Contributions to 

Mineral Resources Research, 1984 

Edited by A.L. BUSH 

Chapter titles are listed on 

volume contents page 

U.S. GEOLOGICAL SURVEY BULLETIN 1694 
DEPARTMENT OF THE INTERIOR 

DONALD PAUL HODEL, Secretary 

U.S. GEOLOGICAL SURVEY 

Dallas L. Peck, Director· 

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON: 1987 

For sale by the 

Books and Open-File Reports Section 

U.S. Geological Survey 

Federal Center 

Box 25425 

Denver, CO 80225 

Library of Congress Cataloging in Publication Data 

Contributions to mineral resources research, 1984. 

(U.S. Geological Survey bulletin ; 1694A-H) 

Includes bibliographies. 

Supt. of Docs. No.: I 19.3:1694 

1. Mines and mineral resources-United States. 2. Mines and mineral 

resources-Saudi Arabia. I. Bush, Alfred Lerner, 1919- 11. Series. 

QE75.B9 no. 1694 A-H 557.3 s 86--600330 

[TN23] [553' .0973] 
CONTENTS 

[Letters designate chapters] 

(A) Authigenic potassium feldspar in the Davis Formation (Upper Cambrian), 

Missouri, by G.A. Desborough, J.J. Connor, and Maya Elrick 1 

(B) Distribution, petrology, and mineral potential of peraluminous granitoid 

plutons in the eastern and southeastern Arabian Shield, Kingdom of Saudi 

Arabia, by Edward A. du Bray 13 

(C) Muscovites from selected aplites and quartz veins of White Pine County, 

Nevada, by D.E. Lee, E.L.M. Brandt, and J.D. Gleason 33 

(D) Chemical and stable-isotope data for micas from meta_clastic rocks of 

eastern White Pine County, Nevada, by D.E. Lee, R.E. Van Loenen, 

E.L.M. Brandt, and J.D. Gleason 41 

(E) Metallogeny of Archean and Proterozoic terranes in the Great Lakes 

region-A brief overview, by Paul K. Sims 55 

(F) Sodium hypochlorite as an aid to the extraction of clay minerals from 

black shales, by Harry C. Starkey 75 

(G) Platinum and palladium in a mafic dike, Sanders County, Montana, by 

R.E. Van Loenen, Joseph Haffty, and R.L. Earhart 81 

(H) Metalliferous oil shale in the Upper Devonian Gibellini facies of the 

Woodruff Formation in the southern Fish Creek Range, Nevada, by 

George A. Desborough, Forrest G. Poole, Richard K. Hose, and Gregory 

N. Green 91 

Volume Contents Ill 
Chapter A 

Authigenic Potassium Feldspar in the 

Davis Formation 

(Upper Cam brian), Missouri 

By G.A. DESBOROUGH, J.J. CONNOR, and MAYA ELRICK 

Chemical and mineralogical studies of the Davis Formation 

show that very fine grained authigenic potassium feldspar, 

a major silicate, formed diagenetically at the expense of 

detrital vitric volcanic ash deposited in a marine environment 

U.S. GEOLOGICAL SURVEY BULLETIN 1694 

CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 
CONTENTS 

Abstract 3 

Introduction 3 

Lithofacies, chemistry, and mineralogy of the Davis Formation 4 

General description 4 

Section near Leadwood, Missouri 4 

X-ray diffraction results 6 

Chemical data 6 

Comparison of the Davis Formation with other potassium feldspar-rich strata 9 

Chemical and mineralogical characteristics of marine strata with a major authigenic 

potassium-feldspar component 9 

Origin of authigenic potassium feldspar in marine strata 10 

References cited 10 

FIGURES 

I. Map showing location of the Rolla I 0 x 2 ° quadrangle in Missouri 4 

2. Maps showing thickness and lithofacies of the Davis Formation in the 

Rolla I 0 X 2 ° quadrangle in Missouri 5 

3-5. Graphs showing concentration of elements versus K20, in samples from 

exposed section of Davis Formation near Leadwood, Missouri: 

3. Barium versus K20 for 26 bulk samples 8 

4. Zirconium versus K20 for 25 bulk samples 8 

5. Rubidium versus K20 for 25 bulk samples 8 

6. Graph showing concentration of boron versus ratio of potassium 

feldspar:illite-smectite for 28 samples at exposed section near Leadwood, 

Missouri 8 

7. Diagram showing concentration of K20 versus K20:Na20 ratio for 

Precambrian, Cambrian, and Ordovician authigenic-feldspathic shales and 

siltstones in the midcontinent and Western United States and for insoluble 

residues from Ordovician carbonates of New York 10 

TABLES 

I. Chemical composition of shale samples of Davis Formation and Deadwood 

Formation 6 

2. X-ray diffraction and chemical data for samples of dolomite, limestone, and 

shale, Davis Formation, near Leadwood, St. Francois County, Missouri 7 

2 Contents 
CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 

Authigenic Potassium Feldspar in the Davis Formation 

(Upper Cambrian), Missouri 

By G.A. Desborough, J.J. Connor, and Maya Elrick 

Abstract 

The chemistry, mineralogy, and texture of the Upper Cambrian Davis Formation are consistent with the concept that its 

silicate minerals originated chiefly from vitric volcanic ash that 

was deposited in a marine environment and then altered during diagenesis to potassium feldspar. Dolomite, calcite, illitesmectite, chlorite, and sparse glauconite and pyrite in the Davis 

also formed diagenetically. Quartz occurs in some beds and 

is largely detrital. Very fine grained authigenic monoclinic 

potassium feldspar is the dominant mineral in all insoluble 

residue samples and constitutes as much as 60 percent of many 

of the residues; only minor amounts of detrital igneous orthoclase and microcline (to 60 J.A.m) are recognized in polished 

grain mounts and thin sections. By contrast, the authigenic 

potassium feldspar crystals are so small (0.03-5 J.A.m) that they 

can be conclusively identified by scanning electron microscopy, 

electron-microprobe, or X-ray diffraction methods. 

That the silicate minerals in the Davis Formation are 

authigenic is also suggested by the composition and distribution of nearby igneous rocks. The Davis is thickest in the region 

surrounding the core of Precambrian igneous rocks that now 

form the St. Francois Mountains. These Precambrian granites 

and rhyolites were also exposed during Davis deposition, yet 

they contributed only a minor local detrital component to the 

Davis Formation. Evidence for this conclusion includes (1) the 

very high K20:Na20 (weight percent) ratio of >55, and (2) 

K20 contents of 10-11 percent for beds in the Davis; these 

values are large compared with those for the local Precambrian 

igneous rocks. In addition, the K20, barium, zirconium, and 

boron concentrations for Davis samples differ significantly from 

the concentrations of these components in local Precambrian 

rocks. Further, the absence of quartz in some of the Davis 

samples, and the euhedral potassium feldspar grains smaller 

than 5 J.A.m, taken with the other evidence preceding, all suggest an authigenic origin for potassium feldspar in many beds 

in the Davis. 

The lower half of the marine Franconia Formation in Minnesota has been reported to contain nearly 50 percent 

authigenic potassium feldspar. These strata are age equivalent 

to the Davis Formation. 

INTRODUCTION 

In the past several years, highly potassic strata in 

which potassium feldspar is the major potassium-bearing 

mineral and also the major silicate mineral have commonly been considered to have originated from the diagenetic 

alteration of volcanogenic or tuffaceous debris. This interpretation has been applied to both lacustrine (Sheppard and Gude, 1965) and marine (Weiss, 1954) 

occurrences. Buyce and Friedman (1975) have emphasized 

these authigenic potassium feldspar occurrences in New 

York and in the upper Mississippi Valley regions for Upper Cambrian and Ordovician marin.e strata. 

Our investigations of the mineralogy and chemistry 

of calcareous and dolomitic shales of the Upper Cambrian Davis Formation and the underlying Upper Cambrian Bonneterre Formation in the Viburnum trend 

mineralized area along the west side of the St. Francois 

Mountains, as well as prior studies of the chemical composition of shales (and other strata) in Missouri by Connor and Ebens (1980), revealed that some of these shales 

are highly potassic, and that the major potassium-bearing 

mineral is potassium feldspar. Samples previously collected by others from the northeastern and central part 

of the Rolla 1 ° X 2 ° quadrangle in southeastern Missouri 

(fig. 1) were studied by us, and regional stratigraphic and 

lithofacies studies of Thacker and Anderson (1979a, b) 

for the entire quadrangle were subsequently useful for 

our interpretations. The bulk samples contain as much 

as 11 weight percent K2

0. Besides dolomite and calcite, 

the major minerals are potassium feldspar (orthoclase), 

disordered illite-smectite, and Fe-rich chlorite. Quartz 

may be absent or may form as much as about 30 percent 

of the noncarbonate fraction. Pyrite is generally detectable by X-ray diffraction in unoxidized samples. Potassium feldspar is the major mineral in most of the insoluble 

residues that were obtained by leaching the bulk samples 

with 20 percent HCl; potassium feldspar is estimated to 

compose as much as 60 percent of the whole rock for 

those samples considered to be shale on the basis of grain 

size. Detrital orthoclase and microcline derived from local 

Precambrian igneous terrane are also recognized by 

chemical, X-ray diffraction, and electron-microprobe 

methods. 

The bulk samples we studied were chiefly collected 

by others in the USGS for trace-element studies of metals 

and general lithologic character studies of the Davis 

Potassium Feldspar, Davis Formation, Missouri 3 
4Qo L 

'-\ 

\ 

I 

'\ )o Kansas City 

I 

I 

MISSOURI 

\ 

\ 

""' ' \ 

vst. Louisi 

\ 

' 

I oRolla JJJJD ~\_ 

I 

L__ _________ .l.. _______ , ~ 

/ : ._ _ __} 

Fi~ure 1. Location of the Rolla 1 ox 2° quadrangle in Missouri, 

w1th study area shown by vertical-lined box. 

Formation. We used these samples collected by others 

because they were considered to be typical of the Davis 

Formation. 

Kurtz (1975) proposed that the detrital component 

of the large volume of strata that makes up the Davis 

Formation in Missouri was principally derived from a 

northwesterly source. We present chemical, mineralogical, and other evidence that precludes as a major local 

source the Precambrian igneous rocks that were exposed 

to weathering and erosion at the time the Davis Formation was deposited. This evidence includes (1) dilution of 

the detrital component by carbonate buildup associated 

with relatively shallow water deposition peripheral to the 

exposed Precambrian igneous rocks of the ancestral St. 

Francois Mountains, (2) differences in chemical composition of the local Precambrian igneous terrane, compared 

with the composition of the silicates in the Davis Formation, (3) no widespread and (or) general increase in 

detrital components in the area proximal to the Precambrian igneous rock~ exposed at the time of deposition of 

the Davis Formation, as shown by isopach and lithofacies 

maps, (4) little, if any, significant grain-size sorting of 

detrital components from the Precambrian igneous rocks 

exposed at the time of deposition of the Davis Formation, and (5) similarities in chemical composition and 

mineralogy with Paleozoic marine strata elsewhere believed to have a primary volcaniclastic origin. Collectively, these aspects led us to consider an exogenous or remote 

source for noncarbonate minerals in the Davis Formation and to examine chemical and mineralogical methods 

for evaluation of local versus remote source areas for the 

detrital material and how it might relate to the authigenic 

potassium feldspar. 

4 Contributions to Mineral Resources Research, 1984 

LITHOFACIES, CHEMISTRY, AND MINERALOGY 

OF THE DAVIS FORMATION 

General Description 

Thacker and Anderson published isopach and 

sand:shale ratio maps (1979a), and carbonate:clastic ratio 

maps (1979b) of the Davis Formation for the Rolla 

1 °X 2 ° sheet, which includes areas adjacent to the exposed, dominantly silicic, Precambrian volcanic and 

granitic rocks (fig. 2). Because much of the presently exposed Precambrian igneous terrane was also exposed during deposition of the Davis Formation, we thought that 

a comparison of the chemistry of the silicate components 

of the sediment with that of the igneous rocks could be 

used to determine whether or not the igneous rocks contributed a major component of clastic material to the 

Davis Formation. If not, then other sources would have 

to be considered. Whole-rock chemical data for more 

than 100 samples of exposed igneous rocks were reported 

by Pratt and others (1980), and 75 of these that were proximal to our study area were used for comparison with 

the Davis Formation chemistry. The ratio K2

0:Na20 

was selected as the principal parameter for chemical comparison. By using K20:Na20, we excluded the variable 

amounts of quartz in both the sedimentary strata and the 

igneous rocks; additionally, the variable amounts of carbonate in the sediment were excluded. Chemical composition of four samples of shale from the Davis Formation 

obtained during th~ studies of Connor and Ebens (1980) 

are given in table 1. (The South Dakota sample is discussed later.) For these four shale samples, the 

K20:Na20 ratios are between 58 and 143. In contrast, 

for the 75 compared samples of Precambrian igneous 

rocks exposed in the Rolla 1 ° X 2 ° quadrangle north of 

lat.370JO' N. (Pratt and others, 1980), the K20:Na20 

ratios average less than 2 and range up to 87. Only nine 

samples (12 percent) of the igneous rocks from that area 

have K20:Na20 ratios exceeding 10; only four samples 

have K2

0:Na20 ratios exceeding 40. 

In addition, the boron content (1 00-150 ppm) in 

the Davis (table 1) is several times that of the mean of 

17 ppm (<10-100) for 75 samples of Precambrian igneous 

rocks (Pratt and others, 1980). The zirconium content 

(70 ppm) in the Davis is less than one-third of the mean 

of 250 ppm (70-500) for the same 75 igneous samples. 

Thus, on the basis of chemical data, it is unlikely that 

the insoluble residue fraction of the Davis was derived 

from the local Precambrian terrane of the ancestral St. 

Francois Mountains. 

Section near Leadwood, Missouri 

An interval of Davis Formation, exposed in a roadcut near Leadwood, was measured and sampled by the 
92 ° 90° 38 o,----------------------------------------------------.~-----------------------. 

A 

PRE-DERBYDOE RUN ROCKS 

38 .-.-~o-:- . ...,..,-..-...,...-,.----,------------------,.:-:-;-..,-----,-_,.,......,----,---,-~~....,..-.----,~--:--...,.-.,-.---,--~90o 

B ·. . . . ~· .. . ' ~ 

CLASTIC: CARBONATE 

RATIOS 

• 

>4 

1-4 

1/16-1 

<1/16 

PRESENT STUDY 

PRE-DAVIS ROCKS 

.. 

, • ' 

•' .· ... 

/ 

37 °~----------------------~----------------------------------------------------~ 15 MILES 

15 KILOMETERS 

Figure 2. Thickness and lithofacies of the Davis Formation in the Rolla 1 ox 2° quadrangle, Missouri. A, Isopach 

map of the Davis Formation. Thickness contours in feet; contour interval 40ft. B, Clastic:carbonate ratios of 

the Davis Formation. Both modified from Thacker and Anderson , 1979a, b. Sample localities: 1, 2, 3. 

Potassium Feldspar, Davis Formation, Missouri 5 
Table 1. Chemical composition of four shale samples of Davis 

Formation, Washington and Ste. Genevieve Counties, Mo., and 

a feldspathic shale from the Deadwood Formation, Lawrence 

County, S. Oak. 

[LOI, loss on ignition at 9000C; n.d., not determined; 0. excluded from 

total, contributed to LOI. All values in weight percent, except as noted] 

Deadwood 

Sample No. Davis Formation Formation 

Laboratory-- DI48004 Dl48015 Dl47998 DI48090 Dl20307W 

Field------- REOSH05 REOSH06 REOSHOI REOSH02 17BBR2 

sw2 45.6 46.6 48.0 47.0 57.2 

Al 2o3 16.8 17.0 16.0 17.0 17.2 

IFe2o3 3.8 2.8 4.8 4.4 5.2 

MgO 4.8 4.7 2.2 2.3 2.4 

CaO 5.7 5.4 6.4 6.4 2.5 

Na 2o o. 17 0.07 0.08 o. 13 o. 14 

KzO 9.8 10.0 10.0 11.0 12.0 

C02 (0.8) (0.2) 1.5 1.5 0.5 

TiOz 0.5 0.7 0.5 o.s 0.7 

c' organic (I. 1) (2. 7) 0.1 0.3 n.d. 

B (ppm) 100 150 150 150 <IO 

Zr (ppm) 70 70 70 70 100 

LOI l1.8 11.4 n.d. n.d. n.d. 

Total--- 99.0 98.7 289.6 29o.5 97.8 

1All iron calculated as Fe 2o3• 

2Low total due to high content of illite-smectite that 

contains hydration water. 

Missouri Department of Natural Resources in 1977. This 

is our principal set of samples that are discussed in the 

following text. The measured stratigraphic interval is 

about 35 m thick, but there is a covered interval of 

unknown thickness between sample numbers 75 and 76 

(table 2). The roadcut is only a few kilometers north of 

the Precambrian terrane that was exposed during deposition of the Davis Formation (fig. 2, locality 3). Thus, if 

Precambrian igneous rocks contributed significant 

detrital material to the Davis Formation, such detritus 

should be a major constituent of this locality. Twentynine composite samples (table 2) were assembled from 

74 field samples of the Davis in the roadcut; each of the 

composite samples represents a lithologic unit recognizable in the field. We then studied these composite samples 

by X-ray diffraction of the bulk, chemical analysis of the 

bulk, and X-ray diffraction of the 20 percent HCl insoluble fraction. 

X-ray Diffraction Results 

Dolomite: calcite ratios for composite bulk samples 

were semiquantitatively determined by measuring the 

relative intensities of the strongest X-ray diffraction peaks 

of each mineral on diffractometer traces. Dolomite is 

dominant over calcite in 21 (72 percent) of the samples; 

dolomite and calcite occur in equal amounts in 2 samples, 

6 Contributions to Mineral Resources Research, 1984 

and 6 of the samples are dominantly calcite (table 2). 

About 60 percent of the interval had dolomite:calcite 

ratios in the range of 2 to 100. 

Silicate mineral contents of bulk samples were 

determined by X-ray diffractometer study of insoluble 

residues, which were obtained by leaching the bulk 

samples with 20 percent HCI. Quartz is absent from seven 

lithologic units, aggregating 24 percent of the interval, 

in the middle and upper part (table 2). Quartz is dominant chiefly in the lower half of the Davis. Ratios of 

potassium feldspar to illite-smectite in insoluble residues 

were estimated on the basis of the X-ray diffractometer 

peak heights; the feldspar 002 peak near 27.5~8 CuKa 

and the illite-smectite 001 peak near 8.9~8 CuKa were 

used. The potassium feldspar:illite-smectite ratios range 

from 1.6 to 17 and indicate that potassium feldspar is 

dominant over illite-smectite in all analyzed samples (lable 

2). Potassium feldspar is the dominant mineral of the insoluble residue in 16 (of 29) samples, aggregating 62 percent of the interval studied (table 2). The potassium 

feldspar in the Davis is chiefly orthoclase, but three 

samples contain detectable microcline. 

The quartz-free samples, except for one with abundant microcline, probably were derived from a source 

remote from the area of deposition, whereas those with 

quartz and abundant microcline were derived from thenexposed Precambrian igneous rocks of the ancestral St. 

Francois Mountains. This interpretation is supported by 

the following evaluation of the chemical data. 

Chemical Data 

To evaluate the possible sources of detrital sediment, the concentrations ··of K2

0, barium, rubidium, 

and zirconium were determined for most of the bulk 

samples by X-ray fluorescence analysis using standards 

with accurately known concentrations of each element; 

results are given in table 2. Values for boron in insoluble 

residues were determined by M.S. Erickson in 1978 using 

semiquantitative emission spectroscopy: results for boron 

in table 2 are the mean for the samples in each interval. 

Figure 3 shows the concentration of barium versus 

K20; on this plot, the linearity of the quartz-free 

samples is remarkable. In other words, the lowtemperature-formed potassium feldspar of authigenic 

origin incorporates much less barium than the igneous 

orthoclase of the Precambrian granites that is of detrital 

origin for the quartz-bearing samples. About 70 percent 

of the samples line along or very close to the "quartzfree" line, whose projected extension to the ~0 content 

of pure potassium feldspar (16.95 percent K20} would 

indicate a concentration of about 450 ± 30 ppm barium 

in the authigenic potassium feldspar of the Davis Formation, if all K20 is in potassium feldspar. It is noteworthy that all samples that have a high barium 
Table 2. X-ray diffraction and chemical data for 29 samples of dolomite, limestone, and shale in the Davis Formation exposed 

along State Highway 8, SWV4NEV4 sec. 3, T. 36 N., R. 4 E., near Leadwood, St. Francois County, Mo. 

[NA, not analyzed; mineral ratios determined by relative X-ray peak intensities; K20, Ba, Rb, and Zr by XRF; analyst M. Elrick; B by semiquantitative emission spectroscopy; analyst M.S. Erickson; q, quartz; K-f, potassium feldspar] 

X-ray diffraction data Chemical data 

Potassium Igneou~ 

Sample Thickness Dolomite: 1 Dominant 1 Quartz2 feldspar: 2 orthoclase in 

No. (m) calcite noncarbonate illi te-smect i.te KzOl Ba 1 Rb 1 zr 1 B2 total orthoclase 

ratio mineral ratio 

7 0.30 100.0 q + 15.0 

8-12 1. 68 100 K-f + 4 

13 .45 100 K-f 0 6.5 

14-17 I. 22 40 K-f 0 4 

18 .45 100 q + NA 

19-23 1.42 100 K-f + 3.4 

24-32 2.9 100 K-f + 8 

33-37 I. 57 33 K-f + 5.7 

38 .30 100 K-f 0 11.5 

39 .25 8 K-f 0 2.5 

40 • 71 20 K-f + 7.5 

41-47 2.29 20 K-f 0 2.4 

48-52 2.44 0. I K-f 0 4.3 

53-58 1.40 30 K-f 0 2.0 

59.:..63 l. 52 100 q + 9.0 

64-66 l. 93 1.0 q + 4.0 

67 • 61 0.3 K-f + I. 6 

68 • 38 100 K-f + 3.3 

69-70 • 71 0.3 q + 2 

71 .20 0.6 q + 7 

72 2. 13 o. 1 q + 17 

73 1. 93 0.1 K-f + 4 

74 1. 22 2.5 q + 10.5 

75 • 71 1.7 q + 12.6 

76 1. 4 7 1.0 + 6.6 

77 1. 75 2.0 q + 7.5 

78 1. 07 1. 2 q + 5.2 

79 l. 98 13 q + II 

80 .33 7 K-f + 5 

1Bulk samples. 2rnsoluble residue; +, yes; o, no. 

concentration-in the range of 300-600 ppm-contain 

quartz, which is inferred to be detrital. The higher concentration of barium in some of the quartz-bearing 

samples is attributed to the detrital igneous orthoclas·e, 

that formed at higher temperatures, and which was locally 

derived from the Precambrian terrane of the St. Francois Mountains (fig. 2). The average content of barium 

in 75 igneous rock samples with more than 5 percent 

K20 is 1,000 ppm (calculated from Pratt and others, 

1980); of the 75 samples, those that are granites average 

625 ppm and range from 300 to 1 ,500 ppm barium. The 

granites with high orthoclase content typically have more 

barium than those with a relatively high content of 

microcline. 

Figure 4 shows the relationship of zirconium to 

K20 concentration for the Davis Formation (table 2). 

For the quartz-free (nonlocal detrital contribution) 

samples, it is obvious that the concentration of zirconium 

versus K2

0 is linear, within the precision of the method. 

However, 44 percent of the Davis samples with inferred 

local detrital contribution contain very high zirconium 

(wt. pet) (ppm) (ppm) (ppm) (ppm) (wt. pet) 

2. I 75 5 15 20 70 

5.5 150 30 40 140 8 

3.6 65 15 15 ISO 0 

7.8 200 NA NA 200 0 

2.8 90 10 25 50 50 

8.1 220 65 55 220 0 

2.9 105 10 35 117 50 

9.6 355 115 lOS 280 20 

2.2 30 15 5 100 0 

10.8 315 ISO 95 500 0 

2. 1 22 <5 <S 200 NA 

10.4 310 135 80 500 0 

1.9 30 5 10 540 0 

6.5 180 so so 230 0 

7.9 500 80 160 110 55 

6.8 220 45 60 170 20 

NA NA NA NA ISO NA 

NA NA NA NA 100 NA 

NA 525 120 145 200 NA 

7. l 325 40 115 100 40 

9.5 590 85 180 100 55 

10.6 600 115 160 100 50 

5.2 350 45 130 100 65 

7.4 425 60 140 100 55 

6.4 190 25 100 70 10 

5.9 190 30 80 100 20 

4.2 110 20 45 100 10 

7.4 275 60 75 100 30 

5.6 160 45 45 200 IS 

3Calculated from Ba content. 

concentrations. The mean for 40 samples of igneous rocks 

in the area that contain more than 5 percent K2

0 is 

270 ppm zirconium, which is more than 2 to 3 times 

higher than the mean for the Davis Formation. The zirconium content of 70 ppm obtained by emission spectroscopy for four samples of Davis (Connor and Ebens, 

1980) shown in table 1, that also contain 10-11 percent 

K20, is in excellent agreement with contents of 80 and 

95 ppm zirconium for the two quartz-free samples from 

the Leadwood section having the highest K20 concentrations (table 2 and fig. 4). The distribution patterns on 

figures 3 and 4 are similar and suggest that samples with 

high concentrations of both barium and zirconium had 

a local detrital source. 

Figure 5 shows the content of rubidium versus 

K20 concentration; the relationship is curvilinear for all 

samples, but the quartz-free samples have a rubidium: 

K20 ratio equal to or higher than those samples that 

contain quartz. The two quartz-free samples with highest 

K20 concentration also have the most rubidium (fig. 5). 

These two samples and the quartz-free sample with 

Potassium Feldspar, Davis Formation, Missouri 7 
800 

z 

0 600 :::::i 

...J 

~ 

a: 

w 

Q.. 

(fJ 

~ 

~ 

~ 

tO 

co 

00 

EXPLANATION 

e QUARTZ ABSENT 

• QUARTZ PRESENT 

m. MICROCLINE 

ABUNDANT 

+ PRECISION 

2 4 6 8 

K20, IN WEIGHT PERCENT 

12 

Figure 3. Concentration of barium versus K20 for 26 bulk 

samples of Davis Formation at exposed section near Leadwood, 

Mo. 

z 

0 

:::::i150 

...J 

~ 

a: 

w 

0..100 

en 

~ 

~ 

~ 

..: 

N 

EXPLANATION 

e QUARTZ ABSENT 

• QUARTZ PRESENT 

m. MICROCLINE 

ABUNDANT 

+ PRECISION 

.m 

.m 

4 6 8 

K20. IN WEIGHT PERCENT 

12 

Figure 4. Concentration of zirconium versus K20 for 25 bulk 

samples of Davis Formation at exposed section near Leadwood, 

Mo. 

200 I I I I I 

z EXPLANATION 

0 • QUARTZ ABSENT :::::; 1501-- • - ...J • QUARTZ PRESENT ~ • a: m. MICAOCLINE 

w ABUNDANT Q.. 1001-- -

~ ~ +PRECISION .. ~ 501- em .. - • .m ..0 

0:: .. .m . ' • .... I I I I 

00 2 4 6 8 10 12 

K20, IN WEIGHT PERCENT 

Figure 5. Concentration of rubidium versus K20 for 25 bulk 

samples of Davis Formation at exposed section near Leadwood, 

Mo. 

8 Contributions to Mineral Resources Research, 1984 

50 ppm rubidium also have low ratios (2.0-2.5) of potassium feldspar:illite-smectite (table 2); therefore much or 

most of the rubidium in these samples is in illite-smectite, 

and the authigenic potassium feldspar contains little if 

any rubidium. This situation accounts for the curvilinear 

relations between rubidium and K20 (fig. 5). The illitesmectite also contains both K20 and rubidium. The 

authigenic potassium feldspar contains less than 70 ppm 

rubidium, but the illite-smectite contains about 200-

300 ppm rubidium, on the basis of calculations using the 

high and low rubidium contents for the quartz-free 

samples and the estimated potassium feldspar:illitesmectite ratios. 

Figure 6 shows a generally inverse relationship between boron concentration in the insoluble residues and 

the ratio of potassium feldspar:illite-smectite. However, 

the three samples having the highest boron content have 

nearly twice as much boron as any of the other samples. 

This is probably due to the reporting intervals (for example, 100, 150, 300, 500 ppm) for the semiquantitative 

emission spectrographic method. In general, the relations 

on figure 6 indicate that boron is chiefly in illite-smectite. 

The amount of boron in illite-smectite is estimated at 

700-1,000 ppm, on the basis of calculations using the 

boron content, potassium feldspar:illite-smectite ratios, 

and K20 content of the bulk samples. Boron substituting for aluminum in potassium feldspar has an effect 

on the a and c cell dimensions when present in concentrations of more than about 600 ppm (Martin, 1971; 

Sheppard and Gude, 1965, 1973). Cell parameter data obtained for potassium feldspar in these samples show that 

no more than about 500 ppm of boron could be in either 

the detrital orthoclase or the authigenic potassium 

feldspar. This contrasts sharply with lacustrine authigenic 

600~--~~--~T~--~~--~~----~~--~~----1~--~~.--~ 

~ 5001--

:::::; 

...J 

~400-

a: 

w 

0.. 3001-

~ ~ 2001--

z 

cci' 1001-

.. 

I 

2 

• 

•• 

• .. 

I I I 

EXPLANATION 

• QUARTZ ABSENT 

• QUARTZ PRESENT 

m. MICROCLINE 

ABUNDANT 

•• • • 

I I I • I 

-

-

-

-

. -

4 6 8 10 12 14 16 18 

potassium feldspar 002 X-ray peak 

intensity illite-smectiteoo1 ratio 

Figure 6. Concentration of boron versus ratio of potassium 

feldspar:illite-smectite for 28 samples at exposed section near 

Leadwood, Mo. 
potassium feldspar in Tertiary strata that commonly 

contains 2,500-4,400 ppm boron (Sheppard and Gude, 

1965, 1973). As·noted, the low boron values (mean= 17 

ppm, range <10-100 ppm) for the Precambrian igneous 

rocks of the St. Francois Mountains (Pratt and others, 

1980) are consistent with the premise that boron in the 

Davis Formation is chiefly in authigenic illite-smectite, 

but some boron is also in authigenic potassium feldspar 

according to the integrated chemical and mineralogical 

data. 

COMPARISON OF THE DAVIS FORMATION 

WITH OTHER POTASSIUM 

FELDSPAR-RICH STRATA 

Several years ago, Buyce and Friedman (1975) 

discussed aspects of the authigenic potassium feldspar occurrences in Upper Cambrian and Lower and Middle Ordovician marine strata of several States including New 

York. In New York they studied insoluble residues of carbonate strata of Late Cambrian, Ordovician, Silurian, 

and Devonian age. They concentrated on the Lower Ordovician Tribes Hill Limestone that contains abundant 

potassium feldspar (12.1-23.9 normative percent) in carbonate beds that contain about 2.5-4.0 percent K2

0; 

Upper Cambrian carbonates contain similar abundances 

(Buyce and Friedman, 1975, p. 814). Tester and Atwater 

(1934) first recognized marine rocks that consist mainly 

of authigenic potassium feldspar in Upper Cambrian and 

Lower Ordovician strata in Minnesota. Subsequent 

studies in Minnesota and Wisconsin of Upper Cambrian 

and Lower and Middle Ordovician marine strata that are 

enriched in potassium feldspar include those of Gruner 

and Theil (1937), Berg (1952), Weiss (1954), Woodard 

(1972, 1974), and Odom (1974). The potassium-rich intervals described by these investigators a~e less than about 

10 m thick for Minnesota (Berg, 1952) and Wisconsin occurrences, although Bowie and others (1966) reported intervals as much as 16 m thick that average 9.5 percent 

K20 in Lower Cambrian marine strata of northwest 

Scotland. 

The Proterozoic Dripping Spring Quartzite in Gila 

County, Ariz., contains marine strata as thick as 34 m 

that consist predominantly of potassium feldspar and 

contain as much as 14 percent K20 (Granger and Raup, 

1964; Nutt, 1981). 

We recently recognized shale beds consisting chiefly 

of authigenic potassium feldspar in the Upper Cambrian 

Deadwood Formation in the Black Hills of South 

Dakota; chemical composition is given in table 1. Barnes 

and Bell ( 1977) 9escribed abundant authigenic potassium 

feldspar that constituted more than 60 percent of the sand 

and silt fraction of the Franconian Point Peak .Member 

of the Wilperns Formation (Upper Cambrian) in westcentral Texas. The Point Peak Member is as much as 

60 m thick, but it is limestone in the upper half (Barnes 

and Bell, 1977, p. 23). 

Thus three widely separated marine basins of Late 

Cambrian age in Missouri, South Dakota, and Texas have 

high contents of authigenic potassium feldspar. We suggest that the potassium feldspar in these strata formed 

by the process suggested by Sheppard and Gude (1973): 

vitric volcanic ash in a diagenetic environment of high 

pH was chemically modified tC? potassium feldspar. 

Determining whether these marine strata with very high 

contents of potassium feldspar represent single or multiple volcanic events requires further detailed investigations, including paleontological, isotopic, mineralogical, 

and chemical studies. 

The Davis Formation, which is more than 60 m 

thick where it is chiefly noncarbonate (Thacker and 

Anderson, 1979a), is certainly the thickest Paleozoic formation on the North American craton that has authigenic 

potassium feldspar as the principal noncarbonate mineral. 

If the precursor were volcanic ash, the volcanic event or 

events might have had significant effects on the biota, 

as well as on sedimentation. 

CHEMICAL AND MINERALOGICAL 

CHARACTERISTICS OF MARINE STRATA 

WITH A MAJOR AUTHIGENIC POTASSIUM 

FELDSPAR COMPONENT 

We have used a simple chemical method for the 

recognition of marine strata that contain more than 

about 50 percent of potassium feldspar that is authigenic. Figure 7 illustrates the chemical method that 

uniquely identifies strata (or their insoluble residues) 

that are rich in authigenic potassium feldspar, by plotting K2

0:Na20 ratios. The six circles for Upper Cambrian Franconian strata on figure 7 include one analysis 

by Berg (1952) for the Franconia Formation in Minnesota 

and five analyses of Upper Cambrian strata of the present study from table 1. Points on figure 7 for the Precambrian Dripping Spring Quartzite in Arizona are from Nutt 

(1981). For Ordovician samples, two are from Minnesota 

and Wisconsin (Gruner and Thiel, 1937; Woodard, 1972). 

Five points for Ordovician strata in New York (fig. 7) 

are from data obtained on insoluble residues (Braun 

and Friedman, 1969, table 2). It is noteworthy that even 

the most potassium-rich illites studied by Weaver (1967) 

have a maximum K20:Na20 ratio of about 7 and a maximum K2

0 content of about 6 percent. Because of the 

presence of analcime (Na-rich) and other zeolites in 

samples consisting mainly of authigenic potassium 

feldspar in lacustrine basins (Sheppard and Gude, 1965), 

the K2

0:Na20 values are 40 or less for bulk samples. 

Data for these are not shown. 

Potassium Feldspar, Davis Formation, Missouri 9 
400 ,---,----,1----.---,----,----.---.- ~ I I • 

100 -

50 r10 '-

7 

5 

Field of most 

Paleozoic 

clay shales 

(Weaver, 1967) 

c§) 

• 0 , 

0 • •• • 0 • •• • 

• • 

EXPLANATION 

• ORDOVICIAN 

• 

0 UPPER CAMBRIAN 

• PRECAMBRIAN 

-

-

-

I I I ~--~---~--~--~----~--- --~--~16 

K20 , IN WEIGHT PERCENT 

Figure 7. Concentration of K20 versus K20: Na20 ratio for 

Precambrian, Cambrian, and Ordovician authigenic-feldspathic 

shales and siltstones in the midcontinent and Western United 

States and for insoluble residues from Ordovician carbonates 

of New York. 

Our studies by X-ray diffraction methods of either 

bulk samples or insoluble residues that consist mainly of 

authigenic potassium feldspar have the pattern of orthoclase (monoclinic potassium feldspar) for marine strata 

that have not been exposed to temperatures above about 

150-200 °C. 

ORIGIN OF AUTHIGENIC POTASSIUM 

FELDSPAR IN MARINE STRATA 

Several mechanisms have been pre, posed for the inplace growth of orthoclase in the Cambrian and Ordovician strata of the midcontinent region of the United 

States. Tester and Atwater (1934) suggested that the 

authigenic potassium feldspar is a normal product of consolidation and cementation in these marine strata. Gruner 

and Theil (1937) considered that it grew at the expense 

of, or during dissolution of, "clayey material." Berg 

10 Contributions to Mineral Resources Research, 1984 

(1952) thought that detrital feldspar served as a nucleus 

for growth of orthoclase overgrowths, but also that 

authigenic crystals formed in pore spaces and filled the 

interstices between the detrital quartz grains. Weiss (1954) 

was the first proponent of an ultimate origin related to 

volcanic ash; he showed orthoclase replacing a 

brachiopod shell and believed that the host Cambrian 

strata were the product of feldspathization of bentonite 

derived from volcanic ash. Hay (1966) suggested reaction 

of the zeolite clinoptilolite to form potassium feldspar 

in Ordovician marine tuffs in Minnesota. Woodard (1972) 

called on direct precipitation of potassium feldspar from 

sea water due to high concentrations of amorbhous silica. 

Odom (1974) believed that the material described by 

Woodard (1972) was the result of authigenic overgrowth 

on detrital feldspar and that the illite (once present) served 

as a major source of potassium, aluminum, and silicon. 

We generally agree with Weiss (1954) that volcanic 

detritus was the main precursor to the development of 

potassium-rich beds in the lower Paleozoic marine strata 

of the midcontinent, and we endorse a similar origin for 

the Proterozoic occurrence in Arizona. 

REFERENCES CITED 

Barnes, V.E., and Bell, W.C., 1977, The Moore Hollow Group 

of central Texas: Bureau of Economic Geology Report of 

Investigations 88, 169 p. 

Berg, R.R., 1952, Feldspathized sandstone: Journal of Sedimentary Petrology, v. 22, p. 221-223. 

Bowie, S.H.U., Dawson, J., Gallagher, M.J., Ostle, D., 

Lambert, R. St. J., and Lawson, R.I., 1966, Potassiumrich sediments in the Cambrian of northwest Scotland: 

Transactions Institute of Mining Metallurgy, v. 75, 

p. B125-B145. 

Braun, Moshe, and Friedman, G.M., 1969, Carbonate lithofacies and environments of the Tribes Hill Formation 

(Lower Ordovician) of the Mohawk Valley, New York: 

Journal of Sedimentary Petrology, v. 39, p. 113-135. 

Buyce, M.R., and Friedman, G.M., 1975, Authigenic K-feldspar 

on the Proto-Atlantic shelf: Journal of Sedimentary 

Petrology, v. 45, p. 808-821. 

Connor, J.J., andEbens, R.J., 1980, Geochemistry of bedrock 

units in Missouri and parts of adjacent states: U.S. Geological Survey Professional Paper 954-F, p. F1-F55. 

Granger, H.C., and Raup, R.B., 1964, Stratigraphy of the Dripping Spring Quartzite, southeastern Arizona: U.S. Geological Survey Bulletin 1168, 119 p. 

Gruner, J.W., and Theil, G.A., 1937, The occurrence of fine 

grained authigenic feldspar in shales and silts: American 

Mineralogist, v. 22, p. 842-846. 

Hay, R.L., 1966, Zeolites and zeolitic reactions in sedimentary 

rocks: Geological Society of America Special Paper 85, 

130 p. 

Kurtz, V.E., 1975, Franconian (Upper Cambrian) trilobite 

faunas from the Elvins Group of southeastern Missouri: 

Journal of Paleontology, v. 49, no 6, p. 1009-1043. 
Martin, R.F., 1971, Disordered authigenic feldspars of the series 

KA1Si30 8 - KBSi 30 8 from southern California: American 

Mineralogist, v. 56, p. 281-291. 

Nutt, C.J., 1981, A model of uranium mineralization in the 

Dripping Spring Quartzite, Gila County, Arizona: U.S. 

Geological Survey Open-File Report 81-524, 49 p. 

Odom, I.E., 1974, Syngenetic sanidine beds from Middle Ordovician St. Peter sandstone, Wisconsin-A discussion: 

Journal of Geology, v. 82, p. 112-116. 

Pratt, W.P., Odland, S.K., Hubert, A.E., Siems, D.F., and 

Viets, J. G., 1980, Spectrographic and chemical analyses 

of exposed Precambrian rocks, Rolla 1 ° by 2 ° quadrangle, 

Missouri: U.S. Geological Survey Open-File Report 80-25, 

26 p. 

Sheppard, R.A., and Gude, A.J., 3rd, 1965, Potash feldspar 

of possible economic value in the Barstow Formation, San 

Bernardino County, California: U.S. Geological Survey 

Circular 500, 7 p. 

___ 1973, Boron-bearing potassium feldspar of authigenic 

origin in closed-basin deposits: U.S. Geological Survey 

Journal of Research, v. 1, p. 377-382. 

Tester, A.C., and Atwater, G.l., 1934, The occurrence of 

authigenic feldspars in sediments: Journal of Sedimentary 

Petrology, v. 4, p. 23-31. 

Thacker, J.L., and Anderson, K.H., 1979a, Preliminary isopach 

and sand/shale ratios of the Cambrian Davis Formation, 

Rolla 1° by 2 ° quadrangle, Missouri: U.S~ Geological Survey 

Miscellaneous Field Studies Map MF-1002-D, scale 1:250,000. 

___ 1979b, Preliminary carbonate lithofacies maps of the 

Cambrian Davis Formation, Rolla 1° by 2 ° quadrangle, 

Missouri: U.S. Geological Survey Miscellaneous Field 

Studies Map MF-1002-E, scale 1:250,000. 

Weaver, C.E., 1967, Potassium, illite and the ocean: Geochimica 

et Cosmochimica Acta, v. 31, p. 2181-2196. 

Weiss, M.P., 1954, Feldspathized shales from Minnesota: Journal of Sedimentary Petrology, v. 24, p. 270-274. 

Woodard, H.H., 1972, Syngenetic sanidine beds from Middle 

Ordovician Saint Peter sandstone, Wisconsin: Journal of 

Geology, v. 80, p. 323-332. 

___ 1974, Syngenetic sanidine beds from Middle Ordovician 

Saint Peter sandstone-A reply: Journal of Geology, v. 82, 

p. 116-119. 

Potassium Feldspar, Davis Formation, Missouri 11 
Chapter B 

Distribution, Petrology, and 

Mineral Potential of Peraluminous 

Granitoid PI utons in the Eastern and 

Southeastern Arabian Shield, 

Kingdom of Saudi Arabia 

By EDWARD A. DU BRAY 

A description of the fundamental characteristics, including 

mineral resource potential, of 13 peraluminous granite plutons in the 

eastern and southeastern Arabian Shield 

U.S. GEOLOGICAL SURVEY BULLETIN 1694 

CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 
CONTENTS 

Abstract 15 

Introduction 15 

Acknowledgments 16 

Age and setting 17 

Petrography 17 

Compositional and textural features 17 

Characteristic minerals 19 

Muscovite and Fe-Li mica 19 

Biotite 19 

Garnet 19 

Geochemistry 20 

Garnets 20 

Muscovite and Fe-Li mica 22 

Major elements 22 

Trace elements 22 

Isotopes 26 

Mineral deposit potential 27 

Petrogenesis 28 

Conclusions 30 

References cited 30 

FIGURES 

1. Map showing locations of peraluminous plutons in eastern and southeastern 

Arabian Shield 16 

2. Ternary diagram showing average modal compositions for 13 peraluminous 

plutons in eastern and southeastern Arabian Shield 19 

3. Graph showing composition of analyzed micas in terms of Si, AI, and octahedral cations (Fe, Mg, Ti, Mn) 22 

4. Harker diagrams showing selected major element-oxide data for samples 

from the peraluminous plutons 23 

5. Ternary ACF diagram showing average compositions of the peraluminous 

plutons 23 

6. Ternary diagram showing average proportions of normative quartz, albite, 

and orthoclase in the peraluminous plutons 23 

7. Trace-element variation diagrams for the peraluminous plutons 26 

TABLES 

1. Petrographic characteristics of peraluminous granitoid plutons in eastern and 

southeastern Arabian Shield 18 

2. Electron microprobe analyses of garnets from three peraluminous plutons in 

eastern and southeastern Arabian Shield 20 

3. Electron microprobe analyses of micas from three peraluminous plutons in 

eastern and southeastern Arabian Shield 21 

4. Average major-element, trace-element, and CIPW compositions for 

peraluminous plutons in eastern and southeastern Arabian Shield 24 

14 Contents 
CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 

Distribution, Petrology, and Mineral Potential of 

Peraluminous Granitoid Plutons in the Eastern and 

Southeastern Arabian Shield, Kingdom of Saudi Arabia 

By Edward A. du Bray 

Abstract 

Thirteen peraluminous granitoid plutons in the eastern 

and southeastern Arabian Shield constitute a suite that 

resembles metallogenically specialized or 5-type granites. Most 

of these plutons are about 1 5 km2 in areal extent, but several 

are as large as 75 km2• Preliminary geochemical results indicate that the plutons contain anomalous concentrations of 

tin, tungsten, and other rare metals. Because of their ore deposit 

potential, these plutons warrant significantly more study. 

The plutons are composed of medium-grained, leucocratic two-feldspar (nonperthitic microcline and albite) granite 

and granodiorite, both having hypidiomorphic- to 

xenomorphic-inequigranular textures. Biotite and muscovite 

or Fe-Li mica are present in varying proportions. Almandinespessartine garnet (as much as 1 modal percent) is present in 

seven of the plutons. Zircon, apatite, ilmenite, topaz, and 

fluorite are the characteristic accessory minerals. 

The major-element chemistry of these plutons is distinctive relative to other granitoid rocks, in that they contain more 

than 72 weight percent Si02 and low total iron, CaO, and 

MgO and have a peraluminous index (molar Al 20 3/ 

(Na20+ Ca0)) approaching or greater than 1.0. In addition, 

the incompatible trace elements Be, Pb, Sn, Li, F, Rb, Y, and 

Nb are significantly enriched in most of these plutons, whereas 

the compatible elements Ba, Cu, La, Sr, and Zr are depleted 

relative to global averages for granites. 

