aplites of Bullrun Creek 0 Granodiorite of Rough Ridge

78 80

FIGURE 11.22. Major-element-oxide concentrations plotted against SiO2 concentrations for samples of Oligocene intrusions from

north end of Ironside Mountain inlier.
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 451

west and the Sparta complex almost 100 km to the 

northeast (Thayer, 1977; Gerlach and others, 1979; 

Phelps, 1979; Phelps and Ave Lallemant, 1980). These 

two complexes have different ages (approximately 260 

Ma and 230 Ma, respectively) (Ave Lallemant and others, 1980) and have distinctive REE patterns (fig. 

11.9). The rock associations of both complexes are 

broadly similar to those of the Ironside Mountain inlier, but the Ironside Mountain rocks correspond in age 

and in their more depleted LREE patterns to those of 

the Canyon Mountain Complex. All three inliers lack 

a basic dike complex, posses little if any pillow basalt, 

and include relatively large proportions of plagiogranite and keratophyre; these characteristics, taken together with their geochemical signatures, have led

previous authors to postulate that the Sparta complex 

and the Canyon Mountain Complex originated as the 

residual roots of an island-arc system (Phelps and Ave 

Lallemant, 1980; Gerlach and others, 1981) rather 

than as a section of mantle and oceanic crust formed 

at some distant midoceanic ridge-rift system (Coleman, 1977). The same arguments apply to the rocks of 

the Ironside Mountain inlier, where they are supported by the date of the first major (Dj) deformation, 

which is almost coincident with the formation of the 

ophiolitic rocks of the Canyon Mountain Complex. 

These rocks were formed and then deformed (presumably in the subduction process, see below) within a 

short period, permitting no time for their transport 

from some distant midoceanic ridge-rift system to the

-\ i i i i i i i i i i i i r

MgO o

K 0 O

A

- Na 2 O

0.4

0.2

o

o A

60 62 64 66 68 70 72 

SiO2 CONTENT, IN WEIGHT PERCENT

EXPLANATION

74 76 78 80

O Granodiorite of Grouse Creek

 Monzodiorite of Bullrun Creek

A Granodiorite of Bullrun Creek

A Aplites of Bullrun Creek

D Dikes associated with the

granodiorite of Bullrun Creek

 Dikes on Rastus Mountain 

() Granodiorite of Rough Ridge

FIGURE 11.22. Continued.
452 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

site of subduction. As argued by Gerlach and others 

(1979) for the Canyon Mountain Complex, the ophiolitic rocks of the Ironside Mountain inlier probably 

represent the roots of a primitive island-arc system, 

an environment in which the large blocks of calc-alkaline andesite are less anomalous than they would be 

in a midocean setting.

The peridotite of Bullrun Mountain and Mine 

Ridge has a restricted dunite-harzburgite mineralogy, 

conspicuously lacking in the plagioclase and well-developed phase layering that would be expected of a 

cumulate formed from a basaltic magma within the 

crust. This mineralogy, reflected in the low CaO and 

A12O3 concentrations, taken together with the coarse 

fabric of elongate refractory (Fo90_92) olivine grains, 

indicates that the peridotite formed as the residuum 

of partial melting in the mantle, not as a crustal cumulate. As in a more typical ophiolitic complex, the 

closely associated basalt, gabbro, and plagiogranite 

are interpreted as the fractionated derivatives of the 

basaltic melt crystallizing within the overlying crust.

The peridotite has undergone at least two distinguishable phases of serpentinization. The first phase 

bears no evidence of slickensides or brecciation, and 

the serpentinite assemblage includes lizardite, normally associated with a low shear-stress environment. 

This event postdates formation of the metamorphic 

fabric of the peridotite, including kink bands. It appears to have occurred at depth and to predate emplacement of the peridotite into the crust. The second 

phase of serpentinization includes the formation of 

talc associated with slickensiding and brecciation, in 

addition to the formation of antigorite, all of which 

imply an environment of considerable shear stress. 

This phase is most obviously associated with the emplacement of the peridotite into the crust, an event 

that culminated in the formation of the melange of 

Mine Ridge.

The three fold phases present in the schistose rocks 

of Mine Ridge are interpreted as the result of progressive, noncoaxial shear deformation and are correlated 

with the D! deformation recognized throughout the 

Baker terrane by Ave Lallemant and others (1980). 

The isoclinal FI folds are associated with a relatively 

high grade (epidote amphibolite facies) metamorphic 

fabric dated at 260 Ma and characterized by a penetrative, near-vertical east- to northeast-striking foliation plane within which the fold axes plunge steeply 

to the north or south. The two later fold phases are 

nonpenetrative and have variable axial planes.

The geometry of isoclinal F! folds whose axes plunge 

consistently down the dip of their axial surfaces is 

consistent with the model advocated by Ave Lallemant 

and others (1980). In this model, folds develop whose

axes are originally horizontal and perpendicular to the 

direction of transport in the planes of variable shear 

parallel to the plane of subduction. These early folds 

are subsequently tightened by flattening and, as the 

degree of extension increases down the subduction 

plane, have their axes rotated into the direction of 

maximum extension (parallel to the direction of subduction). Rotation would be greatly facilitated by the 

incorporation into the shear system of incompetent 

serpentinite between the more competent blocks to 

create the melange of Mine Ridge.

Once the serpentinite was in place, subsequent deformation would be largely absorbed by the serpentinite, the blocks acting only passively as the whole 

melange changed shape. The blocks would rotate but 

would carry little or no internal record of any later 

regional strain. Ave Lallemant and others (1980) 

demonstrated the regionally consistent north-south 

trend of the F! fold axes (and east-striking axial surfaces), which implies a north-south direction of subduction. However, the original direction of subduction 

could have been very different. First, if the area of 

60° clockwise Tertiary rotation of the Wallowa Mountains block (Wilson and Cox, 1980) extended as far as 

the Mine Ridge area, then the original direction of 

compression and subduction resulting in the D! 

strain pattern would have been N. 60° W. Second, the 

SQ-S! plane appears to have been rotated subsequently into parallelism with the Late Jurassic (D2) 

deformation.

The second (D2) episode of regional deformation 

recognized by Ave Lallemant and others (1980) occurred in the Late Jurassic. D2 is responsible for the 

tight chevron folds in the bedded turbidites of the 

Weatherby Formation. The D2 folds have steep eastto northeast-striking axial surfaces and horizontal 

fold axes. Using the asymmetry of the minor folds, 

the bedding/cleavage relations, and the abundant 

sedimentary structures that indicate tops and bottoms of beds, larger fold closures can be recognized 

(fig. 11.14). The D2 deformation appears to represent 

north-south to northwest-southeast shortening with 

maximum extension in a vertical direction.

Although the D2 deformation is relatively severe in 

the Jurassic sedimentary rocks, no trace of it is recognized in the Permian blocks of schistose rocks in the 

melange of Mine Ridge (although it is possible that 

the F3 crenulation folds in these schists are associated with the D2 event). It is assumed that the D2 

strain was absorbed by the incompetent serpentinite 

matrix of the melange, passively rotating the blocks 

to produce steeper S^ axial surfaces and F± fold axes.

The contact between the Jurassic Weatherby Formation and the older Permian rocks is also the
11. GEOLOGY OF THE NORTHERN PART OF THE IRONSIDE MOUNTAIN INLIER, NORTHEASTERN OREGON 453

boundary between the Baker terrane and the Izee 

terrane. All exposures of this boundary appear to be 

faulted, and the original relations must remain somewhat speculative. But as noted by Brooks and Ferns 

(1979) the Weatherby Formation adjacent to the older rocks is often much coarser grained than elsewhere, containing large clasts of serpentinite, chert, 

quartzite, and other rock types clearly derived from 

the adjacent melange. A thin strip of the Weatherby 

Formation also lies along the northern side of the 

melange (fig. 11.1). We conclude that the Weatherby 

was deposited unconformably on the exposed Permian melange and that the contact has been only 

slightly modified by faulting. A significant degree of 

thrusting is not apparent, but the slight asymmetry 

of the cleavage with respect to the fold limbs of the 

D2 folds may imply some small component of thrusting to the northwest.

The source of the Weatherby Formation exposed in 

the Ironside Mountain inlier is predominantly the 

ophiolitic melange and island-arc rocks. To the south 

the Weatherby is derived from the similar island-arc 

rocks of the Olds Ferry terrane (Huntington arc; 

Brooks, 1979b). If the ophiolitic rocks represent the 

roots of the primitive island-arc systems associated 

with the subduction zone that created the melange, 

and the younger Jurassic sediments are derived from 

the same older units, then it is probable that each of 

the four "terranes" of the Blue Mountains province 

were formed in proximity to each other, although 

compressed by the subduction and subsequent accretion processes. Hillhouse and others (1982) demonstrate that the Wallowa terrane was formed offshore 

but in approximately the same latitude relative to 

the North American craton as at present. We can, 

therefore, surmise that the Blue Mountains province 

originated as a complex series of island-arc systems 

lying off the western coast of North America and that 

these were compressed by inter-arc subduction from 

the Permian through the Triassic (Dj) and subsequently by accretion to the craton in the Jurassic 

through the Early Cretaceous (D2) (Fleck and Criss, 

1985). Use of the term "terrane" for these various 

units may be criticized in that although each unit 

does possess discrete stratigraphic sequences, the 

units originated in proximity to each other and are 

not wholly unrelated.

The D2 deformation is cut by undeformed intrusions of Late Jurassic to Cretaceous age (Armstrong 

and others, 1977; Fleck and Criss, 1985). The Late 

Jurassic intrusions often lie across contacts between 

the various "terranes," which must, therefore, have 

been in their present relative positions by the end of 

the Jurassic. Calc-alkaline igneous activity continued

intermittently in northeastern Oregon through the 

Miocene. The Ironside Mountain inlier contains a 

Cretaceous (approximately 120 Ma) intrusion at its 

south end (Thayer and Brown, 1964). At its north 

end are the granodiorites of Bullrun Creek and 

Grouse Creek; their ages 36 to 34 Ma are similar 

to those of the calc-alkaline volcanic rocks of the 

eastern part of the Clarno Formation (Urbanczyk 

and Lilligren, 1990; Walker, 1990), which flank the 

inlier on either side. The eastern part of the Clarno 

Formation is overlain by the Miocene calc-alkaline 

Strawberry Volcanics (Robyn, 1977).

The northern part of the Ironside Mountain inlier is 

intruded by two main stocks the granodiorites of 

Bullrun Creek and Grouse Creek and numerous associated smaller intrusive bodies dated at between 

36 Ma and 34 Ma. The various intrusions are distinguished by their textures and associated hydrothermal 

activity but are mineralogically and chemically very 

similar. Their low-K tonalite-trondhjemite composition 

is cited by Barker (1979) as typical of intrusions in 

accreted terranes on the oceanic side of large calcalkaline batholiths emplaced along cratonic margins. 

The chemical signature of these early Oligocene intrusions is similar in some critical respects to the older 

Permian and Triassic igneous rocks of the inlier, and 

it is possible that they were generated, at least in part, 

by the partial melting of the older accreted rocks.

The younger of the two large stocks and the porphyritic dikes related to it are associated with gold 

and base-metal mineralization in the vicinity of the 

peridotite of Bullrun Creek. The main stock, intruding the Weatherby Formation farther south, forms 

the core of a large northwest-elongated alteration 

zone that is associated with a silver and base-metal 

enrichment. This alteration zone is the first Tertiary 

porphyry system documented in northeastern Oregon.

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3689.

12. PETROLOGY AND DEFORMATION HISTORY

OF THE BURNT RIVER SCHIST AND ASSOCIATED PLUTONIC ROCKS 

IN THE BURNT RIVER CANYON AREA, NORTHEASTERN OREGON

By ROGER P. ASHLEY

CONTENTS

Abstract_______________________________ 457

Introduction 458

General geology 459

LitholOETV ________________________________________________________________________ 461

Deer Creek phyllite unit 461

Metasedimentary rocks 461

Metamorphosed volcanic and volcaniclastic rocks 463

Conditions of deposition, age, and correlation 463

Campbell Gulch phyllite unit 467

Phyllitic rocks and quartzite 467

Marble ___________________________ 467

Metavolcanic rocks 467

Conditions of deposition, age, and correlation 469

Intrusive rocks __________________________ 469

Blue Spring Gulch pluton 469 

Other metamorphosed intrusions in the Deer Creek

phyllite unit _______________________ 470 

Metamorphosed intrusions and serpentinite in the

Campbell Gulch phyllite unit 471

Altered mafic dikes 471

Lamprophyre dikes 471

Pedro Mountain stock 472

Igneous petrology 473

Metavolcanic rocks 473

Blue Spring Gulch pluton 474

Metamorphic petrology 475

Structural geology .- 477

Mesoscopic structural features 477

Deer Creek phyllite unit 477

Blue Spring Gulch pluton and other deformed

intrusive rocks 480

Altered mafic dikes 481

Campbell Gulch phyllite unit - 481

Country rocks of the Pedro Mountain stock 482

Macroscopic structure 483

Deer Creek phyllite unit 483

Campbell Gulch phyllite unit 485

Regional distribution and relations of the Burnt River Schist 486

Cave Creek fault and distribution of the units of the

Burnt River Schist - - - - 486

Relation of the Burnt River Schist to surrounding units - 489

Discussion and conclusions 490

References (jitecj ____________________________ 493

ABSTRACT

The Burnt River Schist is a lithodemic unit that constitutes the 

southeastern part of the Baker terrane in the Blue Mountains region of Oregon. The Burnt River Schist forms an east-northeasttrending belt 6 to 22 km wide that extends for almost 90 km from 

Bald Mountain through Burnt River Canyon into western Idaho. 

The Burnt River Schist in Burnt River Canyon consists of two 

mappable metamorphic rock units that are separated by a highangle fault of pre-Tertiary age that trends east-west through the 

center of the area.

The metasedimentary rocks on the north side of the high-angle 

fault, called the Deer Creek phyllite unit, are thinly layered phyllitic quartzite and quartz phyllite, locally interbedded with tuffs 

and flows of intermediate to mafic composition. Mineral assemblages reflect conditions of low-grade metamorphism. Structural 

analysis of the phyllitic rocks reveals two fold systems. The tight 

folds of the earlier system trend east-northeast and plunge at low 

angles to the west; axial-plane cleavage dips steeply. The later fold 

system is nearly coaxial with the earlier one, but axial surfaces 

dip at low to high angles to the north. Concordant and discordant 

metamorphosed bodies of hornblende gabbro, diorite, quartz diorite, and leucotonalite of Middle Triassic age intrude these metamorphic rocks. North of Burnt River Canyon, numerous intrusions 

form a composite plutonic mass, herein called the Blue Spring 

Gulch pluton, that separates the Burnt River Canyon metamorphic rocks from the Elkhorn Ridge Argillite. The Deer Creek phyllite unit is lithologically similar to parts of the Elkhorn Ridge 

Argillite.

The metamorphic rocks on the south side of the high-angle 

fault, called the Campbell Gulch phyllite unit, are mainly pelitic 

phyllite, greenstone (metamorphosed mafic lavas), and marble. 

Mineral assemblages generally reflect conditions of moderately 

high pressure low-grade metamorphism. Structural analysis reveals two deformations. The folds of the earlier deformation are 

tight to isoclinal and plunge at low to moderate angles to the west. 

Axial-plane cleavage generally dips at high to low angles to the 

south. Effects of the second deformation are limited to local small 

folds accompanied by a crenulation axial-plane cleavage.

The central part of the exposure area of the Burnt River Schist is 

intruded by several unmetamorphosed quartz diorite stocks of inferred Late Jurassic to Early Cretaceous age. In the southeastern 

part of the Burnt River Canyon area, adjacent to the Pedro Mountain 

stock, metabasalt and phyllite have been thermally metamorphosed 

to amphibolite and schist containing mineral assemblages characteristic of medium-grade metamorphism. The foliation planes of the 

schist are refolded, and veinlets and segregations in the amphibolite

457
458 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

are microfolded around steeply plunging axes. These effects postdate 

the two deformations of the Campbell Gulch phyllite unit of the 

Burnt River Schist.

Both phyllite units contain arc-related volcanic rocks, and the 

rocks of these units could have formed in either a fore-arc or an 

intra-arc depositional environment. The Blue Spring Gulch pluton 

and associated predeformation intrusive rocks form an arc-related 

calc-alkaline suite that dates deformation and metamorphism of 

the Deer Creek phyllite unit as younger than Middle Triassic. Middle and Late Triassic faunal ages for the Campbell Gulch phyllite 

unit date its deformation and metamorphism as younger than 

Late Triassic. Contact metamorphism and local deformation 

around the Pedro Mountain stock show that regional metamorphism and both regional deformations affecting the Campbell 

Gulch phyllite unit are older than middle Early Cretaceous.

Most exposures of the Burnt River Schist to the west of Burnt 

River Canyon are assigned to the Deer Creek phyllite unit, whereas most exposures of the Burnt River Schist to the east of Burnt 

River Canyon are assigned to the Campbell Gulch phyllite unit. 

The Burnt River Schist exposure area, as currently recognized, 

does not include any rocks that cannot be correlated with one of 

these two phyllite units. Thus the Burnt River Schist is composed 

of two metamorphic belts. These belts probably represent different 

parts of a subduction complex, on the basis of conditions of metamorphism and character of deformation of the rocks of each unit.

INTRODUCTION

The Baker terrane is a belt of deformed and metamorphosed sedimentary, volcanic, and plutonic rocks 

of late Paleozoic and early Mesozoic age that trends 

eastward through the center of the Blue Mountains 

region (fig. 12.1; Silberling and others, 1987). The 

most extensive formation in the eastern half of the 

Baker terrane, east of Elkhorn Ridge, is the Elkhorn 

Ridge Argillite (Gilluly, 1937). The most common rock 

types in the Elkhorn Ridge Argillite are argillite, 

chert, tuff, and lava flows. Limestone pods are common and widespread but not abundant. The metasedimentary rocks that constitute the southern part of 

the Baker terrane east of Elkhorn Ridge, although 

lithologically similar to the Elkhorn Ridge Argillite, 

have better developed schistosity and appear to be 

more strongly metamorphosed than any other rocks 

in the region. Gilluly (1937) named these rocks the 

Burnt River Schist for exposures in Burnt River Canyon and suggested that they might be older than the

118° 117°30'

FIGURE 12.1. Location of study area, quadrangles, and geographic features referred to in text. Study area, which is also the area of 

figure 12.2, is shaded. Shading on inset map shows area of outcrops of the Baker terrane, modified from Silberling and others, 1987. 

Outline of Blue Mountains physiographic province from Dicken, 1965.
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 459

Elkhorn Ridge Argillite, which was then thought to 

be of Pennsylvania!! age. The Burnt River Schist has 

been geographically extended to the south and east 

of Burnt River Canyon as a result of subsequent investigations (Wolff, 1965; Prostka, 1967; Brooks, 

1978); as currently recognized, it makes up the 

southeastern part of the Baker terrane, forming a 

belt that extends northeastward from Burnt River 

Canyon to the Snake River canyon, and westward 

from Burnt River Canyon to at least Bald Mountain, 

about 22 km south of Baker City.

In the last 20 years several authors have attempted 

to synthesize the pre-Tertiary tectonic and orogenic 

history of northeastern Oregon and western Idaho by 

incorporating modern concepts of plate tectonics and 

tectonostratigraphic terranes (Vallier and others, 

1977; Jones and others, 1977; Brooks and Vallier, 

1978; Brooks, 1979a; Dickinson, 1979; Coward, 1983; 

Mullen, 1985; Silberling and others, 1987; White and 

others, 1992). In these syntheses, the Burnt River 

Schist and Elkhorn Ridge Argillite together have 

been interpreted in various ways. Some investigators 

consider the two units to be oceanic crust terrane; 

others consider them to be fore-arc terrane. Some propose the two units originated a great distance from 

the North American continent, and others believe 

they originated near their present locations. Regarding regional structural features, the two formations 

have been interpreted by some as parts of a subduction zone complex, and by others as parts of a fold 

and thrust belt unrelated to subduction.

In this paper I summarize petrographic and structural information from an earlier detailed study (Ashley, 1966), incorporate petrologic and geochronologic 

information from subsequent regional studies 

(Mullen, 1985; Morris and Wardlaw, 1986; Walker, 

1981, 1986), and present some new analytical data for 

metamorphosed igneous rocks from the Burnt River 

Canyon area. I discuss the possible depositional setting, tectonic setting, and tectonic history of the Burnt 

River Schist, incorporating all information available 

for the area for which it was named (T. 11 and 12 S., 

R. 41 and 42 E., French Gulch and Lost Basin quadrangles, Baker County, Oreg.). I also attempt to clarify 

regional relations and distribution of the unit.

GENERAL GEOLOGY

The Burnt River Schist exposed in Burnt River Canyon consists of two informal units separated by a 

high-angle fault that trends approximately east-west 

across the study area and crosses Burnt River midway 

between the east and west ends of the canyon (fig.

12.2). This structure is called the Cave Creek fault. 

Topographic depressions, springs, and zones of hydrothermal alteration mark the fault zone where it crosses the valley of Cave Creek about 1 km south of Burnt 

River. Minor movement occurred on this fault during 

late Tertiary time, when the rocks on the south side 

were downthrown. The amount and direction of movement prior to late Tertiary time is unknown.

The northernmost unit of the Burnt River Schist is 

here called the Deer Creek phyllite unit. Typical exposures of this unit can be seen in the canyon of 

Deer Creek, a prominent tributary on the north side 

of Burnt River Canyon (fig. 12.2). Quartz phyllite, 

which is the most common rock type in the unit, is 

well exposed along a secondary road that extends 

northwestward up Deer Creek from Burnt River 

Canyon. The southernmost unit of the Burnt River 

Schist is here called the Campbell Gulch phyllite 

unit; it is named for good exposures of phyllite and 

interbedded metavolcanic rocks and marble in Campbell Gulch, a short, high-gradient tributary on the 

south side of Burnt River (fig. 12.2). Campbell Gulch 

is accessible by unmaintained jeep trails extending 

westward from Cave Creek and southward from 

Burnt River Canyon.

Although phyllite is the most common rock type in 

both units, the Deer Creek and Campbell Gulch phyllite units differ significantly in lithology and metamorphic grade. They both show evidence of two 

deformations, but no deformation is closely similar 

between the two units. In both units the first folding 

is the major folding, primarily responsible for the distribution of rock types. The east-northeast trends 

and westerly plunges of the first-deformation fold 

axes are roughly similar in both units, but the orientations of the second-deformation folds differ greatly 

between the two units.

The Deer Creek phyllite unit is intruded by metamorphosed plutonic rocks. The largest intrusive body, 

which bounds the study area on the north (fig. 12.2), 

is here called the Blue Spring Gulch pluton, for good 

exposures in Blue Spring Gulch, a tributary of Deer 

Creek. Mafic dikes cut both the phyllite units. A 

quartz diorite body called the Pedro Mountain stock 

by Wolff (1965) intruded the Campbell Gulch phyllite 

unit and metamorphosed and deformed the metavolcanic rock and phyllite around its periphery. The 

stock belongs to the Late Jurassic to Early Cretaceous plutonic suite that includes the Wallowa and 

Bald Mountain batholiths (Armstrong and others, 

1977).

Large continuous exposures of Tertiary rocks border 

the study area on the west and northeast (Brooks and 

others, 1976). All Tertiary exposures within the area,
460 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 461

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except those in the vicinity of the Cave Creek drainage, 

are small and isolated (fig. 12.2). The oldest Tertiary 

materials consist of fluvial sedimentary deposits in the 

vicinity of Cave Creek and on the spurs 2 to 3 km south 

of Burnt River, and an andesitic tuff breccia exposed 

mainly along the southern boundary of the study area. 

Geologic units of Miocene age include the Dooley Volcanics (previously called the Dooley Rhyolite Breccia 

by Gilluly, 1937), which is exposed on the north and 

west sides of the study area, and overlying porphyritic 

olivine basalt, which is exposed mainly on the north 

side of the study area. The Dooley Volcanics consists 

mainly of rhyolitic welded ash-flow tuff and rhyolite 

domes and related flows (Evans, 1992).

The Tertiary beds strike approximately east-west 

and dip south at low angles. Most of the numerous 

Tertiary faults trend east-west, with south sides downthrown. In the vicinity of Sinker.Creek, the Cave Creek 

fault trends east-northeast under a mantle of Tertiary 

fluvial sedimentary deposits, and several east-west 

Tertiary faults follow this older fault zone in an en 

echelon pattern. In the Durkee and Bridgeport Valleys, 

bordering the study area on the northeast and southwest sides, respectively, the pre-Tertiary rocks have 

been downfaulted and covered with sedimentary deposits of late Miocene and Pliocene age (Brooks and 

others, 1976).

The most important Quaternary deposits are gravels on the south side of Burnt River Canyon up to 

100 m above present river level. These deposits appear to occupy former river channels rather than 

bedrock terraces. Both the Tertiary gravel and the 

Quaternary gravel that overlie bedrock south of 

Burnt River are auriferous.

LITHOLOGY

DEER CREEK PHYLLITE UNIT 

METASEDIMENTARY ROCKS

Metasedimentary rocks of the Deer Creek phyllite 

unit consist of phyllitic quartzite, quartz phyllite, 

quartzofeldspathic phyllite, pelitic phyllite, calc-phyllite, and marble. Phyllitic quartzite and quartz phyllite, intergradational rock types, are the most 

abundant. They consist of quartzite or quartz phyllite 

beds, 1 to 5 cm thick, separated by pelitic phyllite 

partings as much as 3 mm thick. The quartz-rich layers are 85 to 95 percent quartz that has been recrystallized to an interlocking aggregate of grains 5 to 20 um 

in diameter. The phyllosilicate minerals that form the
462 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

pelitic partings and the remaining 5 to 15 percent of 

the quartz-rich layers are mainly muscovite and less 

abundant chlorite and biotite (fig. 12.3). The phyllosilicate flakes within the quartz-rich layers are arranged 

in thin folia spaced 20 to 200 |Lim apart, producing 

slaty cleavage. Microporphyroblasts of clinozoisite, biotite, chlorite, and garnet are sparse constituents of 

the quartz phyllite. Vestiges of radiolarian tests occur 

in a few samples. Before regional metamorphism, the 

phyllitic quartzite and quartz phyllite were likely thinly bedded chert and siliceous mudstone with shale 

partings.

Phyllitic quartzite and quartz phyllite grade into 

pelitic phyllite with few quartzose interbeds. Pelitic 

phyllite is composed mainly of quartz, muscovite, 

albite, chlorite, and minor biotite and clinozoisite, 

with locally abundant fine-grained carbonaceous 

material. The original rocks were probably mudstone 

or claystone.

Quartzofeldspathic phyllite beds show relict clastic 

texture, with monocrystalline grains of quartz and 

albite ranging from silt size to fine sand size. These 

feldspathic siltstones and sandstones are poorly bedded and have moderately developed slaty cleavage. 

Clinozoisite is an important constituent, forming idioblastic microporphyroblasts. Chlorite is more abundant, and muscovite and biotite less abundant, than 

in pelitic phyllite. Polycrystalline clasts, probably 

originally metachert and fine-grained sedimentary or 

felsic volcanic grains, are prominent in some sandstone beds. Both the monocrystalline and polycrystalline clasts are probably recrystallized.

Calcareous rocks occur as thin beds that do not extend for more than 1 km along strike. They range 

from pure calcite marble with less than 1 percent impurities (carbonaceous material, muscovite, and 

quartz) to quartz marble with as much as 50 percent 

quartz, to calc-phyllite with much quartz and carboMetamorphic minerals

Rock type

U

Relict minerals

o

ffi ex 

O &

Phyllitic quartzite

and quartz phyllite 

Pelitic phyllite

Quartzofeldspathic 

phyllite and silicic metavolcaniclastic rocks

Marble and calc-phyllite

Mafic and intermediate 

metavolcaniclastic rocks

Meta-andesite and 

metabasalt

Metagabbro and metadiorite

Metamorphosed quartzbearing diorite

Metamorphosed quartz 

diorite and tonalite

EXPLANATION

Major constituent (>10 percent) 

Minor constituent (1-10 percent) 

Trace constituent (<1 percent)

Note: Bar across entire column indicates 

mineral present in most or all samples; 

bar across half of column indicates mineral present in some samples.

FIGURE 12.3. Mineral assemblages of Deer Creek phyllite unit of Burnt River Schist.
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 463

naceous material. Calc-phyllite bears thin laminae of 

carbonaceous pelitic phyllite and is easily weathered. 

Pure marble forms massive light-gray or bluish-gray 

outcrops.

METAMORPHOSED VOLCANIC AND VOLCANICLASTIC ROCKS

Rocks of volcanic origin form numerous bodies that 

lie mainly in a belt extending from French Gulch to 

the east end of Burnt River Canyon (fig. 12.2). The 

larger bodies are particularly varied and include 

massive lavas, schistose to semi-schistose volcaniclastic rocks, and minor interbedded metasedimentary 

rocks; flows and volcaniclastic rocks also coexist in 

many of the smaller bodies.

The most common rock types are metamorphosed 

coarse andesitic or basaltic tuff and lithic lapilli tuff. 

Larger clasts in the tuffs show original porphyritic 

textures: the phenocrysts are plagioclase that has 

been replaced by albite containing minor muscovite 

and clinozoisite; the matrix consists of fine-grained 

crystalloblastic albite, muscovite, and actinolite, and 

smaller amounts of chlorite, biotite, clinozoisite, leucoxene (fine-grained anatase, rutile, or sphene), and 

quartz (fig. 12.3). A few clasts contain small ragged 

hornblende relics and a very few contain partly altered clinopyroxene. Metamorphosed dacitic volcaniclastic rocks are also abundant. They lack actinolite 

and have more chlorite and muscovite than the andesitic and basaltic rocks. Non-volcanic clasts, including quartz, chert, calcite, and fine-grained 

sedimentary rocks, do not total more than 10 percent 

by volume of any rock examined. The material between clasts is a crystalloblastic aggregate of 5- to 

20-|im clinozoisite, albite, muscovite, chlorite, and biotite grains, concentrated in irregular folia that produce a semischistose texture.

Fine-grained metamorphosed tuffs are crystalloblastic aggregates of albite and quartz and lesser 

amounts of other metamorphic minerals, including 

the assemblages clinozoisite-actinolite-chloritebiotite, actinolite-chlorite-biotite, and muscovite-chlorite-biotite(-clinozoisite). The mineral assemblages 

suggest that the protoliths of these rocks were andesitic and dacitic tuffs.

The protoliths of the metamorphosed lavas were 

basalt, basaltic andesite, and andesite flows and flow 

breccias. Most are porphyritic, containing abundant 

euhedral plagioclase phenocrysts as much as 2 mm 

long, which now consist of secondary albite with 

abundant inclusions of clinozoisite, muscovite, and 

chlorite. The porphyritic rocks have a few pseudomorphs of actinolite after clinopyroxene phenocrysts, 

and their groundmasses are composed of fine-grained

albite, sphene, actinolite, epidote, and chlorite. A few 

flows show intergranular or intersertal texture, consisting of small plagioclase laths with interstitial aggregates of actinolite, epidote, chlorite, and opaque 

minerals. Some rocks contain relics of pumpellyite 

partly replaced by clinozoisite. Some fragments in 

flow breccia have abundant tiny albite amygdules or 

varioles of leucoxene and actinolite.

Chemical analyses and norms of metamorphosed 

and altered igneous rocks from the Burnt River Canyon area are shown in tables 12.1 and 12.2, and sample site locations and descriptions are shown in table 

12.3. The weight percentages of major-element oxides 

have been recalculated to 100 percent volatile-free 

(normalized) to facilitate comparison with unmetamorphosed igneous rocks, but the total volatile content is also shown to indicate the degree to which the 

rocks have been converted to hydroxyl- and carbonate-bearing mineral phases. Samples 14-6B and 77- 

1A are typical metamorphosed lapilli tuffs from the 

Deer Creek phyllite unit. The composition of sample 

14-6B is that of a basalt or basaltic andesite. Sample 

77-1A contains significant amounts of both modal 

and normative quartz. If the silica content represents 

that of the protolith, it is best classified as a dacite 

(Peccerillo and Taylor, 1976). If the potassium content were also representative of the protolith the rock 

would be a low-potassium dacite, but movement of 

alkalis likely occurred during diagenesis and lowgrade metamorphisni.

Chemical analyses for Deer Creek metalavas include samples 76-2, 155-3, 15-22, 78-1B, and 76-3 (table 12.1). Protoliths of all these rocks except sample 

76-3 could be classified either as basalt or basaltic 

andesite on the basis of their silica contents. They 

are probably best classified as basalt, because they 

contain significant amounts of normative olivine, and 

76-2 contains normative nepheline. Sample 76-3 is 

best classified as an andesite rather than a dacite, 

because it lacks modal quartz. Six of the seven metavolcanic rock samples from the Deer Creek phyllite 

unit have relatively high Na2O and low K2O values. 

Sample 155-3, however, has nearly equal amounts of 

Na2O and K2O; both the altered plagioclase phenocrysts and groundmass of this rock contain abundant 

muscovite. The presence of abundant muscovite and 

normative corundum indicate that potassium metasomatism affected this rock.

