Application of Structural Geology to 

Mineral and Energy Resources of the 

Central and Western United States 

U.S. GEOLOGICAL SURVEY BULLETIN 2012 

. . 

Application of Structural Geology to 

Mineral and Energy Resources of the 

Central and Western United States 

Edited by CHARLES H. THORMAN 

Papers presented at a workshop sponsored by the U.S. Geological Survey, 

October 9-10, 1990, Colorado School of Mines, Golden, Colorado 

U.S. GEOLOGICAL SURVEY BULLETIN 2012 
U.S. DEPARTMENT OF THE INTERIOR 

MANUEL LUJAN, JR., Secretary 

U.S. GEOLOGICAL SURVEY 

Dallas L. Peck, Director 

Any use of trade, product, or firm names in this publication is for descriptive purposes only and 

does not imply endorsement by the U.S. Government 

UNITED STATES GOVERNMENT PRINTING OFFICE: 1992 

For sale by 

Book and Open-File Report Sales 

U.S. Geological Survey 

Federal Center, Box 25286 

Denver, CO 80225 

Library of Congress Cataloging-in-Publication Data 

Application of structural geology to mineral and energy resources of the central and 

western United States : papers presented at a workshop sponsored by the U.S. 

Geological Survey, October 9-1 0, 1990, Colorado School of Mines, Golden, 

Colorado I edited by Charles H. Thorman 

p. em.- (U.S. Geological Survey bulletin; 2012) 

Includes bibliographical references. 

Supt. of Docs. no.: I 19.3:2012 

1. Geology, Structural-Southwestern States-Congresses. 2. Geology, 

Structural-West (U.S.)-Congresses. 3. Ore deposits-Southwestern 

States-Congresses. 4. Ore deposits-West (U.S.)-Congresses. I. Thorman, 

Charles H. II. Geological Survey (U.S.) Ill. Series. 

QE75.B9 no. 2012 

[QE627.5.S85l 

557.3 s-dc20 92-1 07 47 

[553'.09781 CIP 
CONTENTS 

[Letters designate chapters] 

(A) Microstructures in the Cambrian Bonneterre Formation, Lamotte Sandstone, and 

basal clastic rocks of southeast Missouri and northeast Arkansas-Implications of 

regional sulfide occurrence in stylolites and extensional veinlets for ore genesis, by 

S.F. Diehl, M.a .. Goldhaber, C.D. Taylor, H.S. Swolfs, and C.A. Gent. 

(B) Tectonic controls on deposition and source-rock properties of the Woodford Shale, 

Anadarko Basin, Oklahoma-Loading, subsidence, and forebulge development, by 

Timothy C. Hester, James W. Schmoker, and Howard L. Sahl. 

(C) Sequential Laramide deformation of the Rocky Mountain foreland of southwestern 

Montana, Wyoming, and north-central Colorado, by William J. Perry, Jr., Douglas 

J. Nichols, Thaddeus S. Dyman, and Christopher J. Haley. 

(D) Influence of crustal structure on the course of the Arkansas River, south-central 

Colorado, by Ken Watson, Daniel H. Knepper, Jr., and Michael W. Webring. 

(E) Implications of gravity and seismic reflection data for Laramide mountain ranges 

and basins of Wyoming and southern Montana, by Stephen L. Robbins and John A. 

Grow. 

(F) Pre-late Eocene structures and their control of gold ore at the Drum Mine, westcentral Utah, by C.J. Nutt and C.H. Thorman. 

(G) High-precision 40 Art39 Ar geochronology, volcanic stratigraphy, and mineral deposits of Keg Mountain, west-central Utah, by Michael A. Shubat and Lawrence W. 

Snee. 

(H) Structural setting of the Chief Mining District, eastern Chief Range, Lincoln 

County, Nevada, by Peter D. Rowley, Lawrence W. Snee, Harald H. Mehnert, R. 

Ernest Anderson, Gary J. Axen, Kelly J. Burke, F. William Simonds, Ralph R. 

Shroba, and Stephen D. Olmore. 

(I) Extensional geometry in the northern Grant Range, east-central Nevada-Implications for oil deposits in Railroad Valley, by Karen Lund and L. Sue Beard. 

(J) Late Paleozoic structure of the southern part of the Uinta Basin, Utah, from seismic 

reflection data, by Christopher J. Potter, Rex Tang, and Timothy J. Hainsworth. 

(K) Audio-magneto-telluric investigation at Turkey Creek caldera, Chiricahua Mountains, southeastern Arizona, by Robert M. Senterfit and Douglas P. Klein. 

(L) Ash-flow calderas as structural controls of ore deposits-Recent work and future 

problems, by Peter W. Lipman. 

Volume Contents Ill 

PREFACE 

This volume includes some of the papers presented at a workshop sponsored by the U.S. Geological Survey and held October 9-10, 1990, at the Green Center, Colorado School of Mines, Golden, Colorado. The workshop presented current research 

by U.S. Geological Survey scientists who are using structural geology in their investigations into the mineral and energy 

resources of the United States. The figure below shows the areal distribution of the 12 chapters included herein; the papers 

are lettered on the figure according to their position in the text of this report. 

One of the most important aspects of this workshop was a demonstration of the integrating of many disciplines in unraveling the structural evolution of an area, regardless of scale. The application of structural geology by itself provides only a 

limited amount of information and leaves many aspects of the evolution of an area unresolved. 

Integration of geophysical techniques (gravity, seismic, and audio-magneto-telluric) with detailed and regional geology in 

sedimentary and volcanic terranes establishes the basis for new approaches to exploration for hydrocarbons in Arizona, Colorado, Utah, Wyoming, and Montana (chapters D, E, J, K). Field studies at detailed to regional scales (chapters A, C, F, G, 

H, I, L) present data that call for moderate to major revision of thinking in areas often thought to be well understood. These 

studies emphasize the continuing need to remap so-called well-understood areas, as well as the need to map those areas for 

which only limited data are available. Detailed petrologic studies in Missouri and Arizona (chapters A, L) of both carbonate 

and volcanic rocks cause us to reconsider our concepts of the origin of well-established rocks units and thus their implications 

regarding the structural evolution of those areas. Depositional patterns can be key indicators of the tectonic control of paleotopographic highs, a major factor in basin evolution, as suggested for the Anadarko Basin (chapter B). The application of 

isotopic dating in structurally complex areas has greatly increased our knowledge of the relative timing and duration of events 

that influenced the migration of mineralizing and hydrocarbon fluids and gases (chapters G, H). With a better grasp of the 

time involved, we can better correlate the various deformational and migrational events in Utah and eastern Nevada. 

I 

i 

! 

\,_ 

Preface V 

Chapter A 

Microstructures in the Cambrian Bonneterre 

Formation, Lamotte Sandstone, and 

Basal Clastic Rocks of Southeast Missouri and 

Northeast Arkansas-Implications of Regional 

Sulfide Occurrence in Stylolites and 

Extensional Vein lets for Ore Genesis 

By S.F. DIEHL, M.B. GOLDHABER, 

C.D. TAYLOR, H.S. SWOLFS, and C.A. GENT 

U.S. GEOLOGICAL SURVEY BULLETIN 2012 

APPLICATION OF STRUCTURAL GEOLOGY TO MINERAL AND ENERGY 

RESOURCES OF THE CENTRAL AND WESTERN UNITED STATES 
CONTENTS 

Abstract 

Introduction 

Methods 

Al 

Al 

A2 

Structural setting of drill holes and description of samples A2 

Reelfoot Rift A2 

Wilson drill hole A2 

Garrigan drill hole A3 

Carbonate platform A 7 

GP-1 0 drill hole A 7 

STH-2 drill hole A8 

317-A drill hole A8 

Ozark Uplift A8 

16-15 drill hole A8 

B-4 and BCC-4 drill holes A9 

Discussion AlO 

Microstructures and mineralization AlO 

Mechanisms for fluid migration All 

Regional fracture patterns Al2 

Summary Al2 

References cited Al2 

FIGURES 

1. Map showing location of drill holes, structural regimes, and ore districts A2 

2. Lithologic logs and stratigraphic correlation of Wilson and Garrigan drill 

holes A3 

3. Diagram comparing sulfur isotope values for samples from Reelfoot Rift and 

samples from various ore districts A6 

4, 5. Photographs showing: 

4. Carbonate-filled veinlets crosscut by mineral front A 7 

5. Deformed stylolite A 7 

6. Diagram showing fracture (veinlet) attitudes in Garrigan drill hole A8 

7. Photograph showing mineral succession in vertical extensional fractures A8 

8, 9. Graphs showing: 

8. Total sulfur contents in mineral fronts versus barren rock AS 

9. Enrichment of lead and cobalt in mineral fronts A9 

10-13. Photographs showing: 

10. Brecciated marcasite in stylolite A9 

11. Layering of mineralization in deformed stylolite A9 

12. Curved cleavage planes in deformed galena along stylolite AlO 

13. Partially dissolved quartz grains and unaffected pyrite along 

stylolite AlO 

14. Map showing flow paths of hot brines due to topographic relief All 

15. Chart showing age dates of authigenic minerals in relation to timing of 

Mississippi Valley-type deposition All 

TABLES 

1. Descriptions of some thin sections of micro fractures and stylolites in core samples 

from the Reelfoot Rift, carbonate platform, and Ozark Uplift structural 

regimes A4 

2. Compilation of sulfur isotope data from the Garrigan and Wilson drill holes A6 

All Contents 
Microstructures in the Cambrian Bonneterre Formation, 

Lamotte Sandstone, and Basal Clastic Rocks of Southeast 

Missouri and Northeast Arkansas-Implications of 

Regional Sulfide Occurrence in Stylolites and 

Extensional Veinl~ts for Ore Genesis 

By S.F. Diehl 1, M.B. Goldhaber2, C. D. Taylor2, H.S. Swolfs 1, and C.A. Gent2 

Abstract 

Petrologic and structural observations of stylolites and 

veinlets in the Cambrian Bonneterre Formation, Lamotte Sandstone, and basal siliciclastic rocks in the Reelfoot Rift indicate 

that episodes of tectonic activity and fluid-flow events related 

to ore deposits in southeast Missouri may be genetically 

related. From the Reelfoot Rift to the Ozark Uplift, these microstructures, including stylolites parallel with stratification and 

small-scale extensional veinlets normal to both bedding and 

stylolites, contain authigenic phases of ore affinities. Sulfidemineral deposition in both horizontal stylolites and vertical 

microfractures is evidence for mineralizing fluid migration 

along microstructures and indicates that the Lamotte Sandstone was not solely a homogeneous conduit as envisioned in 

many models of metals transport. 

Detailed petrologic studies of drill-hole samples from the 

Reelfoot Rift, from the carbonate platform partially buried 

beneath the Mississippi Embayment northwest of the rift, and 

from the Ozark Uplift show that stylolite development began 

early in the diagenetic history of the Cambrian carbonate and 

siliciclastic sequence and probably continued through Paleozoic time. Extensional veinlets sealed by incremental mineral 

growth are commonly associated with the stylolites. Petrographic examination indicates that the microstructures have 

several generations of epigenetic mineral fillings. Multiple 

deformational events are inferred to have induced the flow of 

mineralizing fluids in the microstructures. 

Manuscript approved for publication October 31, 1991. 

1 U.S. Geological Survey, Box 25046, MS 966, Denver, Colorado 

80225. 2 U.S. Geological Survey, Box 25046, MS 973, Denver, Colorado 

80225. 

INTRODUCTION 

The Reelfoot Rift (fig. 1) has only recently been recognized as a major crustal structure of the southern Midcontinent (Ervin and McGinnis, 1975; Kane and others, 1981; 

Hildenbrand, 1985). The existence of the Reelfoot Rift was 

first interpreted from potential-field data as part of studies of 

the New Madrid Seismic Zone (McKeown and Pakiser, 

1982). Recognition that the rift represented a very deep basin 

prompted the oil industry to conduct an exploration program 

that included seismic reflection surveys and drilling of deep 

holes in the upper Mississippi Embayment (Howe and 

Thompson, 1984 ). 

The Reelfoot Rift adjoins the southeast Missouri lead 

belts, which are centered around the Middle Proterozoic St. 

Francois Mountains (fig. 1 ). Ores in the lead belts are hosted 

in the Cambrian Bonneterre Formation and, to a lesser 

extent, the underlying Lamotte Sandstone. Together the lead 

belts, which consist of the Old Lead Belt and New Lead Belt 

or Viburnum Trend, represent the largest known concentration of lead in the crust of the Earth (Kisvarsanyi and others, 

1983) and contain significant amounts of zinc, copper, and 

cobalt. The dominance of lead over zinc and the abundance 

of copper and cobalt are atypical for Mississippi Valley-type 

(MVT) deposits worldwide (Sangster, 1983). Yet this suite 

of metals, particularly copper and cobalt, is commonly associated with rift/red-bed sequences (Mitchell and Garson, 

1981 ), and the source for these elements in southeast Missouri may possibly be in the adjacent red and gray feldspathic siliciclastic rocks such as those penetrated by drill 

holes in the Reelfoot Rift (Hayes, 1985; Kissin, 1988). 

A second major ore district, the Illinois-Kentucky fluorspar district, is also spatially, and perhaps genetically, 

Microstructures in Cambrian Rocks, Missouri and Arkansas A 1 
0 

I I I I r 1 1 1 1 

50 

50 

I 

I 

Cretaceous rocks 

__J__L_ Rift boundary 

• Drill hole 

100 MILES 

I 

where uncertain 

100 KILOMETERS 

e 1. Map showing location of drill holes, structural reg >, and ore districts. 

related to the Reelfoot Rift. Ore deposits in this district are 

structurally controlled and hosted by Mississippian-age sedim, Jtary rocks that overlie rocks in the rift to the northeast 

of the study area (fig. 1 ). The district is spa:· related to 

the Hicks Dome "cryptovolcano," which contams an underlying breccia enriched in Th, Nb, Y, Be, rare-earth elements, 

P, Ti, and F. Both alkalic igneous activity responsible for 

enrichment in this elemental suite and fluorine deposits commonly are rift related (Mitchell and Garson, 1981). 

Analyses of cuttings and core from two of the deepest 

drill holes penetrating into rift-fill sedimentary rocks of the 

upper Mississippi Embayment provide new information on 

processes of rock alteration and potentially on the movement 

of mineralizing fluids. Samples from drill holes in the subsurface carbonate platform northwest of and contiguous with 

the Reelfoot Rift and samples from drill holes within the lead 

belt in the Ozark Uplift were analyzed petrographically and 

geochemically to identify possible flow paths between the 

rift and the southeast Missouri ore districts. The analyses 

provide a regional perspective of changes or similarities in 

the petrologic characteristics of rocks from a basinal-fill 

environment, onto a carbonate platform, and then to depositional pinchouts against structurally high Precambrian crystalline rocks. 

Acknowledgments.-We wish to thank the Missouri 

Division of Geology and Land Survey, the Arkansas Geological Commission, and Mag mont Mines for providing core 

samples. Allison Palmer (Geological Society of America) 

and Michael Taylor (U.S. Geological Survey) identified 

fossils and provided biostratigraphic correlations. Critical 

reviews by George Breit, Paula Hansley, and Frank 

McKeown (U.S. Geological Survey) were very helpful. Sulfur isotope analyses were done in the laboratory ofR.O. Rye 

(U.S. Geological Survey). 

METHODS 

Drill cuttings, core, and junk basket samples used for this 

study were selected from the Bonneterre Formation, 

Lamotte Sandstone, and older stratigraphic units in the Reelfoot Rift. 

Standard-size thin sections were prepared from the sample suite and stained with potassium ferricyanide, Alizarin 

Red-S, and sodium cobaltinitrate to identify ferroan carbonate, nonferroan carbonate, and potassium feldspar, respectively. Authigenic minerals, structural fabric, and the 

paragenetic sequence of minerals were determined by using 

a petrographic microscope and a Cambridge 250 Mark 2' 

scanning electron microscope (SEM) equipped with an 

energy-dispt ' ive X-ray spectrometer. 

Samples trom the Dow Chemical No. 1 Wilson and No. 

1 Garrigan drill holes (fig. 2) were analyzed to determine 

their elemental and isotopic composition. Elemental analyses were made by using inductively coupled plasma spectroscopy and a LECO sulfur analyzer; sulfur isotope 

analyses were made by using gas-source mass spectrometry. 

Results are reported relative to Cafion Diablo troilite (CDT). 

STRUCTURAL SETTING OF DRILL HOLES 

AND DESCRIPTION OF SAMPLES 

Reelfoot Rift 

The Wilson drill hole is off the shoulder of the Reelfoot 

Rift, where modern seismicity is relatively low. In contrast, 

the Garrigan drill hole is in the central part of the rift in the 

seismically active Blytheville Arch. Samples from both drill 

holes record episodic extensional deformation which repeatedly reopened fractures that were then sealed by epigenetic 

mineral cements (Diehl and McKeown, 1989; Diehl and 

Goldhaber, 1990). 

Wilson Drill Hole 

The Wilson drill hole (total depth 14,613 ft, 4,454 m) is 

the deepest drill hole in the Reelfoot Rift that penetrates the 

Precambrian basement. The basal clastic rocks in this hole 

are divided into a basal red feldspathic arenite overlain by a 

gray feldspathic to quartz arenite. From limited 

A2 Application of Structural Geology to Mineral and Energy Resources 
Depth below 

sea level 

(meters! (feet! 

0 0 

Core 

intervals 

Dow Chemical 

No. 1 Garrigan 

Dow Chemical 

No.1 Wilson 

1,000 

5,000 

--~~~--Pa~le-o-zo_i_c-ro-ck-s-----------TOP 

Knox or Arbuckle Group equivalents 

(Lower Ordovician and Upper 

Cambrian) 

EXPLANATION 

Sandstone and shale 

__:_::-:-;-:-....:...:...:_ / (?) Mineral -:-:-; ... ::·. Elvins Group (Upper Cambrian) 

fronts ···-... Davis Formation equivalent 

------ ~.:..:.:..._~ (based on biostratigraphy, 

Veins of sets one __.:_:: :.·: ···. Palmer, 1989) 

2,000 

Shale 

limestone 

and two ···-··· '- ,___:_:...:..:..:__~=--=...!_ ________ _ 

--------=----= " /Bonneterre Formation 

Dolomite 

10,000 Veins of set two~=~~ ", / equivalent (Upper 

==- === ~~~ ~C-'--am_:_:_:::_br-'-"ia_:_n ____________ _ 

3,000 Arkose 

Crystalline rocks ~l. .·Gray arnotte : .. : · .. : ... arkose Basal clastic rocks 

{?) • ... ••.·•· Red (Upper and Middle(?) 

4,000 

. :?:~~;~/; ~~:::lline~o:~~i•nJ -15,000 (Precambrian) 

5,000 

Figure 2. Lithologic logs and stratigraphic correlation of the Dow Chemical No. 1 Wilson and No. 1 Garrigan drill 

holes in the Reelfoot Rift. Location of drill holes shown in figure 1. Modified from McKeown and others (1990) and 

Collins and others (1992). 

biostratigraphic data, A.R. Palmer (written commun., 1989) 

interpreted the Bonneterre-Davis Formations equivalent 

contact to be between 10,800 and 10,920 ft (3,292-3,328 m) 

(Collins and others, 1992). We believe that the underlying 

gray feldspathic to quartz arenite may be the equivalent of 

the Lamotte Sandstone (fig. 2). 

In these arenites, siderite and trace amounts of anhydrite, 

barite, pyrite, and galena are intergranular cements; most of 

these authigenic minerals are also vein fillings. From 12,600 

to 13,900 ft (3,841-4,237 m), an overlapping sequence 

within veins of early fluid inclusion trails is followed by 

quartz-sealed veins, then a succession of siderite, dolomite, 

and ferroan dolomite (table 1). Most of the mineral fillings 

in the fractures are well laminated and have distinct boundaries between carbonate phases. 

Stylolites in this part of the section are parallel with bedding, nonsutured, and wavy and show slickensided clay and 

offset fractures that are normal to the plane of the stylolite. 

The stylolites host trace amounts of barite, pyrite, and euhedral epigenetic monazite. Authigenic monazite is also 

present at Hicks Dome to the northeast in rocks of the fluorspar mineral district of southern Illinois and northwestern 

Kentucky (fig. 1), but a genetic connection has not been 

established. 

Sulfur isotope data on sulfide minerals (dominantly 

FeS2) from the gray arenite are listed in table 2 and plotted 

as a histogram in figure 3. The data in part overlap with values for main-stage ore minerals in the Viburnum Trend of 

the southeast Missouri district. 

Several inferences may be made based on the crystal habits of the epigenetic minerals. The drusy habit of quartz and 

siderite in the veinlets suggests precipitation in a fluid-filled 

fracture (Beach, 1977), and the euhedral monazite suggests 

hydrothermal fluid flow along the stylolites. 

Garrigan Drill Hole 

The Garrigan drill hole (total depth 12,038 ft, 3,669 m) 

bottoms in a quartz and feldspathic arenite that may be 

equivalent to the Lamotte Sandstone. It should be noted, 

however, that the entire siliciclastic section above the basement rock in the rift has been informally divided into the "St. 

Francois formation" and "Reelfoot arkose" by other 

researchers and is considered older than the Lamotte Sandstone (Weaverling, 1987; Houseknecht, 1989). From biostratigraphic data (Taylor and others, 1991; Collins and 

others, 1992) (fig. 2), we infer that an overlying sequence of 

interbedded black shale and siltstone may be equivalent to 

the Bonneterre Formation or, possibly, to the Elvins Group. 

