Geology of the Blue Mountains Region 

of Oregon, Idaho, and Washington: 

The Idaho Batholith and 

Its Border Zone 

U.S. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1436 

Geology of the Blue Mountains Region 

of Oregon, Idaho, and Washington: 

The Idaho Batholith and 

Its Border Zone 

TRACY L. VALLIER and HOWARD C. BROOKS, editors 

U.S. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1436 

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON 1987 
DEPARTMENT OF THE INTERIOR 

DONALD PAUL HODEL, Secretary 

U.S. GEOLOGICAL SURVEY 

Dallas L. Peck, Director 

Library of Congress Cataloging-in-Publication Data 

The Idaho batholith and its border zone. 

(U.S. Geological Survey Professional Paper 1436. Geology of the Blue Mountains region of Oregon, Idaho, and Washington) 

Includes bibliographies. 

Supt. of Docs. No.: I 19.16:1436 

1. Idaho batholith (Idaho and Mont.) I. Vallier, T. L. (Tracy L.) II. Brooks, Howard C. III. Series: U.S. Geological Survey 

Professional Paper 1436. IV. Series: Geology of the Blue Mountains region of Oregon, Idaho, and Washington. 

QE611.5.U6133 1987 552' .3'09796 86-600206 

For sale by the Books and Open-File Reports Section, 

U.S. Geological Survey, Federal Center, Box 25425, Denver, CO 80225 
PREFACE 

This U.S. Geological Survey Professional Paper is one 

volume of a series that focuses on the geology, paleontology, and mineral resources of eastern Oregon, western 

Idaho, and southeastern Washington. The purpose of this 

series is to familiarize readers with the 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. 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 intruded 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 

forearc basins, whereas in others (Tonga-Kermadec) the 

forearc insular slopes descend precipitously to trench 

depths and are only interrupted in places by narrow faultbounded terraces. Moreover, some arcs have erupted 

primarily tholeiitic igneous products throughout their 

histories (Tonga-Kermadec) and others (Aleutian) have a 

long history of both calc-alkaline and tholeiitic eruptive 

activity. Ridge axes of the 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 may have different origins and histories of development. Some 

(Mariana Trough and Lau Basin) have active spreading 

ridges whereas others (Aleutian basin of the Bering Sea) 

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 analogues. 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 province. 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 the 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 about 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. 

III 

CONTENTS 

[Numbers designate the chapters] 

Page 

1. The Idaho batholith and its border zone: a regional perspective 

Tracy L. Vallier and Howard C. Brooks - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - 1 

2. Petrology and origin of the Bitterroot lobe of the Idaho batholith 

Margo I. Toth - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - 9 

3. Structural geology and petrology of a part of the Bitterroot lobe of the Idaho batholith 

Rolland R. Reid - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - 37 

4. Emplacement of the main plutons of the Bitterroot lobe of the Idaho batholith 

C. Gil Wiswall and Donald W. Hyndman ----------------------------- -- 59 

5. Migmatite zones in the Bitterroot lobe of the Idaho batholith 

Enid Bittner- - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - 73 

6. Petrogenesis, geochronology, and hydrothermal systems of the northern Idaho batholith and adjacent areas 

based on 18QJ16Q, D/H, 87Srf86Sr, K-Ar, and 40 Arf39 Ar studies 

Robert E. Criss and Robert J. Fleck - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - 95 

7. Temporal and spatial relations between folding, intrusion, metamorphism, and thrust faulting in the Riggins 

area, west-central Idaho 

Charles M. Onasch - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - 139 

8. Pre-Cenozoic geology of the West Mountain-Council Mountain-New Meadows area, west-central Idaho 

Bill Bonnichsen - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - -151 

9. Lithologic and chemical characteristics of the central and southeastern part of the southern lobe of the Idaho 

batholith 

Reed S. Lewis, Thor H. Kiilsgaard, Earl H. Bennett, and Wayne E. Hall - - - - - - - - - - - - - - 171 

v 

1. THE IDAHO BATHOLITH AND ITS BORDER ZONE: 

A REGIONAL PERSPECTIVE 

By TRACY L. VALLIER and HowARD C. BROOKS1 

ABSTRACT 

The geology of the Idaho batholith and its border zone is interpreted 

by papers in this volume. The geology of the northern (Bitterroot) lobe 

of the batholith is well described in chapters by Toth, Reid, Wiswall and 

Hyndman, and Bittner, who discuss aspects of the petrology, structural 

geology, emplacement history, and zones of migmatization, respectively. 

Criss and Fleck concentrate on chemical and isotopic evidence for the 

origin of the batholith and some of its hydrothermal systems. Onasch 

discusses the structural geology of the western border zone near Riggins; 

Bonnichsen describes the geology of the border zone farther south near 

New Meadows. Lewis and others concentrate their attention on ages, 

chemical compositions, emplacement history, and source regions of 

plutons in the southern (Atlanta) lobe of the batholith. Ages of intrusions in the batholith are well constrained by radiometric dating; 

excluding the complex western border zone, where some plutons from 

the accreted Blue Mountains island arc terrane may be structurally involved, most plutons composing the batholith crystallized between 95 

and 45 Ma. Calc-alkaline volcanism began in central Oregon about 55 Ma, 

at approximately the same time that the last phase of plutonism began 

in the Idaho batholith. Major phases of Late Cretaceous and early 

Tertiary magmatism can be correlated with directions and rates of 

Farallon and Pacific plate motions. 

INTRODUCTION 

In this paper we briefly summarize the major findings 

presented in eight other chapters of this volume and 

subsequently propose a speculative Late Cretaceous and 

early Tertiary geologic history of the Oregon-Idaho region 

that places the Idaho batholith into a time and space setting and that relates its formation in part to movements 

of oceanic plates in the Pacific basin. These speculations 

are not new (for example, Coney, 1978), but on the basis 

of recently published papers of Pacific plate movements 

and the new radiometric ages and interpretations 

presented by authors of chapters in this volume, we can 

offer a more precise correlation scheme. 

10regon Department of Geology and Mineral Industries, Baker, Oregon 97814. 

A review of the literature on the Idaho batholith region 

is beyond the scope of this paper. Important early papers 

on the batholith and its border zone are by Lindgren 

(1904), Ross (1928, 1934, 1952, 1965), Langton (1935), 

Anderson (1942, 1952), Larson and Schmidt (1958), Reid 

(1959), Hamilton (1963), and Hietanen (1963). Significant 

later papers, among many others, are by Chase (1973), 

Greenwood and Morrison (1973), Swanberg and Blackwell 

(1973), Armstrong and others (1977), Chase and Johnson 

(1977), Scholten and Onasch (1977), Hyndman (1980, 

1983, 1984), Criss and others (1982), and Criss and Taylor 

(1983). Reid (chapter 3) reviews several important papers 

in the context of describing the geology of the northern 

part of the batholith. Our major interpretations of the 

geology of the Blue Mountains region west of the Idaho 

batholith can be found in Vallier (1977), Vallier and others 

(1977), and Brooks and Vallier (1978). 

We appreciate the critical reviews of this paper by R.E. 

Criss, R.J. Fleck, R.R. Reid, Bill Bonnichsen, C.G. 

Wiswall, D.W. Hyndman, C.M. Onasch, and D.C. 

Engebretson. Their comments greatly improved both the 

content and style. 

MAJOR CONCLUSIONS OF PAPERS 

IN THIS VOLUME 

This section outlines the major conclusions from papers 

in the volume. The approximate area of study for each 

chapter is shown in figure 1.1, and the presentation order 

follows the table of contents. 

Toth (Chapter 2: Petrology and origin of the Bitterroot 

lobe) presents mineralogic, chemical, and field data on the 

Bitterroot lobe of the batholith. The Bitterroot lobe consists of foliated tonalite and quartz diorite I-type plutons 

of Late Cretaceous age that were intruded by voluminous 

Paleocene monzogranite and granodiorite S-type plutons 

that form the bulk of the lobe. Subsequently, Eocene 

1 
2 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

WASHINGTON 

---·- ---·-·-- --

OREGON 

• P. 

• L.G. 

20 40 MILES 

20 40 KILOMETERS 

120 118 

Z"'MONTANA 

1,..r-) BITTNER -~----f"~) 

LEWIS and OTHERS 

115 

< c: 

....... ) 

I 

46 

FIGURE 1.1.-Index map showing areas discussed by authors in this volume. Each area is outlined and authors are identified by name within 

outlined areas. Area discussed by Lewis and others extends below figure a few miles to about latitude 43° N. Abbreviations, approximately 

from left to right, are M., Mitchell; J.D., John Day; P., Pendelton; L.G., La Grande; B., Baker; H., Huntington; 0., Oxbow; E., Enterprise; 

L., Lewiston; G., Grangeville; R., Riggins; and M.C., McCall. 

syenodiorite and quartz syenite A-type high-level plutons 

intruded the older rocks. The rocks were further injected 

by several sets of dikes. The oldest I -type plutons show 

some chemical variations in the western bodies which apparently reflect the assimilation of some Precambrian 

metasedimentary material. The 8-type granitic rocks that 

form the bulk of the lobe were probably derived from partial melting of metasedimentary material. The A-type 

syenitic rocks may have been generated as a consequence 

of regional metamorphism and anatexsis of continental 

crust. Toth further concludes that most of the oldest 

plutons were emplaced between 100 and 80 Ma, the 

voluminous, middle-stage plutons cooled at about 66 to 

57 Ma, and the youngest plutons formed 51 to 46 Ma. 

Reid (Chapter 3: Structural geology and petrology of 

a part of the Bitterroot lobe) discusses the influence of 

structure on the emplacement of the north-central part 

of the Bitterroot lobe. The granitic massif, which consists 

of tonalite and younger granite and granodiorite, is zoned 

vertically from calcic granodiorite at the base to granite 

at the top. The emplacement of the pluton was synkinematic into an active thrust zone that resulted in a 

metamorphic core complex. Three phases of deformation 

were followed by arching. Uplift of as much as 30 km is 

implied by the data. 

Wiswall and Hyndman (Chapter 4: Emplacement of the 

main plutons of the Bitterroot lobe) compare mineralogic, 

chemical, and field relations to show that the magmas may 

have been emplaced mostly by lateral spreading, rather 

than diapiric rise, which formed a mushroom-shaped body 

at mesozonal to catazonal depths. Evidence suggests that 

the northern part of the Bitterroot lobe had a somewhat 

shallower level of emplacement than the southern part. 

Different styles of emplacement also are indicated; for 

example, in the north forcible injection apparently 

occurred along foliation near the batholith's margins, 
1. THE IDAHO BATHOLITH AND ITS BORDER ZONE: A REGIONAL PERSPECTIVE 3 

whereas in the south the more ductile country rock flowed 

around the laterally spreading magmas. In spite of these 

differences in style of deformation, the country rocks and 

sequence of deformation are similar throughout the 

Bitterroot lobe of the batholith. 

Bittner (Chapter 5: Migmatite zones in the Bitterroot 

lobe) describes selected migmatite screens within the 

Bitterroot lobe. She reports that they are oriented largely east-west and presents evidence that the migmatites 

were derived from the wall and (or) roof rocks. Magma 

flow from west to east apparently caused the east-west 

orientation. The migmatite protolith was injected by 

magmas of several types and, in places, was intricately 

folded. Five deformational events are recorded. 

Criss and Fleck (Chapter 6: Petrogenesis, geochronology, and hydrothermal systems of the northern Idaho 

batholith area) describe the geochemical character of the 

northern Idaho batholith and adjacent areas to the west 

using strontium and oxygen isotopic data, K-Ar ages, 

Rb/Sr and D/H ratios, and major-element chemical data. 

In addition, they discuss Late Cretaceous and early 

Tertiary hydrothermal systems that created many of the 

region's ore deposits. Older plutons in the western 

accreted island arc terrane have some chemical characteristics that are compatible with derivation from mantle 

materials, whereas the main mass of the batholith formed 

from magmas that probably were derived primarily from 

Precambrian lower crust rocks and modified somewhat 

by assimilation of Precambrian wall rocks. The Eocene 

plutons also have Sr-isotopic values that suggest the 

presence of a Precambrian rock component in the melts. 

K-Ar ages of hornblende and biotite range from> 125 Ma 

for typical plutons in the accreted terrane to about 45 Ma 

for the Eocene granitic rocks. Apparent ages decrease, 

in general, from west to east across central Idaho. The 

oldest plutons, discounting those associated with the 

accreted terrane, yield K-Ar ages of 87 to 75 Ma and are 

concentrated along the boundary between the island arc 

rocks of the accreted terrane and the western margin of 

the Precambrian sialic crust. The younger zone of concordant ages of plutons (51 to 44 Ma) is along the IdahoMontana border within the Bitterroot lobe. The age 

pattern in, and adjacent to, the suture zone between the 

accreted terrane and the sialic crust to the east restricts 

the age of accretion to the period between about 135 and 

75 Ma. Several lines of evidence indicate strong uplift of 

the batholith during the Late Cretaceous and early 

Tertiary. 

Onasch (Chapter 7: Temporal and spatial relations 

between folding, intrusion, metamorphism, and thrust 

faulting in the Riggins area, west-central Idaho) studied 

the western border zone that includes the suture between 

the western accreted terrane of ensimatic rocks and the 

ensialic North American continent. It is an extremely 

complex region, not only because of multiple stages of 

metamorphism and thrusting, but also because of the 

original complicated stratigraphy of the ensimatic rocks 

of the accreted terrane. Several fold phases were mapped 

which change along strike of the contact. The time of 

thrust faulting, however, seems to be nearly synchronous 

along the contact region. Intrusions of a probable early 

trondhjemitic phase of the Idaho batholith accompanied 

the peak of metamorphism. Relative to folding and thrust 

faulting, the locus of trondhjemitic activity and metamorphism shifted progressively northward through the succession of fold phases. He agrees with Hamilton (1963) 

that the Rapid River thrust may mark the overlap of two 

orogens. Faulting apparently took place in the Late 

Cretaceous. 

Bonnichsen (Chapter 8: Pre-Cenozoic geology of the 

West Mountain-Council Mountain-New Meadows area, 

west-central Idaho) takes the reader farther south along 

the contact zone to where parts of the Atlanta lobe of the 

batholith and the western accreted terrane are juxtaposed. He discusses the area in light of current ideas on 

both the Blue Mountains region (accreted western 

terrane) and the Idaho batholith. Ultramafic rocks and 

quartz-rich sedimentary rocks (chert protolith?) of the 

western sequence are in contact with metamorphosed 

rocks of the complicated border zone along a major 

tectonic boundary. Plutons in the border zone have K-Ar 

apparent ages of 86 to 80 Ma (biotite) and 120 to 92 Ma 

(hornblende). One of the plutons (No Business Mountain 

pluton) is unique to the region because of its alkalic composition; Bonnichsen speculates that it may be part of the 

continental-margin group of early alkalic plutons that 

have been described from exposures farther south in 

California. The metamorphosed rocks of the western 

terrane probably are part of the ancestral Pacific basin 

rather than the Precambrian North American continent. 

The plutons, then, represent early intrusive phases of the 

Idaho batholith. 

Lewis and others (Chapter 9: Lithologic and chemical 

characteristics of the central and southeastern part of the 

southern lobe) identify several plutonic phases in the 

southern (Atlanta) lobe, determine their ages and chemical 

compositions, and speculate about the depths of emplacement and source regions. In addition, they compare the 

Bitterroot and Atlanta lobes of the Idaho batholith and 

the Idaho and Sierra Nevada batholiths. The Atlanta lobe 

is predominantly biotite granodiorite and muscovitebiotite granite with subordinate amounts of tonalite, 

hornblende-biotite granodiorite, K-feldspar megacryst 

granodiorite, and leucocratic granite. The oldest plutons, 

lying along the western side of the Atlanta lobe, are 

tonalite with emplacement ages of 95 to 85 Ma. The 

youngest plutons are the 75- to 70-Ma leucocratic granite 

plutons. The tonalite plutons yield low radiogenic Sr ratios 
4 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

and probably formed above a subducting slab, whereas 

younger plutons have high ratios, probably reflecting partial melting of the lower continental crust with only a 

minor mantle contribution. The Atlanta and Bitterroot 

lobes are very similar chemically and mineralogically, 

although the Atlanta lobe contains no significant foliation 

and lineation and has no metamorphosed basaltic dikes. 

The Idaho batholith differs from the Sierra Nevada 

batholith by having a lower mafic mineral content and by 

lacking the numerous zoned plutons in sharp contact with 

one another that is characteristic of the Sierra Nevada 

batholith. Most differences between the two batholiths can 

be accounted for by ascribing to differences in water contents of the magmas and in the nature of the source 

regions. 

LATE CRETACEOUS AND 

EARLY TERTIARY HISTORY 

Using mostly the data and interpretations of Engebretson (1982), Vallier and Engebretson (1984), Engebretson 

and others (1984), Robinson and others (1984), Walker and 

Robinson (in press), Criss and Fleck (chapter 6), and Lewis 

and others (chapter 9), we can speculate on major motions 

of plates in the Pacific and their effects on plutonism and 

volcanism in the Oregon-Idaho region during the Cretaceous and early Tertiary. A thorough discussion is beyond 

the scope of this paper, but the correlation of events is 

surprisingly good and some obvious relations can be 

pointed out. Several attempts have been made to correlate 

plate movements with the Laramide orogeny and with 

magmatism in the Pacific Northwest (for example, Coney, 

1978; Ewing, 1980; Engebretson and others, 1984; 

Henderson and others, 1984; Jurdy, 1984; among many 

others), and it is apparent that there are close correlations. No attempt has been made, however, to speculate 

on the relations in the narrow Oregon-Idaho-Montana belt 

that is described in part by the papers in this volume. 

Vallier and Engebretson (1984) stated that the upper 

Paleozoic through Lower Cretaceous rocks of the Blue 

Mountains probably were parts of a single allochthonous 

intra-oceanic island arc composed of several "terranes" 

and that these rocks have been coherent since at least the 

Middle Triassic and perhaps since the Early Permian. The 

island arc may have migrated either from the southwestern Pacific (if it had been part of the Farallon plate) 

or from the equatorial zone of the eastern Pacific (if 

attached to the Kula plate). Accretion probably occurred 

throughout the Early Cretaceous and culminated before 

80 Ma. The early part of the arc's history has been discussed by many authors, but in the present paper we 

discuss only the period between about 135 and 21 Ma. 

Engebretson (1982) and Engebretson and others (1984) 

dealt with all known plates in the ancestral Pacific (for 

example, Farallon, Kula, Izanage, and Phoenix plates) 

during that time frame, but we emphasize the interactions 

of the Farallon, Kula, and Pacific plates as they may relate 

to plutonism and volcanism in the Oregon-Idaho region 

during the Late Cretaceous and early Tertiary. 

The primary ages of plate reorganization are about 135, 

85, 56, and 43 Ma (table 1.1). At these times, there were 

not only changes in directions but also changes in 

velocities (half-rates) of plate movements. The correlation 

between plate movements (table 1.1) and plutonic-volcanic 

sequences in the Idaho-Oregon region (table 1.2) is surprisingly good (table 1.3). 

Engebretson's (1982) conclusions include the following 

major points. (1) Changes in relative plate motion in the 

ancestral Pacific basin occurred at about 135, 85, and 56 

Ma. (2) Spreading rates along the Pacific-Kula ridge 

uniformly decreased between 72 and 56 Ma. (3) Soon after 

56 Ma, a major reorganization in the Pacific-Farallon and 

Pacific-Kula ridge systems occurred, accompanied by an 

abrupt increase in Pacific-Kula ridge spreading. The 

reorganization apparently happened during 56 to 48 Ma. 

(4) At about 43 Ma, Pacific-Kula ridge spreading ceased 

and the Kula-Farallon ridge evolved into alignment with 

the Pacific-Farallon ridge, which increased in spreading 

rate. (5) The Pacific-Farallon ridge relative motion, 

however, remained nearly fixed from 48 to 28 Ma. 

TABLE 1.1.-Changes in plate motions within the Pacific basin 

[Data are from Engebretson (1982) and Engebretson and others (1984)] 

Time of major 

plate change 

(m.y.B.P.) 

135------

85------

56------

Pacific-Farallon 

spreading rate 

(half-rate, in 

degrees per 

million years) 

Increase from 0.31 

to 0.45. 

Decrease from 0.43 

to 0.36. 

Increase from 0.36 

to 0.67. 

Growth or demise of 

plates and other 

changes in plate motion 

Pacific-Izanagi ridge 

also increases in 

half-rate spreading, 

but from 0.43°/m.y. 

to 1.17°/m.y. 

Pacific-Kula ridge 

half-rate spreading 

decreases from 

0.49°/m.y. to 

0.16°/m.y. 

Pacific-Kula ridge 

increases in halfrate spreading from 

0.16°/m.y. to 

0.65°/m.y. Major 

reorganization in 

Pacific, Kula, and 

Farallon plates. 

43------ Increase from 0.74 Demise of the Pacificto 0.99. Kula ridge. 

48-28------ Pacific-Farallon ridge 

relative motion 

fixed. 
1. THE IDAHO BATHOLITH AND ITS BORDER ZONE: A REGIONAL PERSPECTIVE 5 

TABLE 1.2.-Major plutonic and volcanic events in Oregon and Idaho TABLE 1.3.-Possible correlation of plate movements, uplift, and 

magmatic events during the period 135 to 21 Ma 

Age (Ma) 

135-125-----

120-85------

85-70-------

65-62-------

55-40-------

51-46-------

37-21-------

Event 

Demise of plutonism in the Blue Mountains 

is land arc. 

Early phase of Idaho batholith intruded along 

the present western border zone of the 

batholith (and old North American 

continent). 

Intrusion of most plutons of the Idaho 

batholith not associated with the western 

border zone. 

Possible pulse of Idaho batholith plutonism 

(Criss and Fleck, chapter 6) that is not 

reported in the Atlanta lobe (Lewis and 

others, chapter 9). 

Eruption of calc-alkaline volcanic rocks of 

the Clarno Formation. 

Emplacement of granitic plutons in the Idaho 

batholith; emplacement of some shallowlevel alkalic plutons; eruption of Challis 

Volcanics. 

Eruption of rhyolitic tuffs of the John Day 

Formation in Oregon, marking early 

eruptions of Cascade Range volcanoes. 

Criss and Fleck (chapter 6) report three major ages of 

plutonism in the Idaho batholith: 82-75, 65-62, and 51-46 

Ma, with averages (K-Ar apparent ages on hornblende) 

of about 80, 65, and 50 Ma. Lewis and others (chapter 

9) provide data on pluton ages of 95 to 70 Ma. Bonnichsen 

(chapter 8) suggests that there may be an earlier phase 

(120-92 Ma) of Idaho batholith plutonism along the 

southern border. Robinson and others (1984) and Walker 

and Robinson (in press) discuss the Tertiary volcanism in 

Oregon and extreme western Idaho. Periods of major 

volcanic eruptions occurred between 55 and 40 Ma, when 

calc-alkaline volcanic rocks of the Clarno Formation and 

correlative strata were deposited from local vents within 

the Blue Mountains, and between 37 and 21 Ma, when 

the John Day Formation was deposited from vents near 

the present axis of the Cascade Range. Ewing (1980, p. 

622) compiled data that indicate a robust pulse of 

volcanism during 53 to 42 Main the Pacific Northwest. 

The K-Ar apparent ages from the western transition 

(border) zone of the batholith show a broad range, probably because of multiple heating events associated with 

the Idaho batholith intrusions and because older plutons 

were structurally involved in the border zone before 

intrusion of a large part of the batholith. For example, 

U-Pb dating of separated zircons from a tonalite pluton 

within the border zone north of Riggins yielded an Early 

Permian age (N.W. Walker, written commun., 1983), 

which is the same age as several metamorphosed plutons 

within the Blue Mountains (Walker, 1983). In addition, 

Jurassic and Cretaceous (about 165-130 Ma based on 

Age (Ma) Event 

135-85---- Accretion of Blue Mountains island arc to the 

North American continent during increase in 

spreading rates. Most accretion probably 

occurred in the late Early and early Late 

Cretaceous (120-95 Ma). Small plutons emplaced 

along the present western border zone of the 

Idaho batholith during 120-85 Ma (Bonnichsen, 

chapter 8; Toth, chapter 2; Lewis and others, 

chapter 9). Early thrusting in Idaho probably 

associated with accretion of the Blue Mountains 

island arc terrane. 

85-56----- Decrease in spreading rates along both the 

Pacific-Farallon and Pacific-Kula ridges; a 

possible low-angle subduction zone beneath the 

continent may have created the major plutonic 

phases of the Idaho batholith between 85 and 70 

Ma with a possible pulse at 65 to 62 Ma (Criss 

and Fleck, chapter 6). Emplacement of the 

Boulder batholith between 78 and 68 Ma (D.W. 

Hyndman, written commun., 1984). Successively 

younger parts of the Farallon plate were 

subducted beneath North America. Gradual 

uplift of the southern Idaho batholith. Twomica granites ("-'72 Ma) cut major overthrusts in 

eastern Idaho (E.H. Bennett, oral commun, 

1986). 

56-43----- Intrusion of the youngest plutons of the 

batholith (51-46 Ma). Volcanic eruptions 

responsible for the Clarno Formation and 

Challis Volcanics. Major volcanism occurred 

throughout the Pacific Northwest (Ewing, 1980). 

Increase in spreading rates along both the 

Pacific-Kula and Pacific-Farallon ridges; major 

reorganization of plates leading to chaotic 

distribution of eruptive products during 

probable changes of subduction angles and 

directions. Continued and greatly accelerated 

uplift of the Idaho batholith; some faulting in 

the overthrust belt of Idaho, Wyoming, and 

Montana. 

43-37----- Reorganization of plates. Increase in spreading 

rates and the probable demise of the PacificKula spreading ridge. Change in Pacific plate 

motion. Decrease in volcanic eruptions and the 

end of batholith plutonism. 

37-21----- Calc-alkaline volcanism responsible for the John 

Day Formation from eruptive centers near the 

present axis of the Cascade Range; unnamed 

calc-alkaline rocks widespread in central and 

western parts of the Blue Mountains. Probable 

westward jump of subduction zone and (or) 

steepening of the subduction angle; fixed 

Pacific-Farallon relative motion. 

K-Ar ages) plutons also have been identified within the 

border zone (for example, Armstrong and others, 1977). 

These ages suggest that the older plutons probably were 

structurally involved during formation of the border-or 

transitional-zone; thus, the K-Ar ages from this area, 

which previously had been related to the Idaho batholith 

per se, should be suspect. We believe that very few, if any, 

plutons of the Idaho batholith crystallized before 100 Ma. 
6 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

It is tempting to relate not only the ages of major 

changes in plate motions to ages of magmatism, but also 

to faulting events in the overthrust belt of eastern Idaho. 

On the basis of our limited data base, it would be somewhat presumptuous to conclude that the overthrusting in 

eastern Idaho was related directly to plate movements and 

the intrusion and uplift of the Idaho batholith, but the 

correlation-at least locally-cannot be disregarded. In 

fact, as pointed out by Hyndman (1980) there may be some 

cause and effect relations in the Montana-Idaho-Wyoming 

region of the overthrust belt. The dating of overthrusting, 

however, is difficult. Hoffman and others (1976) dated 

K -bentonites from western Montana that they believed 

had been metamorphosed by the additional load of the 

overthrust sheets and reported ages of 72 to 56 Ma for 

the metamorphism. This area was small, and no doubt the 

ages of thrusting vary along the trend of the overthrust 

belt. Most thrusting along the eastern edge of the 

batholith had been completed prior to the intrusion of the 

80- to 70-Ma plutons (E.H. Bennett, oral commun., 1986). 

This means that the thrusting may be related to compressive forces that were exerted during the final accretion of the Blue Mountains island arc terrane to North 

America. We suspect that the accelerated uplift of the 

batholith (table 1.3), particularly during the period 56 to 

43 Ma, may have been responsible for some late-stage 

overthrusting in the region. 

REFERENCES CITED 

Anderson, A.L., 1942, Endomorphism of the Idaho batholith: Geological 

Society of America Bulletin, v. 53, p. 1099-1126. 

__ 1952, Multiple emplacement of the Idaho batholith: Journal of 

Geology, v. 60, p. 255-265. 

Armstrong, R.L., Taubeneck, W.H., and Hales, P.O., 1977, Rb-Sr and 

K-Ar geochronometry of Mesozoic granitic rocks and Sr isotopic composition, Oregon, Washington, and Idaho: Geological Society of 

America Bulletin, v. 88, p. 397-411. 

Brooks, H.C., and Vallier, T.L., 1978, Mesozoic rocks and tectonic evolution of eastern Oregon and western Idaho, in Howell, D.G., and 

MacDougall, K.A., eds., Mesozoic paleogeography of the Western 

United States: Pacific Coast Paleogeography Symposium, 2d, 

Society of Economic Paleontologists and Mineralogists, p. 133-146. 

Chase, R.B., 1973, Petrology of the northeastern border zone of the 

Idaho batholith, Bitterroot Range, Montana: Montana Bureau of 

Mines and Geology Memoir 43, 28 p. 

Chase, R.B., and Johnson, B.R., 1977, Border-zone relationships of the 

northern Idaho batholith: Northwest Geology, v. 6-1, p. 38-50. 

Coney, P.J., 1978, Mesozoic-Cenozoic cordilleran plate tectonics: 

Geological Society of America Memoir 152, p. 33-50. 

Criss, R.E., Lanphere, M.A., and Taylor, H.P., Jr., 1982, Effects of 

regional uplift, deformation, and meteoric-hydrothermal metamorphism on K-Ar ages of biotites in the southern half of the Idaho 

batholith: Journal of Geophysical Research, v. 87, p. 7029-7046. 

Criss, R.E., and Taylor, H.P., Jr., 1983, An 180/160 and D/H study of 

Tertiary hydrothermal systems in the southern half of the Idaho 

batholith: Geological Society of America Bulletin, v. 94, p. 640-663. 

Engebretson, D.C., 1982, Relative motions between oceanic and continental plates in the Pacific basin: Stanford, Calif., Stanford University, Ph.D. dissertation, 211 p. 

Engebretson, D.C., Cox, Allan, and Thompson, G.A., 1984, Correlation 

of plate motions with continental tectonics: Laramide to Basin and 

Range: Tectonics, v. 3, p. 115-119. 

Ewing, T.E., 1980, Paleogene tectonic evolution of the Pacific Northwest: Journal of Geology, v. 88, p. 619-638. 

Greenwood, W.R., and Morrison, D.A., 1973, Reconnaissance geology 

of the Selway-Bitterroot Wilderness Area: Idaho Bureau of Mines 

and Geology Pamphlet 1543, 30 p. 

Hamilton, Warren, 1963, Metamorphism in the Riggins quadrangle, 

western Idaho: U.S. Geological Survey Professional Paper 436, 95 p. 

Henderson, L.J., Gordon, R.G., and Engebretson, D.C., 1984, Mesozoic aseismic ridges on the Farallon plate and southward migration 

of shallow subduction during the Laramide orogeny: Tectonics, v. 3, 

p. 121-132. 

Hietanen, Anna, 1963, Idaho batholith near Pierce and Bungalow, Clearwater County, Idaho: U.S. Geological Survey Professional Paper 

344-D, 42 p. 

Hoffman, Janet, Hower, John, and Aronson, J.L., 1976, Radiometric 

dating of time of thrusting in the disturbed belt of Montana: Geology, 

v. 4, p. 16-20. 

Hyndman, D.W., 1980, Bitterroot dome-Sapphire tectonic block, an 

example of a plutonic-core gneiss-dome complex with its detached 

suprastructure, in Coney, P.J., Crittenden, M.D., Jr., and Davis, 

G.H., eds., Cordilleran metamorphic core complexes: Geological 

Society of America Memoir 153, p. 427-443. 

__ 1983, The Idaho batholith and associated plutons, Idaho and 

western Montana, in Roddick, J., ed., Circum-Pacific plutonic terranes: Geological Society of America Memoir 159, p. 213-240. 

__ 1984, A petrographic and chemical section through the northern 

Idaho batholith: Journal of Geology, v. 92, p. 83-102. 

Jurdy, D.M., 1984, The subduction of the Farallon plate beneath North 

America as derived from relative motions: Tectonics, v. 3, p. 107-113. 

Langton, C.M., 1935, Geology of the northeast part of the Idaho batholith 

and adjacent region in Montana: Journal of Geology, v. 43, p. 27-60. 

Larsen, E.S., and Schmidt, R.G., 1958, A reconnaissance of the Idaho 

batholith and comparison with the southern California batholith: U.S. 

Geological Survey Bulletin 1070-A, p. 1-33. 

Lindgren, W., 1904, A geologic reconnaissance across the Bitterroot 

Range and adjacent region in Montana: Journal of Geology, v. 43, 

p. 27-60. 

Reid, R.R., 1959, Reconnaissance geology of the Elk City region, Idaho: 

Idaho Bureau of Mines and Geology Pamphlet 120, 74 p. 

Robinson, P.T., Brem, G.F., and McKee, E.H., 1984, John Day Formation of Oregon: a distal record of early Cascade volcanism: Geology, 

v. 12, p. 229-232. 

Ross, C.P., 1928, Mesozoic and Tertiary granitic rocks in Idaho: Journal of Geology, v. 36, p. 673-693. 

__ 1934, Geology and ore deposits of the Casto quadrangle, Idaho: 

U.S. Geological Survey Bulletin 854, 135 p. 

__ 1952, The eastern front of the Bitterroot Range, Montana: U.S. 

Geological Survey Bulletin 97 4-E, p. 135-175. 

__ 1965, The Idaho batholith: Geological Society of America Special 

Paper 82, 343 p. 

Scholten, R., and Onasch, C.M., 1977, Genetic relations between the 

Idaho batholith and its deformed eastern and western margins: 

Northwest Geology, v. 6-1, p. 25-37. 

Swanberg, C.A., and Blackwell, D.D., 1973, Areal distribution and 

geophysical significance of heat generation in the Idaho batholith 

and adjacent intrusions in eastern Oregon and western Montana: 

Geological Society of America Bulletin, v. 84, p. 1261-1282. 

Vallier, T.L., 1977, The Permian and Triassic Seven Devils Group, 

western Idaho and northeastern Oregon: U.S. Geological Survey 

Bulletin 1437, 58 p. 
1. THE IDAHO BATHOLITH AND ITS BORDER ZONE: A REGIONAL PERSPECTIVE 7 

Vallier, T.L., Brooks, H.C., and Thayer, T.P., 1977, Paleozoic rocks in 

eastern Oregon and western Idaho, in Stewart, J., and others, eds., 

Paleozoic paleogeography of the Western United States: Pacific 

Coast Paleogeography Symposium, 1st, Society of Economic Paleontologists and Mineralogists, p. 455-466. 

Vallier, T.L., and Engebretson, D.C., 1984, The Blue Mountains island 

arc of Oregon, Idaho, and Washington: an allocthonous terrane from 

the ancestral Pacific Ocean?, in Howell, D.G., and others, eds., Proceedings of the Circum-Pacific Terrane Conference: Stanford 

University Publications in the Geological Sciences, v. 18, p. 197-199. 

Walker, G.W., and Robinson, P.T., in press, Paleocene(?), Eocene, and 

early Oligocene rocks of the Blue Mountains and adjacent areas, 

in Walker, G.W., ed., Geology of the Blue Mountains region of 

Oregon, Idaho, and Washington: Cenozoic geology of the Blue Mountains region: U.S. Geological Survey Professional Paper 1437. 

Walker, N.W., 1983, Pre-Tertiary tectonic evolution of northeastern 

Oregon and west-central Idaho: constraints based on U/Pb ages of 

zircons [ abs. ]: Geological Society of America Abstracts with Programs, v. 15, p. 371. 

2. PETROLOGY AND ORIGIN OF THE BITTERROOT LOBE OF THE 

IDAHO BATHOLITH 

By MARGO I. ToTH 

ABSTRACT 

The Bitterroot lobe of the Idaho batholith was emplaced in Cretaceous 

and Paleocene time into highly deformed Proterozoic metasedimentary 

rocks east of a subduction zone that was active until at least Triassic(?) 

time. The lobe consists of foliated tonalite and quartz diorite plutons 

intruded by voluminous foliated monzogranite and granodiorite plutons. 

Younger epizonal syenogranite and quartz syenite plutons partially encircle the northern, western, and southern sides of the lobe. 

Hornblende-biotite tonalite and quartz diorite were emplaced in 

Cretaceous time into lower mesozonal levels as small isolated plutons 

that now mark the western and northeastern margins of the lobe. They 

are most similar to I-type granite and were probably derived by partial 

melting of metaigneous source rocks involving an upper-mantle component. However, chemical variation is large in some of these plutons, 

reflecting varying degrees of assimilation of Precambrian metasedimentary material. 

Foliated muscovite-biotite granodiorite and monzogranite plutons, 

intruded into mesozonal or upper mesozonal levels during Paleocene 

time, form the bulk of the Bitterroot lobe. They are chemically homogeneous, suggesting that they belong to a single comagmatic suite, but 

they can be distinguished on the basis of mineralogical and textural differences. The plutons show characteristics of S-type granites and were 

most likely derived by partial melting of relatively homogeneous 

metasedimentary rocks of the Middle Proterozoic Belt Supergroup or 

pre-Belt material. Rare-earth-element patterns indicate strong fractionation, possibly reflecting residual hornblende in the source rock, and Eu 

anomalies are generally very small. 

The youngest mesozonal pluton is chemically and mineralogically 

zoned, from a hornblende-biotite granodiorite rim to a hornblende-biotite 

monzogranite core. Biotite, hornblende, and minor plagioclase crystallized early and accreted along the outer margins of the pluton with some 

trapped magma to form the granodiorite. The remaining magma was 

displaced inward to form the monzogranite. This pluton is an I-type 

granite, and its intrusion marked a radical change in the petrology and 

the source region for the mesozonal granitic rocks of the lobe. 

Nonfoliated Eocene syenogranite and quartz syenite plutons, partially 

encircling the lobe, were intruded into epizonallevels as shallow as 1.5 

km and in at least one place intrude their own volcanic ejecta. The plutons 

are characterized chemically by high Si02, high K20 + Na20, high 

K20/Na20, and low CaO. They show strong enrichment of incompatible 

elements such as La, Nb, Sn, and Y and large negative Eu anomalies 

(Eu/Eu* as low as 0.3). These plutons are most similar to A-type granite, 

and the magmas were probably derived by partial melting of relatively 

felsic granulite in previously depleted lower crust. 

Large dike complexes and isolated dikes of diorite, quartz diorite, and 

tonalite are abundant in the granitic rocks in the central and southern 

parts of the lobe. Some of these intrusive bodies show synplutonic relations with the host granitic rocks. However, chemical characteristics 

indicate they are mantle derived and have ingested various amounts 

of crustal material during emplacement. Original emplacement of these 

magmas into the crust may have provided some of the heat necessary 

for partial melting and formation of the granitic magmas. The occurrence of similar dikes in other batholithic terranes may imply a similar 

causal relation to the generation of granitic magmas. 

INTRODUCTION 

In 1978 the U.S. Geological Survey undertook a fiveyear study to assess the mineral resources of the SelwayBitterroot Wilderness, which comprises 1.25 million acres 

in Idaho and western Montana (fig. 2.1) and encompasses 

much of the central and southern parts of the Bitterroot 

lobe of the Idaho batholith. This paper describes the field 

relations and petrology of the granitic rocks of the lobe 

within and adjacent to the wilderness, studied in conjunction with the mineral-resource appraisal. These observations, combined with geochemical and isotopic data and 

with considerations of regional geology, provide a framework for interpreting the origin and evolution of the 

granitic rocks in the Bitterroot lobe. 

The wilderness program utilized extensive helicopter 

support and consisted of geologic mapping at 1:24,000 

scale, sampling of stream sediments and rocks for traceelement and major-element analyses, and a geophysical 

study. Geological and geochemical maps were produced 

at 1:125,000 scale (Coxe and Toth, 1983a; Toth, 1983a; 

Toth and others, 1983) along with geophysical reports 

(Brickey and others, 1980; Bankey and others, 1982) and 

detailed reports on the geology of certain areas within the 

wilderness (Lund, 1980; Bradley, 1981; Motzer, 1981; 

Bittner, chapter 5; Bittner-Gaber, 1983; Garmezy, 1983; 

Toth, 1983b; Reid, chapter 3). Major-element and traceelement analyses of stream sediments and rock samples 

9 
10 

45° 

30' 

Tb 

THE IDAHO BATHOLITH AND ITS BORDER ZONE 

1W30' 

Approximate boundary 

of Selway- Bitterroot 

Wilderness 

Ym 

115°00' 

20 

20 4 0 KILOMETERS 

Missoula 

Ym 

40 MILES 

MONTANA 

Area of map 

IDAHO 

Ym 

EXPLANATION 

Tb 

Ts 

Tv 

Tg 

~~/ .... ~/ .... ~/ 

',:)'(~J{ 

Alluvial and glacial deposits, undivided 

(Quaternary) 

Basalt (Tertiary) 

Sedimentary rocks (Tertiary) 

Volcanic rocks (Tertiary) 

Epizonal granitic rocks (Tertiary) 

Mesozonal granitic rocks of the 

Bitterroot lobe (Paleocene) 

Mesozonal granitic rocks ~"Jf the 

Atlanta lobe (Cretaceous) 

[ Kt I Mesozonal quartz diorite and 

ton a I ite l Cretaceous) 

~ Metamorphic rocks (Middle? Proterozoic) 

~ Migmatite (Middle? Proterozoic) 

----- Contact 

Fault 

FIGURE 2.1.-Simplified geologic map of the Bitterroot lobe of the Idaho batholith and adjacent areas. Modified from Bond (1978), Lund 

and others (1983), and Toth (1983a). 
2. PETROLOGY AND ORIGIN OF THE BITTERROOT LOBE 11 

were done by U.S. Geological Survey laboratories in 

Denver, Colorado, and the geochemical data and methods 

of analyses are presented in Coxe and others (1982), 

Koesterer and others (1982), Coxe and Toth (1983a, 

1983b), and Toth (1983b). Some of the major-oxide data 

were recalculated into mesonorms (Kosinowski, 1982) 

assuming an FeO:Fe203 ratio of 2:1 and are presented 

in Toth (1983b). Microprobe analyses and methods of 

study are also presented in Toth (1983b). 

Granitic rocks in this paper are classified according to 

Streckeisen and others (1973). I have also classified them 

according to their probable source rocks (Chappell and 

White, 1974; Loiselle and Wones, 1981; Collins and 

others, 1982) in an attempt to understand the generation 

of the magmas of a batholith and their evolution through 

time. The granites are designated as "I -type" (igneous), 

"S-type" (sedimentary), or "A-type" (restite), referring 

to their origin by partial melting of igneous source rock, 

pelitic material, or restite material, respectively. The different source rocks are reflected in the mineralogy, 

chemistry, and field occurrence of the resultant plutons 

(table 2.1). Although this classification is a useful tool, 

some granites do not fit within these guidelines, especially 

those with complex histories involving fractionation or 

contamination; many granites also show characteristics 

of both S-and I-type (Lee and Christiansen, 1983). The 

depths of emplacement for the various plutons of the lobe 

were estimated on the basis of structural criteria as outlined by Buddington (1959) and of mineralogical criteria. 

In this report FeO* refers to total iron as FeO. Rareearth-element concentrations were normalized to the 

chondritic abundances of Hanson (1980). 

ACKNOWLEDGMENTS 

Many geologists have made contributions to this work, 

and I would like to acknowledge particularly the assistance and collaboration of Rolland R. Reid, William R. 

Greenwood, Enid Bittner-Gaber, Lawrence Garmezy, 

Dirk S. Hovorka, Mary E. Koesterer, Karen I. Lund, 

Warren M. Rehn, and William E. Motzer. Although this 

paper is partly based on their mapping and petrologic 

work, the interpretations presented here are my sole 

responsibility. This paper benefited from reviews by Don 

E. Lee and Alan R. Wallace of the U.S. Geological Survey. 

PREVIOUS WORK 

The earliest descriptions of the granitic rocks of the 

Bitterroot lobe were based on large-scale reconnaissance 

projects by Lindgren (1900, 1904) and Anderson (1930). 

Williams (1977, p. 8-12) presents a comprehensive review 

of studies done since those early geologists. Because of 

the inaccessibility of the interior of the Bitterroot lobe, 

TABLE 2.1.-Characteristics of S-, I-, and A-type granites 

[Modified from Chappell and White (1974), Collins and others (1982), and 

Loiselle and Wones (1981)] 

S-type granite 

1. Contains muscovite ± garnet, cordierite, sillimanite, 

and andalusite. 

2. Contains ilmenite as an accessory mineral. 

3. Molecular proportion Al2 o3 /Na2 o + K2 o + CaO >1.1. 

4. Relatively low sodium: Na 2 o <3.2 percent in rocks with 

5 percent K2 o; Na2 o <2.2 percent in rocks with 2 percent 

K2 o. 5. Relatively restricted in composition to high-silica 

types with 66-77 percent Sio2 • 

6. Irregular variation diagrams. 

7. Initial Sr ratios >0.708, isochrons show scatter of 

points, possible homogenization of strontium or 

heterogeneous source. 

8. Associated with regional metamorphism up to amphibolite 

grade; migmatites and xenoliths of metasedimentary rocks 

common. 

9. Normative corundum >1 percent in CIPW norm. 

I-type granite 

1. No peraluminous minerals; contains hornblende, sphene. 

2. Contains magnetite as an accessory mineral. 

3. Molecular proportion of Al 2 o3 /Na2 o + K2 0 + CaO <1.1. 

4. High sodium, commonly higher than 3.2 percent in felsic 

varieties, decreasing to <2.2 percent in more mafic 

types. 

5. Broad spectrum of compositions. 

6. Linear variation diagrams. 

7. Low initial Sr ratios, 0.704-0.706. 

8. Rarely associated with regional metamorphism or 

migmatites. 

9. Normative corundum <l percent in CIPW norm. 

A-type granite 

1. Contains Fe-rich minerals such as annite; feldspar is 

mostly alkali feldspar. 

2. Mildly alkaline with low CaO and Al 2 o3 • 

3. High K2 0 and high K2 0/Na2 o. 

4. Enriched in incompatible trace elements but showing 

strong negative Eu anomalies. 

5. Variable initial Sr ratios from 0.703 to 0.712. 

6. F and Cl contents in the magmas are high. 

7. Intruded late in the magmatic cycle; commonly associated 

with tensional tectonic regimes. 

most earlier work was concentrated along the more accessible northeastern and eastern margins of the lobe. 

Prior to the present study, a reconnaissance investigation by the Idaho Bureau of Mines from 1965 to 1968 

(Greenwood and Morrison, 1973) was the only study made 

of the central and southern parts of the Bitterroot lobe. 

GENERAL GEOLOGY 

The Idaho batholith is 390 km long and 80 to 160 km 

wide and extends from north-central Idaho and western 

Montana to southern Idaho, where it is covered by basalt 

of the Snake River Plain. A probable arch of Proterozoic 

metasedimentary rocks divides the batholith into two 

parts (Armstrong, 1975): the Atlanta lobe in the south and 
12 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

the smaller and younger Bitterroot lobe, the subject of 

this paper, in the north (fig. 2.1). Forming a northwesttrending tabular body, the Bitterroot lobe underlies 

approximately 14,000 km2 and intrudes Proterozoic 

metasedimentary rocks that are correlated with the Belt 

Supergroup-correlative Y ellowjacket Formation and 

Lemhi Group to the south and southwest. The western 

margin of the lobe is in contact with allochthonous 

metavolcanic and metasedimentary rocks of probable 

Mesozoic age that were emplaced along a suture between 

a western Paleozoic and Mesozoic volcanic arc terrane and 

an eastern continental terrane (Hamilton, 1976). Metamorphic grade in the country rocks surrounding the 

batholith increases from greenschist facies to upper 

amphibolite facies close to the contact, but this zoning is 

locally truncated by granodiorite and monzogranite 

plutons near the northern end of the batholith (Nold, 

1968). The northeastern contact with the country rock is 

relatively sharp and mostly concordant in outcrop. 

However, the northeast border of the batholith consists 

of a zone as much as 15 to 20 km wide of numerous 

screens of metasedimentary rock hundreds of meters 

thick and showing varying degrees of assimilation and injection by the granitic rocks (Bradley, 1981). 

Cretaceous plutons of foliated hornblende-biotite 

tonalite and quartz diorite form the western border of the 

Bitterroot lobe, and similar rocks also are present along 

the northeastern border as isolated plutons (not shown 

in figure 2.1). Paleocene mesozonal plutons of foliated 

muscovite-biotite granodiorite and monzogranite form the 

major part of the lobe and intrude the tonalite and quartz 

diorite on the west, southwest, and north. Complex 

migmatite terranes sometimes separate different plutons 

(figs. 2.1, 2.2) and involve septa of Middle(?) Proterozoic 

and possibly older metasedimentary rocks. The appearance, chemistry, and mineralogy of the granitic rocks are 

fairly homogeneous, but the recent work of Williams 

(1977), Toth (1983b), and Hyndman (1984) has outlined 

at least ten plutons with different textural and (or) 

chemical variations; more plutons may be present but as 

yet unidentified. 

Large Eocene epizonal plutons of subalkalic syenogranite and quartz syenite intrude the Proterozoic metasedimentary rocks and all the earlier igneous phases of 

the Bitterroot lobe along the northern, southwestern, and 

southern sides (fig. 2.1). Granophyric textures and 

fluorite-filled miarolitic cavities are common in these 

plutons. An extensive network of northeast-trending 

rhyodacite and rhyolite dikes extends outward from all 

of the plutons and crosscuts all older lithologic units. 

Rhyolitic to andesitic tuff and lava flows are preserved 

in two places in down-dropped fault blocks and locally 

overlie parts of the epizonal plutons on the south side of 

the Bitterroot lobe (Lund and others, 1983). 

A 200- to 500-m-thick mylonite zone, which has a 

shallow eastward dip, borders the eastern and southeastern sides of the Bitterroot lobe, adjacent to the Bitterroot River valley, Montana (fig. 2.2). The mylonite zone 

is younger than all other rock units except the epizonal 

plutons. Its formation has been attributed to gravity 

sliding of an overlying block (Hyndman, 1977), to 

emplacement-related phenomena (Garmezy and Greenwood, 1981), and to extensional tectonism or crustal 

thinning (Garmezy, 1983). Mylonitic textures in the granodiorite and monzogranite plutons extend into the center 

of the lobe. 

Two types of faults are widely distributed throughout 

the Bitterroot lobe and are most evident in the higher 

elevations, where vegetation is sparse. The more common 

type consists of normal faults with a consistent trend of 

azimuth 020° that dip dominantly to the southeast and 

crosscut all lithologic units. Offsets average 4 to 10m but 

are as much as 1.5 km along a few of the more extensive 

faults (Lund, 1980). A steep normal fault along the eastern 

and southeastern margins of the Bitterroot lobe separates 

the mesozonal mylonitic rocks of the batholith to the west 

from younger and shallower, upper mesozonal to epizonal 

rocks to the east and southeast. The absence of mesozonal 

granitic rocks south and east of the fault suggests that 

major postbatholithic uplift along this fault exposed the 

more deeply emplaced plutons of the Bitterroot lobe 

(Garmezy, 1983). The second type of faults trends northwestward and consists of thrust faults that have shallow 

dips to the southwest and are pre-Eocene in age. Offsets 

of as much as 10 m were measured, and they may be 

greater because the faults crosscut relatively homogeneous rock units in which offset is difficult to determine. 

TECTONIC SETTING 

The Idaho batholith is one in a chain of probably 

subduction-related granitic batholiths extending from 

Alaska to Mexico along the western cordillera of North 

America and emplaced during the Mesozoic. In age and 

petrology the Idaho batholith is similar to the other 

granitic batholiths, but it is anomalous in its position at 

the eastern border of the Columbia arc and in its emplacement into mostly Proterozoic continental and miogeoclinal 

metasedimentary rocks instead of Phanerozoic eugeoclinal 

rocks. Small Cretaceous tonalite, diorite, and quartz 

diorite plutons along the western margin of the Bitterroot lobe intrude eugeoclinal accretional complexes 

formed in Triassic(?) time (Hamilton, 1963), or possibly 

FIGURE 2.2.-Geologic map of the central part of the Bitterroot lobe 

of the Idaho batholith showing different plutonic units. 
Qu 

? L 

r- I- ,... 'OJ < ,.. 

" .., < " .... .., ,-

TKg 

TKg 

2. PETROLOGY AND ORIGIN OF THE BITTERROCYI' LOBE 

x Grave Peak 

z 

----==:::::::::- vml· 

Y mp...... . ...... --- - /···· -- ---,..._ 

Blodgett 

Mountain X <{ 

Freeman 

Peak 

------................ 

" 

BEAR CREEK PLUTON 

" 

BEAR CREEK PLUTON 

10 

" 

"~Ymi 

.::> 

0 

,_ -\ 

I no/\_ !..~ 

0~=:J 

0 

X El Capitan 

a: 

a: 

w 

,_ 

Trapper 

CD Peak X 

PLuroiV 

15 MILES 

10 15 KILOMETERS 

EXPLANATION 

>-

w 

...... 

...... 

~ 

Ou 

1-

0 

0 

a: 

a: 

w 

~ CD 

Alluvial and glacial deposits (Quaternary) 

Epizonal granitic rocks (Tertiary) 

Mesozonal granitic rocks of the Bitterroot 

lobe, undifferentiated (Paleocene) 

~ Metamorphic rocks (Middle? Proterozoic) 

~ Migmatite (Middle? Proterozoic) 

----- Contact - Dashed where inferred 

Mesozonal quartz diorite and tonalite 

(Cretaceous) 

----- Fault - Dashed where inferred; U. upthrown 

side; D. downthrown side 

13 
14 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

as late as Late Jurassic into Cretaceous time (T.L. Vallier, 

written commun., 1984). According to ,Hamilton (1978), 

these accretional complexes are the youngest direct 

evidence of subduction along the northwest margin of the 

continent. Farther east, the younger and more abundant 

tonalite and quartz diorite plutons of the Bitterroot lobe 

were intruded into Precambrian metasedimentary rocks 

in latest Cretaceous time, probably long after subduction 

had ceased along the western margin. Furthermore, the 

more voluminous granitic rocks were not intruded until 

the Paleocene. Thus, in contrast to many other batholiths 

in the western cordillera, a direct time relationship cannot be established between subduction and plutonism for 

most of the Idaho batholith. 

Extensional tectonism in the Bitterroot lobe as early 

as 51 Ma is documented by the emplacement of the structureless epizonal Whistling Pig pluton (Lund, 1980; K-Ar 

age on hornblende, R.J. Fleck, written commun., 1981), 

but it may have begun as early as 54 Ma (Toth, 1983b) 

with the intrusion of the Paradise pluton. This tectonism 

involved widespread normal faulting and crustal thinning 

(Garmezy, 1983) and continued during emplacement of the 

other epizonal plutons surrounding the lobe. 

AGE OF THE BITTERROOT LOBE 

The study of the geochronology of the Bitterroot lobe, 

as described by Garmezy (1983), has been hampered by 

the effects of several major penecontemporaneous events 

including mylonitization, rapid uplift of the lobe, and the 

thermal effects from the intrusion of Eocene epizonal 

plutons. Nevertheless, U-Pb study of different zircon sizefractions, combined with K-Ar, Ar-Ar, and some Rb-Sr 

work has outlined the major intrusive episodes (fig. 2.3). 

Eptzonal '--I 

plutons 

Mesozonal monzogranite , 

and granodiorite plutons ..---

Tonalite and quartz diorite 

plutons along the west and 

northeastern parts of the lobe 

160 

Isolated tonalite and 

f---------1 quartz diorite plutons 

west of the lobe 

140 120 100 80 60 40 

MILLION YEARS BEFORE PRESENT 

20 

FIGURE 2.3.-Summary of geochronology of the Bitterroot lobe of the 

Idaho batholith. Data from Armstrong and others (1977), Chase and 

others (1978, 1983), Bickford and others (1981), and J.S. Stacey (written commun., 1982). 

The tonalite and quartz diorite plutons along the western 

and northeastern margins of the Bitterroot lobe were 

emplaced between 100 Ma and 80 Ma (Ferguson, 1972, 

1975; Chase and others, 1978; R.J. Fleck, written commun., 1981; J.S. Stacey, written commun., 1982). A few 

satellite quartz diorite, diorite, and tonalite plutons to the 

west of the main body of the lobe are the only rocks that 

give ages greater than 100 Ma (Armstrong and others, 

1977; J.S. Stacey, written commun., 1982). Chase and 

others (1978, 1983), Bickford and others (1981), and J.S. 

Stacey (written commun., 1982) have identified a Proterozoic zircon component in the volumetrically more important granodiorite and monzogranite plutons of the 

Bitterroot lobe and have derived a latest Cretaceous and 

Paleocene emplacement age of 66 to 57 Ma for the bulk 

of the rocks. This major emplacement episode was closely followed by the intrusion of the epizonal plutons 

between 51 and 46 Ma. 

ISOTOPIC DATA 

Work by Armstrong and others (1977) and Criss and 

Fleck (1983) has shown that all the plutonic rocks west 

of the boundary between Phanerozoic eugeoclinal rocks 

and Precambrian crust in Idaho have initial Sr ratios less 

than 0. 7043, and all those east of it have initial ratios 

greater than 0. 7055; there are no known exceptions to 

this observation. The Bitterroot lobe lies to the east of 

that boundary and is surrounded on three sides by Middle 

Proterozoic Belt Supergroup or pre-Belt metasedimentary 

rocks. Chase and others (1978) studied granodiorite and 

monzogranite plutons along the northeastern border of 

the Bitterroot lobe and obtained high initial Sr ratios 

ranging from 0.7103 to 0.7128. R.J. Fleck (written commun., 1978-1982) and Criss and Fleck (1983) found that 

the rocks in the rest of the lobe have more scattered and 

also high initial Sr ratios (0. 7070-0. 7190) and that points 

on Rb-Sr diagrams generally failed to define isochrons, 

results similar to those of Armstrong and others (1977) 

and Chase and others (1978). They also found that initial 

Sr ratios in the tonalite and quartz diorite plutons increase 

in proximity to the screens of Precambrian migmatites. 

TONALITE AND QUARTZ DIORITE PLUTONS 

Minor amounts of tonalite and quartz diorite are present along the western and southwestern border of the 

Bitterroot lobe (fig. 2.1) and along the northeastern 

margin as small isolated plutons (not shown in fig. 2.1). 

Field evidence for individual plutons along the western 

border of the lobe is absent, and it is possible that the 

tonalite and quartz diorite magmas were emplaced 

simultaneously as a single intrusion. The structure of 

these rocks and their relationship to the Precambrian 
2. PETROLOGY AND ORIGIN OF THE BITTERROOT LOBE 15 

country rock are described in detail elsewhere in this 

volume (Reid, chapter 3). Large migmatite screens commonly separate the quartz diorite and tonalite plutons 

from younger monzogranite and granodiorite plutons. 

In hand specimen the tonalite and quartz diorite are 

gneissic to well foliated, gray, equigranular, and medium 

to more rarely fine grained, and contain accessory biotite 

and euhedral hornblende. Tonalite composition is more 

common than quartz diorite (fig. 2.4). In thin section, 

oligoclase (average An27, optical determination) forms 

subhedral to euhedral unzoned grains, which generally 

have sutured boundaries and are bent or fractured. Quartz 

showing undulatory extinction forms granoblastic grains, 

the larger of which are elongate parallel to biotite and 

have developed tails of annealed subgrains. Some of the 

samples contain anhedral microcline (1-5 percent). Brown 

0 

biotite shows minor alteration to pennine and rutile, and 

green hornblende makes up from 0 to 26 percent of the 

rock, increasing in abundance westward. Accessory 

minerals in the tonalite and quartz diorite include sphene, 

epidote, apatite, prismatic zircon, allanite, magnetite, and 

thorite. 

Four analyzed samples of tonalite and quartz diorite 

from the western part of the Bitterroot lobe (table 2.2) 

have compositions ranging from 60 to 69 percent Si02 

and from 2.8 to 3.7 percent Na20. The fairly wide range 

in Na20 + K20 (fig. 2.5) possibly reflects varying 

amounts of assimilated Precambrian rock. One sample 

high in N a20 + K20 was taken adjacent to a screen of 

pelitic migmatite, indicating local assimilation of N a+ Krich Precambrian material. The ratio of molecular 

Al203/CaO+K20+Na20, generally an index of alumina 

I 

/\ \ 

1

1 ~~% 

\\ 

I \ 

' \ II \ I ,\ I \ 60%;----r-----,-------,-----~---7\ 

I 

/ 

Ep1zonal plutons 

20%/L 

A~-~9~0~%------------6~5-%----------3~5-%--------~,0-%--~p 

A COMPOSITE DIAGRAM 

FIGURE 2.4.-Composite triangular diagram and triangular modal plots for some plutonic phases of the Bitterroot 

lobe of the Idaho batholith. A, alkali feldspar; P, plagioclase; Q, quartz. Boundaries from Streckeisen and others 

(1973). Total number of samples, n. Data for tonalite and quartz diorite plutons from R.R. Reid (written commun., 

1982) and for epizonal plutons from W.E. Motzer (written commun., 1982). A, Composite diagram. B, C, Modal plots. 
16 

Q 

65% 

A 

B 

Tonal1te and 

quartz diorite 

plutons 

n=63 

Quartz 

diorite 

10% 

A 

Monzogranite 

. . . 

. . . 

. . 

. . 

.. . . . . . . 

THE IDAHO BATHOLITH AND ITS BORDER ZONE 

p 

Granodiorite 

Monzogranite \ Granodiorite 

65% 

A 

Canyon Lake 

pluton 

n=63 

10% 

p 

MODAL PLOTS 

FIGURE 2.4.-Continued. 

: . . . .·. \. . ~ ... . . . . . . . . . .. . . 

EXPLANATION 

• One sample 

• Two samples 

.. Three samples 

• Four samples 

x Five samples 

Bear Creek 

pluton 

n=268 

10% 

p 
c 

2. PETROLOGY AND ORIGIN OF THE BITTERROOT LOBE 

Monzogranite 

. . 

2t5%/ AG 

65% 

Q 

. 

Paradise pluton 

n=68 

Granodiorite 

. . . .. ... . . . . . . . . . . ...... . . . : .. . . . 

35% 

. . .. 

. 

Quartz monzodiorite 

Epizonal plutons 

n=72 

~20% 

~-~~ 'QQ) %-:;. 

5% 

10% p 

Q 

Monzogranite 

. . . :.· .. 

·. . . . . . . . 

20%~------------------~~~-------------------~ 

Quartz • 

syenite 

Quartz 

monzonite 

AL_ ______ ~ ________ _L ____ L_ __ _L ________ _L ________ -L~P 

90% 65% 35% 

FIGURE 2.4.-Continued. 

17 
18 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

TABLE 2.2.-Representative major-element analyses, normative mineral calculations, 

[A value followed by L indicates that the element was not detected at the indicated lower detection limit; n.d., not determined; LOI, loss of water upon ignition; Fe20 3 is total iron. DI, 

epizonal syenogranite and quartz syenite plutons from R.R. Reid 

Pluton------

Sample------

Lithology---

Sio2--------

Al2o3-------

Fe 2o3-------

Mg 0---------

CaO---------

Na 2 0--------

KzO---------

T102--------

P205--------

Mn0---------

Zro2--------

LOI---------

Total 

Molecular 

Al 2o3/Ca0+ 

K20+Na 2o 

Q-----------

Ab----------

Or----------

An----------

Bi----------

Hb----------

Mt----------

11----------

R-----------

Ap----------

C-----------

D o I o ---------

Quartz------

PlagioclaseK-spar------

Biotite-----

Hornblende--

AccessoriesQuartz diorite and tonalite pluton{s) 

78RR078 78RR095 78RR138 78RR132 

Tonalite Tonalite Tonalite Quartz 

64.11 

16.49 

5.15 

2.19 

3.92 

3.74 

2.07 

.66 

.27 

.05 

.03 

n.d. 

98.83 

1.07 

26.7 

31.7 

3.4 

17.7 

14.3 

o.o 

2.5 

0.6 

-0.6 

0.6 

1.6 

61.8 

25 

57 

0 

18 

0 

0 

60.19 

16.99 

7.78 

2.98 

4.42 

3.48 

2.47 

.92 

.29 

.1 0 

.04 

n.d. 

99.82 

1.04 

21.7 

29.5 

2.0 

20.0 

20.5 

0 

3.8 

0.9 

-0.8 

1.2 

0.7 

73.2 

28 

59 

3 

9 

0 

1 

69.34 

15.87 

2.29 

.67 

1.96 

3.41 

4.04 

.30 

.07 

.03 

• 02 

n.d. 

98.22 

1.17 

29.9 

28.9 

20.6 

9.3 

5.3 

o.o 

1.1 

0.3 

-0.2 

0.2 

2.5 

79.4 

3 

51 

0 

42 

0 

4 

diorite 

65.49 

15.97 

4.68 

1.58 

3.49 

2.77 

1. 78 

.59 

.13 

.05 

• 05 

n.d. 

96.83 

1.25 

35.4 

23.5 

3.4 

16.5 

11.6 

0 

2.3 

0.6 

-o.5 

0.3 

3.4 

62.3 

16 

45 

0 

35 

3 

0 

80MT185 

Monzogranite 

(core) 

Bear Creek pluton 

80MT220 80MT080 

Monzogranite 

(core) 

Granodiorite 

(rim) 

Major oxide analyses 

72.5 

15.4 

1.33 

.30 

1.66 

4.00 

4.00 

.15 

.OIL 

.02L 

n.d • 

.52 

99.86 

1.11 

72.8 

15.7 

1.22 

.40 

2.46 

4.80 

2.16 

.14 

.OIL 

.02L 

n.d. 

.34 

100.02 

1.11 

Mesonorms 

30.4 

33.9 

22.0 

7.9 

2.8 

0 

.6 

.1 

.4 

.1 

1.6 

86.2 

31.3 

40.7 

10.9 

11.9 

3.3 

0 

.6 

.1 

.2 

.1 

1. 1 

82.9 

Modal analyses 

28 

42 

27 

3 

0 

0 

29 

42 

25 

4 

0 

0 

71.4 

14.7 

1. 75 

.74 

1.65 

3.70 

4.41 

.28 

.1o 

.02L 

n.d • 

.59 

99.3 

1.06 

29.6 

31.3 

23.0 

7.5 

4.9 

0 

.8 

.3 

.4 

.2 

1.1 

84.0 

30.5 

42 

20 

7.5 

0 

0 

80MT142 

Granodiorite 

(rim) 

70.5 

15.5 

2.11 

.76 

1.80 

3.80 

4.13 

.29 

.20 

.03 

n.d. 

.41 

99.53 

1.10 

29.2 

32.2 

21.0 

7.6 

5.5 

0 

1.0 

.3 

0 

.5 

2.0 

82.4 

30.5 

43.5 

20.5 

5.5 

0 

0 

80MT254 

Quartz 

diorite 

(zone I) 

63.0 

16.6 

4.55 

2.90 

4.12 

4.30 

2.87 

.68 

.20 

.07 

n.d. 

.55 

99.84 

.94 

17.6 

36.4 

9.0 

16.0 

12.6 

4.8 

2.2 

0.6 

-0.4 

0.5 

0 

63.0 

9.0 

63.5 

5 

12 

9.5 

1 
2. PETROLOGY AND ORIGIN OF THE BITTERROOT LOBE 19 

and modal compositions of rocks from the Bitterroot lobe of the Idaho batholith 

differentiation index Q+Ab+Or. Major oxide analyses and mesonorms in weight percent; modal analyses in volume percent. Modal analyses of tonalite and quartz diorite plutons and 

and W.E. Motzer, respectively (written commun., 1982)] 

Paradise pluton Epizonal syenogranite and quartz syenite plutons 

81MT134 80MT258 80MT017 80MT138 78WM010 78WM018 78WM060 78WM020 

Grano- Grano- Monzo- Tonalite Syeno- Syeno- Syeno- Syenodiorite diorite granite (upper granite granite granite granite 

(zone II) (zone III) (zone IV) contact) 

Hajor oxide analyses 

66.4 70.0 72.5 62.6 77.3 72.0 73.7 76.6 

16.0 15.4 15.2 16.5 12.4 13.9 13.4 12.7 

2.95 2.24 1.28 5.03 1.37 2.58 2.15 1.18 

1.71 .78 .60 3.25 .101 .20 .20 .1 0 

3.06 2.08 1. 77 4.50 .20 .61 • 62 .34 

4.10 4.25 3.79 3.91 3.40 3.30 3.50 3.20 

3.40 3.27 4.35 1.93 4.86 5.91 5.34 5.19 

.49 .41 .32 .71 .12 .25 .20 .09 

• 14 .15 .011 .16 .0101 .101 .101 .101 . 

.03 .06 .011 .09 .021 .02 .02 .021 

n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 

.25 .48 .32 .59 .66 .87 .42 .34 

98.53 99.12 100.13 99.28 99.0 99.7 99.6 99.8 

1.00 1. 07 1.05 .99 1. 36 1.01 1.08 1.11 

Mesonorms 

23.2 26.2 29.4 23.2 38.4 28.8 31.4 38.4 

34.7 36.0 33.0 33.1 28.8 28.0 29.6 28.8 

13.9 14.2 23.2 0.1 27.5 32.6 29.5 27.5 

14.3 11.0 8.5 21.3 1.0 3.0 3.1 1.0 

9.9 8.2 4.0 18.1 2.1 4.1 3.5 2.1 

0 0 o.o o.o o.o o.o o.o o.o 

1. 4 0 0.7 2.4 0.7 1.2 1.0 0.7 

.5 0.4 0.2 0.7 o.l 0.2 0.2 0.1 

-0.2 0.1 0.2 -0.1 0.6 0.7 0.3 0.6 

.3 0.4 0.1 0.4 o.o o.o o.o o.o 

.3 0.4 0.9 0.2 1.2 1.0 0.7 1.2 

71.8 76.4 85.6 56.3 84.7 89.4 90.5 94.7 

Modal analyses 

19 25.5 29 26 33 24 32.5 40 

51 54 39 46.5 52 59 48 46 

16 11.5 27 2.5 13 11 15 12 

8 6 3.5 15 2 1 2.5 2 

6 3 1 10 .3 5 1.5 0 

0 0 0 0 .2 .7 .4 .2 
20 

40% 

20% 

40%-

20% 

THE IDAHO BATHOLITH AND ITS BORDER ZONE 

F 

60% Tonalite an'd quartz 

d1orite plutons 

F 

Paradise 

pluton 

n=55 

n=6 

dzonel 

F 

60% :. 

D1onte 

d1kes 

n=43 

... .. . . ·. . .. . .. :· .. 

40% .· 

F 

60% 

F 

40% Bear Creek 

pluton 

20% ... . .. . . f/J!k.~f'.: • ..... ::;! 

F 

20%:-,·· • 

.. .. · .ti·· 

40% 

20% 

Ep1zonal 

plutons 

n=37 

n=58 

FIGURE 2.5.-AFM diagrams for some plutonic phases of the Bitterroot lobe of the Idaho batholith. A, N~O + K20; 

F, FeO*; M, MgO. Dot, one sample; square, two samples; triangle, three samples; diamond, four samples; n, 

total number of samples. 
2. PETROLOGY AND ORIGIN OF THE BITTERROOT LOBE 21 

saturation, ranges from 1.04 to 1.25 (table 2.2, fig. 2.6) 

and indicates that some of the rocks are peraluminous. 

Three samples with ratios higher than 1.1 were taken adjacent to or at least near migmatite bodies, and their high 

ratios may reflect local assimilation of metasedimentary 

material; samples taken far from the migmatite bodies 

have ratios less than 1.1. Mesonorms indicate a range in 

the differentiation index (D.I. = Q + Ab + Or) from 62.3 

to 79.5 and an average value of 72.1. Normative corundum ranges from 1.2 to 3.4 percent, consistent with a 

peraluminous composition. 

The characteristics of the tonalite and quartz diorite 

plutons along the northern, western, and southwestern 

sides of the Bitterroot lobe are most similar to those of 

1-type granites (table 2.1). They contain no peraluminous 

minerals, have hornblende, sphene, and magnetite, show 

a broad spectrum of compositions, and have very high 

Na20 contents. However, they also show some characteristics typical of S-type granite: they are associated with 

regional metamorphism and migmatites and have more 

than 1 percent normative corundum, molecular Al20s/ 

CaO + K20 + Na20 ratios from 1.04 to 1.25, and initial Sr 

ratios greater than 0.7070 (Criss and Fleck, 1983). The 

initial Sr ratios, molecular Al20s/CaO + K20 + Na20 

ratios, and normative corundum values all increase toward 

migmatite units, and the high values that give these rocks 

some S-type characteristics most likely resulted from the 

assimilation of Precambrian metasedimentary material. 

Primary tonalite and andesite magmas are generated 

at temperatures that are probably too high (about 

1,100 °C) for crustal conditions (Wyllie and others, 1976), 

and therefore they probably do not form within the continental crust. The tonalite and quartz diorite plutons of 

the Bitterroot lobe were most likely derived by partial 

melting of material at upper mantle depths, and this may 

have involved metavolcanic eugeoclinal rocks or sub34~--------------- ----------==----------------------------------------------------, 

18 

22~ 14 

10 

6~ 

H~ 

CXl 

CXl 

c:i 

0 

a> 

c:i 

I I I 

~ 

a> 

c:i 

m 

~ 

CXl 0 N 6 0 q ~ ~ CXl 0 N ~ ~ 

0 q q 

I CTfJ I 

0 

N 

N 

N 

I I I I 

Tonalite and quartz 

diorite plutons 

I I I I 

CXl 0 N M 

n=6 

Epizonal 

plutons 

n=37 

1W1 

FIGURE 2.6.-Percent frequency of the ratio of molecular Al20 3/Ca0 + K20 + N~O for various plutonic phases of the Bitterroot lobe of the Idaho 

batholith. Median value, m; number of samples, n. 
22 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

ducted oceanic lithosphere. Ingestion and assimilation of 

Precambrian rock occurred as the magmas moved 

through and were emplaced in the crust. 

MONZOGRANITE AND GRANODIORITE PLUTONS 

Monzogranite and granodiorite plutons constitute the 

major part of the Bitterroot lobe and are generally well 

exposed in its central and southern parts. Outcrops range 

in elevation from about 900 m in the lower creek bottoms 

to over 3,050 m at Trapper Peak, thereby exposing a 

substantial vertical section through the batholith. Most 

of this part of the batholith was previously thought to be 

chemically and mineralogically homogeneous, but 

Williams (1977) delineated five separate plutons in the 

north-central part of the lobe, and our mapping has outlined at least six mesozonal or lower epizonal plutons in 

the southern and central parts of the lobe (fig. 2.2). These 

plutons are strikingly similar in their major- and traceelement chemistry and have been distinguished and 

mapped on the basis of texture, mineralogy, and crosscutting relationships. Further detailed mapping may show 

that the Bitterroot lobe consists of a composite of possibly 

20 or more different plutons, distinguishable by slight textural, mineralogical, or chemical differences. Geochronologic studies (R.J. Fleck, written commun., 1982; J.S. 

Stacey, written commun., 1982) suggest that the times 

of intrusion of these plutons were fairly close to one 

another. 

BEAR CREEK PLUTON 

The northern boundary of the Bear Creek pluton is 

defined by a string of elongate, mostly east-west-trending 

migmatite units (fig. 2.2) of Proterozoic metasedimentary 

rocks that the pluton intruded in both lit-par-lit and 

strongly disharmonic fashions. Large feldspar phenocrysts (average 3-5 em long) are common in the western 

parts of the pluton, suggesting the pluton grades into a 

border phase or possibly another separate pluton. The 

eastern boundary of the pluton has not been located, but 

the rocks east of the Bitterroot Mountains crest contain 

substantially more accessory muscovite and may be a part 

of another pluton. The location of the southern boundary 

is also unknown, but may lie as far south as the Canyon 

Lake pluton (fig. 2.2). Primary flow foliation in rocks of 

the pluton is weakly to moderately well defined by aligned 

feldspar and biotite grains, compositional layering, and 

rare biotite schlieren. Xenoliths are generally rare and 

consist dominantly of calc-silicate gneiss, with less common quartzofeldspathic schist and quartzite, all correlative 

with Belt or pre-Belt metasedimentary rocks. 

The rocks of the Bear Creek pluton are typically white 

to light gray and slightly porphyritic and contain subhedral to euhedral feldspar phenocrysts 6 to 8 mm long 

in a groundmass of medium-grained feldspar, quartz, and 

biotite, and rare fine-grained muscovite. Medium-grained 

equigranular rocks are less abundant. Modal compositions 

range from granodiorite to monzogranite; the average 

composition of the pluton lies on the granodiorite-monzogranite boundary at about 30 percent quartz (fig. 2.4). 

Variations from the average composition may reflect the 

varying amount of phenocrysts. 

In thin section, the texture of the granitic rocks from 

the Bear Creek pluton is medium-grained hypidiomorphic 

granular to slightly porphyritic. Compositions of oligoclase 

(average Ab76An21 Or3, microprobe determination) show 

very little variation throughout the pluton. Poikilitic 

orthoclase also shows very little compositional variation 

(average Ab17AnoOrg3, microprobe determination), 

although some very slight normal and reverse zoning was 

detected. Biotite (1-4 percent) is characteristically altered 

to a combination of chlorite and subhedral to needle-like 

opaque minerals. All of the main mineral phases show 

varying degrees of deformation and (or) recrystallization 

related to mylonitization. 

Muscovite is present in trace amounts in epitaxial intergrowths with biotite, as an alteration of plagioclase and 

more rarely potassium feldspar, and as small primary 

interstitial grains. Larger primary grains are present only 

in the northernmost part of the pluton adjacent to the 

migmatite unit, where they increase in abundance to 

3 percent of the rock. Using the criteria of Miller and 

others (1981) for distinguishing primary and secondary 

muscovite, unequivocal primary grains were identified 

only in rocks from the northernmost part of the pluton. 

The compositions of primary and secondary grains from 

one sample (table 2.3) show that the muscovite in the 

Bear Creek pluton is far from ideal muscovite 

(KAl13Si30lO(OH)2) and contains substantial celadonite, 

ferrimuscovite, and Ti-muscovite components. Ti02 

ranges from 3.2 to 3.5 percent, MgO from 0.6 to 0.7 percent, and FeO* from 1.1 to 1.8 percent. Secondary 

muscovite is richer in Si02, Al203, and K20, poorer in 

MgO, and contains similar amounts of Ti02 and FeO* 

(table 2.3). 

Accessory minerals in the Bear Creek pluton consist of 

magnetite and minor ilmenite, zircon (with cloudy, 

anhedral cores), monazite, apatite, and rare euhedral palegreen epidote. Sillimanite is absent, although Williams 

(1977) identified fibrolitic sillimanite in similar granite of 

the Tom Beal Park granite in the central part of the 

Bitterroot lobe. 

MAJOR- AND TRACE-ELEMENT CHEMISTRY 

The Si02 content of rocks from the Bear Creek pluton 

ranges from 68 to 75 percent, but more than half of the 

samples contain from 71 to 73 percent Si02. The pluton 
2. PETROLOGY AND ORIGIN OF THE BITTERROOT LOBE 23 

is slightly zoned with respect to Si02 content (table 2.2), 

which increases from less than 72 percent on the margins 

to 72 to 74.5 percent in the center. The zonation is not 

strictly concentric, however, and the distribution of other 

major oxides mimics this variation only erratically. No 

regular variation in mineralogy has been observed. The 

relative homogeneity in the chemistry of the pluton is 

demonstrated by the slight variation on an AFM diagram 

(fig. 2.5) and the slight variation in the mesonorm 

geochemistry. The D.I. of samples from the pluton ranges 

from 90 to 96 and averages 93.3. Apart from those that 

contribute to the D.l., the normative constituents are 

anorthite (4.0-11.9 percent), biotite (1.2-5.4 percent), corundum (0.8-3.4 percent), and minor magnetite, ilmenite, 

and rutile. The molecular ratio of Al203/CaO + K20 + 

Na20, which ranges from 1.05 to 1.25 (fig. 2.6) and has 

a median value of 1.1, indicates the rocks are dominantly 

peraluminous. 

Rare-earth-element (REE) concentrations in rocks from 

the Bear Creek pluton (fig. 2.7) are also markedly similar, 

showing little variation from sample to sample. They are 

characterized by strong fractionation of light REE 

(LREE) relative to the heavy REE (HREE); Ce/Yb ratios 

range from 44.0 to 171.4 and average 95.2. Strongly fractionated REE patterns such as these have commonly been 

interpreted as a reflection of garnet in the source region 

(Arth and Hanson, 1972, 1975), but Arth and Barker 

(1976) have demonstrated that hornblende as a residual 

phase can also cause HREE depletion in a coexisting felsic 

liquid. The depleted HREE patterns of the Bear Creek 

pluton may therefore represent residual hornblende and 

(or) garnet in the source rock. The unusually high abundance of LREE in the pluton is presently unknown, but 

it may reflect a LREE-enriched source. 

All but two of the samples from the pluton show almost 

no europium anomaly, having Eu/Eu* very close to 1 (Eu* 

is the expected Eu concentration obtained by simply interpolating between Sm and Gd; the ratio Eu/Eu * is 

therefore a measure of the Eu anomaly). The absence of 

an Eu anomaly is usually an indication that plagioclase 

fractionation did not occur in the evolution of the magma 

or that plagioclase was not an equilibrated residual phase 

in the source rock. The absence of such an anomaly might 

also reflect oxidation of Eu +2 to Eu +3. One sample near 

the eastern edge of the pluton has an Eu anomaly of 0.5, 

possibly indicating that plagioclase fractionation was 

operative along the margin of the pluton. 

The Bear Creek pluton is an S-type granite body and 

was probably derived by partial melting of metasedimentary crustal material. I infer this because the pluton contains primary muscovite; has a mean molecular ratio of 

Al203/CaO+K20+Na20 of about 1.1, more than 1 percent normative corundum, high Si02 content of 68 to 75 

percent, somewhat irregular variation diagrams, and initial Sr ratios of more than 0. 708 with irregular isochrons; 

is associated with regional metamorphism; and has 

migmatites and xenoliths of metasedimentary origin. The 

only ways in which the pluton does not meet S-type 

criteria are in having some accessory magnetite instead 

of ilmenite (which may reflect the oxygen fugacity), and 

in having relatively high Na20 contents. In rocks with 

5 percent K20, Na20 is about 3.5 percent or greater (versus 3.2 percent for typical S-type granite), and in rocks 

with 2 percent K20, Na20 is about 4.5 percent (versus 

2.2 percent). 

Granitic magma such as that which formed the Bear 

Creek pluton is generated as a consequence of regional 

metamorphism and anatexis of continental crust (Wyllie 

and others, 1976). It can be generated in the presence of 

excess H20 (Piwinskii, 1968; Winkler, 1979) or in dry 

rock containing hydrous minerals such as muscovite, 

biotite, or hornblende (Brown and Fyfe, 1970; Robertson 

and Wyllie, 1971). The presence of H20 substantially 

lowers the solidus, so dry systems need significantly 

higher temperatures to initiate melting. When breakdown 

curves of biotite, muscovite, or hornblende are reached, 

the release of excess H20 lowers the solidus and allows 

melting to begin. As pointed out by both Hyndman (1981a, 

b) and Miller (1981), however, water supplied by the 

breakdown of hydrous minerals in the source rocks is 

probably insufficient for the formation of a large mass of 

granite. Hyndman (1981a, b) suggests that additional 

water released in dehydration reactions in subduction 

zones may rise into the overlying continental crust and 

so contribute to the partial melting of the rocks. 

Anatectically derived magmas commonly contain 

suspended restite material (Winkler, 1979). In the Bear 

Creek pluton, however, there is no evidence of restite 

material other than cores of Precambrian age in zircon 

grains and possibly some rare schlieren layers. Except 

where some assimilation of country rock is observed, the 

composition of the pluton therefore represents the composition of the magma from which it crystallized. If restite 

material did exist, it was removed sometime in the evolution of the pluton before emplacement. 

The major source of the magma for the Bear Creek 

pluton was probably metamorphosed Belt or pre-Belt 

metasedimentary rocks. Melting may have been related 

to subduction, but the timing and mechanics are still 

unclear; subduction may have caused melting by providing 

necessary heat, by supplying necessary water from 

dehydration reactions in the upper mantle, or both. Armstrong and others (1977) indicate that initial Sr ratios in 

Idaho batholith rocks are generally too low to permit 

simple anatectic intracrustal genesis, and as a result they 

suggest some upper mantle contribution. 

Our geochemical studies (Koesterer and others, 1982; 

Coxe and Toth 1983a, b; Toth, 1983b) show that in major-
24 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

TABLE 2.3.-Microprobe analyses and structural formulas 

[Al(iv) is alumina in the tetrahedral- site and AI( vi) is alumina in the octahedral site. n.d., not detected; ---, not determined. Oxides in weight 

Bear Creek pluton Paradise pluton 

MUSCOVITE BIOTITE 

Upper 

Primary Secondary Zone II Zone IV contact 

sio2---- 44.3 44.8 44.5 

Ti02---- 3.5 3.2 3.4 

Al2o3--- 32.1 32.0 33.1 

MgO----- .6 .7 .7 

FeO*---- 1.6 1.1 1.8 

MnO-----

Cao-----

Na 2o---- n.d. n.d. n.d. 

K20----- 10.3 10.1 10.3 

La 2o3---

ce2o3---

Total 92.4 91.9 93.8 

Si------ 6.18 6.24 6.08 

Ti------ .40 .38 .38 

Al(iv)-- 1.82 1.76 1.93 

Al(vi) -- 3.48 3.48 3.42 

Mg------ .12 .14 .14 

Fe------ .18 .10 .16 

Mn------

Ca------

Na------ n.d. n.d. n.d. 

K------- 1.84 1.78 1.78 

La------

Ce------

Fe/Fe+Mg .60 .42 .53 

and trace-element geochemistry the Bear Creek pluton 

is strikingly similar to other granodiorite and monzogranite plutons in the central and southern parts of the 

Bitterroot lobe. Much of the lobe may consequently belong 

to a single comagmatic suite, as suggested by the work 

of Hyndman (1984) and Wiswall and Hyndman (chapter 

4). The plutons of the central and southern parts of the 

lobe were therefore probably derived from a homogeneous 

source rock and underwent very little differentiation or 

assimilation before emplacement. The homogeneous 

chemistry also suggests that the plutons did not travel 

very far from their source rock before they finished 

crystallizing. However, the notable absence of restite 

material does suggest the plutons are at least somewhat 

removed from their source regions. 

GRAVE PEAK PLUTON 

The Grave Peak pluton crops out north of the Bear 

Creek pluton, from which it is separated by a string of 

46.3 38.45 38.78 40.55 39.74 38.15 

3.5 2.57 2.53 2.54 3.11 2.59 

33.9 15.55 14.06 14.05 14.52 14.18 

.2 12.10 12.63 12.55 12.29 14.27 

1.7 14.56 15.04 16.50 16.94 13.57 

.22 .23 .36 .30 .27 

.04 .03 .03 .03 .03 

.1 .06 .o8 .05 .17 .14 

11.1 9.63 9.05 9.76 9.77 9.44 

96.8 93.27 92.43 96.39 96.87 93.64 

6.14 5.85 5.93 5.99 5.87 5-76 

.40 .27 .29 .28 .35 .29 

1.86 2.15 2.06 2.01 2.13 2.24 

3.46 .57 .46 .44 .40 .46 

.04 2.80 2.80 2.04 2.09 3.06 

.16 1.80 2.00 1.76 1.70 1.80 

.03 .03 .04 .04 .03 

.01 .01 0 .01 .01 

.02 .01 .02 .01 .05 .04 

1.88 1.76 1.81 1.84 1.84 1.81 

.8o .39 .41 .43 .44 .37 

elongate migmatite units (fig. 2.2); its other contacts have 

not been mapped. It is similar in all respects to the Bear 

Creek pluton, except that it everywhere contains 3 percent or more of primary muscovite and also contains 

substantially more accessory epidote. This pluton may be 

a part of the Tom Beal Park granite as mapped by 

Williams (1977). It has characteristics most similar to an 

S-type granite, and its origin is presumably closely related 

to that of the Bear Creek pluton. 

CANYON LAKE PLUTON 

The Canyon Lake pluton crops out as an elongate eastwest body underlying approximately 190 km2 along the 

southern border of the Bitterroot lobe (fig. 2.2). It intrudes 

the Bear Creek pluton to the north and is bounded on its 

west and south margins by migmatite units and granitic 

rocks similar in appearance and chemistry to those of the 

Bear Creek pluton. The pluton grades eastward into 

mylonite, which is truncated by the border fault. The 
2. PETROLOGY AND ORIGIN OF THE BITTERROOT LOBE 25 

of minerals from the Bear Creek and Paradise plutons 

percent. Structural formulas of muscovite, biotite, and hornblende based on 22 oxygens; structural formulas of allanite based on 12 oxygens] 

Paradise pluton Paradise pluton 

HORNBLENDE ALLANITE 

Upper Upper 

Zone II Zone IV contact Zone II Zone IV contact 

45.96 45.37 46.55 46.86 47.34 31.31 32.33 30.79 30.83 34.00 

.93 .93 1.12 .94 .83 .84 .69 1.07 .86 .42 

10.40 8.89 10.13 9.02 8.69 16.00 16.72 14.07 14.59 18.46 

10.90 11.15 9.98 10.63 12.86 1.01 .81 1.12 .82 .89 

14.27 14.01 16.42 16.27 12.95 13.82 13.08 14.73 15.09 12.56 

.30 .48 .56 .42 .40 .26 .28 .36 .38 .28 

11.46 11.44 11.60 12.00 12.03 11.41 13.02 11.53 10.85 13.41 

1.50 1.57 1.78 1.60 1.19 

1.28 1.06 1.40 1.21 .82 

5.35 6.08 7.53 5.84 5.63 

9.77 10.03 11.83 11.02 8.49 

97.00 94.90 99.54 99.84 96.63 89.77 93.04 93.03 90.28 94.14 

6.54 6.60 6.53 6.52 6.67 2.98 2.97 2.95 2.99 3.01 

.10 .1 0 .12 .10 .09 .06 .05 .08 .06 .03 

1.46 1.40 1.4 7 1.48 1.33 .02 .03 .os .01 0 

.28 .13 .20 .12 .11 1.78 1.78 1.54 1.66 1.92 

2.32 2.42 2.09 2.36 2.70 .14 .11 .16 .12 .12 

1.69 1.71 1.93 1.82 1.53 1.10 1.01 1.18 1.22 .93 

.04 .06 .07 .os .os .02 .02 .03 .03 .02 

1. 74 1.79 1.74 1. 79 1.82 1.16 1.28 1.18 1.13 1.27 

.43 .44 .so .42 .32 

.23 .20 .30 .21 .15 

.19 

.34 

.42 .43 .48 .43 .36 

Canyon Lake pluton has four synplutonic phases: extremely porphyritic rocks in the west; equigranular rocks 

in the central part; highly deformed, heterolithologic rocks 

in the east; and orbicular rocks along the southern border, 

in contact with Precambrian migmatite. Primary flow 

foliation is generally weak or absent, but a strong phenocryst lineation is commonly present in the porphyritic 

rocks. 

Monzogranite and lesser granodiorite dominate the 

Canyon Lake pluton (fig. 2.4), but minor components 

range from syenogranite to diorite (not shown in fig. 2.4). 

The major phases of the pluton mentioned above are not 

distinguishable chemically and have been mapped on the 

basis of distinctive textures and the relative abundances 

of certain trace minerals. 

Dominating the western part of the pluton is a monzogranite and granodiorite main phase characterized by 

large (as much as 5 em by 9 em), euhedral, poikilitic orthoclase megacrysts, which constitute 15 to 25 percent of the 

rock, and by 1 to 3 percent granular (as large as 1 mm 

.21 .2 7 .21 .18 

.34 .41 .39 .2 7 

.89 .94 .91 .93 .88 

across) magnetite. The matrix consists of medium-grained 

equigranular plagioclase, quartz, and biotite, along with 

minor orthoclase. 

In the central part of the pluton, a fine- to mediumgrained equigranular monzogranite is commonly 

associated with the main phase. Contacts, though sharp 

in hand specimen, are irregular and deformed on a large 

scale and suggest simultaneous emplacement of these 

phases. Fine-grained magnetite constitutes a trace 

amount to 3 percent of the rock and is generally concentrated along joint surfaces. This phase of the Canyon Lake 

pluton is most abundant along its northern contact and 

commonly contains highly folded and deformed biotite 

schlieren. 

In the eastern part of the pluton, large outcrops show 

crosscutting, deformed, and swirled bodies of quartz 

diorite, granodiorite, syenogranite, quartz monzodiorite, 

and diorite, all of which are commonly crosscut by dioritic 

dikes. A few outcrops show as many as twelve texturally 

or mineralogically distinct subphases which were 
26 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

a:: 

0 

z 

0 

:I: 

u 

UJ 

UJ 

a:: 

1000 

100 

Bear Creek 

pluton 

S1000 ~--~--,--TTTTI~~ ~-~-~-l 

~ ~ Ep1zonal - Dionte ~ 

<{ plutons d1kes ~ (f) 

UJ 

~ 100 J 

-1 

~ t 

n=2 

j 

10 t 

~ j 

•,---:~~ . 

-- ----~ --------~,-------~ 

o Ll____L__;__j_ _ _L._L__C~_.L__L J ~ --'-----'---_l____l__l -"- _ _l_l_! ~ La Eu Lu La Eu Lu 

RARE-EARTH ELEMENTS 

FIGURE 2.7.-Chondrite-normalized rare-earth-element (REE) concentrations in some plutonic phases of the Bitterroot lobe of the 

Idaho batholith. Number of samples, n; stippled pattern shows range 

of data. 

apparently emplaced nearly simultaneously. Most of the 

subphases are fine to medium grained, but porphyritic 

granodiorite is locally present. Early intrusive subphases 

also contain minor diorite xenoliths, which are elongate 

and lie parallel to the borders of the enclosing subphase. 

Primary flow foliation defined by biotite is highly variable 

but generally parallels contacts between subphases. 

Magnetite concentrations as much as 4 em thick and 

0.5 m2 in area are present along joint surfaces and foliation planes in the main subphase. No genetic relations 

between subphases were observed in the outcrops where 

intrusive relationships could be determined, and it is not 

known whether these subphases were derived from 

several different sources or resulted from remobilization, 

mixing, or assimilation of earlier intrusions. 

Orbicular rocks are present at six localities along and 

near the southern contact of the Canyon Lake pluton. 

They are present as small (0.5-4 m2), mostly undeformed 

pods of orbicular monzogranite and granodiorite along 

single structural horizons in the main phase of the pluton 

and as individual orbicules in a matrix of medium-grained 

granite. Five types of orbicules are present, ranging from 

2 to 12 em in diameter; some contain hornblende-biotiteplagioclase cores and single or multiple concentric shells 

of plagioclase, biotite, and magnetite oriented in a radial 

or tangential manner. 

In thin section, orthoclase phenocrysts of the main 

phase display Carlsbad twinning and are strongly poikilitic. The groundmass is hypidiomorphic, and matrix 

orthoclase shows well-developed flame perthite. Small 

primary(?) grains of muscovite are present in minor 

amounts. Magnetite is present both as small inclusions 

in feldspar and as large poikilitic interstitial grains. Zircon is a common accessory mineral, and monazite, sphene, 

and secondary calcite are present in trace amounts. 

The major- and trace-element chemistry of the Canyon 

Lake pluton is generally indistinguishable from that of 

the Bear Creek pluton, suggesting a similar origin and 

source rock for both. However, the Bear Creek pluton is 

clearly S-type, whereas the Canyon Lake pluton shows 

some characteristics of I-type granite in that it does contain magnetite and may lack primary muscovite. Because 

it is bordered on the south by a large body of migmatite, 

the Canyon Lake pluton may have assimilated some 

metasedimentary material. 

PARADISE PLUTON 

The Paradise pluton intrudes the middle of the Bear 

Creek pluton in the south-central part of the Bitterroot 

lobe (fig. 2.2), and its intrusion marked major changes in 

structural style of intrusion, mineralogy, and chemistry 

of the mesozonal granitic rocks. The lobe-shaped main 

phase of the pluton underlies about 105 km2 in the west 

and has vertical walls on its northern and western sides. 

Toward the south and east the pluton becomes sheetlike 

and nearly horizontal and has a thickness of approximately 1,450 m (fig. 2.8). This sheet phase extends at least 

5 km south to El Capitan, and possibly as far as 16 km 

south, and it becomes increasingly interlayered with the 

Bear Creek pluton southward, away from the main phase. 

A petrologically distinct, finer grained, equigranular contact rock of varying thickness forms a third phase along 

the vertical walls of the main phase and at the top of the 

sheet phase (contact phase on figure 2.8). 

The lower contact of the Paradise pluton's sheet phase 

and the contacts along the vertical walls of the main phase 

are sharp and locally mylonitized. The upper contact of 

the sheet phase generally has a zone 2 to 100 m thick 

showing degrees of intermixing with the Bear Creek 

pluton and contains abundant xenoliths of calc-silicate 

rock and quartzite along foliation planes; some large 

xenoliths of igneous rock from the Bear Creek pluton are 

also present. This mixed zone is extremely deformed, as 

shown by folding, shearing, and crenulation of the rocks. 

In two places the upper contact is sharp and flat beneath 

large coherent blocks of metasedimentary rock that could 

be remnants of wall or roof material. 
2. PETROLOGY AND ORIGIN OF THE BITTERROOT LOBE 27 

w E 

Bear Creek pluton 

FIGURE 2.8.-Schematic cross section showing structure and main, sheet, 

and upper contact phases of Paradise pluton (stippled). 

Rocks from the main and sheet phases of the Paradise 

pluton are foliated, light to dark gray, and slightly porphyritic. They contain 10 to 30 percent potassium feldspar 

phenocrysts, which are very poikilitic, subhedral to 

euhedral, and 6 to 10 mm long, in a groundmass of 

medium-grained euhedral plagioclase, subhedral biotite 

and hornblende, and anhedral quartz. The rocks are less 

porphyritic toward the core of the main phase. Euhedral 

prisms of sphene are common as grains that average 1 to 

2 mm but are as large as 5 mm across. 

The composition of the main phase of the Paradise 

pluton varies extremely widely (fig. 2.4), and the spatial 

relationships (fig. 2.9) indicate the pluton is compositionally zoned. The main phase ranges from plagioclase-rich 

granodiorite and monzogranite in a rim 0.25 to 3 km wide 

(zone II), through more potassium feldspar-rich granodiorite and monzogranite that forms the bulk of the pluton 

(zone III), to monzogranite in an elongate zone 0.25 to 

0.5 km wide in the core (zone IV). Hornblende and biotite 

decrease from approximately 7 to 10 percent (combined) 

on the outer margins to 3 to 5 percent in the inner core. 

This compositional zonation is rough and only approximately concentric, and compositions vary somewhat 

erratically within each zone. A small, elongate tonalite to 

quartz diorite body (zone I) in the southern part of the 

main phase also complicates the zonal pattern. Rocks in 

the sheet phase have a range in composition similar to 

that in the main phase, but the variation is much less 

systematic. Petrography also suggests that the sheet 

phase has a slight vertical zonation marked by upward 

increase in plagioclase and decrease in potassium feldspar. 

Rocks from the upper contact phase are dark gray and 

equigranular; they have average grain sizes of about 4 mm 

D 

6 KILOMETERS 

6 MILES I 

Zone I 

Zone II 

Zone II 

Zone IV 

FIGURE 2.9.-Simplified lithologic zonation of Paradise pluton as determined from point counts and major-element chemical data. Zone 

I, tonalite and quartz diorite; zone II, granodiorite and quartz monzodiorite; zone III, granodiorite and quartz monzodiorite; zone IV, 

monzogranite. See text for differentiation between zones II and III. 

Dots are sample locations for major-element chemistry. 

and porphyritic varieties are absent. Compositions vary 

widely and include diorite, quartz diorite, tonalite, and 

minor granodiorite. The rocks are characterized by very 

low potassium feldspar content Oess than 5 percent) and 

highly varied amounts of quartz (1-45 percent). Biotite 

ranges from 6 to 56 percent and averages 17 percent, and 

hornblende is locally present. 

Thin sections show that the granitic rocks in the sheet 

and main phases are slightly porphyritic; medium-grained 

equigranular varieties are less common. Albite-twinned 

oligoclase (zone I, Ab7oAn270rg; zone IV, Ab77An2oOrg, 

microprobe determinations) is zoned normally and occurs 

in slightly poikilitic subhedral to euhedrallaths. Orthoclase 

(zone I, Ab1sAn10rgg; zone IV, Ab1s.5Ano.50rs6, microprobe determinations) forms subhedral oikocrysts containing 10 to 80 percent (mostly less than 40 percent) 

inclusions and more rarely occurs as an interstitial phase. 

Green poikilitic hornblende is present in highly varied 

amounts (trace to 7 percent) but rarely exceeds biotite 

in abundance. From zone II to zone IV hornblende composition shows increases in Si02, FeO*, K20, Na20, and 

Ti02 (table 2.3); the Fe/Fe+ Mg ratio also increases from 

0.42 to 0.48. The changes in the chemistry of subhedral 

brown biotite are similar to those for hornblende, and the 

Fe/Fe+ Mg ratio increases from 0.39 to 0.44 (table 2.3). 

Zoned allanite (trace to 1 percent) increases in abundance, 

size, and euhedral form from zone I to zone IV, where 

it reaches an abundance of 0.5 percent and sizes as large 

as 2 mm. The chemistry of the allanite also changes 
28 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

systematically from zone II to zone IV: La203, Ce203, 

FeO*, and MgO increase, whereas Al203, CaO, and Si02 

decrease (table 2.3). Accessory minerals consist of sphene, 

apatite, zircon (some with cloudy, anhedral cores), monazite, magnetite, and rare pyrite; ilmenite and muscovite 

are notably absent. All ofthe main minerals have been 

deformed and (or) recrystallized to varying degrees during mylonitization. 

Thin sections show that the equigranular upper-contact 

rocks have a hypidiomorphic granular texture and average 

grain sizes of 3 to 4 mm. Plagioclase occurs as euhedral, 

rarely poikilitic grains, and sparse orthoclase forms small 

anhedral grains. Olive-green biotite is extremely fresh, 

and hornblende is present in euhedral poikilitic grains that 

contain 10 percent inclusions of round quartz. Strongly 

zoned light-brown allanite is present in very small 

amounts in all samples. Apatite, sphene, magnetite, and 

rare thorite make up the accessory minerals. 

MAJOR- AND TRACE-ELEMENT CHEMISTRY 

The chemistry of the main phase of the pluton (table 

2.2) reflects the mineralogical variation observed in the 

point-count data: an outer, mafic margin grades concentrically inward and upward into more felsic rocks, as 

shown by variations in the weight percent of the major 

oxides. Si02 increases inward from 63 to 72.5 percent, 

Al203 decreases inward from > 17 percent to <15.3 

percent, FeO* decreases inward from 3.6 percent to <2.1 

percent, CaO decreases inward from 4.1 to 1.8 percent, 

and MgO decreases inward from 2.9 to <1 percent. Na20 

and K20 show highly varied concentrations in the pluton; 

K20 is zoned only on the largest scale, increasing slightly inward. From zone I to zone IV, the main phase shows 

an increase in Na20 + K20 and a corresponding decrease 

in FeO* and MgO (fig. 2.5). The variation is continuous 

from zones II through IV, but there is a compositional gap 

between zones I and II, suggesting that these rocks may 

not be directly related. The sheet phase of the Paradise 

pluton may also be zoned, because the major-element 

chemistry shows a general vertical trend of increasing 

Si02 downward and correspondingly decreasing Al20 3, 

FeO*, MgO, CaO, and Na20 {Toth, 1983b). 

In the main phase of the pluton, the mesonorm 

chemistry shows the following trends from zone I to zone 

IV: normative hornblende, biotite, anorthite, and albite 

decrease, and normative orthoclase increases; normative 

corundum increases from 0 to 1.2 percent, although most 

of the samples contain less than 1 percent. The D.l. ranges 

from 54 in zone I (reflecting high quantities of normative 

hornblende and biotite) to 85.6 in zone IV. Molecular 

Al203/CaO + K20 + Na20 ratios vary from 0.85 to 

1.12 (fig. 2.6), indicating that much of the pluton is 

metaluminous. 

The rocks of the Paradise pluton, like those of the Bear 

Creek pluton, typically exhibit an extreme fractionation 

of LREE over HREE (fig. 2. 7); Ce/Yb ratios range from 

38.4 to 120 and average 77 .5. As indicated previously, 

these strongly fractionated patterns may indicate residual 

hornblende or garnet in the source rocks, although the 

high abundance of LREE is still unknown. REE concentrations increase from zone IV to zone I in the main phase; 

this reflects the increase in the amount of hornblende, the 

major site for the REE. The Eu anomalies, which vary 

within each zone and range from 0.64 to 0.93, suggest 

some plagioclase fractionation. 