CONDITIONS OF DEPOSITION, AGE, AND CORRELATION

The siliceous character of the Deer Creek phyllite 

unit and the lack of shallow-water sedimentological
464 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

or faunal features indicate that the depositional environment was most likely hemipelagic, as noted by 

Coward (1983) for the Elkhorn Ridge Argillite. Limestone bodies in the Elkhorn Ridge area, which are

not completely recrystallized, show shallow-water 

features and were interpreted by Coward as olistostromes. In the Burnt River Canyon study area, however, the limestones of the Deer Creek phyllite unit

TABLE 12.1. Chemical analyses of metavolcanic rocks, Burnt River Canyon area

[Major-element oxide contents are in weight percent, recalculated to 100 percent, volatile-free. FeO* is total iron calculated as FeO. I volatiles is the sum of H2O+, H2O~, and CO2, in 

weight percent. Total is the sum of major-element oxides and volatiles before recalculation. CIPW (Cross, Iddings, Pirsson, and Washington) norms were calculated from the majorelement oxides normalized on a volatile-free basis. Major-element oxide contents for samples 76-2, 155-3, 15-22, 14-6B, 78-1B, 76-3, 77-1A, 37-1, and 35-2 were determined in USGS 

laboratories by wavelength-dispersive X-ray spectroscopy; analysts, A.J. Bartel, D.F. Siems, K.C. Stewart, J.E. Taggart, Jr. FeO, H2O+, H2O~, and CO2 contents were determined by 

various conventional analytical methods; analysts, E.L. Brandt, H.G. Neiman, S.T. Fribble. Trace-element contents, in parts per million, were determined by energy-dispersive X-ray 

spectroscopy; analysts, B.-S. W. King, K.O. Dennen, R.G. Johnson. Data for sample BR-59 from Mullen (1985). Data for samples BR-49, BR-52, and BR-59A from E.M. Bishop 

(written commun., 1983). BR-series samples analyzed by PR. Hooper, Washington State University, Pullman, Wash. For sample locations and descriptions, see table 12.3. n.d., not 

determined; - --, absent or value less than 0.01 weight percent]

Deer Creek phyllite unit

Sample 76-2 155-3 15-22 14-6B 78-1B 76-3 77-1A 37-1

Campbell Gulch phyllite unit

35-2 BR-49 BR-59 BR-52 BR-59A

Major-element oxides (weight percent)

A1203 -----

Fe203 -- 

FeO - - - 

FeO* 

MgO---- 

NaoO 

KnO ...

Ti02--- 

P205 ------ 

MnO - 

£ volatiles 

Total 1 - -

51.8 

16.6 

1.54 

8.86 

10.25 

5.62 

8.76 

5.04 

.35 

1.18 

.14 

.17 

5.41 

102.6

52.3 

21.5 

1.45 

6.73 

8.04 

4.65 

4.96 

3.45 

3.56 

1.03 

.15 

.19 

4.98 

101.9

52.8 

17.6 

1.25 

9.20 

10.32 

6.47 

5.35 

5.15 

.63 

1.12 

.15 

.23 

5.49 

100.6

52.3 

16.6 

1.85 

8.27 

9.93 

7.86 

7.04 

4.75 

.03 

1.09 

.10 

.18 

4.18 

102.9

52.8 

16.9 

2.83 

6.62 

9.17 

5.82 

8.97 

4.22 

.63 

.93 

.11 

.17 

2.93 

103.7

62.0 64.3 54.0 

16.7 15.4 15.9 

1.28 1.90 1.05 

5.57 3.94 7.53 

6.71 5.65 8.47 

3.23 2.75 7.04 

3.54 6.78 8.33 

5.61 3.65 4.02 

.88 .27 1.19 

.82 .79 .79 

.19 .17 .07 

.16 .13 .14 

2.74 3.55 3.28 

101.6 99.7 99.9

54.3 

14.9 

1.43 

6.06 

7.35 

7.89 

12.07 

2.53 

.09 

.58 

.05 

.14 

4.31 

100.2

51.3 

18.6 

4.11 

4.71 

8.40 

5.81 

11.21 

3.20 

.01 

.70 

.11 

.25 

n.d. 

99.9

52.5 

17.8 

4.22 

4.82 

8.61 

6.62 

9.84 

3.21 

.02 

.80 

.09 

.13 

n.d.

99.6

53.8 

16.5 

4.63 

5.54 

9.71 

5.84 

8.76 

3.73 

.02 

.91 

.09 

.18 

n.d. 

99.3

56.6 

16.5 

4.21 

4.82 

8.61 

5.72 

6.92 

4.11 

.03 

.90 

.08 

.14 

n.d. 

99.7

0

Normative minerals (weight percent)

Quartz 

Corundum 

Orthoclase 

Albite ----- 

Anorthite 

Nepheline 

Diopside - 

Hypersthene - 

Olivine 

Magnetite ----- 

Hematite 

Ilmenite ------- 

Apatite 

2.1 

37.8 

21.5 

2.6 

17.4

13.8 

2.2

2.2 

.3

3.3 

21.1 

29.2 

23.6

10.2 

8.2 

2.1

2.0 

.3

3.7 

43.6 

23.1

2.0 

5.8 

17.5 

1.8

2.1 

.3

0.2 

40.2 

23.8

8.5 

8.3 

14.1 

2.7

2.1 

.2

3.7 

35.7 

25.2

15.1 

9.4 

5.3 

3.6

1.8 

.3

10.5 23.5 

5

5.2 1.6 7.0 

47.4 30.9 34.0 

16.3 24.7 21.7

6.4 15.6 

16.1 8.3 9.3 

9.1

1.9 2.8 1.5

1.6 1.6 1.5 

.4 .4 .2

4.1

.6 

21.4 

29.0

24.6 

17.0

2.1

1.1 

.1

---

0.1 

27.2 

36.5

15.1 

16.4 

.02 

3.2

1.3 

.3

2.0

.1 

27.2 

34.1

11.6 

19.9

3.4

1.5 

.2

3.1

.1 

31.6 

28.3

12.0 

19.4

3.5

1.7 

.2

7.1

.2 

34.9 

26.4

6.1 

19.9

3.5

1.7 

.2

Trace elements (parts per million)

Rj, ___

Rb

Sr -

Nb 

Y .

75 

<10 

105

22 

65

740 

45 

135

14 

60

101 

6 

106 

6 

20 

62

50 

225

21 

70

170

255 

<10 

<10 

46

270 76 141 

<10 7 26 

200 345 125 

<10 6 5 

23 31 22 

100 184 57

109 

<5 

58 

9 

20 

51

n.d. 

n.d. 

n.d. 

n.d. 

n.d. 

n.d.

n.d. 

n.d. 

n.d. 

n.d. 

n.d. 

n.d.

n.d. 

n.d. 

n.d. 

n.d. 

n.d. 

n.d.

n.d. 

n.d. 

n.d. 

n.d. 

n.d. 

n.d.

^Total is the sum of major-element oxides and volatiles before normalization; it is not the sum of the calculated values that precede it in each column. 

Based on major-element analyses recalculated to 100-percent volatile-free oxides.
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 465

are completely recrystallized to marble and no original textural features are preserved.

No datable fauna have been recovered from the 

Deer Creek phyllite unit. A metadiorite sample collected from the eastern part of the Blue Spring Gulch 

pluton (same locality as sample 105-15, table 12.2) 

yielded a zircon U-Pb age of about 230 Ma (Walker, 

1986). If this determination represents a crystallization age, the host metasedimentary rocks can be no 

younger than Middle Triassic.

The rock types and their relative abundances in the 

Deer Creek phyllite unit are similar to those described 

in the Elkhorn Ridge Argillite by Gilluly (1937), Pardee 

and Hewett (1914), Brooks and others (1976), Coward 

(1983), and Ferns and others (1987), with the possible 

exception that chert is more abundant and argillite

TABLE 12.2. Chemical analyses of meta-intrusive rocks, Burnt River Canyon area 

[All samples determined in USGS laboratories; for methods, analysts, and explanation of symbols, see table 12.1. For sample locations and descriptions, see table 12.3]

Blue Spring Gulch pluton

Samp e 16-3 106-14 16-4 105-15 106- 13B 105-22 77-22

Other meta-intrusive rocks

78-10 78-3B 14-17C

Dikes

78-2 14-29B

Major-element oxides (weight percent)

RiOo ____ ____

Ai203 

Fe203 ----- 

FeO 

FeO* 

MgO - -

cao ____

NaoO - _

K20 -- ---

TiOo _.l\-f_» __________

PoOc ^2^5 ________ 

MnO

__ volatiles 

Totall ____ -

46.6 

17.0 

3.05 

7.49 

10.23 

9.12 

13.34 

1.93 

.14 

1.05 

.09 

.18 

3.75 

104.2

48.9 

14.8 

1.45 

8.83 

10.13 

10.84 

11.89 

1.91 

.17 

.82 

.09 

.21 

4.45 

101.8

48.5 

14.2 

1.93 

8.38 

10.12 

11.77 

11.05 

2.11 

.61 

1.02 

.15 

.20 

3.49 

103.1

58.3 

19.4 

1.01 

5.02 

5.93 

2.87 

6.60 

5.57 

.11 

.82 

.26 

.09 

3.53 

101.0

59.9 62.1 

18.9 18.2 

1.41 4.13 

3.77 0.92 

5.04 4.96 

2.50 2.68 

6.08 5.64 

6.39 4.96 

.08 .36 

.70 .77 

.20 .22 

.08 .08 

3.28 2.92 

100.7 100.6

70.4 

16.0 

.18 

.51 

.67 

1.22 

4.30 

6.16 

.48 

.61 

.19 

.04 

1.36 

99.4

60.7 

19.4 

1.76 

3.14 

4.72 

2.35 

5.94 

5.39 

.30 

.65 

.20 

.08 

3.21

100.9

64.6 

17.5 

1.29 

3.30 

4.47 

2.94 

3.65 

5.79 

.37 

.36 

.10 

.09 

1.77 

99.6

68.0 

15.3 

2.12 

1.93 

3.84 

1.53 

5.61 

4.15 

.48 

.58 

.16 

.07 

1.54 

101.7

53.5 

16.5 

2.83 

4.82 

7.37 

9.29 

7.25 

3.46 

1.37 

.72 

.17 

.14 

4.15 

103.2

54.0 

17.7 

1.89 

4.95 

6.64 

6.57 

8.91 

3.80 

1.13 

.81 

.17 

.13 

4.12

102.0

Normative minerals (weight percent)

Quartz 

Corundum 

Orthoclase 

Albite ____ __ 

Anorthite 

Nepheline 

Hypersthene 

Olivine 

Magnetite 

Hematite 

Ilmenite 

pai e

___

0.9 

16.4 

37.2

23.0 

1.1 

15.5 

3.8

2.0 

.2

___

1.0 

16.1 

31.4

21.8 

13.2 

12.5 

2.1

1.6 

.2

___

3.6 

17.8 

27.6

21.1 

6.9

17.8 

2.8

1.9

.4

5.2

.7 

47.1 

27.5

2.9 

12.9

1.5

1.6 

.6

5.3 13.8

.5 2.1 

54.1 42.0 

22.7 26.2

5.1 .3 

8.5 12.2

2.0 1.3

1.3 1.5 

.5 .5

23.4

2.8 

52.1 

14.5

4.4 

1.0

.003 

.2 

1.2 

.4

11.0

1.8 

45.6 

27.9

.2 

9.2

2.6

1.2

.5

15.6 

1.1 

2.2 

49.0 

17.4

11.9

1.9

.7 

.2

28.9

2.8 

35.1 

21.8

4.0 

2.9

3.1

1.1

.4

8.1 

29.3 

25.5

7.5 

19.8 

4.8 

3.3

1.4 

.4

6.7 

32.1 

27.9

12.3 

15.2 

1.2

2.7

1.5

.4

Trace elements (parts per million)

ga ___ _______ 

Rb -------

Nb __________ 

Y _ . .

2r ________________

50 

<10 

235 

<10 

12 

32

78 

<5 

93 

6 

15 

46

235 

<10 

80 

<10 

11 

50

86 

8 

325 

6 

14 

78

35 278 

11 6 

490 323 

6 5 

9 13 

56 85

310 

6 

202 

9 

31 

185

127 

<5 

231 

<5 

11 

50

311 

7 

140 

9 

16 

89

176 

17 

556 

6 

13 

88

730 

20 

570 

<10 

12 

100

660 

<10 

305 

<10 

14 

95

^Total is the sum of major-element oxides and volatiles before normalization; it is not the sum of the calculated values that precede it in each column. 

Based on major-element analyses recalculated to 100-percent volatile-free oxides.
466 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

TABLE 12.3. Locations and site descriptions of analyzed samples, Burnt River Canyon area

[Locations can be found on U.S. Geological Survey 7J4-minute topographic quadrangles: French Gulch, Oreg.; and Lost Basin, Oreg. For analyses, see tables 12.1 and 12.2]

Sample

76-2

1553

15-22

14-6B

78-1B

765

77-1A

37-1

35-2

BR-49

BR-59, 59A

BR-52

163

106-14

164

105-15

106-13B

105-22

77-22

78-10

78-3B

14-17C

78-2

14-29B

Protolith 

composition

Basalt

Basalt

Basalt

Basalt tuff

Basalt

Andesite

Dacite tuff

Basaltic andesite

Basaltic andesite

Basaltic andesite

Basaltic andesite

Basaltic andesite

Gabbro

Gabbro

Gabbro

Leucodiorite

Leuco-quartz diorite

Leuco-quartz diorite

Leucotonalite

Leuco-quartz diorite

Leuco-quartz diorite

Tonalite

Basaltic andesite

Basaltic andesite

Latitude

(N.)

44°34.9'

44°34.6'

44°33.9'

44°33.3'

44°33.2'

44°35.2'

44°34.0'

44°31.2'

44°32.8'

44°31.0'

44°31.4'

44°31.2'

44°35.6'

44°34.8'

44°35.4'

44°34.5'

44°34.8'

44°34.5'

44°34.3'

44°33.5'

44°33.2'

44°33.6'

44°33.0'

44°33.3'

Longitude 

(W.)

117°36.9'

117°32.1'

117°37.3'

117°39.1'

117°36.8'

117°36.6'

117°36.9'

117°42.7'

117°43.0'

117°41.0'

117°41.0'

117°41.0'

117°39.4'

117°33.0'

117°38.4'

117°33.7'

117°33.2'

117°33.8'

117°35.2'

117°35.4'

117°34.9'

117°37.2'

117°36.2'

117°36.9'

Site description

Large metavolcanic body comprising mostly flows, crest of ridge west

of Hooker Gulch, Deer Creek phyllite unit, sec. 30 (unsurveyed),

T. 11 S., R. 42 E.

Metavolcanic rock adjacent to Blue Spring Gulch pluton, exposed along

Burnt River Canyon road near east end of canyon, Deer Creek

phyllite unit, sec. 26 (unsurveyed), T. 11 S., R. 42 E.

Metavolcanic body exposed on south side of Deer Creek canyon 1 km

west-northwest of confluence of Deer Creek and Burnt River, Deer

Creek phyllite unit, sec. 31 (unsurveyed), T. 11 S., R. 41 E.

Metavolcaniclastic body, south side of Windy Ridge 1.6 km east of

French Gulch, Deer Creek phyllite unit, NE1/4NE1/4 sec. 2, T. 12 S.,

R. 41 E.

Small metavolcaniclastic and metavolcanic body, west side of Cave

Creek canyon near confluence of Cave Creek and Burnt River, Deer

Creek phyllite unit, NE1/4NE1/4 sec. 6, T. 12 S., R. 42 E.

Large metavolcanic body comprising mostly flows, crest of ridge west

of Hooker Gulch, Deer Creek phyllite unit, sec. 19 (unsurveyed),

T. 11 S., R. 42 E.

Small metavolcanic body, north side of Deer Creek canyon near

confluence of Deer Creek and Burnt River, Deer Creek phyllite unit,

sec. 3 (unsurveyed), T. 11 S., R. 42 E.

Greenstone belt, Burnt River Canyon road near west end of canyon,

Campbell Gulch phyllite unit, NE1/4SE1/4 sec. 17, T. 12 S., R. 41 E.

Moderately large metavolcanic body, east side of Dark Canyon,

Campbell Gulch phyllite unit, SW1/4SE1/4 sec. 5, T. 12 S., 41 E.

Greenstone belt, road along crest of Rooster Comb Ridge, Campbell

Gulch phyllite unit, SE1/4SW1/4 sec. 15, T. 12 S., R. 41 E.

Greenstone belt, north side of Rooster Comb Ridge, Campbell Gulch

phyllite unit, SE1/4NW1/4 sec. 15, T. 12 S., R. 41 E.

Greenstone belt, north side of Rooster Comb Ridge, Campbell Gulch

phyllite unit, NE1/4SW1/4 sec. 15, T. 12 S., R. 41 E.

Blue Spring Gulch pluton, Blue Spring Gulch, SW1/4NE1/4 sec. 23,

T. 11 S., R. 41 E.

Blue Spring Gulch pluton, Burnt River Canyon road near east end of

canyon, sec. 27 (unsurveyed), T. 11 S., R. 42 E.

Blue Spring Gulch pluton, ridge crest north of Deer Creek and east of

Blue Spring Gulch, NE1/4SW1/4 sec. 24, T. 11 S., R. 41 E.

Blue Spring Gulch pluton, Burnt River Canyon road near east end of

canyon, sec. 27 (unsurveyed), T. 11 S., R. 42 E.

Blue Spring Gulch pluton, Burnt River Canyon road near east end of

canyon, sec. 27 (unsurveyed), T. 11 S., R. 42 E.

Blue Spring Gulch pluton, Burnt River Canyon road near east end of

canyon, 150 m west of locality 105-15, sec. 28 (unsurveyed), T. 11 S.,

R. 42 E.

Blue Spring Gulch pluton, north side of Burnt River Canyon, 2.3 km

east-northeast of confluence of Deer Creek and Burnt River, sec. 29

(unsurveyed), T. 11 S., R. 42 E.

Small intrusive body in phyllite, ridge crest 1.1 km west of summit of

Sheep Mountain, sec. 32 (unsurveyed), T. 11 S., R. 42 E.

Small intrusive body in phyllite, road to Sheep Mountain manganese

prospect, south side of Sheep Mountain, NE1/4NW1/4 sec. 4, T. 12 S.,

R. 42 E.

Small intrusive body in metavolcanic and metavolcaniclastic rocks,

northwest side of Burnt River Canyon at confluence of Deer Creek

and Burnt River, sec. 31 (unsurveyed), T. 11 S., R. 42 E.

Hornblende lamprophyre dike in phyllite, northeast side of Cave Creek

canyon, SE1/4NW1/4 sec. 5, T. 12 S., R. 42 E.

Hornblende lamprophyre dike in phyllite, Burnt River Canyon road at

confluence of Burnt River and Cave Creek, NW1/4NE1/4 sec. 6,

T. 12 S., R. 42 E.
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 467

less abundant in the Burnt River Canyon area than in 

most parts of the Elkhorn Ridge Argillite. I concluded 

earlier (Ashley, 1966) that the Deer Creek phyllite unit 

could be included as part of the Elkhorn Ridge Argillite 

on the basis of similar lithology. More recent detailed 

mapping has shown, however, that the Elkhorn Ridge 

Argillite is not a simple stratigraphic unit, but it is a 

composite unit consisting of blocks that have different 

lithologic and structural characteristics (see Coward, 

1983; Ferns and others, 1982; Ferns and others, 1983; 

Evans, chap. 8, this volume). In some places these 

blocks are clearly fault- or melange-bounded. Thus the 

Deer Creek phyllite unit cannot be correlated with the 

entire Elkhorn Ridge Argillite. Whether the Deer 

Creek phyllite unit is correlative with a specific subterrane or block within the Elkhorn Ridge cannot be considered until the entire Elkhorn Ridge has been 

mapped in more detail and more age data are available 

for the Deer Creek phyllite unit.

CAMPBELL GULCH PHYLLITE UNIT

PHYLLITIC ROCKS AND QUARTZITE

Pelitic phyllite is an abundant rock type of the 

Campbell Gulch phyllite unit (fig. 12.2). This rock 

type is poorly exposed, and thin interbeds of other 

rock types are relatively scarce; bedding is seen at 

few outcrops.

The pelitic phyllite was originally claystone or shale 

and is now composed of 30 to 50 percent or more finegrained muscovite and chlorite in about equal proportions (fig. 12.4). The phyllosilicates are strongly 

aligned, resulting in slaty cleavage. Most of the remaining material is crystalloblastic quartz in grains 

smaller than 20 jim in diameter. Some carbonaceous 

material is everywhere present, as are scattered siltsize detrital quartz grains that are probably recrystallized. Silty varieties contain as much as 30 percent siltsize quartz grains. Minor minerals include very fine 

grained crystalloblastic albite and clinozoisite.

Thin lenses of quartzofeldspathic phyllite are found 

throughout the pelitic phyllite, but their aggregate 

volume is very small. They were originally beds of 

sandstone and, less commonly, conglomerate. Most of 

these lenses are in the vicinity of the large U-shaped 

bend in Burnt River, east of Dark Canyon (fig. 12.2). 

The clasts of the conglomerates are as much as 5 cm 

long and are mostly mafic and intermediate volcanic 

rocks, mudstone, and chert. Albite grains appear in 

most samples and form as much as 25 percent of 

some of the metasandstones. Lenses with volcanic 

clasts contain minor actinolite.

Phyllitic quartzite is abundant in a few places, for 

example, in the lower part of Dark Canyon. It is lithologically similar to phyllitic quartzite of the Deer 

Creek phyllite unit, containing more than 90 percent 

quartz, numerous quartz veinlets, and scarce relict 

outlines of recrystallized radiolarians.

A few massive quartzite bodies as much 30 m thick 

and 600 m long form prominent exposures in a small 

area on the east side of Cave Creek. Elsewhere, 

quartzite beds generally less than 3 m thick and 30 

m long are scattered throughout the phyllite. Their 

uniform fine grain size and xenoblastic texture indicate that these rocks are metachert.

MARBLE

Marble forms numerous prominently exposed bodies aligned in belts that trend approximately parallel 

to the foliation of the phyllite. Most marble bodies 

are almost entirely calcite, generally with a grain 

size of about 1 mm. Minor quartz and carbonaceous 

material were the only impurities observed (fig. 12.4). 

At some localities there are quartzite interbeds 1 to 3 

cm thick, spaced 5 to 15 cm apart. The fine-grained 

quartz of these interbeds contains minor muscovite 

and scattered rhombs of ferroan carbonate that are 

pigmented with limonite on the weathered surface, 

giving the interbeds a distinctive pale-brown weathered color. A few calcareous quartzite beds containing 

as much as 30 percent calcite are scattered through 

the phyllite.

METAVOLCANIC ROCKS

Greenstone (metamorphosed mafic lava) with 

minor phyllite forms a broad belt that is adjacent to 

the southern boundary of the study area and encompasses Rooster Comb Ridge ("g," fig. 12.2, hereafter 

referred to as the "greenstone belt"). This belt crosses 

the western part of Burnt River Canyon, where it interfingers with phyllite and marble. Relict textures 

and structures are rare, but pillow structures with 

variolitic rims, tuff breccia, and flow breccia are 

present locally. Schistose metatuff is widespread and 

is locally interlayered with greenstone. The pillow 

structures and associated sedimentary rocks indicate 

that the volcanic accumulation is submarine.

Most samples of greenstone are composed of dense 

aggregates of actinolite, albite, pumpellyite, chlorite, 

leucoxene, and minor quartz and muscovite (fig. 

12.4). Scattered through this crystalloblastic groundmass are subhedral relict clinopyroxene phenocrysts
468 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

as much as 1 mm long, which form as much as 20 

percent of the rock. Most rocks contain patches and 

folded, disrupted veinlets of fine-grained pumpellyite. 

In a few rocks, minor granular clinozoisite coexists 

with pumpellyite, and in a very few, subidioblastic 

clinozoisite grains occur instead of pumpellyite in the 

groundmass, accompanied by clinozoisite-albite veinlets instead of pumpellyite veinlets.

The pillows have massive cores with compositions 

and textures like those described above. The pillow 

margins have variolitic rinds as much as several centimeters thick. Individual varioles consist of crudely 

radial bundles of actinolite, dusty granular leucoxene, and minor albite and pumpellyite. Between the 

varioles are actinolite sheaves, patches and stringers 

of fibrous-appearing pumpellyite, and minor albite.

Porphyritic greenstone with as much as 10 percent 

altered plagioclase phenocrysts is another common 

rock type. In some of these rocks the phenocrysts are 

replaced only by twinned albite, but generally they 

are replaced by albite, pumpellyite, minor muscovite, 

and minor calcite. Groundmass textures and mineral 

assemblages are like those of rocks without plagioclase phenocrysts. Porphyritic greenstone also contains slightly altered clinopyroxene phenocrysts. 

Stilpnomelane is largely restricted to these porphyritic rocks; it forms sprays and matted aggregates in 

veins with albite or calcite, and also forms scattered 

plates in the groundmass. The varioles that rim pillows in these rocks are radiating clusters of albite 

laths, with leucoxene granules and tiny chlorite and 

muscovite flakes; the interstices between the varioles

Rock type

Metamorphic minerals Relict minerals

« c/5 "3

?2 <u C

&

_c

0

T3

O

E

(U

I

GH

O

Quartzite 

Pelitic phyllite

Quartzofeldspathic 

phyllite

Marble

Meta-andesite and 

metabasalt

Mafic and intermediate 

intrusions

EXPLANATION

Major constituent (>10 percent) 

Minor constituent (1-10 percent) 

Trace constituent (<1 percent)

Note: Bar across entire column indicates 

mineral present in most or all samples; 

bar across half of column indicates mineral present in some samples.

FIGURE 12.4. Mineral assemblages of Campbell Gulch phyllite unit of Burnt River Schist.
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 469

contain albite, muscovite, chlorite, and minor pumpellyite and sphene. The protoliths of these porphyritic rocks were more felsic than those of the rocks 

free of plagioclase phenocrysts.

The metatuffs range from actinolite-rich to chlorite-muscovite-rich varieties and thus have a range of 

compositions. Some contain scattered clinopyroxene 

crystals. Actinolite and all phyllosilicates are strongly 

oriented, producing slaty cleavage.

Six chemical analyses for greenstones exposed on 

Rooster Comb Ridge and at the west end of Burnt 

River Canyon are shown in table 12.1. The greenstones are derived from basaltic andesites, on the 

basis of silica contents of about 51 to 57 percent and 

small to moderate amounts of normative quartz or 

normative olivine. As is the case with the Deer Creek 

phyllite unit, primary alkali contents probably have 

been modified by secondary processes. I^O contents 

for the BR sample series (last four columns in table 

12.1) seem anomalously low, suggesting possible analytical bias toward low values. Nevertheless, the 

rocks are likely generally low in I^O. Na2O was 

probably added to some rocks during recrystallization 

under very low grade metamorphic conditions.

CONDITIONS OF DEPOSITION, AGE, AND CORRELATION

The fine-grained clastic rocks that dominate the 

Campbell Gulch phyllite unit represent terrigenous 

debris deposited in moderately deep water. Although 

the phyllite is locally siliceous, chert beds are scarce. 

The relatively coarse clastic interbeds may represent 

shallow-water deposits or turbidites or both. The 

sandstones and conglomerates are immature, containing abundant feldspar and volcanic fragments. 

The metavolcanic rocks are derived mainly from basaltic andesite, locally pillowed, and minor tuff. They 

represent products of arc-related submarine volcanism (see section "Igneous Petrology"). Recrystallization has mostly destroyed original textures in the 

marble bodies, but rarely preserved shallow-water 

fauna suggest these were reefal limestones. Whether 

they were formed in shallow water, then transported 

and deposited in deeper water, is problematic.

All the fossils yet recovered from the Burnt River 

Schist are from the Campbell Gulch phyllite unit. A 

bryozoan collected from a marble pod in SW1/4 sec. 

10, T. 12 S., R. 41 E., French Gulch quadrangle, 

yielded a questionable Permian age (Ashley, 1966), 

whereas marble collected by D.A. Bostwick from the 

same or a nearby pod yielded a pentacrinid columnal 

that he considered to be probably Late Triassic in 

age (Brooks and others, 1976). The belt containing

abundant marble bodies narrows and continues eastward into the Durkee quadrangle, where it was 

called the (informal) Nelson marble by Prostka 

(1967). Conodonts from a marble pod in lower French 

Gulch and from the Nelson marble near Durkee are 

of Middle to Late Triassic age at both localities (Morris and Wardlaw, 1986).

Faunal ages from the Campbell Gulch phyllite unit 

overlap those of the Wild Sheep Creek and Doyle 

Creek Formations in the Snake River canyon area 

(Vallier, 1977); the Martin Bridge Limestone, Hurwal 

Formation, and parts of the Clover Creek Greenstone 

(Gilluly, 1937) and Lower Sedimentary Series of 

Smith and Alien (1941) in the Wallowa Mountains 

(Brooks and others, 1976); the Elkhorn Ridge Argillite (Brooks and others, 1976; Blome and others, 

1986); and volcanic rocks exposed near Huntington, 

Oreg. (fig. 12.1; Brooks and. others, 1976; Brooks, 

1979b; Silberling, 1983). Owing to structural complexity, it is not possible to define a stratigraphic succession for the Campbell Gulch phyllite unit (see 

section "Regional Distribution and Relations of the 

Burnt River Schist") or to correlate it directly with 

any other stratigraphic unit in the region.

INTRUSIVE ROCKS

BLUE SPRING GULCH PLUTON

Numerous intrusions into the Deer Creek phyllite 

unit consist of concordant and discordant bodies 

ranging in areal extent from a few hundred square 

meters to several square kilometers. The largest intrusion, a composite body, bounds the mapped area 

on the north (fig. 12.2). Extensive exposures of this 

body continue northward for about 5 km to Alder 

Creek (fig. 12.1; Brooks and others, 1976). This intrusive mass is here called the Blue Spring Gulch pluton. In addition to this pluton, many other small 

bodies occur in a belt extending from Sinker Creek 

westward across Sheep Mountain to Windy Ridge 

(fig. 12.2).

North of Deer Creek and northwest of Stump Spring 

Butte, the Blue Spring Gulch pluton consists mainly of 

metamorphosed hornblende diorite and gabbro (IUGS 

modal classification, Streckeisen, 1976). East of Hooker Gulch, it includes much metamorphosed quartz diorite, leuco-quartz diorite, and leucodiorite, as well as 

metamorphosed diorite and gabbro. The rocks are medium-grained, consisting mainly of hornblende and 

plagioclase, with hypidiomorphic-granular texture. In 

some of the more mafic varieties hornblende is sub-
470 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

poikilitic around plagioclase laths. The hornblende is 

bluish- to brownish-green, possibly of pargasitic composition, and shows only minor alteration to chlorite 

and clinozoisite or epidote (fig. 12.3). The chief mafic 

mineral in the metamorphosed quartz diorite is also 

hornblende, but it is extensively altered to aggregates 

of Fe-Mg chlorite, and chlorite with minor olive-green 

to olive-brown metamorphic biotite replaces scarce primary biotite. Plagioclase in all rocks is completely recrystallized to albite with variable amounts of 

clinozoisite-epidote minerals and muscovite. Metamorphosed quartz diorite has interstitial quartz, and all 

rocks have minor interstitial sphene.

Metamorphosed gabbro and diorite of the Blue 

Spring Gulch pluton locally show mineral lineation 

defined by long axes of hornblende grains. The 

strength and attitude of this lineation both vary 

markedly within tens of meters. The abrupt changes 

in mineral lineation suggest that the lineation is of 

magmatic origin and the pluton is a dike-and-sill 

complex.

The least abundant rocks of the Blue Spring Gulch 

pluton are metamorphosed leucotonalite (also called 

trondhjemite; Streckeisen, 1976) and quartz-albite 

pegmatite (albite granite and silicified gabbro of Gilluly, 1937). A large number of small metamorphosed leucotonalite dikes and plugs intrude metamorphosed 

quartz diorite, leuco-quartz diorite, diorite, and gabbro 

on the north side of Burnt River Canyon east of Hooker 

Gulch, locally forming as much as 20 percent of the 

exposed intrusive rocks. Individual dikes are less than 

100 m long and 10 m wide. Modal compositions are 40 

to 50 percent quartz and 35 to 55 percent albite; other 

minerals total 5 to 15 percent. Albite crystals 1 to 4 mm 

long are set in a matrix of quartz grains 0.5 to 1 mm in 

diameter. Albite crystals are serrate to embayed on 

contacts with quartz. Muscovite and clinozoisite-epidote minerals form inclusions in the albite crystals. 

Chlorite, biotite, sphene, and scarce opaque minerals 

form scattered grains and patches within the quartz 

aggregates. The CaO content of the only analyzed metamorphosed leucotonalite sample (77-22, table 12.2), 

and the presence of clinozoisite-epidote inclusions in 

albite in all the leucotonalite samples indicate that the 

primary plagioclase was oligoclase. The FeO* (total 

iron calculated as FeO) content of sample 77-22 is 

relatively low and the MgO content is relatively high 

compared to those of most trondhjemites (Barker, 

1979), possibly as a result of metamorphism. The sharp 

contacts, lack of mafic or intermediate-rock relics, and 

lack of any transitional rock type suggest that the 

protoliths of the leucotonalite bodies were of magmatic 

origin and were not a product of soda-silica metasomatism of mafic rocks.

Metamorphosed quartz-albite pegmatites with a 

wide range of quartz content (10 to 60 percent) form 

patches as much as 50 m in diameter scattered widely within metamorphosed diorite and gabbro. These 

patches consist of web-like networks apparently filling fractures. In specimens with low quartz percentage, the quartz is restricted to veinlets. Relatively 

quartz-rich rocks have subhedral albite crystals 5 

mm to 2 cm in diameter surrounded by fine-grained 

quartz and cut by many tiny quartz veinlets.

A small, poorly exposed serpentinite body occurs 

within the metamorphosed gabbro and diorite mass 

located south of upper Deer Creek, which is included 

in the Blue Spring Gulch pluton. The serpentinite 

body is penetratively foliated, possibly as a result of 

shearing. It consists of well-oriented chrysotile partly 

recrystallized to randomly oriented chrysotile, and 1 

percent chromite grains surrounded by thin patchy 

halos of fine-grained talc. The shearing and recrystallization have obliterated relict textures, but the presence of chromite suggests that the protolith was an 

olivine-bearing ultramafic rock. This serpentinite 

body is not obviously associated with a fault. It may 

be a small intrusion, on the basis of its metamorphic 

mineral assemblage, which is lower grade than that 

of the surrounding metamorphosed gabbro and diorite (Evans, 1977).

Observed contacts of the Blue Spring Gulch pluton 

are sharp, generally discordant with the foliation of 

metasedimentary host rocks, and devoid of obvious 

contact metamorphic effects. Within the metamorphosed pluton, intrusions of diorite and quartz diorite, including leucocratic varieties, are generally 

younger than gabbro intrusions, and leucotonalite 

dikes and plugs intrude all other plutonic rock types.

Walker (1986) obtained slightly discordant U-Pb 

ages of 230±1 and 233±1 Ma on two zircon fractions 

from a metamorphosed trondhjemite (sample BR-79- 

1; see Vallier, chap. 3, this volume) collected at the 

same locality as sample 105-15 (leucodiorite; tables 

12.2 and 12.3). These are the only radiometric age 

data available for the Blue Spring Gulch pluton.