This drill hole presents a problem of stratigraphic correlation 

because the rock types do not correspond to the stratigraphy 

in the Wilson drill hole or other adjacent drill holes within 

the rift (McKeown and others, 1990). 

A variety of authigenic phases are in horizontal stylolites, 

vertical carbonate-filled fractures, and gray epigenetic mineral fronts in the Garrigan drill hole (fig. 4). Stylolites occur 

in swarms, are parallel with bedding, and are commonly 

deformed (fig. 5). Early diagenetic framboidal pyrite, 

organic matter, clay minerals, and cerium phosphate minerals (monazite?) are concentrated along these stylolites. 

Microstructures in Cambrian Rocks, Missouri and Arkansas A3 
> ,Jlo. 

> ""C' 

"2.. 

g· Table 1. Descriptions of some thin sections of microfractures and stylolites in core samples from the Reelfoot Rift, carbonate platform, and Ozark Uplift 

s· structural regimes :::s [Drill-hole locations shown on figure 1. na indicates not applicable. Fractures are normal to bedding; stylolites are parallel with bedding. Minerals are listed in paragenetic sequence] Q 

~ -

2 Drill hole Number of Stylolite 

~ and sample Structural regime Mineral assemblage fracture amplitude Mineral assemblage c 

.. 

~ depth (ft) and rock unit Host rock in fracture generations or thickness in stylolite Remarks 

C) Wilson Reelfoot Rift; basal F eldspathic arenite Quartz, siderite, dolomite, 5-7? na na Incremental cracking and ~ 

S'" 12,633 clastic rocks (Lamotte ferroan dolomite; clay, resealing of fractures. 

~ 

= 

Sandstone?) hematite 

~ Wilson Reelfoot Rift; basal Quartzarenite Quartz, dolomite, chlorite, 3-4? 0.1-0.4-mm- Minor pyrite, barite, Tension gashes associated :r 

ttl 12,782 clastic rocks minor pyrite thick seam monazite, organic matter with stylolites are offset. 21. 

Dol 

:::s 

Q.. Wilson Reelfoot Rift; basal Feldspathic arenite Dolomite, illitic clay, 3 0.4-0.6-mm- Minor pyrite, monazite, Incremental cracking and rrl 

:::s 13,649 clastic rocks quartz, hematite thick seam organic matter resealing of fractures. ttl 

~ :IIQ Garrigan Reelfoot Rift; Siltstone, black Quartz, siderite, ferroan 4 0. 9-mm-thick Siderite, illitic clay, Stylolite is layered, ttl 

~ 8,000 Bonneterre(?) Formation shale dolomite, ferroan calcite seam thorite, sphalerite, pyrite deformed. ~ l"l 

ttl Garrigan Reelfoot Rift; Black shale, Quartz, chlorite, dolomite, 5 2.6-3.0-mm- Calciwn phosphate, illitic Incremental cracking and Cll 

11,420 Bonneterre(?) Formation siltstone ferroan dolomite, ferroan thick seam clay, minor pyrite resealing of fractures; thrust 

calcite zone. 

GP-10 Carbonate platform; Silty dolostone Marcasite, brecciated 3-4? 0.6-2.0-mm- Marcasite, brecciated Tension gashes associated 

1,817.2 Bonneterre Formation marcasite, pyrite, thick seam marcasite, pyrite, with stylolites; brecciation 

marcasite marcasite between periods of sulfide 

deposition. 

GP-10 Carbonate platform; Litharenite Dolomite 1 0.2-1.0-mm- Illitic clay, pyrite, minor Illitic clay has horizontal 

2,654.4 Lamotte Sandstone thick seam stannite slickensides showing thrust 

movement. 

GP-10 Carbonate platform; Silty limestone, na na 0.6-mm-thick Illitic clay, pyrite, None. 

2,671.1 Lamotte Sandstone litharenite seam monazite, barite, ironmanganese oxides, 

calciwn phosphate 

STH-2 Carbonate platform; Quartzarenite Dolomite, pyrite 2? na na Quartz grains have 

2,700 Lamotte Sandstone undulatory extinction. 
~ r:;· 

0 

~ c 

0. 

c 

;; 

Ill 

:r 

n 

Dl 

3 

~ ;· 

= 

~ ~ 

~ ;;;· 

~ 

5. 

Dl 

&. 

> 

~ 

= 

~ 

> Iori 

Table 1. Descriptions of some thin sections of microfractures and stylolites in core samples from the Reelfoot Rift, carbonate platform, and Ozark Uplift 

structural regimes-continued 

Drill hole Number of Stylolite 

and sample Structural regime Mineral assemblage fracture amplitude Mineral assemblage 

depth (ft) and rock unit Host rock in fracture generations or thickness in stylolite Remarks 

317-A Carbonate platform~ Quartzarenite na na 1.0-2.4-mm- Illitic clay, pyrite, Stylolites are layered, 

2,474 Lamotte Sandstone thick seam carbonate, organic matter deformed. 

B-4 Ozark Uplift~ Lamotte Quartzarenite na na 0.9-3.7-mm- Major pyrite Colloform texture of 

1,349 Sandstone thick seam sulfides in stylolites. 

BCC-4 Ozark Uplift~ Lamotte Quartzarenite na na 0.4-1.0-mm- Major pyrite Colloform texture of 

1,284 Sandstone thick seam sulfides in stylolites. 

BCC-4 Ozark Uplift; Lamotte Quartzarenite na na 2.4-4. 8-mm- Major pyrite Pyrite brecciated along 

1,482.5 Sandstone thick seam stylolite. 

BCC-4 Ozark Uplift; Lamotte Quartzarenite na na <0 .1-mm-thick Pyrite Poikilitic carbonate 

1,509.5 Sandstone seam; 0.2 mm enclosing detrital clay on 

amplitude surfaces of quartz grains 

indicates fluctuation in 

static water level. 

16-15 Ozark Uplift; Bonneterre Silty limestone na na 0.5-1.0 mm Pyrite, chalcopyrite, galena Galena deformed along 

917.9 Formation amplitude stylolites; clay has 

horizontal slickensides. 
Table 2. Compilation of sulfur isotope data from the 

Garrigan and Wilson drill holes 

[Drill-hole locations shown in figure l. o34S in per mil] 

Deeth {ft} o34s Descrietion 

Dow Chemical No. 1 Garrigan 

7,975 26.6 Pyrite in mineral front. 

7,977 25.9 Pyrite in laminated shale associated with 

mineral front. 

7,980.25 30.9 Pyrite in mineral front. 

7,986 12.4 Pyrite in black shale. 

7,992.5 23.8 Pyrite in mineral front. 

7,993.1 27.6 Pyrite in stylolite crosscut by carbonatefilled fracture. 

7,998 28.6 Shale with visible p}rite. 

10,210.6 18.4 Total sulfur from ground core. 

10,225.3 16.2 Total sulfur from ground core. 

11,405.8 0 Total sulfur from ground core. 

11,420 1.2 Total sulfur from ground core. 

11,420.8 6.7 Total sulfur from 8round core. 

Dow Chemical No. 1 Wilson 

12,300-12,600 27.6 Pyrite-rich gray arenite; bulk rock 

sample 

12,300-12,600 17.1 Heavy liquid concentrate containing 

p}rite from gray arenite; bulk rock 

sample 

Sphalerite is present both as a residual mineral in these 

deformed stylolites and as an epigenetic mineral associated 

with late-stage carbonate precipitation. Trace amounts of 

thorite are within the layers of clay minerals and organic 

matter in the deformed stylolites; however, it has not been 

determined if the thorite is residual or epigenetic. 

Carbonate-filled veins are essentially vertical, commonly 

doubly terminated, or bounded against horizontal stylolites 

or shaley interbeds, or dispersed into the mineral fronts (fig. 

4). The lateral extent of these veins is not known, but they 

extend the full 1 0-cm diameter of the core. Azimuths of the 

veins measured by Swolfs and others (1990) in a section of 

oriented core, from 7,973 to 8,002 ft (2,430-2,439 m) depth, 

form two sets (fig. 6). Set one is oriented N. 41° E., subparallel with the rift axis. Set two is oriented N. 35° W., subperpendicular with the rift axis. 

Crosscutting relationships indicate that veins of set two 

are older than those of set one. Set one veins, which are only 

in the siltstone beds, may have formed during a prolonged 

period of extension (Swolfs, 1991). Set two veins, present in 

the siltstone beds and underlying shale sequence, extend 

over a greater vertical interval (3,000 ft, 914 m) and may 

have formed during deformation of regional extent (Swolfs 

and others, 1990). 

From the edges of the set one veins toward their centers, 

the mineral coatings consist successively of drusy quartz, 

siderite, ferroan dolomite, then ferroan calcite (fig. 7). The 

ferroan dolomite and ferroan calcite are optically continuous, in contrast to laminated mineral crusts having distinct 

phase boundaries in fractures in the Wilson and other drill 

holes outside of the seismically active zone. The only sulfide 

mineral recognized in fractures of the Garrigan drill hole is 

trace amounts of cubic pyrite. 

Both set one and set two veins are crosscut by epigenetic 

mineral fronts (fig. 4) that represent reduced zones superimposed on more oxidized host rock as evidenced by elevated 

sulfur content; total sulfur ranges from less than 0.01 weight 

percent sulfur in "barren" host rock to as much as 1.05 

weight percent sulfur in the mineral fronts (figs. 4, 8). The 

pyrite in the mineral fronts is euhedral and coarse grained 

and is unlike early diagenetic framboidal pyrite. 

In addition, the isotopic systematics of epigenetic pyrite 

are unlike those of diagenetic pyrite (table 2, fig. 3). The epigenetic pyrite is isotopically heavy. Several of the values are 

even heavier than main-stage sulfide in the Viburnum Trend; 

however, at isotopic equilibrium, FeS2 is 2-5 percent 

heavier than ZnS/PbS forming from the same H2S source. In 

addition to the sulfur enrichment, metals are also enriched. A 

positive correlation exists between the epigenetic sulfide sulfur and the concentration of lead and cobalt in the mineral 

fronts (fig. 9). Furthermore, several elements of affinity to 

ore-forming processes have been identified in mineral fronts 

ci. 

U5z 

~<t: 

)~ TR::~~n ~ ~ ARKANSA~ I 

° CENTRAL 

MISSOURI 0 D 

GARRIGAN 

WILSON 

I I 

-20 -10 

SOUTHEAST n n MISSOURI 

p o 1ISJ m q 

0 10 20 

o34S, IN PER MILL 

EXPLANATION 

[] Galena 

D Sphalerite 

Pyrite 

Garrigan 

drill hole 

Wilson 

drill hole 

Figure 3. Sulfur isotope values for samples from the Dow 

Chemical No. 1 Garrigan and No. 1 Wilson drill holes in the 

Reelfoot Rift are heavy, similar to those for octahedral mainstage sulfide minerals in various ore districts. Location of drill 

holes shown in figure 1. 

A6 Application of Structural Geology to Mineral and Energy Resources 
Figure 4. Carbonate-filled veinlets are vertical, commonly 

double terminating or bounded against stylolites, shaly interbeds, or epigenetic mineral fronts. Dow Chemical No. 1 Garrigan drill hole, depth 7,998.6 ft (2,438.0 m) (location shown 

in figure 1 ). 

by using SEM-Au, Ag, Cr, Cu, Ni, Pb, Zn, and rare-earth 

elements. Chromium and nickel are associated with authigenic iron-rich clays, gold and silver with authigenic albite, 

and sphalerite and galena as mineral coatings on ferroan calcite. The enrichment of sulfur and metals in the mineral 

fronts suggests a relationship between mineral-front genesis 

and at least local migration of metalliferous fluids in the rift. 

Several generations of diagenetic cement fillings in the 

fractures predate the mineral fronts; however, the introduction of mineralizing solutions may have caused reopening of 

earlier quartz- and siderite-filled fractures because ferroan 

calcite is present in the mineral fronts and as the last diagenetic cement in the fractures (Diehl and Goldhaber, 1990; 

Diehl and others, 1990). The incremental growth of minerals 

in the fractures matches the succession of authigenic minerals that form pore-filling cements. Bulk-rock X-ray 

diffraction of a mineral front versus a barren zone shows that 

siderite and ferroan dolomite pore-filling cements are prevalent in the unaltered barren zones, whereas calcite is more 

abundant in the mineral fronts. 

Carbonate Platform 

Drill-holes GP-10, STH-2, and 317-A are in the 

regional carbonate platform that borders the Reelfoot Rift to 

the northwest (Palmer, 1989; Diehl and Goldhaber, 1990) 

(fig. 1). This platform is partially buried beneath Cretaceous 

sedimentary rocks that fill a southwest-plunging syncline 

known as the Mississippi Embayment (Stearns and Marcher, 

1962). Stylolites and associated fractures form lateral and 

vertical permeable channels in the deeply buried Cambrian 

Bonneterre Formation and underlying Lamotte Sandstone 

(table 1). 

GP-1 0 Drill Hole 

Drill-hole GP-10 penetrates through the Cretaceous 

overburden at the edge of the embayment and into Cambrian-age sedimentary rocks. Stylolites in the Bonneterre 

Formation in this drill hole are layered by successive events 

of mineral precipitation, with intervening periods of dissolution and tectonic movement (table 1). For example, at 

1,817.2 ft (554 m), the genetic sequence can be interpreted 

as follows: (1) early marcasite fills the sockets in a stylolite, 

(2) marcasite is brecciated by a tectonic event, (3) because of 

deformation the stylolite is opened to fluid flow, (4) pyrite 

precipitates from the fluid, and (5) marcasite is again deposited (fig. 10). 

In the underlying Lamotte Sandstone, copper and tin sulfide minerals, probably stannite, precipitated along the 

Figure 5. A reversal of stress orientation has caused the normally ductile stylolite-filling material to break in a brittle manner. This 

deformation, which may be due to overpressure conditions or a paleoearthquake, opened permeable channels for mineralizing 

solutions. Dow Chemical No. 1 Garrigan drill hole, depth 7,999.9 ft (2,438.5 m) (location shown in figure 1 ). 

Microstructures in Cambrian Rocks, Missouri and Arkansas A7 
N 

20 percent 

10 percent 

-- N=509 

Figure 6. Two sets of fracture attitudes are in core from the 

Dow Chemical No. 1 Garrigan drill hole. Set one is parallel 

with the rift axis, and set two is perpendicular to the rift axis. 

These fracture attitudes match fracture orientations in the fluorspar mineral district (T.S. Hayes and L.W. Snee, oral commun., 

1991). 

Figure 7. Closeup view of fracture shown in figure 4. The 

fracture disperses in the mineral front. Ferroan calcite is the last 

mineral to precipitate in both the mineral front and the fracture. 

stylolites. Copper mineralization in this drill hole seems to 

be confined to stylolites in the Lamotte Sandstone. 

STH-2 Drill Hole 

Drill-hole STH-2 is also on the carbonate platform. 

Healed vertical fractures contain pyrite and dolomite but do 

not demonstrate multiple episodes of extensional movement 

(table 1 ). In general, cores from Cambrian strata away from 

the Reelfoot Rift show less evidence of the tectonic episodes 

that produced multiple sets of vertical fractures within the 

rift. 

a: 

UJ 

~ 2 :::> 

z 

EXPLANATION 

Mineral front 

Barren 

0 .2 0.4 0.6 0 .8 1.0 

SULFUR, IN WEIGHT PERCENT 

Figure 8. Total sulfur contents are elevated in the diagenetic mineral fronts, from ~0.01 weight percent sulfur in barren 

host rock to as much as 1.05 weight percent sulfur in the 

mineral fronts. 

317-A Drill Hole 

Although the number of extensional microfracture reactivations decreases northwest from the rift, stylolites 

continue to record multiple precipitation and deformation 

events in the Lamotte Sandstone and in carbonate rocks on 

the platform (fig. 11). In drill-hole 317-A, stylolites are 

zoned with layers of clay, organic matter, and carbonate 

minerals (table 1). Right-angle shearing of normally ductile 

residual material concentrated in stylolites is common. Each 

deformation feature observed in a stylolite may not be due to 

the same tectonic event, but each can be related to the introduction of a mineralizing fluid. 

Ozark Uplift 

Structurally controlled lead-zinc ore deposits surround 

Precambrian granite knobs of the St. Francois Mountains 

(Ohle and Brown, 1954) (fig. 1). Ore deposits of this type are 

present in both the Old and New Lead Belts. Ore minerals, 

particularly in the Old Lead Belt, are concentrated along stylolitic zones, especially in compactional breccias (Tarr, 

1936; Ohle, 1985; Hayes and others, 1989). 

16-15 Drill Hole 

Drill-hole 16-15 is within the Viburnum Trend in the 

Ozark Uplift (fig. 1). Ore mineralization is common along 

clay-filled stylolites throughout the dolomitized Bonneterre 

Formation in this drill hole. Slickensides are abundant in the 

residual clay that fills the stylolites. As clay-lined stylolites 

underwent slippage, galena deformed ductilely along 

A8 Application of Structural Geology to Mineral and Energy Resources 
20,-----------------~----------------~ 

z 

§ 15r--

....J 

~ 

a: 

UJ 

a.. 

(/) h: 10-

~ 

z 

~ ...... ~ -- . ~ 5- •.• --. u , •. - . 

- -

• 

o~----------------.....1----------------~ 50,-----------------,------------------. 

z 

g 

....J 

....J 

~ 

401-

ffi 30 i--

a.. 

(/) 

1--

a: 

~ 

a.. 

z 

;;. 20-

~ 

UJ 

.....J •.. ·· ~ ·~-~/ 

... ///}/ 

10!t- ~-/~ : • 

• 

-

-

O I 0'--------------------0.....l..-5 ________________ _JLO 

SULFUR, IN WEIGHT PERCENT 

Figure 9. Several trace elements, including lead and cobalt, 

positively correlate with sulfide sulfur in mineral fronts in the 

Dow Chemical No. 1 Garrigan drill hole. This correlation may 

imply that these elements were introduced during the sulfidization event. 

cleavage planes rather than undergoing brittle failure 

because it is a softer mineral than pyrite or marcasite. The 

deformed galena encloses and partly replaces pyrite and 

chalcopyrite (fig. 12, table l). 

B-4 and BCC-4 Drill Holes 

In drill-holes B-4 and BCC-4, north of the St. Francois 

Mountains (fig. 1 ), sulfide minerals along stylolites show 

Figure 10. Sequential precipitation and brecciation of marcasite and pyrite along a stylolite indicates that stylolites were 

active sites for both the deposition and dissolution of minerals. 

Drill-hole GP-1 0, depth 1 ,817 .2 ft (553.9 m) (location shown 

in figure 1). 

Figure 11. Layering of mineralization in a deformed stylolite. 

Note the right-angle shearing. Drill-hole 317-A, depth 47 4 ft 

(144.5 m) (location shown in figure 1 ). 

similar alternating mineral precipitation and defonnation 

fabric similar to that observed in other cores (table 1). For 

example, sulfides in stylolites in drill-hole B-4 are brecciated, similar to those in drill-hole GP-10. The breccia fragments are very angular and internally fractured. 

Stylolites or pressure solution seams in drill-hole BCC-4 

are lined with pyrite. Quartz grains along the stylolite boundaries are partly dissolved, whereas remnant intergranular 

pyrite is not {fig. 13). Parallel sets of pyrite-filled stylolitic 

seams are present repetitively throughout the drill hole. 

Well-developed geopetal structures that suggest development in a vadose environment (Walker and others, 1978) are 

detrital clay minerals on the upper surfaces of quartz grains 

Microstructures in Cambr·ian ·Rocks, Missouri and Arkansas A9 
Figure 12. Black triangular pits are curved cleavage planes in 

galena, possibly due to ductile deformation along a stylolite. 

Slickensided clay in stylolites show thrust movement. Drillhole 16-15, depth 915 ft (278.9 m) (location shown in figure 

and pendant quartz cement. Later cementation of these geopetal structures by sparry calcite indicates precipitation in a 

phreatic environment. 

The development of the pyrite-filled stylolitic seams 

probably was related to a fluctuation of static water level. 

The change in relative position of the static water level could 

have occurred through a change in recharge rates that was 

climatically controlled or through tectonic movement such 

as the rotation of fault blocks. 

DISCUSSION 

Microstructures and Mineralization 

Nonsutured low-amplitude stylolites are commonly layered with a variety of minerals. The mineral layers consist of 

(1) residual clay minerals and organic matter, (2) authigenic 

carbonate minerals and quartz, and (3) both residual and epigenetic iron sulfide minerals. These stylolites are weak surfaces, susceptible to deformation during tectonic movement. 

The presence of slickensided clay in stylolite seams and 

brecciated layers is evidence for bedding-parallel slip that 

indicates changes in stress orientation during or subsequent 

to the formation of the stylolite. The breccia is interpreted to 

be tectonic in origin and not a result of simultaneous precipitation and dissolution along the stylolite. Tectonically 

deformed stylolites commonly contain barite, monazite, 

pyrite, marcasite, thorite, and sphalerite. The probable epigenetic origin of these minerals suggests that the stylolites 

are not just pressure solution features but are also active sites 

for fluid flow and precipitation of minerals. 

Vertical, intergranular, mineral-healed veinlets are associated with stylolites and are commonly offset along the 

Figure 13. Quartz grains have been partially dissolved along 

a stylolitic boundary either by pressure solution or introduction of an acidic fluid; however, pyrite has been unaffected. 