Concentric zonation of the main phase of the Paradise 

pluton is demonstrated in silica variation diagrams, AFM 

variation, mesonorm chemistry, REE chemistry, and 

mineral chemistry. Speculations regarding the origin of 

the zonation must take into account the apparent absence 

of any intrusive contacts within the pluton and the 

absence of any xenoliths of country or basement rock. The 

field evidence favors a continuous process rather than 

multiple intrusions of magma, and it seems to preclude 

assimilation as a possible mechanism. Three mechanisms 

for the origin of the zonation should be considered: 

thermogravitative diffusion, fractional crystallization by 

crystal settling, and fractional crystallization by inward 

crystallization and (or) filter pressing. These mechanisms 

are discussed in detail in Toth (1983b), where I utilize data 

for major, trace, and rare-earth elements and modeling 

of Rb-Sr-Ba fractionation to conclude that the origin of 

the zonation involved fractional crystallization of the 

latter type. Biotite, hornblende, and minor plagioclase 

crystallized early and were accreted along the outer 

margins of the pluton together with some trapped magma. 

The remaining magma was in this way displaced inward 

to form the inner zones. 

The intrusion of the Paradise pluton marked a major 

change in the evolution of the Bitterroot lobe, from intrusion of massive amounts of S-type granite to intrusion 

of a single I-type granite. The pluton lacks any 

peraluminous minerals, contains hornblende, sphene, and 

magnetite, has a molecular ratio of Al203/CaO + K20 + 

Na20 of less than 1.1, and has less than 1 percent normative corundum. The Na20 content is greater than 4 percent in most of the rocks, and the Si02 content covers a 

broad range from 60 to 7 4 percent. Variation diagrams 

are typically linear (Toth, 1983b ). 

The main deviation of the Paradise pluton from I -type 

granite in the classification of Chappell and White (197 4) 

is that the initial Sr ratios are greater than 0. 708 (R.J. 

Fleck, personal commun., 1981; Criss and Fleck, chapter 

6). Most I-type granite has low initial Sr ratios (0. 704,_ 

0.706). The relatively higher ratios for the Paradise pluton 

may indicate that either (1) the source rock for the 

Paradise pluton was chemically more primitive and much 
2. PETROLOGY AND ORIGIN OF THE BITTERROOT LOBE 29 

lower in the crust than the source for the Bear Creek 

pluton; ingestion of Precambrian crustal material as the 

magma ascended increased its initial Sr ratios; or (2) the 

source rock for the Paradise pluton was closely related 

to that of the Bear Creek pluton but was of a markedly 

contrasting composition (such as metabasalt or metaandesite) that required higher melting temperatures. The 

younger age of the Paradise pluton would allow time for 

temperature to increase and for melt to be produced from 

the amphibolite. Until more detailed Rb-Sr work is done 

on the Paradise pluton, it is not possible to determine 

which of these origins is more likely. However, all of the 

chemical data appear to require a source rock markedly 

different from that of the Bear Creek pluton. Furthermore, the Paradise pluton was intruded at the start of 

major regional extension, and therefore a different depth 

of origin for the melt might be expected. 

BURNT RIDGE PLUTON 

The Burnt Ridge pluton crops out in the southeastern 

part of the Bitterroot lobe on the east side of the bordering normal fault (fig. 2.2). It is in sharp contact to the 

south with the Piquett Creek pluton and is covered by 

alluvium to the north; its eastern margin has :riot been 

mapped. Along the western border of the pluton, brecciation of the rocks increases in proximity to the fault, and 

a quartz-flooded, sillimanite-bearing marginal phase is 

intermittently present adjacent to the fault. The western 

border also contains various-size inclusions of the Bear 

Creek pluton, which are evidence that the Burnt Ridge 

pluton is younger. 

The pluton consists of medium-grained equigranular 

granodiorite containing 35 to 40 percent quartz. Foliation 

is very weakly defined by aligned mica grains and 

elongate, cigar-shaped quartz crystals. Quartz crystals are 

subhedral to euhedral and rarely terminated, and may 

have been derived from a liquidus phase. Microcline 

poikilitically includes all the other minerals. Euhedral 

oligoclase (An2s-3o, optical determination) is albitetwinned and normally zoned, and the more calcic cores 

are preferentially altered to saussurite. Biotite dominates 

over primary muscovite, and the two constitute 3 to 5 percent of the rock; epitaxial intergrowths between them are 

common. Accessory minerals include anhedral to 

subhedral opaque minerals, zircon, xenotime, apatite, and 

rare sphene. 

Mineralogically the Burnt Ridge pluton is a slightly 

peraluminous S-type granite, but without chemical data 

it is difficult to classify its origin. The presence of possible liquidus quartz suggests that the pluton intruded at 

a relatively shallow level (Tuttle and Bowen, 1958). These 

attributes may indicate a change in the depth of origin 

and possibly in the source rock for the younger plutons 

of the Bitterroot lobe. 

PIQUETT CREEK PLUTON 

The Piquett Creek pluton crops out in an elongate eastwest area in the southeast part of the Bitterroot lobe, east 

of the bordering normal fault (fig. 2.2). It is bounded on 

the north and east by the Burnt Ridge pluton, on the west 

by the border fault, and on the south by Eocene epizonal 

granite. Age relations with the Burnt Ridge pluton are 

unknown, but the two plutons are in mutually sharp 

contact. 

The composition of rocks in the Piquett Creek pluton 

ranges from granodiorite to syenogranite. The pluton contains distinctive euhedral Carlsbad-twinned orthoclase 

phenocrysts 5 to 6 em long (a few are as much as 10 em 

long) in a medium-grained groundmass generally devoid 

of potassium feldspar. Phenocrysts constitute 15 to 20 

percent of the rock and contain zonally arranged inclusions of biotite; many phenocrysts were aligned by flow, 

but others show random orientations. Quartz in these 

rocks is subhedral to euhedral and has circular cross sections 5 to 6 mm in diameter; terminated ends on some 

crystals in a few samples suggest that quartz was possibly 

a liquidus phase. The quartz is strongly aligned and is commonly the only mineral defining the primary flow foliation. 

In thin section, quartz is seen to be strained and almost 

totally recrystallized into mosaics of subgrains averaging 

0.5 mm across. Orthoclase phenocrysts poikilitically include all the other minerals, and euhedral plagioclase 

(An23-25, optical determination) is complexly twinned and 

normally zoned. Biotite is the only mafic mineral and 

makes up 5 to 9 percent of the rock. Primary(?) muscovite 

is present, and secondary muscovite is abundant in altered 

plagioclase. Accessory minerals include euhedral irontitanium oxide minerals, allanite, zircon, thorite, and rare 

apatite. 

Like the Burnt Ridge pluton, the Piquett Creek pluton 

contains possible liquidus quartz, suggesting a shallower 

depth of intrusion than the other monzogranite and 

granodiorite plutons of the lobe. It has mineralogical 

characteristics similar to S-type granite, but definite 

classification is not possible without chemical data. 

SYENOGRANITE AND QUARTZ SYENITE PLUTONS 

Eocene epizonal plutons surround the Bitterroot lobe 

on the south, southwest, and north sides and have been 

studied in some detail by Hall (1980), Holloway (1980), 

Lund (1980), Motzer (1981), Pawlowski (1981), Rehn and 

Lund (1981), Lund and others (1983), Toth (1983a, b), and 

Rehn (1984). Outcrops of these plutons are generally 
30 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

deeply weathered and extremely friable. The plutons were 

intruded to very shallow levels (1.5 km; Rehn and Lund, 

1981) during Eocene time, and they locally intrude their 

own volcanic ejecta. The nonfoliated textures of these 

plutons suggest post-tectonic emplacement; for some of 

the plutons, normal faults apparently controlled the 

emplacement. Miarolitic cavities are abundant and contain crystals of fluorite, smoky quartz, and microcline. The 

cavities average 4 em across but are as large as 15 em 

by 25 em by 30 em. Contacts of the plutons with the country rock are extremely sharp, and hydrothermal alteration effects from the plutons extend as far as 2 km into 

the country rock. A network of northeast-trending 

rhyodacite and rhyolite dikes extends outward from most 

of the plutons, the density of dikes decreasing rapidly 

away from each pluton. 

The plutons consist of medium- to coarse-grained 

syenogranite, quartz syenite, and monzogranite (fig. 2.4), 

with lesser amounts of alkali feldspar granite and quartz 

monzonite (not shown in figure 2.4) (Hall, 1980; Lund, 

1980; Pawlowski, 1981). Thin sections show the texture 

of the rocks to be hypidiomorphic granular to seriate and, 

in places, porphyritic. Orthoclase perthite is the dominant 

mineral and contains as much as 50 percent plagioclase 

in the form of elongate stringers, patches, and irregular 

blebs. Plagioclase also forms subhedral tabular grains that 

are commonly zoned from about An 16 in the core to An4 

at the rim (optical determination, W.E. Motzer, written 

commun., 1983). Biotite is present in all of the samples 

(average 8 percent), and hornblende occurs sporadically 

(average 3 percent). Accessory minerals include allanite, 

apatite, baddeleyite, cassiterite, fluorite, monazite, 

sphene, thorite, zircon (both primary and possibly inherited), and iron-titanium oxide minerals (Motzer, 1981). 

The epizonal plutons are characterized by high Si02, 

high K20+Na20 and K20/Na20, and low CaO (table 2.2). 

The Si02 contents in 37 samples from the Whistling Pig, 

Running Creek, Bad Luck, and Painted Rocks plutons 

range from 70 to 81 percent, and three-quarters of the 

samples contain between 7 4 and 78 percent Si02. The 

sum of Na20+K20 ranges from 7.38 to 9.43 percent, 

K20/Na20 ranges from 1.06 to 2.03, and CaO ranges 

from 0.20 to 0.92 percent. The rocks have abnormally low 

MgO and high but variable amounts of N a20 + K20 

(fig. 2.5). Molecular ratios of Al203/CaO + K20 + Na20 

(fig. 2.6) range from 0.99 to 1.40 and have a median value 

of 1.08. The mesonorm data show a wide D.l. range, from 

85.4 to 95.6 in the more highly differentiated rocks. 

Normative corundum ranges from 0.5 to 2.0 percent. 

The plutons are also very enriched in incompatible 

elements such as Be, La, Nb, Sn, andY (Rehn, 1984); the 

magmas that formed them were saturated with H20 and 

had high F contents. The REE patterns for the plutons 

(fig. 2. 7) show a strong fractionation of LREE over 

HREE, although not as strong as that in the Bear Creek 

and Paradise plutons; Ce/Yb ratios range from 64.2 to 

102.1 and average 79.6. The epizonal granites also have 

substantial Eu anomalies (Eu/Eu* as low as 0.3), suggesting plagioclase fractionation or a strongly depleted 

source rock. 

With the exception of their high H20 content, the 

Eocene epizonal granites have chemical and mineralogical 

signatures typical of A-type granite. They have low MgO 

and CaO, high N a20 + K20, and high F, are enriched in 

the incompatible elements, and have strongly negative Eu 

anomalies. Collins and others (1982) conclude that A-type 

granite is derived in the lower crust by partial melting 

of relatively felsic granulite that had previously been 

melted to produce granite. Vapor-absent melting of the 

granitic source rock at high temperatures generates a 

relatively anhydrous melt containing F. The fluorine 

forms complexes with large highly charged cations, including the REE, thereby producing the high abundances 

of these elements. A depleted source rock would also explain the Eu anomalies, caused by the earlier depletion 

of plagioclase. It is unlikely that the epizonal granitic 

plutons were derived from partial melting of the 

mesozonal granitoids or Precambrian metasedimentary 

material. However, relatively high initial Sr ratios in the 

younger plutons do suggest some assimilation of crustal 

material. 

DIKE ROCKS 

Four types of dike rocks (not shown on figures 2.1 and 

2.2) are present in the central and southern parts of the 

Bitterroot lobe: leucocratic pegmatite; fine-grained biotite 

and hornblende-biotite granite; hornblende-biotite diorite, 

quartz diorite, and tonalite; and hypabyssal rhyolite and 

rhyodacite. Aplite dikes are notably absent. 

PEGMATITE DIKES 

Scattered and minor pegmatite dikes intrude the 

mesozonal granitic rocks and are locally more abundant 

along contacts and near the roofs of the plutons. They 

generally have shallow dips and are randomly discordant 

to the flow foliations of the host rock, but in places they 

lie within the foliation. The dikes are commonly massive, 

but in some places strong foliation is present and is concordant to that in the host rock. Contacts are very sharp, 

and widths of the dikes average 20 to 30 em. 

The pegmatite dikes range compositionally from monzogranite to syenogranite and consist of orthoclase, quartz, 

and lesser amounts of sodic plagioclase. The dikes are 

generally homogeneous in composition, but a few are 

zoned, with feldspars on the margin and gray quartz in 

the core. Grain size averages 3 to 4 em but in a few places 
2. PETROLOGY AND ORIGIN OF THE BITTERROOT LOBE 31 

is as much as 8 em. Mafic minerals are very rare and consist mostly of trace amounts of biotite. Minor amounts 

of muscovite and hornblende are present in one location 

near the northeastern contact of the Bear Creek pluton. 

Garnet-bearing layers are abundant in a pegmatite on the 

northern contact of the Bear Creek pluton, adjacent to 

a pelitic migmatite unit. 

FINE-GRAINED GRANITIC DIKES 

Fine-grained granitic dikes are widely distributed in 

mesozonal rocks throughout the central and southern 

parts of the Bitterroot lobe and constitute 1 to 3 percent 

of the mesozonal plutons. They are locally present in 

greater abundance near the northern and northeastern 

margins of the Bear Creek pluton, near the margins of 

the main phase of the Paradise pluton, and along the upper contact of the sheet phase of the Paradise pluton. They 

range in thickness from 0.2 to 4 m but are typically 0.5 to 

1 m. The dikes generally have a limited lateral extent and 

feather out within 8 to 10m along strike, although a few 

of the thicker dikes extend at least 30 m. Contacts are 

straight-sided to less commonly crenulate and are generally sharp, although in a few places they are diffuse over 

an interval of 2 to 3 em. Chilled borders are absent, and 

biotite selvages occur along the margins of a few of the 

dikes. Many of the granitic dikes lie within the plane of 

flow foliation of the host rock, and weak foliation in the 

dikes is concordant with that of the host rock. Angular 

inclusions of the coarser grained granitic country rock are 

common in the dikes; these are 6 to 10 em long, commonly aligned with the foliation, and show varying degrees 

of assimilation. 

The granitic dike rocks have a sucrosic texture, are light 

to medium gray, fine grained, and equigranular, and contain varying amounts of accessory biotite. They range in 

composition from granodiorite to monzogranite and contain 20 to 35 percent quartz; this quartz content is similar 

to that of the coarse-grained compositional equivalents 

in their respective plutons. In thin section, the texture 

of the dike rocks is seen to be fine-grained xenomorphic 

granular with an average grain size of about 0.5 mm. 

Biotite and minor muscovite are the accessory minerals 

in most of the dikes, but dikes crosscutting the Paradise 

pluton also contain sphene, allanite, and rare hornblende. 

Foliation is weakly to strongly defined by aligned biotite 

grains and, more rarely, by aligned feldspar grains. Few 

deformational features are present. 

The fine-grained granitic dikes intrude the plutons 

generally along foliation planes or at low angles to the 

foliation planes; the dikes also have the same mineralogy 

as their host plutons. They most likely originated late in 

the crystallization history of each pluton and represent 

some of the last compositions of each melt. Dikes that 

crosscut all structures, however, may have originated 

from another pluton at depth. 

DIORITE DIKES 

Diorite, quartz diorite, and tonalite dikes are moderately 

abundant in the central and southern parts of the lobe and 

are collectively referred to herein as diorite dikes. They 

have been described in detail by Toth (1983b) and are present in two modes: large, multiphase complexes and small, 

solitary dikes. 

Several large, multiphase complexes of diorite dikes 

crop out in the Bear Creek, Paradise, and Canyon Lake 

plutons, generally close to the pluton boundaries. The 

complexes extend at an azimuth of 290° as far as 8 km 

and are as thick as 500 m. In each complex, several textural and chemical phases of diorite intruded the host 

plutons concordantly to slightly discordantly. Granitic 

material is commonly present within the complexes as 

partially ingested xenoliths and averages 10 to 15 volume 

percent. Intrusive relationships among the various phases 

of the diorite, and between the granitic rocks and the 

diorite, are complex; they indicate the sequential intrusion of medium- to coarse-grained diorite, fine-grained 

diorite, and porphyritic diorite into the granitic plutons. 

The ages of the various diorite phases overlap, indicating 

that the diorite units are partly contemporaneous. Field 

relationships indicate that, although the granitic rocks are 

mostly older than the diorite, in places they reintrude 

the diorite; this relationship further indicates a general 

contemporaneity. 

Most of the isolated diorite dikes are related to the 

larger diorite dike complexes, and the density of dikes 

decreases rapidly away from the complexes. Fine-grained 

and porphyritic diorite dikes range in thickness from 

0.4 to 10m and have a lateral extent commonly greater 

than 10m. Contacts are generally sharp and well defined, 

but in a few places they are slightly ragged and crenulate. 

Most of the dikes parallel the foliation in the country rock, 

but some crosscut the foliation at low angles {10°-25°). 

Foliation is moderately to well defined by biotite and hornblende, is developed parallel to the contact, and is most 

likely flow related. The isolated diorite dikes generally 

lack inclusions of their host plutons. Partially assimilated 

small inclusions were noted, however, in a few samples, 

and feldspar xenocrysts from the Bear Creek pluton were 

also observed. The dikes commonly crosscut the rocks of 

the pluton in sharp contact, but at one locality the diorite 

shows synplutonic relationships typical of those found in 

the multiphase complexes. 

Fine-grained diorite is equigranular, light gray, and 

rarely aphanitic. The rocks commonly have a porphyritic 

texture containing 5 to 20 percent plagioclase phenocrysts 

that are 1 to 2 mm long and sometimes show flow align-
32 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

ment. The porphyritic samples contain rare glomerophyric 

clots of hornblende and biotite 4 to 5 mm across. Large 

plagioclase xenocrysts, presumably from the granitic 

plutons that the dikes intrude, are intensely corroded and 

altered. Medium-grained diorite is equigranular and is 

either light gray or has a salt-and-pepper appearance. 

Prismatic hornblende dominates over biotite, and together 

they constitute 10 to 15 percent of the rock. Potassium 

feldspar is generally absent, and anhedral quartz is 

present in highly variable amounts. In three localities, 

medium-grained rocks grade into rocks of similar composition that have average grain sizes greater than 1 em. 

Euhedral sphene is present in many of the samples. 

The dike rocks range from diorite to tonalite and quartz 

content varies from virtually zero to almost as much as 

total feldspar (fig. 2.4). Biotite (1 to 46 percent) and hornblende (2 to 52 percent) make up the varietal minerals. 

Major-element chemistry of the diorite dikes demonstrates the widely varying composition of the rocks 

(fig. 2.5), which reflects large amounts of assimilation of 

granitic material. Silica variation diagrams also show a 

relatively high degree of scatter in major-oxide concentrations (Toth, 1983b). The REE concentrations in two 

dikes are similar to those in the Bear Creek and Paradise 

plutons (fig. 2. 7), but the dikes have slightly lower Ce/Yb 

ratios (81.9 and 107.1). The two samples also have small 

Eu anomalies of 0.83 and 0.95. The similarity between 

the dikes and their host plutons in REE concentrations 

and Eu anomalies probably reflects the ingestion of some 

crustal material by both groups of rocks. 

The diorite dikes in the central and southern parts of 

the Bitterroot lobe are mostly younger than the host 

granitic plutons. Because dioritic magma cannot be produced by the differentiation of a granitic magma, the 

diorite dikes must have an origin different from that of 

the granitic plutons. Comparable initial Sr ratios (R.J. 

Fleck, written commun., 1982), however, suggest a 

similar source. Synplutonic diorite dikes like these in the 

Bitterroot lobe have been noted and described in granitic 

batholiths worldwide (see, for example, Grout and Balk, 

1934; Roddick and Armstrong, 1959; Pitcher and Read, 

1960; Gastil, 1975; Pitcher, 1978). The dikes are generally small, have consistent orientation within a pluton, and, 

most important, are intruded over the entire emplacement 

history of a plutonic complex. 

Primary tonalite and andesite magmas are believed to 

be mantle derived (Wyllie and others, 1976). Initial Sr 

ratios of diorite dikes in the Bitterroot lobe indicate that 

the dikes have ingested crustal material with relatively 

high initial Sr ratios, supporting field observations that 

the diorite dikes commonly contain partially assimilated 

inclusions of granitic material. Probably, therefore, the 

diorite dikes in the Bitterroot lobe are mantle derived and 

during their emplacement ingested crustal material. Their 

original emplacement into the crust may have provided 

much of the heat necessary for the partial melting that 

formed the granitic magmas, and their occurrence in other 

granitic batholithic terranes might also indicate a close 

link to the generation of the granitic magmas. The dikes 

may have been emplaced at depth throughout the entire 

formation and crystallization of the granitic plutons, but 

only penetrated into higher levels late in each pluton's 

consolidation. 

RHYOLITE AND RHYODACITE DIKES 

Rhyolite and rhyodacite dikes are associated with 

Eocene epizonal plutons and are most common within and 

in close proximity to these plutons. They are buff, gray, 

or light purple and typically about 1 m thick, and they 

form conspicuous topographic ribs and walls. Commonly 

porphyritic, they contain 15 to 20 percent phenocrysts in 

a matrix that is microcrystalline to cryptocrystalline and 

displays plumose and spherulitic devitrification textures. 

The phenocrysts consist of plagioclase, embayed hipyramidal quartz, and subhedral biotite and hornblende. 

Flow banding and miarolitic cavities are common. Majorand trace-element chemistry is similar to that of the 

epizonal plutons. 

DEPTH OF EMPLACEMENT OF THE 

BITTERROOT LOBE 

Metamorphic grade of the Belt Supergroup surrounding 

the Bitterroot lobe increases toward the lobe from greenschist to upper amphibolite facies (Reid and Greenwood, 

1968). Sillimanite-zone rocks crudely encircle the lobe, 

although the lobe truncates the isograds along its northern 

border (Johnson, 1975). Because the metamorphism was 

probably coincident with or very close in time to batholith 

emplacement (Wehrenberg, 1972; Williams, 1977; Hyndman, 1980), the metamorphic grade constrains the emplacement level of at least part of the batholith to 

mesozonal or possibly upper catazonallevels. Wehrenberg 

(1972) indicates that kyanite is locally present in 

sillimanite-bearing gneiss, and the coexistence of these 

two aluminum silicates constrains depth and pressure 

estimates. The pressure for the kyanite-sillimaniteandalusite invariant point is controversial, but a minimum 

pressure is probably about 3. 75 kbars, equivalent to a 

depth of about 13 km (Holdaway, 1971). This is supported 

by estimates that the thickness of overburden during 

emplacement of the Bitterroot lobe was 13 to 21 km 

(Hyndman, 1983). It is quite likely, however, that the 

metamorphism was directly related to the emplacement 

of only the quartz diorite and tonalite plutons on the north-
2. PETROLOGY AND ORIGIN OF THE BITTERROOT LOBE 33 

ern, western, and southern sides of the Bitterroot lobe. 

Because the more felsic granitic rocks of the Bitterroot 

lobe were intruded 30 to 40 m.y. later, their level of 

emplacement may have been substantially shallower. 

The presence of primary muscovite in granitic rocks has 

been used in the past as a petrogenetic indicator, allowing estimation of a lower limit for the pressure or depth 

of intrusion of a granitic body. The intersection of the 

muscovite stability curve with the water-saturated granite 

solidus was used as an upper limit of intrusion; it has been 

experimentally determined to be at 3 to 4 kbars (11 to 15 

km), depending on which muscovite breakdown reaction 

is used. Recent work by Anderson and Rowley (1981), 

however, has shown that adding a celadonite component 

(phengitic muscovite) to muscovite allows it to coexist with 

liquids at much lower pressures. Their calculations indicate that muscovite can coexist with H20-saturated 

adamellite-granite melts at 2.6 to 3.1 kbars, corresponding to depths of 9.6 to 11.5 km. However, as they also 

point out, granitic rocks formed at depths as shallow as 

5 km contain primary celadonite (see Benoit, 1971; Nelson 

and Sylvester, 1971; Sylvester and others, 1978; Bradfish, 1979), so their estimates may be further decreased 

as better thermodynamic data become available. Because 

the granite solidus is at slightly higher temperatures for 

nonsaturated systems, its intersection with the muscovite 

stability curve will be at slightly higher pressure and 

temperature in such systems. 

The presence of primary celadonite in the Bear Creek 

pluton should put some constraints on the depth of 

crystallization for the granitic rocks of the central part 

of the ~itterroot lobe. According to thermodynamic data, 

minimum depths of crystallization are about 10 km for 

saturated systems and somewhat greater for nonsaturated systems. However, because there are several 

documented cases of celadonite-bearing epizonal granitic 

plutons, depth of crystallization for the granitic rocks of 

the lobe could have been substantially shallower. 

The plutons southeast of the bordering fault (Piquett 

Creek and Burnt Ridge plutons) represent magmas that 

were intruded at depths between mesozonal and epizonal. 

They contain inclusions of mesozonal granite that crop 

out west of the fault, and they are therefore younger than 

the other granitic rocks of the lobe. They show a very 

weak flow foliation, suggesting shallower depths of intrusion, and have no mylonitic textures. Possible liquidus 

quartz in the plutons also suggests relatively shallow 

depths of intrusion (Tuttle and Bowen, 1958). 

The Eocene plutons on the northern, southwestern, and 

southern sides of the Bitterroot lobe were clearly intruded 

into epizonallevels, possibly as shallow as 1.5 km (Rehn 

and Lund, 1981). They are generally structureless, contain abundant miarolitic cavities, and in places intrude 

their own volcanic ejecta. 

CONCLUSIONS 

The history of the Bitterroot lobe is characterized by 

the intrusion of compositionally very variable plutons over 

a relatively short period of time in a changing tectonic 

regime. Early intrusion of minor amounts of tonalite and 

quartz diorite into Mesozoic eugeoclinal rocks was related 

to subduction off the west coast of North America. This 

was followed by the intrusion, farther eastward, of Late 

Cretaceous tonalite and quartz diorite plutons, which had 

probable mantle components but assimilated large 

amounts of Precambrian crustal material as they moved 

upward and were emplaced in the crust. Voluminous 

monzogranite and granodiorite plutons were intruded 

about 30 to 40 m.y.later at mesozonal to upper mesozonal 

levels. These plutons are compositionally homogeneous, 

had a relatively homogeneous source rock, and underwent 

little differentiation before emplacement. They were 

almost immediately uplifted to epizonallevels, concurrently with a change in the tectonic environment to one of 

extension. Large, highly differentiated syenogranite and 

quartz syenite plutons of lower crustal origin were then 

intruded. The granitic rocks of the lobe changed during 

its history from 1-type (with a mantle source and the 

generation of tonalite and quartz diorite magmas) to 

S-type (with a crustal source and the emplacement of the 

large monzogranite and granodiorite plutons that constitute most of the lobe), and finally to A-type (with a 

depleted crustal source and the emplacement of epizonal 

syenogranite and quartz syenite plutons). The transition 

from S-type to A-type was marked by the intrusion of a 

small, compositionally zoned I -type granitic pluton. 

REFERENCES CITED 

Anderson, A.L., 1930, Geology and mineral resources of the region about 

Orofino, Idaho: Idaho Bureau of Mines and Geology Pamphlet 34, 

p. 8-12. 

Anderson, J.L., and Rowley, M.C., 1981, Synkinematic intrusion of 

peraluminous and associated metaluminous granitic magmas, 

Whipple Mountains, California: Canadian Mineralogist, v. 19, 

p. 83-101. 

Armstrong, R.L., 1975, Precambrian (1500 m.y. old) rocks of central 

Idaho-the Salmon River arch and its role in cordilleran sedimentation and tectonics: American Journal of Science, v. 275-A, 

p. 437-467. 

Armstrong, R.L., Taubeneck, W.H., and Hales, P.O., 1977, Rb-Sr and 

K-Ar geochronometry of Mesozoic granitic rocks and their Sr isotopic 

composition, Oregon, Washington, and Idaho: Geological Society 

of America Bulletin, v. 88, p. 397-411. 

Arth, J.G., and Barker, F., 1976, Rare-earth partitioning between hornblende and dacitic liquid and implications for genesis of trondhjemitictonalitic .magmas: Geology, v. 4, p. 534-536. 

Arth, J.G., and Hanson, G.N., 1972, Quartz diorites derived by partial 

melting of eclogite or amphibolite at mantle depths: Contributions 

to Mineralogy and Petrology, v. 3, p. 161-174. 
34 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

__ 1975, Geochemistry and origin of the early Precambrian crust 

of northeastern Minnesota: Geochimica et Cosmochimica Acta, v. 39, 

p. 325-362. 

Bankey, Viki, Brickey, Michael, and Kleinkopf, D.M., 1982, Principal 

facts for gravity stations in the Selway-Bitterroot Wilderness: U.S. 

Geological Survey Open-File Report 82-708, 17 p. 

Benoit, W.R., 1971, Vertical zoning and differentiation in granitic rocks, 

central Flint Creek Range, Montana: Missoula, Mont., University 

of Montana, M.S. thesis, 53 p. 

Bickford, M.E., Chase, R.B., Nelson, B.K., Shuster, R.D., and Arruda, 

E.C., 1981, U-Pb studies of zircon cores and overgrowths, and 

monazite: implications for age and petrogenesis of the northeastern 

Idaho batholith: Journal of Geology, v. 89, no. 4, p. 433-457. 

Bittner-Gaber, Enid, 1983, Geology of selected migmatite zones within 

the Bitterroot lobe of the Idaho batholith, Idaho and Montana: 

Moscow, Idaho, University of Idaho, Ph.D. thesis, 173 p. 

Bond, J.G., compiler, 1978, Geologic map of Idaho: Idaho Department 

of Lands, Bureau of Mines and Geology, with contributions from 

the U.S. Geological Survey, scale 1:500,000, 1 sheet. 

Bradfish, L.J., 1979, Petrogenesis of the Tea Cup granodiorite, Pinal 

County, Arizona: Tucson, Ariz., University of Arizona, M.S. thesis, 

160 p. 

Bradley, M.D., 1981, Geology of the northeastern Selway-Bitterroot 

Wilderness area, Idaho and Montana: Moscow, Idaho, University 

of Idaho, M.S. thesis, 51 p. 

Brickey, Michael, Bankey, Viki, and Kleinkopf, D.M., 1980, Principal 

facts for gravity stations of part of the Selway-Bitterroot Wilderness, 

Idaho and Montana: U.S. Geological Survey Open-File Report 

80-1241, 14 p. 

Brown, G.S., and Fyfe, W.S., 1970, The production of granitic melts 

during ultrametamorphism: Contributions to Mineralogy and 

Petrology, v. 28, p. 310-318. 

Buddington, A.F., 1959, Granite emplacement with special reference 

to North America: Geological Society of America Bulletin, v. 70, 

p. 671-747. 

Chappell, B.W., and White, A.J.R., 1974, Two contrasting granite types: 

Pacific Geology, v. 8, p. 173-174. 

Chase, R.B., Bickford, M.E., and Arruda, C.E., 1983, Tectonic implications of Tertiary intrusion and shearing within the Bitterroot dome, 

northeastern Idaho batholith: Journal of Geology, v. 91, p. 462-470. 

Chase, R.B., Bickford, M.E., and Tripp, S.E., 1978, Rb-Sr and U-Pb 

isotopic studies of the northeastern Idaho batholith and border zone: 

Geological Society of America Bulletin, v. 89, p. 1325-1334. 

Collins, W.J., Beams, S.D., White, A.J.R., and Chappell, B.W., 1982, 

Nature and origin of A-type granites with particular reference to 

southeastern Australia: Contributions to Mineralogy and Petrology, 

v. 80, p. 189-200. 

Coxe, B.W., Mosier, E.L., and McDougal, C.M., 1982, Magnetic tape 

containing analyses of rocks and stream sediments from the SelwayBitterroot Wilderness, Idaho County, Idaho, and Ravalli and 

Missoula Counties, Montana: U.S. Geological Survey Report USGSGD-82-011 (text) and USGS-GD-012 (magnetic tape); available only 

from U.S. Department of Commerce, National Technical Information Service, Springfield, VA 22161, as report PB-82 253 386. 

Coxe, B.W., and Toth, M.l., 1983a, Geochemical maps of the SelwayBitterroot Wilderness, Idaho County, Idaho, and Missoula and 

Ravalli Counties, Montana: U.S. Geological Survey Miscellaneous 

Field Studies Map MF-1495-C, scale 1:125,000. 

__ 1983b, Statistical tables, sample locality maps, and an explanation of data sets for samples from the Selway-Bitterroot Wilderness, 

Idaho County, Idaho, and Missoula and Ravalli Counties, Montana: 

U.S. Geological Survey Open-File Report 83-54, 94 p. 

Criss, R.E., and Fleck, R.J., 1983, Isotopic characteristics of granitic 

rocks from the north half of the Idaho batholith [abs.]: Geological 

Society of America Abstracts with Programs, v. 15, no. 6, p. 550. 

Ferguson, J.A., 1972, Fission track and K-Ar dates on the northeastern 

border zone of the Idaho batholith: Missoula, Mont., University of 

Montana, M.S. thesis, 32 p. 

__ 1975, Tectonic implications of some geochronometric data from 

the northeastern border zone of the Idaho batholith: Northwest 

Geology, v. 4, p. 53-58. 

Frey, F.A., Chappell, B.W., and Roy, S.D., 1978, Fractionation of rareearth elements in the Tuolumne Intrusive Series, Sierra Nevada 

batholith, California: Geology, v. 6, p. 239-242. 

Garmezy, Larry, 1983, Geology and geochronology of the southeast 

border of the Bitterroot dome: implications for structural evolution 

of the mylonite carapace: University Park, Pa., Pennsylvania State 

University, Ph.D. thesis, 269 p. 

Garmezy, L.G., and Greenwood, W.R., 1981, Mylonitization coincident 

with batholith emplacement: Bitterroot lobe, Idaho batholith [abs.]: 

Geological Society of America Abstracts with Programs, v. 13, 

p. 1981. 

Gastil, R. G., 1975, Plutonic zones in the Peninsular Ranges of southern 

California and northern Baja California: Geology, v. 3, p. 361-363. 

Greenwood, W.R., and Morrison, D.A., 1973, Reconnaissance geology 

of the Selway-Bitterroot Wilderness area: Idaho Bureau of Mines 

and Geology Pamphlet 154, 30 p. 

Grout, F.F., and Balk, R., 1934, Internal structures in the Boulder 

batholith: Geological Society of America Bulletin, v. 45, p. 877-896. 

Hall, B. S., 1980, Petrography and geochemistry of a part of the Painted 

Rocks pluton, Idaho County, Idaho: Cheney, Wash., Eastern 

Washington University, M.S. thesis, 94 p. 

Hamilton, Warren, 1963, Metamorphism in the Riggins region, western 

Idaho: U.S. Geological Survey Professional Paper 436, 95 p. 

__ 1976, Tectonic history of west-central Idaho [abs.]: Geological 

Society of America Abstracts with Programs, v. 8, p. 378-379. 

__ 1978, Mesozoic tectonics of the Western United States, in Howell, 

D.G., and McDougall, K.A., eds., Mesozoic paleogeography of the 

Western United States: Society of Economic Paleontologists and 

Mineralogists, Pacific Section, Pacific Coast Paleogeography Symposium, v. 2, p. 33-70. 

Hanson, G.N., 1980, Rare earth elements in petrogenetic studies of 

igneous systems: Annual Review of Earth and Planetary Sciences, 

v. 8, p. 371-406. 

Holdaway, M.J., 1971, Stability of andalusite and the aluminum silicate 

phase diagram: American Journal of Science, v. 271, p. 97-131. 

Holloway, C. D., 1980, Petrology of the Eocene volcanic sequence, Nez 

Perce and Blue Joint Creeks, Southern Bitterroot Mountains, 

Montana: Missoula, Mont., University of Montana, M.S. thesis, 

129 p. 

Hyndman, D.W., 1977, Mylonitic detachment zone and the Sapphire tectonic block, in Geological Society of America Field Guide no. 1: Rocky 

Mountain Section meeting, University of Montana, Missoula, 

Montana, p. 25-31. 

__ 1980, Bitterroot dome-Sapphire tectonic block, an example of a 

plutonic-core gneiss-dome complex with its detached suprastructure, 

in Crittenden, M.D., Jr., Coney, P.J., and Davis, G.H., eds., Cordilleran Metamorphic Core Complexes: Geological Society of America 

Memoir 153, p. 427-433. 

__ 1981a, Controls on source and depth of emplacement of granitic 

magmas: Geology, v. 9, p. 244-249. 

__ 198lb, Reply to comment on controls on source and depth of 

emplacement of granitic magma: Geology, v. 9, p. 503-504. 

__ 1983, The Idaho batholith and associated plutons, Idaho and 

western Montana, in Roddick, J.A., ed., Circum-Pacific plutonic terranes: Geological Society of America Memoir 159, p. 213-240. 

__ 1984, A petrographic and chemical section through the northern 

Idaho batholith: Journal of Geology, v. 92, p. 83-102. 

Johnson, B.R., 1975, Migmatites along the northern border of the Idaho 

batholith: Missoula, Mont., University of Montana, Ph.D. thesis, 
2. PETROLOGY AND ORIGIN OF THE BITTERROOT LOBE 35 

120 p. 

Koesterer, M.E., Bartel, A.J., Elsheimer, H.N., Baker, J.W., King, B.S., 

and Espos, L.F., 1982, Major-element XRF spectroscopy analyses 

from the Selway-Bitterroot Wilderness, Idaho County, Idaho, and 

Missoula and Ravalli Counties, Montana: U.S. Geological Survey 

Open-File Report 82-1094, 36 p. 

Kosinowski, M.H.F ., 1982, MSONRM, a fortran program for the 

improved version of mesonorm calculation: Computers and Geosciences, v. 8, no. 1, p. 11-26. 

Lee, D.E., and Christiansen, E.H., 1983, The granite problem as exposed 

in the southern Snake Range, Nevada: Contributions to Mineralogy 

and Petrology, v. 83, p. 99-116. 

Lindgren, W., 1900, The gold and silver veins of Silver City, De Lamar 

and other mining districts in Idaho: U.S. Geological Survey 20th 

Annual Report, part 3, p. 65-256. 

__ 1904, Geological reconnaissance of Bitterroot Range and Clearwater Mountains: U.S. Geological Survey Professional Paper 27, 

123 p. 

Loiselle, M.C., and Wones, D. R., 1981, Characteristics and origin of 

anorogenic granites [abs.]: Geological Society of America Abstracts 

with Programs, v. 11, no. 7, p. 468. 

Lund, K.I., 1980, Geology of the Whistling Pig pluton, Selway-Bitterroot 

Wilderness, Idaho: Boulder, Colo., University of Colorado, M.S. 

thesis, 115 p. 

Lund, K.I., Rehn, W.M., and Holloway, C.D., 1983, Geologic map of 

the Blue Joint Wilderness Study Area, Ravalli County, Montana, 

and the Blue Joint Roadless Area, Lemhi County, Idaho: U.S. 

Geological Survey Miscellaneous Field Studies Map MF-1557-B, scale 

1:50,000. 

Miller, C.F., 1981, Comment on controls on source and depth of emplacement of granitic magma: Geology, v. 9, p. 503. 

Miller, C.F., Stoddard, E.F., Bradfish, L.J., and Dollase, W.A., 1981, 

Composition of plutonic muscovite: genetic implications: Canadian 

Mineralogist, v. 19, p. 25-34. 

Motzer, W.E., 1981, Tertiary epizonal plutonic rocks of the SelwayBitterroot Wilderness, Idaho [abs.]: Geological Society of America 

Abstracts with Programs, v. 13, p. 210. 

Nelson, C.A., and Sylvester, A.G., 1971, Wall rock decarbonation and 

forcible emplacement of the Birch Creek pluton, southern White 

Mountains, California: Geological Society of America Bulletin, v. 82, 

p. 2891-2904. 

Nold, J.L., 1968, Geology of the northeastern border zone of the Idaho 

batholith, Montana and Idaho: Missoula, Mont., University of 

Montana, Ph.D. thesis, 159 p. 

Pawlowski, M.R., 1981, Geology and exploration geochemistry of the 

Magruder Corridor, Idaho County, Idaho: Moscow, Idaho, University of Idaho, M.S. thesis, 128 p. 

Pitcher, W.S., 1978, The anatomy of a batholith: Journal of the 

Geological Society of London, v. 135, p. 157-182. 

Pitcher, W.S., and Read, H.H., 1960, Early transverse dykes in the Main 

Donegal granite: Geological Magazine, v. 97, p. 53-61. 

Piwinskii, A.J., 1968, Experimental studies of igneous rock series: central Sierra Nevada batholith, California: Journal of Geology, v. 76, 

p. 548-570. 

Rehn, W.M., 1984, Petrology and geochemistry of the Blue Joint Area, 

Idaho and Montana: Golden, Colo., Colorado School of Mines, M.S. 

thesis, 202 p. 

Rehn, W.M., and Lund, K.I., 1981, Eocene extensional plutonism in the 

Idaho batholith region [abs.]: Geological Society of America Abstracts with Programs, v. 13, p. 536. 

Reid, R.R., and Greenwood, W.R., 1968, Multiple deformation and 

associated progressive polymetamorphism in the Beltian rocks north 

of the Idaho batholith, Idaho, U.S.A.: Geology of Precambrian, International Geological Congress, 23rd, Prague, Czechoslovakia, 1968, 

Proceedings, v. 4, p. 75-87. 

Robertson, J.K., and Wyllie, P.J., 1971, Rock-water systems with special 

reference to the water-deficient region: American Journal of Science, 

v. 271, p. 252-277. 

Roddick, J.A., and Armstrong, J.E., 1959, Relict dikes in the Coast 

Mountains near Vancouver, British Columbia: Journal of Geology, 

v. 67, p. 603-613. 

Streckeisen, A.L., chairman, and others, 1973, Plutonic rocks-classification and nomenclature recommended by the lUGS subcommission on the systematics of igneous rocks: Geotimes, v. 18, no. 10, 

p. 26-30. 

Sylvester, A.G., Oertel, G., Nelson, C.A., and Christie, J.M., 1978, 

Papoose Flat pluton: a granitic blister in the In yo Mountains, California: Geological Society of America Bulletin, v. 89, p. 1205-1219. 

Toth, M.I., 1983a, Reconnaissance geologic map of the Selway-Bitterroot 

Wilderness, Idaho County, Idaho, and Missoula and Ravalli Counties, Montana: U.S. Geological Survey Miscellaneous Field Studies 

Map MF-1495-B, scale 1:125,000. 

__ 1983b, Structure, petrochemistry, and origin of the Bear Creek 

and Paradise plutons, Bitterroot lobe of the Idaho batholith: Boulder, 

Colo., University of Colorado, Ph.D. thesis, 337 p. 

Toth, M.I., Coxe, B.W., Zilka, N.T., and Hamilton, M.M., 1983, Mineral 

resource potential of the Selway-Bitterroot Wilderness, Idaho 

County, Idaho, and Missoula and Ravalli Counties, Montana: U.S. 

Geological Survey Miscellaneous Field Studies Map MF-1495-A, scale 

1:125,000. 

Tuttle, O.F., and Bowen, N.L., 1958, Origin of granite in light of experimental studies in the system NaA1Si30 8-KA1Si30 8-Si02-H20: 

Geological Society of America Memoir 7 4, 153 p. 

Wehrenberg, J.P., 1972, Geology of the Lolo Peak area, northern Bitterroot Range, Montana: Northwest Geology, v. 7, p. 25-32. 

Williams, L.D., 1977, Petrology and petrography of a section across the 

Bitterroot lobe of the Idaho batholith: Missoula, Mont., University 

of Montana, Ph.D. thesis, 212 p. 

Winkler, H.G.F., 1979, Petrogenesis of metamorphic rocks, 5th edition: 

New York, Springer-Verlag, 348 p. 

Wyllie, P.J., Huang, W.-L., Stern, C.R., and Maaloe, Sven, 1976, Granitic 

magmas: possible and impossible sources, water contents, and 

crystallization sequences: Canadian Journal of Earth Sciences, v. 13, 

no. 8, p. 1007-1019. 

3. STRUCTURAL GEOLOGY AND PETROLOGY OF A PART OF THE 

BITTERROOT LOBE OF THE IDAHO BATHOLITH 

By RoLLAND R. REID1 

ABSTRACT 

Mapping of an area about 15 km north-south by 85 km east-west across 

the north-central part of the Bitterroot lobe of the Idaho batholith shows 

that the structure of the lobe is asymmetrical. Flow foliation along the 

western margin strikes northwesterly and dips steeply. Within the lobe, 

migmatite screens dip varying amounts and strike at predominantly 

az 090°, near the primary flow lineation trend. Foliation becomes progressively more intense to the east and drapes over a north-south arch 

near the Bitterroot mylonite front. Primary flow foliation is strongly 

overprinted to destroyed by a secondary flow foliation and lineation 

trending near az 290° in a wedge-shaped area that widens to the east. 

The disposition of secondary lineation suggests an original batholith 

thickness in excess of 30 km. Features of gneissic foliation, nonpenetrative blastomylonite, and mylonite are consistent in their indication of 

cover-rock movement toward az 110° throughout three structural 

phases. 

The asymmetric structure began to develop with the rise of tonalite 

magma along a north- to northwest-trending fracture set near the 

western margin of the lobe. This was followed by voluminous mixed 

granite-granodiorite magma, with minor anorthosite and tonalite in 

deeper parts, which spread to the east in a sheet perhaps controlled by 

a low-angle thrust near its base. The granitic part of the massif is vertically zoned, from calcic granodiorite near its base to granite at the 

top. During emplacement, primary flow foliation formed, and migmatite 

screens were rotated into an east-west trend. After consolidation, shearing of still-hot rock created secondary foliation, later blastomylonite, 

and still later mylonite in episodes that continued during arching and 

rise to higher crustal levels. Some east-west extension occurred in conjunction with thrusting during the later stages of cooling. 

INTRODUCTION 

The rocks discussed in this paper form part of the 

Bitterroot lobe of the Idaho batholith (fig. 3.1). They occur 

in north-central Idaho and extend into southwestern 

Montana, in a region with local relief up to 2,000 m (6,000 

ft). Quartzite and calc-silicate gneiss provisionally correlated with units of the Belt Supergroup (its Ravalli 

1Department of Geology, University of Idaho, Moscow, Idaho. 

Group and Wallace Formation, respectively) are the country rocks into which the batholith was intruded. The rocks 

in the western part of the Bitterroot lobe broadly comprise tonalite cut by later granitic rocks, which become 

major to the east. The igneous rock nomenclature used 

herein is that of Streckeisen and others (1973). 

This paper stems from work done on the U.S. Geological 

Survey project (1978-1981) to assess the mineral potential of the Selway-Bitterroot Wilderness Area. I mapped 

the structure in the field and made structural geometric 

analyses and petrographic studies. Terry Hayden mapped 

in my stead during one summer after I broke a leg. I 

studied a total of 1,064 thin sections and did 619 modal 

analyses. Modes were measured in thin section for rocks 

of grain size 2 mm or smaller. For medium-grained rocks 

coarser than 2 mm, modes were measured both in thin 

section and on stained slabs and the results combined. For 

coarsely porphyritic rocks, the percentage of phenocrysts 

was measured at the outcrop using a knotted string 

devised by Bill Greenwood and the results combined with 

thin-section and stained-slab analyses of the matrix. 

Anywhere from 200 to 1,600 points per specimen were 

counted, depending on grain size and inhomogeneity of 

the rocks. Plagioclase composition was determined in thin 

section using the standard a-normal method. 

Detailed data from this work are given in an open-file 

report available from the U.S. Geological Survey (Reid, 

1984). This paper is a shortened version. 

PREVIOUS WORK 

Reconnaissance investigations constituted the early 

work in the Bitterroot lobe of the Idaho batholith and its 

environs. Lindgren (1904) was the earliest worker in the 

area; he wrote of a Mesozoic quartz monzonite mass with 

an eastern schistose slope probably due to a great fault. 

37 
38 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

Qal 

Tb 

Tv 

Tg 

TKg 

Kt 

Kg 

Ymi 

Ym 

/ 

TKg--o 

0 

EXPLANATION 

Alluvium (Quaternary) 

Basalt (Miocene) 

Volcanic rocks (Eocene) 

Epizonal granite (Eocene) 

20 KILOMETERS 

Mesozonal granitic rocks of the Bitterroot 

lobe (Paleocene and Cretaceous) 

Mesozonal tonalite, diorite and quartz 

diorite (Cretaceous) 

Granitic rocks of the Atlanta lobe of the 

Idaho batholith (Cretaceous) 

Migmatite (Middle Proterozoic?) 

Metamorphic rocks (Middle Proterozoic?) 

Major fault 

FIGURE 3.1.-Generalized geologic map of the Bitterroot lobe of the 

Idaho batholith. Dashed lines show approximate boundaries of the 

field of secondary strain, which is the area affected by metamorphic core complex phenomena. A-A', line of section of figure 3.2, 

and the approximate line of quadrangles mapped, shown in figure 

3.2. Other geologic features after Toth (1983). 

Ross (1928) showed that the Idaho batholith is postTriassic and pre-Tertiary and is cut by younger granitic 

masses. Langton (1935) studied the northeastern Idaho 

batholith, showing two deformations in the country rock 

and thrusts related to the emplacement of the batholith. 

Later studies became more detailed. Anderson (1952) 

did further work on the concept of multiple emplacement 

of the batholith, a theme developed in the earlier reconnaissance of Ross (1928). Ross (1952) developed a 

granitization hypothesis to explain the eastern frontal 

gneiss of the Bitterroot Mountains, inferring that emanations from the batholith converted sedimentary rock to 

gneiss. Reconnaissance modal and chemical compositional 

studies by Larsen and Schmidt (1958) showed a considerable variety of rock types (eleven) in the Idaho 

batholith. Reid (1959) discovered three sets of synmetamorphic folds, predating emplacement of the batholith, 

in wall rocks on its west side. Leischner (1959) studied 

the batholith in the vicinity of Lolo Hot Springs, Montana, 

and found an intrusive sequence from early diorite 

through quartz diorite to quartz monzonite. Ross (1965) 

found that the main interior mass of the Idaho batholith 

is largely granodiorite and quartz monzonite and is surrounded by a narrow discontinuous roof and border zone 

of quartz diorite and quartz diorite gneiss. Thrusting 

occurred away from the batholith interior on the eastern 

and western margins, in part preceding and in part following batholith emplacement. Chase (1968) studied country 

rock in the northeast contact zone of the Idaho batholith 

and found metasedimentary units regionally metamorphosed in three phases of synmetamorphic penetrative 

deformation before emplacement of the Idaho batholith, 

the deformation phases perhaps comparable to those 

discovered by Reid (1959). Morrison (1968) also found 

three prebatholithic metamorphic episodes in the Lochsa 

area northwest of the batholith, plus a fourth episode synchronous with batholith emplacement. Further study in 

the country rock northeast of the batholith (Berg, 1969) 

disclosed 20 concordant anorthosite bodies, dominantly 

labradorite, within the gneiss. 

Beginning broader work, the Idaho Bureau of Mines and 

Geology studied the reconnaissance geology and mineral 

potential of the Selway-Bitterroot Wilderness Area, which 

includes much of the Bitterroot lobe of the Idaho batholith 

(Greenwood and Morrison, 1973). In a regional study, 

Ryder and Scholten (1973) showed that the thick 

accumulations of syntectonic conglomerate of the Beaverhead Formation in Montana were derived from the northeast margin of the Idaho batholith in Early Cretaceous 

to middle Paleocene time. Armstrong (197 4) discussed 

relations between tectonic activity and magmatism in the 

cordillera; this work delineates the broad temporal and 

structural framework within which the batholith was 

emplaced. Nold (1974) showed that the Belt Supergroup 
3. STRUCTURAL GEOLOGY AND PETROLOGY OF A PART OF THE BITTERROOT LOBE 39 

and associated orthoamphibolite and orthogneiss units 

were deformed at least four times before they were injected by the quartz diorite to granite of the Idaho 

batholith. This interpretation corresponds well with 

earlier results obtained northwest of the batholith by Morrison (1968) and west of the batholith by Reid (1959). In 

a detailed structural analysis, Cheney (1975) showed that 

anorthosite bodies in the country rock near the northeast 

margin of the batholith have experienced three structural 

events. They are thus closer in age to the country rock 

than to the batholith. Armstrong (1975a) suggested the 

name "Bitterroot lobe" for the northern part of the Idaho 

batholith. Williams (1977) showed that the northern 

Bitterroot lobe consists largely of foliated, medium- to 

coarse-grained, megacrystic, two-mica granite, intruded 

by granodiorite. Wiswall (1979) suggested that the base 

of the Bitterroot lobe was exposed south of my map area, 

where gneiss and migmatite are overlain by granitic rock 

of the batholith. Bradley (1981) studied a gneissic tonalitic 

phase of the Bitterroot lobe along its northeastern border. 

This is a part of Ross' (1965) roof and border zone of 

quartz diorite gneiss. Bradley showed the gneissic tonalite 

to be earlier than the granitic phase of the Bitterroot lobe. 

Toth (1981) defined two major plutons, the Bear Creek 

pluton and the younger Cub Creek pluton, within the main 

Bitterroot lobe. The Bear Creek pluton is a mediumgrained granite having varied amounts of poorly defined 

potassium feldspar phenocrysts and 3 to 5 percent biotite, 

3 to 5 percent hornblende, and as much as 3 percent 

sphene. These units may correspond to those described 

by Williams (1977). The relevant geologic rr..apping was 

published by Toth in 1983. 

As a part of the continuing more detailed study of the 

Bitterroot lobe, work has been done on the extensive 

migmatite within and peripheral to the lobe. Johnson 

(1975) inferred that the migmatite formed by injection, 

metamorphic differentiation, and anatexis. Bittner-Gaber 

(1983) showed that the migmatite screens within the lobe 

are oriented largely east-west (p. 22) and were derived 

from the wall rocks and (or) roof rocks (p. 114). Their eastwest orientation is due to magma flow, postulated to be 

from west to east. The migmatite protolith has been 

strongly injected by magma of several types, and the 

resulting migmatite has in many places been intricately 

folded. 

The continuing detailed studies have resulted in the 

interpretation of a part of the Bitterroot lobe as a metamorphic core complex (Chase and Hyndman, 1977; Hyndman, 1980). Many workers, too numerous ~~o list here, 

have concentrated on the eastern mylonitic detachment 

zone (hearkening back to Lindgren's idea of a major fault 

zone). Core-complex effects extend some distance west 

within the lobe. Toth (1981), for example, showed that 

both the Bear Creek and Cub Creek plutons have 

moderately to well-developed flow foliation that decreases 

in dip upward, suggesting increasing proximity to the roof 

of the batholith. Primary mineral lineations, defined by 

igneous fabric elements, and small fold axes trend at 

az 290° (down-plunge direction) and plunge at shallow 

angles. Secondary mineral lineations, defined by crystalloblastic fabric elements, trend at az 290° and az 110° 

and also plunge at shallow angles, with the az 110° element increasing in abundance toward the eastern bordering mylonite front. Garmezy and Sutter (1983) have done 

the most detailed work on the mylonite zone. Mylonitization began at about 45.5 Ma at a depth probably greater 

than 10 km and continued for about 2 m.y. during a period 

of rapid uplift. Mylonitic fabric indicates eastward tectonic 

transport of the hanging wall throughout the 500- to 

1,000-m-thick mylonite zone. The presence of ductile 

normal faults in this zone, which have a sense of displacement parallel to the mylonitic lineation, suggests that 

mylonitization occurred in an extensional environment in 

which a principal extension direction was aligned at 

az 110°. 

The inferred movement direction of the cover rock 

above the detachment zone has been controversial, but 

increasing evidence is coming to support eastward transport. This view was presented early by Hyndman (1980) 

and has lately been presented with convincing evidence 

for the Okanogan dome of north-central Washington 

(Goodge, 1983). 

Much radiometric dating has been done on the Bitterroot lobe and its country rock. Armstrong (1975b) compiled voluminous data and noted (p. 32) that the ages for 

the Bitterroot lobe are all less than 70 Ma. Chase and 

others (1978) arrived at an age of 66 Ma for the mainstage batholithic emplacement and an age of 46 to 39 Ma 

for a late thermal-plutonic event. Bickford and others 

(1981), in a very detailed study, showed the following intrusive history, based on lower intercept ages (p. 453): 

quartz diorite gneiss at about 73 Ma, medium-grained 

granite at about 55 Ma, and cataclastic granite at 49 to 

46 Ma. The age for cataclastic granite is close to the time 

of beginning of mylonitization, given by Garmezy and 

Sutter (1983) at 45.5 Ma. Upper intercept ages are given 

by Bickford and others (1981, p. 454) in the range 1,700 

to 2,300 Ma. These ages are interpreted as due to countryrock contributions either in partial melting or in contamination. 

WALL ROCK SOUTHWEST OF 

THE BITTERROOT LOBE 

STRUCTURE 

The southwestern contact of the Bitterroot lobe of the 

Idaho batholith consists of metasedimentary and 
40 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

metaigneous rock marginal to quartz diorite and tonalite 

of the batholith. The northern part of this area contains 

granite and granodiorite injected into the tonalite (figs. 

3.1, 3.2). 

115° 

47°r------------------.-------------------. 

0 10 20 KILOMETERS 

Bitterroot front mylonite zone 

Jeanette Mountain mylonite~~ 

123456,~\, A/_,,',...),.,/)/~,./'/~ 7,' ..... _:-~:_,,~A 

EXPLANATION 

TKg Mesozonal granitic rocks of the Bitterroot 

lobe (Paleocene and Cretaceous) 

gr Granitic rocks, part of TKg 

to Older tonalite, part of TKg 

Generalized contact between granitic rocks 

and older tonalite 

Southern limit of secondary strain 

Direction of generally 30° -plunging flow 

lineation, averaged for each quadrangle 

Horizontal flow lineation at the Bitterroot 

arch 

Extension joint trends; the joints average 

to be near perpendicular to flow lineation 

METASEDIMENTARY ROCKS 

Calc-silicate gneiss is the principal country rock exposed 

at the southwestern margin of the lobe, whereas biotite 

quartzite is abundant in migmatite screens inside the 

batholith. The calc-silicate gneiss is deformed in folds of 

steep but varied plunge. Schistosity lies parallel or subparallel to the igneous contact and to original bedding in 

the calc-silicate gneiss; this schistosity has been deformed 

in small passive-flow folds showing both clockwise and 

counterclockwise rotation somewhat randomly through 

the outcrops. The axial planes of these small folds are 

everywhere parallel to the foliation of the enclosing rocks, 

regardless of the orientation of the major foliation. 

Mineral lineation parallels the minor fold axes and consists of elongate biotite clots, diopside and hornblende 

prisms, and quartz rods. 

Pegmatite, alaskite, and aplite have been injected along 

the foliation and are folded with it; in places, the igneous 

material cuts both foliation and minor folds and is itself 

folded. The pegmatite is partly in lozenge boudins, which 

show separations of as much as 1.5 m along the layering. 

The small folds mentioned above are all conical. A set 

of small isoclinal folds, seen in a few places, predates the 

conical folds, as the conical folds deform the foliation 

formed by their axial-plane schistosity. The isoclinal-fold 

foliation stands vertical, and the fold axes plunge 90°. The 

isoclinal folds, the oldest folds recognized in the rock, are 

assigned to a deformation phase D1, and the conical folds 

are assigned to D2. Sequential measurement of the orientation of small conical-fold axes downward along their 

plunge shows that the bearing and plunge change. Thus, 

the folds may be characterized as helical or screwlike. 

The axial planes of the conical folds are deformed in a 

larger synform. The concentric style and low-angle plunge 

of this fold suggest that it is distinctly younger and should 

be assigned a D3 age. Low-angle crinkle lineation in the 

rock is probably related to this fold set. 

In addition to calc-silicate gneiss, three other wall rock 

types occur: biotite quartzite (predominant), biotite gneiss, 

and biotite-muscovite schist (minor). Rare calc-silicate 

FIGURE 3.2.-0utline map showing the thirteen 7.5-minute quadrangles 

entirely or partly mapped across the Bitterroot lobe of the Idaho 

batholith, and along section line A-A' of figure 3.1. Quadrangle 

names: 1, Huckleberry Butte; 2, Greenside Butte; 3, Fish Lake; 4, 

McConnell Mountain; 5, Hungry Rock; 6, Cedar Ridge; 7, Jeanette 

Mountain; 8, Blodgett Mountain; 9, Printz Ridge; 10, Chimney Peak; 

11, Fenn Mountain; 12, Tenmile Lake; 13, Ward Mountain. These 

quadrangle names are used for geographic reference throughout the 

report. Flow directions in granitic rocks, shown as average for each 

quadrangle by solid arrows, form a fan-shaped array. The chloritecoated, microbrecciated shear joints in the several quadrangles are 

believed to be tension fractures initially perpendicular to flow lines 

and reactivated in younger movements. 
3. STRUCTURAL GEOLOGY AND PETROLOGY OF A PART OF THE BITTERROOT LOBE 41 

layers are generally thin but locally reach 65 min thickness. The rock structure was analyzed in a profile section 

perpendicular to the 30° -southeast plunge. The overall 

structure is a nearly recumbent, somewhat refolded 

isoclinal fold. The D1 minor folds are isoclinal and have 

axial-plane schistosity. Amphibolite and tonalite sheets 

occur along the schistosity and are deformed by the D1 

isoclinal folds. The rocks were injected synkinematically 

by concordant and discordant pegmatites, which are 

folded and schistose parallel to the external schistosity. 

Orbicules are flattened in the plane of the schistosity. 

The D2 folds developed a secondary axial-plane 

schistosity (S2) largely restricted to their hinges. Lineated 

biotite and muscovite parallel the fold axes. Boudins produced by stretching in D1 are locally folded in D2; crenulation folds (late D2 or Ds) kink the earlier schistosity and 

the igneous-rock sheets along it. Biotite and muscovite 

are further lineated parallel to kink-fold axes. Late shears 

of diverse orientation contajn sheared-out, streaked 

muscovite and annealed slickensides, all trending nearly 

due east. 

Poles to foliation/bedding locally show a diffuse girdle 

indicative of folding about an axis oriented to the 

southeast; early fold axes are of scattered orientation, as 

are crenulation fold axes. Mineral lineation trends are also 

widely scattered. 

METAIGNEOUS ROCKS 

Tonalitic gneiss occurs as concordant sheets as much 

as 1 or 2 km thick in the country rock southwest of the 

Bitterroot lobe. Attitudes on foliation in the gneiss 

indicate steep dips and a considerable range of strikes 

forming conical folds. Axes of minor folds in the gneiss 

plunge vertically or nearly so, but the axial planes display 

widely varied trends. Mineral lineation consisting of 

elongate biotite streaks and clots commonly parallels the 

axes of the minor folds. The folds show both clockwise 

and counterclockwise rotation more or less randomly 

through the outcrops. Concordant veins of trondhjemite 

were injected along the foliation and are folded with the 

gneiss. 

PETROGRAPHY 

The wall rocks comprise (1) pelitic quartzite and schist, 

from quartz sand and shale, and (2) calc-silieate quartzite 

and gneiss, from dolomitic quartzite and shale-dolomite 

mixtures. 

In the pelitic schist, quartz is present in two populations. 

The first consists of platy grains 2 to 4 mm across that 

are oriented parallel to the schistosity and elongate 

parallel to mineral lineations within it. These grains 

enclose oriented biotite, muscovite, and fibrolite, and are 

cut by a-c extension fractures spaced at intervals of 0.3 to 

0. 7 mm. The second population consists of irregular grains 

0.3 to 0.5 mm across. Most grains in both populations 

show undulatory extinction. 

Plagioclase occurs as unzoned granoblastic grains 0.5 to 

1 mm in diameter and has compositions ranging from 

An2o to Anso· Clay alteration is minor but ubiquitous. 

Biotite is lepidoblastic as grains 0.3 to 2 mm in diameter, 

most of them stretched and elongate or forming clusters 

streaked out in a lineation. Pennine alteration varies from 

weak to strong. Some biotite is in vermiform intergrowths 

with quartz. Muscovite occurs as flakes 0.1 to 0.3 mm 

across, replacing biotite and mostly of random orientation. Some muscovite is also in vermiform intergrowths 

with quartz. Sillimanite is in only a few of the rocks 

sampled. Representative prisms are 0.03 mm by 0.15 mm 

and oriented parallel to other mineral lineations in the 

rock. Prisms are commonly broken and stretched parallel 

to their length. 

Microcline grains are platy and oriented parallel to the 

foliation; within the foliation, they are elongate (for example, 1.5 mm by 4 mm as seen in an a-b section) parallel 

to other mineral lineations. Microcline encloses small, 

oriented biotite flakes. Sillimanite is mostly absent in 

microcline grains, even where it is found elsewhere in the 

same rock, and the few grains that are found in microcline 

are mere shreds that appear to be undergoing replacement. The microcline is cut by a-c extension fractures 

spaced at 0.1- to 0.2-mm intervals. Stained slabs show 

microcline to be concentrated in certain folia. 

Accessory minerals include magnetite (elongate grains 

parallel to other mineral lineations), zircon, baddeleyite, 

thorite, apatite, epidote, and schorlite. 

In calc-silicate quartzite and gneiss, quartz is present 

as granoblastic grains 0.5 mm across and as elongate flattened blebs parallel to other mineral lineation. Plagioclase 

is present as unzoned, granoblastic, 1-mm grains, some 

of which are as elongate plates parallel to other mineral 

lineation. Their composition ranges from An25 to An43· 

Hornblende and diopside form linea ted prisms lying in the 

foliation. Scapolite is in one specimen as granoblastic 

grains 0.5 mm across. 

TONALITE OF THE 

SOUTHWESTERN BITTERROOT LOBE 

STRUCTURE 

A quartz diorite to tonalite intrusive unit invades the 

country rock of the southwestern batholith margin (fig. 

3.2). The intrusive unit shows a strong flow foliation, 

rather swirled and of varied attitudes, that generally has 

a northerly trend and a westerly dip of about 70°. Flow 

lineation consisting of oriented hornblende prisms or 
42 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

biotite grains and clusters plunges down the dip of the 

flow foliation. Rare, dark inclusions are drawn out into 

elongate spindles parallel to the hornblende lineation and 

flattened in the plane of the flow foliation. 

Pegmatite, alaskite, and aplite dikes of shallow to steep 

dip cut the intrusion. The shallow-dipping sheets, mostly 

trondhjemitic, have a northerly trend and easterly dip 

of about 20°, but the steep dikes show no preferred 

orientation. 

At some places, the flow foliation is complexly swirled 

in small folds showing both clockwise and counterclockwise rotation randomly through the outcrop. 

Tonalitic rocks, including pegmatite and alaskite, are 

locally cut by cataclastic shear surfaces on which slickensides are "frozen" or recrystallized, resulting in microrodding and grooving. These features are best shown in 

the pegmatitic rocks, on both near-horizontal and steeply dipping surfaces. 

Migmatite is abundant both here and in the granitic 

rocks described below. It has been discussed by BittnerGaber (1983). 

PETROGRAPHY 

Tonalite occurs in major amounts in the western Bitterroot lobe of the Idaho batholith (fig. 3.1). It is coarsely 

foliated and medium grained, and has a varied flow lineation that is strongest near the margin of the intrusion. 

The flow lineation consists mostly of hornblende prisms 

plunging steeply down the dip of the foliation. 

Hornblende is relatively abundant in the western part 

of the intrusion, but less abundant to absent toward the 

eastern part. It occurs in 2-mm prisms, partly altered to 

biotite. 

Biotite occurs in 0.5- to 2-mm flakes in crude lepidoblastic array in the flow foliation. Grains display olivebrown to strong reddish-brown pleochroism and minor 

pennine-epidote-sericite alteration. Some grains are bent 

cataclastically. 

Plagioclase (An27-37) is euhedral to subhedral in 2- to 

3-mm grains, which have albite twinning and are mostly 

unzoned. Anorthite content of the plagioclase shows no 

systematic areal variation. A few plagioclase grains show 

normal zoning from cores of An2s to rims of An25· Alteration products are minor sericite and epidote, and a few 

grains have bent albite-twin lamellae. 

Quartz occurs in 0.5- to 2-mm grains, most showing undulatory extinction and many having sutured quartzquartz grain boundaries. A trace to 5 percent microcline 

is in a few specimens. Accessory minerals include allanite, 

baddeleyite, magnetite, sphene, thorite, and zircon. 

Rare, dark, spindled and lineated xenoliths of diorite 

in the tonalite perhaps represent cognate inclusions. They 

have a fairly uniform grain size of 0.5 mm and consist of 

hornblende, biotite, and unzoned granoblastic plagioclase 

having a composition of An37· Accessory minerals are 

apatite, baddeleyite, magnetite, and zircon. 

Tonalitic gneiss has intruded the country rock near the 

batholith in concordant sheets at least several hundred 

meters thick. Its mineralogy is like that of the batholithic 

tonalite, from which the rock differs only in its gneissic 

structure. 

Quartz occurs partly as small, equant grains and partly as large, flattened, elongate platelets, typically 1 mm 

by 3 mm by 6 mm, oriented with their long direction 

parallel to the other lineation in the rock. Prisms of hornblende and plagioclase (An37) 4 to 6 mm long also lie 

parallel to the lineation. Biotite in 3- to 4-mm flakes is 

partly stretched into elongate grains and grain clusters 

that parallel the lineation as well. In quartz-rich varieties 

of the tonalite, minor sillimanite occurs in fibers parallel 

to the lineation, and plagioclase is more sodic (An27) than 

in quartz-poor varieties. Muscovite also occurs in some 

specimens. Accessory minerals include apatite, baddeleyite, magnetite, sphene, thorite, and zircon. 

Trondhjemite in the form of lit-par-lit veinlets has intruded the tonalitic gneiss parallel to its foliation. Quartz 

in these veinlets has mild undulatory extinction and forms 

elongate platelets 4-6 mm long that are arranged in a 

crude linear array. Elongate grains and grain aggregates 

of biotite also form a linear array parallel to that of the 

quartz platelets. Plagioclase in 2- to 3-mm grains is more 

sodic (An25) than plagioclase in the enclosing tonalitic 

gneiss. Plagioclase grains are mostly granoblastic and unzoned, but some patchy, irregular relict(?) zoning is seen. 

A trace of muscovite occurs, together ·with chlorite, 

magnetite, thorite, and zircon. 

The tonalite and neighboring rocks are cut by dikes and 

sheets of pegmatite, alaskite, and aplite. Alaskite sheets 

are most common and are representative of the mineralogy of the pegmatite suite; they are mostly trondhjemitic, 

although three sheets of granitic alaskite were found. 

Trondhjemite plagioclase (An31) occurs in 2-mm prisms 

and is mostly unzoned, but a few grains show some 

oscillatory normal zoning. Quartz and biotite are as 1- to 

2-mm grains, and accessories include apatite, epidote, 

muscovite, pennine, and thorite. 

In the granitic alaskite, microcline occurs in grains more 

than 8 mm in diameter. Plagioclase (An2o) is in grains as 

much as 6 mm in diameter, which are unzoned and partly 

myrmekitic where in contact with microcline grains. 

Plagioclase is partly altered to muscovite. Quartz shows 

some undulatory extinction and sutured grain boundaries. 

In both the trondhjemite and the granitic alaskite, some 

shear reduction of quartz and feldspar occurs in cataclastic 

zones with partial strain recovery. Grain size in the 

cataclastic zones averages 0.3 mm. Many cataclastically 
3. STRUCTURAL GEOLOGY AND PETROLOGY OF A PART OF THE BITTERROOT LOBE 43 

reduced grains are free of undulatory extinction, and the 

sutured grain boundaries are annealed. 

GRANITE, GRANODIORITE, AND 

RELATED ROCKS OF THE 

WESTERN, CENTRAL, AND EASTERN SECTORS 

Granite, granodiorite, and smaller amounts of quartz 

monzonite, quartz monzodiorite, tonalite, diorite, and 

anorthosite crop out in a row of nine 7.5-minute 

quadrangles approximately along latitude 46°15' N., 

representing a continuous east-west transect across the 

Bitterroot lobe of the Idaho batholith (figs. 3.1, 3.2). No 

significant modal variations in the main rock types occur 

over this transect, and granite and granodiorite are present in about equal amounts. For convenience and because 

of structural variations, however, the rocks are divided 

for discussion into western, central, and eastern sectors. 

All of these rocks in the three sectors belong to the Bear 

Creek pluton as defined by Toth (1981). 

STRUCTURE 

The western sector covers all or parts of the following 

7.5-minute quadrangles: Fenn Mountain, Chimney Peak, 

Huckleberry Butte, Greenside Butte, and Fish Lake (fig. 

3.2). The sector is discussed in terms of two zones, northern and southern, whose structures are different. 

In this area granite and granodiorite cut the previously discussed tonalite unit of the southwestern batholth (fig. 

3.1). Xenoliths of tonalite occur in the granite. 

The granitic rock of the southern zone has a strong 

primary flow foliation parallel to the tonalite contact. The 

near-vertical orientation of flow lineation is comparable 

to that of the lineation in the tonalite. Thus, a degree of 

flow coupling may have existed between the granitic 

magma and its wall rock during the emplacement of the 

granite. Orientations of pegmatite dikes and sheets in the 

granitic rock are different from those in the tonalite, 

indicating a different system of flow and stretching. 

Farther into the granitic pluton and across a subtle but 

major zone of structural transition, the Greenside Butte 

area is representative of the northern zone of the western 

sector. The rocks here are principally granite and 

granodiorite, which show local flow foliation formed mostly by biotite or microcline bands and less commonly by 

biotite-rich schlieren. The orientation of the flow foliation, 

shown in figure 3.3, is characteristic of both the western 

and central sectors; to a lesser extent, it is also characteristic of the eastern sector. Poles to the flow foliation, 

including some blastomylonite discussed below, lie in a 

diffuse girdle whose axis is oriented west-northwest. 

Reconnaissance mapping suggests that the various 

igneous rocks are rather thoroughly intermixed. In a 

N 

Percent 

6 

per 1 percent 4 

area 2 

FIGURE 3.3.-Poles to 125 foliation planes from the area mapped in the 

Bitterroot lobe of the Idaho batholith. The data define a diffuse girdle 

centering on the dashed great circle. Lineation data (. ). The circled 

point is the pole to the girdle; it lies reasonably close to the average 

plunge of the lineation. Data are plotted on the lower hemisphere 

of the Schmidt equal area stereonet. 

representative traverse for 11 km along a ridge in wellexposed rocks, different rock types were encountered in 

the following order: granite; granite; granodiorite and 

foliated tonalite at the same outcrop; granite; migmatite; 

granite; migmatite; granite and granodiorite at the same 

outcrop; granodiorite; porphyritic granodiorite; tonalite; 

foliated, porphyritic granodiorite, granite, migmatite; 

granodiorite; granodiorite. The rocks, all medium grained 

and gray, look much alike and are not recognized as 

separate varieties in the field; no unit contacts were 

discerned other than those of the migmatite zones. This 

traverse is judged to be indicative of the degree of variation of granitic rock types in the entire mapped area of 

the Bitterroot lobe, because samples from less well exposed parts of the area show similar variation. 

Flow foliation is sporadically present in the rocks of the 

northern zone; about one-third of the specimens collected 

there are coarsely foliated and the rest massive. The two 

types are irregularly mixed over the area among all the 

rock types present. Mostly, the flow foliation is primary, 

expressed not only by rather coarse banding but also by 

aligned schlieren and platy xenoliths. 
44 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

Some of the plutonic rocks of the northern zone show 

subtle foliation expressed principally by flattened quartz 

and feldspar grains. This style of foliation is herein termed 

''secondary foliation''; generally too weakly developed in 

the northern zone to be recognized in weathered outcrops, 

it has been seen principally in stained slabs. The secondary foliation is present in several specimens of granitic 

rock and pegmatite for which a primary foliation was 

measured at the outcrop; therefore, the secondary foliation is at least partly parallel to primary foliation inasmuch 

as it is parallel to it in the hand specimens. This is also 

indicated by figure 3.3, which shows the orientations of 

both foliation sets and which suggests that they formed 

in response to a common stress field. Poles to pegmatite 

dikes and to the foliation in them lie generally in the girdle 

of figure 3.3, and rodding in them plunges at a small angle 

to the west-northwest, parallel to lineation in the granitic 

rock. 

The central sector covers all or parts of the following 

7 .5-minute quadrangle sheets: McConnell Mountain, 

Hungry Rock, and Cedar Ridge (fig. 3.2). This area is 

underlain almost entirely by intermixed granite and 

granodiorite in roughly equal amounts and lies entirely 

within the zone of secondary fabric defined by the wedgeshaped area in figure 3.1. In the western part of this area 

(McConnell Mountain quadrangle), the plutonic rock is 

about two-thirds massive and one-third foliated and 

lineated. In the central part (Hungry Rock quadrangle), 

only about half the rock is massive, and in the eastern part 

(Cedar Ridge quadrangle) massive rock makes up only 

about one-third, the remaining two-thirds being foliated 

and lineated. Foliated granite and granodiorite are 

distributed more or less randomly through the massive 

granite and granodiorite. Both massive and foliated rocks 

are locally porphyritic, with K-feldspar phenocrysts 

making up from 1 to 20 percent of the rock. 

Foliation is expressed in various ways: zones as much 

as half a meter thick of oriented biotite, which may be 

the sole expression of foliation in otherwise massive rock; 

parallel biotite grains; biotite streaks; schlieren; platy 

xenoliths oriented parallel in flow; or parallel K-feldspar 

phenocrysts, the resultant foliation plane of which dips 

more steeply than foliation due to biotite in the same outcrops. The feldspar and biotite planes strongly resemble 

the C and S surfaces defined by Berthe and others (1979, 

p. 33). In many places, a northwest-trending secondary 

foliation expressed principally by flattened quartz and 

feldspar grains crosses and even destroys the primary 

foliation, along whose former trace may remain only some 

schlieren. Secondary lineation is expressed in schlieren 

by linear biotite streaks and in xenoliths of biotite quartzite and schist by minor fold axes and linear biotite arrays. 

Both primary and secondary flow foliation and lineation 

are plotted in figure 3.3. 

In the Cedar Ridge quadrangle, a concordant sheet of 

biotite-hornblende quartz diorite as much as several 

hundred meters thick has been injected more or less along 

the gneissic foliation and extends for several kilometers 

along the strike. Its outer parts are foliated and metamorphosed to about the same degree as the enclosing gneissic 

granite; its central parts are more massive, although the 

fabric is still predominantly metamorphic. 

Injected sheets and dikes of fine-grained biotite granite, 

granodiorite, and tonalite 10 to 70 em thick, both parallel 

and crosscut the foliation in gneissic granitic rocks. These 

sheets and dikes are cut in turn by pegmatite, and the 

secondary foliation cuts all the rocks. 

The eastern sector, which is entirely within the field of 

secondary strain (fig. 3.1), is underlain by commingled 

granite, granodiorite, and minor tonalite and anorthosite. 

Parts of the following 7.5-minute quadrangles were 

mapped: Jeanette Mountain, Blodgett Mountain, Tenmile 

Lake, Printz Ridge, and Ward Mountain (fig. 3.2). Few 

contacts were found, because the various rock types look 

much alike in the field. As in other parts of the pluton 

mapped, granite and granodiorite occur in roughly equal 

amounts and are more or less intermixed. About 20 percent of the rock is porphyritic, more than in the areas to 

the west. The various textural types are commingled in 

random fashion. 

Most of the rock in the eastern sector is foliated. Foliation is expressed by biotite-rich bands, parallel biotite 

grains, flattened or rodded quartz and feldspar grains, 

parallel schlieren, feldspar prisms partly parallel to flow 

planes, and parallel slabby xenoliths of biotite quartzite 

and fine-grained schist derived from biotite diorite, amphibolite, or fine-grained biotite tonalite. 

The relations of foliation and lineation are more complex in the eastern sector than in the two sectors to the 

west, and they require description in three subareas, A, 

B, and C (fig. 3.4). The structure in subarea A displays 

markedly axial symmetry. Much of the foliation appears 

to be secondary and expressed in the texture of mediumgrained rocks. In places it cuts across folded schlieren of 

various trends that appear to reflect an early, primary 

flow foliation; locally it also cuts and distorts pegmatite. 

In other places it parallels primary foliation defined by 

slab by xenoliths and schlieren. The plunge of lineation is 

low to horizontal through this subarea, in contrast to the 

average of about 30° west-northwestward plunge seen in 

the central and western sectors. This reflects a broad arch 

in the lineation (and foliation) across the area (cross section in figure 3.2), from westerly plunges on the west to 

easterly plunges on the east. Lineation includes axes of 

passive flow folds in schlieren, streaked-out biotite, long 

axes of oriented feldspar phenocrysts, and long axes of 

flattened quartz and feldspar grains. In a few places the 

rocks are strongly lineated but not foliated in zones no 
3. STRUCTURAL GEOLOGY AND PETROLOGY OF A PART OF THE BITTERROOT LOBE 45 

7 

'\ 

'\ 

A 

t 

N 

I 

KILOMETERS 

8 '\: 

~·; I 

g 

I I :Ill I ;I 

B I :1 c I 

,'/ I I ;:1 I I ", I ___ , 

" ' I / ' ' .... ---- ' 

\ 

12 .... ~ 13\ 

------- -------: I I 

------- -----------A --------------- : I c 

-------- ----- --------- -- I I 

----- -!I I 

FIGURE 3.4.-Structural sketch map of subareas A, B, and C of the 

eastern sector, Bitterroot lobe of the Idaho batholith. The long dashes 

show foliation trends. Trend differences in subarea A are due to different folia in the axial flow system, all developed in flow along the 

az 290° axis; the lineation trend is the same throughout the subarea. 

Quadrangle numbers as in figure 3.2. 

more than a few meters thick. Furthermore, quartzitesheet xenoliths and related migmatites, with thicknesses 

measured at meter to kilometer scales in screens, have 

orthogonal crossing relations. One sheet stands vertical 

in the foliation and another is perpendicular to it and dips 

at low angles; they intersect in the lineation, and the symmetry is axial. 

In the Tenmile Lake quadrangle, sheets of tonalite as 

much as 20 em thick occur as discontinuous layers along 

the primary flow foliation in the gneissic granitic rock and 

are themselves gneissic. They pass laterally into biotite 

schlieren having a popcorn texture imparted by the 

idiomorphic habit of the remaining small amounts of 

feldspar. Some of the layers are folded with the flow foliation. The tonalite layers show gradual transition to granite 

through zones about 5 mm thick. Tonalite also forms 5- to 

10-cm laminae of lit-par-lit character in biotite quartzite 

sheet xenoliths. 

The eastward transition to the foliation orientation in 

subarea B involves crossing relations. A biotite quartzite 

stringer, oriented in the primary flow foliation and having 

the form of an isoclinal fold extending more than 30 m 

at az 020° along its axial plane, is cut across by a secondary foliation trending az 300° and is refolded in it. This 

has caused the development of crossing, secondary foliation in the fold limbs parallel to the az 300 o trend. A 

related feature is shown by a slabby, 10-cm-thick, biotite 

quartzite xenolith that extends more than 25m at az 020° 

along the relict primary flow direction. Streaky biotite 

lineation in the xenolith plunges 20 o at az 290 o , in the 

general direction of the secondary foliation. 

Foliation in subarea B trends north-northeast and dips 

steeply (fig. 3.4). Angular relations between primary and 

secondary foliations are comparable to those in subarea 

A, except for the absence of extensive flattening and 

mineral rodding. Lineation is rare and is expressed mostly 

by small, northeast-trending fold axes in the gneiss. The 

small folds are part of a conical fold set perhaps caused 

by secondary deformation of the primary flow foliation. 

The rocks in the northern part of subarea Bare more 

complex than those in adjoining areas. Both granite and 

granodiorite occur as early phases, apparently emplaced 

at the same time and recrystallized equally in the secondary strain field. A younger granite injected along the 

foliation remains unaffected by the secondary strain; it 

may equate to or postdate the sheets of fine-grained 

granitic rocks in neighboring areas, no other equivalent 

of which is seen in subarea B. Major migmatite occurs, 

with granodiorite as the principal paleosome. 

The northeast trends die out quickly to the south, in 

areas that were sparsely mapped or bypassed in reconnaissance mapping. If small structures recapitulate the 

large ones, then a major, steep-plunging fold hinge on the 

primary flow foliation should exist near the southern 

margins of Blodgett Mountain and Printz Ridge quadrangles, overprinted and perhaps destroyed by strong 

development of secondary foliation to the south. 

Foliation subarea C is part of the Bitterroot frontalgneiss zone. Foliation dips at shallow angles to the east, 

and its orientations yield an inclined axial symmetry suggesting a tilted dome. Foliation is expressed_ by oriented 

biotite flakes and by flattened quartz and feldspar grains. 

In the transitions, not studied in detail, from subareas A 

and B, the foliation becomes stronger, feldspar grains are 

generally flattened in the foliation, and the rocks in places 

become strongly lineated in the same general trend as the 

rocks to the west. The lineation is expressed by elongate 

biotite grains, trains of small biotite flakes, rodded quartz 

and feldspar grains, and minor fold axes in biotite quartzite xenoliths. Some zones several hundred meters thick 

are lineated but not foliated. 

The remaining structural features to be discussed are 

common to subareas A and C, but are largely or completely absent in subarea B. Feldspar prisms, some of them 

coarse grained to porphyritic in the medium-grained 

matrix, lie with their prism faces mostly parallel to lowangle foliation. This is seen in the rock faces exposing the 

a-c section (inferring the a and c kinematic axes on the 

assumption that rodded quartz and feldspar grains represent along their length the direction of extension and 

therefore the slip line). Viewed along the a-direction 
46 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

(normal to the b-e section), the feldspar-prism cross sections show an axial symmetry, being oriented in all possible directions relative to the foliation. But not all the 

feldspars seen in the a-e section of the outcrop face have 

their e-axes parallel to the foliation. At a particularly 

favorable and representative exposure, all the large 

feldspar grains, whose e-axes are roughly aligned eastsoutheast, were counted in a 37 em by 39 em area on a 

near-vertical face in that direction. Viewed looking northnortheast, seven of the 23 large grains present show 

counterclockwise rotation from the foliation plane, one 

shows clockwise rotation, and 15 are not rotated. None 

of these rotations is large and most of them are less than 

25°. Here, as in the central sector, the resemblance of 

the defined feldspar plane to the S-plane of Berthe and 

others (1979) is remarkable. 

Late-stage secondary shears, developed after the 

pegmatite stage, reduce feldspar phenocrysts to flaser 

structures fairly penetratively, but with intensity varying through gradational zones. The reduced phenocrysts 

are distorted into flattened, somewhat lenticular rods, 

which are parallel to streaky biotite and rodded quartz 

grains and generally also to other lineation. At one 

representative locality, 30 grains were measured, ten in 

each of three orthogonal faces: a-e, b-e, and a-b (kinematic 

axes). The average dimensions of the flattened rods are: 

c, 6. 7 mm; b, 8.6 mm; and a, 13.6 mm; these give an axial 

ratio of 1:1.3:2. Most of the undeformed phenocrysts are 

stubby prisms and are not in linear array, so that stretching seems clearly indicated in rod formation. Thin, quartzfilled extension fractures normal to their length taken as 

~inematic a in the rodded feldspar grains bear out this 

inference. 

A few blocks of anorthositic orbiculite were seen in 

granite talus; sparse orbicules occur strung out and 

rotated along primary flow foliation planes in the granite. 

Anorthosite was found in place at two localities; because 

of its physical resemblance to neighboring granodiorite, 

it was not recognized as anorthosite during field work but 

only when stained slabs were made. 

Xenoliths include biotite quartzite, andesine quartzite, 

calc-silicate gneiss, biotite granodiorite, biotite granodiorite gneiss, tonalite gneiss and schist, quartz diorite, 

quartz diorite gneiss, biotite diorite schist, and amphibolite. They are considerably more plentiful in the eastern 

sector than in the central sector. 

Numerous dikes and sheets of fine-grained granite, 

granodiorite, and tonalite occur in subareas A and C. They 

are injected as sheets along the primary foliation and as 

dikes in fractures containing the secondary lineation or 

perpendicular to it. All are cut by the secondary foliation 

and are intruded by aplite and pegmatite, which are 

themselves partly gneissic because of the secondary 

foliation. 

PETROGRAPHY 

Granitic rocks across the mapped part of the Bitterroot 

lobe are principally an intricate mixture of granite and 

granodiorite, with smaller amounts of quartz monzonite, 

quartz monzodiorite, tonalite, and anorthosite. Modes are 

summarized in figure 3.5A. All these rocks share the same 

pattern of petrographic variation, in which the rocks on 

the west are mildly metamorphosed and weakly foliated, 

if at all, and these grade to rocks on the east that are 

strongly metamorphosed and foliated on three successive 

s planes: gneissic, blastomylonitic, and mylonitic. Because 

of this, the petrography is generalized to cover all rock 

types present and is described mineral by mineral. 

Plagioclase occurs as 0.5- to 4-mm grains. Here and in 

the following descriptions such measures refer to the 

range of average grain sizes determined in the several 

thin sections studied. Plagioclase composition ranges from 

An15 to An33· Both modal amount and composition of 

plagioclase vary more or less randomly over the area, and 

this applies indeed to all mineral modes over the entire 

map area. Virtually all the plagioclase is unzoned; only 

in a few rocks were zoned grains (cores An26-28, rims 

An20-23) seen, and even there they are mostly relics. 

Grains are equant to prismatic, euhedral to subhedral; 

some of the subhedral grains may be crystalloblastic. 

Some plagioclase is myrmekitic against K-feldspar, interpreted as due to strain effects. In three specimens, much 

of the plagioclase is in clusters of 0.5- to 0.7-mm grains 

of irregular granoblastic form with interfacial angles approaching 120 o and with some microcline inclusions. 

These clusters are in part enclosed within K-feldspar. In 

several specimens, plagioclase grains are cut by filled, 

FIGURE 3.5.-Plot of plutonic rock modes from the Bitterroot lobe of 

the Idaho batholith, Idaho-Montana, on a quartz-plagioclaseK-feldspar ternary diagram. Rock names are from Streckeisen and 

others (1973). Modes recalculated to 100 percent. A, Summary plot 

of 509 plutonic rock modes; number of modes for each rock type 

given in parentheses: la, quartzolite (0); lb, quartz-rich granitoids 

(7); 2, alkali-feldspar granite (2); 3, granite (211); 4, granodiorite 

(194); 5, tonalite (68); 6, alkali-feldspar quartz syenite (0); 7, quartz 

syenite to syenite (1); 8, quartz monzonite to monzonite (4); 9, quartz 

monzodiorite/quartz monzogabbro to monzodiorite/ monzogabbro (2); 

10, quartz diorite/quartz gabbro/quartz anorthosite to diorite/gabbro/ 

anorthosite (20). B, Plot of 46 plutonic rock modes from the southwestern Bitterroot lobe; numbers of modes: lb, (4); 2, (1); 3, (6); 4, 

(2); 5, (28); 10, (5). C, Plot of 155 plutonic rock modes from the 

western sector (see figure 3.2) of the Bitterroot lobe; numbers of 

modes: lb, (3); 2, (1); 3, (78); 4, (51); 5, (15); 7, (1); 8, (3); 9, (1); 10, 

(2). D, Plot of 171 plutonic rock modes from the central sector (see 

figure 3.2) of the Bitterroot lobe; numbers of modes: 3, (77); 4, (78); 

5, (11); 8, (1); 9, (1); 10, (3). E, Plot of 137 plutonic rock modes from 

the eastern sector (see figure 3.2) of the Bitterroot lobe; numbers 

of modes: 3, (50); 4, (63); 5, (14); 10, (10). 
/ 

K 

3. STRUCTURAL GEOLOGY AND PETROLOGY OF A PART OF THE BITTERROOT LOBE 

8 

a 

j\ 

/~ \\ I \ 

A 

50 

k ~ 

/ \ ik"" 1\ \ 

50,/i::- ~. 

/;;-

/ 

,<f-

/kD 

50 

a /\.· 

.. -t 

a 

50 

a 

c 

\ 

........... . -:·: · .. ·: .. ::.· .. · : ·:. : :. . 

p K 

a 

E \ 

\ 

47 

p 
48 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

0.1-mm-thick, extension fractures that are perpendicular 

to lineation; filling minerals include K-feldspar, quartz, 

and pennine. Plagioclase in some rocks is in flattened 

grains which in stained slabs impart a distinctly foliated 

aspect to the rock; some of these flattened grains are 

broken into 0.5-mm, strain-free subgrains. These rocks 

contain the strongest development of secondary fabric 

seen in the area. In the plagioclase of one granite, rodding is developed on grains 3 mm by 8 mm or longer, 

segmented into several parts by extension fractures 0.2 to 

0.3 mm thick, which are filled with pennine-muscovitemicrocline-allanite-quartz. Considering the mobility of 

microcline and quartz, the rodding may have developed 

before the granite was completely crystallized. In places, 

plagioclase is broken and recrystallized in 0.3-mm, strainfree, granoblastic subgrains together with quartz and 

microcline of comparable size, in zones as much as 0.5 mm 

across. This is a blastomylonitic texture, and it is overprinted in places by subparallel mylonite. Cataclastic 

effects of microbrecciation along crossing, narrow shears 

include bending of plagioclase twin lamellae, undulatory 

extinction, and breaking of quartz and feldspar grains. 

Quartz occurs as 1.5- to 4-mm grains. The habit of 

quartz in the western sector differs between the northern 

zone and the southern zone. A subtle but major zone of 

structural transition trends approximately east-west 

through the southern half of the Huckleberry Butte and 

Greenside Butte quadrangles and continues east near the 

southern boundary of the Fish Lake quadrangle, well 

within the area of granitic rocks. Beyond the Fish Lake 

quadrangle, this zone continues southeastward (fig. 3.1). 

Quartz of the northern zone, typical also of the central 

and eastern sectors, is altogether broken into 0.5- to 

0.7-mm subgrains, free of strain effects where not 

affected by younger cataclasis. Some rocks contain flattened grains 2 mm by 4 to 6 mm that are also broken into 

strain-free subgrains. Four northern zone specimens show 

blastomylonitic shear zones less than 1 mm thick along 

which 0.2-mm, granoblastic, strain-free grains of quartz 

occur. Quartz of the southern zone is largely as single, 

optically continuous grains free of subgrains, the develop-

-· ·at of which dies out to the north over a zone approximately one kilometer wide. 

In the central and eastern sectors, the quartz grains are 

flattened (from 0. 7 mm by 1.5 mm to 1 mm by 7 mm) 

parallel to foliation, and all have strain-free sub grains. 

Quartz in some rocks is cut by blastomylonitic shear zones 

less than 0.5 mm thick, within which the quartz occurs 

as 0.2-mm, strain-free subgrains. Along restricted shear 

zones less than 0.3 mm wide in the western part of the 

eastern sector, many small grains show undulatory extinction. This may be a result of blastomylonite development 

in which strain outlasted recrystallization, or it may involve late mylonitic overprint. In some of the gneiss in 

the central and eastern parts of the eastern sector, 

blastomylonite zones pass laterally into curved, singlegrain quartz ribbons as large as 0.5 mm by 2 mm. 

Blastomylonite zones and large, platy quartz grains nearly 

wrap some of the feldspar grains, and in part quartz 

masses formed by segregation effects from quartz formerly reduced to fine grain size in blastomylonitic shear. The 

latter view is reinforced by the presence of inclusions of 

blastomylonitic mica within the quartz. 

Many rocks show effects ranging from mild to severe 

of cataclastic deformation in quartz. The milder effects 

include undulatory extinction and subgrain boundary 

suturing in strain-induced boundary migration. Stronger 

effects include breaking into 0.03- to 0.3-mm cataclastic 

subgrains, which show undulatory extinction and more 

or less intricate microsuturing of subgrain boundaries. 

Some of these stronger effects can be related to shear 

zones mapped at the outcrop. 

K-feldspar occurs as 1- to 8-mm grains, and also as 1- to 

4-cm phenocrysts in the porphyritic rocks. Over much of 

the area, the K-feldspar mineral varies: about one-third 

of the rocks contain microcline, another one-third contain 

submicroscopically twinned microcline, and the remaining one-third contain orthoclase. In a few rocks, both 

orthoclase and submicroscopically twinned microcline are 

present. In the eastern sector, only microcline occurs, the 

larger grains of which contain many inclusions of plagioclase, biotite, and quartz. Extension fractures filled by 

quartz and muscovite cut the K-feldspar in several 

specimens. In foliated rocks, the K-feldspar occurs as lenticular grains from 1 mm by 3 mm to 3 mm by 8 mm in 

size and oriented parallel to the foliation. The variety of 

K-feldspar is independent of the degree of flattening. Flattened grains are generally broken into strain-free, 0.5-mm 

subgrains. Part of the K-feldspar in zones of blastomylonite also occurs as 0.5-mm subgrains. 

Biotite occurs as 0.3- to 1-mm grains, crudely lepidoblastic in part of the rock and randomly oriented in the 

rest. Minor pennine alteration is ubiquitous. Fibrolite inclusions occur in some biotite. In blastomylonite zones, 

biotite grain size is near 0.2 mm, and in later mylonite 

zones, 0.02 mm. Muscovite, where present, occurs as 

0.5- to 1-mm grains, but it is absent in many rocks; its 

occurrence is otherwise much like that of biotite. Accessory minerals seen, in approximate order of decreasing 

abundance, include magnetite, zircon, apatite, thorite, 

monazite, baddeleyite, allanite, and sphene. Magnetite is 

more abundant in the western sector than all other accessories together. Averages of all the modes for granitic 

rocks show 0.6 percent accessories in the western sector, 

0.1 percent in the central sector, and 0.4 percent in the 

eastern sector. 

A concordant sheet of biotite-hornblende quartz diorite 

occurs in the Cedar Ridge quadrangle and is the source 
3. STRUCTURAL GEOLOGY AND PETROLOGY OF A PART OF THE BITTERROOT LOBE 49 

of abundant black sands in several creeks in its vicinity. 

It was injected along the secondary foliation in the 

gneissic granitic rocks and then metamorphosed. Plagioclase (An45) as 2.5-mm grains is unzoned, prismatic, and 

subhedral. Quartz as 0.5-mm grains is interstitial and 

largely broken into 0.2-mm, strain-free subgrains. Biotite 

is as 2-mm flakes which partly include and replace 1.5-mm 

grains of green hornblende. Accessories include magnetite 

and apatite. 

One or more small bodies of primary flow-foliated quartz 

diorite, free of secondary foliation, are exposed in the 

Printz Ridge quadrangle. Plagioclase (An35) occurs as 

2-mm, mostly unzoned, prismatic grains; a few relics of 

normal zoning persist. Biotite and hornblende are important constituents and occur as 1-mm grains. Quartz is 

minor and forms 0.5-mm grains free of any subgrains. Accessories include magnetite, apatite, and sphene. 

Anorthosite was found at three locations in the Blodgett 

Mountain quadrangle; the rock closely resembles granodiorite in physical appearance and was not recognized as 

anorthosite until analyzed in the laboratory. I suspect it 

to be synplutonic with the granitic rocks. Plagioclase 

occurs as 4- to 8-mm unzoned grains of composition 

An27-29· These grains are partly converted to granoblastic, pavement-textured grains 0.5 to 1 mm in 

diameter. Quartz occurs as irregular 2-mm grains and contains a few strain-free subgrains 0.5 mm in diameter. 

Microcline is minor and occurs as 1-mm aggregates of 

0.5-mm subgrains with approximately 120° triple junctions. Biotite and muscovite occur in trace amounts. 

Dikes of fine-grained granitic rocks ranging in composition from tonalite to granite intrude the granitic rocks 

in the central and eastern sectors, except in subarea B. 

The dikes lie both along and across the secondary foliation and are themselves cut by it. Plagioclase (An25-35) 

occurs as 0.3- to 0.5-mm, mostly unzoned, equant to 

prismatic grains. Quartz occurs as irregular, 0.5-mm 

grains to platy grains 0.3 mm by 0.5 mm in size, flattened 

parallel to the foliation; some have subgrains and some 

do not. Microcline, submicroscopically twinned, occurs in 

much the same way as quartz. Biotite and muscovite 

(which is absent in some rocks) are lepidoblastic as 0.3- to 

0.5-mm grains. Accessory minerals seen include zircon, 

apatite, allanite, and sphene. 

A 16-m-diameter pipe in the Blodgett Mountain 

quadrangle contains miarolitic intrusive rock identified 

as a younger anorthosite. The plagioclase (An37) is as 

stubby, prismatic, unzoned 3-mm grains, which show mild 

alteration to sericite and clay. Biotite occurs as 0.5-mm 

grains, and a few 0.3-mm grains of quartz occur. A little 

sanidine, intimately intergrown with plagioclase in 

0.3-mm grains, occurs around miarolitic cavities. The pipe 

is cut by a fine-grained, low-angle blastomylonite shear 

zone 10 em thick. The shear zone is filled with pavementtextured, unzoned, strain-free plagioclase grains (An37) 

of 0.5-mm size, showing clay and sericite alteration. A 

comparable anorthosite occurs in the Tenmile Lake quadrangle as a 1-m-thick dike oriented at a high angle to the 

extension lineation. 

The western, central, and eastern sectors show somewhat different average compositions. A total of 463 modes 

of plutonic rocks were determined: 155 in the western 

sector, 171 in the central sector, and 137 in the eastern 

sector. These modes when recalculated to total100 percent for light minerals and then averaged should give a 

reasonable estimate of the bulk composition of the Bitterroot lobe in each sector. For the western sector, the 

resulting average composition is 43 percent plagioclase, 

33 percent quartz, and 24 percent K-feldspar. For the central sector, the average composition is 46 percent plagioclase, 32 percent quartz, and 22 percent K-feldspar; and 

for the eastern sector it is 51 percent plagioclase, 30 percent quartz, and 19 percent K-feldspar. Thus, from east 

to west, the pluton decreases in plagioclase and increases 

in quartz and K-feldspar. The rocks of the western sector are generally granite, those of the central sector 

granodiorite, and those of the eastern sector more calcic 

granodiorite. The dark minerals, principally micas and accessory minerals, average 9 volume percent in each of the 

three sectors. Modal distribution is summarized in figure 

3.5A, C, D, and E, and listed in table 3.1. Modes of the 

southwestern Bitterroot lobe (46, given in fig. 3.5B) are 

not considered here. 

BLASTOMYLONITE 

The granitic rocks, together with their associated dikes 

and mineral-lined or filled primary fractures, are cut in 

places by fine-grained blastomylonitic schistosity or by 

zones of blastomylonite as much as 35 em thick; these 

features are strongest on the east and grade to virtual 

absence on the west. Bell and Etheridge (1973) suggested 

that mylonite is generally the result of ductile rather than 

brittle strain. Their suggestion seems applicable to the 

mylonitic rocks of the Bitterroot lobe. However, their 

definition of mylonite seems too general for discrimination of rocks encountered in this study. For instance, both 

the rocks termed blastomylonite and those termed 

mylonite in this study would qualify as mylonite under 

their definition. Therefore, the definitions in the Glossary 

of Geology (Bates and Jackson, 1980) are used in this 

paper. According to those, blastomylonite is a mylonitic 

rock in which some recrystallization and (or) neomineralization has taken place. In the blastomylonite described 

here, a large amount of recrystallization has occurred, so 

that the small pavement-like grains of the fabric are 

entirely free of strain effects; that is, no undulatory extinction is seen. 
50 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

TABLE 3.1.-Distribution of plutonic rock modes from the western, central, and eastern sectors of the Bitterroot lobe of the Idaho batholith 

Plutonic rock field from 

Streckeisen and others 

(1973) 

Western sector 

(fig. 3.5C) 

Central sector 

(fig. 3.5D) 

Eastern sector 

(fig. 3.5E) 

Number 

present 

Percent 

of total 

Number 

present 

Percent 

of total 

Number 

present 

Percent 

of total 

1a. Quartzolite-------------- 0 

lb. Quartz-rich granitoids--- 3 

2. Alkali-feldspar granite-- 1 

3. Granite------------------ 78 

4. Granodiorite------------- 51 

5. Tonalite----------------- 15 

6. Alkali-feldspar quartz 

syenite to alkali 

feldspar syenite 0 

7. Quartz syenite to 

syenite 

8. Quartz monzonite to 

monzonite 

9. Quartz monzodiorite/ 

quartz monzogabbro to 

monzodiorite/monzogabbro 

10. Quartz diorite/quartz 

gabbro to diorite/gabbro 

3 

2 

0 

1.9 

.6 

50.3 

32.9 

9.7 

0 

.6 

1.9 

.6 

1.3 

The blastomylonitic schistosity partly parallels and 

partly crosscuts flow foliation and is developed independently in massive granitic rock. The blastomylonite 

contains streaky fine-grained biotite lineation, and most 

of it is foliated. Blastomylonite is also developed in 

pegmatite-suite, and other earlier, dikes, either parallel 

to their walls or crossing at various angles. The 

blastomylonitic shears vary widely in their orientation, 

but their poles lie near an inclined girdle whose axis 

plunges west-northwest (fig. 3.3). One gently dipping 

pegmatite sheet 2 em thick is split by a blastomylonitic 

shear parallel to its walls. Two large phenocrysts were 

split by this shear and their top halves transported 5 em 

to the southeast. Steps on exposed parts of the shear surface suggest the same sense of displacement as that 

shown by the offset phenocrysts. 