OTHER METAMORPHOSED INTRUSIONS 

IN THE DEER CREEK PHYLLITE UNIT

Metamorphosed phaneritic intrusive rocks of intermediate composition make up the small bodies exposed on Sheep Mountain, on Windy Ridge, and 1.5 

km southwest of Stump Spring Butte (fig. 12.2). 

These rocks are medium-grained and hypidiomorphic-granular in texture. Their protolith compositions 

are biotite-hornblende diorite, leucodiorite, and tonal-
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 471

ite, and biotite quartz diorite and leuco-quartz diorite; their color indices range from 10 to 30. Quartz, 

in polycrystalline patches and veinlets that locally replace plagioclase, forms as much as 25 percent of the 

rocks. Even the smallest bodies show internal variations in color index, hornblende-biotite ratio, and 

quartz content. Alteration of hornblende and biotite 

is similar to that in the metamorphosed quartz diorite of the Blue Spring Gulch pluton.

METAMORPHOSED INTRUSIONS AND SERPENTINITE 

IN THE CAMPBELL GULCH PHYLLITE UNIT

Some small intrusions are located within or near 

greenstone bodies. The sizes and shapes of the intrusions within the greenstones are difficult to determine, but they are probably dikes. Those in phyllite 

near the greenstones appear to be sills as much as 30 

m long. These intrusions have ophitic to subophitic 

textures, with as much as 40 percent clinopyroxene 

or hornblende partly altered to chlorite or actinolite 

(fig. 12.4). The remainder is albite with abundant inclusions of pumpellyite or clinozoisite and flakes of 

white mica. These dikes and sills are probably subvolcanic intrusions associated with the lavas and 

tuffs of the Campbell Gulch phyllite unit.

Five serpentinite bodies were mapped in the 

Campbell Gulch phyllite unit, none more than 150 m 

long. The serpentine minerals are chrysotile and lizardite, accompanied by minor magnetite. Several 

bodies contain talc and chalcedony veinlets. One body 

shows subhedral serpentine pseudomorphs with 

mesh structure, about 10 percent poikilitic orangebrown hornblende partly altered to Fe-Mg chlorite, 

and the remainder interstitial chlorite, fine-grained 

serpentine, and talc(?). The original rock probably 

contained subhedral olivine and was a hornblendebearing peridotite. All of these bodies are located in 

fault zones. They probably represent small ultramafic 

masses that were intruded along the faults or tectonically emplaced, or both.

At many localities south and east of Burnt River 

Canyon, the Campbell Gulch phyllite unit hosts 

mafic and ultramafic intrusive bodies clearly older 

than the Jurassic to Cretaceous plutons (Wolff, 1965; 

Brooks, 1978). These intrusive rocks have not been 

radiometrically dated. Unlike the Blue Spring Gulch 

pluton, they commonly contain pyroxene and olivine.

ALTERED MAFIC DIKES

Altered mafic dikes intrude the Deer Creek phyllite 

unit within 1.5 km of the Cave Creek fault. They

range from less than 1 m to as much as 3 m in width 

and are less than 50 m long. Even the widest of the 

dikes appears to be homogeneous, and they show no 

recognizable external contact effects. They are concordant with slaty cleavage of phyllitic host rocks, 

but are deformed (see section "Structural Geology").

The dikes contain as much as 50 percent Mg chlorite, which forms relatively coarse-grained patchy 

networks. Other constituents are anhedral quartz 

and albite, radiating sprays of clinozoisite prisms, 

and scattered sphene granules (fig. 12.5). Chlorite is 

the only constituent abundant in all specimens; percentages of the others vary greatly. Calcite is a minor 

constituent in some samples, as is white mica in 

samples with more than 20 percent albite.

All the constituent minerals are typical of lowgrade metamorphic rocks; however, they are coarsegrained compared to those of the surrounding metamorphic rocks. The dikes do not have recognizable 

relict magmatic textures. They probably represent 

basalt that reacted with hydrothermal fluid after 

crystallization and was completely altered to hydrous 

low-grade assemblages.

LAMPROPHYRE DIKES

Lamprophyre dikes, which are much more abundant than the altered mafic dikes, contain 40 to 45 

percent hornblende and are characterized by panidiomorphic-granular texture. They are widely scattered 

through the Deer Creek phyllite unit and occur in 

the Blue Spring Gulch pluton. They also cut the 

Campbell Gulch phyllite unit at five scattered localities. Individual dikes are either concordant or discordant with the slaty cleavage of phyllitic wallrocks 

and are undeformed. Dike margins show relatively 

fine-grained chilled zones, but there are no visible 

contact effects in adjacent wallrocks. Several well-exposed dikes are banded parallel to wallrock contacts. 

Internal bands with either sharp or gradational contacts are marked by relatively large numbers of chlorite amygdules, whereas outer bands generally have 

no amygdules.

The main constituents of these dikes are olivegreen hornblende crystals, as much as 2 mm long, 

and strongly saussuritized subhedral plagioclase 

laths (30 to 40 percent) with irregular overgrowths of 

clear albite (fig. 12.5). Of the 15 dikes examined in 

detail, 5 have former olivine phenocrysts completely 

replaced by smectite and clinopyroxene phenocrysts 

partly to completely replaced by smectite or epidote. 

Interstitial materials besides albite are potassium 

feldspar (1 to 15 percent), quartz (5 to 15 percent),
472 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

and in some specimens, radiating aggregates of epidote or prehnite. All specimens contain scattered euhedral opaque granules. Small amygdules (2 mm or 

less in diameter) composed of Fe-Mg or Fe chlorite 

occupy as much as 15 percent of the rock.

In some of the hornblende lamprophyre dikes, the 

plagioclase has been altered to nearly clear albite 

rather than saussurite. Interstitial material includes 

quartz, sometimes in myrmekitic intergrowths with 

albite, radiating growths of matted chlorite (Fe-Mg, 

Mg-Fe, and, rarely, Mg types), radiating sprays of clinozoisite, and minor muscovite. In one dike, relatively abundant Mg chlorite not only forms interstitial 

aggregates but extensively replaces hornblende. Thus 

these mafic dikes show various degrees of alteration 

to hydrous mineral assemblages.

Before alteration the dikes probably ranged in composition from basalt to andesite. The two analyzed 

samples (78-2 and 14-29B, table 12.2) are basaltic 

andesite, on the basis of silica contents and the presence of small amounts of normative olivine. Barium, 

strontium, and zirconium are more abundant in 

these two samples than in samples of metamorphosed basalt and basaltic andesite from the Deer 

Creek and Campbell Gulch phyllite units and metamorphosed gabbro and diorite from the Blue Spring 

Gulch pluton.

PEDRO MOUNTAIN STOCK

The metamorphic rocks in the study area are cut 

by an unmetamorphosed medium-grained biotiteand hornblende-bearing quartz diorite intrusion (fig. 

12.2). The main mass of the intrusion lies south of 

the area of figure 12.2, and a smaller exposure area 

extends to the east. Various parts of the intrusion 

were mapped by Wolff (1965), Prostka (1967), and 

Kennedy (1956). Wolff (1965) referred to it as the 

Pedro Mountain stock. The entire pluton is delineated on the regional geologic map of Brooks and others 

(1976). This regional map shows that the Pedro 

Mountain stock is the largest of a group of Jurassic 

to Cretaceous plutons that form a 55-km-long northeast-trending belt extending from Cow Valley Butte 

on the southwest to Big Lookout Mountain on the 

northeast (fig. 12.1). These masses belong to a plutonic suite of quartz dioritic and granodioritic composition that is widely exposed in northeastern Oregon 

and includes the Wallowa and Bald Mountain batholiths. Radiometric ages of these plutons range from 

about 120 to 160 Ma (Armstrong and others, 1977; 

Marvin and Dobson, 1979; Walker, 1986). Walker 

(1986) obtained U-Pb ages of 119±1, 120±1, and 

122±1 Ma on three zircon fractions from a tonalite 

sample collected from the southern part of the Pedro

Rock type

Altered mafic dikes

Lamprophyre dikes

Metamorphic and alteration minerals

ii '£«[.£

O - ,. ^ ."ti !SJ ti T-t O -^

 4-J *^"1 ^H O *, ^ H t»J 1 I

 ^.pqoO^^tu ^4^

^ f^3 P "^ CL *7* . ^ ^

<<UUUW^2d,

 H ~

N

:3

3

a

 

. y

o

Hon

'

Relict minerals

t/2

C u

a (L. c 

2 c 6

 2j DH 2 ^ C

CX S 0 CX Cu

< U ffi O t«

^H^ 

EXPLANATION

Major constituent (> 10 percent) 

Minor constituent (1-10 percent) 

Trace constituent (< 1 percent)

Note: Bar across entire column indicates 

mineral present in most or all samples; 

bar across half of column indicates mineral present in some samples.

FIGURE 12.5. Mineral assemblages of altered mafic dikes and lamprophyre dikes in Burnt River Canyon area.
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 473

Mountain stock, in the northeast corner of the 

Bridgeport quadrangle (fig. 12.1).

On the west and south sides of the Pedro Mountain 

stock, the intrusive contact has a simple arcuate outline and the foliation in the wallrocks is strongly deflected around the intrusion, parallel to the contact 

(Wolff, 1965). In contrast, the wallrock foliation on 

the north side of the intrusion is not strongly deflected, and the steeply dipping contact there has a complex outline with several large reentrants (fig. 12.2). 

The wallrock foliation parallels large segments of the 

contact, and mineral foliation in the intrusion, produced by oriented biotite, shows orientations similar 

to those of nearby wallrock foliation. Mineral lineation in the intrusion, produced by oriented hornblende, is generally steeply plunging. These features 

suggest that the northern part of the intrusion, although modally nearly homogeneous, is probably a 

series of connected semiconcordant sheets or lobes 

that pushed upward into the greenstone belt of the 

Campbell Gulch phyllite unit, guided by wallrock foliation.

IGNEOUS PETROLOGY

METAVOLCANIC ROCKS

Because the alkali contents of the metamorphosed 

igneous rocks have probably been modified, alkali-silica variation and the alkali-lime index cannot be 

used to characterize these rocks. Instead variation diagrams of SiO2 versus FeO*/MgO and FeO* versus 

FeO*/MgO can be used to distinguish rocks with affinities to calc-alkaline series from those with affinities to tholeiitic series (Miyashiro, 1975). Analytical 

data for metamorphosed volcanic and intrusive rocks 

from the study area are plotted on an SiO2-versusFeO*/MgO variation diagram in figure 12.6,. with a 

reference line dividing calc-alkaline and tholeiitic series. Below FeO*/MgO=2, trends for calc-alkaline and 

tholeiitic series converge, so the cluster of analyses 

for metamorphosed basalt and basaltic andesite of 

the Campbell Gulch and Deer Creek phyllite units 

cannot be classified as calc-alkaline or tholeiitic.

Titanium, zirconium, niobium, and yttrium are 

thought to be relatively immobile during low-grade 

alteration and metamorphism (Cann, 1970) and can 

provide insight into the tectonic environment in 

which mafic volcanic rocks formed (Pearce and Cann, 

1973). On both Ti-Zr-Y and Ti-Zr discriminant diagrams (fig. 12.7), data for mafic metavolcanic rocks 

from the Deer Creek and Campbell Gulch phyllite

units plot in the field where data from ocean-floor basalts, low-potassium tholeiites, and calc-alkaline basalts overlap (Pearce and Cann, 1973). On Y-Zr, Ti-Zr, 

and Nb-Zr variation diagrams (not shown; see Pearce 

and Norry, 1979), the data fall close to trend lines for 

various volcanic arcs rather than trend lines for 

within-plate volcanoes. The metavolcanic rocks contain moderate amounts of titanium and relatively 

small amounts of zirconium, falling within the compositional range for these elements in island-arc tholeiites (Baker, 1982; Jakes and White, 1971; see table 

12.1 for data). Most analyses fall in the island-arc 

tholeiite field on a TiO2-MnO-P2O5 discriminant diagram (Mullen, 1983a), as shown on figure 12.8. I conclude that the metavolcanic rocks of both the Deer 

Creek and Campbell Gulch phyllite units are derived 

from arc rocks, but whether either group clearly fol75

65

LLI

o

DC 

LLI 

0.

55

45

D

O

D

FeOVMgO

EXPLANATION

Metamorphosed volcanic rocks of the Campbell Gulch 

phyllite unit

Metamorphosed volcanic and volcaniclastic rocks of 

the Deer Creek phyllite unit

Blue Spring Gulch pluton and associated meta-igneous 

bodies

FIGURE 12.6. SiO2 versus FeO*/MgO variation diagram for 

metavolcanic and meta-intrusive rocks of the Burnt River 

Canyon area. FeO* is total iron calculated as FeO. Boundary between calc-alkaline and tholeiitic fields is from Miyashiro (1975).
474 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Ti (x 10-2)

Zr 50 Y(x3)

o tr

HI 

Q.

LU

O 

O

p

1.5

1.0

0.5

B

+ 

O

50 100 150 200 

Zr CONTENT, IN PARTS PER MILLION

EXPLANATION

Metamorphosed volcanic rocks of the Campbell Gulch 

phyllite unit

Metamorphosed volcanic and volcaniclastic rocks of 

the Deer Creek phyllite unit

FIGURE 12.7. Ti, Zr, and Y discriminant diagrams after Pearce 

and Cann (1973) for mafic metavolcanic rocks (45-54 percent 

SiO2) of Burnt River canyon area. A, Ti-Zr-Y discriminant diagram. Scaling factors (TixlO"2, Yx3) used to bring points into 

center of diagram without altering relative positions. Low-K 

tholeiites plot in fields A and B, calc-alkaline basalts in fields B 

and C, ocean-floor basalts in field B, and oceanic-island and continental basalts in field D. B, Ti-Zr discriminant diagram. LowK tholeiites plot in fields A and B, calc-alkaline basalts in fields 

B and C, and ocean-floor basalts in fields B and D.

lows a tholeiitic or a calc-alkaline differentiation 

trend is indeterminate with present data. Mullen 

(1985) also proposed an arc origin for metabasalt 

probably obtained from the Campbell Gulch phyllite 

unit.

BLUE SPRING GULCH PLUTON

Analyses for the Blue Spring Gulch pluton and other 

meta-intrusive rocks that cut the Deer Creek phyllite 

unit (table 12.2) are also shown on the variation diagram (figure 12.6). The six analyses for metamorphosed diorite, quartz diorite, and tonalite samples 

show a calc-alkaline trend. The analysis for sample 77- 

22, a metamorphosed leucotonalite, is not included because its FeO*/MgO is anomalously low. The fact that 

original biotite in metamorphosed quartz diorite samTiO2

MnO(x10) P205 (x10)

EXPLANATION

O 

D

Metamorphosed volcanic rocks of the Campbell Gulch 

phyllite unit

Metamorphosed volcanic and volcaniclastic rocks of 

the Deer Creek phyllite unit

Blue Spring Gulch pluton and associated meta-igneous 

bodies

FIGURE 12.8. MnO-TiO2-P2O5 discriminant diagram for mafic metavolcanic and meta-intrusive rocks (45-54 percent SiO2) of 

Burnt River Canyon area. Scaling factors (MnxlO, P^xlO) 

used to bring points into center of diagram without altering relative positions. Calc-alkaline basalts plot in field A, island-arc 

tholeiites in field B, midocean-ridge basalts in field C, and seamount alkalic basalts and tholeiites in field D. Diagram from 

Mullen (1983a).
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 475

pies is altered to chlorite suggests that potassium may 

be somewhat depleted in these rocks owing to alteration and metamorphism. One metamorphosed leucoquartz diorite sample (78-3B) shows normative corundum, likely the result of alkali depletion and possibly 

calcium depletion as well. All analyzed samples in the 

intrusive series are metaluminous, with the exception 

of sample 78-3B, which is peraluminous. Rocks of basic 

to intermediate calc-alkaline and tholeiitic arc suites 

are typically metaluminous (Brown, 1982).

The mafic minerals in the more felsic members of 

this intrusive series are hornblende and biotite, 

which are typical of calc-alkaline plutonic rocks. The 

only mafic mineral in the metamorphosed diorite and 

gabbro samples, however, is hornblende. Yoder and 

Tilley (1962) showed that magmas of basaltic composition crystallize to form amphibolites at Pn2O greater than 1.4 kbar. Thus these intrusions crystallized 

from hydrous magmas at a minimum depth of about 

5 km. The lack of contact effects indicates that the 

terrane was at elevated temperatures and fluid pressures at the time of intrusion.

Ba, Rb, and Sr contents vary in the analyzed suite 

of intermediate and felsic intrusive rock samples (table 12.2), probably as a result of secondary processes. 

Ba and Sr contents, nevertheless, generally are similar to those found in calc-alkaline andesite and dacite, and island-arc intrusive suites (Jakes and 

White, 1972; Mason and McDonald, 1978; Baker, 

1982). Rb contents are low, providing supporting evidence for K depletion. Zr, which should be relatively 

immobile in secondary processes, is present in the intermediate to felsic intrusive rocks in amounts more 

typical of arc tholeiites than calc-alkaline arc rocks 

(Jakes and White, 1972; Baker, 1982). Nb, which also 

should be relatively immobile, is present in amounts 

characteristic of volcanic-arc magmas (Pearce and 

Gale, 1976; Pearce and Norry, 1979).

Rocks of the Blue Spring Gulch pluton show many 

similarities to intermediate and silicic rocks of the (informal) Sparta complex, about 30 km northeast of 

Burnt River Canyon. Metamorphosed diorite and 

quartz diorite samples from the Burnt River area (table 12.2) are chemically similar to quartz diorite from 

the Sparta complex (Phelps, 1979), except that Na2O 

content is higher in the Burnt River rocks. The metamorphosed leucotonalite of the Burnt River area (sample 77-22, table 12.2) is similar in occurrence, texture, 

and composition to fine-grained low-K2O trondhjemite 

at Sparta described by Phelps (1979), but the Burnt 

River sample has anomalously low FeO* content. The 

southern part of the Sparta complex, consisting of twopyroxene gabbro, hornblende-bearing two-pyroxene 

gabbro, pyroxenite, and peridotite (Prostka, 1962), has

no counterpart in the Blue Spring Gulch pluton. The 

Sparta complex has yielded ages ranging from 213 Ma 

(Vallier and others, 1977) to 253 Ma (Walker, 1981, and 

this volume); this range includes the U-Pb age of approximately 230 Ma determined for a sample from the 

Blue Spring Gulch pluton (Walker, 1986).

The Canyon Mountain Complex consists of harzburgite tectonite overlain successively by plagiogranite intrusions and interlayered variably altered 

basalt, diabase, and keratophyre (Thayer, 1977; 

Gerlach and others, 1981). The only rocks in the 

Canyon Mountain Complex similar to those of the 

Blue Spring Gulch pluton are those of the plagiogranite suite, which includes diorite, quartz diorite, 

tonalite, and trondhjemite (Gerlach and others, 

1981). Thayer (1977) proposed that the Canyon 

Mountain Complex is an ophiolite of oceanic-ridge origin. Himmelberg and Loney (1980) and Gerlach and 

others (1981), however, concluded that it is of islandarc origin. Vallier and others (1977) suggested that 

the complex forms the basement for supracrustal 

sedimentary rocks of the Baker terrane. U-Pb ages of 

279 to 268 Ma for Canyon Mountain intrusive rocks 

(Walker, 1983) indicate that all parts of the Canyon 

Mountain Complex are significantly older than the 

Blue Spring Gulch pluton.

The Blue Spring Gulch pluton is here interpreted 

as an intrusive phase of Middle Triassic arc-related 

calc-alkaline magmatism. Similarly, Phelps and Ave 

Lallemant (1980) concluded that the Sparta complex 

is an intrusive phase of Late Triassic arc-related 

magmatism, on the basis of age data then available, 

which gave an age range of 219 to 213 Ma. Arc magmatism of Middle and Late Triassic age produced the 

younger volcanic-rock-dominated sequences of the 

Wallowa terrane (the upper part of the Clover Creek 

Greenstone (Gilluly, 1937), the informally named 

Gold Creek greenstone of Prostka (1962), the Wild 

Sheep Creek Formation, and the Doyle Creek Formation) and the volcanic rocks of the Huntington Formation of Brooks (1979b) of the Olds Ferry terrane 

(see "Discussion and Conclusions" section; Vallier and 

others, 1977; Brooks and Vallier, 1978; Silberling and 

others, 1987).

METAMORPHIC PETROLOGY

The mineral assemblages of metamorphosed and 

altered rocks of the Burnt River Canyon area are 

summarized in figures 12.3 through 12.5, and 12.9. 

The assemblages for all common rock types of the 

Deer Creek and Campbell Gulch phyllite units are 

summarized in figures 12.3 and 12.4, respectively.
476 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Mineral assemblages in altered mafic dikes and lamprophyre dikes are summarized in figure 12.5. The 

assemblages for five contact metamorphic zones delineated in Campbell Gulch greenstones are summarized in figure 12.9.

Mineral phases were identified mainly by transmitted-light optical microscopy and whole-rock X-ray diffraction. A few mineral separates were prepared from 

metamorphic segregation veinlets and identified by 

powder-camera X-ray diffraction. Several serpentine 

samples were investigated by differential thermal 

analysis. Microanalysis of individual mineral phases 

has not been performed, but the available optical and 

X-ray diffraction data are adequate to define the general conditions of metamorphism.

Mineral assemblages in the Deer Creek phyllite unit 

and in the plutonic rocks that intrude it largely reflect 

conditions of greenschist-facies metamorphism, as

shown by the presence of actinolite in all mafic and 

intermediate metavolcanic rocks and most meta-intrusive rocks and the common occurrence of the biotitemuscovite assemblage in pelitic rocks (Winkler, 1979; 

Liou and others, 1985). In the meta-intrusive rocks, 

the only prominent relict magmatic minerals are hornblende and quartz. In the metamorphosed mafic and 

intermediate volcanic and volcaniclastic rocks, relict 

phases include minor clinopyroxene, hornblende, and 

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two localities. It is probably a relict of lower-temperature metamorphic conditions (prehnite-pumpellyite or 

pumpellyite-actinolite facies) that the Deer Creek 

phyllite unit passed through during prograde metamorphism. Garnet present in some quartz phyllite 

samples in trace amounts is probably spessartine 

formed under low-grade conditions. The pressure

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Major constituent (>10 percent) 

Minor constituent (1-10 percent) 

Trace constituent (<1 percent)

Note: Bar across entire column indicates 

mineral present in most or all samples; 

bar across half of column indicates mineral present in some samples.

FIGURE 12.9. Mineral assemblages of contact-metamorphosed greenstones in the Campbell Gulch phyllite unit of Burnt River Schist.
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 477

range for the greenschist fades is from about 2 kbar or 

less to about 8 kbar, and the temperature range is 

about 320 to 450°C (Liou and others, 1985).

The common presence of pumpellyite and actinolite 

in mafic metavolcanic rocks of the Campbell Gulch 

phyllite unit (fig. 12.4) indicates that mineral assemblages here adjusted to conditions of the pumpellyiteactinolite facies (Liou and others, 1985) and that 

maximum regional temperatures were not as high as 

in the Deer Creek phyllite unit. Lack of reaction of 

stilpnomelane to biotite also indicates lower temperatures of metamorphism in the Campbell Gulch phyllite unit. Relict magmatic clinopyroxene, hornblende, 

and opaque minerals persist in metamorphosed volcanic and intrusive rocks of the Campbell Gulch 

phyllite unit. The pressure range for the pumpellyiteactinolite facies is approximately 3 to 8 kbar, and the 

temperature range is approximately 250 to 350°C.

The regional metamorphic mineral assemblages of 

the Campbell Gulch phyllite unit have been overprinted by contact metamorphic effects of the Pedro 

Mountain stock (fig. 12.9). In greenstone 2 km from 

the intrusive contact, pumpellyite is replaced by clinozoisite, a zone boundary readily seen in the field 

because the characteristic dark green pumpellyite 

veinlets disappear. At about 800 m from the intrusive 

contact, the rocks contain oligoclase instead of albite, 

and all relict clinopyroxene is converted to bluishgreen amphibole, which also replaces fine-grained actinolite in the groundmass. These changes are not 

obvious in the field, but oligoclase-bearing rocks have 

veinlets of plagioclase and clinozoisite that are visible in outcrop. At 600 to 800 m, plagioclase, clinozoisite, and idioblastic bluish-green amphibole begin to 

segregate into lensoid domains parallel to the foliation. Amphibole becomes coarser and more deeply 

colored toward the stock; at about 300 m the rocks 

are gneissose amphibolites with visible amphibole 

grains and abundant plagioclase segregations and 

folded veinlets. Epidote takes the place of clinozoisite 

in these rocks and is prominent to within about 200 

m of the contact, where it disappears and olive-green 

hornblende replaces the blue-green amphibole. A few 

meters from the contact, hornblende is accompanied 

by scattered grains of diopside. The epidote amphibolite- and amphibolite-facies assemblages that occur 

within 800 m of the stock formed at pressures greater than about 2 kbar and temperatures greater than 

about 400°C (Liou and others, 1985).

Pelitic rocks from 1 km to within 300 m of the Pedro 

Mountain stock are mainly muscovite-biotite-quartz 

schists. Spotted schists occur locally within 300 m of 

the contact. The spots, composed of muscovite and 

graphite(?), probably represent altered andalusite.

Serpentinite bodies in both the Campbell Gulch 

phyllite unit and the Blue Spring Gulch pluton are 

composed mainly of chrysotile or lizardite or both. 

Chrysotile and lizardite are stable only under conditions of very low grade metamorphism (Evans, 1977). 

Absence of antigorite suggests that these serpentinized ultramafic bodies postdate the main thermal 

metamorphic event(s).

Altered mafic dikes, although unusual in that they 

contain abundant chlorite and only minor actinolite, 

have low-grade metamorphic mineral assemblages 

that must have formed under temperature and pressure conditions similar to those that affected their 

host rocks (fig. 12.5).

Lamprophyre dikes show little or no alteration of 

apatite, sphene, and opaque minerals, partial alteration of hornblende, partial to complete alteration of 

clinopyroxene, and complete alteration of plagioclase 

and olivine. Although the extent of hydration and recrystallization is variable, alteration assemblages 

suggest that the lamprophyre dikes were exposed to 

conditions of the prehnite-actinolite facies. The temperature range for the prehnite-actinolite facies is 

approximately 250 to 350°C, and the pressure range 

is less than 1 kbar to approximately 2 kbar (Liou and 

others, 1985). Clays in these rocks may have formed 

at temperatures less than 200°C.

STRUCTURAL GEOLOGY

Structural data were collected at a density of 3 or 4 

field stations per square kilometer, for a total of 676 

stations in the study area. For each station, a stereogram was prepared to check the geometric relations 

of the structural data against the relations observed 

in the field.

The notation used here is from Freedman and others (1964): S denotes a planar element, L denotes a 

linear element, F is any fold involving any S as form 

surface, A is a fold axis, and D is a single deformation defined by a set of contemporaneous tectonic 

structural elements. S0 designates bedding, and Sn, 

Ln, Fn, and An refer to planar surfaces, lineations, 

folds, and fold axes related to Dn.

MESOSCOPIC STRUCTURAL FEATURES

DEER CREEK PHYLLITE UNIT

Slaty cleavage (S^ is the dominant S-surface in 

most phyllitic rocks, but bedding (S0) is dominant in
478 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

rocks such as phyllitic quartzites that have interbeds 

with strongly contrasting properties (fig. 12.10A). Sj^ 

is obvious in all rock types except massive metavolcanic and intrusive bodies. It did not form in the 

Blue Spring Gulch pluton, which responded to deformation by irregularly distributed shearing and crushing. A later cleavage (S2) varies greatly in strength 

from place to place, on a scale of tens of meters (fig. 

12.10B). Where S2 is well developed in pelitic phyllite, it is a slaty cleavage, but S2 is never visible in 

quartzite layers. S2 is moderately strong in quartzofeldspathic phyllites and metavolcaniclastic rocks, 

but weak or absent in metalavas.

A strong microcrenulation lineation (L2) pervades 

phyllitic rocks, metavolcaniclastic rocks, and marble 

(fig. 12.10C). It is always parallel to the S1-S2 intersection line and to the axes of minor (F2) folds whose axial 

surfaces are nearly parallel to S2 and whose form surFIGURE 12.10. Mesoscopic structural features of Deer Creek phyllite unit. A, Outcrop of phyllitic quartzite with thin pelitic phyllite interbeds, viewed parallel to Lj and Sj, showing Fj folds. 

Sketch shows trace of bedding (So), which defines Fj folds. 1^ is 

inclined toward observer, and trace of Sj is parallel to L2. Knife

is 8 cm long. B, Outcrop of thinly laminated quartz phyllite, phyllitic quartzite, and pelitic phyllite, viewed parallel to L2, showing 

F2 folds. Sketch shows traces of Sj cleavage, which defines F2 

folds, and S2 cleavage. Bedding is parallel to Sj. Note fanning of 

S2. Hammer is 40 cm long.
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 479

faces are Sj. The So-Sj intersection forms a lineation 

(Lj) that is older than the strong crenulation lineation. 

Wherever S0 is a strong feature, Lj can be distinguished from L2 by careful observation.

FI folds are visible at a few outcrops. Pelitic beds 

show similar-type folds with amplitudes of 15 to 70 

cm, having subsidiary folds with amplitudes as small 

as 1 mm (fig. 12.1QA). Folds with amplitudes from 70 

cm to 5 m are visible in rocks with strong bedding 

features, particularly phyllitic quartzites. Folds are 

tight, with amplitudes ranging from one to four times 

wavelength, and nearly always asymmetrical. Exposures oriented approximately perpendicular to F± 

axes commonly show many small folds with the same 

sense of asymmetry.

F2 folds are locally well developed (fig. 12.10B). Amplitudes range from a few millimeters to 60 m, but for 

most folds, amplitudes do not exceed 1 m, and are 

generally smaller than wavelengths, so the folds are 

open, with no thickening in the axial regions. Locally, 

where F2 folding is very strong, amplitudes exceed 

wavelengths and the folds are similar folds with strong 

S2 cleavage in fans opening outward from fold cores. 

The angle between S2 and the average Sj attitude 

varies within the area and controls the shapes of F2 

folds; the smaller the angle between S2 and average S1? 

the more asymmetrical the fold. Average Sj is the attitude of an envelope containing a given F2-folded Sj 

surface (Turner and Weiss, 1963).

The metavolcaniclastic rocks show moderately to 

well-developed Sj, much like that of the quartzofeldspathic phyllites. S2 is weaker than S1? and locally it is 

absent. Phyllosilicates in the crystalloblastic matrix 

between clasts are concentrated in folia that form S

FIGURE 12.10. Continued. C, Outcrop of quartz phyllite, showing 

strong mullion structure produced by L2 microcrenulation and 

intersecting Sj and S2 cleavages, parallel to L2. Note well-developed cross joints. Hammer, at top of outcrop, is 40 cm long.

and S2. The folia are very poorly developed within the 

clasts, so rocks with coarse protolith textures are semischistose. L2 is well developed, but Lj is difficult to 

observe in most exposures because these rocks are 

poorly bedded, and consequently the line of intersection between S0 and S^ cannot be seen.

Axial ratios of the clasts indicate the amount of apparent strain produced by DI and D2 combined. Where 

L! and L2 are nearly parallel (the angle between them 

less than 20-25°), it is not possible to distinguish 

whether the longest axes of the clasts are parallel to 

L! or L2. The intermediate axes of the clasts lie within 

8} approximately perpendicular to the average of L! 

and L2, and the short axes are perpendicular to Sj_. 

Axial ratios range from 8:2:1 to 3:1.5:1. Where Lj and 

L2 are at a high angle (60° or more), the longest axes of 

the clasts are clearly parallel to LI, the intermediate 

axes perpendicular to Lj within Sj, and the short axes 

perpendicular to Sl5 but the long axes of the clasts 

have been shortened and the short axes lengthened by 

D2 strain, resulting in axial ratios of about 2:1.5:1. 

Since the Lj-L2 angle is generally small and L2 is the 

most prominent lineation throughout the area, the 

clasts at many localities appear to be strongly extended parallel to L2 and the apparent strain they record 

seems to be the result of D2. D1? however, produced 

much of the strain.

The metalavas are generally massive, and they 

show all tectonic structures poorly. Only the originally fine-grained or glassy parts of flow breccias associated with the flows show strongly oriented chlorite 

fabrics that produce conspicuous Sj. Actinolite is 

nearly randomly oriented throughout the rocks.

In phyllitic rocks with weak S2 and no mesoscopic 

F2 folds, marble interbeds are massive, with weak 

Sj. Where S2 is strong and accompanies mesoscopic 

F2 folds in phyllitic rocks, marble beds have an S2 

fabric of flattened calcite grains. L2 is visible only in 

impure marble and calc-phyllite.

Observations of microscopic fabrics, particularly in 

the phyllite and quartz phyllite, show that metamorphic recrystallization continued throughout both Dj 

and D2 deformations. Reorientation of phyllosilicate 

grains from S0 to Sj folia and from both S0 and Si to 

S2 folia was accomplished by recrystallization. 

Throughout the rocks, few phyllosilicate flakes are 

even slightly bent. Biotite microporphyroblasts, seen 

in a few quartz phyllite samples, show weak preferred orientation parallel to Sj, weak to strong preferred orientation parallel to S2, and other 

orientations not clearly related to Dj or D2. All biotite microporphyroblasts are unstrained, regardless 

of size or habit. (See Ashley, 1966, for details of microfabric features.)
480 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

All the quartzose and quartzofeldspathic rocks contain four generations of quartz(-albite) veinlets. PreD! veinlets are parallel to or at various angles to SQ 

and show tight Fj microfolds and refolded fold forms 

produced by D2 (fig. 12.11A). Late-Di veinlets cut the 

rocks at various angles, are gently microfolded with 

axial planes parallel to Si, and show effects of D2 

folding. Post-Dj, pre-D2 veinlets are mostly parallel 

to SQ or Sj_, and they show only F2 microfolds. PostD2 veinlets are parallel to SQ, Sj_, or S2, or have apparently random orientations. They are not folded 

but locally show small offsets along Sj^ or S2 surfaces.