Remnant "arms" of intergranular pore-filling pyrite remain 

above the partially dissolved quartz grains. Drill-hole BCC-4, 

depth 1,284 ft (391 .4 m) (location shown in figure 1 ). 

stylolite surfaces. The veinlets or fractures probably are 

extensional in origin, possibly due to tectonic stretching of 

the layered sedimentary rocks. Structurally controlled fluid 

flow through these veinlets is indicated by repeated opening 

and resealing of fractures by successive generations of 

quartz, barite, siderite, dolomite, clay, calcite, and pyrite. 

Both the number of vein generations and the number of epigenetic minerals filling those veins are greatest in the Reelfoot Rift and decrease in frequency to the northwest across 

the carbonate platform and into the Ozark Uplift. This trend 

may indicate an association between the formation of fractures, fluid movement in the rift, and local seismic activity. 

Sulfide mineralization in the stylolites and microfractures demonstrates that these microstructures existed as 

metal-bearing fluid-flow paths in the Lamotte Sandstone and 

Bonneterre Formation. Although we are not necessarily 

observing the same deformational event along stylolites in 

the Reelfoot Rift and the Ozark Uplift, each deformation 

event recorded in the stylolites and associated fractures can 

be related to an introduction of mineralizing fluids. 

Epigenetic mineral fronts in rift sedimentary rocks demonstrate at least local migration of ore-bearing fluids. That 

fluid migration, combined with similarities in heavy sulfur 

isotopes in sulfide minerals in the rift and Ozark Uplift, suggests that ore-bearing fluids may have passed through the rift 

sedimentary rocks onto the carbonate platform and into the 

Ozark Uplift. This supports the general concept of the rift as 

a flow path for ore-bearing fluids. 

Other lines of evidence also suggest that ore-bearing fluids that formed the ore districts originated in or passed 

through the Reelfoot Rift. 

A 10 Application of Structural Geology to Mineral and Energy Resources 
0 100 

I I I I I I II I I 

200 MILES 

I 

0 100 200 KILOMETERS 

Figure 14. The Appalachian and Ouachita foldbelts may have 

created a hydraulic head through topographic relief in the Late 

Pennsylvanian to Early Permian that influenced the flow of hot 

brines north and northwest (bold arrows) to form the Mississippi 

Valley-type deposits (Leach and Rowan, 1986). Dates for mineralization in the lead-zinc districts overlap with timing of the 

Ouachita tectonism. Heavy lines are province boundaries; 

dashed lines indicate projections of boundaries beneath sediments of Mississippi Embayment. 

1. Lead, zinc, and copper are enriched in acid-insoluble 

residues derived from Cambrian platform carbonate rocks in 

the area between the ore districts and Reelfoot Rift (Erickson 

and others, 1988; M.B. Goldhaber, unpub. data). 

2. Authigenic potassium feldspar is widespread in the 

Lamotte Sandstone. A broad zone of dissolution of these 

overgrowths in the Lamotte Sandstone adjacent to the northwest boundary of the rift suggests that an acidic fluid, spatially related to the rift, was responsible for the dissolution. 

An acidic pore fluid would have also been capable of carrying metals and H2S in solution (Diehl and others, 1989). 

3. Sulfur isotope data for epigenetic sulfide minerals in 

Cambrian arenites in the rift are similar to data for mainstage-mineralized rocks in the Viburnum Trend (fig. 3). 

Mechanisms for Fluid Migration 

Models proposed to induce flow of metal-bearing brine 

fluids include gravity-driven, compaction-driven, and seismic-pumping mechanisms (Bethke, 1985; Leach and 

Rowan, 1986; Clendenin and Duane, 1990). Leach and 

Rowan ( 1986) postulated that the Ouachita orogeny may 

have created a hydraulic gradient through topographic relief 

that influenced the flow of hot brines from the Arkoma Basin 

northward to the Ozark region to form the Mississippi 

K/ Ar glauconite, Viburnum Trend 

Rb/Sr adularia veins, Garland County, 

Arkansas 

Ar/ Ar authigenic potassium feldspar, 

southeast Missouri 

K/ Ar illite, postmineralization, 

St. Francois Mountains vicinity 

Paleomagnetic age, Viburnum Trend 

Paleomagnetic age, northern Arkansas 

Youngest mineralized rocks in Ozark region 

U/Pb basement alteration 

Northern Illinois Basin 

Ouachita convergent tectonics 

• 

I 

I 

• 

• -

-

• 

400 300 200 

AGE, IN MILLIONS OF YEARS 

Figure 15. Authigenic minerals and their ages in relation to 

the timing of Mississippi Valley-type ore deposition (Posey and 

others, 1983; Rothbard, 1983). 

Valley-type deposits (fig. 14). Their interpretation is supported by dates for mineralization in the lead-zinc districts 

that overlap with the timing of Ouachita tectonism (fig. 15). 

Arguments for the hydraulic gradient theory are based primarily on a regional study of fluid inclusions and inferred 

homogenization temperatures as a function of distance from 

the Ouachita front. We do not yet have data on fluid inclusions in samples from the Reelfoot Rift and carbonate platform, but our geochemical, isotopic, petrographic, and 

microstructural data support the concept that ore fluids may 

have originated or migrated through the deep basins associated with the Reelfoot Rift (McKeown and others, 1990). 

Numerical models of compaction-driven flow by Bethke 

(1985) indicate that compaction may induce the migration of 

ore-bearing brines but that velocities were too low to have a 

causal effect on the formation of the Mississippi Valley-type 

ores. Gravity-driven flow is the most likely mechanism but 

requires topographically high areas around the Ozarks. Both 

the Ouachita and Appalachian foldbelts, which were initiated in the Late Mississippian and Early Pennsylvanian and 

reached their greatest development during the Late Pennsylvanian to Permian, have been suggested as topographically 

high areas during the late Paleozoic. Other structurally high 

areas in the late Paleozoic were the Pascola Arch and probably the contiguous Blytheville Arch (McKeown and others, 

1990). On the basis of restoration of the stratigraphic section 

in the Pascola area, Stearns and Marcher (1962) estimated 

that at least 2.5 km and as much as 4.5 km of rocks were 

eroded from the arch before Late Cretaceous subsidence. 

The actual relief of any postulated topographic high area 

during the late Paleozoic cannot be estimated, however, with 

any reasonable validity. The postulated topographic high 

areas are reasonable, but any one or all may have caused separate gravity-driven hydrologic-flow systems. 

The Reelfoot Rift is today a seismically active area and 

was also active at times in the geologic past. Some researchers have advocated seismic pumping (or suction pumping) as 

Microstructures in Cambrian Rocks, Missouri and Arkansas A 11 
an alternative fluid-drive mechanism (Clendenin and Duane, 

1990; Sibson, 1990). Seismic pumping probably had very 

local effects confined to the rift interior as indicated by the 

decrease away from the rift axis of the complexity of fractures and the number of generations of minerals filling fractures in Cambrian deposits (table 1). Events such as 

earthquakes and nuclear explosions produce compressional 

waves that cause fluid- pressure disturbances in aquifers. 

Regional Fracture Patterns 

A key in any argument attempting to relate ore-forming 

processes in southeast Missouri and the structural fabric of 

rift sediments presented here concerns the relative timing of 

mineral precipitation in the two areas. At present, we cannot 

establish a direct timing link, but indirect evidence implies a 

relationship. Fracture azimuths measured in the Garrigan 

drill hole in the Reelfoot Rift (fig. 6) are subparallel with 

those observed in the Illinois-Kentucky fluorspar district, 

northeast of the rift (fig. 1). Mica from peridotite dikes filling 

fractures whose directions are close to set two of figure 6 has 

been dated at 275 Ma (Zartman, 1977). Mineralization filling 

fractures whose orientations are subparallel with set one (fig. 

6) has recently been dated at 200-240 Ma (T.S. Hayes and 

L.W. Snee, oral commun., 1991). If the similarity in fracture 

direction in the two parts of the rift indicates a similarity in 

time of formation, then inspection of figure 15 indicates that 

the epigenetic phases described here could have formed at 

the time of Mississippi Valley-type ore genesis. 

SUMMARY 

Based on our observations in southeast Missouri and 

northeast Arkansas, stylolites and fractures acted as fluidflow conduits throughout the subsurface, and mineralization 

accompanied deformation events recorded in these microstructures. Another implication of our work is that the 

Lamotte Sandstone and the underlying siliciclastic units, 

informally referred to as the Reelfoot arkose in the rift 

(Houseknecht, 1989), may not have acted as homogeneous 

aquifers for migrating fluids. Evidence in these rocks suggests microstructural control of fluid flow, as well as intergranular flow through pore space. 

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Natural Resources Special Report, Proceedings, p. 39-43. 

Swolfs, H.S., McKeown, F.A., and Glick, E.E., 1990, Episodes of 

extension recorded in Cambrian rocks in the Reelfoot Rift, 

Mississippi Embayment: Eos, v. 71, no. 17, p. 638. 

Tarr, W.A., 1936, Origin of the southeastern Missouri lead deposits, part I: Economic Geology, v. 31, p. 712-754. 

Taylor, M.E., Collins, D.S., Repetski, J.E., and Palmer, A.R., 1991, 

Upper Cambrian biostratigraphic correlations in the Reelfoot 

basin, northeastern Arkansas [abs.], in Lowell, G.R., and 

Clendenin, C.W., eds., Louis Unfer, Jr., Conference on Geology of the Mid-Mississippi Valley Region: Southeast Missouri State University, June 13-14, 1991, Cape Girardeau, 

Missouri, Missouri Department of Natural Resources Special 

Report, Proceedings, p. 51-55. 

Walker, T.R., Waugh, B., and Crone, A.J., 1978, Diagenesis in 

first-cycle desert alluvium of Cenozoic Age, southwestern 

United States and northwestern Mexico: Geological Society 

of America Bulletin, v. 89, p. 19-32. 

Weaverling, P.H., 1987, Early Paleozoic tectonic and sedimentary 

evolution of the Reelfoot-Rough Creek rift system, midcontinent, U.S.: Columbia, University of Missouri, M.S. thesis, 

116 p. 

Zartman, R.E., 1977, Geochronology of some alkalic rock 

provinces in eastern and central United States: Annual 

Review of Earth and Planetary Sciences, v. 5, p. 257-286. 

Microstructures in Cambrian Rocks, Missouri and Arkansas A 13 

Chapter B 

Tectonic Controls on Deposition and 

Source-Rock Properties of the Woodford Shale, 

Anadarko Basin, Oklahoma-Loading, 

Subsidence, and Forebulge Development 

By TIMOTHY C. HESTER, JAMES W. SCHMOKER, and 

HOWARD L. SAHL 

U.S. GEOLOGICAL SURVEY BULLETIN 2012 

APPLICATION OF STRUCTURAL GEOLOGY TO MINERAL AND ENERGY 

RESOURCES OF THE CENTRAL AND WESTERN UNITED STATES 
CONTENTS 

Abstract Bt 

Introduction Bt 

Geologic and structural setting Bt 

Depositional patterns B3 

Thermal-maturity patterns BS 

Structural implications BS 

Summary and conclusions BtO 

References cited Btl 

FIGURES 

1. Map showing study area, lines of cross sections, locations of wells, and major structural features, Anadarko Basin, and log suite showing characteristic signatures of 

lower, middle, and upper members of Woodford Shale B2 

2-6. Maps showing: 

2. Thickness of Woodford Shale B3 

3. Depth below sea level of top of Woodford Shale B4 

4. Thickness of lower member of Woodford Shale BS 

5. Thickness of middle member of Woodford Shale B6 

6. Thickness of upper member of Woodford Shale B7 

7. Cross sections of lower, middle, and upper members of Woodford Shale showing 

influence of tectonic evolution of Anadarko Basin on distribution of 

sediments B8 

8, 9. Maps showing: 

8. Vitrinite reflectance of Woodford Shale B9 

9. Thickness of lower member of Woodford Shale, major erosional channels on 

pre-Woodford surface, and distribution of Misener Sandstone BlO 

10. Diagram showing decompacted thickness of middle member of Woodford 

Shale Btl 

811 Contents 
Tectonic Controls on Deposition and Source-Rock 

Properties of the Woodford Shale, Anadarko Basin, 

Oklahoma-Loading, Subsidence, and 

Forebulge Development 

By Timothy C. Hester1

, James W. Schmoker2, and Howard L. SahP 

Abstract 

The Woodford Shale (Late Devonian-Early Mississippian) is an organic-rich, highly compacted black shale that is a 

significant hydrocarbon source rock in the Anadarko Basin of 

Oklahoma. The Woodford can be subdivided into lower, middle, and upper informal members based on geophysical-log 

character. Depositional patterns of these three members were 

shaped by the tectonic evolution of the Southern Oklahoma 

Aulacogen. 

Depositional patterns of the Woodford Shale reveal a 

positive paleotopographic feature, parallel with and about 75 

mi (120 km) north of the Amarillo-Wichita Uplift, that divided 

the Woodford into northeast and southwest depocenters and 

was a hinge line separating areas of differential basement 

movement during Woodford time. Three lines of evidence suggest that the paleotopographic high was tectonically rising 

before and during a Late Devonian (pre-Woodford) episode of 

regional erosion and throughout Woodford time: (1) erosional 

channels on the pre-Woodford unconformity do not cross the 

axis of the paleotopographic high; (2) the Late Devonian Misener Sandstone, which was deposited on the pre-Woodford 

unconformity in topographic lows, is not present along the axis 

of the paleotopographic high; and (3) Woodford sediments 

thicken from the paleotopographic high toward both the southwest and northeast. 

The paleotopographic high can be explained as a forebulge that developed on the basin margin as the direct result of 

loading and subsidence along the central trough of the Southern Oklahoma Aulacogen. The ratio of basin subsidence to 

uplift of the paleotopographic high is in reasonable agreement 

with a theoretical forebulge model. The distance between the 

1

U.S. Geological Survey, Box 25046, MS 971, Denver, Colorado 

80225. 2

U .S. Geological Survey, Box 25046, MS 960, Denver, Colorado 

80225. 3 Advantage Resource, Inc., 1775 Sherman Street, Suite 1375, 

Denver, Colorado 80203. 

paleotopographic high and the load axis of the Southern Oklahoma Aulacogen is comparable to that between the Goodman 

Swell (a forebulge of the Midcontinent Rift System) and its 

nearest load axis. 

INTRODUCTION 

The Woodford Shale is one of several organic-rich black 

shales of Late Devonian and Early Mississippian age in 

basins of the North American Craton. If thermally mature, 

these black shales are economically important as hydrocarbon source rocks. The Woodford Shale is widely regarded as 

a major source of oil and gas in the Anadarko Basin of Oklahoma (Cardott and Lambert, 1985). 

Source-rock properties of the Woodford Shale are dependent on original depositional patterns and subsequent 

thermal-maturation history. This report describes effects of 

the tectonic evolution of the Anadarko Basin on depositional 

and thermal-maturity patterns of the Woodford Shale in the 

Oklahoma part of the Anadarko Basin (fig. 1). The geologic 

and structural setting and the depositional and thermalmaturity patterns of the Woodford are reviewed, and then the 

significance of these patterns and their tectonic implications 

are discussed. 

The data on which conclusions of this report are based 

are mostly from Hester and others ( 1990). The present report 

focuses on relations between tectonic evolution and Woodford sedimentation, rather than on the physical character of 

the formation, as is emphasized by Hester and others (1990). 

GEOLOGIC AND STRUCTURAL SETTING 

The Woodford Shale is a highly radioactive, carbonaceous and siliceous, dark-gray to black shale. In terms of 

geophysical-log character, the Woodford can be generally 

Tectonic Controls, Woodford Shale, Anadarko Basin, Oklahoma B1 
I 

I 

Approximate boundarY! ~ r 

--._Qf Anadarko basin I ' --- _____ . ....__....__ I . -----...J --- l 

I AMARIL::HIT:1 -:PLIF-;--...._........_ -........... 

10 20 30 40 50 60 MILES 

I I I I I I 

I I I I I I 

10 20 30 40 50 60 KILOMETERS 

GAMMA-RAY 

PostWoodford 

PreWoodford 

DENSITY RESISTIVITY 

Upper member 

Lower member 

Figure 1. Map showing study area, lines of cross sections, locations of wells (solid circles) used in this study, and major structural features, Anadarko Basin. Log suite shows characteristic signatures of lower, middle, and upper members of Woodford 

Shale. Modified from Hester and others (1990). 

described as two similar shales separated by a less dense, 

more radioactive, and commonly more resistive middle 

shale member. For this reason, and to better document variations and regional trends within the Woodford, the formation is considered here in terms of three informal members 

(Hester and others, 1988)-the lower, middle, and upper 

members of the Woodford Shale (fig. 1). The primary physical basis for this subdivision is the higher kerogen content 

of the middle member. Total organic carbon (TOC) calculated from log-derived formation density (Hester and others, 

82 Application of Structural Geology to Mineral and Energy Resources 
3r ~------------ OO o __________________ ~gg~o------------------~9~8o __________________ ~g~r 

10 20 30 40 50 60 MILES 

I I I I I I 

I I I I I I 

0 10 20 30 40 50 60 KILOMETERS 

Figure 2. Thickness of Woodford Shale. Heavy line with arrows marks axis of paleotopographic high that separates Woodford 

into northeast and southwest depocenters. Area where Woodford Shale is absent is hachured. Contour interval 25 ft (7.6 m). 

Solid circles show locations of wells used in this study. Modified from Hester and others (1990). 

1990) averages 3.2, 5.5, and 2.7 weight percent within the 

study area (fig. 1) for the lower, middle, and upper members 

of the Woodford Shale, respectively. 

The Woodford Shale was deposited on a major regional 

unconformity developed in Late Devonian time (Amsden, 

1975) and is conformably overlain by shale and limestone of 

Mississippian age. Total thickness of the Woodford ranges 

from near zero to about 125 ft (40 m) on the northern shelf 

areas and increases to more than 700 ft (210m) in limited 

parts of the deep Anadarko Basin (Amsden, 1975). Maximum thickness in the area of this study is about 300 ft (90 m) 

(fig. 2). 

The Anadarko Basin is a two-stage Paleozoic basin 

(Perry, 1989). The Woodford was deposited on the predominantly carbonate sediments of the first-stage Southern Oklahoma Aulacogen (Feinstein, 1981). The present-day 

configuration of the basin, reflected by structure on top of 

the Woodford (fig. 3), developed primarily during a postWoodford second-stage episode of foreland-style downwarping. 

DEPOSITIONAL PATTERNS 

The axis of a positive paleotopographic feature that influenced deposition throughout Woodford time is shown in figure 2. This positive feature, parallel with and about 75 mi 

(120 km) north of the (post-Woodford) Amarillo-Wichita 

Uplift, divided the Woodford into southwest and northeast 

depocenters and was a hinge line separating areas of differential basement movements. 

Northeast of the hinge line, the maximum thickness of 

the lower member of the Woodford is only about 25 ft (8 m). 

Southwest of the hinge line, toward the deep axis of the 

southern Oklahoma aulacogen, the lower member thickens 

to more than 150 ft ( 46 m). Local thickness variations of the 

lower member (fig. 4) result from the filling of topographic 

depressions on the eroded surface of the pre-Woodford 

unconformity. During deposition of the lower member, the 

northeastern part of the study area was quite stable as compared to the central trough of the Southern Oklahoma Aulacogen, which was actively subsiding. 

Tectonic Controls, Woodford Shale, Anadarko Basin, Oklahoma 83 
3r .-------------~10 0o ____________________ gg --------------------~98 ----~-------------- o 

STUDY 

AREA , 

10 20 30 40 50 60 MILES 

I I I I I I 

I I I I I I 

0 10 20 30 40 50 60 KILOMETERS 

Figure 3. Depth below sea level of top of Woodford Shale. Contour interval 1,000 ft (305 m). Solid circles show locations of 

wells used in this study. Modified from Hester and others (1990). 

Because deposition of the lower member covered surface 

irregularities of the pre-Woodford unconformity, the thickness of the middle member of the Woodford (fig. 5) is locally 

more uniform than that of the lower member. Like the lower 

member, the maximum thickness of the middle member in 

the northeastern part of the study area is only about 25 ft (8 

m). The thickness of the middle member increases toward 

the axis of the Southern Oklahoma Aulacogen to more than 

75 ft (23 m). Figure 5 indicates continued stability of the 

northeastern part of the study area relative to the central 

trough of the Southern Oklahoma Aulacogen, which was 

subsiding during middle Woodford time but apparently more 

slowly than during deposition of the lower member. 

The thickness of the upper member of the Woodford 

Shale in the northeastern part of the study area (fig. 6) is 

more than 75 ft (23 m). In contrast to the lower and middle 

members, isopachs of the upper member show significant 

thickening to the northeast toward Kansas. Also in contrast 

to the lower and middle members, the upper member thickens only slightly to the southwest. Figure 6 shows a shift in 

tectonic pattern; during upper Woodford time regional subsidence of the northeastern part of the study area was coupled with continued slowing of subsidence along the axis of 

the Southern Oklahoma Aulacogen. 

Depositional relations between the three members of the 

Woodford Shale, depicted in the regional cross sections of 

figure 7, show that the influence of the first stage of Anadarko Basin tectonic evolution continued at least into Mississippian time. To the southwest, each member of the 

Woodford thickens, reflecting subsidence along the central 

trough of the Southern Oklahoma Aulacogen and relative 

uplift along the axis of the paleotopographic high shown in 

figure 2. To the northeast, the almost uniform thickness of 

the lower and middle members records deposition on a relatively stable shelf, whereas increasing thickness of the upper 

member records downwarping toward the Sedgwick Basin 

of south-central Kansas. 