Blastomylonite shear zones cut the gneiss of the Bitterroot frontal-gneiss zone along dips somewhat shallower 

than those of the gneiss; thus, arching of the secondary 

foliation and therefore doming had already begun when 

the blastomylonite shear zones formed. Stretching 

produced quartz-filled fractures, whose geometry also indicates that tops moved to the southeast. 

MYLONITE 

Late-stage mylonitic shears that cut earlier structures, 

including blastomylonite, were found in several places. 

0 

0 

0 

77 

78 

11 

0 

0 

3 

0 

0 

0 

45.0 

45.6 

6.4 

0 

0 

.6 

.6 

1.7 

0 

0 

0 

50 

63 

14 

0 

0 

0 

0 

10 

0 

0 

0 

36.5 

46.0 

10.2 

0 

0 

0 

0 

7.3 

Mylonite (Bates and Jackson, 1980) "may * * * be 

described as a microbreccia with flow texture." Implicit 

in this definition is the presence of deformation features, 

such as strong undulatory extinction in the mineral particles, and the absence of recrystallization, at least as 

viewed on the microscopic scale. Mylonitic shears are all 

relatively thin, mostly less than 0.5 mm thick, although 

one 8-cm-thick shear was found. They are strongly 

linea ted owing to fine grooving and streaky sericite. The 

shears are all shallow dipping, and their lineations plunge 

near either az 120° or az 300°, depending on the dip direction. Shear steps suggest that the tops moved to the east. 

The mylonitic shears are more numerous in and near the 

gneissic Bitterroot front, where they occur in sheeted 

zones 2 to 5 em thick and spaced 8 to 100 em apart. Some 

mylonite zones are crossed at angles of 15° to 30° to the 

bounding shear surfaces, by shears internal to the zones. 

These are taken to be Riedel shears; the sense of shear 

transport derived from the Riedel shears is the same as 

that shown independently by shear steps. 

STRUCTURAL EVOLUTION OF THE WALL ROCKS 

The calc-silicate gneiss has a flow-foliation fabric containing steeply plunging mineral lineation of diopside, 

hornblende, biotite clots and streaks, and quartz rods. 

Although the order of growth of the minerals in the fabric 

is not entirely clear, the single lineation direction of all 
3. STRUCTURAL GEOLOGY AND PETROLOGY OF A PART OF THE BITTERROCYI' LOBE 51 

the minerals present appears to require a single recrystallization event. Any preexisting metamorphic fabric has 

therefore been completely destroyed in the latest event. 

Deformation associated with the metamorphism formed 

minor passive flow folds. Their axes plunge steeply and 

the axial planes parallel foliation in the enclosing gneiss. 

These small folds show both clockwise and counterclockwise sense of shear, distributed more or less randomly 

through the outcrops. This suggests an orthorhombic symmetry with flattening perpendicular to the x-y plane. 

The parallelism of the mineral lineation and the minor 

fold axes suggests a common origin. One possibility is that 

all the linear elements were generated in a horizontal 

attitude in the plane parallel to the near-vertical flow 

direction, taken as imposed on the wall rock of the batholith by upward-moving magma. In that case, the linear 

elements were rotated toward kinematic a during continuing flow. A second possibility is that the linear 

elements were generated in the direction of kinematic a 

during flow. Available evidence does not permit a choice 

between these two possibilities. 

The structure in the tonalitic gneiss strongly resembles 

that in the calc-silicate gneiss; the two rocks must have 

been subjected to the same structural history. This 

requires that the tonalitic gneiss be injected before or during the metamorphism that produced the steeply plunging 

minor folds and mineral lineation in the rocks. The fabric 

of the tonalitic gneiss does not look metamorphic; rather, 

it has the appearance of an intense primary flow foliation, 

in which the plagioclase retains much of its igneous form. 

Therefore, the tonalitic gneiss is inferred to be a synkinematic intrusion. 

Lit-par-lit stringers of trondhjemite formed by injection 

of late-stage alaskitic material lie along the foliation in 

the tonalitic gneiss and the calc-silicate gneiss. The total 

fabric suggests a single-stage event involving metamorphic recrystallization and igneous injection during nearvertical flow. 

Quartzite is exposed in the southwestern wall rocks. The 

D1 fold set in the quartzite is sharply isoclinal, as shown 

by small outcrops in which minor folds assigned to D1 

have axial-plane foliation. The varied plunges of the fold 

axes, from near the strike line to the dip line, suggest, 

following Hansen (1971), that younger shear in the plane 

of the foliation has rotated D1 folds toward the slip line 

of a younger or continuing deformation. Pegmatite was 

injected into the D1 structures at some time during that 

later event, here taken as D2. 

Folds assigned to Dg appear to be the youngest in the 

section; they are concentric and have varied but relatively 

shallow plunges. They clearly deform lineations assigned 

to D2: the axes of the conical folds and the mineral lineations parallel to them. To analyze the D2 structures, it is 

best to remove the Dg overprint. When the Dg folding has 

been removed, the sequence of metaquartzite is nearly 

vertical. Conical folds with mineral lineation parallel to 

their axes plunge down the dip of the near-vertical 

sequence. 

In the Elk City area (Reid, 1959), the isoclinal D1 folds, 

which have axial-plane foliation, were recumbent before 

later deformation. Perhaps the D1 fold set of that region 

extended into the southwestern wall rocks, only about 

30 miles to the northeast. Near the batholith, the isoclinally folded rocks, which were initially recumbent, 

steeply tilted as the batholith rose diapirically into and 

through superjacent rocks. This resulted in a D2 strain 

axis not parallel to the layering, the general requirement 

for the generation of conical folds (Stauffer, 1964). 

Because the slip line is nearly vertical, the axes of the conical folds generated will also be nearly vertical, plunging 

down the dip of the vertical foliation. The active folding 

of some injected pegmatite in conical folds supports the 

proposed contemporaneity of batholith intrusion and D2 

folding, although pegmatite continued to be injected after 

some conical folds had already developed. The helical or 

screwlike character of the axes of the conical folds is 

attributable to unstable, vertical viscous flow during D2, 

in which a rotational component was imparted to the vertically flowing rocks. Reid and others (1979) discussed 

evidence showing that the wall rocks flowed vertically 

downward (in the absolute sense) along the wall of the 

rising batholith. 

The Dg folds are viewed as a major drag-fold set imposed by the batholith on its wall rocks late in its rise. 

The D2 conical folds are draped over the hinges of the 

concentric D3 folds, roughly at right angles to their axes. 

The complex geometry implicit in this structural array is 

no doubt responsible for the wide scatter of poles to foliation in the wall rocks of the southwestern batholith. 

It is important to note that lineation of biotite and 

muscovite parallels the axes of all three fold sets. Apparently the three fold sets developed during a single 

phase of metamorphism in the hornblende hornfels facies. 

Given the abundant igneous injection into the rocks and 

the metamorphism of much of that igneous material, it 

seems clear that the thermal and mechanical energy for 

the metamorphism was from the emplacement of the 

Idaho batholith. 

STRUCTURE OF THE TONALITE 

Flow foliation in the tonalite intrusive unit of the 

southwestern lobe dips about 70° southwest. Flow lineation plunges about 70° in the plane of the foliation and 

is believed to represent the direction of flow parallel to 

the contact. Steep flow within the intrusion has no doubt 

generated the steep flow structures in the wall rock 

through a mechanism of frictional drag on the walls and 
52 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

resulting passive flow in the strongly heated walls of the 

intrusion. Metamorphic rock in the migmatite screens in 

the batholith behaved similarly. The axial symmetry 

shown by flow elements in these several units supports 

the idea of diapiric intrusion. 

Sheets of trondhjemite and granite alaskite with gentle 

dips are nearly perpendicular to the inferred batholith contact and to the direction of flow lineation. The lineation 

is thus a stretching lineation, and the sheets represent 

primary tensional joints perpendicular to the direction of 

stretching (see Balk, 1937). Similar sheets are found in 

the wall rock, which was stretched in the same way. The 

alaskitic material invaded the primary joints as they 

formed. The varied dike composition, trondhjemite and 

granite, suggests more than one intrusive episode; 

simultaneous emplacement would require synmagmatic 

emplacement of different magmas. Granite synplutonic 

with the dikes was not found, but it may be represented 

by granite xenoliths found in the younger granite discussed below. 

In most thin sections of the biotite tonalite, the plagioclase is euhedral and unzoned, and there is no systematic 

variation in its anorthite content. Plagioclase, the predominant constituent, is free of any strain phenomena 

that might indicate protoclasis has occurred. Therefore, 

it seems that the magma crystallized slowly at approximately equilibrium conditions. 

The foregoing general formulation does not explain a 

number of details. Consider the complex folds in the flow 

foliation: they are arrayed about a steeply plunging axis 

and are not clearly either cylindrical or conical, although 

both kinds are seen. The pattern is one of axial symmetry 

like that within salt stocks, and flow mechanics comparable to those within salt stocks may be postulated. 

Moreover, a twisting or torsional component was present 

in the vertically directed flow in the intrusion, as in the 

wall rock, leading to the complexly twisted or curved axes 

of some of the steeply plunging folds. Associated minor 

folds show both clockwise and counterclockwise rotation 

randomly distributed through individual small outcrops, 

suggesting that the folding was not systematic. 

Dikes of leucocratic biotite tonalite cut the tonalite intrusion. These were sheared while still hot, and they 

recrystallized as fine-grained biotite-plagioclase schist or 

blastomylonite with strong field resemblance to mylonite. 

The shearing, which predated intrusion of the pegmatite, 

suggests faulting along trends parallel to the trans-Idaho 

discontinuity (Yates, 1968). That discontinuity may therefore have been active during or shortly after tonalite 

emplacement. 

The age of the tonalitic rocks presents a problem. Reid 

and others (1973) inferred, on the basis of 207pbf206pb zircon ages, that the tonalitic gneiss and related augen gneiss 

in the country rock of the batholith were Precambrian. 

Field work on this project has led to an alternative idea 

that the tonalite gneiss was derived from the intrusive 

rock emplaced in the Cretaceous as an early precursor of 

the batholith. The zircon ages might then reflect the age 

of the. parent rock from which this earlier tonalite magma 

was derived. Hornblende K-Ar work by R.J. Fleck (oral 

commun., 1979) has shown the age of both the tonalite 

gneiss and the tonalite to be about 65 Ma. The tonalite 

gneiss and the tonalite were apparently intruded nearly 

simultaneously, though the tonalite cuts the tonalite 

gneiss and is thus at least slightly younger. An age of 

about 73 Ma was determined by Bickford and others 

(1981, p. 453) for a quartz diorite gneiss that is found at 

the northeast contact of the batholith and is both structurally and petrographically similar to the tonalite gneiss 

of this study. 

PETROLOGY AND STRUCTURE OF 

THE GRANITIC ROCKS 

MAGMATIC FEATURES 

The granitic rocks of the Bitterroot lobe are clearly 

younger than the tonalite of the southwestern area, inasmuch as there are xenoliths of tonalite in them near the 

contact. The granitic rocks are predominantly granite and 

granodiorite, apparently more or less randomly commingled with each other and with smaller amounts of 

quartz monzonite, quartz monzodiorite, tonalite, and anorthosite. These rocks share many petrographic similarities 

and may have been emplaced synplutonically. 

As shown earlier, the average composition of plutonic 

rocks in the Bitterroot lobe is granite in the western sector 

and changes to calcic granodiorite in the eastern sector. 

Structural analysis shows that the surface rocks of the 

western sector represent the upper part of the batholith, 

grading to the lower batholith in the eastern sector 

(fig. 3.2). The batholith is therefore vertically zoned, from 

granite at the top to calcic granodiorite near the base. This 

would appear to require a differentiation process, but examination of the kind of differentiation is beyond the 

scope of this paper. 

Tabular tonalite bodies along the primary flow foliation 

in the Tenmile Lake quadrangle appear to have been still 

liquid at the time of granite intrusion; this suggests 

synplutonic emplacement, accompanied perhaps by some 

mixing with granite and granodiorite magma. 

The anorthosite bodies in the eastern Bitterroot lobe, 

herein referred to as anorthosite I, require special comment. Anorthosite is generally held to be exclusively of 

Precambrian age; if that were so, the anorthosite here 

would have to be in rafted hypoxenoliths. Because no 

evidence for rafted hypoxenoliths was seen, however, it 

is believed that the conditions for generation of anor-
3. STRUCTURAL GEOLOGY AND PETROLOGY OF A PART OF THE BITTERROOT LOBE 53 

thositic magma were met in this area, which is near the 

base of the batholith and hence near the base of the crust 

at the time of emplacement. The anorthosite I bodies are 

therefore inferred to be magmatic. 

The K-feldspar varies across the pluton: microcline is 

the sole K-feldspar near the eastern margin, whereas 

elsewhere orthoclase is present in addition to microcline. 

Because the batholith cooled more slowly near its base, 

conditions there allowed orthoclase to invert completely 

to microcline. 

The granitic rock of the western sector (upper batholith) 

contains some relict crystalloblastic plagioclase, possibly 

incompletely melted material that has partially retained 

its metamorphic aspect. The source for the granitic rock 

may therefore include older granitic material partly 

melted at depth. 

Steep flow lineation in the granitic rock near its contact with the older tonalite in the southwestern area 

indicates near-vertical flow along a locally east-trending 

contact. Near-horizontal pegmatite sheets are consistent 

with vertical stretching along this contact. The features 

seen along this contact are taken to represent the magma 

tectonics of the granitic rocks; these features have been 

extensively modified or obliterated in areas to the north 

and east. 

Schistosity/foliation in the northern zone of the western 

sector is highly discordant to that of the granitic rock in 

the southern zone, near the contact. The average lineation suggests a dip direction near 30°, az 320°. Orientations of pegmatite and aplite dikes, combined with other 

data, portray an ill-defined girdle (fig. 3.3). The pole to 

this girdle is oriented near 10 o , az 300 ° , not far from the 

slip line for flow in the granitic rock, based on stretching 

lineation. Based on regional structural history, an 

alternative explanation for the girdle involving northeastsouthwest compression and folding is not viable. Solidstate flow evidently continued on about the same axis as 

the liquid-state flow after the granitic rock had consolidated sufficiently to sustain shear. Fractures generated 

in the solid-state flow were injected by pegmatitic magma. 

Blastomylonite schistosity has an average orientation 

of az 320°, 40° SW., and the associated lineation, taken 

as the slip line, averages 30°, az 290°, similar to the orientations of the slip lines postulated above, but rotated 

counterclockwise with respect to them. Therefore, it 

appears that continuing movement along the az 290° 

trend after pegmatite intrusion led to blastomylonite 

development. 

SECONDARY FOLIATION 

U nzoned plagioclase occurs in rocks throughout the 

area. It is possible that cooling under near-equilibrium conditions allowed originally zoned plagioclase to equilibrate. 

Secondary strain may have aided this process, but unzoned plagioclase seems to have developed about as well 

outside the area affected by secondary strain as within 

it. This idea needs to be tested farther away from the field 

of secondary strain. 

Plagioclase composition is roughly the same in granite 

and in granodiorite, ranging from about An1s to An33 

randomly over the whole area. Cores in relict zoned grains 

in granodiorite are somewhat more calcic than those in 

granite, suggesting that the granodiorite magma had the 

potential to form more calcic plagioclase. However, the 

secondary strain has affected both rocks in the same 

pressure-temperature interval during solid-state strain, 

and the plagioclase in both was reequilibrated to lie in the 

same compositional range. 

Evidence of movement in the igneous rock in the lateliquid stage, when the crystal meshwork was sufficiently 

established to undergo strain, is found in quartz-microcline-filled extension fractures in plagioclase. Quartz-filled 

extension fractures in microcline indicate continued movement near the end of the magmatic stage. Extension fractures that developed later during continuing solid-state 

strain are filled with muscovite, quartz, and pennine. 

These fractures all stand more or less perpendicular to 

the mineral lineation and are due to stretching along that 

lineation. 

The fabric elements produced in secondary strain are 

identical in the granite and the granodiorite as well as in 

the less common varieties of granitic rocks. This suggests 

that the several rocks are about the same age (synplutonic), inasmuch as they experienced the same postconsolidation, high-temperature strain event. This strain 

event was severe enough that the granitic rock of the entire area can be considered as essentially a metamorphic 

unit. 

The shallow plunge of secondary elements in granitic 

rocks of northern areas (northern zone of the western sector and areas to the east), compared to the steep plunge 

of linear elements in the older tonalite and younger 

granitic rocks in the southern zone, shows that the rocks 

experienced dramatically different tectonic environments. 

Granite of the southern zone rose vertically along its contact, crystallized, and experienced no further activity. In 

rocks of the northern areas, on the other hand, secondary 

foliation began to form before they were entirely solid, 

as shown by extension fractures in the feldspar that are 

filled with late-stage magmatic material. Strain that 

continued after consolidation formed gneissic foliation, 

flattened feldspar and quartz, and deformed the later, 

fine-grained dikes and the still later pegmatite and aplite. 

The strain is principally axial on the west, but becomes 

more orthorhombic to the east. 

The diffuse boundaries of the secondary strain field in 

the granitic rocks are interpreted as bounding a broad 
54 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

zone of distributed shear. Primary flow foliation formed 

independently of these boundaries. Axial solid-state flow 

in the secondary strain field overprinted and largely 

destroyed the primary foliation within its boundaries. The 

axis of this flow generally plunges at shallow angles to 

the west-northwest, but changes to an east-southeast 

plunge over an arch near the eastern margin of the pluton, 

indicating axial flow symmetry that was later arched. 

The principal petrographic criteria for secondary strain 

include anhedral to flattened quartz grains with strong 

development of annealed subgrains; filled extension fractures in the feldspar perpendicular to the stretching lineation; flattened feldspar grains; and subgrain development 

in the feldspar. 

The eastward-increasing proportion of foliated rocks in 

the area reflects the approach to a zone of more intense 

strain near the base of the pluton. Strain-free subgrains 

in the plagioclase become more common to the east, 

deeper in the pluton, because of higher deformation temperatures and greater concomitant annealing in that 

direction. Mica generally becomes oriented in the rock 

only under more intense strain than that in which quartz 

and feldspar first become flattened. This may be due to 

the type of strain farther to the east: considerably more 

strain may be required to reorient a mica grain in a field 

of flattening strain than in one of rotational strain. 

In many places in the eastern sector, feldspar grains 

are out of parallelism with the main foliation, forming a 

feldspar foliation that dips more steeply than the biotite 

foliation. Thin-section analysis suggests that biotite grains 

lie on shear surfaces; therefore, as suggested above, the 

C and S surfaces of Berthe and others (1979) may be 

present. Analysis based on this assumption and on data 

from three exposures indicates that the upper rocks 

sheared southeast relative to the lower rocks during the 

generation of gneiss early in the tectonic sequence. 

A root zone for thrusting may therefore have cut 

through the pluton near the base during its intrusion; 

alternatively, perhaps the pluton spread laterally along 

a thrust zone and was subsequently cut by continuing 

movement of the thrust. The thickness of the pluton would 

require that this happened deep in or even near its base. 

The flow lines in the Bitterroot lobe have been important in the generation of structures in the late-liquid and 

postconsolidation stages of the batholith (fig. 3.2). The 

granitic rock apparently rose into or crossed a structural 

discontinuity, subsequent movement on which deformed 

it during igneous flow and similarly during solid-state 

flow. The flow is inferred to have been extensional, a conclusion supported by the presence of curving extension 

fractures perpendicular to the flow lines (fig. 3.2). This 

flow pattern implies thinning during deformation. The 

fact that the gneissic foliation in the granitic rock dies 

out quickly to the north is consistent with the 30° northwesterly plunge of the flow lineation and with the concept of its relation to thrusting. Thrusting would appear 

to be inconsistent with the evidence for extension, but this 

may be only apparent. The work of Williams and Chapman (1983) has shown that thrust sheets are subject to 

extension behind the advancing ductile bead associated 

with the thrust tip. Extension is therefore an integral part 

of the thrusting process. 

FLOW TRENDS AND ARCHING 

A compilation of average flow directions in the granitic 

rocks of the project area is given in figure 3.2. These 

average directions are based on both primary and secondary flow features, as well as blastomylonite where it has 

developed. The flow trends form a fan-shaped array opening out to the southeast, within the field of secondary 

strain, and stand more or less perpendicular to fractures 

interpreted to have originated as tensional openings. Flow 

lines plunge northwestward at about 30 o over most of the 

area, but to the east they become horizontal near the 

Bitterroot arch (cross section in fig. 3.2), and then plunge 

about 30 o southeast in the Bitterroot gneissic front, to 

the east of the arch. 

The structure section given in figure 3.2 is based on the 

disposition of flow lineation across the mapped interval. 

As noted above, the intensity of the strain fabric becomes 

less to the west, in areas that the section shows are structurally higher in the pluton. The mylonite zone at the 

eastern margin of the pluton (cross section in fig. 3.2) is 

cut off by erosion and is not certainly seen again farther 

west. If it continued westward as a low-angle zone, it must 

have become highly discordant both to the secondary flow 

lineation and to the blastomylonite. 

As the arch shown in the section appears to be due wholly 

to upward arching of the secondary flow lineation and allied 

features, the batholith must have been 30 km or more 

thick prior to arching. Therefore, the former situation of 

the mylonite zone to the west becomes a major question. 

A mylonite zone about 30 m thick trends north-south 

and dips shallowly to the west in the Jeanette Mountain 

quadrangle. More detailed mapping should investigate the 

possibility that this mylonite is the westward continuation of the frontal mylonite. The structural position seems 

about right (fig. 3.2). If this is the continuation, then the 

mylonite zone appears to be dying out westward within 

the igneous mass. 

Evaluating other work in terms of the cross section 

given in figure 3.2, it is notable that Wiswall's (1979) proposed base of the batholith south of the cross section lies 

along the trend of the arch. That proposed base appears 

to be consistent with the regional structure inferred in 

this paper. On the other hand, Toth (1981) proposed that 

the Bear Creek and Cub Creek plutons south of the cross 
3. STRUCTURAL GEOLOGY AND PETROLOGY OF A PART OF THE BITTERROOT LOBE 55 

section occupy a position near the roof of the batholith. 

This now seems unlikely, as those plutons lie just west 

of the axis of the arch and therefore deep within the 

batholith, possibly near its base. 

The area affected by secondary strain is part of the 

Bitterroot gneiss dome described by Hyndman (1980) as 

a metamorphic core complex. The boundaries of the 

wedge-shaped area defined in this project (fig. 3.1) outline 

the complex; but that is very approximate, as the boundaries are based on extrapolation of a few known points 

on the north and south sides. 

DEVELOPMENT OF SUBAREA B OF THE EASTERN SECTOR 

Flattened mineral grains are scarce and much less 

strongly developed in the rocks of subarea B than in those 

of adjoining areas. At a late stage in consolidation, when 

magma could still move as a partly solidified mush, the 

continuing flow became of secondary character, but strain 

was along directions at only small angles to the earlier 

flow, so that the two flow foliations have much the same 

geometry. Tangential compression seems to have predominated, and flattening was much subdued; what little 

flattening was produced parallels the steep axial planes 

of the folds and is not related to the low-angle flattening 

of adjacent strain fields. 

The primary and secondary flow foliation were deformed in a gigantic fold during early flow, perhaps 

because of impingement of an eastward-moving sheet of 

granitic rock against an eastern buttress or wall-the 

eastern batholith contact. 

THE BITTERROOT GNEISSIC FRONT 

Near the eastern margin of the eastern sector, the foliation dips become shallow and the trends change (fig. 3.4) 

to nearly north-south, cutting abruptly across the trends 

of subareas A and B. Structural analysis yields a neardomical symmetry, which may be of only local significance 

but which raises the possibility of a series of coalescing 

marginal domes concealed in the tilted foliation along the 

eastern side of the Bitterroot lobe. This domical symmetry 

and the easterly trending lineation are features common 

to the metamorphic core complexes of the western 

cordillera (see, for example, Davis and Coney, 1979). 

Hyndman (1980) applied a metamorphic-core-complex interpretation to the Bitterroot lobe, and his discussion is 

particularly germane to the rocks of the gneissic front. 

The possibility of thrusting in this area, mentioned above, 

suggests another way for the lineation to form besides 

the extensional mechanism envisioned by Davis and 

Coney; during the preparation of this paper, Mattauer and 

others (1983) proposed a thrusting model for metamorphic core complexes. 

Foliation is stronger in the gneiss of the Bitterroot frontal zone than in the gneissic granite to the west, raising 

the possibility that they are of different ages. This possibility is supported by Bickford and others (1981, p. 453), 

who show "cataclastic granite" to have an age in the 

range 49 to 46 Ma. This contrasts with their age of 55 Ma, 

mentioned above, for medium-grained granitic rocks of 

the Bitterroot lobe interior. Petrographic data, on the 

other hand, suggest that the rocks are all the same age. 

Both rock types formed earlier than the pegmatite, which 

intruded along their foliation, and the fabric in both types 

is virtually identical. The problem remains unsolved. 

To the west, the dips change with the lineation (fig. 3.2) 

from eastward to westward over an arch and in a narrow transition zone (fig. 3.4) give way either to the steep 

northeast trends of the giant fold or to the northwest 

trends of axial symmetry. This geometry may mirror the 

dual constraints of deep shear zone and eastern contact. 

FINE-GRAINED DIKES 

Dikes and sheets of fine-grained granite, granodiorite, 

and tonalite are present in the central and eastern sectors, but were not seen in the western sector. Except for 

subarea B, where none was seen, they are most abundant 

in the eastern sector, which is deep in the intrusion. Their 

distribution is thus asymmetrical and appears to correlate 

with depth in the intrusion and perhaps with strain intensity. The fine-grained sheets were emplaced into planes 

of secondary shear during secondary strain and were 

metamorphosed during continuing movement. Finegrained dikes were emplaced at the same time and were 

also metamorphosed. These fine-grained igneous rocks 

may have come from the same sources as the main-stage 

batholithic rocks; an alternative possible source was stillunconsolidated batholithic material deep in the magma 

chamber. Some fine-grained dikes were injected into extensional fractures perpendicular to the stretching lineation of the secondary strain field, showing in yet another 

way the extensional character of the secondary strain. 

ANORTHOSITE II 

The miarolitic andesine anorthosite pipe in the Blodgett 

Mountain quadrangle and the miarolitic andesine anorthosite dike in the Tenmile Lake quadrangle are products 

of the late-stage emplacement of anorthosite magma. 

These features may bear the same relation to anorthosite 

I as the fine-grained granitic dikes do to the granitic rocks 

of the Bitterroot lobe. Because the anorthosite is medium 

grained, it must have been emplaced either (1) as a crystal 

mush largely crystallized at a deeper level before emplacement or (2) at a deep level in rocks of high temperature, 

in which crystallization occurred slowly. Like the fine-
56 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

grained granitic dikes, the andesine anorthosite is cut by 

blastomylonitic shears, themselves developed while the 

country rock was hot. The anorthosite bears a strong 

resemblance to the anorthosite I rock of the Blodgett 

Mountain quadrangle and may well be a late product of 

the same magma chamber(s). 

BLASTOMYLONITE 

Blastomylonite is rare in the western sector, becomes 

more common in the central sector, and is abundant in 

the eastern sector. Its abundance thus reflects the intensity of the secondary strain foliation. Petrographic 

features of the blastomylonite suggest simultaneous shear 

and recrystallization along restricted zones of shear. 

Where crystallization outlasted deformation, the fabric 

is strongly annealed. In a few places, deformation 

outlasted crystallization, and polygonal arcs are only 

imperfectly developed. Some blastomylonitic shears in the 

western sector parallel secondary foliation and thus 

mirror the axial symmetry of the secondary strain event. 

Therefore, they represent renewed strain late in that 

event. In some places, blastomylonite plagioclase is slightly more sodic than plagioclase equilibrated in the secondary strain event, indicating slightly lower temperature. 

In most places, however, both kinds of plagioclase have 

the same composition, reflecting little or no change in 

temperature between the two events. Biotite also formed 

in blastomylonite, although its grain size is reduced 

because of shear. 

In the central sector, blastomylonite shears have axial 

symmetry about an axis that plunges toward the northwest, parallel to the strain axis of the secondary strain 

event (fig. 3.3). The two sets of structural features formed 

sequentially in response to the same stress field; they 

evidently shared a common slip line defined by the secondary strain axis, which is also the girdle pole. Rodding 

developed in some blastomylonite because of more intense 

axial strain. Low-angle blastomylonite shears in arrays 

that strike more or less perpendicular to the trend of the 

girdle axis are small, ductile thrust faults. At one of these 

the upper halves of two phenocrysts were thrust 5 em to 

Cie southeast; therefore, the strain has an element of 

monoclinic symmetry. Steeper shears that strike more or 

less parallel to the trend of the girdle axis function as 

small, ductile tear faults. 

Blastomylonite zones are more abundant in the eastern 

sector than in the other sectors, although they were not 

found in subarea B. They are present throughout the other 

subareas, and they resemble those of the other sectors 

in all but two respects. First, in thin section they are seen 

to terminate at large, lenticular grains of quartz or quartz 

ribbons, although fine-grained mica of the blastomylonite 

continues as inclusion trains in the large quartz grains. 

Large quartz grains transected by blastomylonite have 

evidently recrystallized during or after blastomylonite formation in a "quartz-ribbon event," possibly related to 

somewhat higher temperature. Second, the blastomylonite shears are mostly low-angle; blastomylonite appears 

to have formed in low-angle, ductile thrust faults more 

abundantly than to the west, and only a few associated 

ductile tear faults are present. Though the ductile shear 

zones of these thrusts average only 8 em in thickness, they 

can be traced along their strike for kilometers. Their 

generally westward dip, except where arched, is consistent with a thrust mechanism, though the original dips 

may have been substantially changed by the younger 

arching. 

MYLONITE 

Mylonite is rare in the western and central sectors but 

more important in the eastern sector, where it occupies 

low-angle shears. Streaky lineation trends westnorthwest, like lineation in blastomylonite and secondary 

gneiss. Tops moved east for the most part, as indicated 

by shear steps on the shear surfaces. Local conjugate 

shears in the eastern sector show the opposite direction 

of movement. All of these shears have developed mylonite 

that is similar both petrographically and in outcrop. The 

change in stress response, from thrusting farther west 

to conjugate, high-angle reverse shears in the east is 

ascribed to late-stage arching of gneiss in the Bitterroot 

frontal zone. Arching associated with active uplift 

changed the orientation of effective stress across the top 

of the arch. 

CONCLUSIONS 

1. Tonalite and a later granite/granodiorite massif, 

together making up the Bitterroot lobe of the Idaho 

batholith, were emplaced into metasedimentary country 

rock (possibly of Beltian age) as a sheetlike body that 

spread to the east. 

2. The granitic massif, originally more than 30 km 

thick, is vertically zoned, from average calcic granodiorite 

at the base to average granite at the top. 

3. Classic magma tectonics occurred in the southwestern part of the lobe; the resulting steep primary flow 

structures were cut and strongly modified by low-angle 

secondary foliation and lineation within the pluton. 

4. The emplacement of the pluton was synkinematic 

into an active thrust zone, resulting in the formation of 

a metamorphic core complex. 

5. Three phases of deformation yielded gneiss (high 

pressure and temperature), blastomylonite (intermediate 

pressure, high temperature), and mylonite (lower pressure, intermediate temperature). 
3. STRUCTURAL GEOLOGY AND PETROLOGY OF A PART OF THE BITTERROOT LOBE 57 

6. Arching occurred during and following deformation, 

and the base of the massif now lies at and near the surface on the east side. 

7. More than 30 km of relative vertical movement (east 

side up) is implied on the north-south lineament along the 

eastern margin of the Bitterroot lobe. This movement may 

reflect the combined effects of diapirism and arching. 

REFERENCES CITED 

Anderson, A.L., 1952, Multiple emplacement of the Idaho batholith: Journal of Geology, v. 60, p. 255-259. 

Armstrong, R.L., 1974, Magmatism, orogenic timing, and orogenic 

diachronism in the cordillera from Mexico to Canada: Nature, v. 247, 

p. 348-351. 

__ 1975a, Precambrian (1500 m.y. old) rocks of central Idaho-the 

Salmon River arch and its role in cordilleran sedimentation and tectonics: American Journal of Science, v. 275-A, p. 437-467. 

__ 1975b, The geochronometry of Idaho: Isochron West, no. 14, 

p. 1-288. 

Balk, Robert, 1937, Structural behavior of igneous rocks: Geological 

Society of America Memoir 5, 177 p. 

Bates, R.L., and Jackson, J.A., eds., 1980, Glossary of geology (2nd ed.): 

Falls Church, Virginia, American Geological Institute, 751 p. 

Bell, T.H., and Etheridge, M.A., 1973, Microstructure of mylonites and 

their descriptive terminology: Lithos, v. 6, p. 337-348. 

Berg, R.B., 1969, Petrology of anorthosite of the Bitterroot Range, 

Montana, in Origin of anorthosite and related rocks: New York State 

Museum and Science Service Memoir 18, p. 387-398. 

Berthe, D., Choukroune, P., and Jegouzo, P., 1979, Orthogneiss, 

mylonite and non coaxial deformation of granites: the example of 

the South Armorican shear zone: Journal of Structural Geology, v. 1, 

p. 31-42. 

Bickford, M.E., Chase, R.B., Nelson, B.K., Shuster, R.D., and Arruda, 

E.C., 1981, U-Pb studies of zircon cores and overgrowths, and 

monazite: Implications for age and petrogenesis of the northeastern 

Idaho batholith: The Journal of Geology, v. 89, p. 433-457. 

Bittner-Gaber, Enid, 1983, Geology of selected migmatite zones within 

the Bitterroot lobe of the Idaho batholith, Idaho and Montana: 

Moscow, Idaho, University of Idaho, Ph.D. dissertation, 173 p. 

Bradley, M.D., 1981, Geology of northeastern Selway-Bitterroot 

Wilderness Area, Idaho and Montana: Moscow, Idaho, University 

of Idaho, M.S. thesis, 51 p. 

Chase, R.B., 1968, Petrology of the northeast contact zone of the Idaho 

batholith, northern Bitterroot Range, Montana: Missoula, Montana, 

University of Montana, Ph.D. dissertation, 158 p. 

Chase, R.B., Bickford, M.E., and Tripp, S.E., 1978, Rb-Sr and U-Pb 

isotope studies of the northeastern Idaho batholith and border zone: 

Geological Society of America Bulletin, v. 89, 1325-1334. 

Chase, R.B., and Hyndman, D.W., 1977, Mylonite detachment zone, 

eastern flank of the Idaho batholith: Missoula, Montana, Department of Geology, University of Montana, Rocky Mountain Section, 

Geological Society of America Field Guide no. 1, 31 p. 

Cheney, J.T., 1975, Kyanite, sillimanite, phlogopite, cordierite layers 

in the Bass Creek anorthosite, Bitterroot Range, Montana: Northwest Geology, v. 4, p. 77-82. 

Davis, G.H., and Coney, P.J., 1979, Geological development of the cordilleran metamorphic core complexes: Geology, v. 7, p. 120-124. 

Garmezy, Lawrence, and Sutter, J.F., 1983, Mylonitization coincident 

with uplift in an extensional setting, Bitterroot Range, MontanaIdaho [ abs. ]: Geological Society of America Abstracts with Programs, 

v. 15, no. 6, p. 578. 

Goodge, J.W., 1983, Fold reorientation and quartz microfabric in the 

Okanogan dome mylonite zone, Washington: Kinematic and tectonic 

implications: Missoula, Montana, University of Montana, M.S. thesis, 

65 p. 

Greenwood, W.R., and Morrison, D.A., 1973, Reconnaissance geology 

of the Selway-Bitterroot Wilderness Area: Idaho Bureau of Mines 

and Geology Pamphlet 154, 30 p. 

Hansen, Edward, 1971, Strain facies: New York, Springer-Verlag, 207 p. 

Hyndman, D.W., 1980, Bitterroot dome-Sapphire tectonic block, an example of a plutonic-core gneiss-dome complex with its detached 

suprastructure, in Crittenden, M.D., Jr., Coney, P.J., and Davis, 

G.H., eds., Cordilleran metamorphic core complexes: Geological 

Society of America Memoir 153, p. 427-443. 

Johnson, B.R., 1975, Migmatites along the northern border of the Idaho 

batholith: Missoula, Montana, University of Montana, Ph.D. dissertation, 120 p. 

Langton, C.M., 1935, Geology of the northeast part of the Idaho batholith 

and adjacent region in Montana: Journal of Geology, v. 43, p. 27-60. 

Larsen, E.S., and Schmidt, R.G., 1958, A reconnaissance of the Idaho 

batholith and comparison with the southern California batholith: U.S. 

Geological Survey Bulletin 1070-A, 33 p. 

Leischner, L.M., 1959, Border-zone petrology of the Idaho batholith in 

vicinity of Lolo Hot Springs, Montana [abs.]: Geological Society of 

America Bulletin, v. 70, p. 1782. 

Lindgren, Waldemar, 1904, A geological reconnaissance across the 

Bitterroot Range and Clearwater Mountains in Montana and Idaho: 

U.S. Geological Survey Professional Paper 27, 123 p. 

Mattauer, Maurice, Collot, Bernard, and Van den Driessche, Jean, 1983, 

Alpine model for the internal metamorphic zones of the North 

American cordillera: Geology, v. 11, p. 11-15. 

Morrison, D.A., 1968, Reconnaissance geology of the Lochsa area, Idaho 

County, Idaho: Moscow, Idaho, University of Idaho, Ph.D. dissertation, 126 p. 

Nold, J.L., 1974, Geology of the northeastern border zone of the Idaho 

batholith, Montana and Idaho: Northwest Geology, v. 3, p. 47-52. 

Reid, R.R., 1959, Reconnaissance geology of the Elk City region, Idaho: 

Idaho Bureau of Mines and Geology Pamphlet 120, 74 p. 

__ 1984, Structural geology and petrology of a part of the Bitterroot lobe of the Idaho batholith, Idaho County, Idaho, and Missoula 

and Ravalli Counties, Montana: U.S. Geological Survey Open-File 

Report 84-517, 117 p. 

Reid, R.R., Bittner, Enid, Greenwood, W.R., Ludington, Steve, Lund, 

Karen, Motzer, W.E., and Toth, Margo, 1979, Geologic section and 

road log across the Idaho batholith: Idaho Bureau of Mines and 

Geology Information Circular 34, 20 p. 

Reid, R.R., Morrison, D.A., and Greenwood, W.R., 1973, The Clearwater orogenic zone: a relict of Proterozoic orogeny in central and 

northern Idaho, in Belt Symposium 1973: Department of Geology, 

University of Idaho, Moscow, Idaho, v. 1, p. 10-56. 

Ross, C.P., 1928, Mesozoic and Tertiary granitic rocks in Idaho: Journal of Geology, v. 36, p. 673-693. 

__ 1952, The eastern front of the Bitterroot Range, Montana: U.S. 

Geological Survey Bulletin 974-E, p. 135-175. 

__ 1965, Idaho batholith [abs.]: Geological Society of America Special 

Paper 82, p. 343. 

Ryder, R.T., and Scholten, Robert, 1973, Syntectonic conglomerates 

in southwestern Montana: their nature, origin, and tectonic 

significance: Geological Society of America Bulletin, v. 84, 

p. 773-796. 

Stauffer, M.R., 1964, The geometry of conical folds: New Zealand Journal of Geology and Geophysics, v. 7, p. 340-34 7. 

Streckeisen, A.L., and others, 1973, Plutonic rocks: Classification and 

nomenclature recommended by the lUGS Subcommission on the 

Systematics of Igneous Rocks: Geotimes, v. 18, no. 10, p. 26-30. 

Toth, M.l., 1981, Petrology and structure of the east-central Bitterroot 
58 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

lobe of the Idaho batholith [ abs. ]: Geological Society of America Abstracts 

with Programs, v. 13, no. 4, p. 228. 

__ 1983, Reconnaissance geologic map of the Selway-Bitterroot 

Wilderness, Idaho County, Idaho, and Missoula and Ravalli Counties, Montana: U.S. Geological Survey Miscellaneous Field Studies 

Map MF-1495-B, scale 1:125,000. 

Williams, Graham, and Chapman, Timothy, 1983, Strains developed in 

the hangingwalls of thrusts due to their slip/propagation rate: a 

dislocation model: Journal of Structural Geology, v. 5, no. 6, 

p. 563-571. 

Williams, L.D., 1977, Petrology and petrography of a section across the 

Bitterroot lobe of the Idaho batholith: Missoula, Montana, University of Montana, Ph.D. dissertation, 251 p. 

Wiswall, C.G., 1979, The base of the Idaho batholith [abs.]: Geological 

Society of America Abstracts with Programs, v. 11, p. 305. 

Yates, R.G., 1968, The trans-Idaho discontinuity: International 

Geological Congress, 23d, Prague, Czechoslovakia, Proceedings, v. 1, 

p. 117-123. 
4. EMPLACEMENT OF THE MAIN PLUTONS OF THE BITTERROOT LOBE 

OF THE IDAHO BATHOLITH 

By C. GIL WISWALL1 and DoNALD W. HYNDMAN2 

ABSTRACT 

Mineralogic, chemical, and field relationships indicate that the Bitterroot lobe of the Idaho batholith was formed by magmas that underwent 

limited diapiric rise and may have been emplaced by lateral spreading 

to form one or more mushroom-shaped bodies at mesozonal to catazonal 

levels. The main-phase granitoid intrusion exhibits a high degree of 

mineralogic, textural, and chemical homogeneity and so may comprise 

a single comagmatic suite of rocks. Although the country rocks and the 

sequence of deformation are similar throughout the Bitterroot lobe, contact relationships and styles of deformation on the northern and southern 

margins of the batholith show different levels and modes of emplacement. The northern exposures exhibit somewhat shallower levels of 

emplacement near the top of the intrusion. The base of the batholith 

crops out in the south as granitoid cappings on ridge crests underlain 

by migmatites across a near-horizontal contact. Several lines of evidence 

indicate limited rise of the magmas and their emplacement near the 

source area, but conditions evidently varied sufficiently to produce different styles of emplacement. In the north, the country rock was competent enough to cause either lateral or upward forcible injection along 

foliation near the batholith margins, whereas in the south the more ductile country rock flowed around the laterally spreading magmas. The 

different levels of exposure appear to be in sharp structural contact, 

possibly as a result of postemplacement tectonic modification associated 

with the formation of the Bitterroot dome. 

INTRODUCTION 

The Late Cretaceous Idaho batholith occupies much of 

northern Idaho and part of western Montana. The 

39,000-km2 area is divided into a northern (Bitterroot) 

lobe and a southern (Atlanta) lobe by the northwesttrending Salmon River arch (Armstrong, 1975). The 

larger (25,000 km2) Atlanta lobe trends north-south, 

whereas the smaller (14,000 km2) Bitterroot lobe trends 

northwest-southeast. 

The Idaho batholith lies approximately 600 km from the 

present Pacific coastline, in contrast to the more typical 

1Department of Geology and Astronomy, West Chester University, West 

Chester, Pennsylvania 19383. 2Department of Geology, University of Montana, Missoula, Montana 59812. 

distance of 150 to 200 km observed for the Coast Range 

Plutonic Complex of British Columbia, the Sierra Nevada 

batholith of California, and the Andean plutonic arc. The 

batholith lies northeast of the Blue Mountains island arc, 

one of the numerous Mesozoic allochthonous terranes that 

make up the western margin of North America (Coney 

and others, 1980). Although the timing and process of accretion are poorly understood, eastward thrusting of the 

arc sequence against the North American craton was complete prior to the emplacement of the oldest plutons of 

the Idaho batholith (Vallier and Engebretson, 1984). The 

resultant boundary is exceedingly sharp both geologically and isotopically, suggesting the removal of the intervening crust, either by subduction or by lateral transport. 

From the timing of accretion and the spatial relations 

between the Blue Mountains island arc and North America, it can be inferred that a subduction zone existed to 

the west of the allochthonous terrane before intrusion of 

the Idaho batholith began in the Late Cretaceous (Vallier 

and Engebretson, 1984). Therefore, although the clearly 

related pair of accretionary prism and magmatic arc 

recognized in other batholithic terranes (for example, 

Sierra Nevada) has not been established, it seems probable that the Idaho batholith resulted from Mesozoic 

subduction. 

Detailed studies of other major batholiths reveal composite bodies consisting of numerous individual plutons 

belonging to several comagmatic suites. Magmatic suites 

and individual plutons commonly show concentric zoning, 

both compositional and textural, to varying degrees of 

complexity (see, for example Bateman and Chappel, 1979). 

Both modeling experiments (Ramberg, 1970) and field 

studies (Pitcher and Berger, 1972; Green and Kridoharto, 

1975; Sweeny, 1975, 1976; Koller, 1978; Sylvester and 

others, 1978; Wiswall, 1979a) indicate that plutons deeper 

than the brittle/ductile transition rise diapirically and may 

undergo lateral spreading at the level of emplacement to 

form mushroom-shaped bodies. Our studies suggest that 

59 
60 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

most of these general observations apply to the Idaho 

batholith. 

Comparison of the northern Idaho batholith with tectonically similar terranes has been hampered by a lack 

of data. Prior to 1976, detailed work was limited to the 

margins of the batholith. More recently, several detailed 

studies (fig. 4.1) concentrated on the interior of the Bitterroot lobe (Williams, 1977; Wiswall, 1979a, b; Hyndman, 

1984). We summarize these studies here in order to present an overview of the tectonics of emplacement and to 

provide a framework for further study. Perhaps the most 

surprising conclusion from this detailed work is that 

individual plutons are difficult to identify. In fact, petrographic and chemical evidence suggest that approximately 

FIGURE 4.1. -Geologie map of the Bitterroot lobe of the Idaho batholith showing locations of studies referred to in text. Area designated 

Wiswall (1979b) is shown in figure 4.2; section along Loehsa River designated Hyndman (1984) appears in figure 4.5. 
4. EMPLACEMENT OF THE MAIN PLUTONS OF THE BITTERROOT LOBE 61 

40 percent of the volume of Cretaceous rocks in the Bitterroot lobe may belong to one comagmatic suite. Borderzone relationships, structural analyses, and the form of 

the batholith contact indicate that diapiric rise followed 

by lateral spreading at the level of emplacement produced 

one or more mushroom-shaped bodies. However, apparent 

differences in the level of exposure along the margin suggest that the batholith underwent postemplacement 

modification. 

The data presented here are primarily from five detailed 

studies distributed around the lobe (fig. 4.1): Williams 

(1977) and Hyndman (1984) reported on exposures in the 

Lochsa River canyon (herein referred to as the northern 

Bitterroot lobe, northwestern and northeastern border 

zones); Wiswall (1979a, b) studied metamorphic and associated granitoid rocks in the upper Selway River canyon 

(southern Bitterroot lobe, southwestern border zone); and 

Chase (1973) described relationships in the northern 

Bitterroot Range (eastern Bitterroot lobe and border 

zone). 

ACKNOWLEDGMENTS 

Financial support for this work was provided by a 

Penrose grant from the Geological Society of America and 

by National Science Foundation grant 76-84399. We 

thank P.C. Bateman, R.B. Chase, and R.R. Reid for 

reviews of the manuscript. 

EXPLANATION 

~ Valley fill sediments (Tertiary, Quaternary) 

~ Basalt (Miocene) 

~ Felsic volcanic rocks (Eocene) 

~ Granitoid intrusive and volcanic rocks (Tertiary) 

)\~ ~ Granitoid intrusive rocks (Cretaceous) 

~ granodiorite 5~~~~~ quartz diorite 

~ Volcanic rocks, undifferentiated (Mesozoic) 

I p£m I Metasedimentary rocks, paragneiss, and orthogneiss 

(Precambrian) 

I p£bl Sedimentary rocks of the Belt Supergroup 

/ Fault / Thrust fault (sawteeth on 

overthrust side) 

1::§> tjJ.V Mylonite 

FIGURE 4.1.-Continued. 

COUNTRY ROCKS 

Country rocks of the Bitterroot lobe include both 

metasedimentary rocks, generally correlated with the 

Proterozoic Belt Supergroup, and orthogneisses interpreted as pre-Belt basement (Greenwood and Morrison, 

1973; Armstrong, 1975; Hyndman, 1983a). The Belt metasedimentary rocks consist of abundant muscovite-biotite 

quartzofeldspathic schists and gneisses, which locally contain kyanite and almandine, and subordinate quartzite, 

micaceous quartzite, and diopside-hornblende-plagioclase 

gneiss. Pods or boudins of amphibolite are present in some 

places. The lithologic succession suggests correlation with 

the Prichard and Wallace Formations and the Ravalli 

Group (see, for example, Greenwood and Morrison, 1973). 

Belt metasedimentary rocks in the area mapped by 

Wiswall (1979a) are largely anatectic migmatites. Composition of the pre-Belt basement varies widely, from 

felsic to ultramafic. Rock types include hornblende quartz 

monzonite augen gneiss, quartz diorite to diorite orthogneiss, meta-anorthosite, and small bodies of amphibolite, 

metagabbro, and hornblende pyroxenite. The distributions 

of these pre-Belt rocks is poorly known, but they appear 

to form a discontinuous band adjacent to the southern and 

western margins of the Bitterroot lobe (Hyndman and 

Williams, 1977). Layered meta-anorthosites found along 

the northern and eastern margins (Berg, 1964; Cheney, 

1972; Chase, 1973; Jens, 1974) suggest basement exposures there as well. 

Before the formation of the batholith, a regional dynamothermal metamorphic event affected the country rocks 

over an area extending several kilometers beyond the 

present batholith contacts. The rocks adjacent to the 

batholith are in the upper amphibolite facie~ (sillimanite 

zone), but the degree of metamorphism decreases away 

from the batholith margins in a broadly concentric 

fashion. Intense deformation accompanying the metamorphism produced a penetrative schistosity subparallel to 

compositional layering; this schistosity controlled subsequent deformation and, in some areas, emplacement of 

the intrusive phases. Although complexly deformed, the 

schistosity maintains a consistent northwesterly strike 

and steep easterly dip. 

The country rocks around the Bitterroot lobe exhibit 

several generations of mesoscopic structures. Correlation 

between areas, however, is difficult because detailed 

structural analyses are widely scattered. The sequences 

of structural development in the north and southwest are 

similar, but they differ in style. In general, four generations of folds are recognized (Greenwood and Morrison, 

1973; Chase and Johnson, 1977; Wiswall, 1979a). The 

earliest event (F1), which accompanied sillimanite-muscovite to sillimanite-orthoclase regional metamorphism, produced rare, tight to isoclinal, sharp-crested, similar-style 
62 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

folds in compositional layering and a penetrative axialplane schistosity (S1). Because many F 1 structures are 

rootless intrafolial folds, strong transposition is inferred 

to have accompanied deformation (Wiswall, 1979a). The 

second deformation (F 2) occurred after the peak of metamorphism but resulted from movements similar to those 

during F 1· The structures of F 2 are tight shear folds that 

deform S1 and show characteristics of both similar and 

parallel styles. The extremely ductile behavior exhibited 

by the less competent units indicates fairly high temperatures during deformation; however, a new penetrative s 

surface was not produced. Concentric features are confined to more competent units where ductility contrasts 

were low. Where F 1 and F 2 folds are present in the same 

outcrop, they are roughly coaxial. Although there is no 

orientation pattern of the fold axes due to subsequent 

deformation, axial-plane orientations are generally 

reflected in the northwesterly trending structural grain 

as a result of the control exerted by S1 during deformation. Thus, F1 and F2 appear to represent a progressive 

deformation accompanying metamorphism, in which the 

end of F 1 marks the peak of metamorphism. 

The age of this tectonic cycle is uncertain because of 

the complex thermal history of the area. Norwick (1972) 

documented a nondeformational burial metamorphism 

which raised the lower part of the Belt section in Precambrian time to the lower amphibolite facies. The metamorphism associated with F1 and F2 deformation is superimposed on this earlier event. Geochronometric data from 

early tonalite and quartz diorite bodies on the northern 

margin of the batholith indicate an age of 85 to 70 Ma 

(Grauert and Hofmann, 1973; Ferguson, 1975; Chase and 

qthers, 1978). Many of these intrusive bodies exhibit a 

strong foliation parallel to that in the country rocks, and 

their emplacement must have occurred either before or 

synchronous with metamorphism. Reid (chapter 3) has 

concluded that the tonalite bodies on the southern and 

western margins of the Bitterroot lobe are synkinematic. 

Furthermore, the control exerted by S1 on the emplacement of the main-phase granite/granodiorite demonstrates the completion of F1/F2 deformation before the 

onset of major intrusion. Thus, although the timing of 

metamorphism is as yet uncertain, inferred relationship 

between metamorphism and generation of the Idaho 

batholith favors a Cretaceous age. 

In the southwest, Wiswall (1979a) reported a third set 

of structures (F3), which range in scale from microscopic 

to macroscopic and exhibit a variety of forms including 

both similar and parallel styles. The macroscopic structure is revealed by the outcrop pattern of the country rock 

units (fig. 4.2) and the regional foliaton orientation (fig. 

4.3). A large northwesterly plunging antiform straddles 

the Selway River in this area. Mesoscopic structures fall 

into two major categories: (1) earlier, tight, symmetric 

to asymmetric, similar-style fold sets accompanied by 

reorientation of S1 in fold crests and minor recrystallization to form a weak axial-plane schistosity (S3); and (2) 

parallel-style, asymmetric, drag-type folds that consistently show a northeast-over-southwest sense of movement 

away from the batholith and local transposition at the 

crests forming fracture surfaces with fairly consistent 

northeasterly dips of 35 o to 60°. These groupings are not 

mutually exclusive, and individual folds may show characteristics of both groups. Fold orientation data from the 

southwestern mesoscopic F3 structures are shown in 

figure 4.4. Fold axis orientations of F3 form a cluster that 

is nearly coincident with the pole to the great circle defined by poles to foliation (see figs. 4.3, 4.4A), but the 

distribution of poles to F3 axial planes indicates subsequent deformation. Domainal analysis of F3 axial-plane 

orientations demonstrates that deformation was progressive and continued through the cooling of the rocks. 

The earlier similar-style folds, which exhibit ductile 

behavior, were continually reoriented about the axis of 

folding. The drag-type folds fall into domains (based on 

constant axial-plane orientation) that reflect the macroscopic antiform. Thus the F3 deformation in this area 

began at elevated temperature, proceeded through cooling and increasing competence of the country rocks, and 

culminated in the formation of the macroscopic antiform. 

Data presented below suggest that this deformation is 

related to emplacement of the batholith. 

To the north, along the lower Selway and Lochsa 

Rivers, F3 structures are parallel-style tight to open folds 

and fold sets which deform F2 structures (Greenwood and 

Morrison, 1973; Wiswall, 1979a; Hyndman, unpublished 

data). These folds do not exhibit the same diversity of 

form observed to the southwest. In particular, the earlier 

similar-style structures are missing. Granitoid sills are 

commonly concordant with F3 structures; however, no 

causal relationship has been established for this concordance. Data from other areas suggest that F3 is related 

to batholith emplacement. 

The fourth deformation (F 4) was similar in the north 

and southwest and produced open, broad-crested, parallelstyle folds and warps. Recognition of F 4 folds is based 

on warping of F3 fold axes. The F 4 folds are upright, 

horizontal or gently plunging structures that trend northsouth in the north (Greenwood and Morrison, 1973) to 

west-northwest in the southwest (Wiswall, 1979b). The 

orientations and textural evidence indicate that these 

folds formed by vertical movements and cataclasis at low 

temperature and therefore presumably postdate batholith 

emplacement. 

The structural evolution of the northeastern margin, as 

described by Chase (1973, 1977) and Chase and others 

(1978), bears a superficial similarity to that of the 

southwestern and northern borders, but they are not 
4. EMPLACEMENT OF THE MAIN PLUTONS OF THE BITTERROOT LOBE 

N 

0 5 

KILOMETERS 

EXPLANATION 

Cretaceous and Tertiary intrusive rocks Precambrian metamorphic rocks 

/ 

" " " 

,?/ 

/ 

" 

Main-phase granite/granodiorite (Cretaceous) 

Border-zone granodiorite (Cretaceous) 

Granite (Tertiary) 

Contacts A' Foliation 

Mapped 

Inferred \ldaxis with plunge 

Uncertain direction 

From Greenwood and 

Morrison ( 1973) 

II] 

r1 

w 

~ 

Metatonalite orthogneiss 

Migmatite 

Paragneiss 

FIGURE 4.2.-Geologic map of the upper Selway River (from Wiswall, 1979a). See figure 4.1 for location. Structure sections 

A-A' and B-B' appear in figure 4.6. 

63 
64 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

readily correlatable. Chase noted five generations of 

deformation, the youngest of which has been recognized 

only along the eastern margin of the Bitterroot lobe. The 

oldest deformation (F 1) seems to be the same as recognized elsewhere: similar-style isoclines accompanied by 

high pressure (kyanite-sillimanite) prograde metamorphism and transposition. Second-generation (F2) folding 

deformed the schistosity and produced a variety of structures by flexural-slip, which resulted in crenulation and 

a localized axial-plane schistosity. The F3 event produced 

an incipient axial-plane schistosity in some areas and was 

accompanied by transposition, which was strongest adjacent to the batholith and in mechanically weak pelitic 

schists. Chase and others (1978) attributed F1 and F2 to 

preemplacement deformation and F3 to deformation 

either slightly preceding or accompanying intrusion. 

Recognition of F 4 is based on domainal analysis of F3 

structures, although Chase (1973) noted that the variation in orientation of F3 fabric elements may be the result 

of inhomogeneous folding. The youngest deformation (F5) 

is a pronounced mylonitic event primarily affecting the 

easternmost section of the batholith and the adjacent 

country rocks. This deformation is attributed to eastsoutheast movement of the suprastructure (Hyndman and 

others, 1975; Hyndman, 1980) or to mylonitization 

associated with the formation of the Bitterroot dome 

(Chase and others, 1983). Chloritic breccia that overlies 

and is superimposed on the lower, eastern exposures of 

the mylonite may have formed during isostatic rise and 

doming of the mylonite (Hyndman, 1983b). 

N 

EXPLANATION 

• ~4 

~ 3-4 

IIIll1 2-3 

EJ 1-2 

D <I 

FIGURE 4.3.-Lower hemisphere projection of267 poles to foliation in 

country rocks in the upper Selway River area shown in figure 4.2. 

Contours at 1, 2, 3, and 4 percent per 1 percent area (from Wiswall, 

1979a). Bull's-eye marks the pole to the great circle defined by the 

foliation orientations. 

GRANITOID ROCKS 

The mineralogic and chemical variations within 

granitoid rocks of the Bitterroot lobe, especially in the 

interior parts of the batholith, are only now coming to 

light. Most workers identify a more mafic granodioritic 

border phase, which encloses the main-phase granodiorite 

to granite of the batholith interior. However, complex intrusive relationships, gradational contacts, and subtle 

mineralogic and textural variations, coupled with poor exposures, deep weathering and poor access have until 

recently hampered the delineation of individual intrusive 

A N 

N 

EXPLANATION 

• 28 

!:21 6-8 

IIIill 4-6 

E3 2-4 

D <2 

EXPLANATION 

• z2.4 

[2] 1.8-2.4 

[[]]] 1.2-1.8 

E3 0.6-1.2 

D <0.6 

FIGURE 4.4.-Lower hemisphere plots ofF3 structural features from 

the upper Selway River area shown in figure 4.2. A, Total of 95 F 3 

fold axes; contours at 2, 4, 6, and 8 percent per 1 percent area. B, 

Total of 60 poles to F 3 axial planes; contours at 0.6, 1.2, 1.8, and 

2.4 percent per 1 percent area. 
4. EMPLACEMENT OF THE MAIN PLUTONS OF THE BITTERROOT LOBE 65 

bodies. Williams (1977) and Hyndman (1984) recognized 

subunits within both the border phase and the main phase 

of the Bitterroot lobe along the Lochsa River. Toth 

(chapter 2) identified six separate plutons within the 

batholith interior. Available data suggest that the Bitterroot lobe may indeed be a composite body such as is seen 

in other batholithic terranes, but that the plutons are 

much larger and fewer than observed elsewhere. 

Comparison of the available petrographic data from the 

interior of the Bitterroot lobe leads to the conclusion that 

the main-phase granitoid intrusive is fairly homogeneous 

on a regional scale. In hand specimen, the rock is a 

massive to weakly foliated, white to light gray, mediumgrained equigranular hypidiomorphic biotite granodiorite 

to biotite granite. Mineralogically, it averages about 50 

percent plagioclase, 25 percent quartz, 20 percent K-feldspar, 5 percent biotite, and a trace of muscovite (see table 

4.1). Hornblende is absent. All workers report similar textural and compositional variations of an irregular and 

gradational nature: (1) variable grain size, commonly finer 

near the margins; (2) variable abundance (0 to 5 percent 

or more) of late, anhedral to subhedral, pale pink to white 

K-feldspar megacrysts, which are more abundant near the 

margins in some areas (Wiswall, 1979b); (3) localized irregular lenses of biotite granodiorite in otherwise 

homogeneous biotite granite (Williams, 1977); (4) local 

predominance of muscovite over biotite (Williams, 1977); 

(5) moderate to strong foliation, defined by subparallel 

biotite, schlieren, and streaks of biotite, feldspar, and 

megacrysts near the margins, fading to massive rock 

toward the interior; (6) rare isoclinal folds defined by 

schlieren, not necessarily confined to the margin; and (7) 

broad distribution of streaky schlieren, which are somewhat more mafic than the enclosing rock and are especially common near xenoliths or remnants of country rock. 

The origin of the foliation and schlieren is unclear. The 

foliation generally is parallel to and more strongly 

developed adjacent to contacts, and it may therefore have 

a magmatic origin. The diffuse, streaky, and discontinuous 

schlieren suggest incorporation of metasedimentary 

restite from the source area, digested xenoliths, or both. 

From the eastern border of the batholith toward its interior, xenoliths of quartzofeldspathic gneiss and pelitic 

schist become increasingly diffuse and finally are represented only by wispy concentrations of mafic minerals 

(Bickford and others, 1981). Old xenocrystic zircon in the 

granitoid rocks, noted in the same study (Bickford and 

others, 1981), suggests either preservation of refractory 

components from the source or contamination by crustal 

material. Hyndman (1984) concluded on the basis of 

chemical and trace-element studies that assimilation had 

not played a major role in the evolution of the batholith 

as a whole. This conclusion favors the restite origin for 

the schlieren and zircon, and it further suggests a source 

composition close to that of the pelitic schist and quartzofeldspathic gneiss that border the batholith. However, 

both the granodiorite of the border zone and the mainphase granites show microscopic evidence of latemagmatic or postconsolidation alteration, metamorphism, 

and deformation (Williams, 1977; Hyndman, 1984). For 

example, quartz commonly shows strong undulose extinction, subgrain development, and sutured grain boundaries. 

Plagioclase is only weakly twinned, and twin planes are 

commonly bent or broken. Primary biotite grains are bent 

or shredded, and a significant proportion of the biotite 

in the main phase is recrystallized or secondary. Reid 

(chapter 3) distinguished two generations of foliation: a 

primary flow foliation defined by parallel biotite flakes, 

schlieren, and platy xenoliths; and a secondary foliation 

defined by flattened quartz and feldspar grains. The two 

are commonly parallel. Therefore, the apparently conflicting characteristics may be explained by different mechanisms operating during and after crystallization of the 

magma. 

Granitoid rocks from the Bitterroot lobe that are inferred on other grounds to be Cretaceous consistently 

yield K-Ar ages ranging from about 56 to 37 Ma. These 

ages are commonly interpreted to indicate a widespread 

Eocene thermal event (Armstrong, 1974; Criss and 

Taylor, 1983) possibly associated with igneous activity 

that resulted in the formation of the Paleocene (?) and 

Eocene Challis Volcanics. Because Tertiary granite 

plutons and volcanic rocks are widespread throughout the 

Idaho batholith, this thermal event is a possible cause of 

the alteration observed in the Cretaceous intrusive suite. 

However, the distribution and timing of development of 

the secondary foliation (Reid, chapter 3) suggest that it 

is more likely a result of late-magmatic or postconsolidation vertical tectonism associated with formation of the 

Bitterroot dome, as discussed below. 

Subdivision of the main phase is generally based on the 

predominance of one or more of the textural or compositional variations cited above. For example, Hyndman (1984) 

distinguished four subunits of the main phase (M 1 through 

M4) on the basis of textural and compositional variation 

(fig. 4.5). Subunits M 1 to M3 are associated with complexities near the northwest border zone and are volumetrically small compared with M4, which constitutes approximately 60 percent of the main phase along the line of 

Hyndman's section (see fig. 4.5). 'Ibth (chapter 2), noting 

the similar major- and trace-element chemistry throughout 

the lobe, mapped individual plutons on the basis of texture, mineralogy, and crosscutting relationships. The most 

striking feature of her map (fig. 2.2) is the immense size 

of the Bear Creek and Grave Peak plutons. The large areal 

extent and general lack of mineralogic and textural variation in the Bitterroot lobe suggests that the bulk of it is 

formed by a few enormous plutons. 
66 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

TABLE 4.1.-Comparison of modes of major minerals for main-phase granitic rocks from the Bitterroot lobe of the Idaho batholith 

[Modes given in weight percent; anorthite (An) content of plagioclase in cation percent. tr., trace; n.d., no data] 

Area or unit------- Unit M4 SE. border 

Reference---------- Hyndman (1984) Wiswall (l979b) 

Number of analyses- 12 14 

Range Avg Range A VI!, 

Quartz------------- 16.2 -36.5 26.2 .21-31 26.1 

Plagioclase-------- 34.9 -59.2 48.5 32-63 49.8 

An content--------- n.d. n.d. (22)-(46) (33) 

K-fe1dspar--------- 10.8 -29.7 20.0 l-35 16.3 

Biotite------------ 1.9 - 6.8 4.4 3-ll 6 .l 

Muscovite---------- .03- 1.7 .5 0-1 .4 

This conclusion is supported by a striking lack of 

chemical variation. Hyndman (1984) found no clear 

chemical distinctions in either major or trace elements 

between units nor any discontinuities at the contacts 

between units that were defined on the basis of textural 

and mineralogic variation. Analyses of 12 representative 

samples showed only minimal variation in Si02, Al203, 

Na20, K20, and Fe oxides across the whole batholith. 'Ibth 

(chapter 2) arrived at a similar conclusion. Available majorelement analyses reported by other workers for rocks from 

the main body of the Bitterroot lobe exhibit average compositions generally similar to that reported by Hyndman 

(1984). From these sets of analyses (table 4.2), it may be 

noted that (1) average values of Si02 and Al203 vary less 

than 5 percent; (2) average Na20 and K20 values vary less 

than 10 percent; (3) if two analyses of unusually mafic 

rocks are excluded, then the average values of CaO and 

K20 also vary less than 5 percent; and (4) Fe oxides and 

MgO show significant variation. On the basis of available 

data, the chemical uniformity noted by Hyndman (1984) 

appears to extend over a significant part of the Bitterroot 

lobe. 

Tom Beal Park pluton NE. border Bear Creek pluton 

Williams (1977) Chase 0973) Toth (this volume) 

44 33 4 

Range Avg Range Avg Range Avg 

25-44 32.4 15-42 28 28-30.5 29.5 

10-45 32.1 27-69 42 42-43.5 42.4 

(19)-(30) (23) (20)-(36) (27) n.d. n.d. 

21-40 31.4 tr.-49 22 20-27 23.1 

tr.-7 3.0 1-16 5 3- 7.5 5.0 

tr.-3 1 0-7 1 

The internal form of the northern Bitterroot lobe is 

revealed by a weak to moderate foliation (fig. 4.5). Orientations of biotite, plagioclase, diffuse schlieren, and some 

K-feldspar megacrysts define a broad arch plunging 30° to 

50° northward. The arch is flat-topped through the central part of the batholith but has near-vertical limbs at its 

margins. Granodiorite in the western limb of the arch (unit 

M1 of Hyndman, 1984) contains 3 to 4 percent K-feldspar 

megacrysts along with irregular pegmatitic pods having 

rare, large miarolitic cavities. The crest of the arch (unit 

M2 and M3) contains as much as 5 percent K-feldspar 

megacrysts. The eastern limb of the arch consists of the 

biotite granodiorite to quartz monzodiorite (unit M4) that 

constitutes the bulk of the intrusive rocks in this area. 

Xenoliths and schlieren are less prevalent here than farther west. An intrusion breccia 250 m thick cuts this unit 

near the center of its outcrop width. 

Reid (chapter 2) has identified a foliation/lineation arch 

similar to that described above and concluded that it 

formed late in the crystallization history of the batholith 

in response to the formation of the Bitterroot dome. We 

do not believe that the arch along the Lochsa River and 

j.---Border zone---+~----------Main phase-------:~ I 

Lowell ~81+---- 82 I· 83+MI~+-M3-~•+I,,___ ______ M4---------->t4•-84-j 

l3miles b 10 o 10 o 10 o 10 

mileposts:= = ~ ~ !:!::? !:!::? :g: :g: ~rlJ_i~posts- ---

' ' \ : 1\\ . . . ""- + + 't + \ 1 biOtite-plagioclase 

\ + + + +' + + \ + +:I orthogneiss 

+ + \+ + -+ + "' + -+ \ + I + ... 1 ... + + + t >++I 

+ + + + + +'++'+I 

+ -+ + + + + + I 

EXPLANATION 

0 I 2 3 4 5 KILOMETERS Ei!) Granite and granodionte of the Idaho batholith 

0 2 MILES ,-J Trend of mineral foliation or schlieren 

-::._- Subhonzontal set of pegmat1te dikes 

No vertical exaggeration 

FIGURE 4.5.-Geologic cross section through the main-phase granitoid intrusive rocks and border zone of the northern Bitterroot lobe along 

the Lochsa River canyon (from Hyndman, 1984). See figure 4.1 for location. 
4. EMPLACEMENT OF THE MAIN PLUTONS OF THE BITTERROOT LOBE 67 

TABLE 4.2-Comparison of major-mineral analyses of main-phase granitic rocks from the Bitterroot lobe of the Idaho batholith 

[Values given in weight percent. n.d., no data] 

Area or Unit------- Unit M4 Tom Beal Park pluton 

Reference---------- Hyndman (1984) Williams (1977) 

Number of analyses- 9 5 

Range Avg Range Avg 

Sio2 --------------- 69.96-72.21 7l .13 66 .5 -73.7 70 .23 

Ti02 --------------- .18- .25 • 22 .17- .22 .20 

A1 2 o3 -------------- 15.23-16.11 15.73 15 .2 -17 .15 16.01 

Fe 2 o3-------------- .38- • 76 • 61 (as FeO) 

FeO---------------- .58- 1.04 • 74 1.67- 2.12 1.95 

MnO---------------- .02- .05 .025 .03- .05 • 04 

MgO---------------- .12- .47 .40 .34- .43 .JS 

CaO---------------- 1.46- 2. 72 1.92 1.90- 2.29 2.0 

Na 2 0--------------- 4.03- 4.76 4.46 3.82- 4.57 4.38 

K2 o---------------- 2.55- 4.16 3.69 3.05- 3. 93 3.45 

the one identified by Reid are related. Firstly, although 

the trends of the plunges of the two arches are similar, 

the crest of Reid's arch lies approximately 70 km to the 

east of the Lochsa arch. The arches could not, therefore, 

be connected by a continuous crest. More importantly, we 

believe that the arch along the Lochsa River is at least 

partly of primary origin; it is defined by primary flow 

features (primary mineral grains and schlieren). Furthermore, the arch structure is concordant with structure in 

the country rock. Therefore, the Lochsa arch appears to 

be a primary magmatic feature and reflects the original 

gross internal structure of the batholith in this region. 

However, the occurrence of late K-feldspar megacrysts 

oriented parallel to the arch (Hyndman, 1984) may indicate 

accentuation of the arch throughout the crystallization 

history. 

BORDER ZONE AND CONTACT RELATIONSHIPS 

The contact between the Bitterroot lobe and its country rocks varies significantly at different locations. In the 

north the margin is a generally steeply dipping zone of 

interdigitated intrusive and metamorphic rocks that contains numerous dikes, whereas in the south the contact 

lacks dikes and is gently dipping and sharp. Complex 

interaction between the earlier granitoid phases and the 

country rock make the northern margin of the batholith 

difficult to place and necessitates the subdivision of the 

batholith into the main-phase intrusions and a border-zone 

complex. 

The 10-km-thick northwestern border zone of the batholith is characterized by concordant sheets of granodiorite 

(older than the main phase of the batholith) separated by 

remnants of country rock; the granodiorite sheets constitute about 70 percent of the border zone (see Hyndman, 

1984, for a more detailed description of the border zone). 

The metamorphic rocks within the border zone show more 

NE. border Idaho batholith Bear Creek pluton 

Chase (1973) Larsen and Schmidt (1958) Toth (this volume) 

9 1 4 

Range Avg Avg Range Avg 

63.3-73.5 67 .s 71.66 70.5 -72.8 71.8 

.2- .6 .28 .22 .14- .29 .22 

14.4-16.7 15.41 15.48 14.7 -15.7 15 .33 

1.3- 5.3 2.7 .56 (as FeO) 

(as Fe 2 o

3 ) 1.45 1.22- 2.11 1.60 

n.d • n.d. .04 .02- .03 .02 

.6- 1.5 1.0 .47 .30- • 76 .55 

l. 5- 3.5 2.24 1.97 1.65- 2.46 1.89 

3.9- 5.4 4.60 4.34 3. 70- 4.80 4.08 

1.6- 5 .1 3.6 3.24 2.16- 4.41 3.68 

strongly developed gneissic layering than those farther 

west beyond the batholith contact. The granodioritic 

magmas were forcibly injected along foliation planes, and 

in many places they produced intrusion breccias. Xenoliths 

and schlieren are common within the border-zone intrusions, becoming more abundant northeastward toward the 

contact between the border-zone and main-phase granites 

(Wiswall, 1979b). Dikes are present throughout the border 

complex and range in composition from basalt through 

granite pegmatite or aplite. Hyndman (1984) distinguished four periods of dike emplacement that produced: 

(1) early, synmetamorphic, deformed, anatectic pegmatites; (2) tabular pegmatites occurring parallel to planar 

structural features; (3) tabular, subhorizontal pegmatites 

or aplites that intrude without regard for structural grain; 

and (4) contemporaneous or younger dikes, basalt to 

andesite, intruded along fractures and exhibiting chilled 

margins. Other significant features of the northwestern 

batholith margin include a general increase in grain size 

of granitoid rock within the border zone from west to east 

and a concentration of K-feldspar megacrysts near the 

contact between the border complex and the main mass 

of the batholith. Concordance between granitoid phases 

and the country-rock foliation suggests that the northwestern contact dips steeply eastward. 

To the southeast, near the confluence of Moose Creek 

with the Selway River (fig. 4.1), the northwestern border 

zone includes a block of country rock that is underlain by 

commingled Cretaceous and Tertiary granitoid intrusions 

(Wiswall, 1979b). The quartzites, pelitic schists, and 

gneisses of the block are heavily invaded by Cretaceous 

granitic sills injected along steep foliation planes, showing a relationship similar to that farther north along the 

Loch sa River. Angular xenoliths of country rock in the 

sills and apophyses of granite in the country rock imply 

brittle behavior during intrusion. Overprinting relationships indicate that the country rocks underwent three 
68 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

periods of deformation having the same sequence and 

style as F 1 through F 3 noted in the Loch sa area to the 

northwest. The gross structure of the block along the 

Selway River is an arch or dome underlain by granitoid 

rock. On their reconnaissance map, Greenwood and Morrison (1973) showed this block enclosed by Cretaceous 

igneous phases and partly concentric to the Tertiary 

Moose Creek pluton. These relationships suggest a roof 

pendant in the batholith that was deformed into an arch 

during the intrusion of the Tertiary pluton. 

The southwestern margin of the batholith exposed in 

the upper Selway River canyon is distinctly different from 

the northern margin. The intrusive rocks are divisible into 

two mappable units: a main intrusive phase and a 

marginal intrusive phase. The main phase is fairly 

homogeneous granite to granodiorite. Textures vary from 

weakly or moderately foliated at the contact with the 

marginal phase to massive within the main mass. 

Xenoliths, schlieren, pegmatite dikes, and K-feldspar 

megacrysts are rare to absent. The marginal phase is a 

medium-grained, moderately foliated granodiorite to 

quartz monzodiorite that contains K-feldspar megacrysts. 

It forms a shell of rather constant thickness along the 

main batholith contact, contains numerous schlieren, and 

has distinct layering defined by mafic concentrations 

parallel to foliation near the contact with the country rock. 

The contact between marginal phase and country rock 

here is sharp and lacks the country-rock remnants that 

are common to the north. Adjacent to the batholith, the 

contact dips steeply to the east and is concordant with 

the country rock, but it flattens outward perpendicular 

to the strike of the margin. There it becomes discordant 

to the regional foliation, and the granitic rocks, found only 

on ridge crests, overlie the migmatitic rocks exposed in 

the valley bottoms (fig. 4.6). Hamilton (1978) noted similar 

relationships in the Salmon River area within the Atlanta 

lobe. Therefore, the entire southwestern flank of the 

batholith may be in the form of a mushroom: a broad, 

steep feeder stem capped by a tabular mass of granitic 

rock. The floor of this tabular body is subhorizontal but 

not planar. Alternate interpretations include the ideas 

that (1) the southwestern flank may not be part of the 

main-phase plutons but may represent a large granitic sill 

or laccolithic intrusion (P.C. Bateman, written communication, 1984) and that (2) the country rocks may have 

intruded the batholith as a diapir. The consistent 

northeast-over-southwest sense of movement recorded in 

the F 3 meso scopic structures argues against the diapiric 

model. The form of the intrusive contact and the enveloping surface of the F3 antiform described above are 

concordant, indicating a syngenetic relationship. Furthermore, the sense of movement, the progressive nature of 

folding about an axis parallel to the batholith contact, and 

cooling during F3 folding all suggest participation of the 

country rocks during granite emplacement. Although sufficient information is not available to distinguish between 

the remaining two possibilities, we favor the interpretation of the planar body as a subhorizontal extension of 

the main-phase plutons. 

The northeastern border zone described by Chase 

(1973), like that in the Lochsa area, consists of numerous 

interdigitated granitic sills generally concordant with the 

country rock. Xenoliths are common near the contact, 

especially in regions where calcsilicate rocks are the host, 

but become rarer toward the interior of the batholith. 

However, the form of the contact here resembles the 

mushroom shape of that to the southwest. In the easternmost exposure, country rocks dipping 20° to 30° to the 

west concordantly underlie granitic rocks that cap the 

ridges. The contact steepens westward to approximately 

60° adjacent to the main batholith (Chase, 1973). Just to 

the north of the area mapped by Chase, the northeastern 

margin of the main body of the Bitterroot lobe shows a 

major reentrant along the Lochsa River (fig. 4.1). Reconnaissance mapping there shows the same topographic 

distribution noted to the south: intrusive rocks on ridge 

crests overlie country rocks, a relation suggesting a lowangle, west-dipping contact (Williams, 1977; Hyndman, 

unpublished data). 

FORM OF THE BITTERROOT LOBE 

The field relationships at various localities within the 

Bitterroot lobe indicate, as noted above, that it or some 

of its component bodies have the form of a mushroom. 

This shape apparently formed during more-or-less continuous intrusion that emplaced large bodies of magma 

of uniform composition. These interpretations are based 

on (1) the lack of variation in mineralogy and texture 

within the main granitic phase, (2) the lack of significant 

chemical variation across the batholith, (3) recognition of 

a foliation arch in the northern part of the Bitterroot lobe 

(Hyndman, 1984), and (4) exposure of a subhorizontal 

basal contact on the southwestern and northeastern 

borders. Mineralogic and textural relationships within the 

main-phase granitoid rocks indicate that the bulk of the 

magma was emplaced over a limited period of time, 

perhaps in a single protracted intrusive event (see Toth, 

chapter 2; Reid, chapter 3). 

Differences in the contact relationships from place to 

place may be due to the level of exposure. The northern 

Bitterroot lobe forms a central diapiric or mushroomlike 

body which is complexly interdigitated with the country 

rocks near its margins. Border-zone relationships suggest 

intrusion either laterally or from below under conditions 

conducive to parallel-style folding, brecciation, and at least 

in later stages,. fracturing. The preexisting structural 

grain (S1) exerted significant control on the emplacement 
4. EMPLACEMENT OF THE MAIN PLUTONS OF THE BITTERROOT LOBE 69 

of the magmas. Intrusion began with smaller granodiorite 

plutons (Hyndman, 1984) emplaced as sheets throughout 

the area now occupied by the batholith and border zone. 

This was followed by a major intrusive event that emplaced the main-phase granodiorite to granite pluton(s) 

into the center of the previous intrusive activity. The 

shape of the northwest-plunging foliation arch, numerous 

xenoliths near the crest of the arch, abundant injection 

migmatites, and the possible roof pendant along the 

Selway River suggest that the level of exposure in the 

north is near the top of the mushroom cap. Miarolitic 

cavities in late, crosscutting granite pegmatite dikes indicate that crystallization was not complete when the 

magmas attained relatively shallow levels in the crust (see 

Toth, chapter 2). The presence of late-magmatic or 

postconsolidation deformation and alteration suggest accentuation of the foliation arch during diapiric rise or 

uplift late in the crystallization history of the magmas. 

Miarolitic cavities in late pegmatites also support this conclusion. Although the presence of numerous schlieren in 

the granitoid rocks of the main mass suggest deep exposure near the source area, these can also be interpreted 

(/) 

a: 

w 

~ 

w 

~ 

z 

z 

0 

~ 

<( 

> w 

...J 

w 

sw 

A 

1800 

1200 

600 

B 

1800 

1200 

600 

...... 

Selway River 

I 

't~J - _.,._ 

\ 'f' ".) ' ' " .... , 

\ 

~ 

either as restite carried up with the magma from the 

source or as digested xenoliths derived from the overlying 

country rock and stretched out during rise and emplacement of the magma. Thus, exposures in the northwestern 

Bitterroot lobe seem to represent sites near the top of 

the intrusion, which attained shallow levels prior to complete crystallization. 

In contrast, the southwestern margin of the Bitterroot 

lobe along the Selway River exhibits (1) a more ductile 

style of folding characterized by relatively small domains 

of homogeneous deformation and a lack of significant 

brecciation; (2) a sharp boundary between the batholith 

and country rock; and (3) the presence of abundant anatectic migmatites. In addition, the intrusive rocks here clearly 

overlie the country rocks. These characteristics suggest 

a deeper level of exposure, near the bottom of the intrusive mass. The eastern border shows characteristics 

of both northwestern and southwestern borders; data 

from it suggest that the change in style may be gradational over the length of the batholith and that the 

batholith floor has a northerly dip. 

If the interpretations concerning the level of exposure 

--..... ..... 

----.._ -- -- ................. 

NE 

~ 

FIGURE 4.6.-Geologic structure sections across the upper Selway River. See figure 4.2 for location and explanation of symbols. No vertical 

exaggeration. 
70 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

are correct, we may broadly sketch the three-dimensional 

structure of the batholith, at least for the primary level 

of emplacement. The floor marks a distinct break between 

the underlying anatectic migmatites and early igneous 

phases and the overlying main-phase or flanking units. 

The exposed southwest and east edges of the floor are 

subhorizontal and gently undulating. The underlying country rocks are in structural concordance with the floor and 

form an antiform that nestles into the concavity between 

the mushroom cap and stem. At higher levels the form 

is more strongly controlled by the preexisting structural 

grain. Along the sides and roof, the magma injected 

forcibly either laterally or vertically, following planes of 

weakness in the somewhat brittle country rocks and 

producing the characteristics observed in the border zone 

to the north. 

If the batholith is a northerly plunging, mushroomshaped body having its basal part exposed in the south 

and its roof in the north, a section from the southwestern 

basal exposure northward along the Selway River should 

show the same gradational change in the level of exposure 

as is suggested by the intermediate characteristics noted 

on the eastern side. However, reconnaissance mapping 

by Wiswall (1979b) did not show the expected relationship; instead, the change in characteristics appears to be 

abrupt. The batholith floor is exposed for 12 km along the 

Selway River from White Cap Creek, north to North Star 

Creek where it is truncated by a Tertiary pluton (fig. 4.2). 

North of this Tertiary pluton, a Cretaceous quartz monzodiorite crops out for 10 km along the river and is in contact along its northern margin with the possible roof 

pendant described above. The quartz monzodiorite is 

mineralogically and texturally distinct from the intrusive 

phases to the south, and therefore a sharp structural 

discontinuity is apparently masked by the younger intrusive phase. Possible interpretations of this relationship 

are discussed below. 

TECTONICS OF EMPLACEMENT 

Field relationships of the intrusive rocks and the known 

shape of the batholith allow speculation about the tectonics of emplacement. Although the present data are still 

rather sketchy, they show significant similarities in areas 

throughout the Bitterroot lobe. 

Diapirs are classically pictured as teardrop-shaped 

bodies that rise by forcing the overlying rocks laterally, 

causing them to flow down along the margins of the diapir 

and fill in behind as the diapir passes. In a magmatic diapir 

the rising magma thus makes room for itself by displacing 

the overlying rocks, which move by viscous flow around 

the intrusion. The shape of the Bitterroot lobe and the 

structural events recorded in the country rocks support 

such a view of its rise in the crust. Although the lobe is 

presently in the form of a mushroom, this is probably the 

result of its final emplacement rather than its rise. The 

top of a teardrop-shaped pluton will reach its solidus 

temperature before the bottom and hence stop rising. 

Continued rise of the lower parts of the diapir will then 

cause lateral spreading at the level of emplacement, 

thereby altering the shape of the body to that of a mushroom. The movement pattern for F3 deformation in the 

southwest is consistent with this mechanism, for F3 was 

a progressive deformation that continually reoriented 

earlier formed structures about an axis parallel to the 

batholith margin and that proceeded through cooling of 

the rocks. The consistent northeast-over-southwest sense 

of movement recorded in the mesocopic F3 structures in 

the southwest is coincident with the direction of spreading 

there of the magmas at the level of emplacement. On the 

basis of these relationships, Wiswall (1979a) concluded 

that the F3 event records the rise and emplacement of 

the magma. Furthermore, the association of anatectic 

migmatites, which underwent F3 deformation, with the 

floor of the batholith suggests that some of the source 

rocks rose together with the magma. 

Several lines of evidence suggest that the magmas did 

not rise far and were emplaced at depths close to the 

source area in the lower continental crust. Contact relationships indicate mesozonal to catazonal levels of 

emplacement. In addition, Hyndman (1981) presented independent mineralogic and stratigraphic evidence for the 

northeastern border zone that suggested emplacement at 

0.5 to 0.6 MPa (5-6 kilobars) or a corresponding depth 

of 17 to 20 km. Phase considerations, assuming a 

muscovite-bearing parent, would permit only limited rise 

before intersection of the solidus. Finally, the batholith 

margin is generally parallel to regional metamorphic 

isograds in the horizontal plane and occurs within the 

sillimanite zone (Chase and Johnson, 1977). However, the 

batholith contact cuts across the isograds in the third 

dimension (Nold, 1968, 1974), a situation that could reflect 

limited rise of the magma out of the central thermal high 

into the overlying, somewhat lower grade rocks. 

Field relationships suggest that the mechanism of 

batholith emplacement varied from north to south, though 

this change may be a function of the vertical position 

within the intrusion. The main phase(s) throughout the 

lobe generally followed a similar path, rising as a limited 

number (three or four?) of very large diapirs. The path 

was centered in the highest temperature regions and controlled to some extent by the regional northwest-trending 

foliation. However, the differences in the nature of the 

batholith margin and the structural relationships noted 

above suggest a range in physical conditions at the site 

of emplacement. In the north, where structural control 

was strongest, early magmas were constrained to inject 

forcibly along foliation planes, thereby incorporating 

numerous remnants of country rock into the border zone. 

The country rocks were not ductile enough to flow early 
4. EMPLACEMENT OF THE MAIN PLUTONS OF THE BITTERROOT LOBE 71 

in the emplacement history, a conclusion supported by the 

presence of intrusion breccias in the northwestern border 

zone. The main phase was emplaced as a diapir, but the 

steep sides of the foliation arch indicate that the country 

rocks were competent enough to prevent lateral spreading. The change in style of emplacement with time may 

be a function of larger volumes of magma resulting from 

a higher degree of partial melting and more complete 

coalescence, coupled with a higher ductility in the country rocks because of continued intrusion. The numerous 

xenoliths present near the margins suggest at least minor 

stoping, but the absence of large roof pendants (with the 

possible exception cited along the lower Selway River) 

shows that stoping did not play a major role in batholith 

emplacement. 

In the south, on the other hand, the magmas were 

emplaced by lateral spreading. The country rocks were 

ductile enough throughout emplacement to flow passively in response to movement of the magma, producing the 

concordance of the macroscopic country-rock structure 

with the margin of the batholith and preventing forcible 

injection or brecciation. We propose that these observed 

differences in style of emplacement reflect the level of 

exposure within the batholith and deserve further study 

as indicators of relative depth. 

POSTEMPLACEMENT HISTORY 

Any tectonic interpretation of the northern Idaho 

batholith must consider the postemplacement history. 

Throughout this discussion we have emphasized the difference between the apparent levels of exposure in the 

Lochsa Canyon and in the upper Selway Canyon. We have 

noted that the eastern border seems to occupy an intermediate level and thus to suggest a gradational change, 

but an apparently sharp change occurs north along the 

Selway River from the exposed batholith floor. Reconnaissance work shows an equally abrupt change where 

granitic rocks reappear in creek bottoms westward from 

the exposed base (Wiswall, 1979b). The contact between 

the two levels of emplacement is not exposed, but outcrop patterns suggest it has an easterly dip. Thus both 

to the north and west, outcroppings of the batholith floor 

appear to be in sharp contact with batholith exposures 

that were originally emplaced at shallower levels. All of 

these rocks must, however, already have been at shallow 

levels by the Eocene, as shown by the clearly shallow 

Teritiary plutons. 

Chase and Talbot (1973) proposed a domal structure, 

the Bitterroot dome, to explain the eastern exposure of 

the batholith floor, maintaining that this dome is ringed 

by Tertiary plutons. Hyndman (1980) presented additional 

evidence on the extent and timing of dome formation and 

noted that placing the western margin is difficult because 

of the asymmetry in structural fabric. Many workers have 

observed cataclastic deformation that is late, vertically 

directed, and distinct from the mylonite exposed around 

the eastern half of the Bitterroot lobe (Chase, 1973; 

Greenwood and Morrison, 1973; Hyndman, 1980). Such 

deformation features are also present in the southwestern 

exposure of the batholith floor. 

Assuming that the Bitterroot dome is symmetrical and 

that the Bitterroot divide marks its crest, projection 

westward from the crest a distance equal to that of the 

eastern margin (approximately 20 km) would place the 

western margin in the vicinity of the Selway River (fig. 

4.1). The area of abrupt change in level of exposure along 

the Selway River may mark the zone of dislocation formed 

during isostatic rise of the dome following eastward movement of the suprastructure (Hyndman and others, 1975; 

Hyndman, 1980). If this is so, it can explain (1) the apparently abrupt changes in the level of exposure; (2) the 

presence and distribution of the exposures of the batholith 

floor; (3) the F 4 structural event; ( 4) the position of Tertiary plutons; (5) numerous steep fractures with down-dip 

lineations in the eastern half of the lobe (Hyndman, 1980); 

and (6) the course of the Selway River. It should be noted 

that the Tertiary ages obtained by Chase and others (1983) 

from main-phase (?) plutons suggest that the formation 

of the Bitterroot dome is not due to unroofing of the 

eastern Bitterroot lobe but rather may be the result of 

diapiric rise associated with emplacement of those 

plutons. This suggestion underscores the need for continued study and reflects the rapid evolution of geologic 

thought in this tectonically complex area. 

Beyond the scope of this paper, but worth consideration on a broader scale, is the possibility of a major 

prebatholith tectonic suture in the western border zone, 

perhaps in the area of the Selway River. Differences in 

lithology, depth of exposure, and structural style across 

this boundary could mark the edge of a microcontinental 

fragment that impinged on the west. The main-phase 

Idaho batholith might have been emplaced along this 

suture. In such a model, mylonites in the eastern Bitterroot lobe could be related to continued continental collision, in the style outlined by Brown and Read (1983), 

Mattauer and others (1983), and Rhodes and Hyndman 

(1984) for areas farther to the northwest. 

REFERENCES CITED 

Armstrong, R.L., 1974, Geochronometry of the Eocene volcanic-plutonic 

episode in Idaho: Northwest Geology, v. 3, p. 1-15. 

__ 1975, Precambrian (1500 m.y. old) rocks of central Idaho: the 

Salmon River arch and its role in cordilleran sedimentation and tectonics: American Journal of Science, v. 275-A, p. 437-467. 

Bateman, P.C., and Chappell, B.W., 1979, Crystallization, fractionation, 

and solidification of the Tuolumne Intrusive Series, Yosemite 

National Park, California: Geological Society of America Bulletin, 

v. 90, p. 465-482. 
72 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

Berg, R.B., 1964, Petrology of anorthosite bodies, Bitterroot Range, 

Ravalli County, Montana: Missoula, Montana, University of 

Montana, Ph.D. dissertation, 158 p. 

Bickford, M.E., Chase, R.B., Nelson, B.K., Shuster, R.D., and Arruda, 

E.C., 1981, U-Pb studies of zircon cores and overgrowths, and 

monazite: implications for age and petrogenesis of the northeastern 

Idaho batholith: Journal of Geology, v. 89, p. 433-457. 

Brown, R.L., and Read, P.B., 1983, Shuswap terrane of British Columbia: a Mesozoic "core complex": Geology, v. 11, p. 164-168. 

Chase, R.B., 1973, Petrology of the northeastern border zone of the 

Idaho batholith, Bitterroot Range, Montana: Montana Bureau of 

Mines and Geology Memoir 43, 28 p. 

__ 1977, Structural evolution of the Bitterroot dome and zone of 

cataclasis, in Geological Society of America Field Guide No.1: Rocky 

Mountain Section Meeting, Missoula, Montana, University of 

Montana, p. 1-24. 

Chase, R.B., Bickford, M.E., and Arruda, E.C., 1983, Tectonic implications of Tertiary intrusion and shearing within the Bitterroot dome, 

northeastern Idaho batholith: Journal of Geology, v. 91, p. 462-470. 

Chase, R.B., Bickford, M.E., and Tripp, S.E., 1978, Rb-Sr and U-Pb 

isotopic studies of the northeastern Idaho batholith and border zone: 

Geological Society of America Bulletin, v. 89, no. 9, p. 1325-1334. 

Chase, R.B., and Johnson, B.R., 1977, Border-zone relationships of the 

northern Idaho batholith: Northwest Geology, v. 6-1, p. 38-50. 

Chase, R.B., and Talbot, J.L., 1973, Structural evolution of the northeastern border zone of the Idaho batholith, western Montana [ abs.]: 

Geological Society of America Abstracts with Programs, v. 5, no. 6, 

p. 470-471. 

Cheney, J.T., 1972, Petrologic relationships of layered meta-anorthosites 

and associated rocks, Bass Creek, western Montana: Missoula, 

Montana, University of Montana, M.A. thesis, 112 p. 

Coney, P.J., Jones, D.L., and Monger, J.W.H., 1980, Cordilleran suspect 

terranes: Nature, v. 288, p. 329-333. 

Criss, R.E. and Taylor, H.P., Jr., 1983, An 180/160 and D/H study of 

Tertiary hydrothermal systems in the southern half of the Idaho 

batholith: Geological Society of America Bulletin, v. 94, no. 5, 

p. 640-663. 

Ferguson, J.A., 1975, Tectonic implications of some geochronometric 

data from the northeastern border zone of the Idaho batholith: 

Northwest Geology, v. 4, p. 53-58. 

Grauert, B., and Hofmann, A., 1973, Old radiogenic lead components 

in zircons from the Idaho batholith and its metasedimentary aureole: 

Carnegie Institution of Washington Yearbook, no. 72, p. 297-299. 

Green, Ronald, and Kridoharto, P., 1975, New evidence on the form 

of the granitic instrusions in New England, N.S.W.: Journal of the 

Geological Society of Australia, v. 22, p. 51-59. 

Greenwood, W.R., and Morrison, D.A., 1973, Reconnaissance geology 

of the Selway-Bitterroot Wilderness Area: Idaho Bureau of Mines 

and Geology Pamphlet 154, 30 p. 

Hamilton, Warren, 1978, Mesozoic tectonics of the Western United 

States, in Howell, D.G., and McDougall, K.A., eds., Mesozoic 

paleogeography of the Western United States: Society of Economic 

Paleontologists and Mineralogists, Pacific Section, Pacific Coast 

Paleogeography Symposium, no. 2, p. 33-70. 

Hyndman, D.W., 1980, Bitterroot dome-Sapphire tectonic block, an example of a plutonic-core gneiss-dome complex with its detached 

suprastructure, in Crittenden, M.D., Jr., Coney, P.J., and Davis, 

G.H., eds., Cordilleran metamorphic core complexes: Geological 

Society of America Memoir 153, p. 427-443. 

__ 1981, Controls on source and depth of emplacement of granitic 

magma: Geology, v. 9, p. 244-249. 

__ 1983a, The Idaho batholith and associated pluton, Idaho and 

western Montana, in Roddick, J.A., ed., Circum-Pacific plutonic terranes: Geological Society of America Memoir 159, p. 213-240. 

__ 1983b, Structure along the Bitterroot front, western Montana, 

in Hobbs, S.W., ed., Guide to Field Trips, Belt Symposium II: 

Missoula, Montana, University of Montana, p. 115-124. 

__ 1984, A petrographic and chemical section through the northern 

Idaho batholith: Journal of Geology, v. 92, 83-102. 

Hyndman, D.W., and Talbot, J.L., 1976, The Idaho batholith and related 

subduction complex: Geological Society of America Cordilleran Section Field Guide No. 4, 15 p. 

Hyndman, D.W., Talbot, J.T., and Chase, R.B., 1975, Boulder batholith: 

a result of emplacement of a block detached from the Idaho batholith 

infrastructure?: Geology, v. 3, p. 401-405. 

Hyndman, D.W., and Williams, L.D., 1977, The Bitterroot lobe of the 

Idaho batholith: Northwest Geology, v. 6-1, p. 1-16. 

Jens, J.C., 1974, A layered ultramafic intrusion near Lolo Pass, Idaho: 

Northwest Geology, v. 3, p. 38-46. 

Koller, G.R., 1978, Geophysical investigation of the Lexington batholith 

of west-central Maine: a subhorizontal, tabular, composite intrusion 

[abs.]: Geological Society of America Abstracts with Programs, v. 10, 

no. 2, p. 71. 

Larsen, E.S., and Schmidt, R.G., 1958, A reconnaissance of the Idaho 

batholith and comparison with the southern California batholith: U.S. 

Geological Survey Bulletin 1070-A, p. 1-33. 

Mattauer, Maurice, Collot, Bernard, and Van den Driessche, Jean, 1983, 

Alpine model for the internal metamorphic zones of the North 

American cordillera: Geology, v. 11, p. 11-15. 

Nold, J.L., 1968, Geology of the northeastern border zone of the Idaho 

batholith, Montana and Idaho: Missoula, Montana, University of 

Montana, Ph.D. dissertation, 159 p. 

__ 197 4, Geology of the northeastern border zone of the Idaho 

batholith: Northwest Geology, v. 3, p. 47-52. 

Norwick, S.A., 1972, The regional Precambrian metamorphic facies of 

the Prichard Formation of western Montana and northern Idaho: 

Missoula, Montana, University of Montana, Ph.D. dissertation, 

129 p. 

Pitcher, W.S., and Berger, A.R., 1972, The geology of Donegal: a study 

of granite emplacement and unroofing: New York, Wiley Interscience, 435 p. 

Ramberg, Hans, 1970, Model studies in relation to intrusion of plutonic 

bodies, in Newall, Geoffrey, and Rast, Nicholas, eds., Mechanism 

of igneous intrusion: Geological Journal (Liverpool) Special Issue, 

no. 2, p. 261-286. 

Rhodes, B.P., and Hyndman, D.W., 1984, Kinematics of mylonites in 

the Priest River ''metamorphic core complex,'' northern Idaho and 

northeastern Washington: Canadian Journal of Earth Sciences, 

v. 21, no. 10, p. 1161-1170. 

Sweeny, J.F., 1975, Diapiric granite batholith in south-central Maine: 

American Journal of Science, v. 275, p. 1183-1191. 

__ 1976, Subsurface distribution of granitic rocks, south-central 

Maine: Geological Society of America Bulletin, v. 87, p. 241-249. 

Sylvester, A.G., Oertel, Gerhard, Nelson, C.A., and Christie, J.M., 1978, 

Papoose Flat pluton: a granitic blister in the Inyo Mountains, California: Geological Society of America Bulletin, v. 89, p. 1205-1219. 

Vallier, T.L., and Engebretson, D.C., 1984, The Blue Mountains island 

arc of Oregon, Idaho, and Washington: an allochthonous coherent 

terrane from the ancestral western Pacific Ocean?, in Howell, D.G., 

Jones, D.L., Cox, Allan, and Nur, Amos, eds., Proceedings of the 

Circum-Pacific Terrane Conference: Stanford University Publications, Geological Sciences, v. 18, p. 197-199. 

Williams, L.D., 1977, Petrology and petrography of a section across the 

Bitterroot lobe of the Idaho batholith: Missoula, Montana, University of Montana, Ph.D. dissertation, 221 p. 

Wiswall, C.G., 1979a, Field and structural relationships below the Idaho 

batholith: Northwest Geology, v. 8, 18-28. 

__ 1979b, Structure and petrography below the Bitterroot dome, 

Idaho batholith near Paradise, Idaho: Missoula, Montana, University of Montana, Ph.D. dissertation, 129 p. 
5. MIGMATITE ZONES IN THE BITTERROOT LOBE OF 

THE IDAHO BATHOLITH 

By ENID BITTNER1 

ABSTRACT 

The Bitterroot lobe of the Idaho batholith is a complex of migmatite 

zones, plutons, and crosscutting mafic dikes. The migmatite zones are 

concentrated in the inner border of the batholith, where they make up 

about 70 percent of the rock; within the lobe itself they make up 10 to 

20 percent of the rock. Selected migmatite zones were studied on the 

western and southern borders of the lobe and in the central and eastern 

parts of the lobe interior. 

The migmatite formed by the intrusion, metasomatism, and partial 

assimilation of the country rock by tonalite, granodiorite, and granite 

of the batholith. The country rocks included in the migmatite zones consist of tonalite gneiss, quartzite, calc-silicate rocks, schist, quartzofeldspathic gneiss, and amphibolite. They show petrologic and chemical 

affinities to the Belt Supergroup and to other country rocks older than 

the batholith. 

The migmatite zones reflect a sequence of five deformational events: 

Db Jurassic and Early Cretaceous isoclinal and recumbent folding 

accompanied by igneous intrusions and kyanite-zone metamorphism; 

D2, emplacement of the Bitterroot lobe beginning at about 90 Ma and 

resulting in noncylindrical folding, hornblende hornfels contact metamorphism, and the alignment of migmatite screens east-west, parallel 

to the primary flow direction; D3, Late Cretaceous and Early Tertiary 

northwest-southeast extension that produced a secondary axial-symmetric annealing strain fabric, metamorphism of sillimanite-muscovite 

facies, and cylindrical folding; D4, postcrystallization axial-symmetric 

deformation resulting in blastomylonitization and mylonitization; and 

D5, late faulting and arching of the eastern part of the batholith. 