Common lead isotope data suggest that most of these 

plutons include a component generated by partial melting of 

preexisting continental crust. Major-element chemical and 

petrographic data are consistent with crystallization of these 

plutons under conditions of PH2o = 200-400 megapascals 

(2-4 kilobars). Varying degrees of incompatible-element enrichment and compatible-element depletion among the 

peraluminous plutons indicate that they were derived from 

compositionally dissimilar source terranes. Minimum melting 

of four compositionally distinct source terranes could produce 

the trace-element variation observed in the 13 plutons. 

INTRODUCTION 

Geologic investigations conducted throughout the 

world have defined a set of petrologically similar 

granitoid rocks (Streckeisen, 1976) known variously as 

S-type, ilmenite-series, peraluminous, or metallogenically 

specialized granites; these categories overlap in part. 

S-type granites (Chappell and White, 1974) are characterized by 

relatively low Na20 content, 

greater than 1 percent CIPW normative corundum, 

initial 87Sr/86Sr ratios greater than 0.708, 

high silica content, 

irregular interelement variation within plutons, and 

the presence of muscovite or other peraluminous 

minerals. 

In comparison, 1-type granites (Chappell and White, 

1974) are characterized by 

relatively high Na2

0 content, 

less than 1 percent CIPW normative corundum, 

initial 87Sr / 86Sr ratios in the range 0. 704-0.706, 

broad compositional range, 

regular interelement variation within plutons, and, 

commonly, the presence of hornblende. 

Ishihara (1977) has divided granitoid rocks into the 

magnetite series and the ilmenite series. Magnetite-series 

granitoid rocks contain easily recognizable magnetite and 

many are similar to 1-type granites, whereas ilmeniteseries granitoid rocks contain ilmenite but are practically free of opaque oxide minerals (less than 0.1 volume 

percent) and many are similar to S-type granites. Granites 

are peraluminous (Shand, 1949) when the peraluminous 

index, molar (Al20/(Ca0 + Na20 + K20)), exceeds 1. 

Metallogenically specialized granites are characterized by 

incompatible-element (Cox and others, 1979, p. 336; Be, 

Li, F, Pb, Rb, Y, and Nb) enrichment relative to global 

granite averages, compatible-element (Sr, Ba, Zr, and La) 

depletion, the presence of muscovite, and a high silica 

content; and many of these granites are peraluminous 

(Tischendorf, 1977). 

S-type, ilmenite-series, and some peraluminous and 

metallogenically specialized granitoid rocks were derived 

from a source terrane that has undergone previous chemical fractionation, and they may represent partial melt 

Peraluminous Granitoid Plutons, Saudi Arabia 15 
products from sedimentary or evolved igneous rocks. 1-

type, magnetite-series, and nonmetallogenically specialized granitoid rocks are principally first-stage melts 

generated from primitive material, and they have had litde if any interaction with the continental crust. S-type, 

ilmenite-series, peraluminous, and metallogenically 

specialized granitoid rocks share many characteristic 

features. Most of these rocks, similar to those discussed 

here, contain muscovite, but their most significant shared 

trait from the viewpoint of economic geology is their frequent association with deposits of tin and tungsten. The 

association of ore deposits with these rocks has generated 

considerable interest in their genesis and distribution 

throughout the world. 

The· present study area was defined by a reconnaissance geochemical exploration of the eastern and 

southeastern Proterozoic Arabian Shield (du Bray and 

others, 1982). Thirteen peraluminous, metallogenically 

specialized plutons were identified (fig. 1) that strongly 

resemble highly evolved plutons located elsewhere in the 

world that are associated with tin, tungsten, and raremetal deposits. The Arabian Shield is already known for 

its peralkaline granites (agpaitic index, molar 

(K20 + Na20)/ Al20 3>1) (Drysdall, 1979; Stoeser and 

Elliott, 1980). However, both peraluminous and peralkaline plutons are volumetrically minor relative to the vast 

number of calc-alkaline plutons that constitute about 40 

percent of the Arabian Shield. 

Geoscientists working elsewhere in the Arabian 

Shield have reported muscovite-bearing plutons. Elliott 

(1980, 1983), who studied the Gaharra alkali-feldspar 

granite, was perhaps first to realize the significance of 

muscovite-bearing plutons. Cole and others (1981) identified a muscovite-bearing stock in the northeastern Arabian Shield that contains tungsten- and tin-mineralized 

rock. Stacey and others (1983) reported muscovitebearing phases in the Ad Dawadimi batholith, located in 

the central Arabian Shield, northwest of the study area. 

Stoeser (1983) reported muscovite with and without 

garnet in several trondhjemitic plutons located in the 

southern Arabian Shield, and Sable (1982) discussed two 

muscovite-bearing monzogranites that crop out close to 

the Saudi Arabian-North Yemen border. 

Acknowledgments 

The work on which this report is based was performed in accordance with a work agreement between the 

U.S. Geological Survey (USGS) and the Saudi Arabian 

Ministry of Petroleum and Mineral Resources. D.B. 

Stoeser (USGS) performed microprobe analyses of the 

garnets, and Jane Hammarstrom (USGS) performed 

microprobe analyses of the micas. J .E. Quick, J .S. 

Pallister, and J.S. Stuckless (all of USGS), and C.R. 

16 Contributions to Mineral Resources Research, 1984 

0 Bishah 

Tafshah 

8 9 

N 

0 

~ 

Zayd Madha 

10 12

Bwana 

11 

6 

EXPLANATION 

M USCOVITE-8 EARl NG 

PLUTON- Number 

keyed to table 1 

MAJOR FAULT 

EASTERN EDGE OF 

ARABIAN SHIELD 

200 KILOMETERS 

100 MILES 

Figure 1. Locations of peraluminous plutons in the eastern 

and southeastern Arabian Shield. Numbers and names keyed 

to table 1. Plutons numbered 8-11 constitute the Madha 

assemblage. 

Ramsey (Saudi Arabian Directorate General of Mineral 

Resources) provided careful reviews of the manuscript. 

All modal analyses were conducted by Ali Mohammed 

Jabarti. The paper benefited greatly from discussions with 

D.B. Stoeser, J.E. Quick, J.S. Pallister, and C.R. 

Ramsay. 
AGE AND SETTING 

The peraluminous granitoid rocks appear to be 

among the youngest plutonic rocks exposed in the Proterozoic Arabian Shield. They intrude andesitic and 

dacitic volcanic and volcaniclastic rocks and molasse of 

the Balaban and Murdama groups, respectively. Layered 

rocks assigned to the Balaban group, intruded by coeval 

dioritic rocks, were deposited between 800 and 700 Ma 

and are thought to be part of ensimatic-arc terranes that 

were successively accreted to the Arabian plate during its 

cratonization (Greenwood, Anderson, and others, 1980; 

Greenwood, Stoeser, and others, 1982). Rocks of the 

Murdama group, intruded by batholiths composed of 

granite, granodiorite, and tonalite, were deposited following accretionary events at a time coincident with the 

beginning of a collisional event, approximately 670 Ma 

(Schmidt and others, 1979; Greenwood, Anderson, and 

others, 1980; Greenwood, Stoeser, and others, 1982; 

Stacey and Stoeser, 1983). Rocks of the Murdama and 

Balaban groups were metamorphosed at greenschist- and 

amphibolite-facies conditions during pluton emplacement 

and have been deformed several times. No genetic relationship between the peraluminous plutons and rocks of 

the Balaban and Murdama groups is apparent. 

The 13 peraluminous plutons in the study area 

(fig. 1) are between 3 and 75 km2 in areal extent; the 

average area is about 15 km2• The outcrop patterns of 

the plutons range from circular to smoothly elliptical. The 

shapes of the plutons, their intrusive relationships, and 

the fact that they are free of postcrystallization deformational features suggest that they represent small, postorogenic, diapirically emplaced intrusions. 

Geochronologic data also indicate that the plutons 

were emplaced late during the development of the Arabian Shield. Zircon separated from the Sabhah and 

Sitarah monzogranites and from the Madha granodiorite 

and dated by the U/Pb zircon method indicates ages of 

605 ± 5 m.y. (Stacey and others, 1983), 627 ± 30 m.y. (J .S. 

Stacey, oral commun., 1983), and 670± 5 m.y. (Stoeser 

and Stacey, 1983), respectively. 

The peraluminous plutons of the eastern and 

southeastern Arabian Shield are in an intracratonic setting, similar to that of the muscovite-bearing granites that 

crop out in Western North America (Miller and Bradfish, 1980). The latter plutons are located near the 

inferred edge of Proterozoic crust, in an area that was 

intensely deformed coincident with and immediately 

preceding their emplacement. Likewise, the peraluminous plutons in the eastern and southeastern Arabian 

Shield are located in an orogenic zone, and their emplacement followed a hypothesized collision between 

the Arabian plate and an unknown terrane to the east, 

at a time coincident with the Pan-African thermotectonic event (Clifford, 1970). The Bulayfah-Ad 

Dafinah and Nabitah-Bamdah fault zones (to be referred 

to collectively as the BNFZ) (fig. 1), along which are 

numerous outcrops of ultramafic rocks, are the inferred 

sutures; and the area affected by the collision was probably a broad zone that encompassed the peraluminous 

plutons (Schmidt and others, 1979; Schmidt and Brown, 

1982; Stoeser and others, 1982). 

The 13 plutons are in two geographic assemblages. 

One assemblage includes seven plutons in the northern 

part of the study area (plutons 1-7, fig. 1, table 1), and 

the other assemblage of six is located about 400 km to 

the south. For the sake of later discussion, four of the 

plutons in this latter assemblage (plutons 8-11, along or 

west of the HNFZ; fig. 1, table 1) are called the Madha 

assemblage plutons. 

Plutons 1, 2, 4, and 6-13 do not contain xenoliths 

of either wall rock or refractory residuum. Plutons 3 and 

5 have locally stoped and engulfed large blocks of wall 

rock, but there is no indication of chemical interaction 

between these blocks and the enclosing rock, nor between 

the other non-Madha assemblage plutons and the enclosing wall rock. Plutons of the Madha assemblage, however, are characterized by localized mafic schlieren 

layering, which suggests interaction with wall rock, 

possibly including its partial assimilation. 

PETROGRAPHY 

Compositional and Textural Features 

Most of the peraluminous plutons of the eastern 

and southeastern Arabian Shield are alaskitic alkalifeldspar granites, although some samples are syenogranitic, granodioritic, or trondhjemitic. Most of these 

plutons are medium grained, hypidiomorphic to xenomorphic granular, and inequigranular seriate to equigranular, although some are subporphyritic. The 

hypidiomorphic inequigranular texture that characterized 

most of these rocks indicates simultaneous crystallization 

of the feldspars and quartz. The petrographic features 

of these plutons are summarized in table 1, and modal 

data, obtained by counting between 400 and 700 points 

on stained slabs averaging 60 cm2 of surface area, are 

summarized on figure 2. 

The modal compositions of plutons 1, 3-7, and 

12-13 cluster in the alkali-feldspar granite field (fig. 2). 

Modes for plutons 2 and 8 plot in the monzogranite field. 

Plutons 9-11 of the Madha assemblage are distinctly more 

granodioritic than the remaining plutons. 

Most samples contain subequal amounts of quartz, 

plagioclase, and potassium feldspar, and from 0 to 5 percent mafic silicate minerals (although the Madha 

assemblage plutons contain significantly more plagioclase). The potassium feldspar is microcline characterized by gridiron twinning and is amazonitic in some 

Peraluminous Granitoid Plutons, Saudi Arabia 17 
... 

= 

Table 1. Petrographic characteristics of peraluminous granitoid plutons in eastern and southeastern Arabian Shield 

~ [Locality numbers keyed to figures 1 and 2. N, number of samples on which mean (x) and standard deviation (o) are based. Tr., trace; leader(-), not present; a.f.g., alkali-feldspar granite] 0 

~ 0.: Modal mfnerals 1 Muscovite or s. s· Locality number Texture and grain size N Quartz Microcline Plagioclase Biotite Fe-Li mica Garnet :I x x x x x x "' and pluton tS 0' C1' o- (j" 0' 

0 

~ 1 Sabhah a. f. g. Hypidiomorphic, inequigranular, 15 33 3 33 3 30 4 3 :;· 

~ medium grained. 234 llol 

- 2 Hawshah Hypidiomorphic, inequigranular 13 33 6 7 29 6 - - 4 2 :11:1 

., 

~ monzogranite. to subporphyritic, medium ::) 

c:: grained. .. 235 ~ 

~ 3 Si tarah a •. f. g. Hypidiomorphic to xenomorphic, 30 33 7 8 29 5 - - 3 ., inequigranular to subporphyr- :11:1 

1) itic • ., 

~ 

~ 4 Sahah a.f.g. n Hypidiomorphic, equigranular, 15 36 5 38 5 23 5 2 

:r medium to coarse grained • ... 

I.C 3 = 

~ 5 Tarban a.f.g. Hypidiomorphic to xenomorphic, 33 31 9 37 7 30 8 1 2 1 1 Tr. 

inequigranular, medium 

grained. 

6 Kebad a.f.g. Hypidiomorphic, inequigranular, 4 27 5 25 5 44 4 Tr. 4 2 Tr. 

medium grained. 

7 Mahail a. f. g. Hypidiomorphic, equigranular, 4 41 5 17 1 38 6 - - 4 

medium grained. 

8 Tafshah Hypidiomorphic, inequigranular, 6 32 5 31 9 35 7 Tr. 2 1 0.5 

monzogranite. medium grained, gneissic. 

9 Madha ----do------------------------ 5 29 10 13 5 53 11 Tr. 5 3 .2 

granodiorite. 

10 Zayd Hypidiomorphic, equigranular, 9 30 6 21 7 48 8 Tr. 1 1 Tr. 

granodiorite. medium grained. 

11 Habbah Hypidiomorphic, inequigranular, 7 27 9 25 8 47 8 - - 1 1 Tr. 

granodiorite. medium grained, pegmatitic. 234 . 3 12 Bwana a.f.g. Hypidiomorphic, inequigranular 13 6 39 7 23 9 - - 4 2 

to subporphyritic, medium 

grained. 

13 Gaharra a.f.g. Hypidiomorphic, equigranular, 19 32 7 28 5 35 7 Tr. 5 3 

medium grained. 

1Accessory minerals (principally fluorite, ilmenite, and zircon) are <1 percent in most samples. 

2Indicates subporphyritic phase. 

3Amazonite. 
plutons. It is rarely perthitic, which indicates that conditions favoring subsolvus crystallization of two alkali 

feldspars prevailed during crystallization. The plagioclase is unzoned, polysynthetically twinned albite. Zircon, apatite, fluorite, ilmenite, and rarely topaz or 

cassiterite are the accessory minerals identified in these 

plutons. 

Q 

Figure 2. Ternary diagram showing average modal compositions for 13 peraluminous plutons in the eastern and 

southeastern Arabian Shield. Identification numbers refer to 

the 13 plutons; locations shown in figure 1, petrographic data 

in table 1. Composition boundaries after Streckeisen (1976). 

A, Alkali feldspar; Q, quartz; P, plagioclase. 

Characteristic Minerals 

Muscovite and Fe-Li Mica 

Muscovite or Fe-Li mica is characteristic of the 

peraluminous plutons. It is interstitial, occurs as subhedral to anhedral flakes, and appears to have nucleated 

late in the crystaliization process. It forms coarse intergrowths and is in textural equilibrium with biotite. Sericite 

is a common alteration product of plagioclase in some 

of the plutons but is readily distinguishable from primary 

micas. 

Three species of primary mica were distinguished 

on the basis of pleochroism. The most common of these 

is colorless in thin section and is muscovite. This mica 

attains the largest average grain size (2 mm) and has the 

best crystal form of the three species observed. A second 

type of mica, an Fe-Li mica, identified in the Bwana 

alkali-feldspar granite and in the Tarban alkali-feldspar 

granite, is colorless to light green or slightly bluish light 

green. This coloration may be attributable to ferrous iron 

substituting for octahedrally coordinated aluminum (Deer 

and others, 1966, p. 203). A third type of mica, also an 

Fe-Li mica, occurs in samples of the Sabhah and Sitarah 

alkali-feldspar granites and the Hawshah monzogranite, 

and displays colorless. to light-brown pleochroism. It 

typically includes numerous prismatic crystals of zircon 

having distinctive pleochroic halos. The pleochroic 

scheme and paragenesis of this mica most closely match 

those described for zinnwaldite. The coloration is attributable to ferric iron substituting for octahedral 

aluminum (Deer and others, 1966, p. 203). 

Biotite 

Biotite is a nearly ubiquitous constituent of the 

peraluminous plutons. In most samples it is subhedral to 

anhedral, and interstitial. In thin section it displays lightyellow to brown or reddish-brown pleochroism. Numerous included prismatic crystals of zircon characterize the 

biotite. 

Textural relations and mineralogy suggest that a 

reaction relationship between garnet and biotite existed 

in seven of the plutons. The Mahail alkali-feldspar granite 

and Habbah granodiorite, both garnet bearing, contain 

no biotite. In the remaining five garnetiferous plutons, 

biotite is in textural disequilibrium and was being resorbed during the final stages of crystallization. These 

observations indicate that garnet crystallized at the expense of biotite in a generalized reaction such as: biotite 

+ MnO + A120 3 + Si02 (from liquid) go to garnet + 

muscovite. This reaction was first proposed by Miller and 

Stoddard (1981) for the paragenesis of garnets observed 

in leucogranites from the Old Woman-Piute Range in 

southern California. 

Garnet 

Small, subhedral grains of apparently unzoned 

garnet, from 0.10 to 0.50 mm in diameter, were identified 

in seven of the plutons in amounts of as much as 1 modal 

percent (table 1). The garnet has a remarkably similar appearance from one pluton to the next. Completely 

isotropic almandine-spessartine garnet forms light-rosecolored grains containing few inclusions. The garnets 

show no sign of resorption or alteration; however, a reaction between adjacent grains of garnet and biotite, involving the growth of garnet at the expense of biotite, 

is commonly observed. The petrographic observations 

suggest that the garnets nucleated from a silicate melt. 

Garnet-bearing samples and plutons containing more 

than 0.1 weight percent MnO are positively correlated 

(tables 1-3). · 

Peraluminous Granitoid Plutons, Saudi Arabia 19 
Geochemistry 

Garnets 

Major-element analyses of garnet from three of the 

peraluminous plutons were obtained by using an electron 

microprobe (table 2). The garnets analyzed are mixtures 

of the almandine, spessartine, and pyrope end members. 

Garnets from the Mahail alkali-feldspar granite and the 

Tafshah monzogranite are almandine rich, whereas 

garnets from the Madha granodiorite contain approximately equal proportions of almandine and spessartine 

and are substantially enriched in the pyrope component 

relative to garnets from the other two plutons. 

Chemical zoning in the garnets is difficult to 

demonstrate. Results of analyses made at the core and 

Table 2. Electron microprobe analyses of garnets from three peraluminous plutons in eastern and southeastern Arabian Shield 

[Analyses performed using an automated, six~spectrometer electron probe microanalyzer at U.S. Geological Survey, Denver, Colo. All analyses 

made using a focused 5~~-tm beam. Number in parentheses following core/rim designation, number of grains analyzed. Mean and standard 

deviation (±)are listed in cases where more than one grain was analyzed. Leader(-}, nil amount; a.f.g., alkali~feldspar granite] 

Sio2 

Al 2o3 

Fe02 

MnO 

MgO 

CaO 

Ti02 

cr 2o3 

Total--

Mahail a.f.g. 

169584 

Core(2) 

35.91±0.5 

21.47± .os 

31.07± .60 

9.55± .26 

0.30± .03 

.18± • 01 

.02± .02 

.02± .02 

98.52 

169584 

Rim(!) 

36.24 

21.31 

31.53 

9.30 

0.28 

.14 

(.01 

<.01 

98.80 

Tafshah monzogranite 

169620 

Core( 2) 

36.56±0.40 

20.96± .27 

25.79± .43 

13.59± .67 

0.77± .01 

1.11± .04 

.02± .02 

.01± .01 

98.81 

169620 

Rim( 1) 

36.94 

21.67 

24.55 

15.31 

0.65 

1.19 

.OS 

.01 

100.37 

Madha granodiorite 

169635 

Core( 3) 

36.96±0.28 

21.55± .32 

18.39± .36 

19.17± .21 

3.09± .20 

0.52± .08 

.13± .04 

.01± .01 

99.82 

Formula proportions based on 12 oxygens 

Si 

Al 

Fe 

Mn 

Mg 

Ca 

Ti 

Cr 

Almandine-----

Spessartine---

Pyrope--------

Grossularite--

Andradite-----

Uvarovite-----

2.9811 

2.1006 

2.1571 

0.6711 

.0370 

.0157 

.0014 

.001 

74.9 

23.2 

1.3 

.4 

• 1 

• 1 

2.9992 

2.0781 

2.1817 

0.6521 

.0349 

.0125 

75.8 

22.6 

1.2 

.4 

3.0126 

2.0357 

1.7772 

0.9479 

.0944 

.0975 

.0008 

.ooos 

2.9943 

2.0704 

1.6642 

1.0508 

0.0779 

.1032 

.0029 

.0004 

End members, molecular percent 

60.9 

32.5 

3.2 

3.3 

• 1 

57.4 

36.3 

2.7 

3.4 

.2 

2.9746 

2.0455 

1.2353 

1.3174 

0.3704 

.0448 

.0081 

.0014 

41.8 

44.2 

12.5 

1.1 

.4 

• 1 

Precision1 

(percent) 

0.62 

.ss 

.38 

.91 

3.94 

2.50 

47.3 

88.9 

1Analytical precision based on counting statistics given in percent of amount 

indicated. 

2Total iron as FeO. 

20 Contributions to Mineral Resources Research, 1984 
the rim of garnets from the Mahail alkali-feldspar granite 

are very similar, probably within the limits of analytical 

uncertainty. Some small variation between the rim and 

core of garnets from the Tafshah monzogranite is 

suggested. The results suggest that the core is enriched 

in the almandine and pyrope components relative to the 

rim and that the rim is enriChed in the spessartine component relative to the core. 

Table 3. Electron microprobe analyses of micas from three peraluminous plutons in eastern and southeastern Arabian Shield 

[Analyses performed using an automated, six-spectrometer electron probe microanalyzer at U.S. Geological Survey, Reston, Va. All analyses 

made using a defocused 10-15-JLm beam. Number following grain number (Gr.), number of analyses per grain. Mean and standard deviation 

(±)listed when more than one grain analysis was performed. Leader(-), nil amount; a.f.g., alkali-feldspar granite] 

Sio2 

Al 2o3 

2

Fe0 

MgO 

CaO 

Na 2o 

K2o 

Ti02 

MnO 

Sabhah a.f.g. 

169438 

Gr. 1(3) Gr. 2(1) 

36.71±0.39 35.73 

21.59± .03 21.48 

22.51± .42 24.74 

0.07± .01 0.07 

.06± .02 .04 

• 28± • 0 2 • 19 

9.29± .os 9.24 

.23± .05 .23 

.58± .03 .63 

.04± .04 

.09± .02 .09 

3.49± .18 3.43 

2.07± .05 2.09 

Gr. 3(3) 

36.67±0.52 

21.55± .28 

24. 23±1. 09 

0.09± .01 

• 09± • 02 

.25± .01 

9.41± .11 

.21± .01 

• 63± • 01 

.05± .04 

• 08± • 02 

3.28± .16 

2.21± .06 

Bwana a.f.g 

169625 

Gr. 1 (3) Gr. 2(2) 

38.69±2.40 42.86±0.04 

21.49±1.64 22.24± .21 

17.51±3.22 16.38± .06 

0.08± .02 0.08± .01 

.04± .01 .02± .01 

.29± .01 .22± .01 

9.09± .23 10.02± .17 

.10± .02 .14 

.99± .18 .98± .01 

.09± .03 .09± .01 

• 03± • 02 • 03 

4.08± .17 4.02± .21 

1.80± .21 2.09± .12 

Habbaah 

granodiorite 

169609 Precision1 

Gr. 1 (3) 

45.20±0.29 

43.06± .18 

4.98± .36 

1.24± .10 

0.01 

.30± .04 

10.90± .08 

.47± .01 

.05± .02 

.25± .03 

.19± .14 

4.30± .08 

(percent) 

1.44 

0.78 

1. 33 

19.52 

31. 14 

4.87 

1. 86 

16.92 

3.66 

39.92 

26.79 

6.25 

Total4 95.53 96.51 97.35 92.55 97.43 99.48 

Si 

Al 

Al 

Fe 

Mg 

Ti 

Mn 

Ca 

Na 

K 

Ba 

F 

Cl 

3oH 

2.93 

1.07 

.97 

1.50 

.01 

.o 1 

.04 

.04 

.95 

.88 

.01 

1. 11 

2.86 

1.14 

.89 

1.66 

.01 

.o 1 

.04 

.03 

.95 

• 8 7 

.01 

1.12 

Formula proportions based on 12 (0, OH, F) 

2.90 

1.10 

.91 

1. 61 

.01 

.o 1 

.04 

.01 

.04 

.95 

.82 

.01 

1.17 

3.10 

.90 

1. 13 

1.18 

.01 

.o 1 

.07 

.04 

.93 

1.04 

.01 

.96 

3.22 

.78 

1. 19 

1. 03 

.01 

.o 1 

.06 

.03 

.96 

.96 

1.05 

3.08 

.92 

1.66 

.28 

• 13 

.02 

.05 

.95 

.01 

.04 

1. 9 6 

1Analytical precision based on counting statistics given in percent of the amount indicated. 

2rotal iron as FeO. 

3calculated for formula F+Cl+OH=2.0 for a 12 (0, OH, F) formula. 

4corrected for F and Cl. 

Peraluminous Granitoid Plutons, Saudi Arabia 21 
Muscovite and Fe-Li Mica 

Major-element analyses of micas from three of the 

peraluminous plutons were obtained using an electron 

microprobe (table 3). Sums of the oxides and of cation 

occupancy in the octahedral sites in samples of mica from 

the Sabhah and Bwana alkali-feldspar granites are low. 

Li02, which may be a principal component of these 

micas, was not determined and may account for the low 

totals. 

The composition of mica from the Habbah granodiorite plots in the muscovite field (fig. 3) defined by 

Miller and others (1981). As is characteristic of most 

plutonic muscovite, this mica has low contents of total 

Fe, F, and Li02, and high H2

0 content relative to other 

micas (Deer and others, 1962, p. 16-18). Muscovite in 

the Habbah granodiorite, like mica in the Kebad, Mahail, 

Tafshah, Madha, and Zayd plutons, is nonpleochroic; 

the composition of muscovite in the Habbah granodiorite 

is probably representative of the mica in these plutons. 

Samples of mica from the Sabhah and Bwana 

alkali-feldspar granites are compositionally similar to 

zinnwaldite. The high inferred content of Li02, high 

contents of total Fe, MnO, and F, low contents of MgO 

and H20, and intermediate content of Al20 3 relative to 

other micas are all characteristic of zinnwaldite (Deer and 

others, 1962, p. 93). Compared to typical zinnwaldite, 

however, mica from these two plutons is characterized 

by slightly higher occupancy in the octahedral site and 

by a slightly higher ratio of AI to Si (fig. 3). The 

pleochroism observed in mica from these two plutons is 

also characteristic of mica from plutons 2-5 and 13; the 

composition of the zinnwaldite in the Sabhah and Bwana 

alkali-feldspar granites is probably representative of that 

in plutons 2-5 and 13. 

Major Elements 

Forty-three samples of the peraluminous plutons 

were analyzed for their major-element composition. 

These analyses and analyses for 11 other samples 

(Delfour, 1979, 1980; Kanaan, 1979) were used to calculate the average major-element composition of each 

pluton (table 4, figs. 4-6). Data for the 54 samples are 

presented by du Bray (1983a). The Si02 content ranges 

from 72 to 78 weight percent and averages 75 weight percent; the plutons are weakly to strongly peraluminous 

(many samples contain more than 1 percent CIPW normative corundum); total iron, MgO, Ti02, and CaO 

contents are very low. These chemical characteristics are 

similar to those attributed to S-type granites (Chappell 

and White, 1974) and to metallogenically specialized 

granites (Tischendorf, 1977). The muscovite-bearing 

plutons are characterized, however, by subequal Na20 

and K20 contents, whereas S-type and metallogenically 

22 Contributions to Mineral Resources Research, 1984 

Si 

F + 169625-2• + M 169625-1. + A 169438-1-, 169609 -1 

169438-2r-...169438-3 

s 

Figure 3. Composition of analyzed micas with sample and 

grain numbers in terms of Si, AI, and octahedral cations (Fe, 

Mg, Ti, Mn) (modified from Miller and others, 1981). Solid 

circle, zinnwaldite; open circle, muscovite. Zinnwaldites of 

Deer and others (1962, p. 93) are plotted ( + ), without Li 20 

component, for comparison. Locations of mica end-members 

M (muscovite, fluormuscovite, and paragonite-parallelogramshaped field at right edge of triangle); F (ferrimuscovite); C, 

(celadonite); A (annite-phlogopite); and S (siderophylliteeastonite) are also plotted. 

specialized granites are distinctive because they contain 

significantly more K20 than Na20. Either different 

types of source terrane or different conditions of magma 

genesis could account for the alkali element discrepancy. 

As is typical of S-type granites, the within-pluton 

scatter of major-element data on Harker diagrams is 

significant. Average oxide values were calculated for each 

pluton and plotted on Harker and ternary diagrams (table 

4, figs. 4-6). Some between-pluton trends, however, are 

apparent: Al20 3 and CaO decrease and K2

0 increases 

with increasing Si02; Na20, total iron, and the peraluminous index display no consistent covariation with 

Si02 content. Analyses for plutons 8-11 of the Madha 

assemblage plot away from or at the periphery of the data 

fields defined by all other plutons. 

The minor-element compositions of these plutons 

are as follows (in weight percent): Ti02 ranges from 0 

to 0.37 and averages 0.06; MnO and P20 5 range from 

0 to 0.2 and average 0.06 and 0.05, respectively; MgO 

ranges from 0 to 0.4 and averages 0.10; total H20 

ranges from 0 to 1.25 and averages 0.56. None of these 

oxides displays a consistent covariation with Si02 • 

Trace Elements 

Table 4 allows comparison of the trace-element 

compositions of the peraluminous granites with those of 

metallogenically specialized granites, granites of the world, 

and the granitoid rocks of the eastern and southeastern 

Arabian Shield. Note that the trace-element compositions 
..... 16 

z 

w 

~ 15 a: 

w 

0.. 

~ 

M 

0 13 C\1 

< 

....~· ..... 2 <(2 

.-w 

0~ 

..... a: 

0~ (l)z 

LJ..-

2 

..... 

z 

w 

u 

a: 

w 

0.. 

~ 

o" 

~ 

u 

0 

6 

~ w 

u 5 a: 

w 

0.. 

~ 4 

o" C\1 3 ~ 

z 

2 

6 ..... 

z 

w 

~ 5 a: 

w 

0.. 4 

::!: 

o" 3 C\1 

::t: 

2 

~ 1.2 c 

z 

- 1.1 Cl) 

:::;) 

0 

~ 1.0 ~ 

::I 

~ 0.9 

w 

0.. 8.0 

72 73 

.11 

e11 

_,, 

2 •8•4 

• •1 e5 6ex 10i! •• ,3 

7 12 

.s 

74 75 76 

Si02, IN PERCENT 

77 78 

Figure 4. Harker diagrams showing selected major elementoxide data for samples from the peraluminous plutons. Plotted points represent average values for the plutons in weight 

percent (table 3); numbers 1-13 refer to the plutons, whose 

locations are shown in figure 1 and whose petrographic 

characteristics are listed in table 1. X, average value for all 54 

samples. Peraluminous index = AI 20/(Ca0+ K20+ Na20) on 

a molecular basis. 

A 

Figure 5. Ternary ACF diagram showing average compositions 

of the peraluminous plutons. A = AI 20 3-Na20-K20; C = 

CaO; F = FeO+ MgO; all in molecular amounts. Pluton nos. 

1-13 keyed to table 1 (petrographic characteristics) and figure 

1 (map locations). Line, boundary between 1- and S-type compositions, respectively, as suggested by White and Chappell 

(1977). X, average composition for all 54 samples; triangle, 

average composition of 2,327 granites (Krauskopf, 1967); +, 

average composition of 962 metallogenically specialized 

granitoids (Tischendorf, 1977); rectangle, average composition 

of 226 precursor granites (Tischendorf, 1977). 

Q 

Figure 6. Ternary diagram showing average proportions of 

normative quartz (Q), albite (Ab), and orthoclase (Or) in the 

peraluminous plutons. Pluton nos. 1-13 keyed to table 1 

(petrographic characteristics) and figure 1 (map locations). 

Triangle, average composition of 2,327 granites (Krauskopf, 

1967); +, average composition of 962 metallogenically 

specialized granitoids (Tischendorf, 1977); rectangle, average 

composition of 226 precursor granites (Tischendorf, 1977). 

Polybaric minimum melting compositions (minima trend) in 

system Si02-NaAISi 30 8-KAISi 3

0 8-H 20 indicated by open 

circles for PH...20 equal to from 50 MPa (top right circle), 100, 

200, 400, SOu, to 1,000 MPa (bottom left circle) (0.5, 1, 2, 4, 

5, and 10 kbar) (Winkler and others, 1975). 

Peraluminous Granitoi~ Plutons, Saudi Arabia 23 
Table 4. Average major-element, trace-element, and CIPW compositions for peraluminous plutons in eastern and southeastern 

[Number in parentheses following pluton locality name corresponds to number on figure 1 and table 1. N, number of samples on which mean 

anhydrous and corrected for fluorine. Major element analyses from several laboratories, using a combination of X-ray fluorescence, atomic abspectrography; Rb, Sr, Y, Zr, Nb, and Pb by X-ray fluorescence using a 109Cd excitation source; Sn, Ba, La, Ce, and Nd by X-ray fluorescence 

Si02 

Al 2o3 

Fe 2o3 

FeO 

MgO 

Cao 

Na 2o 

K2o 

Ti02 

P2o5 

MnO 

F 

Q 

c 

or 

ab 

an 

wo 

en 

fs 

mt 

hm 

il 

ap 

fr 

Li 

F 

Cu 

Rb 

Sr 

y 

Zr 

Nb 

Pb 

Sn 

Ba 

La 

Ce 

Nd 

Be 

Sabhah(l) 

a.f.g. 

Hawshah(2) Sitarah(3) 

monzogranite a.f.g. 

Sahah(4) 

a. f. g. 

Tarba n( 5) Kebad ( 6) 

a.f.g. a.f.g. 

Major-element analyses, in weight percent 

Mahail(7) Tafshah(8) 

a.f.g. monzogranite 

N = 5 N = 6 N = 6 N = 4 N = 7 N = 2 N = 2 N = 3 

75.4 ±1.69 

13.2 ±0.35 

0.60± .27 

o. 77:1: • 60 

0.11:1: .10 

0.69± .25 

4.43:1: .34 

4.52± .35 

o. 11:1: • 09 

o. 04± • 04 

0.03± .02 

0.21±.13 

31.05 

0.37 

26.73 

37.51 

1. 66 

0 

.28 

.80 

.8 7 

0 

.21 

• 10 

.43 

N = 18 

85± 62 

2035±917 

16:1: 15 

354±106 

22± 35 

181:1: 73 

123:1: 41 

43± 20 

53± 25 

(10:1: -

(8± -

23:1: 12 

55± 21 

(25± -

6± 6 

75.2 ±1.12 

13.6 ±1.00 

0.79±0.10 

1.01± .47 

0.14± .11 

0.81:1: • 50 

·3.63±1.21 

4.53± .98 

0.12± .05 

o. 04± • 01 

0.04± .02 

0.26± .25 

35. 19 

2.01 

26.75 

30.73 

I. 89 

0 

0.35 

I. 07 

1. 15 

0 

.23 

.10 

.53 

N = 13 

105± 79 

1915±1421 

34± 55 

427:1: 188 

46± 45 

96± 50 

100± 37 

19:1: 9 

38± 7 

12± 9 

123:1: 143 

33± 20 

75± 39 

30± 15 

4± 4 

76.9 ±0.97 

12.8 :1: .25 

0.42± .16 

o. 45± • 11 

0.30:1: .64 

o. 55± • 38 

4.25± .49 

4.09± .09 

0.04± .02 

0.04± .07 

0.06± .06 

0.24± .06 

35.53 

I. 10 

24. 14 

35.93 

0.74 

0 

• 75 

.53 

.61 

0 

.08 

• 10 

.49 

75.8 :1:0.83 75.9 ±0.77 

12.9 :1: .31 13.6 :1: .35 

0.66± .23 0.36± .21 

0.76± .27 0.35± .26 

0.11± .04 

0.67± .14 

4.08± .50 

4. 75± • 28 

0.09± .03 

0.08± .09 

0.41± .26 

4.61± .75 

4.40± .15 

0.04± .05 

0.03± .03 0.02± .01 

0.03± .01 0.11± .06 

0.24± .05 0.17± .07 

75.0 ±1.27 

14.5 ±0.57 

0.63± .40 

0.35± .26 

0.04± .01 

o. 42± • 08 

4.60± .31 

4.08± .35 

0.03± -

0.14± .04 

0.04:1: .03 

0.18± .05 

CIPW norms, in weight percent 

32. 79 

0.51 

28.04 

34.54 

1. 40 

0 

0 

• 28 

• 76 

.96 

• 17 

.07 

.49 

31.62 

I. 00 

26.00 

39.02 

0.68 

0 

0 

.20 

.49 

.53 

• 08 

.05 

.35 

Trace-element analyses, in ppm 

32.49 

2.55 

24. 12 

38.95 

0 

0 

0.10 

.15 

.92 

0 

.06 

.33 

.33 

N = 36 N = 15 N = 34 N = 5 

139± 105 

2320±2162 

14:1: 16 

582± 177 

15± 8 

107± 30 

76± 17 

48± 16 

82± 118 

31:1: 23 

30± 46 

13± 6 

33± 14 

(25± -

6± 3 

61± 40 

2083±708 

6:1: 4 

252± 72 

18± 7 

135± 39 

142± 25 

67± 19 

33:!: 8 

<10± -

84± 15 

32± 16 

79± 26 

34± 14 

8± 7 

67± 77 

1316±1139 

15± 22 

488± 205 

32:!: 68 

82± 38 

89:!: 30 

31:1: 20 

61± 39 

13% 10 

193: 217 

15% 10 

38± 26 

(25: -

7± 20 

52% 19 

111 7±366 

12:1: 4 

377± 72 

10± 3 

19: 5 

22:!: 6 

16% 

(11%-

12% 4 

12± 7 

<7: -

(15% -

(12% -

7:!: 5 

74.9 ±0.99 

14.3 ±-

0.46± .11 

0.49± • 11 

0.06± .01 

o. 62± • 08 

5.27± .04 

3.68± .33 

0.03± .01 

0.04± -

0.11± .05 

0.06± .02 

28.67 

0.74 

21. 77 

44.63 

2.42 

0 

N = 5 

o. 15 

• 68 

.67 

0 

.06 

.10 

.12 

80± 21 

361±153 

(5:1: -

288:1: 23 

7:1: 3 

32: 3 

28:!: 4 

10% 3 

<lli: -

(10: -

(8:1: -

<13% -

15: 2 

<12± -

4± 2 

75.5 ±0.40 

13.5 :1: .25 

0.60± .06 

0.20± .16 

0.04± .04 

0. 71± • I 0 

4.44±1.03 

4.82±1.33 

0.03± .01 

0.03± -

0.03± .02 

0.01± .03 

29.89 

0 

28.46 

37.60 

2. 78 

0.21 

.10 

0 

.66 

.15 

.06 

.06 

.02 

N = 6 

5± 3 

188±182 

(5:1: -

31% 9 

40± 30 

46: 16 

43: 11 

5± 5 

<I li: -

<IO± -

681±84 7 

<13% -

15:!: 6 

(12% -

(li: -

1Average of 54 analyses for 13 peraluminous plutons in the eastern and southeastern Arabian Shield. 

Madha(9) 

granodiorite 

N = 2 

74.2 ±0.57 

I 5. I :1: • 21 

0.38± .08 

0.35± .07 

0.23± .01 

1.08± .01 

5.29± .04 

3.08± .03 

0.07± -

0.08:1: .02 

0.07± .03 

0.03± -

29.14 

1. 35 

18.21 

44.79 

4.67 

0 

0.57 

.34 

.55 

0 

.13 

.19 

.05 

N = 5 

11% 22 

265:157 

lli: II 

31: 10 

316±200 

7± 2 

47± 18 

(1:1: -

(21: -

(10: -

1170±503 

16± I 

21:1: 3 

(12% -

(1% -

2Average of 692 samples of granitoid rocks, eastern and southeastern Arabian Shield (du Bray and others, 1982). 

24 Contributions to Mineral Resources Research, 1984 
Arabian Shield 

and standard deviation(±) are based; leader in standard deviation entry(-), zero variance. Pluton element averages normalized to 100 percent, 

sorption and wet chemistry methods. Trace element analyses: Li by atomic absorption; F by selected-ion electrode; Cu and Be by semiquantitative 

using a 241Am excitation source. NO, no data available; a.f.g., alkali-feldspar granite] 

Zayd(IO) Habbah(lf) Bwana(l2) 

granodiorite granodiorite a.f.g. 

Gaharra(l3) 

a.f.g. 

Average 1 Eastern 

Shield2 

Major-element analyses--Continued 

N = 4 N = 4 N = 4 N = 5 N = 54 

74.6 :!:1.05 

14.6 :!:0.41 

0.36:!: .os 

0.23:!: • 10 

0.08:!: .03 

I. 76:!: • 26 

4.49:!: .24 

3. 74:!: .65 

0.05:!: .01 

0.03:!: .01 

0.02:!: .02 

0.03:!: .01 

30.48 

o. 10 

22.00 

37.96 

8.34 

0 

.20 

.OR 

.53 

0 

.10 

.07 

.06 

7 3. 7 :!: I. 09 

15.2 :!:0.52 

0.34:!: .23 

0.42:!: .24 

0. 14:1: .07 

·1. 91:!: • 54 

4.74:!: .76 

3.34±1.60 

0.06:!: .03 

0.05:!: .04 

0.06:!: .03 

0.02:!: -

28.98 

0.47 

19.76 

40.13 

9.05 

0 

0 

.35 

• 51 

.so 

• 12 

• 12 

.03 

75.3 :!:1.29 

14.2 ±l.OR 

0.31±0.011 

0.43± .OS 

0.07± .04 

0.39:!: .29 

5.31± .90 

3.118± .54 

0.02± .01 

0.02± .02 

0.04± .01 

0. 13± .OS 

28.71 

0.9I 

23.01 

44.95 

.8 7 

0 

• 18 

• 58 

.45 

0 

.04 

.05 

.27 

75.4 :!:0.51 

14. 00± • 37 

0.28± .09 

0.41± • 11 

0.03± .01 

0.32± .OS 

5.28± .54 

3.96± .64 

0.02± .01 

0.01± .01 

0.06± .01 

0.42± • 14 

75.2 :!:0. 78 

14.0 :!: .80 

0.48± • 16 

0.48± .23 

o. 11± .08 

0. 79± • 5 I 

4.65± .so 

4.06± .47 

0.05± .OS 

0.04± .03 

0.05± .03 

O. I 5± • 12 

CIPW norms--Continued 

29. 19 

1.02 

23.38 

44.71 

0 

0 

0.08 

.60 

• 4I 

0 

.04 

.02 

.43 

31. 13 

0.99 

24.01 

39.31 

2.60 

0 

0.28 

.50 

.70 

0 

• 10 

.10 

.30 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

Trace-element analyses--Continued 

N = 9 N = 8 N = 16 N = 48 N = 692 

6:!: 3 

161:!:101 

(5:!: -

40:!: 19 

403:!:113 

4:!: -

39:!: 12 

(1:!: -

(11:1: -

<7:!: -

I 170:1:561 

(13:!: -

(15:!: -

(12:!:-

(]:!: -

7:!: 5 

199:!:142 

(5:!: -

36± 12 

446:!:130 

6± 4 

52:!: 35 

(1:!: -

<21:!: -

<7± -

1009±318 

(13:!:-

(15:!: -

(12:!: -

(1:!: -

231±304 

1614±815 

(5:!: -

486±273 

11:!: 17 

90± 57 

89:!: 24 

38:!: 20 

65:!: 8 

35:!: 26 

(8± -

(13:!:-

38± 17 

<25± -

8± 6 

3Tischendorf (1977). 

4Krauskopf (1967, p. 639-640). 

306:!: 91 

3060±1054 

13:!: 9 

603± Ill 

7:1: 6 

97:!: 64 

56:!: 15 

38± 13 

49:!: 19 

40± 26 

(8± -

(13± -

37:!: 13 

<25:!: -

5:!: 2 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

ND 

59±330 

836±962 

31± 36 

158±172 

212±247 

43± 39 

144±140 

23± 15 

28± 30 

(10:!: -

555±492 

39± 28 

ND 

ND 

3:!: 3 

Specialized 

granites3 

N = 962 

74.3 ±1.39 

14.2 ±1.07 

0.81±0.47 

l. 11± .47 

0.48± .56 

o. 76± • 41. 

3.24± .6I 

4. 7 5± • 68 

O.I6± .10 

ND 

0.05± .04 

0.32± .38 

35.42 

3. I6 

28.06 

27.42 

I. 40 

0 

1. 19 

1. 20 

I. 18 

0 

0.31 

0 

.67 

N = 962 

400± 200 

3700±1500 

ND 

580± 200 

ND 

ND 

ND 

ND 

ND 

40:!: 20 

ND 

ND 

ND 

ND 

13:!: 6 

Precursor 

granites3 

N = 226 

73.8 :!:1. I8 

I4.3 ±0.53 

0.49± .I7 

I. 37:!: • 47 

0.46± .36 

I. 06± • 34 

3.27± .so 

4.91± .45 

0.24± .II 

ND 

0.05± .26 

o. 11± .04 

32.57 

1.98 

29.03 

27.66 

4.42 

0 

1. 14 

1. 81 

o. 71 

0 

.46 

0 

.23 

N = 226 

130:!: 50 

810±200 

ND 

250:!: 50 

ND 

ND 

ND 

ND 

ND 

10:!: 5 

ND 

ND 

ND 

ND 

7.5±2 

Granites4 

N = 2,327 

71.4 ±1.41 

14.5 ±0.23 

l. 32± • 29 

1. 79± • 38 

0.82± .23 

I. 91± .40 

3.47± .32 

4.38± .52 

0.34± .08 

ND 

0.06± .03 

0.09 ND 

28.69 

0.79 

25.86 

29.35 

8.79 

0 

2.03 

1. 75 

l. 92 

0 

.65 

0 

.19 

N 2,327 

30 

850 

10 

150 

285 

40 

180 

20 

20 

3 

600 

40 

87 

35 

5 

Peraluminous Granitoid Plutons, Saudi Arabia 25 
of the Madha assemblage plutons, in particular the compositions of plutons 9-11, are distinctly different from 

those of other peraluminous plutons (fig. 7). These 

plutons are depleted in all the lithophile elements except 

barium and strontium. Strontium concentrations are 

slightly higher than those cited for average granites, 

whereas barium is enriched almost twofold relative to 

world averages for granite (Krauskopf, 1967, p. 639-640). 