As a result of the tectonic influence discussed above, 

thickness of the Woodford Shale in the northeastern part of 

the study area is dominated by the upper member and 

84 Application of Structural Geology to Mineral and Energy Resources 
100° 99° 98° 97° 37° .--------------,--------------------~----------------------,---------------------,-, 

I 

·---i 

10 20 30 40 50 60 MILES 

I I I I I I 

I I I I I I 

0 10 20 30 40 50 60 KILOMETERS 

Figure 4. Thickness of lower member of Woodford Shale. Areas where lower member is absent are hachured. Contour interval 

25ft (7.6 m). Solid circles show locations of wells used in this study. Modified from Hester and others (1990). 

thickness of the Woodford Shale in the southwestern part of 

the study area is dominated by the lower and middle members. The effect of these depositional patterns on the sourcerock properties of the Woodford Shale is significant when 

coupled with the thermal-maturity patterns discussed in the 

following section. 

THERMAL-MATURITY PATTERNS 

Depositional patterns of the Woodford Shale were 

mostly shaped by evolution of the first-stage Southern Oklahoma Aulacogen, whereas thermal-maturity patterns were 

shaped by development of the second-stage foreland-style 

Anadarko Basin. Subsidence of the Woodford Shale during 

Mississippian through Permian time produced a large range 

of burial depths (fig. 3) and thus an unusually broad range of 

thermal maturities (Cardott and Lambert, 1985). Vitrinite 

reflectance (Ro) ranges from slightly less than 0.5 percent on 

the northern shelf to well over 2.0 percent in the deep basin 

(fig. 8). 

The Woodford Shale on the northeast shelf, where thickness is dominated by the upper member, is generally immature to marginally mature with respect to oil generation 

(Schmoker, 1986). In contrast, the Woodford Shale in the 

deep Anadarko Basin to the southwest, where thickness is 

dominated by the lower and middle members, is generally 

mature to postmature with respect to oil generation. Thus, 

most hydrocarbons sourced by the Woodford Shale have 

been generated by the lower and middle members. 

STRUCTURAL IMPLICATIONS 

Burial histories of Feinstein (1981) indicate very gradual 

subsidence of the Southern Oklahoma Aulacogen after 

Ordovician time; however, progressive thickening of the 

Hunton Group (Late Ordovician, Silurian, and Early 

Devonian) and the Woodford Shale toward the deep basin 

(Amsden, 1975)(figs. 2, 4-7) shows that differential subsidence between the central trough of the Southern Oklahoma 

Aulacogen and the northern margin of the basin persisted 

Tectonic Controls, Woodford Shale, Anadarko Basin, Oklahoma 85 
3r ~------------,10 0o ____________________ gg --------------------~98~o----~------------ g~r~ 

----l--· ---1 -----~-----

,---~-. ---- • I / ! 

,--- ,/ . I ---~ I / . I ~ '•--" , I I I I , 

I \ -- I : . \• . ~ --------... \ \J . I • , ______ , ,- • ',,_.) . I 

! •• I . r------.---l ~/ h · I I I /'· . r, · --~ ~ · ---~ . I 

I • I • 25 ,----- • 0 " I • ,. I 1!," • -- I • I I (1· l ... ·. 'L _L 

,:) , -~~ - . z6 -- T-- -- - lr--I----~- , I I (\ /i I I 

) · I I 

_ _ __ __ KANSAS 

~ - -OKU\HOMA 

I 

I 

I 

I 

--.,. ,. I I 

, ~------, L 

1 1 .o ••• • ! r-

/ STUDY . AREA / 

/ 

~--~ I r- I 

r--: --1 I ( ~ 

10 20 30 40 50 60 MILES 

I I I I I I 

I I I I I I 

0 10 20 30 40 50 60 KILOMETERS 

Figure 5. Thickness of middle member of Woodford Shale. Area where middle member is absent is hachured. Contour interval 

25ft (7.6 m). Solid circles show locations of wells used in this study. Modified from Hester and others (1990). 

during Hunton and Woodford time. This fact suggests that 

the basin's fundamental loading mechanisms continued to 

operate at least into Mississippian time even as the sedimentary record was being overprinted by pre-Woodford regional 

erosion. 

Several lines of evidence suggest that the paleotopographic high apparent in Woodford isopach maps (figs. 2, 4, 

5) and cross sections (fig. 7) was an active tectonic feature 

that was rising during Late Devonian-Early Mississippian 

time even as the central trough of the aulacogen was subsiding. This evidence, discussed in the following paragraphs, 

falls into three categories: (1) post-Hunton to pre-Woodford 

erosional patterns, (2) distribution of the Misener Sandstone 

(Late Devonian), and (3) Woodford Shale depositional patterns. 

The post-Hunton to pre-Woodford erosional channels 

shown in figure 9 are drawn on the assumption that the lower 

member of the Woodford is thicker where it fills depressions 

on the unconformity surface. The shape and continuity of 

these depressions thus define the erosional channels of figure 

9. Channels mapped by Amsden (1975)(inset map, fig. 9), 

southeast of the present study area, appear to conform with 

the channels mapped here. The erosional channels of this 

report are separated by the paleotopographic high into two 

distinct drainages (fig. 9). To the southwest, the channels 

emanate from the flank of the paleotopographic high and 

trend generally southward. To the northeast, the course of 

the channels deviates roughly 90° to parallel the northeast 

flank. The drainages do not merge within the study area nor 

do the erosional channels cross the axis of the paleotopographic high (fig. 9). These facts suggest that the paleotopographic high was flexing upward before and during the preWoodford episode of regional erosion. 

The Misener Sandstone, a basal facies of the Woodford, 

is thought to have been derived from eroded Simpson Group 

(Middle Ordovician) sandstones and deposited in topographic lows at the beginning of the marine transgression 

that also deposited the Woodford Shale (Amsden and Klapper, 1972). In the study area, the Misener Sandstone is discontinuous across the paleotopographic high and essentially 

B6 Application of Structural Geology to Mineral and Energy Resources 
3r ~----------~1TOO ------------------~gg~o------------------~98~o----~------------~g~r 

------ l--. --

1 

I • ' ' l 

i 76r--..... · II-' · I 

f>O-"-J L . I 

~ • I --.-~ ___j 

~~i I • I 

- . I 

r ' 

---~ : . . 

I 

10 20 30 40 50 60 MILES 

I I I I I I 

I I I I I I 

0 10 20 30 40 50 60 KILOMETERS 

Figure 6. Thickness of upper member of Woodford Shale. Areas where upper member is absent are hachured. Contour interval 

25ft (7.6 m). Solid circles show locations of wells used in this study. Modified from Hester and others (1990). 

absent along its southwest slope (fig. 9). The distribution of 

the Misener Sandstone reflects positive topography at the 

beginning of marine transgression and suggests continued 

tectonic uplift of the paleotopographic high. 

The paleotopographic high significantly influenced 

Woodford deposition through Late Devonian and into Early 

Mississippian time (fig. 7). Cross section C-C' (fig. 7), in 

particular, indicates that the paleotopographic high was rising with respect to both the southwest and northeast depocenters throughout Woodford deposition. This observation 

again favors uninterrupted tectonic uplift of the paleotopographic high. 

The three points of evidence discussed above indicate 

that the paleotopographic high was in fact a tectonic element 

that rose steadily from post-Hunton through Woodford time. 

During the same period, the central trough of the Southern 

Oklahoma Aulacogen was subsiding. It is suggested here 

that the paleotopographic high was a forebulge that developed as a direct result of loading and subsidence along the 

central trough of the Southern Oklahoma Aulacogen. 

As described by Turcotte and Schubert (1982), the 

behavior of the Earth's lithosphere on a geologic time scale 

is that of an elastic plate floating on a viscous fluid. Downward deflection of the plate in response to loading creates an 

upwarp or forebulge adjacent to the depression caused by the 

applied load. Tensional stresses in the upper part of the plate 

and compressional stresses in the lower part of the plate, 

coupled with restoring buoyancy forces, produce the forebulge (Peterman and Sims, 1988). The height and position of 

the fore bulge are dependent on the magnitude of downward 

deflection and the thickness and strength of the lithospheric 

plate. As a rule of thumb, the ratio of downward deflection 

along the load axis to uplift of the fore bulge is about 15 to 1 

(Turcotte and Schubert, 1982). 

In theory, a forebulge is an integral part of basin development in response to loading, but clear examples are not common in the ancient subsurface. A well-documented example 

that is somewhat analogous to the present case is given by 

Peterman and Sims (1988), who recognized the Goodman 

Swell in northern Wisconsin as a forebulge of the 

Tectonic Controls, Woodford Shale, Anadarko Basin, Oklahoma 87 
1--

UJ 

UJ 

LL 

z 

v) 

(/) 

UJ 

z 

~ 

S:? 

I 

INE 

50 

100 

150 

sw NE 

50 

100 

150 

B 

sw 

50 

100 

150 

200 

250 

300 

c 

NW 

F 

NE 

SE 

F' 

5:j~~" 100 

~0 UPPER 

MIDDLE 50 

LOWER 

100 

NW 

G 

SE 

G' 

t=~~±===~----~-----------------TO 

50 50 

100 100 

150 

10 20 30 MILES 

200 

10 20 30 KILOMETERS 

250 

B8 Application of Structural Geology to Mineral and Energy Resources 
Figure 7 (facing page). Cross sections drawn from contoured 

thicknesses (in feet) of lower, middle, and upper members of 

Woodford Shale showing influence of tectonic evolution of 

Anadarko basin on distribution of sediments. Datum is top of 

upper member. Lines of section shown in figure 1. Modified 

from Hester and others (1990). 

Midcontinent Rift System. The distance between the Goodman Swell and the nearest load axis of the Midcontinent Rift 

System is comparable to the distance between the paleo topographic high of this report and the axis of the Southern Oklahoma Aulacogen. The flexural uplift of the Goodman Swell 

is enhanced by a tightly curved arrangement of rift axes and 

is therefore anomalously large compared to uplift of the 

paleotopographic high identified here. 

Measuring flexural uplift of the paleo topographic high is 

not straightforward. An estimate of flexural uplift during a 

limited span of geologic time is based on the decompacted 

thickness of the middle member of the Woodford Shale at 

the crest of the paleotopographic high and at the northeast 

depocenter (fig. 10). Because the thickness of the lower 

member is affected by topographic relief on the pre-Woodford unconformity and the thickness of the upper member is 

affected by subsidence of the Sedgwick Basin north of the 

study area, only the thickness of the middle member was 

used to estimate flexural uplift. 

The thickness of the middle member of the Woodford is 

controlled by a combination of three penecontemporaneous 

processes, as shown in figure 10. Thickness due to compactional subsidence (Schmoker and Gautier, I989) of the 

underlying Woodford member and to regional sediment 

accumulation is independent of forebulge development. 

Figure I 0 indicates that the paleotopographic high, which 

is interpreted here as a possible forebulge, was uplifted about 

30 ft (9 m) during middle Woodford time. Based on a subsidence to uplift ratio of 15 to I (Turcotte and Sphubert, 

I982), the central trough of the Southern Oklahoma Aulacogen would then be expected to have subsided roughly 450 ft 

100° ggo 98° 9r 3r ,-------------~--------------------~----------------------~----~--------------~ 

---------1----T-----~-----

,' '---------- 0.50 I I I 

I I 

~------j I I 

----------!----< STUDY 

AREA// 

I 

10 20 30 40 50 60 MILES 

I I I I I I 

I I I I I I 

0 10 20 30 40 50 60 KILOMETERS 

Figure 8. Vitrinite reflectance (Ro, in percent) of Woodford Shale contoured using data of Cardott and Lambert (1985) and B.J. 

Cardott (written commun., 1987). Contour interval 0.10 percent. Modified from Hester and others (1990). 

Tectonic Controls, Woodford Shale, Anadarko Basin, Oklahoma 89 
1 00° ggo 98° 9r 3r .- -----------.----------------------~--------------------,---------------------_, 

I 

I 

I 

j AMARILLO-WICHITA 

UPLIFT 

STUDY 

AREA 

10 20 30 40 50 60 MILES 

I I I I I I 

I I I I I I 

0 10 20 30 40 50 60 KILOMETERS 

Figure 9. Thickness of lower member of Woodford Shale, major erosional channels on pre-Woodford surface inferred from 

thickness of lower member (medium lines with arrows), and distribution of Misener Sandstone (screened) as depicted by Amsden (1975). Heavy line with arrows marks axis of paleotopographic high. Areas where Woodford Shale is absent are hachured. 

Contour interval 25ft (7.6 m). Inset map of study area shows stream channels of Amsden (1975) (heavier lines to southeast). 

(140 m) during the same time interval. The thickness of the 

middle member of the Woodford projected south of the 

study area along cross section C-C' (fig. 7) to the approximate axis of the aulacogen suggests that the central trough 

probably subsided at least 430 ft (130 m) during middle 

Woodford time. Thus the ratio of basin subsidence to uplift 

of the paleotopographic high during middle Woodford time 

is in reasonable agreement with that predicted by a theoretical forebulge model. 

In view of the tectonic interpretations presented here, 

sedimentation of the Woodford Shale and thus Woodford 

source-rock properties have been directly influenced by 

crustal loading during the terminal stages of the Southern 

Oklahoma Aulacogen. Subsidence along the axis of the load 

and the resulting upward flexure of the fore bulge at the basin 

margin explain the observed regional depositional patterns 

of the Woodford Shale. 

SUMMARY AND CONCLUSIONS 

Source-rock properties of the Woodford Shale are in 

large part controlled by original depositional patterns shaped 

by evolution of the first-stage Southern Oklahoma Aulacogen and by thermal-maturation patterns shaped by development of the second-stage Anadarko Basin. This report 

emphasizes the influence of subsidence along the axis of the 

Southern Oklahoma Aulacogen and uplift at the basin margin upon the depositional patterns of the Woodford Shale. 

A positive paleotopographic feature parallel with and 

about 75 mi (120 km) north of the Amarillo-Wichita Uplift 

divided the Woodford Shale into northeast and southwest 

depocenters and was a hinge line separating areas of regional 

basement subsidence and flexure during Woodford time. 

Three lines of evidence suggest that this positive feature was 

tectonically rising during post-Hunton through Woodford 

810 Application of Structural Geology to Mineral and Energy Resources 
sw 

(CREST OF 

PALEOTOPOGRAPHIC 

HIGH) TOP OF MIDDLE MEMBER 

NE 

(NORTHEAST 

DEPOCENTER) 

3J ft REGIONAL SEDIMENT ACCUMULATION 30ft 

j15 ft CQ _____ IJ_P_LIFT OF PALEOTOPOGRAPHIC' 

----- MPACTtON-0-------_HIGH 30ft ---- FLo ----- • BASE-DFM/o-~!~ MEMBER ---t 

DLE MEi?a_____ 22 tt 

ER -----l 

17 MILES 

(27.4 KILOMETERS) 

Figure 10. Decompacted thickness of middle member of 

Woodford Shale along line of section C-C' (fig. 7) at crest of 

paleotopographic high and at northeast depocenter. Subdivisions do not represent stratigraphic units or time-sequential 

events but show contributions of three penecontemporaneous 

processes to middle-member thickness. Decompaction based 

on shale-compaction curves of Baldwin and Butler (1985). 

time: (1) erosional channels on the pre-Woodford surface do 

not cross the axis of the paleotopographic high; (2) the Misener Sandstone, which was deposited in topographic lows, is 

not present along the axis of the paleotopographic high; and 

(3) Woodford sediments thicken from the paleo topographic 

high toward the southwest and northeast depocenters. 

In view of this evidence, it is suggested here that the 

paleotopographic high is a forebulge that developed on the 

basin margin as a direct result of loading and subsidence 

along the central trough of the Southern Oklahoma Aulacogen. Subsidence along the rift axis coupled with 

simultaneous development of a fore bulge can account for the 

Woodford Shale depositional patterns described here. 

The Woodford Shale was deposited near the end of the 

evolution of the Southern Oklahoma Aulacogen, during a 

relatively short interval of geologic time. The evidence 

developed in this study regarding a possible fore bulge associated with the aulacogen is thus limited to a very narrow 

phase of the aulacogen's development. An investigation of 

selected pre-Woodford sediments, analogous to that reported 

here for the Woodford, is suggested as an approach to further 

elucidate the nature and development through time of the 

paleotopographic high interpreted here as a forebulge. 

REFERENCES CITED 

Amsden, T.W., 1975, Hunton Group (Late Ordovician, Silurian, 

and Early Devonian) in the Anadarko Basin of Oklahoma: 

Oklahoma Geological Survey Bulletin 121, 214 p. 

Amsden, T.W., and Klapper, Gilbert, 1972, Misener Sandstone 

(Middle-Upper Devonian), north-central Oklahoma: American Association of Petroleum Geologists Bulletin, v. 56, no. 

12, p. 2323-2334. 

Baldwin, Brewster, and Butler, C.O., 1985, Compaction curves: 

American Association of Petroleum Geologists Bulletin, v. 

69,no.4,p.622-626. 

Cardott, B.J., and Lambert, M.W., 1985, Thermal maturation by 

vitrinite reflectance of Woodford Shale, Anadarko Basin, 

Oklahoma: American Association of Petroleum Geologists 

Bulletin, v. 69, no. 11, p. 1982-1998. 

Feinstein, Shimon, 1981, Subsidence and thermal history of Southern Oklahoma Aulacogen-Implications for petroleum 

exploration: American Association of Petroleum Geologists 

Bulletin, v. 65, no. 12, p. 2521-2533. 

Hester, T.C., Sahl, H.L., and Schmoker, J.W., 1988, Cross sections 

based on gamma-ray, density, and resistivity logs showing 

stratigraphic units of the Woodford Shale, Anadarko Basin, 

Oklahoma: U.S. Geological Survey Miscellaneous Field 

Studies Map MF-2054, 2 plates. 

Hester, T.C., Schmoker, J.W., and Sahl, H.L., 1990, Log-derived, 

regional source-rock characteristics of the Woodford Shale, 

Anadarko Basin, Oklahoma: U.S. Geological Survey Bulletin 1866-D, 38 p. 

Perry, W.J., Jr., 1989, Tectonic evolution of the Anadarko Basin 

region, Oklahoma: U.S. Geological Survey Bulletin 1866-A, 

19 p. 

Peterman, Z.E., and Sims, P.K., 1988, The Goodman Swell-A 

lithospheric flexure caused by crustal loading along the Midcontinent Rift System: Tectonics, v. 7, no. 5, p. 1077-1090. 

Schmoker, J.W., 1986, Oil generation in the Anadarko Basin, Oklahoma and Texas-Modeling using Lopatin's method: Oklahoma Geological Survey Special Publication 86-3, 40 p. 

Schmoker, J.W., and Gautier, D.L., 1989, Compaction of basin sediments-Modeling based on time-temperature history: Journal of Geophysical research, v. 94, no. B6, p. 7379-7386. 

Turcotte D.L., and Schubert, Gerald, 1982, Geodynamics, applications of continuum physics to geological problems: John 

Wiley, New York, 450 p. 

Tectonic Controls, Woodford Shale, Anadarko Basin, Oklahoma 811 

Chapter C 

Sequential Laramide Deformation of the 

Rocky Mountain Foreland of 

Southwestern Montana, Wyoming, and 

North-Central Colorado 

By WILLIAM J. PERRY, JR., DOUGLAS J. NICHOLS, 

THADDEUS S. DYMAN, and CHRISTOPHER J. HALEY 

U.S. GEOLOGICAL SURVEY BULLETIN 2012 

APPLICATION OF STRUCTURAL GEOLOGY TO MINERAL AND 

ENERGY RESOURCES OF THE CENTRAL AND WESTERN UNITED STATES 
CONTENTS 

Abstract C1 

Introduction C1 

Contrasting models of Laramide sequence of deformation C3 

Blacktail-Snowcrest Uplift CS 

Madison-Gravelly Uplift C8 

Ancestral Teton-Gros Ventre and Wind River Uplifts C8 

Eastward progression of Laramide deformation in Wyoming C1 0 

Wind River Basin C1 0 

Hanna Basin region C1 0 

Colorado Front Range C11 

Economic implications C11 

Summary C11 

References cited C11 

FIGURES 

1, 2. Maps showing: 

1. Principal Laramide basins and uplifts, Rocky Mountain foreland 

province C2 

2. Principal Laramide uplifts and other structural features, northwestern Wyoming, 

western Montana, and east -central Idaho C3 

3. Chart showing duration of intense deformation ofLaramide uplifts keyed to age and 

to palynomorph zones C4 

4-6. Maps showing areas of active Laramide uplifts during period: 

4. 80-90 Ma C6 

5. 70-80 Ma C7 

6. 60-70 Ma C9 

C II Contents 
Sequential Laramide Deformation of the Rocky Mountain 

Foreland of Southwestern Montana, Wyoming, and 

North-Central Colorado 

By Wi II iam J. Perry, Jr. 1, Douglas J. N ichols2, Thaddeus S. Dyman1, and Christopher J. Haley1 

Abstract 

Basement-involved compressional and transpressional 

deformation of the Rocky Mountain foreland began at about 90 

Ma in the Lima region of southwestern Montana. In the Lima 

region, Laramide-style foreland deformation culminated prior 

to mid-Campanian time (approximately 80 Ma), based on 

palynostratigraphically dated sediments shed from the early 

Laramide Blacktaii-Snowcrest Uplift. The Madison-Gravelly 

Uplift, located immediately northwest of Yellowstone Park, is 

the next youngest Laramide uplift of the northern Rocky Mountain foreland in Montana. Laramide deformation in this area 

began soon after 79 Ma and was apparently completed by 

68-69 Ma, before the close of Cretaceous time. Laramide 

deformation of the Wind River Uplift, which may have been 

initiated about 85 Ma based on independent fission-track evidence, culminated in late Paleocene to Eocene time. Available 

data on timing from Montana and Wyoming suggests that Laramide deformation proceeded eastward and southward from 

southwestern Montana, reaching the Colorado Front Range by 

latest Maastrichtian time. 