The migmatite units have retained their original east-west flow alignment and thus form good strain indicators. 

INTRODUCTION 

The Selway-Bitterroot Wilderness Area is located in 

north-central Idaho and western Montana and contains 

most of the Bitterroot lobe of the Idaho batholith (fig. 5.1). 

Migmatitic screens and remnants have been mapped 

around the border of the lobe and to a lesser extent in 

its interior. The screens predominantly consist of metasedimentary rocks that resemble the country rock, and 

they define specific structural trends. 

1Auburn University, Auburn, Alabama 36849. 

The main objectives of this study were: (1) to map the 

major migmatite zones within the Selway-Bitterroot 

Wilderness Area on a reconnaissance basis; (2) to describe 

the petrography, chemistry, and structure of the mimnatite zones and their associated igneous rocks; (3) to determine the relationship of the migmatite to the country rock; 

and (4) to propose a model to explain frequency of migmatite occurrence and the consistent east-west trend. 

BITTERROOT Cr e1r 

LOBE 0 

~ J -. v 

---~-STUDY \--i;:-.~--~) __ AR~A __ / _, 

Idaho 

0 

\ 

\ 

\ 

\ 

\ 

\ 

\ 

19 ' .. 20-..... 30 

Kilomet~rs / . / 

, ... "' 

FIGURE 5.1.-Location map, simplified from Toth (1982), for the SelwayBitterroot Wilderness (outlined by the heavy black line). Areas of 

reconnaissance mapping are designated P (perimeter country rock), 

North (country rock parallel to the Lochsa River), West, South and 

Central. 

73 
74 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

The Bitterroot lobe of the Idaho batholith extends 

100 km east-west and 50 km north-south. The SelwayBitterroot Wilderness Area lies within the Bitterroot lobe 

and covers an area of approximately 5,000 km2 (1.24 

million acres) in Idaho County, Idaho, and Missoula and 

Ravalli Counties, Montana. 

ACKNOWLEDGMENTS 

This study was condensed and revised from my doctoral 

dissertation done at the University of Idaho. Financial 

support for the dissertation was provided by Health 

Education and Welfare Fellowship, a Sigma Xi grant, and 

an Idaho Bureau of Mines grant. I thank R.R. Reid for 

review of the manuscript. 

PREVIOUS WORK 

Reconnaissance mapping of the Bitterroot and Clearwater Mountains was first done by Lindgren (1904), who 

identified metasedimentary rocks intruded by plutonic 

rocks. Anderson (1952) and Toth (1982) recognized the 

plutonic event as one of multiple intrusions. More complete mapping of the wilderness area by Greenwood and 

Morrison (1973) and more recently by the U.S. Geological 

Survey Wilderness Evaluation study (Toth, 1983) has 

delimited the eastern and western borders of the Bitterroot lobe and established major structural elements. 

Ross (1928) first suggested that the igneous complex 

of the Bitterroot lobe was of Mesozoic age. Subsequently, Tripp and Chase (1976), Chase and others (1978), and 

J.S. Stacey (oral commun., 1979) narrowed the age of the 

main stage of batholith intrusion to between 75 and 55 Ma 

by U-Pb dating methods. A later (Eocene) thermal event, 

identified by Armstrong (197 4, 1976), Ferguson (1975), 

and Armstrong and others (1977), has reset many of the 

K-Ar, fission tract, and Rb-Sr isotope ages, and so 

accurate dating of the rock by these methods is problematical. 

Reid and Greenwood (1968) and Reid and others (1973) 

defined the Clearwater orogenic zone as a zone of metasedimentary and metaigneous rocks enclosing the Bitterroot lobe. This zone extends north to the St. Joe River 

and east to the Bitterroot River valley but has undefined 

western and southern borders. Studies of the metasedimentary rocks by Reid (1959), Hietanen (1961b, 1963b, 

1968), Chase (1968), Nold (1968), and Harrison (1972) 

show regional low-grade metamorphism, accompanied by 

multiple preintrusion deformations, which increases 

through greenschist to amphibolite and hornblende hornfels facies nearer the contact zone of the batholith. 

Hietanen (1963a), Morrison (1968), and Bradley (1981) 

identified rocks of the border contact zone of the intrusion as diorite to quartz diorite. The initial diorite-quartz 

diorite phase was intruded by a later two-mica granite 

mapped by Williams (1977) and a granodiorite-quartz 

monzonite phase described by Leischner (1959), Ross 

(1963a, 1965), Wehrenberg (1968), and Toth (1982). 

Hamilton and Meyers (1967), Hamilton (1969), and Hyndman and Williams (1977) related the granodiorite-quartz 

monzonite phase of the Bitterroot lobe to Late Cretaceous 

subduction of the Pacific plate, which induced melting in 

the lower crust and subsequent plutonism. More recent 

works by Coney (1980) and Hyndman (1980) have related 

the formation of the eastern zone of the Bitterroot lobe 

to the 55- to 15-Ma event that produced the core complexes (Davis, 1975) along the western cordillera of the 

United States. Coney (1980) suggested that the two major belts along which these core complexes lie diverge at 

the Idaho batholith. 

Migmatite zones were mapped around the northern, 

northwestern, and southern contact zones of the Bitterroot lobe by Johnson (1975), Reid and others (1979), 

Wiswall (1979), and Allen (1980). Kell and Chase (1976) 

noted migmatite screens extending 8 to 10 km outward 

from the batholith. Johnson (1975) interpreted the 

migmatite zones as open systems of metasomatism and 

injection and inferred that only minor amounts of metamorphic differentiation and anatexis took place. 

METHODS OF INVESTIGATION 

This investigation was part of the United States Geological Survey Mineral Evaluation study of the SelwayBitterroot Wilderness Area. Three field seasons, from 

1978 to 1980, were spent in reconnaissance mapping on 

19 U.S. Geological Survey 7112-minute (1:24,000) topographic sheets. Special attention was paid to migmatitic 

terrain within designated study areas. Information on 

areas not mapped by the author was gathered from 

coworkers. 

Field work included chip and hand-specimen sampling, 

stream-sediment sampling and reconnaissance mapping. 

Modal analyses were done by point counting both thin sections, where grain size allowed, and slabs. Because of the 

remarkable heterogeneity of grain size and mineral content, thin-section modal analyses are considered to have 

little statistical viability. The modal analyses were plotted 

on a ternary quartz-potassium feldspar-plagioclase 

diagram, and the rocks were named according to 

Streckeisen and others' (1973) plutonic rock classification. 

Structural data were plotted on the lower hemisphere of 

the Wulff stereonet. 

Chemical analyses were done on 119 selected samples, 

including migmatite and Cretaceous and Tertiary intrusive rocks. Whole-rock analyses were run on 108 samples 

and trace-element analyses on 33. 
5. MIGMATITE ZONES IN THE BITTERROOT LOBE 75 

The whole-rock analyses were done at Washington 

State University on a PW 1410-70 Phillips X-ray spectrometer and at the U.S. Geological Survey, Denver. The 

samples were analyzed for 9 elements: Si, Al, Fe (total), 

Mn, Ca, Mg, K, Na, and P. The standard deviations for 

error in the whole-rock analyses range from 0 to 0.61 and 

average 0.13 for 8 samples and replicates. Silica yields 

the greatest deviations (between 0.13 and 0.61) and 

Fe203, MgO, CaO, Ti02, P20s, and MnO yield the lowest 

deviations (between 0 and 0.6). 

Trace elements were determined by neutron activation 

analysis at the Washington State University's nuclear 

reactor. The samples were run for 35 elements. 

GEOLOGIC AND TECTONIC SETTING 

In the Northwestern United States, sediments of the 

Middle Proterozoic Belt Supergroup were deposited on 

an older crystalline basement in a marine trough between 

1,500 and 1,100 Ma. After the deposition of the Belt 

Supergroup, orogenic events produced multiple episodes 

of metamorphism, deformation, and magmatic intrusion. 

A major period of deformation and plutonism in the 

Western United States began in the Mesozoic at about 

180 Ma, and recurring episodes of deformation and 

plutonism continued into the Cenozoic (Coney, 1972, 1981; 

Roberts, 1972; Drewes, 1978). The Idaho batholith was 

intruded during the mid-Sevier orogeny of the Late 

Cretaceous. Local plutonism continued, producing the 

"Tertiary pink granites," which intruded the Idaho 

batholith. A regional metamorphism (55 to 15 Ma), 

recognized in northwest Washington and Idaho (Coney, 

1980; Hyndman, 1980), formed metamorphic core complexes during post-Cretaceous extensional tectonics. Two 

core-complex belts formed, one developed on oceanic crust 

adjacent to the cordilleran batholiths, the other closer to 

the eastern cordillera on the original North American 

craton. These belts merge in the region of the Idaho 

batholith (Coney, 1980). The period of extension was 

overlapped and followed by volcanism, which formed the 

basalts of the Columbia and Snake River regions during 

the Pliocene (Hietanen, 1960). 

The Idaho batholith is separated into two lobes, the 

northern (Bitterroot) lobe and the southern (Atlanta) lobe 

(Armstrong, 1975). The Bitterroot lobe extends from the 

North Fork of the Clearwater River south to the crystalline Salmon River arch, which divides the batholith into 

the two lobes. 

The Bitterroot lobe is bounded on the north, east, and 

west by Precambrian crystalline basement rocks, Belt 

rocks, and the Early Mesozoic Riggins and Seven Devils 

Groups. The eastern part of the lobe was uplifted along 

the north-south-trending Bitterroot front fault zone. On 

the south the lobe is bordered by metasedimentary rocks 

and Eocene plutons across northwest-trending fault 

zones. 

The first phase of batholith intrusion induced an amphibolite grade of metamorphism in the adjacent Belt 

Supergroup. This first intrusive phase consisted of 

tonalite-quartz diorite, which now crops out on the northeastern, western, and southern borders of the lobe. Most 

of the lobe, however, is made up of gray granodiorite to 

monzogranite, which has been subdivided into separate 

plutons (table 5.1). Late Cretaceous and the Paleocene 

plutons are considered to be mesozonal in origin, and 

depths of emplacement from 11 to 19 km have been 

calculated by Williams (1977), W.E. Motzer (unpub. data, 

1979), and J.E. Selverstone (oral commun., 1979). Migmatite zones make up 20 to 70 percent of the border phase 

and inner boundary of the lobe but decrease to 0 to 15 

percent in the inner granodiorite. The major phase of the 

batholith emplacement was followed by intrusion of 

granodiorite and monzogranite, in the Paleocene and 

Eocene, and epizonal syenogranite to monzogranite in the 

Eocene (table 5.1). 

The long axis of the Bitterroot lobe lies along a northwest trend, concordant to the foliation of the surrounding 

country rocks. Steeply dipping structures were formed 

on the western border of the lobe during the emplacement 

of the batholith. Farther east, a secondary strain field of 

penetrative shearing has overprinted the primary fabric 

and developed axial symmetry about a shallowly plunging west-northwest axis. Late-stage uplift has formed a 

north-south-trending domal structure through the eastern 

section of the lobe (Reid, chapter 3; Wiswall and Hyndman, chapter 4). The eastern front is characterized by a 

zone of mylonitizaton, trending north-northeast, in which 

zones of ultramylonite to pseudotachylyte are interlaminated with mylonite. Faults that also trend northnortheast cut both the Cretaceous and Eocene granitic 

rocks. 

METAMORPHIC ROCKS 

AND MIGMATITE COMPLEXES 

GENERAL STATEMENT 

The wall rock of the Idaho batholith has been tentatively 

correlated by numerous workers (for example, Nold 1968; 

Greenwood and Morrison, 1973; Allen, 1980) as units of 

the Belt Supergroup. Wall rock on the northeastern and 

northern borders of the Bitterroot lobe are correlated with 

the Wallace Formation, the Ravalli Group, and the Belt's 

basal unit, the Prichard Formation (table 5.2). The 

Wallace Formation consists primarily of diopside gneiss 

intercalated with pelitic schist and gneiss, the Ravalli 

Group is noted for thick sequences of quartzite and pelite, 
76 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

TABLE 5.1.-Mapped individual plutons within the Bitterroot lobe of the Idaho batholith and their sequence of intrusion 

(Informal names and descriptions of plutons from Bradley (1981) and Toth (1983) except where otherwise noted. Sr rations (analysed by R.J. Fleck) from Reid and others (1978); 

radiometric ages from Ferguson (1975), Chase and others (1978), R.J. Fleck (written commun., 1982), and J.S. Stacey (oral commun., 1982). Order of emplacement mainly 

from intrusive contacts] 

Geologic age 

Eocene 

Paleocene 

and Eocene 

Late Cretaceous 

and Paleocene 

Radiometric 

age (Ma) 

51.3±10 

(K-Ar, Fleck) 

57±10 

(U-Pb, Stacey) 

66±10 

(U-Pb, Chase) 

81 

(Ferguson) 

0.7094 

.7089 

Rock description 

syenogranite-monzogranite 

granodiorite 

biotite, microcline 

An28-30 

granodiorite-monzogranite 

biotite, orthoclase 

An23-25 

diorite-gabbro 

biotite, augite 

An 57 

quartz diorite-granodiorite 

biotite, orthoclase 

An24-38 

tonalite-monzodiorite 

biotite+hornblende 

porphyritic 

granodiorite-monzogranite 

biotite, muscovite megacrysts 

aplite, pegmatite dikes 

monzogranite-granodiorite 

biotite+muscovite 

magnetite 

granite-granodiorite 

biotite, muscovite megacrysts 

(orthoclase) 

granite 

biotite, muscovite megacrysts 

aplite, pegmatite 

granite-granodiorite 

diorite 

biotite, hornblende 

An20-45 

Plutons mapped 

Running Creek 

Whistling Pig 

Fire Mountain 

Painted Rocks 

Lolo Springs1 

Burnt Ridge 

Piquett Creek 

White Sand Creek1 

Bushy Fork1 

Paradise 

Canyon Creek 

Bear Creek 

Canyon Lake 

Tom Beal Park 2 

Boulder Creek2 

Skookum Butte1 

migmatites, Black Canyon complex2 

Cretaceous (~;~~1, Id. 

( U-Pb, Stacey) 

82±10 

NE. border 

( U-Pb, Chase) 

Precambrian 

.715 

quartzofeldspathic gneiss, 

schist, amphibolite, 

quartzite, calc-silicate rocks, 

tonalite-monzogranite 

0.7130, 0.7196 tonalite-quartz diorite 

biotite, hornblende 

0.7147 metasedimentary rocks 

metaigneous rocks 

1rnformal name and description of pluton from Nold (1968). 2 Informal name and description of pluton from Williams (1977, 1979). 

Elongation 

of pluton 

east-west 

east-west 

northeast 

northeast 

east-west 

northeast? 

east-west 
5. MIGMATITE ZONES IN THE BITTERROOT LOBE 77 

TABLE 5.2.-Lithology of Belt Supergroup and Belt-correlative rocks from Coeur d'Alene to Clearwater and Idaho Counties, Idaho 

[Lithologies from Hietanen (1962), Harrison and Campbell (1963), Harrison (1972), Greenwood and Morrison (1973), Allen (1980), and Reid and others (1981)] 

Stratigraphic unit Lithology, relict structures, metamorphic grade 

Wallace Formation 

(carbonaceous, shallowwater marine, and littoral sediments) 

Ravalli Group 

(guartzitic and pelitic, 

shallow-water marine 

sediments) 

Revett Formation 

St. Regis and Burke 

Formations 

Prichard Formation 

(argillaceous, thinly 

interbedded, turbidite 

sediments) 

Coeur d'Alene area 

Interlayered clastic rocks, 

black limestone, and 

dolomite 

Oolites, stromatolites 

Chlorite-sericite zone 

greenschist facies 

Quartzite 

( 60-90 percent) 

Interlayered siltite, 

argillite, and sericitic 

quartzite; Na=K 

Cross beds 

Biotite zone 

greenschist facies 

Argillaceous upper unit, 

black pyritic lower unit, 

interbedded argillite and 

quartzite 

Biotite zone 

greenschist facies 

and the Prichard Formation is characterized by massive 

white quartzite intercalated with pelitic schist and gneiss. 

Remnants of older rocks included within the batholith 

closely resemble the wall rock. They are concentrated in 

migmatitic terrains near the margins of the plutons and 

occur as swarms of dismembered rafts and concordant 

slabs ranging from one meter to more than two kilometers 

in extent. Different degrees of assimilation have resulted 

in indistinct xenolith contacts, diffuse lamination, and 

mafic schlieren. Inclusions of tonalite and orthoamphibolite, as well as strongly sheared and recrystallized 

gneiss, also constitute part of the migmatitic terrain. 

PERIMETER OF THE BATHOLITH 

The northern, western, and southern borders of the 

Bitterroot lobe contain the highest concentrations of inSt. Joe area 

Argillite with interlayered dolomitic 

quartzite; minor 

scapolite 

Ripples, mudcracks, 

slumps, graded beds 

Greenschist to amphibolite facies; Na 

increases with grade 

Thick beds of 

quartzite 

Thin siliceous beds, 

dark argillite, sericitic quartzite and 

schist 

Greenschist to 

amphibolite facies 

Argillite, quartzite, 

and garnet schist, 

muscovite quartzite 

Amphibolite facies 

Clearwater and Idaho Counties 

Diopside-plagioclase gneiss interbedded with pelitic schist; 

biotite-garnet-sillimanite schist 

Sillimanite facies, hornblendehornfels facies 

Quartzite and pelite 

Quartzite interbedded 

with pelitic schist and gneiss, 

biotite-feldspar-garnetkyanite schist 

Amphibolite facies 

Pelitic quartzite, biotitemuscovite schist, white 

quartzite with kyanite, garnetmica schist, aluminosilicate schist 

Amphibolite facies 

eluded material, ranging from 20 to 70 percent. Screens 

of metamorphic rock are intruded by both concordant and 

discordant sheets of diorite, tonalite, and granite. The 

migmatite consists of these intrusive rocks and included 

screens and megabreccia blocks of quartzite, pelitic gneiss, 

calc-silicate rock, and amphibolite. The inclusions generally display sharp to partially diffuse contacts, but these are 

more disseminated and diffuse where assimilation has 

been more complete. 

In the northeastern corner of the lobe, the inclusions 

of country rock consist of calc-silicate, biotite quartzite, 

and minor amphibolite. Inclusions make up more than 20 

percent of the total rock and are predominantly mixed 

with tonalite gneiss (M.D. Bradley, oral commun., 1981). 

At lower elevations in this northeastern area, the percentage of inclusions is less, declining to less than 10 percent 

in massive gray tonalite gneiss. 
78 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

The migmatite screens of the western margin are made 

up mainly of two-mica schist, white quartzite, biotite 

gneiss, and minor calc-silicate inclusions, which occur as 

blocks or stretched xenoliths that have diffuse borders and 

show complex migmatitic folding. The screens are flat 

lying, as much as several kilometers long, and one-half 

to one kilometer wide. Extensive veining has broken up 

the sheets of metasedimentary rocks. The blocks of the 

brecciated rocks display sharp contacts and are typically 

separated by a few meters of igneous material. 

Migmatite screens near the southern margin of the 

Bitterroot lobe have been truncated on the southsouthwest by the Running Creek pluton and by other 

Eocene granite. 

Overall, igneous rocks constitute a slightly greater 

percentage of the volume of the migmatite. Mediumgrained granitic rock makes up as much as 35 percent of 

the total rock, pegmatite as much as 20 percent, and aplite 

as much as 10 percent. The biotite quartzite and gneiss 

range from 20 to 50 percent of the total rock, and schist 

forms 15 to 20 percent. 

CENTRAL AREA AND EASTERN FRONT 

From the central part of the lobe to the eastern front 

of the Bitterroot lVIountains, xenolith concentrations are 

low. The xenoliths are generally large blocks or 

disseminated rafts, which lie on an east-west trend and 

are as long as four kilometers. 

Along the eastern front, blocks and elongate xenoliths 

of quartz diorite and tonalite gneiss, biotite gneiss, biotite 

quartzite, amphibolite, and calc-silicate rocks occur in 

granodiorite and granite. The xenoliths make up 5 to 40 

percent of the total rock volume. 

QUARTZITE 

Massive white quartzite to finely laminated biotite 

quartzite contain relict sedimentary features which have 

withstood metamorphism; these features include crossbeds, pebbles, and graded bedding. The relict crossbeds 

are in a finer grained sugary-textured quartzite. The grain 

size generally ranges from less than 1 mm to 5 mm, but 

pebble conglomerates are interbedded with the white 

quartzite in the western border zone. These conglomerates contain quartzite pebbles, which have been orthogonally stretched and average 5 em in length, in a matrix 

of medium-grained quartz and mica. 

In the northeastern comer of the Bitterroot lobe, biotite 

quartzite in layers 0.5 to 1 em thick are interlayered with 

granodiorite to quartz monzonite layers as much as 1 m 

wide. The competent, fine-grained (1 to 2 mm) quartzite 

remnants have sharp contacts and show only minor 

assimilation. Where assimilation has been more complete, 

the igneous rock shows an increase in the concentrations 

of quartz and sillimanite (M.D. Bradley, oral commun., 

1980). 

QUARTZOFELDSPATHIC GNEISS 

The quartzofeldspathic gneiss units are granodioritic to 

tonalitic in composition. They were formed either by 

lit-par-lit injection of biotite quartzite and schist or by 

shearing-induced metamorphic differentiation and 

equilibration. 

The alternating leucosome amd melanosome layers in 

the gneiss range in width from 1 to 10 em. The layers consist of the following ranges of minerals (in volume 

percent): 

Quartz Plagioclase K-feldspar Biotite Muscovite 

Leucosome 29-30 

Melanosorne 41-60 

17-50 

10-31 

0-47 

0-11 

0-20 

17-26 

0-6 

1-10 

The tonalite gneiss, like the enveloping granitic rock, 

has high percentages of quartz and feldspar (An30-34). 

The gneiss is finely laminated. The neosome contains 

lepidoblastic biotite, muscovite, and sillimanite, minor 

amounts of sericite, and granoblastic quartz and feldspar. 

Biotite forms a mafic selvage parallel to the leucosome. 

The leucosome contains quartz, feldspar, and minor 

amounts of muscovite and biotite. 

SCHIST 

Biotite-muscovite-sillimanite schist containing varying 

amounts of garnet are interbedded with quartzite, injected 

lit-par-lit by granodiorite, and crosscut by aplite and 

pegmatite dikes. The schist units occur as blocks within 

the igneous host. Partial assimilation has affected the 

margins of the blocks, and in some cases near-total 

assimilation has left only remnant schlieren. 

A complex sequence of events is evident in the schist. 

The original schistose rock, composed of biotite, muscovite, and minor feldspar and garnet, was intruded by 

granodiorite parallel to schistosity. This rock was subjected to a second metamorphic event which concentrated 

sillimanite along shear zones and within muscovite and 

feldspar. Metamorphism was followed by cataclasis and 

the formation of sericite. During or after the sillimanitegrade metamorphism, feldspar was introduced or remobilized along and into fractures. The leucocratic layers 

of the schist generally consist of quartz, plagioclase, and 

pennine. The garnet, which is relict from the earlier 

metamorphic event, is rounded, fractured, and partly 

altered to biotite. 
5. MIGMATITE ZONES IN THE BITTERROOT LOBE 79 

CALC-SILICATE ROCKS 

The calc-silicate rocks occur as blocky inclusions, 1 to 

10 min length, interlayered with biotite quartzite and with 

granodiorite, tonalite, or quartz monzonite. Large screens 

commonly are broken and brecciated. 

Layers of plagioclase and quartz generally alternate 

with layers of amphibole or pyroxene in the calc-silicate 

rocks. Individual bands range from 1 to 10 em in width 

and are defined by composition, color, textural changes, 

and sills parallel to the layering. Two metamorphic events 

are indicated by the occurrence of diopside and younger 

amphibole and epidote. 

HORNBLENDE GNEISS AND AMPHIBOLITE 

Amphibolite occurs as rounded or teardrop-shaped 

xenoliths along the western border zone and in the northeast corner of the Bitterroot lobe. The inclusions range 

from 0.5 to 15 min length and from 0.25 to 10 min width. 

Smaller inclusons in the form of lenticular pods or boudins 

locally parallel the gneissic foliation. The rocks are garnetrich orthoamphibolite that developed reaction rims as 

much as 6 em wide adjacent to the host. 

Plagioclase ranges from An24 to An65 depending upon 

locality. Amphibole occurs as hornblende or tremolite or 

as hornblende cores with rims of actinolite. The zoned 

hornblende may indicate a second, weaker metamorphism, which induced the growth of actinolite. 

CHEMICAL TRENDS IN COUNTRY ROCKS AND INCLUSIONS 

The metamorphic mineral assemblages and major-oxide 

chemical data for the country rock are summarized in 

table 5.3. The mineral assemblages indicate metamorphism in the sillimanite-muscovite zone of the amphibolite 

facies (fig. 5.2A). R.R. Reid (oral commun., 1980) noted 

in the western country rock the presence of minor cordierite and andalusite, which may represent a contact 

metamorphic zone of hornblende hornfels facies (fig. 

5.2B). Kyanite occurs in the metamorphic rocks to the 

west and north of the batholith, as observed by Hietanen 

(1961, 1962) and by R.R. Reid (unpub. data, 1979), who 

described kyanite altered to muscovite in the rocks closest 

to the western contact. The relict kyanite indicates an 

earlier metamorphic event, possibly before emplacement 

of the batholith, under pressures greater than 400 MPa 

(Turner, 1981). Sillimanite and muscovite are found in all 

of the country rock, and their formation represents the 

last event in the amphibolite-facies metamorphism. 

The major-oxide chemical data and metamorphic 

mineral assemblages for the country-rock inclusions are 

given in table 5.4. The mineral assemblages represent a 

recrystallization event in the sillimanite-muscovite zone 

of the amphibolite facies (fig. 5.3). A previous metamorphic event is indicated by partly altered garnet in a matrix 

of recrystallized plagioclase, biotite, and quartz. 

Slight variations are present in the K20 and Na20 contents of the migmatite units. Near the southern border 

there is a slight decrease in Na20 and an increase in K20. 

The variation in N a20 in the metasedimentary rocks is 

similar to that in the granodiorite where the inclusions 

are found. 

Chondrite-normalized rare-earth-element (REE) data 

for the country rock and migmatite units are plotted in 

figure 5.4. The country-rock data show a linear trend with 

a substantial negative Eu anomaly (fig. 5.4A). The 

migmatite units show a similar linear trend (fig. 5.4B). 

The REE trends for the granite and tonalite gneisses in 

the migmatite correspond closely to the trends for the 

country rock. They differ, however, from those in the 

tonalite and granite units of the batholith, and two 

separate intrusive events are thus indicated. 

IGNEOUS ROCKS 

The rocks of the batholith fall mostly within the granite, 

granodiorite, and tonalite-trondhjemite fields of the 

Streckeisen and others (1973) classification (fig. 5.5). They 

are not distinguishable in the field. 

WESTERN BORDER 

Near the western border of the Bitterroot lobe, the 

granite, granodiorite, and tonalite of the batholith are 

partly intermixed and laterally gradational. Concordant 

sheets of country-rock tonalite gneiss were intruded by 

younger quartz diorite and tonalite. The older tonalite 

gneiss and other country rock occur in migmatite screens 

and as xenoliths. 

Granite and granodiorite vary from massive, weakly 

foliated rock which contain sparse but large plagioclase 

and potassium feldspar phenocrysts to sheared gneissic 

rocks. Biotite, feldspar, and a later generation of muscovite and sericite lie parallel to the shear planes. Most of 

the younger muscovite is an alteration product of biotite 

and plagioclase. Muscovite also fills fractures cutting 

across lepidoblastic biotite at right angles. Sericite is the 

predominant mica in highly sheared zones. Extension 

fractures are commonly filled with epidote and remobilized quartz. 

Included in the granite are xenoliths of granodiorite 

gneiss. These range from well-defined blocky inclusions 

to more intensely assimilated stretched lenses. The 

coherent nature of the xenoliths suggests at least partial 

consolidation prior to granite intrusion. In other areas the 

granodiorite-granite contacts are gradational. 
80 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

TABLE 5.3.-Chemical analyses for country rock ofthe Bitterroot lobe ofthe Idaho batholith and corresponding metamorphic mineral 

assemblages 

[Analyses given in weight percent. N (north), from Hietanen (1962); P (perimeter), from Reid (1980); LS, from Ross (1963b); H, from Hobbs and others (1965). Each analysis 

represents one sample unless other number (n) given, in which case the average is given] 

Rock type Calc-silicate Schist Quartzite Gneiss Amphibolite 

Sample area N p LS HH N N p N p p 

(n=2) 

Weight percent 

Sio2------- 63.2 64.7 58.8 47.9 67.3 90.6 78.9 75.5 70.3 50.4 

Al2 o3 ------ 11.2 10.9 9.0 7.0 15.6 5.6 8.6 11.7 ll.5 14.2 

Fe2o3------ .24 3.4 .28 .5 .9 .04 3.1 .21 6.7 13.9 

FeO-------- 3.4 1.9 3.9 6.8 .13 1.7 

MgO-------- 5.3 4.9 4.5 15.7 1.7 .15 1.8 2.4 2.7 6.1 

GaO-------- 14.8 8.5 21.1 21.6 .41 .5 .55 1.4 .89 9.1 

Na2o------- .76 .?. • 7 1.3 1.0 .85 2.5 1.4 3.6 1.2 2.6 

KzO-------- .12 2.0 2.3 1.6 3.2 .26 2.4 1.7 3.5 1.1 

T102------- .44 .43 .4 .19 2.0 .19 .47 .57 .91 1.8 

P 2 °5------- .09 .ll .08 .09 .81 .05 .06 .05 .15 .21 

MnO-------- .o1 .11 .04 .22 .04 .os .04 .02 .07 .17 --- Total---- 99.56 97.75 99.70 99.70 99.61 100.07 97.32 98.86 97.92 99.58 

Metamorphic mineral assemblages 

I I aJ aJ I aJ "0 I 4-J aJ 4-J t:: I aJ ·M t:: 'M +I I I aJ aJ 4-J t:: •M t:: aJ 1-1 aJ aJ r-l "0 •M ell r-l ell 4-J ell 4-J 4-J 

..0 •M (.) s (.) S ·M p.. •M •M 

t:: CJJ r-l ·M 0 •M 1-1 CJJ 4-J 4-J 1-1 p.. ell r-l 1-1 r-l 0"0 0 0 0 0 (.) r-l (.) r-l r-l r-l ·M •M ,.c:: ·M I •M oM •M ,.C:: aJ ..0 ,.0 

I "0 aJ CJJ s CJJ (.)~ I I aJ aJ I "0 I I I I 4-J 1-1 4-J 1-1 4-J 4-J aJ •M aJ aJ aJ aJ s aJ ell aJ ell 'M •M CJJ CJJ 4-J 4-J 4-J+J ::l t:: P..t:: p.. 

4-J r-l ell p.. •M 'M ·M •M •M 1-1 CJJ 1-1 CJJ 

0 0.--l 0 :> :> :> t:: CJJ ell "0 ell "0 'M P..CJ •M 0 0 0 ell CJJ 00 r-l 00 r-l 

..0 ell 0 "0 (.) (.) (.) s ell +I aJ I aJ I CJ,.C:: I 11) 11) CJJ •M 4-J aJ ~ aJ ~ 

aJ 11) 4-J aJ § § ;:lr-l 0 4-J 4-J 

t:: I 1-1 CJJ Sr-l P..•M S •M s 'M aJ 0 aJ ell I I I •M I t:: ::l t:: ::l r-l "0 I 4-Jr-l aJ aJ aJ 11) (!) 

~ •M 

~ 'M (.) ·M aJ ·M (.) 4-J 4-J 4-J I 4-J CJJ CJJ 0 11) 4-J (.) 0 ·ri ·ri •M aJ ·M •M CJJ •M CJJ 1-1 P..•M r-l ,.C:: 4-J 4-J 4-J+J+Jr-l ellr-l ell (.) 0..0 ell 4-J 0 0 0 ·ri 0.--l 4-Jr-l .... 'M ·M r-l (.) 1-1 oM •M ·M :> ·M ·ri 0 ·M 0 s "0 ell +I 0 ..0 ..0 ..0 0..0 CJJ P..CJJ p.. 

I I +I aJ I I I I (.) I I I I +I aJ aJ aJ 4-J aJ aJ aJ aJ CJJ aJ aJ aJ aJ aJ aJ 11) aJ CJJ CJJ ·M 11) CJJ 11) aJ CJJ § CJJ 4-J CJJ 4-J Ulr-l 

ell"' ell ellr-l ell ell ell 4-J ell ell •M ell ·M ell 0 r-l •M r-l r-l 0.--l aJ r-l 4-J r-l ·M r-l I r-l :> r-l :> r-l..O (.) 11) (.) (.) t:: (.) 4-J (.) aJ (.) 1-1 (.) aJ (.) 0 (.) 0 CJ •M 

0 P..O 0 •M 0 ·M 0 t:: 0 0 0 4-J 0 (.) 0 (.) o..c:: 'M O·M·M+J·ri..O •M 1-1 •M r-l ·M •M ·M CJJ ·M 11) •M p.. 

00 ·ri 0000(.) 00.--l OOell oo..c:: 004-J OO::l OO::l oos ell"' ell ell ell ell ell ell 00 ell (.) ell 0 ell s ell s ell ell r-l r-l r-l r-l r-l r-l r-l ·ri r-l r-l r-l p.. p..p.. p.. p.. p.. p..p::)p.. p.. p.. 

CENTRAL REGION TO EASTERN BORDER 

In the central regions of the Bitterroot lobe and east 

to its eastern border, the granitic rock is both intruded 

by and synkinematic with tonalite, quartz diorite, diorite, 

and pegmatitic and aplitic dikes. The granitic rock is porphyritic and has tabular to augen-shaped phenocrysts of 

potassium feldspar and plagioclase, which range in size 

from 1.5 to 7 em and totallO to 40 percent of the rock 

mode. The granitic rocks are massive to medium grained, 

gray, and lineated and have an average color index of 15. 

The fabric changes from the central regions, where the 
5. MIGMATITE ZONES IN THE BITTERROOT LOBE 81 

Muscovite 

8 Andalusite Andalusite Muse 

F C Wollastonite Oiopside 

FIGURE 5.2.-AKF and ACF diagrams applicable to the country rock 

of the Bitterroot lobe of the Idaho batholith (from Turner, 1981). 

A, Diagrams representing the sillimanite-muscovite zone of metamorphism. B, Diagrams representing the hornblende hornfels facies 

of metamorphism. 

rocks are generally massive and lineated, eastward 

toward the eastern border, where they are locally foliated. 

At higher elevations the rocks are layered and consist 

of pegmatite interfingered with medium-grained granitic 

rock and finely ground cataclastic rock. Layers vary from 

1 em to 2m in thickness. At lower elevations the rocks 

are more massive and lineated, and pegmatite dikes are 

absent. The metamorphic inclusions are in varying stages 

of assimilation from undeformed blocks to very stretched 

and folded incompetent bodies, and they become more 

abundant eastward and toward higher elevations. 

East from the Bitterroot crest the intensity of cataclasis 

increases. The medium-grained, massive granitic rocks 

become more equigranular and have alternating concentrations of quartz, feldspar, and biotite. The grain size is 

generally 3 to 5 mm; phenocrysts are as large as 5 em 

but are mostly smaller. Foliation is better defined and 

coarser than it is to the west. Layers are rich in lepidoblastic and microschlieren biotite, which makes up as 

much as 10 to 20 percent of the finer grained granulated 

rock. Locally the shearing was so intense that sheets of 

mylonite and ultramylonite were formed. The ultramylonite occurs as zones 7 to 14 em wide and contains 

porphyroblasts as long as 1 em. The shears average 1 em 

in width. Chlorite and epidote are concentrated along 

fractures. 

Compositionally, the granitic rocks differ in plagioclase 

composition and percentage, in mafic content and 

minerals, and in potassium feldspar percentage. Plagioclase in the granodiorite ranges from An21 to An34, but 

is mostly around An25-27. The mafic minerals are biotite 

and minor actinolite, which has altered to biotite and 

chlorite and includes some apatite and zircon. Plagioclase 

in the tonalite has a higher An range, from Ans2 to Ans6· 

The mafic mineral of the tonalite is biotite. Quartz monzonite and quartz monzodiorite have plagioclase that is 

slightly more sodic, from An23 to An27. Mafic minerals 

here are biotite and hornblende. The hornblende is 

subhedral and averages 1.25 mm in size, is partly altered 

to biotite, and includes apatite and zircon. 

CHEMISTRY OF IGNEOUS ROCKS 

Oxide and normative mineral percentages for the igneous rocks are shown in table 5.5. The uneven distribution 

of data among the various areas representing the major 

batholithic rocks makes interpretation difficult. In spite 

of the probably insufficient data, however, trend analyses 

and interpretations have been attempted. 

The diagrams of oxides versus Si02 (fig. 5.6) clearly 

demonstrate the absence of a linear fractionation trend 

between the tonalites and the granites and granodiorites, 

the absence of chemical distinctions between the granites 

and granodiorites, and an irregular distribution of composition within the several rock types. 

Concentrations of N a20 do not noticeably vary between 

areas or between rock types. This uniformity may indicate 

a redistribution of N a during shearing or an original irregular distribution of N a among the different magma 

types. The K20/Na20 ratios fall between 0.45 and 2.11 

(average of 0.99) for the granites, between 0.39 and 1.54 

(average of 0.88) for the granodiorites, and between 0.29 

and 0.97 (average of 0.51) for the tonalites. When compared to the data of Cox and others (1979), these ratios 

are within the range for granites (0.7) and tonalites 

(0.3 to 0.5) but are high for the granodiorite range (0.5 

to 0.7). The high ratios for granodiorite indicate a higher 

K20-concentration from biotite or feldspar. As a result, 

there is little or no difference between the granites and 

granodiorites. 

From the composition of the rocks it should be possible 

to infer the chemistry and origin of the magma from which 

they were derived. Table 5.6 gives data for those oxides 

and minerals which are used to determine the source of 

the igneous rocks in Chappell and White's (1974) classifi-
82 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

TABLE 5.4.-Chemical analyses of metamorphic rocks in migmatite of the Bitterroot lobe of the Idaho batholith and corresponding 

metamorphic mineral assemblages 

[Analyses given in weight percent. Areas: C, central; S, southern; W, western. Each analysis represents one sample unless other number (n) given, in which case the average 

is given] 

Rock type--------- Calc-silicate Schist Quartzite Gneiss Amphibolite 

Sample area------- c s s w s w s w c 

(n=3) ( n=3) (n=5) (n=5) (n=5) 

Sio2-------------- 64.3 69.3 46.6 66.1 87.5 86.4 67.8 73.8 66.2 

Al2o3------------- 10.2 12.0 25.4 18.2 5.2 6.7 16.3 12.7 17.3 

Fe2o3------------- 1.4 1.3 11.8 3.3 1.4 .95 4.2 1.5 1.6 

FeO--------------- 1.6 1.5 3.8 1.2 1.1 1.5 1.8 

MgO--------------- 5.6 4.9 2.8 1.7 .88 1.1 .95 1.9 1.7 

cao--------------- 10.8 3.9 .26 .06 .07 .73 .07 2.3 3.9 

Na2o-------------- 3.1 2.4 .24 • 78 .74 .88 .83 2.2 4.3 

~!~~============== 2.4 4.2 4.3 4.8 2.0 1.6 6.0 2.8 2.3 

.4 .23 2.0 1.0 .46 .33 .59 .49 .63 

P2os-------------- .13 .04 .05 .05 .03 .04 .05 .ll .17 

MnO--------------- .09 ~ .14 .06 .03 ~ .02 .07 ~ 

Total----------- 100.02 99.72 93.95 99.85 99.51 99.87 96.81 99.37 99.95 

Metamorphic mineral assemblages 

+I (I) (I) +I "0 +.1 (I) I (I) 

s:: ·r-t +.1 (I) (I) +.1 (I) ...-l ·r-t +.1 +.1 0 ...-l 0 s:: ·r-t •r-t "0 ,.c s:: ~ :> :> +.1 •r-t 

s:: ·r-t +.1 0 0 (I) p. ~ (I) +.1 ·r-t (I) () () s:: (I) 

0 +.1 () ...-l s:: C/) C/) ~ I ,..c: 0 m ...-l ~ § § m Q) 

I "0 I ·r-t m bl) +.1 (I) ·r-t (I) C/) (I) bl) I I +I ·r-1 "0 P."' I +.1 +I (I) Q) Q) .-1 ·r-t (I) ·r-t (I) ·r-t Q) +.I +.I +.1 0 C/) I C/) +.1 s:: +.1 ·r-t •r-t +.1 ·r-t s:: P.+-1 p. ·r-t ~ ·~ +.1 +.1 Q) +.1 ·r-t 

0 (I) 0 :> 0 0 s:: 0 +.1 ·r-t s:: ·r-t 0 ·r-t 0 •r-t ·r-t ~ •r-t () 

"0 ~"0 () ...-l () .D.D m..c m I m I C/) ...-l C/) I I bOt I (I) bi)Q) § ·r-t § Q) Q) +I Q) Q) C/) I C/) C/) s:: s:: Q) s:: Q)"' m (I) m I I +I ·r-t •r-t +.1 •r-t +.1 s:: ...-4"0...-l (I) (I) Q) .-1 ...-4 0.-1 •r-t Q) 

() ·r-1 () +.1 +.I +.I () ()"0 () .-I.-I 0 C/) 0 ·r-t ·r-t ·r-t 0 0 •r-t 0 ,..c: P.,..C: +.1 +.I +.I ~ ~ p.~ ~.g +.1 0 +.1 0 0 0 () () Q) () Q) ~ ~ ·r-t ~ •r-t ·r-t ~ •r-t •r-t +I~ ~ 0 0"0 0 ,.c ,.c,.c a a Q) a +.I,..C: I I I I +I +I I I +.1 +I I I Q) (I) Q) (I) Q) (I) Q) Q) •r-t Q) Q) Q) C/) C/) C/) C/) C/) C/) C/) C/) s:: C/) C/) C/) 

m m m Q) m m m m m ~~ m m ...-4...-4...-4+.1 ...-4 +.1 ...-4.-1 .-1...-4 .-I.-I (.) (.) () ·r-t () Q) () () () () •r-t () () () 

0 0 0 (.) 0 s:: 0 0 0 0.-1 0 0 0 •r-t ·r-t ·r-t ...-4 ·r-1 ~ •r-t ·r-t •r-t ~ .-1 •r-t •r-t •r-t bObObOtll bl)(l! bl)bl) bl) bO •r-t bl) bl)bl) 

m m m () m bl) m m m m C/) m Cll m ...-4 ...-4 ...-4 ...-4 ...-4...-1 ...-4.-1 ...-4 .-I.-I p..,p..,p.., p.., p..,p.., p..,p.., p.., p..,p.., 

cation, as well as for those minerals which may contribute 

to REE fractionation trends. 

The southern granites, granodiorites, and tonalites differ from the central and western rocks mainly in their 

higher values of normative enstatite, forsterite, magnetite, and ilmenite, which are characteristic of 1-type rocks. 

The high Alz03/KzO + NazO + CaO ratio and the high 

normative corundum values for the granites and granodiorites support an interpretation as S-type rocks. 

However, the values of these parameters are also high 

for the tonalites. 

S-type granite contains metasedimentary restite 

material (White and Chappell, 1977). Where xenoliths are 

present in the granitic rocks of the Bitterroot lobe the in-
5. MIGMATITE ZONES IN THE BITTERROOT LOBE 83 

elusions are predominantly metasedimentary rocks, 

although some are orthogneiss. The xenolith material 

could be restite in granitic rocks that represent minimum 

melts, or the xenoliths could be stoped and brecciated 

blocks of the country rock. The presence of these metasedimentary inclusions supports an interpretation of the 

rocks as S-type granite. 

The southern granites, like the others, include a high 

percentage of metasedimentary and tonalitic xenoliths. 

Their 1-type characteristics, therefore, may be the result 

of assimilating more igneous material. 

Data on RE E in rocks from the western and southern 

areas show two curves on the chondrite-normalized 

diagrams (fig. 5.7): (1) a general linear trend of LREE 

(light rare-earth-element) enrichment and a negative Eu 

anomaly and, (2) a saddle-shaped pattern of HREE (heavy 

rare-earth-element) abundances accompanied by a small 

positive Eu anomaly and LREE enrichment. Minimum 

melting of a metasedimentary source, which would involve 

a high degree of fractionation of amphiboles, garnet, and 

feldspars (Albuquerque, 1977; Hanson, 1978), would produce a linear trend like that seen in the granites. 

Compositional data for the tonalites and granodiorites 

are varied and complex. Tonalites can be derived from 

graywacke (Albuquerque, 1977) or from an igneous source 

(Arth and Hanson, 1972; Wyllie and others, 1976). The 

high Al203 and N a20 values in the tonalites here support 

a graywacke source, but they could also be the result of 

contamination or assimilation of metasedimentary 

material. The tonalites are rich in apatite and contain 

hornblende; as a result they have high abundances of both 

LREE and HREE. 

The close association between the granites and granodiorites in the field and their chemical similarity suggest 

they are synplutonic. The characteristic REE patterns for 

the granodiorites indicate partial fractionation of plagioclase, hornblende, and pyroxene (Albuquerque, 1977; Hanson, 1978); possibly a higher content of garnet in the melt 

is indicated by the trough-shaped HREE pattern. 

Sillimanite Muscovite 

K F Calcite Diopside Actinolite F 

FIGURE 5.3.-AKF and ACF diagrams applicable to the prebatholithic 

rocks in the migmatite of the Bitterroot lobe o.f the Idaho batholith. 

The diagrams represent the sillimanite-muscovite zone of the amphibolite facies of metamorphism (from Turner, 1981). 

A 

100 

10 

w 

1-

-

a: 

c 

z 

0 

:3: La Ce SmEu Tb Yb 

u 

~ ' 8 0 

0 

a: 

100 

10 

LaCe SmEu Tb Yb 

REE 

FIGURE 5.4.-Chondrite-normalized rare-earth-element abundances in 

metamorphic rocks associated with the Bitterroot lobe of the Idaho 

batholith. Chondrite data from Masuda and others (1973). A, 

Country-rock gneiss (solid lines) and schist (dashed lines). B, 

Prebatholithic migmatite rocks: quartzite and calc-silicate quartzite 

(solid lines); gneiss (dashed lines). 
84 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

INFERRED INTRUSIVE 

AND METAMORPHIC HISTORY 

Several deformational and metamorphic episodes prior 

to batholith emplacement are indicated by the chemical 

trends, mineralogy, and field relationships of the rocks. 

The following sequence of intrusion is represented by the 

field relationships: (1) Tonalite gneiss including plagioclase 

quartzite; (2) granite gneiss including tonalite gneiss; (3) 

pegmatite including granite gneiss; and ( 4) the batholith 

rocks. A prebatholithic event is indicated by the tonalite 

and granite gneiss and by evidence for earlier kyanitegrade metamorphism (500 oc and 10-30 km depth) in 

other rocks. The negative anomalies for Eu in the gneisses 

indicate derivation from melts depleted in Eu, either from 

the basement (Albuquerque, 1977; Haskin, 1977) or by 

anatexis from the enclosing metasedimentary rocks. 

Original metamorphism of the tonalite and granite 

gneisses in the migmatite was prior to batholith emplacement, as indicated by crosscutting relationships, and these 

rocks are distinctly different from the batholithic rocks 

in their REE trends. 

The emplacement of the batholithic tonalites along the 

northern, western, and southern borders, as early as approximately 90 Ma (table 5.1), caused lit-par-lit intrusion, 

assimilation, rafting, and contact metamorphism of the 

country rock to hornblende hornfels facies. 

Intrusion of the magma amd metasomatism were the 

major contributors to migmatization. Partial melting and 

1\ I \ I \ 

/ \ / \ / \ 

I A I I .. ,. 

I I . , .... "' 

,;;_.__ . :: ~ I I • ., • 

K~ 

FIGURE 5.5.-Classification of the batholithic rocks of the Bitterroot 

lobe of the Idaho batholith using Streckeisen and others' (1973) 

plutonic rock classification. Data from modal analyses. Fields are: 

A, granite; B, granodiorite; C, tonalite; D, quartz monzonite; E, 

quartz monzodiorite; F, quartz diorite. 

assimilation of the older metamorphosed units did, 

however, play some role, producing variable oxide data, 

high alkali and AI values, and the high 87Srf86Sr values 

of 0. 7196 (table 5.1). 

Later intrusions of the batholithic granites and granodiorites included xenoliths of tonalite and migmatized rock 

as well as the country rock. The rafts and xenoliths 

became aligned during the upward and eastward flow of 

the granitic crystalline mush. 

Blastomylonitic shearing and annealing strain resulted 

in sillimanite-muscovite zone recrystallization and a strong 

axial symmetry within the plutonic, migmatitic, and adjacent country rocks. After solidification and the major 

metamorphic event, the rocks underwent sericitization 

and low-temperature alteration and shearing. 

STRUCTURE 

The western, central, and southern areas can be divided 

into a western and an eastern structural domain. Lineation in the western domain, which is defined by fold axes 

and oriented minerals, has a southerly trend and plunges 

steeply. Foliation mostly strikes northwest to west and 

dips northward from 10° to 90°, but some strike northeast to east and dip steeply (fig. 5.8). 

The eastern structural domain includes most of the 

Bitterroot lobe and extends southward and westward to 

the inner western margin of the batholith (fig. 5.8). It is 

characterized by axial symmetry about a shallowly plunging northwest axis. The rocks generally have a linear 

fabric, which parallels fold axes, or locally a linear-planar 

fabric. Lineation trends predominantly west-northwest 

and has a relatively shallow plunge. Foliation strikes 

roughly northwest and varies in dip from vertical to 

horizontal. Eastward toward the Bitterroot front mylonite 

zone (fig. 5.8), which extends along the east border of the 

lobe, the fabric of the rock develops into a linear-planar 

fabric with a strong east to northeast trend. The mylonite 

zone is bordered on the west by the Bitterroot Mountain 

front and on the east by the Bitterroot River valley. Lineation in the mylonite zone plunges gently to the southeast 

and a pervasive, gently dipping planar fabric is present. 

The southern border of the lobe lies within the eastern 

domain and has a strong linear-planar fabric with a suggestion of axial symmetry about a northwest axis (fig. 5.8). 

Four separate fold events (F1 to F 4) are in the western 

and eastern domains. 

The earliest folding (F 1) was prior to intrusion and produced rootless isoclinal and recumbent folds in country 

rock xenoliths and migmatite. These folds occur locally 

along the margins of plutons along the western and 

southern borders and in the northeast corner of the lobe. 

Dikes intruded along existing axial-plane foliation (81). 
5. MIGMATITE ZONES IN THE BITTERROOT LOBE 85 

TABLE 5.5.-Average chemical and normative-mineral compositions for granitic rocks of the Bitterroot lobe of the Idaho batholith 

[Analyses and norms in weight percent. Areas: C, central; S, southern; W, western; N, northern] 

Rock type---------------- Granite Granodiorite Quartz diorite Tonalite 

Sampling area------------ c s w c 

Number of samples-------- 12 4 5 11 

Weight 

Sio2--------------------- 72.1 69.9 72.6 70.2 

Al2o3-------------------- 15.5 16.2 15.8 15.9 

Fe2o3-------------------- 1.32 1.27 .46 1.43 

FeO---------------------- .16 .62 .53 .15 

Mgo---------------------- .ss .96 .46 .62 

GaO---------------------- 1.98 2.27 1.46 2.19 

Na2o--------------------- 4.16 4.34 3. 77 4.43 

~t~~===================== 3.38 3.82 4.6 2.95 

.18 .28 .15 .19 

P205--------------------- .06 .11 .06 .06 

MnO---------------------- .02 .02 .02 .02 

Total-------------- 98.41 99.49 99.91 98.14 

CIPW 

Q------------------------ 28.2 23.7 28.1 26.1 

c------------------------ 1.24 1.11 2.6 .68 

or----------------------- 23.1 22.6 28.8 11.3 

ab----------------------- 29.7 36.8 31.4 24.8 

an----------------------- 7. 76 10.4 6.6 8.4 

en----------------------- 1.28 2.4 1.2 1.6 

fs----------------------- .43 .45 

mt----------------------- .79 .77 

il----------------------- .03 .36 .32 .os 

ap----------------------- .14 .26 .13 .18 

The second folding (F 2) formed steep concordant structures within plutonic rocks and prebatholithic screens during continued deformation and batholith emplacement. 

Along the western, southern, and northeastern borders 

of the lobe, the F 1 folds are refolded by conical and cylindrical folds. These folds developed axial-plane schistosity (82) having axes and lineations plunging steeply to the 

south-southwest, structures characteristic for the western 

structural domain. Primary folia are steep and concordant 

with the country-rock contact. Structures of F2 along the 

southern border include reclined isoclinal folds and tight 

folds that plunge more gently to the south-southwest. 

Granitic plutonism that progressed eastward from the 

steep western border aligned the country rock parallel to 

its flow and formed a primary flow foliation. Pegmatites 

and aplites intruding along F2 axial planes were synkinematically folded. 

A third folding (F 3) was caused by continued flow and 

deformation. The F2 structures were folded and the 

primary east-west foliation was overprinted by a secondary annealing strain fabric, which defines the eastern 

s w N w N s w 

6 5 2 1 3 7 11 

percent 

70.9 72.8 59.7 65.4 68.8 65.4 65.5 

14.9 14.8 18.2 16.1 16.7 16.4 16.5 

1.44 .98 1.42 2.27 .36 2.45 2.33 

1.4 1.12 3.85 3.30 2.13 1.09 2.44 

.95 .73 2.98 1.83 1.06 2.55 1. 74 

2.05 1.91 6.21 4.1 4.1 s.o 4.15 

3.60 3.57 3.90 3.75 4.06 3.83 3. 77 

3.67 3.30 1.37 2.29 1.37 1.97 2.18 

.so .11 .79 .59 .44 .so .71 

.14 .32 .28 .21 .11 .14 .20 

.03 .02 .10 .06 .02 ____:Q2_ ~ 

99.58 99.66 98.71 99.90 99.15 99.38 98.58 

norms 

30.2 29.5 12.8 28.6 22 24.6 

1.33 

21.7 20.3 8.12 8.1 11.7 11.9 

30.5 26.5 33.0 34.3 32.6 33.8 

9.29 6.34 28.0 19.6 20.2 17.4 

2.38 1. 71 7.43 2.64 6.39 3.58 

.84 .72 4. 76 2.93 .72 1.6 

1. 78 1.39 2.06 .52 1.39 2.4 

.92 .61 1.5 .84 .33 .46 

.33 .08 .67 .27 .33 .46 

structural domain. The secondary strain sheared the 

rocks, thus flattening and annealing feldspar and quartz, 

and developing lepidoblastic mica and quartz-filled tension fractures. The secondary strain field does not extend 

north of the Lochsa River nor into the northeast corner 

of the lobe. The northern border and the southeast extension of the secondary strain outlines the wedge-shaped 

eastern structural domain (R.R. Reid, oral commun., 

1980). North of the western domain, eastward to the 

Bitterroot River valley and southeast to the southern 

border, F3 is the predominant event. The F2 folds are 

refolded by open to gentle folds having axes plunging 

shallowly to the northwest. 

Protoclasis accompanied F3 folding and formed a 

penetrative foliation that obliterated the primary foliation. Migmatite zones and prebatholithic inclusions also 

developed a secondary strain fabric, although they retained their primary east-west alignment. Mineral lineations produced by the secondary strain have a shallow 

plunge to the northwest; the foliation is axially symmetric 

about this trend. Deformation continued after intrusion 
86 

..... 

z 

w 

0 

0:: 

w 

Q. 

..... 

:z: 

(!) 

w 

31: 

~ 

ui 

w 

0 

x 

0 

Ill 

0 

C\1 ~ 

C\1 

0 

i= 

THE IDAHO BATHOLITH AND ITS BORDER ZONE 

0.5 I 

0.41-- • -

o.3r . : -

0.2 ~ -

..... + 0 0 0.1 

.~S~t~o -

• o.o 

• 

•• 

• ' 0 + II f. • <t, + ~"t oit5 + o o • ~0~~-ot 0 

+ 0 + 

+01+ 0 0 

"'.p~ 

<o '691= ()1-o + + 

• o 0 oo 0 

I"J.• 

• • 0 

0 

+ • 

• + 

+ 

• 

• + \.t· 0 

• 't.;>~+ 0 • 

'§' +~~*it. + CfJ "¥0 

60 65 70 75 

Si02, IN WEIGHT PERCENT 

~~------~2-------3L-----~4~----~ 

Na20, IN WEIGHT PERCENT 

and solidification, as indicated by discordant relationships 

with plutonic contacts, and resulted in coaxial blastomylonitization and later mylonitization. 

The latest folding (F 4) is represented by the gentle 

warping of the northwest axis about a northeast axis, 

resulting in linear trends plunging both northwest and 

southeast . 

PREBATHOLITHIC STRUCTURES IN THE WESTERN DOMAIN 

RELICT STRUCTURES 

The prebatholithic rocks in the western domain are included in the granitic rock as xenoliths and large remnants 

and have formed large migmatitic screens having westerly 

trends. The structures in the metasedimentary rocks 

reflect multiple folding but also include relict sedimentary 

structures. Among relict structures in the metasedimentary units are orthogonally stretched pebbles, relict 

bedding defined by differences in grain size and mineral 

composition, and heavy-mineral concentrations that 

define crossbed sets. Locally, bedding trends are discordant to the foliation, either because of rotation of blocks 

or from overprinting of the more stable bedding planes 

by foliation and schistosity planes. 

FOLIATION 

The prebatholithic rocks have developed a fabric concordant with the local trend, which strikes variously from 

northeast through west to northwest and dips steeply 

(fig. 5.8). 

The northeast corner of the Bitterroot lobe structurally resembles the western border zone and therefore 

reflects the western domain (fig. 5.8). Country-rock inclusions make up more than 30 percent of the total rock. 

The layering and foliation strike east to northeast and 

have moderate dips . 

LINEATIONS 

Rodded quartz grains, stretched biotite, and sillimanite 

are all aligned with their long dimensions parallel to the 

F 2 fold axes and define a lineation. This lineation generally plunges steeply to the south (fig. 5.8). 

The fold axes in the northeastern corner of the lobe 

plunge steeply to the southwest; they show evidence of 

FIGURE 5.6.-Variation diagrams of major oxides in the batholithic 

granitic rocks of the Bitterroot lobe of the Idaho batholith based 

on whole-rock chemical data. Filled squares, tonalite; open circles, 

granodiorite; crosses, granite. 
5. MIGMATITE ZONES IN THE BITTERROCYr LOBE 87 

TABLE 5.6.-Selected compositional data for granitic rocks of the Bitterroot lobe of the Idaho batholith compared with characteristics of 

I- and S-type granites 

[Criteria for granite types from Chappell and White (1974). Areas: C, central; S, southern; W, western. HREE, heavy rare earth elements. Sr-isotope ratios from table 5.2. M, 

muscovite; B, biotite; H, hornblende; A, amphibole; Px, pyroxene; Sp, sphene; Ep, epidote; Gr, garnet; S, sillimanite; Ap, apatite; An, anorthite; Cd, cordierite; Ep, epidote; metased., 

metasedimentary] 

Rock type-------- Granite Granodiorite Tonalite I-type S-type 

granite granite 

Sample area------ c s w c s w N s w 

Weight percent 

Na2o------------- 4.2 4.3 3.7 4.4 3.6 3.6 4.1 3.8 3.7 3.2 3.2 

K2o-------------- 3.4 3.8 4.6 2.9 3.7 3.3 1.4 1.9 2.2 5 or 2 

Al 2o3/K20+Na2+Ca0 1.6 1.6 1.6 1.7 1.6 1.7 1.8 1.5 1.6 1.1 1.1 

Sio2------------- 72 70 73 70 71 73 69 65 66 high 

Norms 

c---------------- 1.3 1.1 2.6 1.2 1.7 3.0 1.4 0.7 1.3 1.0 1.0 

il--------------- .04 .4 .3 .05 .9 .6 .8 .7 1.0 il 

mt--------------- .o .8 .8 .o 1.8 1.4 .5 1.4 2.4 mt 

en--------------- 1.3 2.4 1.2 1.6 2.4 1.7 2.6 6.4 3.6 di 

fs--------------- .o .4 .4 .o .8 .7 2.9 .7 1.6 

Minerals present for HREE concentration 

M+B H+B+M B+M B+M+Px B+M+An 

Gr s Gr+S 

Accessory minerals Ap Ap Ap Ap Ap 

Sp Ep 

87sr;86sr ratios--

Xenolith types--- metased. metased. 

H-bearing 

a later folding about a northeast-trending axis. Biotite 

lineation is not parallel to the prevailing fold-axis direction but plunges steeply to the northwest. This northwest 

trend may be a reflection of the eastern structural domain. 

FOLDS 

Three distinct fold sets are well displayed in the migmatite in the western domain and reflect the first three of 

the fold events discussed earlier. The Frfolds consist of 

overturned to recumbent folds, from 0.5 to 5 m wide, and 

rootless folds showing axial-plane schistosity (81). 

B+M H+B H+B B+M H B+M 

Gr+Cd 

Ap Ap Ap Ap Ap Ap 

Ep Sp Sp+Ep Sp+Ep Sp Ep 

0.7103, 0. 7196 0.704 0.708 

• 706 

metased. H-bearing metased. 

H-bearing 

The F 1 folds were refolded about a south-southwest 

axis by conical and cylindrical upright to overturned folds 

(F2), which have developed axial-plane schistosity. The 

axial planes, defined by lepidoblastic biotite, muscovite, 

and sillimanite, have random strike and dips ranging from 

50° to vertical; they are commonly invaded by dikes. 

The third fold event (F3), weakly developed on the 

western margin, is a reflection of the eastern domain overprinting the western. It is indicated by open to gentle 

parallel and asymmetrical folds which refold F2 about a 

northwest-trending axis. The F3 folds are as wide as 3m 

from crest to trough. 
88 

LIJ 

t: 

a: 

Q 

z 

0 

:::1: 

u 

::.:: ' 

(.) 

0 

a: 

10 

LaCe 

THE IDAHO BATHOLITH AND ITS BORDER ZONE 

8 

c 

SmEu Tb Yb 

REE 

BATHOLITHIC STRUCTURES IN THE WESTERN DOMAIN 

Along the western contact zone, the rocks of the batholith show the same northwesterly structural trends as the 

prebatholithic rocks. 

FOLIATION 

The igneous rocks range from flow foliated to gneissic. 

Biotite aligned in the plane of flow defines the foliation, 

which strikes from northwest through north to northeast 

and dips moderately (fig. 5.8). 

LINEATION 

The lineation trends southwesterly, has a shallow to 

steep plunge (fig. 5.8), and cuts across the foliation. 

Stretched biotite, elongate feldspar phenocrysts, and 

sillimanite needles define the lineation; they lie in the 

plane of maximum extension and are aligned parallel to 

F 2 fold axes. 

PROTOCLASIS 

Protoclastic recrystallization and shearing are common 

throughout the western area. Sillimanite, which is an indicator of the intensity of these effects, lies in the plane 

of shear and defines a lineation for the protoclastic event. 

Quartz grains are extended parallel to the lineated 

sillimanite. Lepidoblastic biotite and muscovite are concentrated along the shear planes. 

Subsequent shearing possibly linked to overprinting by 

the eastern domain has truncated the primary schistosity in the batholithic rocks and has reoriented fold axes 

and rodded grains to the northwest (az 300°-330°) in the 

migmatitic terrain. 

The principal ductile shear planes were later crosscut 

by northeast-trending shears, which granulated the 

batholithic rocks. These shears are only 1 em wide and 

cut the protoclastic fabric and the pegmatites. Subgrains 

of feldspar and strained quartz and semialigned secondary 

sericite and muscovite have developed along these narrow shears. 

FOLDS 

Large, open, cylindrical Fz folds affect both the prebatholithic and batholithic rocks. The folds plunge to the 

southwest (az 250°) and southeast (az 150°). 

FIGURE 5.7.-Chondrite-normalized rare-earth-element abundances in 

batholithic rocks of the Bitterroot lobe of the Idaho batholith. 

Chondrite data from Masuda and others (1973). A, Tonalitic rocks. 

B, Granodioritic rocks. C, Granitic rocks. 
5. MIGMATITE ZONES IN THE BITTERROOT LOBE 89 

DIKES 

Pegmatite and aplite have intruded primary joints, foliation surfaces, and shear surfaces. Poles to the pegmatite 

dikes have a crude axial symmetry like that of the poles 

to the foliations (82). 

PREBATHOLITHIC STRUCTURES IN THE EASTERN DOMAIN 

Prebatholithic rocks in the eastern domain have notretained relict sedimentary structures nor depositional 

fabric. The rocks are present as xenoliths and remnants 

and in rare migmatite zones. The inclusions trend predominantly eastward through the central part of the 

domain, but are at various orientations along the southern 

border. The fabric of the included rock is axially symmetric about a shallow northwest-trending axis (az 297°, 

plunge 36 ° ). 

Along the southern border of the Bitterroot lobe is a 

migmatitic contact zone formed by screens of metasedimentary rock and tonalite gneiss separated by concordant 

sheets of granodiorite and granite. The migmatite screens 

have been truncated on the south-southwest by Eocene 

granitic intrusions. 

46° 

oo• 

WESTERN 

STRUCTURAL DOMAIN 

WESTERN AREA 

10 20 30 

Kilometers 

EASTERN STRUCTURAL 

DOMAIN 

FIGURE 5.8.-Stereonet plots of structural data from areas assigned to 

the western and eastern structural domains of the Bitterroot lobe 

of the Idaho batholith. Plots are on the lower hemispheres of the 

Wulff stereonet. Dots are lineations and fold axes; great circles are 

derived from poles to foliation (girdle from scattered poles in northeastern area). Rock units: Tg; Tertiary granite; TKg, Late Cretaceous and early Tertiary granite and granodiorite; Kt, Cretaceous 

tonalite; Kmi, Cretaceous migmatite. 

Xenoliths and remnants of prebatholithic rock in the 

central area trend east-west. They range from 1 m to 3 km 

in length, and rotated blocks are locally present. 

FOLIATION 

Rotation of xenoliths during emplacement has resulted 

in foliation and schistosity that are not everywhere concordant with the prevailing trend of the inclusions. The 

initial flow and rotation were followed by protoclasis, 

which developed a penetrative foliation that strikes from 

northwest through north to northeast and dips shallowly 

to steeply (fig. 5.8). 

LINEATION 

Biotite lineation along the foliation planes plunges 

gently to the northwest (az 270° to 305°; fig. 5.8). 

Elongate feldspar grains commonly are spindled and have 

been rotated clockwise or counterclockwise to become 

perpendicular to the lineation. 

FOLDING 

In the migmatitic rocks of the eastern domain the predominant folds represent the F2, Fs, and F 4 events. 

The F2 folds are identified by fold axes plunging 

variously to the southwest and northeast. Near the southern border a quartzite schistosity has been recumbently 

folded about a shallow southwest-trending axis. An axialplane schistosity formed by lepidoblastic biotite and 

muscovite (82) has developed. In the central area, F2 folds 

are symmetric to isoclinal and have axes plunging 

moderately to the southwest. 

The F 3 folds are generally open cylindrical and conical 

folds with axes plunging shallowly to the northwest and 

southeast. Small symmetric and ptygmatic parasitic and 

flow folds have formed in the inclusions and in the granitic 

rocks. These folds were formed during the same refolding 

event and have axes which also plunge shallowly to the 

northwest and southeast. The varying directions of rotation and symmetry shown by the folds indicate formation 

in an environment of flattening or pure shear. 

Uplift and doming during F 4 have refolded F2 and F3 

about a northeast-trending axis. 

BATHOLITHIC STRUCTURES IN THE EASTERN DOMAIN 

A single structural pattern of axial symmetry about a 

lineation plunging shallowly to the northwest characterizes the batholithic rocks in the eastern domain. 
90 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

FOLIATION MYLONITIZATION 

Shearing that represents the Fa event has produced a 

nonplanar foliation, which overprints a weak primary 

foliation at a small angle. The primary foliation is defined 

by aligned biotite and undeformed feldspar crystals 

having normal zonation. An alignment of xenoliths and 

migmatitic zones in the east-west direction suggests that 

that was the primary flow direction. 

The flow structure of the secondary strain event is 

characterized by protoclastic texture resulting from solidstate flow and recrystallization. The flow structure 

becomes better developed and more pervasive eastward 

to the mylonite zone. Poles to the shear planes are concentrated in a broad girdle that strikes northeast and dips 

moderately to the southeast (fig. 5.8). 

LINEATION 

Preferred orientations of stretched lepidoblastic biotite 

and rodded quartz grains, both parallel to the direction 

of shear, define a linear fabric trending northwest (fig. 

5.8). Phenocrysts of feldspar form elongate clumps along 

the strike of the foliations; in the plane of shear they have 

been rotated clockwise and counterclockwise to positions 

perpendicular to the lineation. Crenulation and fold axes 

parallel the predominant linear trend. 

FOLDS 

Folds in the gneiss and migmatite correspond to the Fa 

and F 4 events. The Fa folds are large, similar, reclined 

isoclinal to tight folds and have axes plunging shallowly 

to the southwest. Parasitic and shear folds that formed 

during the Fa event developed an axial-plane schistosity, 

which cuts the F2 schistosity. Later doming (F 4) refolded 

the lineations and fold axes about a northeast-trending 

axis. The refolding has produced a second concentration 

of lineations trending southeast at az 110°. 

BLASTOMYLONITIZATION 

Axial deformation continued beyond protoclasis to produce blastomylonitic shears. The shears are nonpenetrative and have biotite covering the shear planes. These 

biotite-coated shears, which cut the secondary penetrative 

schistosity, strike predominantly east-west and have 

moderate dips. The preferred orientation of biotite on the 

shear surfaces defines a lineation plunging to the northwest or southeast. The shear planes decrease in dip and 

become more concentrated in the mylonite front eastward 

from the Bitterroot crest. 

The mylonite zone trends north-south along the Bitterroot front (fig. 5.8). The rock grades from a mylonite 

gneiss to a protomylonite-mylonite interlayered with 

sheets of ultramylonite that contain pseudotachylyte. The 

ultramylonite sheets are 12 to 14 em wide and have porphyroclasts as much as 1 em long. As the intensity of 

shearing increased, prominent planes of relative movement formed. Smeared-out grains of chlorite and mica are 

parallel to striations on the shear surface. 

DIKE INTRUSION 

Pegmatite and aplite dikes invaded fractures during the 

late stages of cooling before the imposition of secondary 

strain. The dikes have been cut and offset by the 

penetrative secondary shear event. Secondary shear folds 

(Fa) also formed during rotation of the dikes into 

parallelism with the foliation. Dikes synkinematic with the 

penetrative deformation intruded along the schistosity as 

veins and sheets and were subsequently deformed. 

SUMMARY OF DEFORMATIONAL EVENTS 

The events in the structural history of the Bitterroot 

lobe can be grouped into five stages of deformation. 

(1) Kyanite-zone metamorphism and folding of the 

country rock before intrusion of the batholith. 

(2) Injection of the batholithic tonalite to granodiorite, 

followed by pulses of the main-phase granite and granodiorite; steep concordant flow foliations and lineations 

along the western border and shallow flow foliations and 

lineations in the central, eastern, and southern areas of 

the lobe; migmatite screens and xenoliths oriented in eastwest alignment and conical and cylindrical folding that 

refolded F 1 around a southwest-trending axis; hornblende 

hornfels contact metamorphism. 

(3) Development of a secondary annealing strain fabric 

at the sillimanite-muscovite facies of metamorphism; overprinting by the secondary fabric of the primary foliation 

of stage 2; formation of shallow axial-symmetric fabrics 

and refolding of F2 about a shallow northwest axis by 

open gentle cylindrical folds. 

(4) Blastomylonitic shearing and intense mylonitization 

after crystallization of the batholith; formation of the 

eastern-front mylonite zone defined by bands of mylonite, 

ultramylonite, and pseudotachylyte that have varied 

strikes and lineations plunging shallowly to the southeast. 

(5) Refolding of Fa and of the lineations of stages 3 and 

4 by faulting and arching of the eastern batholith. 
5. MIGMATITE ZONES IN THE BITTERROOT LOBE 91 

CONCLUSIONS 

Significant deformation and metamorphism occurred in 

the area before the batholith was intruded. The structural 

setting at the time included thrust blocks that were east 

directed and strike-slip faults and lineaments that trended 

northwest. Plutonism occurring at about 90 Ma invaded 

the country rocks and thrust blocks in a lit-par-lit manner. 

The plutonism was accompanied by deformation and 

migmatization, and it resulted in high-grade metamorphism of the kyanite zone. Assimilation of the country 

rock or an anatectic origin for the magma may account 

for the high 87Srf86Sr ratios (0. 7103 to 0. 7196) in the 

tonalite gneisses and their similarity to those of the country rock. The similar REE patterns in country rock and 

tonalite gneiss may support an anatectic origin for the 

migmatitic rocks. 

Major structural amd metamorphic events accompanied 

batholithic intrusion of the Bitterroot lobe. Emplacement 

of the lobe was by lower crustal melting and thrusting. 

The shape of the lobe, elongate in a west-northwest direction, suggests emplacement along a major lineament of 

that trend. Yates (1968) and Taubeneck and Armstrong 

(1974) have inferred the existence of a northwesttrending transform, the trans-Idaho discontinuity, in the 

vicinity of the lobe. Such a preexisting transform would 

have provided a pathway in the crustal rocks for the 

ascending melt. 

The major plutonism that produced the lobe occurred 

during the mid-Sevier orogenic period. The magma intruded the country rock lit-par-lit and in the cores of folds, 

and locally intrusion was forceful enough to break up the 

wall rock. Blocks and screens of country rock were incorporated and aligned parallel to the flow direction of the 

intruding material. The west to northwest alignment of 

the migmatized country rock could represent the primary 

flow direction of the intruding magma. Some of the 

screens may represent dividers between separate plutonic 

intrusions. Plutonism continued through the Late Cretaceous into early Tertiary. Chemical analyses suggest a 

lower crustal "S-type" paragenesis which would support 

melting in the lower crust. 

Migmatite zones formed around the periphery of the 

lobe and sparsely within its interior. The last plutonic 

phases of the batholith incorporated early batholithic 

phases and migmatized material. Complex zones were 

formed by intrusion and assimilation of prebatholithic 

metasedimentary rocks and orthogneiss by the final 

phases of batholithic plutonism. The screens and blocks 

are concentrated mostly at higher elevations, capping 

ridges and peaks. In the western and southern areas, 

zones of migmatization are at lower elevations and have 

a more assimilated appearance. The high concentration 

of migmatized rock at lower elevations in the western and 

southern areas may represent exposures at greater depth 

within the batholith, near the basal contact. The sparsely 

concentrated migmatite zones in the batholith interior 

may have resulted from breaking up, stoping, or rafting 

of the roof rocks into ascending magma. Wall-rock and 

plutonic fabrics indicate an initially steep flow and a 

folding about a southwest-trending axis during the intrusion of the batholith. 

Deformation continued through the late stages of batholith emplacement and afterward. During the final stages 

of plutonism a penetrative axial-symmetric annealing 

strain fabric overprinted the plutonic primary foliations 

and migmatite schistosities. Continued plutonism and 

protoclasis of the cooling magma against an unyielding 

roof may have caused development of strain. Other possible causes are crustal extension, perhaps the driving force 

in the Cenozoic development of the west coast metamorphic core complexes (Coney, 1980), and renewed low-angle 

faulting and movement along the proposed trans-Idaho 

transform. 

Axial-symmetric deformation developed blastomylonitic 

shear zones after crystallization of the batholith. The 

strain regime remained in effect after protoclasis and produced nonpenetrative biotite-covered shears subparallel 

to the secondary foliation. The blastomylonitic shearing 

granulated and recrystallized the rock. On the eastern 

front of the Bitterroot Range the protoclastic fabric and 

blastomylonitic shears are cut by a third zone of cataclasis, 

the eastern front mylonite zone, which is characterized 

by bands of mylonite, ultramylonite, and pseudotachylyte. 

Shear lineations in the zone plunge shallowly to the southeast. Recent work by Garmezy (1983) and by Chase and 

others (1983) has suggested that the cause of mylonitization was an Eocene (45.5 Ma) episode of extension and 

domal uplift along a fault parallel to the Bitterroot River 

valley. The principal direction of extension was in a 

southeast (az 110°) direction (Garmezy, 1983). The episode 

of doming refolded the F3 structures and the lineations 

of deformation stages 3 and 4. 

REFERENCES CITED 

Albuquerque, C.A.R. de, 1977, Geochemistry of the tonalitic and granitic 

rocks of the Nova Scotia southern plutons: Geochimica et Cosmochimica Acta, v. 41, p. 1-13. 

Allen, C.C., 1980, Petrologic and chemical study of a portion of the northeastern border zone of the Idaho batholith, east-central Idaho: 

Williamstown, Mass., Williams College, independent study, 43 p. 

Anderson, A.L., 1952, Multiple emplacement of the Idaho batholith: Journal of Geology, v. 60, p. 255-265. 

Armstrong, R.L., 1974, Magmatism, orogenic timing, and orogenic 

diachronism in the cordillera from Mexico to Canada: Nature, v. 247, 

p. 348-351. 

__ 1975, Precambrian (1500 m.y. old) rocks of central Idaho-the 

Salmon River arch and its role in cordilleran sedimentation and tee-
92 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

tronics: American Journal of Science, v. 275A, p. 437-467. 

__ 1976, Geochronometry of the Idaho batholith and nearby rocks 

[abs.]: Geological Society of America Abstracts with Programs, v. 8, 

no. 3, p. 350-351. 

Armstrong, R.L., Taubeneck, W.H., and Hales, P.O., 1977, Rb-Sr and 

K-Ar geochronometry of Mesozoic granite rocks and their Sr isotopic 

composition, Oregon, Washington and Idaho: Geological Society of 

America Bulletin, v. 88, p. 397-411. 

Arth, J.G., and Hanson, G.N., 1972, Quartz diorites derived by partial 

melting of eclogite or amphibolite at mantle depths: Contributions 

to Mineralogy and Petrology, v. 37, p. 161-174. 

Bence, A.E., and Albee, A.L., 1968, Empirical correction factors for 

the electron microanalysis of silicates and oxides: Journal of Geology, 

V. 76, p. 382-403. 

Bradley, M.D., 1981, Geology of northeastern Selway-Bitterroot 

Wilderness Area, Idaho and Montana: Moscow, Idaho, University 

of Idaho, Masters thesis, 51 p. 

Chappell, B.W., and White, A.J.R., 1974, Two contrasting granite types: 

Pacific Geology, v. 8, p. 173-174. 

Chase, R.B., 1968, Petrology of the northeast contact zone of the Idaho 

batholith, northern Bitterroot Range, Montana: Missoula, Montana, 

University of Montana, Ph.D. dissertation, 158 p. 

Chase, R.B., Bickford, M.E., and Arruda, E.C., 1983, Tectonic implications of Tertiary intrusion and shearing within the Bitterroot dome, 

northeastern Idaho batholith: Journal of Geology, v. 91, p. 462-470. 

Chase, R.B., Bickford, M.E., and Tripp, S.E., 1978, Rb-Sr and Pb-Pb 

isotopic studies of the northeastern Idaho batholith and border zone: 

Geological Society of America Bulletin, v. 89, p. 1325-1334. 

Coney, P.J., 1972, Cordilleran tectonics and North America plate motion: 

American Journal of Science, v. 272, p. 603-628. 

__ 1980, Cordilleran metamorphic core complexes: an overview: 

Geological Society of America Memoir 153, p. 7-31. 

__ 1981, Accretionary tectonics in western North America, in Dickinson, W.R., and Payne, W.D., eds., Relations of tectonics to ore 

deposits in the southern cordillera: Arizona Geological Society 

Digest, v. 14, p. 23-39. 

Cox, K.G., Bell, J.D., and Pankhurst, R.J., 1979, The interpretation of 

igneous rocks: London, George Allen and Unwin, 445 p. 

Davis, G.H., 1975, Gravity-induced folding of a gneiss dome complex, 

Rincon Mountains, Arizona: Geological Society of America Bulletin, 

v. 86, p. 979-990. 

Drewes, Harald, 1978, The cordilleran orogenic belt between Nevada 

and Chihuahua: Geological Society of America Bulletin, v. 89, 

p. 641-657. 

Ferguson, J.A., 1975, Tectonic implications of some geochronometric 

data from the northeastern border zone of the Idaho batholith: 

Northwest Geology, v. 4, p. 53-58. 

Garmezy, Lawrence, 1983, Mylonitization coincident with uplift in an 

extensional setting, Bitterroot Range, Montana-Idaho [abs.]: 

Geological Society of America Abstracts with Programs, v. 15, no. 6, 

p. 578. 

Greenwood, W.R., and Morrison, D.A., 1973, Reconnaissance geology 

of the Selway-Bitterroot Wilderness Area: Idaho Bureau of Mines 

and Geology Pamphlet 154, 30 p. 

Hamilton, Warren, 1969, The volcanic central Andes-a modern model 

for the Cretaceous batholiths and tectonics of western North 

America, in Proceedings of the Andesite Conference: Oregon 

Department of Geology and Mineral Industries Bulletin, 65, 

p. 175-184. 

Hamilton, Warren, and Meyers, W.B., 1967, The nature of batholiths: 

U. S. Geological Survey Professional Paper 554-C, p. C1-C30. 

Hanson, G.N., 1978, The application of trace elements to the petrogenesis of igneous rocks of granitic composition: Earth and Planetary 

Science Letters, v. 38, p. 26-43. 

Harrison, J.E., 1972, Precambrian Belt basin of Northwestern United 

States: its geometry, sedimentation, and copper occurrences: 

Geological Society of America Bulletin, v. 83, p. 1215-1240. 

Harrison, J.E., and Campbell, A.B., 1963, Correlations and problems 

in Belt Series stratigraphy, northern Idaho and western Montana: 

Geological Society of America Bulletin, v. 74, p. 1413-P1427. 

Haskin, L.A., 1977, On rare-earth element behavior in igneous rocks, 

in Ahrens, L.H., ed., Origin and distribution of the elements: New 

York, Pergamon Press, p. 175-189. 

Hietanen, Anna, 1960, Superposed deformations northwest of the Idaho 

batholith: International Geological Congress, 21st, Copenhagen, Proceedings, v. 26, p. 87-102. 

__ 1961, Relation between deformation, metamorphism, metasomatism, and intrusion along the northwest border zone of the 

Idaho batholith, Idaho: U.S. Geological Survey Professional Paper 

424-D, p. D161-D164. 

__ 1962, Metasomatic metamorphism in western Clearwater County, 

Idaho: U.S. Geological Survey Professional Paper 344-A, p. A1-A116. 

__ 1963a, Idaho batholith near Pierce and Bungalow, Clearwater 

County, Idaho: U.S. Geological Survey Professional Paper 344-D, 

p. D1-D42. 

__ 1963b, Metamorphism of the Belt Series in the Elk River-Clarkia 

area, Idaho: U.S. Geological Survey Professional Paper 344-C, 

p. C1-C49. 

__ 1968, Belt Series in the region around Snow Peak and Mallard 

Peak, Idaho: U.S. Geological Survey Professional Paper 344-E, 

p. E1-E34. 

Hobbs, S.W., Griggs, A.B., Wallace, R.E., and Campbell, A.B., 1965, 

Geology of the Coeur d'Alene district, Shoshone County, Idaho: U.S. 

Geological Survey Professional Paper 478, 139 p. 

Hyndman, D.W., 1980, Bitterroot dome-Sapphire tectonic block, in Cordilleran metamorphic core complexes: Geological Society of America 

Memoir 153, p. 427-444. 

Hyndman, D.W., and Williams, L.D., 1977, The Bitterroot lobe of the 

Idaho batholith: Northwest Geology, V. 6, p. 1-16. 

Johnson, B.R., 1975, Migmatites along the northern border of the Idaho 

batholith: Missoula, Montana, University of Montana, Ph.D. dissertation, 120 p. 

Kell, R.E., and Chase, R.B., 1976, Multiphase deformation and polymetamorphism along the north-central border zone of the Idaho 

batholith, Kelly Fork district, Clearwater County, Idaho [abs.]: 

Geological Society of America Abstracts with Programs, v. 8, no. 

4, p. 485-486. 

Leischner, L.M., 1959, Border-zone petrology of the Idaho batholith in 

vicinity of Lolo Hot Springs, Montana [abs.]: Geological Society of 

America Bulletin, v. 70, p. 1782. 

Lindgren, Waldemar, 1904, A geological reconnaissance across the 

Bitterroot Range and Clearwater Mountains in Montana and Idaho: 

U.S. Geological Survey Professional Paper 27, 123 p. 

Masuda, Akimasa, Nakamura, Noboru, and Tanaka, Tsuyoshi, 1973, Fine 

structures of mutually normalized rare-earth patterns of chondrites: 

Geochimica et Cosmochimica Acta, v. 37, p. 239-248. 

Morrison, D.A., 1968, Reconnaissance geology of the Lochsa area, Idaho 

County, Idaho: Moscow, Idaho, University of Idaho, Ph.D. dissertation, 126 p. 

Nold, J.L., 1968, Geology of the northeastern border zone of the Idaho 

batholith, Montana and Idaho: Missoula, Montana, University of 

Montana, Ph.D. dissertation, 159 p. 

Reid, R.R., 1959, Reconnaissance geology of the Elk City region, Idaho: 

Idaho Bureau of Mines and Geology Pamphlet 120, 7 4 p. 

Reid, R.R., Bittner, Enid, Greenwood, W.R., Ludington, Steve, Lund, 

Karen, Motzer, W.E., and Toth, Margo, 1979, Geologic section and 

road log across the Idaho batholith: Idaho Bureau of Mines and 

Geology, Information Circular 34, 20 p. 

__ 1980, U.S. Geological Survey Selway-Bitterroot Mineral Evaluation Study: U.S. Geological Survey Open-File Report. 
5. MIGMATITE ZONES IN THE BITTERROOT LOBE 93 

Reid, R.R., and Greenwood, W.R., 1968, Multiple deformation and 

associated progressive polymetamorphism in the Beltian rocks north 

of the Idaho batholith, Idaho, U.S.A.: International Geological Congress, 23rd, Prague, Proceedings, v. 4, p. 75-87. 

Reid, R.R., Morrison, D.A., and Greenwood, W.R., 1973, The Clearwater orogenic zone: a relict of Proterozoic orogeny in central and 

northern Idaho: Idaho Bureau of Mines Belt Symposium 1973, v. 1, 

p. 10-56. 

Reid, R.R., Greenwood, W.R., Nord, G.L., 1981, Metamorphic petrology 

and structure of the St. Joe area, Idaho: Geological Society of 

America Bulletin, pt. II, v. 92, p. 94-205. 

Roberts, R.J., 1972, Evolution of the cordilleran fold belt: Geological 

Society of America Bulletin, v. 83 p. 1989-2003. 

Ross, C.P., 1928, Mesozoic and Tertiary granitic rocks in Idaho: Journal of Geology, v. 36, p. 673-693. 

__ 1963a, Modal composition of the Idaho batholith: U.S. Geological 

Survey Professional Paper 475-C, p. C86-C90. 

__ 1963b, The Belt series in Montana: U.S. Geological Survey Professional Paper 346, 122 p. 

__ 1965, Idaho batholith [abs.]: Geological Society of America Special 

Paper 82, p. 343. 

Streckeisen, A.L., and others, 1973, Plutonic rocks: classification and 

nomenclature recommended by the lUGS Subcommission on the 

Systematics of Igneous Rocks: Geotimes, v. 18, no. 10, p. 26-30. 

Taubeneck, W.H., and Armstrong, R.L., 1974, The trans-Idaho discontinuity viewed as a transform fault [abs.]: Geological Society of 

America Abstracts with Programs, v. 6, no. 5, p. 478. 

Toth, M.I., 1982, Reconnaissance geologic map of the Selway-Bitterroot 

Wilderness, Idaho County, Idaho, and Missoula and Ravalli Counties, Montana: U.S. Geological Survey Miscellaneous Field Studies 

Map MF-1495-B, scale 1:125,000. 

Tripp, S.E., and Chase, R.B., 1976, Rb/Sr and U/Pb geochronology of 

the northeastern border of the Idaho batholith, Bitterroot Range, 

Montana [abs.]: Geological Society of America Abstracts with Programs, v. 8, no. 4, p. 514-515. 

Turner, F.J., 1981, Metamorphic petrology; mineralogical, field, and 

tectonic aspects (2nd ed.): New York, McGraw-Hill, 524 p. 

Wehrenberg, J.P., 1968, Structural development of the northern Bitterroot Range, western Montana [abs.]: Geological Society of America 

Special Paper 115, 456 p. 

White, A.J.R., and Chappell, B.W., 1977, Ultrametamorphism and 

granitoid genesis: Tectonophysics, v. 43, p. 7-22. 

Williams, L.D., 1977, Petrology and petrography of a section across the 

Bitterroot lobe of the Idaho batholith: Missoula, Montana, University of Montana, Ph.D. dissertation, 189 p. 

__ 1979, General geology of a section across the Bitterroot lobe of 

the Idaho batholith: Northwest Geology, v. 8, p. 29-39. 

Wiswall, C.G., 1979, The structure and petrography below the Bitterroot dome, Idaho batholith, near Paradise, Idaho: Missoula, Montana, 

University of Montana, Ph.D. dissertation, 129 p. 

Wyllie, P.J., Huang, W.L., Stern, C. R., and Maaloe, Sven, 1976, Granitic 

magmas: possible and impossible sources, water contents, and 

crystallization sequences: Canadian Journal of Earth Sciences, v. 13, 

p. 1007-1019. 

Yates, R.G., 1968, The trans-Idaho discontinuity: International Geological Congress, 23rd, Prague, Proceedings, v. 1, p. 117-123. 

6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS 

OF THE NORTHERN IDAHO BATHOLITH AND ADJACENT AREAS 

BASED ON 180/ 160, D/H, 87SR/ 86SR, K-AR, AND 40AR/ 39AR STUDIES 

By RoBERT E. CRiss and RoBERT J. FLECK 

ABSTRACT 

Geochemical, geological, and geochronological variations among the 

granitoid rocks of north-central Idaho and adjacent areas indicate that 

the region is composed of two essentially different terranes. Jurassic 

and Cretaceous plutons intrude Permian to Jurassic volcanic and 

sedimentary rocks of the Wallowa-Seven Devils terrane, a province 

devoid of Precambrian rocks that probably represents an accreted islandarc complex. These Jurassic and Cretaceous plutons are distinguished 

by low initial 87 Sr/86Sr ratios (rd of 0.7030 to 0.7042, Rb/Sr ratios less 

than 0.08, 6180 values of + 5.5 to + 10.5 per mil, and oD values of -41 

to - 82 per mil, all of which are compatible with derivation, at least in 

part, from mantle materials. These rocks are separated from the main 

mass of the Idaho batholith by a 5- to 25-km-wide suture zone containing deformed, Late Cretaceous transition zone plutons. Steep gradients 

in ri, 6180, dD, and major element concentrations of the plutons occur 

across this zone and reflect the contrasting characteristics of the bordering source regions. The main mass or Bitterroot lobe of the northern 

Idaho batholith intrudes the Belt-Y ellowjacket terrane, which contains 

abundant Precambrian rocks. These granitic plutons can be divided into 

a Cretaceous-early Tertiary main-phase suite and an Eocene epizonal 

suite. The main-phase plutons have high ri values of ~0. 708, Rb/Sr 

ratios of 0.04 to 0.24, and high primary 6180 values of +8.0 to + 12.4 

per mil. These magmas apparently were derived primarily from Precambrian lower crust, but they also display evidence of substantial assimilation of Precambrian wall rock. The Eocene epizonal plutons all have 

ri values ~0.704 (mean value of 0.7088), suggesting the presence of an 

ancient component, but several values between 0.704 and 0.708 indicate 

minimal wall-rock involvement. The Rb/Sr ratios of these plutons are 

very large (as much as 49.0), probably due to differentiation. 

Extensive hydrothermal circulation systems developed during the 

emplacement of both the Cretaceous-early Tertiary plutons and the 

Eocene plutons. Belt Supergroup (Wallace Formation) metasedimentary rocks around the northern Idaho batholith show a progressive 

decrease from 6180 ~ + 15± 1 per mil in low-grade metamorphic zones 

more than 60 km from mapped pluton contacts to 6180 ~ + 11.5± 1 per 

mil in the highest grade metamorphic zones (sillimanite ± kyanite) 

adjacent to the batholith. These data reflect large-scale and deep circulation of high-6180 metamorphic-hydrothermal fluids during Cretaceous-early Tertiary time. In contrast, meteoric-hydrothermal interactions involving heated, low-dD, low-6180 fluids occurred around the 

Eocene epizonal plutons and caused large reductions in dD values ( -120 

to -148 per mil) over an approximately 7,000-km2 area of the Bitterroot lobe. Major 6180 depletions (as low as -0.3 per mil) of the rocks 

occur only in restricted areas, however, indicating that the overall 

water/rock ratios were not very large. Most ore deposits occur near the 

perimeter of the Bitterroot lobe and therefore are probably related to 

the hydrothermal systems established during emplacement of the main 

batholith mass rather than to the Eocene meteoric-hydrothermal 

systems, in marked contrast to relations in the southern Idaho batholith. 

Although the whole-rock Rb-Sr isochron systematics of the Cretaceousearly Tertiary plutons and particularly the Eocene plutons were relatively unaffected by the hydrothermal systems, the metamorphic-hydrothermal activity appears to have significantly altered the isochrons of 

Precambrian Belt metasedimentary rocks. 

K-Ar ages of hornblende range from ~160 Ma for some Jurassic 

plutons in the Wallowa-Seven Devils terrane to less than 51 Ma for 

Eocene granitic rocks in the Bitteroot lobe. Some low-ri plutons in the 

Wallowa-Seven Devils terrane have concordant hornblende and biotite 

K-Ar ages of 135 Ma and show no geochemical indication of proximity 

to the Precambrian craton. These data, together with K-Ar ages of 87 

to 75 Ma for transition zone plutons, indicate that the tectonic suture 

between the Wallowa-Seven Devils terrane and the Belt-Y ellowjacket 

terrane formed between 135 and 75 Ma, and probably between 95 and 

80 Ma. Most of the Bitterroot lobe was emplaced after this probable 

accretionary event. However, the K-Ar apparent ages of biotite and hornblende become systematically younger toward the east-central part of 

the Bitterroot lobe, where biotite data define a 10,000-km2 area where 

apparent ages range from 51 to 44 Ma. The relations are thought to 

reflect temporal migration of magmatism as well as rapid uplift and 

cooling of the region. 

INTRODUCTION 

Stable- and radiogenic-isotope studies have contributed 

much new knowledge to the geology of the Idaho batholith. K-Ar studies by McDowell and Kulp (1969), Armstrong (197 4), Armstrong and others (1977), and Criss and 

others (1982) helped establish the regional importance of 

Tertiary plutonism and placed constraints on the emplacement age and uplift history of the batholith. Sr-isotope 

data of Armstrong and others (1977), Chase and others 

(1978), and Fleck and Criss (1985) also constrained age 

relations, helped identify the source materials from which 

the batholith was made, and were used to map a fundamental terrane boundary that coincides with the edge of 

95 
96 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

the Precambrian continental crust in the Pacific Northwest. U-Pb studies by Reid and others (1970, 1973), 

Grauert and Hofmann (1973), and Bickford and others 

(1981) provided some information on the emplacement 

ages of the plutons and presented unequivocal evidence 

that a significant Precambrian component is present in 

most of them. The hydrogen- and oxygen-isotope data of 

Taylor and Magaritz (1978) and Criss and Taylor (1983) 

demonstrated the importance of Tertiary hydrothermal 

metamorphism in the batholith and were used to map 

individual hydrothermal systems and to relate those 

systems to ore deposits. 

This paper presents considerable new o18Q, dD, 87Sr/ 

86Sr, K-Ar, 40 Ar/ 39 Ar, and chemical data from northern 

Idaho and adjacent areas. We attempt to summarize the 

salient results of pertinent previous investigations, provide more detailed maps of isotopic variations in the 

batholith, and augment present understanding of regional 

geologic relations. Data tabulated in this paper include 

results presented graphically in a companion study by 

Fleck and Criss (1985). 

ACKNOWLEDGMENTS 

We thank J.F. Sutter, M.A. Lanphere, H.P. Taylor, 

T.L. Vallier, and J.R. O'Neil for discussion and criticism, 

and W.R. Greenwood for providing assistance and invaluable information in the field. Elliot Kollman and Anne 

Tapay assisted in Rb-Sr analyses. Jarel Von Essen, James 

Saburomaru, Suzanne Makita, and Elliot Kollman performed argon analyses and assisted in sample preparation. Potassium analyses were made by Paul Klock, Sarah 

Neil, and David Vivit. 

GEOLOGIC SETTING 

The Idaho batholith includes granitic plutons intruded 

into Precambrian crystalline and metasedimentary rocks. 

However, a contiguous but geologically distinct group of 

plutons west of the main batholithic mass is in a region 

entirely devoid of Precambrian rocks. In this report we 

will use the terms "plutons of the Bitterroot lobe" to refer 

to the northern part of the Idaho batholith where 

Precambrian-hosted plutons occur, the "Jurassic and 

Cretaceous plutons in the Wallowa-Seven Devils terrane" 

(WSD) to refer to the distinct group of plutons in the accreted terrane west of the Bitterroot lobe, and "plutons 

of north-central Idaho" to refer to both groups of plutons 

in a general sense. The term "transition zone plutons" 

is used to refer to a distinct group of geochemically 

variable, Late Cretaceous granitoids that occur only 

within and near the boundary between the WSD and Belty ellowjacket terranes. 

The Bitterroot lobe underlies a nearly 15,000-km2 area 

of north-central Idaho and adjacent parts of Montana (fig. 

6.1). The plutons of the Bitterroot lobe can be divided into 

a Cretaceous-early Tertiary main-phase group and an 

Eocene epizonal group (Greenwood and Morrison, 1973; 

Hyndman and Williams, 1977). The main-phase plutons 

generally comprise leucocratic granodiorite and granite, 

are metaluminous to weakly peraluminous, and define a 

calcic to calc-alkaline trend (Criss, 1981; Hyndman, 1984). 

Hyndman (1983) has likened the main-phase plutons to 

the sodic series of the Boulder batholith, Montana, 

described by Tilling (1973). The main-phase plutons were 

emplaced at relatively deep crustal levels of generally 

~10 km (Swanberg and Blackwell, 1973; Hyndman, 1981, 

1983). No definitely comagmatic volcanic rocks are 

known, although volcanic detritus occurs in several 

Cretaceous formations in eastern Idaho, Montana, and 

Wyoming. The emplacement of the main-phase plutons 

was attended by high-grade regional metamorphism of 

sedimentary rocks of the Belt Supergroup of Middle Proterozoic age (Hietanen, 1962, 1963, 1984). However, a 

regional discordance between some of the metamorphic 

isograds and the margin of the batholith (Chase and 

Johnson, 1977; Lang and Rice, 1985) suggests that this 

metamorphism in part predated batholith emplacement 

or that large areas of high-grade metamorphic rocks may 

be underlain by part of the batholith. Numerous inliers 

and screens of metamorphic rock occur within the Bitterroot lobe (Larsen and Schmidt, 1958; see fig. 6.1). 

A significant part of the east-central Bitterroot lobe 

includes Eocene epizonal plutons (fig. 6.1). These calcalkaline plutons consist predominantly of pink granite, 

although granodiorite and quartz diorite are also present, 

and tend to be metaluminous to weakly peraluminous. 

Distinctive characteristics of many of these Eocene 

plutons include high heat-generation values (Swanberg 

and Blackwell, 1973), smoky quartz, and common miarolitic cavities (Bennett, 1980); additional properties are 

described by Ross (1934), Williams (1979), Criss (1981), 

Criss and Champion (1984), and Motzer (1985). Although 

some deep-seated plutons are Eocene in age (Bickford and 

others, 1981; Chase and others, 1983) most plutons of 

Eocene age had an epizonal emplacement, and some intrude approximately coeval rocks of the Challis Volcanics 

(Ross, 1934; Cater and others, 1973; Swanberg and 

Blackwell, 1973; Armstrong, 1974). 

The Bitterroot lobe is bounded on the east by a 100-kmlong mylonite zone (fig. 6.1) that dips approximately 25° 

east along the prominent front of the Bitterroot Mountains (Lindgren, 1904; Langton, 1935; Chase, 1973). 

Various models have been advanced to explain this major 

structure, which may be related in some way to a domelike 

uplift-termed the "Bitterroot dome"- of the batholith 

interior (Chase and Talbot, 1973; Hyndman and others, 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 97 

1975; Chase and others, 1983). Major Cretaceous and 

early Tertiary uplift of the greater Idaho batholith region 

is required to explain (1) tectonic features (Chase and 

Talbot, 1973; Scholten and Onash, 1977), (2) the juxtaposition of Cretaceous-early Tertiary and Eocene plutons 

having markedly different inferred emplacement depths 

(Swanberg and Blackwell, 1973; Chase and others, 1983), 

(3) fission-track and K-Ar apparent ages (Ferguson, 1975; 

Criss and others, 1982; Garmezy and Sutter, 1983; Sutter 

and others, 1984), and (4) the record of clastic sedimentation (Ryder and Ames, 1970; Ryder and Scholten, 1973). 

However, the uplift documented by Criss and others 

(1982) for the southern Idaho batholith may have started 

earlier and was much more gradual than that inferred by 

Garmezy and Sutter (1983) for the Bitterroot lobe. 

In western Idaho, the Bitterroot lobe is bounded by 

another zone of intense deformation. This 5- to 25-kmwide zone coincides with an important tectonostratigraphic boundary between the cratonal Belt-Y ellowjacket 

terrane on the east and the Wallowa-Seven Devils terrane 

on the west; the latter may be an accreted microplate 

(Hamilton, 1976; Jones and others, 1977). Stratified rocks 

in the Wallowa-Seven Devils terrane are predominantly 

Permian and Triassic volcanic and volcaniclastic rocks, 

in places overlain by Jurassic flyschlike sedimentary rocks 

(Vallier, 1977; Walker, 1983). These rocks were affected 

by low-grade metamorphism and west-directed thrusting, 

followed by and in places associated with intrusion of 

Jurassic and Cretaceous quartz diorite gneiss, tonalite, 

and trondhjemite (Hamilton, 1963). These plutons are 

generally calcic and metaluminous (see Hietanen, 1963), 

and in many places were also emplaced at deep levels 

(Swanberg and Blackwell, 1973; Zen and Hammarstrom, 

1984). All the above rocks are thought to represent an 

island arc (Hamilton, 1976; Vallier, 1977). 

Juxtaposition of the Wallowa-Seven Devils terrane and 

the Belt-Y ellowjacket terrane apparently occurred during the Late Cretaceous (Fleck and Criss, 1985), insofar 

as their suture zone is intruded by deformed transition 

zone plutons having K-Ar ages of 87 to 75 Ma and 

characteristics intermediate between those of the WSD 

and the Bitterroot lobe. Most plutons in the Bitterroot 

lobe were emplaced after this period of accretion. Initial 

intrusion of the main phase of the Idaho batholith apparently coincided with submergence and renewed marine 

sedimentation in the epicontinental Cretaceous seaway 

to the east, which extended from the Gulf Coast to the 

Arctic Ocean (McGookey and others, 1972). 

METHODS 

The 87Srf86Sr ratios, Rb and Sr concentrations, and 

a180 and aD values were made on crushed (- 200 mesh) 

whole-rock powders of samples. Standard techniques were 

used in all cases; details pertinent to the Sr-isotope study 

are given by Fleck and Criss (1985). Rb and Sr concentrations were measured by isotope dilution or by X-ray 

fluorescence, and strontium-isotope ratios were measured 

on a thermal-ionization mass spectrometer. Calculations 

of initial S7SrfS6Sr ratios (ri) utilized the measured 

isotopic ratios of Sr, the Rb/Sr ratios, and assumed ages 

as discussed by Fleck and Criss (1985). However, in cases 

where independent age data were available, these ages 

were used in the calculations. Because Rb/Sr ratios are 

low in all but the Eocene epizonal plutons, the calculated 

ri values are generally insensitive to large uncertainties 

in age. 