The other nine plutons are depleted in Ba, Cu, La, Sr, 

and Zr relative to average granites, but they are enriched 

in combinations of Be, Pb, Sn, Li, F, Rb, Y, and Nb and 

have Rb/Sr ratios that range from 9 to about 85. In addition, some samples contain unusually high concentrations of Bi, W, Ag, Mo, or Zn. The trace-element 

composition of the Tafshah monzogranite is transitional 

between the composition of the two groups of plutons. 

It is important to note that the various plutons are enriched in incompatible elements to substantially different 

degrees. 

Distinctive trace-element covariation exists among 

the plutons studied (fig. 7). In particular, zirconium and 

yttrium concentrations increase in a nearly 1: 1 ratio (fig. 

7A). Yttrium and niobium concentrations display similar 

covariation (fig. 7C), except that the yttrium concentration increases at more than twice the rate of the niobium 

concentration. The plot of the agpaitic index versus zirconium content (fig. 7D) demonstrates that there is no 

relation between agpaitic ratio and zirconium content in 

nonperalkaline granites, which is consistent with the findings of Watson (1979) that the zirconium concentration 

of nonperalkaline granites is buffered at about 100 ppm. 

Variation of zirconium concentration among the 

peraluminous plutons may be related to the degree of 

magmatic evolution, the composition of the source terrane, or a combination of these factors. 

Rare-earth element (REE) contents of non-Madha 

assemblage plutons distinguish them from Madha 

assemblage plutons and are typical of highly evolved intrusive rocks. Rare-earth element patterns for plutons 

1-8, 12, and 13 are flat, have large negative europium 

anomalies, and are LREE (light REE) depleted; total REE 

content is high (J .S. Pallister and 1 .S. Stuckless, written 

commun., 1983). REE patterns for Madha assemblage 

plutons are positively sloping and are characterized by 

moderate to small Eu anomalies; total REE content is 

significantly less than in the other plutons. 

Isotopes 

Stacey and Stoeser (1983) have determined in a 

study of feldspar and whole-rock common lead compositions that the plutons in the northeastern and extreme 

eastern and southeastern parts of the Arabian Shield have 

elevated 208Pb/204Pb and 207Pbf204Pb ratios relative to 

their 206Pb/204Pb ratios and average radiogenic lead 

26 Contributions to Mineral Resources Research, 1984 

Figure 7. Trace-element variation diagrams for the 

peraluminous plutons. A, zirconium versus yttrium; 8, K/Rb 

ratio versus Rb/Sr ratio; C, yttrium versus niobium; 0, zirconium 

versus the agpaitic index ((Na2

0+ K20)/AI20 3) on a molecular 

basis. Pluton numbers 1-13 keyed to figure 1 (map locations) 

and table 1 (petrographic characteristics). Data are average 

trace-element values from table 3. Open circle, average composition of 2,327 granites (Krauskopf, 1967); circle with dot, 

average composition of 692 granitoid samples from eastern and 

southeastern Arabian Shield (du Bray and others, 1982). 
growth curves for the mantle and oceanic crust (Doe and 

Zartman, I979). These plutons are located east of the 

HNFZ (fig. I). The Sabhah alkali-feldspar granite is 

characterized by anomalously radiogenic common lead 

(2°8Pbf204Pb = 37.339, 207Pb/204Pb = I5.5II, 206Pb/204Pb 

= I7. 726), whereas common lead in the Madha granodiorite (208Pb/204Pb = 37.045, 207Pb/204Pb = I5.470, 206Pb/ 204Pb = I7 .697) and in the Gaharra alkali-feldspar 

granite (2°8Pb/204Pb = 37 .I82, 207Pb/204Pb = I5 .480, 206Pb/204Pb = I7 .89I) is significantly less radiogenic. 

MINERAL DEPOSIT POTENTIAL 

Peraluminous plutons are potential hosts to 

deposits of tin and tungsten. For example, the lode tin 

and tungsten deposits of southwest England (Jackson, 

I979) and the alluvial tin deposits of Thailand and 

Malaysia (Ishihara and others, I980) are characteristically associated with muscovite-bearing peraluminous 

granitoid rocks. A quartz-vein stockwork associated 

with a small, greisenized peraluminous stock and containing O.IO percent wo3 and O.OI percent tin was 

recently located near Baid al Jimalah (Cole and others, 

I98I; Lofts, I982). Identification of this deposit suggests that other peraluminous plutons in the eastern 

and southeastern Arabian Shield might contain ore 

deposits. 

The average compositions of the peraluminous 

plutons, as indicated on the ACF ternary diagram 

(fig. 5), suggest that only three or four of the plutons 

areS-type as defined by White and Chappell (I977); yet 

preliminary geochemical data (du Bray and others, I982) 

suggest that tin deposits may exist associated with most 

of these plutons. Takahashi and others (1980) have also 

noted that some tin deposits are associated with granites 

whose compositions plot in the 1-type field on the ACF 

diagram. The compositional distinction made on the ACF 

diagram (White and Chappell, I977) between S-and 1-

type granitoid rocks-and therefore on a pluton's 

favorability for associated tin deposits-needs reconsideration. Perhaps, as indicated by Ishihara and others 

(I980), identification of ilmenite-series granitoid rocks 

will expedite discovery of tin deposits. 

Pan concentrates of wadi sediment collected from 

wadis draining the peraluminous plutons in the eastern 

and southeastern Arabian Shield contained I ,000 ppm 

tin (du Bray and others, I982), or up to IOO times the 

tin concentration in associated granites. Because the mica 

minerals are removed from the samples during panning, 

tin must not be concentrated in micas but in a heavy 

mineral such as cassiterite. Tungsten was also detected 

in several of these samples but did not exceed concentrations of IOO ppm. In addition, some of these samples were 

enriched in combinations of the elements Ag, Bi, Zn, Pb, 

and Mo relative to average values determined for felsic 

plutonic rocks in the eastern and southeastern Arabian 

Shield. 

Despite the geochemical anomalies in tin and other 

elements, detailed ground traverses failed to identify any 

mineralized zones. Quartz veins and host granites were 

thoroughly sampled, but analyses indicate that, although 

concentrations of Sn, Mo, Bi, and Ag are well above 

regional averages (3, 2, <I, and <I ppm, respectively), 

surface outcrops are not strongly mineralized (du Bray, 

I983b). A small wolframite-bearing quartz vein was 

found cutting the Tarban alkali-feldspar granite, and a 

single sample of a quartz vein cutting the Sitarah alkalifeldspar granite contained I ,000 ppm molybdenum, but 

these are exceptional occurrences. Plutons 3, 5, I2, and 

I3 probably warrant further study because the concentrations of tin and other trace elements in average 

specimens were so high (table 4). 

The failure to locate either mineralized quartz-vein 

stockworks or greisens, despite the appropriateness of the 

geologic environment, should be discussed. Formation 

of a tin deposit requires a manyfold concentration of tin 

in a late-stage magmatic fluid, possibly through volatile 

complexing, and its subsequent precipitation in a stanniferous phase that is included either in quartz veins or 

in a greisen. A fluid phase must be present in order for 

tin to be concentrated in late-stage processes. The occurrence of primary white mica and of two independently 

nucleated alkali feldspars suggests that most plutons 

studied here were emplaced at relatively great depths. The 

depths may preclude early evolution of a fluid phase during crystallization, so that tin was partitioned not into 

a fluid phase but into the melt. To a limited extent, tin 

may have been incorporated in micas and oxide minerals, 

but the tin-saturation maximum was probably reached 

in melts of various high-potential plutons. Tin saturation 

may have resulted in nucleation of a tin-bearing phase, 

probably cassiterite, that was then widely disseminated 

throughout the plutons. When a fluid phase ultimately 

did exsolve, tin was no longer available for concentration, and relatively barren quartz veins were deposited 

in fractures. 

Deposits of tin, tungsten, molybdenum, and other 

rare metals are generally found in the apical part of 

plutons emplaced at fairly shallow levels (Groves 

and McCarthy, I978). Regional studies of the Arabian 

Shield (D.B. Stoeser, oral commun., I982; J.W. Whitney, 

oral commun., I982) suggest that erosion has extended 

much deeper in the southeastern Arabian Shield than 

in the northeastern shield and that the tops of highmineral-potential plutons have been removed by erosion. 

Consequently, mineral potential associated with felsic 

plutonic rocks is greater in plutons of the northeastern 

region. The discoveries in the northeastern part of the 

Arabian Shield of both a tungsten and tin quartz-vein 

stockwork associated with the peraluminous granite 

at Baid al Jimalah (Cole and others, I98I) and a 

Peraluminous Granitoid Plutons, Saudi Arabia 27 
tin-anomalous muscovite-topaz monzogranite with 

associated tin greisens in a subcaldera ring fracture (du 

Bray, 1982c) corroborate this finding. 

PETROGENESIS 

Some constraints can be placed on the genesis and 

crystallization of the peraluminous plutons by comparing their chemistry and petrology with experimental data. 

The highly differentiated major-element chemistry, incompatible trace-element-enrichment patterns, and radiogenic lead-enriched compositions of these plutons all 

suggest that they were generated by partial melting of 

preexisting continental material. The compositions of the 

peraluminous plutons are compared on the Q-Ab-Or ternary diagram (fig.()), which also shows minimum melting 

compositions at several pressures (polybaric) in the experimental system Si02-NaA1Si30 8

-KA1Si30 8-H20 

(Winkler and others, 1975). The average composition of 

the plutons is coincident with the composition of the 

200 MPa (megapascal) (or 2 kbar (kilobar)) minimum 

melt for conditions of PH2

o =Ptotal· Steiner and others 

(1975) indicated that determinations of intensive 

parameters such as temperature and pressure of crystallization based on comparisons of this sort may be tenuous 

and are based on certain petrogenetic assumptions. 

However, the convergence of chemical and petrographic 

data for these peraluminous plutons suggests that the full 

compositional range represented by averages for the 

peraluminous plutons of the eastern and southeastern 

Arabian Shield could be explained by crystallization 

under conditions of PH2

o ranging from 200 to 400 MPa 

(2 to 4 kbar), equivalent to between 7 and 14 km of overburden. Figure 6 shows that the compositions of the 

plutons depart a little from the trend line of polybaric 

minimum melting compositions (minima trend) of the experimental system, but the differences are within the 

analytical precision of the method. In all cases, onestandard deviation bars (omitted for clarity) intersect the 

minima trend. 

The 200- to 400-MPa PH2

o range also meets the 

requirements necessary for crystallization of muscovite 

from a granitic melt and for crystallization of two independently nucleated (nonperthitic) alkali feldspars. 

Muscovite crystallizes from a eranite magma when PH2

o 

is greater than or equal to 300 MPa (3 kbar) at about 670° 

C (Yoder and Eugst~r, 1954). Two discrete alkali 

feldspars crystalliz({::·"in the experimental system 

NaA1Si30 8-KA1Si30 8-Si02-H20 in favor of a single alkali 

feldspar when PH2o is greater than 360 MPa (3.6 kbar) 

{Tuttle and Bowen, 1958; Morse, 1970). 

The 200- to 400-MPa pressure range must be 

regarded as a maximum, however. Manning (1981) ha~ 

demonstrated that the minimum melt composition in the 

28 Contributions to Mineral Resources Research, 1984 

granite system is displaced toward compositions richer 

in Ab by the presence of several percent fluorine. The 

high fluorine content of most plutons studied suggests 

that the magmas from which they crystallized may have 

contained up to several percent fluor_ine and that their 

eutectic compositions were affected accordingly. The 

pressure necessary for nucleation of two alkali feldspars 

may be similarly reduced by the presence of fluorine. 

Thermodynamic considerations regarding the pressuredependent stability of muscovite in granitic melts indicate 

that these micas can be stable at anomalously low 

pressures. Anderson and Rowley (1981) suggested that 

muscovite is stable at magmatic temperatures and 

pressures of 200 MPa because of its non-ideal composition. No experimental data are available for the stability 

of zinnwaldite, but its occurrence in several of the plutons 

studied may not necessitate great pressure. 

The presence of garnet in seven of the plutons probably reflects bulk composition variations between plutons 

rather than high pressure. Hsu (1968) and Green (1978) 

indicated that high manganese concentrations enhance 

garnet stability in magmas at pressures as low as 300 MPa 

(3 kbar). This fact is in good agreement with the correlation between high manganese concentrations and 

garnetiferous plutons in the study area (tables 1, 4). Miller 

and Stoddard (1981) suggested that an elevated manganese concentration in melts may be the critical factor 

relative to garnet nucleation and growth. Weak zoning 

preserved in garnets from the Tafshah monzogranite may 

indicate either that manganese was concentrated during 

crystallization or that garnet composition in equilibrium 

with the melt became richer in manganese during 

magmatic ascent. 

Preliminary common lead isotope data for the 

Sabhah alkali-feldspar granite indicate that its genesis involved partial melting that included a component of 

preexisting lower continental crust. Stacey and Stoeser 

(1983) suggested that the common lead composition of 

the Sabhah alkali-feldspar granite lies on a mixing line 

between the lead composition field of more radiogenic 

lower continental crust and the lead composition field of 

less radiogenic oceanic crust, as defined by Doe and Zartman (1979). Common lead data for the Gaharra alkalifeldspar granite and the Madha granodiorite are consistent with their derivation from oceanic crust, as indicated 

by Doe and Zartman (1979); however, the lead isotope 

data are at odds with other petrologic and geochemical 

evidence regarding the genesis of the Gaharra alkalifeldspar granite. In particular, the chemistry, modal 

mineralogy, and setting of this pluton are very similar 

to those of the other peraluminous plutons that crop out 

east of the HNFZ in the eastern and southeastern Arabian 

Shield. In addition, the Gaharra alkali-feldspar granite 

crops out near Najran (fig. 1), where a quartz monzonite 

crops out that has a lead isotopic composition consistent 
with derivation from or incorporation of older, evolved 

material (Stacey and Stoeser, 1983). The complexities of 

the common lead composition of the Gaharra alkalifeldspar granite are beyond the scope of this study, but 

it is important to note that most of the peraluminous 

'plutons crop out near and east of the HNFZ in terrane 

in which common lead compositions indicate contribution 

from older continental material (Stacey and Stoeser, 

1983). The isotopic composition of lead in the Madha 

granodiorite and probably in the other plutons of the 

Madha assemblage may be attributable to their position 

west of the HNFZ in a terrane in which common lead 

compositions suggest derivation from oceanic crust or 

depleted granulite facies material. 

The peraluminous plutons, except those of the 

Madha assemblage, are enriched in the incompatible 

elements relative to global averages for granites and 

depleted in the compatible elements; however, the degree 

of enrichment and depletion varies from one pluton to 

the next (table 4; fig. 7). Enrichment was probably not 

achieved by assimilation and contamination during 

magma ascent because the peraluminous plutons neither 

contain xenoliths nor preserve assimilation textures. Only 

the Madha assemblage plutons (nonenriched) display 

evidence of wall-rock assimilation. 

The Madha assemblage plutons are best considered 

separately because their tectonic position and chemical 

data indicate that, although peraluminous, they are distinctly different from the other plutons. The Tafshah 

monzogranite is problematic because it shares characteristics both of Madha assemblage plutons and of the other 

peraluminous plutons. The factors that controlled genesis 

of this pluton were probably transitional between the two 

groups, but in light of the position of the pluton west 

of the HNFZ-and its trace-element composition-it 

seems most appropriate to consider genesis of the Tafshah 

monzogranite with the other three Madha assemblage 

plutons. 

The non-Madha assemblage plutons were probably 

derived by partial melting of old continental material, a 

fact that precludes derivation of the Madha plutons by 

advanced progressive melting of the same source terrane. 

The non-Madha assemblage plutons could not be derived 

by partial melting of noncontinental material, except for 

very small amounts, because the extreme incompatibleelement enrichment and the anomalously radiogenic common lead compositions characteristic of the non-Madha 

plutons could not be obtained. The non-Madha assemblage magmas could have been obtained by extreme fractionation of less evolved melts, which could yield 

incompatible-element-enriched melts, but the cumulate 

rocks that would also result are not in evidence. To account for the incompatible-element depletion observed 

for the Madha assemblage plutons, an unreasonably large 

extent of melting of common continental source material 

would be required. 

The barium content of these rocks also suggests that 

Madha assemblage and non-Madha assemblage plutons 

were derived from very different sources. The principal 

mineralogic site of barium is potassium feldspar, whose 

mineral-melt distribution coefficient for barium favors 

concentration of barium in residual potassium feldspar 

(Hanson, 1978). The high barium content of the Madha 

assemblage rocks probably resulted from advanced partial melting of a source terrane that contained potassium 

feldspar, whereas the pronounced barium depletion 

characteristic of the other plutons probably resulted from 

a small amount of partial melting of material, such as 

graywacke, in which potassium feldspar was a minor 

constituent. 

Trace-element plots-particularly the K/Rb ratio 

versus Rb/Sr ratio plot (fig. 7 B)-emphasize the 

dissimilarity that exists between the Madha assemblage 

plutons and the others. The compositions of the Madha 

assemblage plutons plot in fields (fig. 7) that are widely 

separated from fields and trends defined by the other 

plutons. The K/Rb ratio versus Rb/Sr ratio plot provides 

information concerning melting and crystallization events 

that might have produced the muscovite-bearing plutons. 

A trend involving an increasing K/Rb ratio and a decreasing Rb/Sr ratio would suggest that the plutons could have 

been generated by varying amounts of progressive melting 

of a common source. As partial melting proceeds, 

potassium and strontium concentrations should increase 

relative to rubidium concentration. Instead, the K/Rb 

ratio is nearly constant for non-Madha assemblage 

plutons, and the Rb/Sr ratio is almost independent of 

the K/Rb ratio. These relationships suggest that nonMadha assemblage plutons are crystallized partial melts 

from different source terranes. 

Variation in degrees of incompatible-element 

enrichment between the plutons substantiates the relations 

indicated by the K/Rb versus Rb/Sr ratio plot. Concentrations of the incompatible elements Zr, Nb, and Y 

should be highest in those magmas generated by the 

smallest amount of partial melting (and subjected to the 

greatest amount of fractional crystallization), and as 

melting progresses the relative proportion-that is, 

concentration-of these elements should decrease. If the 

same source terrane is presumed for all plutons, then 

plutons 3, 6, 7, 12, and 13 (fig. 7B) represent the greatest 

degree of chemical evolution; however, plutons 1, 2, 4, 

and 5 (representing less evolved magma) contain greater 

concentrations of Zr, Nb, andY (fig. 7A, C). The extent 

of incompatible-element enrichment is not a function of 

the amount of progressive melting (or subsequent fractional crystallization) of a common source. Figure 7 B 

and, less convincingly, figure 7A, C, and D suggest that 

partial melting in four geochemically distinct source 

Peraluminous Granitoid Plutons, Saudi Arabia 29 
terranes could account for the observed trace-element 

variation. The Madha assemblage plutons (8-11) were 

probably generated in one source terrane, plutons 1, 2, 

4, and 5 in a second, plutons 3, 6, 7, and 12 in a third, 

and pluton 13 in a fourth. 

Crystal fractionation processes may account for 

within-source terrane group variation and enhancement 

of incompatible-element enrichment. Concentrations of 

yttrium and zirconium in plutons numbered 1, 2, 4, and 

5 and those numbered 3, 6, 7, and 12 are consistent with 

plagioclase fractionation (Walsh and others, 1979). 

Large negative europium anomalies and flat 

chondrite-normalized REB patterns are characteristic of 

high silica intrusions (Miller and Mittlefehldt, 1982). 

LREE depletion in felsic magmas has been explained 

(Miller and Mittlefehldt, 1982) as resulting from crystallization and removal of very small grains of an LREEenriched phase such as monazite or allanite. Alternatively, 

these minerals may reside in the residue upon completion of partial melting. Similar processes may account for 

LREE depletion of some peraluminous plutons in the 

eastern and southeastern Arabian Shield. The large 

negative europium anomalies characteristic of nonMadha assemblage plutons are consistent with plagioclase 

fractionation. The high total REB content of the nonMadha assemblage plutons argues in favor of an evolved 

source for these rocks. REE data for the Madha 

assemblage plutons suggest an unevolved source. 

CONCLUSIONS 

Thirteen plutons of peraluminous monzogranite 

and granodiorite, similar to S-type, metallogenically 

specialized, ilmenite-series, and peraluminous granitoid 

rocks described elsewhere in the world, were identified 

in the eastern and southeastern Arabian Shield. These 

medium-grained, inequigranular rocks contain subequal 

amounts of quartz, microcline, and albite. Biotite or FeLi mica is the principal mafic phase, muscovite occurs 

in six of the plutons, and trace amounts of garnet were 

identified in seven of the plutons. Geochemical studies 

indicate that at least four of the plutons could host 

associated tin-tungsten deposits. 

Major-element and petrographic data for the peraluminous granitoids of the eastern and southeastern 

Arabian Shield indicate that these rocks were generated 

by partial melting of older continental material and subsequently crystallized under conditions of PH2

o = 200-400 

MPa (2-4 kbar). The trace-element and common lead 

isotope data indicate that most of the peraluminous 

plutons were produced by partial melting involving a 

component of old continental crust. The Madha 

assemblage plutons, however, must have been produced 

as partial melts of geochemically primitive material. 

Magma genesis may have been initiated in response to 

30 Contributions to Mineral Resources Research, 1984 

a collisional event that mobilized the crust adjacent to 

the Hulayfah-Ad Dafinah and Nabitah-Hamdah fault 

zones, the supposed suture zone. Compositional variations in the peraluminous plutons across the fault zone 

indicate that the two colliding plates were compositionally 

distinct; the eastern plate was evolved continental crust, 

whereas the western plate was geochemically primitive, 

oceanic crust or depleted granulite facies material. The 

distinctive composition of the plutons suggests that significant regional or temporal differences related to 

magma genesis, including compositional variation and 

variation of physical conditions within the source crust, 

may have existed within the developing Arabian Shield. 

The peraluminous plutons of the eastern and southeastern 

Arabian Shield owe their gross petrologic similarities not 

to a common source terrane but to a set of relatively 

similar physical conditions that prevailed during their individual genesis. 

REFERENCES CITED 

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. 

Chappell, B.W., and White, A.J.R., 1974, Two contrasting 

granite types: Pacific Geology, v. 8, p. 173-174. 

Clifford, T.N., 1970, The structural framework of Arabia, in 

Clifford, T.N., and Gass, I.G., eds., African magmatism 

and tectonics: Darien, Conn., Hafner, p. 1-26. 

Cole, J.C., Smith, C.W., and Fenton, M.D., 1981, Preliminary 

investigation of the Baid al Jimalah tungsten deposit, 

Kingdom of Saudi Arabia: U.S. Geological Survey Saudi 

Arabian Mission Technical Record 20 (Interagency Report 

377), 26 p. 

Cox, K.G., Bell, J.D., and Pankhurst, R.J., 1979, The interpretation of igneous rocks: London, Allen and Unwin, 450 p. 

Deer, W.A., Howie, R.A., and Zussman, J., 1962, Rockforming minerals: London, Longmans, v. 3, 270 p. 

___ 1966, An introduction to the rock-forming minerals: 

London, Longmans, 528 p. 

Delfour, Jacques, 1979, Geologic map of the Halaban 

quadrangle, sheet 23G, Kingdom of Saudi Arabia (with 

topographic base): Saudi Arabian Directorate General of 

Mineral Resources Geologic Map GM-46-A, 32 p., scale 

1:250,000. 

___ 1980, Geologic map of the Wadi ar Rika quadrangle, 

sheet 22G, Kingdom of Saudi Arabia: Saudi Arabian Directorate General of Mineral Resources Geologic Map 

GM-51-A, 34 p., scale 1:250,000. 

Doe, B.R., and Zartman, R.E., 1979, Plumbotectonics I, the 

Phanerozoic, in Barnes, H.L., ed., Geochemistry of hydrothermal ore deposits: New York, John Wiley, p. 22-70. 

du Bray, E.A., 1983a, Petrology of muscovite-bearing granitoid 

plutons in the eastern and southeastern Arabian Shield, 

Kingdom of Saudi Arabia: Saudi Arabian Deputy Ministry 

for Mineral Resources Open-File Report USGS-OF-03-10, 

36 p. 
___ 1983b, Mineral potential of selected felsic plutons in the 

eastern and southeastern Arabian Shield, Kingdom of 

Saudi Arabia: Saudi Arabian Deputy Ministry for Mineral 

Resources Open-File Report USGS-OF-03-39, ~8 p. 

___ 1983c, Reconnaissance geology of the Jabal as Silsilah 

quadrangle, sheet 26/42 D, Kingdom of Saudi Arabia: 

Saudi Arabian Deputy Ministry for Mineral Resources 

Technical Record USGS-TR-03-4, 52 p., scale 1:100,000. 

du Bray, E.A., Elliott, J.E., and Stoeser, D.B., 1982, Geochemical evaluation of felsic plutonic rocks in the eastern 

and southeastern Arabian Shield, Kingdom of Saudi 

Arabia: Saudi Arabian Deputy Ministry for Mineral 

Resources Technical Record USGS-TR-02-2, 53 p. 

Drysdall, A.R., 1979, Petrogenesis of two peralkaline microgranites from N. W. Hijaz; some considerations relevant 

to further prospecting: Saudi Arabian Directorate General of Mineral Resources Open-File Report DGMR-705, 

51 p. 

Elliott, J.E., 1980, Tin-bearing granite of Jabal al Gaharra in 

the southern Arabian Shield, Kingdom of Saudi Arabia: 

U.S. Geological Survey Saudi Arabian Mission TechniCal 

Record 4 (Interagency Report 299), 28 p. 

___ 1983, Peralkaline and peraluminous granites and related 

mineral deposits of the Arabian Shield, Kingdom of Saudi 

Arabia: Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report USGS-OF-03-56, 37 p. 

Green, T.H., 1978, Garnet in silicic liquids and its possible use 

as a P-T indicator: Contributions to Mineralogy and Petrology, v. 65, p. 59-67. 

Greenwood, W .R., Anderson, R.E., Fleck, R.J ., and Roberts, 

R.J., 1980, Precambrian geologic history and plate tectonic evolution of the Arabian Shield: Saudi Arabian 

Directorate General of Mineral Resources Bulletin 24, 

35 p.; also 1977, U.S. Geological Survey Saudi Arabian 

Project Report 222, 97 p.; both include colored geologic 

map, scale 1 :4,000,000. , 

Greenwood, W.R., Stoeser, D.B., Fleck, R.J., and Stacey, J.S., 

1982, Late Proterozoic Island-arc complexes and tectonic 

belts in the southern part of the Arabian Shield, Kingdom 

of Saudi Arabia: Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report USGS-OF-02-8, 46 p.; 

also 1983, U.S. Geological Survey Open-File Report 

83-296, 49 p. 

Groves, D.I., and McCarthy, T.S., 1978, Fractional crystallization and the origin of tin deposits in granitoids: Mineralium 

Deposita, v. 13, p. 11-26. · 

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. 

Hsu, L.C., 1968, Selected phase relationships in the system AlMn-Fe-Si-0-H, a model for garnet equilibria: Journal of 

Petrology, v. 9, p. 40-83. 

Ishihara, Shunso, 1977, The magnetite-series and ilmenite-series 

granitic rocks: Mining Geology, v. 27, p. 293-305. 

Ishihara, Shunso, Sawata, Hideho, Shibata, Ken, Terashima, 

Shigera, Arrykul, Surapon, and Kohei, Sato, 1980, 

Granites and Sn-W deposits of peninsular Thailand: Mining Geology Special Issue 8, p. 223-241. 

Jackson, N.J., 1979, Geology of the Cornubian tin field: A 

review: Geological Society of Malaysia Bulletin 11, 

p. 209-237. 

Kanaan, F.M., 1979, The geology, petrology, and geochemistry 

of the granitic rocks of Jabal al Hawshah and vicinity, Jabal 

al Hawshah quadrangle, Kingdom of Saudi Arabia: Saudi 

Arabian Directorate General of Mineral Resources Bulletin 

23, 114 p. 

Krauskopf, K.B., 1967, Introduction to geochemistry: New York, 

McGraw-Hill, 721 p. 

Lofts, P.G., 1982, A preliminary evaluation of the Baid al Jimalah 

tungsten prospect: Saudi Arabian Deputy Ministry for 

Mineral Resources Open-File Report RF-OF-02-21, 223 p. 

Manning, D.A.C., 1981, The effect of fluorine on liquidus phase 

relationships in the system Qz-Ab-Or with excess water at 

1 kb: Contributions to Mineralogy and Petrology, v. 76, 

p. 206-215. 

Miller, C.F., and Bradfish, L.J., 1980, An inner Cordilleran belt 

of muscovite-bearing plutons: Geology, v. 8, p. 412-416. 

Miller, C.F., and Mittlefehldt, D.W., 1982, Depletion of light rareearth elements in felsic magmas: Geology, v. 10, p. 129-133. 

Miller, C.F., and Stoddard, E.F., 1981, The role of manganese 

in the paragenesis of magmatic garnet-An example from 

the Old Woman-Piute Range, California: Journal of 

Geology, v. 89, p. 233-246. 

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. 

Morse, S.A., 1970, Alkali feldspars with water at 5 kb pressure: 

Journal of Petrology, v. 11, p. 221-251. 

Sable, E.G., 1982, Reconnaissance geology of part of the Najran 

quadrangle, sheet 17 I 44 C, Kingdom of Saudi Arabia: Saudi 

Arabian Deputy Ministry for Mineral Resources Open-File 

Report USGS-OF-03-7, 25 p., scale 1:100,000. 

Schmidt, D.L., and Brown, G.F., 1982, Major-element chemical 

evolution of the late Proterozoic shield of Saudi Arabia: 

Saudi Arabian Deputy Ministry for Mineral Resources 

Open-File Report USGS-OF-02-88, 34 p.; also in press, Jiddah, King Abdulaziz University, Faculty of Earth Sciences, 

First Symposium of Pan-African Project 164, Proceedings. 

Schmidt, D.L., Hadley, D.G., and Stoeser, D.B., 1979, Late Proterozoic crustal history of the Arabian Shield, southern 

Najd province, Kingdom of Saudi Arabia: U.S. Geological 

Survey Saudi Arabian Project Report 251, 44 p.; also 1979, 

in Thhoun, S.A., ed., Evolution and mineralization of the 

Arabian-Nubian Shield: Oxford-New York, Pergamon 

Press, Jiddah, King Abdulaziz University, Institute of Applied Geology Bulletin 3, v. 2, p. 41-58. 

Shand, S.J., 1949, Eruptive rocks, their genesis, composition, classification, and their relation to ore deposits, with a chapter 

on Meteorites, 3rd edition: London, Thomas Murby, 488 p. 

Stacey, J.S., and Stoeser, D.B., 1983, Distribution of oceanic and 

continental leads in the Arabian-Nubian Shield: Saudi Arabian Deputy Ministry for Mineral Resources Open-File 

Report USGS-OF-03-55, 36 p. 

Stacey, J.S., Stoeser, D.B., and Greenwood, W.R., 1983, U/Pb 

zircon geochronology and geologic evolution of the 

Halaban-Al Amar region of the eastern Arabian Shield, 

Kingdom of Saudi Arabia: Unpublished report available 

from the U.S. Geological Survey, Jiddah, Saudi Arabia. 

Steiner, J.C., Jahns, R.H., and Luth, W.C., 1975, Crystallization 

of alkali feldspar and quartz in the haplogranite system 

NaA1Si30 8-KA1Si30 8-Si02-H20 at 4 kb: Geological Society of America Bulletin, v. 86, p. 83-98. 

Peraluminous Granitoid Plutons, Saudi Arabia 31 
Stoeser, D.B., 1983, Reconnaissance geology of the Wadi Tarib 

quadrangle, sheet 18/43 C, Kingdom of Saudi Arabia: Unpublished report available from the U.S. Geological Survey, 

Jiddah, Saudi Arabia. 

Stoeser, D.B., and Elliott, J.E., 1980, Post-orogenic peralkaline 

and calc-alkaline granites and associated mineralization of 

the Arabian Shield, Kingdom of Saudi Arabia, in Cooray, 

P.G., and Tahoun, S.A., eds., Evolution and mineralization of the Arabian Nubian Shield, volume 4: Oxford and 

New York, Pergamon Press; Jiddah, King Abdulaziz 

University, Institute of Applied Geology Bulletin 3, p. 1-23. 

Stoeser, D.B., Fleck, R.J., and Stacey, J.S., 1982, Geochronology and origin of an early tonalite gneiss of the Wadi Tarib 

batholith and the formation of syntectonic gneiss complexes 

in the southeastern Arabian Shield, Kingdom of Saudi 

Arabia: Saudi Arabian Deputy Ministry for Mineral 

Resources Open-File Report USGS-OF-02-62, 27 p.; also 

in press Jiddah, King Abdulaziz University, Faculty of 

Earth Sciences, First Symposium on Pan-African crustal 

evolution in the Arabian-Nubian Shield, IGCP Project 164, 

Proceedings. 

Stoeser, D.B., and Stacey, J.S., 1983, Pan-African age continental collision in and plutonic rock U-Pb zircon geochronology of the southeastern Arabian Shield, Kingdom 

of Saudi Arabia: Unpublished data available from the U.S. 

Geological Survey, Jiddah, Saudi Arabia. 

Streckeisen, A.L., 1976, To each plutonic rock its proper name: 

Earth-Science Reviews, v. 12, p. 1-33. 

32 Contributions to Mineral Resources Research, 1984 

Takahashi, Masaki, Aramaki, Shigeo, and Ishihara, Shunso, 

1980, Magnetite-series-ilmenite-series versus !-type-S-type 

granitoids: Mining Geology Special Issue 8, p. 13-28. 

Tischendorf, G., 1977, Geochemical and petrographic characteristics of silicic magmatic rocks associated with rareelement mineralization, in Stemprok, M., and others, eds., 

Metallization associated with acid magmatism: Prague, 

Czechoslovakia Geological Survey, v. 2, p. 41-96. 

Thttle, O.F., and Bowen, N.L., 1958, Origin of granite in the light 

of experimental studies in the system NaA1Si30 8-

KA1Si308-Si02-H20: Geological Society of America 

Memoir 74, 152 p. 

Walsh, J.N., Beckinsdale, R.D., Skelhorn, R.R., and Thorpe, 

R.S., 1979, Geochemistry and petrogenesis of Tertiary 

granitic rocks from the island of Mull, northwest Scotland: 

Contributions to Mineralogy and Petrology, v. 71, p. 99-166. 

Watson, E.B., 1979, Zircon saturation in felsic liquidsExperimental results and applications to trace element 

geochemistry: Contributions to Mineralogy and Petrology, 

v. 70, p. 407-419. 

White, A.J.R., and Chappell, B.W., 1977, Ultrametamorphism 

and granitoid genesis: Tectonophysics, v. 43, p. 7-22. 

Winkler, H.G.F., Boese, M., and Marcopoulos, T., 1975, Low 

temperature granitic melts: Neues Jahrbuch fur Mineralogie 

Monatshefte, v. 6, p. 245-268. 

Yoder, H.S., Jr., and Eugster, H.P., 1954, Phlogopite synthesis 

and stability range: Geochimica et Cosmochimica Acta, v. 

6, p. 157-185. 
· Chapter C 

Muscovites from Selected Aplites and 

Quartz Veins of 

White Pine County, Nevada 

By D.E. LEE, E.L.M. BRANDT, and R.E. VAN LOENEN 

Selected muscovite data indicate both jurassic and 

Cretaceous hydrothermal activity associated with quartz veins 

U.S. GEOLOGICAL SURVEY BULLETIN 1694 

CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 
CONTENTS 

Abstract 35 

Introduction 35 

Muscovite analyses 36 

Muscovites from aplites 39 

Muscovites from quartz veins 39 

References cited 39 

FIGURE 

1. Map showing localities of analyzed muscovites 35 

TABLES 

1. Analytical data for muscovites from White Pine County, Nevada 36 

2. Additional data for analyzed muscovites 38 

34 Contents 
CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 

Muscovites from Selected Aplites and 

Quartz Veins of White Pine County, Nevada 

By D.E. Lee, E.L.M. Brandt, and R.E. Van Loenen 

Abstract 

Chemical analyses, physical properties, stable-isotope 

data, K-Ar ages, and coexisting minerals are listed for 

muscovites from selected aplites and quartz veins of White Pine 

County, Nevada. 

Two muscovites were recovered from a set of quartz veins 

in an area of wolframite mineralization in the southern Snake 

Range. Data obtained for these muscovites indicate that there 

may have been both Jurassic and Cretaceous hydrothermal activity associated with the quartz veins, corresponding to known 

episodes of magmatic activity in the area. 

One of the quartz-vein muscovites, present with minor 

amounts of wolframite, contains 1.90 weight percent fluorine. 

This may be the most fluorine-rich naturally occurring 

muscovite reported to 1985. 

INTRODUCTION 

This is one of a series of reports resulting from work 

on the granitoid rocks of eastern Nevada during the past 

20 years. In our efforts to understand these rocks we have 

selected particular samples for intensive study. For example, where possible we have tried to obtain (1) data 

on coexisting minerals and (2) different kinds of analytical 

data for splits from a given mineral fraction. Here we 

bring together chemical, radiometric-age, and stableisotope data for muscovites from selected quartz veins 

and aplites of eastern Nevada (fig. 1). We also summarize 

previously reported field data and information on 

minerals coexisting with these muscovites. A given sample number refers to the whole rock and to each of the 

minerals recovered from that rock. We emphasize that 

the different types of data reported for each of these 

muscovites were obtained on splits taken from the same 

mineral fraction. 

The muscovites analyzed were recovered from the 

respective samples after grinding each rock to -150 mesh 

and washing the powder in tap water to remove material 

smaller than about 40 llm (micrometers). The final 

purification of each of the muscovites was by repeated 

centrifuging in density-adjusted mixtures of methylene 

iodide and .bromoform and by repeated passes through 

. Area of 

figure 1 

Tertiary aplite~119-MW-60 

392-DL-62 

SNAKE CREEK-WILLIAMS 

CANYON PLUTON (155 Ma) 

~MW-61 

~121-MW-60 

POLE CANYON-CAN YOUNG 

CANYON PLUTON (79 Ma) 

38°50' ._ _________________ _j 

10 KILOMETERS 

Figure 1. Localities of analyzed muscovites. Outlines of intrusions modified from map of Hose and Blake (1976). Several 

other intrusions within map area are not shown for the sake 

of simplicity. Sample 245-MW-60 is from about 83 km north 

of locality of samples 119-MW-60 and 119-MW-61. Exact coordinates of all samples are listed in table 1. 

the Frantz isodynamic separator .1 Densities of the 

analyzed muscovites were determined by the suspension 

method during centrifuging. Each of these muscovites 

fractionated over a very narrow density range, indicating 

a similarly small range of composition. 

Major elements were determined by the chemical 

methods described by Peck (1964) and calculated on the 

basis of 24 {O,OH,F,Cl) to the general mica formula 

X2Y4_

6Z8

0 20 (OH,F,Cl)4 , using the computer program 

described by Jackson and others (1967). Calculated values 

for density and cell volume were determined by the same 

program. The observed cell parameters were obtained by 

1

Any use of trade names in this report is for descriptive purposes 

only and does not imply endorsement by the U.S. Geological Survey. 

Muscovites of White Pine County, Nevada 35 
least squares refmement of powder diffractometer data, with 

an internal standard of CaF2 and a self-indexing computer 

program developed by Evans and others (1963). Quantitative 

barium, rubidium, and strontium values were determined 

by X-ray fluorescence (Lee and Doering, 1974, barium; and 

1980, rubidium and strontium). Results of semiquantitative 

spectrographic analyses are based on their identity with 

geometric brackets whose boundaries are 1.2, 0.83, 0.56, 

0.38, 0.26, 0.18, 0.12, and so forth, and are reported arbitrarily as midpoints of these brackets: 1, 0.7, o.s', 0.3, 0.2, 0.15, 

and 0.1, respectively. The precision of a reported value is 

approximately plus or minus one bracket at 68 percent confidence or two brackets at 95 percent confidence. 

Stable-isotope data are based on the SMOW (standard mean ocean water) standard (Friedman and O'Neil, 

1977). 

MUSCOVITE ANALYSES 

All of the analyzed muscovites (table 1; fig. 1) are 

2M1 micas, as determined from X-ray diffraction patterns. Ionic ratios are close to the ideals for dioctahedral 

mica, and rather good agreement exists between the 

observed and calculated values for density and cell volume 

(table 2). Coexisting minerals present in "sparse" or 

''trace'' amounts were detected only during mineral 

separation work. Those listed as sparse make up an 

estimated 0.02-0.05 weight percent of the whole rock. 

Each trace mineral makes up less than about 0.005 weight 

percent of the rock. We emphasize that such minerals as 

gahnite and columbite-tantalite are not recognizable in 

the field. 

Table 1. Analytical data for muscovites from White Pine County, Nev. 

[Chemical analyses by E.L. Munson Brandt. Quantitative Ba, Rb, and Sr analyses by W.P. Doering. Semiquantitative spectrographic analyses 

by L.A. Bradley. Leaders ( ---), not determined] 

Muscovites from aplites Muscovites from quartz veins 

Sample No.-- 119-MW-60 1199-MW-61 121-MW-60 245-MW-60 392-DL-67 396-DL-67 

Sample locality: 

Lat N.------ 38°59'30" 38°58'11" 38°58'31" 39°35'55" 38°58'36" 38°57'04" 

Long W.----- 114°12'32" 114°13'12" 114°11'57" 114°07'55" 114°21'17" 114°20'36" 

Chemical analyses (weight percent) 

Sio2 46.08 45.96 49.95 46.27 46.36 47.54 

Al 2o3 30.10 32.52 30.30 33.57 32.71 32.40 

Fe 2o3 3.29 3.15 2.84 1.81 1.54 .59 

FeO 2.47 1. 89 1.11 1.61 1.51 1.77 

MgO 1.38 .28 .53 .54 1.45 .94 

CaO .oo .oo .08 .oo .00 .oo 

Na 2o 1.02 .54 1.29 .54 .49 .23 

K2o 10.45 10.86 9.53 10.79 10.82 11.18 

H o<+) 3.84 4.09 4.15 4.20 3.65 2 H (-) 2o .06 .03 .05 .05 .02 .01 

Ti02 .60 .11 .20 .15 .39 .20 

P2o5 .01 .01 .01 .04 .01 

MnO .26 .13 .06 .06 .29 

Cl .oo .oo .01 .01 .00 .oo 

F .71 .42 .45 .58 .24 1.90 

Subtotal- 100.27 99.99 100.14 99.83 100.71 

Less 0 .30 .18 .24 .10 .80 

Total 99.97 99.81 99.90 99.73 99.91 

Quantitative Ba, Rb, Sr (weight percent) 

Ba 0.0071 0.0011 0.2717 0.0720 

Rb .1844 .3110 .0666 .0957 

Sr .00003 .00002 .0044 .00117 

36 Contributions to Mineral Resources Research, 1984 
Table 1. Analytical data for muscovites from White Pine County, Nev.-Continued 

Sample No.-- 119-MW-60 

Sample locality: 

Lat N.------ 38°59'30" 

Long W. ----- 114 ° 1 2 '3 2" 

Muscovites from aplites 

1199-MW-61 121-MW-60 

38°58'11" 38°58'31" 

114°13'12" 114°11'57" 

245-MW-60 

39°35'55" 

114°07'55" 

Muscovites from quartz veins 

392-DL-67 396-DL-67 

38°58'36" 38°57'04" 

114°21'17" 114°20'36" 

Number of ions on basis of 24(0,0H,F,Cl) 

{

p5+ 

Z Si 

-!1~~~ Fe3+ 

Y Ti 

Mg2+ Fe 

Mn 

l

ea 

Na 

X K 

Ba 

Rb 

OH 

B 

Ba 

Be 

Cr 

Cu 

Ga 

Li 

Nb 

Ni 

Pb 

Sc 

Sn 

Sr 

v 

Zn 

Zr 

F 

Cl 

6.28} 8.00 

1.72 

3.34 12l 

.06 4.11 

.28 

• 28 

.03 

.27ool 

1.82 2.11 

.oo 

.02 

3.49} .31 3.80 

.oo 

6.23} 8.00 . 1. 77 

3.32 42l 

.01 4.04 

.06 

0 21 . 

.02 

.00114 

1.88 2.05 

.oo 

.03 

3.70} .18 3.88 

.oo 

6.62} 8.00 

1.38 

3.35}28 

.02 3.87 

.10 

.12 

.33021 

~= 61 1. 96 

6.22} 8.00 

1.78 

3.54] .18 

.02 4.04 

.11 

0 18 

.01 

.ool14 

~=85 1.99 

3.72} .25 3.97 

.oo 

0.01} 6.24 8.00 

1.75 

3.44] .16 

.04 4.11 

.29 

.17 

.01 

.ool13 

1.86 2.02 

.02 

.01 

3.771 .10 3.87 

. .00 

Semiquantitative spectrographic analyses (weight percent) 

0 

(2) 

.0015 

.0003 

.0007 

.015 

.15 

.02 

.0005 

0 

.015 

.01.5 

(2) 

.007 

0 

0 

~.005 

( ) 

.003 

0 

.0002 

.03 

.03 

.02 

0 

.0015 

.0005 

.05 

(~) 

.0007 

.07 

.0015 

0.003 

.003 

.002 

0 

0 

0 

0 

0 

0 

.0005 

.02 

.07 

.015 

.002 

.07 

.005 

0 

0 

.015 

.0005 

.0001 

.0005 

.015 

.02 

.005 

0 

.001 

.003 

0 

.002 

.001 

.003 

0.005 

(2) 

0 

.001 

.001 

.0002 

.01 

.02 

.002 

.002 

.0015 

.01 

(2) 

.015 

0 

.07 

6.37} 8.00 

1. 63 

3.49l .06 

.02 3.99 

.19 

.20 

.• 03 

.00]6 

1.91 1.98 

.oo 

.01 

3.26} .81 4.07 

.oo 

0.007 

(2) 

0 

.002 

.0003 

.0003 

.005 

.OS 

.001 

0 

.0007 

.003 

(2) 

.005 

0 

0 

1Structural formula for 121-MW-60 calculated on anhydrous basis of 22(0). See Ernst, 

1963, p. 1359). 