Economic implications of this newly recognized 

sequence of deformation of the Rocky Mountain foreland 

include progressive blockage by structural disruption of eastward migration paths for hydrocarbons generated from Paleozoic source rocks in east-central and southeastern Idaho, 

southwestern Montana, and Wyoming. 

INTRODUCTION 

Beginning about 1978, the thinking and structural concepts of many structural geologists working in the Rocky 

Mountains began to converge on regional compression and 

1

U.S. Geological Survey, Box 25046, MS 940, Denver, Colorado 

80225 2

U.S. Geological Survey, Box 25046, MS 919, Denver, Colorado 

80225 

crustal shortening as the primary mode for Laramide deformation of the Rocky Mountain foreland (fig. 1) (Smithson 

and others, 1978, 1979; Brewer and others, 1980; Gries, 

1981, 1983a, b; Hamilton, 1981, 1988; Petersen, 1983; 

Schmidt and Garihan, 1983; Brown, 1984, 1988; Kulik and 

Schmidt, 1988). This concept was first championed by 

Chamberlin (1919, 1923) and later by Berg (1962) and 

Blackstone (1963). 

The relationship between plate tectonics and Laramide 

deformation was first explored by Lowell (1974), whose 

analysis incorporated Lipman and others' (1971) concept of 

low-angle subduction. The low-angle subduction model was 

further developed by Dickinson and Snyder (1978), who 

believed that sufficient shear developed between the base of 

the continental lithosphere and the top of the underlying subhorizontal "Laramide-style" oceanic crustal slab to facilitate 

Laramide crustal shortening in the Rocky Mountain foreland. How and when Laramide crustal shortening occurred is 

still controversial (see, for example, conflicting views by 

Hamilton, 1981, 1988; Gries, 1983b; Kulik and Schmidt, 

1988). 

The theme espoused by Armstrong (1968) that the Laramide orogeny did not begin until latest Cretaceous time has 

led to a virtual disregard by some workers of early Laramide 

events in the Rocky Mountain foreland (for example, Dickinson and others, 1988). In the Lima region of southwestern 

Montana (fig. 2), Nichols and others (1985) dated beds in the 

fining-upward sequence near the top of the foreland-derived 

conglomerates as middle Campanian (78-81 Ma). The tectonic implications of this date were discussed by Perry and 

others (1988) and are expanded upon in this report. Numerous other examples of pre-latest Cretaceous Laramide-style 

deformation of the Rocky Mountain foreland exist, as mentioned by Wilson (1936), Tweto (1975), Schmidt and Garihan (1983), and Tysdal (1986, 1988). In this report we 

demonstrate that there is a definite sequence of development 

of Laramide deformation in the Rocky Mountain foreland of 

Sequential Laramide Deformation of Rocky Mountain Foreland Cl 
- 1120- 1100 1080 1060 1040 ___ l102,:__ __ 

--r --r----,---,---r 

MONTANA 

200 

I 

I 

I 

I 

NORTH 

DAKOTA 

SOUTH 

DAKOTA 

NEBRASKA 

J ____ __, 

COLORADO I 

I 

\----

1 

I 

\ 

400 KILOMETERS 

EXPLANATION 

Area of major uplift 

Area of subdued uplift 

--0-- Sea-level contour at surface 

of Precambrian basement 

---t-•111 Basin axis-Showing 

direction of plunge 

C2 Application of Structural Geology to Mineral and Energy Resources 

--+-•111 Arch axis-Showing 

direction of plunge 

A A A Foreland thrust-fault zoneSawteeth on upthrown block 

+ + + Front of Cordilleran thrust 

belt-Sawteeth on upthrown block 

-'//// //// Left-slip fault zone-Obscured by 

younger cover 

Strike-slip fault zone 
Figure 1 (facing page). Principal Laramide basins and uplifts, 

Rocky Mountain foreland province. Major uplifts are shown in 

closely spaced dot pattern; broad positive areas in more widely 

spaced dot pattern. Sawteeth on thrust faults point into upper 

plates. AU, Axial Uplift; BB, Bighorn Basin; BHU, Black Hills 

Uplift; BSU, Blacktaii-Snowcrest Uplift; BU, Bighorn Uplift; 

CA, Casper Arch; CMB, Crazy Mountains Basin; FRU, Front 

Range Uplift; GHB, Grand Hogback; GRB, Green River Basin; 

GR-SU, Granite Mountains-Shirley Mountains (Sweetwater) 

Uplift; HB, Hanna Basin; HU, Hartville Uplift; LB, Laramie Basin; LU, Laramie Uplift; MBU, Medicine Bow Uplift; OCU, 

Owl Creek Uplift; PB, North and Middle Parks Basin; PCB, 

Piceance Creek Basin; PRB, Powder River Basin; PSU, Park-Sierra Madre Uplift; RU, Rock Springs Uplift; SCU, Sangre de 

Cristo Uplift; SSU, Sawatch-San Luis Uplift; SWB, Sand Wash 

Basin; UB, Uinta Basin; UMU, Uncompahgre Mountains Uplift; UU, Uinta Mountains Uplift; WB, Washakie Basin; WHU, 

White River Uplift; WMU, Wet Mountains Uplift; WRB, Wind 

River Basin; WRU, Wind River Uplift. Modified from Bayley 

and Muehlberger (1968). 

Montana and Wyoming, from northwest to southeast. We 

focus on those Laramide uplifts that have a Cretaceous 

growth history (fig. 3); other uplifts that show little evidence 

of growth until the late Laramide (late Paleocene and early 

to middle Eocene) are not discussed. Welts of a few tens of 

meters of inferred structural relief such as discussed by 

Merewether and Cobban (1986}, defined chiefly by minor 

unconformities, are also not discussed. We consider the 

amount of tectonism represented by the minor partitioning of 

the Early and early Late Cretaceous foreland basin 

(Schwartz, 1982; DeCelles, 1986; Schwartz and DeCelles, 

1988) to be modest, not a source of significant synorogenic 

sediments or uplift, and not of the scale discussed here. 

Contrasting Models of Laramide 

Sequence of Deformation 

Tweto (1975) marshalled available evidence concerning 

the age and sequence of Laramide deformation in Colorado 

and adjacent areas showing that the early Laramide uplifts of 

central Colorado-the Front Range, Medicine Bow, 

Park-Sierra Madre, and Wet Mountains Uplifts (fig. 1)-are 

on the site of the late Paleozoic Front Range highland. These 

uplifts began in late to very late Maastrichtian time, as indicated by closely dated synorogenic debris deposited on their 

flanks (Baltz, 1972; Tweto, 1975; Kluth and Nelson, 1988). 

Tweto (1975) concluded that the Sawatch-San Luis (fig. 1) 

and Wet Mountains Uplifts are more than 70 million years 

old, slightly older than the other Colorado Laramide uplifts. 

He also concluded that the Axial, Grand Hogback, and 

White River Uplifts (fig. 1) were the latest to develop, in 

Eocene time, whereas the Laramie and Sangre de Cristo 

100 KILOMETERS 

EXPLANATION 

,, > l A<ea of maio< uplift 

Area of subdued uplift 

---+l-- Basin axis 

l IJJ Basin axis-Showing directifn of plunge 

-----rl -- Arch axis 

I 

+ + + Front of Cordilleran thrust ieit- Sawteeth on upthrown block 

6 +~ Front of Cordilleran thrust "elt with major right-slip 

component I, 

6 6 6 Foreland thrust fault zone-\t5awteeth on upthrown block 

---- Strike-slip fault zone · 

left-slip fault zone-Obscure\:! by younger cover 

Queried relations-Obscured •by Cenozoic volcanic rocks 

Figure 2. Principal Laramide uplifts (screen pattern), northwestern Wyoming, western Montana, and east-central Idaho. 

BHB, Bighorn Basin; BSU, Blacktaii~Snowcrest Uplift; BTU, 

Beartooth Uplift; CC, Cache Creek Thrust; CMB, Crazy 

Mountains Basin; GRB, Green River Basin; GU, Gallatin Uplift; H, Hilgard Thrust System; LR, Lima region; MA, Moxa 

Arch; MGU, Madison-Gravelly Uplift; PB, Pinyon Basin; RB, 

Ruby Basin; S, Snowcrest structural terrane and thrust, southeastern margin of Blacktaii-Snowcrest Uplift; SB, Sphinx Basin; TGVU, Ancestral Teton-Gros Ventre Uplift; TRU, 

Tobacco Root Uplift; WR, Wind River Thrust; WRB, Wind 

River Basin; WRU, Wind River Uplift, WU, Washakie Uplift. 

Sequential Laramide Deformation of Rocky Mountain Foreland CJ 
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biozones 

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--

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Perry and others ( 1 988) 

r-----?- _pg_--

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PALEOCENE _r------?--p;--

r------7-Pl=-

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sprnata ~ ~ ffi_U[IjD]OJij QJ "C ~ ? 

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Tysdal and 

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Law and Johnson 

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Bryant and others ( 1 981 ); 

Kluth and Nelson (1988) 

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Figure 3. Chart showing duration of intense deformation of Laramide uplifts keyed to age and to palynomorph zones. x•s mark times during which 

Precambrian rocks were exposed on up I ifts. 
Uplifts (fig. 1) began to develop in Paleocene time. The 

varying trends ofthese uplifts require first east-west shortening in late Maastrichtian and Paleocene time, followed by 

northeast-southwest shortening in early Eocene time, assuming that shortening was dominantly perpendicular to the 

trends of the ranges. 

Hamilton ( 1981) offered the hypothesis that Laramide 

compressive deformation was caused by a 2°-4° clockwise 

rotation of the Colorado Plateau relative to the continental 

interior during late Late Cretaceous and early Tertiary time. 

This rotation can be readily eliminated as a cause for all but 

the latest Laramide events because structures bounding and 

defining the northern and northeastern margin of the Colorado Plateau (the south flank thrust of the Uinta Mountains, 

Axial Arch, Grand Hogback, and White River Uplift) are latest Laramide, middle to late Eocene in age (Tweto, 1975). 

Hamilton's mechanism may explain latest Laramide deformation in the southern and central Rocky Mountains. 

Gries (1983b) proposed a multi phase model for Laramide deformation in Colorado and Wyoming. The first 

phase is characterized by east-west compression in latest 

Cretaceous time, which she believed was caused by the westward movement of the North American plate during opening 

of the North Atlantic ocean. The second is characterized by 

northeast-southwest compression in early Tertiary (chiefly 

Paleocene) time. The third is characterized by north-south 

compression in late Laramide (Eocene) time, during the 

opening of the Arctic Ocean. This is consistent with very 

young Laramide thrusting along the south flanks of the 

Granite Mountains and Shirley Mountains in southeastern 

Wyoming (see also Love, 1971; Blackstone, 1983 ), the Medicine Bow Mountains in Colorado (Independence Mountain 

Thrust) (see also Blackstone, 1977; Park, 1977), and the 

Uinta Mountains in northeastern Utah (see also Ritzma, 

1971; Bryant and Nichols, 1988). 

Schmidt and Garihan (1983) added another dimension to 

this controversy by showing evidence of oblique left-reverse 

Laramide motions along northwest-southeast-trending fault 

zones. This evidence is consistent with Brown's (1984) 

observations that a variety of fault styles would be generated 

with overall east-west shortening if older fault zones of varying trends were reactivated during the Laramide. Kulik and 

Schmidt (1988) stressed the spatial and temporal overlap 

between Sevier (thin-skin) thrusting and Laramide 

basement-involved thrust uplifts. Few of these workers, 

however, have considered the oldest Laramide events in 

southwestern Montana, the sequence of events, and the 

implications of these events, discussed below. 

Acknowledgments. -Careful reviews by Chris Potter and 

Ben Law are appreciated, particularly the insistence by Law 

that we appreciate the complexity of the Cretaceous rocks in 

western Wyoming and by Potter that we carefully distinguish between uplift and denudation. The assistance of 

Marge MacLachlan in reformatting the charts (fig. 3) is most 

appreciated. 

BLACKTAIL-SNOWCREST UPLIFT 

Eardley (1960) recognized the Blacktail-Snowcrest Anticline of Scholten and others (1955) as a major Laramide 

uplift (figs. 1, 2) and, following the mapping by Klepper 

(1950), showed southeast-directed thrusting along the southeastern margin of this uplift. Eardley (1960, fig. 3, p. 88), 

however, mistakenly placed the Blacktail-Snowcrest Uplift 

in his "Paleocene phase (mid-Laramide)." 

Ryder (1967), Wilson (1970), and Ryder and Scholten 

(1973) reexamined the Cretaceous synorogenic conglomerates and related siliciclastic rocks of the Beaverhead Group 

(Ryder's Beaverhead Formation) of the Lima region and 

showed that both foreland-sourced and thrust-belt-sourced 

conglomerates are present (figs. 2, 3). The Lima Conglomerate (fig. 3), the basal conglomerate of the Beaverhead Group, 

is mostly made up of limestone clasts derived from unroofing of the Blacktail-Snowcrest Uplift. The overlying 

sequence, characterized by the almost ubiquitous presence 

of Proterozoic quartzite clasts, was derived from the fold and 

thrust belt to the west. Based on palynologic work by A.T. 

Cross, Wilson (1970) showed that the Blacktail-Snowcrest 

Uplift was a highland during Campanian time. More 

recently, Nichols and others (1985) determined that the 

Blacktail-Snowcrest Uplift developed between Coniacian(?) 

and middle Campanian time. Perry and others (1988) 

showed that the Blacktail-Snowcrest Uplift had reached its 

present structural relief prior to middle Campanian time. The 

middle Campanian palynologic date (78-81 Ma) of Nichols 

and others (1985) was in the fining-upward sequence near 

the top of the Lima Conglomerate, which was derived from 

the Blacktail-Snowcrest Uplift. 

Recent field work by Perry, Dyman, and Haley revealed 

the presence of Mesozoic limestone clasts in conglomerates 

of the upper part of the Frontier Formation (fig. 3) southeast 

of Lima Peaks. These clasts are as large as cobble size (20 

em diameter) near the peaks and fine southeastward away 

from the projected trace of the Blacktail-Snowcrest Uplift. 

They are composed of Lower Cretaceous gastropod limestone of the Kootenai Formation, micritic limestone of the 

lower part of the Kootenai, and oolitic limestone of the 

Jurassic Ellis Group. These clasts within the upper part of the 

Frontier become less abundant as well as smaller southeastward and generally occur within 17 km southeast of Lima 

Peaks. The thickness of the limestone-conglomerate-bearing 

stratigraphic interval also decreases markedly southeastward 

(Haley and others, 1991). The Frontier-Beaverhead contact 

is an angular unconformity near Lima Peaks; however, 

because the two units become conformable to the southeast, 

the contact is placed at the top of a green porcellanite unit in 

the Frontier-Beaverhead transition zone, the stratigraphically highest porcellanite (Dyman and others, 1991). 

Although the top of the Frontier Formation has not been 

precisely dated southeast of Lima Peaks, it is unlikely to be 

younger than Turonian; a carbonaceous siltstone sample 

Sequential Laramide Deformation of Rocky Mountain Foreland CS 
__ lJW -- mo 1W 10Bo 106o 104o ---l102,:__ __ 100o 

,---r----~---,---r-

\ 

\ 

~ ~ ., 

' ' 

1 

I 

\ 

MONTANA 

NORTH 

DAKOTA 

I I 

/ / -- ~,~~ ~ \ I J_BLACKTAIL-SNOWCREST I 

~\ ~~~~~~-------- WYOMING__jl 

\t --- \ I 

\ I 1 

~ I ,~, I 

SOUTH 

DAKOTA 

IDAHO '\.>,, I " 'yiNFERRED PRE-WIND 

'-, ',RIVER FOREBULGE I 

)...1 \ \ L-------- \ \ }\ \ \ I I ~ \--~ --r------11 i 

: I I -&-OLDEST PART OF I 

NEBRASKA 

f MOXA ARCH _j ____ __, I /IL_ __T ____________ COLORADO \ 

I / I I 

I "/ I \----

1 SAN RAFA~L I \ : SWELL 1 \ 

I / I I

. UTAH I 

(/ I 

I 

I \ 

I I 

I I _l ,_ I ----

/ -----------t ,----------------c ----- 1 ARIZONA I NEW MEXICO I ~ I 

200 400 KILOMETERS 

Figure 4. Areas of active Laramide uplifts during period 80-90 Ma. Dark pattern shows area of thrust-bounded 

uplift with more than 1,500 m of structural relief; light pattern marks area of less than 1,500 but more than 300m 

of structural relief generated during this time period. 

collected from near the top of the Frontier south of Lima 

Peaks (Dyman and others, 1991) contained primarily Cenomanian to Turonian palynomorphs typical of the Frontier 

Formation of Wyoming (Nichols and others, 1982). These 

rocks near the top of the Frontier are lithologically similar to 

other Frontier rocks of the region, the lower part of which are 

definitely Cenomanian. On this basis, Laramide uplift 

intense enough to have locally sourced conglomerates began 

about 90 Ma, or slightly earlier in the Lima region of southwestern Montana (fig. 4). The abnormal thickness (more 

than 2.1 km) of Frontier Formation in the vicinity of Lima 

Peaks suggests the possibility of a foredeep localized by tectonic loading due to thrusting of the adjacent southeastern 

margin of the Blacktail-Snowcrest Uplift as well as thin-skin 

thrusting in the Cordilleran thrust belt to the west (Dyman 

and others, 1989). 

Archean clasts (gneiss, schist, and marble) have been 

observed in the upper part of the Lima Conglomerate 

C6 Application of Structural Geology to Mineral and Energy Resources 
~ \ 

~ MONTANA I 

NORTH 

DAKOTA 

'~7 ~ \ \ 1 1 

;J T I ------

v\ (j{ ~MADISON-GRAVELLY ~ 'L d r---~:.~~----------___.l ' I WYOMING \ 

l --- :.J 

~ \ I 

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IDAHO 

SOUTH 

DAKOTA 

\J"-"" VENTRE UPLIFT 

h,~ <:RIVERUPLIFT r-------

\ \ 1 NEBRASKA r-------j 1 \\---MOXAARCH ? 

1 

I ---p---T---------- COLORADO \ 

: )li/ ~ j ___ __, 

: , : \-----

/ I 1 SAN RAFAEL (\ 

I 

1 

I SWELL-............._~ I 

I

. UTAH /~/ 

I I \.._/ 

\ 

I 

I 

I I 

I I I '--- I _ _L---- / --------,-------------[ ____ _ I ARIZONA I NEW MEXICO I r I 

200 400 KILOMETERS 

Figure 5. Areas of active Laramide uplifts during period 70-80 Ma. Dark pattern shows area of thrust-bounded 

uplift with more than 1,500 m of structural relief; light pattern marks area of less than 1,500 but more than 300m 

of structural relief generated during this time period. (d) indicates previously active uplift area that is dormant during time period;(?) indicates queried uplift extension. 

(Wilson, 1970; Perry and others, 1988) in association with 

Phanerozoic clasts derived from the uplift. Therefore it is 

reasonable to conclude that the entire Phanerozoic cover was 

stripped off at least part of the Blacktail-Snowcrest Uplift by 

78-81 Ma (median of 79.5). The median palynologic date 

for cessation of deposition of Laramide synorogenic sediments eroded from the uplift (79 .5 Ma) subtracted from 90 

Ma provides an estimate ( 10.5 Ma) of duration of uplift and 

erosion. 

Based on well data southeast of the uplift (Perry, 1986, 

drill-holes 1 and 2), a minimum of 2 km of pre-Laramide 

Cretaceous cover is estimated to have been stripped from the 

uplift. Similarly, 360m of lower Mesozoic rocks, 1 km of 

Upper Paleozoic rocks (the Snowcrest trough sequence, 

Perry, 1986, fig. 3), and 800 m of Mississippian Madison 

Formation and older Phanerozoic rocks (Perry, 1986, drillhole 2) would also have been stripped, a total of 4.16 km of 

Phanerozoic rocks during approximately 10.5 Ma. From 

Sequential Laramide Deformation of Rocky Mountain Foreland C7 
these data, we calculated an exhumation rate (England and 

Molnar, 1990) of 0.4 m/1 03 years. England and Molnar 

(1990) restricted the term "uplift" to mean uplift in terms of 

the geoid, usage not followed here because of the virtual 

impossibility of making accurate and precise estimates of the 

position of the pre-Quaternary geoid. By uplift we mean the 

uplift of a body of rock with respect to the surrounding land 

surface. Our calculated exhumation rate obviously underestimates actual uplift rate because it makes no allowance for 

the topographic relief between crest of uplift and adjacent 

basin or the depth to which the Precambrian basement was 

eroded, only the thickness of Phanerozoic strata eroded. To 

compensate for the second problem, we tried to record the 

time when basement was first exposed in an uplift, through 

biostratigraphic data, by bracketing the time when basement 

clasts, arkose, or metamorphic rock fragments first appear in 

the sediments shed from that uplift. Because uplift likely 

occurred chiefly during episodes of stick-slip uplift-margin 

faulting, a quite variable short-term uplift rate probably 

existed; we do not intend to imply continuous constant-rate 

uplift. Variations in rainfall also affect the exhumation rate. 