BLUE SPRING GULCH PLUTON 

AND OTHER DEFORMED INTRUSIVE ROCKS

The intrusive rocks are penetratively fractured and 

brecciated. Plagioclase and hornblende show bent 

and broken twin lamellae and deformed cleavage surfaces on the microscopic scale throughout the intrusive bodies. Cataclasite and microbreccia sheets 

(nomenclature of Higgins, 1971) are the only mesoscopic tectonic features visible in the metamorphosed 

intrusions (fig. 12.11B). Individual sheets are typically several millimeters thick, spaced several centimeters apart, and do not persist for more than 3 m. 

They cut and offset early epidote veinlets as much as 

10 cm. The finely crushed matrix of the sheets is 

completely recrystallized to crystalloblastic aggregates of clinozoisite or epidote, albite, chlorite, actinolite, and leucoxene. Most exposures show one or two 

sets of parallel cataclasite sheets, but as many as 

four sets may be visible in a single outcrop. At many 

localities one or two sets of cataclasite sheets strike 

east to northeast and dip steeply, generally similar to 

Sj in the metasedimentary rocks, but direct correlation of these shear surfaces with S-surfaces in the 

phyllite is not possible because no suitable metasedimentary rock exposures were found at the external 

contacts of the large intrusive bodies.

At the margins of the small metamorphosed diorite 

and quartz diorite bodies exposed on Sheep Mountain 

and Windy Ridge, intense pervasive cataclasis was 

followed by development of phyllosilicates, which 

form wispy, irregular folia within the cataclasite 

sheets and have a common orientation throughout 

the rock that may diverge from the orientations of 

the sheets. Near the margins of these intrusive bodies, extremely strong cataclasis plus recrystallization 

of phyllosilicates into folia produced a semischistose 

texture parallel to the slaty cleavage (Sj) of the surrounding metasedimentary rocks. F2 microfolds are 

weakly developed in the semischistose quartz diorites, but no other feature in the metamorphosed intrusions can be definitely attributed to D2 .

1 mm

B

1 mm

FIGURE 12.11. Photomicrographs showing structural features. A, 

Pelitic phyllite of the Deer Creek phyllite unit showing both Fl 

and F2 microfolds. White bands are quartz veinlets, many of 

which pre-date Dj and show refolded fold forms, with folds 

symmetrical to Sj planes refolded about axial planes parallel to 

82- Traces of Sj and S2 and examples of refolded quartz veinlets are shown. B, Metamorphosed leucogabbro. Main constituent is albite with muscovite inclusions (ab); also contains 

hornblende grains (hb), sphene (sp), and epidote veinlets (ep). 

Intensity of deformation increases from right side of view area 

to the left. To right of center, albite and hornblende are fractured; to the left, they are crushed, with a matrix of crystalloblastic epidote, actinolite, and albite. On the extreme left is a 

cataclasite band consisting of fine-grained crystalloblastic epidote, albite, actinolite, and leucoxene, with scattered albite and 

hornblende porphyroclasts.
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 481

ALTERED MAFIC DIKES

The altered mafic dikes are always concordant 

with the slaty cleavage (Sj) of their host rocks. The 

thinner dikes and the margins of the thicker ones 

have weak schistosity parallel to Sj_, produced by 

preferred orientation of chlorite. Where Sj in the 

host rocks forms F2 folds, the dikes are also folded 

and Si in the dikes remains parallel to the dike 

margins and to Sj of the host rocks. Thus the dikes 

were emplaced before D2. Although they could be 

pre-Dj sills, more likely they were emplaced after Sj 

formed in the metasedimentary rocks, then were foliated by renewed compressive stress before Dj ended.

CAMPBELL GULCH PHYLLITE UNIT

The phyllitic rocks have a well-developed slaty 

cleavage, Sj_, generally parallel to bedding. At a few 

outcrops of pelitic phyllite, thin marble or quartzite 

(metasandstone) beds outline Fj folds that show 

greatly thickened hinges and pinched-out limbs (fig. 

12.12A). The Fj folds are tight to isoclinal, and Sj is 

an axial-plane cleavage. Phyllosilicates strongly 

aligned parallel to Sj are evenly distributed through 

most phyllite, which lacks strong bedding features. 

Quartz and plagioclase grains in metamorphosed 

conglomerate and lithic sandstone beds form domains 

nearly free of cleavage folia. Feldspathic metasandstone, which contains abundant quartz and plagioclase grains, is semischistose.

L! is well defined in some exposures by one of several 

features, including intersection of prominent interbeds 

with S1? boudinage structure in thin marble beds interlayered with phyllite, quartz rods that are probably 

remnants of disrupted folded veinlets, stretched clasts 

and calcareous concretions, and crenulations developed near fold hinges in competent beds. More commonly, poorly developed spaced cleavages intersect Sj 

at low angles to produce Lj_. The origin of these spaced 

cleavages is unknown. In some cases they may be parallel to subtle bedding features.

The long axes of stretched clasts are parallel to L1? 

their intermediate axes are perpendicular to Lj in 

the plane of Sj_, and their short axes are perpendicular to Sj. Stretching ratios in the plane of Sj are 

1.5:1 to 3:1, indicating moderate apparent stretching 

parallel to L1} and flattening ratios are 2:1 to 15:1, 

indicating strong apparent flattening in Sj_.

Adjacent to the southern boundary of the study 

area, a second lineation, L2, is strongly developed in 

the phyllitic rocks and rarely developed in metatuff. 

Si is folded to form F2 folds that have a new axialplane cleavage, S2, which intersects Sj to form L2

(fig. 12.12B). S2 is a spaced crenulation cleavage 

produced by thin bands of aligned phyllosilicates 

(fig. 12.12C). Both zonal and discrete cleavages are 

seen in individual samples, but discrete cleavages 

are more common where S2 is relatively well developed (nomenclature of Gray, 1979). Because the 

zonal type of S2 crenulation cleavage is most common, the rock usually breaks along Sj folia, and L2 

thus appears as a series of strongly asymmetric 

wrinkles on Sj surfaces. Mesoscopic F2 folds are 

open flexural folds with amplitudes less than wavelengths (fig. 12.12B). Apparent offsets on S2 surfaces 

consistently move the fold limbs away from the 

hinges enough to produce slight thinning of the 

limbs and thickening of the hinges. These apparent 

offsets are probably the result of dissolution and removal of quartz and feldspar from the phyllosilicate 

domains (Gray, 1979).

North of the greenstone belt C'g," fig. 12.2), no mesoscopic F2 folds exist and no S2 is present, but L2 persists. Sj surfaces are deformed into microcrenulations 

with a consistent sense of asymmetry. Near the Cave 

Creek fault, D2 features are absent. In general, within 

the Campbell Gulch phyllite unit, rocks with F2 and S2 

grade northward into rocks with only L2j and these in 

turn grade into rocks with no D2 features.

Marble bodies generally show vague foliation. Very 

small bodies and the margins of larger bodies are 

moderately to strongly foliated, and the foliation is 

parallel to Sj in the surrounding phyllite. The degree 

to which the constituent calcite grains are flattened 

controls the strength of the foliation. The smaller, 

more highly foliated bodies have an Lj lineation produced by elongation of the grains. The extremely 

tight F! folding and moderate apparent stretching 

parallel to Lj observed in the phyllite suggest that 

the podlike and lensoid shapes of the marble bodies 

are at least partly tectonic features rather than entirely original sedimentary features.

The more quartz-rich rocks, including phyllitic 

quartzite and massive metasandstone, contain many 

quartz veinlets, including strongly folded pre-Dj 

veinlets, weakly folded late-Dj veinlets, and post-Dj, 

pre-D2 veinlets.

Strength of Sj varies greatly in the metavolcanic 

rocks. This variation is related primarily to original 

differences in texture. Lj is generally weak. It results 

from small-scale folding of pre-Di veinlets and of primary structures such as flow layering. Lj is not visible in the most massive metalava, whereas in the 

metatuff it is as strongly developed as it is in the 

phyllite. L2 in metalava is very rare, even though F2 

affects phyllite interbedded with greenstone on the 

south side of the greenstone belt.
482 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

COUNTRY ROCKS OF THE PEDRO MOUNTAIN STOCK

In addition to thermal contact effects, the Pedro 

Mountain stock produced wallrock structures that

postdate all structural features associated with regional deformation and metamorphism. The most 

commonly observed structural effect is folding of veinlets and segregations in the amphibolite of the contact

B

\ Phyllitic 

partings

FIGURE 12.12. Mesoscopic structural features of Campbell Gulch phyllite unit. A, Small isoclinal FI folds shown by marble beds in phyllite. Sketch shows contact between marble and 

phyllite, trace of Sj, and phyllitic partings within marble. Partings mark cores of Fj folds, or 

in some cases bound sheared-off fold limbs. Knife is 8 cm long. B, F2 folds with S2 axialplane cleavage. Sketch shows traces of Sj and S2 . Bedding (S0) is parallel to Sj. Knife at 

lower center is 8 cm long.
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 483

aureole. Networks of post-Dj veinlets appear in greenstone about 1 km from the intrusion. At 600 to 800 m 

the veinlets are microfolded strongly enough to form a 

new lineation. Since the veinlets are segregation features that resulted from heating by the intrusion, this 

lineation must have been produced by mechanical effects of the intruding mass. The lineation persists in 

gneissose amphibolite close to the stock and generally 

plunges steeply.

Biotite-muscovite schist 600 to 800 m from the intrusion has a weak spaced S3 cleavage produced by 

discontinuous mica folia. Sj in schist less than 300 m 

from the intrusion is minutely folded. Much of the 

biotite and muscovite in these rocks is reoriented to 

form S3 , a closely spaced discrete crenulation cleavage parallel to the axial planes of the new folds. The 

intersection of Sj and the new axial-plane cleavage 

(S3) forms a new lineation (L3). Locally the schist has 

three lineations; two of the lineations, interpreted as 

Lj and L2, are folded around late mesoscopic folds 

(F3) with axes equivalent to the new lineation (L3). 

Thus the intrusion locally produced a third deformation as it was emplaced.

MACROSCOPIC STRUCTURE

DEER CREEK PHYLLITE UNIT

Orientations of S1? S2, L1? L2, and L3 in the study 

area are shown in figure 12.13. In the Deer Creek 

phyllite unit (subareas 2 and 3, fig. 12.13), Sj generally

1 mm

FIGURE 12.12. Continued. C, Photomicrograph of pelitic phyllite 

showing spaced S2 cleavage consisting of asymmetric microcrenulations confined to phyllosilicate-rich layers. Thin section 

is perpendicular to L2.

strikes east-northeast and dips steeply. From west to 

east Sj changes from mainly south dips through vertical to north dips, and back to south dips; the area with 

north dips overlaps subareas 2 and 3. Lj plunges west 

to west-southwest at low to moderate angles throughout the area, generally steepening from west to east.

FI folding is characterized throughout the unit by 

tight, upright folds. Because Sj is an axial-plane 

cleavage and Lj is parallel to Fj fold axes, the attitudes of Sj and Lj together show the attitudes of FI 

folds. At many outcrops Fj folds are asymmetric and 

the enveloping surface enclosing S0 forms a small 

angle with Sj. Average orientations of S0 and S1? 

however, are not significantly different throughout 

the Deer Creek phyllite unit. Observations of relatively well exposed areas in the lower part of French 

Gulch and the central part of Deer Creek suggest 

that most macroscopic Fj folds have wavelengths 

from 15 to 300 m. Judging from the size and geometry of the FI folds, the distribution of metavolcaniclastic beds within the Deer Creek phyllite unit must 

primarily reflect stratigraphic variations rather than 

repetition of beds by macroscopic FI folds (fig. 12.2). 

In restricted areas, however, there may be some repetition of units owing to FI folding. Neither Sj nor S2 

shows significant postfolding shearing, and there is 

no evidence of penetrative post-D2 shearing, so Fj 

folds are not significantly disrupted.

S2 in the Deer Creek phyllite unit is quite variable, 

but it generally dips to the north (subareas 2 and 3, 

fig. 12.13). Locally in the center of the study area, S2 

shows horizontal to low dips to the south. At the 

west edge of the area, L2 is approximately horizontal, 

shown by the point density contours on the subarea 2 

diagram of figure 12.13. To the east it becomes steeper, locally dipping as much as 40° W. (subarea 3, fig. 

12.13). F2 folds are not large enough (wavelengths 

generally less than 60 m) to affect the distribution of 

lithologic units depicted at the scale of figure 12.2. Lj 

orientation is generally near that of L2 throughout 

the Deer Creek phyllite unit, but tends to plunge 

more steeply to the west (fig. 12.13). Because the 

angle between Lj and L2 is generally relatively small 

and L2 is parallel to F2 fold axes, the geometry of S0 

approaches that of two coaxially superposed fold systems (Turner and Weiss, 1963).

Variations in the angles between all the S-surfaces 

produce differences in the appearance of outcrops in 

different parts of the area, even in a single rock type. 

Where S0 and Sj are nearly parallel and at a high 

angle to S2, a strong cleavage-mullion structure results (the "pencil" structure described by Gilluly, 

1937; fig. 12.14A). Where bedding in thinly-bedded 

rocks diverges notably from Sj, the mullions have
484

SUBAREA 

1

PETROLOGY AND TECTONIC

EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

EXPLANATION

+ LI ° Pole to 81 

A Pole to ^2

Contact 

Fault

0 1 2 3 KILOMETERS 

117°35'
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 485

oblique terminations, forming "torpedo" shapes (Gilluly, 1937; fig. 12.14B). Where Sx and S2 are relatively close together but S0 and Sj^ diverge, so that 

phyllitic quartzite beds with pelitic phyllite interlayers are cut obliquely by strong cleavage, the rocks 

resemble those in the vicinity of Sumpter called 

"pseudoconglomerates" by Pardee and Hewett (1914; 

fig. 12.14C). They attribute this structure to extensive shearing of fragments of competent beds into incompetent layers, and Coward (1983) states that it 

results from cataclastic deformation where shear 

zones cut thick chert sequences. Rocks in both areas 

consist of rhomboid quartz-rich microlithons bounded 

by phyllitic partings, but observations in the Burnt 

River Canyon area indicate that such rocks can form 

without notable shearing. All prominent S-surfaces 

and lineations show local variations in orientation 

that result from multiple folding. Detailed structural 

analysis by Ashley (1966) shows that all geometrical 

relations can be explained by D! and D2, and there is 

no evidence for significant subsequent penetrative 

deformation.

CAMPBELL GULCH PHYLLITE UNIT

In the Campbell Gulch phyllite unit, Sx strikes 

east to northeast and dips mainly south at moderate 

to steep angles throughout the area north of the

FIGURE 12.13. Orientations of structural features of Burnt River 

Canyon area, shown by summary structural diagrams. Map 

shows boundaries of numbered subareas as heavy lines, solid 

where they follow lithologic contacts, dashed within lithologic 

units. Shaded areas of the map are post-deformation cover 

rocks and intrusive rocks. Diagram for subarea 1 shows contoured density of poles to 162 sets of cataclasite and microbreccia bands in Blue Spring Gulch pluton; contours at point 

densities of 0.5, 1.5, 2.5, 3.5, 4.5, and 5.5 percent. Areas having point densities greater than 3.5 percent are shaded. Diagrams for subareas 2, 3, and 4 show contoured densities of 

147, 61, and 35 orientation measurements of L2, respectively; 

contours at point densities of 0.5, 2.0, 4.0, 6.0, 8.0, and 10.0 

percent. Areas having L2 point densities greater than 6.0 percent are shaded. Measurements for other structural elements 

shown by symbols. Diagram for subarea 5 shows contoured 

density of 55 measurements of L2, contours at point densities 

of 0.5, 2.0, 4.0, and 6.0 percent. Areas having L2 point densities greater than 4.0 percent are shaded. Diagram for subarea 

6 shows contoured density of 29 measurements of L3, contours 

at point densities of 0.5, 2.0, 4.0, 6.0, and 8.0 percent. Areas 

having L3 point densities greater than 4.0 percent are shaded. 

All diagrams are lower-hemisphere Schmidt equal-area projections oriented with north at the top. Point densities for contouring were calculated using a spherical Gaussian weighting 

function (Robin and Jowett, 1986).

greenstone belt (subarea 4, fig. 12.13). On the south 

side of the greenstone belt (southern parts of subareas 5 and 6, fig. 12.13), dips are generally to the 

north. L! generally plunges west to southwest at low 

to moderate angles, with little change along strike 

from one end of the study area to the other. Detailed 

structural analysis (Ashley, 1966) shows that bedding 

is generally parallel to Sj_, as anticipated from the 

isoclinal style of F^ folding commonly observed at 

mesoscopic scale. Structural analysis also indicates 

that fanning of Sx related to Fx folding is important 

on the macroscopic scale. The wide greenstone belt 

may be a complex west-plunging antiform with Sj 

fanning upward and outward from its core (subarea 

5, fig. 12.13).

Many marble exposures in the steep-walled valleys 

of tributaries entering Burnt River from the south 

outline the noses of mesoscopic and macroscopic isoclinal folds. It is clear from the outcrop pattern, however, that the marble horizons are discontinuous 

along strike, probably owing both to depositional 

variations and tectonic stretching. The exposures of 

marble in the walls of Burnt River Canyon just west 

of the confluence of French Gulch and Burnt River 

reveal a westward-plunging complex antiform that 

repeats one or two marble-bearing horizons. At many 

localities, however, it is impossible to determine 

whether numerous marble pods reflect different marble-bearing stratigraphic intervals, or structural repetition of one or a few marble beds.

Along the south side of the area, S2, although visible only in phyllite outcrops, consistently strikes 

about N. 60° E. and dips steeply (subarea 5, fig. 

12.13). Mesoscopic F2 folds are found only in this 

part of the area, also restricted to the patches of 

phyllite. In the northern part of the Campbell Gulch 

phyllite unit (subarea 4, fig. 12.13), S2 is developed 

at only a few localities but L2 is visible at many localities. L2 here has various orientations that tend to 

lie in a plane that strikes east-northeast to northeast 

and dips at a low angle to the southeast, shown by 

the girdle pattern of L2 point density contours on the 

subarea 4 diagram (fig. 12.13). This plane is interpreted to represent an average S2 plane for that part 

of the Campbell Gulch phyllite unit north of the 

greenstone belt. Structural analysis indicates that D2 

had only minor geometric effects on Dj features, as 

might be expected from the restricted occurrence of 

F2 folds (Ashley, 1966).

L3, seen around the Pedro Mountain stock, varies 

greatly in orientation, but generally plunges moderately to steeply (see point density contours on subarea 6 diagram, fig. 12.13).
486 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

REGIONAL DISTRIBUTION AND RELATIONS 

OF THE BURNT RIVER SCHIST

CAVE CREEK FAULT AND DISTRIBUTION OF THE UNITS 

OF THE BURNT RIVER SCHIST

The Cave Creek fault marks a discontinuity in lithology, metamorphic grade, and style and orientation 

of structural elements, although not all these features 

appear different everywhere along the fault. Although

the fault is nowhere well exposed, field work has 

shown that it is a narrow zone with no known associated mylonitization or retrograde alteration.

Thermal metamorphism persisted through Dj and 

D2 in both the Deer Creek and Campbell Gulch phyllite units, and involved higher temperatures during 

both D! and D2 in the Deer Creek phyllite unit than 

in the Campbell Gulch phyllite unit, as described 

earlier. The difference in metamorphic grade across 

the fault is visible in the field only in greenstone outcrops, where it is indicated by the presence of pumpellyite in mafic and intermediate rocks of the 

Campbell Gulch phyllite unit.

The strike of Sj and the plunge of LI are generally 

similar on both sides of the fault, but locally there is 

an abrupt change in the dip of S]_. D2 features in the 

Deer Creek phyllite unit are truncated by the fault, 

but D2 features in the Campbell Gulch phyllite unit 

die out south of the fault. Local D3 features produced 

by the Pedro Mountain stock also die out south of the 

fault, so in the Burnt River Canyon area the age of 

displacement cannot be fixed relative to the Jurassic 

to Cretaceous plutonism. Wherever Tertiary rocks of 

Miocene age or older cover the fault zone, they show 

small displacements, generally south side down. 

Northwest-trending Tertiary faults truncate the Cave 

Creek fault at the west edge of Durkee Valley (fig. 

12.15), and the Pliocene sedimentary rocks of Durkee 

Valley completely cover the fault zone.

An interpretation of (1) the geographic distribution 

of the two units of the Burnt River Schist in the 

eastern part of the Baker terrane and (2) the location 

of the Cave Creek fault to the east and west of the

FIGURE 12.14. Effects of variations in relative orientations of Ssurfaces, Deer Creek phyllite unit. A, Debris from outcrop of 

phyllite, fragments showing strong cleavage-mullion or "pencil" 

structure, produced by S2 at a high angle to S1( and S0 nearly 

parallel to Sj. Hammer is 40 cm long. See also figure 12.10C.

B, Outcrop of quartz phyllite showing obliquely terminated 

mullions bounded by Sj and SQ surfaces, or "torpedo" structure, 

produced by S2 at a high angle to Sj, and S0 at a moderate angle 

to Sj. Sketch shows traces of S0 and Sj. Surface of outcrop is approximately parallel to S2. Knife is 8 cm long.
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 487

Burnt River Canyon area is shown in figure 12.15. 

Tracing the Cave Creek fault along strike to the east 

is made difficult by lack of marker rock types and 

detailed structural information. Existing information 

from the geologic maps of Prostka (1967) and Brooks 

(1978) is not adequate to detect discontinuities in Si 

or L! attitudes or various D2 features.

In the Durkee quadrangle, between the Burnt 

River Canyon area and Big Lookout Mountain (fig. 

12.1), relations are obscured by Jurassic to Cretaceous intrusive bodies and extensive cover of Tertiary 

sedimentary and volcanic rocks. Prostka (1967) 

states that the southwest-trending lobe of the Big 

Lookout Mountain pluton invades a major northeasttrending fault zone. I infer that this zone is the eastward extension of the Cave Creek fault. If this inference is correct, juxtaposition of the Deer Creek and 

Campbell Gulch phyllite units along the Cave Creek 

fault predates the Big Lookout Mountain pluton, 

dated at about 124 Ma (Walker, 1986). The area of 

the Deer Creek phyllite unit shown north of this 

fault on figure 12.15 is mapped by Prostka (1967) as 

his greenschist unit. Rock types described by Prostka 

for this map unit match closely those seen in the 

Deer Creek phyllite unit at the east end of Burnt 

River Canyon, between Sheep Mountain and Durkee 

Valley.

East of the study area, the marble bodies of the 

Campbell Gulch phyllite unit form a single belt that 

is easily traced eastward across Burnt River at Nelson, south of Durkee Valley (equivalent to the informally named Nelson marble of Prostka, 1967), and 

thence east-northeast to the Snake River canyon

FIGURE 12.14. Continued. C, Hand specimen of interbedded phyllitic quartzite and pelitic phyllite cut approximately perpendicular 

to Sj and 82, which are nearly parallel. Shows "pseudoconglomerate" structure, produced by phyllitic quartzite beds (Sg) cut obliquely by strong combined Sj and S2 cleavages, which form 

phyllitic partings.

(Brooks and others, 1976; Brooks, 1978). The Nelson 

marble and the phyllite and greenschist associated 

with it in the Durkee and Mineral quadrangles constitute the eastern extension of the Campbell Gulch 

phyllite unit from the Burnt River Canyon area.

East of Big Lookout Mountain and west of the 

Snake River, Brooks (1978) did not find any well-defined subunit in the Burnt River Schist other than 

the Nelson marble, and he did not recognize any 

major fault within the Burnt River Schist such as 

the Cave Creek fault. Mitchell and Bennett (1979) do 

not show any internal subdivisions or structures in 

the Burnt River Schist on the east side of the Snake 

River. The Deer Creek phyllite unit, therefore, is not 

recognized and may not exist east of Big Lookout 

Mountain. The Burnt River Schist is covered by flows 

of the Columbia River Basalt Group 8 to 12 km east 

of the Snake River.

South of Burnt River Canyon, Wolff (1965) described the Burnt River Schist as a northeast-dipping 

block that was pushed upward and southwestward 

by the Pedro Mountain stock. In this area, Si generally strikes west-northwest and dips north-northeast, 

and L! plunges to the northwest. Major rock types of 

the Burnt River Schist west of Pedro Mountain (fig. 

12.15) consist of metamorphosed sedimentary and 

volcanic rocks in about equal proportions. Metasedimentary rocks include quartz phyllite, pelitic phyllite 

and schist, quartzite (metachert), and minor marble. 

Metavolcanic rocks include metalava and metatuff; 

flows are locally pillowed. On the basis of similar lithology, structural style, and orientation of structures, I conclude that the Campbell Gulch phyllite 

unit continues south into the Bridgeport quadrangle, 

around the south side of the Pedro Mountain stock 

(fig. 12.15; Brooks and others, 1976). In the Bridgeport quadrangle, the metamorphic grade of the Burnt 

River Schist is generally higher than it is in the 

Burnt River Canyon area, owing to thermal contact 

metamorphism by the Pedro Mountain stock and 

other intrusions southwest of Pedro Mountain.

To the west of Burnt River Canyon, extensive Tertiary rocks of the Dooley Volcanics cover the Cave 

Creek fault. Exposures in the southeastern part of 

the Dooley Mountain quadrangle include slate, metamorphosed andesitic tuff and flows, and minor marble (Evans, 1992) that I interpret as part of the 

Campbell Gulch phyllite unit. Exposures extending 

westward from the west-central part of the Dooley 

Mountain quadrangle to the vicinity of Bald Mountain in the Brannan Gulch quadrangle are mainly 

quartz phyllite, phyllite, argillite, and minor marble. 

On the basis of mapping by Evans (1994) and reconnaissance examination of mesoscopic structural fea-
488 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 489

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tures, I include these exposures in the Deer Creek 

phyllite unit. The Cave Creek fault must continue 

westward under Tertiary rocks, passing south of the 

Bald Mountain area. Farther west, Pardee (1941) 

mapped metasedimentary rocks on strike with the 

Deer Creek phyllite unit of the Brannan Gulch quadrangle and included them in the Elkhorn Ridge Argillite. The western limit of the Burnt River Schist 

cannot be denned with the geologic information that 

is currently available.

Geologic maps of the Durkee quadrangle (Prostka, 

1967), the Oregon part of the Mineral quadrangle 

(Brooks, 1978), and the Bridgeport quadrangle (Wolff, 

1965) show that bedding and Sj in the Campbell 

Gulch phyllite unit generally dip steeply to the north. 

The Campbell Gulch exposures in Burnt River Canyon and the Dooley Mountain quadrangle (Evans, 

1992), however, include an appreciable area with 

southerly dips, and therefore one cannot assume that 

stratigraphic tops are to the north. The tight to isoclinal folding and lack of indicators of stratigraphic tops 

in the Burnt River Canyon area appear to be characteristic of all rocks of the Burnt River Schist that are 

assigned to the Campbell Gulch phyllite unit. Consequently, there is no basis for a consistent interpretation of the internal stratigraphy of the Campbell 

Gulch phyllite unit, and correlations with specific 

units elsewhere in the region are not possible.

RELATION OF THE BURNT RIVER SCHIST 

TO SURROUNDING UNITS

Structural and stratigraphic relations between the 

Deer Creek phyllite unit of Burnt River Canyon and 

exposures of the Elkhorn Ridge Argillite to the northwest, north, and northeast are obscured by intrusive 

rocks of the Blue Spring Gulch pluton and Tertiary 

volcanic and sedimentary cover rocks. East of Durkee 

Valley, the exposures of Prostka's (1967) greenschist 

unit that I consider to be part of the Deer Creek phyllite unit are separated from the Elkhorn Ridge Argillite to the north by a mass of metagabbro (fig. 12.15).

Between the Big Lookout Mountain pluton and the 

Snake River, the contact between the Burnt River 

Schist and the Elkhorn Ridge Argillite shown on figure 12.15 is that of Brooks and others (1976). In the 

center of the Mineral quadrangle, Brooks (1978) 

shows a fault at this boundary and shows the same 

fault cutting and displacing Tertiary basalt. Elsewhere in the Mineral quadrangle the boundary is 

shown as a depositional contact. In the Snake River 

canyon the contact is located between a section of 

quartz phyllite with interbedded metavolcaniclastic
490 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIAKY ROCKS, BLUE MOUNTAINS REGION

rocks and pillowed greenstone to the north and a section of interbedded pelitic phyllite and quartz phyllite to the south. The schistosity consistently dips 

steeply northward throughout both sections. Mullen 

(1983b) stated that the contact here is gradational. 

Whether the contact is affected only by minor faulting of Tertiary age or whether significant pre-Tertiary shear displacement is distributed through the 

gradational zone is unknown. On the east side of the 

Snake River, Mitchell and Bennett (1979) show the 

contact between the Elkhorn Ridge Argillite and the 

Burnt River Schist as a high-angle fault (fig. 12.15).

The south side of the exposure belt of the Burnt 

River Schist is in contact with volcanic sandstone 

and siltstone of the Weatherby Formation of Brooks 

(1979b), which is included in the Olds Ferry terrane 

by Silberling and others (1987) and in the Izee terrane by Vallier (chap. 3, this volume) and Ave Lallemant (chap. 7, this volume). These rocks are much 

less strongly folded and metamorphosed than are the 

rocks of the Burnt River Schist, and they are of 

Early and Middle Jurassic age (Brooks and others, 

1976; Brooks, 1979b). Between Burnt River and the 

Snake River canyon, the contact is a north-dipping 

high-angle reverse fault locally called the Connor 

Creek fault (Brooks, 1978, 1979a, 1979b; Roure, 

1982). In the vicinity of Pedro Mountain and south of 

Bridgeport, Wolff (1965) shows the contact as a series 

of high-angle faults with various orientations (fig. 

12.15). West of Bridgeport, the contact is obscured by 

Tertiary deposits. Because the Connor Creek fault is 

cut by a quartz diorite mass that is probably related 

to the Big Lookout Mountain pluton (Prostka, 1967; 

Brooks and others, 1976), the Burnt River Schist and 

the Weatherby Formation must have been juxtaposed 

between Middle Jurassic and Early Cretaceous time.

DISCUSSION AND CONCLUSIONS

The Blue Mountains region has been rotated approximately 60° clockwise since Early Cretaceous 

time (Wilson and Cox, 1980), so the present east to 

northeast trends of the pre-Tertiary formations of the 

region were originally northeast to north. In the following discussion, regional tectonic features are described in their original orientations. Local features 

are described in their present orientations, and their 

original orientations are also indicated.

The available minor-element data show that both 

the Deer Creek and Campbell Gulch phyllite units 

contain arc-related volcanic rocks. Both units could 

have formed in either a fore-arc or an intra-arc depositional environment. The Deer Creek phyllite unit,

which contains abundant siliceous sedimentary materials, more likely formed in a fore-arc environment, 

as Mullen (1985) suggested for the lithologically similar Elkhorn Ridge Argillite. Deposition of the Deer 

Creek phyllite unit in a fore-arc setting is also compatible with the idea presented by White and others 

(1992) that the Wallowa and Baker terranes formed 

as an arc/fore-arc pair over a west-dipping subduction zone (fig. 12.1). The only volcanic rocks in northeastern Oregon that could be contemporaneous with 

the Deer Creek phyllite unit are those in the Windy 

Ridge and Hunsaker Creek Formations of Permian 

age (Vallier, 1977) and the Cougar Creek Complex of 

Vallier (1968) of late Paleozoic age (see Vallier, chap. 

3, this volume; fig. 12.16). All these units are found 

in the Wallowa terrane.

The Campbell Gulch phyllite unit, which contains 

abundant clastic and volcanic rocks, more likely 

formed in an intra-arc rather than a fore-arc environment, as Mullen (1985) suggested for the Burnt River 

Schist (undivided). White and others (1992) proposed 

that westward subduction continued into Middle Triassic time, followed by possibly eastward subduction 

in Late Triassic to Early Jurassic time. The Late Triassic to Early Jurassic subduction produced the arcrelated volcanic rocks of the Olds Ferry terrane. 

Based on the Middle and Late Triassic faunal ages 

for the Campbell Gulch phyllite unit, Campbell 

Gulch volcanism could be contemporaneous with the 

Wild Sheep Creek Formation, Doyle Creek Formation, upper part of the Clover Creek Greenstone, or 

Gold Creek greenstone of the Wallowa terrane, or the 

Huntington Formation of the Olds Ferry terrane (fig. 

12.16). It could therefore be related either to westward or eastward subduction, or both, and its apparent intra-arc setting is explained by its position 

between the volcanic belts of the Wallowa and Olds 

Ferry terranes, even though arc volcanism may not 

have occurred simultaneously both to the east and 

west (Silberling, 1983).

The Blue Spring Gulch pluton and associated rocks 

form an arc-related calc-alkaline suite; they intrude 

the Deer Creek phyllite unit and do not represent 

ophiolitic basement rocks of oceanic origin. The Sparta complex is also an arc-related pluton rather than 

an inlier of oceanic basement (Phelps, 1979). The late 

Paleozoic through Late Triassic age span of volcanic 

rocks in the Wallowa terrane and the range of age 

determinations for arc-related plutons in the Wallowa 

terrane (265-225 Ma) are similar to the range of age 

determinations for plutons in the eastern part of the 

Baker terrane (277-213 Ma; Vallier and others, 1977; 

Ave Lallemant and others, 1980; Walker, 1981, 1982, 

1983, 1986, chap. 6 of this volume), whereas the age
w

Blue Spring Gulch pluton

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Deer Creek 

sedimentation, volcanism

Deer Creek D!,D2, 

metamorphism

Campbell Gulch 

sedimentation, volcanism

Campbell Gulch 

Dj, metamorphism

Basaltic lamprophyre 

P

Big L

Campbell Gulch D2

Deer Creek faulted against Campbell 

Huntington Fm. Weatherby Fm. 

volcanism, sedimentation sedimentation

Campbell Gulch faulted 

against Weatherby Fm.

Windy Ridge and Hunsaker Creek Fms. 

sedimentation, volcanism

Cougar Creek Complex 

volcanism

Age I 

(Ma) 320 300

I L I 

CARBONIFEROUS

I I 

280 260

E I L 

PERMIAN

Wild Sheep Creek and Doyle Creek Fms. 

sedimentation, volcanism

I I I 

240 220 200

|El M I L I E 

TRIASSIC

180 160 

I M I L 

JURASSIC

140 

FIGURE 12.16. Chronology of selected late Paleozoic and Mesozoic geologic events 

in the Blue Mountains region, northeastern Oregon, showing age relations of 

events recorded in the Burnt River Canyon area. Solid horizontal lines indicate 

direct paleontologic or radiometric evidence for age within range shown; dashed 

lines indicate indirect evidence for age within range shown. Heavy vertical bar 

indicates age limit constrained by radiometric data. Arrow indicates approximate 

or poorly constrained age limit; arrow with query indicates age limit unknown. 