2see quantitative Ba, Rb, Sr heading above for values. 

Muscovites of White Pine County, Nevada 37 
Table 2. Additional data for analyzed muscovites 

[Leaders ( ---), not determined 

Sample No.-----

Optical properties: 

f3 index 

2V (-) 

Density: 

Observed (±0.01) 

Calculated 

119-MW-60 

2.89 

2.897 

Cell parameters observed: 

a (±0.0054) 

b (±0.010~) 

c (::t:0.010A) 

fj (±5.0') 0 

Volume (::t:1.0A3) 

5.212 

9.050 

20.034 

95°45' 

940.2 

Cell volume calculated: 

Volume (A3 ) 945.68 

Stable isotope data (SMOW): 

d18o, permil 

bD, permil -135.3 

K-Ar age (m.y.)1 __ 

Coexisting minerals 

34±1 

Quartz 

Feldspar 

Garnet 

Fluorite 

(spars~) 

Apatite 

(sparse) 

Muscovites from aplites 

199-MW-61 

1.600 

41° 

2.88 

2.889 

5.201 

9.031 

20.056 

95°49' 

937.3 

939.75 

9.9 

-123.3 

53.9±1.5 

Quartz 

Feldspar 

Garnet 

Fluorite 

(sparse) 

Columbitetantalite 

(trace) 

121-MW-60 

2.88 

31±2 

Quartz 

Feldspar 

Garnet 

Traces of 

gahn~tej 

apat1te , 

columbitetantalite 

1K-Ar age data from Lee and others (1970). 

Muscovites from quartz veins 

245-MW-60 

1. 586 

42° 

2. 8 7 

8.0 

-61.5 

247.9±2.0 

Quartz 

Feldspar 

Garnet 

Traces of 

gahnite, 

monazite, 

columbitetantalite, 

zircon 

392-DL-67 

1. 598 

39° 

2.86 

2.865 

5.207 

9.030 

20.050 

95 °48 I 

938.0 

933.01 

9.9 

-64.6 

156±6 

Quartz 

Apatite 

(sparse) 

396-DL-67 

1. 582 

44° 

2. 86 

2.856 

5.200 

9.031 

20.053 

95°45' 

937.0 

929.16 

11.9 

-66.8 

72.1±2.2 

Quartz 

Wolframite 

(sparse) 

2K-Ar age data from Lee and others (1980); uranium-lead isotopic age of coexisting zircon is 

75±9 Ma (unpublished data). 

3Analytical data for apatite given by Lee and others (1973). 

38 Contributions to Mineral Resources Research, 1984 
Muscovites from Aplites 

Muscovite 119-MW-60 is from an isolated mass of 

Tertiary aplite apparently squeezed up between two faults 

by stresses related to late movement on the Snake Range 

decollement (Lee and Van Loenen, 1971). This aplitic 

muscovite is notably high in Fe2

0 3 , FeO, and MgO 

(table 1), and as such it represents solid solution between 

muscovite and celadonite and is phengitic (Foster, 1956). 

Muscovites from the two-mica granites of the area are 

also phengitic (Lee and others, 1981). 

Three of the analyzed muscovites were recovered 

from aplites that intrude the unusual Cretaceous two-mica 

granites of northeastern Nevada that bear muscovite 

phenocrysts (Lee and others, 1981). Samples 199-MW-61 

and 121-MW-60 are from the Pole Canyon-Can Young 

Canyon area of the southern Snake Range, and sample 

245-MW -60 is from the Kern Mountains, about 83 km 

north of the other two. These three muscovites are 

chemically similar except that number 121-MW-60 is 

rather rich in Na20 (1.29 weight percent), equivalent to 

about 16.8 mol percent of paragonite. The K-Ar ages 

determined for these muscovites are spuriously low for 

reasons discussed by Lee and others (1970, 1980). 

Muscovites from Quartz Veins 

Samples 392-DL-67 and 396-DL-67 were collected 

about 3 km apart, from a set of quartz veins that strikes 

N. 60°-75° E. and dips 60°-90° SE., just north of 

Williams Canyon on the west side of the southern Snake 

Range (Lee and Van Loenen, 1971, pl. 1). These quartz 

veins are present within the Ju~assic (155 Ma) Snake 

Creek-Williams Canyon intrusion (Lee and Christiansen, 

1983). Some of the larger veins in the set are more than 

3 m wide and are continuously exposed over a horizontal distance of 300m or more. A few (one in particular) 

of the largest of these quartz veins were worked for 

wolframite at the Hub mine during World War I. During 1911-1916 the Hub district produced 9,058 tons of 

wolframite concentrate valued at $672,353 (Couch and 

Carpenter, 1943). 

Sample 392-DL-67 is from a quartz vein about 

10 em wide, and sample 396-DL-67 is from a bull quartz 

vein about 1 m wide. The two muscovites gave K-Ar ages 

of 156±6 and 72.1 ±2.2 Ma, respectively. The 156-Ma 

date is in good agreement with the age (155 Ma) determined for the host intrusion. It was suggested (Lee and 

others, 1970, p. 097) that the 72.1-Ma date determined 

for 396-DL-67 might result from loss of argon due to late 

slippage along the quartz vein from which the sample was 

collected. However, more recently (Lee and Christiansen, 

1983), it has become apparent that there was also a 

Cretaceous (79 Ma) magmatic event in the area. Thus it 

is possible that muscovites 392-DL-67 and 396-DL-67 

actually formed at very different times, even though the 

host veins have the same attitude. Therefore the 

wolframite mineralization at the Hub mine may be Late 

Cretaceous in age, much younger than the Late Jurassic 

host intrusion. 

The stable-isotope data (table 2) are consistent with 

the idea that formation of the veins sampled might have 

been related to one or more of the magmatic episodes 

recognized in the area. The o180 values of the muscovites are near the o180 value for magmatic water postulated by Taylor (1977), and their oD values are near the 

estimated oD of "juvenile" water (Sheppard and Epstein, 

1970). 

Muscovite 392-DL-67 contains somewhat more 

Fe2

0 3, MgO, N~O, and BaO and much less fluorine 

than does 396-DL-67. Indeed, to our knowledge, 

muscovite 396-DL-67 contains more fluorine (1.90 weight 

percent) than any natural muscovite mentioned in the 

literature. This fluorine-rich muscovite is present with 

minor amounts of wolframite. 

The high fluorine content of muscovite 396-DL-67 

is not readily apparent from the physical properties listed 

in table 2, but it does seem to have an effect on the color 

of the mineral. In grains of 0.10-mm size, the bulk 

mineral has a chalky white color. A bulk fraction of the 

same size grains of muscovite 392-DL-67 is brownish gray 

in color, but as noted this muscovite also has slightly 

higher Fe2

0 3 and MgO contents than the fluorine-rich 

396-DL-67. 

REFERENCES CITED 

Couch, B.F., and Carpenter, J.A., 1943, Nevada's metal and 

mineral production (1859-1940, inclusive): University of 

Nevada Bulletin, Geology and Mining Series 38, 159 p. 

Ernst, W.G., 1963, Significance of phengitic micas from lowgrade schists: American Mineralogist, v. 48, p. 1357-1373. 

Evans, H.T., Jr., Appleman, D.E., and Handwerker, D.S., 

1963, The least squares refinement of crystal unit cells with 

powder diffraction data by an automatic computer indexing method: American Crystallographic Association Program and Abstracts 1963 Annual Meeting, p. 42-43. 

Foster, M.D., 1956, Correlation of dioctahedral potassium 

micas on the basis of their charge relations: U.S. Geological 

Survey Bulletin 1036-D, p. 57-67. 

Friedman, Irving, and O'Neil, J .R., 1977, Compilation of stable 

isotope fractionation factors of geochemical interest, in 

Fleischer, Michael, ed., Data of geochemistry, 6th edition: 

U.S. Geological Survey Professional Paper 440-KK, 12 p. 

Hose, R.K., and Blake, M.C., 1976, Geology of White Pine 

County, Nevada: Nevada Bureau of Mines and Geology 

Bulletin 85, pt. 1, p. 1-35. 

Jackson, E.D., Stevens, R.E., and Bowen, R.W., 1967, A 

computer-based procedure for deriving mineral formulas 

from mineral analyses, in Geological Survey research 1967: 

U.S. Geological Survey Professional Paper 575-C, 

p. C23-C31. 

Muscovites of White Pine County, Nevada 39 
Lee, D.E., and Christiansen, R.L., 1983, The granite problem 

as exposed in the southern Snake Range, Nevada: Contributions to Mineralogy and Petrology, v. 83, p. 99-116. 

Lee, D.E., and Doering, W.P., 1974, Barium in hybrid granitoid 

rocks of the southern Snake Range, Nevada: U.S. Geological Survey Journal of Research, v. 2, p. 671-675. 

___ 1980, Rubidium and strontium in hybrid granitoid rocks, 

southern Snake Range, Nevada: U.S. Geological Survey Professional Paper 1158-A, p. 1-7. 

Lee, D.E., Kistler, RW., Friedman, Irving, and Van Loenen, R.E., 

1981, 1\vo-mica granites of northeastern Nevada: Journal 

of Geophysical Research, v. 86, no. B11, p. 10607-10616. 

Lee, D.E., Marvin, R.F., and Mehnert, H.H., 1980, A radiometric 

age study of Mesozoic-Cenozoic metamorphism in eastern 

White Pine County, Nevada, and nearby Utah: U.S. 

Geological Survey Professional Paper 1158-C, p. 17-28. 

Lee, D.E., Marvin, R.F., Stern, TW., and Peterman, Z.E., 1970, 

Modification of potassium-argon ages by Tertiary thrusting 

40 Contributions to Mineral Resources Research, 1984 

in the Snake Range, White Pine County, Nevada, in 

Geological Survey research 1970: U.S. Geological Survey 

Professional Paper 700-D, p. D196-D206. 

Lee, D.E., and Van Loenen, R.E., 1971, Hybrid granitoid rocks 

of the southern Snake Range, Nevada: U.S. Geological 

Survey Professional Paper 668, 48 p. 

Lee, D.E., Van Loenen, R.E., and Mays, R.E., 1973, Accessory 

apatite from hybrid granitoid rocks of the southern Snake 

Range, Nevada: U.S. Geological Survey Journal of Research, 

v. 1, no. 1, p. 89-98. 

Peck, L.C., 1964, Systematic analysis of silicates: U.S. Geological Survey Bulletin 1170, 89 p. 

Sheppard, S.M.F., and Epstein, Samuel, 1970, D/H and 180/160 

ratios of minerals of possible mantle or lower crustal origin: 

Earth and Planetary Science Letters, v. 9, p. 232-239. 

Thy lor, H.P., Jr., 1977, Water-rock interactions and the origin 

of H20 in granitic batholiths: Journal of the Geological 

Society of London, v. 133, p. 509-558. 
Chapter D 

Chemical and Stable-Isotope Data for 

Micas from Metaclastic Rocks of 

Eastern White Pine County, Nevada 

By D.E. LEE, R.E. VAN LOENEN, 

E.L.M. BRANDT, and J.D. GLEASON 

Micas related to movement on thrust faults 

have oD values well below the "normal" range 

U.S. GEOLOGICAL SURVEY BULLETIN 1694 

CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 
CONTENTS 

Abstract 43 

Introduction 43 

Analytical techniques 43 

Samples 44 

Chemical analyses 44 

Stable-isotope data 49 

Conclusions 51 

References cited 53 

FIGURES 

1. Index map showing localities of mica samples analyzed 45 

2. Camera Iucida drawings of four quartzite samples from which recovered 

micas were analyzed 48 

3. Plot of o180 versus oD values for analyzed micas from metaclastic rocks of 

eastern White Pine County, Nevada 52 

TABLES 

1. Analytical data for white micas recovered from late Precambrian and Early 

Cambrian metasedimentary rocks 46 

2. Mineral content of the rocks from which the analyzed micas were 

recovered 49 

3. Properties of analyzed white micas 50 

4. Analytical data for biotites recovered from late Precambrian and Early Cambrian metasedimentary rocks 51 

5. Stable-isotope data and K-Ar ages for micas recovered from late Precambrian and Early Cambrian metasedimentary rocks 52 

42 Contributions to Mineral Resources Research, 1984 
CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 

Chemical and Stable-Isotope Data for Micas from 

Metaclastic Rocks of Eastern White Pine County, Nevada 

By D.E. Lee, R.E. Van Loenen, E.L.M. Brandt, and J.D. Gleason 

Abstract 

Regional thrust faults in eastern White Pine County, 

Nevada, have moved Paleozoic sedimentary rocks over late 

Precambrian and Early Cambrian metaclastic rocks. The lower 

plate metaclastic rocks are exposed over thousands of square 

kilometers and were derived from a deeply weathered Proterozoic crystalline terrane. 

Metamorphic micas recovered from the late Precambrian 

and Early Cambrian lower plate rocks have yielded Tertiary KAr ages that generally relate to the time of latest movement 

on the overlying decollement. Chemical and stable-isotope data 

are presented for these same metamorphic micas. The most 

remarkable of the white micas analyzed contains 2.91 weight 

percent Na2

0, equivalent to 38 mol percent paragonite. Many 

of the white micas and biotites studied have oD values well 

below the "normal" range of -50 to -85 permil, possibly as 

a result of isotopic exchange with hydrothermal-meteoric 

waters circulating near the regional thrust faults. However, the 

fairly "normal" o180 values of the analyzed micas imply rather 

low water:rock ratios during the isotopic exchange. 

INTRODUCTION 

In eastern White Pine County, Nev., regional thrust 

faults have moved Paleozoic sedimentary rocks over late 

Precambrian and Early Cambrian metaclastic rocks. 

Plutons ranging in age from Jurassic to Tertiary are also 

present in lower plate rocks. 

Lower plate quartzitic rocks are exposed over 

thousands of square kilometers. Previous studies have 

focused on the chemical composition of these quartzites 

(Lee, Van Loenen, and others, 1980) and on the K-Ar ages 

of their constituent micas (Lee and others, 1970; Lee, Marvin, and Mehnert, 1980). We have also obtained U-Th-Pb 

isotopic ages (78±9 Ma) for metamorphic monazite and 

recrystallized detrital zircon from lower plate metaclastic 

rocks that have been subjected to staurolite-kyanite-grade 

metamorphism (Lee and others, 1981, and unpub. data). 

1\vo main conclusions have been drawn from these 

earlier studies. First, the original miogeosynclinal shelf 

sediments were derived from a deeply weathered, Early 

Proterozoic crystalline terrane. These sediments were mixtures of (1) sand with a fair degree of mineralogical 

maturity and (2) argillaceous materials. In most cases the 

bulk composition of the argillaceous material apparently was appropriate for the crystallization of metamorphic 

micas under conditions of metamorphism that were essentially isochemical, except for depletion of pore fluids. 

The second main conclusion drawn from the earlier 

studies is that the K-Ar ages (commonly Thrtiary) obtained 

for the metamorphic micas recovered from lower plate 

(late Precambrian and Early Cambrian) metasedimentary 

rocks generally relate to the time of latest movement on 

the overlying decollement (Lee and others, 1970; Lee, Marvin, and Mehnert, 1980). In other words, these metamorphic micas were partially to totally degassed by strain 

heating related to late movement on the overlying 

decollement and therefore give a maximum age for the 

time of that movement. Samples collected just below the 

sole of the decollement have been degassed most completely and so give K-Ar ages most representative of the 

time of latest movement. 

Despite past efforts, however, two essential questions remain with regard to any one of the micas recovered from the lower plate metasedimentary rocks: (1) If 

the mica was only degassed at, say, 20 Ma, then when did 

it first crystallize from the argillaceous material in the 

original clastic sediment? (2) Did the mica re-equilibrate 

chemically to the conditions of the most recent event during which it was degassed? 

Here we present chemical and stable-isotope data 

for selected micas recovered from lower plate metaclastic 

rocks of eastern White Pine County, Nevada. We have previously reported K-Ar ages for most of these same micas 

(Lee and others, 1970; Lee, Marvin, and Mehnert, 1980), 

and we emphasize that all data reported for a given sample were determined on splits of the same mineral fraction. 

ANALYTICAL TECHNIQUES 

The micas were prepared for analysis by methods 

described by Lee and Van Loenen (1970, p. D198). Briefly, 

Data for Micas from Nevada 43 
the final purification was by centrifuging in densityadjusted mixtures of methylene iodide and bromoform 

and by repeated passes through the magnetic separator. 

Generally the muscovites fractionated over a very narrow density range, indicating a similarly small range of 

composition of the mineral in the rock. On the other 

hand, each biotite fractionated over a rather large density range, as indicated in table 3, and the recovery of a 

homogeneous sample for analysis was difficult. Similar 

difficulties were encountered in the purification of the 

biotites discussed by Lee and Van Loenen (1970), who 

have described the problem in some detail. 

Major elements were determined by the chemical 

methods described by Peck (1964) and calculated on the 

basis of 24 (O,OH,F,Cl) to the general mica formula 

X2Y4_

6Z8

0 20 (OH,F,Cl)4 , using the computer program 

described by Jackson and others (1967). Calculated values 

for density and cell volume were determined by the same 

program. 

Quantitative barium, rubidium, and strontium 

values were determined by X-ray fluorescence. The 

methods for barium and for rubidium and strontium are 

described by Lee and Doering (1974 and 1980, respectively). Results of semiquantitative spectrographic 

analyses are based on their identity with geometric 

brackets whose boundaries are 1.2, 0.83, 0.56, 0.38, 0.26, 

0.18, 0.12, and so forth, and are reported arbitrarily as 

midpoints of these brackets: 1, 0.7, 0.5, 0.3, 0.2, 0.15, 

and 0.1, respectively. The precision of a reported value 

is approximately plus or minus one bracket at 68 percent 

confidence or two brackets at 95 percent confidence. 

Stable-isotope data are based on SMOW, standard 

mean ocean water (Friedman and O'Neil, 1977). 

A given sample number refers to the whole rock 

and to each of the mineral fractions recovered from that 

rock. 

SAMPLES 

Sample localities are shown in figure 1. Exact coordinates are listed and the field setting of each of these 

samples is described by Lee, Marvin, and Mehnert (1980). 

Some representative textures are shown in figure 

2. Sample 503 was collected from lower plate rocks on 

the northwestern flank of the northern Snake Range. The 

quartz grains show evidence of pronounced strain and 

recrystallization, with sutured contacts except where 

shielded by mica. The orientation of the micas is determined mainly by the outlines of the original detrital 

quartz grains. Samples 505 and 506 (not shown) are 

similar to 503, but in sample 506 a directional element 

is evident in thin section. 

Sample 555 (K-Ar age, 141 ± 5 Ma) is unique among 

those analyzed in that it does not appear to be spatially 

44 Contributions to Mineral Resources Research, 1984 

related to any thrust fault. It was collected at an elevation of 3,440 m in the Wheeler Peak area of the southern 

Snake Range. K-Ar ages determined on samples from the 

Wheeler Peak area show no relation to present altitude, 

position within the stratigraphic section, or distance 

below any hypothetical but now-eroded thrust fault. The 

original argillaceous cement in sample 555 has crystallized to an extremely fine grained muscovite surrounding 

the detrital quartz, which shows no evidence of strain or 

recrystallization. These textural features may be the result 

of a very low grade static or load metamorphism (Lee, 

Marvin, and Mehnert, 1980). 

Sample site 561 probably is not more than a few 

meters below the sole of the Snake Range decollement. 

The quartz is completely recrystallized and drawn out to 

express a planar element approximately parallel with the 

present surface and presumably with the overlying thrust 

fault. The micas are aligned in the same plane. 

Sample 569 was collected from a shaly unit within 

quartzite. About 23 percent of the rock is a very 

homogeneous, well-crystallized metamorphic muscovite 

that shows no preferred orientation. 

CHEMICAL ANALYSES 

All the white micas analyzed (table 1) are 2M1 

micas. Many are high in silica, low in alumina, and contain appreciable amounts of FeO and MgO. As such, they 

represent solid solution between muscovite and celadonite 

and are phengites (Foster, 1956). Most of the white micas 

that contain notable amounts of Fe20 3 (table 1) coexist 

with magnetite or hematite. Those highest in Fe20 3 

(numbers 547 and 548) coexist with a mineral that is part 

of the metamorphic fabric and is tentatively identified 

as maghemite. The mineral content of the rocks from 

which the analyzed micas were recovered is summarized 

in table 2. 

One of the micas in the literature most similar to 

some of those in table 1 comes from a very similar 

geologic setting. It is the "iron-rich muscovitic mica" 

from the Grandfather Mountain area, North Carolina 

(Foster and others, 1960). This same mica was referred 

to as a "phengite" by Ernst (1963, p. 1358) and as a 

"phengitic mica" by Velde (1965, p. 906). The analysis, 

as corrected for impurities and adjusted to 100 weight 

percent (Foster and others, 1960, p. 845), is Si02, 47.62; 

A120 3, 24.74; Fe20 3, 8.11; FeO, 2.55; MgO, 1.88; MnO, 

0.08; Na2

0, 0.11; K20, 10.72; F, 0.14; and H20, 4.05. 

The mica analyzed was recovered from a late Precambrian or Early Cambrian micaceous arkosic quartzite exposed below the thrust surface in the Grandfather 

Mountain area. Similar mica is present in most of the 

metamorphosed arkoses and arkosic quartzites exposed 

both above and below the thrust (Foster and others, 1960, 
15' 1WOO' 

JUAB 

0 506 L.S;Q. 

I 

I 

39°30' I 

I 

I 

I 

<I 

<< QIJ: 

w >lLIJ:J (!) ZJ 

z I 

<( 

I a: 

~ 

I w 

w 

a: I u 

>- I ...J LLJ 

'>-

....J ...J ....J w 

:I: ~ IUJ u J....J en ....J 

~ 

10 15 20 25 KILOMETERS 

Figure 1. Localities of mica samples analyzed. Precise latitudes and longitudes of localities are listed in Lee, Marvin, and Mehnert (1980). Sample locality 511 (not shown) is from the Toana Range about 150 km north of map area. 

Data for Micas from Nevada 45 
Table 1. Analytical data for white micas recovered from late Precambrian and Early Cambrian metasedimentary rocks 

[Chemical analyses by E.L. Munson Brandt. Quantitative Ba, Rb, and Sr analyses by W.P. Doering. Semiquantitative spectrographic analyses 

Sample No.-- 1427 1428 1476 478 501 503 505 

Chemical analyses (weight percent) 

SiOz 46.40 45.86 49.42 47.23 46.97 45.99 46.45 

Al 2o3 29.60 33.82 29.62 28.52 30.10 30.24 31.04 

Fe 2o3 4.35 1. 69 2.72 5.41 2.24 4.31 2.39 

FeO 1.48 1.45 1. 8 7 .94 1. R5 1.62 2.57 

MgO 1. 53 .76 1. 52 I. 62 1.88 1.02 1. 17 

CaO .05 0 01 .00 .oo .oo .oo .00 

Na 2o .42 .55 .16 .39 .37 .30 

KzO 10.50 10.36 10.99 10.79 10.86 lJ 0 03 

HzO+ 4.37 4.35 4.16 4.05 4.07 4.07 

HzO .05 .as .06 .02 .03 .01 .01 

TiOz .75 .33 0 58 .76 1.40 1.10 .50 

Pzos .05 .05 .00 .oo .oo .00 

MnO .02 .01 .02 .02 .02 .02 

Cl .00 .01 .01 .oo .01 .00 

F .07 .13 .07 .16 .16 0 38 

Subtotal 99.64 99.43 85.79 99.91 99.88 99.78 99.93 

Less 0 .03 .as .03 .07 .07 .16 

Total 99.61 99.38 99.88 99.81 99.71 99.77 

Quantitative X-ray fluorescence analyses (weight percent) 

Ba 0.3636 0.4300 o. 0867 0.1040 0.1584 .1370 

Rb .0297 .0217 .0312 .0362 .0375 .0409 

Sr .0064 .0094 .0019 .0015 .0055 .0021 

Numbers of ions on basis of 24(0, OH, F, Cl) 

r+ .01} .01) Z Si 6.30 8.00 6.17 8.00 6.42 J 8.00 6. 36} 8.00 6. 27 } 8.00 6. 31} 8.00 

(~~ 1. 69 1.82 I. 58 1.64 1.73 1. 69 3.04} 3.53} 2.99 J 3.16} 3.13} 3.28} Fe 3+ .44 • 17 .55 .23 .44 .24 

Y Fe 2+ .17 4.04 .16 4.04 0 11 4.06 0 21 4. 12 • 19 4.08 .29 4.10 

Mg .31 .15 .33 .38 0 21 .24 

Ti .08 .03 .08 0 14 .11 .05 

r .01) .o] .0] .ool Na 0 11 0 14 .04 .10 .00)10 .oo) .OR 

X K 1.82 1. 96 1.78 1. 9 5 1.91 1.96 1.86 1. 97 I.R9 2.00 1. 91 2.00 

Ba .02 .03 .01 .01 0 01 .01 

Rb .oo .oo .oo .oo .00 .oo 

Sr .oo .oo .00 .oo .oo .oo 

(~" 3.96} 3.90} 3.77} 3.66} 3. 70} 3. 69} .03 3.99 .06 3.96 .03 3.80 .07 3.73 .07 3. 77 .16 3.85 

Cl .oo .oo .oo .oo .00 .oo 

Semiquantitative spectrographic analyses (weight percent) 

B (.002 .002 .002 .002 .002 <.002 

Be .0002 .0003 .0003 .0005 .0007 .0007 

Co .0005 .ooos .002 .0007 .002 .0007 

Cr .015 .015 .007 .015 .02 .007 

Cu .002 .0005 .0003 .0002 .0005 .0003 

Ga .005 .01 .007 .007 .005 .007 

Li .01 0 0 .007 .03 .03 

Mo 0 0 .0003 0 (.0003 <.0003 

Nb 0 <.001 0 .001 .001 .001 

Ni .001 .001 .007 .001 .005 .001 

Pb .001 .005 .0015 .0015 .0015 .001 

Sc .001 .003 .003 .007 0 0015 .0015 

v .05 .as .02 .03 .007 .007 

y (.001 .0015 .001 .0015 0 0 

Yb .0001 .0002 .0001 .0001 (.0001 <.0001 

Zr .005 .01 .005 .005 .002 .002 

1samples 427' 428' 476 recovered from Precambrian Z metasedimentary rocks; all others recovered 

from Early Cambrian metasedimentary rocks. 

46 Contributions to Mineral Resources Research, 1984 
by L.A. Bradley. Leaders(---), not determined] 

506 511 547 548 549 2sss 561 569 

Chemical analyses--Continued 

46.98 45.76 45.71 45.56 46.07 51.89 46.40 47 .OS 

31.21 29.26 30.38 30.44 36.20 27.31 30.53 27.04 

1.63 5.23 6.44 6.49 1. 38 4. 38 3.55 5.87 

I. 70 I.80 .56 • 68 • 58 .56 1.21 1. 15 

I. 70 1.17 .84 .66 .62 .80 1.42 2.03 

.oo .oo .03 .01 .01 .00 .oo .00 

.33 .36 .42 .32 2.91 .19 .39 .27 

II. 02 10.89 10.69 10.80 7. 12 10. 11 10.69 10.89 

4.02 4.08 4.27 4.25 4.43 3.62 4.22 4.02 

.oo .00 .02 .03 .os .OS .01 .04 

.85 1.02 .36 .44 .35 • 67 1.05 1.30 

.oo .02 .01 .01 .01 .03 

.01 .02 .02 .oo .OI .03 

.01 .00 .01 .oo .01 .oo .oo .01 

.48 .35 .06 .04 .04 .OS .os .08 

99.94 99.96 99.82 99.72 99.78 99.63 99.54 99.81 

.20 • 15 .03 .02 .02 .02 .02 .03 

99.74 99.81 99.79 99.70 99.76 99.61 99.52 99.78 

Quantitative X-ray fluorescence analyses--Continued 

• I353 • IOI7 .0984 .0843 • 1134 .2140 .0855 

.0384 .0495 .0277 .0089 .0292 .0248 .0338 

.0093 .OOI8 .0058 .0456 .0022 • 0274 .0017 

Numbers of ions on basis of 24(0, OH, F, Cl)--Continued 

6. 34 } 

I.66 

8.00 6. 26 } 

I.74 8.00 6. 22 } I. 78 8.00 6.24 } 8.00 1. 76 

6.08} 

1. 92 8.00 6.48} 1. 52 8.00 6. 29} 1. 71 8.00 6.44} 

1.56 8.00 

3.31} 2.98} 3.09 J 3.11} 3. 71} 3. 23} 3.17} 2.80} • I7 .54 .66 .67 • 14 .49 .36 • 61 

.I9 4.IO • 2I 4.08 .06 4.02 .08 4.05 .06 4.07 .07 4.05 • 14 4.07 • 13 4.08 

.34 • 24 .17 • 14 .12 .18 .29 .41 

.09 .II .04 .OS .04 .08 • 11 .13 

.00)9 .00)10 .oo) .I1 .oo) .09 .oo) .74 .oo) .06 .1000) .oo) .07 

I.90 2.00 I.90 2.02 1.86 I. 98 

=~t 1. 98 1.20 I.9S I.90 1. 97 1.85 1.96 1.90 1. 98 .01 .01 .01 .01 .01 .01 • 01 

.oo .01 .oo .oo .oo .oo .oo 

.oo .oo .oo .oo .oo .oo .oo 

3.62} 3. 72 } 3.R8} 3.88} 3.90} 3.57} 3.82} 3.67} • 21 3.83 • I5 3. 8 7 .03 3.9I .02 3.90 .02 3.92 .02 3.59 .02 3.84 .04 3. 71 

.oo .oo .oo .oo .oo .oo .00 .oo 

Semiquantitative spectrographic analyses--Continued 

<.002 (.002 <.002 <.002 <.002 .003 .002 .003 

.001 .OOI5 .0002 .0002 .0003 .0005 .0005 .0003 

.OOI .0007 .0007 .001 .0003 .001 .0007 .0015 

.007 .015 .003 .oos .015 .005 .OIS .007 

.0002 .0003 .0002 .0002 .0003 .0003 .0001 .0002 

.oos .007 .007 .005 .007 .003 .003 .005 

.02 .02 0 0 0 0 .015 .oos 

0 (.0003 0 0 0 0 0 0 

.001 .0015 <.001 <.001 <.001 .015 <.001 <.001 

.0007 .002 .003 .003 <.0005 .002 .002 .003 

.0015 .OOIS .0015 .001 .005 .001 .0015 0 

.0015 .003 .001 .0015 .0015 .0015 .002 .002 

.015 .015 .007 .007 .02 .007 .015 .015 

0 <.001 <.001 <.001 .(.001 .003 (.001 .001 

<.0001 <.OOOI <.0001 <.0001 <.0001 .0003 <.000 1 .0001 

0 .003 .001 .0015 .005 .015 .003 .007 

2Phengite 555 analyzed contained 8 percent admixed quartz. Structural formula computed on quartzfree basis. 

Data for Micas from Nevada 47 
p. 840). Thus, given the appropriate bulk composition, 

it appears that a thrust-fault environment is conducive 

to the crystallization of such micas. 

A c 

B D 

1 mm 

Figure 2. Camera Iucida drawings of four quartzite samples 

from which recovered micas were analyzed. A, sample 503: 

quartz and muscovite; 8, sample 555: quartz, very fine grained 

muscovite, and three grains offeldspar; C, sample 561: quartz, 

muscovite, and a trace of biotite; 0, sample 569: quartz, 

muscovite, and a few grains of potassium feldspar. 

In calculation of the structural formulas (table 1), 

the X-ray fluorescence values for Ba and Rb were used, 

as neither BaO nor Rb2

0 was sought by the chemist. 

The amounts of BaO present probably were lost during 

the gravimetric analyses, and in each case Rb20 probably was weighed with K2

0 (L.C. Peck, oral commun., 

1970). The very small amounts of SrO present either were 

weighed with CaO or lost. 

The structural formulas (table 1) are in rather good 

agreement with the ideal for dioctahedral mica: In all 

cases the octahedral occupancy is slightly greater than 

4.00, a feature common to most analyzed phengites 

(Mather, 1970, p. 258). The alkali contents tend to be 

slightly low, but agreement with the ideal 2.00 total is 

close considering the fact that a deficiency in the X-group 

totals is a persistent feature of muscovite analyses (Butler, 

1967, p. 253). 

The physical properties of the white micas analyzed 

are listed in table 3. Annealed fluorite was used as an internal standard for X-ray diffraction patterns which were 

used to give unit-cell parameters by least squares refine48 Contributions to Mineral Resources Research, 1984 

ment. The standard error listed for each unit-cell 

parameter varies slightly from sample to sample, apparently because of slight differences in the quality of the 

X-ray diffraction patterns. Agreement is good between 

the observed and calculated values for density and unitcell volume. 

One of the most important results of this work is 

the discovery of a sodic muscovite (549, table 1}, which 

contains 2.91 weight percent N~O, equivalent to 38 mol 

percent paragonite. This sodic muscovite coexists only 

with quartz and biotite, except for trace amounts of 

chlorite, garnet, and staurolite. Paragonite as a discrete 

mineral is absent, and this sodic muscovite is present 

where one might expect the pair paragonite and lowNa20 muscovite. The simplicity of this assemblage and 

the chemical data available for rock (Lee, Van Loenen, 

Brandt, and Doering, 1980}, muscovite (table 1), and 

biotite (table 4) may help to direct future experimental 

work on the stability of sodic muscovite. 

Biotites coexisting with four of the white micas in 

table 1 were recovered for analysis. As emphasized earlier, 

each of these biotites fractionated over a range of density when centrifuged in heavy liquids. Thus each of the 

analyses listed (table 4) applies to a bulk sample with the 

indicated density range. During work on these biotites, 

and on igneous biotites from the southern Snake Range, 

Nevada (Lee and Van Loenen, 1970), we have had chemical analyses made on the heavier and lighter biotite fractions present in different rock samples of about 1,000 g 

each. Our data show that in each sample the most significant compositional differences between the lighter and 

heavier biotite fractions involve the FeO and Fe20 3 contents. Invariably the heavier biotite fraction contains 

more FeO and less Fe20 3 than the lighter. For example, 

the Fe20 3 in sample 428 (table 4) is listed as 6.11 weight 

percent; the range might be about 5.2-7.0 weight percent 

Fe20 3• The equivalent range of FeO (22. 72 weight percent) might be about 23.7-21.7 weight percent, on the 

basis of our experience. 

X-ray powder patterns were obtained for a split of 

each of the analyzed biotites in table 4. Oriented and 

disoriented mounts of each sample were scanned from 

8 ° 29 to 65 ° 29 at a speed of 1 I 4 ° per minute. However, 

we were unable to refine the measured 29 values into 

meaningful unit-cell parameters, even by the method 

described by Dodge and others (1969, p. 263). 

The d06fJ peak provides a measure of b (Wones, 

1963). The (060) interplanar spacings of samples 428, 501, 

506, and 549 are 1.5428, 1.5433, 1.5424, and 1.5410 A, 

respectively. These results are in good agreement with 

those of Lee and Van Loenen (1970) and Dodge and 

others (1969}, who found the (060) interplanar spacings 

of biotites to be rather constant within a narrow range 

regardless of composition. 

The·basal (005) interplanar spacings of samples 428, 

501, 506, and 549 are 2.0145, 2.0077, 2.0085, and 2.0107 
Table 2. Mineral content of the rocks from which analyzed micas were recovered 1 

[Measured in weight percent; Tr., trace] 

Sample Quartz 

No. 2 Plagio- K-feld- Musco- ~iotite Chlorite Other 

clase spar vite 

427 92 2 0 4 2 Tr. 0.05 magnetite. 

428 69 2 0 21 8 Tr. 0.25 ilmenite. 

476 20 3 0 70 0 Tr. Illite?, chlorite, 0.05 hematite. 

478 70 <1 4 26 0 (3) 0.05 magnetite. 

501 35 

503 

505 

506 

511 

547 

548 

549 

555 

561 

569 

89 

69 

90 

82 

89 

88 

82 

62 

94 

63 

30 6 22 

0 

(1 

(1 

0 

(1 

0 

0 

3 

0 

(1 

0 

5 

0 

10 

(1 

0 

0 

5 

0 

10 

11 

25 

9 

8 

10 

12 

10 

30 

6 

27 

7 

0 

(1 

(1 

Tr. 

0 

0 

8 

0 

Tr. 

0 

Tr. 

None noted. 

Do. 

Do. 

Do. 

0.05 magnetite. 

<1 "maghemite." 

(1 "maghemite." 

Garnet, staurolite. 

None noted. 

Do. 

0.05 magnetite. 

1Muscovite, biotite, chlorite, garnet, staurolite, and probably "maghemite" are 

metamorphic in these rocks. All other minerals are detrital, though some quartz is 

recrystallized. 

2Data for samples 427, 501, 511, 555, and 561 from unpublished notes. Data for all 

other samples from Lee, Van Loerien, Brandt, and Doering (1980). 

3Blank, no data. 

A, respectively. Muller (1966), Dodge and others (1969), 

and Lee and Van Loenen (1970) all found the (005) spacing to decrease with increase of fluorine. Data from the 

present study are too few for a good check of this 

relation. 

STABLE-ISOTOPE DATA 

Stable-isotope data and previously reported K-Ar 

ages are presented (table 5) for most of the chemically 

analyzed biotites. Taylor (1977) has emphasized that 

metamorphic rocks with oD values outside the "normal" 

range of -50 to -85 permil tend to be confined to unique 

geologic environments. Taylor has also noted that lowering of oD values commonly correlates with resetting of 

K-Ar ages. The K-Ar ages of our micas have been reset, 

and many of them have oD well below the "normal" 

range (table 5; fig. 3). 

Frey and others (1976, p. 173) and Taylor (1977, 

p. 511) found ~D muscovite-biotite values in the range 

15-35 permil, compatible with an equilibrium interpretation (Suzuoka and Epstein, 1976). Of the four white 

mica-biotite pairs included in our study (table 5; fig. 3), 

three have ~D values exceeding 50 permil, and one (428) 

has the very high value of 64 permil. On the other hand, 

our results are similar to those in the studies just cited 

in that white mica always has heavier isotopic values than 

biotite where the two coexist. Also the slopes of the lines 

connecting the white mica-biotite pairs (fig. 3) are similar 

to the slopes of such line~ based on the data of Frey and 

others (1976, p. 173). Thus, although the isotopic compositions of our samples vary over a wide range of values, 

apparently some sort of equilibrium relation was reached 

between each of the white mica-biotite pairs. 

The deuterium contents of natural waters can vary 

much from place to place and from time to time (Friedman, 1953; Friedman and others, 1964). Thus the pore 

waters present in the argillaceous sands may have differed 

from one body of sediment to another. Also, the isotopic 

composition of the hydrous minerals present in the 

original argillaceous materials may have differed from 

place to place. Thus we have no reason to assume that 

the original unmetamorphosed shelf sediments were pervaded by isotopically homogeneous waters. 

Data for Micas from Nevada 49 
1.11 

= 

n 

0 

~ s: 

s. 

cs· 

:I 

<ll 

0 

~ Table 3. Properties of analyzed white micas 

:r [Leaders (---), not determined] . ~ 

~ 

;:IC 

~ 

~ Sample No.-- 427 428 476 478 501 503 505 506 511 547 548 549 555 561 569 5; 

1"1 

~ 

<ll Optical properties: ;:IC 11.598 1 ~ 8 index 1. 608 1.594 1.607 1.603 1.610 1. 600 1. 598 1.612 1.606 1.610 1.596 1.600 1. 606 1.615 <ll 

~ 1390 2:!.._(-) 33° 43° --- 39° 29° 31° 37° 35° 32° 39° 39° 44° 31° 30° ~ 1"1 

~=- Density (g/cm3): ... 

~ Observed (±.01) 2.89 2.88 2.88 2.89 2.88 2.89 2.89 2.88 2.90 2.90 2.90 2.87 2.88 2.88 2.89 

= 

~ Calculated 2.893 2.875 --- 2.889 2.883 2.887 2.891 2.875 2.896 2.891 2.883 2.870 2.882 2.879 2.892 

Cell parameters observed: 

~ (A) 5.203 5.188 5.197 5.208 5.206 5.215 5.196 5. 195 5.212 5.204 5.206 5.175 5.210 5.204 5.219 

± .005 .002 .004 .002 .006 .003 .003 .002 .002 .003 .002 .004 .004 .002 .003 

..!?. (A) 9.032 9.007 9.021 9.040 9.024 9.055 9.046 9.036 9.054 9.025 9.047 8.973 9.048 9.040 9.066 

± .009 .004 .009 .004 .010 .008 .006 .005 .004 .005 .005 .006 .008 .003 .005 

.£. (A) 20.018 20.042 20.053 20.035 20.036 20.037 20.035 20.037 20.047 20.086 20.069 19.811 20.040 20.036 20.028 

± .010 .005 .010 .005 .010 .006 .007 .006 .005 .008 .006 .007 .010 .004 .007 

/3 95°44.8' 95°50.0' 95°44.3' 95°42.3' 95°43.8 I 95°47.3' 95°46. 5' 95°43.6' 95°51.8' 95°46.7' 95°46.6' 95°34.0' 95°45.8' 95°49.5' 95°44.9' 

Volume (A3) 936.0 931.7 935.5 938.6 936.5 941.4 936.9 936.0 941.7 938.5 940.5 915.6 940.1 937.9 942.9 

± 1. 0 0.5 1.0 0.5 1.0 0.8 0.6 0.6 0.5 0.5 0.5 0.7 0.9 0.4 0.6 

Cell volume calculated: 

Volume (A3) 936.97 930.10 --- 938.25 937.41 940.27 937.13 934.34 940.49 935.40 934.91 915.70 938.25 937.49 943.48 

-

1Value uncertain because the mineral is very finely crystalline. 
Table 4. Analytical data for biotites recovered from late Precambrian and Early Cambrian metasedimentary rocks 

[Chemical analyses by E.L. Munson Brandt. Quantitative Ba, Rb, and Sr analyses by W .P. Doering. Semiquantitative spectrographic analyses 

by L.A. Bradley. Leaders(---), not determined] 

Sample No. 501 549 

Physical properties of analyzed biotites 

Density: 

Average---

Range-----

Refractive 

index (b)-

2:!_(- )-----=--

3.15 3.07 3.06 3.05 

3.14-3.16 3.03-3.11 3.02-3.10 3.01-3.09 

Chemical analyses (weight percent) 

Sio2 

Al 2o3 

Fe 2o3 

FeO 

MgO 

CaO 

Na 2o 

K2o 

H2o+ 

H2o 

no 2 

P2os 

HnO 

Cl 

F 

33.80 

17.91 

6. 11 

22.72 

3.90 

.06 

• 21 

8.83 

3.80 

. 12 

2. 13 

.04 

.04 

.04 

.30 

36.50 

16.48 

3.18 

18.81 

8.43 

.02 

0.20 

9.41 

3.30 

• 16 

2.97 

.03 

• 11 

.02 

.49 

Subtotal-- 100.01 100.11 

Less 0-- .14 .21 

Total-- 99.87 99.90 

36.16 

16.36 

6.43 

16.59 

7. 71 

.29 

.27 

8.70 

.20 

2.59 

1. 626 

so 

36.36 

19.22 

2. 11 

16.82 

10.42 

.oo 

.34 

9.07 

3.78 

.08 

1.48 

.00 

.os 

.01 

.20 

99.94 

.08 

99.86 

Quantitative X-ray fluorescence analyses (weight percent) 

Ba 

Rb 

Sr 

o. 1783 

.0558 

.0012 

0.0431 

.0690 

• 0011 

0.0295 

.0622 

.0008 

Inasmuch as most of the sample sites are spatially 

related to regional thrust faults, it is likely that the original 

oD values of many of the sedimentary rocks were lowered 

through isotopic exchange with meteoric waters circulating along the fault plane. Meteoric waters are the 

only abundant natural materials that are depleted in 

deuterium. However, meteoric waters are also depleted 

in 180, and the o180 values of most of our micas (table 

5, fig. 3) are fairly "normal." It is generally easier to 

change o0 than o180 in a system because of the oxygenbuffering capacity of rocks. Thus, the stable isotope 

results in table 5 imply rather low water:rock ratios during the isotopic exchange (Taylor, 1977). 

Sample No. 501 549 

Numbers of ions on basis of 24 (0, OH, F, Cl) 

{ 

pS+ 

Z Si 

AliV 

!

AlVI 

Fe3+ 

y Ti 

Mg 

Fe 2+ 

Mn 

{

Ca 

Na 

X K 

Ba 

Rb 

~H 

Cl 

0.01 } 

5.27 

2.72 

.57) .72 

.25 

.91 

2.96 

.o 1 

.06 

8.00 

5.42 

.01} 

1. 76 1.85 

.01 

.01 

3. 95} 

.15 4.11 

.01 

s.s7} 8.oo s.s3} 8.oo s.4s} 8.oo 

2.43 2.47 2.55 

.54) .48} .85) .37 .74 .24 

.34 5.58 .30 5.40 .17 5.71 

1.92 1.76 2.33 

2.40 2.12 2.11 

.01 --- .o 1 

.06oo} .08o7} .oo)10 

1.83 1.90 1.71 1.86 1.74 1.85 

.oo --- .00 

.01 --- .01 

3.36l --- 3. 78} 

.24 3.61 --- .10 3.88 

.01 --- .00 

Semiquantitative spectrographic analyses (weight percent) 

Be 

Co 

Cr 

Cu 

Ga 

Li 

Mo 

Nb 

Ni 

Pb 

Sc 

v 

y 

Yb 

Zn 

Zr 

0 

.oos 

.01 

.002 

.003 

.02 

.0005 

.002 

.015 

.003 

.001 

.02 

.003 

.0005 

0 

.01 

0.0002 

.003 

.015 

.002 

.005 

0 

.0007 

.001 

.01 

.002 

.002 

.015 

.0015 

.0003 

.OS 

.005 

0.0002 

.007 

.007 

.002 

.002 

0 

0 

.0015 

.01 

0 

.0005 

. 01 

0 

0 

0 

.002 

0.0002 

.007 

.015 

.003 

.003 

0 

0 

.002 

0 

0 

0 

.01 

.0015 

.001 

.02 

.003 

1sample recovered from Late Proterozoic schistose 

quartzite. All other samples recovered from Early Cambrian 

metasedimentary rocks. 