Determinations of alSO values on the whole-rock samples were made by techniques similar to those described 

by Clayton and Mayeda (1963) and by Borthwick and 

Harmon (1982); the techniques used for aD determination 

are similar to those described by Friedman (1953) and by 

Godfrey (1962). All aD and a1SO determinations are 

reported in the ordinary notation as per mil deviations 

from the SMOW standard (Craig, 1961); precision is 

generally better than ± 0.2 per mil for alSO and ± 1.0 per 

mil for aD. K-Ar ages reported here were determined 

utilizing conventional isotope-dilution techniques (Dalrymple and Lanphere, 1969), with selected samples studied 

by 40 Arf39 Ar methods (Dalrymple and Lanphere, 1971). 

All data are reported in tables 6.1-6.5; refer to table 6.1 

for sample locations. 

GEOCHEMICAL CHARACTER OF 

THE PLUTONIC ROCKS 

REGIONAL VARIATIONS OF Rn RB/SR, AND 

MAJOR-ELEMENT CHEMISTRY 

Kistler and Peterman (1973) discovered that the ri 

values of Sierra Nevada plutons vary with geographic 

position relative to the Paleozoic miogeosyncline and 

eugeosyncline of the cordillera. Armstrong and others 

(1977) extended this concept to the Idaho batholith region 

and demonstrated that the ri values changed abruptly 

(from ~0.704 to ~0.706) across the geologic transition 

between the Phanerozoic eugeosyncline to the west and 

Precambrian crust to the east. They attributed these 

variations to contamination of mantle-derived magmas by 

assimilation of enclosing crustal rocks having different 

chemical and isotopic characteristics. These mechanisms, 

together with large-scale anatexis, were confirmed by the 

results of Fleck and Criss (1985). 

New determinations of r i in the Idaho batholith region 

(table 6.1; fig. 6.2) are in large part consistent with the 

observations of Armstrong and others (1977). Whereas 

the WSD plutons in the accreted terranes all have r i 

values less than 0. 7042, most Bitterroot lobe plutons in 
98 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

the Belt-Yellowjacket terrane have ri values greater than 

0.7080. The highest ri values (as high as 0.732) occur in 

plutons adjacent to screens and pendants of Precambrian 

rocks (compare figs. 6.1 and 6.2), indicating significant 

incorporation of such material in the magmas (Fleck and 

Criss, 1985). Away from such zones, the ri values of 

plutons in the Bitterroot lobe generally range from 0.7080 

to 0.7100. 

Deformed plutons having ri values between 0.704 and 

0.708 generally occur within the 5- to 25-km-wide tectonic 

zone along the boundary between the Wallowa-Seven 

Devils terrane and the Belt-Yellowjacket terrane. These 

transition zone plutons contain materials derived from 

both the accreted and Precambrian terranes, and thereFIGURE 6.1.-Generalized geologic map of north-central Idaho and adjacent areas, modified from Ross and others (1955), Bond (1978), Chase 

and others (1978), and Toth and others (1983). Heavy line is the 0. 706 

initial-Sr contour (see figure 6.2), which approximately delineates a 

major suture zone between the Precambrian craton to the east and 

fore provide important clues regarding the significance 

of ri variations in Idaho as well as on the geochronology 

of the tectonic suture zone they intrude. Fleck and Criss 

(1985) provide a detailed discussion of the ri, d180, and 

Rb and Sr relations of these rocks and of the neighboring Bitterroot lobe and WSD plutons. 

The Rb and Sr concentration data also distinguish the 

plutonic series that comprises the northern Idaho batholith region (fig. 6.3). The low Rb concentrations and low 

Rb/Sr ratios of the WSD plutons contrast markedly with 

the high Rb contents and high Rb/Sr ratios of the Eocene 

epizonal plutons, whereas the Rb contents and Rb/Sr 

ratios of the main-phase granitoids of the Bitterroot lobe 

are intermediate. This variation is clearly paralleled by 

the accreted Wallowa-Seven Devils terrane. Abbreviations of Eocene 

epizonal plutons are BP (Bungalow pluton), LHS (Lolo Hot Springs 

pluton), WP (Whistling Pig pluton), RC (Running Creek pluton), PR 

(Painted Rocks pluton), CG (Crags pluton), and CP (Casto pluton). 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 99 

variations in K, with the alkalis generally increasing 

transverse to orogenic structures, as has bee_n documented in the Sierra Nevada batholith by Bateman and 

Dodge (1970). As pointed out by Hyndman (1984), however, no clearcut eastward trend appears if the mainphase plutons are considered alone. 

Consistent differences between the various plutonic 

groups are also revealed by bulk chemical analyses (table 

6.2). A variation diagram (fig. 6.4) illustrates that plutons 

in the Wallowa-Seven Devils terrane are significantly 

more calcic than the calcic to calc-alkaline plutons in the 

Bitterroot lobe. Figure 6.5 shows that the Eocene epizonal 

plutons lie along the calc-alkaline trend defined by the 

Challis Volcanics. 

Temporal trends also appear in the chemical data (see 

tables 6.1 and 6.2). Aside from being more calcalkaline, 

younger plutons tend to have higher Si02 concentrations 

and higher degrees of aluminum saturation than older 

plutons. Exceptions are common, however, probably indicating that shorter differentiation cycles or other effects 

are superimposed on these overall chemical trends. 

d180 AND dD DATA 

The o180 and dD values for the Mesozoic and Tertiary 

granitoids of north-central Idaho exhibit great variations. 

Whole-rock o180 values range from -0.3 to + 12.4 per 

mil, and the dD values range from -148 to -41 per mil 

(table 6.1). As discussed below, the large also variations 

can be partly attributed to variable source characteristics 

of the plutons. However, the lowest o180 and dD values 

result from hydrothermal alteration processes, and these 

secondary effects must be quantified before discussion of 

the primary isotopic variations is possible. 

~ 

~ 

EXPLANATION 

Quaternary and Tertiary deposits 

Tertiary epizonal intrusive rocks 

Tertiary Challis Volcanics 

Mylonite zone 

Tertiary and Cretaceous intrusive rocks 

Intrusive rocks (Mesozoic) of the accreted terranes -- mostly Jurassic and Cretaceous plutons; probably includes some Triassic plutons 

Jurassic to Permian metasedimentary and 

metavolcanic rocks of accreted terranes 

Precambrian rocks, undifferentiated 

Geologic contact 

FIGURE 6.1.-Continued. 

Taylor and Magaritz (1978) proposed that hydrothermal circulation systems involving fluids derived from low 

alBO, low dD meteoric waters developed in the Idaho 

batholith during the intrusion of the Eocene epizonal 

plutons. Isotopic exchange between this meteoric-hydrothermal fluid and the rocks produced the major alSO and 

dD depletions in the rocks within and peripheral to the 

Eocene plutons. Criss and Taylor (1983) estimated that 

the Eocene meteoric waters in the southern Idaho batholith had d180 and dD values of -16 and -120 per mil, 

respectively; the Bitterroot lobe hydrothermal fluids were 

apparently derived from waters with similarly low values. 

A graph of o180 versus dD values (fig. 6.6) shows that 

most Bitterroot lobe samples with 80~ + 8.0 per mil 

have extremely low dD values ( -120 per mil or less). As 

discussed by Taylor (1977), the coupled isotopic depletions 

in such rocks are inferred to have been produced by 

meteoric-hydrothermal alteration at high water/rock 

ratios. On the other hand, the great range in dD observed 

in rocks with alSO~+ 8.0 per mil is attributed to weak 

Qow water/rock ratios) hydrothermal interactions that, for 

material balance reasons described by Taylor (1977), can 

markedly lower the dD values without significantly modifying the o180 values. 

According to the above interpretation, the primary 

(igneous) whole-rock d180 values of the plutons in the 

Bitterroot lobe and the northern Atlanta lobe range from 

+ 8.0 to + 12.4 per mil. This result is similar to the range 

of+ 8 to + 11 per mil obtained by Criss and Taylor (1983) 

for the bulk of the southern (Atlanta) lobe of the batholith. 

On the other hand, the low o180 values of some of the 

WSD plutons cannot be attributed to these meteorichydrothermal interactions because they are far from the 

Eocene epizonal plutons and because their dD values are 

invariably ~- 82 per mil. Although it is possible that the 

WSD plutons underwent interactions with hydrothermal 

fluids having different isotopic characteristics (for example, sea water), the d180 and dD values of the rocks are 

not correlated in a systematic way that would be expected 

if this were the case. We prefer the interpretation that 

the alSO range of + 5.5 to + 10.5 per mil observed in 

WSD plutons is primary. 

Because of the sensitivity of dD values to alteration 

processes, the primary dD range of the plutons of northcentral Idaho is more difficult to establish. Taylor and 

Magaritz (1978) suggested that primary dD values of 

major cordilleran batholiths range from about -50 to - 85 

per mil, which is similar to the range observed in most 

deep-seated, isotopically normal igneous rocks throughout 

the world. Similar dD values (>- 85 per mil) occur within 

the WSD plutons and in parts of the Bitterroot lobe that 

are remote from mapped Eocene epizonal plutons. Thus, 

the dD values in these zones are probably primary, an 

interpretation supported by dD measurements of biotite-
TABLE 6.1.-Results of Rb-Sr, cY

80, and dD analyses 1---' 

0 

0 

North West Rb Sr 

Sample Latitude Longitude 87Rb/86Sr 87Sr/86Sr (87Sr/86Sr) 0180 (ppm) (ppm) Age (Ma) oD 0 

787-18A 46°14.0 1 115°24 • 6 I 87.8 986 0.25766 0.70940 0.7092 60 8.2 -88 

188 46°15.1 I 115°24.0 1 115.5 647 .51661 .71077 .7104 60 9.6 -86 

18C 46°17 • 1 I 115°23.1 I 80 488 .47453 .71471 .7143 60 10.3 -95 

18D 46°21.5 1 115°17.2 I 95.6 535 .51743 .71221 .7118 60 10.8 -107 

18E 46°24.3 1 115°10.2 I 75.5 864 .25272 .70940 .7092 60 9.2 -111 

18F 46°27 • 1 I 115°05.3 1 82.5 944 .25291 .70892 .7087 60 9.7 -98 

18G 46°28.5 1 114°53.1 I 93.9 500 .54378 .71045 .7100 60 10.7 -111 

18H 46°34.9 1 114°36.5 1 89.2 695 .37143 .70923 .7089 60 6.0 -145 

181 46°34.9 1 114°36 • 5 I 83.6 822 .29419 .70943 .7092 60 9.3 -127 

18J 46°40.6 1 114°34.3 1 184.4 55.9 9.548 .71361 .7072 45 7.2 -126 

1-3 

19A 46°09.5 1 115°21.7 1 53.1 542 .2835 .71169 .7114 70 8.8 -98 ::q 

198 46°09.1 I 115°21.7 I 42.2 594 .2056 .71103 .7108 70 8.5 t:z:::l -- 1--4 

19C 46°08.9 1 115°21.6 I 59.4 510 .3374 .71108 .7107 70 8.5 -B !Z 19E 46°10.6 1 115°13.8 I 52.5 567 .2682 .71183 .7116 70 7.5 -127 ::q 19F 46°14.9 1 115°12.6 1 131 469 .8102 .71332 .7126 60 9.3 -59 0 

46°14.9 1 115°12.6 I 

t:d 

19G 157 437 1. 0361 .71213 .7114 60 10.0 -76 ~ 20A 46°10.3 1 115°11.6 1 59.4 524 .32775 .71219 .7119 60 8.3 -109 ::I: 

20C 46°12.8 1 115°09.6 I 106 492 .62453 .71965 .7191 60 10.3 -103 0 

20D 46°04.9 1 115°08. ')I 44.7 599 .21604 .71404 .7138 70 8.8 -73 t""' 

1--4 

20E 46°04.9 1 115°08.5 1 42.0 607 .20000 .71475 .7146 70 8.0 -78 1-3 

::I: 

21A 46°03.3 1 114°57.2 1 157 364 1.158 .70818 .7073 45 9.3 -14') > z 

218 45°59.0 1 114°57.9 1 237 8.2 84.063 .76360 .7065 48 9.0 -- t::1 

21C 45°58.9 1 114°57 • 9 I 220 26.5 24.043 .72294 .7075 48 8.2 -128 :::3 

21D 45°58.4 1 115°00.1 I 224 148 4.3786 .71149 .7085 48 9.0 -126 w. 

21E 46°01 • 1 I 114°50 • 7 I 50.3 1033 • 14098 .70937 .7092 60 9.1 -- t:d 

0 

46°01.1 I 114°50.7 1 

~ 

21F 59.6 1055 .16340 .70967 .7095 60 8.8 -131 t::1 

24A 46°10.2 1 114°38.5 1 71.6 1014 .20429 .70901 .7088 60 9.2 -100 t:z:::l 

~ 248 46°10.2 1 114°38.6 1 74 1039 .20624 70904 .7089 60 8.8 -- N 

24C 46°10.2 1 114°38.6 1 70 1224 .16555 .70829 .7081 60 9.0 -81 0 

24D 46°08.1 I 114°36.9 1 68 864 .22751 .70987 .7097 60 9.0 -- z 

24E 46°08.1 I 114°36.8 1 <1 806 <.00124 .71032 .7103 60 8.8 -117 t:z:::l 

24F 46°08.2 1 114°36.9 1 78.2 759 .29805 .71000 .7097 60 8.8 

24G 46°08.2 1 114°36.9 1 57.8 561 .29805 .70648 .7063 50 8.5 -108 

24H 46°05.3 1 114°32.6 1 97.7 432 .65442 .70944 .7090 50 8.8 -85 

241 46°05.3 1 114°32.7' 105 542 .56022 .70942 .7090 50 9.5 -80 

25A 46°07.5 I 114°56.0 I 139 90.7 4.4226 .71206 .7089 50 1.2 -141 

258 46°04.8 1 114°58.4 1 177 206 2.4898 .71068 .7089 50 2.3 -141 

25C 46°04.8 1 114°25.5 1 102 746 .39433 .70926 .7090 50 10.8 -91 

26A 46°13.2 1 115°31.9 1 87.6 883 .28726 .70888 .7086 70 9.3 -72 

268 46°13.2 1 115°31. 9' 59.2 898 .19073 .70937 .7092 70 7.2 -136 

26C 46°13.2 I 115°31.9' ')4.2 814 .19273 .70924 .7090 70 8.7 -129 

26D 46°12.0 1 115°32 • 5 I 43.6 889 .14186 .71360 .7135 70 9.9 -109 

26E 46°12.0 1 115°32.5 I 32.9 682 .13963 .71492 .7148 70 9.1 -85 

26F 46°12.0 1 115°32.5 I 56.7 435 .37747 .71388 .7135 70 8.4 -141 
797-18A 45°33.3 1 114°32.9 1 223.6 210.2 3.0782 0.70745 0.7053 50 6.9 -131 

188 45°33.4 1 114°32 • 8 I 221.2 83.63 7.6638 .72402 .7153 80 

18C 45°38.4 I 114°25.1 I 288.4 46.86 17.826 .71850 .7058 50 8.4 -118 

18D 45°39.6 1 114°30.1 I 205.6 117.9 5.04708 .71089 .7073 50 -- -- ~ 

18E 45°33.8 1 114°34.1 I 206 219.7 2.7132 .70859 .7067 50 -- -- ""0 

tz:j 

19A 45°43.1 1 114°28.0 1 215.4 68.17 9.1496 .71530 .7088 50 t-3 -- -- ~ 

198 45°43.1 I 114°28.0 1 232.7 61.51 10.953 .71574 .7087 45 8.1 -- 0 

19C 45°41.5 1 114°28.9 1 221.4 226.5 2.8288 .70978 .7079 45 -- -- Q 

tz:j 

19D 45°32.7 I 114°24.4 1 36.63 142.5 .74731 .70752 .7067 80 3.4 -157 z 

19E 45°32 o 5 I 114°24.8 1 31.27 135.9 .66627 .72124 .7205 80 6.3 -145 tz:j 

en 

1--i 

45°32.5 1 114°24.8 1 en 19F 31.9 138.2 .66906 .72452 .7238 80 4.4 -- ~ 

Q 19G 45°32.6 1 114°26.6 1 251.1 5Q. 91 12.137 .71800 .7094 50 9.8 -126 tz:j 

20A 45°44.8 I 114°26.8 1 272.9 19.90 39.804 .74003 .7080 50 -- -- 0 

208 45°42.4 1 114°36.0 I 335.9 6.728 145.88 .80904 .7054 50 5.8 -133 n 

20C 45°42.1 I 114°38.3 1 174.1 257.9 1.9539 .71260 .7112 50 ::I: -- -- ~ 

0 

20D 45°41.6 1 114°39.3 1 220.1 214.2 2. 9744 .71193 .7098 50 -- -- z 

0 

20E 45°44.0 1 114°44.8 1 106.1 318.1 .96552 .71076 .7101 50 -- -- ~ 

20F 45°44.1 1 114°45.0 1 151.6 82.40 5. 3263 .71140 .7076 50 -- -152 0 

20G 45°39.0 1 114°35.7 1 214.2 54.18 11.452 .71797 .7098 50 -0.3 -133 ~~ 21A 45°43.2 1 114°16.6 1 325.8 14.46 65.469 .75402 .7074 50 7.8 -140 > z 

22A 46°09.7 1 114°47.2 1 82.40 845.6 .28196 .70882 .7086 60 10.2 -95 t::l 

228 46°13.2 1 114°27.1 I 61.88 904.9 • 19788 .70923 .7091 60 9.4 -90 ::I: 

22C 46°13.3 1 114°27.0 1 60.35 1003 • 17584 .70882 .7087 60 9.2 -89 ~ 

t::l 

23A 46°08.4 1 115°36.1 I 29.61 704.3 • 12164 .70866 .7085 93 -- -- ~ 

238 46°20.4 I 115°06.7 I 115.4 808.7 .41286 .71144 • 7111 60 9.9 -123 ~ ::I: 

23C 46°18.1 I 115°01.4 I 92.15 816.6 .32655 .70950 .7092 60 3.4 -131 tz:j 

24A 45°56.2 I 115°03.0 1 203.9 32.44 18.211 .72093 .7080 50 ~ 

-- -- ~ 248 45°54.6 1 114°43.9 1 124.6 122.6 2.9410 .71090 .7088 50 -- -131 > 

24C 45°57.4 I 114°37.4 1 79.26 1013 .22647 .70855 .7083 60 -- -141 ~ 

25A 46°08.2 1 115°37 o 3 I 51.29 946.1 .15688 .70948 .7093 90 -- -- en 

~ 

en 

258 46°08.2 1 115°37.3 1 44.16 • 14155 .70941 .7092 90 10.1 t-3 902.7 -- tz:j 

25C 46°08.3 1 115°39.4 1 -8.31 111.7 • 21756 .81459 .8144 75 -- -- ~ 

25D 46°08.2 1 115°45.1 I 76.70 483.2 .45Q23 .70798 .7075 75 -- -- en 

25E 46°08.2 1 115°45.1 I 48.28 653.9 .21361 .70753 .7073 80 9.4 -71 0 

"'%j 

25F 46°08.5 I 115°51.8 I 23.05 615.3 • 10837 .70349 .7034 80 6.7 -78 t-3 

46°01.8 1 

::I: 

25G 115°58.4 I .533 304.9 .00505 .70408 .7041 80 -- -- tz:j 

25H 46°01.5 1 115°58.2 I 16.46 261.7 • 18191 .70424 .7046 80 -- -71 z 

25I 46°00.4 1 115°57.5 1 9.48 333.9 .08210 • 70398 .7039 80 n.6 -5R 0 

26A 46°15.3 1 116°03.8 I 7.142 269.2 .07673 .70378 .7037 80 5.5 -64 ~ 268 45°57.0 1 115°58.7 I 15.35 635 .06988 .70325 .7032 80 7.8 -51 ::I: 

tz:j 

26C 45°57.0 1 115°58.7 I 5.52 379 .04214 .70391 .7039 80 5.9 -46 ~ 

z 

26D 45°53.3 1 116°02.1 I 26.19 524.6 .14439 .70386 .7037 80 9.0 -132 1--i 

26E 45°50.5 1 115°59.9 I 26.25 805.5 .09425 .70387 .7038 80 8.5 -66 ~ 26F 45°49.7 1 115°33.3 I 96.89 88.18 3.2032 .78462 .7810 80 11.3 -- ::I: 

26G 45°47.6 1 115°45.1 I 69.95 368.4 .54996 .71751 .7170 60 11.2 -79 0 

26H 45°49.2" 115°49.7 1 114 562.3 .58696 .71654 .7159 80 10.2 -R1 to 

26I 45°49.6 1 115°54.1 I 64.46 900.5 .20718 .70808 .7078 80 9.6 -- ~ ::I: 

0 

~ 

:::3 

::I: 

~ 

0 

~ 
TABLE 6.1.-Results of Rb-Sr, 

6180, and oD analyses-Continued """"' 0 

~ 

North West Rb Sr 

Sample Latitude Longitude (ppm) (ppm) 87Rb/86Sr 87Sr/86Sr (87Sr/86Sr) Age 0180 (Ma) oD 0 

807-27A 45°50 • 1 I 115°58.9 1 23.2 832 0.08065 0.70375 0.7037 86 10.4 

27B 45°49.8 I 115°58.4 I 68.6 189 1.05002 .70624 .7050 86 

27C 45°49.8 1 I 115°55 • 9 I 38.2 852 .12969 .70493 .7048 86 9.1 

28A 45°50.1 I 115°57.9 I 51.9 513 .29261 .70420 .7038 86 -- 288 4'5°50.2 1 115°56.0 I 33.3 1228 .07844 .70476 .7047 86 8.5 

28C 45°48.2 1 1.15°58.6 I 9.71 303 .09269 .70488 .7048 80 

280 45°48.? 1 115°58.6 1 17.2 357 .13936 .70497 • 7048 80 

28E 45°48.2 1 115°58.6 1 34 107 .91930 .70695 .7059 80 

28F 45°49.8 1 115°55 • 8 I 54.8 988 .16045 .70588 .7057 86 9.8 

28G 45°49.7 I 115°55.5 1 45.6 842 .15666 .70577 .7056 86 10.2 

45°49.6 1 t-3 

28H 115°55 • 5 I 29.9 1406 .06152 .70657 • 7065 86 10.4 -- :::0 

28! 45°49.4 1 115°55.0 1 41.7 709 .01702 .70748 .7073 86 9.2 trl -- ~ 29A 45°49.4 1 115°51.6 1 110 461 .69079 .71377 .7130 80 11.2 -- 298 45°49.4 1 115°51.5 I 104 529 .56905 .71180 .7112 80 10.2 -- 29C 45°49.3 1 115°51.3 I 96.8 599 .46774 .71142 .7109 80 11.6 :::0 -- 0 

45°49.3 1 115°50.7 I 

t:d 

29E 

290 

45°49.0 1 115°48.9 1 184 

65.5 726 .26110 .71017 .7099 80 9.9 -- ~ 14.2 37.5833 • 73292 .7089 45 -- -- :::0 29F 45°52.0 1 115°48.5 I 98.5 449 .63268 .71645 .7157 80 12.1 -- 0 

29G 45°48.5 1 115°48.0 1 89.1 513 .50290 .71529 .7147 80 11.4 -- t"' 

~ 

29H 45°48.0 1 115°47.0 I 62.9 672 .27123 .72338 .7231 80 11.5 -- t-3 

:::0 

29! 45°47.8 1 115°46.2 1 108.4 407 .77148 .71929 .7184 RO 10.9 > -- z 29J 45°47.7 1 115°45.5 I 97.8 710 .39869 .71142 .7110 80 10.4 -- t;j 

30A 45°47.8 1 115°44.7 1 72.7 570 .36924 • 71376 .7133 80 10.9 -- ~ 

t-3 308 45°47.7 1 115°43.5 I 66.0 517 .36927 .71512 .7147 80 10.0 -- r:n 

30C 45°47.7 1 115°43.5 1 88.3 479 .53373 .71483 .7142 80 10.9 -- t:d 

0 

45°47.9 1 115°42.7 I 

~ 30D 59.2 610 .28101 .71556 .7152 80 11.5 -- t;j 

30E 45°48.7 1 115°39.5 1 117 163 2.08641 • 75502 • 7526 80 10.8 -- trl 

30F 45°48.7 1 115°39.7 I 113.7 167 1. 97866 .75328 .7510 80 10.9 ~ 

-- N 30G 45°48.7 1 115°39.7 1 67.3 189 1.03416 .75248 • 7513 80 12.0 -- 0 

30H 45°48.7 I 115°39.5 I 106.1 192 1.60481 .74575 .7439 80 9.5 -- z 

trl 

30! 45°48.4 1 115°38.4 I 113.1 326 1. 00628 • 73313 • 7320 80 9.1 

30J 45°48.8 I 115°37.8 1 83.7 433 .55998 .72046 .7198 80 10.4 

30K 45°48.7 1 115°37 • 8 I 284.5 94.3 8.93133 • 94462 .9345 80 8.9 

31A 45°49.7 1 115°37.3 1 167.9 277.1 1. 75735 .73241 .7304 82 11.0 

318 45°49.7 1 115°37 o 3 I 92.9 356.4 • 75520 .72151 .7206 82 10.7 

31C 45°49.7 I 115°36.7 I 60.1 426.9 .40863 .74081 .7404 65 10.6 

31D 45°49.5 1 115°35.7 1 57.9 397 .42239 .71787 .7175 65 11 • 1 

31E 45°49.4 I 115°35.0 1 49.7 805 .17881 .71820 .7180 65 10.1 

31F 45°49.8 1 115°34.9 1 50.9 1262 .11680 .71683 .7167 65 10.4 

31G 45°49.4 I 115°34.1 I 59.8 651 .26612 .72092 .7207 65 10.9 

31H 45°49.4 1 115°34.1 I 35.6 620 .16635 .72137 .7212 65 10.9 

311 45°49.9 I 115°33.4 I 42.5 490 .25130 .72207 .7218 65 10.6 

31J 45°49.8 1 115°33.4 1 84.1 298 • 81862 .73385 .7329 80 10.3 

31K 45°49.8 1 115°33.4 1 66.6 91.2 2.11955 .73995 .7374 85 8.7 
808-01A 45°49.5 1 115°31.6 I 77.8 898 0.25107 0.72425 0.7240 80 1?.4 

018 45°49.5 1 115°30.4 I 125.6 415 .88122 .77252 .7717 65 9.5 

01C 45°49.7 1 115°30,1 I 62.8 60.6 3.01601 .76790 .7645 80 9.4 

010 45°49,2 I 115°29.5 1 123 203 1,76165 .75762 .7556 80 9.8 -- 0 

01E 45°48.4 1 115°28.6 I 196 124 4.61352 • 79761 .7924 80 10,0 -- ""d 

M 

45°48.2 1 115°24.7 I 6.06141 .82884 .8220 80 1-3 01F 194 93.7 9.9 -- ~ 

01G 45°47.7 1 115°24.0 1 178 88.3 5.89950 .82513 .8184 80 9.3 -- 0 

C1 01H 45°48.4 I 115°28,3 I 143 176 2.36352 .76295 • 7603 80 9.7 -- M 

on 45°48.2 1 115°28.3 I 35 144 .70411 .72045 • 7196 80 7.7 -- z 

01J 45°47.8' 115°27.9 1 154 130 3.45392 .78662 .7827 80 9.2 -- M 

r:n 

45°47.8 1 115°26.4 I 194 .74658 .7446 80 

_w 01K 117 1. 75157 10.3 -- C1 01L 45°47.7 1 115°26.3 1 158 247 1. 85547 .73357 .7315 80 -- -- M 

04A 45°58.4 1 114°08.2 1 68.2 523 • 37748 .71286 .7125 65 7.8 -111 0 

048 45°58.1 I 114°08.4 1 74.5 538 .40087 .71342 .7130 65 8.4 -- CJ 

p:: 04C 45°56.8 1 114°12.3 1 76.6 114 1. 94976 .73747 .7361 50 -- -- ~ 

0 

040 45°55.8 1 114°13.4 I 75 563 .38535 .70562 .7053 65 z -- -- 0 

04E 45°55.1 1 114°09.9 1 70.5 594 .34361 .71401 .7137 55 -- -141 r 

04F 45°55.2 I 114°09:7 1 77.4 536 .41803 .71325 .7129 55 0 -- -- 04G 45°55.3 -~ 1 114°09.6 1 224 87.1 7.44514 .71353 .7093 40 -- -- 05A 46°17.6 1 114°16.1 I 102 700 .42175 .71137 .7110 55 10.5 -76 > z 

058 46°17.4 1 114°20.4 I 128 574 • 64546 • 71191 .7114 55 10.4 -- t:J 

05C 46°17.4 1 114°24.6 1 118 585 .58382 .71147 .7110 55 10.6 -72 p:: 

050 46°17.0 1 114°28.2 1 34 1381 .07127 .71243 .7124 55 10.3 ~ 

-- t:J 05E 46°17.8 1 114°36.3 1 92 638 .41736 .71107 .7107 55 10.6 -60 ~ 

~ 05F 46°08.7 1 114°26 • 2 I 62.3 861 .20941 .71023 .7101 55 9.3 -- p:: 

06A 46°08.6 1 114°28.3 1 73.4 727 .29220 .71070 .7105 55 9.3 M -- ~ 068 46°07.1 I 114°29.6 1 77.6 870 .25814 .71034 .7101 55 8.9 -- ~ 

06C 46°08.5 1 114°30.1 I 78.4 834 .27207 .71090 .7107 55 9.5 -- > 

060 46°08.8 1 114°30.1 I 62 872 .20576 .71002 .7099 55 r -- -- r:n 

46°08.8 1 114°30.1 I 

~ 

06E 67.6 901 .21713 .71008 .7099 55 -- -- r:n 

1-3 06F 46°08.9 1 114°30.0 1 85.6 751 .32988 .71066 .7104 55 -- -- M 

06G 46°08.4 1 114°25.9 1 65.8 1062 .17931 .71013 .7100 55 9.3 -- ~ 

06H 46°08.3 1 114°25.2 1 51.7 1006 .14873 .71005 .7099 55 r:n -- -- 0 061 46°08.3 1 114°24.2 1 61.2 853 .20763 .70981 .7096 55 8.8 -61 l-%j 

46°48.6 1 117°00.4 1 1-3 

818-09A 72 494 .42195 .71379 .7134 65 9.8 -93 p:: 

098 46°41.4 1 116°58.3 I 177 724 .70742 .70887 .7082 65 -- -- M 

09C 46°41.5 1 116°58.6 1 220 662 .96169 .70943 .7085 65 -- -- z 

0 090 46°57.6 1 116°19.2 I 70.2 1070 .18990 .71167 .7115 80 -- -- 09E 46°46.2 1 116°14.3 I 94.9 371 .74033 .71090 .7101 80 -- -- ~ :::c: 

46°46.2 I 116°14.3 1 M 09F 55.7 949 .16985 .70982 .7096 80 11.7 -71 ~ 

09G 46°50.7 1 116°10 

8 1 89.1 344.7 .74780 • 70660 .7059 65 -- -- z 

09H 46°53.5 1 116°08:5 1 

65.8 599 .31452 .70989 .7095 80 -- -- ""'"' 

09I 46°53.3 1 116°08.5 I 69.8 588 .34353 .70975 .7094 80 -- -116 ~ 09J 46°51.9 I 116°11.0 I 110.1 260 1.22370 .70688 .7058 65 -- -- p:: 

0 

46°29.7 1 116°27.7 1 b:i 10A 

108 46°28.2 I 116°25.2 I 33.5 

<1 229 

461 

<.00437 

.21017 .70383 

.70297 

.7034 

.7030 160 

160 -- -- ~ -- -- p:: 10C 46°30.0 1 116°23.2 23.4 577 .11729 .70357 .7033 160 -- -- 0 

100 46°30.1 I 116°21.1 I 2.5 896 .00807 .70345 .7034 160 -- -- r 

10E 46°29.7 1 116°16.4 1 ""'"' 17.23 905 .05506 .70379 .7037 160 -- -- 1-3 

:::c: 

jo--J. 

0 

co 
TABLE 6.1.-Results of Rb-Sr, c5180, and c5D analyses-Continued 1--l 

0 

~ 

North West Rb Sr 

Sample Latitude Longitude (ppm) (ppm) 87Rb/86Sr 87Sr/86Sr (87Sr/86Sr) Age 0180 (Ma) on 0 

818-lOF 46°30.8 1 116°17.9 1 23.7 756 0.09067 0.70411 o. 7039 160 

lOG 46°30.5 1 116°18,8 I 46.6 587 .22061 .70400 • 7035 160 

11A 46°36.5 1 116°10.8 1 44.4 811 • 15839 .70723 .7071 65 

11B 46°36.0 1 116°10,4 1 53.8 795 .19588 .71199 .7118 65 11. 1 11C 46°26,1 I 116°12.8 1 3.683 661 ,01611 .70342 .7034 160 6.2 -41 

110 46°24,7 I 116°11,7 I 12.8 682 ,05428 .70344 .7033 160 

11E 46°22,6 1 116°10.6 1 8.3 864 ,02778 .70349 .7034 160 

11F 46°19.4 1 116°08.4 1 25.7 629 ,11817 .70357 .7033 160 -- -56 11G 46°17,4 1 115°58,5 I 21,1 773 .07894 .70355 .7034 160 

12A 46°19.9 1 115°49.4 1 38.8 623 ,18023 .70987 .7095 130 

t-3 12B 46°19.4 1 115°49,2 1 56 624 .25974 .71091 .7104 130 -- -- ~:I: 

12C 46°20.6 1 M 115°43.8 1 76.6 833 .26613 .71039 .7099 65 -- -- 120 46°14.7 - 1 115°46.6 1 57.2 610 .27150 .71501 .7145 130 9.5 -52 12E 46°08.6 1 115°50,7 I 1.1 347 .00917 .70912 • 7091 130 7.0 -63 ~ 12F 46°08.3 1 ~:I: 115°47.8 1 30.9 616 • 14508 .70389 .7037 85 -- -- 0 

ttl 

13B 

13A 

45°35.1 

45°35,1 I 

I 115°40.8

115°40,8 1 I 

178 

61.7 

150 

859 ,20836 .71067 .7104 80 10,0 -72 ~ 3.43631 .71638 .7125 80 -- -- ~:I: 13C 45°31,1 I 115°40,1 I 107 451 .68682 .71330 .7125 80 10.8 -52 0 

t"" 130 45°37.3 1 115°46,4 1 124 278 1,29138 .71434 .7129 80 -- -- 13E 45°20,6 1 :::3 115°52,5 I 122 437 .80805 • 71158 .7107 80 10.7 -62 ~:I: 

13F 45°32.4 1 115°27,8 I 76.8 435 .51133 .71785 .7173 80 10,0 -62 ~ 14A 45°30,7 I 115°02,0 1 307 7.602 117.62028 .77563 .7043 43 9.7 -105 t1 

14B 45°30.5 1 115°00,9 1 231 216 3.09502 .71030 .7083 45 -- --

-t-3 14C 45°38,7 I 115°11,3 I 204 49.3 11.98380 .71749 .7090 45 -- -- r:n 15A 45°47.9 1 115°57.2 1 44.2 1093 ,11697 .70506 • 7049 80 -- -- ttl 

0 

~ 158 45°38.9 1 115°59.0 1 71.1 826 .24906 .70830 .7080 80 -- -71 t1 

15C 45°38.7 1 115°59,3 I 66.4 949 .20246 .70852 .7083 80 -- M 

150 -- ~ 45°37.8 1 116°04,5 I 3.055 1556 ,00568 .70402 .7040 80 -- -- N 15E 45°37,9 I 116°04.5 1 19.7 986 .05779 .70387 .7038 80 8.9 -67 0 

15F 45°38.2 1 116°06.3 1 28,1 144 .56444 .70488 .7042 80 -- -- z 

M 

16A 45°24,0 1 116°13,4 I .3 610 ,00142 • 70401 • 7040 160 

16B 45°24,0 1 116°12,6 I 11.5 546 ,06092 .70391 .7038 160 

16C 45°24 •. 

0 1 116°12.6 1 39.9 85.4 1. 35187 .70826 • 7033 260 

160 45°25,1 I 116°10.2 1 2 254 ,02277 .70389 .7039 80 6,4 -63 16E 45°25.6 1 116°08,1 I 21,2 857 .07155 .70411 • 7040 80 

16F 45°24.6 1 116°07,5 I 48.6 693 ,20290 .70750 .7073 80 

16G 45°24,4 1 116 °07,5 I 84 486 .50007 .70745 • 7069 80 

16H 45°24.2 1 116°05.5 1 142 52.1 7.90246 .72929 .7203 80 

16! 45°25.4 1 116 °02,5 I 65 644 • 29203 .70782 .7075 80 9.1 -62 16J 45°27.2 1 115°59.2 1 103 537 .55498 .70805 .7074 80 

16K 45°25.8 1 115°56,5 I 87.1 450 .56014 .70986 .7092 80 11.4 -70 
827-26A 46°15.7 1 117°36.8 I 9.466 752.9 0.03636 0.70324 0.7032 160 

268 46°15.7 1 117°36.8 1 6.948 709.9 .03636 .70317 .7032 160 

26C 46°35.4 I 116°43.4 I 39.78 394.4 .29216 .71939 • 7191 80 -- -54 

260 46°35.2 1 116°43.0 1 59.68 206.3 .84321 .78356 .7826 80 -- -- ~ 

27A 46°27.8 I 116°07.2 I 33.65 569.7 .17083 .70392 .7035 160 8.3 -- '"d 

t_%j 

278 46°28.7 1 116°06.0 1 17.39 662.8 .07588 .70381 .7036 160 -- -63 1-3 

~ 27C 46°29.2 I 116°05.1 I 55.47 657.1 .24425 .70794 • 7077 80 9.1 -- 0 

270 46°29.1 1 116°05.5 I 6.198 356.6 .05028 .70534 .7052 160 -- -- 0 

t_%j 

27E 46°29.1 I 116°05.5 I 5.828 1653 .01020 • 70560 .7056 160 -- -- z 

27F 46°28.9 1 116°06.0 1 2. 735 1111 .00712 .70398 • 7040 160 -- -- trj 

w 

...... 

28A 46°29.4 I 116°11.7 I 3. 732 542.3 .01991 .70486 .7048 160 -- -- ~w 

288 46°29.2 1 116°03.0 1 42.50 420.8 .29214 .70503 .7047 80 -- -- 0 

trj 

28C 46°29.2 I 116°14.7 I 35.12 635.9 .15974 .70431 .7039 160 -- -- 0 

280 46°28.3 1 116°14.8 1 .6681 280.1 .00690 .70316 .7031 160 -- -- Cl 

::r: 28E 46°23.2 I 116°10.4 I 2.755 763.7 .01043 .70337 .7033 160 6.4 -- ~ 

0 

28F 46°17.2 1 116°06.8 1 32.26 467.2 • 19971 .70376 .7033 160 -- -- z 

0 28G 46°11.5 I 116°00.5 I 42.76 535.4 .23099 .70390 .7034 160 -- -- t"'4 

29A 45°38.1 I 116°14.2 1 15.91 504.5 .09121 .70390 .7037 160 -- -- 0 

298 45°38.0 1 116°13.8 1 .188 104.7 .00519 .70391 .7039 160 -- -- ~ 29C 45°38.2 1 116°08.6 1 16.88 691.6 .07059 .70387 .7037 160 -- -- > z 

290 45°38.4 1 116°07.3 1 20.27 459." .12768 .70455 .7044 80 -- -- t:::::l 

29E 45°38.0 1 116°00.7 1 52.60 1013 .15024 .70829 .7081 80 -- -- ::r: 

29F 45°37 • 7 I 116°00.9 1 59.72 916.1 .18863 .70864 .7084 80 -- -- ~ 

t:::::l 29G 45°26.9 1 116°10.6 1 9.618 127.4 .21836 .70457 .7043 80 -- -- ~ 

29H 45°25.4 I 116°09.4 I 46.65 950.4 .14199 .70567 .7055 80 -- -- ~ ::r: 

29I 45°25.7 1 116°08.6 1 14.25 682.4 .06040 .70417 .7041 80 -- -- trj 

29J 45°25.7 1 116°08.6 I 25.96 379.9 • 19766 .70530 .7051 80 ~ -- -- a= 29K 45°25.4 1 116°07.8 1 42.31 919.7 .13310 .70726 .7071 80 -- -- > 

29L 45°25.5 I 116°07.8 1 51.66 932.5 .16024 .70477 .7046 80 -- -- t"'4 

30A 45°10.0 1 116°18.0 1 6.905 1078 .01853 .70393 .7039 80 -- -- w 

~ 

w 308 45°07.8 I 116°17.1 I 20.94 911.1 .06647 .70372 .7036 80 1-3 -- -- trj 

30C 44°57.5 1 116°12.5 1 11.32 475.2 .06890 .70403 • 7040 80 -- -- a= 

300 44°57.5 1 116°12.5 I 34.08 705.5 .13072 .70417 .7040 80 -- -- w 

30E 44°57.5 1 116°12.5 1 9.448 688.2 .03971 .70395 .7039 80 -- -- 0 

"'%j 30F 44°58.0 I 116°11.1 I 63.56 764.8 .24038 .70461 .7043 80 -- -- 1-3 

44°58.0 1 116°10.9 1 ::r: 

30G 37.49 984.9 .11011 • 70587 .7057 80 -- -- trj 

30H 44°57.8 I 116°10.4 I 33.88 890.5 .11001: .70545 .7053 80 -- -- z 

30I 44°55.0 1 116°00.2 1 168.7 310.9 1. 57048 .71108 .7093 80 -- -- 0 

30J 44°55.6 1 116°01.2 I 187.2 139.2 3.89342 .71404 • 7096 80 -- -- ~ 30K 44°58.1 1 116°03.2 1 65.52 727.3 .26069 .70923 .7089 80 -- -- ::r: 

t_%j 30L 44°57 0 7 I 116°05.5 I 114.5 473.4 .69991 .70927 .7085 80 -- -- ~ 

z 

837-23A 46°35.2 1 116°48.5 1 287.5 690.7 1. 20498 .71337 .7123 80 -- -- ...... 

238 46°35.5 1 116°46.4 I 25.46 923.6 .07976 .70772 .7076 80 -- -- ~ 23C 46°35.3 1 116°18.3 1 92 621.2 • 42853 .70835 .7079 80 -- -- ::r: 

24A 46°30.7 I 116°17.4 I 30.46 698.1 .12619 .70363 .7035 80 -- -- 0 

248 46°30.6 1 116°16.7 1 45.22 433.8 .30149 .70380 .7015 80 -- -- Oj 

~ ::r: 

0 

t"'4 

...... 

1-3 

::r: 

~ 

0 

c.n 
TABLE 6.1.-Results of Rb-Sr, 

6180, and oD a,nalyses-Continued ~ 

0 

~ 

North West Rb Sr 

87Rb/86Sr 87Sr/86Sr (87Sr/86Sr) 0180 Sample Latitude Longitude (ppm) (ppm) Age (Ma) oD 0 

837-24C 46°31.7 1 116°18.4 1 93.6 663.9 0.40796 0.70888 0.7084 80 

240 46°30.6 1 116°16.0 1 3.04 743.2 .01183 .70487 .7049 80 

24E 46°33.3 1 116°15.1 I 37.68 966.9 .11274 .70654 .7064 80 

24F 46°29.8 I 116°15.4 1 32.73 113.2 .83640 .70578 .7048 80 

25A 46°21.9 1 116°05.6 1 3.638 1284 .00819 .70356 .7035 80 

25B 46°21.9 1 116°04.7' 11,19 867.1 • 03733 .70407 .7040 80 7.4 

25C 46°17.8 1 115°58.5 1 14.39 903.7 • 04606 .70423 .7042 80 

25D 46°18.5 1 116°02.8 1 20.65 557 .10723 .70467 .7046 80 

25E 46°16.3 1 115°58.9 1 15.05 803.5 .05417 .70345 .7034 80 

25F 46°14.5 1 115°59.9 1 12.42 453.5 .07921 .70335 .7033 80 

46°15.1 I 115°57 • 6 I 

>-3 

25G 2.537 610.2 .01202 .70356 .7035 80 -- -- ::q 

26A 46°30.7. 116°19.0 I 36.49 584.1 .18068 • 70378 .7036 80 -- -- t?=:l 

26B 46°33.0 1 116°17.7 1 """'" .5 248.7 .00582 .70546 .7054 80 -- -- 26C 46°32.9 1 116°18.4 1 31.77 580.9 .15819 .70484 .7047 80 10.3 -- ~ 260 46°32.8 1 116°18.4 1 44.48 638.2 .20160 .70504 .7048 M -- ::q -- 0 

46°34.9 1 115°45.3 1 ttl 27A 30.59 728.3 .12155 .70952 .7094 65 -- -79 ~ 27B 46°34.9 1 115°45.3 1 20.77 741.7 .08104 .70940 .7093 65 -- -- ::q 

27C 46°34.5 1 115°39.1 I 59.55 832.2 .20707 .70911 .7089 65 -- -- 0 

270 46°35.4 1 115°35.3 I 58.32 744.4 .22671 .70930 .7091 65 ,..._ -148 t'"-1 

27E 46°36.9 I 115°30.3 I 272.9 64.53 12.24150 .71235 .7045 45 -- -131 ~ ::q 

27F 46°42.9 1 115°33.6 1 222.8 22.45 28.75622 .72272 .7043 45 -111 > - z 

27G 46°32.3 1 115°37.9 1 60.85 595.8 .29563 .71225 .7120 65 -- -79 t::1 

27H 46°27.2 1 115°36.3 I 64.05 1091 • 16994 .71268 .7125 60 -- -66 """'" >-3 

271 46°25.2 1 115°24.9 1 64.29 866.6 .21468 .70961 .7094 60 -- -124 rn 

27J 46°27 • 1 I 115°14.7 1 100.6 64.91 4.48600 .71187 .7087 50 -- -102 ttl 

0 

46°29.6 1 115°09.1 I 

~ 

27K 170 202.3 2.43158 .70866 .7069 50 -- -- t::1 

27L 46°30.5' 115°05.7 I 175.1 172.8 2.93218 .70896 .7069 50 -- -132 t:z:l 

27M 46°28.4' 115°05.3 1 123.1 397.9 .89503 .70669 • 7060 50 -- -- ~ 

N 28A 45°57.2 1 115°55.8 1 148.4 271.3 1.58308 .71063 .7088 80 -- -- 0 

28B 45°58.8 I 115°53.0 1 94.47 574.5 .47578 .70796 .7074 80 -- -- z 

t:z:l 

28C 45°58.8 1 115°52.4 I 72.36 778.7 .26887 .70832 .7080 80 

280 45°58.7 I 115°51.5 I 139.3 343.7 1.17351 .71527 .7139 80 

28E 45°58.5 1 115°52.0 1 115.3 498.9 .66886 .71056 .7098 80 

28F 45°57.8 1 115°54.6 1 79.88 702.3 .32906 .70694 .7066 80 

28G 45°49.5 1 115°53.2 1 181.5 547.8 ,96020 .72467 .7236 80 

28H 45°49.5 I 115°53.4 I 96.95 723.8 • 38760 .70904 .7086 80 

29A 45°24.1 1 116°06.8 1 103.2 365.7 .81662 .70937 .7084 80 

29B 45°24.6' 116°04.7 I 16.68 50.37 .95831 .70976 .7097 80 

29C 45°25.6 1 116°00.4 1 179.4 225.3 2.30614 .71779 .7152 80 

290 45°25.5 I 116°01.5 I 145.6 544.8 .77337 .70935 .7085 80 

29E 45°24.1 1 116°08.6 1 27.26 1056 .07466 .70398 .7039 80 

29F 45°24.1 I 116°09.1 I 37.89 890.8 .12302 .70408 .7039 80 

29G 45°17.5 1 116°14.6 1 26.15 603.4 • 12534 .70395 .7038 80 

29H 45°17.5 I 116°14.6 1 33.24 633.9 .15166 .70401 .7038 80 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 107 

muscovite pairs from the Atlanta lobe (Criss and Taylor, 

1983). However, Masi and others (1981) discovered a relation between dD values and water contents in the Sierra 

Nevada batholith and proposed that this trend is primary. 

If the latter is true, granitic rocks having low water contents may have primary dD values as low as -110 per 

mil. Such low primary values can be excluded in most of 

the Idaho batholith, but may apply in certain cases, such 

as in the Eocene epizonal plutons. Unfortunately, the intimate association of these epizonal plutons with the 

meteoric-hydrothermal systems will complicate resolution 

of this possibility. 

RELATION OF R1 AND 6180 

Values of ri and d18Q are governed by fundamentally 

different processes, and graphs of these two parameters 

have provided much insight into the petrogenesis of igneous rocks (for example, Magaritz and others, 1978; 

Taylor, 1981). Taylor and Silver (1978) constructed a 

graph of ri versus d18Q for the Peninsular Ranges batholith (PRB) and discovered that most data points comprise 

a linear band extending from mantlelike igneous rocks 

(fig. 6. 7). They suggested that the source materials for 

the PRB were dominated by two end members, one 

derived from upper mantle materials and the other from 

high-180, high 87Srf86Sr materials probably related to 

sedimentary or metasedimentary rocks. Evidence for a 

third end member was found in the rocks of the San 

Jacinto block, which have high ri values relative to the 

PRB trend (fig. 6.7); Taylor and Silver (1978) suggested 

that involvement of ancient cratonal rocks may be implicated. In contrast, Masi and others (1981) found that 

1160 1150 1140 47°r----n--------------------~~------------------~~---------------,----,-------~ 

\ 

( 0 

', 

/ 

(_ 

WASHINGTON\ 

46° ----------\ 

OREGON ·•

1 

•. 

EXPLANATION ·---Q ·. 

( 87Sr/86Sr)o 

I 

\ 

• >0. 712 ) a o.7o8-0.712 :" 

() o. 7065-0.708 ( G 0.704-0.7065 

0 <0.704 ( 

/ "1 

( i 0 10 Ml LES '( 1---r-'-r-----' I 

0 10 KILOMETERS I "'· 

) 

f 

{ 

: 

45° ( 

FIGURE 6.2.-Contour map showing the distribution of initial 87Sr/ 86Sr (ri) for the Jurassic to Eocene plutons of north-central Idaho. Note the 

well-defined 0.704 and 0.708 contours that separate the high-ri, Precambrian-hosted plutons of the Belt-Yellowjacket terrane on the east 

from the low-ri plutons of the accreted Wallowa-Seven Devils terrane (compare with figure 6.1). Patterl)ed areas same as in figure 6.1. 
108 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

only about half the data points for Sierra Nevada plutons 

conform to the PRB trend. The remaining analyses also 

have high ri values relative to the PRB trend, similar to 

the San Jacinto block as well as to most Mojave Desert 

granitoids (fig. 6. 7). Masi and others (1981) also showed 

that Klamath Mountains plutons are distinguished by ri 

values that are uniform and low (~0. 705), independent of 

the o18Q values. 

The Idaho batholith is unique among the major cordilleran batholiths in that very few plutons have ri and alSO 

values that conform to the trend of PRB granitoids (fig. 

6.8). Practically all Bitterroot lobe samples, including both 

the Cretaceous-early Tertiary main-phase plutons and the 

Eocene epizonal plutons, have high ri values relative to 

the PRB trend. These plutons are most similar to 

117° 

47° 

I 

I 

i i() 

j i~oscow 

.u 

i~ 

I 

( 

\~ 

\ 

( 

,f 

c_' 

WASHINGTON\ 

46° --------··\ 

OREGON \. 

'· ·,__ 

EXPLANATION '·. 

l 

Rb/Sr \ 

) • >1.0 a o.25-1.o 

() 0.1-0.25 

Q 0.05-0.1 

0 <0.05 

,..-' 

~0 MILES ( 

0 10 KILOMETERS / 

i 

) 

/ 

./ 

( 

( 

116° 

Q 

() 

I 

~ 

granitoids of the San Jacinto block, the Mojave Desert, 

and other areas that are known or inferred to contain 

Precambrian rocks. This observation could possibly indicate that mantle-derived plutons have assimilated a 

great deal of ancient crystalline rocks in all these areas; 

such a model was proposed by Armstrong and others 

(1977) for the radiogenic rocks of the Idaho batholith. 

Although Fleck and Criss (1985) demonstrated that 

assimilation of upper crustal Precambrian wall rock was 

important in the Bitterroot lobe magmas, the mixing 

trends for the main-phase granitoids indicate that the 

magmatic end member has ri values (about 0.708) and 

Rb/Sr ratios that are much too high to characterize 

materials derived from depleted mantle. Fleck and Criss 

(1985) proposed that the main-phase magmas primarily 

115° 114° 

45°~----~~r_·------------L-----~----~~--------------~~~~----------------~L-----~ 

FIGURE 6.3.-Contour map showing the distribution of Rb/Sr concentration (weight) ratios for the plutons of north-central Idaho. The Jurassic 

and Cretaceous plutons in the Wallowa-Seven Devils terrane typically have Rb/Sr ratios ~0.05, with most values being <0.08, whereas those 

for the Eocene epizonal plutons ( + pattern) are typically> 1. Ratios for the main-phase plutons of the Bitterroot and Atlanta lobes are intermediate, ranging from 0.06 to 0.24 in most samples. 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 109 

TABLE 6.2.-Chemical analyses of granitic plutons from the northern Idaho batholith 

[Rapid rock analyses by Z.A. Hamlin and Nancy Skinner, U.S. Geological Survey] 

Sample-- 7a7-1aA 7a7-1a8 1a1-1ac 7a 7-1ao 7a7-1aE 7a7-1aF 1a1-1aG 7a7-1aH 7a7-1a1 7a7-1aJ 

Si02---- 71 .1 71.0 72.3 73.3 71.3 71.4 71.4 69.1 ~7 .4 75.2 

A1 2o3--- 15 .o 15.4 14.9 14.4 15 .1 15 .2 15 .o 15.2 16.6 12.7 

Fe2o3--- .96 .a1 .97 .52 .64 .67 .sa .ao 1.0 .76 Feo----- .96 .a4 .68 .64 .64 .76 .96 1.6 1.5 .72 

MgO----- .43 .39 .36 .30 .31 .36 • 75 1.1 .62 .17 cao----- 2.7 2.2 2.1 2.3 2.3 2.4 2.8 3.3 3.8 1.1 Na2o---- 4.0 4.3 4.5 4.6 4.4 4.4 4.1 4.1 4.6 4.0 

K2o;---- 3.6 4.2 3.9 3.6 4.1 3.7 3.0 2.5 2.6 4.a 

HzO_ ---- .43 .4a .43 .34 .40 .s1 • 79 .aa .54 .41 

H2o ---- .10 .08 .10 .09 .06 .02 .09 .10 .12 .10 

Tl.02---- .27 .22 .13 .12 .18 .22 .27 .Ja .29 .1a 

P2os---- .14 .09 .11 .os .15 .09 .11 .15 .1a .06 Mno----- .03 .03 .02 .02 .01 .02 .03 .03 .os .04 

co2----- ____:.Ql __!.!!.!. __!.!!.!. _:.E _:.E __.!21. __!.!!.!. __!.!!.!. _:.E _!2.l. 

Total- 100 100 100 100 100 100 100 99 99 100 

Sample-- 787-19C 7a7-19E 7a7-19G 7a7-20A 7a7-208* 7a7-20C 7a7-200 787-20E 787-21A 7a7-218 

Sio2---- 62.1 57.6 12.7 63.7 73.0 70.9 65.8 64.8 68.7 76.3 

A1 2o3--- 16.2 17.3 13.7 16.1 11.0 15 .1 17.1 17.0 15 .2 12 .s 

Fe2o3--- 1.6 2.5 .97 2.8 .49 .42 1.1 1.6 .as .64 

Feo----- 3.4 s.o .56 3.5 .sa 1.3 2.2 2.6 1.7 .40 

MgO----- 2.6 3.1 .25 2.0 4.2 .65 1.5 1.7 .96 .01 

Cao----- 5.4 6.2 1.5 4.4 4.8 1.3 4.1 4.3 2.5 .29 

Na2o---- 3.6 3.7 3.1 4.0 3.7 2.8 4.2 4.0 3.8 3.6 

K2o;---- 2.2 2.3 5.4 2.1 .52 5.4 2.0 2.0 4.0 5.1 

HzO_ ---- 1.1 1 .1 .52 .94 .36 .87 • 74 .87 .so .64 

~0 ---- .17 .os .16 .12 .04 .10 .oa .os .07 .18 

T1.02---- .70 .96 .16 .79 .so .26 .54 .64 .55 .09 

P205---- .23 .30 .16 .40 .17 .17 .29 .27 .24 .10 

MnO----- .os .09 .01 .os .02 .02 .os .07 .04 .01 

co2 _____ _:.E ~ __!.!!.!. ~ _!2.l. ~ ~ __!.9.!. ~ ~ 

Total- 99 100 99 101 100 99 100 100 99 100 

Sample-- 787-21C 787-210 787-21E 787-21F 787-24A 787-248 787-24C 787-240 787-24E 787-24F 

Sio2---- 75 .s 73.8 68.2 69.3 70.9 70.1 69.7 72.1 71.9 71.5 

A1 2o3--- 12.4 13.2 17.1 16.1 15 .s 15.9 16.2 14.9 15.1 15.1 

Fe2o3--- • 78 0.45 .91 .98 .6a .71 • 79 0.56 .56 .69 

Feo----- .64 1.1 .92 1.0 .84 .a4 .96 0.84 .68 .68 

Mgo----- .03 .27 .45 .47 .40 .43 .43 0.31 .28 .26 

Cao----- .59 1.2 2.9 2. 7 2.4 2.9 2.6 2.4 2.5 2.5 

Na2o---- J.a 3.6 5.1 4.6 4.5 4.9 4.6 4.3 4.5 4.3 

K2o;---- 4.9 4.7 3.0 3.8 3.8 3.1 4.0 3.8 J.a 4.3 

HzO_ ---- .51 .ss .61 .52 .sa .62 .52 .52 .55 .45 

H

2

o ---- .09 .09 .03 .07 .os .14 .10 .11 .11 .11 

T102---- .19 .23 .29 .26 .23 .23 0.26 .21 .20 .22 

P2os---- .10 .14 .14 .18 .16 .20 0.18 .15 .12 .15 

MnO----- .02 .02 .02 .03 .01 .02 0.02 .01 .02 .01 

co2----- _:.E _:.E _:.E __!.!!.!. __!.!!.!. __:..Q.!. ~ __:..Q.!. __!.!!.!. ~ 

Total- 100 99 100 100 100 100 100 100 100 100 

Sample-- 787-24G 787-241 787-25A 7a7-25C 787-26A 787-268 787-26C 7a7-260 787-26E 7a7-26F 

Sio2---- 54.9 70.8 72.1 66.5 67 .o 67.6 70.9 66.1 66.9 64.0 

A1 2o3--- 17 .o 15.2 14.1 16.9 15.3 14 .s 14.9 17.4 15.4 16.0 

Fe2o3--- 1. 7 • 74 1.6 .77 1.4 1.9 1.2 .68 1.5 1.6 

Feo----- 6.7 1.3 .92 1.8 2.3 2.5 1.9 2.2 2.2 3.7 

MgO----- 3.4 .92 .11 1.4 1.4 1.6 0.97 1.5 1.4 2.7 

Cao----- 7.1 2. 7 1.5 3.5 3.8 4.4 4.5 4.7 3.8 4.9 

Na2o---- 3.7 4.3 4.0 4.8 4.1 4.2 4.0 4.a 4.0 3.5 

K2o;---- 2.3 3.6 5.7 3.4 3.9 2.4 1.3 1.5 4.1 2.3 

HzO_ ---- 1.2 .51 .61 .68 .66 .77 • 75 .66 .61 1.0 

H2o ---- .14 .07 .14 .10 .09 .10 .10 .14 .09 .09 

T102---- 2.4 .31 .27 .46 .53 .63 .44 .ss .45 .82 

P205---- .45 .15 .12 .19 .26 .29 .23 .26 .22 .32 

MnO----- .09 .02 .os .03 .06 .09 .02 .03 .06 .06 

C02----- ~ __:Q!. _:.E __:..Q.!. ~ _:.E ---!.!1. __!.9.!. __.:& ~ 

Total- 101 101 101 100 101 101 101 100 101 101 

*Precambrian ( ?) metasedimentary rock. 
110 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

originated from melting of ancient continental basement. 

Available data also indicate a significant Precambrian 

component in the Eocene plutons (Armstrong, 1974; 

Chase and others, 1978; Fleck and Criss, 1985), but do 

not completely deny the existence of a mantle component. 

12 

1-

z 

UJ 

() 

a: 10 UJ 

a.. 

1-

J: 

CJ 8 iii 

s: 

~ 

~ 6 

~ 

+ 

q. 4 

nl 

z 

0 2 

0 

nl 

() 

0 

44 48 52 60 64 68 72 76 80 

Si0 2 , IN WEIGHT PERCENT 

FIGURE 6.4.-Variation diagram of Mesozoic and Tertiary plutons of 

north-central Idaho showing total alkalis (Na20 + K20, solid symbols) and CaO (open symbols) plotted against Si02• The calcic trend 

of the low-ri WSD plutons (triangles) contrasts with the calcic to 

calc-alkaline trend of the Precambrian-hosted plutons in the Bitterroot lobe, including both the main-phase (circles) and Eocene epizonal 

plutons (squares). Analyses are from Lindgren (1904), Hietanen 

(1962, 1963), Larsen and Schmidt (1958), Ross (1934), and table 6.2. 

12 1-

z 

UJ 

() 

ffi 10 

a.. 

1-

J: 

Q 8 UJ 

s: 

z 

-_ 6 

q. 

~ 

+ 4 

q. 

nl 

z 

.... 2 0 

0 

nl 

() 0 

44 48 52 56 60 64 68 72 76 80 

Si0 2 , IN WEIGHT PERCENT 

FIGURE 6.5.-Variation diagram showing that Eocene epizonal granitoids conform to the calc-alkaline trend defined by the Challis 

Volcanics. Volcanic-rock analyses (circles) are from Ross (1962) and 

Mcintyre and others (1982); granitoid analyses (squares) are from 

the same references listed in figure 6.4. Solid symbols, Na20 + 

K20; open symbols, CaO. 

In fact, some values near 0.704 (table 6.1) suggest only 

minor contamination by either upper or lower sialic crust. 

The ri and d18Q relations of the WSD samples are 

distinctly different from those of the Bitterroot lobe (fig. 

6.8). The data points for these samples generally fall 

within the PRB band, but their flat trend is most similar 

to that of the Klamath Mountains granitoids. The relatively high d18Q values of some samples (as high as + 10.5 per 

mil; sample 807-27 A) may reflect either incorporation of 

or exchange with high-18Q materials, or low-temperature 

alteration, but the uniformly low ri values preclude 

significant participation of ancient sialic material. The 

isotopic relations strongly suggest that the WSD plutons 

were derived from island-arc or other mantle components. 

-40 EXPLANATION 

• EOCENE EPIZONAL 

e MAIN PHASE 

-50 & WSD 

X PRECAMBRIAN 

-60 

-70 

-80 

...J 

~ 

a:: -90 w 

a. 

z 

c:i -100 

...0 

-110 

-120 

-130 • • 

-140 • • 

-150 

• • 

• 

• • • -

• • 

• '·. • 

•• 

•• 

o 18Q, IN PER MIL 

FIGURE 6.6.-Relation of o180 to oD for plutonic rocks of north-central 

Idaho. Rocks with the highest oD values are unaltered; these are 

low-ri plutons in the Wallowa-Seven Devils terrane and main-phase 

plutons in the Bitterroot lobe that are remote from Eocene epizonal 

plutons. Because of interactions with meteoric-hydrothermal fluids, 

the epizonal plutons and their adjacent country rocks are significantly depleted in oD. Where oD is most depleted (oD~ -120 per 

mil) the o180 values are lowered as well. 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 111 

CRETACEOUS-EARLY TERTIARY AND EOCENE 

HYDROTHERMAL SYSTEMS 

REGIONAL VARIATIONS OF dD AND d180 VALUES 

Geographic variations of dD and o180 values provide 

important information on the location and geologic 

characteristics of hydrothermal systems (for example, 

Taylor, 1971, 1977; Taylor and Forester, 1979). Taylor 

and Magaritz (1978) made a reconnaissance map of dD 

values of the Idaho batholith and demonstrated the 

regional extent of Eocene meteoric-hydrothermal activity. 

Criss and Taylor (1983) used d180 data to map several 

meteoric-hydrothermal systems in the southern Idaho 

batholith, and Criss and others (1984, 1985) extended this 

method to the Challis volcanic field. As discussed below, 

shallow-level, Eocene meteoric-hydrothermal systems as 

well as deep-level, Cretaceous-early Tertiary metamorphic-hydrothermal systems operated in north-central 

Idaho. 

A contour map of whole-rock dD values illustrates 

strong geographic variations (fig. 6;9). All WSD plutons, 

as well as the main-phase granitoids remote from Eocene 

epizonal plutons, have isotopically normal dD values of 

~ - 82 per mil. The primary dD character of these plutons 

0.709 

0.708 

0 0.707 

-lo.. 

(J) 

(0 

~ 0.706 lo.. 

(J) 

.... 

Q) 

- 0.705 

0.704 

0.703 

0.702 

5 6 7 8 9 10 11 12 13 

0 18Q, IN PER MIL 

FIGURE 6.7.-Generalized graph of initial s7Sr/s6Sr (ri) versus d1so 

delineating granitoids of the Peninsular Ranges batholith (PRB), 

the San Jacinto block (SJ), the Klamath Mountains (KM), the Mojave 

Desert (MD), and the Sierra Nevada batholith (SN) from Taylor and 

Silver (1978) and Masi and others (1981). The San Jacinto, Mojave 

Desert, and many Sierra Nevada plutons occur in regions underlain 

by Precambrian wall rock and tend to have the highest ri values 

at a given a1so content. 

is further supported because the dD contours make the 

same right-angle bend as do the 0. 704 and 0. 708 initial-Sr 

contours of figure 6.2. Another zone of normal dD values 

occurs in the east-central Bitterroot lobe, immediately 

west of Hamilton (fig. 6.9). Included in this zone are 

several mylonized samples, indicating that deep-level 

penetration of meteoric-hydrothermal fluid was not essential to the development of the mylonite zone. In contrast, 

extremely low dD values (~ -120) occur within several 

zones in the Bitterroot lobe, together constituting more 

than 7,000 km2. All low values occur within or near the 

Eocene epizonal plutons and are indicative of meteorichydrothermal systems established during their emplacement, as originally pointed out by Taylor and Magaritz 

(1978). The most depleted samples (<-140 per mil) occur 

along the perimeter of these epizonal plutons, which 

is where the most intense fluid circulation would be 

expected. 

0.720 

0.718 

0.716 • 

BITTERROOT 

• 0. 714 • • • • 

~ 0.712 

en 

"' 

• • • 

• LOBE • 10 • • ....... 

.... 

~ 0. 710 

~ 

• • -:. 

• 

I • 

• • • ·- 818-12E ~ . • 0.708 

0.706 

0.704 

0.702 5 6 7 8 9 10 11 12 

o 18Q, IN PER MIL 

FIGURE 6.8.-Graph of ri versus a1so for the plutons of north-central 

Idaho. Only essentially unaltered rocks with aD;;?; -120 per mil are 

shown. Note the significant differences between the low-ri plutons 

in the Wallowa-Seven Devils terrane (WSD), the high-ri plutons in 

the Bitterroot lobe (including several samples from the northern 

Atlanta lobe), and the Peninsular Ranges batholith (PRB) trend from 

figure 6. 7. These differences are explicable in terms of the source 

rocks from which the plutons were derived. Transition-zone samples 

collected along detailed traverses across the suture-zone boundary 

(0. 704-0.708 line) have been discussed in detail by Fleck and Criss 

(1985) and are not included here. 
112 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

Variations of o180 values in the plutons of north-central 

Idaho are illustrated in figure 6.10. With the exception 

of a few WSD plutons whose rather low a18Q values are 

probably primary, all a18Q values less than + 8.0 per mil 

occur within the zones of extreme deuterium depletion 

in the Bitterroot lobe and are related to the meteorichydrothermal systems. High water/rock ratios of unity 

or greater are required to produce the most extreme a18Q 

values (as low as -0.3 per mil). As discussed in the subsection "o18Q and dD data," the high water-rock ratios that 

are required to produce significant 18Q/ 16Q changes in 

rocks demand complete reequilibration of D/H ratios, thus 

accounting for the observed spatial correspondence 

between low a18Q and low dD values. 

117° 

47° 

1 

I 

! 

jQ -ao j ;~oscow 

.tl 

--J----60 

! 0 

i 

' ( 

'\ ~ 

\ 

\ 

·-/ 

< 

WASHINGTON\ 

46° ---·--·-·-\ 

OREGON ·•

1 

•. 

EXPLANATION ·,___ 

' 

60, IN PER MIL l 

e -140 to -160 \ 

a -120 to -140 ) 

() -100 to -120 I 

.• 

~ -80 to -100 ( 0 -60 to -80 

0 -40 to -60 ( 

/ 0 10 MILES ( f----r-1-r----' 

0 10 KILOMETERS / 

j 

j 

I 

{ 

: 

45° ( 

FIGURE 6.9.-Contour map of whole-rock dD values for the plutons of 

north-central Idaho. Relatively high(~- 82 per mil) dD values of the 

plutons in the Wallowa-Seven Devils terrane are thought to be primary; 

note that the contours make a similar right-angle bend as do the 0. 704 

and 0.708 initial strontium contours shown in figure 6.2. High dD 

The zones of 18Q depletion in the plutons of the Bitterroot lobe are much smaller than the meteoric-hydrothermal systems mapped by Criss and Taylor (1983) in 

the southern Idaho batholith, many of which encompass 

several thousand square kilometers. Furthermore, the 

most extreme o180 values in the Bitterroot lobe are not 

as low as some of those observed farther south, several 

of which are -4.0 or less (Criss and Taylor, 1983). 

Petrographic examination of thin sections reveals that the 

propylitic alteration that generally attends moderatetemperature water-rock interactions is not as widespread 

or intense in the Bitterroot lobe as that observed in the 

southern Idaho batholith. These observations cannot 

simply be explained by a reduced volume of Eocene 

t.--'"'. 

( 

!...--

) 

\ ': 

( 

\ 

'"·. 

values also occur in the Bitterroot lobe at positions remote from the 

Eocene epizonal plutons (pattern). In contrast, marked dD depletions 

(~ -120 per mil) caused by meteoric-hydrothermal alteration processes 

occur within and near the epizonal intrusions. 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 113 

epizonal plutons in the north. In fact, recent mapping by 

Toth and others (1983) shows that the epizonal plutons 

are at least as areally abundant in the Bitterroot lobe as 

in the Atlanta lobe. 

On the other hand, the age and temperature contrasts 

between host rocks and epizonal plutons in the Bitterroot 

lobe were almost certainly less than those of the Atlanta 

lobe. McDowell and Kulp (1969), Armstrong and others 

(1977), and Criss and others (1982) report biotite K-Ar 

ages of 70 to 60 Ma for most of the main-phase plutons 

of the Atlanta lobe, with some ages older than 80 Ma. 

These authors argue that much of the Atlanta lobe is 

Cretaceous. In contrast, K-Ar ages of main-phase plutons 

in the Bitterroot lobe where intruded by the belt of 

epizonal granites are only slightly older than or equal to 

those of the granites (62 to 45 Ma as compared to 51 to 

44 Ma). Whereas temperature anomalies associated with 

the older plutonic rocks in the Atlanta lobe would have 

diminished to the geothermal gradient before emplacement of the Eocene plutons, host-rock temperatures for 

these epizonal plutons in the Bitterroot lobe would have 

been significantly higher. We suggest that smaller lateral 

temperature gradients and lower permeabilities in the 

Bitterroot lobe resulted in less intense fluid circulation 

than occurred in the Atlanta lobe. 

Evidence for earlier, deep-level metamorphic-hydro1170 

47or----:r---------------------~11~6---------------------~1~1~5o-----------------,----1 1_4° ______ __, 

I 

I 

i 

i() l t~oscow 

:n 

' 1 i----

1--

\ 

( 

.r 

<' 

WASHINGTON\ 

46° ---·----·-\ 

OREGON \. 

'· 

EXPLANATION 

olBQ, IN PER MIL 

• >11.0 

~ 10-11 

() 9-10 

Q 8-9 

0 7-8 

0 <7.0 

~0 MILES 1 

0 10 KILOMETERS 1 

i 

j 

/ 

./ 

'·. I 

\ 

) 

,,~~' 

( 

( 

•. / 45°~----~--L-----------~--~--~----~~L_~~~-------L~~~L-----~~~----~~------~ 

FIGURE 6.10.-Contour map of whole-rock d180 values for the plutons 

of north-central Idaho. Although the d180 values for some of the 

plutons in the Wallowa-Seven Devils terrane are rather low (as low 

as + 5.5 per mil), these values are thought to be primary, partly because 

the dD values for the rocks are high. Note that the d180 contours in 

this zone make a similar right-angle bend as do the 0. 704 and 0. 708 

initial strontium contours (fig. 6.2) and the -80 per mil dD contour 

(fig. 6.9). The d180 values from the Bitterroot lobe tend to be high 

(;;?; + 8 per mil) except in the zones of intense meteoric-hydrothermal 

alteration (hachured contours) within and adjacent to the Eocene 

epizonal plutons (pattern). Only representative samples are shown 

along a closely sampled traverse near 45°50' N., 116° W. (see Fleck 

and Criss, 1985). 
114 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

thermal activity is provided by d180 analyses of Belt 

Supergroup (Wallace Formation) pelitic metasedimentary 

rocks (table 6.3) collected around the north margin of the 

Bitterroot lobe. The d180 values progressively decrease 

toward the Idaho batholith (fig. 6.11), in the general direction of increasing metamorphic grade (Hietanen, 1962, 

1984). Thus, the whole-rock d180 values decrease from 

about + 15 ± 1 per mil in the low-grade metamorphic zones 

at large distances from the batholith outcrop (>60 km) to 

generally + 11.5 ± 1 per mil in the highest grade, sillimanite ± kyanite zones adjacent to the batholith (fig. 6.11; 

table 6.3). The latter values are similar to those typical 

of the Idaho batholith ( + 10 ± 1.5 per mil), with plutons 

near metasedimentary wall rock generally having the 

highest c5180 values. 

Hietanen (1984) and Lang and Rice (1985) recognize 

both synkinematic and postkinematic metamorphism in 

the region north of the Bitterroot lobe, with the later, 

higher temperature metamorphism and associated metasomatic effects being related to batholith intrusion. It is 

probable that the 180 reductions in the metasedimentary 

rocks were principally produced during this later metamorphism, apparently in Cretaceous-early Tertiary time. 

Note that Fleck and Criss (1985) observed a similar 

approach to 180 homogenization between plutons and 

adjacent wall rock along the west side of the Idaho batholith and argued that most of the indicated 18Q exchange 

occurred during emplacement of the Late Cretaceous 

plutons. 

The large scale of 180 exchange observed in the 

Wallace Formation metasedimentary rocks clearly indicates fluid migration rather than diffusion as the dominant transport process. The observations are consistent 

with convective circulation of high-180 fluid, probably 

derived from metamorphic dehydration, connate sources, 

and magmatic water, mostly during batholith emplacement and associated postkinematic regional metamorphism. General correlations between c5180 values and 

metamorphic grade have been observed in other areas by 

Garlick and Epstein (1967) and Dontsova (1970); 180 exchange between igneous plutons and metasedimentary 

wall rock has been documented by Shieh and Taylor (1969) 

and Turi and Taylor (1971); and large-scale homogenization of oxygen isotopes among various rocks of high-grade 

metamorphic terranes has been reported by Taylor (1970), 

Shieh and Schwarcz (1974), and Hoernes and Friedrichsen 

(1980). The oxygen isotopic exchange was probably effected by high-temperature, high-180 aqueous fluid in all 

of these cases. 

COMPARISON OF dD AND AEROMAGNETIC MAPS 

Several workers have pointed out that the Eocene 

epizonal plutons in the Idaho batholith region are commonly associated with positive magnetic anomalies 

(Kiilsgaard and others, 1970; Cater and others, 1973; 

Tschanz and others, 1974; Mabey and others, 1978). Criss 

and Champion (1984) showed that these Eocene plutons 

generally have much higher magnetic susceptibilities and 

remanent magnetic intensities than typical pre-Eocene 

rocks of the southern Idaho batholith interior, although 

they also outlined some areas that were exceptional. Thus, 

the aeromagnetic intensity maps of Zietz and others (1978, 

1980) can greatly aid deduction of the distribution of the 

epizonal rocks, particularly in regions that have been 

mapped only in reconnaissance. Furthermore, in certain 

cases magnetic intensity maps provide unique interpretative data, such as when subsurface plutons are present. 

Criss and Taylor (1983) and Criss and Champion (1984) 

demonstrated a close correlation between the fossil hydrothermal systems (low d18Q and dD zones) and positive 

magnetic anomalies in the southern Idaho batholith; this 

correlation was in some areas better than that between 

the isotopic data and geologic maps. 

As regards the present investigations, however, a good 

correlation between aeromagnetic anomalies (Zietz and 

others, 1978, 1980) and dD values was found only in the 

southern part of the Bitterroot lobe (fig. 6.12). Most of 

the lowest dD values occur along and within the large, 

northwest-trending positive aeromagnetic anomaly (fig. 

6.12, features A and B). Feature A clearly overlaps the 

outcrop of the epizonal Painted Rocks pluton. The Running Creek pluton lies along the same magnetic trend immediately to the northwest, between features A and B. 

Feature B is not clearly related to a mapped Eocene 

pluton, and a normal dD value of - 73 per mil that occurs 

nearby is enigmatic. However, the only extremely low dD 

values in the western Bitterroot lobe lie on the northwest 

shoulder of feature B, on the Lochsa River near Lowell. 

It is possible that Eocene material, possibly present only 

in the subsurface, is at least partly responsible for feature 

B. The Whistling Pig and Paradise plutons (see map by 

Toth and others, 1983) occur along the north margin of 

the large aeromagnetic anomaly (A and B), but are not 

clearly related to it. Extremely low dD values occur within 

and near the Whistling Pig pluton. Pronounced hydrothermal effects are not associated with the Paradise pluton, 

but this Eocene intrusion was emplaced at a relatively 

deep level and probably is best referred to the main-phase 

group. 

On the other hand, the correlation between positive 

magnetic anomalies and dD values is extremely poor in 

the main north-central mass of the Bitterroot lobe. Two 

large Eocene plutons, the Lolo Hot Springs and the 

Bungalow (see fig. 6.1), have been mapped in this region 

and both are associated with low dD values. Neither of 

these plutons is associated with a significant magnetic 

anomaly, and a rather intense positive anomaly nearby 

(feature C) is not clearly associated with a mapped Eocene 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 115 

TABLE 6.3.-Rb-Sr and 6180 analyses of Wallace Formation metasedimentary rocks 

[Rb and Sr concentrations in parentheses are Kevex determinations; all others are by isotope dilution; t, diorite dike] 

North West Rb 

Sample latitude longitude (ppm) 

68-9-10B 46°35.3 1 ll5°47 e8 I 153.1 

-10C 46°35.3 1 115°47.8 I 176.6 

-10D 46°35.3 1 ll5°47.8 1 159.6 

-11B 46°47.0 1 115°49 o3 I 105.0 

-llC 46°47.0 1 115°49.3 1 (113.0) 

-liD 46°47.0 1 115°49 e6 I 129.0 

-12A 46°49.8 1 ll6°06.2 1 153.5 

-12B 46°54.2 1 116°06.9 1 183.3 

-12C 46°54.3 1 116°07.0 1 213.5 

-12D 46°54.7 1 116°09.6 1 (164 .8) 

-12E 47°00.3 1 ll6°ll.l 1 179.8 

-12F 47°00.5 1 116°13 oO I 178.2 

-12G 47°05.2 1 116°21.0 I 200.0 

-12H 47°28.0 1 116°34 e6 I 207.1 

-13A 47°28.4 1 115°55.3 1 (271.6) 

-13B 47°25.8 1 115°53 .s t (151.2) 

-13C 47°25.9' 115°53.7 1 (154.6) 

-13D 47°26.0' 115°54.3 1 (190 .1) 

69-7-lA 47°00.7 1 116°18.8 1 98.8 

-lB 47°00.7 1 116°19.1' ( 75 .4) 

-lC 47°00.4 1 116°20.8 1 205.6 

-lD 47°00.2 1 116°21.0' 187.6 

-IE 47°03.3 1 116°19.4 1 160.5 

-2A 47°12.1' 115°12.0 1 171.0 

-2B 47°ll .9 1 115°12.0 1 146.7 

-2ct 47°11.7' 115°11.9 1 (34.0) 

-3A 47°04.1' 115°21.3 1 165.5 

-3B 47°04.9' 115°21.5 1 181.1 

-3C 47°04.6 1 115°21.8 1 175.2 

-3D 47°04.4 1 115°22.2 I 105.0 

-3E 47°04.2 1 ll5°22.5 1 (179 .4) 

-3F 47°04.2 1 115°23 oS I (112.8) 

-3G 47°04.1 1 ll5°24.0 1 (167 .0) 

-3H 47°04.0' 115°24.5 1 (169.5) 

-31 47°04.0 1 115°25.3 1 (164 .3) 

-3J 47°03.9 1 115°25.9 1 (180.4) 

-3K 47°03.8 1 ll5°26.5 1 (166 .4) 

-31 47°03.8 1 115°26.9 1 ( 204. 7) 

-3M 47°03.5 1 115°27 o 7 I 053 .s) 

-4A 47°06.0 1 115°30 o9 I (149 .8) 

-4B 47°06.5 1 ll5°30 • 7 I 151.5 

-4C 47°08.1 1 115°29.2 t 171.8 

-4D 47°09.0 1 ll5°27 .9 1 (193.5) 

-4E 47°05.9 1 115°30.8 1 ( 207 .2) 

-4F 47°05.5 1 115°32.1 1 (166.6) 

-SA 47°06.5 1 115°35.9 1 (156.4) 

-SB 47°17.4 1 ll5°46 .4 1 (165. 7) 

-sc 47°16.3 1 115 °46.3 t (138.2) 

-SD 47°16.1 1 115°46.4 1 

epizonal pluton or with low dD values. In addition, the 

large positive anomaly immediately west of the Bitterroot lobe (feature D) generally coincides with exposures 

of the Tertiary Columbia River Basalt Group; such rocks 

would be expected to have high magnetic susceptibilities 

and high remanent magnetizations. Several small anomalies south of feature D apparently reflect varying 

volcanic and plutonic lithologies of the WSD terrane and 

the transition zone. 

Many of the above inconsistencies can be resolved by 

comparing the aeromagnetic anomalies to limited magSr a18o 

(ppm) 87 Rb/86sr 87 Sr /86sr (per mil) 

59.24 7.6016 0.88137 13.2 

31.14 16.6532 .86190 12.1 

27.15 17.2654 .86493 12.6 

45.86 6.6694 • 77983 11.8 

32.74 10.0371 .75960 10.5 

79.45 4. 7066 • 72518 10.8 

58.68 4.6088 .76146 11.2 

84.05 6.3750 .81509 12.0 

68.91 9.04 72 .80484 12.2 

49.90 9.6844 .84696 11.3 

(64.61) 12.7 

(14.19) 12.7 

48.57 12.1844 .93915 9.4 

70.13 8.6346 .81470 12.7 

14.43 56.5910 1.10938 15.0 

29.93 14.8983 .90732 14.5 

19.26 24.1304 1.10367 15.9 

20.97 27.0060 1.01210 14 .s 

23.16 12.4778 .818ll 12.5 

(66.55) 11.0 

219.2 2.7356 • 79272 12 .o 

51.84 10.6583 .88874 11.8 

97.73 4.8018 .81628 11.7 

(32.25) 15.34 .93207 13.0 

31.59 13 .8030 .98966 13.4 

(905.65) .1108 • 70689 8.6 

12.8 

46.40 11.5007 .89554 12.7 

49.80 10.4310 .96188 12.6 

110.20 2.7835 .80377 12.4 

58.85 6.9685 .87788 12.1 

45.31 7.2868 .82518 12.9 

56.84 8.6070 .83663 12 .s 

73.25 6. 7723 .82608 11.3 

40.56 12.3187 .91093 12.4 

64.05 8.2573 .84073 12 .s 

(113.91) 12.6 

( 71.69) 8.3664 .83696 12.0 

99.61 4.4998 .80009 12.5 

84.66 5.1640 • 79654 12.5 

73.56 6.0188 .8ll53 12.8 

(71.41) 12.8 

(102.09) 12.4 

115.69 5.2369 .81582 12.0 

61.24 7.9781 .84686 12 .s 

82.59 5.5264 .79407 13.3 

(24.85) 15.5 

76.31 5.2889 .80654 15.2 

14.3 

netic susceptibility data for the plutons (fig. 6.13). Although these data are preliminary, it is evident that 

extremely low magnetic susceptibilities are typical of the 

western and northern parts of the Bitterroot lobe. The 

absence of positive anomalies and the extremely low 

magnetic relief of such areas indicates that the lowsusceptibility rocks extend to great depths. Most of these 

low-susceptibility samples represent main-phase plutons, 

but a surprising number come from Eocene plutons. Much 

higher susceptibilities occur in the east-central part of the 

Bitterroot lobe, within a zone partly associated with the 
116 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

northwest-trending anomaly A-B (see fig. 6.12). An important and unexpected feature is that this high-susceptibility population is primarily composed of main-phase 

plutons, with the Eocene epizonal plutons generally 

having only moderate to low susceptibilities. Possibly, 

many of these main-phase rocks have enhanced susceptibilities due to Eocene alteration (see Criss and Champion, 1984), but further study is needed. In any case, no 

distinct susceptibility contrast exists between Eocene and 

older plutons in the Bitterroot lobe, in contrast to overall 

relations in the Atlanta lobe. Positive aeromagnetic 

anomalies and high magnetic susceptibilities are therefore 

not clearcut indicators of Eocene epizonal plutons in the 

northern Idaho batholith. 

RELATION OF ORE DEPOSITS TO THE 

HYDROTHERMAL SYSTEMS 

Following the discovery by Ross (1928, 1934) that Tertiary as well as Mesozoic plutonism was important in cen01-l --'--,.,.....Jl,L-0 -..L..i ,---_.2,0 KILOMETERS 

0 5 1

1

0 MILES 

tral Idaho, most workers have recognized that significant 

ore deposits were formed during both magmatic episodes 

(for example, Ross, 1931; Finch, 1932; Bergendahl, 1964; 

Mcintyre and others, 1976). Ross (1931) believed that 

most central Idaho deposits are Mesozoic and related to 

the Idaho batholith; however, Anderson (1951) argued 

that most ores in the region are Tertiary and that the 

Idaho batholith was metallogenically insignificant. On the 

other hand, Finch (1932) concluded that most gold 

mineralization in north-central Idaho (for example, near 

Elk City and Warren) is related to the Idaho batholith, 

insofar as the deposits are mesothermal and generally occur near roof pendants, whereas most gold deposits in 

south-central Idaho are epithermal and related to younger 

magmas. 

Stable-isotope data have considerable bearing on the 

genesis and probable ages of ore deposits in central Idaho. 

Criss and Taylor (1983) argued that most ore deposits in 

the southern half of the Idaho batholith are Tertiary 

because they are spatially associated with large zones of 

anomalous d18Q values, generally related to shallow-level, 

EXPLANATION 

Plutonic rocks related 

to the Idaho batholith 

....--- Metamorphic isograd 

0, in per mil 

• 15- 16 

a 14- 15 

() 13 - 14 

~ 12 - 13 

0 11 - 12 

0 10 - 11 

Metamorphic Zones 

B Biotite 

G Garnet 

St Staurolite 

St+K Staurolite + kyanite 

K Kyanite 

K+S Kyanite + sillimanite 

S+M Sillimanite + muscovite 

A+K+S Andalusite + kyanite + sillimanite 

FIGURE 6.11.-Whole-rock o180 analyses of Belt Supergroup (Wallace Formation) pelitic metasedimentary rocks plotted on the metamorphic 

isograd map of Hietanen (1984). The o180 values decrease progressively toward the north margin of the Idaho batholith. These data indicate 

widespread and pervasive exchange of the metasedimentary rocks with the plutonic rocks and with high-180 fluids during Cretaceous-early 

Tertiary time. 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 117 

Tertiary meteoric-hydrothermal systems. In this regard, 

the deposits are similar to epithermal gold- and silverbearing veins that occur in the Challis Volcanics (Criss 

and others 1984, 1985) as well as in even younger volcanic 

rocks at Silver City (Criss, unpub. data, 1983), all of which 

are irrefutably Tertiary in age. The following discussion, 

however, demonstrates that only minor mineralization is 

associated with the relatively weak Eocene meteorichydrothermal systems in the Bitterroot lobe, but that 

significant mineralization is more likely related to deeplevel hydrothermal activity related to emplacement of the 

main-phase Cretaceous-early Tertiary plutons. The overall 

FIGURE 6.12.-Relation of positive aeromagnetic anomalies to the -120 

per mil dD contour (hachured line, from fig. 6.9) in the northern Idaho 

batholith and adjacent areas. The 11,000-gamma contour of Zeitz and 

others (1978, 1980) is shown as the 0-gamma contour on this map. 

In places where the magnetic anomalies and the low-dD values are 

proximal (features A and B) the magnetic anomalies are probably 

situation may therefore be similar to that envisaged by 

Finch (1932). 

In order to evaluate mineral controls in the Idaho batholith, figure 6.14 was prepared to show the distribution of 

mines and prospects in the region, based on data compiled 

by Mitchell and others (1981a, b) in Idaho. In constructing this figure an attempt has been made to delete placer 

deposits, stone quarries, sedimentary deposits, and a few 

pegmatitic and metamorphic deposits, but the data are 

otherwise unedited. Very few of the deposits shown in 

figure 6.14 were productive and even fewer were significant, but no attempt is made here to evaluate the imporrelated to Eocene granitoids with high magnetic susceptibilities. 

However, positive aeromagnetic anomalies are not associated with 

the low-dD Bungalow and Lolo Hot Springs plutons (see fig. 6.1). Other 

positive aeromagnetic anomalies in the region (C, D) are probably 

related to different sources (Columbia River basalt for D?). 
118 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

tance of individual deposits. It can be assumed, however, 

that the areas around the most important mines were the 

most actively explored, so the density of points on figure 

6.14 does give some indication of mineral potential. 

Lindgren (1904) originally emphasized that the Bitterroot lobe is remarkably devoid of economic mineral 

deposits. Figure 6.14 clearly shows that most of the ore 

deposits are near the margins of the Bitterroot lobe, principally within highly metamorphosed Precambrian host 

rocks or within areas where screens and roof pendants 

of such rocks are closely interspersed with the batholith. 

This relation, combined with the nearly complete absence 

of correlation of the deposits to the low dD zones, suggests that most of the mineralization is related to emplacement of the main phase of the batholith rather than to 

the Eocene hydrothermal systems. This observation further suggests that most of the ore deposits in fact predate 

emplacement of the Eocene plutons; note that Snee and 

others (1985) have recently reported 40 Arf39 Ar ages of 

78 to 71 Ma for muscovites from gold-bearing quartz veins 

from several mining districts in north-central Idaho. As 

discussed above, the significant d180 depletions in the 

metasedimentary wall rock surrounding the batholith provide evidence that major deep-level metamorphic-hydrothermal systems existed in the region during Cretaceousearly Tertiary time. 

On the other hand, Eocene meteoric-hydrothermal 

systems are probably responsible for at least a few of the 

1160 1150 1140 

47°r----rr----------------------,r---------------------~~----------------,---~---------, 

I 

t 

! 

i i ;~oscow .n 

i 

I 

i 

I 

Lewi-;tnn 

( 

'\ 

\ 

( 

.? 

<' 

WASHINGTON\ 46° ---·-··-··\ 

OREGON \. 

' L-, EXPLANATION .,__ 

MAGNETIC 

SUSCEPTIBILITY 

• >100 

a 5o-1oo 

() 25-50 

Q 10-25 

0 <10 

/ 

I 

0 10 MILES : 

f--r-J-,--' / 10 KILOMETERS j 

j 

/ 

' t 

\ 

? 

( 

r 

( 

0 

oo 

a 

00 ~ 0 

0 

0 Q 

0 

~ 

f 

oOCID@ ~ '\ 

mo .,/ 0 45°~----~~~----------~------~----~L----------------L--~~L---------------~~------~ 

FIGURE 6.13.-Magnetic-susceptibility data for plutons of north-central Idaho. Low susceptibilities are typical of the west-central and northern 

parts of the Bitterroot lobe, but relatively high susceptibilities occur within a broad northwest-trending zone that contributes to high aeromagnetic 

intensities (see fig. 6.12). These rapid determinations were obtained on unmodified hand specimens using a hand-held magnetic susceptibility 

device and are reported as x x 10- 5 cgs (see Criss and Champion, 1984). Patterned areas same as in figure 6.1. 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 119 

ore deposits shown on figure 6.14; the commodities of 

interest are indicated for several of these. Obvious candidates are the Pb-Ag, Fe-Cu, and Mo deposits present 

within the Lolo Hot Springs pluton. Toth and others 

(1983) have additionally emphasized that many of the 

other Eocene epizonal plutons have Mo potential which 

could be of importance in the future. Furthermore, a few 

workings described by Toth and others (1983) in westernmost Montana occur in one of the low-dD zones and could 

very well be Eocene. Another possibility is a zone of SbAu-Ag mineralization near Lowell, which is near a small 

low-dD zone and near a positive aeromagnetic anomaly 

as well. Last, several significant deposits in the lowermost 

part of figure 6.14 (Edwardsburg and Profile districts) 

occur along the northwest side of the giant d180 anomaly mapped by Criss and others (1984) in the Challis 

volcanic field. However, prebatholithic rocks occur in the 

latter areas, so the origin of these deposits is unclear. 

In summary, Eocene meteoric-hydrothermal activity 

was not as important to mineralization in north-central 

Idaho as in the Atlanta lobe; conversely, Cretaceous-early 

Tertiary metamorphic-hydrothermal activity was not as 

important in the Atlanta lobe as it was in the Bitterroot 

lobe. This difference between the Bitterroot and Atlanta 

lobes is almost certainly related to the rather low intensity of the Eocene hydrothermal systems in the north and 

to the scarcity of prebatholithic wall rock and inliers in 

the south. An important corollary is that future exploration for Eocene ore deposits in the Bitterroot lobe should 

be concentrated in the relatively small areas of intense 

meteoric-hydrothermal alteration which occur within and 

near areas having the lowest d180 values. Additional 

study of Cretaceous-early Tertiary hydrothermal activity 

and its relation to ore deposits is needed. 

INFLUENCE OF HYDROTHERMAL SYSTEMS ON 

WHOLE-ROCK RB-SR ISOCHRONS 

Hydrothermal transport processes have the potential 

of disturbing the whole-rock Rb-Sr isochron systematics 

of rocks by introducing subsolidus changes in either the 

87SrJ86Sr ratio or the Rb/Sr ratio, or both. For example, 

any nonuniform modification of the 87SrJ86Sr ratios of 

rocks in a suite would contradict the necessary conditions 

of the isochron method, regardless of when the modification occurred. Additionally, hydrothermal alteration that 

significantly postdates crystallization of a rock suite could 

produce Rb/Sr ratios that are either much too low or too 

high to be consistent with observed 87SrJ86Sr ratios and 

the probable ages of the rocks. 

An important aspect of the Bitterroot lobe data are that 

the hydrothermal alteration systems do not seem to have 

profoundly influenced ri, even in cases where major 180 

depletions were produced. Thus, the altered (dD~ -120 

per mil) rocks have ri and d180 values either similar to 

those of unaltered Bitterroot lobe rocks or simply display 

lower d180 values. More importantly, some suites of 

Eocene rocks define good Rb-Sr isochrons (fig. 6.15) in 

spite of the fact that many of the samples underwent 

major 180 depletions. In any case, we submit that varying source characteristics, and not hydrothermal processes, were primarily responsible for the confused Rb-Sr 

isochrons observed by Armstrong and others (1977) and 

by Fleck and Criss (1985) in the main-phase granitoids. 

Note that the above results differ from those observed 

in submarine hydrothermal systems, where pervasive seawater interactions can produce significant changes in the 

87 Sr/ 86Sr ratios of the rocks (for example, McCulloch and 

others, 1980). This difference between the Eocene 

hydrothermal systems in Idaho and submarine systems 

can partly be attributed to the low Sr concentration of 

typical meteoric-hydrothermal fluids compared to sea 

water. Furthermore, the Sr that existed in the Bitterroot 

lobe systems would have been largely derived from the 

Cretaceous and Tertiary rocks undergoing alteration, and 

not from some external reservoir having a markedly different 87SrJ86Sr ratio. The Precambrian rocks provide an 

exception to the last statement, but their high 87 Sr/ 86Sr 

ratios are moderated by their low Sr concentrations. 

As regards any possible hydrothermal alteration of 

Rb-Sr ratios, it is important in the case at hand that the 

Bitterroot lobe plutons are not drastically older than the 

hydrothermal systems responsible for the alteration. Even 

if Eocene hydrothermal transport processes changed the 

Rb-Sr ratios from the values at emplacement, such 

changes would introduce only moderate complications into 

the isochron systematics of Late Cretaceous and early 

Tertiary plutons, and the whole-rock Rb-Sr isochrons of 

the Eocene plutons would be essentially unaffected. 

On the other hand, major complications in Precambrian 

isochrons would result from such nonclosed system 

behavior. An example is provided by Belt Supergroup 

(Wallace Formation) metasedimentary rocks collected 

peripheral to the northern margin of the Bitterroot lobe 

(fig. 6.16). Although a Precambrian age is still suggested 

by the data, the absence of conformity to a simple isochron 

is notable, especially when these data are compared to the 

good to excellent isochrons obtained on Belt sedimentary 

rocks elsewhere (Obradovich and Peterman, 1968). The 

effects shown in figure 6.16 probably reflect the highgrade regional metamorphism of the Belt rocks in the 

vicinity of the Bitterroot lobe, as well as the pervasive 

interactions with metamorphic-hydrothermal fluid which 

apparently occurred in Cretaceous-early Tertiary time. 

It is alternatively possible that the results indicate large 

spatial variations in the initial 87SrJ86Sr ratios of the 

sedimentary rocks in the region of study, but this is considered to be less likely. 
120 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

# • 

.. , ... 

: • c·· . . . . . . . . 

ATLANTA LOBE 

{/. 

. . .. • • • 

'r 

10 20 MILES 

r' I, I 

10 20 KILOMETERS 

EXPLANATION 

Number of mines and prospects 

• 2:.10 

• 3-9 

/ 

r_ ........ 

;l 

',) 

) 

\ 

i.-....... \ 

; ........ \ 

~ -120 per mil 60 contour from figure 6.9 

- Margin of Idaho batholith 

Alunite 

,.....___.- ./. 

• •• 

.. 

( 

/~ 

:-!. 

•• 

\ 

.... 

• 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 121 

K-AR RELATIONS IN PLUTONS OF 

NORTH-CENTRAL IDAHO 

REGIONAL VARIATIONS IN K-AR AND 40AR/ 39AR 

APPARENT AGES OF BIOTITE 

Excluding Precambrian rocks, the K-Ar apparent ages 

of biotite from plutons of north-central Idaho range from 

136 to 39 Ma (fig. 6.17). These apparent ages are a function of both rock type and geographic position, and in 

many cases are significantly younger than the true emplacement ages of the rocks. The oldest K-Ar ages from 

biotites (136 to 78 Ma) are obtained from plutons in the 

Wallowa-Seven Devils terrane, west of the 0.704line for 

87Srf86Sr initial ratios (compare figs. 6.17 and 6.2). Apparent ages of biotite in transition zone plutons that 

intrude the eastern and northern margins of the WSD 

terrane range from 82 to 76 Ma, similar to the youngest 

apparent ages of biotite of the WSD plutons (table 6.4; 

fig. 6.17). Such values occur only near the terrane boundary with Precambrian sialic crust. 

Biotite K-Ar apparent ages of plutons of the Bitterroot 

lobe of the Idaho batholith are in most cases younger than 

80 Ma. These apparent ages decrease systematically from 

about 80 Ma near the suture with the WSD terrane to 

less than 50 Ma in the broad central part of the Bitterroot lobe (fig. 6.17). The region of low (<51 Ma) K-Ar apparent ages is more than 10,000 km2 in extent, as 

originally pointed out by Armstrong (197 4). This unusual 

region completely encompasses the Bitterroot dome of 

Hyndman and others (1975). The youngest K-Ar apparent 

ages ( 45 to 39 Ma) occur either within and immediately 

adjacent to the Eocene plutons or within the extensive 

mylonite zone near the Idaho-Montana state line (see section "Geochronology and Structural Development of the 

Idaho Batholith"). 

As discussed by Armstrong (1974), the K-Ar apparent 

ages (51 to 44 Ma) of the Eocene plutons are essentially 

the same as those typical of the Challis Volcanics (also 

see Axelrod, 1966; Mcintyre and others, 1982). This 

volcanic sequence contains an Eocene paleoflora (Axelrod, 

1966) and is in places crosscut by the Eocene plutons 

FIGURE 6.14.-Distribution of mines and prospects (dots; predominantly 

representing hydrothermal ore bodies) in north-central Idaho 

(modified from Mitchell and others, 1981a, b). Mineral deposits are 

scarce in the Bitterroot lobe, except along its margins and near 

inliers of Precambrian rocks, where metamorphic-hydrothermal activity related to the main-phase plutons occurred. Only a few deposits 

appear to be related to the Eocene meteoric-hydrothermal systems, 

whose positions are indicated by the -120 per mil oD contour 

(hachured lines). The commodities of interest are indicated for these 

Eocene deposits. Mineral deposits in Montana are not shown, except one alunite deposit (Eocene?) described 'by Toth and others 

(1983) is illustrated. 

(Ross, 1934; Cater and others, 1973). In the area of the 

Painted Rocks pluton (fig. 6.1), tuffs of the volcanic sequence yield biotite and sanidine K-Ar and 40 Arf39 Ar 

ages ranging from 49 to 45 Ma (table 6.4, 6.5, sample 

797 -19C) and rest nonconformably on the epizonal plutonic 

rocks. Initial 87Srf86Sr ratios of the volcanic rocks are indistinguishable from those of the Painted Rocks pluton 

(797-20E). 

REGIONAL VARIATIONS IN K-AR AND 40AR/ 39AR 

APPARENT AGES OF HORNBLENDE 

The K-Ar and 40 Arf39 Ar apparent ages of plutonic 

hornblende exhibit a range from 192 to 46 Ma within the 

central Idaho-westernmost Montana area (figs. 6.18, 6.19; 

tables 6.4, 6.5). Where zircon U-Pb ages are available for 

plutons of central Idaho, discordance with K-Ar ages on 

hornblende is generally small (Armstrong and others, 

1977; J.S. Stacey, personal commun., 1983). Ages of 

plutonic hornblendes drop abruptly across the boundary 

between the WSD and Belt-Y ellowjacket terranes and 

continue to decrease in steps eastward across the Idaho 

batholith (figs. 6.18, 6.19). The two terranes are sutured 

by transition zone plutons that intrude both crustal types 

and yield biotite, hornblende, and muscovite K-Ar ages 

of 87 to 75 Ma, many of which are concordant. 

Excluding plutons of the transition zone, the youngest 

plutons intruding the WSD terrane are Early Cretaceous 

in age, commonly greater than 125 Ma. Sample 827 -28G 

has concordant biotite and hornblende K-Ar ages at 135 

Ma, and sample 818-10B is nearly concordant, with hornblende at 135 Ma and biotite at 129 Ma (table 6.4). Jurassic 

bodies, yielding K-Ar ages of 170 to 140 Ma on hornblende, are also common in the WSD terrane (Armstrong 

and others, 1977). Although Lower Jurassic to Upper 

Permian intrusive rocks are reported in eastern Oregon 

and south of 45° N.latitude in western Idaho (Armstrong 

and others, 1977; Ave Lallemant and others, 1980; 

Walker, 1983), only one of our samples yields K-Ar ages 

on hornblende greater than 170 Ma, and that sample is 

poorly reproducible, ranging from 199 to 170 Ma (tables 

6.4, 6.5). These Jurassic and Early Cretaceous plutons are 

notably absent in the Belt-Y ellowjacket terrane, north and 

east of the suture zone. 

Hornblende K -Ar apparent ages in the Bitterroot lobe 

are oldest along the western margin and generally 

decrease to a minimum along the Bitterroot front. The 

hornblende ages define three groups of rather uniform 

ages-87 to 75 Ma, 66 to 62 Ma, and 51 to 46 Ma (fig. 