Sequential events that share a mutual age limit are connected by a light vertical

line. Analytical ranges of radiometric age determinatio

es. Unit terminology shown in figure includes the follo

phyllite unit (of Burnt River Schist); Campbell Gulch, 

(of Burnt River Schist); Weatherby Formation of B

Complex of Vallier (1968); Huntington Formation of 

include Brooks (1979b), Brooks and others (1976), Im

Walker (1986). Geologic time scale is from Palmer (1

Late.
492 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

span of volcanic rocks in the Olds Ferry terrane is 

much more restricted (fig. 12.16). It therefore seems 

likely that the plutons of the eastern part of the 

Baker terrane, including the Sparta complex and 

Blue Spring Gulch pluton, are related to arc magmatism of the Wallowa terrane.

The extension parallel to fold axes that is associated 

with DI and D2 deformations in the Deer Creek phyllite unit and DI deformation in the Campbell Gulch 

phyllite unit probably represents approximately orogen-parallel extension. Such extension is a feature 

commonly seen in belts of deformed low- to mediumgrade metamorphic rocks (Hobbs and others, 1976). 

The relation between stretching lineations and relative 

plate motions in subduction orogens, however, is not 

necessarily simple (Ellis and Watkinson, 1987).

If the Deer Creek phyllite unit is part of a fold belt 

deformed as a result of simple regional compression, 

the entire exposure area of the Deer Creek phyllite 

unit studied here constitutes the south (originally 

southeast) side of a major F± antiform. Alternatively, 

if it is part of a fold-and-thrust belt, north- (originally 

northwest-) directed overthrusting could account for 

the asymmetric folding. Ave Lallemant (chap. 7, this 

volume), infers a major northwest-directed component 

of tectonic transport for all penetratively deformed 

rocks in the eastern part of the Baker terrane, on the 

basis of microfabric and microtextural kinematic indicators. The greatly varied dips of axial planes of F2 

folds in the Deer Creek phyllite unit probably cannot 

be reliably interpreted as related to any simple regime of thrust faulting.

Mesoscopic and inferred macroscopic D± structures 

in the part of the Campbell Gulch phyllite unit in the 

Burnt River Canyon area are not consistently compatible with any simple structural regime involving 

regional thrusting. DI fabric elements in the Campbell Gulch phyllite unit do not show a consistent direction of rotation or asymmetry. The northwestdirected thrusting inferred by Ave Lallemant (chap. 

7, this volume) for DI in the eastern part of the 

Baker terrane thus best applies to the Elkhorn Ridge 

Argillite and the Deer Creek phyllite unit. As in the 

Deer Creek phyllite unit, the varied dips of axial 

planes of F2 folds in the Campbell Gulch phyllite 

unit are not easily related to regional thrusting.

The ages of the deformations in both units of the 

Burnt River Schist are not closely constrained, but 

deformation could have begun earlier in the Deer 

Creek phyllite unit (fig. 12.16). Deformation in the 

Blue Spring Gulch pluton is correlative with DI in 

the Deer Creek phyllite unit, so the age of the pluton 

dates DI in the Deer Creek phyllite unit as late Middle Triassic at the earliest. Similarly, faunal ages for

the Campbell Gulch phyllite unit constrain the time 

of deformation there to later than Late Triassic. The 

ages of undeformed plutons of the Pedro MountainBig Lookout Mountain belt (about 125-120 Ma; 

Walker, 1986) provide evidence that no regional deformation seen in the Burnt River Canyon area is 

younger than Early Cretaceous, because structures 

and contact metamorphism around the Pedro Mountain stock postdate D2 in the Campbell Gulch phyllite unit, and contact metamorphism around the Big 

Lookout Mountain pluton probably postdates D2 in 

the Deer Creek phyllite unit. Ave Lallemant (chap. 7, 

this volume) concludes that DI in the Baker terrane 

is probably pre-Early Jurassic in age.

The actinolite- and biotite-bearing greenschist mineral assemblages of the Deer Creek phyllite unit 

show the highest temperatures (350-450°C) of regional metamorphism yet recognized in the eastern 

part of the Baker terrane, whereas the pumpellyiteactinolite assemblage of the Campbell Gulch phyllite 

unit records relatively high pressures at relatively 

low temperatures (about 3-8 kbar at 250-350°C; 

Liou and others, 1985). High-pressure/low-temperature metamorphic assemblages are most easily 

formed in subduction zones (Coleman, 1972). Although the greenschist-facies metamorphism of the 

Deer Creek phyllite unit is not diagnostic of tectonic 

environment, presence of pumpellyite relicts in the 

Deer Creek rocks suggests that the prograde pressure-temperature path included a pumpellyite-bearing facies and thus may have involved moderately 

high pressures. Models of thermal relaxation accompanying uplift and erosion of subducted sediment 

wedges predict that high-pressure assemblages can 

be destroyed by retrograde reactions under conditions of the greenschist facies (England and Richardson, 1977). Ernst (1988) showed that the majority of 

exposed blueschist metamorphic belts have undergone pervasive retrograde reaction to greenschistand amphibolite-facies assemblages, but some have 

retained high-pressure assemblages because they migrated slowly back up near their subduction zones, 

retracing their prograde pressure-temperature paths. 

Therefore the prograde metamorphism of both the 

Deer Creek and Campbell Gulch phyllite units probably occurred in a subduction environment, but because the Deer Creek rocks underwent extensive 

retrograde alteration to greenschist-facies assemblages, whereas the Campbell Gulch rocks did not, the 

retrograde metamorphic history of each unit is significantly different.

Differences in style of DI and D2 deformation and 

metamorphic history suggest that the Deer Creek 

and Campbell Gulch phyllite units were deformed in
12. PETROLOGY AND DEFORMATION OF BURNT RIVER SCHIST, NORTHEASTERN OREGON 493

different erogenic episodes, or if they were deformed 

in the same erogenic episode(s), they must have been 

located in different parts of the deformation zone. 

The older age of the Deer Creek rocks permits the 

speculation that the Deer Creek phyllite unit was 

metamorphosed and deformed (D± and possibly also 

D2) during Late Triassic subduction, possibly subjacent to Campbell Gulch strata accumulating at the 

same time (fig. 12.16). Just as it is not possible to 

determine whether volcanic rocks of the Campbell 

Gulch phyllite unit are related to Wallowa terrane 

subduction or Olds Ferry terrane subduction, it is not 

possible to determine whether deformation of the 

Deer Creek phyllite unit, if indeed it occurred in Late 

Triassic time, is related to Wallowa terrane or Olds 

Ferry terrane subduction. The fact that Deer Creek 

D! structures are compatible with north (originally 

northwest)-directed overthrusting or south (originally 

southeast)-directed underthrusting favors eastward 

subduction related to Olds Ferry terrane arc volcanism (White and others, 1992).

The Campbell Gulch phyllite unit was metamorphosed and deformed in Early to Middle Jurassic 

time, probably subjacent to the northwestern (originally western) part of the Weatherby Formation, 

which was accumulating at the same time (fig. 

12.16). This metamorphism and deformation could 

have been produced by eastward subduction related 

to Olds Ferry terrane arc volcanism, especially if the 

subduction zone migrated eastward. However, the 

retrograde metamorphism of the Deer Creek phyllite 

unit, which implies relatively rapid uplift, could better be accommodated by westward subduction beneath the Baker terrane. White and others (1992) 

suggested that westward subduction occurred in Middle Jurassic time, producing renewed arc volcanism 

in the Wallowa terrane.

The Deer Creek phyllite unit was faulted against 

the Campbell Gulch phyllite unit on the Cave Creek 

fault and the Campbell Gulch phyllite unit was faulted against the Weatherby Formation on the Connor 

Creek fault between Middle Jurassic and Early Cretaceous time (fig. 12.16). The maximum age of faulting 

of the Deer Creek phyllite unit against the Campbell 

Gulch phyllite unit is shown on fig. 12.16 as somewhat greater than the maximum age of faulting of the 

Campbell Gulch phyllite unit against the Weatherby 

Formation, mainly to indicate that the Connor Creek 

and Cave Creek faults are not necessarily the same 

age. Although the upper age limit of the Connor Creek 

fault is constrained by the age of the Weatherby Formation, the age of the Cave Creek fault is not similarly constrained, and it could be older or younger than 

the Connor Creek fault. The fact that the Connor

Creek fault is a high-angle reverse fault suggests that 

compressive deformation occurred during the Middle 

Jurassic to Early Cretaceous.

The lamprophyre dikes of basaltic composition in 

the Burnt River Canyon area were emplaced after Dj 

in the Campbell Gulch phyllite unit, but were subjected to conditions of the prehnite-actinolite facies, 

suggesting that they were emplaced before Tertiary 

uplift. Thus extension must have occurred at some 

time between the Middle Jurassic and the early Tertiary. Since the Middle Jurassic, serpentinite has 

been intruded locally along faults in the Baker terrane, including faults that may be of Tertiary age.

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13. GRAVITY STUDIES OF AN ISLAND-ARC/CONTINENT SUTURE ZONE IN 

WEST-CENTRAL IDAHO AND SOUTHEASTERN WASHINGTON

By GREGORY B. MoHL1 '2 and RICHARD L. TniESSEN1

CONTENTS

	Page

_______________________________ 497

Introduction 497

Acknowledgments - 499

Methods 499

Geologic and tectonic elements 501

North American assemblage 501

Accreted-terrane assemblage 501

Boundary zone 501

Columbia River Basalt Group 502

West-central Idaho 504

Southeastern Washington 504

Gravity modeling 508

Interpretation 511

OrtrjfMii sion s --»« -_- ._......_......._......_......_......._...... 512

References cited 513

ABSTRACT

Gravity anomalies in west-central Idaho and southeastern 

Washington correlate well with an island-arc/continent suture zone 

previously defined by initial 87Sr/86Sr ratios and detailed field 

mapping. Using a regional compilation of gravity data supplemented with data collected in southeastern Washington, we correlated 

the observed anomalies with geologic and tectonic elements associated with the suture zone. These correlations are supported by 

theoretical models that establish the gravity signature of geologic 

elements suspected to be causing anomalies in the study area.

The character of the observed gravity anomalies varies within 

the study area. In the northern part, the gravity signature consists 

of a regional high-low paired anomaly with local highs superimposed near the suture zone. In the central part, this signature 

does not correlate well with the suture zone defined by initial 

87Sr/86Sr ratios, and the locally superimposed highs are absent. 

The gravity signature observed in the southern part of the study 

area returns to the regional paired anomaly with a broad gradient 

that steepens in the suture zone. Isolated highs outboard of the 

suture zone were also observed.

Washington State University, Pullman, WA 99164-2812 

2Present address: Conoco, Inc., Casper, WY 82601

Correlation of these anomalies to geologic features, magnetic 

anomalies, and theoretical models indicates that the primary source 

of the lows in the northern part of the study area is probably the 

Idaho batholith. The highs are likely to be related to near-surface 

mafic and ultramafic bodies caught in or near the suture zone. The 

lack of correlation between gravity anomalies and initial 87Sr/86Sr 

ratios in the central part of the study area may be due to the 

relativity coarse level of gravity data available, or it may be related 

to geologic elements that are not reflected in the isotopic analyses. 

Gravity anomalies in the southern part of the study area appear to 

be related to differences in crustal composition or thickness that 

have been overprinted by the effects of the Idaho batholith.

In Washington, geologic control is limited to a few isolated exposures, but detailed gravity data closely define the suture zone by 

extending the anomaly signature observed in Idaho. This investigation provides a basis for better understanding the geometry of 

the suture zone and accurately delineating its location to the west 

where thick sections of the Columbia River Basalt Group prevent 

detection by traditional geochemical and field-mapping techniques.

INTRODUCTION

The recognition of exotic terranes in the northwestern United States has led to the definition of a suture zone between these terranes and the North 

American craton. Where exposed in west-central Idaho, the suture zone is characterized by a steep gradient in initial 87Sr/86Sr ratios (Armstrong and others, 

1977; Fleck and Criss, 1985; Criss and Fleck, 1987) 

(fig. 13.1) and discontinuities in petrologic, metamorphic, and structural features (Hamilton, 1963; Myers, 

1982; Onasch, 1979, 1987; Aliberti and Manduca, 

1988a, b; Davidson, 1988; Lund, 1988; Lund and 

Snee, 1988; Strayer, 1988; Strayer and others, 1989, 

1990; Lund and others, 1990; Blake, 1991; Selverstone and others, 1992; Manduca and others, 1992, 

1993; Blake and Thiessen, in press). In adjacent 

Washington, westward-thickening flood basalts of the 

Columbia River Basalt Group obscure the suture 

zone's surface expression, and the currently used 

techniques are of little value in defining its subsurface location. Accurately locating the suture zone

497
498 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

46°15

10 0 10 20 30 40 50 60 70 KILOMETERS

Line of equal initial 87Sr/86Sr ratio

FIGURE 13.1. Index map of study area showing suture zone defined by lines of equal initial 87Sr/86Sr ratios (modified from 

Armstrong and others, 1977, and Fleck and 

Criss, 1985). Accreted terrane assemblage is 

defined as material south and west of these 

lines, North American assemblage is defined 

as material north and east of the lines, and 

boundary zone is material of undetermined 

affinity between the lines. BM, Blue Mountains; PB, Pasco basin.

- 45°

44 C
13. GRAVITY STUDIES OF AN ISLAND-ARC/CONTINENT SUTURE ZONE IN IDAHO AND WASHINGTON 499

under the basalt is essential to understanding the geometry of the suture zone as well as to developing 

models that explain accretion of exotic terranes.

Detailed mapping of gravity variations is one method by which the suture zone in southeastern Washington may be denned. This approach utilizes differences 

in bulk density that result from juxtaposition of crust 

of differing composition and thickness across the suture zone. The purpose of this paper is to use geologic 

and tectonic elements associated with the suture zone 

to define its gravity signature in west-central Idaho. 

This signature is then used to delineate the suture 

zone in southeastern Washington where the suture 

zone's location is poorly constrained by exposures of 

rock underlying the basalt.

ACKNOWLEDGMENTS

We are grateful for critical reviews by Peter Hooper 

and John Watkinson as well as helpful discussions 

with Paul Myers, David Blake, and Kent Johnson.

METHODS

From various published sources we compiled information about geologic and tectonic elements in the 

study area that may affect gravity measurements. 

Then we correlated these elements with the suture 

zone in west-central Idaho that had been defined previously by geologic and geochemical analyses (Armstrong and others, 1977; Vallier, 1977; Vallier and 

others, 1977; Brooks and Vallier, 1978; Brooks, 1979; 

Myers, 1982; Fleck and Criss, 1985; Criss and Fleck, 

1987; Onasch, 1987; Aliberti and Manduca, 1988a, b; 

Davidson, 1988; Lund, 1988; Lund and Snee, 1988; 

Strayer, 1988; Strayer and others, 1989, 1990; Lund 

and others, 1990; Blake, 1991; Selverstone and others, 

1992; Manduca and others, 1992, 1993; Blake and 

Thiessen, in press). These correlations were supplemented with observations of aeromagnetic data (Zietz 

and others, 1971, 1978). Due to the regional nature of 

this investigation, the geology of the study area has 

been simplified to four lithotectonic assemblages: the 

North American assemblage, the accreted-terranes assemblage, the boundary zone, and the Columbia River 

Basalt Group.

A gravity signature of the suture zone in west-central Idaho was defined by comparing available gravity 

data (Bankey and others, 1984) with the geologic and 

tectonic elements and magnetic anomalies observed in 

the area. This comparison was supplemented with numeric models generated to examine the gravity signature of the various tectonic elements observed in the 

region. The data are directly from Bankey and others 

(1984) and were collected from approximately 1,000 

stations distributed as shown in figure 13.2. Also 

shown in figure 13.2 are approximately 900 additional 

stations where data were collected in southeastern 

Washington by Mohl, 1987. Bankey and others (1984) 

used the International Gravity Standardization Net 

1971 (IGSN 71) datum (Morelli, 1974), the 1967 gravity formula (International Association of Geodesy, 

1967), and a Bouguer reduction density of 2.67 g/cm3. 

Terrain corrections were applied radially from each 

station to a distance of 167 km (Bankey and others, 

1984).

When we began this study, gravity data in southeastern Washington were insufficient to define the suture zone. The thick basalt cover in southeastern 

Washington made it necessary to use a higher density 

of gravity stations there than in Idaho to determine 

the location of the suture zone. Approximately 900 

stations were added to the existing network in the 

Washington portion of the study area by Mohl (1987; 

fig. 13.2). Additional stations where Phole (1979) collected data were added to the network, as were stations where Gregory and Jackson (1976) took 

measurements. The data gathered by Mohl (1987) 

were collected with a Worden gravimeter at bench 

marks, triangulation stations, and other points of 

known elevation. All the measurements were tied to 

the IGSN 71 datum and corrected for drift, latitude 

(formula from International Association of Geodesy, 

1967), free air, and a Bouguer anomaly (density of 2.67 

g/cm3) (Mohl, 1987). The data are currently on file 

with the Defense Mapping Agency (U.S. Department 

of Defense), which has recalculated them using a 1984 

latitude-correction formula. Differences between the 

1967 and 1984 formulas are slight and result in changes that are well within the accuracy of the investigation. Terrain corrections were not applied to these data 

because the amount of correction that would result is 

not significant on the scale of the anomalies observed.

This report discusses the regional map patterns of 

gravity anomalies in the western Idaho suture zone 

(WISZ) in western Idaho and southeastern Washington as first defined in print by Strayer and others 

(1987, 1989, 1990), Strayer (1988), and Fleck and 

Criss (1988). Subsequent publications by Lund (1988 

and chap. 14, this volume), Lund and Snee (1988), 

Lund and others (1990), and Snee and others 

(chap. 10, this volume) called the western Idaho suture zone the "Salmon River suture zone" (see discussion and reply in Lund and others, 1990, and Strayer 

and others, 1990). In a companion study (Johnson and 

others, 1988; Mohl, 1989), we collected detailed data
500 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

117°45' 117° 116 C

118°30' 

46°45'

46°15'

TT

','/ ' .Yv' '

/ v-:-iA: MS

115°30' 

47 C

GRAVITY STATION 

DENSITY

EXPLANATION 

:. ;!: Gravity stations

I0 0 10 20 30 40 50 60 MILES

FIGURE 13.2. Gravity-station density of study 

area. Washington gravity data from Mohl 

(1987), Idaho gravity data from Bankey and 

others (1984). See figure 13.1 for labels of 

towns and 87Sr/86Sr contours.

< ' ,'

/ 

I

/ 

N ' ;. .

V ^ "" ' ' 

U

- 46<

- 45 C

44.
13. GRAVITY STUDIES OF AN ISLAND-ARC/CONTINENT SUTURE ZONE IN IDAHO AND WASHINGTON 501

along a series of four gravity traverses across the western Idaho suture zone straddling the major bend in the 

suture zone near Orofino, Idaho. Detailed gravity models were prepared for these traverses on the basis of 

local geologic and density observations.

GEOLOGIC AND TECTONIC ELEMENTS

The geologic and tectonic units associated with the 

western margin of the North American craton and the 

accreted terranes are well studied. Significant works 

include, but are not limited to, Armstrong and others 

(1977), Vallier (1977), Vallier and others (1977), Bond 

and Wood (1978), Brooks and Vallier (1978), Brooks 

(1979), Hyndman (1979, 1983), Myers (1982), Fleck 

and Criss (1985), Criss and Fleck (1987), Onasch 

(1987), Wiswal and Hyndman (1987), Aliberti and 

Manduca (1988a, b), Davidson (1988), Lund (1988), 

Strayer (1988), Strayer and others (1989, 1990), Lund 

and others (1990), Blake (1991), Selverstone and others (1992), Manduca and others (1992, 1993), and 

Blake and Thiessen (in press). The generalized assemblages presented herein are based on the work of these 

and other investigators.

NORTH AMERICAN ASSEMBLAGE

The North American assemblage is defined as the 

suite of rocks exposed north and east of the 87Sr/ 

86Sr=0.708 and 87Sr/86Sr=0.706 lines (fig. 13.1; Armstrong and others, 1977; Fleck and Criss, 1985). 

Termed the "Belt-Yellowjacket terrane" by Fleck and 

Criss (1985) and Criss and Fleck (1987), this assemblage consists of granodiorite and monzogranite of the 

Idaho batholith, Proterozoic metasedimentary rocks of 

the Belt Supergroup (Hyndman, 1979), and graniticgneissic pre-Belt basement rocks (Armstrong, 1975; 

Fleck and Criss, 1985). Seismic-refraction studies 

(Smith, 1978; Allenby and Schnetzler, 1983; Sheriff 

and Stickney, 1984) indicate that the crust in central 

Idaho is between 33 and 45 km thick.

The dominant quartzofeldspathic composition and 

relative thickness of the crust formed by the North 

American assemblage result in smooth and consistent long-wavelength gravity anomalies. Density 

measurements (Criss and Champion, 1984; Mohl, 

1989) indicate that the densities of the batholith and 

later epithermal plutons range from 2.43 to 2.81 g/ 

cm3 and average 2.56 g/cm3 and 2.58 g/cm3, respectively. The Precambrian units have average densities 

of 2.67 g/cm3 for gneissic units, 2.71 g/cm3 for schists 

and marbles, and 2.56 g/cm3 for quartzites (Mohl,

1989). These densities are all within the published 

ranges for metasedimentary rocks (Sharma, 1976; 

Telford and others, 1976). The density of possible 

pre-Belt basement rocks (Armstrong, 1975) averages 

2.66 g/cm3 , which is approximately the often quoted 

value of 2.67 g/cm3 for continental crust. Magnetic 

anomalies observed over the North American assemblage are also relatively smooth and consistent (Mabey and others, 1978).

ACCRETED-TERRANE ASSEMBLAGE

The accreted-terrane assemblage consists of the material south and west of the 87Sr/86Sr=0.704 line (fig. 

13.1). More diverse and structurally complex than the 

North American assemblage, this suite of rocks includes thick clastic and carbonate sequences, islandarc volcanic rocks, dismembered oceanic crust, and 

granitic plutons. Basement rocks include quartz diorite, plagiogranite, gabbro, and pyroxenite (Brooks, 

1979; Hyndman, 1979). Seismic-refraction and other 

studies indicate that the crust is 20 to 25 km thick in 

central Washington and 25 to 30 km thick in the area 

of this assemblage (Hill, 1972, 1978; Smith, 1978; Allenby and Schnetzler, 1983; Zervas and Crosson, 1986; 

Catchings and Mooney, 1988; Sobczyk, 1994).

Geophysical characteristics of the accreted-terrane 

assemblage are not as well defined as their counterparts in the North American assemblage. The magnetic signature is chaotic and typified by relatively shortwavelength anomalies (Mabey and others, 1978). 

Gravity anomalies are similar owing to the relatively 

thin crust and disjointed nature of the assemblage. 

Densities of the island-arc volcanic rocks (Mohl, 1989) 

average 2.72 g/cm3. Basement rocks, as well as overlying clastic and carbonate rocks, were poorly exposed in 

Mohl's (1989) study area and so were not sampled for 

density measurements. Characterization based on 

lithologic association indicates that the densities of the 

basement rocks in the accreted-terrane assemblage 

are generally higher than those observed in the North 

American assemblage, owing to the higher proportion 

of denser mafic components.

BOUNDARY ZONE

The width and limits of the boundary zone between 

the accreted-terrane and North American assemblages are different depending the data used to define 

them. Locations of suture-zone plutons, deformation, 

metamorphism, lithologic changes, isotopic changes, 

and geophysical characteristics all define somewhat
502 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

different boundary zones. However, the limits of the 

boundary zone are best defined by the 87Sr/ 

86Sr=0.708 and 87Sr/86Sr=0.706 lines on the north 

and east and by the 87Sr/86Sr=0.704 line on the south 

and west. The boundary zone consists of a series of 

geochemically and mineralogically distinct lensshaped plutons (Kuntz, M.A., Alien, C.C., and LaFortune, J.R., unpub. data, 1984; Fleck and Criss, 1985; 

Criss and Fleck, 1987; Aliberti and Manduca, 1988a, 

b; Lund, 1988), ultramafic pods (Hietanen, 1962; 

Hamilton, 1963; Bonnichsen, 1987), and highly deformed schistose metasedimentary rocks (Hamilton, 

1963, 1969; Myers, 1982; Onasch, 1979, 1987; Aliberti and Manduca, 1988a, b; Davidson, 1988; Lund, 

1988; Lund and Snee, 1988; Strayer, 1988; Strayer 

and others, 1989, 1990; Lund and others, 1990; 

Blake, 1991; Selverstone and others, 1992; Manduca 

and others, 1992, 1993; Blake and Thiessen, in 

press). This zone ranges from 5 to 25 km wide (Criss 

and Fleck, 1987) in west-central Idaho and is of undefined dimensions in southeastern Washington. Plutons of the boundary zone contain material derived 

from both the accreted-terrane and the North American assemblages (Criss and Fleck, 1987), and they 

tend to be compositionally and texturally similar to 

the plutons of the accreted-terrane assemblage (Fleck 

and Criss, 1985).

Geophysical characteristics of this assemblage vary 

with the width and lithology of the boundary zone 

itself. Ultramafic pods or mafic plutons composed of 

large percentages of ferromagnesian minerals have 

relatively high densities that result in gravity highs. 

Conversely, absence of these ferromagnesian minerals in the metasedimentary rocks results in gravity 

lows. Anomalies associated with this assemblage are 

likely to be isolated due to the discontinuous nature 

of the bodies observed. Detailed analysis of the gravity anomalies within the boundary zone must be done 

with much higher resolution gravity data sets than 

those currently available (Mohl, 1989), particularly 

because over part of its length, the boundary zone is 

covered with basalt.

COLUMBIA RIVER BASALT GROUP

The Columbia River Basalt Group covers much of 

the surface expression of the suture zone. This sequence of tholeiitic basalts originated in a series of 

dike swarms near the western Idaho border (Hooper, 

1982). In the Idaho part of the study area, the Clearwater and Weiser embayments represent the eastern 

limit of the basalt (fig. 13.1; Swanson and others, 

1980). In these areas, deep drainages were invaded

and filled with as much as 100 m of basalt (P.R. Hooper, oral commun., 1988). Near the margins, nonmarine 

clastic sediments shed from the surrounding highlands intertongue with the basalt flows to create a 

complex stratigraphic sequence (Cavin, 1964; Lin, 

1967; Myers and Price, 1979). In south-central Washington, the basalt thickens to the west to more than 

3,000 m in the central part of the Pasco basin (PB on 

fig. 13.1; Hooper, 1982), and the relative proportion of 

interbedded sedimentary rocks diminishes (Cavin, 

1964).

The problem of defining the gravity signature of 

the Columbia River Basalt Group is in determining 

the density and thickness of the basalt section. Near 

the margins, interbedded sedimentary rocks of varying proportions could reduce the density of the basalt stratigraphic section to below that of the 

surrounding country rocks. This would result in a 

gravity low of varying proportion. Toward the center 

of the basin the interbedded sedimentary rocks thin, 

but the relative proportion of highly fractured and 

vesicular basalt, which also would reduce the density of the section, is unknown. Furthermore, the basalt is known to have inundated a deeply dissected 

preexisting topographic surface that is poorly defined 

(P.R. Hooper, oral commun., 1988), resulting in highly variable thickness. Where measured from basalt 

samples taken across the plateau (Yost and Steele, 

1978; Robbins and others, 1979; Smith, 1979; 

Schmidt and others, 1980), densities range from 1.8 

g/cm3 to 2.8 g/cm3 as a function of vesicularity and 

degree of fracturing in a given sample. Samples 

from west-central Idaho (Mohl, 1989) range from 

2.49 g/cm3 to 2.91 g/cm3 with an average of 2.74 g/ 

cm3 . These measured densities are significantly 

lower than those commonly reported for basalts 

(Sharma, 1976; Telford and others, 1976), probably 

because textbook examples do not take into account 

the vesicularity and fracturing observed. Near the 

margins of the basins, the overall density of the basalt section is further reduced by the interbedded 

sedimentary rocks, whose densities range from 0.95 

to 2.4 g/cm3 (Yost and Steele, 1978; Mohl, 1989).

GRAVITY SIGNATURE

Gravity studies over several suture zones exposed 

elsewhere in the world imply that suture zones are 

characterized by a paired high-low gravity anomaly 

(Gibb and Thomas, 1976; Fountain and Salisbury, 

1981; Hutchinson and others, 1983; Earner and 

Watts, 1983; Price and Hatcher, 1983; Thomas, 1983; 

Cook, 1984; Dainty and Frazier, 1984; Fisher and von
13. GRAVITY STUDIES OF AN ISLAND-ARC/CONTINENT SUTURE ZONE IN IDAHO AND WASHINGTON 503

Huene, 1984; Johnson and others, 1984; Keller and 

others, 1985; Thomas and others, 1988). In most of 

these studies, a regional high is observed over the 

accreted terranes and a regional low over the continental material. The suture zone itself is often characterized by isolated highs either in the suture zone 

or just outboard, and a steep gradient that marks the 

change from the regional high of the accreted terranes to the regional low of the continent. As shown 

in figure 13.3, the total difference in gravity values 

and the distance between the regional high and low

anomalies vary from suture zone to suture zone. Relatively few models have been developed to explain 

the attributes of the paired anomalies. In these investigations most of the regional high and low gravity anomalies have been conceptually modeled as 

products of either obduction of oceanic crust (fig. 

13.3, models A and B) or crustal thickening (model 

C). Local anomalies, such as the highs directly associated with the suture zone and those produced by 

relatively light upper-crustal bodies, are shown in 

models D and E.

Report

Gibb and Thomas (1976)

Fountain and Salisbury (1981)

Hutchinson and others (1983)

Thomas (1983)

Karner and Watts (1983)

Johnson and others (1984)

Cook (1984)

Dainty and Frasier (1984)

Fisher and von Huene (1984)

Keller and others (1985)

Region

Canadian Shield

5 around world

Southern Appl.

Southern Appl.

Alps, Appl.

Wyoming

Southern Appl.

Southern Appl.

Southern Alaska

Virginia

A Grav.

(milligals)

66

30-140

60-150

60-150

80-200

50-100

60-140

60

60-70

<15

Dist.

(kilometers)

80

30-60

60-120

80

70-200

90

50-140

60

50

40

Model

A

B

C, D

A

F

C, D, E

C

C, D

B

C, D, E

isis< Zi 

>=!

DISTANCE, IN KILOMETERS

B

D

EXPLANATION

Medium-density material

Oceanic sedimentary rocks

Low-density material

High-density material

FIGURE 13.3. Gravity anomalies and conceptual models (cross sections A-E) tor suture 

zones in various regions. Graph shows general shape of all anomalies; specific values 

given in table. Appl., Appalachian Mountains in North America; Dist., distance; A Grav., 

total change in gravity over specified distance. Arrows in models A and B show direction 

of relative movement.
504 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

WEST-CENTRAL IDAHO

West-central Idaho gravity data (Bankey and others, 1984) display a paired high-low anomaly that is 

similar to the signature observed in other suture 

zones (fig. 13.4). Gravity values in the accreted-terrane assemblage tend to be relatively higher than 

their continental counterparts. With respect to the 

anomalies that characterize the suture zone itself, a 

series of isolated highs are observed immediately 

south and west of the 87Sr/86Sr=0.704 line. A gradient 

with gravity values decreasing to the north and east 

is also seen. The relative magnitude of the gradient is 

25 to 80 mGal over a distance of 15 to 70 km. The 

highs are smaller (5-15 mGal) and of much more limited extent.

In the northern part of the study area, lat 46°-470 

N., the suture zone appears to have two distinct 

anomaly signatures. The east-west segment from 

Orofino, Idaho, to the Washington border is typified by 

a small but uniform gradient with gravity values decreasing to the north and by well-defined highs outboard of the 87Sr/86Sr=.704 line. In this area, highs 

appear to coincide with local mafic Mesozoic intrusions emplaced near the suture zone. The north-south 

segment displays a steep gradient with gravity values 

decreasing to the east. The gradient and paired anomaly have higher amplitudes than in the east-west segment, and highs are poorly defined. On the whole the 

gradient defines the northern and eastern extent of 

the Columbia River Basalt Group in the Clearwater 

embayment (fig. 13.1). It also approximately coincides 

with the western edge of the Idaho batholith.

In the central part of the study area, lat 45°-460 N., 

the gradient swings to the west and crosses the suture 

zone defined by 87Sr/86Sr ratios. The gradient also 

crosses the southern limit of the Clearwater embayment (fig. 13.1) and places metasedimentary rocks of 

the Riggins Group (Hamilton, 1963; Myers, 1982) as 

well as at least one major boundary pluton on the east 

side of the gradient (see area A, fig. 13.4). In the area 

of the (informal) Little Goose Creek and Hazard 

Creek complexes of Aliberti and Manduca (1988a, b) 

and Manduca and others (1993), indicated with the 

letter B on fig. 13.4, the gradient appears to diverge 

and surround these plutonic bodies. The local highs 

are absent in this part of the study area.

In the southern part of the study area, lat 44°-45° 

N., the gravity contours once again form a single, 

steep gradient. The gradient is steep in the vicinity 

of the 87Sr/86Sr=0.706 line and roughly coincides 

with the western border of the Idaho batholith, or 

the eastern limit of the Columbia River Basalt Group 

in the Weiser embayment (fig. 13.1). Several poorly

defined highs are present outboard of the suture 

zone, but these, unlike those found in the northern 

part of the study area, do not appear to be related to 

exposures of Mesozoic intrusive rocks.

Comparison of the gravity signatures with aeromagnetic data from the study area (Mabey and others, 1978; Zietz and others, 1978) shows that several 

of the gravity features appear to be associated with 

magnetic anomalies (fig. 13.5). The magnetic anomalies exhibit a major change in character from relatively smooth, long-wavelength anomalies over the 

North American assemblage to more chaotic, shortwavelength anomalies over the accreted-terrane assemblage (Mabey and others, 1978). In the northern 

part of the study area, this change in character corresponds with the line of gravity highs observed at 

the suture zone. Near Grangeville, Idaho, the gravity 

gradient that crosses from east to west of the suture 

zone is reflected by a similar magnetic gradient. In 

the southern part of the study area, the change in 

character of the magnetic anomalies coincides with 

the steep slope of the gravity gradient as well as the 

suture zone defined by 87Sr/86Sr ratios.