Despite these drawbacks, the exhumation rate is a useful 

comparative measure because it can be calculated for other 

uplifts farther to the east and southeast, for which first 

appearance of basement clasts in the material eroded from 

the uplift can be dated. Therefore the exhumation rate provides a basis for comparison of relative uplift rates, if we are 

allowed to assume that topographic relief was not vastly different between uplifts and that climatic differences were not 

so extreme as to overwhelm topographic controls on erosion 

rates. 

MADISON-GRAVELLY UPLIFT 

_ S~holten (1967) recognized that the Madison and Gravelly Ranges comprise the east and west flanks, respectively, 

of a Laramide uplift that he termed the Madison-Gravelly 

Uplift (fig. 2). The core of this uplift subsided during Tertiary and Quaternary Basin-Range extension to become the 

Madison Valley. Tysdal (1986) described the thrust faults 

and backthrusts of the Hilgard Fault System (fig. 2) along the 

complexly deformed eastern side of this uplift. This 

deformed terrane is similar to that along the southeastern 

limb of the Blacktail-Snowcrest Uplift described by Perry 

and others (1983, 1988). K-Ar and 40Arf39Ardating of hornblende from dacitic laccolithic rocks of the same suite as 

dacite dikes that intruded the Hilgard Fault System yielded 

ages of 68-69 Ma (Tysdal and others, 1986). These ages 

indicated to Tysdal and others that Laramide deformation in 

this area was virtually completed by the close of Cretaceous 

time, possibly 15 m.y. later than the cessation of growth of 

the Snowcrest structural terrane to the west (fig. 5). Tysdal 

and others ( 1986) reported a palynological age of approximately 79 Ma (early Campanian) from the lower part of the 

Livingston Formation (fig. 3), just below the oldest locally 

derived conglomerates (upper part of the Livingston) 

beneath the Sphinx Conglomerate (fig. 3), the youngest 

Laramide synorogenic deposit recognized in the Madison 

Range. If we accept that locally the major episode of Laramide deformation was over by 68-69 Ma and also accept 

that erosion proceeded synchronous with thrusting to expose 

Precambrian basement by this time, then primary Laramide 

growth and exhumation of the Madison-Gravelly Uplift took 

place during an interval of 10--11 m.y., between 79 Ma and 

68-69 Ma. During this time 3.35 km of Phanerozoic cover 

would have been eroded from the uplift (thickness measured 

from Tysdal, 1990, section D-D'), yielding an estimated 

exhumation rate of 0.3 m/1 o3 years, very close to that for the 

Blacktail-Snowcrest Uplift. In any case, this estimated rate is 

only a first approximation because the time of first exposure 

of the Precambrian basement is unknown for the MadisonGravelly Uplift. 

The contact between the Sphinx Conglomerate and 

underlying Livingston is conformable almost everywhere. 

Only the upper member of the Livingston is conglomeratic, 

consisting of interstratified sandstone and conglomerate 

(Tysdal and others, 1986). This lithofacies is similar to the 

much older Cenomanian to Turonian upper part of the Frontier Formation to the west, south of Lima Peaks. The base of 

the 200-m-thick upper (conglomeratic) member of the Livingston is 30~00 m above the interval dated as 79 Ma 

(Tysdal and others, 1986, p. 861). Laramide deformation in 

the Madison Range began probably no more than 76 Ma, 

about 14 m.y. later than that of the Blacktail-Snowcrest 

Uplift, based on estimated rapid sedimentation rates in the 

Livingston. The east-west shortening represented by the Hilgard Thrust System during very late Cretaceous time is consistent with the earliest phase of Laramide deformation in 

Colorado, described by Tweto (1975). 

ANCESTRAL TETON-GROS VENTRE AND 

WIND RIVER UPLIFTS 

Dorr and others (1977) were among the first to point out 

that the Late Cretaceous rise of the Ancestral Teton-Gros 

Ventre Uplift (fig. 2) provided a high-standing buttress of 

Precambrian rocks that deflected later structures in the frontal part of the Cordilleran thrust belt. Wiltschko and Dorr 

(1983, fig. 2) estimated the age of initial growth of the 

Ancestral Teton-Gros Ventre-Wind River Arch as about 73 

Ma (figs. 3, 5). They showed that this early phase of uplift 

continued until early Maastrichtian time. In his section nearest the northern flank of the Wind River Range in the northwestern Wind River Basin (fig. 2), southwest of the axis of 

the basin, Keefer (1965) showed conglomerates in the basal 

Lance Formation (fig. 3) that he dated only as Late Cretaceous. If this basal Lance is correlative with that described 

C8 Application of Structural Geology to Mineral and Energy Resources 
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Figure 6. Areas of active Laramide uplifts during period 60-70 Ma. Dark pattern shows area of thrust-bounded 

uplift with more than 1,500 m of structural relief; light pattern marks area of less than 1,500 but more than 300m 

of structural relief generated during this time period. (d) indicates previously active uplift area that is dormant during time period;(?) indicates queried uplift extension. 

by Law and Johnson (1979) from the subsurface west of the 

Wind River Uplift, then the presence of these conglomerates 

is compatible with the 73-Ma estimate of initiation of uplift 

by Wiltschko and Dorr (1983). Keefer (1965) also showed 

the Lance pinching out along the northeastern flank of the 

Ancestral Teton-Gros Ventre and Wind River Uplifts. 

history curves, sandstone provenance studies, and inferred 

thrust-loading along the northeastern margin of the Green 

River Basin (adjacent to the Wind River Uplift). They concluded that a "subsidence event at approximately 90 Ma may 

indicate initial uplift of the Wind River block." They also 

identified a very rapid subsidence event in the area of the 

Pinedale Anticline (adjacent to the Wind River Uplift) during Maastrichtian time, which they considered to be 

More recently, Shuster and Steidtmann (1988) developed 

a growth history of the Wind River Uplift based on burial 

Sequential Laramide Deformation of Rocky Mountain Foreland C9 
probably due to tectonic loading by the growing Wind River 

Thrust. Apatite fission-track studies by Shuster (1986) indicate "that the Wind River uplift began between 80 and 90 Ma 

and moved upward about 8 km." 

More extensive, recent apatite fission-track studies by 

Cerveny (1990) indicate that growth of the Wind River 

Uplift may have begun by 85 Ma but was clearly underway 

by 7 5 Ma, in reasonably good agreement with Wiltschko and 

Dorr (1983). Cores of the Late Cretaceous Lance Formation 

from wells drilled in the northeastern Green River Basin 

contain detrital limestone and dolomite grains (Law and others, 1986) probably derived from Paleozoic limestones being 

eroded from the adjacent Wind River Uplift during Maastrichtian time, consistent with the results of Cerveny (1990). 

Ages of uplift in the 80-90-Ma time period in the Wind 

River Uplift area may relate to a broad fore bulge that encompassed the present area of the Rock Springs Uplift as well as 

the Ancestral Teton-Gros Ventre and Wind River Uplifts 

and not to the time of initiation of these Laramide structures. 

In contrast to the Madison-Gravelly Uplift, which apparently lacks a Tertiary growth history, Cerveny (1990) 

showed that the northern part of the Wind River Uplift 

underwent most rapid cooling (presumably uplift) between 

60 and 57 Ma, soon after Precambrian basement was 

exposed based on the first appearance of arkose derived from 

the Wind River Uplift in the northern Green River Basin 

(Prensky, 1989). The arkose is in an unnamed lower Tertiary 

unit (fig. 3) dated palynologically as early to middle 

Paleocene in age (Law and Johnson, 1989). 

EASTWARD PROGRESSION OF LARAMIDE 

DEFORMATION IN WYOMING 

Wind River Basin 

The Wind River Basin in west-central Wyoming (figs. 1, 

6) occupies a critical position with respect to the sequential 

development of Laramide structure in Wyoming. Our discussion of the tectonic development of the Wind River Basin 

is summarized from Keefer ( 1965) unless otherwise stated. 

Conglomerate in the Late Cretaceous Lance Formation in the 

northwestern part of the Wind River Basin, nearest the Wind 

River Uplift, contains granule-size fragments and scattered 

pebbles of chert, silicious shale, and porcellanite. Here the 

Lance is about 350m thick (Keefer, 1965, p. A17), and only 

the lower part is conglomeratic. Keefer found no definite 

evidence for uplift of the Wind River Range during Cretaceous time, but his control was inadequate along the southwestern margin of the basin (contours dashed, no control 

points within 48 km of the northeastern flank of the Wind 

River Mountains, see his fig. 9). Perry suggests that the 

above described conglomerate in the lower part of the Lance 

was probably eroded from Frontier and older Mesozoic 

rocks exposed on the growing Wind River Uplift. 

Murphy and Love (1958) inferred that a broad, domal 

uplift developed in latest Cretaceous time on the southeastem flank of the Wind River Basin in the area of the present 

Granite Mountains. Keefer ( 1965) reached similar conclusions. In a summary of the Laramide history of the Granite 

Mountains area, Love (1971) indicated that uplift of this area 

did not begin until latest Cretaceous time and culminated in 

earliest Eocene time. He suggested that the early phase of 

this uplift may have been coextensive with that of the southcentral part of the Wind River Range. 

Keefer (1965) concluded that the Cretaceous-Tertiary 

boundary in the Wind River Basin is generally conformable, 

but that extensive downwarping occurred at this time along 

the present-day northern margin of the basin. Along the 

northeastern margin of the Wind River Basin, Keefer ( 1965) 

observed that the oldest conglomerate zones are in the lower 

Eocene Indian Meadows Formation. The oldest arkosic conglomerate in this part of the basin is at the base of the Lost 

Cabin Member of the overlying Eocene Wind River Formation. The presence of extensive lacustrine sediments, which 

first appeared in the Wind River Basin in late Paleocene time 

(Nichols and Ott, 1978; Phillips, 1983), is indicative of internal drainage that likely reflects initial growth of the Casper 

Arch and Owl Creek Uplift (figs. 1, 6) that closed the outlets 

of the basin. 

Keefer ( 1965) estimated 2. 7 km of middle to late 

Paleocene uplift in the Owl Creek Mountains and 3.2 km of 

subsidence in the adjacent Wind River Basin; these amounts 

indicate a cumulative vertical separation (uplift+subsidence) 

rate of about 1.3 m/103 years using the time scale shown in 

figure 3. Keefer estimated an additional2.6 km of uplift and 

an additional 1. 7 km of subsidence in early Eocene time, a 

cumulative vertical separation rate of about 1.2 m/1 03 years. 

He showed that thrust faults of the Casper Arch cut the lower 

Eocene Indian Meadows Formation and that the rocks 

deformed by this thrusting are erosionally truncated by the 

overlying lower Eocene Wind River Formation. These relations date the cessation of major Laramide deformation in 

the area as early Eocene. For comparability with exhumation 

rates calculated for uplifts in southwestern Montana 

(approximately 0.4 m/103 years for the Blacktail-Snowcrest 

Uplift and 0.3 m/103 years for Madison-Gravelly Uplift 

derived above), the middle to late Paleocene uplift rate for 

the Owl Creek Mountains is 0.46 m/103 years and the early 

Eocene rate 0.6 m/103 years, based on Keefer's estimates of 

topographic relief and exhumation and our time scale (fig. 

3). We find it remarkable how similar these generalized rates 

are, calculated from different data sets for different Laramide uplifts. The amount of surface relief shown by Keefer 

(1965) is merely a guess, of course. 

Hanna Basin Region 

Blackstone (1983) and Hansen (1986) summarized, 

respectively, the complex structural geometry and tectonic 

Cl 0 Application of Structural Geology to Mineral and Energy Resources 
history of the Hanna Basin region southeast of the Granite 

Mountains-Sweetwater Uplift (fig. 1) and suggested that the 

major west-northwest-trending thrust fault bounding the 

northern flank of the Hanna Basin is late Paleocene to early 

Eocene in age. A major unconformity between the Late Cretaceous Pierre Shale and middle Paleocene part of the Coalmont Formation in the North Park Basin south of the 

Medicine Bow Uplift indicates that the entire area from the 

Medicine Bow Mountains (and Park Range to the south) 

eastward to the Colorado Front Range may have been a single broad highland in latest Cretaceous time (Hail, 1965; 

Wellborn, 1977; Madden, 1979; S.B. Roberts, U.S. Geological Survey, oral commun., 1990). 

COLORADO FRONT RANGE 

Laramide basement faulting in the Front Range began at 

69.3±1.1 Ma (Wallace, 1988). Kluth and Nelson (1988) 

detailed the uplift history of the Colorado Front Range based 

on proprietary fission-track and palynologic investigations 

that are compatible with unpublished data of Nichols. They 

stated that the basement was exposed in the core of the Front 

Range by 68 Ma, as indicated by the first appearance of arkosic sediments and basement clasts in the Upper Cretaceous to 

Eocene Dawson Arkose (fig. 3). Three kilometers of Phanerozoic cover (estimated by Wallace, 1988) was removed from 

the From Range during a strikingly brief period of one (or at 

most two) million years. Based on these values an exhumation rate of3 m/103 years (or minimally 1.5 m/103 years) can 

be calculated. These rates are much greater than Laramide 

exhumation rates calculated for the Owl Creek Mountains in 

south-central Wyoming and almost an order of magnitude 

greater than rates calculated for the older Laramide uplifts of 

southwestern Montana. 

ECONIDMIC IMPLICATIONS 

A relatively simple unpartitioned Cretaceous foreland 

basin was envisioned by Dickinson and Snyder (1978, fig. 1) 

and Cross (1986), although Schwartz (1982), DeCelles 

(1986), ~v1erewether and Cobban (1986), and Schwartz and 

DeCelles ( 1988) indicated that broad tectonic welts of probably low amplitude were present in the Rocky Mountain 

foreland during late Early and early Late Cretaceous time. 

The rise of the Blacktail-Snowcrest Uplift, beginning about 

90 Ma, would have provided a barrier to eastward updip 

migration of hydrocarbons generated in the rapidly subsiding foredeep developed in front of the eastward-advancing 

Helena and Wyoming Salients of the Cordilleran Thrust Belt 

(fig. 4). Beginning about 75 Ma, a perhaps continuous structural barrier to eastward migration of hydrocarbons would 

have developed from the northeastern tip of the BlacktailSnowcrest Uplift, southward through the Madison-Gravelly 

Uplift, and southeastward through the Ancestral Teton-Gros 

Ventre-Wind River Uplifts. As a consequence of this structural barrier, hydrocarbons generated farther to the west may 

have migrated into the southern Green River Basin-Hanna 

Trough region of southern Wyoming, a major fairway of 

hydrocarbon production in Wyoming. 

SUMMARY 

Intense Laramide deformation of the Rocky Mountain 

foreland began in southwestern Montana at about 90 Ma. 

The Madison-Gravelly Uplift to the east and Ancestral 

Teton-Gros Ventre-Wind River Uplifts in western Wyoming began intense growth about 75 Ma. These two uplifts 

may have been originally coextensive (fig. 2). Near the close 

of Cretaceous time, intense Laramide deformation spread 

eastward to the Granite Mountains and Colorado Front 

Range. This sequence of events cannot be related directly to 

either Eocene rotation or displacement of the Colorado Plateau but may be related to a southeastward-advancing zone 

of decoupling in the lower to middle crust that was initiated 

in southwestern Montana about 90 Ma. 

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65-77. 

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foreiand basins and uplifts: Rocky Mountain Association of 

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anticline area, Wyoming, and at the Multiwell Experiment 

site, Colorado: U.S. Geological Survey Bulletin 1886, p. 

Hl-H21. 

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and Jarre, G.A., eds., Symposium on Wyoming tectonics and 

their economic significance: Wyoming Geological Association 23rd Annual Field Conference Guidebook, p. 145-150. 

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p. 773-796. 

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Rocky Mountain Association of Geologists, p. 271-294. 

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1, p. 159-183. 

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C14 Application of Structural Geology to Mineral and Energy Resources 
Chapter D 

Influence of Crustal Structure on the 

Course of the Arkansas River, 

SoiJth-Central Colorado 

By J(EN WATSON, DANIEL H. KNEPPER, JR., and 

MIC:HAEL W. WEBRING 

U.S. ~GEOLOGICAL SURVEY BULLETIN 2012 

APPLICATION OF STRUCTURAL GEOLOGY TO MINERAL AND 

ENERGY RESOURCES OF THE CENTRAL AND WESTERN UNITED STATES 
CONTENTS 

Abstract Dl 

Introduction Dl 

General geology D2 

Near-surface expression D4 

Crustal expression D6 

Aeromagnetic data D6 

Gravity data D9 

Discussion D9 

References cited Dll 

FIGURES 

1. Three-dimensional and plan views of study area, south-central Colorado D2 

2. Digital terrain image and location map of study area D4 

3,4. Maps showing: 

3. Northeast-trending zone of northeast-trending faults in study area D6 

4. Generalized lithology of Proterozoic rocks along Arkansas River between 

Coaldale and Parkdale, south-central Colorado D7 

5. Images of magnetic data for western half of Pueblo, Colorado, 1 °X2° 

Quadrangle D8 

6. Gravity map of western half of Pueblo, Colorado, 1 °X2° Quadrangle D9 

7. Map showing structural features, mineral deposits, and magnetic anomalies, 

Colorado and southeastern Wyoming DlO 

011 Contents 
Influence of Crustal Structure on the Course of the 

Arkansas River, South-Central Colorado 

By Ken Watson 1

, Daniel H. Knepper, Jr. 1, and Michael W. Webring 1 

Abstract 

On small-scale aerial photos, satellite images, topographic maps, and digital terrain images, the linear northeasttrending segment of the Arkansas River between Coaldale and 

Parkdale, south-central Colorado, and the somewhat longer 

associated topographic lineament suggest to even the casual 

observer an element of geologic control. This segment of the 

river is generally considered to be a superposed stream established during regional uplift and incised in late Pliocene and 

Pleistocene time. Although joints in Proterozoic crystalline 

rocks probably exert control on the path of the river on a local 

scale, evaluation of geologic maps indicates that the regional 

trend and location of the river and its associated topographic 

lineament are not controlled by a single structural feature 

(fault, shear zone, broad-scale folding). 

Examination of published geophysical maps and processing and analysis of digital geophysical data for the western half 

of the Pueblo, Colorado, 1 °X2° Quadrangle indicate that the 

Arkansas River and its topographic lineament spatially correspond approximately to a northeast-trending boundary 

between contrasting crustal blocks of the southern Front Range 

to the north and the Wet Mountains and DeWeese Plateau to 

the south. Exposed Cambrian alkalic plutonic rocks and associated carbonatites, which are restricted to the Wet Mountains-DeWeese Plateau block, and granodiorites account for 

some of the observed geophysical contrasts; however, a geophysical contrast in the vicinity of the Arkansas River topographic lineament is difficult to explain based on exposed rock 

types. 

Possible surface expression of the block boundary 

includes northeast-trending geologic features within an anomalous zone extending from the Pikes Peak batholith southwestward to the Sangre de Cristo Mountains. The dimensions of this 

zone are controlled by an anomalous northeast-trending zone 

of northeast- trending faults. Within the zone a northeast-trending contact between granodiorite of the Routt Plutonic Suite 

and metamorphic rocks follows within 2 km of the Arkansas 

River for much of the distance between Coaldale and Parkdale, 

Colorado. 

1

U.S. Geological Survey, Box 25046, MS964, Denver, Colorado 

80225. 

The northeast-trending course of the Arkansas River and 

its topographic lineament are interpreted as being loosely controlled by the northeast-trending zone of faults and the Precambrian lithologic contact that reflects the boundary between the 

southern Front Range and Wet Mountains-DeWeese Plateau 

crustal blocks. Extensive vertical and lateral erosion of the 

Arkansas River canyon during late Pliocene and Pleistocene 

time has removed any physical evidence of the topography of 

the exhumed late Eocene erosion surface upon which the original course of the river was established; however, some combination of differential uplift or erosion along the northeasttrending zone of faults and the granodiorite-metamorphic contact could have influenced establishment of the linear northeast trend of the river before downcutting. Indirectly, then, the 

modern course of the Arkansas River between Coaldale and 

Parkdale and its associated topographic lineament reflect the 

northeast-trending boundary between the southern Front 

Range and Wet Mountains-DeWeese Plateau crustal blocks. 

Understanding structural controls in the vicinity of this 

northeast-trending boundary is important because the boundary is near mineral deposits that are conspicuously outside the 

Colorado Mineral Belt. 

INTRODUCTION 

The Arkansas River in south-central Colorado includes a 

northeast-trending segment, approximately 40 km long, 

between Coaldale and Parkdale. At a regional scale the topographic lineament associated with the river extends southwest to Hayden Pass in the Sangre de Cristo Mountains. 

Hayden Pass is the lowest point (3,567 m) along the crest of 

the rugged Sangre de Cristo Mountains between Methodist 

Mountain 32 km to the northwest and Mendano Pass 72 km 

to the southeast (fig. lA). In addition, Hayden Pass marks a 

conspicuous northwestward narrowing of the Sangre de 

Cristo Mountains (fig. lB). On digital terrain data of the 

State (fig. 2A) a somewhat longer (60 km) topographic lineament is apparent. 