6.19). These age groupings are based on too few data to 

be regarded as clearly discrete events and additional data 

must be obtained. However, the geographic pattern is consistent with eastward migration of plutonism across central Idaho during Late Cretaceous and early Tertiary 

time. 
122 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

GEOCHRONOLOGY AND STRUCTURAL 

DEVELOPMENT OF THE 

IDAHO BATHOLITH REGION 

GENERAL GEOCHRONOLOGIC RELATIONS 

Potassium-argon ages of biotite, hornblende, and 

muscovite provide a pattern of decreasing apparent age 

from west to east across central Idaho, including plutons 

in the Wallowa-Seven Devils terrane, plutons on the 

western margins of both the Bitterroot and Atlanta lobes 

of the Idaho batholith, the main-phase plutons of the 

Bitterroot lobe, and the late-phase epizonal plutons. 

Jurassic-Early Cretaceous plutons were emplaced into 

older units of the WSD magmatic-sedimentary arc at a 

time when it was spatially remote from its present juxtaposition with Precambrian sialic crust. Plutons of this 

group are confined to the WSD terrane and yield ages 

from 192 Ma to 135 Ma (fig. 6.19). Late Cretaceous transition zone plutons yield K-Ar mineral ages of 87 to 75 

Ma and define a belt of strongly deformed bodies centered 

along the terrane boundary (Armstrong and others, 1977; 

Fleck and Criss, 1985). Whereas some of the apparent 

ages may represent resetting during recrystallization, zircon U-Pb ages (J.S. Stacey, written commun., 1983) of 

some plutons are also concordant with the K-Ar determinations, and most intrusive rocks within the suture zone 

are probably syntectonic. As discussed by Hyndman 

(1983) and Zen and Hammarstrom (1984), depths of 

crystallization of these plutons were probably greater than 

10 to 15 km. The near concordance of mineral ages virtually requires a tectonic emplacement, because slow 

uplift from these depths would yield extremely discordant 

ages in the order: zircon U-Pb > hornblende K-Ar > 

muscovite K-Ar ~biotite K-Ar. This mode of emplacement is consistent with the strongly deformed fabrics of 

the plutons and with high-grade metamorphism of Precambrian wall rocks. 

A younger zone of nearly concordant ages occurs within 

the Bitterroot mylonite (figs. 6.1, 6.19 at about 100 km), 

a shallow-dipping zone of deformation in the youngest 

plutons of the main phase of the Bitterroot lobe. Zircon 

U-Pb ages (Chase and others, 1978, 1983; Bickford and 

others, 1981) and K-Ar ages on hornblende and biotite 

(table 6.4) are nearly concordant (52 to 45 Ma) for the 

youngest deformed plutons. Concordant biotite and hornblende K-Ar ages of mylonite sample 787-25C (table 6.4; 

Fleck and Criss, 1985) indicate an age of mylonization of 

45.7 Ma, consistent with Garmezy and Sutter's (1983) 

determination of 45.5 to 43.5 Ma. As with the suture zone, 

.780 TERTIARY PLUTONS 

,....,. 

L. 

en .760 

co 

co 

.......... 

L. 

en .740 EXPLANATION 

I'-

co 

'-' o Painted Rocks pluton 

.720 • Running Creek pluton 

• Whistling Pig pluton 

.700~----~----~------~----~------~----~----~------~----~----~ 

0 1 0 20 30 40 50 60 70 80 90 100 

FIGURE 6.15.-Whole-rock Rb-Sr isochron diagram for epizonal granites from the most highly altered part of the Bitterroot lobe, namely, the 

Painted Rocks, Running Creek, and Whistling Pig plutons. These rocks define reasonable isochrons in spite of the high degree of alteration 

of several samples. Sample 797-20B is off scale, having an 87Rb/ 86Sr ratio of 145.9 and a present-day 87Sr/ 86Sr ratio of 0.8090, but this 

highly altered sample (d180 = 5.8 per mil; dD = -133 per mil) would also plot close to the 50-Ma reference isochron. 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 123 

the plutons with mylonitic fabrics clearly crystallized at 

least in part at great depth, but rose abruptly (see subsection on "Bitterroot Lobe Uplift") to levels sufficiently 

shallow that biotite could retain argon (<250 to 300 °C). 

Except for the two elongate zones discussed above and 

less common, undeformed, Early Cretaceous plutons in 

the accreted terrane (table 6.4), plutonic rocks of northcentral Idaho generally yield discordant K-Ar apparent 

ages (fig. 6.19). The patterns of these age variations appear to be regular (figs. 6.17, 6.18), but ages tend to fall 

into several groups rather than being evenly distributed 

in time. Discordant ages of Late Cretaceous and early Tertiary plutons suggest uplift and cooling of an episodically 

emplaced batholith after about 75 Ma and before 45 Ma, 

with biotite apparent ages generally being 4 to 14 m.y. 

younger than hornblende. The youngest plutonic event 

in the region was represented by widespread volcanism, 

caldera formation, hydrothermal activity, and intrusion 

of the highly differentiated epizonal Eocene plutons at 

51 to 44 Ma. 

TIMING OF ACCRETION 

Age patterns in and adjacent to the suture zone of the 

WSD terrane with the sialic crust of the North American 

continent restrict the time of accretion to the period 

between 135 and 75 Ma. West and south of the narrow 

zone of highly deformed transition zone plutons having 

nearly concordant mineral ages of 87 to 75 Ma, K-Ar apparent ages on biotite and hornblende from some tonalite 

U) '"" \() 

00 

t::; 

U) 

..... 

1.100 

1.000 

00 0.900 

0.800 

• 

0.700 '------~--~--~---~-----'-----___j 

10 15 20 25 

FIGURE 6.16.-Whole-rock Rb-Sr isochron diagram for Belt Supergroup 

(Wallace Formation) metasedimentary rocks collected north of the 

Bitterroot lobe. In contrast to a 1,100-Ma isochron (Obradovich 

and Peterman, 1968), data scatter widely about an 800-Ma reference 

isochron, perhaps reflecting hydrothermal metamorphism in 

Cretaceous-early Tertiary time. 

and quartz diorite plutons are concordant at 135 Ma. 

These Early Cretaceous plutons have flat 40 Arf39 Ar age 

spectra (hornblende) and have initial 87Srf86Sr ratios of 

0. 7033 to 0. 7035 that show no indication of proximity to 

a sialic craton of Precambrian age (fig. 6.20). Less than 

7 km from these ensimatic plutons, Late Cretaceous 

syntectonic tonalite and trondhjemite bodies of the transition zone intrude the WSD terrane and yield flat hornblende age spectra and the nearly concordant mineral 

ages discussed above. Zircon fractions studied by J.S. 

Stacey (written commun., 1983) from one of these low-ri 

trondhjemites yield a lower intercept U-Ph age concordant 

with muscovite K-Ar ages of 87 Ma (table 6.4; sample 

797 -26D), but contain a significant component of inherited 

Precambrian zircon. The initial87Srf86Sr ratio from this 

body is low (0.7037) but significantly above values from 

the Early Cretaceous plutons. Some transition zone 

plutons also intrude Precambrian crust, and their high ri 

values (0. 705 to 0. 709) demonstrate a significant Precambrian component in these magmas (Fleck and Criss, 1985; 

table 6.1). All lines of evidence indicate at least some involvement of Precambrian sialic crust in the magmatic 

history of the transition zone plutons. 

Early Cretaceous tonalite (sample 797-25F) 1 to 2 km 

west of the suture zone near Kooskia, Idaho, has an initial87Srf86Sr ratio of 0.7034 (table 6.1) and a 40Arf39Ar 

plateau age on hornblende of 130.4 ± 5.3 Ma (table 6.5), 

but biotite K-Ar apparent ages are about 78 Ma. Biotite 

may have been reset by deformation and intrusion within 

the suture zone during the Late Cretaceous event. On the 

other hand, hornblende in this sample, which yields a partially disturbed 40 Arf39 Ar age spectrum (table 6.5) and 

conventional K-Ar ages of 118 Ma, still retains evidence 

of its Early Cretaceous age. The absence of Early Cretaceous intrusive bodies within the Precambrian terrane 

east of the suture zone and the evidence of a major 

magmatic and tectonic event at 87 to 75 Ma argue persuasively for a Late Cretaceous age for final accretion 

(fig. 6.20). 

BITTERROOT LOBE UPLIFT 

In addition to indicating the timing of the regional igneous events, K-Ar data provide important information 

on the uplift of the Bitterroot lobe. In the discussion 

below, we elaborate on Ferguson's (1975) suggestion that 

the K-Ar apparent ages of Bitterroot lobe plutons reflect 

early Tertiary uplift and cooling of the terrane. Uplift 

models have been advanced to explain apparent age patterns in several orogenic terranes, including the Alps 

(Jager, 1962), the Ruby Mountains of Nevada (Kistler and 

O'Neil, 1975), the Peninsular Ranges batholith (Krummenacher and others, 1975), the Sandpoint quadrangle 

of Washington and Idaho (Harms and Price, 1983; com-
124 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

pare with Miller and Engels, 1975), and the southern 

Idaho batholith (Criss and others, 1982). 

Several independent lines of evidence are in accord with 

marked Late Cretaceous and early Tertiary uplift of the 

Idaho batholith, and these have been summarized earlier 

in this report. Structural evidence is also generally consistent with such an uplift history, although certain 

disagreements exist at present and many details remain 

to be elucidated. Chase and Talbot (1973) originally proposed that the Bitterroot lobe has a domelike structure 

that formed by an uplift mechanism. Hyndman and others 

(1975) and Hyndman (1980) suggested that the Bitterroot 

dome formed by isostatic rebound, following detachment 

of the Sapphire tectonic block along the major mylonite 

zone and the removal of this formerly superincumbent 

block by gravitational gliding. Hyndman proposed that 

the mylonite zone is therefore genetically connected to 

78-Ma thrust faults more than 100 km to the east. However, Chase and others (1983) obtained U-Pb ages of 52 to 

48 Ma for three sheared plutons and pointed out the age 

discrepancy between this result and the genetic implications of the early model. The timing of mylonite formation has been confirmed by Garmezy and Sutter (1983), 

who indicate an age range of 45.5 to 43.5 Ma for the 

mylonization event, significantly younger than the age of 

crystallization of the youngest mylonized plutons. 

1160 1150 1140 

47°~---:r----------------------,r-~------------------~~--------------~r----.---------, 

( 

·\~ 

\ 

( 

.> 

< 

WASHINGTON\ 46° ----------\ 

OREGON \. 

'· 

& 

124 

l.-,, 

EXPLANATION ·._. 

e Data from tables6.4.6.5t ' A Data from Armstrong (1975·)\. 

~ 40Ar/ 39Ar total fusion age } 

:' 47 Conventional K-Ar age ( 

( ~0 MILES / 

0 10 KILOMETERS ( 

/ 

/ 

/119 

... 

/ 124 

.

/ 45°~----L-~~----------~--~~~-----2~~------~--~-L~~~L-~~--------~~~------~ 

FIGURE 6.17.-Contour map of K-Ar and 40Ar/ 39Ar apparent ages of biotites from plutons of north-central Idaho. Note the large area of 

apparent age of 50 to 45 Ma that reflects Eocene intrusives as well as rapid uplift and cooling of the east-central Bitterroot lobe. 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 125 

TABLE 6.4.-K-Ar apparent ages of rocks of central Idaho and western Montana 

[Age uncertainties represent larger of (1) standard deviation of replicates, or (2) pooled estimates of standard deviations, using formula of Cox and 

Dalrymple (1967). Sample locations are given in table 6.1. 40Ar refers to radiogenic argon. t, see table 6.5 for 40Ar/ 39Ar analyses. E, exponential 

to the negative value given] 

K2o Percent Age 

Sample Mineral (weight percent) mol40 Ar/g 40Ar (Ma) 

Plutonic rocks of the Wallowa-Seven Devils terrane 

797-25Ft Biotite 9.21' 9.23 1.058 E-9 73.1 77. 9±0 .5 

do. 1.056 E-9 77.7 

Hornblende .799, .818 1.424 E-10 77.7 ll8 .6±0 .8 

797-25 It 

do. 1.429 E-10 70.3 

Hornblende .319, .320 8.906 E-ll 79 .5 191. 7±11.0 

do. 9.700 E-ll 65.9 

807-280 Biotite 8. 76' 8.76 1.005 E-9 82.1 78.0±0.5 

( metavo1 canic) 

Hornblende .858, .863 1.000 E-10 55.9 78.2±1.1 

818-10Bt 

do. 9.796 E-ll 53.5 

Biotite 9.68' 9.64 1.810 E-9 82.3 129.2±4.8 

do. 1.914 E-9 84.1 

Hornblende .805, .816 1.625 E-10 74.1 134.6±0.8 

do. 1.638 E-10 70.3 

818-16C Biotite 9.88' 9.88 1.205 E-9 62.7 82.8±0.7 

Muscovite 10.27, 10.30 1.217 E-9 78.5 80.4±0.6 

827-28G t Biotite 9.51, 9.54 1.887 E-9 88.0 135. 6±4 .3 

do. 1.976 E-9 85.7 

Hornblende 1.021, 1.021 2.049 E-10 86.9 135.4±1.6 

do. 2.085 E-10 67.9 

Plutonic rocks of the transition zone 

797-250 Biotite 9.34, 9.37 8.169 E-10 87.4 59.8±0.3 

do. 8.210 E-10 77.8 

Hornblende 1. 750' 1. 777 1.919 E-10 69.7 74.4±0.5 

797-260 t 

do. 1.935 E-10 83.9 

Muscovite 10.75, 10.76 1.382 E-9 67.1 87.2±0.5 

do. 1.385 E-9 85.1 

797-26E Biotite 9.60, 9.65 1.117 E-9 70.8 79.2±0.6 

do. 1.128 E-9 83.2 

797-26H Biotite 9.62, 9.64 1.020 E-9 89.5 72.1 ±0 .4 

797-26It 

Muscovite 10.74' 10.76 1.159 E-9 86.4 73.4±0.4 

Biotite 9.62, 9.63 1.078 E-9 51.7 75.1±1.6 

do. 1.046 E-9 89.2 

Hornblende 1.594, 1.598 1. 900 E-10 61.3 79.8±1.5 

do. 1.851 E-10 59.4 

807-27A Biotite 9.28, 9.28 1 .063 E-9 a8.2 77. 6±0 .4 

do. 1.057 E-9 79 .3 

807-28A Biotite 9.17, 9.23 1.067 E-9 87.6 7a.6±0.4 

do. 1.061 E-9 83.8 

Muscovite 9.91, 9.93 1.134 E-9 72.6 78.1±0.9 

do. 1.147 E-9 18 .1 

807-28H Biotite 9.34, 9.42 1.059 E-9 93.8 76.6±0.4 

do. 1.054 E-9 73.0 

Muscovite 9.95' 9.97 1.133 E-9 82.9 76.7±0.9 

807-2ait 

do. 1.113 E-9 ao.9 

Biotite 9.43' 9.45 1.057 E-9 76 .a 75.a±0.5 

do. 1.046 E-9 71.5 

Hornblende 1.165, 1.16a 1.355 E-10 12 .a 7a.5±0.6 

do. 1.341 E-10 49.a 
126 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

TABLE 6.4.-K-Ar apparent ages of rocks of central Idaho and western Montana-Continued 

[Age uncertainties represent larger of (1) standard deviation of replicates, or (2) pooled estimates of standard deviations, using formula of Cox and 

Dalrymple (1967). Sample locations are given in table 6.1. 40Ar refers to radiogenic argon. t. see table 6.5 for 40Ar/ 39Ar analyses. E, exponential 

to the negative value given] 

Sample 

818-llC 

818-15D 

818-16D 

818-16E 

818-16Ft 

818-16 I 

818-16J 

827-29F 

827-29J 

787-18A 

787-18B 

787-19At 

787-19G 

787-20A t 

787-20Et 

787-24A 

787-24G 

Mineral 

K2 0 

(weight percent) 

Percent 

40Ar 

Plutonic rocks of the transition zone--Continued 

Hornblende 

do. 

Biotite 

do. 

Hornblende 

do. 

Biotite 

do. 

Hornblende 

do. 

Biotite 

do. 

Biotite 

do. 

Hornblende 

do. 

Biotite 

do. 

Hornblende 

Biotite 

do. 

Hornblende 

Biotite 

Biotite 

do. 

Hornblende 

do. 

.385' .388 

8.67, 8.70 

.898, .901 

9.01, 9.03 

.830' .839 

9.16, 9.16 

9.22, 9.24 

.982, .985 

9.55, 9.68 

1.312, 1.316 

9.56, 9.56 

1 .4ll' 1 .413 

9.70, 9.65 

9.33, 9.31 

1 .455' 1 .446 

4.091 E-ll 

4.332 E-ll 

1.027 E-9 

1.048 E-9 

1.045 E-10 

1.063 E-10 

1.078 E-9 

1.081 E-9 

9.293 E-ll 

9.628 E-ll 

1.104 E-9 

1 .068 E-9 

1 .080 E-9 

1 .122 E-9 

1.178 E-10 

1.360 E-10 

1.086 E-9 

1.081 E-9 

1.547 E-10 

1 .066 E-9 

1.094 E-9 

1. 700 E-10 

1.117 E-9 

1 .133 E-9 

1.123 E-9 

1. 834 E-10 

1. 705 E-10 

21.1 

28.4 

75.5 

89.4 

78.8 

69.3 

91.5 

78.9 

70.8 

65.0 

79.5 

89.9 

73.2 

92.8 

70.3 

85.6 

82.7 

89.1 

80.5 

78.4 

92.9 

82.9 

82.6 

83.6 

79 .2 

67.2 

68 .1 

Plutonic rocks of .the Bitterroot lobe, undivided 

Biotite 

do. 

Biotite 

do. 

Muscovite 

do. 

Biotite 

do. 

Hornblende 

do. 

Biotite 

do. 

Muscovite 

do. 

Biotite 

do. 

Hornblende 

Biotite 

do. 

Biotite 

do. 

Biotite 

do. 

8.98, 9.04 

9.62' 9.64 

10.74, 10.74 

9.33, 9.34 

1.463, 1.465 

9.23, 9.25 

10.48, 10.50 

9.32, 9.32 

1.611, 1.642 

9.69, 9.71 

9.43, 9.45 

8.90, 8.95 

6.392 E-10 

6.410 E-10 

6.943 E-10 

6. 820 E-10 

8.078 E-10 

8.026 E-10 

7.226 E-10 

7.441 E-10 

1.399 E-10 

1.275 E-10 

6.581 E-10 

6.690 E-10 

7.840 E-10 

7.927 E-10 

7.201 E-10 

7.268 E-10 

1.482 E-10 

7.331 E-10 

7.397 E-10 

6.263 E-10 

6.263 E-10 

5.850 E-10 

5.864 E-10 

71.5 

78.2 

81.7 

84.4 

87.2 

72 .o 

83.9 

78.4 

72.2 

72.9 

83.3 

69.9 

81.4 

73.7 

78.9 

73.3 

72.5 

76.1 

75.6 

72.1 

77.8 

66.6 

86.6 

Age 

(Ma) 

74.2.±2.9 

81.1.±1.1 

79.6.±0.9 

81.3.±0.4 

77.1.±1.9 

80 .5.±1. 9 

81.0.±2.1 

87.5.±8.7 

76.6.±0.5 

80.0.±0.6 

76.8±1.4 

81. 7±0. 6 

78.4.±0.6 

82.2±0.5 

82.9±4.2 

48.7.±0.3 

49.0.±0.6 

51.3.±0.3 

53.8.±1.1 

62.4.±4.0 

49.2.±0.6 

51.5.±0 .4 

53.1.±0.3 

62.2.±0.4 

52.0±0.3 

45.5±0.2 

45.0±0.2 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 127 

TABLE 6.4.-K-Ar apparent ages of rocks of central Idaho and western Montana-Continued 

[Age uncertainties represent larger of (1) standard deviation of replicates, or (2) pooled estimates of standard deviations, using formula of Cox and 

Dalrymple (1967). Sample locations are given in table 6.1. 40Ar refers to radiogenic argon. t, see table 6.5 for 40Ar/ 39Ar analyses. E, exponential 

to the negative value given] 

K20 Percent Age 

Sample Mineral (weight percent) mo1 40 Ar/g 40Ar (Ma) 

Plutonic rocks of the Bitterroot lobe, undivided--Continued 

Hornblende 1.238, 1.240 8.933 E-ll 74.6 49.6±0.3 

do. 9.007 E-ll 78.1 

787-24Ht Biotite 9.21, 9.23 6.181 E-10 62.3 45.9±0.2 

do. 6.151 E-10 90.9 

Hornblende 9.414 E-ll 73.3 47.7±0.4 

787-25C Biotite 8.57, 8.59 5.703 E-10 73.7 45.3±0.4 

do. 5.630 E-10 74.4 

Hornblende 1.211' 1.215 8.183 E-ll 63.7 45.7±0.8 

787-26At 

do. 7.986 E-ll 44.8 

Biotite 9 .C9, 9.10 7.042 E-10 79.4 53.0±0.3 

Hornblende 1.373' 1.376 1.277 E-10 80.3 64.9±2.0 

78 7-26Ft 

do. 1.335 E-10 68 .1 

Biotite 9.34, 9.35 7.302 E-10 86.1 52.4±1.4 

do. 7.018 E-10 83.4 

Hornblende 1.256, 1.260 1.138 E-10 79.8 61.7±0.3 

do. 1.136 E-10 63.6 

797-23A Biotite 9.27, 9.28 7. 735 E-10 18.9 57.3±0.6 

Muscovite 10 .15' 10.16 8.148 E-10 77.5 55.0±0.2 

do. 8.190 E-10 79.5 

808-04E Biotite 9.24, 9.25 6.722 E-10 70.0 49.9±0.3 

do. 6.743 E-10 67.7 

Muscovite 10 .31' 10.45 7.883 E-10 69.5 52.3±0.4 

do. 7.973 E-10 78.1 

808-04F Biotite 9.54' 9.59 6.733 E-10 67.2 48.2±0.3 

Muscovite 10.43, 10.37 8.347 E-10 74.9 53 .8±1.6 

do. 7.993 E-10 80.9 

818-12D Biotite 9.56, 9.55 7.859 E-10 83.3 56.2±0.3 

818-13A Biotite 9.82, 9.86 1.083 E-9 82.2 74 .0±1.2 

do. 1.058 E-9 86.4 

68-9-11 A Biotite 9.06, 9.08 7.489 E-10 81.6 56.1±0.5 

do. 7.398 E-10 82.1 

Hornblende 1.413, 1.414 1.302 E-10 77.2 62. 7±0 .3 

do. 1.424, 1.434 1.308 E-10 53.6 

80-MT-246t Biotite 9.12, 9.12 5.917 E-10 56.6 44.5±0.3 

Hornblende 1.160' 1.173 7.441 E-ll 27.0 46.6±4.0 

do. 8.425 E-ll 40.0 

Eocene epizona 1 plutonic and associated volcanic rocks 

787-18Jt Biotite 8. 71' 8. 73 5.956 E-10 90.7 47.0±0.2 

do. 5.989 E-10 79.7 

787-21 A Biotite 8.89' 8.90 6.063 E-10 82.3 46.9±0.3 

do. 6.096 E-10 82.0 

787-21C Biotite 7.18, 7.19 4.816 E-10 76.8 45.9±0.2 

do. 4.800 E-10 79.8 

797-19B Biotite 8.60, 8.61 5.670 E-10 59.1 44.8±0.6 

do. 5.563 E-10 77.8 

Sanidine 11.15, ll .15 7.665 E-10 89.8 46.3±1.2 

d,o. 7.376 E-10 78.8 

797-19Ct Biotite 8.65 8.66 5.914 E-10 72.3 46.6±0.3 

(tuff) do. 5.850 E-10 78.5 

Sanidine 11 .51' ll.53 7.788 E-10 82.3 44.6±2.5 

do. 7.189 E-10 94.5 
128 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

TABLE 6.4.-K-Ar apparent ages of rocks of central Idaho and western Montana-Continued 

[Age uncertainties represent larger of (1) standard deviation of replicates, or (2) pooled estimates of standard deviations, using formula of Cox and 

Dalrymple (1967). Sample locations are given in table !1.1. 40Ar refers to radiogenic argon. t, see table 6.5 for 40Ar/ 39Ar analyses. E, exponential 

to the negative value given] 

K20 Percent Age 

Mineral (weight percent) mol Sample 40 Ar/g 40Ar (Ma) 

Eocene epizonal plutonic and associated volcanic rocks--Continued 

797-20E Biotite 5 .so' 5.53 3.604 E-10 47.8 46.0±1.6 

do. 3.794 E-10 64 .s 

Hornblende 1.121, 1.128 7.635 E-ll 73.0 46.6±0.4 

Metamorphosed Precambrian rocks 

797-25C Biotite 8. 72' 8.81 7.196 E-10 28.9 58 .2±2. 9 

do. 7. 723 E-10 72 .o 

Muscovite 9.35, 9.44 8.412 E-10 37.5 61.3±0.4 

do. 8.456 E-10 81.8 

807-31J Biotite 9.19, 9.30 8.624 E-10 75.0 64.1 ±0. 6 

807-31K t 

do. 8.736 E-10 86.8 

Biotite 8.20, 8.23 7.910 E-10 82.0 65 .0±1 .o 

do. 7.734 E-10 78.0 

Hornblende 0.964 1.082 E-10 57.2 74.6±2.4 

do. 1.033 E-10 59.6 

808-01 D Biotite 9 .ss' 9.57 9.121 E-10 81.5 65.0±0.4 

do. 9.091 E-10 77.3 

Muscovite 10.69' 10.72 1.015 E-9 88.4 65.7±1.4 

808-01Gt 

do. 

Biotite 9.47, 

do. 

808-01 I Hornblende 1.521, 

do. 

Several unusual features of the K-Ar data in the Bitterroot lobe are consistent with the tectonic explanation. 

Principal among these is the large central zone of low 

(51 to 44 Ma) K-Ar mineral ages (figs. 6.17 and 6.18). This 

10,000 km2 area and the surrounding belt of high apparent-age gradients and high age discordance is much too 

large to be solely attributed to the thermal aureoles of 

the crosscutting Eocene plutons and is also much larger 

than the zones of intense meteoric-hydrothermal alteration (compare figs. 6.9 and 6.10). However, the observed 

age pattern could have been produced by relatively rapid 

uplift and cooling of the rocks in this broad area, in which 

case the terrane could constitute an infrastructure. This 

uplift and cooling was apparently sufficiently rapid that 

vertical gradients in the K-Ar apparent ages are not 

discernible given the topographic relief of the area. 

Garmezy and Sutter (1983) used 40 Arf39 Ar data from 

hornblende, muscovite, biotite, and orthoclase to calculate 

1.047 E-9 83.5 

9.54 9.510 E-10 79.9 72 .4±6 .o 

1.071 E-9 91.9 

1.535 1.765 E-10 74.2 79.1±0.9 

1.793 E-10 84.8 

that the uplift rate was about 20 ± 10 mm/yr from 45.5 

to 43.5 Ma, approaching that observed in the Himalaya 

mountains. Chase and Talbot (1973) suggested that the 

Bitterroot dome plutons were emplaced during the uplift 

event in a semiconsolidated condition, and Garmezy and 

Greenwood (1981) proposed that the mylonite zone formed 

by injection and flowage of crystal mush. In any case, the 

high temperature conditions implied by these workers and 

the subsequent rapid cooling after uplift can explain the 

51- to 44-Ma K-Ar mineral ages of these rocks. 

CONCLUSIONS 

North-central Idaho is composed of two essentially 

different terranes, each having distinct geologic and 

geochemical characteristics that are reflected by igneous 

intrusions. The western or Wallowa-Seven Devils terrane 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 129 

TABLE 6.5.-40Ar/ 39 Ar ages of rocks of the Bitterroot lobe of the Idaho batholith 

[Sample locations are given in table 6.1. 40 Ar refers to radiogenic argon. TF, total fusion of separate aliquant of sample; +, indicates temperature step omitted from plateau; 

J, measure of the integrated fast neutron flux determined using a monitor biotite, SB-3] 

Sample 

797-25F------

797-25F( IH)--

797-25I------

797-25l(IH)--

797-26A------

797-26A(IH)--

818-lOB(IH)--

827-28G------

827-28G------

Temperature 

step 

( oC) Mineral 

Percent 

of total 

39 Ar 

Percent 

40Ar 

Percent 

36Arca 40Ar/39Ar 

Plutonic rocks of the Wallowa-Seven Devils terrane 

TF Hornblende 100 

750+ do. 6.53 

850+d do. 11.43 

900 + do. 19 .48 

930 do. 15.89 

960 do. 5.78 

1000 do. 6.73 

1040 do. 29.25 

1100 do. 4.33 

Fusion do. .588 

TF 

750 + 

850+ 

900 

930 

960 

1000 

1040 

1100 

Fusion 

Hornblende 100 

do. .63 

do. 4.33 

do. 43.78 

do. 14.12 

do. 

do. 

do. 

do. 

do. 

3.38 

5.53 

7.89 

13.97 

6.37 

TF Hornblende 100 

750+ do. 5.17 

900 do. 57.01 

940 do. 9.22 

960 do. 6.26 

1000 do. 12.25 

1050 do. 8.44 

1100+ do. 1.24 

Fusion+ do. .40 

750+ Hornblende 3.76 

850+ do. 8.59 

900 do. 19.56 

925 do. 34.75 

950 do. 9.85 

975 do. 3.93 

1000 do. 5.60 

1025 do. 7.66 

Fusion do. 6.32 

TF 

750+ 

850+ 

900 

925 

950 

975 

1000 

1025 

Fusion 

Hornblende 100 

do. 3. 71 

do. 4.62 

do. 57.64 

do. 13.53 

do. 

do. 

do. 

do. 

do. 

2.46 

5. 59 

6.30 

5.47 

.67 

71.0 

22.0 

36.1 

83.3 

87.4 

89 .2 

82.6 

95.7 

93.5 

46.4 

52.4 

22. 7 

35.7 

81.8 

84.8 

56.7 

74.3 

69.6 

81.3 

70.8 

63.7 

.01 

42.3 

47.9 

41.1 

52.3 

44.8 

20.7 

9.1 

8.0 

53.3 

62.0 

88.7 

92.9 

91.5 

93.6 

89.6 

90.3 

75.4 

20.7 

13.3 

58.8 

95.6 

70.7 

91.7 

93.3 

92.6 

46.0 

10.11 

.36 

3.52 

21.68 

24.29 

26.82 

19.50 

51 .68 

40.05 

4.29 

7.04 

.1 7 

2.35 

23.30 

27.01 

9.02 

16.37 

lJ .64 

23.34 

14.47 

15.42 

52.00 

21.88 

12.02 

12.88 

12.69 

13.35 

11 .86 

12.46 

23.13 

28.42 

690.8 

87.92 

17.39 

17. 7l 

23.96 

21.18 

21.85 

17.78 

20.20 

12.48 24 .o1 

.13 800.4 

5.78 35.09 

7.28 31 .46 

4. 90 42.54 

8.08 30.46 

6.26 34.86 

1.37 119.9 

.24 606.3 

.4 

6.3 

8.2 

29.7 

42.4 

37.8 

45.0 

32.1 

34.1 

10.32 

.34 

.47 

5.38 

46.37 

9.01 

32.24 

37 .1 7 

38.65 

5.39 

90.24 

17.87 

15.82 

11 .06 

10.52 

10.74 

10.53 

11.00 

10.88 

14.05 

59.82 

97.59 

17.79 

10.75 

14.88 

11.29 

11.18 

11 .15 

21 .66 

6.259 

1.845 

6.335 

6.895 

6.470 

6.230 

6.989 

6.668 

6.637 

6.910 

12.74 

ll. 55 

16.96 

12.59 

12.37 

12.78 

13.25 

l3 .04 

12.59 

12.465 

15.45. 

12.72 

15.47 

15.98 

16.06 

15.90 

15.99 

16.46 

16.75 

4.593 

7.009 

6. 729 

6.565 

6. 733 

6.853 

6. 774 

6.698 

6. 762 

4.946 

2.008 

4.997 

5.172 

4.987 

5.366 

5.537 

5.503 

6.398 

8.143 

0.0168 

.1378 

.0490 

.0087 

.0072 

.0063 

.0097 

.0035 

.0045 

.0438 

J 

0.0066059 

.0066059 

.0066059 

.0066059 

.0066059 

.0066059 

.0066059 

.0066059 

.0066059 

.0066059 

Age 

(Ma) 

126.3±1.5 

131.4±3 

92.1±1.8 

116 .0±1.5 

129.9±1.7 

130.6±3.3 

127.4±3.0 

130.9±1.4 

134.3±4.3 

124.2±31.1 

Recalculated total----- 123.4 

Plateau age------------ 130.4±5.3 

0.0492 

1.810 

.1960 

.0141 

.0124 

.0385 

.0220 

.0260 

.0147 

.0234 

0.0066059 

.0066059 

.0066059 

.0066059 

.0066059 

.0066059 

.0066059 

.0066059 

.0066059 

.0066059 

170.6±2.2 

1289. 7±131 .l 

343.2±11.6 

163.4±2.0 

171.9±3.9 

156.3±15.1 

179.8±9'.3 

174.1 ±6. 7 

165.8±4.0 

164.7±8.3 

Recalculated total----- 184.4 

Plateau age------------ 169.9±3.6 

0.0337 

2. 712 

.0728 

.0598 

.0891 

.05351 

.06947 

.3260 

1.869 

0.005898 

.005898 

.005898 

.005898 

.005898 

.005898 

.005898 

.005898 

.005898 

157.3±2.7 

0.9±108 

152.7±3.4 

155.2±5.9 

178.8±8.8 

163.4±5.4 

160.5±6.7 

249.2±30.9 

514.9±238.3 

Recalculated total----- 151.9 

Plateau age------------ 157.9±2.8 

.2821 

.0301 

.0222 

.0060 

.0043 

.0049 

.0041 

.005 7 

.0054 

0.007852 

.007852 

.007852 

.007852 

.007852 

.007852 

.007852 

.007852 

.007852 

100.0±2.5 

130.6±1.4 

134.3±1.3 

134.4±1.3 

133. 9±1 .4 

134.6±1.6 

135.0±1.4 

134.4±1.4 

134.6±1.4 

Recalculated total----- 132.8 

Plateau age------------ 134.4±0.5 

0.0130 

.01610 

.2883 

.0262 

.0029 

.0162 

.0047 

.0040 

.00465 

.0411 

0.007851 

.007851 

.007851 

.007851 

.007851 

.007851 

.007851 

.007851 

.007851 

.007851 

144.6±2.1 

167.7±2.3 

173.0±2.9 

143.0±1.4 

140.5±1.4 

143.7±2.2 

141.5±1.6 

142.4±1.5 

141.1±1.6 

139 .2±6 .1 

Recalculated total----- 144.7 

Plateau age------------ 141.9±1.0 
130 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

TABLE 6.5.-40Ar/ 39Ar ages of rocks of the Bitterroot lobe of the Idaho batholith-Continued 

[Sample locations are given in table 6.1. 40 Ar refers to radiogenic argon. TF, total fusion of separate aliquant of sample; +, indicates temperature step omitted from plateau; 

J, measure of the integrated fast neutron flux determined using a monitor biotite, SB·3] 

Temperature Percent 

step of tota 1 Percent Percent Age 

Sample (oC) Mineral 39 Ar 40Ar 36Ar 40 Ar ;39 Ar 37 Ar/39Ar 36 Ar ;39 Ar J (Ma) Ca 

Plutonic rocks of the transition zone 

797-26D------ TF Muscovite 100 94.7 .03 7.892 0.0013 .0014 0.0066059 86.9±0.9 

797-261( 1H)-- 750+ Hornblende .59 .4 .04 789.2 3.523 2.661 .005898 35.7±114.2 

900 do. 35.92 34.3 1. 78 21.80 3.232 .0493 .005898 78.0±2.0 

940 do. 38.59 44.0 2.55 17.59 3.204 .0342 .005898 80. 7±1.8 

960 do. 7.24 38.7 2.19 19.88 3.393 .0421 .005898 80.2±2.4 

1000 do. 16.55 49.4 3.16 15.96 3.272 .0282 .005898 82.2±1.7 

1050+ do. .80 5.6 .17 164.8 3.248 .5272 .005898 96.2±24.7 

1100+ do. .26 5.4 .20 142.8 3.293 .4582 .005898 80.2±46.5 

Fusion + do. .04 4.7 .03 1696 6.036 5.468 .005898 703.5±914.6 

Recalculated total----- 80.1 

Plateau age------------ 80.5±1.0 

807-281------ TF Hornblende 100 64.6 7.93 11.53 4.369 0.0150 0.006233 82.1±1.1 

TF do. 100 65.0 9.9 9.110 4.350 .0120 .007852 82.1±1.4 

807-281( 1H)-- 750+ do. 5.26 24.6 1.6 26.31 4.083 .0682 .007852 89 .8±1.1 

850 do. 29.88 70.9 13.4 8.125 4.553 .0092 .007852 80.0±0.8 

900 do. 28.61 90.2 36.9 6.448 4.531 .0033 .007852 80.8±0.8 

925 do. 10.28 70.2 12.7 8.252 4.439 .0095 .007852 80.4±0.8 

950 do. 17.00 95.7 57.8 6.126 4.418 .0021 .007852 81.4±0 .8 

975 do. 7.61 91.0 37.7 6.448 4.345 .0031 .007852 81.4±0.9 

1000+ do. .42 41.8 4.4 13.40 4.454 .0276 .007852 77 .8±5 .8 

1025 + do. .57 35.6 3.4 15.86 4.506 .0358 .007852 78.4±4.3 

Fusion + do. .36 16.6 .98 44.06 4.519 .0156 .007852 100.7±6.9 

Recalculated total----- 81.2 

Plateau age------------ 80.8±0.4 

Plutonic rocks of the Bitterroot lobe, undivided 

818-16F------ TF Hornblende 100 39.6 2.73 24.27 5.120 .0510 0.0050695 86.6±1.1 

818-16F(1H)-- 750+ Hornblende .61 1.6 .01 653.1 1.031 2.174 .0050695 95.1±81.1 

840+ do. .43 3.0 .02 749.4 2.202 2.461 .0050695 194.0±125.0 

890+ do. .20 2.5 .02 1241 3.106 4.098 .0050695 261.8±414.9 

930+ do. .32 .4 .02 1779 3.632 6.043 .0050695 -59.2±448.3 

960+ do. .64 3.3 .22 162.8 4.371 .5342 .0050695 47.9±31.2 

1000 do. 15.85 59.7 6.31 15.17 5.119 .0220 .0050695 81.2±1.3 

1040 do. 61.43 83.7 19.23 1.87 5.243 .0074 .0050695 81.5±0 .9 

1100 do. 20.33 74.1 11 .53 12.75 5.352 .0126 .0050695 84 .6±1.2 

Fusion + do. .19 7.6 .09 475.5 5.145 1.1488 .0050695 305.6±176.0 

Recalculated total----- 82.9 

Plateau age------------ 82.3±0.6 

787-19A------ TF Hornblende 100 70.9 11.41 7.572 3.521 0.0084 0.006930 66.0±1.0 

787-19A(1H)-- 750+ do. 2.74 19.8 .30 68.74 2.044 .1872 .00693 162.6±8.1 

890+ do. 12.40 21.3 2.17 15.73 3.418 .0428 .00693 41.5±2.4 

925 do. 14.41 55.7 5.97 9.988 3.489 .0159 .00693 68.4±1.3 

960 do. 16.45 55.6 6.21 9.717 3.547 .0156 .00693 66. 5±1.2 

1000 do. 15.73 50.5 5.48 10.03 3.575 .0177 .00693 62 .4±1.3 

1075 do. 28.64 70.3 11 .41 7.450 3.538 .0084 .00693 64.5±1.0 

Fusion do. 9.63 48.4 4.83 11 .12 3.169 .0204 .00693 66.2±1.6 

Recalculated total----- 65.1 

Plateau age------------ 65.4±0.6 

787-20A------ TF Hornblende 100 55.7 4.76 10.97 3.020 .0173 0.005723 62.1±1.2 

787-20E------ TF Biotite 100 79.1 .02 5.360 .0029 .0038 .006979 52.6±0.8 

78 7-24H( IH) -- 750+ Hornblende 3.55 7.8 .03 127.9 .4398 .3992 .0062025 107.9±14.8 

755+ do. .3 7 7.0 .or 565.7 .4838 1.782 .0062025 394.±149 

835+ do. 1.12 4.6 .11 154.7 2.053 .5002 .0062025 77.6±24.8 

900 do. 3.52 22.3 1.29 26.52 3.346 .0706 .0062025 65.3±7.2 

960 do. 14.24 48.1 5.96 9. 777 3.995 .0183 .0062025 52.0±2.0 

1029 do. 36.74 71.3 14.81 6.174 3.816 .0070 .0062025 48. 7±1.0 

1100 do. 38.09 80.3 22.16 5.496 3.812 .0047 .0062025 48.8±0.8 

Fusion + do. 2.36 20.8 1.02 38.43 3.915 .1041 .0062025 87 .5±10 .8 

Recalculated total----- 54.6 

Plateau age------------ 49.2±1.9 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 131 

TABLE 6.5.-4°Ar/ 39Ar ages of rocks ojthe Bitterroot lobe of the Idaho batholith-Continued 

[Sample locations are given in table 6.1. 40 Ar refers to radiogenic argon. TF, total fusion of separate aliquant of sample; +, indicates temperature step omitted from plateau; 

J, measure of the integrated fast neutron flux determined using a monitor biotite, SB-3] 

Temperature Percent 

step of total Percent Percent Age 

Sample (oC) Mineral 39Ar 40Ar 36Arca 40 Ar ;39 Ar 37Ar/39Ar 36Ar/39Ar J (Ma) 

Plutonic rocks of the Bitterroot lobe, undivided--Continued 

787-26A------ TF Hornblende 100 17 .o .96 36.34 3.617 0.0103 0.005723 62.8±3.8 

787-26F------ TF do. 100 77.6 15.25 8.120 4.060 .0072 .005 723 64 .1±1.0 

787-26F(IH)-- 750+ do. 2. 73 33.1 .75 40.01 2.501 .0913 .005723 131. 9±8 .4 

890+ do. 1.34 23.5 .94 28.64 2.576 .0749 .005 723 68.2±7.5 

925+ do. 2.43 43.9 3.59 15.60 4.048 .0307 .005 723 69. 6±4 .8 

960 do. 12.66 85.3 23.30 7.396 4.088 .0048 .005723 64 .1±1.1 

1000 do. 18.90 90.6 34.09 6.678 4.019 .0032 .005 723 61.5±0.9 

1050 do. 32.27 92.6 41.25 6.425 4.078 .0027 .005 723 60.6±0.8 

Fusion do. 28.67 88.4 28 .8·2 7.163 4.160 .0039 .005723 64.4±0.9 

Recalculated total----- 64.6 

Plateau age------------ 62.4±0.5 

80MT246------ TF Hornblende 100 31.0 3. 72 12.95 4.292 .0314 0.006287 47.4±0.6 

TF do. 100 52.1 8.16 8.086 4.271 .0142 .0066124 47.3±0.6 

80MT246( IH) -- 750+ Hornblende 1.52 3.9 .08 143.3 1.448 .4664 .0066124 65.2±11.4 

820+ do. 4.21 3.4 .35 85.72 3.648 .2813 .0066124 34.2±3.9 

875 do. 12.10 18.2 1.85 22.52 4.320 .0635 .0066124 42.5±1.4 

950 do. 26.64 54.0 9.57 7.278 4.404 .0125 .0066124 46.4±0.7 

1050 do. 53.86 61.8 12.73 6.464 4.470 .00956 .0066124 47 .2±0 .6 

Fusion + do. 1.67 11.2 1.02 40.54 4.613 .1231 .0066124 '53.4±8.9 

Recalculated total----- 46.9 

Plateau age------------ 47.0±0.5 

Eocene epizonal plutons and associated volcanic rocks 

787-l8J------ TF Biotite 100 74.7 .10 5.179 0.0164 0.0044 0.006979 48 .1±0. 7 

787-25A------ TF Hornblende 100 72.4 13.85 6. 792 3.736 .0073 .006930 50.2±0.7 

TF do. 100 56.0 8.50 7.423 3.767 .0121 .005723 51.3±1.0 

787-25B------ TF Biotite 100 36.4 4.03 11.35 3.767 .0254 .006930 51.1±1.4 

797-l9C------ TF Sanidine 100 95.2 .35 4.89 .0100 .0078 .005898 48.9±0.6 

807-29E------ TF Sanidine 100 42.3 1.60 8. 794 1.023 .0174 .0050695 34.0±0.7 

(dike) TF do. 100 9.1 .21 44.12 1.042 .1360 .004 771 34 .2±1.1 

Metamorphosed Precambrian rocks 

807-31K------ TF Hornblende 100 18.2 1.27 41.32 5.391 .1159 0.006233 82.7±1.3 

808-01G------ TF Muscovite 100 87.2 .06 9.412 .0090 .0039 .004771 69.8±0.7 

is devoid of Precambrian rocks and probably represents 

an accreted island-arc complex. Jurassic and Cretaceous 

granitic intrusions reflect the character of this zone in 

their low ri values (~0. 7042), low Rb/Sr ratios (~0.08), 

d18Q values of + 5.5 to + 10.5 per mil, dD values of -41 

to - 82 per mil, and in their calcic and metaluminous 

major-element chemistry. The eastern or Belt-Yellowjacket terrane contains abundant Precambrian rocks and 

is intruded by both Cretaceous-early Tertiary main-phase 

granitoids and by Eocene epizonal plutons. The mainphase plutons reflect their Precambrian source region in 

their high ri values (~0. 708), Rb/Sr ratios (0.04 to 0.24), 

high primary d18Q values ( + 8.0 to + 12.4 per mil), and 

in their calcic to calc-alkaline, metaluminous to peraluminous chemistry. The Eocene epizonal plutons all have 

ri values of ~0.704, suggesting the presence of an ancient 

component, but several values between 0.704 and 0.708 

indicate limited wall-rock involvement. Rb/Sr ratios of 

these epizonal plutons are very large (as much as 49.0), 

probably due to differentiation. These calc-alkaline plutons 

are chemically similar to the Challis Volcanics. 

Extensive hydrothermal circulation systems developed 

in the region during both Cretaceous-early Tertiary and 

Eocene times. The Cretaceous-early Tertiary hydrothermal activity was most pronounced near the perimeter of 

the Bitterroot lobe, where most of the ore deposits in the 

region are located. This deep hydrothermal system was 

probably contemporaneous with regional metamorphism 
132 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

of the Precambrian wall rock. Major shallow hydrothermal 

systems involving fluids derived from low-180 meteoric 

waters developed during the emplacement of the Eocene 

epizonal plutons. Water-rock interactions produced large 

reductions in the dD values (as low as -148 per mil) and 

d18Q values (as low as -0.3 per mil) of the granitic rocks 

in the affected areas. Only a few ore deposits in the Bitterroot lobe appear to be related to these Eocene hydrothermal systems, however. The hydrothermal systems did not 

greatly disturb the Rb-Sr systematics of Bitterroot lobe 

plutons, partly because large intervals do not separate the 

times of emplacement and alteration, but the isochrons 

of Precambrian rocks appear to have been profoundly 

modified by later hydrothermal metamorphism. 

The K-Ar apparent ages of hornblende range from ~ 160 

Ma for some Jurassic plutons in the Wallowa-Seven Devils 

terrane to as low as 44 Ma for Eocene granitic rocks. 

Some low-ri plutons of the Wallowa-Seven Devils terrane 

have concordant biotite and hornblende ages of 135 Ma 

and show no geochemical indication of proximity to the 

Precambrian craton. Transition zone plutons having K-Ar 

mineral ages of 87 to 75 Ma intrude both the WallowaSeven Devils terrane and the Precambrian sialic rocks. 

These results indicate that the tectonic suture between 

these terranes formed between approximately 135 and 

75 Ma, probably between 95 and 80 Ma. Most of the 

Bitterroot lobe was emplaced after this probable accretionary event. However, the hornblende and biotite K-Ar 

1160 1150 1140 

47°r---~r-----------------------~--------------------~~~----------------~----r---------, 

Moscow 

n 

( 

\ 

\ 

( 

,f 

< 

.. 144 

46° -~~~2.~~~~\ )-

OREGON \. 

'· EXPLANATION \... '· 

·..__ • Data from tables 6.4.6.5 ... '. 

.6. Data from Armstrong (1975) l. 

[47] 4 uAr/ 39Ar plateau age \._ 

4oA f39 ) .±Z. r Ar total fusion age /.• 

4 7 Conventional K-Ar age r' 

( ~0 MILES / 

0 10 KILOMETERS 1 

/ 

/140 

) ~30 / 

..!2.!.. 1921 

GrangPvil~~ 

n 

[R1l, 80 

tl' 

[81]. _!g, 79 

• 80 

........ 

. '-.. 

l MONTANA 

\ ~~ .. 

<.. ....... 

( 

i..,.... .. 

) 

\ 

'! 

t 

\ 

'""·. 

.

/ 45°L-----~--L-------------~----~--~~~-----------------L------L-----------------~~----~ 

FIGURE 6.18.-Contour map of K-Ar and 40 Ar/ 39 Ar apparent ages of hornblende. Most of these apparent ages are thought to be within 10 m.y. 

of the emplacement ages and indicate that magmatism migrated progressively toward the east-central part of the Bitterroot lobe. The ages 

of the transition-2one plutons, which intrude the suture zone between the Precambrian craton and the accreted terrane, provide an absolute 

lower limit of 75 Ma for the accretionary event. 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 133 

apparent ages become systematically younger toward the 

mylonitic front of the Bitterroot lobe, with biotite data 

defining a 10,000-km2 area of apparent ages of 51 to 44 

Ma. The detailed relations of hornblende and biotite K-Ar 

200 

apparent ages in the Bitterroot lobe are thought to reflect 

temporal migration of magmatism as well as rapid uplift 

and cooling of the region. 

EXPLANATION 

180 I o Biotite K-Ar or 40 ArP9 Ar total fusion age 

I 

I 

I t::,. Muscovite K-Ar or 40 ArP9 Ar total fusion age 

~ 160 

-- I Cl:l 

~ I • Hornblende K-Ar age 

ro I <1) 

~ 140 ------------r----j ~ 

0 

120 ~!-------------! : Cl:l 

~ 

0 

·~ ~ 

~ 

·~ 

100~ ...._, s 

~ 

Cj 

~ 80~ ---•---r 

I 

60 I 

I 

I 

40~ I 

I 

WSD I 

I 

-e- Hornblende 40 ArP9 Ar plateau age 

+ Concordant biotite-hornblende K-Ar age 

• Concordant biotite-muscovite K-Ar age 

CD ------,--------

--~ -----~?---

Belt - Yellowjacket 

~------l-- ------------11----------

0 ___ 2,-~6-----~-~-

I 20~----~----~------~----~----~------~----~----~----~ 

-40 -20 0 20 40 60 80 100 120 140 

DISTANCE (kilometers from terrane boundary) 

FIGURE 6.19.-K-Ar and 40Ar/ 39Ar ages of plutonic rocks plotted 

against their distance from the WSD/Belt-Y ellowjacket terrane 

boundary. Vertical tielines connect ages determined from the same 

sample. Plutons yielding Jurassic and Early Cretaceous ages are 

confined to the WSD terrane. This abrupt termination at the terrane boundary suggests that the Jurassic-Early Cretaceous 

magmatic arc was truncated tectonically and juxtaposed with 

Precambrian sialic crust subsequent to 135 Ma. Ages determined 

on hornblende and muscovite from transition-zone plutons, which 

intrude both terranes in the 5- to 25-km zone on either side of the> 

suture, range from 87 to 75 Ma. Within the Belt-Y ellowjacket te"-

rane, two early Tertiary plutonic episodes are defined by rather 

abrupt steps in hornblende ages: 62 to 66 Ma and 46 to 51 Ma. Notf• 

that no hornblende-bearing samples were found in the central par! 

of the Bitterroot lobe (55 to 90 km from the terrane boundary) 

Biotite ages exhibit a smooth decrease in age normal to the terrane 

boundary, reflecting uplift and (or) argon loss due to younger i::-: 

trusive events. 
134 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

C/) 

a: 

<( 

w 

>-

z 

0 

.....J 

.....J 

~ 

z 

w 

(!) 

<( 

.... 

<( 

CJ) 

C') 

....... 

.... 

<( 

0 

'lit 

150 

140 -

130 - I 

"' 

120 -

110 -

100 -

90 ~ 

80 r--.l-

~ 

70 

0 

I 

"' I 

I 

I 

_l 

I I I I r I I I 

827-28G <rt 0.7034) 

I ~ 

818-108 (r= 0.7034) 

- I 

-

-

-

-

818-16F (r=0.7073) 

J 797-261 (r.= 0.7078) 

J I 807-281 <r.=0-7073) ~ 

I I I I I I I I 

20 40 60 80 100 

PERCENT 39Ar RELEASED 

FIGURE 6.20.-40 Ar/ 39 Ar age spectra of hornblende from granitic rocks 

within and near the tectonic suture between the Wallowa-Seven Devils 

terrane and the Precambrian craton. Samples 827-28G and 818-10B 

have flat Early Cretaceous age spectra, concordant or nearly concordant biotite and hornblende K-Ar ages, and low initial strontium (ri) 

ratios that give no indication of assimilation of Precambrian rocks. 

REFERENCES CITED 

Anderson, A.L., 1951, Metallogenic epochs in Idaho: Economic Geology, 

v. 46, p. 592-607. 

Armstrong, R.L., 197 4, Geochronometry of the Eocene volcanic-plutonic 

episode in Idaho: Northwest Geology, v. 3, p. 1-15. 

__ 1975, The geochronometry of Idaho: Isochron!West, v.14, p. 1-50. 

Armstrong, R.L., Taubeneck, W.H., and Hales, P.O., 1977, Rb-Sr and 

K-Ar geochronometry of Mesozoic granitic rocks and their Sr isotopic 

composition, Oregon, Washington, and Idaho: Geological Society 

of America Bulletin, v. 88, 397-411. 

Ave Lallemant, H.G., Phelps, D.W., and Sutter, J.F., 1980, 40Ar-39Ar 

ages of some pre-Tertiary plutonic and metamorphic rocks of eastern 

Oregon and their geologic relationships: Geology, v. 8, p. 371-374. 

Samples 818-16F, 807-28I, and 797-26I have flat Late Cretaceous age 

spectra, nearly concordant biotite and hornblende K-Ar ages, and 

higher initial strontium ratios that indicate involvement of both Precambrian and arc-derived materials in the magmas. Taken together, 

these data establish that the arc-continent suture zone formed between 

135 and 80 Ma. 

Axelrod, D.I., 1966, Potassium-argon ages of some western Tertiary 

floras: American Journal of Science, v. 264, p. 497-506. 

Bateman, P.C., and Dodge, F.C.W., 1970, Variations of major chemical 

constituents across the central Sierra Nevada batholith: Geological 

Society of America Bulletin, v. 81, p. 409-420. 

Bennett, E.H., 1980, Granitic rocks of Tertiary age in the Idaho batholith 

and their relation to mineralization: Economic Geology, v. 75, 

278-288. 

Bergendahl, M.H., 1964, Gold, in Mineral and water resources of Idaho: 

U.S. Geological Survey Special Report No. 1, p. 93-101. 

Bickford, M.E., Chase, R.B., Nelson, B.K., Shuster, R.D., and Arruda, 

E.C., 1981, U-Pb studies of zircon cores and overgrowths, and 

monazite: Implications for age and petrogenesis of the northeastern 

Idaho batholith: Journal of Geology, v. 89, p. 433-457. 
6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 135 

Bond, J.G., 1978, Geologic map of Idaho: Idaho Bureau of Mines and 

Geology. 

Borthwick, J., and Harmon, R.S., 1982, A note regarding ClF3 as an 

alternative to BrF5 for oxygen isotope analysis: Geochimica et 

Cosmochimica Acta, v. 46, p. 1665-1668. 

Cater, F.W., Pinckney, D.M., Hamilton, W.B., Parker, R.L., Weldin, 

R.D., Close, T.J., Zilca, N.T., Leonard, B.F., and Davis, W.E., 1973, 

Mineral resources of the Idaho primitive area and vicinity, Idaho: 

U.S. Geological Survey Bulletin 1304, 431 pp. 

Chase, R.B., 1973, Petrology of the northeastern border zone of the 

Idaho batholith, Bitterroot Range, Montana: Montana Bureau of 

Mines and Geology Memoir 43, 28 p. 

Chase, R.B., Bickford, M.E., and Arruda, E.C., 1983, Tectonic implications of Tertiary intrusion and shearing within the Bitterroot dome, 

northeastern Idaho batholith: Journal of Geology, v. 91, p. 462-470. 

Chase, R.B., Bickford, M.E., and Tripp, S.E., 1978, Rb-Sr and U-Pb 

isotopic studies of the northeastern Idaho batholith and border zone: 

Geological Society of America Bulletin, v. 89, p. 1325-1334. 

Chase, R.B., and Johnson, B.R., 1977, Border-zone relationships ofthe 

northern Idaho batholith: Northwest Geology, v. 6, no. 1, p. 38-50. 

Chase, R.B., and Talbot, J.L., 1973, Structural evolution of the northeastern border zone of the Idaho batholith, western Montana [abs.]: 

Geological Society of America Abstracts with Programs, v. 5, 

p. 470-471. 

Clayton, R.N., and Mayeda, T.K., 1963, The use of bromine pentafluoride 

in the extraction of oxygen from oxide and silicates for isotopic 

analysis: Geochimica et Cosmochimica Acta, v. 27, p. 43-52. 

Cox, A., and Dalrymple, G.B., 1967, Statistical analysis of geomagnetic 

reversal data and the precision of potassium-argon dating: Journal 

of Geophysical Research, v. 72, p. 2603-2614. 

Craig, H., 1961, Standard for reporting concentrations of deuterium 

and oxygen-18 in natural waters: Science, v. 133, 1833-1834. 

Criss, R.E., 1981, An 180/ 160, D/H and K-Ar study of the southern half 

of the Idaho batholith: Pasadena, California Institute of Technology, 

Ph.D. dissertation, 401 pp. 

Criss, R.E., and Champion, D.E., 1984, Magnetic properties of granitic 

rocks from the southern half of the Idaho batholith: influences of 

hydrothermal alteration and implications for aeromagnetic interpretation: Journal of Geophysical Research, v. 89, p. 7061-7076. 

Criss, R.E., Champion, D.E., and Mcintyre, D.H., 1985, o180, aeromagnetic and gravity anomalies associated with hydrothermally 

altered zones in the Yankee Fork mining district, Custer County, 

Idaho: Economic Geology, v. 80, p. 1277-1296. 

Criss, R.E., Ekren, E.B., and Hardyman, R.F., 1984, Casto ring zone: 

a 4,500-km2 fossil hydrothermal system in the Challis volcanic field, 

central Idaho: Geology, v. 12, p. 331-334. 

Criss, R.E., Lanphere, M.A., and Taylor, H.P., Jr., 1982, Effects of 

regional uplift, deformation, and meteoric-hydrothermal metamorphism on K-Ar ages of biotites in the southern half of the Idaho 

batholith: Journal of Geophysical Research, v. 87, p. 7029-7046. 

Criss, R.E., and Taylor, H.P., Jr., 1983, An 180/ 160 and D/H study of 

Tertiary hydrothermal systems in the southern half of the Idaho 

batholith: Geological Society of America Bulletin, v. 94, p. 640-663. 

Dalrymple, G.B., and Lanphere, M.A., 1969, Potassium-argon dating, 

principles, techniques, and applications to geochronology: San Francisco, California, W.H. Freeman, 258p. 

__ 1971, 40 Ar/ 39 Ar technique of K-Ar dating: a comparison with the 

conventional technique: Earth and Planetary Science Letters, v. 12, 

p. 300-308. 

Dontsova, Ye I., 1970, Oxygen isotope exchange in rock-forming processes: Geochemistry International, v. 7, p. 624-636. 

Ferguson, J.A., 1975, Tectonic implications of some geochronometric 

data from the northeastern border zone of the Idaho batholith: 

Northwest Geology, v. 4, p. 53-58. 

Finch, J.W., 1932, Prospecting for gold ores: Idaho Bureau of Mines 

and Geology Pamphlet, v. 36, p. 1-33. 

Fleck, R.J. and Criss, R.E., 1985, Strontium and oxygen isotopic variations in Mesozoic and Tertiary plutons of central Idaho: Contributions to Mineralogy and Petrology, v, 90, p. 291-308. 

Friedman, I., 1953, Deuterium content of natural waters and other 

substances: Geochimica et Cosmochimica Acta, v. 4, p. 89-103. 

Garlick, G.D., and Epstein, S., 1967, Oxygen isotope ratios in coexisting 

minerals of regionally metamorphosed rocks: Geochimica et Cosmochimica Acta, v. 31, p. 181-214. 

Garmezy, L., and Greenwood, W.R., 1981, Mylonitization coincident with 

batholith emplacement: Bitterroot lobe, Idaho batholith [ abs ]: 

Geological Society of America Abstracts with Programs, v. 13, 

p. 197. 

Garmezy, L., and Sutter, J.F., 1983, Mylonization coincident with uplift 

in an extensional setting, Bitterroot Range, Montana-Idaho [abs.]: 

Geological Society of America Abstracts with Programs, v. 15, 

p. 578. 

Godfrey, J.D., 1962, The deuterium content of hydrous minerals from 

the east-central Sierra Nevada and Yosemite National Park: Geochimica et Cosmochimica Acta, v. 26, p. 1215-1245. 

Grauert, B., and Hofmann, A., 1973, Old radiogenic lead components 

in zircons from the Idaho batholith and its metasedimentary aureole: 

Carnegie Institution of Washington Year Book, v. 72, p. 297-299. 

Greenwood, W.R., and Morrison, D.A., 1973, Reconnaissance geology 

of the Selway-Bitterroot Wilderness Area: Idaho Bureau of Mines 

and Geology Pamphlet, v. 154, 30 p. 

Hamilton, W.B., 1963, Metamorphism in the Riggins region, western 

Idaho: U.S. Geological Survey Professional Paper 436, 95p. 

Hamilton, W., 1976, Tectonic history of west-central Idaho [abs.]: 

Geological Society of America Abstracts with Programs, v. 8, p. 378. 

Harms, T.A., and Price, R.A., 1983, The Newport fault: Eocene crustal 

stretching, necking and listric normal faulting in NE Washington 

and NW Idaho [abs.]: Geological Society of America Abstracts with 

Programs, v. 15, p. 309. 

Hietanen, Anna, 1962, Metasomatic metamorphism in western Clearwater County, Idaho: U.S. Geological Survey Professional Paper 

344-A, 116 p. 

__ 1963, Idaho batholith near Pierce and Bungalow, Clearwater 

County, Idhao: U.S. Geological Survey Professional Paper 344-D, 

p. 1-42. 

__ 1984, Geology along the northwest border zone of the Idaho 

batholith, northern Idaho: U.S. Geological Survey Bulletin 1608, 

17 p. 

Hoernes, S., and Friedrichsen, H., 1980, Oxygen and hydrogen isotope 

composition of Alpine and pre-Alpine minerals of the Swiss central 

Alps: Contributions to Mineralogy and Petrology, v. 72, p. 19-32. 

Hyndman, D.W., 1980, Bitterroot dome-Sapphire tectonic block, an 

example of a plutonic gneiss-dome complex with its detached suprastructure, in Coney, P.J., Crittenden, Max, Jr., and Davis, G.H., 

eds., Cordilleran metamorphic core complexes: Geological Society 

of America Memoir 153, p. 427-443. 

Hyndman, D.W., 1981, Controls on source and depth of emplacement 

of granitic magma: Geology, v. 9, p. 244-249. 

__ 1983, The Idaho batholith and associated plutons, Idaho and 

western Montana: Geological Society of America Memoir 159, 

p. 213-240. 

__ 1984, A petrographic and chemical section through the northern 

Idaho batholith: Journal of Geology, v. 92, p. 83-102. 

Hyndman, D.W., Talbot, J.L., and Chase, R.B., 1975, Boulder batholith: 

a result of emplacement of a block detached from the Idaho batholith 

infrastructure?: Geology, v. 3, p. 401-404. 

Hyndman, D.W., and Williams, L.D., 1977, The Bitterroot lobe of the 

Idaho batholith: Northwest Geology, v. 6., no. 1, p. 1-16. 

Jager, E., 1962, Rb-Sr age determination on micas and total rocks from 

the Alps: Journal of Geophysical Research, v. 67, p. 5293-5306. 
136 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

Jones, D.L., Silberling, N.J., and Hillhouse, J., 1977, Wrangellia-a 

displaced terrane in northwestern North America: Canadian Journal of Earth Sciences, v. 14, p. 2565-2577. 

Kiilsgaard, T.H., Freeman, V.L., and Coffman, J.S., 1970, Mineral 

resources of the Sawtooth Primitive Area, Idaho: U.S. Geological 

Survey Bulletin 1319D, 17 4 pp. 

Kistler, R.W., and O'Neil, J.R., 1975, Fossil geothermal gradients in 

crystalline rocks of the Ruby Mountains, Nevada as indicated by 

radiogenic and stable isotopes [abs.]: Geological Society of America 

Abstracts with Programs, v. 7, p. 334-335. 

Kistler, R.W., and Peterman, Z.E., 1973, Variations in Sr, Rb, K, Na, 

and initial 87Sr/ 86Sr in Mesozoic granitic rocks and intruded wall 

rocks in central California: Geological Society of America Bulletin, 

v. 84, p. 3489-3512. 

Krummenacher, Daniel, Gastil, R.G., Bushee, J., and Doupont, J., 1975, 

K-Ar apparent ages, Peninsular Ranges batholith, southern California and Baja California: Geological Society of America Bulletin, v. 86, 

p. 760-768. 

Lang, H.M., and Rice, J.M., 1985, Metamorphism of pelitic rocks in the 

Snow Peak area, northern Idaho: sequence of events and regional 

implications: Geological Society of America Bulletin, v. 96, 

p. 731-736. 

Langton, C.M., 1935, Geology of the northeastern part of the Idaho 

batholith and adjacent region in Montana: Journal of Geology, v. 43, 

p. 27-60. 

Larsen, E.S., Jr., and Schmidt, R.G., 1958, A reconnaissance of the Idaho 

batholith and comparison with the southern California batholith: U.S. 

Geological Survey Bulletin 1070-A, p. 1-33. 

Lindgren, Waldemar, 1904, A geologic reconnaissance across the Bitterroot Range and Clearwater Mountains in Montana and Idaho: U.S. 

Geological Survey Professional Paper 27, 123 p. 

Mabey, D.R., Zietz, I. Eaton, G.P., and Kleinkopf, M.D., 1978, Regional 

magnetic patterns in part of the cordillera in the Western United 

States, in Smith, R.B. and Eaton, G.P., eds., Cenozoic tectonics and 

regional geophysics of the Western Cordillera: Geological Society 

of America Memoir 152, p. 93-106. 

Magaritz, M., Whitford, D.J., and James, D.E., 1978, Oxygen isotopes 

and the origin of high 87Sr/ 86Sr andesites: Earth and Planetary 

Science Letters, v. 40, p. 220-230. 

Masi, U., O'Neil, J.R., and Kistler, R.W., 1981, Stable isotope 

systematics in Mesozoic granites of central and northern California and southwestern Oregon: Contributions to Mineralogy and 

Petrology, v. 67, p. 116-126. 

McCulloch, M.T., Gregory, R.T., Wasserburg, G.J., and Taylor, H.P., 

Jr., 1980, A neodymium, strontium, and oxygen isotopic study of 

the Cretaceous Samail ophiolite and implications for the petrogenesis 

and seawater-hydrothermal alteration of oceanic crust: Earth and 

Planetary Science Letters, v. 46, p. 201-211. 

McDowell, F.W., and Kulp, J.L., 1969, Potassium-argon dating of the 

Idaho batholith: Geological Society of America Bulletin, v. 80, 

p. 2379-2382. 

McGookey, D.P., Haun, J.D., Hale, L.A., Goodell, H.G., McCubbin, D.G., 

Weimer, J.R., and Wulf, G.R., 1972, Cretaceous System, in Mallory, 

W.M., ed., Geologic Atlas of the Rocky Mountain region: Rocky 

Mountain Association of Geologists, A.B. Hirschfeld Press, 

p. 190-228. 

Mcintyre, D.H., Ekren, E.B., and Hardyman, R.F., 1982, Stratigraphic 

and structural framework of the Challis Volcanics in the east half 

of the Challis 1 o by 2° quadrangle, in Bonnichsen, B., and Breckenridge, R., eds., Cenozoic geology of Idaho: Idaho Bureau of Mines 

and Geology Bulletin 26, p. 3-33. 

Mcintyre, D.H., Hobbs, S.W., Marvin, R.F., and Mehnert, H.H., 1976, 

Late Cretaceous and Eocene ages for hydrothermal alteration and 

mineralization, Bayhorse district and vicinity, Custer County, Idaho: 

Isochron/West, v. 16, p. 11-12. 

Miller, F.K., and Engels, J.C., 1975, Distribution and trends of discordant ages of the plutonic rocks of northeastern Washington and 

northern Idaho: Geological Society of America Bulletin, v. 86, 

p. 517-528. 

Mitchell, V.E., Strowd, W.B., Hustedde, G.S., and Bennett, E.H.,l~l;-;la. 

Mines and prospects of the Elk City quadrangle, Idaho: Idaho Bureau 

of Mines and Geology, Mines and Prospects Map Series. 

__ 1981b. Mines and prospects of the Hamilton quadrangle, Idaho: 

Idaho Bureau of Mines and Geology, Mines and Prospects Map 

Series. 

Motzer, W.E., 1985, Anorogenic granite plutons of the Selway-Bitterroot 

Wilderness (SBW), Idaho [abs.]: Geological Society of America 

Abstracts with Programs, v. 17, p. 257. 

Obradovich, J.D., and Peterman, Z.E., 1968, Geochronology of the Belt 

Series, Montana: Canadian Journal of Earth Sciences, v. 5, 

p. 737-747. 

Reid, R.R., Greenwood, W.R., and Morrison, D.A., 1970, Precambrian 

metamorphism of the Belt Supergroup in Idaho: Geological Society 

of America Bulletin, v. 81, p. 915-918. 

Reid, R.R., Morrison, D.A., and Greenwood, W.R., 1973, The Clearwater orogenic zone: a relic of Proterozoic orogeny in central and 

northern Idaho: Belt Symposium, Moscow, University of Idaho, 

Department of Geology, p. 10-56. 

Ross, C.P., 1928, Mesozoic and Tertiary granitic rocks in Idaho: Journal of Geology, v. 36, p. 673-693. 

__ 1931, A classification of the lode deposits of south-central Idaho: 

Economic Geology, v. 26, p. 169-185. 

__ 1934, Geology and ore deposits of the Casto quadrangle, Idaho: 

U.S. Geological Survey Bulletin 854, 135 pp. 

__ 1962, Stratified rocks in south-central Idaho: Idaho Bureau of 

Mines and Geology Pamphlet 125, 126 p. 

Ross, C.P., Andrews, D.A., and Witkind, I.J., 1955, Geologic map of 

Montana, scale 1:500,000: U.S. Geological Survey. 

Ryder, R.T., and Ames, H.T., 1970, Palynology and age of Beaverhead 

Formation and their paleotectonic implications in Lima region, 

Montana-Idaho: American Association of Petroleum Geologists 

Bulletin, v. 54, p. 1155-1171. 

Ryder, R.T., and Scholten, R., 1973, Syntectonic conglomerates in 

southwestern Montana: their nature, origin, and tectonic significance: Geological Society of America Bulletin, v. 84, p. 773-796. 

Scholten, R., and Onash, C.M., 1977, Genetic relations between the Idaho 

batholith and its deformed eastern and western margins: Northwest 

Geology, v. 6, no. 1, p. 25-37. 

Shieh, Y.N., and Schwarcz, H.P., 1974, Oxygen isotope studies of granite 

and migmatite, Grenville province of Ontario, Canada: Geochimica 

et Cosmochimica Acta, v. 38, p. 21-45. 

Shieh, Y.N., and Taylor, H.P., 1969, Oxygen and hydrogen isotope 

studies of contact metamorphism in the Santa Rosa Range, Nevada 

and other areas: Contributions to Mineralogy and Petrology, v. 20, 

p. 306-356. 

Snee, L.W., Lund, K., and Gammons, C.H., 1985, Mineralization history 

of the central Idaho batholith: importance of fracture-controlled 

Cretaceous activity [abs.]: Geological Society of America Abstracts 

with Programs, v. 17, p. 265. 

Sutter, J.F., Snee, L.W ., and Lund, K., 1984, Metamorphic, plutonic, 

and uplift history of a continent-island arc suture zone, west-central 

Idaho [abs.]: Geological Society of America Abstracts with Programs, 

v. 16, p. 670. 

Swanberg, C.A., and Blackwell, D.D., 1973, Areal distribution and 

geophysical significance of heat generation in the Idaho batholith 

and adjacent intrusions in eastern Oregon and western Montana: 

Geological Society of America Bulletin, v. 84, 1261-1282. 

Taylor, H.P., 1970, Oxygen isotope studies of anorthosites, with particular reference to the origin of bodies in the Adirondack Mountains, New York, in Origin of anorthosites: New York State Museum 
_ 6. PETROGENESIS, GEOCHRONOLOGY, AND HYDROTHERMAL SYSTEMS OF THE NORTHERN IDAHO BATHOLITH 137 

and Science Service Memoir, v. 18, p. 111-134. 

__ 1971, Oxygen isotope evidence for large-scale interaction between 

meteoric ground waters and Tertiary granodiorite intrusions, 

western Cascade Range, Oregon: Journal of Geophysical Research, 

v. 76, p. 7855-7874. 

__ 1977, Water-rock interactions and the origin of H20 in granitic 

batholiths: Geological Society of London Journal, v. 133, p. 509-558. 

__ 1981, The effects of assimilation of country rocks by magmas on 180/ 160 and 87Sr/ 86Sr systematics in igneous rocks: Earth and 

Planetary Science Letters, v. 47, p. 243-254. 

Taylor, H.P., Jr., and Forester, R.W., 1979, An oxygen and hydrogen 

isotope study of the Skaergaard intrusion and its country rocks: a 

description of a 55-m.y.-old fossil hydrothermal system: Journal of 

Petrology, v. 20, p. 355-419. 

Taylor, H.P., Jr., and Magaritz, M., 1978, Oxygen and hydrogen isotope 

studies of the cordilleran batholiths of western North America, in 

Robinson, B.W., ed., Stable isotopes in the earth sciences: Wellington, New Zealand, New Zealand Department of Scientific and 

Industrial Research Bulletin, v. 220, p. 151-173. 

Taylor, H.P., Jr., and Silver, L.T., 1978, Oxygen isotope relationships 

in plutonic igneous rocks of the Peninsular Ranges batholith, 

southern and Baja California, in Short papers of the Fourth International Conference on Geochronology, Cosmochronology, and 

Isotope Geology: U.S. Geological Survey Open-File Report 78-701, 

p. 423-426. 

Tilling, R.I., 1973, The Boulder batholith, Montana: A product of two 

contemporaneous but chemically distinct magma series: Geological 

Society of America Bulletin, v. 84, p. 3879-3900. 

Toth, M.I., Coxe, B.W., Zilka, N.T., and Hamilton, M.M., 1983, Mineral 

resource potential of the Selway-Bitterroot Wilderness, Idaho 

County, Idaho, and Missoula and Ravalli Counties, Montana: U.S. 

Geological Survey Miscellaneous Field Investigations Map MF 

1495-A, scale 1:125,000. 

Turi, B., and Taylor, H.P., 1971, An oxygen and hydrogen isotope study 

of a granodiorite pluton from the southern California batholith: 

Geochimica et Cosmochimica Acta, v. 35, p. 383-406. 

Tschanz, C.M., and others, 1974, Mineral resources of the eastern part 

of the Sawtooth National Recreation Area, Custer and Blaine Counties, Idaho: U.S. Geological Survey Open-File Report, 648 p. 

Vallier, T.L., 1977, The Permian and Triassic Seven Devils Group, 

western Idaho and northeastern Oregon: U.S. Geological Survey 

Bulletin 1437, 58 p. 

Walker, N. W., 1983, Pre-Tertiary tectonic evolution of northeastern 

Oregon and west-central Idaho: constraints based on U/Pb ages of 

zircons [abs.]: Geological Society of America Abstracts with Programs, v. 15, p. 371. 

Williams, L.D., 1979, General geology of a section across the Bitterroot lobe of the Idaho batholith: Northwest Geology, v. 8, p. 29-39. 

Zen, E-an, and Hammarstrom, J.J., 1984, Magmatic epidote and its 

petrologic significance: Geology, v. 12, p. 515-518. 

Zietz, I., Gilbert, F.P., and Kirby, J.R., Jr., 1978, Aeromagnetic map 

of Idaho: Color coded intensities: U.S. Geological Survey Geophysical 

Investigations Map GP-920, scale, 1:1,000,000. 

Zietz, I., Gilbert, F.P., and Snyder, S.L., 1980, Aeromagnetic map of 

Montana: in color: U.S. Geological Survey Geophysical Investigations Map GP-934, scale, 1:1,000,000. 

7. TEMPORAL AND SPATIAL RELATIONS BETWEEN .FOLDING, 

INTRUSION, METAMORPHISM, AND THRUST FAULTING 

IN THE RIGGINS AREA, WEST-CENTRAL IDAHO 

By CHARLES M. 0NASCH 1 

ABSTRACT 

Two pre-Cenozoic, largely volcanogenic rock sequences are exposed 

along the western margin of the Idaho batholith in the Riggins area, 

west-central Idaho. The eastern, upper sequence, the Riggins Group, 

is in contact with the Idaho batholith. It is separated from the unnamed, 

lower sequence to the west by an east-dipping imbricate thrust fault 

system, the Rapid River thrust. 

The temporal and spatial relations between folding, intrusive activity, 

metamorphism, and thrust faulting in the two sequences were determined from textural data and crosscutting relations. In the Riggins 

Group, emplacement of associated peridotite and intrusion of quartz 

diorite and quartz monzonite predate all folding in the southern and central parts of the area. Metamorphism and intrusion of trondhjemite 

associated with the main phase of the batholith occurred with the first 

fold phase in the south, between the third and fourth fold phases in the 

central part, and followed the fifth fold phase in the north. In the lower 

sequence, intrusion of quartz diorite preceded the first phase of folding 

and associated metamorphism. Juxtaposition of the two sequences along 

the Rapid River thrust occurred just prior to or synchronous with the 

last phase of folding in each sequence, which is the only fold phase common to both. Following thrust faulting, both sequences were metamorphosed to the same grade and intruded by trondhjemite in the northern 

part of the area. 

INTRODUCTION 

The Riggins area in west-central Idaho contains the best 

exposures of pre-orogenic (pre-Late Cretaceous) rocks 

along the western margin of the Idaho batholith. This area 

has long been recognized as important to understanding 

the evolution of the western margin of North America 

during the Mesozoic. The western edge of the continental basement, defined on the basis of strontium isotopic 

data (Armstrong and others, 1977; Kistler, 1978), contains 

a deep embayment in the Riggins area (Davis and others, 

1978). Eugeosynclinal additions to the craton were accreted during the Mesozoic construction of the cordilleran 

"collage" (Davis and others, 1978). Rocks in the Riggins 

1Department of Geology, Bowling Green State University, Bowling Green, OH 

43403. 

area belong to distinctly different petrotectonic divisions 

that owe their present positions to structural juxtaposition (Brooks and Vallier, 1978). Lithologic similarities with 

the Seven Devils Group may support correlation of some 

Riggins area rocks with the far-traveled Wrangellia terrane (Jones and others, 1977; Hillhouse and others, 1982). 

Many authors have suggested that the tectonic development of the area was dominated by island-arc processes 

including subduction, which occurred either in the Riggins 

area or immediately to the west (Talbot and Hyndman, 

1975; Hamilton, 1976; Onasch, 1976; Scholten and 

Onasch, 1976). 

Various aspects of the geology in the Riggins area have 

been studied for many years. Folding in the area was 

discussed by Onasch (1977, 1979); thrust faulting was 

described by Hamilton (1963a, b) and Onasch (1977, 1980); 

and the metamorphism and intrusive history were studied 

by Hamilton (1963a, c). The purpose of this paper is to 

summarize the data on folding, thrust faulting, metamorphism, and the intrusive history and to examine the temporal and spatial relations between these events in order 

to establish a framework in which broader scale tectonic 

models may be constructed. Only the pre-Cenozoic history 

will be considered here. For an account of the Cenozoic 

block faulting, crustal warping, and eruption of the 

voluminous Columbia River Basalt Group, the reader is 

referred to Hamilton (1962, 1963d). 

ACKNOWLEDGMENTS 

Versions of this manuscript were reviewed by Robert 

Scholten, Karl Evans, Ferrell McCollough, Tracy Vallier, 

and Howard Brooks; their comments are greatly appreciated. Discussions in the field with Paul Myers and 

Edward Price proved helpful. Support for the field work 

was provided by the Geological Society of America, Sigma 

Xi, and the Paul D. Krynine Fund of The Pennsylvania 

State University. 

139 
140 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

GEOLOGIC SETTING 

The Riggins area lies along the western margin of the 

Idaho batholith (fig. 7.1). The traditional view of the 

batholith as a massive igneous body surrounded by a 

gneissic shell has proved too simple. The batholith is now 

known to consist of two lobes: the Atlanta lobe in the 

south and the Bitterroot lobe in the north. Both lobes are 

a composite of numerous smaller plutons rather than a 

[ill] 

-

~ 

.. 

D 

-0 0 

-

EXPLANATION 

Volcanic and sedimentary rocks (Cenozoic) 

Granite intrusions (early Cenozoic) 

Granitic rocks ( Mesozoic) 

Eugeosynclinal rocks (Mesozoic and Paleozoic) 

Miogeosynclinal rocks (Mesozoic and 

Paleozoic) and Belt Supergroup (Middle 

Proterozoic), undivided 

Crystalline basement rocks of platform 

(pre-Belt Supergroup) 

Gneisses and intrusive rocks of Salmon 

River arch (pre- Belt Supergroup) and 

other gneisses (age uncertain), undivided 

FIGURE 7.1.-Regional geologic map of Idaho batholith and vicinity 

modified from Armstrong (1975). Riggins area is indicated with 

circled R. 

single body. In general, the age of the batholith is Late 

Cretaceous. The Atlanta lobe is slightly older (43 to 100 

Ma) than the Bitterroot lobe (38 to 80 Ma) (McDowell and 

Kulp, 1969; Armstrong, 1975; Armstrong and others, 

1977). Numerous Eocene granitic plutons have resulted 

in widespread resetting of isotopic ages to 50 Ma or less 

(Armstrong, 197 4). 

Armstrong (1975) ascribed the separation of the 

batholith into lobes to the presence of the Salmon River 

arch (fig. 7.1). This arch of older gneisses and intrusive 

rocks, some of which may predate the Belt Supergroup 

(pre-1,500 Ma) is postulated to have influenced the tectonics and paleotopography of central Idaho throughout 

its existence. However, Evans and Lund (1981) found that 

the area occupied by the Salmon River arch in east-central 

Idaho was a depositional basin that received thick accumulations of sediment during the Middle Proterozoic. 

In light of this work, the concept of the Salmon River arch, 

at least in east-central Idaho, may need to be re-evaluated. 

Because the Columbia River Basalt Group forms an extensive cover in this region, prebatholithic eugeosynclinal 

rocks are only exposed in several isolated areas along the 

western margin of the Atlanta lobe. Where exposed, the 

nature of the border zone rocks changes markedly from 

north to south. Along the northwest corner of the Atlanta 

lobe, strongly deformed Belt and pre-Belt(?) to Late 

Triassic rocks are intruded by the batholith (Myers, 1968, 

1972, 1976). The older rocks represent continental basement and presumably mark the northwest end of the 

Salmon River arch. Intrusive rocks of the batholith were 

injected along east-dipping thrust faults and high-angle 

reverse faults. Many of the country rocks were assimilated 

in the contact zone, leaving only schlieren and xenoliths 

of calc-silicate gneiss and quartzite. 

Southward along the western margin, younger rocks 

are in contact with the batholith. The Belt and pre-Belt 

rocks found in the north appear to be absent in the Riggins area. Here, the batholith is in direct contact with 

volcanic rocks of Paleozoic and (or) Mesozoic age. 

South of the Riggins area, the border zone consists of 

varied granitic gneisses with minor calc-silicate gneiss and 

quartzite of unknown age (Hamilton, 1969). These rocks 

are lithologically unlike those in the Riggins area and may 

mark the reappearance of Belt or older rocks in the 

marginal zone. The Riggins area may then be anomalous 

in that it is the only place along the western margin of 

the batholith where eugeosynclinal rocks are in direct contact with the batholith and not separated from it by older 

continental crust (Onasch, 1977). 

ROCK UNITS 

The pre-Cenozoic rock units in the Riggins area consist 

of two largely volcanogenic rock sequences of eugeo-
7. RELATIONS BETWEEN FOLDING, INTRUSION, METAMORPHISM, AND THRUST FAULTING IN THE RIGGINS AREA 141 

synclinal affinity separated by faults of the Rapid River 

thrust. The Riggins'Group and the lower sequence-have 

been variably metamorphosed and intruded by rocks 

related to the batholith. The metamorphic grade and 

amount of intrusive material in both sequences increases 

to the east, and to a lesser extent, to the south and north. 

RIGGINS GROUP 

Hamilton (1963a) named the thick sequence of metamorphosed volcanic and sedimentary rocks lying above and 

to the east of the Rapid River thrust the Riggins Group. 

He defined four formally named units that he assigned 

to the group. In addition, he recognized an informal 

ultramafic rock unit (mainly metaperidotite) which he 

noted as associated (fault emplaced) with the group. From 

lowest to highest, the formal units are: the Fiddle Creek 

Schist, the Lightning Creek Schist, the Berg Creek Amphibolite, and the Squaw Creek Schist (Hamilton, 1963a, 

p. 16-17). This succession, he pointed out, may be partly 

or wholly structural rather than stratigraphic. 

During the present investigation, several members were 

defined in the units of the Riggins Group which aided in 

expanding the mapped extent of the group, especially to 

the east (Onasch, 1977). The other modification to 

Hamilton's original stratigraphy was the reduction of the 

Berg Creek Amphibolite to a member of the Lightning 

Creek Schist (Onasch, 1977, 1979, 1980). The amphibolite 

grades along strike into other lithologies and is believed 

to be a stratigraphic or metamorphic facies of the Lightning Creek Schist. 

Abundant lithologic variations, due to the strong 

regional metamorphic gradient, occur in the area, not only 

between units of the Riggins Group but within each unit. 

Within any unit, the texture and mineralogy vary considerably in response to the changes in metamorphic 

grade (see Hamilton, 1963a, for a detailed description of 

the petrography and metamorphism). 

To the east, the Riggins Group is in contact with intrusive rocks related to the batholith. The contact zone is not 

sharp; rather, it consists of an 8- to 10-km-wide border 

zone dominated by intrusive rocks that increase in abundance to the east. Subunits of the Riggins Group were 

traced well into this transition zone, and it appears that 

they persist until the rocks are entirely plutonic. 

The exposed thickness of the Riggins Group, no doubt 

thickened structurally, is approximately 8 km. The age 

is unknown other than it must be older than the middle 

to Late Cretaceous rocks of the batholith. Brooks and 

Vallier (1978) have tentatively correlated the lower part 

of the Riggins Group with rocks ranging in age from 

Devonian to Middle Triassic, and the upper part with 

Jurassic rocks. 

LOWER SEQUENCE 

For convenience, the three units exposed below and to 

the west of the Rapid River thrust have been informally 

designated the lower sequence by Onasch (1977, 1979, 

1980). They consist, from oldest to youngest, of the Early 

Permian(?) to Late Triassic Seven Devils Group, the Upper Triassic Martin Bridge Formation, and the Triassic 

(?)Lucile Formation. Restrictive lithologic names for the 

Martin Bridge and Lucile Formations were dropped in 

favor of "formation" because both units contain numerous 

lithologies. At low metamorphic grades, typical lithologies 

for the three lower sequence units are greenstone, marble, and phyllite, respectively. In the northern part of the 

area, where rocks of the lower sequence are intruded by 

trondhjemite related to the batholith, schist and marble 

predominate. For more lithologic detail, see Hamilton 

(1963a) and Vallier (1974, 1977). 

CORRELATIONS BETWEEN THE RIGGINS GROUP 

AND THE LOWER SEQUENCE 

The Riggins Group and the lower sequence display 

many lithologic similarities yet also differ in some important ways. This situation has led to a wide range of opinion 

regarding the original stratigraphic relations between the 

two rock sequences. Both sequences consist largely of 

volcanic and volcaniclastic rocks of similar bulk compositions, and both include admixtures of terrigenous and carbonate rocks. They differ in their gross stratigraphy and 

in the absence of peridotite in the lower sequence. Vallier 

(1974, 1977) and Onasch (1976) suggested that the two 

sequences may be partly correlative. Hamilton (1963a, 

1978) believes that the two are unrelated other than many 

of the rock types in both are the product of intraoceanic 

arc volcanism. Any possible correlations will be important 

in some of the structural interpretations developed later. 

FOLDING 

Folding in both rock sequences is polyphase. In either 

sequence, the style of folding and type of associated foliation or lineation are strongly influenced by lithology and 

metamorphic grade. The folding in each rock sequence 

will be discussed separately. An additional subscript will 

be added to the standard structural shorthand notation 

to differentiate between the two sequences: R corresponds to the Riggins Group and B to the lower sequence. 

For example, F2R corresponds to the folds of the second 

phase in the Riggins Group. 

RIGGINS GROUP 

Five phases of mesoscopic folding were identified in the 

Riggins Group. Although the style and degree of develop-
142 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

ment for any phase varies widely over the area, the orientations vary considerably less due to the relative simplicity 

of the macroscopic geometry (fig. 7.2). 

The FIR folds are isoclinal and are the only folds that 

deform compositional layering, SoR, but not the regional 

foliation, SIR· SIR is generally parallel to SoR except in 

the hinges of F1R folds. F2R folds are strongly asymmetric and indicate a rotation sense of east over west. 

F3R and F 4R folds are upright, symmetrical, and the most 

widespread. F5R folds are very open folds, or more rarely, kink-band folds. The average orientations of the fold 

axes and associated S-surfaces for all phases except the 

first are shown in figure 7.3A. First-phase structures have 

highly variable orientations. 

The type of associated foliation and the style of most 

fold phases are strongly dependent on the metamorphic 

grade. Pressure solution and crenulation cleavages are 

typical at low metamorphic grades, whereas schistosity 

and gneissic layering predominate at higher grades. Correspondingly, open, concentric folds become tight and flattened and approach a similar style. The main exception 

to these generalizations are fifth-phase structures. This 

phase postdates metamorphism in most of the area and 

remains relatively constant in style and type of associated 

foliation and lineation. 

As previously mentioned, the complexity seen on a 

mesoscopic scale is not evident on a macroscopic scale. 

The areal distribution of units in the Riggins Group is controlled mainly by two southeast-plunging folds (fig. 7.2), 

the Riggins synform and the Lake Creek antiform 

(Onasch, 1977, 1979). A second synform in the extreme 

eastern part of the area is defined by SIR attitudes, but 

the homogeneous nature of the stratigraphy in this area 

precludes the tracing of individual units with any degree 

of certainty. All these folds correlate with F 4R. Macroscopic examples of F3R and F5R folds can be identified 

from SIR attitudes in the southeast part of the area, but 

the absence of mappable units in this area prevents their 

recognition by map pattern. 

Tectonic transport directions were determined for all 

fold phases, except F lR, using separation angles (Hansen, 

1971) and folded lineations (Ramsay, 1960). For F3R, F 4R, 

and F5R they are subvertical across the area. If the effects of subsequent deformations are removed, F2R 

transport directions plunge to the east at progressively 

steeper angles. 

LOWER SEQUENCE 

Three phases of mesoscopic folding were identified in 

the lower sequence. The style of folding in the lower 

sequence is much more uniform than in the Riggins 

Group, owing to a smaller variation in metamorphic grade. 

Variations in style are due almost entirely to lithologic 

differences. 

The FIB folds are tight to isoclinal and deform compositional layering, SoB· SIB is parallel to the axial surfaces 

of these folds and is a poorly developed foliation in the 

marble defined by flattened calcite grains. In the phyllite, 

SIB is a well developed schistosity. In the greenstone, slB 

consists of anastomosing surfaces defined by the preferred orientation of layer silicates. F2B folds are strongly 

asymmetric with a west over east rotation sense. F3B 

folds are open folds in all lithologies, or more rarely, kinkband folds in phyllite. The orientation of second- and thirdphase structures is shown in figure 7.3B. First-phase 

structures are omitted due to their variable orientations. 

The macroscopic geometry of the lower sequence is even 

less complex than that of the Riggins Group. SIB and SoB 

dip consistently at low angles to the east or southeast. 

The dominant control on the areal distribution of units 

in the lower sequence is the faults of the Rapid River 

thrust (fig. 7.2). Movement directions determined from 

separation angles for F2B plunge gently to the east; for 

F 3B they are vertical. 

Just north of the Riggins area, five phases of folding 

were recognized in the lower sequence (McCollough, 

1984). The two additional phases occurred between F2B 

and F3B as defined in the Riggins area. Thus, the higher 

grade of metamorphism that affected the lower sequence 

north of the Riggins area is apparently accompanied by 

an increase in the structural complexity. 

INTRUSIVE HISTORY 

The intrusive history of both sequences is long and complex. Intrusion occurred prior to, synchronous with, and 

after folding in both sequences. 

RIGGINS GROUP 

The oldest intrusive rocks that are found within the 

mapped extent of the Riggins Group are lensoid bodies 

of metaperidotite. At low metamorphic grade, schists with 

varying proportions of talc, magnesite, antigorite, and 

chlorite are most common. At higher grades, these lowgrade schists are replaced by anthophyllite-garnet schists. 

On the basis of their chemical composition, Hamilton 

(1963a) believes the parent rock to be an enstatite-bearing 

peridotite. 

The metaperidotite bodies are typical of alpine-type 

ultramafic bodies. Their contacts with the enclosing 

schists are conformable and show no effects of contact 

metamorphism. The bodies are highly sheared and lack 

any internal stratigraphy. The origin of alpine-type 

ultramafic bodies is debatable, but most workers agree 

that their present position is a result of tectonic emplacement, commonly along faults. Because SIR can be identified in most of the bodies, emplacement, presumably by 
7. RELATIONS BETWEEN FOLDING, INTRUSION, METAMORPHISM, AND THRUST FAULTING IN THE RIGGINS AREA 143 

CJ Columbia River Basalt Group ·· · (Miocene) 

Lower sequence, undivided 

(Early Permian- Triassic ? ) 

[-<:i:: H Intrusive rocks, undivided -· -- - (Cretaceous ) 

- Metaperidotite (age uncertain) 

EXPLANATION 

Riggins Group 

(age uncertain ) 

D Squaw Creek Schist 

- Lightning Creek Schist 

l:i3~W§ Fiddle Creek Schist 

8' 

/ Contact 

/ Thrust fault 

~ Plunging antiform 

_x Plunging synform 

s---- 8

• Line of section 

FIGURE 7.2.-Generalized geologic map and cross sections of the Riggins area. 

km 
144 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

0 Fold axis 

D 5-surface pole 

faulting, must have occurred prior to D1R, the first phase 

of deformation in the Riggins Group. 

The oldest intrusive rocks related to the Idaho batholith 

are quartz diorite and quartz monzonite. Both occur as 

semi concordant stocks in the eastern part of the area (fig. 

7 .2). They are younger than the metaperidotite because 

they have not been everywhere metamorphosed to the 

same grade as the country rocks, as is the case with the 

metaperidotite. South of the Riggins area, the quartz 

diorite is intruded by the quartz monzonite (Hamilton, 

1963c). The quartz diorite is an early mafic border phase 

of the batholith, whereas the quartz monzonite is nearly 

contemporaneous with the main phase of the batholith 

(Hyndman and Talbot, 1976). Both were intruded prior 

to D1R, as deformed S1R surfaces can be seen in each. 

Trondhjemite associated with the main phase of the 

batholith is the most abundant granitic rock in the Riggins 

area. It was intruded at different times with respect to 

folding at different places. In the southern part of the 

area, trondhjemite dikes predate SlR· In the central part 

of the area, a large composite pluton of variably foliated 

trondhjemite is confined to the axial region of the Lake 

Creek antiform (fig. 7.2), an F 4R fold, implying coeval 

formation. Ten kilometers north of the Riggins area, in 

the Slate Creek area, trondhjemite intrudes across the 

Rapid River thrust and into the lower sequence. The foliation parallels the fault, and it appears that intrusion was 

synchronous with faulting (Hamilton, 1963c). The foliation presumably is related to movements during faulting. 

Although local exceptions are numerous, the intrusion 

of trondhjemite was apparently earlier than D1R in the 

south, approximately synchronous with D4R in the middle 

of the area, and possibly later in the north. The only intrusive rocks younger than trondhjemites are pegmatites 

and aplite dikes, most of which are unfoliated. 

LOWER SEQUENCE 

The intrusive history of the lower sequence is considerably less complex than that of the Riggins Group. 

The oldest intrusive rocks are small stocks of sheared and 

altered quartz diorite (Hamilton, 1963a, b). The shear surfaces are parallel to S1B; therefore, intrusion occurred 

prior to DlB· These quartz diorite bodies and others of 

similar composition are common throughout the Seven 

Devils Mountains, Snake River canyon, and eastern 

FIGURE 7.3.-Summary plot of fold axis and S-surface orientations for 

all deformational phases except the first. Great circle segments show 

separation angles determined from fold asymmetry. A, Riggins 

Group. B, Lower sequence. Numbers refer to phases discussed in 

text. 
7. RELATIONS BETWEEN FOLDING, INTRUSION, METAMORPHISM, AND THRUST FAULTING IN THE RIGGINS AREA145 

Oregon. They range in age from Permian to Early 

Cretaceous, all predating the Idaho batholith (White, 

1973; Vallier, 1974, 1977). 

In the Slate Creek area, trondhjemite related to the 

batholith intrudes the lower sequence. This trondhjemite 

appears to crosscut all fold structures; hence, it is 

post-D3B· 

METAMORPHISM 

With the exception of the younger intrusive rocks, all 

rocks in the Riggins area have been metamorphosed. The 

grade of metamorphism varies from lower greenschist to 

upper amphibolite facies in the Riggins Group and from 

lower to upper greenschist facies in the lower sequence. 

The metamorphic grade in both sequences increases from 

west to east, most markedly so in the Riggins Group. A 

noticeable northward increase in grade is also evident in 

the lower sequence just north of the Riggins area 

(McCollough, 1984). 

The metamorphism in the Riggins Group and most of 

it in the lower sequence is related both spatially and temporally to intrusion of the Idaho batholith. This relation 

is expressed well in Hamilton's (1963a) reference to the 

metamorphism as contact metamorphism on a regional 

scale. An older, low-grade metamorphism was identified 

in the lower sequence by Hamilton and is believed to be 

part of a Late Jurassic event recognized over much of 

eastern Oregon and western Idaho (Armstrong and 

others, 1977). Any possible effects on the Riggins Group 

have been obliterated by the more intense Cretaceous 

everit. The polymetamorphic history indicated by textures 

in some samples of both rock sequences is interpreted in 

most cases to be the result of multiple stages of metamorphism during a single event, because the textures are 

developed irregularly across the area. Where a polymetamorphic history is indicated, reference to the 

metamorphism will correspond to the thermal peak of 

metamorphism. 

Metamorphic isograds in the Riggins area (fig. 7.4) have 

been defined by Hamilton (1963a) and Onasch (1977). The 

relation between the isograds and the faults of the Rapid 

River thrust changes progressively from south to north. 

In the southern part of the Riggins area, the isograds are 

truncated by the faults and metamorphism predates 

faulting. In the north, near Lucile, the isograds cross the 

faults undeflected, and both sequences were metamorphosed to the same grade after faulting. Somewhere in 

the central part of the area, beneath the Columbia River 

Basalt Group, thrust faulting must have been synchronous 

with metamorphism. 

Metamorphism is also related intimately with folding. 

Rock textures show that the temporal relations between 

r-"'1 a......J 

• ~ 

~ 

D 

li.1..J 

Kilometers 

EXPLANATION 

Columbia River Basalt Group (Miocene] 

Lower sequence, undivided (Early Permian-Triassic?] 

Intrusive rocks, undivided !Cretaceous I 

Riggins Group, undivided (age uncertain I 

.,....... Thrust fault 

-·-· lsograd 

AP·alumina prochlorite 

B · bio.tite 

CZ·clinozoisite 

G ·garnet 

0 · oligoclase 

A· andesine 

CP·clinopyroxene 

Z · zoisite 

FIGURE 7.4.-Metamorphic isograds in the Riggins area defined by 

Hamilton (1963a) in the area around Riggins and near Lucile by 

Onasch (1977) 

folding and metamorphism vary with latitude but are 

more constant with changes in longitude. 

RIGGINS GROUP 

In the extreme southern part of the Riggins area, the 

highest grade minerals in the Riggins Group define the 

S1R surface, suggesting that the peak of metamorphism 

was synchronous with D1R· Approximately 10 to 15 km 

to the north, clinozoisite porphyroblasts, the highest grade 
146 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

mineral, grow across S1R and S3R, indicating that the 

metamorphic peak followed D3R· Farther west, at the 

same approximate latitude, S4R is defined by chlorite, indicative of a lower metamorphic grade than the oligoclase, 

garnet, and biotite assemblage that makes up the rest of 

the rock. From these two observations, it can be seen that 

the metamorphic peak in the central part of the area occurred between D3R and D4R· In the extreme northern 

part of the area, metamorphism followed D 4R, as garnet 

porphyroblasts grow across S4R surfaces. 

LOWER SEQUENCE 

Shear surfaces and associated alteration in the quartz 

diorite bodies are believed to be the product of the Late 

Jurassic event (Hamilton, 1963b). As the S1B surfaces in 

adjacent rocks are parallel to these shear surfaces, D1B 

is interpreted to have been synchronous with this metamorphic event. In the northern part of the area, the lower 

sequence is metamorphosed to the same grade as the 

Riggins Group. Here, chloritoid in the Lucile Formation 

grows across F3B kink bands, indicating that metamorphism followed D3B· 

RAPID RIVER THRUST 

The Rapid River thrust extends the entire length of the 

Riggins area, making it the most obvious structural 

feature. It separates the two major rock sequences, and 

in some places, juxtaposes rocks displaying different 

grades of metamorphism. For these reasons, Hamilton 

(1963b) considered it to be a major westward-directed 

overthrust with as much as 24 km of displacement. The 

Rapid River thrust takes on added significance if the lower 

sequence is stratigraphically unrelated to the Riggins 

Group. Brooks and Vallier (1978) suggested that the 

thrust marks the suture between the Seven Devils 

volcanic arc and the lower part of the Riggins Group, 

which they included in their dismembered oceanic crust 

terrane. Correlation of the Seven Devils terrane by Jones 

and others (1977) to Wrangellia implies that the thrusting 

is a middle to Late Cretaceous event that coincides with 

the accretion of Wrangellia (Coney and others, 1981) and 

may represent many hundreds of kilometers of closure. 

The character of the fault changes several times along 

its length (fig. 7 .2). In the extreme southwest corner of 

the area, it is a single surface separating the two sequences. In the central part of the area, just before it 

disappears beneath the Columbia River Basalt Group, the 

fault splays into several surfaces that produce intercalation of the two sequences. Where it emerges from beneath 

the basalt farther north it is marked by at least two surfaces which split and reconnect, creating lens-shaped 

bodies. 

In addition to the main belt of faults just described, at 

least one additional thrust fault was mapped approximately 3 km to the west of the main belt of thrusts in the 

central part of the area (fig. 7.2). Hamilton (1963a) 

recognized the possibility of additional faults in this area 

in his mention of intercalations of marble (Martin Bridge 

Limestone) in the Seven Devils Group. Other low-angle 

faults could be drawn in the lower sequence based on 

anomalous formation thicknesses and local stratigraphic 

inconsistencies. 

The complex distribution of units along the fault is 

puzzling, and no single model can explain all the data. 

Hamilton (1963a, 1969) interpreted the fault to be imbricated southwest of Riggins. Based on reconnaissance 

mapping father north, near Lucile, he ascribed the distribution of units there to be the result of recumbent folding 

of the thrust. Detailed mapping along the entire length 

of the fault by this author has shown the geology to be 

more complex than indicated by Hamilton's reconnaissance map. This complexity is especially apparent in 

the Lucile area, where numerous intercalations of the 

lower sequence with itself and with the Riggins Group 

were found (Onasch, 1977, 1980). Recent work by 

McCollough (1984) just north of the Lucile area argues 

persuasively that intercalations of the lower sequence 

units are stratigraphic in nature. Intercalations of the 

lower sequence and the Riggins Group are, however, 

interpreted by McCollough and this author to everywhere 

be the result of imbricate thrust faulting. 