2structural formula calculated on basis of total 

cationic charge of 44. 

CONCLUSIONS 

We·still do not understand the timing and conditions of the metamorphic events that have affected the 

lower plate metaclastic rocks of eastern White Pine County, Nev., despite the many age, chemical, and stableisotope data that have been assembled for these rocks and 

their constituent micas. Nonetheless, these data may help 

to direct and place in perspective future studies of these 

rocks. Such future studies might include: (1) 40Ar/39Ar 

dating of selected micas from lower plate metaclastic 

rocks of eastern White Pine County; (2) additional UTh-Pb work on the metamorphic monazite that appears 

Data for Micas from Nevada 51 
to be widespread in the lower plate metaclastic rocks; (3) 

additional work such as the intensive study of metamorphic and structural features of the Hampton Creek area 

of the northern Snake Range (Rowles, 1982); and (4) 

fluid-inclusion work on the recrystallized quartz in 

samples such as 503 and 561 (fig. 2) to provide evidence 

bearing on pressures and temperatures of metamorphism. 

Table 5. Stable-isotope data and K-Ar ages for micas recovered from late Precambrian and Early Cambrian metasedimentary rocks 

[o180 and oD determinations by J.D. Gleason; leaders (---), not 

determined] 

Sample Rock type Mineral K-Ar age2 

No. (symbol) Ma.:!:_2o 

427 

428 

428 

476 

478 

501 

SOl 

503 

505 

506 

511 

511 

547 

548 

549 

549 

555 

561 

569 

p6 

p6 

p6 

p6str 

p66pm 

6p 

6p 

p66pm 

p66pm 

p66pm 

p66pm 

p66pm 

p66pm 

p66pm 

p66pm 

p66pm 

p66pm 

p6Spm 

p66pm 

White 63.6.:!:_2.4 

mica. 

Biotite 45.8+1.9 

White 61.7~2.3 

mica. 

--do--- 105+4 

--do--- 8 7 .4+2. 2 

Biotite 43.3+1.6 

White 51.1~1.4 

mica. 

--do--- 57.1+1.7 

--do--- 54.t+t. 7 

--do--- ss. 2::!1. s 

Biotite 

White 72.S..:t_2.0 

mica. 

--do--- 20.8+0.5 

--do--- 22. 7"+0. 5 

Biotite 4 7. t::!t. 6 

White 30.5.:!:_1.0 

mica. 

--do--- 141+5 

--do--- 36.9+0.8 

--do--- 66. 9::f). 5 

~18o, permil dD, permil 

(SMOW) (SMOW) 

+ 3.2 

+10.8 

+ 8.0 

+ 7.6 

+10.3 

+ 8.8 

+ 8.8 

+ 4.5 

+ 8.3 

+ 9.3 

+ 7.1 

+10.5 

+10.3 

+ 8.4 

+ 8.6 

-182.3 

-118.0 

-150.5 

-109.3 

- 58.9 

- 98.6 

-101.9 

-124.9 

- 62.5 

-167.4 

- 77.2 

- 51.1 

- 80.6 

-131.3 

- 77.4 

1samples 427 and 428 are the quartzite and mica schist, 

respectively, of Misch and Hazzard's (1962, p. 308, 309) A1 

of unit A. p€~m, Precambrian and Cambrian Prospect Mountain 

Quartzite. ~p, Cambrian Pioche Shale; p6str, Precambrian 

Strawberry Creek Formation of Misch and Hazzard (1962). 

2K-Ar ages from Lee, Marvin, and Mehnert (1980). 

52 Contributions to Mineral Resources Research, 1984 

-50 

-60 

-70 

-80 

-90 

§' o-1oo 

~ 

en 

~-110 

~ 

a: -120 w 

CL. 

ci -130 CQ 

-140 

-150 

-160 

-170 

8180, PER MIL (SMOW) 

Figure 3. Plot of c5 180 versus oD values for analyzed micas 

from metaclastic rocks of eastern White Pine County, Nevada. 

The large range of generally below "normal" oD values of these 

micas together with their fairly "normal" c5 180 values implies 

rather low water: rock ratios during isotopic exchange (Taylor, 

1977). Slopes of lines connecting coexisting white micas (circles) 

and biotites (X) are similar to slopes of such lines based on data 

of Frey and others (1976, p. 173). 
REFERENCES CITED 

Butler, B.C.M., 1967, Chemical study of minerals from the 

Moine schists of the Ardnamurchan area, Argyllshire, 

Scotland: Journal of Petrology, v. 8, p. 233-267. 

Dodge, F.C.W., Smith, V.C., and Mays, R.E., 1969, Biotites 

from granitic rocks of the central Sierra Nevada batholith, 

California: Journal of Petrology, v. 10, p. 250-271. 

Ernst, W.O., 1963, Significance of phengitic micas from lowgrade schists: American Mineralogist, v. 48, p. 1357-1373. 

Foster, M.D., 1956, Correlation of dioctahedral potassium 

micas on the basis of their charge relations: U.S. Geological 

Survey Bulletin 1036-D, p. 57-67. 

Foster, M.D., Bryant, B., and Hathaway, J., 1960, Iron-rich 

muscovitic mica from the Grandfather Mountain area, 

North Carolina: American Mineralogist, v. 45, p. 839-851. 

Frey, M., Hunziker, J.C., O'Neil, J.R., and Schwander, H.W., 

1976, Equilibrium-disequilibrium relations in the Monte 

Rosa granite, western Alps: Contributions to Mineralogy 

and Petrology, v. 55, p. 147-179. 

Friedman, Irving, 1953, Deuterium content of natural waters 

and other substances: Geochimica et Cosmochimica Acta, 

v. 4, p. 89-103. 

Friedman, Irving, and O'Neil, J .R., 1977, Compilation of stable 

isotope fractionation factors of geochemical interest, 

Chapter KK in Fleischer, Michael, ed., Data of 

geochemistry (6th edition): U.S. Geological Survey Professional Paper 440-KK, 12 p. 

Friedman, Irving, Redfield, A.C., Schoen, B., and Harris, 1., 

1964, The variation of the deuterium content of natural 

waters in the hydrologic cycle: Reviews of Geophysics, v. 

2, p. 177-224. 

Jackson, E.D., Stevens, R.E., and Bowen, R.W., 1967, A 

computer-based procedure for deriving mineral formulas 

from mineral analyses, in Geological Survey research 1967: 

U.S. Geological Survey Professional Paper 575-C, 

p. C23-C31. 

Lee, D.E., and Doering, W.P., 1974, Barium in hybrid granitoid 

rocks of the southern Snake Range, Nevada: U.S. Geological Survey Journal of Research, v. 2, no. 6, p. 671-675. 

___ 1980, Rubidium and strontium in hybrid granitoid rocks, 

southern Snake Range, Nevada: U.S. Geological Survey 

Professional Paper 1158-A, p. 1-7. 

Lee, D.E., Marvin, R.F., and Mehnert, H.H., 1980, A radiometric age study of Mesozoic-Cenozoic metamorphism in 

eastern White Pine County, Nevada, and nearby Utah: 

U.S. Geological Survey Professional Paper 1158-C, 

p. 17-28. 

Lee, D.E., Marvin, R.F., Stern, T.W., and Peterman, Z.E., 

1970, Modification of potassium-argon ages by Tertiary 

thrusting in the Snake Range, White Pine County, Nevada, 

in Geological Survey research 1970: U.S. Geological Survey 

Professional Paper 700-D, p. D92-D102. 

Lee, D.E., Stern, T.W., and Marvin, R.F., 1981, Uraniumthorium-lead isotopic ages of metamorphic monazite from 

the northern Snake Range, Nevada: ISOCHRON/WEST, 

no. 31, p. 23. 

Lee, D.E., and Van Loenen, R.E., 1970, Biotites from hybrid 

granitoid rocks of the southern Snake Range, Nevada, in 

Geological Survey research 1970: U.S. Geological Survey 

Professional Paper 700-D, p. Dl96-D206. 

Lee, D.E., Van Loenen, R.E., Brandt, E.L.M., and Doering, 

W.P., 1980, Composition of calcium-poor quartzites of 

late Precambrian and Early Cambrian age from eastern 

White Pine County, Nevada: U.S. Geological Survey Professional Paper 1158-B, p. 9-16. 

Mather, J.D., 1970, The biotite isograd and the lower 

_greenschist facies in the Dalradian rocks of Scotland: Journal of Petrology, v. 11, p. 253-275. 

Misch, Peter, and Hazzard, J.C., 1962, Stratigraphy and 

metamorphism of late Precambrian rocks in central northeastern Nevada and adjacent Utah: American Association 

of Petroleum Geologists Bulletin, v. 46, p. 289-343. 

Muller, G., 1966, Der Einfluss verschiedener Substitutionen auf 

die Gitterkonstanten von koexistierenden Biotiten und 

Muskowiten: Contributions to Mineralogy and Petrology, 

v. 13, p. 59-74. 

Peck, L.C., 1964, Systematic analysis of silicates: U.S. Geological Survey Bulletin 1170, 89 p. 

Rowles, L.D., 1982, Deformational history of the Hampton 

Creek Canyon area, northern Snake Range, Nevada: Stanford, Calif., Stanford University M.S. thesis, 80 p. 

Suzuoka, T., and Epstein, S., 1976, Hydrogen isotope fractionation between OH-bearing minerals and water: Geochimica 

et Cosmochimica Acta, v. 40, no. 10, p. 1229-1240. 

Taylor, H. P., Jr., 1977, Water I rock interactions and the origin 

of H20 in gr~nitic batholiths: Journal of the Geological 

Society of London, v. 133, p. 509-558. 

Velde, B., 1965, Phengitic micas-Synthesis, stability, and 

natural occurrence: American Journal of Science, v. 263, 

p. 886-913. 

Wones, D.R., 1963, Physical properties of synthetic biotites on 

the join phlogopite-annite: American Mineralogist, v. 48, 

nos. 11-12, p. 1300-1321. 

Data for Micas from Nevada 53 
Chapter E 

Metallogeny of Archean and 

Proterozoic Terranes in the 

Great Lakes Region-A Brief Overview 

By PAUL K. SIMS 

A description of major mineral deposits 

of known or potential future importance 

U.S. GEOLOGICAL SURVEY BULLETIN 1694 

CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 
CONTENTS 

Abstract 57 

Introduction 57 

General geology 57 

Archean gneiss terrane 58 

Archean greenstone-granite terrane 58 

Early Proterozoic epicratonic sequence 59 

Early Proterozoic magmatic terrane 60 

Ladysmith-Pembine belt 61 

Chippewa-Wausau belt 63 

Marshfield-Stevens Point belt 64 

Middle Proterozoic Midcontinent rift system 64 

Summary of tectonic evolution 64 

Metallogeny 65 

Deposits in island-arc crustal terranes 66 

Massive sulfide deposits 66 

Algoma-type iron-formations 67 

Nickel-copper deposits 68 

Gold deposits 68 

Deposits in continental crustal terranes 68 

Superior-type iron-formations 68 

Carbonate-hosted copper deposits 69 

Conglomerate-hosted uranium deposits 69 

Volcanic- and shale-hosted copper deposits 69 

Gabbro-hosted copper-nickel deposits 70 

Algoma-type iron-formation 70 

Summary of tectonic evolution and mineralization 70 

References cited 71 

FIGURES 

1. Map of southern margin of Canadian Shield showing Midcontinent rift 

system and edge of Phanerozoic platform cover 58 

2. Generalized geologic map of major Precambrian tectono-stratigraphic 

terranes in Lake Superior region 59 

3. Paleogeologic maps of Archean terranes in Great Lakes region 60 

4. Paleogeologic maps of Proterozoic terranes in Great Lakes region 61 

5. Generalized geologic map of northern Wi.sconsin, showing belts of magmatic 

terrane 62 

6. Map of Lake Superior region, showing major known and potential metallic 

mineral districts in Precambrian rocks 67 

TABLES 

1. Mineral deposits in island-arc crust, Lake Superior region 66 

2. Characteristics of massive sulfide deposits in the Wisconsin magmatic 

terrane 66 

3. Mineral deposits in continental crust, Lake Superior and Lake Huron 

regions 69 

56 Contributions to Mineral Resources Research, 1984 
CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 

Metallogeny of Archean and Proterozoic Terranes in the 

Great Lakes Region-A Brief Overview 

By Paul K. Sims 

Abstract 

Four major Precambrian tectono-stratigraphic terranes 

have been delineated in the Great Lakes region of the United 

States and Canada. Two of these-an Archean greenstonegranite terrane and an Archean gneiss terrane-have a partial 

epicratonic cover, which is a continental margin sequence of 

Early Proterozoic miogeoclinal rocks. The third is an Early Proterozoic eugeoclinal oceanic-arc assemblage (Wisconsin 

magmatic terrane) that mainly lacks an Archean basement, and 

the fourth is a Middle Proterozoic rift assemblage (Midcontinent rift system). The two Archean terranes were sutured 

together in Late Archean time, and the Early Proterozoic 

oceanic-arc assemblage was accreted to the North American 

continent about 1,850 Ma. The Midcontinent rift system is a 

major intracontinental rift (about 1,100 Ma) that aborted before 

significant crustal separation took place. 

Major known or potential ore deposits occur in all terranes except the Archean gneiss. These deposits can be 

grouped into two classes according to the tectonic environment in which they formed: island-arc crust and continental 

crust. 

Of the four types of major mineral deposits in island-arc 

crust, only one is common to both the Early Proterozoic 

Wisconsin magmatic terrane and the Archean greenstonegranite terrane: volcanic-hosted Cu-Zn type massive sulfide 

deposits. The massive sulfide deposits are in calc-alkaline and 

tholeiitic volcanic rocks having island-arc chemical affinities. 

The other deposits in island-arc crust-Algoma-type ironformations, ultramafic- and gabbro-hosted Cu-Ni deposits, and 

stratabound- and intrusion-hosted Au deposits-are virtually 

restricted to Archean rocks and, except for the ironformations, 

occur mainly in Canada. 

With respect to deposit types in continental crust, the Early 

Proterozoic epicratonic rocks of the Lake Superior region contain the vast Superior-type iron-formations and apparently 

subeconomic carbonate-hosted copper deposits. In the Lake 

Huron area of Canada, major conglomerate-hosted uranium 

deposits occur in the Huronian Supergroup, a continentalmargin accumulation like the miogeoclinal sequence in the 

Lake Superior region. The Middle Proterozoic rocks of the Midcontinent rift system have deposits containing chalcophile 

elements, which ultimately were derived from magmas ascending from the mantle. The principal developed deposits are 

the famous volcanic- and sedimentary-hosted copper deposits 

on the Keweenaw Peninsula. A major resource for the future 

is the large, marginal-grade gabbro-hosted Cu-Ni deposits of 

the Duluth Complex, northern Minnesota. 

INTRODUCTION 

The Great Lakes region is well known for its diverse 

Precambrian geology and for its vast iron, copper, and 

uranium resources. Not as well known is the presence of 

important-and as yet untapped-base-metal deposits, 

such as the gabbro-hosted copper-nickel sulfide deposits 

in the Duluth Complex, northern Minnesota, of Middle 

Proterozoic age, and the volcanic-hosted massive sulfide 

deposits in northern Wisconsin, of Early Proterozoic age. 

Knowledge of the geology of the region has advanced considerably during the past few years, and it is 

now possible to relate the major mineral deposits to 

specific tectono-stratigraphic terranes and, to a certain 

extent, to postulated plate tectonic regimes. A terrane, 

as defined in the region, has a unique stratigraphy, structure, metamorphism, plutonism, and age, and at least in 

part is fault-bounded. 

The purpose of this report is to discuss, in brief, 

the metallogeny of each of the terranes that has been 

delineated in the Great Lakes region, and to relate the 

mineral resources to inferred plate tectonic models. Only 

major mineral deposits of known or potential future importance are discussed, and emphasis is given to deposits 

in the United States. A few deposits known to occur in 

adjacent Ontario, Canada, but not in the United States 

are mentioned because of the possibility that such deposits 

could exist in this country. This report is a revision of 

an earlier one (Sims, 1976), as well as an elaboration of 

an oral presentation to a Society of Economic Geologists 

symposium at the 113th annual meeting of SME-AIME 

in Los Angeles, Calif., February 1984. 

GENERAL GEOLOGY 

The Great Lakes region is on the southern margin 

of the Superior province of the Canadian Shield (fig. 1). 

Metallogeny in the Great Lakes Region 57 
SUPERIOR 

Figure 1. Map of southern margin of Canadian Shield, showing location of Lake Superior and adjacent Lake Huron regions. 

Horizontal line pattern, Midcontinent rift system; dashed line, 

southern edge of Superior province; dot pattern, edge of 

Phanerozoic platform cover. Modified from King and Edmonston (1972) . 

The Precambrian rocks in the region are overlapped by 

platform strata of Paleozoic and Mesozoic age on the 

west, south, and east, and are irregularly covered by 

Pleistocene glacial deposits. 

Four major geologic terranes have been delineated 

in the Lake Superior region (fig. 2): (1) an Archean gneiss 

terrane; (2) an Archean greenstone-granite terrane; (3) 

an Early Proterozoic magmatic terrane (a eugeoclinal 

volcanic-plutonic sequence), most of which apparently 

lacks an Archean basement; and (4) a Middle Proterozoic 

intracontinental rift assemblage (Midcontinent rift 

system). An Early Proterozoic epicratonic, dominantly 

miogeoclinal sequence forms a partial cover overlapping 

the composite Archean basement. 

It is ,useful to discuss the sequential development 

of the individual terranes and their inferred tectonic environments by means of a series of simplified regional 

paleogeologic maps. 

Archean Gneiss Terrane 

The oldest crustal segment in the region is the 

Archean gneiss terrane (Morey and Sims, 1976). The inferred position of this terrane during the Archean, as 

reconstructed prior to Keweenawan rifting and based on 

its present-day relative position, is shown in figure 3A. 

Rocks of this terrane are exposed in southern Minnesota, 

principally in the Minnesota River Valley, and in the cores 

58 Contributions to Mineral Resources Research, 1984 

of mantled gneiss domes and uplifts in east-central 

Minnesota and northern Michigan; they are inferred to 

extend eastward into the southern part of the Lake 

Huron region (Sims and others, 1981). The westward 

extension into South Dakota and the southern limit during Archean time are uncertain. The terrane is composed mainly of varied migmatitic gneisses and amphibolite, of amphibolite- and granulite-metamorphic 

facies, and includes metapelites and granitic plutons 

(Grant, 1972). Migmatitic gneisses in the Minnesota River 

Valley record a long history of rock-forming events. They 

have U-Pb zircon and Rb-Sr isochron ages of about 

3,500 Ma (Goldich and Wooden, 1980), they were 

modified by addition of a granitic neosome about 3,000 

Ma, and they were metamorphosed at 3,050 and 2,600 

Ma, followed by a low-grade thermal event at 1,800 Ma. 

Granitic bodies of batholithic dimensions were intruded 

at about 2,600 Ma. Gneisses of comparable age and complexity are exposed in the Watersmeet dome in northern 

Michigan. A tonalitic augen gneiss and associated biotite 

gneiss have U-Pb zircon ages of :::::3,560 Ma, and a 

younger succession of interlayered amphibolite and 

biotite gneiss has a zircon age of :::::2,640 Ma (Sims and 

others, 1984). These rock types are intruded by a 

:::::2,600-m.y.-old leucogranite. Intense deformation and 

metamorphism during the Early Proterozoic orogenic 

event reset whole-rock and mineral Rb-Sr systems in the 

interval 1,800-1,750 Ma. 

Archean Greenstone-Granite Terrane 

Voluminous mafic to felsic volcanic rocks and 

related volcanogenic graywacke accumulated adjacent to 

the north side (present-day relative position) of the gneiss 

terrane (fig. 3B) in Late Archean time (:::::2, 700 Ma). They 

were intruded by abundant tonalite-granodiorite and 

granite bodies, many of batholithic dimensions. The 

depositional overlap of the Late Archean volcanic rocks 

onto the adjacent gneiss terrane, to the south, and the 

tectonism and plutonic stitching in the boundary zone 

(Great Lakes tectonic zone; Sims and others, 1980) 

sutured the greenstone-granite terrane to the older gneiss 

terrane. The greenstone-granite terrane (Superior province of Canadian usage) represents the single greatest addition of volcanic-plutonic rocks to the North American 

crust in Precambrian time. The mafic volcanic rocks are 

dominantly tholeiitic basalt; the lesser felsic volcanic 

rocks have both tholeiitic and calc-alkaline affinities 

(Campbell and others, 1982; Schulz, 1980). The tectonic 

environment in which the greenstone-granite complexes 

were formed is still debated, but the rocks have oceanic 

and island-arc compositional affinities. The granitoid 

rocks are mainly calc-alkaline, but relatively late tectonic and post-tectonic rocks typically are alkaline. The 
LAKE SUPERIOR 

100 200 Kl LOM ETEAS 

EXPLANATION 

... ~ .. .. 

·· . 

Phanerozoic 

Sedimentary rocks (= 570 Ma) 

Middle Proterozoic 

Midcontinent rift system (= 1,100 Ma) 

Anorogenic granite and syenite (= 1,500 Ma) 

Early Proterozoic 

Granitoid rocks (= 1,80Q-1,850 Ma) 

Wisconsin magmatic terrane (= 1,83Q-1,890 Ma) 

Epicratonic sequence (= 1,90Q-2,100 Ma) 

I Late Archean 

Greenstone-granite terrane (2,65Q-2,750 Ma) 

Late to Early Archean 

Gneiss terrane Gneiss terrane (2 ,600-3,550 Ma) 

-----~-- Reverse fault-Bar and ball on downthrown side 

High-angle fault 

NOTE: Archean rocks in central Wisconsin are 

possibly derived from an exotic terrane 

Figure 2. Generalized geologic map of Lake Superior region, showing major Precambrian tectono-stratigraphic terranes. Modified 

from Morey and others (1982). 

granitoid rocks have initial 87Sr/86Sr of less than 0.701 

(Sims and Peterman, 1981) and appear to have been derived either by fractional melting of basic rocks or by partial melting of graywacke having a short crustal residence 

time (Arth and Hanson, 1975). This segment of the crust 

stabilized by the end of Archean time (2,500 Ma). 

Early Proterozoic Epicratonic Sequence 

Following stabilization of the crust, two apparently separate asymmetrical sequences of sediments accumulated in the Great Lakes region along the southern 

margin of the Archean craton (fig. 4A): an older sequence 

Metallogeny in the Great lakes Region 59 
A 

NORTH 

DAKOTA 

300 KILOMETERS 

EXPLANATION 

Archean 

Greenstone-granite terrane 

Gneiss terrane 

- -- Known or inferred terrane boundary 

Figure 3. Paleogeologic maps of Archean terranes in the Great 

Lakes region (present coordinates). A, Archean gneiss 

(2,600-3,550 Ma); 8, Archean continent (:::::::2,500 Ma). Extent 

of terranes queried where uncertain. 

in the Lake Huron area, the Huronian Supergroup of 

Canada, and a somewhat younger sequence in the Lake 

Superior region, the Marquette Range Supergroup and 

its equivalents in Minnesota (Sims and others, 1981; 

Morey, 1983; Cannon and Gair, 1970). Volcanic rocks 

at the base of the Huronian have U-Pb zircon ages of 

about 2,500 Ma (K.D. Card, oral commun., 1984); 

sedimentation terminated by about 2.1 Ga (Sims and 

others, 1981). The time of deposition of the Marquette 

Range Supergroup is poorly known, but it probably 

spanned the interval2,100-1,900 Ma. Depositional patterns of rocks in the Marquette Range Supergroup, as 

well as in the Huronian Supergroup, reflect contrasting 

tectonic conditions in the northern and southern segments 

of the depositional basin, whose axis was approximately 

over and parallel to the Great Lakes tectonic zone (fig. 

4A). A relatively thin (<2 km) succession of sedimentary 

rocks, represented especially by the Animikie Group in 

60 Contributions to Mineral Resources Research, 1984 

northern Minnesota (Morey, 1973), was deposited over 

relatively nonstretched (stable) Archean greenstonegranite complexes on the north, whereas a much thicker 

(:::::6 km) succession of interlayered sedimentary and 

volcanic rocks accumulated on stretched Archean gneiss 

basement on the south. Sedimentation began with deposition of quartz-rich and carbonate rocks and continued 

with precipitation of major iron-formations. South of the 

Great Lakes tectonic zone, in the area of stretched crust, 

these miogeoclinal rocks were deposited in rather narrow, 

fault-bounded basins only tens of kilometers long and on 

intervening platforms (Larue and Sloss, 1980); north of 

the tectonic zone sedimentation occurred within a single 

(Animikie) basin several hundred square kilometers in 

area, which composed a southward-facing shelf. Later, 

the shelf foundered into deeper water, and a thick, 

southward-facing, flysch-like clastic wedge was deposited 

by southward-flowing turbidity currents. In Michigan, 

abundant volcanic rocks were intermittently extruded 

through stretched crust during accumulation of the 

deeper water sediments. Chemically, the volcanic rocks 

in northern Michigan are largely bimodal, with abundant 

tholeiitic basalt and minor high-K20 rhyolite. The 

basalts show strong iron enrichment and high Ti02 

and incompatible-element contents (Fox, 1983); they are 

compositionally similar to continental rift basalts, 

such as those of the Keweenawan in Minnesota (Green, 

1983). 

There is now a general consensus that the Early 

Proterozoic rocks in the Lake Superior region accumulated on a south-facing, rifted, passive continental 

(Atlantic-type) margin undergoing north-south extension 

(Sims and Peterman, 1984), as proposed earlier by Cambray (1978) and Larue and Sloss (1980) (fig. 4A). 

Early Proterozoic Magmatic Terrane 

The relation of the Early Proterozoic sequence in 

Wisconsin, dominated by volcanic and plutonic rocks, 

to the epicratonic sequence (Marquette Range Supergroup) in Michigan has been uncertain, because of poor 

exposures, sparse geologic mapping, and lack of chemical 

data on the volcanic and plutonic rocks of Wisconsin. 

A recently completed regional geologic map (Morey and 

others, 1982) and new chemical data (Greenberg and 

Brown, 1983; Schulz, 1984), however, support an 

oceanic-arc complex-origin for the magmatic terrane and 

the subsequent accretion to the continent on the north 

during the Penokean orogeny (:::::1 ,850 Ma) (fig. 4B). This 

terrane is informally called the Wisconsin magmatic terrane in this report. 

The Wisconsin magmatic terrane consists of abundant mafic to felsic volcanic rocks and granitoid rocks 

ranging from tonalite to granite, and lesser, intercalated 

sedimentary rocks. An Archean basement is present in 
NORTH 

DAKOTA 

300 KILOMETERS 

r-.-----.-< .. :-:>: ·- 1 .·· . .. ·._. ·,·.·. 

' ' 

' 

EXPlANATION 

Middle Proterozoic 

Midcontinent rift system 

Early Proterozoic 

Wisconsin magmatic terrane 

Marquette Range Supergroup and correlative rock; 

Archean basement exposed locally 

Late Archean 

Greenstone-granite terrane 

Late to Early Archean 

Gneiss terrane 

- --- Contact-Approximately located; dashed where now 

covered by water; queried where uncertain 

---- Fault-Dashed where approximate; queried where uncertain 

~ Direction of tectonic transport 

_.;;.;GL"""T-.Z - Great Lakes tectonic zone; boundary between Archean 

terranes; extension beneath Marquette Range 

Supergroup shown as dashed line 

Figure 4. Paleogeologic maps of Proterozoic terranes in the Great Lakes region (present coordinates). A, Early Proterozoic 

epicratonic sequences (1,900-2,500 Ma) deposited on a rifted continental margin; 8, Penokean orogeny (::::::1,850 Ma) resulting 

from collision between Wisconsin magmatic terrane and continen,tal margin; C, Midcontinent rift system (::::::1, 100-1,225 Ma). 

central Wisconsin (fig. 5). Whether this Archean crustal 

segment was rifted from the continental margin and later 

returned or whether it originated elsewhere is not known. 

The terrane is bounded on the north by the Niagara fault. 

It appears to extend westward across the Midcontinent 

rift system into Minnesota and northwestern Iowa 

(Anderson and Black, 1983), as shown diagrammatically in figure 4B. 

The Wisconsin magmatic terrane is tentatively 

divided into three sub-terranes (fig. 5), or belts, each of 

which has internal characteristics that distinguish it from 

adjacent terranes. The distinguishing characteristics are 

rock types, stratigraphy, structure, metamorphism, 

plutonism and, to some extent, age. 

Ladysmith-Pembine Belt 

The Ladysmith-Pembine belt, the northernmost 

sub-terrane in Wisconsin, is arc-shaped, about 80 km 

wide, and more than 300 km long. Geophysical data 

indicate that the belt extends at least another 100 km 

southeastward. The belt is composed mainly of (1) 

metamorphosed basalt-andesite-rhyolite, (2) an apparently continuous basement of Early Proterozoic gneiss having a combined volcanic-plutonic protolith, and (3) 

intrusive rocks ranging in composition from gabbro to 

granite. In the best exposed area, in northeastern Wisconsin (Sims and others, 1985), a thick (:::::1 km) succession 

of metamorphosed shallow-water dolomite, quartzite, 

carbonaceous shale, and calc-silicate rocks underlies the 

volcanic rocks. Except for the 1,835-m.y.-old Athelstane 

batholith in northeastern Wisconsin (fig. 5) and a few 

smaller plutons in northwestern Wisconsin, the granitoid 

rocks are closely associated with the Early Proterozoic 

gneiss in the cores of domes, having been intruded mainly 

along the core-cover contact in the domes. Geologic and 

geochronologic data from northeastern Wisconsin indicate that the gneiss and overlying volcanic rocks are 

:::::1,865 Ma old; the intrusive rocks range in age from 

:::::1,860 to 1,835 Ma. Younger epizonal granite, dated at 

1, 760 Ma, locally intrudes the older granitoid rocks in 

the domes. 

Metallogeny in the Great Lakes Region 61 
= 

. 

. 

. 

. 

50 

EXPlANATION 

LADYSMITH-PEMBINE BELT 

Early Proterozoic 

Granite (1,760 Ma) 

Granodiorite-granite (= 1,835 Ma) 

Dominantly metavolcanic rocks (= 1,865 Ma) 

Gneiss, amphibolite, and tonalite-granite (1,835-

1,865 Ma) 

CHIPPEWA-WAUSAU BELT 

Middle Proterozoic 

Anorogenic syenite (1,518 Ma) and granite (1,485 Ma) 

Early Proterozoic 

Granitic gneiss (= 1,835 Ma) 

Granodiorite-granite (= 1,840-1,860 Ma) 

Dominantly metavolcanic rocks; includes tonalite 

gneiss (= 1,840-1,860 Ma) 

X 

. . . 0." . • .. o_ ~ , . g •• • • • Q •• o .. 

0 . Q 

. . • 0 .· 

... a·. a. 

100 KILOMETERS 

MARSHFIELD-STEVENS POINT BELT 

Early Proterozoic 

Granite (= 1,835 Ma) 

Metamorphosed volcanic rocks, carbonaceous 

schist, quartzite, and chert 

Late Archean 

Gneiss and amphibolite (= 2,515-2,800 Ma) cut by 

foliated tonalite (= 1,830-1,890 Ma) and granitic 

gneiss (= 1,870 Ma) 

Inferred boundary of belt-Short dashed where concealed 

High-angle reverse fault-Hachures on down thrown side 

Fault-Showing apparent relative horizontal movement 

High-angle fault-Relative movement not known 

Massive sulfide deposit 

1, Thornapple 

2, Flambeau 

3, Pelican River 

4,Crandon 

Figure 5. Generalized geologic map of northern Wisconsin, showing three sub-terranes (belts) of magmatic terrane. Modified 

from Morey and others (1982); age data mainly from W. R. Van Schmus, written commun., 1984. 

62 Contributions to Mineral Resources Research, 1984 
The volcanic rocks in the Ladysmith-Pembine belt 

(Schulz, 1984) are dominantly calc-alkaline and are 

enriched in LIL (large-ion lithophile) elements 

([La/Yb]N = 2.5-9.4) but are depleted in high-fieldstrength elements (Hf, Zr, Ta, and others). They are 

similar to volcanic sequences found in recent island-arcs 

(such as N.ew Hebrides and Japanese arcs). In contrast, 

basalt (Quinnesec Formation) from the same belt in 

northeastern Wisconsin is tholeiitic and strongly depleted 

in LIL elements ([La/Yb]N=0.09-0.89). This basalt is 

compositionally similar to recent back-arc basin basalts 

(for example, Lau Basin) and island-arc tholeiites (for example, Scotia arc). 

The Early Proterozoic gneiss and volcanic rocks are 

polydeformed and metamorphosed. The major deformation, regional in scope, produced tight folds whose axial 

planes form a gentle, arcuate pattern that is subparallel 

to the trace of the Niagara fault (fig. 5). The folds have 

steep limbs and steep plunges. This deformation was accompanied by greenschist facies metamorphism and, 

locally in the western part of the belt, by low-amphibolite 

facies. At places, younger northeast-trending tight to 

open folds produced fold-interference domal structures 

cored by the Early Proterozoic gneisses; these domal 

structures were further modified by diapirism, which produced mylonite and a steep-stretching lineation in dome 

boundaries. The gneiss domes have a surrounding annular, thermal metamorphic aureole, which reached middle amphibolite facies (Bayley and others, 1966). 

The position of the southern margin of the 

Ladysmith-Pembine volcanic-plutonic belt is ill defined 

and is partly based on interpretation of aeromagnetic and 

gravity data; presumably, this margin is a major fault. 

Chippewa-Wausau Belt 

An arcuate belt of volcanic and pb~tonic rocks to 

the south of the Ladysmith-Pembine belt (fig. 5) is comparable in width to the northern belt but is truncated on 

the east by the Wolf River batholith (:::::1 ,500 Ma; Van 

Schmus and others, 1975). The belt differs from the belt 

on the north in that it consists of two distinct, major 

volcanic successions of different lithology, metamorphic 

grade, and structure (LaBerge and Myers, 1984). A possible third, younger volcanic succession is present locally 

in the Wausau area. 

The older volcanic succession consists of interlayered quartzofeldspathic gneiss and schist (felsic 

volcanic rocks) and amphibolite (mafic volcanic rocks), 

which are intruded by abundant tonalite-granodioriteprobably entirely subvolcanic intrusions. Intrusion breccias are commonly associated with the intrusions. Chemically, the rocks are tholeiitic and calc-alkaline, having 

overall island-arc affinity. This succession apparently 

forms the basement in this volcanic belt. It is metamorphosed mainly to amphibolite facies and is deformed on 

west- to northwest-trending fold axes; the folds are tight, 

plunge steeply, and have steep limbs. 

The second, younger succession consists of basalt, 

andesite, and rhyolite, and has large subvolcanic intrusions of quartz monzonite composition and some intercalated conglomerate, graywacke, arkose, and quartzite 

(LaBerge and Myers, 1983). The conglomerate clasts include granitoid rocks. The rocks are widely distributed 

and now apparently compose scattered patches in fault 

contact with the older succession. The rocks are deformed on northeast fold. axes that plunge moderately to the 

northeast or southwest; they have a penetrative schistosity 

and are mainly metamorphosed to middle or upper 

greenschist facies. Although regional in scope, the deformation was remarkably inhomogeneous. It was intense 

in certain northeast-trending ductile deformation zones, 

such as the Milan shear zone (Sims and Peterman, 1983, 

fig. 3), where strain indicators show as much as 40 percent shortening (Palmer, 1980), whereas it was relatively 

mild in other areas. Metamorphism reaches low amphibolite facies in these deformation zones. 

The third, youngest succession consists mainly of 

rhyolite and volcanogenic graywacke that occupies an 

area of about 12 km2 in the vicinity of Wausau. 

LaBerge and Myers (1984) described welded tuff, flowbanded rhyolite and rhyolite breccia, lahars, and interbedded sandstone and conglomerate-indicative of 

probable subaerial accumulation-from the succession; 

it is possibly an erosional remnant of a rhyolite plateau 

or caldera complex. The rocks are mildly metamorphosed 

(low greenschist facies), lack penetrative structures, and 

are deformed on gently plunging northeast-trending fold 

axes. 

In addition to the subvolcanic intrusive rocks 

associated with successions one and two in this belt, a 

quartz monzonite pluton characterized by dotty biotite 

intrudes the two older successions in an area west of 

Wausau (Cherokee Granite of LaBerge and Myers, 1983). 

The most abundant intrusive rocks, however, are pink 

leucocratic granite bodies that intrude and apparently are 

coeval with succession three. Judged from the mapping 

of LaBerge and Myers (1983), the granite occurs as several 

separate(?) plutons, each of which is a few tens of square 

kilometers in area. The granite has moderately high 

K20/Na20 as compared with the intrusive rocks in the 

Ladysmith-Pembine belt. 

The southern margin of the belt is a ductile deformation (shear) zone along the Eau Pleine River, south 

of Wausau; this zone is marked by local mylonite, small 

bodies of ultramafic rocks, and a prominent northwestto west-trending aeromagnetic anomaly (LaBerge and 

Myers, 1983). 

Metallogeny in the Great Lakes Region 63 
Marshfield-Stevens Point Belt 

The third, most southerly exposed sub-terrane, the 

Marshfield-Stevens Point belt (fig. 5), differs from the 

other two belts in possibly having a continuous Archean 

gneiss basement; it also has a distinctive deformation style 

(Maass, 1983) characterized by steeply plunging fold axes 

and mineral lineations. Maass has recognized two main 

generations of folds: isoclinal F1 folds and tight to very 

open F3 folds, which are coaxial. 

Archean gneiss and amphibolite are exposed in the 

southern and western parts of the terrane, in major river 

valleys. The gneisses have U-Pb zircon ages of about 

2,800 Ma (Van Schmus and Anderson, 1977; DuBois and 

Van Schmus, 1978). The Archean gneisses are intruded 

by abundant, generally small bodies of tonalite, granodiorite, and granite, and all but the granite are deformed 

by F1 and F3• A tonalite near Stevens Point has a U-Pb 

zircon age of 1,892±9 Ma (W. R. Van Schmus, written 

commun., 1984). East of Marshfield (fig. 5), the Archean 

gneiss is overlain by a succession of mafic to felsic 

volcanic rocks that have numerous interbeds of ferruginous chert, white quartzite, conglomerate, and carbonaceous shale (tuff?). The rocks are middle greenschist 

facies and are folded into a steeply inclined syncline. 

Clearly, they are Early Proterozoic in age because they 

have the same steep penetrative lineation as the dated 

tonalite. Maass, Medaris, and Van Schmus (1980) have 

demonstrated that the deformation took place between 

about 1 ,880 and 1 ,825 Ma. During the deformation, older 

structures in the Archean gneisses were partly transposed 

(Maass, 1983), but the extent to which the structures were 

transposed is equivocal. 

Middle Proterozoic Midcontinent Rift System 

The youngest major geologic terrane in the region, 

the Keweenawan (formed 1,100 Ma) Midcontinent rift 

system, is an intracontinental assemblage of igneous and 

sedimentary rocks that formed in a rift that aborted 

before significant crustal separation was achieved (fig. 

4C). The rocks are dominantly (1) bimodal basalt and 

rhyolite, which occupy a central horst; (2) gabbroanorthosite complexes (Weiblen and Morey, 1980; 

Weiblen, 1982) which were intruded along the unconformity along the margins of the rift between Archean 

and (or) Early Proterozoic rock and the younger 

Keweenawan lavas; and (3) red-bed sedimentary rocks. 

The basalt is an aluminum-rich olivine tholeiite having 

a composition typical of large plateau lava provinces 

(Green, 1983). Chemically, the olivine basalt resembles 

ocean ridge basalt, but it differs in that it has slightly lightrare-earth-enriched REE patterns and contains high 

alumina. Large cumulate bodies, such as the Duluth 

64 Contributions to Mineral Resources Research, 1984 

Complex in Minnesota, are the dominant intrusive rocks 

associated with the lavas; but basaltic dike swarms, 

diabase sills, granophyre, and granite are also present. 

The sedimentary rocks consist of an older, well-indurated 

succession in basins within the medial part of the rift 

system, and a younger, less indurated succession along 

the rift margins, in half-grabens. The well-indurated 

sedimentary rocks include the Copper Harbor Conglomerate, which formed as volcanism waned, and the 

succeeding Nonesuch Shale, which was deposited in a 

large anoxygenic lake. 

SUMMARY OF TECTONIC EVOLUTION 

The geologic history of the region during the 

Archean is speculative. Knowledge of the original size and 

extent of the 3,500 Ma gneiss terrane is fragmental, but 

clearly it formed a protocontinental mass of at least 

several hundred square kilometers. Quite possibly, it was 

continuous with the ancient gneisses in Labrador and 

West Greenland. The principal protolith of the gneisses 

in the Lake Superior region probably was volcanic rocks 

and associated intrusive rocks (Sims and others, 1984a), 

but pelitic rocks are a component of the gneisses in the 

Minnesota River Valley (Grant, 1972) and, if the gneisses 

in West Greenland were a part of the terrane, ironformation also was a locally significant component 

(Windley, 1977). Thus, except perhaps for tectonic style, 

it is probable that geologic processes operative in Early 

Archean time were not appreciably different from those 

of today. 

About 800 m.y. later, vast quantities of submarine 

mafic to felsic lavas and associated pyroclastics and 

volcanogenic graywackes-with attendant voluminous 

granitoid plutons-were added in repeated cycles along 

the northern margin (present coordinates) of the older 

sialic mass. The volcanic-sedimentary pile and associated 

granitoid rocks were developed in 50-100 m.y., about 

2, 700 Ma. Evidence is lacking for a sialic crust beneath 

the greenstone belts in the Great Lakes region, and the 

rocks are presumed to have accumulated on oceanic crust 

(Sims and Peterman, 1981). The greenstone-granite complexes were formed in an orogenic environment, probably 

analogous to modern oceanic-arc complexes. Something 

of the nature of the boundary (suture) between the two 

Archean crustal segments (Sims and others, 1980) has 

been determined by a recent seismic reflection study 

(Gibbs and others, 1984). The study showed that a 

shallow north- to northwest-dipping reflection zone, 

among many others, projects to the surface near the trace 

of the boundary, as mapped previously (Sims and others, 

1980). The shallow-dipping reflection zones are interpreted by Gibbs and colleagues as probably being thrust 

faults resulting from a Late Archean collision between 
the greenstone-granite terrane (Superior province) and the 

older gneiss terrane on the south; possibly they are 

mylonite zones. 

During the interval between stabilization of the 

crust-termination of the Archean, about 2,500 Maand deposition of the Early Proterozoic rocks of the Marquette Range Supergroup and its equivalents in 

Minnesota-probably beginning about 2,100 Ma-the 

tectonic regime in the Lake Superior region was 

dominantly extensional. Abundant northwest-trending 

gabbroic dikes of probably two ages transect the Archean 

basement in northern Minnesota (Southwick and Day, 

1983) and predate the Animikie Group and, possibly, all 

of the Early Proterozoic sedimentary succession. An older 

set is about 2,200 m.y. old (K-Ar method; Hanson and 

Malhotra, 1971 ), and a younger set has a Rb/Sr isochron 

age of 2,120±67 Ma (Beck and Murthy, 1982). In the 

Lake Huron area, mafic dikes about 2.1 Ga cut the Huronian Supergroup, indicating that this sequence is mostly, if not entirely, older than the Marquette Range 

Supergroup. 

Extension culminated in the Lake Superior region 

with deposition of a thick (in part >6 km) wedge-shaped 

succession (Marquette Range Supergroup) of dominantly craton-derived clastic sediments (fig. 4A). The Early 

Proterozoic basin formed over and along the profound 

structural zone (Great Lakes tectonic zone) marked by 

the suture between the two Archean basement terranes; 

a clastic wedge was also deposited over the rifted continental margin to the south as a result of passive-margin 

subsidence. In Michigan and northwest Wisconsin, abundant bimodal volcanic rocks accumulated with the clasticwedge sediments over rifted gneissic crust. 

A recent tectonic model of Schulz and others (1984) 

suggests the following. The breakup of the continent as 

a consequence of Early Proterozoic rifting led to oceanic 

spreading, northward-facing subduction along the 

southern margin of the rifted fragment, and the formation of a volcanic arc (Chippewa-Wausau volcanic belt, 

fig. 5). Continuing consumption of oceanic crust led to 

collision of the arc with an Archean microcontinent(?) 

on the south (now within the Marshfield-Stevens Point 

belt). Upon reversal in the polarity of subduction, a 

second arc terrane (Ladysmith-Pembine belt) was developed above the subduction zone, and subsequent closing of the ocean basin resulted in collision of the arc with 

the continental margin sequence and Archean crust on 

the north. As noted earlier, it is not known whether the 

Archean microcontinent(?) associated with the 

Marshfield-Stevens Point belt is a rifted fragment from 

the North American continent or an exotic block. 

The sequence of Early Proterozoic events outlined 

above is based on age relations established in the field. 

The actual time span from crustal breakup to suturing 

of the ocean-arc complexes with the continental margin 

is estimated to exceed 50 m.y., and it may well have been 

longer. 

Because of moderately good exposures of rocks 

composing the Midcontinent rift system in the vicinity 

of Lake Superior, its pronounced gravity and magnetic 

signatures, and a lengthy period of intensive study (see 

Wold and Hinze, 1982), this rift system has been well 

documented as a major intracratonic rift. The actual 

nature and stages of development of the rift are controversial, but in the main, eruption of voluminous flood 

basalts-apparently in nine separate plateaus-was 

followed by felsic volcanism, subsidence, and sedimentation. At some stage during sedimentation, compressional tectonism commenced, producing major high-angle 

strike faults with reverse displacement on either side of 

a central horst. Sedimentation and faulting then continued concurrently for some time (Green, 1983). Van 

Schmus and others (1982) concluded that rifting and 

associated igneous activity began 1,200-1,225 Ma, peaked 

at 1,100 Ma, and ceased shortly thereafter; but widely 

disparate published radiometric ages suggest that evolution of the rift system may have spanned a longer interval of time. 