The Arkansas River between Salida and Canon City, 

Colorado (fig. 2B), has been interpreted as a superposed 

Influence of Crustal Structure on Course of Arkansas River, Colorado 01 
A 10 KILOMETERS 

Figure 1 (above and facing page). Study area in south-central Colorado. A, Three-dimensional view of the Sangre de 

Cristo Mountains (looking northeast) showing Hayden Pass and Arkansas River drainage. Vertical exaggeration about 

6:1 . Data from 1:250,000 topographic map of Pueblo 1 °X2° Quadrangle. 8, Plan view of area shown in A 

drainage, although the covering rocks from which it was 

superposed are not preserved (Thornbury, 1965, p. 349). The 

present course of the river was probably established in late 

Oligocene or early Miocene time on the extensive exhumed 

late Eocene erosion surface of low relief in south-central 

Colorado (Epis and Chapin, 1975). Paleochannels containing gravel deposits and remnants of Oligocene ash-flow tuffs 

are preserved on the late Eocene erosion surface immediately north and south of the present Arkansas River canyon, 

but extensive lateral and vertical erosion by the Arkansas 

River and its tributaries have removed any physical evidence 

of the early history of the Arkansas River itself. 

Steeply dipping Precambrian metamorphic and igneous 

rocks exposed in and along the Arkansas River canyon, 

although complexly deformed, do not indicate that a geologic structure (shear zone?) controls the exceptionally 

straight course of the river;' indeed, the regional metamorphic foliation is at a high acute angle to theN. 60° E. trend 

of the river, such that the canyon exposes an extensive cross 

section of these steeply dipping rocks. 

This segment of the Arkansas River is in an area established as an airborne geophysics demonstration area for testing new geophysical instrumentation and methods and 

extending integrated interpretation techniques (Watson and 

others, 1989). The Geophysics Environmental and Minerals 

(GEM) demonstration area (fig. 2B) includes most of the 

western half of the Pueblo 1 °X2° Quadrangle. Several digital 

data sets have been compiled for the GEM area, including 

gravity, aeromagnetic, radiometric, geologic, remote sensing 

(Landsat Thematic Mapper; Heat Capacity Mapping Mission), and digital elevation model terrain data. Numerous 

other geologic and geophysical maps provide additional 

local and regional information. 

This report summarizes the results of a preliminary 

investigation to determine whether available data might provide some insights to the nature and cause of the linearity of 

the Arkansas River between Coaldale and Parkdale, Colorado. These results are preliminary but provide some observations that suggest possible structural and lithologic 

controls of the river, as well as broad-scale crust-mantle 

influence on the geologic evolution in this area of the Front 

Range-VIet Mountains junction. 

Acknowledgment.-The authors would like to acknowledge the assistance of Adel Zohdy in providing illustrations 

of the three-dimensional and plan views of the Sangre de 

Cristo Mountains. 

GENERAL GEOLOGY 

The GEM area is roughly centered on Canon City, Colorado, at the junction between the southern Front Range and 

the northern Wet Mountains (fig. 2B). These Precambriancored uplifts are flanked by a well-exposed section of Phanerozoic sedimentary strata deformed by diverse types of 

mountain-flank structures. Precambrian rocks include a variety of Proterozoic igneous and metamorphic rocks. In the 

northern Wet Mountains, Precambrian crystalline rocks 

were intruded by Cambrian alkalic and carbonatite rocks 

prior to deposition of the Phanerozoic sedimentary cover 

rocks. Oligocene volcanic rocks in various parts of the study 

area are related to local eruptive centers or are distal deposits 

D2 Application of Structural Geology to Mineral and Energy Resources 
in early Tertiary paleochannels. Miocene and Pliocene sediments and sedimentary rocks occupy several north- to northwest-trending grabens formed during the evolution of the 

Rio Grande Rift Zone, and gravels locally lie on pediment 

surfaces and in terraces formed during various stages of the 

Pleistocene. 

The largest and most important mineral deposit in the 

GEM area is the world-class gold-telluride deposit associated with hydrothermally altered Miocene alkalic volcanic 

rocks at Cripple Creek, Colorado (Lovering and Goddard, 

1950, p. 289). Smaller precious- and base-metal deposits are 

present at several Oligocene volcanic centers in the region. 

Lead, zinc, and copper sulfide ores and minor gold and silver 

have been reported in Precambrian rocks at various places; 

copper sulfide ore was mined from Pennsylvanian redbeds at 

ELEVATION 

lJN FEET) 

7,000 

8,000 

9,000 

10,000 

11,000 

12,000 

13,000 

Cotopaxi (Vanderwilt, 1947, p. 83-87). Small amounts of 

placer gold have been produced from the Arkansas River 

between Salida and Florence. Iron and small quantities of 

rare-earth elements are associated with Cambrian alkalic and 

carbonatite rocks in the northern Wet Mountains and 

DeWeese Plateau (Shawe and Parker, 1967; Armbrustmacher, 1988), and a small uranium deposit is in sediments 

deposited at the base of a paleovalley at Tallahassee Creek. 

Precambrian pegmatites in the area have been mined for 

beryl, columbite-tantalite, muscovite, and feldspar, and gypsum and fire clay have been produced in the area (Vanderwilt, 1947). Cretaceous coals have been mined south of 

Canon City, and a producing oil field is located near 

Florence. Because of normal faulting during the evolution of 

the Rio Grande Rift Zone, the area may have undiscovered 

Influence of Crustal Structure on Course of Arkansas River, Colorado 03 
Figure 2 (above and facing page). Digital terrain image (A), illumination from northwest, and location map 

(8) showing features mentioned in text and boundaries of GEM study area, Colorado. 

mineral deposits on downdropped blocks that are now concealed by upper Tertiary sedimentary rocks and Quaternary 

alluvial materials. 

NEAR-SURFACE EXPRESSION 

The courses of modem drainages are commonly controlled by geologic phenomena such as geomorphic features 

reflecting differential erosion along structures or lithologic 

units of varying resistance and active faulting and uplift. 

Detailed and reconnaissance geologic maps ( 1 :62,500 and 

1:250,000 scale) of the southern Front Range, South Park, 

Wet Mountains, and DeWeese Plateau were examined to 

determine if geologic features exposed at the surface could 

be related to the trend of the Arkansas River between 

Coaldale and Parkdale and the associated topographic lineament. Both the pattern of faulting and the distribution of Precambrian granodiorite and metamorphic rocks in the vicinity 

of the Arkansas River suggest possible surficial controls on 

the course of the river. 

Detailed and reconnaissance geologic mapping shows 

that the present course of the Arkansas River between 

Coaldale and Parkdale does not follow a fault or shear zone 

(Scott and others, 1978; Taylor, Scott, Wobus, and Epis, 

1975a); however, the pattern of faults in the area suggests 

that faulting may have played a major part in establishing the 

straight course of the river. The major faults in the Wet 

Mountains-DeWeese Plateau block trend northwesterly 

northward to the vicinity of Hillside and the Cambrian 

McClure Mountain Complex, an alkalic complex, where a 

northeast-trending zone of northeast-trending faults subparallel with the Arkansas River signals a rather abrupt change 

in structural style (fig. 3). The northeast-trending fault zone 

extends northeastward across the Canon City embayment 

and the southern Front Range to the Pikes Peak batholith. 

Rocks of Precambrian, Paleozoic, and Mesozoic age are cut 

by the northeast-trending faults; lower Eocene sediments of 

the Echo Park Alluvium, rocks of the Oligocene Gribbles 

Park Tuff(Taylor, Scott, and Wobus, 1975), and sodic granite ofthe Miocene Slide Rock Mountain stock (Taylor, Scott, 

and Wobus, 1975) are the youngest rocks cut by northeasttrending faults in the zone. 

Prior to Neogene movements on the Ilse fault and the 

uplift of the Wet Mountains and southern Front Range 

blocks, Oligocene to Miocene low-gradient streams flowed 

generally eastward across an exhumed, relatively flat Eocene 

erosion surface (Taylor, 1975). Most of the faults having 

demonstrated Neogene movement trend north or northwesterly (Epis and Chapin, 1975, fig. 19; Taylor, 1975, fig. 1); 

04 Application of Structural Geology to Mineral and Energy Resources 
109° 

41°~----------

1 

I 

I 

__ WYO~I~~~B_B.~------~ a COLORADO I 

0 DENVER 

I 

\ 

I 

I 

L 

100 200 KILOMETERS 

I B I 

37°L-------------------- COL_9BADO --------------r------' 

NEW MEXICO 

the northeast-trending faults in the northeast-trending fault 

zone may have also had minor movement in the Neogene, 

although they were probably in existence prior to the major 

Neogene tectonism that segmented the paleochannels. The 

northeasterly trend of Oligocene to Miocene paleochannels 

identified within or near the northeast-trending fault zone 

(Goat Creek-Hillside and Oak Creek, Taylor, Scott, Wobus, 

and Epis, 1975a; Tanner Peak, Taylor, Scott, and Wobus, 

1975; Howard, Taylor, Scott, and Wobus, 1975) suggests 

either a uniform gradient direction on the late Eocene erosion surface or a northeast-trending grain to the subtle topography on the erosion surface. The presence of so many 

northeast-trending faults within this area suggests the possibility that low ridge and valley topography, perhaps reflecting differential erosion or movement along the abundant 

northeast-trending faults, may have influenced the course of 

the paleovalleys. The Arkansas River, now deeply incised 

through the late Eocene erosion surface, may have originated 

as another paleochannel before uplift and deep incision in 

the Pliocene and Pleistocene; however, subsequent downcutting and lateral erosion by the Arkansas River and its tributaries have removed any evidence of the original shape and 

trend of the early Arkansas River drainage channel and any 

deposits that may have accumulated in the channel. 

Two major Precambrian rock types are exposed in the 

area of the southern Front Range and northern Wet Mountains and De Weese Plateau-metamorphic rocks and granodiorite (Scott and others, 1978; Tweto, 1979). The oldest and 

most widespread rocks are quartzite, biotite schist, gneiss 

and migmatite, and other metavolcanic and metasedimentary 

rocks older than about 1,700 Ma (fig. 4). These rocks were 

intruded by abundant granodioritic magmas at catazonal to 

mesozonal depths (Hutchinson, 1976) about 1,700 Ma 

(Routt Plutonic Suite of Tweto (1987)). 

Smaller granitic and quartz monzonitic plutons emplaced 

about 1,400 Ma (Berthoud Plutonic Suite age) at mesozonal 

depths are exposed in a few areas of the southern Front 

Range-northern Wet Mountains area. Granitic rocks of the 

extensive Pikes Peak batholith (about 1,000 Ma) make up 

much of the southern Front Range north of the study area. 

Along the 35-km-long, northeast-trending segment of the 

Arkansas River between Coaldale and Parkdale, Colorado, 

granodioritic rocks of Routt Plutonic Suite age to the north 

are in contact with older Proterozoic metasedimentary and 

metavolcanic rocks to the south (Taylor, Scott, Wobus, and 

Epis, 1975a). The regional trend of this contact is essentially 

parallel with the Arkansas River and for 25 km along this 

stretch of the river the contact is within 2 km of the river and 

commonly within 1 km. This spatial relationship strongly 

suggests that the granite-metamorphic contact may have 

influenced the ultimate course of the river, although the 

nature of this influence is unknown. 

Influence of Crustal Structure on Course of Arkansas River, Colorado 05 
\ 

20 40 KILOMETERS 

Figure 3. Northeast-trending zone of northeast-trending faults (dashed where inferred) (after Tweto, 1979) in study area, southcentral Colorado. 

CRUSTAL EXPRESSION 

Because geologic maps of the southern Front Range-Wet 

Mountains/DeWeese Plateau areas (Taylor, Scott, and 

Wobus, 1975; Taylor, Scott, Wobus, and Epis, 1975a, b; 

Scott and others, 1978) clearly show that the course of the 

Arkansas River between Coaldale and Parkdale is not controlled by a fault or shear zone, potential-field data were 

examined to see if there are any magnetic or gravity characteristics in the area that might suggest a possible controlling 

influence. 

Aeromagnetic Data 

Residual total-field magnetic data (1-mi and some 3-mi 

line spacing, acquired 400 ft above mean terrain) were digitally processed for display in image format at 1:250,000 

scale to examine geologic structure. The data were then 

reduced to the pole using a 0.25-km grid size (fig. 5A) to 

remove inclination errors and place anomalies above their 

sources. Two distinct areas (areas I and II) having similar 

high spatial frequency pattern signatures were identified. 

These two areas are characterized by anomalies having max06 Application of Structural Geology to Mineral and Energy Resources 
0 Hayden Pass 

10 15 KI LOMETERS 

EXPLANATION 

8 Phanerozoic rocks 

Berthoud Plutonic Suite granitic rocks 

' 

Routt Plutonic Suite granodiorite 

Precambrian metamorphic rocks 

Figure 4. Generalized lithology of Proterozoic rocks along Arkansas River between Coaldale and Parkdale, south-central Colorado. Modified from Scott and others (1978). 

imum dimensions of 8 km or less (Boler and Klein, 1990a), 

although the eastern boundary of area I and the southern 

boundary of area II differ somewhat from the Front Range 

area as mapped by Boler and Klein (1990a). 

Area I, covering the Wet Mountains and DeWeese Plateau, corresponds at the surface to predominantly migmatitic 

gneiss older than Routt Plutonic Suite. Most of the magnetic 

anomalies have been interpreted to be caused by a Cambrian 

alkalic intrusive complex having mafic derivatives (Boler 

and Klein, 1990a). Anomalies on the northeastern side of 

area I, along the crest of the Wet Mountains, are caused by 

Precambrian granodiorite (Boler and Klein, 1990a). The 

boundaries of area I generally coincide with the Wet Mountains Fault on the east, the WestcliffFault on the west, and a 

zone of northeast-trending faults near the Arkansas River 

canyon on the north. The cause of the northern truncation of 

area I is not readily apparent. Rocks to the north of area I, in 

a magnetic quiet zone, are mostly Precambrian granodiorite 

(Routt Plutonic Suite) and are the source of magnetic anomaly highs in both areas I and II. Because the zone is similar 

in outcrop pattern, elevation, and topography to magnetic 

rocks of the DeWeese Plateau and because there is no evidence for alteration, Boler and Klein were unable to identify 

the cause of this magnetically quiet area. To the north a 

series of magnetic highs are associated with Precambrian 

basement highs covered by Tertiary sedimentary and volcanic rocks (Boler and Klein, 1990a). The quiet area, which 

contains a number of northeast-trending faults, also includes 

the northeast-trending segment of the Arkansas River 

described previously. 

The second area of magnetic anomalies (area II, fig. SA) 

is northeast of area I. The surface rocks are predominantly 

granodiorite of the Routt Plutonic Suite. Area II is bounded 

on the east by the Pikes Peak batholith (figs. 2B, SA), on the 

southeast by Paleozoic sediments, and on the west by the 

Elkhorn-Current Creek-lise Fault Zone. Area II corresponds approximately to Boler and Klein's (1990a) feature 

H. Although volcanic cover rocks are present west of area II, 

their influence on the boundary is not evident. The sources 

of the magnetic anomalies in area II are Precambrian 

Influence of Crustal Structure on Course of Arkansas River, Colorado 07 
granodiorite (Routt Plutonic Suite) and Precambrian quartz 

monzonite (Berthoud Plutonic Suite). To the east a pronounced magnetic low is associated with Precambrian Pikes 

Peak Granite (1,000 Ma); this association is based on the low 

magnetization of the granite, not on reverse magnetization 

(Boler and Klein, 1990a). 

106° 105° 

39° 

Figure 5 (above and facing column). Magnetic data for 

western half of Pueblo, Colorado, 1 °X2° Quadrangle. A, Image of reduced-to-pole magnetic data. Original grid was 

produced using minimum curvature (Briggs, 1974) with elliptical weights to enhance N. 45° W. direction (M.W. Webring, U.S. Geological Survey, 1990, oral commun.). Areas 

of magnetic highs are bright on image. B, Image of horizontal gradient of reduced-to-pole magnetic data (upward continued 200 m). Aeromagnetic highs are bright and 

correspond to edges of bodies. C, Map showing magnetic 

signature patterns (see text) (dotted line patterns) and areas 

(1,11) of magnetic anomalies. 

To examine the boundaries of areas I and II more fully, 

the horizontal gradient of the reduced to the pole field was 

computed and upward continued 200 m to smooth highfrequency artifacts (fig. 5B). In this representation, magnetic 

gradient highs are associated with the edges of bodies thus 

making it easier to infer subtle trends and structural alignments. Two prominent zones are evident. One, in the northcentral part of the image, corresponds to the edge of the Pike 

Peak batholith. The other, a north-trending magnetic gradient high at approximately long 105.5° W., corresponds to the 

western edges of both areas I and II. From south to north this 

high coincides with the northwest-trending Westcliffe Fault, 

passes through a regional grain of northwest-trending faults 

(locally coinciding with numerous northeast-trending 

faults), then to the Current Creek-lise Fault Zone, and then 

north-northwest along the fault zone to the Guffey volcanic 

center. 

Although the northern boundary of area I and the southem boundary of area II suggest some spatial coincidence 

D8 Application of Structural Geology to Mineral and Energy Resources 
Figure 6. Gravity map of western half of Pueblo 1 °X2° 

Quadrangle. Areas of gravity highs are light on image. 

Map area is same as figure 5. 

with the straight segment of the Arkansas River drainage, the 

reasons are not directly evident. Why are the granodiorites of 

the quiet zone less magnetic? Hydrothermal alteration could 

result in destruction of magnetite, but Boler and Klein 

(1990a) found no evidence for alteration. One hypothesis is 

that the magma source here contained less magnetite than 

did sources to the south and to the east. Another explanation 

is that during the cooling of this batholith reverse magnetism 

could have been captured and this remanence would have 

reduced the total magnetic field. A last hypothesis is that the 

granodiorite in the quiet zone is much thinner than elsewhere, thereby reducing the mass of magnetic minerals and 

thus the total field. Field sampling and extensive drilling are 

required to establish the correct cause with a high confidence 

level. At this time we are left with fairly strong evidence that 

anomalies along the boundary between area I and the quiet 

zone are fault controlled, that Cambrian intrusive rocks are 

not at the surface north of area I, and that possible explanations of the magnetic quiet zone are lithologic or tectonic differences, or possibly both. 

Gravity Data 

Bouguer gravity data for the Pueblo Quadrangle (Boler 

and Klein, 1990b) were examined in the vicinity of the 

straight segment of the Arkansas River between Coaldale 

and Parkdale (fig. 6). Area I identified on the magnetic data 

is a gravity high, the boundaries of which closely approximate the -220-mGal contour and the steepest gradient. Adjacent to and east of area I a northwest-trending gravity low is 

associated with the Florence basin that cuts across the southwestern end of area II. The northeastern boundary of area II 

is controlled by the Pikes Peak batholith, which appears as a 

weak, but very broad, gravity high. The western edge of area 

II corresponds to a gravity gradient (roughly along the -237-

mGal contour). Southeast of Pikes Peak a broad northeast 

trend to the gravity contours parallels the southeast boundary 

of area II due to variations in crustal thickness. These contours strike northeast along a 100-km-long segment before 

swinging to the north parallel with the front of the Front 

Range. 

The gravity data confirm that the areas identified on the 

basis of the magnetic data correspond to crustal differences, 

but they provide no additional evidence as to the nature of 

the boundary between the Wet Mountains-DeWeese Plateau 

and southern Front Range blocks. 

DISCUSSION 

The zone of northeast-trending faults and the northeasttrending granodiorite-metamorphic contact are possible elements of geologic control of the location and trend of the 

Arkansas River between Coaldale and Parkdale, Colorado. It 

is clear, however, that these features are not related to a simple shear zone such as the shear zones ancestral to the Colorado Mineral Belt to the north in central Colorado (Tweto 

and Sims, 1963; Warner, 1978). Aeromagnetic and gravity 

data, as well as the restricted distribution of Cambrian alkalic 

rocks, indicate that the anomalous northeast-trending zone, 

including the northeast-trending faults, the granodioritemetamorphic contact, and the Arkansas River topographic 

lineament, may be the relatively subtle expression of a 

deeper seated boundary between contrasting blocks of 

Proterozoic crust. 

The Proterozoic crust of Colorado and northern New 

Mexico was accreted to the southern margin of the Archean 

craton in Wyoming between 1,790 and 1,660 Ma (Reed and 

others, 1987). The suture between Archean and Proterozoic 

crust is well marked by the northeast-trending Proterozoic 

Cheyenne Belt of Proterozoic crustal blocks bounded by 

shear zones (Duebendorfer and Houston, 1987) in southern 

Wyoming (fig. 7). Southward into Colorado, numerous 

northeast-trending shear zones and faults of Early Proterozoic ancestry mark conspicuous changes in regional structure (Tweto and Sims, 1963; Reed and others, 1987) and 

controlled the emplacement of Laramide granitic intrusive 

bodies of the Colorado Mineral Belt (Tweto and Sims, 

1963). Bowring and Karlstrom (1990) suggested that these 

structures may reflect the boundaries between relatively thin 

lithospheric fragments assembled into the Proterozoic crust 

that maintained their identity through subsequent orogenic 

Influence of Crustal Structure on Course of Arkansas River, Colorado 09 
COLORADO 

PROVINCE 

CR 

BN 

northeast-trending 

faults 

I 

I 

l 

I 

I 

---_WYOM~.Q I NEB!!_A_§KA ~~ ---- 20 

COLORADO 1 1 

~ I 

~ I 

0 DENVER 

I 

I 

l 

I 

r I 

37oL _____ l_ ____ _j _____ L COLQ.R~DO l ___ __ _l _____ --r _____ j 

NEW MEXICO I 

100 200 KILOMETERS 

Figure 7. Map showing structural features, mineral deposits, and magnetic anomal ies, Colorado and southeastern Wyoming. Double-wavy lines indicate shear zones: H, Homestake; BP, Berthoud Pass; IS, Idaho Springs. Mineral deposits: CC, 

Cripple Creek; BN, Bonanza; CR, Creede. Dotted lines indicate magnetic anomalies. Compiled from Reed and others 

(1987) and Duebendorfer and Houston (1987). 

plutonism (Routt Plutonic Suite, 1.75-1.6 Ga) and anorogenic plutonism and differential uplift and erosion (Berthoud 

Plutonic Suite and Pikes Peak Granite, 1.4-1.1 Ga) and were 

reactivated repeatedly through Tertiary time. It is likely that 

such structures would be variously modified by metamorphic and plutonic activity but might still maintain a subtle 

influence on subsequent tectonic events. 