SOUTHEASTERN WASHINGTON

Increased detail afforded by the higher density of 

gravity stations in southeastern Washington (Mohl, 

1987) permits more accurate definition of gravity 

anomalies there than in the Idaho portion of the 

study area. Contoured on an interval of 2 mGal rather than the interval of 5 mGal used in Idaho, the 

Washington data accurately define several major 

gravity features that may be associated with the suture zone. Correlation of these detailed gravity anomalies to geologic and tectonic features is difficult, 

because outcrops of pre-basalt rocks are limited to a 

few isolated exposures near the Idaho-Washington 

border and in the northern Blue Mountains south of 

Pomeroy, Wash. (fig. 13.1). Given this lack of exposure, extrapolation of the geologic and tectonic associations expressed in Idaho and correlation with 

structures within the basalt and with physiographic 

features provide the primary means of interpreting 

the data.

Gravity anomalies observed in southeastern Washington are labeled A through F on figure 13.6. The 

regional paired anomaly signature observed in Idaho 

is not as strongly defined in these data. However, the 

signature observed from Orofino, Idaho, to the Washington border (fig. 13.4) does extend into southeastern Washington, where it is expressed as an eastwest-trending series of highs (Bl, B2, B3 on fig. 13.6)
13. GRAVITY STUDIES OF AN ISLAND-ARC/CONTINENT SUTURE ZONE IN IDAHO AND WASHINGTON 505

117°45' 117

118°30' 

46°45'

46°15'

BOUGUER GRAVITY

EXPLANATION

Bouguer gravity contours 

Hachured in direction of 

gravity low. Contour 

interval, 5 milligals. 

Reduction density, 2.67 gm/cm3

0 0 10 20 SO 40 50 60 70 KILOMETERS ,

FIGURE 13.4. Bouguer gravity map of study 

area. Idaho portion from Bankey and others 

(1984). Washington portion modified from 

Mohl (1985). 87Sr/86Sr contours modified 

from Armstrong and others (1977) and Fleck 

and Criss (1985). See figure 13.1 for labels of 

towns and 87Sr/86Sr contours.
506 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

117°45' 117° 116°

46°15

AREOMAGNETIC

EXPLANATION

Magnetic contours Hachured in 

direction of magnetic low. 

Contour intervals, 20 and 100 

gammas

0 20 30 40 50 60 70 KILOMETERS

^~^^f*=-*'/rS^'6~\ V ))) J

FIGURE 13.5. Aeromagnetic map of study area 

showing change in character of anomalies 

across suture zone defined by initial 87Sr/ 

86Sr ratios. From Zietz and others (1971, 

1978). 87Sr/86Sr contours modified from Armstrong and others (1977) and Fleck and Criss 

(1985). See figure 13.1 for labels of towns 

and 87Sr/86Sr contours.
117°45' 

47

BOUGUER GRAVITY ANOMALY MAP

SOUTHEASTERN WASHINGTON

EXPLANATION

Bouguer gravity contours Observation station 

Hachured in direction of Base station 

gravity low. Contour + Gravity high 

interval, 2 milligals Gravity low

46°15'

FIGURE 13.6. Detailed Bouguer gravity map of Washington portion of study area showing relative location and magnitude of anoma

description of origin of anomalies A through F. Modified from Mohl (1985).
508 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

and a parallel gradient sloping down to the north 

and northeast (Dl). This signature is complicated by 

a crosscutting linear feature (F) and two subcircular 

features (A, E). An isolated subcircular feature (C) 

corresponds to the westernmost exposure of the 

North American assemblage in southeastern Washington (Armstrong and others, 1977).

The east-west-trending series of highs (Bl, B2, B3) 

correlates to several local geologic and physiographic 

features shown on regional maps by Myers and Price 

(1979) and Swanson and others (1980). Anomaly Bl 

coincides with a complex set of folds, faults, and 

faulted folds called the Lewiston structure (Camp, 

1976). This structure is characterized by anomalous 

paleomagnetic rotations of basalts (Lim, 1986; Thiessen and others, 1986). The gravity anomaly corresponds to the structure for several kilometers, then 

continues to the west (B2), following a series of small 

monoclines along the northern edge of the Blue 

Mountains near Pomeroy (fig. 13.1). Near the northwest corner of the Blue Mountains, the folds die out, 

and the anomaly abruptly changes character from a 

high-amplitude feature (B2) to a more subdued, lowamplitude feature (B3). This series of highs is approximately 5 to 10 km south of the extrapolated location of the 87Sr/86Sr=0.704 line and is on the 

accreted-terrane assemblage side of the suture zone.

The gradient (Dl) and low (D2) also correlate to 

local geologic and physiographic features. The gradient roughly defines the southern extent of the Palouse ridge, which is a quartzitic and granitic 

topographic high buried beneath the Columbia River 

Basalt Group (Strand, 1949; Cavin, 1964; Lin, 1967). 

North and west of Pullman the gradient weakens, 

possibly divides, then joins a major regional northeast-trending gradient discussed by Cady and Fox 

(1984), Johnson and others (1989), and Thiessen and 

others (1992). The subcircular low (C) located between the gradient and the high trend appears to be 

related to a subbasalt ridge similar to that correlated 

with the Dl and D2 set of anomalies. On the basis of 

87gr/86gr ratiOSj this material is unquestionably of 

North American assemblage affinity (Armstrong and 

others, 1977).

Other significant gravity features include subcircular lows (A and E) and a linear anomaly (F). Anomaly A corresponds to the Blue Mountains highland, 

which includes several isolated exposures that resemble the rocks of the accreted-terrane assemblage 

(Armstrong and others, 1977; Swanson and others, 

1980). The other subcircular anomaly, E, does not 

correlate to any surface geologic or physiographic element. However, it does correlate to a zone of anomalous paleomagnetic rotations as observed in a latestage basalt flow (Lim, 1986; Thiessen and others, 

1986). The linear anomaly, F, is tentatively correlated 

to an extension of Riddihough and others' (1986) 

Klamath-Blue Mountain lineament (Mohl and Thiessen, 1985; Mohl, 1985). This feature truncates the 

high-amplitude high (B2), changing the character of 

the anomaly to a much more subdued feature. It 

aligns with the Kite Fault, a normal fault that drops 

the basalt down to the west. It also correlates with a 

major break in the gradient (Dl) and forms the western edge of anomaly A.

When compared with aeromagnetic anomalies observed in southeastern Washington (Zietz and others, 

1978), three of the gravity features are seen to have 

corresponding magnetic expressions. These features 

include the east-west-trending series of gravity highs 

(Bl, B2, B3) that corresponds to a similar series of 

magnetic highs, and gravity lows (A and C) that coincide with magnetic lows. The linear gravity anomaly 

(F) coincides with a reduction of the magnetic intensities from relatively high-magnitude features on the 

east side to more subdued features on the west side.

GRAVITY MODELING

The origin and significance of gravity anomalies observed in this investigation may be addressed with 

simple two-dimensional computer models. The program used in this analysis (Thiessen, 1986) is based 

on Talwani and others' (1959) algorithm, which determines the expected gravity signature of isolated 

geologic elements of known dimension and density.

The effects of crustal differences on gravity anomalies were modeled on the basis of three variables: 

crustal thickness, crustal composition, and suturezone geometry (fig. 13.7A). In models A through G on 

figure 13.7, the density of cratonic crust was assumed 

to be 2.67 g/cm3, on the basis of the limited samples of 

Criss and Champion (1984) and Mohl (1989). This is 

the density commonly assumed for continental crust. 

The density of the accreted block was assumed in 

models B, D, and F to be equal to that of the continental block, reflecting an assumption that the two blocks 

are of similar composition. In models C, E, and G, the 

density of the accreted block was assumed to be 2.70 

g/cm3, 0.03 g/cm3 denser than the continental block, 

reflecting an assumption that the accreted block consists of slightly more mafic rocks. The effect of varying 

this density contrast by various intervals, from 0.02 to 

0.1 g/cm3 ' is indicated by dashed lines in the graph at 

the upper left of figure 13.7A. The density of the upper 

mantle is assumed to be 3.0 g/cm3 . The geometry of 

the suture zone was varied to simulate a vertical con-
13. GRAVITY STUDIES OF AN ISLAND-ARC/CONTINENT SUTURE ZONE IN IDAHO AND WASHINGTON 509

>

< > tr li CD <

= ^ ui O

CD **

Z> 

O

CQ

CO

< 

CD

A

200

100

2.71-- -__

-2.70 A~ 

2.69 --

200

100

0

100 0 

DISTANCE, IN KILOMETERS

100

DEPTH, KILOMETERS IN 

jO N)COfOCO

oo o c

^ZXfr&Tt* 

I (2.70) : I

2.67

B

2.67

2.67

C

ZffQ , ;

2.67

0 100 

DISTANCE, IN KILOMETERS

B

50

-50

50

0 100 

DISTANCE, IN KILOMETERS

H-J

100

H 2.60 

I 2.70 

J 2.80

2.67

tr CD 

tr

LU

=>

CD

Z> 

O

CQ

Z -50 

- 0

K, L

K 2.72 

L 2.62

2.67

100

50

0

M

-50

0 100 

DISTANCE, IN KILOMETERS

2.90

N

2.67

FIGURE 13.7. Generalized models of 

Bouguer gravity anomalies produced by various geologic elements 

observed or inferred in study area. 

Numbers to two decimal places are 

densities in g/cm3. A, Models showing effects of variations in crustal 

thickness, crustal density, and suture-zone orientation. Accreted 

crust shown on left, continental 

crust on right. In two graphs at top 

of figure 13.7A, solid lines labeled 

with letters show anomalies that 

correspond to lettered cross sections. Dashed lines are anomalies 

produced by varying differences in 

density between accreted crust and 

continental crust in model A; 

dashed lines are labeled with densities postulated for accreted crust. 

Heavy solid line in each graph 

shows representative anomaly 

across suture zone in west-central 

Idaho. B, Models showing effects of 

near-surface geologic bodies. Models H through J show effects of 

variations in density of slab of Columbia River Basalt Group 30 km 

wide by 1 km thick. Models K and 

L represent suture-zone plutons of 

same size (20 km wide at top, 10 

km wide at bottom, and 8 km 

thick) but different densities. 

Model M represents vertical highdensity body such as sliver of mafic 

rocks 1 km wide by 5 km deep; 

model N represents low-density 

body 50 km wide by 10 km thick 

emplaced in upper crust, such as 

Idaho batholith. In cross sections, 

density values on left are for nearsurface geologic bodies and those 

on right for enclosing continental 

crust.

o 2.67
510 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

tact (models A C), inclined subduction (models D and 

E), and inclined obduction (models F and G). The 

heavy line on each of the two graphs at the top of 

figure 13.7A shows a representative section across the 

suture zone in west-central Idaho. The effects of an 

abrupt contrast in crustal thickness across the suture 

zone were studied on all three geometries using crustal thicknesses as indicated in seismic-refraction studies (Smith, 1978; Allenby and Schnetzler, 1983; 

Sheriff and Stickney, 1984): 20 km for the accretedterrane assemblage and 30 km for the continent. As 

summarized in figure 13.7A, the overall shape of the 

gravity anomaly varies with each change in variables, 

but the long wavelength and the total change in gravity remain similar. If a simple density contrast of 0.03 

g/cm3 is applied (model A), the resulting anomaly is 

about 30 mGal with a wavelength of more than 50 

km. If a corresponding 10-km change in crustal thickness is applied (model C), the magnitude of the gradient is increased to approximately 100 mGal and the 

wavelength to 80 km. Increasing either the density 

contrast between the two blocks to 0.1 g/cm3 or the 

density of the upper mantle component to 3.4 g/cm3 

(simulating more mafic mantle) causes the anomaly to 

increase by another 100 mGal. Reducing the thickness 

contrast across the suture would reduce the gravity 

anomaly accordingly.

This series of models demonstrates that several 

combinations of differences in crustal composition and 

thickness and in underlying components associated 

with the suture zone will result in gravity anomalies 

that are about 100 mGal or greater. Such configurations are characterized by the long wavelength of the 

anomalies observed in the study area.

Near-surface geologic bodies are characterized by 

sharp, short-wavelength anomalies whose magnitude 

is directly proportional to the density contrast between the body and the surrounding background material. Near-surface geologic bodies in west-central 

Idaho include the Columbia River Basalt Group, suture-zone plutons, and the Idaho batholith; models of 

these bodies are shown in figure 13.IB.

Several models were developed to estimate the influence of the rocks of the Columbia River Basalt 

Group in the study area. For these models, H through 

J on figure 7.13.B, we assumed the presence of a tabular basalt body 1,000 m thick and varied the density 

of the body to examine the influence of changes in 

density in the basalt section. Three densities were 

used for the basalt body: 2.60 g/cm3 (model H) to simulate basalt with numerous, less dense sedimentary 

interbeds; 2.70 g/cm3 (model I) to simulate basalt with 

a few sedimentary interbeds; and 2.80 g/cm3 (model J) 

to simulate basalt with no sedimentary interbeds. A

density of 2.67 g/cm3 was assumed for the surrounding rock. As shown in figure 13.7B, gravity anomalies 

generated by these parameters are of negligible significance on the scale of the anomalies observed in this 

study. If the basalt body is assumed to be wedgeshaped the generated anomaly is even less significant.

Suture-zone plutons, such as the Kamiah pluton 

or Blacktail stock, are represented by models K and 

L on figure 13.7B. The sign of the anomaly is positive if the body is more dense and negative if the 

body is less dense than the background material. 

These models assume that the suture-zone pluton is 

limited in lateral and vertical extent and that the 

background material has a density of 2.67 g/cm3. 

The gravity anomaly produced by a body that is 0.05 

g/cm3 more dense (2.72 g/cm3) or 0.05 g/cm3 less 

dense (2.62 g/cm3) than the background material has 

a magnitude of about 10 mGal..

Pods of mafic or ultramafic rocks caught in the accreted-terrane assemblage, or in the suture zone itself, may be responsible for some of the gravity highs 

outboard of the suture zone. Such a feature is represented by model M on figure 13.IB, where the pod is 

assumed to be a body 1 km wide by 5 km deep with a 

density of 2.90 g/cm3 set in a matrix of country rock 

with a density of 2.67 g/cm3 . This configuration produces a very sharp gravity anomaly of 8 mGal, which 

is very similar to the anomalies observed in the 

northern portion of the study area. This anomaly is 

much sharper than those produced by the deeper 

crustal differences across the suture zone (fig. 13.7A).

Anomalies generated by relatively thick uppercrustal bodies of low-density material, such as the 

Idaho batholith, are more like the anomalies observed in west-central Idaho. A model of such an 

anomaly, N on figure 13.IB, assumes the same background density as continental crust (2.67 g/cm3), and 

places a body 10 km thick and 0.11 g/cm3 less dense 

(2.56 g/cm3) into the upper crust. The geometry of 

the lighter body was not varied because such variation would only affect the shape of the anomaly, and 

not its magnitude. The importance of this model is 

that it might explain why the gradient observed in 

west-central Idaho follows the western margin of the 

Idaho batholith. It does not, however, explain why 

the gradient crosses to the west of the suture zone as 

defined by strontium ratios in the central part of the 

study area. This model also does not explain the 

magnitude of the anomaly in the southern part of the 

study area unless one assumes an unrealistically 

thick batholith.

The approach of identifying possible anomaly 

sources is appropriate for areas where the gravity 

data are relatively coarse and there is geologic con-
13. GRAVITY STUDIES OF AN ISLAND-ARC/CONTINENT SUTURE ZONE IN IDAHO AND WASHINGTON 511

trol to guide the interpretation. As applied to the 

anomalies observed in the study area, these models 

represent a qualitative analysis of the relation between the anomalies and the geologic elements to 

which they are tentatively correlated. Where gravity 

anomalies are more tightly constrained, the general 

models may be supplemented with a specific model 

generated to match the anomalies in the field data. 

Johnson and others (1988) and Mohl (1989) generated such models for four detailed traverses in the 

north-central portion of the study area.

A detailed model (fig. 13.8) developed on a northsouth profile at a longitude of 117° W. near the 

Idaho-Washington border utilizes the high-resolution 

Washington gravity data. The geology of the model is 

constrained by exposures of the Idaho batholith at 

the north end of the profile and the Columbia River 

Basalt Group in the central and southern parts. It is 

also weakly constrained by the projection, near the 

center of the profile, of the suture zone delineated by 

initial 87Sr/86Sr ratios. The anomalies observed along 

this profile may be generated with the combination of 

a relatively light body (model N, fig. 13.7) and two 

north-dipping dense bodies (model M, fig. 13.7). Variations in crustal thickness or composition are poorly 

constrained and not consistent with the gravity 

anomalies observed. Thus, they are not included in 

this model. The basalt section was not included due 

to the relative insignificance of the anomaly it generates. Note that any gravity model matched to field 

data, such as that shown in figure 13.8, provides only 

a first approximation of the anomaly-producing bodies and is not a unique solution.

INTERPRETATION

The models in figure 13.7 indicate that there are 

two general classes of gravity anomalies that may be 

associated with the geologic features of the suture 

zone. The first group consists of relatively narrow and 

sharp anomalies that are limited to less than 30 mGal 

(fig. 13.7JB). These types of anomalies can only be created by near-surface bodies that are isolated and have 

distinctly different densities than the surrounding 

country rock. The second group consists of broad, 

long-wavelength anomalies that may display a total 

change in gravity of greater than 30 mGal (fig. 13.7A). 

The magnitude of the anomaly produced by this type 

of feature depends on the size of the density contrast, 

the difference in crustal thickness, and the density of 

the underlying crustal material.

The suture zone in the northern part of the study 

area is characterized by two distinct regional gravity

> 

< :E 

o

i«>

-70

-80

-90

-100

-110

-120

Observed

20 40 60 

DISTANCE, IN KILOMETERS

80

87Sr/ 86Sr ratios:

5 -,

20 40 60 

DISTANCE, IN KILOMETERS

EXPLANATION

Medium-density material 

Low-density material 

High-density material

FIGURE 13.8. Model of Bouguer gravity along Idaho-Washington 

border. Field observations shown with dashed line; calculated 

gravity anomaly with solid line. The model includes high- and 

medium-density slivers along the boundary zone, which lies between the initial 87Sr/86Sr ratio values of 0.704 and 0.708 near 

the center of the profile. It also includes a less dense body in the 

upper crust associated with outlier plutons of the Idaho batholith.

signatures: (1) a relatively steep gradient of 20 mGal 

in the east-west part of the suture zone west of Orofino, Idaho, and (2) a true paired-anomaly gravity signature with a regional high and a regional low 

separated by a gradient in the north-south part of 

the suture zone immediately south of Orofino. The location, wavelength, and magnitude of the gradient 

between the regional high and low suggest that this 

gradient is due in part to differences in composition 

and (or) thickness of the crust on either side of the 

suture zone. The gradient is also due in part to largescale, near-surface features such as the Idaho batholith. Several anomalies superimposed on the gradient
512 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

appear to originate in the near surface on the basis 

of their relatively small size and sharp features. 

They are explained as representing isolated bodies of 

rock emplaced in the upper part of the crust within 

or near the suture zone, such as the Kamiah pluton 

and amphibole-rich plutons. The relatively steep gradient of 20 mGal in the east-west part of the suture 

zone west of Orofino appears to be due entirely to 

near-surface bodies. The magnitude and shape of this 

anomaly indicate that it does not represent differences on the crustal scale, unless those differences are 

offset by other geologic effects such as crustal thinning of continental North America, as discussed by 

Mohl (1989).

Interpretation of the gravity signature of the central part of the study area is difficult because the 

gradient crosses the suture zone defined by initial 

87Sr/86Sr ratios. This apparent lack of correlation between the gravity signature and the suture zone 

leads to several possible interpretations: (1) geologic 

elements overprint the signature of the suture zone,

(2) the geometry of the suture zone changes at depth,

(3) the geophysical characteristics of the material 

within the suture zone vary along the suture zone, or

(4) the apparent intersection of gradient and suture 

zone is simply a relict of the relatively coarse level of 

data available for this area. The magnitude and 

shape of the gradient suggest that the gravity signature is due at least in part to crustal differences on 

either side of the gradient. However, the complex geology that has been observed on the surface in this 

part of the study area indicates that contributions 

from near-surface elements are to be expected. Significantly more geophysical and density data are 

needed to evaluate interpretations.

The gravity signature in the southern part of the 

study area is a long-wavelength anomaly overprinted 

by a sharper gradient near the suture zone defined 

by initial 87Sr/86Sr ratios. The long wavelength and 

large magnitude of the gradient indicate a significant 

contribution from crustal-scale anomaly sources as 

demonstrated by the theoretical models in figure 

13.7A. Probable upper-crustal sources include contrasts in thickness and density that are larger than 

the gravity data indicate farther to the north. The 

superimposed gradient is most likely related to the 

western edge of the Idaho batholith. This sharper 

anomaly also coincides with the abrupt change in 

character of the magnetic anomalies. To the west, 

several isolated gravity highs that do not correlate 

with any surface features probably define other nearsurface elements, but more detailed gravity data 

are needed before these highs can be adequately 

analyzed.

CONCLUSIONS

Gravity studies in west-central Idaho and southeastern Washington delineate the suture zone between an assemblage of accreted terranes and 

continental North America as a marked change in regional gravity values and closely associated suturezone anomalies. The character of the observed gravity anomalies varies within the study area. In the 

northern part, the signature is a classic regional 

high-low paired anomaly with local highs superimposed near the suture zone. The gradient between 

the regional high and low does not always coincide 

with the suture zone defined by initial 87Sr/ Sr ratios, as is evident in the central part of the study 

area where the gradient crosses the suture zone. In 

the southern part of the study area, the broad gradient that separates the high and low in the regional 

paired anomaly does coincide with the suture zone 

and steepens there. Isolated highs outboard of the 

suture zone are also observed in this part of the 

study area.

Correlation of these anomalies to geologic features, 

magnetic anomalies, and theoretical models indicates 

that the primary source of the lows in the northern 

part of the study area is probably the Idaho batholith, whose effects are superimposed on contrasts in 

crustal thickness and (or) composition. The highs are 

probably related to near-surface mafic and ultramafic 

bodies caught in or near the suture zone. Lack of correlation in the central part of the study area may be 

due simply to the relatively coarse level of gravity 

data available, or it may be related to geologic elements that are not reflected in the Sr analyses. Gravity anomalies in the southern part of the study area 

appear to be related to differences in crustal composition or thickness that have been overprinted by the 

effects of the Idaho batholith.

In southeastern Washington, geologic control is 

limited to a few isolated exposures, but detailed gravity data closely define the suture zone by extending 

the anomaly signature observed in Idaho. This signature can be traced approximately 60 km west of the 

Idaho-Washington border, where it becomes more 

poorly defined due to thicker basalt cover. It is still 

identifiable to the north in the east-central Columbia 

plateau, approximately 25 to 35 km north of Washtucna (fig. 13.1; Johnson and others, 1989, Thiessen 

and others, 1992). However, rocks that appear to 

have a continental affinity have been sampled with a 

drill hole 23 km west of the southwest corner of figure 13.6 (Reidel and others, 1994; Sobczyk, 1994), indicating that the boundary zone may extend to the 

southwest from anomaly B2 through the southern
13. GRAVITY STUDIES OF AN ISLAND-ARC/CONTINENT SUTURE ZONE IN IDAHO AND WASHINGTON 513

anomaly F (fig. 13.6). In northeastern Washington 

and British Columbia, accreted terranes are thought 

to have been thrust over a wedge of continental material. This relationship may extend directly to the 

south under the Columbia River Basalt Group. The 

gravity anomaly discussed by Johnson and others 

(1989) and Thiessen and others (1992) aligns with 

the eastern limit of the accreted terranes in northeastern Washington and British Columbia. The eratonic margin defined by Reidel and others (1994) and 

Sobczyk (1994) corresponds to the western limit of 

the wedge of continental crust.

Further definition of the tectonic significance of 

these gravity anomalies in west-central Idaho and 

southeastern Washington requires that significantly 

higher resolution data be acquired. In west-central 

Idaho, detailed gravity traverses across the suture 

zone have been studied and modelled (Mohl, 1989) in 

order to more accurately define the signature where 

it is constrained by better exposure. In Washington, 

higher resolution gravity data are needed to further 

define the westward extension of the suture zone 

(Johnson and others, 1989; Thiessen and others, 

1992). This definition is particularly crucial for determining the point where the suture zone returns to a 

north-south orientation; such a determination is vital 

to tectonic models developed for the region.

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14. METAMORPHIC AND STRUCTURAL DEVELOPMENT OF 

ISLAND-ARC ROCKS IN THE SLATE CREEK-JOHN DAY CREEK AREA,

WEST-CENTRAL IDAHO

By KAREN LUND

CONTENTS

Abstract 517

IntrodUCtiOH. ._..__.____.__.___ _ __..._._...__.________ _..___--____________. ._ ^) 1 7

Acknowledgment 519

Rocks of the Blue Mountains island arc 519

Rocks of the Wallowa terrane 521

Seven Devils Group 521

Upper Triassic Martin Bridge Formation 521

Stratigraphy - 521

Correlation 523

Metamorphosed rocks of uncertain age 525

Rapid River plate 525

Description of rocks 525

Proposed correlation 526

North Fork block - - T 526

Description of rocks 526

Proposed correlation 527

Late Cretaceous crosscutting plutonic rocks 527

Rock units ____________________________ 528

Isotoi)ic a£Tes ___.-.-- - _»«-»««....__«««___...___« 528

Metamorphic history 528

Description -._.__.........._.____.._._....__ _ _____.._ ^52Q

Ages of metamorphism 530

Structure . .... ...... ....... 530

Fabrics and geometries 530

Ages 534

___ ___ _._ .__.._____._____ _ .....-._...__ ._____.. _____._.______- .__ ^5 ^^5

S _._.._..... ____.______ ._____...,._____..__»._..______....___. « ....__ _____. ^5^6

..............__..............__ ...._._____........ 538

ABSTRACT

Metamorphosed volcanic island-arc rocks of the Slate CreekJohn Day Creek area of western Idaho are bounded on the east by 

the Salmon River suture, the structure that separates these oceanic rocks from North American continental rocks to the east. The 

island-arc rocks are characterized by multiple overprints of medium- to high-grade metamorphism and multiple phases of deformation, all of which were caused by late Early to Late Cretaceous 

movement along the Salmon River suture as the oceanic rocks 

were being accreted to the North American continent. The islandarc rocks occur as three imbricated tectonostratigraphic units that 

probably originated as part of the Blue Mountains island arc but

are now part of a complex, inverted metamorphic sequence. At the 

bottom of the sequence are the lowest grade metamorphic rocks in 

the study area; these are part of the Wallowa terrane (the lower 

plate). This lower plate is composed of Lower Permian and Upper 

Triassic rocks of the Seven Devils Group as well as the Upper Triassic Martin Bridge Formation (here made up primarily of calcareous slate to schist containing marble lenses, minor metavolcanic 

rocks, and possible marble clast conglomerate). The middle tectonostratigraphic unit or plate, which is herein also called the Rapid 

River plate, generally is correlative with rocks previously called 

the Riggins Group; in this paper, rocks of the Rapid River plate 

are tentatively considered to be a tectonic slice of the Wallowa terrane. The uppermost tectonostratigraphic unit or plate, herein 

called the North Fork block, is multiply metamorphosed, deformed, and intruded and cannot definitely be correlated with 

known units.

Metamorphism of rocks in the study area began about 120 Ma 

and continued during deformation until about 90 Ma. Rocks in the 

North Fork block show conclusive evidence of overprinting amphibolite-facies events. The earliest observed deformation, which produced foliation, occurred with metamorphism about 120 Ma; 

multiple phases of megascopic deformation, which imbricated and 

folded the rocks, occurred from 109 to 90 Ma. A 90-Ma dynamothermal event caused the last observed metamorphic/deformational fabric in rocks in the North Fork block and juxtaposed the deep-seated 

rocks of this upper plate against shallower, earlier metamorphosed 

and deformed rocks of other plates. The earliest nonallochthonous 

plutons intruded across the Salmon River suture starting 93 m.y. 

ago; these are the first plutons of the Idaho batholith. Plutonic 

rocks in the North Fork block were strongly deformed together with 

the older wallrocks during the 90 Ma event; away from the Salmon 

River suture, these plutons (as well as younger plutons) are not 

strongly deformed.

INTRODUCTION

As recognized early (Lindgren, 1904), west-central 

Idaho is underlain by metamorphosed oceanic volcanic 

and volcaniclastic rocks that are directly juxtaposed 

against metamorphosed older terrigenous metasedimentary rocks. The Salmon River suture between these 

crustal blocks is everywhere intruded by Late Cretaceous plutons of the Idaho batholith and is covered by 

Miocene basalt along much of its length (fig. 14.1).

517
518 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

118°00' 117°00'

0 50 MILES 

J

I ' ' ' ' I

0 50 KILOMETERS

EXPLANATION

Quaternary to Tertiary volcanic 

rocks and alluvium

Eocene to Cretaceous igneous rocks 

Jurassic to Permian island-arc rocks

Paleozoic to Middle Proterozoic 

North American continental rocks

Initial strontium isotope ratio 

isopleth
14. METAMORPHIC AND STRUCTURAL DEVELOPMENT OF ISLAND-ARC ROCKS, WEST-CENTRAL IDAHO 519

The Slate Creek-John Day Creek area of western 

Idaho (fig. 14.2) is underlain by oceanic rocks west of 

an exposed segment of the Salmon River suture zone. 

This area contains the major rock types and deformational features commonly found along the suture. It 

has been mapped in detail (Lund, 1984; McCollough, 

1984; Lund and others, 1993) and has been the subject of detailed dating studies (Snee and others, 1985; 

Snee and others, 1987; Snee and others, chap. 10, this 

volume). Data from these studies are the basis for this 

paper and the reader is referred to them for further 

details. Tectonic models to explain the metamorphic 

and structural geometry on both sides of the Salmon 

River suture (Lund, 1984) and the timing of these 

events were published by Lund and Snee (1988).

Several mechanisms for accretion of the oceanic 

rocks to North American continental rocks have been 

proposed: (1) subduction (Hamilton, 1976; Hyndman 

and Talbot, 1976; Onasch, 1977; Davis and others, 

1978; Hillhouse and others, 1982), (2) dextral transpression along the Salmon River suture involving no 

subduction at this site of final suturing (Lund, 1984; 

Lund and Snee, 1988), and (3) collision penetration 

with subduction in western Idaho and tectonic escape 

to north (Wernicke and Klepacki, 1988; Selverstone 

and others, 1992). Prior to isotopic dating studies, estimates for the time of accretion were based on geologic constraints and range from Triassic to Cretaceous 

(Hamilton, 1963a; Brooks and Vallier, 1978; Dickinson 

and Thayer, 1978; Hillhouse and others, 1982; Lund, 

1984; Wernicke and Klepacki, 1988). Dating of minerals from metamorphic rocks, which formed during suturing, indicates that orogenic events in the Salmon 

River suture zone occurred during the interval from 

130 to 75 Ma (Sutter and others, 1984; Lund and 

Snee, 1988; Davidson, 1990; Getty and others, 1991).

ACKNOWLEDGMENT

This work was possible because of the guidance 

and inspiration of W.R. Greenwood, Robert Scholten, 

and Anna Hietanen.

FIGURE 14.1. General geology of Salmon River suture zone. Geology modified from Bond (1978), Brooks and Vallier (1978), Lund 

(1984), and McCollough (1984). Initial strontium isotope ratio 

line divides oceanic values of <0.704 (generally west of line) 

from continental values of >0.706 (generally east of line); isopleth was approximated from data published by Armstrong and 

others (1977) and by Fleck and Criss (1985). Ba, Baker City; 

BC, Big Creek; Bo, Boise; E, Elk City; G, Grangeville; Ha, 

Hamilton; Hr, Harpster; Hu, Huntington; L, Lewiston; Lu, Lucile; M, McCall; N, New Meadows; O, Orofino; Ox, Oxbow; R, 

Riggins; SC, Slate Creek; SM, South Mountain; YP, Yellow Pine.

ROCKS OF THE 

BLUE MOUNTAINS ISLAND ARC

Hamilton (1963a, 1969) provided the first comprehensive study of the oceanic rocks in western Idaho. 

He subdivided oceanic rocks near Riggins (fig. 14.1) 

into three stratigraphic sequences similar to the three 

tectonostratigraphic units used in the present study 

(fig. 14.3). The metamorphic, structural, and intrusive 

history of the tectonostratigraphic units is different 

enough to make it difficult to correlate map units 

among the three. The structurally lowest tectonostratigraphic unit is in stratigraphic continuity with Permian and Triassic rocks of the Wallowa terrane 

(Silberling and others, 1992) that is equivalent to the 

so-called Seven Devils island arc (Brooks and Vallier, 

1978; Dickinson, 1979; Hillhouse and others, 1982) 

and that is one of four terranes grouped together and 

named the Blue Mountains island arc (Vallier and 

Brooks, 1986) of eastern Oregon and western Idaho. 

Thus, a large portion of the island-arc rocks of western Idaho are known to be part of the Wallowa terrane. The middle tectonostratigraphic unit is herein 

called the Rapid River plate; it contains rocks that are 

along strike with the Riggins Group (Hamilton, 

1963a). The structurally uppermost tectonostratigraphic unit, herein called the North Fork block, consists of a complex of metavolcanic, metavolcaniclastic, 

and plutonic rocks; the premetamorphic-predeformation stratigraphy is not well preserved.

Rocks of the Riggins Group have been correlated in 

the past with rocks of the Baker and Izee terranes 

(Brooks and Vallier, 1978; Silberling and others, 1992) 

or with the Olds Ferry terrane (White and others, 

1992), which are all part of the Blue Mountains island 

arc of eastern Oregon and westernmost Idaho. Rocks of 

the North Fork block have not been previously correlated with any particular terrane of eastern Oregon because these metavolcanic rocks were not previously 

mapped separately from younger plutonic rock (Hamilton, 1963a, c; 1969). These two structurally highest 

tectonostratigraphic units were also tentatively correlated to the Wallowa terrane (Lund, 1984; Lund and 

Snee, 1988; Lund and others, 1993) on the basis of the 

model that these metamorphosed and deformed rocks 

were the leading edge of the Wallowa terrane that 

underwent the most severe effects of accretion to North 

America. This correlation with the Wallowa terrane 

was also suggested by Onasch (1977) and Vallier (1977).