Since sometime in the Late Proterozoic, the region 

has been stable, and except for the effects of Pleistocene 

glaciation, it apparently has been in isostatic equilibrium. 

METALLOGENY 

Major mineral deposits occur in each of the terranes 

except the Archean gneiss. Iron-formation has overwhelmingly been the principal source of metallic ore in 

the Lake Superior region, the total production having a 

value of about $20 billion. Copper has also been an important metal produced in this region, the output coming entirely from the Keweenaw Peninsula, Michigan. 

Other major types of deposits, as yet unmined, are 

volcanic-hosted massive sulfide deposits in the Early Proterozoic Wisconsin magmatic terrane and gabbro-hosted 

copper-nickel deposits in the Middle Proterozoic 

(Keweenawan) Duluth Complex, Minnesota. Uranium 

has been by far the principal ore produced in the Lake 

Huron region. Without exception, the major deposits 

were formed synchronously with, or shortly after, formation of the enclosing rocks, although secondary processes of enrichment were important in deposits such as 

the iron-formations. 

The known or potential ore deposits can be grouped 

into two classes according to the tectonic environment 

in which they were formed: island-arc crust and continental crust. This format is used in describing the mineraldeposit types below. 

Two terranes in the Lake Superior region are underlain by former island-arc, or simatic, crust: the Early 

Metallogeny in the Great Lakes Region 65 
Proterozoic Wisconsin magmatic terrane and the Archean 

greenstone-granite terrane. The terranes underlain by continental, or sialic, crust are the Archean gneiss terrane 

and its partial cover, the Early Proterozoic epicratonic 

sequence, and the Midcontinent rift system. The Midcontinent rift system is a major intracratonic feature, 

bounded by continental crust, but its contained mineral 

deposits are mainly chalcophile elements ultimately derived from the Earth's mantle. The deposits, therefore, 

have at'finities with those formed in island-arc and oceanic 

crusts. Except for a rather small iron-formation deposit 

in central Wisconsin, the Archean gneisses are virtually 

devoid of mineral deposits of economic value. 

Deposits in Island-Arc Crustal Terranes 

Four types of major mineral deposits occur in 

island-arc crust, as noted in table 1, only one of which 

is common to both the Early Proterozoic Wisconsin 

magmatic zone and the Archean greenstone-granite terrane: volcanic-hosted Cu-Zn type massive sulfide 

deposits. The other major deposits apparently are 

restricted to Archean rocks. The metallogeny is dominated by siderophile elements of the greenstone 

association. 

Table 1. Mineral deposits in island-arc crust, Lake Superior 

region 

[do, same as above entry] 

Deposit type Terrane Age 

Volanic-hosted Cu-Zn Greenstone-granite Late Archean. 

type massive sulfide. Wisconsin magmatic Early Proterozoic. 

Algoma-type iron- Greenstone-granite Late Archean. 

formation. 

Ultramafi~- and gabbro- ----do------------ Do. 

hosted Cu-NL 

Stratabound- and in- ----do------------ Do. 

trusion-hosted gold. 

Massive Sulfide Deposits 

At least four significant massive sulfide deposits are 

known from the Wisconsin magmatic terrane (Afifi and 

others, 1984); no potentially minable massive sulfide 

deposits have been reported from the United States segment of the Archean greenstone-granite terrane (Superior 

province), but several world-class deposits are known in 

adjacent Ontario, Canada (Franklin and Thorp, 1982), 

most notably the incredibly rich Kidd Creek deposit. 

The known massive sulfide deposits in northern 

Wisconsin are in the Ladysmith-Pembine belt of the 

Wisconsin magmatic terrane (fig. 5). The deposits belong 

66 Contributions to Mineral Resources Research, 1984 

to the primitive type of Zn-Cu massive sulfide deposits 

of Hutchinson (1980). Characteristics of the deposits are 

listed in table 2. The Crandon deposit is by far the largest 

and most extensively studied of the four deposits (May 

and Schmidt, 1982). This deposit is a steeply dipping, 

tabular, stratabound body composed of massive sulfide 

ore and an underlying zone of stringer ore. The tabular, 

massive sulfide ore contains abundant pyrite and lesser 

sphalerite, chalcopyrite, and galena, as well as trace 

amounts of tetrahedrite-tennantite and arsenopyrite. It 

occurs within a unit of fine-grained pyritic tuff and 

associated sedimentary rocks. The footwall stringer ore, 

which makes up about half of the reserves, is composed 

of pyrite and chalcopyrite, and has subordinate 

sphalerite, galena, and arsenopyrite. It occurs in brecciated, silicified, and sericitized dacite and andesite 

pyroclastic rocks. The hanging-wall rocks are tuffs, mainly of dacitic composition. Within the oxidized zone of 

the subcrop, the primary sulfides have been altered by 

supergene processes to limonite. In contrast, the 

Flambeau deposit at Ladysmith has an extensive 

supergene-enriched zone, which is preserved under a local 

capping of Cambrian sandstone (May, 1977). The host 

rocks for all the deposits are calc-alkaline volcanic rocks 

having overall island-arc chemical affinities. 

Table 2. Characteristics of massive sulfide deposits in the 

Wisconsin magmatic terrane 

[Data modified from Mudrey, 1979; na, not available; Tr., trace] 

Reserves 

Deposit (millions Zn Cu 

of tons) 

(percent) 

Flam beau 

and 

Thornapple. 4-6 Tr. 35-40 

Crandon- 83 1. 1 

Pelican 

River. 2.2 4.5 

Estimated 

Pb Au Ag life 

in years 

(oz/ton) 

Tr. 0.05 0.53 10-20 

0.4 .03 .01 25 

Tr. Tr. Tr. na 

Although exploration has been extensive and is continuing in the Archean greenstone-granite terrane of 

northern Minnesota, potentially minable deposits have 

not been discovered, so far as known. The deposits can 

be expected to have the general characteristics listed by 

Franklin and Thorpe (1982) for massive sulfide deposits 

in the Superior province of Canada. A major attempt to 

characterize REE (rare-earth-element) patterns of felsic 

volcanic rocks associated with Cu-Zn type massive sulfide 

deposits in the Superior province, Canada, has indicated 

that some of these rhyolites are of tholeiitic affinity; they 

form the felsic members of a bimodal volcanic sequence, 

and are commonly underlain by subvolcanic intrusions. 
Other ore-associated rhyolites are calc-alkaline (Campbell and others, 1982). The rhyolites of tholeiitic affinity 

are characterized by relatively flat chrondite-normalized 

REE patterns, strong Eu anomalies, and high 

incompatible-trace-element contents (Lesher and others, 

1981). This technique for determining the chemical affinity of ore-associated rhyolites is a potentially powerful region.al prospecting tool in greenstone-granite 

terranes. 

Algoma-Type l.ron-Formations 

Algoma-type iron-formations, originally defined by 

Gross (1965), have been important sources of iron ore 

from the Archean greenstone-granite terrane in northern 

ONTARIO 

--\._.-..1/"--........, 

'""' v-' ....... , / 

VERMILION ~ 

DISTRICT X 0 Ely 

MINNESOTA 

Minnesota (Sims, 1972) and adjacent Canada (Franklin 

and Thorpe, 1982). Several deposits have been mined in 

the Vermilion district of Minnesota from near Ely (fig. 

6). The deposits are smaller, less continuous, and more 

irregular in thickness and iron content than the Superiortype iron-formations of Early Proterozoic age in the 

region. The ore at the major (Soudan) mine in the 

Vermilion district is of ''lump'' grade, being a massive, hard, bluish-gray, nearly pure hematite. The 

hematite ore is an enriched iron ore formed by hydrothermal processes (Gruner, 1926) from dominantly 

magnetite-bearing banded iron-formation of cherty 

oxide type. The iron-formations in the Vermilion district 

are associated with volcanic rocks-both basalt flows and 

pyroclastic deposits-as well as with epiclastic sedimentary rocks, some of which are carbonaceous; and thus 

I:AKE MICHIGAN 

300 KILOMETERS 

Figure 6. Map of Lake Superior region, showing major known and potential metallic mineral districts in Precambrian 

rocks. Horizontal rule, known mining district; vertical rule, potential mining district; black pattern, iron-formation. 

Metallogeny in the Great Lakes Region 67 
these iron-formations are clearly distinguished from the 

Superior-type iron-formations, which lack any obvious 

volcanic relation. As pointed out by Gole and Klein 

(1981 ), however, the two types of deposits otherwise have 

many striking similarities in textures and structures. 

Algoma-type iron-formations are exhalative-type deposits 

that were formed during lulls in the volcanism, commonly 

at the completion of a volcanic cycle. The iron-formations 

in northern Minnesota are complexly deformed (Sims and 

Southwick, 1980; 1985) and are metamorphosed to 

greenschist facies. A few iron-formations in the region 

are hosted by graywacke turbidites; generally these are 

smaller and less suitable for mining than the volcanichosted deposits. 

Nickel-Copper Deposits 

Nickel-copper deposits in Archean greenstone belts 

of the Superior province are common (Franklin and 

Thorpe, 1982) and contribute substantially to Canada's 

mineral economy, but deposits of economic significance 

have not been identified in the United States. The deposits 

are grossly of two types, one being associated with 

tholeiitic rocks and the other with komatiitic rocks; both 

rock types are mantle derived. Komatiitic (magnesiumrich) basalts are rather uncommon in Minnesota but occur locally in the Newton Lake Formation (Schulz, 1980). 

Gold Deposits 

The Archean greenstone-granite terrane of the 

Superior province in Canada has yielded more than 4.3 

million kg of gold-a major world producer-from 

several different types of deposits (Franklin and Thorpe, 

1982), but production from this terrane in the United 

States has been rather small. At least 1,250 kg of gold 

has been produced from the Ropes gold mine in northern 

Michigan (Rossell and Kalliokoski, 1983), and a small· 

amount of gold was produced from a mine on an island 

in Rainy Lake (Sims, 1972). 

The recent discovery of the Hemlo gold deposit at 

Heron Bay on the north shore of Lake Superior, in Ontario (Muir, 1983), has spurred a major new exploration 

phase for gold in the Archean rocks of the region. The 

Hemlo deposit is a large stratabound(?) deposit of exhalative type in felsic to intermediate pyroclastic volcanic 

host rocks. Reserves of at least 50 t (metric tons) of gold 

have been reported; the tenor of the ore is reported to 

be 0.28 oz gold/ton. Molybdenite occurs in approximately equal amounts as the gold in the ore. 

With respect to Early Proterozoic island-arc rocks, 

anomalous gold values have been reported (G. L. 

LaBerge, oral commun., 1984) from ferruginous chert 

and associated rocks in a volcanic-sedimentary basin 

68 Contributions to Mineral Resources Research, 1984 

within the Marshfield-Stevens Point belt (fig. 5) in central Wisconsin; the extent and potential significance of 

the showings, however, are not known. 

Deposits in Continental Crustal Terranes 

Early Proterozoic epicratonic rocks, which overlie 

the composite Archean basement, contain the vast ironformations for which the region is famed (Bayley and 

James, 1973; Morey, 1983) as well as small, probably 

uneconomic deposits of carbonate-hosted copper (table 

3). In adjacent Canada, major conglomerate-hosted 

uranium deposits (Robertson, 1976) occur in the Huronian Supergroup, which is a continental margin accumulation (Young, 1984) similar to the Marquette Range 

Supergroup. 

Middle Proterozoic rocks of the Midcontinent rift 

system contain volcanic- and sedimentary-hosted copper 

deposits (table 3), which have been mined for more than 

a century in the Keweenaw Peninsula of Michigan, and 

gabbro-hosted Cu-Ni deposits in the Duluth Complex 

(Ely-Hoyt Lakes district, fig. 6), a complex which constitutes this country's principal resource of nickel and 

cobalt for the future. 

A small Algoma-type iron-formation has been 

mined at one locality from the Archean gneiss microcontinent(?) associated with the Wisconsin magmatic terrane in central Wisconsin. 

Superior-Type Iron-Formations 

Superior-type iron-formations of Early Proterozoic 

age have provided more than 96 percent of the 4.6 billion 

t of iron ore shipped from the Lake Superior region since 

1848 (Morey, 1983). The production has come from 

separate districts, or iron ranges, within the Animikie 

basin, a depositional basin having an area of more than 

200,000 km2 , and several other, separate, possibly 

isolated smaller basins, in Michigan. James (1983) has 

estimated that the initial amount of Superior-type ironformation in the Lake Superior region was ·1013 t. 

Superior-type iron-formations are characteristically 

thin banded, cherty rocks with iron-rich layers, containing granules or oolites of both chert and iron minerals, 

and generally lacking clastic material (Gross, 1965). They 

are associated with quartzite and carbonaceous shale and 

commonly with dolomite, argillite, and conglomerate as 

well. Major iron-formations tend to be continuous and 

commonly extend for several hundreds of kilometers. 

They differ from Algoma-type iron-formations in lacking any obvious volcanic relation. In nearly all the depositional basins, iron-formation precipitation persisted, at 

least intermittently, throughout depositional history; the 
Table 3. Mineral deposits in continental crust, Lake Superior and Lake Huron regions 

[do, same as above entry) 

Terrane 

Midcontinent rift 

system. 

Epicratonic rocks 

Deposit type 

Volcanic-hosted 

copper. 

Host rock 

Portage Lake 

Volcanics of 

Keweenawan 

Supergroup. 

Shale-hosted copper Nonesuch Shale of 

Keweenawan 

Supergroup. 

Gabbro-hosted copper- Duluth Complex--

nickel. 

Superior-type ironformation. Animikie Group--

Menominee Group 

of Marquette 

Range Supergroup. 

Age 

Middle Proterozoic 

----do-------------

----do-------------

Early Proterozoic--· 

----do-------------

Locality 

Keweenaw Peninsula, 

Michigan. 

Do. 

Ely-Hoyt Lakes 

district, Minnesota. 

Mesabi and Cuyuna 

ranges, Minnesota. 

Gogebic, Marquette, 

and Menominee ranges, 

Michigan and Wisconsin. 

Carbonate-hosted 

copper. 

Kona Dolomite ----do------------- Northern Michigan. 

(Chocolay Group) 

of Marquette 

Range Supergroup. 

Conglomerate-hosted 

uranium. 

Elliott Lake Group ----do------------- North shore, Lake Huron 

region. 

Gneiss---------- Algoma-type ironformation. 

of Huronian 

Supergroup. 

Metavolcanic(?) 

rocks. 

facies type of iron-formation (James, 1954) appears to 

be related to local conditions in the basins. There is now 

widespread acceptance that the source of the iron was the 

ocean waters, but the mechanism by which iron and silica 

were withdrawn from the reservoirs is poorly understood. 

Carbonate-Hosted Copper Deposits 

The Kona Dolomite, near the base of the Marquette 

Range Supergroup, contains local stratabound copper 

deposits. The largest deposit, so far as known, is in the 

vicinity of Marquette, Mich., but it appears to be too low 

grade and too cut up by faults to be considered for 

mining. 

Conglomerate-Hosted Uranium Deposits 

Quartz-pebble conglomerates and associated 

coarse-grained arenites from near the base of the Huronian Supergroup in the Blind River-Elliot Lake-Agnew 

Lake area, Ontario, Canada, have been a major source 

of uranium (Robertson, 1976). Between 1955 and 1979, 

uranium oxide totalling 120,000 twas recovered from the 

milling of approximately 100 million t of ore. Thirteen 

mines have been worked at different times .. The 

uraniferous conglomerates are in the Matinenda Formation, and the pay shoots are typically broad paleovalley 

fillings. Sedimentological studies have shown that the 

uraninite and associated monazite, pyrite, and zircon are 

Late Archean------- Black River Falls, 

Wisconsin. 

detrital and were transported down the paleoslope from 

Arch(!an source areas to the northwest. Anoxygenic conditions, which terminated by the end of the Early Proterozoic, were required for the transport and deposition 

of the uranium-bearing placer minerals. 

Volcanic- and Shale-Hosted Copper Deposits 

The copper deposits in the Middle Proterozoic 

(Keweenawan) rocks of the Keweenaw Peninsula are of 

two broad types: (1) native copper in amygdaloidal flow 

tops and conglomerate beds of the Portage Lake 

Volcanics (White, 1968), and (2) chalcocite and native 

copper deposits in the Nonesuch Shale at White Pine 

(Ensign and others, 1968). The former has yielded about 

1.5 million tons and the latter about 1 million tons of 

copper. The principal producing areas for both deposit 

types are shown in figure 6. 

The native copper deposits in the Portage Lake 

Volcanics occur primarily as open-space fillings and 

replacements in flow tops and conglomerate. Both 

primary permeability and tectonic fracturing controlled 

the distribution of the copper ore shoots. White (1968) 

proposed that the copper-bearing solutions that formed 

the deposits were ascending and were derived either from 

a large concealed intrusion or, more likely, from water 

driven out of the porous tops of buried lava flows when 

the rocks were compacted and metamorphosed (to 

prehnite-pumpellyite facies). Concentrations of copper 

have been found associated with lava flows in other areas 

Metallogeny in the Great Lakes Region 69 
in the Lake Superior region, but apparently none of these 

are of economic significance. 

The shale-hosted copper deposit in the Nonesuch 

Shale is stratabound, although slightly cross cutting on 

a regional scale, is continuous over an area of several tens 

of square kilometers, and contains chalcocite as the dominant copper mineral (Ensign and others, 1968). It is grossly similar to the Kupferschiefer deposits of Paleozoic age 

in Germany and Poland. A major factor controlling the 

form and location of ore bodies is a favorable host (darkgray or black, laminated shale). Apparently organic matter had a direct role in localizing ore. White (1971) has 

proposed an elegant paleohydrologic model to account 

for the mineralization in the White Pine deposit. This 

model involves the lateral migration of fluids through the 

subjacent Copper Harbor Conglomerate to the site of the 

deposit and the precipitation of copper from these solutions where they percolated upward through the carbonaceous Nonesuch Shale. 

Another large area of Nonesuch Shale in northwestern Wisconsin (Morey and others, 1982) has been unsuccessfully explored for copper deposits. 

Gabbro-Hosted Copper-Nickel Deposits 

The Duluth Complex, of Keweenawan age, in 

northeastern Minnesota (Weiblen and Morey, 1980), constitutes an enormous untapped nickel resource and a 

substantial copper resource. Cobalt and platinum-group 

elements are important potential byproducts. Estimates 

of the magnitude of the resource indicate the existence 

of several billion tons of marginal- and low-grade copperand nickel-bearing material (see Sims, in press, for a summary). An unfavorable metal market and environmental concerns have hampered development of any of the 

deposits. 

Copper- and nickel-bearing sulfides are dispersed 

widely in the lower, western part of the Duluth Complex 

and are concentrated locally into bodies of potential 

economic significance. The major deposits that have been 

delineated by drilling are in the Ely-Hoyt Lakes district, 

outlined in figure 6. The principal concentrations are in 

troctolite, commonly in the vicinity of inclusions of older 

argillite. Basinlike depressions and growth faults are 

favorable structures that have localized a part of the oremineral concentrations. The ore minerals-mainly pyrrhotite, chalcopyrite, cubanite, and pentlandite-are 

disseminated in the host rock. Semimassive ore has been 

delineated in one potential ore body (Listerud and 

Meineke, 1977). The magma generally is interpreted as 

the source of the metals; most of the sulfur is believed 

to have been derived from a crustal source, probably the 

Early Proterozoic Virginia Formation (Mainwaring and 

Naldrett, 1977). Naldrett (1981) assigned the deposit to 

his class of "Intrusions feeding flood basalt activity 

70 Contributions to Mineral Resources Research, 1984 

associated with intra-continental rift zones." The magma 

was derived directly from the mantle. 

Algoma-Type I ron-Formation 

In central Wisconsin, Algoma-type iron-formation 

of presumed Archean age has been mined at one locality 

(Maass and Van Schmus, 1980), the Jackson County iron 

mine, east of Black River Falls (fig. 6). The ironformation is a banded magnetite-hematite rock that differs greatly from iron-formations in the greenstonegranite terrane of northern Minnesota in being metamorphosed to middle amphibolite grade (Jones, 1978). It is 

interbedded with granitic and migmatitic gneisses, 

schistose greenstone, and metasedimentary rocks, which 

are intruded by massive granite, aplite, and metadiabase 

dikes (Harry Klemic, written commun., 1970). 

SUMMARY OF TECTONIC EVOLUTION 

AND MINERALIZATION 

A tectono-stratigraphic .approach to metallogeny 

provides a framework within which to relate existing and 

possible undiscovered mineral deposits to their enclosing rocks and, with a proper understanding of regional 

geologic evolution, to possible plate tectonic regimes. 

Knowledge of the construction and assembly of the 

several Archean and Proterozoic terranes in the Lake 

Superior region-and thus of the environment of 

mineralization-decreases with increasing age of rocks. 

The Middle Proterozoic (:::::1,225-1,100 Ma) Midcontinent 

rift system is recognized as a classic example of an aborted 

continental rift, but the actual process of development 

of the rift and its affinity with tectonic analogs elsewhere 

are still debated (Green, 1983). As a consequence of the 

rifting, mantle-derived mafic magma ascended into the 

upper crust, where it reacted with crustal concentrations 

of sulfur-to give rise to magmatic concentrations of nickel 

and copper sulfides. Naldrett (1981) has likened the CuNi deposit in the Duluth Complex to the Norilsk-Talnakh 

Cu-Ni deposit in the U.S.S.R., which also lies within a 

major continental rift along the northwestern margin of 

the Siberian platform. The volcanic- and sedimentaryhosted copper deposits of the Keweenaw Peninsula are 

secondary concentrations resulting from the leaching of 

country rocks and the transport and deposition of the 

metal in favorable lithologic or structural traps (White, 

1968; 1971). 

Major aspects of the evolution of the Early Proterozoic rocks (:::::2,100-1,850 Ma) in the region have been 

controversial (Sims, 1976; Van Schmus, 1976), but there 

now is a convergence of opinion, if not general agreement, that these rocks formed in a plate tectonic regime, 

probably analogous to modern tectonic environments. 
The recent model proposed by Schulz and others (1984), 

a refinement of an earlier one by Cambray (1978), suggests early rifting of the Archean continental margin, 

spreading, subsequent subduction and formation of a 

complex volcanic arc, and collision of the arc, first with 

an Arche~n microcontinent(?) on the south and then with 

the continental margin sequence and Archean crust (of 

northern Michigan) on the north. The collision and 

related tectonism produced the Penokean orogen at about 

1,850 Ma. 

Superior-type iron-formations were deposited on 

the passive, rifted margin of the Archean craton under 

stable conditions, before the beginning of passive-margin 

subsidence. They were deposited near a transgressing and 

regressing strandline (White, 1954), probably on a shelf 

located seaward from the subtidal sandstone facies of the 

stratigraphically older sandstone (quartzite) deposits 

(Ojakangas, 1983). 

The Cu-Zn-type massive sulfide deposits in the 

volcanic-plutonic complexes of the Wisconsin magmatic 

zone accumulated in felsic-intermediate volcanic host 

rocks that are chemically similar to modern island-arcs, 

like those of the western Pacific (Schulz, 1984). The local 

coexistence of tectonically emplaced ultramafic rocks with 

basalts of mid-ocean ridge chemical affinities further supports such an environment of formation. All the known, 

potentially economic, massive sulfide deposits occur in 

the northern, Ladysmith-Pembine volcanic belt, perhaps 

because of a favorable chemical environment. This belt 

of volcanic rocks is tentatively interpreted as the younger 

of the two arc complexes in the Wisconsin magmatic terrane, and it is believed to have formed over a southwarddipping subduction zone. 

Still more uncertainty exists in interpreting the environments in which the rocks and mineralization of the 

Archean terranes respectively were formed and took 

place. The ancient gneiss terrane is especially enigmatic. 

Apparently it was part of a protocontinent of substantial size that was metamorphosed at moderately deep 

crustal levels, to produce the ubiquitous amphibolite and 

granulite metamorphic facies, then rose vertically(?) to 

its approximate present level by Late Archean time. The 

sparsity of economic mineral deposits in this terrane could 

mean that they were removed by erosion, but it could also 

mean that few metal deposits were formed or that they 

were obliterated by subsequent crustal reworking and 

metamorphism. 

The tectonic environment of the Late Archean 

greenstone-granite complexes is a perplexing problem, but 

they probably formed in repetitive cycles adjacent to the 

older gneiss terrane (present coordinates) to the south as 

continental or oceanic arcs. There was some overlap of 

deposition onto the sialic gneiss terrane (Sims and others, 

1984), but for the most part they probably were deposited 

on an oceanic floor. Low initial 87Sr / 86Sr values (Sims 

and Peterman, 1981) are interpreted as indicating massive 

additions of juvenile, mantle-derived material to the crust 

by ways of melting, assimilation, and mobilization. A 

common current postulate is that plate tectonic processes 

occurred during the Late Archean. The requirements for 

greater rates of dissipation of mantle heat can be accommodated merely by increasing the number of plates and 

the rates of subduction. The volcanic rocks have both 

tholeiitic and calc-alkaline compositions. Most of the 

felsic volcanic rocks are calc-alkaline, but recently it has 

been shown (Lesher and others, 1981) that some of the 

felsic volcanic rocks associated with Cu-Zn-type massive 

sulfide deposits in the Superior province are of tholeiitic 

affinity; they form the felsic members of bimodal volcanic 

assemblages. These rocks, which have rare-earth-element 

patterns that differ from those of felsic rocks in apparently barren terranes, are inferred to have been derived from 

partial melting of associated tholeiitic basalts during the 

intrusion of and the subsequent fractional crystallization 

of magmas in high-level chambers (Lesher and others, 

1981). 

REFERENCES CITED 

Afifi, Afifa, Doe, B.R., Sims, P.K., and Delevaux, M.H., 1984, 

U-Th-Pb isotope chronology of sulfide ores and rocks in 

the Early Proterozoic metavolcanic belt of northern 

Wisconsin: Economic Geology, v. 79, p. 338-353. 

Anderson, R.R., and Black, R.A., 1983, Early Proterozoic 

development of the southern Archean boundary of the 

Superior province in the Lake Superior region: Geological 

Society of America Abstracts with Programs, v. 15, no. 6, 

p. 515. 

Arth, J.G., and Hanson, G.N., 1975, Geochemistry and origin 

of the early Precambrian crust of northeastern Minnesota: 

Geochimica et Cosmochimica Acta, v. 39, p. 325-362. 

Bayley, R.W., Dutton, C.E., and Lamey, C.A., 1966, Geology 

of the Menominee iron-bearing district, Dickinson County, Michigan, and Florence and Marinette Counties, 

Wisconsin: U.S. Geological Survey Professional Paper 513, 

96 p. 

Bayley, R.W., and James, H.L., 1973, Precambrian ironformation of the United States: Economic Geology, v. 68, 

p. 934-959. 

Beck, Warren, and Murthy, V. Rama, 1982, Rb-Sr and 

Sm-Nd isotopic studies of Proterozoic mafic dikes in 

northeastern Minnesota [abs.]: 28th Annual Institute on 

Lake Superior Geology, International Falls, Minnesota, 

p. 5. 

Cambray, F.W., 1978, Plate tectonics as a model for the environment of deposition and deformation of the early Proterozoic (Proterozoic X) of northern Michigan: Geological 

Society of America Abstracts with Programs, v. 10, no. 

7, p. 376. 

Campbell, I.H., Coad, P., Franklin, J.M., Gorton, M.P., Scott, 

S.D., Sowa, J.M., and Thurston, P.C., 1982, Rare earth 

elements in volcanic rocks associated with Cu-Zn massive 

sulphide mineralization-A preliminary report: Canadian 

Journal of Earth Sciences, v. 19, p. 619-623. 

Metallogeny in the Great Lakes Region 71 
Cannon, W.F., and Gair, J.E., 1970, A revision of stratigraphic 

nomenclature for middle Precambrian rocks in northern 

Michigan: Geological Society of America Bulletin, v. 81, 

p. 2843-2846. 

DuBois, J.F., and Van Schmus, W.R.,1978, Petrology and 

geochronology of Archean gneiss in the Lake Arbutus area, 

. west-central Wisconsin [abs.]: Abstracts and Proceedings, 

24th Annual Institute on Lake Superior Geology, 

Milwaukee, Wis., p. 11. 

Ensign, C.O., Jr., White, W.S., Wright, J.C., Patrick, J.L., Leone, 

R.J., Hathaway, D.J., Trammell, HV., Fritts, J.J., and Wright, 

T.L., 1968, Copper deposits in the Nonesuch Shale, White 

Pine, Michigan, in Ridge, J.D., ed., Ore deposits of the 

United States, 1933-1967: American Institute of Mining, 

Metallurgical, and Petroleum Engineers, Inc., p. 460-488. 

Fox, T.P., 1983, Geochemistry of the Hemlock Metabasalt and 

Kiernan sills, Iron County, Michigan: East Lansing, Mich., 

Michigan St~te University M.S. thesis, 81 p. 

Franklin, J.M., and Thorpe, R.I., 1982, Comparative metallogeny 

of the Superior, Slave and Churchill provinces, in Hutchinson, RW., Spence, C.D., and Franklin, J.M., eds., Precam- · 

brian sulphide deposits (H.S. Robinson Memorial Volume): 

Geological Association of Canada Special Paper 25, p. 

3-90. 

Gibbs, A.K., Payne, Barton, Setzer, Thomas, Brown, L.D., Oliver, 

J.E., and Kaufman, Sidney, 1984, Seismic-reflection study 

of the Precambrian crust of central Minnesota: Geological 

Society of America Bulletin, v. 95, p. 280-294. 

Goldich, S.S., and Wooden, J.L., 1980, Origin of the Morton 

Gneiss, southwestern Minnesota, Part 3, Geochronology, 

in Morey, G.B., and Hanson, G.N., eds., Selected studies 

of Archean gneisses and lower Proterozoic rocks, southern 

Canadian Shield: Geological Society of America Special 

Paper 182, p. 77-94. 

Gole, M.J., and Klein, Cornelius, 1981, Banded iron-formations 

through much of Precambrian time: Journal of Geology, 

v. 89, p. 169-183. 

Grant, J.A., 1972, Minnesota River valley, southwestern Minnesota, in Sims, P.K., and Morey, G.B., eds., Geology of 

Minnesota-A centennial volume: Minnesota Geological 

Survey, p. 177-196. 

Green, J.C., 1983, Geologic and geochemical evidence for the 

nature and development of the Middle Proterozoic 

(Keweenawan) Midcontinent rift of North America: Tectonophysics, v. 94, p. 413-437. 

Greenberg, J.K., and Brown, B.A., 1983, Lower Proterozoic 

volcanic rocks and their setting in the southern Lake 

Superior district, in Medaris, L.G., Jr., ed., Early Proterozoic 

geology of the Great Lakes region: Geological Society of 

America Memoir 160, p. 67-84. 

Gross, G.A., 1965, Geology of iron deposits in Canada, Volume 

1, General geology and evaluation of iron deposits: 

Geological Survey of Canada Economic Geology Report 

22, 181 p. 

Gruner, J.W., 1926, The Soudan formation and a new suggestion as to the origin of the Vermilion iron ores: Economic 

Geology, v. 21, p. 629-644. 

Hanson, G.N., and Malhotra, R., 1971, K-Ar ages of mafic dikes 

and evidence for low-grade metamorphism in northeastern 

Minnesota: Geological Society of America Bulletin, v. 82, 

p. 1107-1114. 

72 Contributions to Mineral Resources Research, 1984 

Hutchinson, R.W., 1980, Massive base metal sulphide deposits 

as guides to tectonic evolution, in Strangway, DW., ed., The 

continental crust and its mineral deposits: Geological 

Association of Canada Special Paper 20, p. 659-684. 

James, H.L., 1954, Sedimentary facies of iron-formation: 

Economic Geology, v. 49, p. 235-293. 

___ 1983, Distribution of banded iron-formation in space and 

time, in Trendall, A.F., and Morris, R.C., eds., Ironformation-Facts and problems: Amsterdam and New 

York, Elsevier, p. 471-490. 

Jones, D.O., 1978, Geology of the iron formation and associated 

rocks of the Jackson County iron mine, Jackson County, 

Wisconsin: Madison, Wis., University of Wisconsin M.S. 

thesis, 117 p. 

King, P.B., and Edmonston, G.J., 1972, Generalized tectonic map 

of North America: U.S. Geological Survey Miscellaneous 

Geologic Investigations Map I-688, scale 1:15,000,000. 

LaBerge, G.L., and Myers, P.E., 1983, Precambrian geology of 

Marathon County, Wisconsin: Wisconsin Geological and 

Natural History Survey Information Circular 45, 88 p. 

___ 1984, 1\vo Early Proterozoic successions in central 

Wisconsin and their tectonic significance: Geological Society of America Bulletin, v. 95, p. 246-253. 

Larue, D.K., and Sloss, L.L., 1980, Early Proterozoic sedimentary basins of the Lake Superior region: Geological Society of America Bulletin, Pt. I, v. 91, p. 450-452. 

Lesher, C.M., Campbell, I.H., Coad, P., Franklin, J.M., Gibson, 

H.L., Goodwin, A.M., Gorton, M.P., Hart, T.R., Sowa, 

J.M., and Thurston, P.C., 1981, Rare earth elements in felsic 

volcanic rocks associated with Cu-Zn massive sulphide 

mineralization: Ontario Geological Survey Miscellaneous 

Paper 113, p. 127-134. 

Listerud, W.H., and Meineke, D.O., 1977, Mineral resources 

of a portion of Duluth Complex and adjacent rocks in St. 

Louis and Lake Counties, northeastern Minnesota: 

Minnesota Department of Natural Resources Report 93, 

73 p. 

Maass, R.S., 1983, Early Proterozoic tectonic style in central 

Wisconsin, in Medaris, L.G., Jr., ed., Early Proterozoic 

geology of the Great Lakes region: Geological Society of 

America Memoir 160, p. 85-95. 

Maass, R.S., Medaris, L.G., Jr., and Van Schmus, W.R., 1980, 

Penokean deformation in central Wisconsin: Geological 

Society of America Special Paper 182, p. 147-157. 

Maass, R.S., and Van Schmus, W.R., 1980, Precambrian tectonic history of the Black River Valley: Field trip guidebook 

2, 26th Annual Institute on Lake Superior Geology, Eau 

Claire, Wis., 43 p. 

Mainwaring, P.R., and Naldrett, A.J., 1977, Country rock 

assimilation and the genesis of Cu-Ni sulfides in the Water 

Hen intrusion, Duluth Complex: Economic Geology, v. 

72, p. 1269-1284. 

May, E.R., 1977, Flambeau-a Precambrian supergene enriched 

massive sulfide deposit: Geoscience Wisconsin, v. 1, 

p. 1-26. 

May, E.R., and Schmidt, P.O., 1982, The discovery, geology 

and mineralogy of the Crandon Precambrian massive 

sulphide deposit, Wisconsin, in Hutchinson, R.W., Spence, 

C.D., and Franklin, J.M., eds., Precambrian sulphide 

deposits (H.S. Robinson Memorial Volume): Geological 

Association of Canada Special Paper 25, p. 447-480. 
Morey, G.B., 1973, Stratigraphic framework of Middle Precambrian rocks in Minnesota, in Young, G.M., ed., Huronian 

stratigraphy and sedimentation: Geological Association of 

Canada Special Paper 12, p. 211-249. 

___ 1983, Animikie basin, Lake Superior region, U.S.A., in 

Trendall, A.F., and Morris, R.C., eds., Iron formationFacts and problems: Amsterdam and New York, Elsevier, 

p. 13-68. 

Morey, G.B., and Sims, P.K., 1976, Boundary between two 

Precambrian W terranes in Minnesota and its geologic 

significance: Geological Society of America Bulletin, v. 87, 

p. 141-152. 

Morey, G.B., Sims, P.K., Cannon, W.F., Mudrey, M.G., Jr., and 

Southwick, D.L., 1982, Geologic map of the Lake Superior 

region, Minnesota, Wisconsin, and northern Michigan: 

Minnesota Geological Survey State Map Series S-13, scale 

1:1,000,000. 

Mudrey, M.G., Jr., 1979, The massive sulfide occurrences in 

Wisconsin: Wisconsin Geological and Natural History 

Survey Paper 79-2, 20 p. 

Muir, T.L., 1983, Geology of the Hemlo-Huron Bay area, in 

Col vine, A.C., ed., The geology of gold in Ontario: Ontario 

Geological Survey Miscellaneous Paper 110, p. 230-239. 

Naldrett, A.J., 1981, Nickel sulfide deposits-Classification, composition, and genesis, in Sims, P.K., and Skinner, B.J., eds., 

75th Anniversary Volume: Economic Geology, p. 628-685. 

Ojakangas, R.W., 1983, Tidal deposits in the early Proterozoic 

basin of the Lake Superior region-The Palms and 

Pokegama Formations: Evidence for subtidal-shelf deposition of Superior-type banded iron-formation, in Medaris, 

L.G., Jr., ed., Early Proterozoic geology of the Great Lakes 

region: Geological Society of America Memoir 160, 

p. 49-66. 

Palmer, E.A., 1980, The structure and petrology of Pre~ambrian 

metamorphic-rock units, northwestern Marathon County, 

Wisconsin: Duluth, Minn., University of Minnesota-Duluth 

M.S. thesis, 127 p. 

Robertson, J.A., 1976, The Blind River uranium deposits-The 

ores and their setting: Ontario Division of Mines Miscellaneous Paper 65, 45 p. 

Rossell, Dean, and Kalliokoski, Jo, 1983, The Ropes gold mine 

and its geologic setting: 29th Annual Institute on Lake 

Superior Geology, Houghton, Michigan, p. 44-74. 

Schulz, K.J., 1980, The magmatic evolution of the Vermilion 

greenstone belt, northeastern Minnesota: Precambrian 

Research, v. 11, p. 215-245. 

___ 1984, Volcanic rocks of northeastern Wisconsin, in Sims, 

P.K., Schulz, K.J., and Peterman, Z.E., Guide to the geology 

of the Early Proterozoic rocks in northeastern Wisconsin: 

30th Annual Institute on Lake Superior Geology, Field trip 

1, p. 51-80. 

Schulz, K.J., LaBerge, G.L., Sims, P.K., Peterman, Z.E., and 

Klasner, J.S., 1984, The volcanic-plutonic terrane of northern Wisconsin-Implications for Early Proterozoic tectonism, Lake Superior region: Geological Association of 

Canada-Mineralogical Association of Canada, London, 

Ontario, Canada, Program with abstracts, p. 103. 

Sims, P.K., 1972, Mineral deposits in lower Precambrian rocks, 

northern Minnesota, in Sims, P.K., and Morey, G.B., eds., 

Geology of Minnesota-A Centennial Volume: Minnesota 

Geological Survey, p. 172-176. 

___ 1976, Precambrian tectonics and mineral deposits, Lake 

Superior region: Economic Geology, v. 71, p. 1092-1118. 

___in press, History of discovery of copper-nickel deposits 

in the Duluth Complex, northeast Minnesota, in Hollister, 

V.F., ed., Case histories of mineral discoveries: Society of 

Mining Engineers. 

Sims, P.K., Card, K.D., and Lumbers, S.B., 1981, Evolution of 

early Proterozoic basins of the Great Lakes region, in 

Campbell, F.H.A., ed., Proterozoic basins of Canada: Geological Survey of Canada Special Paper 81-10, p. 379-397. 

Sims, P.K., Card, K.D., Morey, G.B., and Peterman, Z.E., 1980, 

The Great Lakes tectonic zone-A major crustal structure 

in central North America: Geological Society of America 

Bulletin, Pt. I, v. 91, p. 690-698. 

Sims, P.K., and Peterman, Z.E., 1981, Archean rocks in the 

southern part of the Canadian Shield-A review: Geological Society of Australia Special Publication 7, p. 85-98. 

__ 1983, Evolution of Penokean foldbelt, Lake Superior 

region, and its tectonic environment, in Medaris, L.G., Jr., 

ed., Early Proterozoic geology of the Great Lakes region: 

Geological Society of America Memoir 160, p. 3-14. 

___ 1984, A partisan review of the Early Proterozoic geology 

of Wisconsin and adjacent Michigan [abs.]: 30th Annual 

Institute on Lake Superior Geology, Wausau, Wis., 

p. 73-76. 

Sims, P.K., Peterman, Z.E., Prinz, W.C., and Benedict, F.C., 

1984, Geology, geochemistry, and age of Archean and Early 

Proterozoic rocks in the Marenisco-Watersmeet area, northern Michigan: U.S. Geological Survey Professional Paper 

1292-A, 41 p. 

Sims, P.K., Peterman, Z.E., and Schulz, K.J., 1985, The Dunbar 

Gneiss-granitoid dome-Implications for early Proterozoic 

tectonic evolution of northern Wisconsin: Geological Society of America Bulletin, v. 96, p. 1101-1112. 

Sims, P.K., and Southwick, D.L., 1980, Geologic map of the 

Soudan quadrangle, St. Louis County, Minnesota: U.S. 

Geological Survey Geologic Quadrangle Map GQ-1540, 

scale 1:24,000. 

___ 1985, Geologic map of Archean rocks, western Vermilion 

district, northern Minnesota: U.S. Geological Survey 

Miscellaneous Investigations Map 1-1527, scale 1:48,000. 

Southwick, D.L., and Day, W.C., 1983, Geology and petrology 

of Proterozoic mafic dikes, north-central Minnesota and 

western Ontario: Canadian Journal of Earth Sciences, v. 20, 

p. 622-638. 

Van Schmus, W.R., 1976, Early and middle Proterozoic history 

of the Great Lakes area, North America: Royal Society of 

London Philosophical Transactions, ser. A280, no. 1298, 

p. 605-628. 

Van Schmus, W.R., and Anderson, J.L., 1977, Gneiss and 

migmatite of Archean age in the Precambrian basement of 

central Wisconsin: Geology, v. 5, p. 45-48. 

Van Schmus, W.R., Medaris, L.G., Jr., and Banks, P.O., 1975, 

Geology and age of the Wolf River batholith, Wisconsin: 

Geological Society of America Bulletin, v. 86, 

p. 907-914. 

Van Schmus, W.R., Green, J.C., and Halls, H.C., 1982, 

Geochronology of Keweenawan rocks of the Lake Superior 

region, in Wold, R.J., and Hinze, W.J., eds., Geology and 

tectonics of the Lake Superior region: Geological Society 

of America Memoir 156, p. 165-172. 

Metallogeny in the Great Lakes Region 73 
Weiblen, P.W., 1982, Keweenawan intrusive igneous rocks, in 

Wold, R.J., and Hinze, W.J., eds., Geology and tectonics 

of the Lake Superior region: Geological Society of America 

Memoir 156, p. 57-82. 

Weiblen, P.W., and Morey, G.B., 1980, A summary of the 

stratigraphy, petrology and structure of the Duluth Complex: American Journal of Science, v. 280-A, p. 88-133. 

White, D.A., 1954, Stratigraphy and structure of the Mesabi 

range, Minnesota: Minnesota Geological Survey Bulletin 

38, 92 p. 

White, W.S., 1968, The native copper deposits of northern 

Michigan, in Ridge, J.D., ed., Ore deposits of the United 

States, 1933-1967: American Institute of Mining, Metallurgical, and Petroleum Engineers, Inc., p. 303-326. 

74 Contributions to Mineral Resources Research, 1984 

___ 1971, A paleohydrologic model for mineralization of the 

White Pine copper deposit, northern Michigan: Economic 

Geology, v. 66, p. 1-13. 

Windley, B.F., 1977, The evolving continents: New York, John 

Wiley, 385 p. 

Wold, R.J., and Hinze, W.J., 1982, Introduction, in Wold, R.J., 

and Hinze, W.J., eds., Geology and tectonics of the Lake 

Superior region: Geological Society of America Memoir 

156, p. 1-4. 

Young, G.M., 1984, The Huronian Supergroup-An example of 

an Early Proterozoic passive margin sequence [abs.]: 30th 

Annual Institute on Lake Superior Geology, Wausau, Wis., 

p. 80. 
Chapter F 

Sodium Hypochlorite as an Aid to the 

Extraction of Clay Minerals from 

Black Shales 

By HARRY C. STARKEY 

Disaggregation of black shales is increased by 

sodium hypochlorite dissolution of cementing organic materials 

U.S. GEOLOGICAL SURVEY BULLETIN 1694 

CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 
CONTENTS 

Abstract 77 

Introduction 77 

Experimental method 78 

Results and discussion 79 

References cited 80 

FIGURE 

1. Diagram showing weight loss of sodium hypochlorite- and distilled-watertreated black shale samples compared to number of treatments 80 

TABLES 

1. Weight loss from sodium hypochlorite versus distilled-water treatments compared with partial data on total organic carbon content 78 

2. Comparison of particle-size fractionation of black shales treated with sodium 

hypochlorite and with distilled water 79 

76 Contents 
CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 

Sodium Hypochlorite as an Aid to the 

Extraction of_ Clay Minerals from Black Shales 

By Harry C. Starkey 

Abstract 

Black shales are important as petroleum source rocks and 

also may contain heavy minerals. Clay minerals, which may 

indicate depositional and postdepositional environments, are 

important shale constituents. Identifying these clay minerals is 

difficult because the organic matter in the black shales may act 

as a cementing agent, which makes disaggregating the samples 

difficult. 

A 5.25 percent solution of sodium hypochlorite, adjusted 

to a pH of 9.50, was used as an aid to the disaggregation and 

dispersion of 17 samples of black shale by removing organic 

matter prior to the removal of the clay-size fraction. Controls 

of the same samples were run using distilled water instead of 

the sodium hypochlorite. In every sample, the amount of sandsize particles was reduced and, in all but one sample, the 

amount of clay-size particles was increased when the sodium 

hypochlorite was used. The sodium hypochlorite-treated 

samples lost an average of 1.7 percent more weight than did 

the samples treated with distilled water; some of the increased 

weight loss was probably due to the destruction of organic matter by the sodium hypochlorite. A lightening in the dark color 

of most samples treated with sodium hypochlorite substantiated 

this loss of organic matter. 

All samples but one showed a greater weight loss with the 

sodium hypochlorite solution than with the distilled water; the 

exception, sample 1145, contained gypsum. The differences in 

the final weights of the samples treated with sodium hypochlorite and those treated with distilled water are less than the 

experimental error in any particular sample. As all the experimental error is unlikely to be in the same direction, the differences all indicate that more organic matter was removed by 

the sodium hypochlorite. 