Several lines of evidence suggest that the Rapid River 

thrust near Lucile is imbricated rather than recumbently 

folded. Graded bedding in the Lucile Formation indicates 

that the rocks are nowhere overturned. Compositional 

layering and schistosity in this area dip consistently to 

the east and southeast. No mesoscopic or macroscopic 

folding with an attitude or style compatible with recumbent folding was identified in either the lower sequence 

or the Riggins Group. 

Just to the north of the Riggins area, McCollough (1984) 

identified open folds in a thrust within the Seven Devils 

Group which he believes is related to the Rapid River 

thrust. Approximately 5 km farther north, along Slate 

Creek, where both sequences are metamorphosed to the 

same grade, recumbent folds occur in both sequences and 

possibly the thrust itself (E. Price, oral commun., 1980). 

The metamorphic grade here is much higher, and the 

structural setting appears to be different. It is possible 

that thrusting is gradually replaced by folding and that, 

along Slate Creek, the two sequences are juxtaposed not 

by thrusting but by nappelike folding. 

Interpreting the Riggins area in terms of an imbricate 

thrust fault model rather than recumbently folded thrust 

faults has important implications with regard to the structural history of the area. If recumbent folding were the 
7. RELATIONS BETWEEN FOLDING, INTRUSION, METAMORPHISM, AND THRUST FAULTING IN THE RIGGINS AREA147 

norm, a major post-thrust episode of folding would have 

had to occur. The imbrication model not only fits the data 

better but simplifies the structural history considerably. 

TIMING OF THRUST FAULTING 

The time of thrust faulting relative to folding can be 

determined by examining possible correlations between 

fold phases in the Riggins Group and those in the lower 

sequence. Using the criteria of style, asymmetry, orientation, and relative age as tools for correlation, only one 

fold phase appears to be common to both sequences. Both 

F5R and F3B are northwest- or southeast-plunging, open, 

upright folds with poorly developed axial plane foliations 

that developed late in their respective deformational sequences. All other fold phases show differences in style, 

asymmetry, orientation, or a combination thereof with 

respect to one another. 

Correlation of only the youngest phase of folding in both 

sequences indicates that movement of the Rapid River 

thrust occurred after D4R and D2B and prior to D5R and 

D3B· The F5R and F3B folds are possibly a product of the 

same stresses responsible for the thrusting, in which case 

folding would be synchronous with faulting. Relative to 

metamorphism and intrusive activity, thrust faulting 

postdates both in the south and predates both in the north. 

DISPLACEMENT ON THE RAPID RIVER THRUST 

The direction of relative motion on the thrusts, based 

on slickenside striae, is west or northwest, approximately perpendicular to the trace of the fault. The sense of 

relative motion, based on the asymmetry of folds in faultrelated shear zones, is reverse, as the east block has 

moved relatively westward. 

An accurate estimate of the lateral shortening associated with the Rapid River thrust requires that the 

original stratigraphic relations between the two sequences 

be known. Based entirely on the discordance in metamorphic grade at the southern end of the fault and assuming 

a metamorphic gradient similar to that in the hanging 

wall, a minimum of 5 to 6 km of displacement is required. 

Detailed versions of the cross sections in figure 7.2 require a minimum of 6 to 8 km of displacement (Onasch, 

1980). These estimates would approach the true displacement if the two sequences were stratigraphically equivalent and originally separated by a rapid facies change. If 

the two sequences are entirely unrelated, the actual 

amount of movement would most likely be much higher. 

The possibility of additional faults in the lower sequence 

increases the estimates in either case. 

The structural histories of the two sequences differ in 

that only the youngest folding in each is correlative. Prior 

to this common fold phase, the two sequences must have 

been separated by a distance great enough to allow 

separate structural evolutions. This distance need not be 

extremely great but would undoubtedly be more than the 

6 to 8 km minimum estimates determined from the cross 

sections. 

SUMMARY 

The structural, igneous, and metamorphic evolution of 

the Riggins area is complex, because the temporal relations between different events change from south to 

north. The relative ages of folding, igneous activity, 

metamorphism, and thrust faulting for southern, central, 

and northern parts of the area are summarized in figure 

7.5. Changes in the relative age of events from one place 

to another are presumed to be gradual and continuous. 

In the southern part of the area, the earliest events after 

deposition and eruption of the rocks of the Riggins Group 

were the intrusion of quartz diorite, quartz monzonite, and 

trondhjemite of the Idaho batholith. The onset of folding 

and the metamorphic peak followed the igneous activity. 

Movement on the Rapid River thrust preceded or accompanied the last stage of folding. In the lower sequence, 

intrusion of quartz diorite occurred prior to D1B and 

associated low-grade metamorphism. The relative ages of 

the first two fold phases in the lower sequence and the 

first four fold phases in the Riggins Group are unknown; 

they may overlap, or early folding in the lower sequence 

may predate all folding and even the intrusive activity and 

metamorphism in the Riggins Group. 

In the central part of the area, intrusion of quartz diorite 

and quartz monzonite and emplacement of peridotite all 

predate folding and the thermal peak in the Riggins 

Group. Intrusion of large volumes of trondhjemite and the 

metamorphic peak occurred between D3R and D 4R· These 

events were followed by thrust faulting and D5R and D3B· 

Just to the north of the Riggins area, along Slate Creek, 

intrusion of trondhjemite and metamorphic peak followed 

thrust faulting and most folding in both sequences. 

Thrusting was synchronous with or slightly older than the 

phase of folding common to both sequences (F5R and 

F3B). 

Some temporal relations remain constant throughout 

the area, whereas others vary from north to south. Thrust 

faulting occurred at the same time relative to folding in 

both sequences everywhere in the Riggins area as did the 

intrusion of trondhjemite relative to the metamorphic 

peak. Relative to folding and thrust faulting, the locus of 

trondhjemite activity and metamorphism shifted progressively northward through the succession of fold phases. 

The changes in the relative age relations could result from 

folding and thrust faulting having affected the entire area 

at one time, while the intrusion of trondhjemite and the 

metamorphism became progressively younger toward the 

north. These changes could also result from intrusion and 
148 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

Rapid River thrust ~ 

R 

.c Metamorphism 

- B ? ? ... 

0 R Tr z Intrusion 

B ?-Od-? 

Folding R F1R F2R F3R F4R FsR 

B ?-F1B F2B ? ? F39 

Rapid River thrust --- R 

'! Metamorphism B ?---? -

s::::: 

G) R Mp, Od, Om ~Tr~ Intrusion 0 B ?-Od-? 

Folding R F1R F2R F3R F4R FsR 

B ?-F1B F2B ? ? F39 

Rapid River thrust --- Metamorphism R 

.c 

- B ? ? :::::1 

0 R Od, Om, Tr U) Intrusion 

B 

Folding R F1R F2R F3R F4R FsR 

B ?-F1B F2B ? ? F39 

Old ~ ~~------------------------------------~· Young Relative Age 

FIGURE 7.5.-Time-space diagram showing relative ages of various events in southern, central, and northern parts of the Riggins area. R, Riggins 

Group; B, lower sequence; Qd, quartz diorite; Mp, metaperidotite; Qm, quartz monzonite; Tr, trondhjemite; F, fold phase. Arrows show range 

of age, question marks show where age extent is uncertain. Significant hiatus may exist between F 1B and F ZB· Relative age of trondhjemite 

in north determined from outcrops in Slate Creek, north of the Riggins area. 

metamorphism occurring everywhere at one time and 

folding and thrust faulting becoming progressively 

younger toward the south. Crosscutting relations in trondhjemite in the central part of the area suggest that the 

former is the case and that intrusion and metamorphism 

are the dynamic elements. The various generations of 

trondhjemite show a general trend of becoming younger 

toward the north. 

The temporal relations between events in the Riggins 

Group and those in the lower sequence older than thrust 

faulting cannot be determined with certainty. In the lower 

sequence, the intrusion of quartz diorite, the low-grade 

metamorphism, D1B, and possibly D2B are all apparently related to a Late Jurassic event for which there is no 

record in the Riggins Group. The absence of this Jurassic 

event in the Riggins Group could have resulted from the 

lower sequence being unrelated stratigraphically to the 

Riggins Group or from the lower sequence being in a different location at that time, or both. As Hamilton (1963b) 

pointed out, the Rapid River thrust may mark the overlap 

of two orogens: a Late Jurassic one to the west and a Late 

Cretaceous one to the east. Following the juxtaposition 

of the two sequences during the Late Cretaceous event, 

they were both affected by the waning stages of this 

event. Polyphase folding in the lower sequence would then 

be the result of two distinctly different orogenic events 

rather than phases of a single event, as was the case in 

the Riggins Group. 

REFERENCES CITED 

Armstrong, R.L., 1974, Geochronology of the Eocene volcanic-plutonic 

episode in Idaho: Northwest Geology, v. 3, p. 1-15. 

__ 1975, Precambrian (1500 m.y. old) rocks of central Idaho-The 

Salmon River arch and its role in cordilleran tectonics and sedimentation: American Journal of Science, v. 275-A, p. 437-467. 

Armstrong, R.L., Taubeneck, W.H, and Hales, P., 1977, Rb-Sr and K-Ar 

geochronometry of Mesozoic granitic rocks and their Sr isotopic composition, Oregon, Washington, and Idaho: Geological Society of 

America Bulletin, v. 88, p. 397-411. 

Brooks, H.C., and Vallier, T.L., 1978, Mesozoic rocks and tectonic evolu-
7. RELATIONS BETWEEN FOLDING, INTRUSION, METAMORPHISM, AND THRUST FAULTING IN THE RIGGINS AREA 149 

tion of eastern Oregon and western Idaho, in Howell, D.G., and 

McDougall, K.A., eds., Mesozoic paleogeography of the Western 

United States: Society of Economic Paleontologists and Mineralogists, Pacific Coast Paleogeography Symposium, 2d, p. 133-146. 

Coney, P.J., Jones, D.L., and Monger, J.W.H., 1981, Cordillera suspect 

terranes: Nature, v. 288, p. 329-333. 

Davis, G.A., Monger, J.W.H., and Burchfiel, B.C., 1978, Mesozoic construction of the cordilleran "collage", central British Columbia to 

central California, in Howell, D.G., and McDougall, K.A., eds., 

Mesozoic paleogeography of the Western United States: Society of 

Economic Paleontologists and Mineralogists, Pacific Coast Paleogeography Symposium, 2d, p. 1-32. 

Evans, K.V., and Lund, K., 1981, The Salmon River "arch"? [abs.]: 

Geological Society of America Abstracts with Programs, v. 13, 

p. 448. 

Hamilton, W., 1962, Late Cenozoic structure of west-central Idaho: 

Geological Society of America Bulletin, v. 73, p. 511-516. 

__ 1963a, Metamorphism in the Riggins region, western Idaho: U.S. 

Geological Survey Professional Paper 436, 95 p. 

__ 1963b, Overlapping of Late Mesozoic orogens in western Idaho: 

Geological Society of America Bulletin, v. 74, p. 779. 

__ 1963c, Trondhjemite in the Riggins quadrangle, western Idaho: 

U.S. Geological Survey Professional Paper 450-E, p. E98-E101. 

__ 1963d, Columbia River basalt in the Riggins quadrangle, western 

Idaho: U.S. Geological Survey Bulletin 1141-L, 37 p. 

__ 1969, Reconnaissance geologic map of the Riggins quadrangle, 

west-central Idaho: U.S. Geological Survey Miscellaneous Geologic 

Investigations Map I-579, 1:125,000. 

__ 1976, Tectonic history of west-central Idaho [abs.]: Geological 

Society of America Abstracts with Programs, v. 8, p. 378-379. 

__ 1978, Mesozoic tectonics of the Western United States, in Howell, 

D.G., and McDougall, K.A., eds., Mesozoic paleogeography of the 

Western United States: Society of Economic Paleontologists and 

Mineralogists, Pacific Coast Paleogeography Symposium, 2d, 

p. 33-70. 

Hansen, E., 1971, Strain facies: New York, Springer-Verlag, 207 p. 

Hillhouse, J.W., Gromme, C.S., and Vallier, T.L., 1982, Paleomagnetism 

and Mesozoic tectonics of the Seven Devils volcanic arc in northeastern Oregon: Journal of Geophysical Research, v. 87, p. 3777-

3794. 

Hyndman, D.H., and Talbot, J.L., 1976, The Idaho batholith and related 

subduction complex: Geological Society of America Annual Meeting, 

Cordilleran Section, 72d, Field Guide No.4, Pullman, Washington, 

14 p. 

Jones, D.L., Siberling, N.J., and Hillhouse, J.W., 1977, Wrangellia-a 

displaced terrane in northwestern North America: Canadian Journal of Earth Sciences, v. 14, p. 2565-2577. 

Kistler, R.W., 1978, Mesozoic paleogeography of California: A viewpoint from isotope geology, in Howell, D.G., and McDougall, K.A., 

eds., Mesozoic paleogeography of the Western United States: Society 

of Economic Paleontologists and Mineralogists, Pacific Coast 

Paleogeography Symposium, 2d, p. 75-84. 

McCollough, W.F., 1984, Stratigraphy, structure, and metamorphism 

of Permo-Triassic rocks along the western margin of the Idaho 

batholith, John Day Creek, Idaho: University Park, Pennsylvania 

State University, M.S. thesis, 140 p. 

McDowell, F.W., and Kulp, J.L., 1969, Potassium-argon dating of the 

Idaho batholith: Geological Society of America Bulletin, v. 80, p. 

2379-2382. 

Myers, P.E., 1968, The geology of the Harpster quadrangle and vicinity, Idaho: Ann Arbor, University of Michigan, Ph.D. dissertation, 

130 p. 

__ 1972, Batholithic emplacement in the Harpster region, Idaho 

[abs.]: Geological Society of America Abstracts with Programs, v. 4, 

p. 206. 

__ 1976, Batholithic emplacement in the Harpster region, Idaho 

County, Idaho: Idaho Bureau of Mines and Geology Pamphlet 80, 

25 p. 

Onasch, C.M., 1976, Infrastructure-suprastructure relations and structural mechanics along the western margin of the Idaho batholith 

[abs.]: Geological Society of America Abstracts with Programs, v. 8, 

p. 407-408. 

__ 1977, Structural evolution of the western margin of the Idaho 

batholith in the Riggins, Idaho, area: University Park, Pennsylvania 

State University, Ph.D. dissertation, 196 p. 

__ 1979, Multiple folding along the western margin of the Idaho 

batholith in the Riggins, Idaho, area: Northwest Geology, v. 8, 

p. 94-100. 

__ 1980, A model for evolution of the Rapid River thrust in the 

Riggins area, Idaho: Northwest Geology, v. 9, p. 19-25. 

Ramsay, J.G., 1960, The deformation of early linear structures in areas 

of repeated folding: Journal of Geology, v. 68, p. 79-93. 

Scholten, R., and Onasch, C.M., 1976, Genetic relations between the 

Idaho batholith and its deformed eastern and western margins [abs.]: 

Geological Society of America Abstracts with Programs, v. 8, p. 

407-408. 

Talbot, J.L., and Hyndman, D.W., 1975, Consequence of subduction 

along the Mesozoic continental margin of the Idaho batholith [abs.]: 

Geological Society of America Abstracts with Programs, v. 7, 

p. 1290. 

Vallier, T.L., 1967, Geology of part of the Snake River canyon and adjacent areas in northwestern Oregon and western Idaho: Corvallis, 

Oregon State University, Ph.D. dissertation, 267 p. 

__ 1974, A preliminary report on the geology of part of the Snake 

River canyon, Oregon and Idaho: Oregon Department of Geology 

and Mineral Industries Geologic Map Series GM-6, 15 p. 

__ 1977, The Permian and Triassic Seven Devils Group, western 

Idaho and northeastern Oregon: U.S. Geological Survey Bulletin 

1437, 58 p. 

White, W.H., 1973, Flow structure and form of the Deep Creek stock, 

southern Seven Devils Mountains, Idaho: Geological Society of 

America Bulletin, v. 84, p. 199-210. 

8. PRE-CENOZOIC GEOLOGY OF THE WEST MOUNTAIN-COUNCIL 

MOUNTAIN-NEW MEADOWS AREA, WEST-CENTRAL IDAHO 

By BILL BoNNICHSEN1 

ABSTRACT 

Pre-Tertiary rocks are exposed in two areas west of McCall and Donnelly and south of New Meadows in west-central Idaho. The southern 

area, called the West Mountain-Council Mountain block, contains the 

West Mountain, Council Mountain, and Deserette dioritic to tonalitic 

plutons and the No Business Mountain monzodioritic to syenitic pluton. 

Five provisional metamorphic rock units are recognized in that block: 

(1) metavolcanic rocks northeast of No Business Mountain, (2) a 

plagioclase-quartz-biotite gneiss unit, (3) metavolcanic rocks on Council Mountain, ( 4) a unit containing abundant metasedimentary rocks, 

and (5) a small marble-calc-silicate rock-hornfelsed amphibolite unit adjacent to the No Business Mountain pluton. Both the plutons and the 

metamorphic units are elongate north to south and both show moderately 

to steeply eastward-dipping north-south foliations. 

The northern area, called the New Meadows block, contains preTertiary rocks that are mostly peridotite accompanied by metachert and 

metagraywacke with local manganese-enriched beds and layers of possible volcaniclastic derivation. The peridotite is mainly unaltered dunite 

that locally is severely deformed. The dunite masses contain harzburgite 

and local chromitite segregations, and serpentinite and chloriteamphibole schist zones at their margins. The deformed olivine and pyroxene grains in the peridotite commonly are cut by undeformed acicular 

anthophyllite crystals that have largely been replaced by talc. The growth 

of these hydrous minerals clearly occurred after the deformation of the 

olivine and pyroxene. Their presence implies that a thermal metamorphic and hydration event affected the peridotite masses after they were 

emplaced into their present tectonic setting. The associated metachert 

and other metasedimentary rocks in the New Meadows block show a 

similar history. Locally their layering is strongly deformed, yet these 

rocks have hornfelsic textures indicative of reheating and annealing after 

emplacement into their present tectonic setting. 

The lithologies, structures, and metamorphic grades of the rocks in 

the New Meadows block differ so much from those in the nearby West 

Mountain-Council Mountain block that the two areas must have had quite 

different geologic histories and probably are separated by a zone (or 

zones) of major tectonic displacement. The metavolcanic and associated 

metasedimentary rocks in the West Mountain-Council Mountain block 

may be correlative with the Lightning Creek and Fiddle Creek Schists 

of the Riggins Group to the north. The peridotite and associated metasedimentary rocks of the New Meadows block are postulated to be a 

melange assemblage that was mixed together and deformed in an oceanic 

trench environment. The rocks in the New Meadows block probably are 

part of the dismembered oceanic crust terrane that is exposed in northeastern Oregon. 

1ldaho Geological Survey, Moscow, Idaho 83843. 

INTRODUCTION 

Pre-Cenozoic rocks west of the Idaho batholith are exposed in the West Mountain-Council Mountain area and 

in the New Meadows area, both of west-central Idaho. In 

this paper these two areas are referred to as the West 

Mountain-Council Mountain block and the New Meadows 

block. The pre-Cenozoic rocks in the West MountainCouncil Mountain block consist of metamorphosed plutons 

of probable Cretaceous age. Previous authors generally 

have included these metamorphic rocks as part of the 

western border zone of the Cretaceous Idaho batholith. 

The pre-Cenozoic rocks in the New Meadows block, approximately 20 kilometers north-northwest of the West 

Mountain-Council Mountain area, include dunite and 

peridotite of unknown age and associated strongly 

deformed quartz-rich rocks, interpreted as metachert, and 

intercalated metagraywacke. A strongly sheared tectonic 

boundary exposed along the east side of the New 

Meadows block separates this small block from a terrane 

to the east that contains metamorphic rocks generally 

resembling those in the West Mountain-Council Mountain 

block. The purpose of this paper is to briefly describe the 

pre-Cenozoic rock units in the West Mountain-Council -

Mountain block and in the New Meadows block. At the 

end of the paper, brief comments regarding how these 

rock units may be related to other geologic units in the 

region are included. 

The West Mountain-Council Mountain-New Meadows 

area consists of north- to north-northwest-trending ridges 

and intervening valleys displaced along late Cenozoic 

faults. The pre-Cenozoic metamorphic rocks comprising 

these blocks are exposed primarily along the east sides 

of the fault-block ridges. The relief in this area is about 

1,300 meters from the low point at New Meadows to the 

top of Council Mountain, which is at an elevation of 2,4 77 

meters. Exposures are poor due to trees and heavy brush 

at lower elevations. Extensive fresh bedrock exposures 

are present only above about 2,100 meters elevation, 

where Pleistocene glaciation has occurred. Such areas are 

151 
152 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

along parts of the West Mountain and Council Mountain 

ridges. Sufficient exposures along streams and roads 

allowed preparation of the preliminary geologic map included in this report. To date, my mapping has been of 

a reconnaissance nature designed to identify the major 

rock units, to gain an understanding of their gross 

lithologic and general structural relations, and to identify more clearly the geologic problems that need to be 

resolved in order to better elucidate the geologic history. 

The information used for this report was obtained from 

29 days of field work during 1978, 1979, and 1983 and 

from the examination of about 100 thin sections. 

The lUGS classification and nomenclature for plutonic 

rocks (Streckeisen and others, 1973) and volcanic rocks 

(Streckeisen, 1979) are used in this paper. For the names 

of most metamorphic rocks, all abundant minerals 

(greater than 5 or 10 percent) are listed in decreasing 

order of abundance. 

ACKNOWLEDGMENTS 

I am pleased to thank Cathy Allen, Earl Bennett, 

Howard Brooks, and Tracy Vallier for reviewing this 

paper and providing many useful suggestions for improving it. In addition, I would like to acknowledge the 

assistance of Ruth Barrett, Charles Knowles, Robert 

Jones, and James Zarubica for help with my research. 

Finally, I would like to thank Serena Smith and Melinda 

Nichols for typing various versions of my manuscript, 

Roger Stewart for his valuable editorial comments, and 

Margaret Jenks for drafting the figures. 

GEOLOGIC SETTING AND PREVIOUS WORK 

The West Mountain-Council Mountain-New Meadows 

area lies between the Cretaceous Idaho batholith 

(Schmidt, 1964; Fisher and others, 1983), which was intruded into a crystalline Precambrian basement and 

Paleozoic marine strata to the east, and the collage of 

Paleozoic and Mesozoic volcanic arc and dismembered 

oceanic crust terranes to the west (Brooks and others, 

1976; Brooks and Vallier, 1978; Hamilton, 1978; Brooks, 

1979a, b), which was accreted onto the North American 

craton during the Mesozoic. The West Mountain-Council 

Mountain-New Meadows area lies astride the postulated 

western limit of Precambrian rocks in North America, as 

deduced from the change in initial87SrJ86Sr ratios from 

greater than 0. 706 to the east to less than 0. 704 to the 

west (Armstrong and others, 1977; see fig. 8.1). This 

isotopic boundary may denote the paleomargin of the 

North American plate after it was truncated but before 

the accretion of the volcanic-arc and dismembered 

oceanic-crust terranes to the west .(see Armstrong and 

others, 1977; Hamilton, 1978; Davis and others, 1978). 

WSVA 

c. ."""') 

HVA j 

0 50 100 

Kilometers 

FIGURE 8.1.-Index map showing the geologic setting of the West 

Mountain-Council Mountain-New Meadows area in west-central 

Idaho. The terrane boundaries are adapted from Brooks (1979a); 

the Riggins area geology is simplified from Hamilton (1969); and 

the 0. 704-0.706 initial strontium isotope line is from Armstrong and 

others (1977). WSVA, Wallowa Mountains-Seven Devils volcanic arc 

terrane; DOC, dismembered oceanic-crust terrane; FB, forearc basin 

terrane; HV A, Huntington volcanic arc terrane; SC, Squaw Creek 

Schist, Riggins Group; LF, Lightning Creek and Fiddle Creek 

Schists, undivided, Riggins Group; IB, Idaho batholith. 

Immediately north of the West Mountain-Council 

Mountain-New Meadows area, the Riggins Group is present within the tectonic boundary zone between the Permian and Triassic volcanic-arc rocks to the west and the 

Cretaceous Idaho batholith and remnants of Precambrian 

rocks to the east (fig. 8.1). The Riggins Group consists 

of metamorphosed sedimentary and volcanic rocks; 

associated with the Riggins Group are enclosed blocks of 

serpentinized ultramafic rocks (Hamilton, 1963, 1969, 

1978). Myers (1982) mapped the Riggins Group north of 
8. PRE-CENOZOIC GEOLOGY OF THE WEST MOUNTAIN-COUNCIL MOUNTAIN-NEW MEADOWS AREA 153 

the Riggins area but still within this tectonic boundary 

zone. Previous workers have not identified the Riggins 

Group in the West Mountain-Council Mountain-New 

Meadows area. Brooks and Vallier (1978) and Brooks 

(1979a) have suggested that the Riggins Group is the 

northeastward extension of late Paleozoic to middle 

Mesozoic dismembered oceanic-crust and forearc-basin 

terranes that lie between the Wallowa Mountains-Seven 

Devils volcanic arc terrane and the Huntington volcanic 

arc terrane in northeastern Oregon (fig. 8.1). 

Palmer (1963) subdivided the southern part of the preCenozoic crystalline rock area at Council Mountain into 

a migmatite unit and an undifferentiated metamorphic 

suite. Schmidt (1964) subdivided the pre-Cenozoic rocks 

in the West Mountain-Council Mountain area, from west 

to east, into the following three units: (1) the gneiss of 

Council Mountain, (2) the migmatite of McCall, and (3) 

the quartz dioritic gneiss of Donnelly. Taubeneck (1971) 

prepared a reconnaissance geologic map of the West 

Mountain-Council Mountain area in which he subdivided 

the pre-Cenozoic rocks, from west to east, into five units: 

(1) the quartz diorite of Deserette, (2) the quartz diorite 

of Council Mountain, (3) a zone of schist, gneiss, and minor 

pegmatite, ( 4) a zone of gneiss, migmatite, pegmatite, and 

granitic rocks, and (5) the quartz diorite of Donnelly. Both 

Schmidt (1964) and Taubeneck (1971) published modal 

analyses of the units that they recognized. Armstrong and 

others (1977) determined the initial strontium isotope 

ratios for the quartz diorites of Deserette, Council Mountain, and Donnelly, and for the nearby units to the north 

and east. 

Very little descriptive information, previous to this 

study, exists for the pre-Cenozoic units in the New 

Meadows area. In a study of the Columbia River Basalt 

Group, Breeser (1972) included the pre-Cenozoic rocks 

south of New Meadows as part of a map unit consisting 

of metasedimentary schist and gneiss injected by quartz 

diorite and trondhjemite. She did not comment on the extensive occurrence of ultramafic rocks in that area ' although she did report that chromite is present south of 

New Meadows. Thayer and Brown (1964) first mentioned 

that ultramafic rocks and associated chromitite deposits 

are present south of New Meadows, but they did not 

describe the rocks. 

The Columbia River Basalt Group is present throughout 

the West Mountain-Council Mountain-New Meadows area 

and essentially covers the pre-Cenozoic rocks for many 

kilometers to the southwest and west, so the relation 

between the rocks discussed in this report and those in 

the Huntington volcanic arc terrane to the west (fig. 8.1) 

are nowhere exposed. The Columbia River Basalt Group 

has been described in the Council Mountain area by 

Palmer (1963), near New Meadows by Breeser (1972), and 

throughout the region by Fitzgerald (1982). 

THE WEST MOUNTAIN-COUNCIL 

MOUNTAIN BLOCK 

The West Mountain-Council Mountain block, in the 

southern part of the study area (fig. 8.2), contains the 

larger of the two exposures of pre-Cenozoic rocks. This 

block contains several metamorphosed plutons, which are 

outlined on figure 8.2, and metamorphosed volcanic and 

sedimentary rocks that I have subdivided into several provisional map units on the basis of differences in lithologic 

assemblages. Foliation and lineation attitudes throughout 

the area are essentially the same in the metamorphosed 

volcanic and sedimentary rocks as in the metamorphosed 

plutons (fig. 8.3). Foliations generally trend north-south 

and dip moderately to steeply east (fig. 8.2). Where present, the lineations plunge moderately to steeply eastward 

in the plane of foliation. The lineations commonly consist 

of mineral streaking or elongation (fig. 8.4). Where compositional layering is present in the metavolcanic and 

metasedimentary rocks and in the plutons, it is parallel 

to the foliation in most instances. Small-scale isoclinal 

folds, which are a common feature in many metamorphic 

terranes, were not observed in the metamorphosed 

plutons and were rarely seen in the metavolcanic and 

metasedimentary rocks in the West Mountain-Council 

Mountain block. Ptygmatic injections of thin aplite dikes 

are present locally, however, as are boudinage structures 

(figs. 8.5, 8.6) in some parts of the metavolcanic and 

metasedimentary sequences. 

Obvious major late Cenozoic faults that offset the 

Columbia River Basalt Group and that have had an important influence on the physiographic development are 

shown in figure 8.2. Undoubtedly, significant faults of preCenozoic age, as well as numerous additional faults of 

Cenozoic age, are present in the area, but careful mapping will be required to find them. 

PLUTONIC ROCKS 

Several felsic to intermediate plutons are present in the 

West Mountain-Council Mountain block. The characteristics of four of the largest plutons are noted in table 8.1. 

Other plutonic bodies appear to have been injected into 

the,area, but due to their poor exposures, smaller sizes, 

and structural dismemberment they have not been outlined on figure 8.2. All the plutons were metamorphosed 

to a greater or lesser degree, depending on their location, 

so they have been modified by recrystallization and the 

development of foliation (fig. 8.3). Where deformation was 

intense, a lineation developed and compositional banding 

formed. 

The West Mountain pluton has been called the quartz 

dioritic gneiss of Donnelly (Schmidt, 1964) and the quartz 

diorite of Donnelly (Taubeneck, 1971). Schmidt's (1964) 

map shows that this unit extends more than 30 kilometers 
154 

116° 

22'30" 

THE IDAHO BATHOLITH AND ITS BORDER ZONE 

0 10 Kilometers ~--L---L-__ L_ __ L_~ 

north-northeast of West Mountain. The Council Mountain 

and Deserette plutons were referred to as the quartz 

diorites of Council Mountain and Deserette by Taubeneck 

(1971), and they form an undifferentiated part of 

Schmidt's (1964) gneiss of Council Mountain. Palmer 

(1963) recognized that the Council Mountain pluton probably had an igneous origin and outlined the southern part 

of it as a migmatite complex. No previous workers 

recognized the relatively deformed No Business Mountain 

pluton as an igneous body. Inasmuch as most of the rock 

exposed in these plutonic bodies is something other than 

quartz diorite, according to the widely accepted lUGS 

classification system (Streckeisen and other, 1973), and 

because the bodies vary somewhat in composition, I have 

dropped the rock-type designation from the names of the 

plutons. 

COUNCIL MOUNTAIN AND DESERETTE PLUTONS 

The least metamorphosed and least deformed plutons 

in the West Mountain-Council Mountain block are the 

Council Mountain and Deserette plutons on the west side 

of the area. The Council Mountain pluton is exposed in 

two areas adjacent to the region covered by later basalt 

flows to the west (fig. 8.2). Although it is uncertain if both 

exposures are part of the same pluton, their great 

similarities in composition and appearance, as well as in 

their north-south alignments, suggest they are. 

The Deserette pluton, which adjoins and extends northwestward from the Council Mountain pluton, is much 

FIGURE 8.2.-Preliminary geologic map of the West Mountain-Council 

Mountain-New Meadows area. Compiled principally from the reconnaissance geologic mapping of the writer, with minor additions from 

Palmer (1963), Schmidt (1964), Schmidt and Mackin (1970), 

Taubeneck (1971), Breeser (1972), Mitchell and Bennett (1979), and 

Fitzgerald (1982). The New Meadows block consists of the peridotite 

and metachert (PC) and the intensely sheared schistose rock (SZ) 

zones in the northern part of the area. The West Mountain-Council 

Mountain block includes the various plutonic and metamorphic rock 

units in the central and southern parts of the area. Plutonic rock 

units: CM, Council Mountain pluton; D, Deserette pluton; WM, West 

Mountain pluton; NBM, No Business Mountain pluton; LV, quartz 

diorite of Little Valley. Metamorphic rock units: 1, metavolcanic 

rocks northeast of No Business Mountain; 2, plagioclase-quartzbiotite gneiss; 3, metavolcanic rocks on Council Mountain; 4, zone 

with abundant metasedimentary rocks; 5, zone of marble, calc-silicate 

rocks, and hornfelsed amphibolite. Other units: PC, peridotite and 

metachert along west side of New Meadows block; SZ, zone consisting mainly of intensely sheared schistose rocks; GC, Goose Creek 

mylonitic gneiss (C. Allen and M. Kuntz, oral commun., 1984). Symbols: heavy lines, faults, bar and ball on downthrown side; light lines, 

geologic contacts, dashed where approximately located; foliation 

symbols, strikes and dips of representative foliations; patterned area, 

Cenozoic materials: consist mainly of the Columbia River Basalt 

Group and unconsolidated sediments. 
8. PRE-CENOZOIC GEOLOGY OF THE WEST MOUNTAIN-COUNCIL MOUNTAIN-NEW MEADOWS AREA 155 

FIGURE 8.3.-View of eastward-dipping foliation and shearing within the West Mountain pluton. Note hammer for scale. 

more leucocratic than the Council Mountain pluton. It 

generally contains less than 10 percent mafic minerals, 

whereas the Council Mountain pluton typically has nearly 

30 percent (table 8.1). Although these two plutons adjoin, 

their contact is not exposed (Taubeneck, 1971). Their 

proximity and nearly identical initial87Srf86Sr ratios and 

K-Ar biotite ages (table 8.1) suggest that they are 

genetically related. Such a relation is also suggested by 

the occurrences of fairly leucocratic zones and leucocratic 

tonalite dikes in the Council Mountain pluton. These two 

plutons may have formed from the same parent magma, 

with the more leucocratic Deserette pluton crystallizing 

later and intruding parts of the earlier crystallized, more 

mafic Council Mountain pluton. 

WEST MOUNTAIN PLUTON 

The West Mountain pluton, commonly having slightly 

more than 30 percent mafic minerals (table 8.1), generally 

is similar in appearance and composition to the Council 

Mountain pluton. It differs by containing augite, which 

the Council Mountain pluton evidently lacks, and by 

having much less epidote than can be found in the Council Mountain pluton (table 8.1). The West Mountain pluton 

also has a considerably higher initial 87Srf86Sr ratio and 

a younger K-Ar biotite age than the Council Mountain 

pluton and shows a greater degree of gneissic structure 

and foliation disrupting its original plutonic fabric 

(fig. 8.3). 

NO BUSINESS MOUNTAIN PLUTON 

The No Business Mountain pluton covers 10 square 

kilometers or more at the top and on the south and west 

slopes of No Business Mountain. It lies between the Council Mountain and Deserette plutons to the west and the 

West Mountain pluton to the east (fig. 8.2). Due to the 

paucity of exposures and its deformed character, the 

southern extent of this body is poorly known. Its contact 

with metavolcanic and metasedimentary rocks is exposed 

at elevation point 7,340 feet on the top of No Business 

Mountain (sec. 34, T. 17 N., R. 2 E.). 

The No Business Mountain pluton varies greatly in composition from one location to another; overall, it ranges 

from diorite to syenite (table 8.1). Most samples are 

enriched in potassium feldspar, in contrast to the West 
156 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

FIGURE 8.4.- Well-developed lineation in the form of mineral streaking developed on a foliation plane in the Council Mountain pluton. 

Mountain and Council Mountain plutons. The No Business 

Mountain pluton apparently has much less quartz than the 

others. Furthermore, parts of the No Business Mountain 

pluton contain relatively high quantities of augite, sphene 

(indicating titanium enrichment), and zeolites, whereas 

none of the samples has more than a trace of biotite. In 

contrast, the tonalitic rocks of the nearby plutons are rich 

in biotite. 

The potassium feldspar in the No Business Mountain 

pluton typically is perthitic, with small to abundant quantities of exsolved plagioclase. Where strongly deformed, 

the potassium feldspars show well-developed microclinegrid twinning. Aplitic dikes which cut the calc-silicate 

rocks on the west side of the No Business Mountain pluton 

(fig. 8.2) and which very likely are apophyses of the pluton 

consist mainly of a single alkali feldspar. These habits of 

potassium feldspar are generally quite different from 

those in the other plutonic rocks or in the nearby metamorphosed volcanic and sedimentary rocks. 

The petrographic features outlined above suggest that 

the No Business Mountain pluton may be petrogenetically different from the other plutons in the area, even 

though samples from all of them are superficially similar. 

The range of rock types, from quartz monzodiorite to 

syenite (table 8.1), indicates that the No Business Mountain pluton is more alkalic than the nearby tonalites and 

quartz diorites. Thus, it may be one of the group of early 

Mesozoic alkalic intrusions that occur near the Mesozoic 

continental margin, from southern California to the 

middle of British Columbia (Miller, 1978). These plutons 

typically have high strontium contents compared to the 

calc-alkaline plutons with which they are associated. X-ray 

fluorescence analyses of one sample from the No Business 

Mountain pluton and one from the West Mountain pluton 

indicate that the No Business Mountain pluton sample 

contains 3 to 4 times more strontium than the sample from 

the West Mountain pluton (C. Knowles, oral commun., 

1984). This contrast, along with the high abundance of 
8. PRE-CENOZOIC GEOLOGY OF THE WEST MOUNTAIN-COUNCIL MOUNTAIN-NEW MEADOWS AREA 157 

FIGURE 8.5.-Boudinage of an amphibolitic layer within a more ductile matrix of quartzofeldspathic gneiss, within the unit of metavolcanic 

rocks on Council Mountain. 

sphene within the No Business Mountain pluton, again 

suggests that the No Business Mountain pluton may be 

part of the continental margin group of early Mesozoic 

(Triassic?) alkalic plutons. In this regard, it is interesting 

to note that the No Business Mountain pluton actually lies 

exactly on the old continental margin, as inferred from 

the initial strontium isotope boundary determined by 

Armstrong and others (1977). 

OTHER PLUTONIC ROCKS 

Additional plutonic rocks are present in the West Mountain-Council Mountain area but are not yet well enough 

understood to show as map units. These include scattered 

occurrences of metamorphosed hornblende-bearing quartz 

diorite or tonalite within the two metavolcanic rock units 

and within the plagioclase-quartz-biotite gneiss map unit 

on the west side of the West Mountain pluton (fig. 8.2). 

These metaplutonic rocks are similar to the metamorphosed tonalites in the Council Mountain and West Mountain plutons. Taubeneck (1971) suggested that part of this 

belt of plagioclase-quartz-biotite gneiss includes maficpoor plutonic rocks of tonalitic composition. He referred 

to these rocks as the quartz diorite west of the Donnelly 

(West Mountain) pluton but did not indicate their distribution on his map, although he included several modal 

analyses. 

Additional plutonic rocks are present immediately south 

of the area shown in figure 8.2 and are believed to extend into the southeast corner of the map area. Rocks 

collected on reconnaissance traverses in sees. 7 and 8, T. 

13 S., R. 3 E., near the crest of West Mountain ridge, 

about 2 kilometers south of the south edge of the area 

shown in figure 8.2, are largely foliated quartz diorites 

or tonalites that contain biotite, muscovite, and garnet, 

but no hornblende. Additional two-mica plutonic rocks 

were observed locally farther west, about 2 kilometers 

south of the south edge of the area shown in figure 8.2, 
158 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

FIGURE 8.6.-Pods or boudins of amphibolite, interpreted as a dismembered, originally continuous, amphibolite layer in the metavolcanic 

rock unit northeast of No Business Mountain. 

in sec. 7, T. 13 S., R. 2 E. These two-mica plutonic rocks 

have only about half the mafic minerals as the tonalitic 

plutons on Council and West Mountains. Their general 

compositions and their locations on the map compiled by 

Mitchell and Bennett (1979) suggest that they are part 

of the leucocratic quartz diorite of Little Valley (Schmidt, 

1964), which was mapped by Fisher and others (1983) as 

the sphene-enriched variety (Kgds) of biotite granodiorite, 

a widespread interior unit of the Idaho batholith. Fisher 

and others (1983) considered this unit to be younger than 

the tonalitic rocks in the West Mountain-Council Mountain block. 

METAVOLCANIC AND METASEDIMENTARY ROCKS 

Five provisional map units within the area of metavolcanic and metasedimentary rocks that surrounds the 

metamorphosed plutons have been outlined within the 

West Mountain-Council Mountain block (fig. 8.2). These 

units are: (1) a zone interpreted as mostly metavolcanic 

rocks northeast of No Business Mountain, (2) a zone of 

mostly plagioclase-quartz-biotite gneiss west and north 

of the West Mountain pluton, (3) a second zone of probable metavolcanic rocks along the western part of the area 

and best exposed on Council Mountain, (4) a zone having 

abundant metasedimentary rocks in the southern part of 

the area, and (5) a small area of marble, skarnlike calcsilicate rocks, and hornfelsed amphibolite next to the No 

Business Mountain pluton. 

The respective ages of these metamorphic rock units 

are unknown, and it is not yet known if their boundaries 

are depositional or tectonic. The first two units noted 

above generally correspond to Schmidt's (1964) migmatite 

of McCall and to the unit that Taubeneck (1971) 

designated as gneiss, migmatite, pegmatite, and granitic 

rocks. The other three generally correspond to Schmidt's 

(1964) gneiss of Council Mountain and to the unit that 

Taubeneck (1971) referred to as schist, gneiss, and minor 

pegmatite. 
8. PRE-CENOZOIC GEOLOGY OF THE WEST MOUNTAIN~COUNCIL MOUNTAIN-NEW MEADOWS AREA 159 

TABLE 8.1.-Characteristics of plutons in the West Mountain-Council Mountain area 

[Rock type nomenclature from Streckeisen and others (1973). K-Ar ages were recalculated for revised decay constants using the method of 

Dalrymple (1979)_ Table includes data from Palmer (1963); Schmidt (1964); Taubeneck (1971); Armstrong (1975); Armstrong and others (1977). 

tr, trace] 

Number of modes 

averaged-------

Quartz-----------

K-feldspar-------

Plagioclase------

Augite-----------

Hornblende-------

Biotite----------

Epidote----------

Muscovite--------

Black opaques----

Apatite----------

Zircon-----------

Sphene-----------

Zeolites---------

Pyrite-----------

An content of 

plagioclase----

Council 

Mountain 

pluton 

6 

16.2 

.4 

54.5 

13.9 

10 .5 

3.4 

tr. 

tr .• 

tr. 

30 (?) 

tonalite 

quartz diorite 

diorite 

Deserette 

pluton 

Modal composition 

7 

29.7 

2.6 

59.5 

5.9 

1.1 

.7 

tr. 

tr. 

tr. 

tr. 

Rock types 

tonalite 

quartz diorite 

granodiorite 

West 

Mountain 

(Donnelly) 

pluton 

8 

21.0 

1.3 

48.4 

.4 

13.0 

15.5 

tr. 

tr. 

tr. 

tr. 

tr. 

34-39 

No Business 

Mountain 

pluton 

range 

noted 

0-10 

15-50 

5-50 

0-15 

10-40 

tr. 

tr.-5 

tr. 

tr. 

tr.-3 

tr. 

tr. 

tonalite syenite 

quartz diorite monzonite 

quartz 

monzonite 

quartz 

monzodiorite 

diorite 

Average initial 87 srJ86 sr ratio 

Number of locations 

averaged---------

Biotite-----------

Hornblende--------

5 

.7039 

84.1 ±2 

119 .9±4 

4 

.0739 

K-Ar ages (Ma) 

86.1 ±3 

4 

• 7086 

77. 9±2 

92.3±3 
160 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

METAVOLCANIC ROCKS NORTHEAST OF 

NO BUSINESS MOUNTAIN 

The zone of probable metavolcanic rocks northeast of 

No Business Mountain (fig. 8.2) consists predominantly 

of strongly foliated, mostly homogeneous to weakly 

layered schist typically containing 50 to 80 percent plagioclase, as much as 20 percent quartz, and from 5 to 15 percent biotite plus hornblende. Biotite commonly is more 

abundant than hornblende. Potassium feldspar generally accounts for less than 10 percent of the schist but locally 

is as abundant as 30 percent. Minor minerals include 

muscovite, apatite, zircon, sphene, opaque oxides, 

clinozoisite, traces of chlorite replacing biotite, and local 

garnet adjacent to the No Business Mountain pluton. 

The rocks in this zone have mineral compositions that 

suggest they originally might have been felsic andesites 

or dacites. Locally the schist contains uniformly distributed feldspar porphyroblasts in a textural arrangement 

similar to that in metamorphosed volcanic rocks of many 

other regions. The rocks in this map unit may have 

originated as a combination of lava flows, welded or 

nonwelded pyroclastic rocks or sedimentary tuffs, and immature sediments derived from a volcanic terrane. 

Foliated fine-grained layers of amphibolite and plagioclase-hornblende-augite-potassium feldspar gneiss are 

intercalated with the felsic metavolcanic rocks and 

generally show sharp contacts. These rocks consist of 

about equal parts of felsic and mafic constituents. 

Plagioclase is the principal felsic mineral. It is accompanied by 0 to 5 percent quartz and minor to locally as 

much as 15 percent potassium feldspar. Hornblende is the 

predominant mafic mineral, but augite locally may 

account for as much as 20 percent of the rock. Minor constituents include biotite, chlorite, epidote, opaque oxides, 

apatite, sphene, and calcic scapolite as a local vein-filling 

mineral. The mineralogy of these amphibolitic rocks suggests that the parent rocks were lava flows or tuffs of 

basalt to basaltic andesite compositions, carbonatebearing volcanogenic sedimentary rocks, or perhaps 

spillitized basalt or tuff. Rocks having compositions intermediate between the felsic and mafic varieties are rare 

in this unit. 

The amphibolite layers commonly are only a few meters 

in maximum thickness and some are less than a meter 

thick. During deformation, many of the mafic layers were 

broken into boudins (fig. 8.6) that are partially to completely surrounded by the more ductile feldspar- and 

quartz-rich schist matrix. The thin nature of many of the 

amphibolite layers suggests that they originated as 

sedimentary or tuffaceous layers rather than as lava 

flows, and the local presence of augite and scapolite suggests that these were carbonate bearing. 

Locally, along the west side of this map unit, the schist 

is medium to coarse grained and gneissic and resembles 

some of the more strongly deformed gneissic parts of the 

metamorphosed plutons on West Mountain and Council 

Mountain. This relation indicates that some metamorphosed intrusive rocks are present within this metavolcanic unit. 

At many localities the rocks in this unit are cut by thin 

alkali feldspar- and quartz-rich pegmatitic dikes. Many of 

the dikes have sharp contacts and are generally concordant with the layering and foliation of the enclosing 

schists. These dikes are probably the result of incipient 

partial melting of their enclosing host rocks during 

metamorphism. 

PLAGIOCLASE-QUARTZ-BIOTITE GNEISS 

The plagioclase-quartz-biotite gneiss north of and immediately west of the West Mountain pluton (fig. 8.2) is 

largely felsic and fine to medium grained, and typically 

contains 65 to 75 percent plagioclase, 10 to 25 percent 

quartz, as much as 10 percent potassium feldspar, and 

generally only 5 to 15 percent mafic minerals, mostly 

biotite. As much as 10 percent hornblende locally is 

present but more commonly is absent. Trace amounts of 

apatite, zircon, and opaque oxide minerals generally are 

present, and minor amounts of augite, sphene, epidote, 

muscovite, tourmaline, and graphite have been noted 

locally. The premetamorphic textural features have generally been obliterated by deformation and recrystallization. 

The rock compositions in this unit typically consist of 

diorites, quartz diorites, and tonalites, or their volcanic 

equivalents. The augite-, hornblende-, and sphene-bearing 

rocks may largely be plutonic in origin, whereas at least 

part of the finer grained gneiss with biotite as the main 

mafic mineral may be volcanic in origin. The local 

homogeneity of most exposures also suggests an igneous 

origin for most of this unit. 

Rocks suspected of having a volcanic origin are present 

principally in the northern part of this unit near the 

metavolcanic rocks northeast of No Business Mountain 

(fig. 8.2), whereas most of the rocks in the southern part 

of this unit are probably of plutonic origin. Alkali feldsparand quartz-rich aplitic to pegmatitic dikes as much as one 

meter thick are present in many exposures in the southern 

part of this unit. The unit which Taubeneck (1971) 

designated as the quartz diorite west of the Donnelly 

(West Mountain) pluton probably corresponds to the 

metaplutonic southern part of this unit. 

Rocks that are compositionally heterogeneous locally 

or rocks that contain minerals whose presence would infer 

a probable sedimentary origin, such as garnet, muscovite, 

diopsidic pyroxene, calcite, or scapolite are notably sparse. 

To date, the only area identified in this unit that might 

contain significant quantities of metasedimentary rocks 

is in the east half of sees. 15 and 22, T. 16 N., R. 2 E., 
8. PRE-CENOZOIC GEOLOGY OF THE WEST MOUNTAIN-COUNCIL MOUNTAIN-NEW MEADOWS AREA 161 

where traces of graphite occur in a medium-grained, compositionally layered plagioclase-quartz-biotite mixture of 

gneiss and schist. 

Sparse to abundant potassium feldspar megacrysts are 

present to conspicuous in the northern part of the 

plagioclase-quartz-biotite gneiss and perhaps elsewhere. 

Where these megacrysts occur the gneiss is very similar 

in appearance to the Goose Creek mylonitic gneiss (C. 

Allen and M. Kuntz, oral commun., 1984) located in the 

northeastern corner of the area shown in figure 8.2. The 

Goose Creek mylonitic gneiss sits astride the initial strontium isotope line in its occurrence east of New Meadows 

in the same fashion as can be inferred for the plagioclasequartz-biotite gneiss in the West Mountain-Council Mountain block. 

METAVOLCANIC ROCKS ON COUNCIL MOUNTAIN 

The zone of predominantly metavolcanic rocks on Council Mountain, along the west side of the pre-Cenozoic 

exposure belt (fig. 8.2), contains gneisses and schistose 

rocks very similar to those in the metavolcanic rock unit 

northeast of No Business Mountain. As in the No Business 

Mountain area, two rock types, felsic and mafic, dominate 

in the metavolcanic rock unit on Council Mountain. The 

more felsic rocks consist mostly of plagioclase and quartz, 

commonly having 5 to 15 percent biotite plus hornblende, 

with biotite generally predominating. These rocks typically are strongly foliated and fine grained, and locally 

display textures resembling those of metavolcanic rocks 

of other regions around the world. These felsic rocks 

probably were initially dacites to felsic andesites. 

Most of the more mafic rocks are moderately to strongly 

lineated amphibolites having about equal amounts of 

plagioclase and hornblende. The abundance and thickness 

of the amphibolite layers in the Council Mountain area 

typically are much greater than in the No Business Mountain area; this relation implies that many of the mafic 

layers of the Council Mountain area may initially have 

been mafic andesite or basaltic flows or tuffs. Also supporting a metavolcanic rather than metasedimentary 

origin for most of the Council Mountain amphibolites is 

the apparent absence of diopsidic pyroxene segregations 

within most amphibolite masses, in contrast to the 

behavior described below for some of the hornblende-rich 

rocks within the map unit of abundant metasedimentary 

rocks. 

Locally within the Council Mountain area, thin amphibolite layers display boudinage structure within the 

more ductile felsic matrix (fig. 8.5). Aplitic and pegmatitic 

dikes cutting the metavolcanic rocks are quite common 

throughout the Council Mountain area. Most pegmatites 

contain either albite or potassium feldspar and quartz, and 

minor biotite or muscovite. Rarely, small garnets are present in the pegmatites and aplites. Most dikes appear to 

have been metamorphosed, as they commonly display 

foliation concordant to the regional foliation. Quartz veins 

are present locally in the metavolcanic rocks around Council Mountain. Most were recrystallized and segmented 

during the deformation of their host rocks. 

Felsic rocks that may originally have been plutonic are 

present locally in the metavolcanic rocks on Council Mountain. Such rocks are less abundant than the rocks of probable volcanic origin. A few metasedimentary rocks that 

locally contain garnets and thin quartz- and muscoviterich layers are also present in this area. These layers may 

have formed from thin felsic-tuff layers. 

MAP UNIT WITH ABUNDANT METASEDIMENTARY ROCKS 

A zone having abundant metasedimentary rocks is 

present southwest of the plagioclase-quartz-biotite gneiss 

in the southern part of the area shown in figure 8.2. Two 

main rock types are in this unit: (1) Ca- and Mg-rich rocks 

composed largely of clinopyroxene and hornblende which 

appear to have originated as carbonate-rich sediments, 

and (2) quartzofeldspathic gneisses and schists that commonly contain garnets and two micas, but which vary in 

minor mineral content. This second group predominates 

within the map unit and encloses the Ca- and Mg-rich 

rocks. The quartzofeldspathic gneisses and schists include 

rocks that were probably deposited as sediments and 

others that may have been intermediate to silicic volcanic 

rocks. Metaplutonic rocks are scarce or absent from this 

unit, as are rocks with quartz-rich or pelitic compositions. 

The Ca- and Mg-rich rocks include compositionally 

layered types consisting mostly of diopsidic clinopyroxene, tremolitic amphibole, plagioclase, and quartz or calcic 

scapolite and calcite, accompanied by minor potassium 

felspar, apatite, sphene, and opaque minerals. Palmer 

(1963) also noted the occurrence of similar interlayered 

marble, amphibolite, and quartz-orthoclase-diopsidescapolite gneiss within this unit, in either sec. 6, T. 14 N., 

R. 2 E. or sec. 31, T. 15 N., R. 2 E. 

Other Ca- and Mg-rich rocks are present as groups of 

irregular to streamlined pods similar to the boudin shown 

in figure 8.6. These pods are as much as a few meters long 

and are present within some of the quartzofeldspathic 

gneisses. The cores of the larger pods commonly consist 

of augite or diospide that is partly replaced by hornblende 

and that is accompanied by small percentages of calcite, 

epidote, opaque minerals, apatite, and sphene. The 

clinopyroxene-rich cores commonly are surrounded by 

hornblende-rich coronas as much as several centimeters 

thick. These pods are believed to be boudins which initially 

were thin layers rich in carbonate minerals of the 

dolomite-ankerite series. 
162 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

The appearance, overall compositions, and absence of 

muscovite, garnet, graphite, or other minerals that infer 

a sedimentary origin suggest that the gneisses enclosing 

the clinopyroxene-rich pods probably were andesitic to 

dacitic flows or tuffs. Closely associated with the local 

metavolcanic rocks that enclose the clinopyroxene-rich 

pods are sparse occurrences of leucocratic schistose to 

gneissic rocks that are rich in plagioclase, potassium 

feldspar, and quartz. These rocks have compositions and 

textures that suggest a rhyolitic tuff origin. 

The quartzofeldspathic gneisses and schists of this map 

unit are predominantly medium grained and foliated, and 

contain 40 to 80 percent plagioclase, 10 to 40 percent 

quartz, and 0 to 15 percent mafic minerals. Potassium 

feldspar generally is absent or is present in trace amounts. 

Biotite is the principal mafic mineral. Some of the gneisses 

are leucocratic but contain little or no potassium feldspar. 

These rocks contain traces to a few percent muscovite or 

garnet (almandine?), or both. Traces of apatite, zircon, 

opaque minerals, and chlorite (replacing biotite) commonly 

are present, but clinopyroxene, hornblende, epidote, 

sphene, and carbonate minerals have not been noted. The 

presence of muscovite and garnet, coupled with the 

absence or paucity of potassium feldspar, suggests that 

these gneisses are metamorphosed sediments rather than 

volcanic rocks. The common appearance of thin biotiterich layers and streaks within the gneiss also suggests a 

metasedimentary origin. 

MARBLE, CALC-SILICATE ROCKS, AND 

HORNFELSED AMPHIBOLITE 

A small area of Ca- and Mg-rich rocks, including marble 

and gneissic- to hornfelsic-textured rocks with varying 

amounts of hornblende, clinopyroxene, and plagioclase, 

is present near the corner of sees. 4, 5, 8, and 9, T. 16 

N., R. 2 E., between the No Business Mountain pluton 

and the metavolcanic rocks on Council Mountain (fig. 8.2). 

A marble body that may be as much as several tens of 

meters thick is present in that area. It varies from approximately 50 percent quartz to nearly pure calcite. Accessory minerals include tremolite, phlogopite(?), graphite, 

and magnetite or pyrite. As noted above, Palmer (1963) 

reported an additional marble occurrence at a location 

about 18 kilometers to the south in the map unit having 

abundant metasedimentary rocks (fig. 8.2). Leucocratic 

dikes and irregular masses of aplite locally cut the marble 

and the associated calc-silicate rocks. 

The hornfelsic- to gneissic-textured rocks vary widely 

in their mineral proportions. Overall, these rocks resemble the amphibolites in the region but locally include calcsilicate skarn. The textures in some of these rocks suggest 

that reheating partially annealed a preexisting gneissic 

fabric to a hornfelsic arrangement. Less abundant 

minerals include grossularite to andradite garnet, quartz, 

potassium feldspar, calcite, epidote, scapolite, opaque 

minerals, chalcopyrite, apatite, sphene, and idocrase. The 

minor-mineral assemblage suggests that these rocks were 

derived from carbonate-bearing to carbonate-rich 

(dolomitic to ankeritic) sedimentary material, rather than 

mafic volcanic rocks. 

The composition of the aplitic dikes and masses and 

their proximity to the No Business Mountain pluton, suggest they are apophyses of the No Business Mountain 

pluton that were injected into and thermally metamorphosed the sedimentary rocks in this small area. All the 

rocks in this unit, including the skarnlike mass, were 

deformed and metamorphosed again, along with the rest 

of the rocks in the West Mountain-Council Mountain 

block. Although this small area of metasedimentary rocks 

is approximately on line with the map unit containing 

abundant metasedimentary rocks to the south (fig. 8.2), 

it is not yet clear if this area represents a northward extension of that unit. 

THE NEW MEADOWS BLOCK 

The New Meadows block is in the northern part of the 

West Mountain-Council Mountain-New Meadows area 

(fig. 8.2). The western part of the block consists of a northnortheast-trending, 7 -kilometer-long zone of mainly 

peridotite and impure quartzite. Both rock types are cut 

by aplitic to pegmatitic dikes and masses having granitic 

to trondhjemitic compositions. Scattered outcrops in the 

eastern part of the block are a mixture of amphibolite, 

quartzofeldspathic schist, and local metapyroxenite and 

serpentinite, all cut by aplites and pegmatites. 

The foliation in the New Meadows block trends north 

to north-northeast and dips moderately to steeply eastward, as in the West Mountain-Council Mountain block. 

Lineations in the western part of the block, however, are 

subhorizontal within the foliation plane. They are commonly in the form of elongate mineral streaks and shallow 

elongate grooves that kink the wispy to thin layers in the 

quartz-rich rocks. Their attitude is in marked contrast to 

the eastward-plunging, down-dip lineations in the West 

Mountain-Council Mountain block. 

The strong deformation and juxtapositioning of various 

rock types in this area, along with many straight stream 

courses, escarpments, and other topographic features, 

indicate a complex history of pre- and post-Columbia River 

basalt faulting. Aside from noting on figure 8.2 the obvious major postbasalt faults that control topographic 

features and the distribution of pre-Cenozoic rock 

exposures no attempt has been made to decipher the 

faulting history of the New Meadows block. 
8. PRE-CENOZOIC GEOLOGY OF THE WEST MOUNTAIN-COUNCIL MOUNTAIN-NEW MEADOWS AREA 163 

ULTRAMAFIC ROCKS 

Ultramafic rocks are abundant in a 4-kilometer-long 

zone along the west side of the New Meadows block in 

sees. 2, 3, 10, and 15, T. 18 N., R. 1 E. Reconnaissance 

traverses indicate the presence of several different ultramafic rock bodies separated by quartzose, and locally 

schistose quartzofeldspathic to amphibolitic, rocks. The 

ultramafic rocks vary from nearly unaltered peridotite to 

schists composed of various proportions of anthophyllite, 

tremolite, talc, serpentine (mainly antigorite and 

lizardite), and chlorite (probably corundophilite). These 

schistose rocks generally occupy the peripheral parts of 

the ultramafic masses and occur as zones that crosscut 

fresh peridotite. 

Forsteritic olivine, enstatitic orthopyroxene, and 

chromite are the principal minerals in the unaltered 

peridotite. Clinopyroxene has not been seen, but small 

quantities were likely present judging from locally abundant tremolite in hydrated zones. Olivine is much more 

abundant than orthopyroxene, so most of the ultramafic 

rock is dunite. None of the peridotite examined to date 

contains plagioclase. Float samples of chromitite have 

been found near prospect pits in the SW1

/4 NW1

/4 sec. 2, 

T. 18 N., R. 1 E. The chromitite probably originated as 

a magmatic segregation when the peridotite crystallized. 

The peridotites vary from cumulates with equant olivine 

grains packed loosely in an interstitial orthopyroxene 

matrix (fig. 8. 7) to severely tectonized rocks in which large 

(about 1 em or more) olivine grains have been strained 

into sets of subparallel, curved lamellae (fig. 8.8). The 

small percentage of chromite consists mainly of equant 

grains that are enclosed by orthopyroxene but not by 

olivine, suggesting that chromite is also a cumulus 

mineral. The overall petrologic, mineralogic, and structural characteristics of the peridotite bodies resemble 

those of alpine peridotites or fragments of ophiolite complexes, including the Canyon Mountain Complex (Thayer, 

1977; Ave Lallemant, 1976) in northeastern Oregon 

within the dismembered oceanic crust terrane (fig. 8.1) 

described by Brooks and Vallier (1978) and Brooks 

(1979a). 

Peridotite that has been partially converted to hydrous 

minerals shows evidence of having undergone at least two 

hydration episodes, both postdating the straining of the 

olivine. During the first episode, numerous thin veins and 

local patches of very fine grained serpentine developed 

throughout the rocks, cutting across olivine and pyroxene (fig. 8. 7). Locally forming with this early serpentine 

were veins and patches of chlorite and traces of finegrained magnetite. No amphiboles and little, if any, talc 

appeared during this early serpentinization episode. 

During the second hydrous alteration episode, anthophyllite and talc formed and were accompanied locally by 

tremolite and additional serpentine, chlorite, and magnetite. The anthophyllite developed throughout the peridotite masses, both within olivine and pyroxene and cutting 

the earlier serpentine veins (fig. 8.9). Tremolite developed 

concurrently with anthophyllite and very likely consumed 

calcium released by the alteration of clinopyroxene. 

Following the formation of anthophyllite and tremolite, 

abundant talc formed, both in veins and as masses replacing anthophyllite (fig. 8.9). Talc formation was an 

early retrograde stage of the second hydration episode. 

Accompanying and following the talc, additional serpentine, chlorite, and magnetite formed as a late retrograde 

stage. 

Although anthophyllite and talc are present throughout 

the peridotite masses, they are much more conspicuous 

in some areas. Their maximum development is on the 

northwest side of Vick Creek in NE1

/4 sec. 15, T. 18 N., 

R. 1 E., where the anthophyllite occurs as sheaves and 

rosettes of acicular crystals (fig. 8.10), some longer than 

20 centimeters. After they formed, the anthophyllite 

sheaves and rosettes were largely replaced by talc 

(fig. 8.9). 

There is a marked contrast between the nondeformed 

anthophyllite and talc and the strongly deformed olivine 

grains in the peridotite (compare figs. 8.8 and 8.10). The 

nondeformed character of the anthophyllite and talc suggests that the peridotite bodies were near their present 

structural positions when they were reheated, allowing 

the anthophyllite and talc to grow statically. The strong 

olivine deformation probably occurred during the emplacement of the peridotite masses into their present tectonic settings, or it might be inherited from an earlier 

stage. 

Locally cutting the peridotite are trondhjemitic to 

granitic dikes of aplitic to pegmatitic texture. Their injection probably was contemporaneous with the heating 

event that formed the anthophyllite, tremolite, and talc. 

The thoroughly hydrated ultramafic rocks that are 

marginal to or that cut the peridotite masses range from 

almost pure anthophyllite, tremolite, talc, chlorite, or 

serpentine to mixtures of these minerals. Their textures 

vary from schistose to decussate. Inasmuch as these rocks 

are entirely hydrated, their evolution is not as easily 

deciphered as that of rocks containing olivine and pyroxene. Nevertheless, assuming the amphiboles and talc 

formed in response to heating after the ultramafic rocks 

were in place, it is probable that the peripheral concentrations of these minerals developed concurrently, probably as water migrated from adjacent rocks into the 

peridotite. Whether the serpentinite and chlorite schist 

adjacent to the peridotite bodies formed at this time, or 

earlier, is not known. 
164 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

FIGURE 8.7.-Photomicrograph of remnants of equant olivine grains, about a millimeter across, set in a talc-rich matrix developed mostly from 

orthopyroxene that was interstitial to the olivine. This texture is interpreted as a relict cumulate texture. Note the abundant thin veins of 

serpentine cutting the olivine grains. Thin section in plane light; width of view is about 7 millimeters. 

METAMORPHIC ROCKS 

Mixed with the peridotite is an assortment of thinly 

layered quartz-rich metasedimentary rocks that range 

from gneissic to hornfelsic in texture. Present with these 

quartz-rich rocks, mainly to the south, are layered rocks 

having quartzofeldspathic to amphibolitic compositions 

that originally may have been graywackes or similar 

clastic sedimentary rocks. Also associated with the quartzrich rocks are bedded manganese-oxide occurrences (SE1f4 

sec. 3, T. 18 N., R. 1 E., and vicinity). Only the quartzrich rocks have been studied in thin section. 

The quartz-rich rocks in the western part of the New 

Meadows block consist of alternating quartz-rich layers 

and thinner, strongly deformed layers and pods of Ca- and 

Mg-rich to Fe- orAl-bearing rocks that are less abundant. 

The quartz-rich layers commonly are 90 percent or more 

quartz and have a few percent hornblende, biotite, or 

magnetite, and minor amounts of garnet (almandine?), 

epidote, and apatite. Local traces of sphene, calcite, 

muscovite, tourmaline, chlorite, and rutile(?) have been 

noted. The layers with little or no quartz (typically less 

than 20 percent) consist largely of hornblende (20-70 percent), epidote (as much as 50 percent), clinopyroxene (0-20 

percent), plagioclase (0-20 percent), and zoisite (as much 

as 15 percent), accompanied by sparse magnetite, sphene, 

and hematite(?) or rutile(?). Minerals occurring locally include as much as a few percent scapolite, minor amounts 

of xanthophyllite(?) or clintonite(?), and traces of apatite. 

No potassium feldspar has been found in either the quartzrich or quartz-poor layers. These quartz-rich metasedimentary rocks may have originated as pelagic marine 
8. PRE-CENOZOIC GEOLOGY OF THE WEST MOUNTAIN-COUNCIL MOUNTAIN-NEW MEADOWS AREA 165 

FIGURE 8.8.-Photomicrograph of an olivine crystal that initially was about 5 millimeters across that has been reorganized into parallel, but 

bent, lamellae. Note the acicular anthophyllite crystals cutting the olivine in the middle of the view. Thin section in plane light; width of 

view is about 6 millimeters. 

cherts, in which the intercalated quartz-poor layers were 

marl or other carbonate-rich and clay-bearing material 

having local concentrations of ferruginous material. 

Structural features in these quartzose rocks are complex. Tight folds and fragmented layers are common in 

hand samples and small outcrops. Although poorly exposed, part of the interlayered quartz-rich and quartz-poor 

material was apparently so strongly deformed as to locally 

rupture and scramble the layering. At other places, 

however, the layering is much less deformed, or even 

planar. 

Well after the disruption of their layering, the quartzrich and other metasedimentary rocks were hornfelsed. 

This sequence of deformation and recrystallization is 

similar to that noted for the peridotite. The deformation 

of the layering may have occurred during or before the 

emplacement of these metasedimentary rocks into their 

present tectonic position, whereas the development of the 

hornfelsic texture, similar to the growth of the undeformed anthophyllite crystals in the peridotite, probably 

occurred during reheating. Aplitic to pegmatitic dikes, 

generally having granitic to trondhjemitic compositions, 

cut the metasedimentary rocks as they cut the periodotite. 

Metamorphic rocks, mostly of quartzofeldspathic and 

amphibolitic compositions, are present to the east of the 

metasedimentary rocks and peridotite in sees. 5 and 18, 

T. 18 N., R. 2 E. (fig. 8.2). Most of these rocks are strongly 

sheared but appear to have been mostly plutonic, ranging from gabbro to tonalite. They are thought to be within 

a major zone of faulting. Locally present in the zone are 

detached pyroxene- and hornblende-rich boudins of gabbroic and pyroxenitic compositions and tectonic slivers as 
166 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

FIGURE 8.9.-Texture of acicular crystals of anthophyllite, now almost entirely replaced by talc (light colored), developed at the expense of olivine. 

Note the fine-grained magnetite (black grains) within the talc and the numerous thin serpentine veins cutting the olivine but truncated by 

the talc masses. Thin section in plane light; width of view is about 9.5 millimeters. 

much as a few meters across composed of chlorite, serpentine, and talc, similar in composition to the peridotite to 

the west. The shear zone is cut by many pegmatitic dikes, 

consisting of albite, quartz, white mica, and traces of redbrown garnet. The pegmatites are foliated locally but are 

essentially intact. Evidently their injection largely postdated the shearing and dismemberment of their host 

rocks. These sheared rocks on the eastern side of the New 

Meadows block are probably either part of a tectonic zone 

between accreted terranes or within the major tectonic 

zone between the North American craton and the accreted 

terranes to the west. 

DISCUSSION 

As is common in geology, the presence of the thoroughly 

deformed and metamorphosed sedimentary, volcanic, 

ultramafic, and plutonic rocks in the West MountainCouncil Mountain-New Meadows area raises more questions than answers about the evolution of this part of 

Idaho and of the Pacific Northwest. By taking into account the regional geology as deciphered by various 

geologists in the papers cited and by attempting to interpret the premetamorphic nature of the rock units, it seems 

that the rocks may have originated somewhere in the 

Pacific Ocean basin during the Paleozoic and Mesozoic. 

No available evidence suggests that any of the metasedimentary or metavolcanic rocks in the West MountainCouncil Mountain-New Meadows area were part of the 

early or middle Precambrian North American craton, or 

part of the supracrustal cover of Middle Proterozoic (Belt 

Supergroup or its lateral equivalents) and Paleozoic 

miogeosynclinal sedimentary rocks. 
8. PRE-CENOZOIC GEOLOGY OF THE WEST MOUNTAIN-COUNCIL MOUNTAIN-NEW MEADOWS AREA 167 

FIGURE 8.10.-Radiating sheaves of talc that replaced earlier formed 

anthophyllite in a matrix consisting mostly of olivine. Note the 

nondeformed habit of the hydrous minerals. Pocket knife is 8.5 centimeters long. 

My interpretation that these rocks are not part of the 

Precambrian craton agrees with the observation of Armstrong and others (1977) that the 0.704-0.706 initial strontium isotope ratio line, which demarks the western margin 

of the Precambrian North American craton, passes 

through the area. This boundary, which separates rocks 

having initial87Srf86Sr ratios of greater than 0.706 to the 

east from those having less than 0.704 to the west, passes 

between the West Mountain pluton and the Deserette and 

Council Mountain plutons on Council Mountain (table 8.1). 

The West Mountain pluton, as part of the Idaho batholith, 

probably was intruded after the continental accretion 

process was completed. By the reasoning of Armstrong 

and others (1977), its higher strontium isotope ratio is due 

to contamination by old crustal material. The plutons on 

Council Mountain, however, do not contain excess 

radiogenic strontium and thus were not contaminated by 

assimilating older crustal rocks. 

Lithologic similarities suggest that the metavolcanic and 

metasedimentary rocks in the West Mountain-Council 

Mountain block might be the southern extension of the 

Lightning Creek and Fiddle Creek Schists. These 

dominantly volcanic units form the structurally lower part 

of the Paleozoic or Mesozoic Riggins Group (Hamilton, 

1963). Brooks and Vallier (1978) and Brooks (1979a) have 

suggested that the Fiddle Creek and Lightning Creek 

Schists are part of their dismembered oceanic crust terrane (fig. 8.1), and, at least in part, are correlative with 

the Burnt River Schist and the Elkhorn Ridge Argillite 

in northeastern Oregon. If this correlation is correct, then 

the metavolcanic and associated metasedimentary rocks 

in the West Mountain-Council Mountain block also should 

be considered as part of the dismembered oceanic crust 

between the two volcanic arc terranes. 

As an alternative, the metavolcanic and metasedimentary rocks in the West Mountain-Council Mountain block 

may be unrelated to the Riggins Group or to the other 

terranes to the west noted in figure 8.1. If so, the rocks 

in the West Mountain-Council Mountain block could be 

either substantially younger or older than the rocks in 

either the Riggins Group or the Huntington volcanic arc 

terrane. 

The peridotite bodies in the New Meadows block probably are related to other ultramafic occurrences in northeastern Oregon and western Idaho. In Oregon these 

ultramafic rocks include the Canyon Mountain Complex 

(Ave Lallemant, 1976; Thayer, 1977) and small ultramafic 

rock outcrops near Sparta, in the Virtue Hills, along the 

Connor Creek fault (Brooks, 1979a), and in the Greenhorn 

Mountains (T. Vallier, oral commun., 1984). In western 

Idaho there are several serpentinized peridotite bodies 

near Riggins along the contact between the Squaw Creek 

Schist and the Lightning Creek Schist (Hamilton, 1963, 

1969, 1978; Onasch, 1977; Sarewitz, 1982), some occurrences at Asbestos Peak about 20 kilometers east of 

Whitebird (in sees. 26, 35, and vicinity, T. 28 N., R. 3 E.), 

an occurrence near West Fork Grouse Creek (sec. 31, T. 

29 N., R. 4 E.) (A. Hoover, oral commun., 1984), metagabbro and metaperidotite at Blacktail Butte near Harpster 

(Anderson, 1930; Myers, 1982), several occurrences of 

dunite extensively replaced by anthophyllite in the Glenwood area and near the mouth of Pete King Creek east 

of Kamiah (Anderson, 1930, 1931), and additional occurrences mentioned by Anderson (1930) near Ahsahka and 

Teakean in western Clearwater County. Among this 

western Idaho group, the occurrences that Anderson 

(1930, 1931) described in the Glenwood area most resemble the peridotite in the New Meadows block. He shows 

photographs of anthophyllite rosettes similar to those in 

figure 8.10. Furthermore, his description of the sequence 

in which the hydrous minerals formed at the expense of 

original olivine, first involving serpentinization and then 

the formation of anthophyllite, is identical to that which 

is described above for the New Meadows peridotite bodies. 

The writer knows of no other occurrences in western 

Idaho of quartz-rich rocks having accompanying quartzpoor layers quite like those mixed with the peridotite 

masses that have been interpreted as metacherts and 

associated pelagic metasedimentary rocks. However, 

chert is a common lithology in the dismembered oceanic 

crust terrane in northeastern Oregon (Brooks and Vallier, 

1978; Brooks, 1979a). These cherts probably formed in 

an oceanic environment far removed from a source of 

clastic sediments. The metacherts in the New Meadows 
168 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

block might be equivalent to part of the Squaw -Creek 

Schist, although neither Hamilton (1963) nor Myers (1982) 

indicates quartz-rich rocks to be an important part of that 

formation. Around Riggins, the Squaw Creek Schist is 

dominated by volcanic rocks and sedimentary rocks derived from volcanic rocks, although Hamilton (1963, p. 32) 

briefly mentions a local type of schist typically having 75 

to 95 percent quartz within the formation. 

The structural disruption of the metacherts in the New 

Meadows block suggests that, after deposition, they may 

have been tectonically incorporated, along with the 

peridotite, into a melange wedge that accumulated between an active trench and an island arc or continental 

margin. I believe this explanation is the most plausible 

for the association of metachert and peridotite, and I suggest that the rocks in the western part of the New 

Meadows block are dismembered oceanic crust-probably 

an eastward extension of the oceanic crust terrane of 

Brooks and Vallier (1978) and Brooks (1979a) (fig. 8.1). 

The plutons in the West Mountain-Council Mountain 

block resemble other tonalite and quartz diorite plutons 

along the western margin of the Atlanta lobe of the Idaho 

batholith and those plutonic outliers west of the batholith 

(Schmidt, 1964; Taubeneck, 1971; Myers, 1982; Fisher 

and others, 1983; and E.H. Bennett, oral commun., 1984), 

and thus are assumed to be related to the early development phase of the composite Atlanta lobe. The Council 

Mountain, Deserette, and West Mountain plutons clearly are part of this group. However, because the No 

Business Mountain pluton has a somewhat alkalic 

character, a relative abundance of clinopyroxene (table 

8.1), and an apparent enrichment in Sr and Ti in comparison to nearby, more typical, quartz diorite and tonalite 

plutons, it is not yet clear whether the No Business Mountain pluton is a normal pluton of the group along the 

western margin of the Idaho batholith that has assimilated 

considerable calcareous material or a part of the group 

of early Mesozoic plutons that constitutes the alkalic 

magmatic belt in western North America (Miller, 1978). 

After intrusion, the plutons in the West MountainCouncil Mountain block were metamorphosed to the 

amphibolite facies, and a fairly pervasive foliation and 

lineation, along with abundant but somewhat localized 

shears, were formed within the plutons. Foliation, lineation, and shear zones that developed concurrently in the 

enclosing metavolcanic and metasedimentary rocks 

generally obliterated whatever earlier metamorphic 

features they may have contained. The complex tectonic 

history of the area and the composite nature of the nearby Idaho batholith imply that the metamorphism may have 

lasted for several million years or more and may have 

involved more than one intense pulse of thermal or structural activity. A better understanding of the metamorphic history will require more detailed mapping, 

radiometric ages and chemical characterization of the 

rocks and minerals. 

The latter, most obvious part of the metamorphism may 

have partly been in response to the intrusion of plutons, 

including the West Mountain pluton, within the Atlanta 

lobe of the Idaho batholith. Forces accompanying this 

event could have helped form the moderately to steeply 

eastward-dipping foliaton and shear zones and the 

eastward-plunging lineation in the West Mountain-Council 

Mountain block. Such attitudes could have occurred in 

response to a diapiric rise and westward spreading of 

parts of the Idaho batholith, in the same fashion envisioned by Myers (1982, fig. 28) for the northern tip of 

the Atlanta lobe. 

Emplacement of the Atlanta lobe may have been accompanied by east-west-directed compressive stress that 

resulted from east-directed subduction of oceanic crust 

in an active trench somewhere to the west or by tectonic 

pushing or crumpling of a forearc basin or island arc 

against the North American plate. Vallier and Engebretson (1983) have suggested that such large-scale structural 

events may have occurred in western Idaho during about 

95 to 100 Ma, when the Blue Mountains island arc completed its accretion to the North American craton and 

when the first plutons of the Atlanta lobe were emplaced. 

Accordingly, the east-dipping structural and metamorphic 

features may have resulted from a shear couple operating 

around a north-south locus of horizontal axes, as the continental margin was compressed and the Idaho batholith 

was emplaced (looking northward parallel to this axial 

direction, the rotation sense during the shearing would 

have been counterclockwise). 

The deformation and metamorphism of the New 

Meadows block differ from that in the West MountainCouncil Mountain block. Although the foliations are 

similar in attitude for the two blocks (fig. 8.2), the lineations in the New Meadows block dip shallowly rather than 

steeply and are approximately at right angles to those in 

the West Mountain-Council Mountain block. These differences indicate that the structural histories of the two 

blocks were quite different, even though the reasons are 

not entirely clear. The metamorphism in the West 

Mountain-Council Mountain block appears to have 

culminated with recrystallization under amphibolite-facies 

conditions, accompanied by strong shearing but little thermal readjustment of the textures after deformation. In 

the New Meadows block, however, the metamorphic 

grade was probably within the epidote-amphibolite facies. 

Additionally, the static growth habit of undeformed anthophyllite rosettes in the peridotite and the hornfelsed 

nature of the previously strongly deformed metacherts 

are convincing evidence that the latter part of the 

metamorphism in that block was primarily thermal. The 

presence of pegmatitic to aplitic dikes cutting the 
8. PRE-CENOZOIC GEOLOGY OF THE WEST MOUNTAIN-COUNCIL MOUNTAIN-NEW MEADOWS AREA 169 

ultramafic and quartz-rich rocks (these cannot be interpreted as due to partial melting of their host rocks) suggests that the thermal metamorphism in the western part 

of the New Meadows block may be due to a hidden granitic 

pluton rather than to a regional heating event. 

Cooling to subamphibolite-facies temperatures probably 

occurred in the metamorphosed volcanic and sedimentary 

rocks and plutons after about 84 to 86 Ma, in the western 

part of the West Mountain-Council Mountain block, as 

these are the K-Ar ages of biotite from the Council Mountain and Deserette plutons (table 8.1). Cooling of the West 

Mountain pluton, to the east, to similar temperatures may 

have been more prolonged, as suggested by the biotite 

K-Ar age there of about 78 million years (table 8.1). The 

K-Ar age of approximately 120 Ma (table 8.1) on hornblende from the Council Mountain pluton indicates that 

the pluton is at least this old, according to Armstrong and 

others (1977). Similarly, the approximate 92-Ma K-Ar age 

on hornblende from the West Mountain pluton (table 8.1) 

should also be considered as a minimum age, because of 

the possible loss of radiogenic argon from hornblende between the time of its original crystallization in the pluton 

and the cooling near the end of the metamorphism. 

Because of the metamorphism, U-Pb ages on zircons will 

probably be required to obtain the true igneous ages of 

the metamorphosed quartz diorites and tonalites in the 

West Mountain-Council Mountain block. 

REFERENCES CITED 

Anderson, A.L., 1930, The geology and mineral resources of the region 

about Orofino, Idaho: Idaho Bureau of Mines and Geology Pamphlet 

34, 63 p. 

__ 1931, Genesis of the anthophyllite deposits near Kamiah, Idaho: 

Journal of Geology, v. XXXIX, no. 1, p. 68-81. 

Armstrong, R.L., 1975, The geochronometry of Idaho: Isochron/West, 

no. 14, p. 1-50. 

Armstrong, R.L., Taubeneck, W.H. and Hales, P.O., 1977, Rb-Sr and 

K-Ar geochronomtry of Mesozoic granitic rocks and their Sr isotopic 

composition, Oregon, Washington, and Idaho: Geological Society 

of America Bulletin, v. 88, p. 397-411. 

Ave Lallemant, H. G., 1976, Structure of the Canyon Mountain (Oregon) 

ophiolite complex and its implication for sea-floor spreading: 

Geological Society of America Special Paper 173, 49 p. 

Breeser, P.J., 1972, General geology and highway realignment considerations in the Pinehurst-New Meadows-Tamarack area, west-central 

Idaho: Moscow, Idaho, University of Idaho, M.S. thesis, 103 p. 

Brooks, H. C., 1979a, Plate tectonics and the geologic history of the Blue 

Mountains: Oregon Geology, v. 41, no. 5, p. 71-80. 

__ 1979b, Geologic map of Huntington and part of Olds Ferry 

quadrangles, Oregon: Oregon Department of Geology and Mineral 

Industries Geologic Map Series GMS-13, scale 1:62,500. 

Brooks, H.C., Mcintyre, J.R., and Walker, G.W., 1976, Geology of the 

Oregon part of the Baker 1 o by 2° quadrangle: Oregon Department 

of Geology and Mineral Industries Geologic Map Series GMS-7, scale 

1:250,000, 25 p. 

Brooks, H.C., and Vallier, T.L., 1978, Mesozoic rocks and tectonic evolution of eastern Oregon and western Idaho, in Howell, D.G., and 

McDougall, K.A., eds., Mesozoic paleogeography of the Western 

United States: Society of Economic Paleontologists and Mineralogists, Pacific Section, Pacific Coast Paleogeography Symposium, 

2d, p. 133-145. 

Dalrymple, G.B., 1979, Critical tables for conversion of K-Ar ages from 

old to new constants: Geology, v. 7, p. 558-560. 

Davis, G.A., Monger, J.W.H., and Burchfiel, B.C., 1978, Mesozoic construction of the cordilleran "collage," central British Columbia to 

central California, in Howell, D.G. and McDougall, K.A., eds., 

Mesozoic paleogeography of the Western United States: Society of 

Economic Paleontologists and Mineralogists, Pacific Section, Pacific 

Coast Paleogeography Symposium, 2d, p. 1-32. 

Fisher, F.S., Mcintyre, D.H., and Johnson, K.M., compilers, 1983, 

Geologic map of the Challis 1° x 2° quadrangle, Idaho: U.S. 

Geological Survey Open-File Report 83-523, scale 1:250,000, 60 p. 

Fitzgerald, J.F., 1982, Geology and basalt stratigraphy of the Weiser 

embayment, west-central Idaho, in Bonnichsen, Bill, and Breckenridge, R.M., eds., Cenozoic Geology of Idaho: Idaho Bureau of Mines 

and Geology Bulletin 26, p. 103-128. 

Hamilton, Warren, 1963, Metamorphism in the Riggins region, western 

Idaho: U.S. Geological Survey Professional Paper 436, 95 p. 

__ 1969, Reconnaissance geologic map of the Riggins quadrangle, 

west-central Idaho: U.S. Geological Survey Miscellaneous Geologic 

Investigations Map I-579, scale 1:125,000. 

__ 1978, Mesozoic tectonics of the Western United States, in Howell, 

D.G., and McDougall, K.A., eds., Mesozoic paleogeography of the 

Western United States: Society of Economic Paleontologists and 

Mineralogists, Pacific Section, Pacific Coast Paleogeography Symposium, 2d, p. 33-70. 

Miller, C.F., 1978, An early Mesozoic alkalic magmatic belt in western 

North America, in Howell, D.G., and McDougall, K.A., eds., 

Mesozoic paleogeography of the Western United States: Society of 

Economic Paleontologists and Mineralogists, Pacific Section, Pacific 

Coast Paleogeography Symposium, 2d, p. 163-173. 

Mitchell, V.E., and Bennett, E.H., compilers, 1979, Geologic map of 

the Baker quadrangle, Idaho: Idaho Bureau of Mines and Geology 

Geologic Map Series, Baker 1 o x 2° quadrangle, scale 1:250,000. 

Myers, P.E., 1982, Geology of the Harpster area, Idaho County, Idaho: 

Idaho Bureau of Mines and Geology Bulletin 25, 46 p. 

Onasch, C.M., 1977, Structural evolution of the western margin of the 

Idaho batholith in the Riggins, Idaho area: University Park, Pennsylvania, Pennsylvania State University, Ph.D. dissertation, 296 p. 

Palmer, I.F., Jr., 1963, Geology of the Council Mountain area, Adams 

County, Idaho: Moscow, Idaho, University of Idaho, M.S. thesis, 

54 p. 

Sarewitz, D.R., 1982, Geology of a part of the Heavens Gate quadrangle, 

Seven Devils Mountains, western Idaho: Corvallis, Oregon, Oregon 

State University, M.S. thesis, 144 p. 

Schmidt, D.L., 1964, Reconnaissance petrographic cross section of the 

Idaho batholith in Adams and Valley Counties, Idaho: U.S. Geological 

Survey Bulletin 1181-G, 50 p. 

Schmidt, D.L., and Mackin, J.H., 1970, Quaternary geology of Long 

and Bear Valleys, west-central Idaho: U.S. Geological Survey 

Bulletin 1311-A, 22 p. 

Streckeisen, Albert, 1979, Classification and nomenclature of volcanic 

rocks, lamprophyres, carbonatites, and melilitic rocks: Recommendations and suggestions of the lUGS Subcommission on the 

Systematics of Igneous Rocks: Geology, v. 7, p. 331-335. 

Streckeisen, Albert (chairman), and others, 1973, Plutonic rocks, 

classification and nomenclature recommended by the lUGS Subcommission on the Systematics of Igneous Rocks: Geotimes, v. 18, no. 

10, p. 26-30. 

Taubeneck, W.H., 1971, Idaho batholith and its southern extension: 

Geological Society of America Bulletin, v. 82, p. 1899-1928. 

Thayer, T.P., 1977, The Canyon Mountain Complex, Oregon, and some 
170 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

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. 

Thayer, T.P., and Brown, C.E., 1964, Pre-Tertiary orogenic and plutonic 

intrusive activity in central and northeastern Oregon: Geological 

Society of America Bulletin, v. 75, p. 1255-1262. 

Vallier, T.L., and Engebretson, D.C., 1983, The Blue Mountains island 

arc of Oregon, Idaho, and Washington: an allochthonous coherent 

terrane from the ancestral western Pacific Ocean?, in Howell, D. G., 

Jones, D.L., Cox, Allen, and Nur, Amos, eds., Proceedings of the 

Circum-Pacific Terrane Conference: Stanford University Publications in Geological Sciences, v. 18, p. 197-199. 
9. LITHOLOGIC AND CHEMICAL CHARACTERISTICS OF THE 

CENTRAL AND SOUTHEASTERN PART OF 

THE SOUTHERN LOBE OF THE IDAHO BATHOLITH 

By REEDS. LEwis, THOR H. KnLSGAARD, EARL H. BENNETT, and WAYNE E. HALL 

ABSTRACT 

The Late Cretaceous Idaho batholith consists of the Bitterroot and 

Atlanta lobes, which are separated by Precambrian metamorphic rocks. 

Mapping in the southern (Atlanta) lobe indicates the presence of several 

plutonic phases. The oldest unit is tonalite, intruded by biotite granodiorite, which is transitional into a core of muscovite-biotite granite. 

The youngest phase is fine-grained leucocratic granite. Locally, potassium metasomatism resulted in the formation of porphyritic (K-feldspar 

megacryst-bearing) granodiorite. Seven outlying stocks intrude Paleozoic 

metasedimentary rocks east of the Atlanta lobe and, with the exception of the pyroxene-bearing quartz monzodiorite of the Croesus stock, 

they are lithologically similar to the main batholithic mass to the west. 

Major-element data from 105 samples document the remarkable 

homogeneity of the Atlanta lobe; a tonalitic west margin gives way 

eastward to biotite granodiorite and muscovite-biotite granite that show 

minimal variation along a 90-km transect. Tonalite along the westernmost margin of the batholith is low in radiogenic strontium and was 

probably formed from magma generated in or above a subducting slab 

of oceanic lithosphere. Rocks to the east have high initial 87Sr/ 86Sr 

ratios and probably formed by partial melting of pre-Belt Supergroup 

basement in the lower crust, with only a minor mantle contribution. 

Rocks in the Bitterroot and Atlanta lobes of the Idaho batholith are 

similar both chemically and mineralogically. The Idaho batholith has a 

lower mafic mineral content and is higher in Si02 content than the 

Sierra Nevada batholith, and it lacks zoned plutons in sharp contact with 

one another that characterize the Sierra Nevada batholith. These differences result from the Sierra Nevada batholith being largely underlain 

by oceanic material, whereas the Idaho batholith is largely underlain 

by continental crust. In addition, the Idaho batholith may have had a 

higher water content during emplacement, which inhibited the rise of 

magma and resulted in slow cooling and diffuse contacts. 

INTRODUCTION 

The Idaho batholith, a large granitoid body that is widely exposed in central Idaho, was emplaced during the Late 

Cretaceous. It is composed of several plutonic phases, 

most being of great areal extent, that range in composition from granite to tonalite. A northwest-trending belt 

of Precambrian metamorphic rocks, part of the Salmon 

River arch as defined by Armstrong (1975a), separates 

the northern (Bitterroot) and southern (Atlanta) lobes of 

the batholith (fig. 9.1). Both lobes were intruded by 

Eocene plutons having ages of 49 to 44 Ma (Bennett, 

1980). 

Atlanta lobe 

0 100 KILOMETERS 

FIGURE 9.1.-Index map of the northern (Bitterroot) and southern 

(Atlanta) lobes of the Idaho batholith and the area of study. 

171 
172 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

This paper outlines the lithologic and chemical characteristics of Late Cretaceous plutonic rocks in an area 

south of the 45th parallel (fig. 9.1). Chemical emphasis 

is on major-element data from 105 samples collected in 

the central and southeastern parts of the Atlanta lobe. 

The samples were collected during geologic mapping of 

the Challis 1 o by 2° quadrangle by the U.S. Geological 

Survey and the Idaho Geological Survey (Fisher and 

others, 1983), the Ten Mile West Roadless area (Kiilsgaard, 1983), and part of the Hailey 1 o by 2° quadrangle 

(Hall and others, 1978; unpublished mapping by the authors). Eocene plutonic rocks have not been included in this 

study, so the reader is referred to Bennett (1980) and Bennett and Knowles (1985) for discussion of those rocks. 

Early work on the Atlanta lobe of the batholith was 

done chiefly by Ross (1928, 1934) and Anderson (1947, 

1952), who recognized several phases of Mesozoic rocks 

as well as Tertiary plutonism. Subsequent mapping by 

Schmidt (1964) in the western part of the Atlanta lobe 

near Cascade delineated four Late Cretaceous rock units, 

and Reid (1963) and Kiilsgaard and others (1970) 

separated Late Cretaceous from Tertiary rocks in the 

Sawtooth Range southwest of Stanley. Swanberg and 

Blackwell (1973), using gamma-ray spectrometry to subdivide rocks of the Atlanta lobe, recognized the extensive 

nature of the rock units that form the core of the batholith. 

Reconnaissance geology of the southwest corner of the 

batholith was performed by Taubeneck (1971). Mapping 

in the southeast part of the batholith was done by Schmidt 

(1962) and by Hall and others (1978). A comparison of the 

Idaho batholith with the southern California batholith was 

made by Larsen and Schmidt (1958); their study included 

a compilation of available chemical analyses. 

Numerous radiometric ages of the Idaho batholith, using 

mostly K-Ar dating techniques, have been published, 

including those by McDowell and Kulp (1969) and Armstrong (1975b). Criss and others (1982) outlined the complex nature of K-Ar systematics in the Idaho batholith, 

and Criss and Taylor (1983) used oxygen-isotope data to 

document the effect of Eocene hydrothermal systems on 

rocks of the Late Cretaceous batholith. Criss and Fleck 

(chapter 6) review their earlier work and interpret new 

data. An extensive review of research on the batholith 

was made by Hyndman (1983). 

GEOLOGIC SETTING 

Along its eastern margin (fig. 9.2) the Idaho batholith 

intruded Paleozoic sedimentary rocks that were metamorphosed to hornfels and calc-hornfels (Ross, 1934; Hall and 

others, 1978; Hall, 1985); some of the eastern contact is 

now covered by the Eocene Challis Volcanics (Ross, 1937; 

Mcintyre and others, 1982). 

The northern part of the Atlanta lobe intrudes various 

types of metasedimentary and metaigneous rocks, the majority of which have yet to be adequately mapped. These 

rocks are in a structural setting referred to as the Salmon 

River arch by Armstrong (1975a), who postulated an approximate age of 1,500 Ma, based on Rb-Sr systematics, 

for granitic and gneissic rocks near Shoup. Granitic rocks 

of the Salmon River arch intrude argillaceous quartzite 

of theY ellowjacket Formation (Ross, 1934; Ruppel, 1975; 

Lopez, 1982). Armstrong (1975a) suggested that the 

Y ellowjacket Formation is pre-Belt Supergroup in age, 

though most workers consider it time-correlative with the 

Prichard Formation of the lower part of the Belt Supergroup (Evans and Zartman, 1981; Lopez, 1982; Hahn and 

Hughes, 1984). Although the base of the Yellowjacket is 

not exposed, high-grade gneiss and schist found west of 

Shoup at the mouth of the Middle Fork of the Salmon 

River may represent Precambrian basement rocks older 

than the Yellowjacket Formation (Cater and others, 1973). 

Paragneiss and orthogneiss immediately north of the main 

Salmon River were tentatively assigned a pre-Belt Supergroup age (W eis and others, 1972) and may be part of the 

same basement complex. 

A northwest-trending belt of isolated metasedimentary 

roof pendants extends across the northeastern part of the 

study area (Fisher and others, 1983). These rocks are exposed chiefly on high ridge tops and are probably near 

or at the intrusive roof of the batholith. Most of these roof 

pendants are probably Paleozoic in age, but correlation 

with known formations has not been established. 

Rocks adjacent to the western margin of the batholith 

are part of an allochthonous Permian and Triassic 

volcanic-arc complex with oceanic affinities (Hamilton, 

1963; Brooks and Vallier, 1978; Myers, 1982; Lund, 1984). 

This volcanic-arc complex includes the Wallowa-Seven 

Devils terrane, composed of metavolcanic and metasedimentary rocks, and the structurally overlying, higher 

grade meta volcaniclastic rocks of the Riggins Group. This 

complex was believed to have been accreted to the old continental margin during the Triassic or Jurassic (Hamilton, 

1976; Jones and others, 1977; Brooks and Vallier, 1978). 

This assumption was based primarily on the presence of 

Triassic and Jurassic plutons intrusive into the volcanicarc complex. However, recent mapping in the suture area, 

coupled with 40 Arf39 Ar age-spectrum dating of a 118-Ma 

metamorphic event in the Riggins Group, has shown that 

juxtaposition of the arc terrane with continental rocks, 

at least in this part of Idaho, may have begun as late as 

118 Ma along a right-lateral transcurrent fault (Lund, 

1984; Sutter and others, 1984). Thus, the Triassic and 

Jurassic plutons were accreted along with the volcanic and 

sedimentary arc rocks. 

The western margin of the Idaho batholith is at the edge 

of the old continental margin, as shown by Armstrong and 
9. LITHOLOGIC AND CHEMICAL CHARACTERISTICS OF THE SOUTHERN LOBE OF THE IDAHO BATHOLITH 173 

I 

0.704 

~ ,~ 

I ' 

0.706 0 

I 

EXPLANATION 

D Volcanic rocks of southwest Idaho and 

Snake River Plain (Miocene to Recent) 

[::::)./·?/} Columbia River Basalt Group (Miocene) 

~ Challis Volcanics (Eocene) 

• Plutonic rocks (Eocene) 

~ Plutonic rocks (Late Cretaceous) 

~~ Volcanic-arc complex (Mesozoic) 

~ Sedimentary rocks 

~ (Mesozoic and Paleozoic) 

Metasedimentary rocks of Belt 

Supergroup (Middle Proterozoic) 

and basement rocks (pre-Belt age) 

/ Sr 0.704-o.706 line 

100 KILOMETERS 

I 

FIGURE 9.2.-Generalized geologic map of Idaho showing the geologic setting of the Idaho batholith. Compiled from Armstrong and others 

(1977), Bond (1978), Bennett (1980), Fisher and others (1983), and Kiilsgaard (1983). 
174 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

others (1977), who delineated a sharp break in initial 

87Srf86Sr ratios of igneous rocks in this area. A northsouth-trending line extends along the westernmost edge 

of the batholith, with initial Sr ratios of greater than 

0.7055 to the east and ratios less than 0.7043 to the west 

(0.704-0.706line, fig. 9.2). More detailed work by Fleck 

and Criss (1985) shows a continuum of initial 87Srf86Sr 

ratios across this line, but the transition zone is nevertheless narrow, typically no more than several hundred 

meters wide. 

Intrusive relations in the vicinity of the western margin 

of the Idaho batholith are in large part obscured by extensive flows of the Miocene Columbia River Basalt 

Group. Miocene and younger volcanic and sedimentary 

rocks of the Snake River Plain overlap the entire southern 

margin of the Atlanta lobe, masking intrusive features 

in this area. 

LITHOLOGY OF THE CENTRAL ATLANTA LOBE 

A peculiarity of the Atlanta lobe of the Idaho batholith 

is the absence of numerous small, separate granitic 

plutons that typify other batholiths. Instead, the plutons 

are very large, approximately 1,000 to 10,000 km2 in 

area, and generally display gradational contacts between 

each other. Even where exposures are good, sharp contacts are rare. 

Mapping in the central part of the Atlanta lobe (Fisher 

and others, 1983) indicates that the batholith can be subdivided into six major rock types of Late Cretaceous age. 

The rock types are tonalite, hornblende-biotite granodiorite, porphyritic granodiorite, biotite granodiorite, 

muscovite-biotite granite, and leucocratic granite. A brief 

summary of the field relations and lithologies of these 

units is given in the following subsections, and a discussion of the geochronology is presented in the section 

"Geochronology of the Central Atlanta Lobe." More 

detailed rock descriptions can be found in Kiilsgaard and 

Lewis (1985). Granitic rocks are classified in this report 

according to recommendations of Streckeisen and others 

(1973). 

TONALITE 

Tonalite grading to granodiorite has been mapped along 

the northwestern margin of the Atlanta lobe and, to a 

lesser extent, in the eastern part of the lobe (fig. 9.3). A 

large mass of unmapped tonalitic rock is present west of 

long 116 ° W. and is well exposed near Banks. Tonalitic 

and gneissic rocks in the McCall area have been studied 

by Kuntz and Allen (1984) and are presently under investigation by them (oral commun., 1985). 

Field evidence and radiometric age determinations indicate that tonalite is the oldest of the batholithic rocks. 

Northeast of Banks tonalite xenoliths are enclosed by 

younger biotite granodiorite. The tonalite is typically dark 

gray, medium grained, and equigranular, with a varying 

foliation defined by biotite. Hornblende content is generally 2 to 8 percent, and biotite, 5 to 15 percent. Sphene is 

a conspicuous accessory mineral and magnetite is common. Modally the rock plots as tonalite grading to granodiorite (fig. 9.4), with a mean composition, excluding mafic 

and opaque constituents, of 29 percent quartz, 8 percent 

potassium feldspar, and 63 percent plagioclase. 

HORNBLENDE-BIOO'ITE GRANODIORITE 

Small exposures of hornblende-biotite granodiorite are 

found on the high ridges in the central part of the Atlanta 

lobe (fig. 9.3). This granodiorite is typically gray, medium 

grained, and slightly foliated. It has more potassium 

feldspar and less hornblende than the tonalite, but in 

general appearance the rocks are similar. Modally it is a 

granodiorite (fig. 9.4), with a mean composition, excluding 

mafic and opaque minerals, of 28 percent quartz, 14 percent K-feldspar, and 58 percent plagioclase. Hornblende 

content is generally 1 to 4 percent, and biotite, 8 to 15 

percent. Intrusive relations are ambiguous, but this unit 

probably represents a discontinuous roof phase that is 

slightly older than the biotite granodiorite that lies 

beneath it. No radiometric ages have been obtained for 

the hornblende-biotite granodiorite. 

PORPHYRITIC GRANODIORITE 

A discontinuous, northwest-trending belt of porphyritic 

(K-feldspar megacryst-bearing) granodiorite 125 km long 

extends across the northeastern part of the study area 

(fig. 9.3). This unit crops out primarily on ridge tops, 

similar to the hornblende-biotite granodiorite. The porphyritic granodiorite is gray, medium to coarse grained, 

and characterized by pink or light-pinkish-gray microcline 

megacrysts as long as 10 em. Locally biotite defines a 

foliation. Modally the rocks plot as granodiorite grading 

to granite (fig. 9.4), with a mean composition, excluding 

mafic and opaque constituents, of 30 percent quartz, 22 

percent potassium feldspar, and 48 percent plagioclase. 

Biotite is the primary mafic constituent, but as much as 

5 percent hornblende is present in many places, particularly north and east of Stanley. The porphyritic granodiorite that contains hornblende is similar to the tonalite 

described earlier, differing mainly in that it contains 

megacrysts. Where hornblende is absent the rock more 

closely resembles biotite granodiorite. Ross (1963) 

postulated formation of the megacrysts late in the origin 

of the rock, probably as a result of fluid circulating 

through the rock after it had consolidated but while it remained hot. Olson (1968) and Criss (1981) thought the 

megacrysts may have resulted from similar subsolidus 
9. LITHOLOGIC AND CHEMICAL CHARACTERISTICS OF THE SOUTHERN LOBE OF THE IDAHO BATHOLITH 175 

deuteric porphyroblastic growth. Lewis (1984) and 

Kiilsgaard and Lewis (1985) described the megacrysts as 

resulting from potassium metasomatism. Numerous 

potassium feldspar-rich veinlets cut the rock and appear 

0 

to have been a source of potassium-rich fluid for the 

metasomatism. This metasomatism affected both the 

tonalite and the biotite granodiorite and was most pronounced near the roof of the batholith. 