The oceanic terranes in Oregon and Idaho have 

been variously correlated to regional superterranes. 

In particular, the Wallowa terrane first was suggested 

to be a piece of the western British Columbian Wrangellia terrane (Jones and others, 1977; Hillhouse and
Miocene Columbia River Basalt Group 

Late Cretaceous granite

Late Cretaceous tonalite to granodiorite

Late Cretaceous tonalite to granodiorite 

tectonite

EXPLANATION 

Oceanic Rocks

Triassic to Permian rocks of Wallowa terrane 

Mesozoic to Late Paleozoic(?) rocks of Rapid River plate 

Mesozoic to Late Paleozoic(?) rocks of North Fork block

Continental rocks

Paleozoic(?) to Late Proterozoic(?) metasedimentary rocks

Middle Proterozoic(?) metasedimentary rocks

-AA Rapid River 

concealed

A Thrust fault 

_ Normal fault 

side

-> Axis of antic

plunge

 Axis of sync

FIGURE 14.2. Simplified geology of Slate Creek-John Day Creek area (modified from Lund, 1984; McCollough, 1984; Lund and 
14. METAMORPHIC AND STRUCTURAL DEVELOPMENT OF ISLAND-ARC ROCKS, WEST-CENTRAL IDAHO 521

others, 1982) but more recently was argued to be correlative with the eastern British Columbian Stikinia 

terrane (Mortimer, 1986).

ROCKS OF THE WALLOWA TERRANE

SEVEN DEVILS GROUP

Volcanic rocks of the Seven Devils Group (included 

in "Triassic and Permian rocks of Wallowa terrane" 

unit in fig. 14.2; Vallier, 1977) crop out along the Salmon River near the mouth of John Day Creek and possibly as a small klippe in the Slate Creek drainage 

(Lund and others, 1993). Lithologically, these rocks 

consist of orange-weathering (green where fresh), massive metamorphosed tuffs and flows and metamorphosed volcanic conglomerate that consists of 

stretched volcanic clasts in a tuffaceous matrix. Recrystallized quartz, untwinned plagioclase, chlorite, 

tremolite, epidote, biotite, and muscovite make up the 

mineralogically variable fine-grained groundmass. 

Garnet and chlorite porphyroblasts are common; plagioclase occurs as both relict phenocrysts and metamorphic porphyroblasts. Marble layers less than 10 m 

thick in the upper part of the Seven Devils Group are 

known both in this area and in the Riggins area to the 

south (Hamilton, 1963a; Onasch, 1977).

A tonalitic to quartz dioritic pluton crops out in the 

Salmon River canyon (Hamilton, 1963a; McCollough, 

1984; Lund and others, 1993). The mineralogy of the 

pluton is plagioclase, quartz, altered hornblende, and 

chloritized biotite. The isotopic age of the pluton is 

about 259 Ma as determined by U-Pb zircon methods 

on two splits (Walker, 1986). Foliations in the pluton 

and in the metavolcanic country rock are parallel, 

but the contact between the two units is discordant 

to the foliation. This pluton is therefore interpreted 

to have intruded one of the two Permian units of the 

Seven Devils Group prior to the metamorphic event 

that formed the foliation probably during construction of the Blue Mountains island arc.

Although all of the four formations, which make up 

the Seven Devils Group elsewhere (fig. 14.3), have 

not been mapped in the Slate Creek-John Day Creek 

area, both the lower and upper parts of the group 

presumably crop out in the area. Because of the 259- 

Ma (Permian) age for the crosscutting pluton, the 

rocks of the Seven Devils Group that crop out along 

the Salmon River in this area (fig. 14.2) are presumably Early Permian. Either the Hunsaker Creek or 

Windy Ridge Formation could have been intruded by 

a 259-Ma pluton. Gradational stratigraphic contacts 

between the Seven Devils Group and the overlying 

calcareous phyllite (see section "Upper Triassic Martin Bridge Formation") indicate that the uppermost 

unit of the Seven Devils Group, the Upper Triassic 

Doyle Creek Formation, is also present in the study 

area. Others have reported a gradational upper contact for the Doyle Creek Formation in the Wallowa 

Mountains of eastern Oregon (Folio, 1994) and along 

the Snake River on the Oregon-Idaho boundary (fig. 

14.1; Vallier, 1977). The base of the Seven Devils 

Group is not exposed in the study area.

UPPER TRIASSIC MARTIN BRIDGE FORMATION

About one-third of the Slate Creek-John Day Creek 

area is underlain by a sequence of Upper Triassic 

metasedimentary rocks. The top of this sequence, 

which is truncated by faulting, is not known to crop 

out in western Idaho. The lower part of the section 

crops out between Slate Creek and John Day Creek; 

here greenschist-facies metasedimentary rocks grade 

into the underlying metavolcanic rocks of the Seven 

Devils Group. Also included in this unit are amphibolite-facies rocks in the core of the northeast-plunging 

Slate Creek antiform that crosses Slate Creek in the 

study area (fig. 14.2).

STRATIGRAPHY

There are two major and two minor rock types in 

this metasedimentary section. The most common 

rock type is calcareous phyllite to schist with which 

the less common rock types are interbedded. The calcareous phyllite to schist is composed of calcite, plagioclase, biotite or phlogopite, muscovite, chlorite, 

rounded fine-grained quartz, epidote, and siderite 

cubes. It also contains disseminated fine graphite 

grains. At higher metamorphic grades, the schist contains abundant sphene, amphibole, and garnet; in 

the John Day Creek drainage, andalusite, kyanite, 

sillimanite, chloritoid, staurolite, and prochlorite are 

reported (E.H. Price, 1985, written commun.). Gray 

and white, fine- to coarse-grained calcite marble, 

which forms layers from 5 to 500 m thick, is the second most common lithology. At high metamorphic 

grade, the marble usually contains a few percent 

phlogopite and a trace of fine-grained graphite.

The two minor rock types are metavolcanic and 

conglomeratic rocks. Green, chloritic metavolcanic 

layers occur at several stratigraphic horizons and are 

composed of fine, interlocking grains of plagioclase 

that form the groundmass for porphyroblastic intergrowths of actinolite and chlorite. Fine grains of 

opaque minerals, quartz, sphene, and apatite are 

also present. Lenses of calcareous schist containing
522 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

multilithologic clasts (fig. 14.4) crop out in a few 

places. Although first interpreted to be fault breccia 

(Lund, 1984), the presence of more than one carbonate-clast rock type suggests that these diamictites 

were originally limestone-conglomerate lenses similar 

to those found in the Martin Bridge Limestone in the 

Wallowa Mountains, eastern Oregon (Folio, 1994).

As seen between the mouth of Slate Creek and the 

mouth of John Day Creek, the lower contact between 

the Upper Triassic metasedimentary section and the 

metavolcanic rocks of the Seven Devils Group is interbedded through about 10 m of stratigraphic section (fig. 14.5). This contact zone is characterized by 

centimeter-scale internngering among layers of finely

laminated tuffaceous chloritic greenstone, layers of 

laminated gray calcareous phyllite, and lenses of 

quartz-carbonate rock (McCollough, 1984); phyllite 

predominates upward. Marble layers or lenses 5 to 

500 m thick occur at several stratigraphic horizons in 

the calcareous phyllite to schist section (fig. 14.6). At 

the intraformational contacts with the marble, interbedding of phyllitic and marble layers or gradation 

from calcareous phyllite to graphitic-quartzitic-micaceous marble to fairly pure marble occurs across 

about 2 m. A thick marble band (possibly as much as 

1000 m thick, probably technically thickened) forms 

the structurally truncated top of the Upper Triassic 

metasedimentary section from south of John Day

Wallowa terrane 

(lower plate)

1,2

\

_ _ Top of section _'

:^|lr|i^ti|rj|-:: ::::: ;'||:"::::;:

S^^lMiP

Rapid River plate 

(middle plate)

7s,~:' 7^<)f£!e<$0^^

North Fork plate 

(lower plate)

1 Vallier, 1977.

2 Folio, 1994.

3 Hamilton, 1963a. 

4 Onasch, 1977.

FIGURE 14.3. Generalized correlation diagram for oceanic rocks of western Idaho.
14. METAMORPHIC AND STRUCTURAL DEVELOPMENT OF ISLAND-ARC ROCKS, WEST-CENTRAL IDAHO 523

FIGURE 14.4. Multilithologic, calcareous-cobble conglomerate with calcareousschist matrix in amphibolite-facies Upper Triassic Martin Bridge Formation. 

Lens cap is 5.25 cm in diameter.

FIGURE 14.5. Detail of interbedding between metavolcanic-rock 

(light) and calcareous-shale (dark) layers at gradational top 

of greenschist-facies rocks of the Seven Devils Group. Lens 

cap is 5.25 cm in diameter.

Creek to Slate Creek (Onasch, 1977). Toward the 

north and west, this single, thick marble unit splits 

into several thinner marble layers that gradually rejoin to form a single layer again near the mouth of 

Slate Creek.

Stratigraphic relationships such as interbedding 

are difficult to recognize in the deformed, medium- to 

high-grade metamorphic rocks along Slate Creek and 

upper John Day Creek; commonly, stratigraphic layering cannot be distinguished from structural intercalation. In addition, transposition of bedding can be 

observed in the low-grade rocks (McCollough, 1984) 

but is difficult to identify in higher grade, multiply 

deformed rocks. However, despite local structural 

complexities, the majority of foliation measurements 

are consistently parallel to lithologic contacts. This 

parallelism, together with the apparently gradational 

contacts, indicates that the major lithologic contacts 

are original stratigraphic features.

CORRELATION

Correlation of the metasedimentary rocks is problematic. On the basis of reconnaissance mapping between the Snake and Salmon Rivers, the name 

"Lucile Series" (Wagner, 1945) was first applied to a 

sequence of metasedimentary and subordinate metavolcanic rocks near the Slate Creek-John Day Creek 

area. Because probable Triassic fossils were found in 

marble beds within metasedimentary rocks continuous with rocks of the "Lucile Series" near Riggins, 

Idaho, and because of metamorphic and structural 

complexity within the upper part of the "Lucile Series," Hamilton (1963a) reevaluated the stratigraphy
524 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

of these rocks. Along the east flank of the Seven Devils Mountains southwest of the study area, a thick 

marble layer that overlies the Seven Devils Group 

was correlated with the Martin Bridge Formation of 

the Wallowa Mountains and named the "Martin 

Bridge Limestone" (Hamilton, 1963a). South of Riggins, Hamilton (1963a) thought that the lower contact 

with the volcanic rocks was depositional although 

he cited no supporting evidence whereas north of 

Riggins, the contact was described as tectonic (Hamilton, 1963a). Calcareous slate to schist, which overlies 

the marble layer along the east flank of the Seven 

Devils Mountains, was named the "Lucile Slate" 

(Hamilton, 1963a). Although Hamilton did not do detailed mapping of the type locality of the "Lucile 

Slate" north of Riggins near the town of Lucile, Idaho, 

he described lithologic and structural characteristics 

of rocks between the town of Lucile and the mouth of 

Slate Creek: "Near Lucile, the [Lucile] slate is repeated structurally several times and lies between sheets 

of Martin Bridge Limestone and Seven Devils" Group 

(Hamilton, 1963a). The calcareous schist described at 

Hamilton's type locality for the "Lucile Slate" is in the 

present study area. This calcareous schist is less than 

5 km south of and along strike (in continuously exposed outcrops) with the calcareous phyllite in this 

study area that has gradational contacts with both

the underlying Seven Devils Group and the interbedded and overlying marble layers (Onasch, 1977; McCollough; 1984; Lund and others, 1993). Much of the 

complexity that Hamilton (1963a) described previously as structural is due to stratigraphic interfingering 

(McCollough, 1984).

Folio (1986, 1994) has described complex meterscale interfingering of limestone, organic-rich calcareous and noncalcareous shale, and argillite in the 

Upper Triassic Martin Bridge Formation (Limestone) 

where it was originally described in the southern 

Wallowa Mountains of Oregon (Ross, 1938). The Martin Bridge Formation is in gradational contact with 

and laterally interfingers with the overlying argillaceous sedimentary rocks, graywacke, calcareous argillite, and limestone of the Upper Triassic and 

Lower Jurassic Hurwal Formation (fig. 14.3; Folio, 

1986, 1992).

Detailed studies of unmetamorphosed and only 

slightly deformed rocks in the Wallowa Mountains indicate that interfingered units and limestone lenses 

in the Slate Creek-John Day Creek area of western 

Idaho are common in this complex island-arc setting. 

Although inadequate fossil control makes correlation 

difficult, the probable Triassic fossils found in the 

section south of Riggins, Idaho (Hamilton, 1963a), 

and mapping completed in other parts of the region

FIGURE 14.6. Marble layers and lenses interbedded with calcareous phyllite in greenschistfacies rocks of the Martin Bridge Formation. Intraformational marble layer outlined in 

photograph is 10 to 20 m thick. "Em, Upper Triassic Martin Bridge Formation; TiPs, Permian and Triassic Seven Devils Group.
14. METAMORPHIC AND STRUCTURAL DEVELOPMENT OF ISLAND-ARC ROCKS, WEST-CENTRAL IDAHO 525

(Vallier, 1977) confirms Hamilton's (1963a) conclusion 

that the metasedimentary section that overlies the 

Seven Devils Group in western Idaho is Upper Triassic. However, the entire metasedimentary section is 

most probably correlative with the Upper Triassic 

Martin Bridge Limestone of the Wallowa Mountains. 

Herein the unit is called the Martin Bridge Formation both because of prior use of this name for the 

unit in Oregon (Ross, 1936) and along the Snake 

River in Oregon and Idaho (Vallier, 1977) and because "limestone" describes neither the stratigraphic 

nor metamorphic character of the unit in western 

Idaho.

METAMORPHOSED ROCKS OF UNCERTAIN AGE

RAPID RIVER PLATE

Structurally overlying the rocks that are known to 

be part of the Wallowa terrane is a group of rocks 

that had both volcanic and volcaniclastic-sedimentary 

protoliths (fig. 14.3). Hamilton (1963a) described 

these rocks near Riggins, named them the Riggins 

Group, and subdivided them into formations (Fiddle 

Creek Schist, Lightning Creek Schist, Berg Creek 

Amphibolite, and Squaw Creek Schist); many aspects 

of the stratigraphy and structure of the Riggins 

Group near Riggins were refined by Onasch (1977, 

1987). These rocks are also found along strike in the 

Slate Creek-John Day Creek area, where metamorphic grade is generally higher than in the Riggins 

area. Metamorphic grade also increases eastward 

across the Slate Creek-John Day Creek area. Metamorphic variation across the study area complicates 

description of the formations of the Riggins Group; 

likewise, metamorphic variation across the region 

complicates correlations. Both the base and top of the 

Rapid River plate are faulted contacts. Therefore, 

rocks in the same structural position and along 

strike with the Riggins Group are herein treated as a 

single lithodemic unit and referred to as the Rapid 

River plate. Rocks of the Rapid River plate in the 

Slate Creek-John Day Creek area are named according to their lithology; problems and possible correlations are discussed separately.

DESCRIPTION OF ROCKS

The dominant rock type in the Rapid River plate is 

green-gray to dark-gray chlorite-biotite schist and 

hornblende-biotite gneiss (fig. 14.3). The mineralogy 

of these metavolcanic rocks changes with metamorphic grade. Plagioclase and subequal amounts of 

quartz are the most abundant minerals. Chlorite and

biotite are the most common mafic minerals in greenschist- to lower-amphibolite-facies metamorphic rocks 

on the western side of the area, whereas biotite and 

hornblende are the mafic minerals in the amphibolite-facies metamorphic rocks on the east side of the 

Rapid River plate. Retrograde chlorite is found in 

some of the hornblende-biotite gneiss in the Rapid 

River plate. The common accessory minerals throughout are garnet, opaque minerals, epidote, muscovite, 

and apatite. Interlayered in the chlorite-biotite schist 

and hornblende-biotite gneiss unit are two minor rock 

types of special note. A subordinate but important 

rock type is hornblende-biotite conglomeratic gneiss 

that has relatively felsic multilithologic clasts; this 

gneiss forms layers that probably represent metamorphosed volcaniclastic conglomerate beds. Another 

minor rock type is a more feldspathic gneiss with 

fine-grained plagioclase, quartz, and biotite as the 

common minerals. The protolith for this gneiss may 

have been fine-grained plutonic rocks or a more felsic 

volcanic flow.

The second most common metavolcanic rock type 

in the Rapid River plate is greenish muscovite-chlorite schist or orange-weathering muscovite schist (fig. 

14.3). The major minerals in the groundmass are 

muscovite, plagioclase, quartz, and chlorite. These 

schists commonly contain porphyroblastic garnet, 

chlorite, biotite, or quartz; many of the garnet porphyroblasts have chloritic rims (McCollough, 1984). 

Accessory minerals include common plagioclase, epidote, and opaque grains as well as sparse carbonate 

minerals.

Epidote-garnet-muscovite-biotite leucotonalite to 

quartz diorite in the chlorite-biotite schist and biotite 

gneiss unit along lower and middle Slate Creek (Lund 

and others, 1993) is metamorphosed and deformed. 

Plagioclase is the most common mineral in the felsic 

metaplutonic rocks; it forms untwinned, sericitized 

intergrowths with other minerals. Other primary 

minerals in these rocks are fine-grained and strained 

quartz, chloritized biotite, and accessory apatite. As 

much as 5 percent of fine- to medium-grained secondary pyrite and calcite are present; garnet, muscovite, 

and epidote are metamorphic minerals. Although no 

radiometric ages are available for these rocks, foliation in the pluton and associated dikes is concordant 

with that in the metavolcanic country rocks whereas 

the contact between the rocks is commonly discordant 

to the foliation. The metamorphic overprint on the 

pluton and the parallel foliation in the two rock types 

indicate that, although possibly the pluton intruded 

the volcanic host rocks during formation of the island 

arc and may be Permian to Triassic, (see "Proposed 

Correlation" section under "Rapid River Plate"), the
526 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

pluton was definitely emplaced prior to the metamorphic event that formed the foliation.

Garnet-muscovite-chlorite schist or muscovite schist 

(depending on metamorphic grade) is commonly found 

near the structural base of the plate. In upper John Day 

Creek, the thickness of this schist is estimated to be 

2,700 m (McCollough, 1984), whereas in Slate Creek, it 

is common to have less than 50 m preserved at the base 

of the Rapid River plate. Chlorite-biotite schist and 

hornblende-biotite gneiss compose the structurally 

upper unit (fig. 14.3). The original contact between 

these rock types was not observed; the type of contact 

and the relative age of these units are unknown. The 

only evidence of primary depositional features in the 

metavolcanic rocks is volcaniclastic conglomerate that 

forms layers in the hornblende-biotite gneiss. In general, the foliation parallels the major lithologic breaks 

and relict conglomeratic layers. Therefore, except in 

areas of local structural disturbance, foliation in these 

rocks generally parallels bedding.

PROPOSED CORRELATION

Physical continuity indicates that garnet-muscovitechlorite schist and muscovite schist in the John Day 

Creek drainage (McCollough, 1984) are equivalent to 

the Fiddle Creek Schist (the lowest Riggins Group 

unit, fig. 14.3) that has been traced into the south side 

of John Day Creek (Onasch, 1977, 1987). On the basis 

of mineralogy and chemistry of rocks of the Riggins 

Group south of this area (Hamilton, 1963a), muscovite 

schist in the Slate Creek area is probably also part of 

the Fiddle Creek Schist. However, direct tracing of the 

Lightning Creek Schist, the lowest unit of the Riggins 

Group, into the Slate Creek-John Day Creek area is 

not possible and the mafic schist and gneiss in the 

study area cannot be confidently correlated because 

rocks of this composition are found in both the Fiddle 

Creek and Lightning Creek Schists (Hamilton, 1963a; 

Onasch, 1977). Metavolcaniclastic rocks similar to the 

Squaw Creek Schist of the upper Riggins Group do 

not form part of the Rapid River plate in the Slate 

Creek-John Day Creek area.

No direct age information is available for rocks of 

the Rapid River plate. Although undated rocks that 

are part of the Rapid River plate in the John Day 

Creek drainage were previously considered to be correlative, in part, with the undivided Riggins Group 

(Hamilton, 1963a), and physical connection between 

one formation of the Riggins Group (Fiddle Creek 

Schist) and schists of the Rapid River plate is documented on the south side of John Day Creek (Onasch, 

1977, 1987; McCollough, 1984), the age of the Riggins 

Group is also uncertain (Hamilton, 1963a). In light of

suggestions that the Riggins Group is possibly a higher grade, more deformed equivalent of part of the 

Wallowa terrane (Onasch, 1977; Vallier, 1977; Lund, 

1984; also see "Discussion" section), rocks of the Rapid River plate are also probably Triassic to Permian. 

If the Riggins Group is correlative with units in the 

Baker terrane of eastern Oregon (Elkhorn Ridge Argillite and Burnt River Schist) as has also been suggested (Vallier, 1977; Brooks and Vallier, 1978), the 

age is still uncertain.

NORTH FORK BLOCK

The North Fork reverse fault separates multiply 

deformed, complexly intruded amphibolite-facies 

metavolcanic rocks of the North Fork block from 

rocks of the Rapid River plate that lie structurally 

below and to the west; the Salmon River suture 

forms the east boundary of the plate (fig. 14.2). Metavolcanic rocks of the North Fork block are preserved as screens, roof pendants, and xenoliths 

within deformed plutonic rocks. Because of this, previous workers generally did not differentiate between the metamorphic and igneous rocks that crop 

out in this complex structural position. However, it 

is generally accepted that the metamorphic rocks 

had an island-arc protolith, and some workers have 

suggested that similar rocks to the south may be 

equivalent to the Riggins Group (Hamilton, 1963a, 

b; Onasch, 1977, 1987) and therefore equivalent to 

rocks of the Rapid River plate.

DESCRIPTION OF ROCKS

The primary rock type in the North Fork block is 

garnet-biotite-hornblende feldspathic gneiss and garnet amphibolite gneiss. Most of the rocks contain distinct mafic and quartzofeldspathic layers. Rolled 

porphyroblasts of garnet and hornblende are common. 

A complex paragenetic sequence of metamorphic minerals is preserved. In many places, two phases of the 

same mineral with different orientations and (or) characteristics are present; for instance, early formed hornblende porphyroblasts are wrapped by oriented, 

medium-grained hornblende grains (fig. 14.7). A less 

common rock type in the North Fork block is biotitecalc-silicate gneiss; centimeter-scale quartz-rich layers 

are common and marble was noted in this plate near 

the headwaters of John Day Creek (E.H. Price, oral 

commun., 1987). These calc-silicate gneiss and calcareous rocks occur along the east side of the North Fork 

block. An uncommon rock type is garnet-muscovitebiotite schist that is in a few isolated localities near the 

west edge of the North Fork block.
14. METAMORPHIC AND STRUCTURAL DEVELOPMENT OF ISLAND-ARC ROCKS, WEST-CENTRAL IDAHO 527

Rocks in the North Fork block contain no known 

original sedimentary features. They are structurally 

discontinuous with rocks in the Wallowa terrane or 

Rapid River plate; the rocks are multiply deformed 

and complexly intruded by tonalite plutons. Exposures of the metamorphic rock are isolated in the 

dominant plutonic rock, and, after intrusion, metamorphic and plutonic rock were concordantly deformed. As a result, lithologic units commonly lack 

mappable continuity. For rocks of the North Fork 

block, layering probably does not parallel bedding 

and original stratigraphic relationships among lithologic units are unknown.

PROPOSED CORRELATION

The metavolcanic rocks in the North Fork block previously were not separately mapped or correlated to 

specific units mapped elsewhere. This igneous-metamorphic complex (or "gneiss complex" of Hamilton, 

1963a) was referred to as the "western border zone of 

the Idaho batholith" (Hamilton, 1963b, c; Vallier and 

Brooks, 1987). Hamilton (1963a) suggested that metamorphic rocks within the complex are at least partly 

correlative with the Riggins Group (fig. 14.3).

There is no direct information available on the age 

of metamorphic rocks in the North Fork block. Based 

on the Slate Creek area, it is apparent from both

structural studies and dating of the metamorphic 

rocks (Lund, 1984; Lund and Snee, 1988) that there is 

no direct stratigraphic continuity between the Rapid 

River and North Fork blocks. Therefore, all that can 

be concluded about the correlation of these rocks is 

that they are metavolcanic and metavolcaniclastic 

rocks of probable island-arc origin (Hamilton, 1963a; 

Fleck and Criss, 1985) and that, on the basis of lithology, they are probably higher grade, structurally more 

complex equivalents of rocks in the Wallowa terrane 

and (or) the Rapid River plate (fig. 14.3). Accordingly, 

garnet-biotite-hornblende gneiss and garnet amphibolite gneiss are possibly equivalent to metavolcanic 

rocks of the Fiddle Creek and Lightning Creek Schists 

of the Riggins Group (see Hamilton, 1963a) and (or) to 

metavolcanic rocks of the Seven Devils Group. The 

rocks with calc-silicate- and quartz-bearing layers are 

possibly equivalent to the Squaw Creek Schist of the 

Riggins Group described by Hamilton (1963a) and (or) 

to the Martin Bridge Formation.

LATE CRETACEOUS CROSSCUTTING 

PLUTONIC ROCKS

Quartz diorite, tonalite, and granodiorite plutons intruded metamorphosed and deformed rocks of the

1 MILLIMETER 

J

FIGURE 14.7. Photomicrograph showing evidence of overprinted dynamothermal events 

in hornblende gneiss from the North Fork plate. Hornblende porphyroblast of S^ fabric 

is wrapped by hornblende of S4 fabric; postkinematic garnet postdates both. Hbi, earlier hornblende; Hb4, later hornblende; Gt, garnet.
528 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

Wallowa terrane, Rapid River plate, and North Fork 

block as well as the North American continent to the 

east; these plutonic rocks are the early phases of the 

Idaho batholith and are undivided on figure 14.2, where 

plutonic rocks of Late Cretaceous through Eocene age 

are shown as one unit. These rocks are progressively 

more common eastward across the area; they occur over 

only a small part of the middle of the area, whereas the 

eastern part of the area is dominated by these plutonic 

rocks. Thus, in the North Fork block, metamorphic 

rocks are found primarily as roof pendants and xenoliths within these plutonic rocks (fig. 14.2). These plutonic rocks are grouped separately from those thought 

to be island-arc related or premetamorphic because 

they (1) do not contain retrograde-metamorphic mineral assemblages, (2) cut metamorphic fabrics and 

postmetamorphic structures, (3) are affected by late 

deformation that cuts earlier features, (4) have isotopic 

ages that are younger than ages of metamorphism 

affecting other plutonic rocks (see "Isotopic Ages" section and Snee and others, chap. 10, this volume), and 

(or) (5) can be shown to stitch island arc to continental 

rocks by intruding across the Salmon River suture. In 

this chapter, only these and younger, autochthonous or 

semiautochthonous plutons are considered part of the 

Idaho batholith.

ROCK UNITS

The prevalent postmetamorphic plutonic rocks in 

Slate Creek-John Day Creek area are hornblende-biotite and biotite tonalite. Primary epidote (Zen and 

Hammarstom, 1984; Zen, 1985) and sphene are common accessory minerals and, if present, are usually 

coarse enough to be seen in hand specimen. Textures 

vary according to structural position across the area 

but generally may be divided into three major textural groups: (1) Plutonic rocks that cut continental 

rocks to the east (Lund, 1984) are mostly homogeneous in texture although faint, inconsistently oriented 

foliation is present in places. (2) Plutonic rocks that 

intruded Wallowa terrane and Rapid River plate 

rocks in the center of the Slate Creek-John Day 

Creek area cut metamorphic fabric, the Rapid river 

thrust fault, and the Slate Creek antiform but were 

intruded into and deformed along reverse faults related to pervasive deformation and emplacement of 

the North Fork block. These plutonic rocks are in 

places foliated and contain secondary muscovite. (3) 

Plutonic rocks that intruded the North Fork block 

have ubiquitous strong fabrics wherein postplutonic 

foliation and lineation have been superimposed on 

both the plutonic rocks and the metamorphic country

rocks (see "Structure" section). Foliation, where 

present, is defined by biotite; lineation is defined by 

hornblende. Quartz grains are strained.

Less dominant quartz diorites and granodiorites 

that are similar in mineralogy and texture have been 

noted in the field but not mapped separately. It is not 

known if these are local phases of one pluton or if 

several different plutons are present. Plutonic boundaries have not been found, and any that existed may 

have been obscured by subsequent deformation.

ISOTOPIC AGES

Age information for these plutonic rocks comes 

from field relationships and from dating of both metamorphic events and plutonic rocks. In the central 

part of the area, these intrusive rocks cut across the 

metamorphic fabric and postmetamorphic structures 

that deform rocks of the Wallowa terrane and Rapid 

River plate (Lund, 1984). 40Ar/39Ar-age-spectrum 

analysis indicates that metamorphic events in rocks 

of the Wallowa terrane and Rapid River plate occurred before 101 Ma. Radiometric dates for the 

crosscutting plutonic rocks are 93 to 85 Ma. In addition, U-Pb data for sphene-epidote-hornblende-biotite 

tonalite on the east side of the area give a concordant 

age (on four zircon splits and one sphene split) of 

90.4±3.0 Ma (Snee and others, chap. 10, this volume). 

In the North Fork block, 40Ar/39Ar data shows that 

the record of an approximately 100-Ma event was destroyed by another event that caused argon loss and 

reequilibration about 90 Ma.

Because these Late Cretaceous plutons intruded 

across the suture and are unmetamorphosed, they 

are considered to be the earliest plutons of the Idaho 

batholith; some workers have even referred to them 

as a specific compositional phase of the Idaho batholith that occupies a peculiar structural setting and 

have called them the "western border zone of the 

Idaho batholith" (Hamilton, 1963a; Vallier and 

Brooks, 1987). However, chemically identical plutons 

of the same age can be found across the top and on 

the east side of the central Idaho batholith (Lund, 

1984; Lund and others, 1986; Lewis and others, 

1987; Toth, 1987). These tonalitic plutons seem to 

have been widespread across the area of the batholith rather than limited to the west side.

METAMORPHIC HISTORY

The Slate Creek-John Day Creek area has a complex metamorphic history involving several apparent
14. METAMORPHIC AND STRUCTURAL DEVELOPMENT OF ISLAND-ARC ROCKS, WEST-CENTRAL IDAHO 529

episodes of metamorphism and disruption of metamorphic patterns caused both by overlapping metamorphic events and by deformation. The pervasive, 

prograde metamorphic events, which are very strongly represented in the Slate Creek-John Day Creek 

area, are younger than metamorphic events documented in the island-arc rocks about 35 km outboard 

from the suture (Hotz and others, 1977; Ave Lallemant and others, 1980, 1985). Information from outside of the study area provides the only evidence of 

earlier metamorphic history in these rocks because, 

near the Salmon river suture, evidence of earlier 

events has been largely obliterated by complex higher grade metamorphism.

Although not manifested in the Slate Creek-John 

Day Creek area, the island-arc rocks apparently underwent low-grade metamorphism and reaction to 

seawater during and immediately following deposition 

or emplacement. The resulting spilite, keratophyre, 

and slate are preserved in parts of the Wallowa terrane that crop out in the Seven Devils Mountains and 

along the Snake River in Idaho as well as in Oregon 

(Hamilton, 1963a; Vallier, 1977).

It is probable that earlier Mesozoic metamorphism 

and deformation affected rocks in the Slate CreekJohn Day Creek area because outboard rocks exhibit 

effects of several earlier dynamothermal events: (1) 

The Baker terrane (Silberling and others, 1992), 

which is one suggested correlative for metavolcanic 

rocks of the Rapid River plate (Brooks and Vallier, 

1978), underwent Triassic metamorphism and deformation (Hotz and others, 1977; Ave Lallemant and 

others, 1980, 1985). (2) Rocks of the Wallowa terrane 

in the Wallowa Mountains of Oregon (probable correlatives of most metamorphic rocks in the study 

area) underwent pre-Late Jurassic deformation 

(Prostka, 1962; Folio, 1986). (3) On the basis of tectonic models, it is possible that rocks along the leading edge of the oceanic terrane underwent 

metamorphism and deformation related to collision 

with North America at some other location prior to 

final accretion in what is now western Idaho. What is 

inferred about the early Mesozoic history depends on 

the preferred correlation between high-grade rocks of 

the Rapid River plate and North Fork block, on how 

widespread earlier orogenic activity was, as well as 

on tectonic models for accretion of these rocks to the 

continent.

DESCRIPTION

In the Slate Creek-John Day Creek area, earlier 

metamorphic effects are overprinted and, in most

cases, obliterated by prograde regional metamorphism. As elsewhere in western Idaho (Hamilton, 

1963a; Onasch, 1977, 1987; Myers, 1982), metamorphic grade and complexity increase toward the Salmon River suture from greenschist facies on the west 

side of the area (fig. 14.2) to upper amphibolite facies 

near the Salmon River suture on the east. This occurs mostly by means of structural shortening whereby several major thrust or reverse faults juxtapose 

rocks having different metamorphic histories and 

bring higher grade rocks over lower grade rocks.

Rocks of the Wallowa terrane generally record a 

fairly simple metamorphic history. Only one foliation 

is seen in these rocks. At low metamorphic grade, 

porphyroblasts (especially hornblende) are randomly 

oriented in the plane of foliation; at higher metamorphic grades, porphyroblasts are well aligned 

within the single foliation. These single fabric rocks 

contain evidence for only a single dynamothermal 

event. Mineral growth after the metamorphic peak 

or related to a late-stage, nondeformational metamorphic event is indicated by common unoriented 

biotite grains in the lower- to middle-amphibolitefacies rocks.