Compared to the clay-size fraction after the sodium hypochlorite treatment, only one sample showed a higher clay-size 

fraction after distilled-water treatment; the difference was 

0.4 percent. As this difference is well within the experimental 

error, the clay-size fraction present after both treatments of this 

sample is considered to be the same. 

INTRODUCTION 

Black shales are widespread; more than 200 formations containing black shales were investigated in the 

United States between 1944 and 1956 (Swanson, 1961). 

These shales are organic-rich, fine-grained rocks which 

may contain heavy metals, and they are also important 

as petroleum source rocks. Because they are composed 

of claystones and mudstones, the shales contain abundant 

clay minerals. These clay minerals may be important indicators of depositional and postdepositional environments. Identification of minerals in shales, especially clay 

minerals, is often hindered by difficulty in disaggregating 

the shales in which organic matter as well as silica and 

carbonates may act as cementing agents (van Langeweld 

and others, 1978; Sterne and others, 1982). To study the 

clay minerals, enough clay must be separated to make randomly oriented powder mounts and oriented aggregates 

for X-ray powder diffraction analysis. Removing cementing agents such as organic matter from shales facilitates 

the disaggregation necessary for particle-size analysis, for 

heavy-mineral extraction, and for the extraction of clay 

minerals for identification. 

This report discusses a method of removing organic 

material from black shales to increase the proportion of 

the clay-size fraction resulting from disaggregation and 

thus the amount of clay minerals extracted. Removing silica or carbonate cements may further increase the amount 

of clay minerals extracted, but the chemicals used for 

removing these cements may attack the clay minerals also. 

Hydrogen peroxide was often used in the past to 

remove organic material from geologic samples, but it produced some undesirable side effects such as the formation of calcium oxalate phases (Martin, 1954; Jones and 

Beavers, 1963), the dissolution of carbonates, and the partial dissolution of the clay minerals. Douglas and Feissinger (1971) demonstrated that the dissolution of clay 

minerals during hydrogen peroxide treatment could be 

reduced somewhat by using a sodium acetate buffer. 

Cassidy and Mankin (1960) proposed using sodium 

hypochlorite to remove organic matter from black shales. 

Their method consisted of boiling the sample for as much 

as 3 hours in a 5 percent solution of sodium hypochlorite 

(pH 12.1). This method effectively removed most of the 

organic matter from the shales, but some of the clay 

Extraction of Clay Minerals from Black Shales 77 
minerals were attacked, especially when the treatment was 

prolonged. Kaolins were subject to attack by the highpH solutions, particularly if the solutions became concentrated through evaporation by boiling. The solutions 

were so strong that Cassidy and Mankin cautioned against 

the use of glass beakers because the concentrated solutions would attack them. 

More recently, sodium hypochlorite has been used 

to destroy organic matter in soils (Anderson, 1963; 

Omueti, 1980). In a comparative study of the effectiveness 

of the sodium hypochlorite and the hydrogen peroxide 

methods of organic matter destruction, Lavkulich and 

Wiens (1970) found that the sodium hypochlorite, when 

used according to Anderson's methoq, destroyed more 

organic material while removing fewer inorganic cations 

than did the hydrogen peroxide. 

Anderson (1963) used calcium- and phosphorusfree sodium hypochlorite and Omueti (1980) used sodium 

hypochlorite (commercial bleach) adjusted to pH 9.5 to 

remove organic matter from soil samples. The pH of that 

solution is less destructive of the silica and alumina in 

the clay minerals than is the higher pH (12.1) of the unadjusted sodium hypochlorite. They used shorter reaction 

times and lower temperatures than did Cassidy and 

Mankin (1960): they heated the samples for only 15 min 

over a steam bath, after which they removed the spent 

sodium hypochlorite solution and replaced it with fresh 

solution. They then repeated the process twice. 

EXPERIMENTAL METHOD 

Seventeen black shales were prepared by grinding 

to -100 mesh. The shales used in this study were very fine 

grained, and this amount of grinding should not have produced much crystal fracturing. Grinding to such a small 

size was deemed desirable in order to allow better penetration of the sample by the solutions and thus reduce the 

number of treatments. Duplicate 1-g (gram) samples were 

taken; one to be treated with sodium hypochlorite, and 

the other to be treated, as a control, with distilled water 

in the same manner. 

A 1-g sample of each shale was placed in 40-mL 

(milliliter) centrifuge tubes with 5-mL sodium 

hypochlorite solution. The sodium hypochlorite solution 

used was commercial household laundry bleach containing 5.25 percent sodium hypochlorite and freshly adjusted 

to a pH of 9.5 with hydrochloric acid. The centrifuge tube 

was then placed in a steam bath and heated for 15 min. 

After removal from the steam bath, ~he sample was 

treated ultrasonically for 5 min, allowed to settle for 10 

min, and the suspended material was decanted and retained. Five milliliters of fresh solution was added to the 

sample, and the procedure was repeated several times until the supernatant was clear after the 10-min settling time. 

All the supernatants from each sample were combined, 

78 Contributions to Mineral Resources Research, 1984 

and the number of treatments recorded. The sodium 

hypochlorite solution was removed from the suspended 

material in the combined supernatants by vacuum 

through bacteriological filter candles (Starkey and others, 

1984). The suspended material was washed with distilled 

water, which then was removed by filter candles. After 

washing, this suspended material was recombined with 

the material that had settled out. 

This method did not remove all the organic matter 

(table 1), nor did it completely disaggregate all of the 

shales, but it did free enough of the clay minerals for 

study. For this study, the samples were separated after 

treatment into sand, silt, and clay-size fractions to 

demonstrate the changes in particle-size distribution. The 

method of particle-size fractionation was that described 

by Starkey and others (1984) in which the sand is removed 

by wet sieving through a 230-mesh sieve and the clay and 

silt were separated by centrifugation. 

Table 1. Weight loss from sodium hypochlorite versus 

distilled-water treatments compared with partial data on total 

organic carbon content 

[All values in weight percent; n.d., not determined] 

Sample NaOCl HzO Total 

No. treatment treatment organic carbon 1 

29S 3.2 1. 4 n.d. 

30S 4.2 0.8 n.d. 

33S 3. 1 1. 9 n.d. 

35S 1.0 1.2 n.d. 

46S 3.6 2.3 n.d. 

71S 3.8 2.5 n.d. 

90S 3.2 1. 2 n.d. 

lOOF 4.2 3.2 15.41 

lOOS 4.0 3.9 14.86 

110S 4.4 2. 1 6.84 

114S 4.8 7.4 23.43 

119S 4.2 2.9 12.03 

132S 3.8 2.0 13.27 

135S 4.4 1. 6 6.08 

303S 3.6 0.8 n.d. 

311S 4.4 2.2 n.d. 

562S 3.0 2.2 2.9 

1Total organic carbon data furnished by 

F .G. Poole, u.s. Geological Survey. 

To ascertain how much of the disaggregation 

achieved was caused by the sodium hypochlorite treatment rather than the ultrasonic treatment, duplicate 

samples were treated in the same manner, but distilled 

water was used in place of the sodium hypochlorite. The 

number and duration of the treatments were repeated 

with the distilled water, and the size fractions were 

separated by the same method already referred to 

(Starkey and others, 1984). 
RESULTS AND DISCUSSION 

The results obtained from this study show that 

sodium hypochlorite is an effective aid for removing 

organic matter from black shales. Although the sodium 

hypochlorite does not remove all the organic matter, it 

removes enough to increase the quantity of clay-size particles obtained without adverse effects on the clay 

minerals. 

The results of the particle-size fractionation of the 

samples after treatment with sodium hypochlorite and 

with distilled water are given in table 2. For all samples 

except one, the clay-size fractions were greater in the 

samples treated with sodium hypochlorite than in those 

samples treated with distilled water. The one exception 

was sample 33S, which produced 0.4 percent more claysize fraction in the sample treated with distilled water. 

This increase is so slight that it can be attributed to experimental error, and the amount of the clay-size fraction is considered the same for both samples. 

Not only was the percentage of the clay-size fraction generally greater after the sodium hypochlorite treatment, but the percentage of the sand-size fraction was 

decreased. Some of the samples produced more silt with 

the sodium hypochlorite treatment and some produced 

less. The percentage of silt produced depended upon 

whether the sodium hypochlorite was more effective in 

breaking down the large particles, thus producing silt, or 

in breaking down the silt, thus producing more clay-size 

particles. 

Samples treated with sodium hypochlorite had a 

greater overall weight loss than those treated with distilled water (fig. 1). All but one of the sodium 

hypochlorite-treated samples lost 3.0 weight percent or 

more, and all but three of the distilled-water-treated 

samples lost less than 3.0 weight percent. Some weight 

loss in all samples could be caused by handling and 

transferring samples. Other weight loss may have been 

caused by soluble salts that went into solution, or by 

variations in amounts of adsorbed water. 

The greater weight loss in the sodium hypochloritetreated samples was caused by the removal of the organic 

matter. The difference in weight loss between the sodium 

hypochlorite-treated and distilled-water-treated (control) 

samples averaged 1. 7 percent, if sample 114S is eliminated. This weight loss was accompanied by a general 

lightening of the color of the black shales that were 

treated with sodium hypochlorite. 

Two samples did not show greater weight loss with 

the sodium hypochlorite treatment. Sample 35S showed 

a 0.2 percent greater loss with distilled water than with 

the sodium hypochlorite treatment, but this amount is 

negligible, and so the weight loss is basically the same 

for both treatments. Any color change in sample 35S was 

Table 2. Comparison of particle-size fractionation of black shales treated with sodium hypochlorite and with distilled water 

[All values in weight percent] 

Sample Number of 5.25 pet NaOCl solution Distilled water 

No. treatments Sand Silt Clay Total Sand Silt Clay Total 

(percent) (percent) 

29S 6 30.9 46.8 19.1 96.8 31.5 66.5 0.6 98.6 

30S 13 20.9 53.2 21. 7 95.8 31.8 50.0 17.4 99.2 

33S 11 15.0 66.7 15.2 96.9 18.8 63.7 15.6 98. 1 

35S 8 14.0 61.0 24.0 99.0 17.8 58.4 22.6 98.8 

46S 9 3.2 77.5 15.7 96.4 21.2 65.3 11.2 97.7 

71S 13 2.4 76.4 17.4 96.2 28.9 62.0 6.6 97.5 

90S 7 23.2 62.8 10.8 96.8 27.1 71.3 0.4 98.8 

100F 3 17.7 52.4 25.7 95.8 21.2 51.8 23.8 96.8 

100S 4 2.6 44.4 49.0 96.0 11. 1 47.1 37.9 96.1 

110S 8 1.4 62.1 32. 1 95.6 14.6 58.5 24.8 97.9 

114S 5 27.6 40.8 26.8 95.2 30.9 39.6 22.1 92.6 

119S 4 21.7 42.7 31.4 95.8 25.8 44.3 27.0 97.1 

132S 11 13.6 68.8 13.8 96.2 28. 1 66.5 3.4 98.0 

135S 9 7.6 43.2 44.8 95.6 24.2 43.8 30.4 98.4 

303S 8 0.2 52.8 43.4 96.4 18.8 54.3 26.1 99.2 

311S 11 2.6 52.8 40.2 95.6 29.2 49.0 19.6 97.8 

562S 10 6.6 50.0 40.4 97.0 19.9 49.8 28. 1 97.8 

Extraction of .Clay Minerals from Black Shales 79 
indiscernible. Sample 114S had a greater weight loss in 

distilled water because the sample contained gypsum that 

went into solution. Although it produced a lighter color 

in the sample, the sodium hypochlorite treatment of sample 114S did not result in as great a weight loss because 

the commercial sodium hypochlorite solution used in this 

work contained carbonate anions. These anions combined with the calcium dissolved from the gypsum to form 

calcite, which was precipitated and weighed with the claysize fraction. If gypsum is present in a sample, and if 

particle-size distribution is to be determined, it may be 

necessary to use a purer sodium hypochlorite solution. 

Because handling and transferring samples may produce weight loss, a positive correlation between the 

number of treatments and the total weight loss was expected. Such a correlation was not evident (fig. 1). Apparently the loss of weight due to handling of the samples 

was minimal. 

Sodium hypochlorite, although niore efficient than 

hydrogen peroxide in removing organic matter, must only be considered as an aid to the disaggregation of black 

shales. Available data show that sodium hypochlorite 

removed only about one-fourth of the organic material 

from some organic-rich samples (table 1), which increased the amount of clay-size particles by f-11 percent (table 

2). Repetition of the treatment together with removal of 

8r--r-l-r-l-r-l-lr-~lr-~l~~~--~~--~l--,l--,l--,1--,l~ 

7~ 

6-

..... 

z 

w 

u 5- a: 

w 

a.. 

~ 

u) 4-

en 

0 

...J 

CY"contained gypsum 

EXPLANATION -

6 Sample treated with 5.25 

percent NaOCI 

0 Sample treated with 

distilledHzO 

-

-

-

6 

..... 0 6 6 6 

-----o-------~---­

w 

~ 0 

0 0 0 0 8 2- -

0 

0 

0 0 

1- 6 -

0 0 

o~~~~~~~~~--~~--~~--~1--~I--~1--~I--~I--~1--~J--~J~ 

2 3 4 5 6 7 8 9 10 11 12 13 

NUMBER OF TREATMENTS 

Figure 1. Weight loss of sodium hypochlorite- and distilledwater-treated samples of black shale compared to the number 

of treatments. Line at 3.0 percent, weight-loss break by 

treatment. 

80 Contributions to Mineral Resources Research, 1984 

other cementing agents can be expected to free more claysize particles. 

The sodium hypochlorite, and possibly solutions of 

other hypohalous salts, offers several advantages over 

hydrogen peroxide. When commercial household laundry 

bleach is used as the sodium hypochlorite solution, the 

cost is quite reasonable. The sodium hypochlorite is more 

stable and safer to use than is the hydrogen peroxide, and 

it is less destructive of clay minerals, an important consideration because the chemical character of the clay 

minerals may be affected if a part of them is destroyed 

by the pretreatment. An added advantage to the sodium 

hypochlorite treatment is that the exchange positions of 

the clay minerals become sodium saturated, which 

prevents flocculation and aids in dispersion of the particles (Omueti, 1980). 

REFERENCES CITED 

Anderson, J.U., 1963, An improved pretreatment for 

mineralogical analysis of samples containing organic matter: Clays and Clay Minerals, v. 10, p. 380-388. 

Cassidy, M.M., and Mankin, C.J., 1960, Clorox used in 

preparation of black shale for mineral analysis: Oklahoma 

Geology Notes, v. 20, no. 11, p. 275-281. 

Douglas, L.A., and Feissinger, Francois, 1971, Degradation of 

clay minerals by H20 2 treatment to oxidize organic matter: Clays and Clay Minerals, v. 19, no. 1, p. 67-68. 

Jones, R.L., and Beavers, A.H., 1963, Silt size whewellite and 

waddellite by soil peroxidation: Soil Science Society of 

America Proceedings, v. 27, no. 3, p. 356-357. 

Lavkulich, L.M., and Wiens, J.H., 1970, Comparison of 

organic matter destruction by hydrogen peroxide and 

sodium hypochlorite and its effects on selected mineral constituents: Soil Science Society of America Proceedings, v. 

34, no. 5, p. 755-758. -

Martin, R.T., 1954, Calcium oxalate formation in soil from 

hydrogen peroxide treatment: Soil Science, v. 77, no. 2, 

p. 143-145. 

Omueti, J.A.I., 1980, Sodium hypochlorite treatment for 

organic matter destruction in tropical soils of Nigeria: Soil 

Science Society of America Journal, v. 44, no. 4, 

p. 878-880. 

Starkey, H.C., Blackmon, P.D., and Hauff, P.L., 1984, The 

routine mineralogical analysis of clay-bearing samples: U.S. 

Geological Survey Bulletin 1563, 32 p. 

Sterne, E.J., Reynolds, R.C., Jr., and Zantop, Half, 1982, 

Natural ammonium illites from black shales hosting a 

stratiform base metal deposit, DeLong Mountains, northern Alaska: Clays and Clay Minerals, v. 30, no. 3, 

p. 161-166. 

Swanson, V.E., 1961, Geology and geochemistry of uranium 

in marine black shales, a review: U.S. Geological Survey 

Professional Paper 356-C, p. 67-112. 

van Langeveld, A.D., van der Graast, S.J., and Eisma, D., 

1978, A comparison of the effectiveness of eight methods 

for the removal of organic matter from clay: Clays and 

Clay Minerals, v. 26, no. 5, p. 361-364. 
Chapter G 

Platinum and Palladium in a 

Mafic Dike, Sanders County, Montana 

By R.E. VAN LOENEN, JOSEPH HAFFTY, and R.L. EARHART 

Hydrothermal activity has concentrated platinum 

and palladium in a mafic intrusive body of Late 

Proterozoic age 

U.S. GEOLOGICAL SURVEY BULLETIN 1694 

CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 
CONTENTS 

Abstract 83 

Introduction 83 

Acknowledgments 83 

Geologic setting 83 

Sampling and analysis 85 

Dike composition 85 

Mineralization 88 

Metal distribution 88 

Nature of platinum in dike 89 

Summary 89 

References cited 89 

FIGURES 

1. Index map showing mines and sample localities and outcrop pattern of 

Green Mountain dike 84 

2. Histograms showing the nu~ber of samples at various concentrations of 

metals 88 

TABLES 

1. Analyses of intrusive rocks from Green Mountain dike 86 

82 Contents 
CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 

Platinum and Palladium in a 

Mafic Dike, Sanders County, Montana 

By R.E. Van Loenen, Joseph Haffty, and R.L. Earhart 

Abstract 

Platinum and palladium have been concentrated by 

hydrothermal activity in altered rocks in and along fractures 

in the Green Mountain dike, Sanders County, Montana. The 

Green Mountain dike, a mafic intrusive of Late Proterozoic age, 

is intruded into metasediments of the Middle Proterozoic Belt 

Supergroup. Unaltered dike rocks contain slightly more 

platinum than palladium; however, palladium, the most mobile 

element of the platinum-group, is enricheo in altered rock over 

platinum by a factor of 2:1. Copper, which follows a similar 

pattern of enrichment in altered rocks, is closely associated to 

platinum and palladium. Sperrylite (PtAs2

) was identified in a 

sample of altered rock from the dike. 

INTRODUCTION 

The Green Mountain dike, of Late Proterozoic age, 

crops out south of the Flathead River in northwest Montana. Minor base metals (mainly copper), gold, and silver 

were produced from vein deposits in the dike. Since the 

early mining days the dike was known to contain local 

occurrences of platinum-group elements, but little has 

previously been known of their nature of occurrence or 

mineralogy. 

Vein deposits in the dike were mined during a 

40-year period from 1910 through 1949. During this time, 

recorded production amounted to 9,000 tons of ore that 

contained 1,275 oz of gold, 5, 750 oz of silver and 

1 ,400,000 lbs of copper (Crowley, 1963). The amount of 

platinum and palladium produced is not known, but these 

metals were probably recovered with gold and were included with the gold. 

Several mines were developed along the dike; 

however, most of the metals produced were from the 

Green Mountain (Drake) mine (fig. 1). 

The Green Mountain dike is about 5 mi south of 

the Flathead River along the east side of Revais Creek 

in the southeast part of Sanders County, Montana 

(fig. 1). Part of the dike is located on the Flathead 

Indian Reservation. Data for this report are partly from 

unpublished data in a 1980 U.S. Geological Survey 

administrative report to the Bureau of Indian Affairs, 

concerning the mineral and geothermal resources of the 

Flathead Indian Reservation, by R.L. Earhart, R.E. Van 

Loenen, and J. W. Whipple. The purpose of the present 

report is to present analytical data for 99 samples collected from both underground and surface exposures and 

to discuss the distribution and mode of occurrence of the 

platinum group metals. 

Acknowledgments 

We thank Zane Kelly of Polson, Mont., for assistance in collecting the samples, and G. A. Des borough 

of the U.S. Geological Survey for assistance with the electron microprobe analysis. 

GEOLOGIC SETTING 

The Green Mountain dike was intruded into rocks 

of the Burke Formation, the lowermost unit of the Ravalli 

Group of the Belt Supergroup. The Belt Supergroup in 

this region is more than 40,000 ft thick and consists of 

fine-grained clastic and carbonate rocks that were 

deposited rather continuously during Middle Proterozoic 

time. Regional metamorphism to greenschist to epidoteamphibolite facies converted the strata to argillite, siltite, 

and quartzite. Mafic sills that intruded into the Belt rocks 

during Late Proterozoic time are especially abundant in 

lower parts of the Belt Supergroup of this area. Mafic 

intrusions are more often discordant higher up in the 

stratigraphic succession. Local hornfels-facies metamorphism accompanied the emplacement of many of the 

larger sills. Major tectonic deformation that began during the Early Jurassic Epoch and continued throughout 

the Laramide orogeny resulted in a structurally complex 

folded and faulted terrane. 

In the region of the Green Mountain dike, the 

Burke Formation is folded into a large, open, northerly 

plunging anticline; several faults occur parallel to the fold 

axis; and the dike is located on the east limb of the 

Platinum and Palladium, Montana 83 
25' 24' 23' 22' 21' 

00 ~ToPerma==================================~==~~==~~================~ 

....--Location of 

figure 1 

MONTANA 

tf 

II 

II 

II v~/1 

;/II 

I! II II II 

II II 

II 

II 

II 

II ,. -- -f~mple• fmm 

surface outcrops 

)...__ 45 ', {.: f ' 89 Underground samples 

LUCKY STRIKE 

OH 1 MINE 

92<?99i t I Samples from surface 

1 - -x10 and drill core 

f GREEN MOUNTAIN 

40 (DRAKE) MINE 

15J_j25 

Samples from surface 

Yb 

-;:::====::::::---Extent 

uncertain 

91 x 90 

........-s;,:;;les from 

prospect pits 

EXPlANATION 

MIDDLE PROTEROZOIC BURKE FORMA· 

TION- Ybl, upper member; Ybu, lower 

member 

---- CONTACT 

_!_P_ STRIKE AND DIP OF BEDS 

>- ADIT 1 

DRILL HOLE 

x-J _______ l_O-x SAMPLE TRAVERSE 

15' 4r L_ __ _L _______ L_ ______ ~--------L-------~------~ 

2 MILES 

2 KILOMETERS 

Figure 1. Sample localities and mines, and outcrop pattern of the Green Mountain dike, Sanders County, Mont. 

84 Contributions to Mineral Resources Research, 1984 
anticline. The Burke Formation is divided into lower 

and upper members. The lower member is about 3,300 

ft thick (R.L. Earhart and others, unpub. data, 1980) and 

consists of gray-green argillites and siltites, some of which 

are calcareous. This member is host to low-grade stratabound copper deposits about 3 mi west of the Green 

Mountain dike. The upper member is about 4,000 ft 

thick (R.L. Earhart and others, unpub. data, 1980) and 

consists mainly of light-gray magnetite-bearing siltite, 

argillite, and quartzite. The northernmost outcrops of the 

Green Mountain dike are in the lower member of the 

Burke Formation; to the south the dike cuts up stratigraphic section into the upper member (fig. 1). 

The dike occurs in an arcuate pattern over its 3.5 mi 

of exposed length, and it is at least 335 ft thick as 

measured in a cross-cutting adit (Crowley, 1963). The 

radiometric age of the dike has not been determined, but 

the dike is probably related to the other mafic dikes and 

sills intruded into Belt terrane during Late Proterozoic 

time. Three metadiorite sills in the region that are possibly 

related in time to the Green Mountain dike were dated 

by potassium-argon methods at 803, 854, and 770 m.y. 

(Marvin and others, 1984). 

SAMPLING AND ANALYSIS 

Sampling of the Green Mountain dike was primarily designed to study the distribution of metals with respect 

to structures and to compositional variations within the 

dike. Sources of samples include outcrops, prospect pits, 

underground workings, and drill core. The approximate 

localities of the samples are shown in figure 1. Of the 

underground workings, only the Blue Ox mine was accessible at the time of the investigations. Samples were 

collected systematically at intervals from west to east 

across the width of the dike (fig. 1). Unless otherwise 

specified, all samples are single grab samples weighing 

from 1 to 2 lbs. 

Samples 1 through 10 (channel samples), collected 

from exposures in an open pit near the Lucky Strike mil)e, 

represent a 27-ft section of the lower (western) part of 

the dike. Eight samples (92-99) are from drill core provided by the U.S. Bureau of Mines in 1979. The inclined 

drill hole intersected the dike below the Lucky Strike 

mine (fig. 1). It was drilled at a 45 ° angle to a depth of 

350 ft. The samples are composites of 1- to 2-ft intervals 

of material that was selected because of megascopically 

visible copper minerals. Sample sets 11 through 25 (near 

the Green Mountain (Drake) mine) and 26 through 44 

(about 1 mi north of the Blue Ox mine) are samples from 

surface exposures spaced at large intervals across the 

width of the dike. Samples 45 through 89 were collected 

from underground workings of the Blue Ox mine. The 

mine adit, with a bearing perpendicular to strike of the 

dike, cuts the entire dike width of 340 ft. These samples 

were collected at 10-ft intervals, and represent a complete 

subsurface cross section of the dike. Samples 57 through 

62 of this set are channel samples taken from a northsouth-trending fracture zone. Ore-bearing rock from this 

zone (which is 150ft from the mine portal) was depleted 

by mining. Samples 90 and 91 are from small prospect 

pits near the northernmost exposures of the dike. 

Samples were analyzed for 50 elements by a six-step 

semiquantitative spectrographic method (table 1). Results 

of this analysis are identified with respect to geometric 

brackets whose boundaries are 1.2, 0.83, 0.56, 0.38, 0.26, 

0.18, 0.12, etc., and are reported arbitrarily as midpoints 

of these brackets: 1, 0.7, 0.5, 0.3, 0.2, 0.15, and 0.1, etc. 

The precision of a reported value is approximately plus 

or minus one bracket at 68 percent, or two brackets at 

95 percent confidence. 

Elements determined by this method and shown in 

table 1 include Co, Cr, Cu, Mo, Ni, Pb, and Ag. Other 

elements determined are not included because either they 

have no recognizable significance, or they were below 

their detection limits. Palladium, platinum, and rhodium 

(also table 1) were determined by combined fire assay and 

emission spectrography (Haffty and Riley, 1968). Gold 

was determined by fire assay and atomic absorption 

spectrography. 

DIKE COMPOSITION 

The dike rock composition, although somewhat 

obscured by alteration, ranges from dioritic to gabbroic. 

Modal analysis of a sample of relatively unaltered hornblende diorite collected underground from the Blue Ox 

adit near the center of the dike gave the following composition: quartz 6 percent, feldspar (andesine) 28 percent, 

amphiboles (hornblende and cummingtonite) 42 percent, 

ilmenite 6 percent, diopside 2 percent, chlorite and illite 

10 percent, and a trace Df apatite and malachite. Olivine 

(fayalite), although not found in this sample, is abundant 

in others. 

The Green Mountain dike is strongly altered both 

deuterically and by later hydrothermal activity along fractures. Mafic minerals are commonly altered to chlorite, 

which imparts a greenish tint to the dike. The host rocks 

(Burke Formation) adjacent to the dike are not visibly 

altered by contact metamorphism. 

Platinum and Palladium, Montana 85 
Table 1. Analyses of intrusive rocks from the Green Mountain dike 

[For sample localities see figure 1. All data in parts per million. Numbers in parentheses below elements indicate lower limits of 

determination. Leaders (---), element not looked for; N, not detected at limit of detection; L, detected but below limit of determination. Co, Cr, Cu, Mo, Ni, and Pb by semiquantitative spectrographic analysis by M.J. Malcolm and Nancy Conklin; Pd, Pt, 

and Rh determined by fire assay-emission spectrography by Joseph Haffty and A. W. Haubert; Au determined by fire assay-atomic 

adsorption spectrography by J.G. Crock, A.W. Haubert, and Joseph Haffty] 

Sample Co 

No. (5) 

1 20 

2 70 

3 50 

4 50 

5 30 

6 70 

7 70 

8 70 

9 100 

10 3 

11 70 

12 70 

13 100 

14 70 

15 70 

16 70 

17 70 

18 50 

19 70 

20 50 

21 70 

22 70 

23 70 

24 50 

25 50 

26 70 

27 70 

28 30 

29 70 

30 30 

31 70 

32 50 

33 30 

34 20 

35 20 

36 70 

37 30 

38 20 

39 70 

40 70 

41 15 

42 15 

43 70 

44 7 

45 L 

46 20 

47 15 

48 L 

49 50 

50 70 

Cr 

(1) 

30 

50 

70 

70 

30 

70 

70 

70 

100 

30 

50 

70 

20 

1 

70 

70 

70 

70 

70 

70 

70 

70 

70 

70 

70 

70 

70 

70 

70 

50 

70 

70 

30 

30 

30 

70 

30 

30 

30 

70 

15 

30 

30 

30 

50 

70 

50 

20 

70 

50 

Cu Mo 

(1) (3) 

300 N 

700 N 

7,000 N 

15,000 N 

5,000 N 

3,000 N 

1,500 N 

5,000 N 

10,000 N 

150 N 

10,000 30 

15,000 30 

2,000 N 

300 N 

500 N 

300 N 

300 N 

300 N 

300 N 

300 N 

300 N 

200 N 

300 N 

300 N 

300 N 

300 N 

300 N 

200 N 

300 N 

300 N 

200 N 

200 N 

150 N 

200 N 

300 N 

200 N 

300 N 

300 N 

500 N 

300 N 

300 N 

150 N 

300 N 

20 N 

70 7 

200 N 

50 N 

70 N 

300 N 

300 N 

Ni 

(5) 

70 

70 

300 

300 

70 

150 

100 

100 

500 

10 

200 

200 

200 

20 

100 

100 

100 

100 

100 

100 

100 

70 

70 

70 

70 

70 

150 

70 

150 

70 

70 

70 

50 

30 

30 

70 

30 

30 

70 

70 

20 

30 

70 

15 

5 

20 

20 

L 

100 

70 

Pb 

(10) 

N 

N 

50 

15 

N 

N 

N 

N 

N 

20 

N 

N 

100 

N 

20 

N 

N 

N 

N 

L 

L 

N 

N 

N 

N 

N 

N 

L 

10 

L 

N 

N 

N 

N 

L 

N 

N 

N 

N 

N 

N 

N 

N 

10 

N 

15 

50 

N 

N 

N 

Pd 

(. 004) 

0.02 

.024 

.66 

3.2 

.25 

.11 

.084 

.075 

1. 8 

.99 

.79 

.026 

.009 

.013 

.023 

.012 

.026 

.010 

.013 

.009 

.045 

.033 

.009 

.027 

.011 

.007 

.009 

.040 

.008 

.008 

.016 

.018 

.052 

.040 

.035 

.035 

.031 

.Oll 

.010 

.019 

.012 

.012 

N 

.008 

.009 

.005 

.009 

.018 

.017 

1silver determined by fire assay-atomic adsorption. 

86 Contributions to Mineral Resources Research, 1984 

Pt 

(.010) 

0.037 

.028 

.99 

.80 

.075 

.056 

.067 

.043 

• 70 

.75 

1.00 

.024 

N 

.029 

.033 

.026 

.040 

.027 

.020 

.016 

.040 

.039 

.027 

.019 

.025 

.029 

.020 

• 028 

.010 

.027 

.023 

.040 

.043 

.074 

.068 

.062 

.053 

.031 

.021 

.021 

.034 

.026 

N 

N 

N 

N 

N 

.010 

.028 

Rh 

(.005) 

N 

N 

0.023 

.062 

N 

N 

N 

N 

.026 

.018 

.035 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

Au 

(.05) 

2.9 

.73 

.07 

.06 

N 

N 

.64 

.30 

.28 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

Ag 

(.5) 

N 

N 

1. 5 

3 

N 

N 

N 

N 

1. 5 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 
Table 1. Analyses of intrusive rocks-Continued 

Sample 

No. 

51 

52 

53 

54 

55 

56 

57 

58 

59 

60 

61 

62 

63 

64 

65 

66 

67 

68 

69 

70 

7 1 

72 

73 

74 

75 

76 

77 

78 

79 

80 

81 

82 

83 

84 

85 

86 

87 

88 

89 

90 

91 

92 

93 

94 

95 

96 

97 

98 

99 

Co 

(S) 

50 

50 

70 

50 

50 

50 

70 

70 

30 

30 

30 

30 

so 

50 

50 

50 

50 

50 

50 

70 

70 

70 

30 

50 

70 

70 

50 

70 

70 

70 

so 

70 

so 

so 

so 

50 

30 

5 

30 

30 

so 

30 

20 

50 

70 

30 

30 

15 

15 

Cr 

(l) 

50 

70 

70 

70 

50 

70 

70 

150 

100 

150 

70 

70 

50 

70 

50 

50 

50 

50 

70 

50 

50 

70 

70 

50 

50 

30 

30 

50 

30 

50 

15 

20 

20 

20 

50 

30 

70 

30 

30 

20 

50 

50 

50 

100 

20 

50 

70 

30 

50 

Cu Mo 

(1) (3) 

200 N 

200 N 

700 N 

300 N 

300 N 

500 N 

30,000 30 

15,000 N 

7,000 N 

2,000 N 

2,000 N 

1,500 N 

300 N 

300 N 

300 N 

500 N 

500 N 

500 N 

500 N 

500 N 

1,000 N 

1, 000 N 

1, 000 N 

1,000 N 

1,000 N 

500 N 

500 N 

700 N 

2,000 N 

2,000 N 

300 N 

2,000 N 

500 N 

500 N 

200 N 

500 N 

150 N 

20 N 

500 N 

700 N 

20,000 N 

700 N 

300 N 

3,000 N 

30,000 N 

2,000 N 

3,000 N 

2,000 N 

1, 000 N 

Ni 

(5) 

70 

70 

70 

70 

70 

70 

700 

200 

ISO 

100 

100 

70 

70 

70 

so 

70 

70 

70 

70 

70 

70 

100 

50 

100 

100 

70 

70 

100 

100 

100 

so 

70 

70 

50 

70 

70 

50 

I5 

so 

50 

300 

70 

20 

150 

300 

50 

100 

30 

70 

Pb 

(10) 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

L 

L 

N 

N 

N 

N 

N 

150 

N 

N 

Pd 

(.004) 

0.022 

.033 

.024 

.006 

.006 

.029 

2.9 

2.2 

1. 6 

.27 

.24 

.34 

.026 

.018 

.024 

.029 

.066 

.063 

.055 

.036 

.13 

• 12 

.089 

.15 

.057 

.049 

.044 

.069 

• 15 

• 16 

.021 

.094 

.042 

.021 

.037 

.090 

.024 

.008 

.023 

.026 

1.7 

.062 

.016 

.57 

I. 0 

.20 

• 12 

.46 

.10 

Pt 

(.OIO) 

0.029 

.024 

.038 

.010 

N 

.028 

1. 2 

.91 

.34 

.13 

• 11 

.053 

.038 

.025 

.034 

.033 

.050 

.053 

.052 

.056 

.078 

.087 

.057 

.093 

.067 

.062 

.044 

.062 

.077 

.073 

.024 

.095 

.036 

.03I 

.049 

.084 

.021 

N 

.027 

.043 

.81 

.032 

.02 

.19 

.86 

.084 

.046 

.18 

.08 

Rh 

(.005) 

N 

N 

N 

N 

N 

N 

.039 

.021 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

.018 

N 

N 

N 

• 11 

N 

N 

N 

N 

Au 

(.OS) 

N 

N 

N 

N 

N 

N 

.55 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

• 12 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

.67 

N 

N 

• 12 

• 15 

.07 

N 

.13 

N 

Ag 

(.5) 

N 

N 

N 

N 

N 

N 

2 

7 

5 

1 

1 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

1 

N 

N 

• 7 

N 

N 

N 

63 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

N 

11.5 

N 

N 

L 

3 

N 

1 

N 

N 

Platinum and Palladium, Montana 87 
MINERALIZATION 

Mineralization within the dike was apparently confined to fractures, fissures and breccia zones, which 

provided the conduits for movement of hydrothermal 

fluids. These faults and shear zones, as well as altered 

rock adjacent to these structures, are enriched in 

platinum, palladium, copper, gold, silver, and nickel. 

Through the effects of weathering and hydrothermal activity, Pd, Pt, and probably Cu have been remobilized 

from the mafic dike rocks and concentrated in fractures. 

The dike has been broken and shattered by cross-cutting 

faults and shear zones, and by movement along the dikehost rock contact. Most fractured rock contains copper 

minerals. 

The most important ore production was from a 

deposit along the lower contact of the dike at the Green 

Mountain (Drake) mine (Crowley, 1963). The ore there 

was mined from fractures, quartz veins, and altered rock, 

with the richest ore in an oxide zone that consisted 

primarily of hematite and clay. The ore body is zoned vertically, according to 1976 unpublished data by the U.S. 

Geological Survey and the U.S. Bureau of Mines (on the 

status of mineral resource information for the Flathead 

Indian Reservation), which stated: ''The enriched zone 

is made up of both an oxide zone and a zone of secondary chalcocite. The oxide zone continues from the surface to a depth of about 80 ft and contains malachite, 

chrysocolla, cuprite, and tenorite. The oxide zone is 

underlain by the secondary chalcocite zone which is about 

30 ft thick. The enriched zone grades into a zone of 

primary mineralization containing chalcopyrite, chalcocite, and bornite. An appreciable amount of silver, gold, 

and platinum-palladium was recovered from both oxide 

and sulfide ores." 

The principal copper sulfide in samples collected 

for this report is chalcopyrite, which occurs in variable 

amounts and ranges from massive in fractures to fine 

disseminations in wall rocks. Other associated copper 

minerals include chalcocite, bornite, and secondary 

malachite. Pyrite and pyrrhotite closely associated with 

the enriched rock are partly oxidized to goethite. 

METAL DISTRIBUTION 

Histograms that plot element concentrations 

against numbers of samples were constructed for copper 

and the sum of the platinum-group metals (Pt + Pd + Rh) 

(fig. 2). The threshold concentration values, indicated by 

arrows on the histograms, were chosen by plotting the 

cumulative percent frequency distributions as described 

by Sinclair (1974). The threshold, the concentration value 

at which a rock is considered anomalous, was chosen at 

1,000 ppm for copper and 0.1 ppm for the platinumgroup metals. Using these somewhat arbitrary threshold 

88 Contributions to Mineral Resources Research, 1984 

values, 28 percent of all samples are enriched in both copper and the platinum-group metals, 13 percent are enriched in only the platinum-group elements and only one 

sample contains anomalous copper without enrichment 

of the platinum-group. These results clearly illustrate a 

close geochemical association of copper and platinumgroup metals. Considering platinum and palladium 

separately, each with a threshold of 0.1 ppm, palladium 

is enriched in 25 percent of the samples, whereas platinum 

is enriched in only 15 percent. 

McCallum and others (1976), in a study of Pt-PdCu ores of the New Rambler mine (Wyoming), stated that 

the most significant evidence for a strictly hydrothermal 

origin is the predominance of the more readily soluble 

platinum-group elements Pd and Pt (Pd being most 

mobile), and a relative lack of the more inert platinumgroup elements lr, Os, Ru, and Rh. Ores from shear 

zones in metagabbroic rocks at the New Rambler mine 

have a Pd:Pt ratio of 18:1 and a Pd:Pt:Ir+Ru+Os 

(minor platinum-group elements) of nearly 1800:100:1 

(McCallum and others, 1976). The distribution of 

platinum-group elements in the Green Mountain dike 

follows similar patterns, even though concentrations of 

the metals in the dike are much lower. The Pd:Pt ratio 

for samples of enriched rock averages 2:1 with one value 

of 6:1. Samples of unaltered gabbroic and dioritic dike 

rocks have a Pd:Pt ratio which averages 0.8: 1. The 

average Pd:Pt:Rh ratio for enriched rock is 60:30:1. 

25 

~ :E 20 

~ 

l:5 15 

r::r::: 

..... 

~ 10 ::I 

z 5 

0 o.n o.n c -.-NC"')~,..... &.n 

dddddd-,...;NMo.n,...co.ncc V .--NM 

COPPER, IN PPM x 1000 

TOTAL Pt-GROUP ELEMENTS, IN PPM 

Figure 2. Histograms showing number of samples at various 

concentrations of metals. Arrow indicates threshold value. 

Samples with metal content at arrow and above are considered 

anomalous (enriched). 
These ratios, although not as high as those of the New 

Rambler mine, are inferred to reflect migration and concentration of the more soluble platinum-group metals by 

hydrothermal fluids. 

Three samples with relatively high concentrations 

of Pd and Pt (Nos. 11, 12 and 57) were further analyzed 

for iridium and ruthenium; however, neither one was 

detected at the respective lower limits of detection of these 

elements (0.03 ppm and 0.1 ppm). 

Chromium and cobalt are present in very low 

amounts (trace to 100 ppm) and show no correlation or 

increased concentrations in the platinum-palladiumenriched rock. 

Nickel, although present in very small amounts, can 

be correlated with enriched rock. Values ranged from 

10 ppm in unaltered rock to 700 ppm in enriched rock. 

Gold and silver show a similar association to enriched rock in that none was detected in unaltered rock. 

Lead and zinc do not appear to be associated with 

precious-metal mineralization. Zinc did not exceed the 

detection limit of the method (200 and 500 ppm-a higher 

detection limit was required by some samples) in any of 

the samples. Samples 92-99 included from 100 to 

200 ppm zinc (determined by atomic absorption 

methods). 

Ten samples contained detectable lead that ranged 

from <10 ppm to 100 ppm, but these samples apparently were not enriched in platinum-group metals. Only a 

few pounds of lead reportedly were shipped from the 

district during mining operations (Crowley, 1963). 

NATURE OF PLATINUM IN DIKE 

Several polished sections of dike rock that contained the highest amounts of platinum were examined 

microscopically. The highly reflective grains were further 

analyzed using an electron microprobe in attempts to 

identify platinum minerals. A single grain of sperrylite 

(PtAs2

) was found in sample 91. The sample was collected from a small prospect pit near the north end of 

the dike (see fig. 1). The sperrylite grain is pear-shaped, 

about 0.004 mm by 0.002 mm, and occurs in fractured 

gabbro that contains 2 percent copper as disseminated 

chalcopyrite. Heavy concentrates, separated from several 

metal-enriched samples, were analyzed with the electron 

microprobe; however, no platinum minerals were found. 

This paucity of microscopically visible platinumgroup minerals suggests that at least some of these 

elements may be incorporated in the lattice of other 

minerals. 

SUMMARY 

The distribution of platinum-group metals and 

associated copper-gold-silver minerals in the Green 

Mountain dike is controlled by faults and shear zones that 

cut across the dike or parallel the dike contacts. These 

faults and shear zones, as well as altered rock adjacent 

to these structures, are enriched in platinum, palladium, 

copper, gold, silver, and nickel. Through the effects of 

weathering and hydrothermal activity, Pd, Pt, and probably Cu have been remobilized from the mafic dike rocks 

and concentrated in fractures. Anomalous rocks, defined 

as those containing 1 ,000 ppm or more copper, are enriched in both Pd and Pt. Furthermore, Pd, the most 

mobile of the Pt-group elements, is enriched over Pt by 

a factor of 2:1. Unaltered dike rocks contain slightly more 

platinum than palladium. 

Analytical results of samples collected across the 

dike indicate a lack of metal association with any layered 

compositional differentiate that might be present in the 

dike. Trace amounts of the same metals found in fractured zones are apparently evenly distributed throughout 

unfractured dike rock. 

REFERENCES CITED 

Crowley, F.A., 1963, Mines and mineral deposits (except fuels), 

Sanders County, Montana: Montana Bureau of Mines and 

Geology Bulletin 34, p. 37. 

Haffty, Joseph, and Riley, L.B., 1968, Determination of 

palladium, platinum and rhodium in geologic materials by 

fire assay and emission spectrography: Talanta, v. 15, 

p. 111-117. 

Marvin, R.F., Zartman, R.E., Obradovich, J.D., and Harrison, 

J.E., 1984, Geochronometric and lead isotope data on 

samples from the Wallace 1 °X 2 ° quadrangle, Montana 

and Idaho: U.S. Geological Survey Map MF-1354-G, scale 

1:250,000. 

McCallum, M.K, Loucks, R.R., Carlson, R.R., Cooley, E.F., 

and Doerge, T.A., 1976, Platinum metals associated with 

hydrothermal copper ores of the New Rambler mine, 

Medicine Bow Mountains, Wyoming: Economic Geology, 

v. 71, no. 7, p. 1443. 

Sinclair, A.J., 1974, Selection of threshold values in geochemical 

data using probability graphs: Journal of Geochemical Exploration, v. 3, no. 2, p. 129-149. 

Platinum and Palladium, Montana 89 
Chapter H 

Metalliferous Oil Shale in the 

Upper Devonian Gibellini Facies of the 

Woodruff Formation, Southern Fish 

Creek Range, Nevada 

By GEORGE A. DESBOROUGH, FORREST G. POOLE, 

RICHARD K. HOSE, and GREGORY N. GREEN 

The Upper Devonian Gibellini facies of the Woodruff Formation 

consists of marine oil shale that contains anomalous amounts 

of vanadium, zinc, selenium, molybdenum, and nickel 

U.S. GEOLOGICAL SURVEY BULLETIN 1694 

CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 
CONTENTS 

Abstract 93 

Introduction 93 

Acknowledgments 94 

Methods of study 94 

Stratigraphic and tectonic relations 98 

Rock constituents 98 

Organic constituents and fossils 98 

Distribution and concentration of metals 102 

Conclusions 103 

References cited 104 

FIGURES 

1. Map locating Gibellini and Bisoni properties and Bisoni-McKay property, 

Nevada 94 

2. Geologic map and block diagram of the Gibellini and Bisoni properties, 

southern Eureka County, Nevada 96 

3. Photomicrographs of samples, showing texture and fabric of organic 

fragments, opaline silica, and sulfides 100 

TABLES 

1. Concentration of selected elements for composite samples from drill holes, 

Gibellini vanadium property 95 

2. Oil yield and bulk mineralogy for composite samples from drill holes, 

Gibellini vanadium property 99 

3. Concentration of vanadium, zinc, selenium, molybdenum, and silver in composite samples from drill holes, Gibellini and Bisoni-McKay properties 102 

: .. 
CONTRIBUTIONS TO MINERAL RESOURCES RESEARCH, 1984 

Metalliferous Oil Shale in the 

Upper Devonian Gibellini Facies of the 

Woodruff Formation, Southern Fish Creek Range, Nevada 

By George A. Desborough, Forrest G. Poole, Richard K. Hose, 1 and Gregory N. Green 

Abstract 

The Gibellini facies of the Woodruff Formation is a 

kerogen-rich marine sequ.ence, as much as 60 m thick, of Upper Devonian siliceous mudstone, siltstone, and chert. It has 

been evaluated in two areas in Nevada for its potential 

resources of vanadium, zinc, selenium, molybdenum, and sync rude oil. In one of the areas studied, an exotic block 

(>1.0x 106 m3) contains strongly deformed allochthonous 

eugeosynclinal Devonian siliceous rocks of the Gibellini facies 

structurally overlying deformed allochthonous eugeosynclinallike Early Mississippian siliceous rocks (Bisoni facies of the 

Webb Formation). Both siliceous facies within this exotic block 

structurally overlie less deformed autochthonous Mississippian 

Antelope Range Formation (lower part) and are overlapped 

locally by autochthonous Mississippian conglomerates of the 

upper Antelope Range Formation. 