EXPLANATION 

D Pre- and post-batholithic 

rocks 

1::::::::::1 ~~~~:~~:~~t~~~ekds 

~Muscovite-biotite granite 

~ Biotite granodiorite 

[{~{;) Porphyritic granodiorite 

• 

Hornblende-biotite 

granodiorite 

• Thompson Creek stock 

ilclayton 

White Cloud stock 

Little Boulder 

Creek stock 

Rook's Creek stock 

Summit 

FIGURE 9.3.-Simplified geologic map of Cretaceous rocks in the Atlanta lobe of the Idaho batholith. Also shown is the location of transect 

A-A'. Compiled from Bond (1978), Fisher and others (1983), Kiilsgaard (1983), and unpublished mapping by the authors. 
176 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

BIOTITE GRANODIORITE 

Biotite granodiorite is the most extensive rock unit in 

the Atlanta lobe (fig. 9.3). It is light gray, medium to 

coarse grained, and locally porphyritic, containing K-feldspar phenocrysts 2 em or less in length. Foliation is rarely present within this unit, and where seen may have 

resulted from the movement of one mass of granodiorite 

magma into another. Modally the unit consists of granodiorite grading to granite (fig. 9.4), with a mean composition, excluding mafic and opaque constituents, of 30 

percent quartz, 22 percent potassium feldspar, and 48 percent plagioclase. The biotite content is generally 2 to 8 

percent, and hornblende and primary muscovite are rarely 

present. Field relations indicate that the biotite granodiorite is younger than the tonalite and the hornblendebiotite granodiorite. The biotite granodiorite is gradational into muscovite-biotite granite, and together these 

units comprise what is consjdered the main phase of the 

Atlanta lobe. 

MUSCOVITE-BIOTITE GRANITE 

A discontinuous, north-south-trending body of muscovite-biotite (2-mica) granite is present in the core of the 

southern lobe of the batholith (fig. 9.3). It grades into 

biotite granodiorite over a zone as wide as 2 km or more. 

The 2-mica granite is massive, light gray, medium to 

coarse grained, and equigranular to slightly porphyritic. 

The presence of books of primary muscovite was the principal field criterion used in mapping this unit. Muscovite 

comprises as much as 5 percent of the rock, but generally averages about 2 percent, as does biotite. Garnet is 

present locally. Modally the unit consists of granite and 

granodiorite (fig. 9.4), with a mean composition, excluding 

mafic minerals, of 32 percent quartz, 23 percent K-feldspar, and 45 percent plagioclase. The radiometric ages 

and transitional contact with biotite granodiorite suggest 

that this unit is slightly younger than the biotite 

granodiorite. 

LEUCOCRATIC GRANITE 

Leucocratic granite (not indicated on fig. 9.3, for locations refer to Fisher and others, 1983) is present as dikes, 

sills, and irregular small stocks that cut sharply across 

all earlier phases of the batholith. The granite is light gray, 

fine grained, and equigranular and is generally considered 

to be a late differentiate of the batholith (Anderson, 1947; 

Reid, 1963; Luthy, 1981). The leucocratic granite was 

mapped as alaskite by Cater and others (1973). Modally 

the rock plots as granite (fig. 9.4), with a mean composition of 31 percent quartz, 33 percent K-feldspar, and 36 

percent plagioclase. The leucocratic granite generally contains 2 percent or less biotite, and trace amounts of 

primary muscovite. Locally, on the east side of the batholith, the granite contains as much as 10 percent biotite. 

This biotite is very fine grained, however, so the rock retains its light color. Garnet grains as large as 2 mm in 

diameter are present locally. 

GEOCHRONOLOGY OF 

THE CENTRAL ATLANTA LOBE 

Application of radiometric dating techniques in the central Atlanta lobe has met with only limited success. Most 

of the radiometric ages available were done by the K-Ar 

method, and they give only a minimum age of emplacement. Some of the K-Ar ages are too young due to argon 

loss during Eocene reheating (Armstrong, 197 4), and all 

of them record only the time since the rock cooled past 

the argon-blocking temperature of the mineral dated. 

Criss and others (1982) studied the effects of both Eocene 

plutonism and regional uplift on K-Ar ages of biotite and 

found that argon loss was primarily confined to areas close 

to Tertiary plutons, and that the rest of the "ages" could 

be accounted for by slow (about 0.1-0.14 mm/yr) uplift 

of the batholith. Uplift rates of much higher magnitude 

have been postulated by Snee and others (1985) for Late 

Cretaceous batholithic rocks in the Elk City area in the 

northern part of the Atlanta lobe. 

The K-Ar biotite ages that have been determined for 

tonalitic rocks on the west side of the Atlanta lobe range 

from 82 to 67 Ma (table 9.1; Armstrong, 1975b; Criss and 

others, 1982). Tonalite on the east side of the lobe 

southeast of Stanley yielded an age of 81.2 ± 2.9 Ma on 

biotite (table 9.1). This wide age spread is typical in the 

Idaho batholith, and it illustrates the difficulty encountered with using the K-Ar method to date these rocks. 

Even the age of 82 Ma represents cooling below the blocking temperature of biotite (about 280 ± 40 °C, Harrison 

and McDougall, 1980; Snee, 1982), so true emplacement 

ages are older, probably in the range of 85 to 95 Ma. 

Available geochronologic data for the tonalite and the 

other five rock units in the central Atlanta lobe are summarized in figure 9. 5. 

A sample of porphyritic (K-feldspar megacryst-bearing) 

granodiorite from the east side of the Atlanta lobe east 

of Stanley yielded a U-Pb zircon age of 88 ± 6 Ma (Lynn 

Fischer, written commun., 1983). Another sample from 

the same outcrop gave a K-Ar hornblende age of 84.7 ±2.9 

Ma and a K-Ar biotite age of 79.8 ± 2.7 Ma (table 9.1). This 

granodiorite, sampled 13 km east of Stanley at milepost 

201, U.S. Highway 75, is a metasomatically altered 

tonalite. A sample of the same unit south of Yell ow Pine 

yielded a K-Ar biotite age of 77.9 ± 2.8 Ma (table 9.1). The 

88 ± 6 Ma U-Pb zircon age approximates the age of 

emplacement of the porphyritic granodiorite (fig. 9.5). 

K-Ar ages on biotite from the biotite granodiorite unit 

span a wide range (table 9.1; Armstrong, 1975b; Criss and 
9. LITHOLOGIC AND CHEMICAL CHARACTERISTICS OF THE SOUTHERN LOBE OF THE IDAHO BATHOLITH 177 

others, 1982). Samples from the central and western part 

of the Atlanta lobe give apparent ages ranging from 79 

to 54 Ma. Many of these rocks have probably undergone 

partial Ar loss. An area of intense Ar loss (greater than 

75 percent) surrounds the Eocene plutons, and K-Ar ages 

there range from 46 to 37 Ma (Criss and others, 1982). 

The actual age of the biotite granodiorite is probably between 85 and 75 Ma (fig. 9.5), but the present data do not 

adequately constrain the age of this unit. Intrusion of the 

biotite granodiorite may have directly followed the 

tonalite magmatism, or several million years may have 

passed between intrusive events. 

K-Ar ages on biotite and muscovite in the 2-mica granite 

are generally concordant, with muscovite giving apparent 

ages as much as 4 m.y. older than biotite in the same rocks 

(table 9.1). Dated samples in this study have muscovite 

ages of 76 to 69 Ma. 40 Arf39 Ar determinations on 2-mica 

granite west of Landmark and west of Yell ow Pine gave 

ages of 72 and 70 Ma, respectively, but geologic relations 

in these areas are not known well enough to interpret 

these as emplacement or cooling ages (L.W. Snee, oral 

commun., 1985). In addition, muscovite-bearing biotite 

granite 14 km east of Yellow Pine at Stibnite is cut by 

mineralized veins dated by Snee at 78 Ma. Excluding the 

rocks at Stibnite, the actual emplacement age of the 

muscovite-biotite granite is probably between 78 and 72 

Ma, though it may be slightly older (fig. 9.5). 

K-Ar ages on biotite in the leucocratic granite range 

from 76 to 64 Ma, and most are from 73 to 70 Ma (table 

9.1; Criss, 1981). The young (65 to 64 Ma) ages are from 

samples collected near Eocene plutons. Emplacement age 

is probably 75 to 70 Ma (fig. 9.5). 

LITHOLOGY OF THE SOUTHEASTERN PART 

OF THE ATLANTA LOBE 

The batholith immediately north and west of Dollarhide 

Summit (fig. 9.3) contains biotite granodiorite and subordinate hornblende-biotite granodiorite, but more mapping 

is necessary to differentiate these units. Both appear 

similar to previously described granodiorites in the central part of the study area. A sample of the biotite granodiorite (H22, table 9.1) yielded a K-Ar age on biotite of 

60.7 ± 1.8 Ma. This rock almost certainly underwent Ar 

QUARTZ 

GRANITE 

K-FELDSPAR 

EXPLANATION 

IJ. Leucocratic granite 

• Muscovite-biotite 

granite 

0 Biotite granodiorite 

+ Porphyritic granodiorite 

• Hornblende-biotite 

granodiorite 

D Tonalite 

PLAGIOCLASE 

FIGURE 9.4.-Modal plot of rock samples from the central part of the Atlanta lobe of the Idaho 

batholith. Classification is that of Streckeisen and others (1973). 
178 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

TABLE 9.1.-K-Ar ages of Cretaceous rocks of the Idaho batholith 

[All samples were collected by the authors except BL1-BL4, which were collected by B.F. Leonard, U.S. Geological Survey, Denver, Colo. Rock unit symbols: Kt, tonalite; Kgdp, 

porphyritic (K-feldspar megacryst-bearing) granodiorite; Kgdh, hornblende-biotite granodiorite; Kgd, biotite granodiorite; Kg, muscovite-biotite granite; Klg, leucocratic granite; 

Peg, pegmatite. Minerals analyzed: B, biotite; M, muscovite; H, hornblende. Analysts: RM, R.F. Marvin; HM, H.H. Mehnert; VM, V.M. Merritt; KF, K. Futa; JK, J.J. Kenny; 

JG, J. Graves; EB, E.L. Brandt; RF, R.J. Fleck; JVE, J.C. Von Essen] 

sample Rock 

unit 

Age, 

Ma 

Mineral K20, 

weight 

ercent 

BOl 

Bl9 

BLl 

BL2 

K08 

K09 

B20 

B21 

H09 

H03 

B03 

B04 

B05 

B22 

B23 

H22 

R04 

B24 

B25 

B26 

BL3 

BL4 

K52 

B29 

B30 

B31 

B32 

K53 

B35 

B33 

Bl6 

B34 

H23 

Kt 

Kt 

Kt 

Kt 

Kt 

Kt 

81.6 ±. 2.9 

81.2 ±. 2.9 

78.9 ±. 2.6 

71.9 ±. 2.5 

73.6 ±. 3.0 

72.1 ±. 1.6 

Kgdp 79.8 ±_ 2.7 

84.7 ±. 2.9 

Kgdp 77.9 ±_ 2.8 

Kgdh 76.7 ±_ 1.5 

114 + 2 

Kgdh 84.0 ±_ 2.5 

Kgd 73.0 ±_ 2.5 

Kgd 72.7 ±. 2.6 

Kgd 78.9 ±_ 2.8 

Kgd 76.3 ±_ 2.6 

Kgd 74.9 ±. 2.5 

Kgd 60.7 ±_ 1.8 

Kgd 81.3 ±_ 2.9 

Kg 72.1 ±_ 2.6 

Kg 

Kg 

Kg 

Kg 

Kg 

72.2 ±. 1.7 

73.1 + 2.6 

74.7 ±. 2. 7 

73.9 ±. 2.7 

75.8 ±. 2.7 

69.8 ±. 2.3 

74.1+1.7 

69.5 ±. 2.3 

70.7 ±. 2.3 

66.2 ±. 0.9 

68.8 ±. 0.2 

Klg 70.3 ~ 2.5 

Klg 

Klg 

Klg 

Klg 

72.6 ~ 2.5 

63.6 ±. 1.4 

72.1 ~ 1.7 

65.1 ~ o.8 

Klg 75.0 ~ 2.7 

Klg 86.6 ±. 3.1 

Kgdp 84.7 ±. 1.9 

Klg 83.2 ±_ 2.8 

Peg 76.4 ±. 1.5 

B 

B 

B 

B 

B 

B 

B 

H 

B 

B 

H 

B 

B 

B 

B 

B 

B 

B 

B 

B 

M 

B 

M 

B 

M 

B 

M 

B 

M 

B 

M 

B 

B 

B 

B 

B 

M 

B 

B 

B 

M 

7.635 

8.41 

8.635 

7.555 

9.445 

9.510 

9.46 

1.075 

7.39 

8.57 

0.941 

7.71 

8.35 

8.35 

8.32 

7.035 

8. 745 

8.26 

8.215 

7.495 

10.05 

8.60 

8.555 

5.585 

8.495 

8.13 

10.38 

7.69 

8.27 

8.945 

10.59 

6.385 

8.32 

7.019 

5.046 

6.500 

9.43 

7.955 

7.445 

7.575 

10.62 

40Arrad, Analysts 

percent 

Latitude 

(North) 

9.175 

10.06 

10.03 

7.977 

10.516 

9.9004 

9.9205 

84 

86 

80 

77 

73 

84 

67 

10.233 22 

11.12 

1.342 

8.471 

9.641 

1.600 

9.547 

8.959 

8.914 

9.667 

7.899 

9.635 

7.339 

9.832 

7.933 

10.67 

9.243 

9.398 

6.066 

9.472 

8.333 

11.30 

7.847 

8.585 

8.6049 

83 

88 

88 

79 

77 

81 

82 

78 

91 

92 

87 

74 

89 

86 

85 

81 

81 

86 

80 

87 

83 

75 

86 

78 

8.7764 80 

10.0715 95 

10.0655 90 

6.591 

8.876 

6.546 

5.343 

6.1478 

83 

89 

84 

49 

74 

6.2575 32 

10.40 

10.16 

9.290 

9.285 

11.91 

84 

91 

88 

85 

90 

RM,HM,EB 

RM 

RM,HM,VM 

RM,HM,VM 

RF 

RF 

RF 

RM,HM,VM,JK 

RM 

JVE 

JVE 

RM,HM,VM,KF 

RM,HM,JG 

RM,HM,EB 

RM,HM,VM,KF 

RM,HM,JK,VM 

JVE 

RM 

RM,HM,EB 

RM,HM,EB 

RM,HM,EB 

RM,-HM,VM 

RM,HM,VM 

RF 

RM,HM,JG 

44°40 1 55" 

44°10 1 2P 

44°55 1 25" 

44°56 1 00" 

44°05.6 1 

44°47 1 18" 

43°31.8 I 

43°32 o5 I 

44°42 1 17" 

44°43 1 36" 

44°35 1 4P 

44 °34 1 19" 

44°36 1 49" 

43°38.5 1 

44°17 1 13" 

44°57 1 25" 

~M,HM,VM,JK 44°41 1 00" 

RM,HM,JK,KF,VM 44°11 1 02" 

RM,HM,JK,KF,VM 44°36 1 52" 

RF 44°08'10" 

RM 

RM,HM,EB 

RM,HM,VM,KF 

RM,HM,VM,KM 

JVE 

44°35'54" 

44°19'05" 

44°05 '45" 

44°06'26" 

43°36.6' 

Longitude 

(West) 

115°59 1 2P 

114°48'26" 

116°01'00" 

115°56 '40" 

116°06.5. 

115°31'23" 

114°31.8' 

114°38.5' 

115 °49. 30" 

115°38 1 2P 

115°57'10" 

115°01'30" 

115°23'46" 

114°51.0 1 

114°43'24" 

115°38'00" 

115°53'24" 

115°30'07" 

115°05'57" 

115°23 '42" 

115°18 '20" 

114°52'49" 

114°32 '47" 

114°35'26" 

114°36'23" 

114 °38 .1' 

Location 

NE. of cascade. 

SE. of Stanley. 

E. of Mccall. 

E. of Mccall. 

E. of Banks. 

E. of Banks. 

E. of Stanley 

(Milepost 201). 

s. of Yellow Pine. 

w. of Ketchum. 

w. of Ketchum. 

E. of cascade. 

NW. of Landmark. 

sw. of Landmark. 

NW. of stanley. 

E. of Landmark. 

w. of Ketchum. 

E. of stanley. 

w. of Yellow Pine. 

NE. of cascade. 

w. of Landmark. 

w. of Yellow Pine. 

w. of Yellow Pine. 

s. of Landmark. 

NE. of cascade. 

NW. of Landmark. 

sw. of stanley. 

E. of Landmark. 

E. of Lowman. 

Lost Packer Mine. 

Thompson creek 

stock. 

White cloud stock. 

White Cloud stock. 

w. of Ketchum. 
9. LITHOLOGIC AND CHEMICAL CHARACTERISTICS OF THE SOUTHERN LOBE OF THE IDAHO BATHOLITH 179 

loss during Eocene reheating. Biotite granodiorite and 

hornblende-biotite granodiorite southeast of Dollarhide 

Summit, termed the Hailey granodiorite unit by Schmidt 

(1962), are separated from the main batholith mass by the 

Challis Volcanics. These rocks are light gray, medium to 

coarse grained, equigranular to porphyritic, and probably 

continuous with the rest of the batholith. 

K-Ar ages of 76.7 ± 1.5 Ma (biotite) and 114 ± 2 Ma (hornblende), as well as 84.0 ± 2.5 Ma (biotite), were obtained 

from two samples southeast of Dollarhide Summit (H09 

and H03, table 9.1). The 114-Ma age on hornblende is 

unusually high and may be the result of incorporation of 

additional argon into the hornblende. If correct, it 

represents the oldest (by 17 m.y.) K-Ar age yet determined for the main batholith. Additional geologic mapping is needed to determine the exact relation between 

these southeastern batholithic rocks and those in the central and western part of the Atlanta lobe. 

SATELLITE PLUTONS 

Seven outlying stocks intrude Paleozoic metasedimentary rocks east of the main Atlanta lobe (fig. 9.3). Modes 

EXPLANATION 

Mineral dated 

• Muscovite 

-Emplacement age 

____ Range of 

apparent ages 

• Biotite > 

this study 

• Hornblende 

eZircon 

o Biotite (previous studies) 

-;---;. ~---; -.--?·----? K t 

-----?·----

0 0 • 

-------------?----

0 00 «liD co • •• • • 

I 

60 

----?-? 

• • . • • 6. Klg 

I 

70 80 

AGE,Ma 

• 

Kgd 

I 

90 

Kgdp 

FIGURE 9.5.-Radiometric ages for tonalite (Kt), porphyritic granodiorite 

(Kgdp ), biotite granodiorite (Kgd), 2-mica granite (Kg), and leucocratic 

granite (Klg) in the central part of the Atlanta lobe of the Idaho 

batholith. Data from this study and Criss and others (1982, tables 1 

and 2). Biotite, muscovite, and hornblende were dated by the K-Ar 

method. Zircon age was determined by U-Pb systematics. 

of most of these stocks are given in figure 9.6. The 

descriptions that follow are of the stocks, from north to 

south. 

THOMPSON CREEK STOCK 

The Thompson Creek stock consists of medium-grained 

equigranular biotite granodiorite that is intruded by equigranular to porphyritic biotite granite (Hall and others, 

1984). Both of these units are cut by leucocratic granite. 

Potassic alteration is imprinted on all rock types. The 

Thompson Creek stock contains ore-grade molybdenite 

mineralization and is the site of the open-pit Thompson 

Creek mine. A K-Ar age of 86.6 ± 3.1 Ma (biotite) was obtained as part of this study from a sample of the leucocratic granite (table 9.1). D.L. Schmidt (oral commun., 

1980) reported an age of 88.4 ± 3 Ma on the biotite 

granodiorite. Muscovite in the ore body yielded 40 Arf39 Ar 

ages of 87.35 ± 0.43 Ma and 87.58 ± 0.31 Ma (L. Snee, written commun., 1985). 

WHITE CLOUD STOCK 

The White Cloud stock is a zoned intrusion with a porphyritic (K-feldspar megacryst-bearing) granodiorite core 

and an outer rim of equigranular biotite granodiorite. The 

megacrysts are similar to, though not as large as, those 

found in the porphyritic granodiorite unit of the main 

batholith. Locally the stock is cut by dikes of leucocratic 

granite. K-Ar (biotite) ages of 83.2±2.8 Ma and 84.7 ± 1.9 

Ma were obtained on two samples of the White Cloud 

stock collected during this study (table 9.1). 

LITTLE BOULDER CREEK STOCK 

The Little Boulder Creek stock is a small, poorly exposed granitoid body that lies near the head of Little 

Boulder Creek 300 m east of the White Cloud stock. The 

stock is exposed for 600 m in a north-south direction and 

about 240 min an east-west direction. The stock is sheared 

and intensely silicified and has undergone potassic alteration. The least altered rock is a medium-grained, equigranular leucocratic granitoid that contains 40 to 50 

percent plagioclase, 40 to 45 percent quartz, minor 

potassium feldspar, and about 5 percent amphibole. Locally the plagioclase has been replaced nearly entirely by 

potassium feldspar. The stock is associated with disseminated molybdenite (Little Boulder Creek deposit) in 

adjacent hornfelsed Paleozoic quartzite. The Little 

Boulder Creek stock is probably a satellitic body of the 

equigranular border facies of the White Cloud stock, but 

it has been so sheared, silicified, and altered that its 

original composition is uncertain. 
180 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

ROOK'S CREEK STOCK 

The Rook's Creek stock ranges in composition from 

biotite granodiorite at the margins to biotite granite in 

the interior. The stock is light gray, medium to fine 

grained, and generally equigranular; locally it is porphyritic. Marvin and Dobson (1979) obtained a K-Ar age 

of 92.2 ± 2.2 Ma on biotite from this stock. The western 

and central parts of the Rook's Creek stock have undergone intense argillic and potassic alteration associated 

with the intrusion of dikes and the intrusion of a finegrained, medium-gray, hornblende-biotite granodiorite 

stock of probable Eocene age. 

DEER CREEK STOCK 

The Deer Creek stock was first delineated by Lindgren 

(1900). The stock ranges in composition from fine-grained, 

equigranular biotite-hornblende tonalite at the margin to 

coarse-grained, porphyritic biotite granite in the interior. 

Equal amounts of biotite and hornblende, and very minor 

augite, are present in the marginal rocks. Euhedral 

K-feldspar phenocrysts in the interior of the pluton are 

6 to 8 mm long. 

CROESUS STOCK 

Originally described by Lindgren (1900) and later by 

Anderson and others (1950), the rocks of the Croesus 

stock do not resemble any other plutonic rocks in this part 

of Idaho. The Croesus stock consists of medium- to darkgray quartz monzodiorite that contains hypersthene, 

augite, hornblende, and biotite. It is intruded by the 

McCoy Mine stock. The Croesus stock may actually be 

older than, and unrelated to, the Late Cretaceous Idaho 

batholith. 

McCOY MINE STOCK 

The McCoy Mine stock, consisting of light-gray, 

medium- to coarse-grained, equigranular to porphyritic 

biotite granodiorite, intrudes the Croesus stock. The 

McCoy Mine stock is part of what Schmidt (1962) termed 

the Hailey granodiorite unit. Described by Lindgren 

(1900) and Anderson and others (1950) as a quartz monzonite, this stock is very similar to biotite granodiorite 

found in the main batholith to the west. A K-Ar age on 

biotite of 83.8 ± 2.5 Ma (recalculated from Berry and 

others, 1976) was obtained from a sample collected at the 

McCoy mine. 

EMPLACEMENT DEPTH 

The depth of emplacement of the Idaho batholith has 

been discussed by Hyndman (1981, 1983), who suggested 

a depth of 17 to 20 km for the base of the Bitterroot lobe. 

He based this estimate on: (1) the presence of muscovitebearing granite that had to form at pressures greater than 

the intersection of the muscovite breakdown curve and 

the granite solidus (rv5 kbar), (2) an estimated 17- to 

21-km-thick cover of Belt metasedimentary rocks, and (3) 

emplacement shortly after culmination of regional highgrade metamorphism. The composition of the muscovite 

in the Idaho batholith has not been well established, but 

samples from the Bitterroot lobe have a high celadonite 

component (M. Toth, oral commun., 1985). Experimental work by Centanni (1985) has shown that muscovite 

with a high celadonite component has an extended stability field when compared to ideal muscovite. The presence 

of the celadonite component reduces the minimum 

pressure restriction on the equilibrium between muscovite 

and a granitic melt and requires a re-evaluation of depth 

estimates of the Idaho batholith. Although sillimanitegrade metasedimentary rocks are present as roof pendants in the Atlanta lobe, Smitherman (1985) found only 

contact metamorphism associated with the batholith in 

the Yell ow Pine area, and he attributed higher grade 

mineral assemblages to a prebatholithic event. A depth 

of emplacement of 12 ± 3 km was estimated for most of 

the Atlanta lobe, using heat-generation values (Swanberg 

and Blackwell, 1973), but a recent 40 Arf39 Ar agespectrum analysis indicates a depth of emplacement of 

between 4 and 9 km for 2-mica granite north of the study 

area in the Buffalo Hump region (Lund and others, 1986). 

Resolution of the emplacement depth problem requires 

more study. 

The Idaho batholith may at least locally be thinner along 

the east side than along the west. Tonalitic gneiss, migmatite, and other rocks interpreted to lie beneath the 

batholith have been described in the deep canyon of the 

Middle Fork of the Salmon River (Hamilton, 1978), in the 

south-central part of the Bitteroot lobe (Wiswall, 1979), 

and along the northeastern margin of the Bitteroot lobe 

(Chase, 1973). 

GEOCHEMISTRY 

Over 140 rock samples were analyzed for ten major 

elements, and limited trace-element determinations were 

made on 25 of these samples. Of the original analyses, 

105 are reported in table 9.2. The remaining samples had 

unusually low or high total oxides, had undergone 

moderate to extreme alteration, or had large errors in 

replicate samples. The latter case involved many samples 

of porphyritic (megacryst-bearing) rock, where the size 

and varied distribution of 5- to 10-cm-long microcline 

megacrysts made representative sampling extremely difficult. Also listed in table 9.2 are CIPW normative calculations; sample locations are shown in figure 9.7. 
9. LITHOLOGIC AND CHEMICAL CHARACTERISTICS OF THE SOUTHERN LOBE OF THE IDAHO BATHOLITH 181 

Major-element oxides were determined mostly by X-ray 

fluorescence spectroscopy (XRF) at the U.S. Geological 

Survey laboratory in Denver, Colorado. Twenty samples 

were analyzed by the XRF method at the Washington 

State University laboratory in Pullman, Washington. 

Samples H02 to H22 were analyzed in the U.S. Geological 

Survey laboratory in Reston, Virginia, using spectrophotometry and atomic-absorption spectrometry as 

described by Shapiro (1975). Trace elements were determined in the U.S. Geological Survey laboratory in Denver, 

Colorado, using inductively coupled plasma spectroscopy 

(for Ba, Sr, Y, Zr) and atomic-absorption spectrometry 

(for Rb), except for samples H02, H04, H09, H21, and 

H22, which were analysed by XRF methods in the U.S. 

Geological Survey laboratory in Reston, Virginia. Samples 

with trace-element data are limited in number and widely distributed. Thus, attempts at interpreting the results 

will have to await further sampling and analysis. 

The chemical analyses show that rocks of the tonalite 

unit vary the most in SiOz content, ranging from 60.6 to 

68.7 percent with a standard deviation (o) of 2.4 percent. 

The muscovite-biotite granite and biotite granodiorite 

units exhibit minimal variation between samples as shown 

by the low standard deviations among the various oxides. 

This homogeneity is a significant feature of the main 

phase of the southern lobe of the batholith, and it reflects 

the unusually large and chemically monotonous units. The 

limited chemical variation is compatible with geophysical 

data, which also indicate that the main phase of the 

southern lobe of the batholith is a remarkably uniform 

rock (Mabey, 1982). Large parts of the northern lobe of 

the batholith also are homogeneous, both chemically and 

mineralogically (Hyndman, 1984). 

To illustrate the variation of major-element oxides in 

rocks of the Atlanta lobe a chemical transect was made 

across the central part of the study area (A to A', fig. 9.7) 

from west to east. Oxide contents from samples 10 km 

north and south of line A-A' (fig. 9. 7) were plotted against 

distance, as shown in figure 9.8. Sample locality elevations range from a high of 2,910 m (9,550 ft) to a low of 

975 m (3,200 ft) along this transect. Leucocratic granite 

was omitted from the projection because the samples of 

QUARTZ EXPLANATION 

K-FELDSPAR 

GRANITE 

EB 

QUARTZ 

MONZONITE 

• Biotite granodiorite 

• Hornblende:biotite granodionte 

0 Thompson Creek stock 

6. White Cloud stock 

Ell Little Boulder Creek 

stock 

D Rook's Creek stock 

.A. McCoy Mine stock 

PLAGIOCLASE 

FIGURE 9.6.-Modal plot of rock samples from the southeastern part of the Atlanta lobe of the Idaho batholith 

and satellite plutons along its eastern margin. Plot is only of those rocks analyzed for major elements 

in this study. Classification is that of Streckeisen and others (1973). 
TABLE 9.2.-MaJ·or- and trace-element analyses and CIPW norms of selected samples from study area 

[Analysts: J.S. Walhberg, J.E. Taggart, J.D. Baker, A.J. Bartel, K. Stewart, L.L. Jackson, G.R. Mason, H.G. Neiman, E. Engleman, and D.B. Hatfield, U.S. Geological Survey, Denver, Colo.; H. Smith, G. Holmes, Z.A. Hamlin, N. Skinner, 

K. Coates, H.J. Rose, J. Lindsay, B. McCall, G. Sellers, and R. Johnson, U.S. Geological Survey, Reston, Va.; J. Amistoso, F.J. Moye, S.T. Luthy, W.B. Strowd, and D. Johnson, Washington State University, Pullman, Wash. Leaders 

( ---), not determined] 

sample 

number 

Si02 

Al203 

FeT03l 

Fe203 

FeO 

MgO 

cao 

Na2o 

K20 

Ti02 

P205 

MnO 

H20+ 

H20-

C02 

LOI 2 

Total 

Ba 

Rb 

sr 

y 

zr 

Q 

or 

Ab 

An 

c 

Di 

Hy 

Mt 

Il 

Ap 

KOl 

63.2 

16.7 

l. 31 

3.a3 

l. 75 

4.38 

3.57 

l.a9 

1.19 

.H 

.l 

_d! 

98.7 

22.8 

11.4 

30.7 

19.4 

l.a 

8.7 

1.9 

2.3 

1.0 

K02 

62.6 

18.0 

1.14 

2.3a 

1.51 

5.14 

3.65 

2.70 

.73 

.22 

.07 

.41 

9a.6 

17.6 

16.3 

31.5 

24.5 

.3 

6.2 

1.7 

1.4 

.5 

K03 

63.9 

17.4 

1.51 

1.96 

1.44 

4.96 

3.a2 

2.41 

.57 

.22 

.07 

~ 

9a.9 

19.9 

14.5 

32.9 

23.6 

5.2 

2.2 

1.1 

.s 

K04 

66.7 

16.5 

1.00 

l. 74 

1.03 

3.79 

3.41 

3.10 

.6a 

.17 

.04 

.37 

9a.5 

25.7 

la.7 

29.4 

la.o 

1.1 

4.0 

1.5 

1.3 

.4 

K05 

62.6 

16.5 

2.2a 

3.34 

l. 76 

4.a8 

3.48 

2.27 

.91 

.26 

.a a 

.30 

98.7 

20.3 

13.6 

29.9 

22.9 

.1 

7.4 

3.4 

La 

.6 

K06 

63.5 

17.3 

2.02 

2.54 

1.24 

4.54 

3.84 

2.33 

.64 

.23 

.05 

.44 

9a.7 

20.a 

14.0 

33.1 

21.4 

.8 

5.2 

3.0 

1.2 

.6 

Tonalite 

K07 K08 K09 ROl KlO SOl 

Major-element oxides in weight percent 

65.4 

16.9 

.94 

2.15 

1.18 

3.64 

3.72 

2.95 

.62 

.20 

.as 

.69 

63.6 6a.7 

17.4 16.5 

1.16 .92 

3.33 l.a9 

1.9 1.0 

4.56 4.07 

3.9 3.9 

2.44 2.06 

.a7 .50 

.3 .2 

.06 .03 

.51 .3a 

.07 .oa 

.03 .07 

60.6 63.0 66.56 

15.9 17.3 16.76 

4.26 

1.55 1.07 

3.90 3.05 

3.23 1.99 1.37 

5.58 5.48 4.40 

3.05 3.92 3.57 

2.2a 1.53 2.45 

1.02 .76 .ao 

.31 .32 .25 

.1 .07 .05 

1.14 .59 

.15 .10 

<.01 .01 

801 

6a.49 

17.41 

3.73 

1.02 

3.99 

3.70 

2.a5 

.60 

.21 

.04 

802 

65.1 

16.3 

1.19 

2.22 

1.49 

5.07 

3.48 

2.04 

.64 

.25 

.06 

.69 

.08 

.17 

9a.4 100.1 100.3 9a.a 99.2 100.47 102.04 9a.a 

23.0 

17.a 

32.2 

17.1 

1.5 

5.3 

1.4 

1.2 

.5 

Trace elements in parts per million 

1340 1500 

ao 46 

7 51 910 

21 12 

205 190 

CIPW norms in weight percent 

la.2 

14.5 

33.2 

20.a 

.a 

a.6 

1.7 

1.7 

.7 

25.5 

15.3 

32.9 

la.a 

.3 

4.4 

1.3 

.9 

.5 

17.3 

13.8 

26.5 

23.5 

2.2 

11.7 

2.3 

2.0 

.a 

19.2 

9.2 

33.7 

25.5 

a.7 

1.6 

1.5 

.a 

24.9 

14.4 

30.1 

20.2 

.a 

5.0 

2.5 

1.5 

.6 

25.4 24.4 

16.5 12.3 

30.a 30.1 

la.l 23.3 

1.5 

.6 

4.0 5.7 

2.1 1.8 

1.1 1.2 

.5 .6 

Mean cl 

N=l4 

64.6 

16.9 

4.29 

1.6 

4.61 

3.6 

2.3a 

• 75 

.25 

.06 

2.4 

.6 

1.02 

.6 

.61 

.2 

.43 

.19 

.06 

.02 

Hornblende-biotite 

granodiorite 

R02 

66.5 

15.7 

3.76 

1.53 

4.20 

3.41 

2.84 

.66 

.22 

.06 

~ 

99.4 

24.9 

17.0 

29.2 

19.4 

.2 

5.3 

2.2 

1.3 

.5 

R03 

66.9 

15.5 

l.la 

2.30 

1.61 

3.85 

3.55 

2.58 

.53 

.19 

.08 

.a5 

.10 

.01 

99.2 

EOl 

66.0 

15.2 

3.aa 

1.95 

3.90 

3.01 

3.50 

.69 

.21 

.a a 

.76 

99.2 

25.6 24.2 

15.5 21.1 

30.6 25.9 

la.2 la.o 

.3 

.3 

6.6 6.4 

1.7 2.3 

1.0 1.3 

.5 .5 

H03 

64.0 

15.3 

1.3 

3.2 

2.4 

4.2 

3.7 

2.3 

.73 

.23 

.09 

1.1 

.22 

.03 

9a.a 

20.a 

13.9 

32.1 

la.a 

.a 

9.6 

1.9 

1.4 

.6 

1--l 

00 

t\:) 

~ tzj 

p:: 

~ 

0 

td 

~ p:: 

0 

~ 

-

~ 

p:: 

~ t:; 

:::3 

U1 

td 

0 

~ 

t:; 

tzj 

~ 

N 

0 

z 

tzj 
sample 

number 

Si02 

Al203 

FeT03l 

Fe203 

FeO 

MgO 

cao 

Na2o 

K20 

Ti02 

P205 

·MnO 

H20+ 

H20-

C02 

101 2 

Total 

sa 

~b 

sr 

y 

zr 

Q 

or 

Ab 

An 

c 

Di 

Hy 

Mt 

Il 

AP 

Hornblende-biotite granodiorite 

Mean3 d 

H06 H07 H09 Hl3 N=8 

68.3 

15.1 

1.4 

1.6 

1.3 

3.6 

2.8 

3.6 

.49 

.23 

.06 

.85 

.04 

.01 

99.4 

29.2 

21.6 

24.1 

16.6 

.6 

4.4 

2.1 

.9 

.6 

68.4 

15.1 

1.5 

1.8 

1.6 

3.6 

3.4 

3.3 

.60 

.17 

.03 

.57 

.14 

.02 

100.2 

26.1 

19.6 

28.9 

16.3 

.4 

4.9 

2.2 

1.1 

.4 

69.1 

15.1 

.79 

2.0 

1.3 

3.5 

3.5 

3.6 

.45 

.22 

.12 

• 72 

.03 

.01 

100.4 

733 

147 

427 

13 

108 

25.3 

21.3 

29.7 

14.9 

.9 

5.3 

1.1 

.9 

.5 

70.4 

15.3 

1.8 

1.0 

1.0 

2.9 

3.2 

3.2 

.58 

.16 

.06 

1.2 

.22 

.02 

101.0 

32.4 

19.0 

27.2 

13.4 

1.7 

2.5 

1.7 

1.1 

.4 

67.5 

15.3 

3.61 

1.6 

3.7 

3.3 

3.1 

.59 

.20 

.07 

2.0 

.2 

.62 

.4 

.4 

.3 

.5 

.10 

.03 

.03 

K11 

69.2 

16.0 

.92 

1.50 

.81 

3.26 

4.0 

2.71 

.43 

.2 

.03 

.49 

.05 

<.01 

99.6 

1300 

85 

455 

4 

50 

27.6 

16.2 

34.2 

15.0 

1.0 

3.4 

1.3 

.8 

.5 

Biotite granodiorite 

Kl2 K13 Kl4 Kl5 Kl6 Kl7 

Major-element oxides in weight percent 

72.8 71.4 

15.1 16 .2 

.49 

1.01 

.52 

2.16 

3.5 

4.27 

.26 

<.1 

<.02 

.28 

.04 

<.01 

.63 

.65 

.5 

1.36 

4.9 

3.04 

.20 

<.1 

<.02 

.72 

.12 

.04 

70.74 72.7 71.4 

15.15 15.1 15.6 

l. 74 

.29 

2.08 

3.44 

4.13 

.28 

.09 

.03 

.57 

.73 

1.28 

.4 .41 

1.66 l. 79 

4.5 4.46 

2.89 3.26 

.16 .14 

<.1 <.05 

< .02 .03 

.60 

.02 

.34 

___£ 

74.6 

14.5 

.42 

.91 

.3 

1.49 

4.2 

3.40 

.15 

< .1 

.04 

.33 

.04 

< .01 

100.4 99.8 97.97 99.7 99.0 100.4 

Trace elements in parts per million 

1300 

94 

697 

13 

97 

1800 

51 

736 

8 

79 

1300 

57 

588 

6 

62 

CIPW norms in weight percent 

30.4 

25.2 

29.5 

10.0 

1.0 

2.3 

.7 

.5 

.2 

28.0 

18.1 

41.9 

6.2 

2.6 

1.6 

.9 

.4 

.2 

30.6 

24.9 

29.7 

9.9 

1.5 

1.5 

1.0 

.5 

.2 

31.7 

17.3 

38.5 

7.7 

1.8 

1.7 

.8 

.3 

.2 

28.9 

19.6 

38.4 

8.7 

1.6 

1.7 

.8 

.3 

.1 

1360 

68 

477 

5 

53 

33.2 

20.1 

35.5 

6.7 

1.4 

1.9 

.6 

.3 

.2 

Kl8 Kl9 

72.3 73.5 

15.2 14.8 

.97 .98 

.29 .26 

1.52 1.59 

3.98 3.84 

3.96 3.04 

.10 .13 

<.05 <.05 

.03 .02 

--=.!?.. ~ 

98.8 98.8 

30.7 

-23.8 

34.2 

7.3 

1.8 

1.3 

.6 

.2 

.1 

36.3 

18.3 

33.1 

7.7 

2.5 

1.1 

.6 

.3 

.1 

K20 

70.4 

15.9 

1.90 

.59 

2.40 

4.25 

2.45 

.36 

.09 

< .02 

.44 

98.8 

30.8 

14.7 

36.6 

11.5 

2.1 

2.2 

1.1 

.7 

.2 

K2l 

72.2 

14.5 

1.67 

.54 

1.94 

3.43 

3.76 

.29 

.09 

.04 

--=12. 

98.8 

33.4 

22.6 

29.5 

9.2 

1.5 

2.1 

1.0 

.6 

.2 

K22 

72.4 

15.3 

1.11 

.33 

l. 72 

4.73 

2.61 

.15 

< .05 

.04 

.47 

98.8 

31.1 

15.7 

40.7 

8.3 

1.7 

1.4 

0.7 

.3 

.1 

K23 

72.9 

14.6 

1.58 

.31 

1.29 

4.09 

3.62 

.17 

< .05 

.04 

.49 

99.1 

32.3 

21.7 

35.1 

6.2 

1.8 

1.6 

.9 

.3 

.1 

K24 

71.0 

15.7 

1.50 

.34 

l. 76 

4.48 

3.37 

.18 

.06 

.04 

__:..!2.. 

98.9 

28.1 

20.3 

38.5 

8.5 

1.7 

1.6 

.9 

.3 

.1 

~ 

t"" 

~ 0 

t"" 

0 

C':) 

(1 

-

~ t::1 

(1 

::q 

~ 

~ 

~ t"" 

(1 

~ ~ 

~ 1-3 

~ 

~ 

~ -(1 

r:n 

0 

l"%j 

1-3 

::q 

~ 

r:n 

0 

c:: 

1-3 

::q 

~ 

~ 

z 

t"" 

0 

ttl 

~ 

0 

l"%j 

1-3 

::q 

~ 

~ ::q 

0 

ttl 

~ ::q 

0 

t"" 

-

1-3 

::q 

....... 

00 

~ 
TABLE 9.2.-MaJ·or- and trace-element analyses and CIPW norms of selected samples from study area-Continued 

[Analysts: J.S. Walhberg, J.E. Taggart, J.D. Baker, A.J. Bartel, K. Stewart, L.L. Jackson, G.R. Mason, H.G. Neiman, E. Engleman, and D.B. Hatfield, U.S. Geological Survey, Denver, Colo.; H. Smith, G. Holmes, Z.A. Hamlin, N. Skinner, 

K. Coates, H.J. Rose, J. Lindsay, B. McCall, G. Sellers, and R. Johnson, U.S. Geological Survey, Reston, Va.; J. Amistoso, F.J. Moye, S.T. Luthy, W.B. Strowd, and D. Johnson, Washington State University, Pullman, Wash. Leaders 

(---),not determined] 

sample 

number 

Si02 

Al203 

FeT03l 

Fe203 

FeO 

MgO 

cao 

Na2o 

K20 

Ti02 

P205 

MnO 

H20+ 

H20-

C02 

LOI 2 

Total 

sa 

Rb 

Sr 

y 

zr 

Q 

or 

Ab 

An 

c 

Di 

Hy 

Mt 

Il 

Ap 

K25 K26 K27 K28 R03 

73.1 71.0 72.1 73.5 72.1 

14.8 15.2 15.1 14.4 14.6 

.43 .92 .78 .43 .64 

.82 1.08 • 76 .62 .87 

.31 .46 .40 .14 .48 

1.97 2.19 1.73 .78 2.14 

4.19 3.74 4.02 3.93 3.85 

2.74 3.35 3.30 4.56 3.20 

.14 .27 .22 .09 .26 

< .05 .07 .08 < .05 .09 

.03 .04 .03 .03 .02 

~ .44 .49 .34 

.33 

.11 

.10 

98.8 98.8 99.0 98.8 98.8 

33.7 

16.4 

36.0 

9.6 

1.5 

1.8 

.6 

.3 

.1 

31.4 

20.1 

32.2 

10.6 

1.6 

2.0 

1.4 

.5 

.2 

32.3 

19.8 

34.5 

8.2 

2.0 

1.5 

1.1 

.4 

.2 

31.6 

27.3 

33.8 

3.6 

1.7 

1.1 

.6 

.2 

.1 

32.7 

19.2 

33.2 

10.2 

1.1 

1.9 

.9 

.5 

.2 

R04 

70.4 

15.1 

.86 

1.11 

.66 

2.05 

3.96 

3.63 

.31 

.11 

.04 

.33 

.10 

.20 

98.9 

28.5 

21.8 

34.1 

9.6 

1.2 

2.6 

1.3 

.6 

.3 

R05 

71.22 

14.98 

1.88 

.38 

2.12 

3.44 

3.59 

.35 

.11 

.03 

98.10 

32.4 

21.2 

29.1 

9.8 

1.8 

1.7 

1.1 

.7 

.3 

Biotite granodiorite 

R06 WOl K29 Lol4 L024 803 

Major-element oxides in weight percent 

72.0 70.9 

15.2 15.3 

1.63 1.90 

.45 .56 

2.17 2.57 

3.75 3.58 

3.25 3.21 

.22 .26 

.11 .08 

<.02 .03 

.41 .48 

99.2 98.9 

71.4 

16.0 

.39 

.90 

.4 

2.29 

4.7 

2.99 

.18 

<.1 

<.02 

.17 

.06 

.01 

99.5 

73.84 

14.91 

1.30 

.48 

1.60 

3.79 

2.82 

.16 

.04 

.03 

98.97 

73.06 

15.18 

.85 

.31 

1.20 

3.76 

4.03 

.13 

.03 

.03 

98.58 

72.32 

14.82 

1.33 

.49 

1.83 

3.09 

3.18 

.21 

.05 

.03 

97.35 

Trace elements in parts per million 

32.8 

19.5 

32.1 

10.2 

1.9 

1.9 

1.0 

.4 

.3 

2260 

38 

834 

3 

69 

CIPW norms in weight percent 

31.8 

19.3 

30.8 

12.4 

1.5 

2.3 

1.1 

.5 

.2 

27.1 

17.8 

40.0 

10.8 

1.1 

2.1 

.6 

.3 

.2 

33.1 

16.9 

32.4 

7.8 

2.8 

1.9 

.8 

.3 

.1 

37.1 

24.2 

32.3 

5.8 

2.6 

1.2 

.5 

.3 

.1 

38.5 

19.3 

26.9 

9.0 

3.2 

1.8 

.8 

.4 

.1 

804 805 806 

72.68 69.48 71.85 

15.16 17.04 16.41 

2.08 2.07 2.50 

.63 .56 1.07 

2.54 2.95 3.19 

3.47 3.39 4.22 

2.56 4.54 2.32 

.29 .41 .38 

.05 .13 .09 

.03 .02 .03 

99.49 100.59 102.6 

36.3 

15.2 

29.5 

12.4 

2.2 

2.5 

1.2 

.6 

.1 

25.2 

16.7 

28.6 

13.7 

1.5 

2.1 

1.2 

.8 

.3 

29.2 

13.5 

35.0 

15.0 

1.3 

3.7 

1.4 

.7 

.2 

807 

71.63 

15.01 

1.63 

.63 

2.49 

3.61 

3.63 

.34 

.10 

.10 

99.17 

30.7 

21.7 

30.8 

11.8 

.9 

2.3 

1.0 

.7 

.2 

BOB 

73.4 

14.6 

.49 

.29 

.16 

1.04 

4.09 

4.16 

.08 

<.05 

<.02 

.32 

.10 

<.01 

98.7 

31.9 

25.0 

35.2 

4.9 

1.6 

.4 

.7 

.2 

.1 

H11 Hl2 

69.6 70.3 

14.4 15.3 

1.3 1.4 

1.2 1.0 

1.1 .92 

2.4 2.1 

3.2 3.6 

4.6 4.0 

.39 .39 

.13 .13 

.03 .06 

.so .88 

.21 .17 

.04 .02 

99.1 100.3 

27.2 

27.6 

27.5 

11.2 

.1 

3.3 

1.9 

.8 

.3 

29.3 

23.8 

30.7 

7.9 

2.2 

2.9 

2.0 

.2 

.9 

f--1. 

00 

~ 

1-3 

::q 

t_:rj 

::q 

~ 

0 

to 

~ ::q 

0 

t"' 

~ ::q 

> z 

t:1 

~ 

1-3 

r:n 

to 

0 

~ 

t:1 

t_:rj 

~ 

N 

0 

z 

t_:rj 
sample 

number 

Si02 

Al203 

FeT03l 

Fe203 

FeO 

MgO 

cao 

Na2o 

K20 

Ti02 

P2o5 

MnO 

H20+ 

H20-

C02 

LOI 2 

Total 

sa 

Rb 

sr 

y 

zr 

Q 

or 

Ab 

An 

c 

Di 

Hy 

Mt 

Il 

Ap 

Biotite granodiorite 

Mean3 o' 

Hl6 H22 N=36 

71.1 

14.9 

.82 

.62 

.23 

1.5 

5.0 

3.4 

.11 

.06 

.09 

.55 

.17 

.03 

98.6 

68.5 

15.6 

2.1 

1.4 

1.0 

3.2 

4.0 

3.2 

.51 

.17 

.05 

.76 

.36 

100.9 

1564 

80 

800 

12 

220 

26.0 25.1 

20.5 19.0 

43.2 33.9 

7.2 14.8 

.4 .1 

1.0 2 .6 

1.2 3.1 

.2 1.0 

.1 .4 

71.8 

15.2 

1.69 

.5 

2.0 

4.0 

3.4 

.24 

.07 

.03 

1.4 

.6 

.58 

.2 

.6 

.5 

.6 

.11 

.04 

.02 

K30 K31 

73.5 74.6 

14 .a 14 .l 

1.04 • 72 

.27 .16 

1.63 .53 

3.85 3.78 

3.24 4.41 

.12 .as < .. 05 <.05 

< .02 .02 

~ ----!2.! 

98.9 99.0 

35.3 34.9 

19.4 26.5 

33.1 32.5 

7.9 2.3 

2.1 2.3 

1.2 .8 

.6 .4 

.2 .1 

.1 .l 

Muscovite-biotite granite 

K32 K33 K34 K35 K36 K37 K38 

Major-element oxides in weight percent 

73.3 

14.9 

1.04 

.28 

1.42 

3.83 

3.50 

.10 

.0'7 

.04 

71.3 

16.1 

1.20 

.43 

2.57 

4.35 

2.23 

.18 

.07 

.02 

74.6 

14.2 

.33 

.4 7 

.19 

.54 

3.47 

4.15 

.06 

<.05 

.l 

73.5 74.7 

14.7 14.1 

.24 .34 

.66 .22 

.24 .13 

1.48 .88 

4.01 3.75 

3.44 4.06 

.11 .04 

<.05 <.05 

.04 .02 

.62 ~~_ill~ 

99.1 98.9 98.9 98.8 98.8 

74.9 

14.0 

.29 

.38 

.17 

1.21 

3.87 

3.41 

.l 

<.05 

.04 

.39 

98.8 

Trace elements in parts per million 

34.6 

21.0 

32.9 

6.7 

2.4 

1.3 

.6 

.2 

.2 

CIPW norms in weight percent 

31.8 

13.4 

37.4 

12.5 

2.1 

1.6 

.7 

.3 

.2 

37.6 

25.0 

29.9 

2.4 

3.2 

1.2 

.5 

.1 

.l 

33.7 

20.6 

34.5 

7.1 

1.8 

1.5 

.4 

.2 

.1 

36.0 

24.4 

32.3 

4.1 

2.1 

.4 

.5 

.l 

.1 

37.1 

20.5 

33.3 

5.8 

1.9 

.8 

.4 

.2 

.1 

73.6 

14.7 

.23 

.65 

.25 

1.64 

3.76 

3.50 

.13 

< .05 

.02 

.34 

98.8 

34.7 

21.0 

32.3 

7.9 

1.9 

1.5 

.3 

.3 

.1 

K39 

74.3 

14.4 

.33 

.42 

.24 

1. 79 

3. 7l 

2.87 

.09 

< .05 

< .02 

:l.§. 

98.5 

38.2 

17.3 

32.0 

8.7 

2.1 

1.0 

.5 

.2 

.1 

K40 

72.9 

14.7 

1.16 

.31 

1.62 

3.89 

3.23 

.11 

< .05 

.04 

1.02 

99.0 

34.5 

19.5 

33.6 

7.9 

2.0 

1.4 

.7 

.2 

.l 

W02 

74.5 

14.6 

.40 

.54 

.3 

1.66 

3.7 

3.62 

.l 

< .l 

< .02 

.30 

.17 

< .01 

99.9 

1400 

100 

886 

20 

228 

35.3 

21.5 

31.5 

7.6 

1.8 

1.3 

.6 

.2 

.2 

B09 

73.7 

15.0 

.58 

.67 

.30 

1.59 

3.63 

2.74 

.14 

< .05 

< .02 

• 74 

.24 

.04 

99.4 

38.7 

16.5 

31.2 

7.7 

3.3 

1.3 

.9 

.3 

.1 

BlO 

75.03 

14.37 

.90 

.55 

1. 39 

3.28 

4.17 

.12 

.02 

.03 

99.86 

36.1 

24.7 

27.8 

6.8 

2.0 

1.8 

.5 

.2 

.l 

Mean3 

N=l4 

73.9 

14.6 

.95 

.3 

1.43 

3.8 

3.47 

.10 

.04 

.03 

o' 

1.0 

.5 

.20 

.1 

.53 

.2 

.60 

.04 

.02 

.02 

:.0 

t""' 

H 

~ 

p:: 

0 

t""' 

0 

CJ 

H 

C":l 

> z 

t:1 

C":l 

p:: 

t:zj 

~ 

~ t""' 

C":l 

p:: 

> 

~ 

f; 

~ 

t:zj 

~ 

w 

~ H 

C":l 

w 

0 

~ 

~ 

p:: 

t:zj 

w 

0 

~ 

~ 

p:: 

t:zj 

~ t""' 

0 

to 

t:zj 

0 

~ 

~ 

p:: 

t:zj 

p:: 

~ 

0 

to 

~ p:: 

0 

t""' 

H 

~ 

p:: 

1--l 

00 

Ol 
TABLE 9.2.-Major- and trace-element analyses and CIPW norms of selected samples from study area-Continued 

[Analysts: J.S. Walhberg, J.E. Taggart, J.D. Baker, A.J. Bartel, K. Stewart, L.L. Jackson, G.R. Mason, H.G. Neiman, E. Engleman, and D.B. Hatfield, U.S. Geological Survey, Denver, Colo.; H. Smith, G. Holmes, Z.A. Hamlin, N. Skinner, 

K. Coates, H.J. Rose, J. Lindsay, B. McCall, G. Sellers, and R. Johnson, U.S. Geological Survey, Reston, Va.; J. Amistoso, F.J. Moye, S.T. Luthy, W.B. Strowd, and D. Johnson, Washington State University, Pullman, Wash. Leaders 

(···), not determined] 

sample 

Number 

Si02 

Al203 1 FeT03 

Fe203 

FeO 

MgO 

cao 

Na2o 

K20 

Ti02 

P205 

MnO 

H20+ 

H20-

C02 

101 2 

Total 

sa 

Rb 

sr 

y 

zr 

Q 

or 

Ab 

An 

c 

Di 

Hy 

Mt 

Il 

Ap 

K41 

74.9 

l3 .8 

.23 

.65 

.12 

.65 

4.09 

4.44 

.08 

.03 

.04 

.37 

.13 

99.5 

K42 

76.8 

13.3 

.22 

.23 

.1 

.90 

3.6 

4.31 

<.02 

<·1 

<.02 

.25 

.06 

.01 

99.8 

1420 

92 

440 

5 

21 

32.5 37.6 

26.5 25.6 

34.9 30.6 

3.1 3.8 

1.2 1.3 

1.3 .5 

.3 .3 

.2 .0 

.1 .2 

K43 

73.9 

14.8 

.15 

.68 

.2 

1.12 

4.4 

3.82 

.07 

<.1 

.04 

.30 

.07 

<.01 

99.6 

1520 

70 

393 

3 

46 

K44 

75.1 

l3 .8 

.87 

<.10 

.57 

4.77 

3.71 

<.02 

<.05 

<.02 

.31 

99.1 

K45 

75.6 

14.0 

.32 

.36 

.2 

.62 

4.5 

3.98 

.05 

<·1 

<.02 

.24 

.09 

.02 

100.0 

880 

84 

252 

7 

15 

31.0 32.0 32.8 

22.7 22.2 23.6 

37.5 40.8 38.2 

4.9 2.5 2.4 

1.6 1.0 1.4 

1.6 .8 .9 

.2 .5 .5 

.1 .o .1 

.2 .1 .2 

K46 

74.1 

14.7 

.31 

.43 

.07 

1.12 

4.78 

4.10 

.09 

.03 

.01 

.42 

.13 

100.3 

K47 

76.7 

13.8 

.28 

.27 

.1 

• 72 

4.9 

3.39 

<.02 

<•1 

<.02 

.06 

.04 

<.01 

100.3 

150 

79 

75 

5 

45 

28.1 33.6 

24.3 20.0 

40.6 41.3 

5.4 2.9 

.4 1.0 

.6 .5 

.5 .4 

.2 .o 

.1 .2 

1eucocratic granite 

K48 K49 K50 K51 1034 1044 

Major-element oxides in weight percent 

71.9 

16.0 

• 71 

.20 

.2 

.84 

5.2 

4.10 

.09 

.1 

<-02 

.28 

.14 

<.01 

99.8 

77.5 

12.0 

.40 

.23 

<.10 

.68 

3.41 

3.96 

.04 

<.05 

<.02 

~ 

98.4 

75.3 

14.0 

.4 7 

.38 

.2 

.77 

4.1 

4.74 

.09 

<.1 

.03 

.17 

.04 

<.01 

76.7 

13.9 

.03 

.18 

.1 

.58 

4.0 

4.65 

<.02 

<·1 

<.02 

.30 

.07 

.01 

100.3 100.5 

75.51 76.06 

14.59 15.07 

.12 .12 

.30 .35 

.64 1.03 

3.38 3.58 

4.10 3.95 

.05 .04 

.01 .02 

.01 .00 

98.71 100.22 

Trace elements in parts per million 

2080 

83 

700 

10 

52 

24.3 

24.4 

44.3 

3.5 

1.7 

.5 

.5 

.2 

.2 

740 

71 

221 

6 

8 

680 

126 

168 

16 

1 

CIPW norms in weight percent 

41.7 

23.8 

29.3 

3.1 

1.0 

.3 

.6 

.1 

.1 

31.5 

28.0 

34.6 

3.1 

1.0 

.7 

.7 

.2 

.2 

34.3 

27.4 

33.8 

2.2 

1.5 

.6 

.o 

.o 

.2 

38.9 

24.5 

29.0 

3.2 

3.5 

.8 

.1 

.1 

.o 

37.4 

23.3 

30.2 

5.0 

3.1 

.9 

.1 

.1 

.o 

Bll 

73.30 

14.72 

.02 

.29 

1.35 

3.76 

4.75 

.05 

.05 

.01 

98.30 

30.6 

28.6 

32.4 

6.5 

1.1 

.7 

.o 

.o 

.1 

812 

74.77 

14.53 

.59 

.45 

1.08 

2.51 

4.53 

.03 

.02 

.03 

98.54 

40.3 

27.2 

21.6 

5.3 

3.6 

1.5 

.4 

.1 

.o 

Bl3 

73.45 

14.47 

.87 

.27 

1.32 

3.68 

4.34 

.11 

.02 

.03 

98.56 

32.5 

26.0 

31.6 

6.5 

1.4 

1.1 

.5 

.2 

.o 

814 

74.93 

14.82 

.34 

.33 

1.10 

4.31 

3.97 

.06 

.01 

.01 

99.88 

31.8 

23.5 

36.5 

5.4 

1.5 

1.0 

.2 

.1 

.o 

815 

71.9 

15.5 

.70 

.23 

.16 

.78 

5.08 

3.49 

.11 

<.05 

<-02 

.44 

.12 

.03 

98.5 

27.8 

21.0 

43.9 

3.6 

2.1 

.4 

.5 

.2 

.1 

Mean3 oN=l8 

74.9 

14.3 

.59 

.2 

.88 

4.1 

4.13 

.06 

.04 

.02 

1.6 

.9 

.32 

.1 

.25 

.7 

.40 

.03 

.02 

.01 

........ 

00 

0":1 

.-3 

~ 

tr:1 

~ ~ 

0 

to 

~ ~ 

0 

~ 

-.-3 

~ 

~ t:;j 

-.-3 

r:n 

to 

0 

~ 

t:;j 

t_:tj 

~ 

N 

0 

z 

tr:1 
ThomEson creek stock White Cloud stock Little Boulder Rook's creek McCoy Mine Croesus 

sample creek stock stock stock stock 

numl:;>er Hl9 H215 817 818 ROB Hl7 616 Hl56 H05 Hl8 H04 H02 

~ 

Major-element oxides in weight percent ~ ~ 

t-3 

~ 

0 

SiOz 72.9 77.8 73.28 72.45 70.5 71.2 71.49 70.8 68.0 72.1 67.3 57.8 ~ 

0 

Alz03 14.3 12.8 15,45 15.29 15.3 15.7 14.90 15.3 16.5 14.7 14.7 18.0 8 

FeT03l -- -- 1. 73 1.04 -- -- 1. 74 -- -- -- -- -- 0 

Fe2o3 .53 .27 -- -- .93 .65 -- .56 • 72 .53 .56 2.1 > 

FeO .76 .84 -- -- .86 1.1 -- .60 2.6 1.6 2.7 4.6 z 

MgO .33 .28 .73 .69 .58 .46 .68 .38 .95 .66 2,0 2.8 t:::) 

cao 1.1 .94 2.29 1. 74 2.07 2.2 1. 77 2.2 3.1 2.0 3.1 7.7 0 

~ Na2o 3.0 2.4 2.95 2.85 4.15 4.1 4.16 3.1 3.4 2.9 3.1 3.8 t:zj 

KzO 4.5 4.8 3,07 4.27 3.59 3.2 3.60 6.5 3.1 4.2 3.6 1.3 s:: 

~ 

TiOz .24 .15 .30 .28 .26 .35 .27 .35 .56 .45 .65 1.2 

P2os .07 .08 .10 .10 .11 .11 .09 .11 .18 .14 .09 .52 ~ ~ 

MnO .03 .02 .03 .03 .03 .03 .04 .02 .06 .03 .02 .OS 0 

HzO+ .84 .76 -- -- .30 .43 -- .43 .65 ,62 .90 .95 ~ 

HzO- .12 .28 -- -- .07 .08 -- .28 .65 .19 .16 .17 > 

~ co2 .03 .01 -- -- <.01 .02 -- .10 .30 .as .10 .10 

LOI -- -- -- -- -- -- -- -- -- -- -- -- ~ t-3 

t:zj 

Total 98.8 101.4 99.93 99,74 98.75 99.6 98.74 100.7 100 .a 100.2 99.0 101.1 ~ ~ 

r:n 

t-3 

~ 

Trace elements in parts per million 0 

r:n 

0 

~ 

sa -- 1034 -.. -- -- 776 -- 939 -- 1092 702 720 t-3 

Nb -- 14 -- -- -- 18 -- 17 -- 21 25 11 ~ 

t:zj 

Rb -- 117 -- -- -- 80 -- 200 -- 171 168 35 r:n Sr -- 633 -- -- -- 800 -- 735 -- 557 550 938 0 

":( -- 15 -- -- -- 11 -- 17 -- 14 30 24 c:: 

t-3 zr -- 148 -- -- -- 190 -- 198 -- 171 314 214 ~ t:zj 

~ 

z 

CIPW norms in weight percent ~ 

0 

ttl 

29.8 29.0 23.2 33.7 25.1 10.7 t:zj 

Q 36.1 42.6 38.9 34,8 27.9 27.5 0 or 27.2 28.3 17.5 25.3 21.6 19.1 21.6 38.4 18.5 25.0 21.7 7.7 ~ 

Ab 26.0 20.2 25.0 24.2 35.7 35.0 35.7 26.3 29.0 24.7 26.8 32.2 t-3 

An 5.1 4.1 10.7 8,0 9.7 10 • .3 8.3 8.6 14.3 9.1 15.1 28.2 ~ 

c 2.7 2.1 3.5 3.1 1.1 1.8 1.2 -- 2.4 2.1 .3 -- t:zj 

Di 1.2 ~ 

-- -- -- -- -- -- -- -- -- -- 5.4 

Hy 1.5 1.8 2.5 2.7 1.9 2.1 2.5 .5 5.8 3.5 8.6 9.2 ~ ~ Mt .a .4 1.0 1.2 1.4 1.0 1.0 .a 1.1 .a .a 3.0 0 

Il .5 .3 .6 .5 .5 .7 .5 .7 1.1 .9 1.3 2.3 

Ap .2 .2 .2 .2 .3 .3 .2 .3 .4 .3 .2 1.2 ~ ~ 

~ 

0 

!Total iron reported as Fe2o3 • Fe2o3/Feo ratio of .75 assigned for CIPW norm calculation. t'"4 ~ 

2Loss on ignition at 900° c. t-3 

3Assigned values for calculation of mean and standard deviation are as follows: <.1=.05; <.05-.03; <.02=.01. ~ 

4oata from Luthy (1981). 

5sample has undergone silicification. f.-A 

6sample has undergone K-feldspar alteration. 00 

~ 
188 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

leucocratic granite along the transect were taken from 

dikes. A sharp increase in Si02 content and a corresponding decrease in total Fe and CaO contents are evident over 

a distance of 10 km from the western edge of the transect, 

.821 

where the east margin of the tonalite gives way to the 

core of biotite granodiorite and 2-mica granite. Variation 

in concentrations of Si02, CaO, and total Fe is minimal 

eastward across the next 90 km of the batholith; CaO and 

•8o4 o 40 

Boise® 

•8o3 •83o 

•• ®Landmark 826 810 

•8o7 

•wo1 •Ro2 

•Ro3 

•K32 •K52 •E01 

•Ro7 •K30 

•K19 •K31 

•K40 •K18 eK28 

•K23 

®Idaho City 

oV~ ®Atlanta 

L---~----~----~--~ 

Thompson Creek stock 

H19 H21 I 817 

~~~· 

KILOMETERS 

H18t:-._Ho5 

H13 \) 

®Ketchum 

FIGURE 9.7.-Locations of samples collected for this study, and location of transect A-A'. 
9. LITHOLOGIC AND CHEMICAL CHARACTERISTICS OF THE SOUTHERN LOBE OF THE IDAHO BATHOLITH 189 

total Fe contents then increase and Si02 content 

decreases. At the extreme eastern edge Si02 content 

again rises and concentrations of total Fe and CaO 

decrease. Relations along the eastern margin are complex, 

owing to the discontinuous nature of the tonalitic rocks 

in this area. The K20 and Na20 contents display very 

little systematic variation along the transect, though there 

is an increase in K20 in the nontonalitic phases of the 

batholith. The minimal chemical variation across the central 90-km part of the batholith suggests that little or no 

differentiation took place as these rocks were emplaced. 

A similar conclusion was reached by Hyndman (1983) for 

the Bitterroot lobe of the batholith. Fleck and Criss (1985) 

have shown that assimilation-fractional crystallization 

models permit only insignificant amounts of crystal fractionation during anatexis and mixing for the majority of 

the plutons in the Bitterroot lobe. 

The break between the western tonalite and more felsic 

rocks to the east corresponds to the quartz diorite line 

of Moore (1959). The initial 87Srf86Sr 0. 704-0.706 line is 

not coincident with the quartz diorite line, but lies in the 

tonalitic rocks to the west (Armstrong and others, 1977; 

Fleck and Criss, 1985). 

An AFM plot of the Atlanta lobe rocks (fig. 9.9) shows 

a calc-alkaline alkali-enrichment trend. Harker diagrams 

with Al20s, total Fe, CaO, and K20 concentrations 

plotted against Si02 concentrations (fig. 9.10) also show 

typical trends for a granitic suite of rocks. Both CaO and 

total Fe contents have excellent inverse correlations with 

Si02 content, a relation also seen in figure 9.8. Considerable scatter, however, occurs in the Al203 and K20 

concentration plots. Although rock suites that show 

smooth variation on such diagrams are interpreted by 

some authors as being comagmatic (for example, Hyndman, 1972), the distribution of data in figure 9.10 is probably not a good indicator of, or evidence against, a 

comagmatic suite in the Idaho batholith because of the 

size of the study area and the scatter seen for some of 

the major-element oxides. 

A K20-CaO-Na20 plot (fig. 9.11) shows the range and 

average value of rocks in the study area compared with 

the range and average values of some other granitoids 

in the Western United States. The Tuolumne Intrusive 

Suite (Bateman and Chappell, 1979) of the Sierra Nevada 

batholith and the quartz monzonite and granodiorite 

plutons in the southern California batholith (Baird and 

others, 1974) have virtually identical average K20-CaON a20 values. Both are more calcic than the average value 

of Idaho batholith rocks in this study, though there is conA It siderable overlap among individual samples. Hyndman 

.---------------------. (1983) has shown that rocks of the Idaho batholith are 

5 

ui 

~ 4 

>< 0 

0 10 20 30 40 50 60 70 80 90 100 110 

DISTANCE, IN KILOMETERS 

FIGURE 9.8.-Chemical variation of selected oxides from samples across 

the central part of the Atlanta lobe of the Idaho batholith. Fe203 is 

total Fe as Fe203• Locations of sample sites and transect A-A' are 

given in figure 9.7. 

similar chemically to Tilling's (1973) "sodic series" in the 

Boulder batholith of Montana; the "main series" of the 

Boulder batholith is appreciably richer in K20. Data from 

this study supports Hyndman's correlation, as seen in 

figure 9 .11. 

Rocks of the Atlanta lobe of the Idaho batholith are predominantly peraluminous according to the definition of 

Shand (1949), with molecular Al20s/(CaO + Na20 + K20) 

ratios (A/CNK) greater than 1.0 (fig. 9.12). Some of the 

tonalite and hornblende-biotite granodiorite is metaluminous by this classification. 

SOURCE REGION 

Chappell and White (197 4) recognized two major orogenic granite types (I and S) in the Lachlan fold belt of 

Australia, which they separated on the basis of mineralogy 

and chemistry. The 1-type granite was thought to have 

formed from partial melting of gabbroic to granitic igneous rocks and the S-type granite, from partial melting 

of sedimentary rocks. A late-formed group of felsic 

anorogenic granitoids (A-type) has since been recognized 

(Collins and others, 1982; White andChappell, 1983). Application of this classification system in the Idaho batholith 

meets with only limited success. The tonalite has I -type 

characteristics, such as the presence of biotite and horn-
190 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

blende, high CaO content, and low A/CNK ratios; the 

muscovite-biotite granite has many S-type features, such 

as high A/CNK ratios, the presence of muscovite, and low 

CaO content. However, the biotite granodiorite, which is 

the most extensive unit in the southern lobe, has both Iand S-type characteristics. The A/CNK ratios (fig. 9.12) 

of the biotite granodiorite best illustrate the transitional 

nature of this unit. The ratios range from 1.0 to 1.25; 

Chappelland White (1974) indicate that S-type granites 

have ratios of greater than 1.1. Hyndman (1984) has found 

that the central part of the northern lobe of the Idaho 

batholith is neither clearly 1- nor S-type, and he concluded 

that such a distinction may not be appropriate for these 

rocks. 

The I -type granitoids of Australia are high in sodium, 

whereas the S-type granitoids are low in sodium. Chappell and White (197 4) proposed that the low sodium content in the S-type rocks reflected a sedimentary source 

region that had been through a chemical weathering cycle. 

During this cycle potassium was incorporated into clays, 

A 

F 

and sodium was removed into sea water, resulting in 

sodium-depleted sediment. The S-type (2-mica) granite of 

the Idaho batholith does not show a sodium depletion, 

however, and is indistinguishable from the other batholith 

rock types on a Na20 versus K20 plot (fig. 9.13). The 

peraluminous nature of the Idaho two-mica granite suggests that these rocks may have had a sedimentary source 

region, hence a Na20 versus K20 plot is not adequate for 

distinguishing granitoid source regions in the Idaho 

batholith. A similar conclusion has been reached by Armstrong and Boudette (oral commun., 1985), who found no 

overlap between the western European 2-mica granitoids 

and the Australian S-type granitoids on a Na20 versus 

K20 diagram, even though these European rocks are 

thought to have been derived from a sedimentary source 

region. 

Because large bodies of biotite granodiorite in the Atlanta lobe cannot be categorized as either I- or S-type granite, 

this classification scheme is of limited use in the Idaho 

batholith. The usefulness of the I and S classification 

M 

FIGURE 9.9.-AFM plot of nonsatellitic granitic rocks of the Atlanta lobe of the Idaho batholith showing 

alkali-enrichment trend. Solid line is an iron-enrichment (tholeiitic) trend and shaded area is a composite 

of several calc-alkaline trends. A = N~O + K20; F = total Fe as FeO; M = MgO. 
9. LITHOLOGIC AND CHEMICAL CHARACTERISTICS OF THE SOUTHERN LOBE OF THE IDAHO BATHOLITH 191 

system has been questioned by Czamanske and others 

(1981), and Barker (1981) has pointed out that although 

the S-type designation of Chappell and White (197 4) is 

often applicable, the 1-type category has not met with 

similar widespread application. 

The tonalitic rocks along the westernmost margin of the 

batholith are low in radiogenic strontium (initial Sr-isotope 

ratios of less than 0. 704, fig. 9.2). Armstrong and others 

18 -

16 -n 

14 -

12 -

0 

5-

0 l3 3 -

60 

0 

o!SD 

0 0 

0 

• 

I 

0 0 

0 0 

• 

• • 

~ 0 0 

11:11 0 0 

• • 

I 

• r:1 0 

65 

I I I 

0 

0 

I I 

75 80 

(1977) have suggested that these tonalites formed from 

mantle-derived magmas that assimilated young crustal 

rocks not greatly enriched in radiogenic Sr. They postulate that the remainder of the batholith, with initial ratios 

greater than 0.706, was formed by similar magma that 

encountered and assimilated large quantities of Precambrian (largely pre-Belt Supergroup) rocks rich in radiogenic Sr. Hyndman (1983) has pointed out that the low 

I 

0 

0 

l 

60 

0 

0 

o• 

0 

0 0 

0 

ofR 

0 

0 

I 

0 

0 ••• 

0 

• 

0 

0 

65 

0 

0 

• 

• 0 

I I 

• 0 

0 

70 75 

- 6 

- 4 

- 2 

- 0 

- 5 

- 4 

-3 

-2 

80 

0 

C\1 

~ 

EXPLANATION 

o Tonalite 

• Hornblende-biotite granodiorite 

• Muscovite-biotite granite 

o Biotite granodiorite 

~ Leucocratic granite 

FIGURE 9.10.-Harker diagrams showing Si02 plotted against Al20 3, Fe20 3 (total Fe as Fe20 3), CaO, and K20 (all in weight percent) for 

nonsatellitic granitic rocks in the Atlanta lobe of the Idaho batholith. 
192 

• 

THE IDAHO BATHOLITH AND ITS BORDER ZONE 

FIGURE 9.11.-K20-Ca0-Na20 plot for nonsatellitic granitic rocks of the Atlanta lobe (shown as open circles 

with the average value given by the solid circle) of the Idaho batholith. Also shown are average values 

for the Tuolumne Intrusive Suite of the Sierra Nevada batholith (Bateman and Chappell, 1979) and the 

quartz monzonite and granodiorite plutons in the southern California batholith (Baird and others, 197 4), 

which are coincident and indicated by the solid triangle. The "Main Series" and "Sodic Series" fields 

of the Boulder batholith of Montana (Tilling, 1973) are also shown for comparison. 

Kt 

I I~~~~ 

8++., Kgd 

II I I dJ tttttJ ITIIJ 0 D 

Metaluminous +Peraluminous Mm Kg 

1.0 1.1 1.2 

FIGURE 9.12.-Ratios of molecular Al20 3/(Na20+K20+Ca0) for 

tonalite (Kt), biotite granodiorite (Kgd), and 2-mica granite (Kg) in 

the Atlanta lobe of the Idaho batholith. 

initial87Srf86Sr rocks along the west side of the batholith 

could have formed by melting of the volcanic rocks of the 

Seven Devils Group, subducted oceanic crust, or upper 

mantle. He considers the main body of the batholith to 

be from crustal-derived magma and not the product of extensive assimilation by mantle-derived magmas. This idea 

is supported by Fleck and Criss (1985), whose oxygen- and 

strontium-isotopic study of the Bitterroot lobe documents 

plutons that are mixtures of Precambrian wall rock and 

melts dominated by continental lower crust (initial 

87Srf86Sr greater than 0. 708) rather than mantle components. The immense volume of the batholith and the 

peraluminous nature of the central and eastern part suggest, at the very least, a high crustal component. Initial 

Sr ratios are not high enough, however, to permit a simple anatectic, intracrustal genesis (Armstrong and others, 
9. LITHOLOGIC AND CHEMICAL CHARACTERISTICS OF THE SOUTHERN LOBE OF THE IDAHO BATHOLITH 193 

1-

z 

w 

(.) 

a: 

~ 

II 

(!) 

w 

~ 

Cf 

C\1 

«S 

z 

• 

4 

2 

2 4 

K20, WEIGHT PERCENT 

FIGURE 9.13.-Na20 versus K20 plot of nonsatellitic granitic rocks of 

the Atlanta lobe of the Idaho batholith. Muscovite-biotite (2-mica) 

granite shown as squares, all other units shown as circles. I and S 

fields of White and Chappell (1983) shown for comparison. 

1977). Whether the crustal component was from metasedimentary rocks such as those of the Belt Supergroup and 

(or) the Yellowjacket Formation, or from pre-Belt Supergroup basement rocks is unknown. However, owing to the 

low metamorphic grade of the bulk of the Belt Supergroup 

and Y ellowjacket lithologies, the pre-Belt Supergroup 

basement is a more likely source. The sodic nature of the 

batholith indicates that it originated in rocks enriched in 

sodium relative to calcium and potassium. 

COMPARISON OF THE ATLANTA LOBE WITH 

THE BITTERROOT LOBE 

Rocks of the Atlanta lobe of the Idaho batholith are 

more massive and contain less schlieren than those of the 

Bitterroot lobe. Particularly lacking in the Atlanta lobe 

is the penetrative foliation and lineation that is present 

in a 0.5-km-wide mylonitic shear zone on the east flank 

of the Bitterroot lobe (Chase, 1977; Hyndman, 1979; 

Bittner, chapter 5). K-Ar ages are younger in the Bitterroot lobe than in the Atlanta lobe, and some authors (for 

example, Toth, 1985) consider the main mass of the Bitterroot lobe to be Paleocene in age. The Atlanta lobe also 

differs in that it lacks metamorphosed mafic dikes, which 

are found in the Bitterroot lobe and which were apparently emplaced into fractures in the granite just before 

crystallization of the last pegmatite fluids. Hyndman 

(1984) has interpreted these dikes as the later stages of 

rising basalts that supplied heat for melting of the crust, 

producing the batholith magmas. 

Biotite granodiorite and muscovite-biotite granite are 

the most abundant rock types in both lobes of the batholith 

and together can be considered to be the main phase of 

the batholith. A comparison of averaged major-element 

oxides from samples of the main phase of each lobe (table 

9.3) shows remarkable similarities. The Atlanta lobe is 

slightly higher in Si02 content and lower in Al203 and 

K20 contents, but otherwise the average concentration 

values are virtually identical. This relation suggests that 

both lobes of the batholith formed under similar conditions from a source material of nearly the same composition. The main phase of the Bitterroot lobe shows minimal 

chemical variation, as does the main phase in the Atlanta 

lobe. Both are relatively sodic compared to the main series 

of the Boulder batholith. Although K20 varies more in 

the Atlanta lobe, there is no eastward increase in either 

lobe. 

COMPARISON OF THE IDAHO BATHOLITH 

WITH THE SIERRA NEVADA BATHOLITH 

As already noted, one characteristic of the Idaho batholith is its lack of numerous separate granitic plutons in 

sharp contact with one another. This distinction is in contrast to the Sierra Nevada batholith (Bateman, 1983), as 

well as other large batholiths (Pitcher, 1978), where a wide 

variety of plutons are present, many of which are texturally and compositionally zoned in a concentric fashion 

(Bateman and Chappell, 1979). The unusual homogeneity 

of the Idaho batholith compared to the Sierra Nevada 

batholith suggests generation by partial melting of broadly similar rocks, coupled with only minimal amounts of 

differentiation. Hyndman (1981) proposed that the Idaho 

batholith was formed by relatively ''wet'' magma that was 

unable to rise far above its source region in the crust 

before reaching the solidus. If so, the differences between 

the Idaho and Sierra Nevada batholiths may in part be 

due to a higher water content and thus deeper emplacement level of the Idaho batholith. The deep emplacement 

level and resultant slow cooling would produce diffuse contacts such as those present in most of the Idaho batholith. 

Hamilton (1983) pointed out the possible water content 

differences between the two batholiths and further suggested that a batholith similar to the one in Idaho may 

underlie the Sierra Nevada batholith. He also suggested 

that a Sierra-type batholith once existed above the Idaho 

batholith, a proposal that seems unlikely, as metasedimentary roof pendants are present across the top of the 

Atlanta lobe. An exception to the "wet" magma proposal 

for the Idaho batholith is found with the Croesus stock, 

which lies southeast of the main batholith 20 km south 

of Ketchum (fig. 9.3). In this pluton pyroxene and biotite 

are major mafic constituents, whereas amphibole is a 

secondary mineral. Thus, the magma that formed the 
194 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

TABLE 9.3.-Major-elernent oxide averages of the Atlanta and Bitterroot 

lobes of the Idaho batholith and of the Tuolumne Intrusive Suite of the 

Sierra Nevada batholith 

[Idaho batholith averages are only for the main-phase (biotite granodiorite and 2-mica granite) 

rocks; values given are in weight percent. Atlanta lobe values from this study; Bitterroot 

lobe values from Coxe and Toth (1983) and Hyndman (1984); Tuolumne Intrusive Suite 

values from Bateman and Chappell (1979). n, number of samples. Total Fe as Fe2

0 3] 

Atlanta Bitterroot Tuolumne 

lobe lobe Intrusive 

n=50 n=81 Suite 

n=22 

Sio2------ 72.4 71.9 68.2 

Al2o3----- 15.0 15.4 15.7 

K20------- 3.4 3.8 3.6 

Na2o------ 3.9 4.1 4.0 

Total Fe-- 1.48 1.37 3.0 

MgO------- .4 .4 1 .1 

Cao------- 1.84 1.80 3.2 

Tio2------ .20 .18 .5 

quartz monzodiorite of the Croesus stock was relatively 

dry. 

The principle difference between the Idaho and Sierra 

Nevada batholiths is composition. The Idaho batholith has 

a lower mafic-mineral content and is dominated by granite 

and granodiorite. Granodiorite and tonalite are the dominant rock types in the Sierra Nevada batholith. The Idaho 

batholith is predominantly muscovite and biotite bearing 

with subordinate hornblende, whereas biotite and hornblende predominate in the Sierra Nevada rocks. Chemically, the Idaho batholith is higher in Si02 content and lower 

in CaO and total Fe contents than many of the Sierra 

Nevada rocks (table 9.3). The K20 content does not increase at constant Si02 content eastward across the 

Idaho batholith as it does in the Sierra Nevada batholith. 

The initial 87Srf86Sr 0.704 line extends down the 

western side of both batholiths (Kistler, 1974; Armstrong 

and others, 1977). However, the 0.706 line extends 

through the central part of the Sierra Nevada, whereas 

it is nearly coincident with the 0. 704 line in Idaho. This 

represents a fundamental difference in the nature of the 

lithosphere below each batholith. Much of the Sierra 

Nevada batholith is underlain by oceanic (island arc) crust, 

but the Idaho batholith is largely underlain by continental crust. These crustal differences greatly contributed 

to the contrasts between the two granitoid bodies. 

SUMMARY 

The Atlanta lobe of the Idaho batholith consists primarily of biotite granodiorite and muscovite-biotite granite, 

with subordinate amounts of tonalite, hornblende-biotite 

granodiorite, K-feldspar megacryst granodiorite, and 

leucocratic granite. The oldest rocks are tonalite, which 

have emplacement ages of approximately 95 to 85 Ma. 

The youngest rocks are fine-grained leucocratic granite, 

emplaced at about 75 to 70 Ma. Major-element chemistry 

documents the remarkable homogeneity of the batholith, 

with minimal east-west variation across 90 km of the central part of the Atlanta lobe. The bulk of the batholithic 

rocks are peraluminous, and they show a calc-alkaline 

alkali-enrichment trend. 

Available Sr-isotope data show that tonalite along the 

extreme western margin of the Idaho batholith formed 

from magma depleted in radiogenic strontium. These 

rocks were probably generated by partial melting in or 

above a subducting slab of oceanic lithosphere. Tonalite, 

granodiorite, and granite to the east have high initial Sr 

ratios and probably formed by partial melting of pre-Belt 

Supergroup basement rocks in the lower crust. The 

mantle component in these rocks is minor. 

Chemical and mineralogical comparisons show the 

Bitterroot and Atlanta lobes of the Idaho batholith to be 

similar in composition. Significant foliation and lineation, 

as well as metamorphosed basaltic dikes, are absent in 

the Atlanta lobe but are present in the Bitterroot lobe. 

The Idaho batholith differs from the Sierra Nevada batholith in two significant ways. First, the Idaho batholith has 

a lower mafic-mineral content than the Sierra Nevada 

batholith, and second, it lacks the numerous zoned plutons 

in sharp contact with one another that is characteristic 

of the Sierra Nevada batholith. Biotite and muscovite 

predominate in the Idaho batholith, whereas hornblende 

and biotite are characteristic of the Sierra Nevada rocks. 

Chemically, the Idaho batholith tends to be higher in Si02 

content and lower in CaO and total Fe contents. Higher 

water content in the magma during emplacement of the 

Idaho batholith may have inhibited the rise of magma 

there relative to the Sierra Nevada batholith. The nature 

of the lithosphere, as reflected in initial87Srf86Sr ratios, 

is different below the two batholiths. The Sierra Nevada 

batholith is largely underlain by oceanic material, and the 

Idaho batholith is underlain almost entirely by continental crust, thus resulting in numerous contrasts between 

the two batholiths. 

REFERENCES CITED 

Anderson, A.L., 194 7, Geology and ore deposits of Boise Basin, Idaho: 

U.S. Geological Survey Bulletin 944-C, 319 p. __ 1952, Multiple emplacement of the Idaho batholith: Journal of 

Geology, v. 60, no. 3, p. 255-265. 

Anderson, A.L., Kiilsgaard, T.H., and Fryklund, V.C., Jr., 1950, Detailed 

geology of certain areas in the Mineral Hill and Warm Springs 

mining districts, Blaine County, Idaho: Idaho Bureau of Mines and 

Geology Pamphlet 90, 73 p. 

Armstrong, R.L., 1974, Geochronometry of the Eocene volcanic-plutonic 

episode in Idaho: Northwest Geology, v. 3, p. 1-15. __ 1975a, Precambrian (1500 m.y. old) rocks of central Idaho-the 

Salmon River arch and its role in cordilleran sedimentation and tectonics: American Journal of Science, v. 275-A, p. 437-467. 
9. LITHOLOGIC AND CHEMICAL CHARACTERISTICS OF THE SOUTHERN LOBE OF THE IDAHO BATHOLITH 195 

__ 1975b, The geochronometry of Idaho: Isochron West, no. 14, 50 p. 

Armstrong, R.L., Taubeneck, W.H., and Hales, P.O., 1977, Rb-Sr and 

K-Ar geochronometry of Mesozoic granitic rocks and their Sr isotopic 

composition, Oregon, Washington, and Idaho: Geological Society 

of America Bulletin, v. 88, p. 397-411. 

Baird, A.K., Baird, K.W., and Welday, E.E., 1974, Chemical trends 

across Cretaceous batholithic rocks of southern California: Geology, 

v. 2, no. 10, p. 493-495. 

Barker, F., 1981, Introduction to special issue on granites and rhyolites: 

a commentary for the nonspecialist: Journal of Geophysical 

Research, v. 86, no. Bll, p. 10131-10135. 

Bateman, P.C., 1983, A summary of critical relations in the central part 

of the Sierra Nevada batholith, California, U.S.A.: Geological Society 

of America Memoir 159, p. 241-254. 

Bateman, P.C., and Chappell, B.W., 1979, Crystallization, fractionation, 

and solidification of the Tuolumne intrusive series, Yosemite 

National Park, California: Geological Society of America Bulletin, 

Part I, v. 90, p. 465-482. 

Bennett, E.H., 1980, Granitic rocks of Tertiary age in the Idaho batholith and their relation to mineralization: Economic Geology, v. 75, 

p. 278-288. 

Bennett, E.H., and Knowles, C.R., 1985, Tertiary plutons and related 

rocks in central Idaho, in Mcintyre, D.H., ed., Symposium on the 

geology and mineral deposits of the Challis 1 o x 2° quadrangle, 

Idaho: U.S. Geological Survey Bulletin 1658 A-S. 

Berry, A.L., Dalrymple, G.B., Lanphere, M.A., and Von Essen, J.C., 

compilers, 1976, Summary of miscellaneous potassium-argon age 

measurements, U.S. Geological Survey, Menlo Park, California, for 

the years 1972-1974: U.S. Geological Survey Circular 727, 13 p. 

Bond, J.G., compiler, 1978, Geologic map of Idaho: Idaho Bureau of 

Mines and Geology, scale 1:500,000. 

Brooks, H.C., and Vallier, T.C., 1978, Mesozoic rocks and tectonic evolution of eastern Oregon and western Idaho, in Howell, D.G., and 

McDougall, K.A., eds., Mesozoic paleogeography of the Western 

United States: Society of Economic Paleontologists and 

Mineralogists, Pacific Coast Paleogeography Symposium II, Pacific 

Section, p. 133-145. 

Cater, F.W., Pinckney, D.M., Hamilton, W.B., Parker, R.L., Weldin, 

R.D., Close, T.J., and Zilka, N.T., 1973, Mineral resources of the 

Idaho Primitive Area and vicinity, Idaho: U.S. Geological Survey 

Bulletin 1304, 431 p. 

Centanni, J.P., 1985, Composition, synthesis and thermal stability of 

celadonitic white mica: implications for plutonic muscovite [abs.]: 

Geological Society of America Abstracts with Programs, v. 17, no. 

6, p. 347. 

Chappt>ll, B.W., and White, A.J.R., 1974, Two contrasting granite types: 

Pacific Geology, v. 8, p. 173-174. 

Chase, R.B., 1973, Petrology of the northeastern border zone of the 

Idaho batholith, Bitterroot Range, Montana: Montana Bureau of 

Mines and Geology Memoir 43, 28 p. 

__ 1977, Structural evolution of the Bitterroot dome and zone of 

cataclasis: Geological Society of America, Annual Meeting, Rocky 

Mountain Section, Missoula, Montana, Guidebook, no. 1. 

Collins, W.J., Beams, S.D., White, A.J.R., and Chappel, B.W., 1982, 

Nature and origin of A-type granites with particular reference to 

southeastern Australia: Contributions to Mineralogy and Petrology, 

v. 80, p. 189-200. 

Coxe, B.W., and Toth, M.l., 1983, Geochemical maps of the SelwayBitterroot Wilderness, Idaho County, Idaho, Missoula and Ravalli 

Counties, Montana: U.S. Geological Survey Miscellaneous Field 

Studies Map MF-1495-C, scale 1:125,000 (2 sheets). 

Criss, R.E., 1981, An 180/ 160, D/H and K-Ar study of the southern half 

of the Idaho batholith: Pasadena, California Institute of Technology, 

Ph.D. dissertation, 401 p. 

Criss, R.E., Lanphere, M.A., and Taylor, H.P., Jr., 1982, Effects of 

regional uplift, deformation, and meteoric-hydrothermal metamorphism on K-Ar ages of biotites in the southern half of the Idaho 

batholith: Journal of Geophysical Research, v. 87, no. B8, p. 7029-

7046. 

Criss, R.E., and Taylor, H.P., Jr., 1983, An 180/ 160 and D/H study of 

Tertiary hydrothermal systems in the southern half of the Idaho 

batholith: Geological Society of America Bulletin, v. 94, p. 640-663. 

Czamanske, G.K., Ishihara, S., and Atkin, S.A., 1981, Chemistry of rockforming minerals of the Cretaceous-Paleocene batholith in southwestern Japan and implications for magma genesis: Journal of 

Geophysical Research, v. 86, no. Bll, p. 10431-10469. 

Evans, K.V., and Zartman, R.E., 1981, U-Th-Pb zircon geochronology 

of Proterozoic Y granite intrusions in the Salmon area, east-central 

Idaho [ abs. ]: Geological Society of America Abstracts with Programs, 

v. 13, no. 4, p. 195. 

Fisher, F.S., Mcintyre, D.H., and Johnson, K.M., 1983, Geologic map 

of the Challis 1 o x 2° quadrangle, Idaho: U.S. Geological Survey 

Open-File Report 83-523, 41 p., 2 maps, scale 1:250,000. 

Fleck, R.J., and Criss, R.E., 1985, Strontium and oxygen isotopic variations in Mesozoic and Tertiary plutons of central Idaho: Contributions to Mineralogy and Petrology, v. 90, p. 291-308. 

Hahn, G.A., and Hughes, G.J., Jr., 1984, Sedimentation, tectonism, and 

associated magmatism of the Y ellowjacket Formation in the Idaho 

cobalt belt, Lemhi County, Idaho, in Hobbs, S.W., ed., Abstracts 

and summaries, Belt Symposium II: Montana Bureau of Mines and 

Geology Special Publication 90, p. 65-67. 

Hall, W.E., 1985, Stratigraphy and mineral deposits in Middle and Upper Paleozoic rocks of the black-shale mineral belt, central Idaho, 

in Mcintyre, D.H., ed., Symposium on the geology and mineral 

deposits of the Challis 1 o x 2° quadrangle, Idaho: U.S. Geological 

Survey Bulletin 1658 A-S. 

Hall, W.E., Rye, R.O., and Doe, B.R., 1978, Wood River mining district, 

Idaho-intrusion-related lead-silver deposits derived from country 

rock source: U.S. Geological Survey Journal of Research, v. 6, no. 

5, p. 579-592. 

Hall, W.E., Schmidt, E.A., Howe, S.S., and Broch, M.J., 1984, The 

Thompson Creek, Idaho, porphyry molybdenum deposit-an example 

of a fluorine-deficient molybdenum granodiorite system: Proceedings of the Sixth Quadrennial International Association on the 

Genesis of Ore Deposits Symposium, Stuttgart, Germany, p. 349-

357. 

Hamilton, Warren, 1963, Metamorphism in the Riggins region, western 

Idaho: U.S. Geological Survey Professional Paper 436, 95 p. 

__ 1976, Tectonic history of west-central Idaho [abs.]: Geological 

Society of America Abstracts with Programs, v. 8, no. 3, p. 378-

379. 

__ 1978, Mesozoic tectonics of the Western United States, in Howell, 

D.G., and McDougall, K.A., eds., Mesozoic paleogeography of the 

Western United States: Society of Economic Paleontologists and 

Mineralogists, Pacific Coast, Paleogeography Symposium II, Pacific 

Section, p. 33-70. 

__ 1983, Depth-related contrasts between Idaho and Sierra Nevada 

batholiths [abs.]: Geological Society of America Abstracts with Programs, v. 15, no. 5., p. 334. 

Harrison, T.M., and McDougall, 1., 1980, Investigations of an intrusive 

contact, northwest Nelson, New Zealand-I. Thermal, chronological, 

and isotopic constraints: Geochimica et Cosmochimica Acta, v. 44, 

p. 1985-2003. 

Hyndman, D.W., 1972, Petrology of igneous and metamorphic rocks: 

New York, McGraw-Hill Book Co., p. 86. 

__ 1979, Major tectonic elements and tectonic problems along the 

line of section from northeastern Oregon to west-central Montana: 

Geological Society of America Map and Chart Series MC-28C, 11 

p., scale 1:250,000. 

__ 1981, Controls on source and depth of emplacement of granitic 
196 THE IDAHO BATHOLITH AND ITS BORDER ZONE 

magma: Geology, v. 9, p. 244-249. 

__ 1983, The Idaho batholith and associated plutons, Idaho and 

western Montana: Geological Society of America Memoir 159, p. 

213-240. 

__ 1984, A petrographic and chemical section through the northern 

Idaho batholith: Journal of Geology, v. 92, p. 83-102. 

Jones, D.L., Silberling, N.J., and Hillhouse, J., 1977, Wrangellia-a 

displaced terrane in northwestern North America: Canadian Journal of Earth Sciences, v. 14, p. 2565-2577. 

Kiilsgaard, T.H., 1983, Geologic map of the Ten Mile West Roadless 

area, Boise and Elmore Counties, Idaho: U.S. Geological Survey 

Miscellaneous Field Studies Map MF-1500-A, scale 1:62,500. 

Kiilsgaard, T.H., Freeman, V.L., and Coffman, J.S., 1970, Mineral 

resources of the Sawtooth Primitive Area, Idaho: U.S. Geological 

Survey Bulletin 1319-D, 17 4 p. 

Kiilsgaard, T.H., and Lewis, R.S., 1985, Plutonic rocks of Cretaceous 

age and faults in the Atlanta lobe of the Idaho batholith, Challis 

quadrangle, in Mcintyre, D.H., ed., Symposium on the geology and 

mineral deposits of the Challis 1 o x 2° quadrangle, Idaho: U.S. 

Geological Survey Bulletin 1658 A-S. 

Kistler, R.W., 1974, Phanerozoic batholiths in western North America: 

Annual Reviews, Earth and Planetary Science Letters, v. 2, p. 

403-418. 

Kuntz, M.A., and Allen, C. C., 1984, Rock units and deformation on the 

western margin of the Idaho batholith near McCall, Idaho [abs.]: 

Geological Society of America Abstracts with Programs, v. 16, no. 

4, p. 227. 

Larsen, E.S., Jr., and Schmidt, R.G., 1958, A reconnaissance of the Idaho 

batholith and comparison with the southern California batholith: U.S. 

Geological Survey Bulletin 1070-A, 33 p. 

Lewis, R.S., 1984, Geology of the Cape Horn Lakes quadrangle, southcentral Idaho: University of Washington, Seattle, M.S. thesis, 91 p. 

Lindgren, W., 1900, The gold and silver veins of Silver City, De Lamar, 

and other mining districts in Idaho: U.S. Geological Survey 20th 

Annual Report, part 3, p. 65-256. 

Lopez, D.A., 1982, Constraints on the shape and position of theY ellowjacket (Proterozoic Y) depositional basin [abs.]: Geological Society 

of America Abstracts with Programs, v. 14, no. 6, p. 320. 

Lund, K., 1984, The continent-island arc juncture in west-central 

Idaho-a missing link in cordilleran tectonics [abs.]: Geological 

Society of America Abstracts with Programs, v. 16, no. 6, p. 580. 

Lund, K., Snee, L.W., and Evans, K.V., 1986, Age and genesis of 

precious metals deposits, Buffalo Hump district, central Idaho: Implications for depth of emplacement of quartz veins: Economic 

Geology, v. 81, no. 4, p. 990-996. 

Luthy, S.T., 1981, Petrology of Cretaceous and Tertiary intrusive rocks, 

Red Mountain-Bull Trout Point area, Boise, Valley, and Custer 

Counties, Idaho: Missoula, University of Montana, M.S. thesis, 109 p. 

Mabey, D.R., 1982, Gravity and magnetic features along the eastern 

margin of the Idaho batholith in central Idaho [abs.]: Geological 

Society of America Abstracts with Programs, v. 14, no. 6, p. 320. 

Marvin, R.F., and Dobson, S.W., 1979, Radiometric ages: Compilation 

B, U.S. Geological Survey: Isochron/West, no. 26, p. 17. 

McDowell, F.W., and Kulp, J.L., 1969, Potassium-argon dating of the 

Idaho batholith: Geological Society of America Bulletin, v. 80, p. 

2379-2382. 

Mcintyre, D.H., Ekren, E.B., and Hardyman, R.F., 1982, Stratigraphic 

and structural framework of the Challis Volcanics in the eastern 

half of the Challis 1 o x 2° quadrangle, Idaho, in Bonnichsen, B., 

and Breckenridge, R.M., eds., Cenozoic geology of Idaho: Idaho 

Bureau of Mines and Geology Bulletin 26, p. 3-22. 

Moore, J.G., 1959, The quartz diorite boundary line in the Western 

United States: Journal of Geology, v. 67, no. 2, p. 198-210. 

Myers, P.E., 1982, Geology of the Harpster area, Idaho County, Idaho: 

Idaho Bureau of Mines and Geology Bulletin 25, 46 p. 

Olson, H.J., 1968, The geology and tectonics of the Idaho porphyry belt 

from the Boise Basin to the Casto quadrangle: Tucson, University 

of Arizona, Ph.D. dissertation, 154 p. 

Pitcher, W.S., 1978, The anatomy of a batholith: Journal of the 

Geological Society of London, v. 135, p. 157-182. 

Reid, R.R., 1963, Reconnaissance geology of the Sawtooth Range: Idaho 

Bureau of Mines and Geology Pamphlet 129, 37 p. 

Ross, C.P ., 1928, Mesozoic and Tertiary granitic rocks in Idaho: Journal of Geology, v. 36, no. 8, p. 673-693. 

__ 1934, Geology and ore deposits of the Casto quadrangle, Idaho: 

U.S. Geological Survey Bulletin 854, 135 p. 

__ 1937, Geology and ore deposits of the Bayhorse region, Custer 

County, Idaho: U.S. Geological Survey Bulletin 877, 161 p. 

__ 1963, Geology along U.S. Highway 93 in Idaho: Idaho Bureau 

of Mines and Geology Pamphlet 130, 98 p. 

Ruppel, E.T., 1975, Precambrian and Lower Ordovician rocks in eastcentral Idaho: U.S. Geological Survey Professional Paper 889, 34 p. 

Schmidt, D.L., 1962, Quaternary geology of the Bellevue area in Blaine 

and Camas Counties, Idaho: U.S. Geological Survey Open-File 

Report 62-120, 92 p. 

__ 1964, Reconnaissance petrographic cross section of the Idaho 

batholith in Adams and Valley Counties, Idaho: U.S. Geological 

Survey Bulletin 1181-G, 50 p. 

Shand, S.J., 1949, Eruptive rocks: New York, John Wiley and Sons, 

Inc., 3d ed., 488 p. 

Shapiro, L., 1975, Rapid analysis of silicate, carbonate, and phosphate 

rocks-revised edition: U.S. Geological Survey Bulletin 1401, 76 p. 

Smitherman, J.R., 1985, Geology of the Stibnite roof pendant, Valley 

County, Idaho: Moscow, University of Idaho, M.S. thesis. 

Snee, L.W., 1982, Emplacement and cooling of Pioneer batholith, southwestern Montana: Columbus, Ohio State University, Ph.D. dissertation, 320 p. 

Snee, L.W., Lund, K., and Hoover, A.L., 1985, Discrimination of Idaho 

batholith plutons by structure and age [abs.]: Geological Society of 

America Abstracts with Programs, v. 17, no. 6, p. 409. 

Streckeisen, A.L., and others, 1973, Plutonic rocks-classification and 

nomenclature recommended by the lUGS Subcommission on the 

systematics of igneous rocks: Geotimes, v. 18, no. 10, p. 26-30. 

Sutter, J.F., Snee, L.W., and Lund, K., 1984, Metamorphic, plutonic, 

and uplift history of a continent-island arc suture zone, west-central 

Idaho [ abs. ]: Geological Society of America Abstracts with Programs, 

v. 16, no. 6, p. 670-671. 

Swanberg, C.A., and Blackwell, D.D., 1973, Areal distribution and geophysical significance of heat generation in the Idaho batholith and 

adjacent intrusions in eastern Oregon and western Montana: Geological Society of America Bulletin, v. 84, p. 1261-1282. 

Taubeneck, W.H., 1971, Idaho batholith and its southern extension: 

Geological Society of America Bulletin, v. 82, p. 1899-1928. 

Tilling, R.I., 1973, Boulder batholith, Montana: a product of two contemporaneous but chemically distinct magma series: Geological 

Society of America Bulltein, v. 84, p. 3879-3900. 

Toth, M.I., 1985, Petrology and evolution of the Bitterroot lobe of the 

Idaho batholith [abs.]: Geological Society of America Abstracts with 

Programs, v. 17, no. 4, p. 269. 

Weis, P.L., Schmitt, L.J., Tuchek, E.T., and Davis, W.E., 1972, Mineral 

resources of the Salmon River Breaks Primitive Area, Idaho: U.S. 

Geological Survey Bulletin 1353-C, 91 p. 

White, A.J.R., and Chappell, B.W., 1983, Granitoid types and their 

distribution in the Lachlan fold belt, southeastern Australia: Geological Society of America Memoir 159, p. 21-34. 

Wiswall, C. G., 1979, Structure and petrology below the Bitterroot lobe 

of the Idaho batholith: Northwest Geology, v. 8, p. 18-28. 

GPO 785-048/48039