Fabrics in rocks of the Rapid River plate are more 

complex. Along upper John Day Creek, garnet- and 

biotite-porphyroblastic muscovite schists contain evidence of two high-grade events. In the first event, 

helicitic-garnet and biotite porphyroblasts were 

formed and the muscovite-rich fabric was wrapped 

around the porphyroblasts; after deformation, euhedral rims grew over the syntectonic-garnet cores 

(McCollough, 1984). A lower grade dynamothermal 

overprint in these rocks is demonstrated by deformed chlorite porphyroblasts that grew across the 

preexisting fabric (McCollough, 1984). In addition, 

evidence of low-grade metamorphic overprinting of 

higher grade events is found near Riggins south of 

this study area, where oriented chlorite flakes grew 

in the axial-planar cleavage of a postmetamorphic 

megascopic fold in middle-amphibolite-facies rocks 

(Onasch, 1977, 1987).

Metavolcanic rocks of the North Fork block contain 

clear evidence of more than one dynamothermal 

event. Two nearly equal magnitude metamorphic 

events are indicated in many places by the presence 

of two generations of the same mineral that have different orientations and characteristics. In some of 

these rocks, deformed metamorphic hornblende porphyroblasts form augen; these augen are wrapped by 

a finer grained groundmass that has a secondary 

schistosity containing a second-growth hornblende. 

Postkinematic garnets commonly overprint these oriented fabrics (fig. 14.7).
530 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

AGES OF METAMORPHISM

As just described, it is common, except in the lowest grade rocks, to observe evidence of two or more 

metamorphic events in most rocks in the region. This 

preserved overprinting is characterized by (1) common unoriented biotite flakes superimposed on the 

earlier upper-greenschist- to lower-amphibolite-facies 

directional fabric; (2) garnet porphyroblasts, formed 

in response to prograde dynamothermal amphibolitefacies metamorphism and subsequently cut by a metamorphic fabric of slightly lower grade; and (3) in 

most middle to upper amphibolite-facies rocks along 

the east side of the area, multiple generations of amphibole growth punctuated by episodes of deformation. 

Detailed 40Ar/39Ar-age-spectrum dating of metamorphism in these rocks (Snee and others, 1987, chap. 

10, this volume; Lund and Snee, 1988) revealed additional information on the metamorphic history. Where 

the appropriate mineralogy is present, it is possible to 

date either the prograde metamorphic event or the 

retrograde metamorphic event, and, in some cases it 

is even possible to date both the major metamorphic 

event and the overprinting event.

40Ar/39Ar-age-spectrum dating of metamorphic 

minerals from the lowest grade, simplest island-arc 

metamorphic rocks shows that the closest approximation of time of formation of foliation in the greenschist-facies rocks is about 120 Ma. This age is 

constrained by data from foliated but nonlineated 

hornblende at its lower stability limit in greenschistfacies metavolcanic rocks (Snee and others, chap. 10, 

this volume). This is the best information on the 

onset of metamorphism and deformation that can be 

gained from rocks in the study area. An apparent age 

for the overprinting, randomly oriented biotite porphyroblasts from lower-amphibolite-facies calcareous 

schist of the Martin Bridge Formation (Wallowa terrane) of about 84 Ma indicates local late reheating at 

that time.

Metamorphic ages of hornblende from higher 

grade, more complex rocks of the Rapid River plate 

form two age groups about 109 and 101 Ma. Age data 

for some of the rock samples, which contain overprinting chlorite porphyroblasts, indicate initial cooling at about 109 Ma and also an argon-loss event at 

about 94 Ma (Snee and others, chap. 10, this volume); these ages represent both cooling after prograde metamorphism that formed the hornblende 

fabric and a later reheating event that caused the 

growth of overprinting chlorite porphyroblasts.

40Ar/39Ar-age-spectrum data for hornblende from 

the paragenetically complex metavolcanic rocks of the 

North Fork block commonly exhibit hybrid age spectra

indicating argon loss from about 101 to 90 Ma or reset 

ages of about 90 Ma. These data indicate that the 

metavolcanic rocks cooled to below 530°C by 101 Ma 

and were subsequently reheated and deformed before 

final cooling about 90 Ma.

The age data show clearly that metamorphism 

spanned from about 120 to 90 Ma. Careful age-spectrum work documents the eastward-increasing 

complexities of the paragenetic sequences and metamorphic overprinting as well as documenting the 

timing of events. The apparent eastward decrease in 

metamorphic ages may be in part related to simple 

cooling of rocks at deeper levels within the erogenic 

belt or to greater eastward structural disruption that 

brought deeper rocks up (from the east) to cool at 

later times. However, on the basis of age data and 

the mineralogic and deformational overprinting revealed by examination of the metamorphic textures, 

it is apparent that some aspects of the eastward increasing complexity of metamorphic fabrics and eastward-decreasing ages are due to distinct cycles of 

metamorphism and deformation.

STRUCTURE

Deformation as well as metamorphism becomes 

more intense to the east across the Slate Creek-John 

Day Creek area, and structural features differ accordingly. For this reason, mesoscopic (outcrop-scale) 

structures described in low-grade rocks to the south 

(Onasch, 1977,1987) and on the west side of this study 

area commonly take a different form from and are 

difficult to correlate with mesoscopic structures in the 

high-grade rocks in eastern John Day Creek or along 

Slate Creek. In addition, there are many structural 

domains caused by overprinting events in these rocks 

as well as several structural plates, and same-age features commonly have different orientations. Tracing 

macroscopic (map-scale) structures from area to area 

and interpreting the associated local mesoscopic structures in terms of the macroscopic structures are the 

best ways to piece together the deformational history 

of this area.

FABRICS AND GEOMETRIES

According to McCollough (1984), the low-grade calcareous slate to schist of the Wallowa terrane near 

the mouth of John Day Creek contains evidence of 

folding that produced schistosity. The regional foliation (Si), which is a demonstrable axial-planar schistosity or cleavage in this part of the study area, is
14. METAMORPHIC AND STRUCTURAL DEVELOPMENT OF ISLAND-ARC ROCKS, WEST-CENTRAL IDAHO 531

defined by the preferred orientation of muscovite or 

chlorite in pelitic rocks and of carbonate grains in 

calcareous rocks where pressure solution may have 

been the major process in forming directional fabric. 

Except in fold hinges, the axial-planar fabric (S^ 

parallels bedding (S0; McCollough, 1984). The first 

generation folds (F^, in bedding (So), are tight to isoclinal folds that plunge shallowly to the north, south, 

and east. Upper-greenschist- to amphibolite-facies 

rocks along Slate Creek and upper John Day Creek 

do not retain details of S0 (bedding), and F± (early 

folds) that produced Si (foliation) are not found because of more intense dynamothermal processes.

The low-grade calcareous slate to phyllite of the 

Martin Bridge Formation contains a particularly 

spectacular array of mesoscopic structures presumably because the composition and grade are conducive to the formation and preservation of minor local 

structures that cannot be identified in the massive 

metavolcanic rocks or marble that make up the rest 

of the section. Although these structures attest to the 

complexity of the structural history, they commonly 

cannot be correlated with macroscopic features. In

EXPLANATION

Fold axis 

X F2

 F3 

O F4 

l^> Sense of rotation

A

A

A 

Pole to fold axial plane

F2

F'

F4

7i F3 at 21°,041°

FIGURE 14.8. Lower-hemisphere, equal-area projection of 

mesoscopic structures in rocks of the Wallowa terrane in 

Slate Creek area, n pole (©) is fold axis of F3 fold determined from fig. 14.11.

this greenschist-facies calcareous slate to phyllite, 

the Si foliation is locally deformed by later northwest-verging asymmetric folds of variable orientation, and these as well as the early isoclinal FI folds 

are locally overprinted by northeast-plunging open 

folds with northwest-dipping closely spaced fracture 

cleavage (McCollough, 1984). The higher grade rocks 

of the Wallowa terrane and Rapid River plate contain 

evidence for only one phase of early folding F2 (figs. 

14.8 and 14.9) that deforms the foliation. In the higher grade rocks, the F2 folds are tight to isoclinal folds 

of Sl5 the principle schistosity. F2 fold axes plunge at 

a low angle toward the northeast or southwest, but 

the dip of axial planes varies from place to place.

A thrust fault juxtaposes amphibolite-facies rocks 

of the Rapid River plate over upper-greenschist- to 

amphibolite-facies rocks of the Wallowa terrane (fig. 

14.10). This fault can be traced discontinuously 

across the Slate Creek-John Day Creek area. In the 

John Day Creek area, the trace of this fault connects 

with the Rapid River thrust fault (Onasch, 1977, 

1987; McCollough, 1984). Mesoscopic structures in 

the narrow Rapid River thrust fault zone verge westEXPLANATION

Fold axis 

X F2

 F3 

O F4 

l<~^ Sense of rotation

A

AA 

Pole to fold axial plane

F2

F'

F4

7i F3 at 21°,041°

FIGURE 14.9. Lower-hemisphere, equal-area projection of 

mesoscopic structures in rocks from the Rapid River 

plate in Slate Creek area, n pole (ffi) is fold axis of F3 

fold determined from fig. 14.11.
532 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

FIGURE 14.10. Rapid River thrust fault along Slate Creek. Geologist is pointing to 

fault surface. Hanging wall consists of middle-amphibolite-facies metavolcaniclastic 

conglomerate of Rapid River plate (mcgl); footwall consists of lower-amphibolite-facies calcareous schist of the Upper Triassic Martin Bridge Formation (~Rm) of the 

Wallowa terrane.

" *%/*& 

_ * * V

EXPLANATION 

 Pole to Sx 

 Trace of F3 axial plane 

© TC F3 at 21°,041°

FIGURE 14.11. Lower-hemisphere, equal-area projection 

showing poles to Sj foliation in rocks of the Wallowa 

terrane and Rapid River plate. Girdle represents axial 

plane to F3 fold, n is pole to F3 axial plane.

northwest. The tight F2 folds of foliation (figs. 14.8 

and 14.9) may have formed during movement along 

the thrust fault.

On a macroscopic scale, the Rapid River thrust 

fault is folded into the gently northeast-plunging 

Slate Creek antiform (figs. 14.2 and 14.11) that has 

been subsequently deformed on the east limb. Related mesoscopic symmetric and asymmetric F% folds 

(figs. 14.8 and 14.9) that refold earlier mesoscopic 

folds are found throughout the area. These folds 

plunge consistently northeast at a low angle. Folds 

on the east limb are commonly asymmetric with an 

west-verging sense of rotation.

Evidence of the presumably latest phase of folding 

is found in the high-grade rocks on the east limb of 

the Slate Creek antiform. Macroscopic reverse faults 

disrupt the trace of the Rapid River thrust fault and 

the Slate Creek antiform (fig. 14.2). These faults cut 

the core of the Slate Creek antiform and rotated the 

east limb to vertical or overturned. Klippen of higher 

grade rocks on lower grade rocks along lower John 

Day and Slate Creeks may also be the result of the 

same deformation. Additionally this event is manifested by F4 upright, open mesoscopic folds and 

crenulations that plunge north-northeast at a low 

angle with steeply dipping axial planes and consistent west-vergent rotation (figs. 14.8 and 14.9). The 

synclines are commonly cut out by steep reverse
14. METAMORPHIC AND STRUCTURAL DEVELOPMENT OF ISLAND-ARC ROCKS, WEST-CENTRAL IDAHO 533

faults that formed during the late stages of F4 folding. Boudinage structures formed along many of the 

exposed faults by the deformation of minor leucocratic tonalite dikes that were emplaced along the faults 

(fig. 14.12). Thus, this complex stage of deformation 

began with ductile folding and continued into brittle 

faulting, then magma was emplaced and the resulting dikes were deformed during the last stages of the 

brittle faulting.

Rocks of the Wallowa terrane and Rapid River 

plate that contain sequential evidence of (1) early 

folding, (2) west-directed movement along the Rapid 

River thrust fault, (3) folding of the Slate Creek antiform, and (4) west-vergent folding and reverse faulting form the footwall to the steeply east-dipping 

North Fork reverse fault. These lower-plate rocks, 

which retain evidence of multiple deformations of the 

primary schistosity (Sj), are juxtaposed against deformed metamorphic and plutonic rocks of the North

FIGURE 14.12. Detail of F4 folds in calcareous schist of the Upper 

Triassic Martin Bridge Formation (~Em) of the Wallowa terrane. 

East-vergent, mesoscopic folds were faulted and intruded by 

Late Cretaceous tonalite (Kt); tonalite shows boudinage structure, which formed along fault during final movement. Shaft of 

hammer is approximately 27 cm long.

Fork block (upper plate). Metavolcanic rocks of the 

North Fork block retain only vestiges of the earlier 

fabric and amphibolite-facies mineralogy but are almost completely transformed by younger amphibolite-facies minerals of an overprinted younger fabric. 

The paragenetic sequence of metamorphic minerals 

and their structural characteristics (described earlier 

in this section) indicates that the early amphibolitefacies minerals, which are probably contemporaneous 

with the primary schistosity (still preserved in less 

complex rocks of the footwall), were deformed prior 

to the growth of younger amphibolite-facies minerals 

(fig. 14.7). Second-stage hornblendes are within a 

younger foliation (84) or define a secondary lineation 

(1/4) that has an orientation markedly different from 

the orientation of earlier hornblendes. The dominant 

fabric in both the deformed metamorphic and igneous 

rocks of the North Fork block is characterized by 

steeply inclined north-northeast-trending 84 foliation 

and (or) steeply southeast-plunging L4 mineral lineation (fig. 14.13). The fabric is steepest along the suture and becomes more shallow at higher elevations 

on the west side of the North Fork block (Lund and 

others, 1993). Although most of the rocks are L- and 

(or) S-tectonites indicative of plane strain, the occuro

n

EXPLANATION 

Pole to S4 foliation

Metamorphic rocks 

Intrusive rocks 

L4 lineation

FIGURE 14.13. Lower-hemisphere, equal-area projection 

showing poles to S4 foliation and L4 lineation in rocks of 

North Fork block.
534 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

rence of L-tectonites indicates a component of intense 

unidirectional vertical flow (pure shear) during deformation. Evidence for rotational deformation (simple 

shear) has not been found in the North Fork block in 

the Slate Creek-John Day Creek area.

Although obscured by intrusive rocks, the North 

Fork reverse fault cuts off the Rapid River thrust 

fault and the nose of the Slate Creek antiform and is 

therefore younger (fig. 14.2; Lund and others, 1993). 

The pervasive S4 schistosity and L4 lineation in the 

North Fork block formed in response to extreme 

strain. Formation of the new fabric destroyed evidence of bedding and transposed lithologic and plutonic contacts. Also, the imposed changes in 

mineralogy and fabric effectively cut off or transposed 

early structures. However, the multiple generations of 

some minerals and the microfabrics are good evidence 

of earlier metamorphic and structural events. Therefore, the overprinting fabric in rocks of the North Fork 

block is younger than both the regional Sj foliation 

and the F2 and F3 structures in the Wallowa terrane 

and Rapid River plate. Movement along the North 

Fork reverse fault and the young deformation (S4 and 

L4) in the North Fork block are probably contemporaneous with (1) formation of the F4 west-vergent folds 

and reverse faults that are on the east side of the 

lower plate (Martin Bridge Formation) near the North 

Fork reverse fault and (2) formation of late, moderately plunging crenulation lineations in rocks of upper 

John Day Creek (McCollough, 1984).

Crosscutting structures and metamorphic-mineral 

paragenesis evident in the regional-metamorphic fabrics indicate progressively younger deformation toward the Salmon River suture. Characteristically, the 

younger deformation in the higher grade rocks is 

more pervasive than older deformation in lower 

grade rocks and resulted in virtual obliteration of 

early structures.

AGES

A 40Ar/39Ar-age-spectrum study of metamorphic 

rocks in the Salmon River suture zone provides information on the age of specific structures and of the 

duration and complexity of dynamothermal processes 

in the area. A minimum of four important successive 

deformational events are recognized from progressive 

overprinting in rocks of the Slate Creek-John Day 

Creek area: (1) formation of the Sj primary regional 

foliation, (2) regional juxtaposition of higher grade 

rocks over lower grade rocks along the Rapid River 

thrust fault and simultaneous formation of F2 folds, 

(3) macroscopic F3 folding of the regionally juxtaposed rocks represented by the Slate Creek antiform

(Lund, 1984) and possibly the Riggins synform and 

Lake Creek antiform (Hamilton, 1963a; Onasch, 

1977), and (4) initiation of the North Fork reverse 

fault, dynamothermal overprinting S4 schistosity and 

L4 lineation in rocks of the North Fork block, and 

disruption of the Slate Creek antiform by reverse 

faulting coincident with F4 folding.

Dating of poorly formed, nonlineated, metamorphic 

hornblende in greenschist-facies, foliated metavolcanic 

rocks shows that the Sj foliation formed about 120 Ma, 

and this age dates the beginning of dynamothermal 

metamorphism. Metamorphic ages for minerals that 

delineate the Sj foliation in amphibolite-facies rocks of 

the Wallowa terrane and Rapid River plate form two 

groups, about 109 and 101 Ma, and indicate either 

that these higher grade rocks cooled more slowly from 

the foliation-forming event that began about 120 Ma 

or that these ages represent successive dynamothermal events. Metamorphic ages that date foliation in 

the North Fork block are complex. Some hornblende 

samples have age spectra that indicate minor argon 

loss; compared to primary cooling ages of about 

100 Ma, reset ages are 89 and 93 Ma, most samples 

having been completely reset to about 90 Ma. Therefore, the pervasively overprinting S4 and L4 fabric in 

rocks of the North Fork block formed during a dynamothermal event about 90 Ma.

The Rapid River thrust fault cuts rocks having metamorphic ages in the range of 118 through 109 Ma, 

and the Slate Creek antiform folds foliation that is of 

the same age. Therefore, the best estimate for the 

maximum age of these structures is 109 Ma. The 

minimum age for the thrust and antiform is approximately 93 Ma, the age of crosscutting tonalitic rocks. 

Thus, the Rapid River thrust fault and the Slate 

Creek antiform formed between 109 and 93 Ma. The 

age of the North Fork reverse fault that cuts both the 

Rapid River thrust fault and the Slate Creek antiform is also a limiting factor in determining relative 

ages of structures.

Because the North Fork reverse fault cuts approximately 93-Ma plutonic rocks and cuts off rocks in the 

North Fork block that have an approximately 90 Ma 

metamorphic overprint the fault probably formed no 

earlier than about 90 Ma. The 90-Ma metamorphic 

fabric in the North Fork block formed during a dynamothermal event that was distinct from earlier 

events and was restricted to a narrow zone along the 

Salmon River suture. It is not known whether the 

steep to vertical tectonic flow that formed this fabric 

was caused by a distinctly, late-stage transpressional 

event along the suture or whether the ductile effects 

constitute an early nonbrittle response of deep-seated 

rock to uplift parallel to the suture.
14. METAMORPHIC AND STRUCTURAL DEVELOPMENT OF ISLAND-ARC ROCKS, WEST-CENTRAL IDAHO 535

DISCUSSION

The general structural geometry of the Slate 

Creek-John Day Creek area is that of stacked thrust 

plates. In each successively higher and more easterly 

plate, the stratigraphy is less well understood because of eastward-increasing complexity of metamorphic overprinting and structure; this gives rise to 

several major, interrelated problems. First, it is unclear how many stratigraphically distinct allochthons 

(terranes) occur in this area; there may be as many 

as three such terranes, or a single terrane may have 

been imbricated. Secondly, the mechanism for the 

eastward-increasing complexity of metamorphism 

and deformation is problematic; causes that have 

been suggested include plate interactions or dynamic 

emplacement of the Idaho batholith.

Questions about the correlation of stratigraphically 

coherent sequences arise from the absence of dates 

on the units, medium- to high-grade metamorphism 

of the rocks, and transposition of bedding. Physical 

continuity between dated rocks along the Snake 

River and rocks along the Salmon River makes it 

possible to correlate rocks of the Wallowa terrane of 

Oregon with exposures in the western part of the 

Slate Creek-John Day Creek area. However, the sedimentology and stratigraphy of the Martin Bridge 

Formation in western Idaho is not well documented 

because of metamorphism and deformation.

The progressive increase in metamorphic grade is 

also a problem in correlation of rocks of the Rapid River 

plate. Greenschist-facies rocks are found near Riggins, 

where these rocks were first described (Hamilton, 

1963a), whereas in the Slate Creek-John Day Creek 

area, rocks thought to be along strike are of the upper 

amphibolite facies (Onasch, 1977; Lund, 1984; McCollough, 1984). Therefore it is difficult to correlate units 

of this plate across structural breaks. Most importantly, because there is little lithologic or chemical difference between units of the Wallowa terrane and Rapid 

River plate (Hamilton, 1963a), it is virtually impossible 

to distinguish them at like metamorphic grades if the 

units cannot be traced along strike.

This problem is even more significant for rocks of 

the North Fork block where even the gross lithologic 

boundaries between units have been transposed and 

intruded. Furthermore, no known differences in composition or previous history conclusively separate 

these rocks from either the Wallowa terrane or the 

Rapid River plate.

Probably only through detailed chemical studies or 

through dating the extrusive age of the volcanic protolith can any differences among these allochthonous 

sequences be conclusively determined. Because all

the rock sequences are compositionally similar and 

the only known differences among them were superimposed during Cretaceous orogenic events, it seems 

most plausible to consider these rocks to be imbricated parts of a single island-arc terrane.

On the basis of successively overprinting fabrics and 

of age data, deformation of the rocks in the Slate CreekJohn Day Creek area apparently occurred in several 

discrete steps. At least two stages of metamorphism are 

documented: Early low-grade metamorphism was associated with dynamic processes that produced a foliation, and the latest high-grade metamorphism was 

associated with processes that formed a pervasive, 

crosscutting foliation in rocks of the North Fork block 

(including Late Cretaceous tonalite of the Idaho batholith). Age data indicate several additional groupings 

among the medium- to high-grade rocks. These groupings of the data may be related to multiple discrete 

thermal pulses and (or) to intermittent nonthermal 

tectonism that periodically exposed hotter rocks from 

deeper crustal levels to cooling.

The earliest plutonic rocks of the Idaho batholith 

intruded both island-arc and continental rocks and 

were involved in the deformation that affected the 

North Fork block. A zone of intense deformation clearly 

crosscuts earlier dynamothermal fabrics and structures. Analyzing this same structural zone east of Riggins and south of the Slate Creek-John Day Creek area, 

Hamilton (1963b, c) interpreted this complex mixture 

of deformed metamorphic and plutonic rocks as a zone 

of vertical "plutonic flow" and called it part of "the 

western border zone of the Idaho batholith." He concluded that deformation and metamorphism in western 

Idaho were the result of westward thrusting of the 

Idaho batholith over the low-grade metavolcanic rocks 

of the Wallowa terrane; that the area of the North Fork 

block was the zone of movement between the batholithic rocks and the island-arc rocks; and that the metamorphism and deformation to the west of the zone were 

caused by the same event, thought to have occurred in 

the middle Cretaceous.

However, evidence that strain was greater toward 

the Salmon River suture is found in the increasingly 

complex dynamothermal overprinting in island-arc 

rocks all along the oceanic-continental boundary from 

north of Harpster to near Riggins as one approaches 

the suture (fig. 14.1; Hamilton, 1963a; Myers, 1968, 

1982; Onasch, 1977). In addition, the metamorphic intensity and amount of strain recorded in continental 

rocks are also greater toward the boundary irrespective of the relative amount of plutonic rock present. 

Except for the zone within 5 km of the suture, plutonic rocks cut earlier dynamothermal fabrics. Therefore, 

metamorphism and deformation in the Slate Creek-
536 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

John Day Creek area predated most plutonic rocks of 

the Idaho batholith, dynamothermal processes were 

localized along the Salmon River suture and were independent of intrusive boundaries, and the earliest 

plutons were deformed in the last stages of the late 

Early to Late Cretaceous regional dynamothermal 

event. This orogenic event that formed the Salmon 

River suture and caused the associated metamorphism and deformation was probably also responsible 

for formation of the Idaho batholith.

Age data show that metamorphic and deformational 

features described herein and in mapping of nearby 

areas (Hamilton, 1963a; Onasch, 1977, 1987; McCollough, 1984) are all Cretaceous. The dynamothermal 

events lasted about 30 Ma and included several major 

stages. In amphibolite-facies rocks, any record of possible earlier (preCretaceous) dynamothermal events 

was destroyed by the >500°C temperatures and high 

strains associated with the late Early to Late Cretaceous deformation. In addition, the area affected by 

any earlier event localized along the east edge of these 

island-arc rocks apparently was no wider than the 

present orogenic belt because the only record of dynamothermal events in the lower-greenschist-facies 

rocks indicates that deformation in them also dates 

from the Cretaceous.

CONCLUSIONS

Even in the lowest grade metamorphic rocks of the 

Wallowa terrane in west-central Idaho, the stratigraphy is not adequately established. Metamorphic complications, structural complexity, and lack of isotopic 

age data all contribute to the problem. In spite of 

generally inadequate stratigraphic data, detailed 

study of the Slate Creek-John Day Creek area has 

provided new information on and insight into the 

stratigraphy of the Wallowa terrane in western Idaho. Although formations have not been delineated for 

all such rocks in the study area, both Permian metavolcanic rocks (either Hunsaker Creek or Windy 

Ridge Formation) and the Triassic Doyle Creek Formation of the Seven Devils Group are in this area 

(Lund and others, 1993). Detailed stratigraphic observations indicate that the upper part of the Seven 

Devils Group is gradational with overlying calcareous slate to phyllite, which is in turn interfingered 

with marble lenses. Stratigraphic data from the 

Slate Creek-John Day Creek area and from the Riggins area to the south show that the Triassic rocks 

of the Wallowa terrane are composed of irregularly 

stacked lithologic lenses and should not be divided 

into lithostratigraphic units and that the Triassic

rocks are remarkably like the sequence of unmetamorphosed rocks in the type section in the Wallowa 

Mountains (Folio, 1992, 1994). Therefore, probably 

all the metasedimentary rocks of the Wallowa terrane in western Idaho are correlative with the Triassic Martin Bridge Limestone (Vallier, 1977; Folio, 

1994). The name "Lucile Slate" is thought to be inappropriate for stratigraphic reasons and its use 

should be discontinued; the name Martin Bridge Formation should be adopted for the Triassic section of 

interfingered calcareous phyllite-schist, biotite-chlorite schist, calc-silicate gneiss, and marble gradationally overlying the Seven Devils Group (Lund and 

others, 1993).

Despite the fact that rocks on the south side of 

John Day Creek are along strike and in physical continuity with units of the Rapid River plate (the Riggins Group as defined by Hamilton, 1963a), lateral 

tracing of the units across structures and metamorphic-facies boundaries has proven to be impossible in 

many places. Where rocks of the Wallowa terrane 

consist of medium-grade metamorphic rocks, the rock 

types are identical to those originally described in 

medium-grade rocks of the Rapid River plate. Likewise, amphibolite-facies rocks of the North Fork 

block are lithologically similar to those of amphibolite-facies Rapid River plate. Rocks of the Rapid 

River plate were originally separated from metavolcanic rocks of the Wallowa terrane because of higher 

metamorphic grade, because distinct map units had 

been recognized in these rocks but not (at that time) 

in the lower grade volcanic rocks, and despite chemical similarity (Hamilton, 1963a). It is now known 

that distinct units exist in the Seven Devils Group of 

the Wallowa terrane (Vallier, 1977) and that, at like 

grades of metamorphism, there are more similarities 

than differences among rocks of the three main 

structural units that are described in this study.

The combination of similar stratigraphy in all three 

plates, characteristically complex original stratigraphy, changes in metamorphic grade, and complexly 

overprinted macroscopic structures makes it difficult 

to separate rocks of the different plates on a regional 

basis. Additionally, because the stacked plates had 

similar histories throughout the metamorphic and deformational events that affected them and because the 

structures that juxtaposed the plates are late Early to 

Late Cretaceous, it is proposed that rocks of the Rapid 

River plate (previously called Riggins Group) and the 

North Fork block might best be considered the leading 

edge of the Wallowa terrane. These rocks may be highgrade, multiply deformed portions of the Wallowa 

terrane that were more strongly affected by dynamothermal processes with their increasing involvement
14. METAMORPHIC AND STRUCTURAL DEVELOPMENT OF ISLAND-ARC ROCKS, WEST-CENTRAL IDAHO 537

in the suturing process and that were ultimately imbricated. Correlation of all three tectonic slices with 

the Wallowa terrane simplifies regional stratigraphic 

problems as well as models for the structural and microplate tectonic histories.

Metamorphic grade increases eastward across the 

Slate Creek-John Day Creek area. The simplest metamorphic fabrics are found in the lowest grade rocks in 

the west part of the area; complex, multiple metamorphic overgrowths and complications from later cycles 

of deformation occur farther east in the study area. 

Onset of prograde metamorphism occurred about 

120 Ma in the study area and as early as 130 Ma in 

other parts of the suture (Davidson, 1990). A separate 

90-Ma event recorded in high-grade metamorphic 

rocks within 5 km of the suture is the youngest dynamothermal stage. Intermediate ages of 109 and 101 Ma 

in medium- to high-grade rocks resulted either from 

separate cycles of metamorphism or from disruption of 

metamorphic patterns by macroscopic structures that 

juxtaposed higher grade rocks over lower grade rocks 

and thus produced the regional inverted metamorphic 

gradient.

Structural complexity also increases eastward. Evidence of early phases of mesoscopic deformation is 

found in lower grade rocks in the west side of the 

area. In eastern parts of the area, intense metamorphism and deformation obliterated most early mesoscopic features. The earliest preserved mesoscopic 

structures formed in the Wallowa terrane during prograde regional metamorphism about 120 Ma. Early 

structures in rocks of the Rapid River plate cannot 

be directly correlated with those in Wallowa terrane 

rocks because early phases are not consistent in 

style, appression, or orientation. However, mesoscopic 

features that formed in the juxtaposed Rapid River 

plate and Wallowa terrane after movement on the 

Rapid River thrust fault are correlative. After formation of the Rapid River thrust fault, both the thrust 

fault and rocks of the Wallowa terrane and Rapid 

River plate were folded together to form the Slate 

Creek antiform. The best limits on the timing of 

Rapid River thrust faulting and formation of the 

Slate Creek antiform range between 109 and 93 Ma. 

The first postmetamorphic plutons were emplaced at 

93 Ma. The last pervasive deformation event at 

90 Ma formed the North Fork reverse fault and a 

steep L- and (or) S-tectonite fabric in plutonic and 

metamorphic rocks of the North Fork block; this 

90 Ma event cut off and (or) overprinted earlier 

structures in the North Fork block.

The best estimate for direction of relative movement of the upper plate along the Rapid River thrust 

fault is west-northwest (Lund, 1984); this is in agreement with that determined on the same fault near 

Riggins (Hamilton, 1963a; Onasch, 1977, 1987). The 

fold axes of the Slate Creek antiform and associated 

mesoscopic folds are perpendicular to the determined 

direction of movement along the Rapid River thrust 

fault. Folds in late reverse faults, which formed after 

the Slate Creek antiform and may signify the lower 

plate's response to deformation in the North Fork 

block, are coaxial with the Slate Creek antiform and 

indicate that movement along the thrust faults was 

westward, generally parallel to movement directions 

determined for the Rapid River thrust fault. Thus, 

regional stresses seem to have maintained a consistent orientation for several discrete events related to 

and following formation of the Rapid River thrust 

fault.

Prograde metamorphic fabric in the island-arc metamorphic rocks of the Slate Creek-John Day Creek 

area developed and was multiply deformed between 

120 and 93 Ma prior to intrusion of earliest plutons 

of the Idaho batholith. These plutons intruded both 

island-arc and continental rocks after 93 Ma and 

thereby sealed the join between the North American 

continent and the accreted island-arc rocks. Although 

metamorphic grade increases eastward toward the 

main mass of plutonic rocks, pluton contacts regionally crosscut metamorphic patterns. More importantly, in continental metamorphic rocks that form roof 

pendants in the main mass of the Idaho batholith, 

metamorphic grade also increases westward into the 

suture zone; metamorphic patterns and structures in 

continental metamorphic rocks also are cut by plutonic rocks (Lund, 1984). The age information confirms this time gap between metamorphism and 

plutonism. The only deformed Idaho batholith plutonic rocks are the 93 to 85 Ma tonalite suite, where 

emplaced along the suture. Therefore instead of causing the regional dynamothermal events, the Idaho 

batholith probably formed as a result of the same 

processes that caused the Salmon River suture.

On basis of this work on island-arc rocks of the 

Slate Creek-John Day Creek area and of work on 

continental rocks in the upper Slate Creek drainage 

and to the east, metamorphism and deformation of 

rocks in the vicinity of the Salmon River suture apparently occurred during formation of and movement 

along the accretionary boundary. Age data indicate 

that the metamorphism and deformation occurred 

from 120 to 88 Ma. Geochemical data (Armstrong 

and others, 1977; Fleck and Criss, 1985; Hoover, 

1986), geophysical data (Bankey, 1992; McCafferty, 

1992), and geological constraints (Lund, 1984; 

Hoover, 1986; Manduca, 1988) indicate that the suture is steep to vertical. The structural geometry fits
538 PETROLOGY AND TECTONIC EVOLUTION OF PRE-TERTIARY ROCKS, BLUE MOUNTAINS REGION

best with a model of transpressive (oblique-transcurrent) movement (Lund, 1984). The contact between 

island-arc and continental rocks is obscured by postaccretionary plutons, and the evidence for timing of 

motion between them indicates that the motion was 

late-stage and postplutonic (after 93 Ma). Additionally, and as expected from models (Wilcox and others, 

1973; Harding, 1976; Sylvester and Smith, 1976; 

Odonne and Vialon, 1983; Sylvester, 1988), shear indicators show the rock underwent dominantly pure 

shear resulting in vertical motion up and out of the 

suture. However, during the part of the Cretaceous 

when transpressive movement was occurring along 

the Salmon River suture, displacement of the Farallon plate with respect to North America had a major 

right-lateral component (Engebretson and others, 

1985; Mammerickx and others, 1988) and major contemporaneous transcurrent faults in Washington and 

Canada were also right lateral (Monger, 1984; Gabrielse, 1985). Thus, metamorphism and deformation in 

the Slate Creek-John Day Creek area as well as 

along both sides of the rest of the Salmon River suture occurred during the interval 120 to 88 Ma and 

was probably due to right-lateral transpressive motion that juxtaposed island-arc rocks against continental rocks.

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U.S. GOVERNMENT PRINTING OFFICE: 1995 - 673-046 / 20010 REGION NO. 8
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