The Gibellini facies in both study areas contains 

disseminated grains and framboids of pyrite which may be 

primary (early diagenetic), but pyrite and marcasite in 

chalcedonic quartz vein lets appear to have formed during later 

diagenesis. Organic matter in the Gibellini facies consists mainly 

of amorphous, flaky and stringy sapropel. Recognizable organic 

remnants include radiolarian tests, sponge spicules, conodonts, 

brachiopod shells, algae, and humic debris. 

Most fresh black rock of the Gibellini facies is low-grade 

oil shale. Black solid bitumen and liquid oil commonly fill voids 

and microfractures. Gas chromatographic analysis of the 

saturated hydrocarbon fraction indicates a complex mixture 

dominated by branched and cyclic compounds. Conodont and 

palynomorph color alteration, vitrinite reflectance, and RockEva! pyrolysis data suggest that the organic matter in the rock 

is thermally immature and has not been subjected to 

temperatures greater than 60 °C since deposition in Devonian 

time. All these characteristics are consistent with the interpretation of a low-temperature and a shallow-burial history for the 

Gibellini facies on ridge 7129 in the Cockalorum Wash 

15-minute quadrangle, southern Fish Creek Range, Eureka 

County, Nevada. 

Vanadium in fresh black rocks of the Gibellini facies obtained from drill holes and fresh exposures in trenches and 

1

ARCO Exploration Company, Box 5540, Denver, CO 80217. 

roadcuts occurs chiefly in organic matter. Metahewettite is the 

principal vanadium mineral in the oxidized zone, but it also 

occurs sparsely as red and pink nodules and fillings of 

microfractures in unweathered strata. High concentrations of 

vanadium occur in unoxidized subsurface samples and in oxidized and bleached rock that is prevalent at the surface. An 

observed tendency toward enrichment is due to surficial 

weathering and ground-water movement. Zinc which occurs 

as sphalerite is in high concentrations in unoxidized rock, but 

is depleted in oxidized rock due to leaching and weathering. 

Selenium occurs in organic matter. Concentrations of selenium 

in unoxidized rock are low, whereas in oxidized and leached 

rock they are enriched by oxidation. Molybdenum occurs principally in organic matter, but molybdenite is present. Concentrations of molybdenum are higher in unoxidized rock than 

in oxidized rock, indicating depletion upon weathering. Nickel 

concentrations are high in unoxidized strata but are low in the 

oxidized zone. 

INTRODUCTION 

The particular objective of this study was to determine the concentration and residence of certain metals 

relative to organic and inorganic components of a lowtemperature shale and chert sequence in central Nevada. 

Many Paleozoic dark-colored marine mudstones 

and siltstones (black shales) contain significant 

syngenetic-early diagenetic metal concentrations as well 

as sufficient syncrude oil to qualify as oil shales. In addition, they commonly host epigenetic metal deposits, and 

serve as significant source rocks for petroleum. As black 

shales commonly form several-meter-thick laterally persistent units, they potentially could serve as vast low-grade 

metal resources. Additionally, the rock masses are either 

low-grade oil shales or have already generated and expelled petroleum that has migrated elsewhere. For these 

reasons, it is important to identify the mineralogical, 

chemical, depositional, diagenetic, and tectonic factors 

that collectively govern the size, grade, and distribution 

of the metalliferous and petroliferous deposits, as these 

Metalliferous Oil Shale, Nevada 93 
factors will influence the choice of extractive process or 

processes necessary to recover these resources. 

The economic potential of marine black shales in 

the Western United States is difficult to evaluate for a 

number of reasons: 

1. The strata commonly are in allochthonous blocks 

above regional thrust faults, and generally are intensely deformed and broken. Both vertical and lateral continuity is virtually unpredictable without subsurface 

data, and reliable resource projections are generally 

limited to small, well-studied areas. Tectonic fracturing promotes accelerated weathering arid erosion so 

that the shales are generally poorly exposed; in addition, the broken rocks are extensively leached by 

ground water and tend to be severely altered to substantial depths. 

2. The rocks are generally fine to very fine grained 

(microclastic to microcrystalline), and megascopic 

field methods generally are ineffective in mineral or 

chemical identification. Even petrographic examination of standard thin sections is of limited value. 

3. The sapropelic organic matter in these sediments is 

widely dispersed, and appears virtually opaque in 

standard thin sections. 

4. The elements are so extensively dispersed that 

chemical and mineralogical studies must be conducted 

on a microscale. 

Vanadium, zinc, selenium, and molybdenum occur 

in anomalously high concentrations in the Upper Devonian kerogenous marine Gibellini facies at the Gibellini 

and Bisoni properties on ridge 7129 in the Cockalorum 

Wash 15-minute quadrangle, southern Fish Creek Range, 

Eureka County, Nev. (fig. 1). The ridge is located 

0.48 km west of the Gibellini manganese mine and about 

4 km northwest of Eightmile well. The Gibellini facies 

at the Bisoni-McKay property 14.5 km to the south also 

is enriched in these metals. 

Acknowledgments 

We are particularly indebted to Geoffrey Snow of 

Noranda Exploration Inc. for providing drill-cutting 

samples, vanadium analytical data, maps, and diagrams 

that were all essential for our study and interpretations. 

We also thank Richard Schoenfeld of the U.S. Geological 

Survey for drafting the geologic map and block diagram 

(fig. 2). Ronald Keil of Geolabs Inc. provided quantitative 

analyses of zinc, selenium, and silver. Claudia True, 

formerly of the U.S. Geological Survey, provided data 

and interpretations on vitrinite reflectance of kerogen in 

sample 122. 

94 Contributions to Mineral Resources Research, 1984 

EUREKA COUNTY 

I 

\ 

I 

\ , 

J 

I 

( WHm: PINE §j 

Gibellini and Bisoni vanadium 

properties on ridge 7129 

I 

-......!_. ... 

NYE 

COUNTY 

81son1-McKay .. ~ 

vanadium 

property 

Figure 1. Location of Gibellini and Bisoni vanadium properties on ridge 7129, southern Eureka County, Nev., and BisoniMcKay vanadium property, northern Nye County, Nev. 

METHODS OF STUDY 

Vanadium has been known in high concentrations 

in the Gibellini facies in the southern Fish Creek Range 

for many years; in addition, analyses of drill cuttings and 

fresh outcrop samples in U.S. Geological Survey and 

other laboratories have revealed anomalously high concentrations of zinc, selenium, molybdenum, and nickel 

(Desborough and others, 1979; Bohlke and others, 1981). 

Extensive drilling and sampling of the Gibellini 

property by Noranda Exploration Inc. provided the subsurface samples, and Dr. Geoffrey Snow of Noranda 

generously provided analytical data for vanadium. Splits 

of samples from two drill holes, NG-28 and NG-12 

(fig. 2), have been quantitatively analyzed for several 

metals (table 1). The Fischer oil-assay yield and bulk 

mineralogy of these samples have also been determined. 

Samples of twelve 5-ft (1.5-m)2 composites from 

three widely spaced drill holes on the Bisoni McKay property and three 5-ft (1.5-m) composites from two drill 

2Measurements originally in feet, converted to metric for this 

study. 
Table 1. Concentration of selected elements for composite samples from drill holes NG-28 and NG-12, Gibellini 

vanadium property, Eureka County, Nev. 

[Quantitative X-ray fluorescence analysis by G.N. Green; values in parts per million. Original depth measurements were in feet; 1 

ft=0.3048 m] 

Depth v Fe Ni Zn (f t) 

Drill 

75-85 1 '960 6,320 (100 670 

85-95 2,350 10,300 210 3,360 

95-105 2' 180 11 '000 220 6,700 

105-115 2,410 11,800 610 8,260 

115-125 2,300 11,800 340 4,580 

125-135 ·2,460 13,400 260 3,920 

135-145 1,850 20,500 300 4,530 

145-155 2,240 20,600 475 7,290 

155-165 2,350 11,100 525 6,910 

165-175 1,500 23,700 235 3,299 

Drill 

100-110 4,200 3,950 <100 <100 

110-120 3_,860 5,530 (100 <100 

120-130 4,420 10,300 <100 (100 

130-140* 3,250 6,320 <100 135 

140-150* 3,700 7,900 135 1 '430 

150-160* 3,360 15,000 410 4,140 

160-170 2,300 22' 100 1,330 18,800 

170-180 1,620 13,400 480 8,830 

180-190 2,130 14,200 1,030 13,900 

190-200 1,900 16,600 575 8,690 

* Zone of less intense oxidation. 

holes on the Gibellini property also were analyzed for 

vanadium, zinc, selenium, molybdenum, and silver. 

Rock chips of subsurface samples from the Gibellini 

and Bisoni properties have been polished and studied by 

reflected light microscopy and electron microprobe 

methods. Fresh outcrop samples were obtained from 

rocks exposed by blasting and bulldozing (fig. 2, locality 

122). Incipiently oxidized surface samples were obtained 

from rock outcrops (fig. 2, localities 116 and 120). These 

surface samples provided fragments large enough to 

evaluate gross texture and mineralogy, as well as material 

sufficient for organic-geochemical, microprobe, X-raydiffraction, palynologic, micropaleontologic, vitrinitereflectance, fluorescence-microscopic, Rock-Eval 

pyrolysis (Espitalie and others, 1977), oil yield, and 

maturation studies. 

Methods for elemental analysis included atomic absorption and X-ray fluorescence. Vanadium, zinc, and 

As Se Rb Sr Zr Mo u 

hole NG-28 

20 70 30 120 40 80 20 

40 60 40 180 60 110 40 

40 55 35 120 so 105 20 

40 55 30 145 45 125 so 

so so 30 120 45 115 25 

40 65 45 190 65 105 45 

30 45 70 275 70 70 25 

55 55 60 210 65 90 25 

45 so 30 170 so 120 35 

35 20 90 170 100 60 20 

hole NG-12 

<20 275 <20 140 20 20 60 

20 455 25 130 35 15 so 

30 265 40 90 40 65 45 

20 400 25 100 25 35 40 

40 415 20 130 25 90 45 

55 190 6p 230 60 130 35 

105 85 65 175 100 700 55 

45 55 25 90 35 215 40 

55 60 30 80 so 415 so 

80 60 35 125 55 195 35 

selenium were analyzed by both methods, and good 

agreement was obtained. Iron, nickel, arsenic, rubidium, 

strontium, zirconium, molybdenum, and uranium were 

· analyzed using bulk samples, an energy dispersive detector with a Cd 109 excitation source, and prepared 

standards. The concentration of the following elements 

is less than the following values, in ppm: Mn 150, Cu 

150, Y 40, Pb 25. Silver was analyzed by atomic absorption analysis. 

Owing to the microscopic grain size and microfabric 

of the siliceous sapropelic mudstones, siltstones, and 

cherts of the Gibellini facies, most megascopic methods 

of mineral and organic-matter identifications are inadequate. The microscopic grain size also limits the use of 

standard petrographic thin sections in identifying mineral 

and organic matter. Most of the organic material is amorphous and cannot be identified even with high 

magnification. 

Metalliferous Oil Shale, Nevada 95 
1.0 

"" 

(""') 

0 

~ c;.: 

s. 

o· 

~ 

0 

~ 

:;· 

ltl 

~ 

~ 

ltl 

"' 0 

:; 

l"l 

~ 

~ 

~ ltl 

~ l"l 

~ 

... 

1.0 

co 

"'" 

39° 

12' 

30" 

0 0.5 KILOMETERS 

A 

EXPlANATION 

Unconsolidated colluvium 

~ and alluvium (Quaternary) 

ALLOCHTHONOUS AUTOCHTHONOUS 

~ Upper member . Lower member 

Webb Formation Antelope Range Formation 

Bisoni facies (Mississippian) 

(Mississippian) 

~ 

~ 

UNCONFORMITY 

Woodruff Formation 

Gibellini facies 

(Devonian) 

Devils Gate Limestone 

(Devonian) 

u 

D 

.....___... 

122. 

~ 

NG12 0 

Contact 

Bay State Dolomite 

(Devonian) 

Fault-Dashed where inferred, dotted where 

concealed; letters show relative vertical displacement 

Thrust fault-Dashed where inferred; dotted 

where concealed: queried where depth or 

extent uncertain; sawteeth on upper plate 

Claim boundary-Approximately located 

Road 

Sample locality 

Inclined shaft 

Drill hole 

Figure 2 (above and facing page). Gibellini and Bisoni properties and vicinity, southern Fish Creek Range, Eureka County, Nev. A, Geologic map (modified 

from Hose, 1983). B, Exploded block diagram showing subsurface detail of Gibellini and Bisoni vanadium and zinc deposit, ridge 7129. 
N 

X 

J)~ z 1- 0 w i= w 

u.. < 1%: 

~ en ~ w 

(!I 

(!I ~ < X 0 w 

...J 

< 

~#Ill 

(.) 

i= 

1%: 

w ~ 

> 

Metalliferous Oil Shale, Nevada 97 
STRATIGRAPHIC AND TECTONIC RELATIONS 

The Upper Devonian Gibellini facies of the 

Woodruff Formation at the Gibellini and Bisoni properties occurs as a northwest-trending allochthonous mass 

of organic-rich siliceous mudstone, siltstone, and chert, 

whose approximate dimensions are 700 m long, 300m 

wide, and a minimum of 60 m thick (fig. 2). This mass, 

ridge 7129, has a maximum relief of about 90 m. The 

maximum thickness of the rock sequence is not known 

because the stratigraphic base and top have not been 

established, owing to structural complexities. Surface 

geologic relations, however, indicate that locally the maximum thickness may be as much as 100m. On ridge 7129, 

the Gibellini facies is underlain structurally by the Bisoni 

facies of the Lower Mississippian Webb Formation, 

which consists chiefly of dolomitic and (or) argillaceous 

siltstone, siliceous mudstone, chert, and subordinate 

limestone, sandstone, and grit. Field studies and drill-hole 

information indicate that the vanadiferous Gibellini facies 

of the Woodruff Formation caps ridge 7129 and is 

separated by a low-angle thrust fault from the underlying nonmineralized Bisoni facies of the Webb Formation 

(fig. 2). Rocks of the Gibellini and Bisoni facies rest structurally on flysch deposits of the autochthonous lower 

Antelope Range Formation of Mississippian age. The 

Antelope Range Formation in the southern Fish Creek 

Range is underlain by the Devonian Devils Gate 

Limestone (fig. 2; Hose, 1983). On the southeast side of 

ridge 7129, a depositional contact between the 

allochthonous Bisoni facies and overlapping molasse conglomerate beds of the upper Antelope Range Formation 

(fig. 2) suggests that the combined Gibellini and Bisoni 

facies form an exotic block (>LOX 106 m3) that detached from the eastern front of the moving Roberts 

Mountains allochthon in Early Mississippian time and slid 

eastward onto deposits of the lower Antelope Range Formation in the flysch trough (Poole and others, 1983). This 

exotic block subsequently was overlapped by coarse 

molasse deposits of the upper Antelope Range 

Formation. 

At the Bisoni-McKay property (fig. 1) about 14 km 

to the south, the stratigraphic and structural relations between the Gibellini facies of the Woodruff Formation and 

adjacent rocks are less clear. Moreover, the Gibellini 

facies is more strongly deformed and consequently 

resource evaluation is more difficult. Here the Gibellini 

facies occurs as a north-northeast-trending, cigar-shaped, 

allochthonous mass of organic-rich siliceous mudstone, 

siltstone, and chert approximately 2,400 m long, 70 m 

wide, and 80 m thick. 

98 Contributions to Mineral Resources Research, 1984 

ROCK CONSTITUENTS 

The bedded siliceous sediments in the Gibellini 

facies of the Woodruff Formation on ridge 7129 consist 

largely of siliceous mudstone, siliceous siltstone, and 

banded chert layers. Unaltered strata are olive black, olive 

gray, brownish black, and brownish gray, whereas altered 

(weathered) strata are generally lighter shades of gray and 

brown, and locally (where intensely oxidized) yellow and 

red owing to the presence of secondary vanadium, selenium, and iron minerals. Fresh rocks in excavations oxidize 

rapidly as a result of surficial weathering, and a pinktinted coating of metahewettite (CaO · 3V20 5 • 9H20) 

forms at the surface and usually persists for several years. 

Rocks that have been exposed for longer periods lose this 

coating as the vanadium is dissolved and carried away. 

The bedded siliceous mudstones, siltstones, and 

cherts of the Gibellini facies are irregularly laminated and 

color banded, and many beds appear to be diagenetically silicified fine-grained turbidites. Pyrite, marcasite, and 

sapropel are common constituents. Most fresh rock is 

fractured and veined with bluish-white chalcedonic quartz 

(fig. 3G) and sparse black solid bitumen. Some of the 

beds contain thin (<2.5 em) layers of healed diagenetic 

breccia. Dark-colored laminae contain amorphous 

organic material. Sparse, poorly preserved algal cysts, 

radiolarians, and sponge spicules are the characteristic 

fossils seen in these siliceous rocks. The silica in the rocks 

is mostly microcrystalline, but it occurs also as fillings 

of algal cysts (fig. 3A-D). 

The Gibellini facies in the southern Fish Creek 

Range contains no calcareous fossils and little carbonate 

of any type except minor carbonate-fluorapatite (table 

2). The siliceous mudstones and siltstones are transitional 

into cherts by increase of silica and decrease in clay and 

sapropelic material. Most chert beds have wavy layering 

and locally are nodular; some look like diagenetically 

silicified very fine grained turbidites. Except for the lOA 

mica and some quartz, all the minerals listed in table 2 

are considered to be authigenic. In addition to 

chalcedonic quartz, the major authigenic minerals in the 

siliceous sediments on ridge 7129 are sphalerite, apatite, 

potassium feldspar, pyrite, and iron-chlorite (table 2). 

Marcasite and molybdenite are also present, but in small 

amounts, and were not detected by X-ray diffraction. 

Organic Constituents and Fossils 

The dark color of the siliceous mudstone, siltstone, 

and chert of the Gibellini facies results from black, / 
Table 2. Oil yield obtained by modified Fischer assay analysis, 

and bulk mineralogy obtained by X-ray diffraction for composite 

samples from drill holes NG-28 and NG-12, Gibellini vanadium 

property, Eureka County, Nev. 

[K, potassium; Fe, iron; original measurements were in feet; 1 ft = 0.3048 

m] 

Depth Oil 

( f t) (wt.pct) 

75-85 2.8 

85-95 3.0 

95-105 3.6 

105-115 3.3 

115-125 3.2 

125-135 3. 1 

135-145 2. 1 

145-155 2.4 

155-165 2.8 

165-175 1.3 

100-110 (.5 

110-120 (.5 

120-130 (.5 

130-140* 1.6 

140-150* 3.5 

150-160* 2.3 

160-170 1.6 

170-180 3. 1 

180-190 2.4 

190-200 2.5 

Minerals in addition to quartz 

(listed in order of abundance) 

Drill hole NG-28 

lOA mica, K-feldspar, apatite, 

Fe-chlorite. 

Sphalerite, K-feldspar, Fe-chlorit e , 1 OA mica. 

Sphalerite, K-feldspar, 0pyrite, apatite, Fe-chlorite, lOA mica. 

Sphalerite, K-feldspar, pyrite, apatite, lOA mica, Fe-chlorite. 

Sphalerite, K-feldspar, pyrite, apatite, lOA mica. 

Sphalerite, K-feldspar, lOA mica, 

pyrite, apatite, Fe-chlorite. 

lOA mica, sphalerite, K-feldspar, 

pyrite, apatite, calcite, Fechlorite. 

Sphalerite, lOA mica, pyrite, Kfeldspar, apatite, Fe-chlorite. 

Sphalerite, apatite, pyrite, Fechlorite, lOA mica, calcite. 

lOA mica, sphalerite, K-feldspar, 

Fe-chlorite, marcasite. 

Drill hole NG-12 

0 

Apatite, lOA mica. 

Apatite, lOA mica. 

lOA mica. 

Apatite, lOA mica. 

Apatite, lOA mica, sphalerite. 

lOA mica, Fe-chlorite, sphalerite, pyrite, calcite. 

K-feldspar, sphalerite, lOA mica, 

pyrite, Fe-chlorite. 

Apatite, sphalerite, lOA mica, 

Fe-chlorite, 0pyrite. 

Sphalerite, lOA mica, Fechlorite, pyrite, K-feldspar, 

apatite. 

Sphalerite, pyrite, K-feldspar, 

lOA mica, Fe-chlorite, apatite. 

* Zone of less intense oxidation. 

sapropelic (principally algal), organic material. In fresh 

rock the organic matter is dominantly very fine grained 

kerogen consisting of flaky ·and stringy organism 

fragments less than 15 J.tm (micrometers) in length, 

preferentially oriented paraJlel to bedding. Most of the 

organic matter examined microscopically is translucent 

and yellowish brown. Most of the kerogen is amorphous, 

probably owing to microbial alteration during sedimentation and early diagenesis, and only sparse organic remains can be recognized. Organic matter is generally 

abundant and forms as much as 12 weight percent of the 

rock. Round and flattened algal cyst fragments commonly are filled with opaline quartz (fig. 3A -D). Liquid oil 

was seen on some fracture surfaces in freshly broken 

unoxidized rock. 

Surface sample 122 from ridge 7129 (figs. 2, 3) contains 11.5 weight percent organic carbon; 0. 7 percent of 

the rock consists of chloroform-extractable bitumen 

(7,300 ppm), and about 10.8 percent consists of carbon 

in kerogen. Less than 10 percent of the total organic matter consists of vitrinite, liptinite (mostly sporinites), and 

inertinite. Although most vegetal components are generally less than 15 J.tm in length, algal cyst fragments about 

50-150 J.tm in diameter were found scattered throughout 

all sections studied (fig. 3A-D). Rock-Eval pyrolysis 

results on sample 122 were as follows: 13.9 weight percent organic carbon, 62.05 mg/g S1 + S2, 0.05 S/S1 + S2, 

423 hydrogen index, 16 oxygen index, and 416 °C maximum pyrolytic temperature. Six additional unoxidized 

samples, including two from bulldozer cuts and four from 

drill cuttings, from both the Gibellini and Bisoni and 

Bisoni-McKay properties, were analyzed by Rock-Eval 

pyrolysis in order to characterize and evaluate the degree 

of thermal maturity of the organic matter. Poole and 

Claypool (1984, p. 209, table 1) reported these data; they 

concluded that these strata were not subjected to thermal regimes above 60 °C since deposition. The bitumen 

in sample 122 is 59 percent hydrocarbon with the balance 

being nonhydrocarbon asphaltic material. The sample 

contains 1,690 ppm saturated hydrocarbons and 2,620 

ppm aromatic hydrocarbons with a saturated-to-aromatic 

ratio of 0. 7 and a hydrocarbon-to-organic carbon ratio 

(HC:C) of 3.76. The saturated-to-aromatic hydrocarbon 

ratio of 0. 7 is consistent with the dominantly asphaltic 

content of the extract. Generally, HC:C values greater 

than 1 suggest either thermochemical conversion of solid 

organic matter to hydrocarbons or the presence of nonindigenous hydrocarbons. According to G .E. Claypool of 

the U.S. Geological Survey (written commun., 1978), gas 

chromatographic analysis of the saturated hydrocarbon 

fraction indicates a complex mixture dominated by 

branched and cyclic compounds, with the bulk of the 

material in the range of n-C15 to n-C25 • The bulk rock 

yielded 51.7 L/t (liters/metric ton) of syncrude oil by 

Fischer assay. Total sulfur in sample 122 is 2.16 weight 

Metalliferous Oil Shale, Nevada 99 
100 Contributions to Mineral Resources Research, 1984 
percent. Conodonts in the sample are pale yellow indicating a conodont color alteration index (CAl) of 1, 

which suggests less than 60 percent fixed carbon and a 

maximum rock temperature of less than 60 °C since 

deposition (Epstein and others, 1977). The vitrinite reflectance mean value of 0.43 (n=23) for surface sample 122 

is compatible with the conodont and palynomorph color 

alteration and other organic geochemical data (Poole and 

Claypoql, 1984, table 1, loc. 23). These data indicate that 

the organic matter in the rock is immature and has not 

reached the postdiagenetic stage of catagenesis. G.E. 

Claypool (written commun., 1979) noted that red components of the hydrocarbon extract from sample 122 were 

retained on a silica-gel column during elution with n-heptane, but were partially removed during benzene elution. 

He considered the red material to be a vanadyl prophyrin 

or a selenium organic complex. 

A composite sample of cuttings from drill hole 

NG-12 on ridge 7129 (fig. 2) taken 35.5 to 58 m below 

the surface averages 7.9 weight percent organic carbon 

and 1.5 weight percent sulfur. Rock-Eval pyrolysis results 

on the composite sample were as follows: 7.5 weight percent organic carbon, 35.32 mg/g S1 +S2, 0.04 S/S1 +S2, 

454 hydrogen index, 13 oxygen index, and 417 °C maximum pyrolytic temperature. The composite sample contains 5,000-5,600 ppm extractable hydrocarbons 

(saturated and aromatic fractions). The organic matter 

seen microscopically in cuttings is very fine grained, 

averaging 2-6 J.tm in maximum dimensions, although 

silica-filled cyst fragments 50-150 #'m in diameter were 

found in polished sections of the cuttings. The organic 

material occurs principally as pervasively disseminated 

amorphous debris that is intimately admixed with the 

small mineral grains. Unoxidized mudstone yields 1.6-3.6 

weight percent (about 17-37 L/t) of oil (table 2) by 

Fischer-assay analysis. During Fischer-assay pyrolysis, a 

high content of water was derived from both dehydration of "opaline" silica and kerogen. Silica dehydration 

is verified by electron microprobe studies, during which 

the electron beam leaves a decomposition hole in the 

Figure 3 (facing page). Photomicrographs of samples 121 and 

122 showing texture and fabric of organic fragments, opaline 

silica, and sulfides. Bedding is parallel to bar scale on photos; 

all but G in reflected light, bar scale is 100 JLm. A, Core of 

opaline silica filling cyst fragment; 8, Fluorescence photo of 

cyst in A; C, Opaline silica filling flattened cyst fragment; 0, 

Fluorescence photo of cyst fragment in C; E, Sphalerite (s), 

pyrite (p), and marcasite (m) in opaline silica veinlet; F, Finegrained sphalerite disseminated in opaline silica bleb (black 

areas are voids due to plucking); G, Opaline silica vein let containing euhedral pyrite (p) (transmitted polarized light, crossed 

polars, polished thin section); H, Marcasite (m) framboid and 

sphalerite (gray area in lower part of photo). 

Si02 • nH20 phase, which does not happen with quartz. 

This phenomenon is the reason for using the term silica, 

although the X-ray diffraction pattern is that of quartz. 

Surface sample 120 on ridge 7129 (fig. 2) is oxidized 

and most of the organic material has been removed (0.21 

weight percent organic carbon). The sample is olive-gray 

nodular chert that weathers yellowish to pinkish gray and 

light brownish gray. The rock contains both marcasite 

and pyrite, and is strongly fractured, with most of the 

fractures healed with silica. In reflected-light examination under crossed nicols, round to subhedral grains of 

marcasite and pyrite can be seen that range in diameter 

from about 50 to 500 J.tm. Also, unoriented bluish-white 

veinlets of opaline-like silica (chalcedonic quartz) transect 

the rock. Red and yellow grains of metahewettite as long 

as 50 J.tm are scattered through the chert, and metahewettite also was seen in a veinlet 25 J.tm wide. 

Organic geochemical data on two outcrop samples 

from the Bisoni-McKay property are given in table 1 

(locality 24) of Poole and Claypool (1984). Their RockEva! results are comparable to those we obtained for the 

Gibellini facies at ridge 7129. 

Radiolarian tests and sponge spicules are locally 

abundant in the siliceous rocks of the Gibellini facies. 

Sparse conodonts, phosphatic linguloid brachiopods, 

algal cysts, and small pieces of plant material were seen 

in some mudstone specimens. Radiolarians composed of 

cryptocrystalline silica are largely resorbed, and most 

sponge spicules are obscure and appear to have been partially or marginally resorbed. Some beds, however, contain well-preserved radiolarians and sponge spicules that 

show little or no compaction and whose peripheries are 

sharply defined; an example is surface sample 116 (silicified siltstone) collected near drill hole NO 12 (fig. 2). The 

sample yielded radiolarians of Late Devonian age according to D.L. Jones of the U.S. Geological Survey (oral 

commun., 1978). Another sample collected from a chert 

bed near sample locality 122 (fig. 2) also yielded 

radiolarians of Late Devonian age according to Brian 

Holdsworth of the University of Keele, England (oral 

commun., 1978). 

Conodonts on bedding plane surfaces collected 

from a 1-m-thick mudstone unit in the Gibellini facies 

at locality 122 by Poole and Green (August 27, 1981) were 

identified by C.A. Sandberg of the U.S. Geological 

Survey (written commun., 1981). They include the following taxa (and number of specimens): Palmatolepis rugosa 

rugosa (8), P. perlobata postera (9), Bispathodus stabi/is 

(2), ·Meh/ina strigosa (1), Branmehla inornata (2), 

Ozarkodina sp. {1), Neoprioniodus sp. (2), and 

Hindeode//a sp. (2). The conodonts have a CAl of 1, indicating a maximum roc.k temperature of less than 60°C 

(Epstein and others, 1977). The association of P. rugosa 

Metalliferous Oil Shale, Nevada 101 
rugosa and P. perlobata postera occurs worldwide only 

in the Late Devonian (late Famennian) Upper 

Polygnathus styriacus and Lower Bispathodus costatus 

Zones. Sandberg's knowledge of the many occurrences 

of these subspecies in the Western United States suggested 

to him that this collection more likely represents the Upper P. styriacus Zone. 

The marine algal cyst Tasmanites sinuosus Winslow 

was identified by R. M. Kosanke of the U.S. Geological 

Survey (written commun., 1981) in a mudstone bed at 

sample locality 122 on ridge 7129 (figs. 2, 3). In addition 

to Tasmanites in the mudstone sample, Kosanke recognized Hystrichosphaeridium cf. H. ohioensis Winslow 

and Punctatisporites sp. The middle Paleozoic range of 

Tasmanites sinuosus is compatible with the conodont age 

determination. 

Conodont, foraminiferal, and radiolarian faunas 

from the Bisoni facies of the Webb Formation indicate 

an Early Mississippian (Kinderhookian to Osagean) age 

(Hose and others, 1979). Conodonts in thin (<1 m) 

limestone beds in the Bisoni facies near the southern 

margin of the map area shown in figure 2A have a CAl 

of 1, which is similar to that of the structurally overlying 

Gibellini facies. 

Fragments of detrital vascular plant stems are 

common in the underlying autochthonous Mississippian Antelope Range Formation. A boulder of fossil 

wood was collected from the conglomeratic part of the 

upper Antelope Range Formation just beyond the 

southeast corner of the map area shown in figure 2. The 

wood was identified by R.A. Scott of the U.S. Geological Survey (written commun., 1979) as Callixylon of 

Late Devonian to Early Mississippian age. The fossilwood clast suggests that large trees were growing on or 

along the eastern margin of the Antler orogenic highland. 

A western source for the silicified-wood boulder is 

indicated by associated gravels composed of quartzite 

and chert similar to lower and middle Paleozoic units 

exposed in the Antler orogen. The wood may have been 

silicified prior to its eastward transport by streams, 

whereby the boulder was incorporated in Upper Mississippian molasse deposits (Poole, 1981; Poole and Claypool, 

1984, fig. 6B and C). 

Distribution and Concentration of Metals 

Metal concentrations and mineral assemblages in 

samples from drill hole NQ-12 on ridge 7129 suggest that 

oxidized rock extends to a depth of approximately 49 m, 

whereas the oil content suggests that the base of the 

oxidized zone may be at a depth of approximately 40 m 

(tables 1, 2). Hence, an intermediate zone of approximately 9 m in drill hole NG-12 represents a transition 

from intense oxidation above 40 m to fresh rock below 

49 m. 

102 Contributions to Mineral Resources Research, 1984 

In order to accurately determine the lower contact 

of the metal-rich Devonian strata with underlying 

Mississippian siltstone and mudstone on ridge 7129, cuttings were obtained from the deepest penetrations in drill 

holes NG-35 and NG-51 (fig. 2; table 3). Concentrations 

of vanadium, zinc, selenium, molybdenum, and silver are 

reported for these samples in table 3. In contrast with 

the values for the vanadiferous Devonian strata in drill 

holes NG-35 and NG-51 shown in table 3, the underlying Mississippian siltstones contained less than 150 ppm 

vanadium, 100 ppm zinc, 20 ppm selenium, and 20 ppm 

molybdenum. 

Table 3. Concentration of vanadium, zinc, selenium, 

molybdenum, and silver in composite samples from drill holes 

on the Gibellini (NG) and Bisoni-McKay (N-BM) properties, 

Eureka and Nye Counties, Nev. 

[Values in parts per million. Original measurements were in feet; 1 

ft=0.3048 m] 

Depth v Zn Se Mo Ag (ft) 

Drill hole NG-35: 

150-155 1,180 1,100 20 80 2.5 

Drill hole NG-51: 

140-145 3,140 500 120 20 3.0 

155-160 1,290 1,800 30 40 3.0 

Drill hole N-BM-1: 

150-155 780 300 440 <10 8.0 

185-190 500 300 <10 <10 3.5 

. 240-245 450 200 <10 <10 3.5 

Drill hole N-BM-2: 

90-95 '3, 640 800 <10 80 4.5 

150-155 3,140 600 60 20 4.0 

190-195 3,360 2,700 60 150 3.5 

225-230 2,460 1,300 <10 140 3.0 

275-279.5 2,860 2,200 70 110 4.0 

Drill hole N-BM-3: 

95-100 620 200 (10 10 5.0 

125-130 560 200 30 <10 6.0 

155-160 500 300 <10 10 6.5 

240-245 1,120 400 <10 100 2.0 

The vanadium in fresh (unoxidized) black rocks of 

the Gibellini facies on ridge 7129 occurs chiefly in organic 

material, with concentrations ranging from I ,500 to 

3,700 ppm. Vanadium in oxidized and leached rocks 

ranges from 850 to 9,600 ppm. Average concentration 

of vanadium for thirteen 1.5-m samples in the upper oxidized part (depth 3-23m) of drill hole NG-28 (not shown 

in table 1) is 1,840 ppm. Average concentration of 

vanadium in eighteen 1.5-m samples in the upper oxidized 

part of drill hole NG-12 (not shown in table 1) is 

2,600 ppm. Because we cannot establish the initial 
distribution of vanadium in the oxidized strata, we cannot quantitatively assess the influence of oxidation and 

(or) ground-water leaching on the present distribution of 

vanadium. However, the highest concentrations of 

vanadium are in samples that are oxidized, and one 6-mthick zone (NG-12, depth 24-30 m) averages 7,140 ppm, 

·.which is more than twice the amount found in any sample of unoxidized strata. 

Zinc in the Gibellini facies occurs in sphalerite, and 

selenium is chiefly in organic matter; molybdenum is 

mostly in organic material, but molybdenite has been 

identified. Concentrations of zinc in unoxidized rocks 

range from 4,000 to 18,800 ppm, whereas in oxidized rock 

they range from 30 to 100 ppm, showing strong depletion due to weathering. Concentrations of selenium in 

unleached rocks range from 20 to 85 ppm, but in oxidized and leached rocks they range from 190 to 455 ppm, 

indicating enrichment upon weathering. Concentrations 

of molybdenum in unoxidized rocks range from 60 to 

700 ppm, whereas in oxidized and leached rocks they 

range from 20 to 60 ppm, showing strong depletion by 

weathering. Unoxidized surface sample 122 contains the 

following, in ppm: V 3,140, Zn 2,570, Ni 315, Se 68, and 

Mo 252. The concentration of iron is usually greater in 

the unoxidized rocks, but oxidative leaching removes 

much of it from iron sulfides. Nickel concentrations in 

unoxidized rocks range as high as 1,330 ppm, but they 

are reduced to less than 100 ppm in the oxidized zone. 

As much as 105 ppm of arsenic is present in fresh rocks, 

but the arsenic concentration seems to be somewhat lower 

in the oxidized zone. No marked concentrations of Rb, 

Sr or Zr were encountered. Uranium is moderately enriched in these sapropelic sediments, with concentrations 

from 20 to 60 ppm. Silver contents are in the range of 

2-8 ppm, and seem to be highest in the oxidized rock 

(table 3). 

Sparse, microcrystalline red metahewettite in 

microfractures in subsurface samples and in nodules of 

fresh surface samples of the unoxidized zone has been 

identified by X-ray diffraction methods. Electron 

microprobe analysis for all elements heavier than oxygen 

revealed that calcium and vanadium were the only major constituents in the metahewettite. 

Microscopic, electron microprobe, and X-ray diffraction studies of unoxidized drill cuttings and of the 

unoxidized samples obtained from excavations show that 

vanadium minerals are sparse. Most of the vanadium in 

unoxidized mudstone occurs in the organic matter, according to electron microprobe studies using both 

wavelength- and energy-dispersive spectrometers 

simultaneously. The VK{j peak was used for both 

qualitative and quantitative electron microprobe analyses 

because of the serious interference of the TiK and VKa 

peaks. 

Concentration of vanadium in organic matter is in 

the range of about 0.3-3.0 weight percent. The exact 

amount cannot be determined by electron microprobe 

analysis because even under lowest intensity beam conditions, the kerogen volatilizes immediately. Vanadium 

and sulfur are the chief components (heavier than oxygen) 

in the kerogen. Thus, virtually all the vanadium in the 

unoxidized and unleached mudstone is dispersed in 

kerogen at low concentrations. 

Zinc occurs in widely dispersed grains of sphalerite 

that are generally less than about 40 J.Lm in diameter (fig. 

3E, F, and H). Microprobe analysis shows that all 

sphalerite grains contain less than 0.1 weight percent of 

manganese and iron, but that cadmium is present locally. Cadmium constitutes as much as 6.0 weight percent 

of some grains, but generally constitutes 0.2 weight percent or less of the sphalerite. Pyrite and marcasite are 

also present as small, widely dispersed grains in the unoxidized mudstone; these grains are distributed completely 

independently of the sphalerite grains. Although 

volumetrically insignificant, the pyrite and marcasite are 

conspicuous microscopically owing to their high reflectance in comparison with other minerals in the rock. 

Concentrations of selenium in kerogen are probably 

greater than 0.04 weight percent, but accurate determinations cannot be made because of decomposition of kerogen in the microprobe beam. The selenium dispersed in 

kerogen is concentrated during oxidation in the weathered 

zone where concentrations range from 190 to 455 ppm 

(table 1). Minor amounts of selenium and arsenic were 

detected in iron sulfides concentrated from sample 122. 

Molybdenum occurs in widely dispersed grains of 

molybdenite less than 5 J.Lm in diameter. It appears that 

most of the molybdenite grains were plucked from the 

polished sections during preparation. Molybdenum· probably constitutes more than 0.01 weight percent of the 

kerogen, but like selenium, accurate determination cannot be made owing to decomposition of kerogen in the 

microprobe beam. 

Unofficial reports indicate that the Gibellini facies 

at the Bisoni-McKay property may contain several million 

tons of rock with an estimated grade of as much as 0.44 

percent V20s- Concentrations of vanadium, zinc, 

selenium, molybdenum, and silver in composite samples 

from three drill holes on the Bisoni-McKay property are 

listed in table 3. 

CONCLUSIONS 

Siliceous marine sediments that contain abundant 

sapropel dominate the Gibellini facies of the Woodruff 

Formation. These silica-rich sediments of the Gibellini 

facies originally consisted of siliceous silts and muds, 

slimes, and oozes whose high organic content is derived 

mainly from microscopic phytoplanktonic organisms. In 

many beds, outlines of radiolarian tests and sponge 

spic.ules tend to be blurred and obscure, indicating some 

Metalliferous Oil Shale, Nevada 103 
marginal resorption and some masking by kerogen or 

bitumen. During diagenesis, silica probably dissolved, 

reached saturation, and reprecipitated; at the same time, 

authigenic sulfide minerals of iron, zinc, molybdenum, 

and other metals formed. Pyrite and marcasite occur both 

in the matrix and in opaline-like silica veinlets, which indicates that their principal time of formation was during 

diagenesis. 

Within the siliceous sediments, abundant amorphous, flaky and stringy organic matter is present, and 

probably includes remains of bacteria, phytoplankton 

(algae), zooplankton, and a minor amount of higher 

plants. The organic matter (kerogen and bitumen) includes, and is the source of, additional hydrocarbons produced by heating. The early-phase hydrocarbons (liquid 

oil and solid bitumen) present in the rock probably 

formed without any major thermal degradation of the 

kerogen. These hydrocarbons were probably released during diagenesis from complex nonhydrocarbon molecular 

structures originating from living organisms. 

These kerogen-rich sediments were deposited in an 

area receiving organic debris from large volumes of nearsurface marine waters that contained abundant phytoplankton. Fine-grained mineral particles of terrigenous 

origin were sparse, or were chemically degraded during 

diagenesis, and organic sediments were not effectively 

diluted by hemipelagic sediments. Abundant kerogen of 

marine origin indicates anoxic bottom conditions. 

The Gibellini facies of the Woodruff Formation appears to have been deposited in a continental-rise setting 

on the east side of a Devonian inner-arc basin (Poole and 

others, 1977) that periodically was anoxic, which allowed 

deposition and preservation of organic-rich sapropelic 

sediments. Anoxic bottom water (that is, minimal mixing with near-surface water) probably accounted for the 

preservation of the organic matter. Water depth was well 

below wave base-influence, and may have exceeded 

100m. Although most of the organic matter is kerogen 

of marine origin, traces of vegetal macerals derived from 

land plants have been recognized. 

The occurrence of opaline-silica veinlets, kerogen 

that yields significant amounts of syncrude oil (as much 

as 50 L/t), no significant thermal alteration of the 

kerogen (Poole and Claypool, 1984, p. 209, table 1), low 

vitrinite reflectance values (<0.5 percent), conodont color alteration index .of 1, and high ratio of hydrocarbon:organic carbon, collectively demonstrate that the 

Upper Devonian Gibellini facies of the Woodruff 

104 Contributions to Mineral Resources Research, 1984 

u U.S. GOVERNMENT PRINTING OFFICE: 1987-773-047/46,057 REGION NO.8 

Formation in the southern Fish Creek Range has not been 

subjected to thermal regimes above 60 °C since 

deposition. 

REFERENCES CITED 

Bohlke, J.K., Radtke, A.S., Heropoulos, Chris, and Lamothe, 

P.J., 1981, Spectroscopic geochemical study of vanadiferous 

marine sediments of the Gibellini claims, southern Fish 

Creek Range, Eureka County, Nevada: U.S. Geological 

Survey Open-File Report 81-32, 37 p. 

Desborough, G.A., Poole, F.G., Hose, R.K., and Radtke, A.S., 

1979, Metals in Devonian kerogenous marine strata at 

Gibellini and Bisoni properties in southern Fish Creek 

Range, Eureka County, Nevada: U.S. Geological Survey 

Open-File Report 79-530, 31 p. 

Epstein, A.G., Epstein, J.B., and Harris, L.D., 1977, Conodont 

color alteration-an index to organic metamorphism: U.S. 

Geological Survey Professional Paper 995, 27 p. 

Espitalie, J., Laporte, J.L., Madec, M., Marquis, F., Leplat, P., 

Paulet, J., and Boutefeu, A., 1977, Methode rapide de 

caracterisation des roches meres, de leur potential petrolier 

et de leur degre d'evolution: Revue de l'Institut Francais 

du Petrole, v. 32, no. ·1, ·p. 23-42. 

Hose, R.K., 1983, Geologic map of the Cockalorum Wash 

quadrangle, Eureka and Nye Counties, Nevada: U.S. 

Geological Survey Miscellaneous Investigations Series Map 

1-1410, scale 1:31,680. 

Hose, R.K., Wrucke, C.T., and Armstrong, A.K., 1979, Mixed 

Devonian 'and Mississippian conodont and foraminiferal 

faunas and their bearing on the Roberts Mountains thrust, 

Nevada: Geological Society of America Abstracts with Programs, v. 11, no. 7, p. 446. 

Poole, F.G., 1981, Molasse deposits of the Antler foreland basin 

in Nevada and Utah: Geological Society of America 

Abstracts with Programs, v. 13, no. 7, p. 530-531. 

Poole, F.G., and Claypool, G.B., 1984, Petroleum source-rock 

potential and crude oil correlation in the Great Basin, in 

Woodward, J., Meissner, F.F., and Clayton, J.L., eds., 

Hydrocarbon source rocks of the Greater Rocky Mountain 

Region: Rocky Mountain Association of Geologists, 

p. 179-230. 

Poole,F.G., Sandberg, C.A., and Boucot, A.J., 1977, Silurian 

and Devonian paleogeography of the Western United States, 

in Stewart, J.H., Stevens, C.H., and Fritsche, A.B., eds., 

Paleozoic paleogeography of the western United States: 

Society of Economic Paleontologists and Mineralogists, 

Pacific Sec., Pacific Coast Paleogeography Symposium 1, 

p. 39-65. 

Poole, F.G., Sandberg, C.A., and Green, G.N., 1983, 

Allochthonous Devonian eugeosynclinal rocks in southern 

Fish Creek Range of central Nevada: Geological Society 

of America Abstracts with Programs, v. 15, no. 5, p. 304.