Geophysical expression of the Proterozoic crust in southcentral Colorado suggests a boundary between contrasting 

blocks in the vicinity of the Arkansas River topographic lineament. Possible surface expression of this boundary is limited to a northeast-trending zone of northeast-trending faults, 

a northeast-trending contact between granodiorite of the 

Routt Plutonic Suite and older metamorphic rocks, and the 

010 Application of Structural Geology to Mineral and Energy Resources 
Arkansas River topographic lineament itself. If this boundary originated during the Early Proterozoic assembly of heterogeneous crustal blocks, it was either irregular and 

complex or it has been greatly modified by subsequent plutonism and metamorphism. Nevertheless, this crustal boundary may account for the en echelon pattern of uplift during 

orogenesis in late Paleozoic (Ancestral Rockies) and Late 

Cretaceous-Eocene (Laramide) times, as well as during 

uplift of the Front Range and Wet Mountains in Neogene 

time (fig. 7). 

Based on observations over an admittedly limited geographic area, the modem course of the Arkansas River 

between Coaldale and Parkdale, Colorado, and the associated Arkansas River topographic lineament probably are 

indirectly controlled by geologic features expressing a northeast-trending boundary between contrasting Proterozoic 

crustal blocks. This interpretation is preliminary at best, and 

additional geological and geophysical studies are necessary 

to understand the nature of the block boundary, particularly 

its lateral extent, and the effects it may have had on Precambrian and Phanerozoic events including the formation of 

mineral deposits. The location of several major mineral 

deposits (Cripple Creek, Bonanza, Creede) conspicuously 

outside the Colorado Mineral Belt (fig. 7) provides additional motivation for understanding crustal structure in this 

region. 

REFERENCES CITED 

Annbrustmacher, T.J., 1988, Geology and resources of thorium and 

associated elements in the Wet Mountains area, Fremont and 

Custer Counties, Colorado: U.S. Geological Survey Professional Paper 1049-F, 34 p. 

Briggs, I.C., 1974, Machine contouring using minimum curvature: 

Geophysics, v. 39, p. 39--48. 

Boler, F.M., and Klein, D.P., 1990a, Residual aeromagnetic and 

interpretative maps of the Pueblo 1 °X2° Quadrangle, southcentral Colorado: U.S. Geological Survey Miscellaneous 

Investigations Series Map 1-1419-C, scale 1:250,000. 

___ 1990b, Residual Bouguer gravity and interpretative maps of 

the Pueblo 1 °X2° Quadrangle, south-central Colorado: U.S. 

Geological Survey Miscellaneous Investigations Series Map 

1-1419-B, scale 1:250,000. 

Bowring, S.A., and Karlstrom, K.E., 1990, Growth, stabilization, 

and reactivation of Proterozoic lithosphere in the southwestem United States: Geology, v. 8, p. 1203-1206. 

Duebendorfer, E.M., and Houston, R.S., 1987, Proterozoic accretionary tectonics at the southern margin of the Archean Wyoming craton: Geological Society of America Bulletin, v. 98, 

p. 554--568. 

Epis, R.C., and Chapin, C.E., 1975, Geomorphic and tectonic 

implications of the post-Laramide, Late Eocene erosion surface in the southern Rocky Mountains: Geological Society of 

America Memoir 144, p. 45-74. 

Hutchinson, R.M., 1976, Precambrian geochronology of western 

and central Colorado and southern Wyoming, in Epis, R.C., 

and Weimer, R.J., eds., Studies in Colorado field geology: 

Professional Contributions of Colorado School of Mines, no. 

8, p. 73-77. 

Lovering, T.S., and Goddard, E.N., 1950, Geology and ore deposits 

of the Front Range, Colorado: U.S. Geological Survey Professional Paper 223, 319 p. 

Reed, J.C., Jr., Bickford, M.E., Premo, W.R., Aleinikoff, J.N., and 

Pallister, J.S., 1987. Evolution of the early Proterozoic Colorado province-Constraints from U-Pb geochronology: 

Geology, v. 15, p. 861-865. 

Scott, G.R., Taylor, R.B., Epis, R.C., and Wobus, R.A., 1978, Geologic map of the Pueblo 1 °X2° Quadrangle, south-central 

Colorado: U.S. Geological Survey Miscellaneous Investigations Map 1-1022, scale 1:250,000. 

Shawe, D.R. and Parker, R.L., 1967, Mafic-ultramafic layered 

intrusion at Iron Mountain, Fremont County, Colorado: 

U.S. Geological Survey Bulletin 1251-A, p. A1-A28. 

Taylor, R.B., 1975, Neogene tectonism in south-central Colorado: 

Geological Society of America Memoir 144, p. 211-226. 

Taylor, R.B.. Scott, G.R., and Wobus, R.A., 1975, Reconnaissance 

geologic map of the Howard Quadrangle, central Colorado: 

U.S. Geological Survey Miscellaneous Investigations Map 

1-892, scale 1 :62,500. 

Taylor, R.B., Scott, G.R., Wobus, R.A., and Epis, R.C., 1975a, 

Reconnaissance geologic map of the Cotopaxi 15-minute 

Quadrangle, Fremont and Custer Counties, Colorado: U.S. 

Geological Survey Miscellaneous Investigations Map 1-900, 

scale 1:62,500. 

___ 1975b, Reconnaissance geologic map of the Royal Gorge 

15-Minute Quadrangle, Fremont and Custer Counties, Colorado: U.S. Geological Survey Miscellaneous Investigations 

Map 1-869, scale 1:62,500. 

Thornbury, W.D., 1965. Regional geomorphology of the United 

States: New York, John Wiley and Sons, 609 p. 

Tweto, Ogden, 1979, Geologic map of Colorado: U.S. Geological 

Survey, scale 1:500,000. 

___ 1987, Rock units of the Precambrian basement of Colorado: 

U.S. Geological Survey Professional Paper 1321-A, 54 p. 

Tweto, Ogden, and Sims, P.K., 1963, Precambrian ancestry of the 

Colorado mineral belt: Geological Society of America Bulletin, v. 74, no. 8, p. 991-1014. 

Vanderwilt, J.W., 1947, Mineral resources of Colorado: State of 

Colorado, Mineral Resources Board, 547 p. 

Warner, L.A., 1978, The Colorado lineament-A middle Precambrian wrench fault system: Geological Society of America 

Bulletin, v. 89, p.161-171. 

Watson, Ken, Knepper, D.H., Jr., Clark, R.N., Pitkin, J.A., Labson 

Vic, Campbell, D.L., Rowan, L.C., Kruse, F.A., Lee. 

Keenan, and Livo, K.E., 1989. Plans for integrated airborne 

geophysical study of the Geophysics Environmental and 

Minerals demonstration area, south-central Colorado, in 

USGS Research on Mineral Resources-1989 Program and 

Abstracts, Fifth Annual V.E. McKelvey Forum on Mineral 

and Energy Resources: U.S. Geological Survey Circular 

1035, p. 77-78. 

Influence of Crustal Structure on Course of Arkansas River, Colorado 011 

Chapter E 

Implications of Gravity and Seismic Reflection 

Data for Laramide Mountain Ranges and 

Basins of Wyoming and Southern Montana 

By STEPHEN L. ROBBINS and JOHN A. GROW 

U.S. GEOLOGICAL SURVEY BULLETIN 2012 

APPLICATIONS OF STRUCTURAL GEOLOGY TO MINERAL AND 

ENERGY RESOURCES OF THE CENTRAL AND WESTERN UNITED STATES 
CONTENTS 

Abstract El 

Introduction El 

Geologic setting E2 

Gravity data E2 

Previous geophysical studies E3 

Wind River Mountains-Green River Basin E3 

Owl Creek Mountains/Casper Arch-Wind River Basin-Granite Mountains-Great 

Divide Basin E4 

Beartooth Mountains-Bighorn Basin E4 

Bighorn Mountains-Powder River Basin-Black Hills 

Other geophysical studies supporting low-angle thrusting 

ElO 

ElO 

Horizontal gravity gradient El 0 

Isostatic compensation El3 

Rock density El5 

Structural synthesis El6 

Conclusions El7 

References cited El8 

PLATES 

{Plates are in pocket] 

1-5. Maps of Wyoming and southern Montana showing: 

1. Topography and lines of geophysical profiles used in study. 

2. Structure of the top of the Precambrian basement. 

3. Complete Bouguer gravity anomaly. 

4. Isostatic residual anomaly gravity. 

5. Horizontal gradient anomaly gravity. 

FIGURES 

1. Seismic line across Wind River Thrust E4 

2. Deep-crustal seismic-reflection line across Wind River Mountains E5 

3, 4. Gravity profiles across: 

3. Wind River Mountains E5 

4. Wind River Basin E6 

5. Interpreted cross section across Casper Arch Thrust E8 

6. Gravity profile across Beartooth Mountain front E9 

7. Seismic line across east flank of Bighorn Mountains 

8. Gravity and magnetic profiles across Powder River Basin 

9-11. Diagrams showing: 

Ell 

9. Gravity anomaly-fault dip comparison model El4 

El2 

10. Airy-Heiskanen isostatic compensation model across Wyoming El5 

11. Sediment density contrast model across Wind River Thrust El6 

12. Interpreted cross section across Wyoming El7 

Ell Contents 
Implications of Gravity and Seismic Reflection Data for 

Laramide Mountain Ranges and Basins of 

Wyoming and Southern Montana 

By Stephen L. Robbins 1 and John A. Grow2 

Abstract 

Gravity profiles and seismic reflection profiles indicate 

low-angle thrusting of several Laramide mountain ranges onto 

basins in Wyoming and southern Montana. Examples include 

thrusting of the Bighorn Mountains northeastward onto the 

western Powder River Basin, the Wind River Mountains southwestward onto the Green River Basin, the Owl Creek Mountains southward onto the Wind River Basin, the Shirley 

Mountains southward onto the Hanna Basin, the Granite 

Mountains southwestward toward the Great Divide Basin, and 

the Beartooth Mountains northeastward onto the Bighorn 

Basin. Interpretations of the profiles contain several similarities: (1) thrust angles of 20°-30°; (2) thrust distances onto the 

basins of 10-20 km; and (3) thrusts that appear to sole in the 

middle or lower crust. 

Isostatic residual anomaly (IRA) values range from -77 

mGal over the Hanna Basin to +75 mGal over the Laramie 

Mountains but average near zero. Large positive IRA values 

(+40 to +70 mGal) over all of the Laramide mountain ranges 

indicate that local roots are absent beneath these mountain 

ranges. IRA lows of -20 to -70 mGal over the deeper basins are 

primarily a reflection of low-density basin-fill deposits. The fact 

that the IRA values average near zero indicates that the longer 

wavelength elevation changes (greater than 1 00 km) must be 

isostatically compensated; that is, the crustal thickness compensates for regional elevation changes but not for local mountain ranges. 

Positive gravity anomalies within some of the basins in 

areas of little structural relief (areas such as the central Powder 

River Basin where some borehole and seismic data are available) suggest that density varies laterally significantly within 

Precambrian basement rocks, probably because of differences 

in basement rock types. Aeromagnetic patterns within these 

basins also support this interpretation. 

1

U.S. Geological Survey, Box 25046, MS939, Denver, Colorado 

80225. 2

U.S. Geological Survey, Box 25046, MS960, Denver, Colorado 

80225. 

A better understanding of the geometry of thrust faults 

that have carried mountain ranges over the edges of adjacent 

sedimentary basins, and of the structural relief within these 

basins, is needed in order to effectively explore for future 

hydrocarbon resources. 

INTRODUCTION 

The Laramide orogeny uplifted Precambrian basement 

blocks throughout the western United States during Late 

Cretaceous and early Cenozoic time (Hamilton, 1988; Perry 

and others, this volume). Vertical uplifts of 13 km and 

crustal shortening of more than 20 krn have been demonstrated in the Wind River Mountains (plates 1, 2) (Smithson 

and others, 1978; Sharry and others, 1986; Blackstone, 

1990). Within northern Wyoming and southern Montana, the 

basement blocks are composed of Archean granite, gneiss, 

amphibolite, and minor mafic dikes (Love and Christensen, 

1986). Cratonic sedimentary rocks of Cambrian through 

Jurassic age unconformably overlie the basement and thin 

from 2I/2 km in western Wyoming (Lamerson, 1982) to less 

than 1 km in northeastern Wyoming (Rocky Mountain Association of Geologists, 1972) with no evidence of any significant pre-Laramide relief on the basement. Although thinskinned sedimentary thrusts formed along the UtahWyoming border area during the Cretaceous Sevier orogeny, 

the Laramide orogeny is the only significant tectonic event 

during Phanerozoic time in northern Wyoming that involved 

major disruptions in the basement. Therefore, northern Wyoming offers a unique opportunity to study the processes 

involved in compressional deformation of the continental 

crust and lithosphere in an area relatively uncomplicated by 

either earlier or later Phanerozoic tectonic events. 

Debate over whether Laramide mountain ranges are bordered by thrust faults or vertical faults goes back almost a 

century (Darton, 1906). Berg (1963) published a seismic line 

over the west flank of the Wind River Mountains that shows 

Gravity and Seismic Reflection Data, Wyoming and Southern Montana E1 
a thrust fault flattening to the northeast. Subsequently, however, Prucha and others (1965) introduced the concept of 

concave-downward (upthrusting) faults for which the dip 

increases downward to near vertical at basement level. V ariations on the vertical fault theme were dominant into the late 

1970's (Matthews, 1978). A combination of multifold seismic reflection profiles, gravity models, and drill holes 

through Precambrian overhangs, especially in the last dozen 

years, confirms that thrust faulting was the dominant mode 

of deformation along Laramide mountain fronts (Berg, 

1963; Berg and Romberg, 1966; Smithson and others, 1978; 

Blackstone, 1981; Gries, 1981, 1983; Bally, 1983; Lowell, 

1983; Skeen and Ray, 1983; Gries and Dyer, 1985). Furthermore, a gravity model and a reprocessed version of the 

COCORP (Consortium for Continental Reflection Profiling) 

seismic profile over the Wind River Mountains indicates that 

the Wind River Thrust flattens within the lower crust at 

depths of about 22 km (Hurich and Smithson, 1982; Sharry 

and others, 1986). 

Gravity data recently collected in the Powder River Basin 

and the Bighorn Mountains of northeastern Wyoming (Robbins and Grow, 1990) and seismic profiles along the east 

flank of the Bighorn Mountains (Grow and others, 1988) 

indicate that low-angle faults (20°-30° W. dip) that flatten 

within the middle or lower crust are also dominant in northeastern Wyoming. Therefore, thrusts that flatten within the 

middle and lower crust probably are dominant across most of 

the Laramide province of Wyoming from the Wind River 

Mountains eastward for a distance of 300 km to the Bighorn 

Mountains. 

In this paper, we present new compilations of complete 

Bouguer gravity anomaly, isostatic gravity anomaly, and 

horizontal gravity gradient maps for Wyoming and southern 

Montana in color at a scale of 1:1,500,000. Although data are 

missing for some parts of this region, these are the first color 

gravity maps at a large enough scale to allow discussion of 

individual Laramide mountain fronts as well as regional 

trends across the entire province. We also review some of the 

previously published key seismic reflection lines and gravity 

models for the area. Finally, we discuss the implications of 

the new gravity maps for tectonic models of the Wyoming 

Laramide province and for applications of these types of 

maps for improved mineral, oil, and gas exploration techniques. 

GEOLOGIC SETTING 

Elevations in southern Montana and Wyoming range 

from about 1,100 m in the northeast to almost 4,200 min the 

Wind River Mountains (plate 1). The Bighorn Mountains 

(4,000 m), Beartooth Mountains (3,800 m), and Medicine 

Bow Mountains (3,700 m) are the other major ranges in the 

area. The intermontane basins have a systematic topographic 

E2 Application of Structural Geology to Mineral and Energy Resources 

gradient from 1,100 m in the high plains of the northeast to 

more than 2,000 m in the Green River Basin of western 

Wyoming. 

Structural relief on the Precambrian basement surface 

extends from more than 9 km below sea level to more than 4 

km above sea level along the northeast flank of the Green 

River Basin and the west flank of the Wind River Mountains 

(plate 2) (Blackstone, 1990). 

Paleozoic cratonic shallow-marine sedimentary rocks 

that lie unconformably on Archean basement rocks are 1 km 

thick in western Wyoming and thin to less than 1 km in 

northeastern Wyoming; there is no evidence for any significant pre-Laramide relief on the basement surface (Rocky 

Mountain Association of Geologists, 1972; Lamerson, 1982; 

Peterson, 1983). Cratonic deposition continued in Triassic 

and Jurassic time; about 1.5 km of rock was deposited in the 

western part of Wyoming and about 0.3 km in the eastern 

part. In the Early Cretaceous, Sevier thrusting along the 

Utah-Wyoming border resulted in debris being shed to the 

east; as much as 1.6 km of rock was deposited in the western 

part of Wyoming and 0.2 km in the eastern part. Amounts of 

synorogenic clastic debris increased rapidly in the Late Cretaceous; this debris was fed into a broad marine seaway that 

connected the Gulf of Mexico with the Arctic Ocean 

(McGookey, 1972). Upper Cretaceous sedimentary rocks 

range from 3.6 km thick in western Wyoming to about 1.5 

km thick in eastern Wyoming except along the flanks of the 

Wind River Mountains, which began to rise in the Late Cretaceous, where Upper Cretaceous rocks are only 1.2 km 

thick (McGookey, 1972, fig. 22). The rapid rise of Laramide 

uplifts resulted in thick lower Cenozoic nonmarine deposits 

of the Fort Union (Paleocene) and Wasatch (Paleocene and 

Eocene) Formations that have combined thicknesses of more 

than 3 km in western Wyoming and as much as 2 km in 

northeastern Wyoming. 

Central and northern Wyoming and southern Montana 

have been subjected to limited basement-involved deformation since Precambrian time, whereas southern Wyoming 

and Colorado were subjected to more intense late Paleozoic 

tectonic activity; we focus our attention in this paper on 

northern Wyoming where Laramide deformation is far more 

significant. 

GRAVITY DATA 

Complete Bouguer anomaly (CBA), isostatic residual 

anomaly (IRA), and horizontal gradient anomaly (HGA) 

gravity maps of Wyoming and southern Montana are presented on plates 3-5; 21,206 data points were used to generate these maps. Most of these data are from U.S. Defense 

Mapping Agency files (St. Louis, Missouri), but 1,386 stations in eastern Wyoming and southeastern Montana were 

collected by us. Principal facts for 1,040 of these gravity 
stations are in Williamson and Robbins (1991) and were 

used together with the Defense Mapping Agency data to 

compile CBA gravity maps of the Sheridan and Gillette 

1 °X2° quadrangles (Robbins, in press; Robbins and Williamson, in press). Principal facts for our 346 stations in the 

Arminto, Newcastle, and Torrington 1 °X 2° quadrangles are 

unpublished. 

All gravity values are based on the IGSN-1971 datum 

(Morelli, 1974), and the data were reduced using the 

GRS-1967 formulas (International Association of Geodesy, 

1971) and an assumed average crustal density of 2.67 g/cm3. 

The terrain effect to a distance 167 km from each station was 

removed using a computer program based on Plouff (1977). 

The terrain effect for the stations established by the authors 

was removed by manually making corrections to a distance 

of 0.59 km and using the computer program from 0.59 to 167 

km. Isostatic corrections were made using a computer program by Simpson and others (1983) assuming an AiryHeiskanen compensation model where the parameters used 

are (1) density ofthe topographic load, 2.72 g/cm3, (2) depth 

of compensation below sea level, 40 km, and (3) density contrast at the crust-mantle boundary, +0.35 g/cm3. The CBA 

and IRA gravity maps were generated by first gridding the 

scattered data using a computer program by We bring (1981) 

and then plotting the gridded data using the program 

COLOR by Godson ( 1980). A 2.5-km grid interval was used 

to plot the CBA and IRA maps (plates 3, 4). The HGA map 

(plate 5) was generated from the IRA grid file using a package of programs called BOUNDARY based on Blakely and 

Simpson (1986). 

The maps are plotted at a scale of 1:1,500,000. Station 

spacing is quite variable, and the blank areas are where the 

station spacing is greater than 10 km. The CBA and IRA 

maps are contoured at 5 mGal, and the HGA map is contoured at 1 mGal/km. 

The CBA values range from -75 mGal in southeastern 

Montana to -285 mGal in the Green River Basin in western 

Wyoming (plate 3). The most dominant feature on the CBA 

map is the regional gradient from northeast to southwest, 

which is a direct reflection of the topographic gradient (plate 

1). The mountain ranges are associated with relative highs 

that are 30-50 mGal above the regional trends and the basins 

with relative lows that are 30-50 mGal below the regional 

trend. 

The IRA values are as much as + 75 mGal in the Laramie 

Mountains, the deepest low is -77 mGal in the Hanna Basin 

(plate 4). The average IRA value of near zero indicates that 

the region is isostatically compensated at longer wavelengths (> 100 km). The basins have low IRA values of from 

-35 mGal in the Powder River Basin to -77 mGal in the 

Hanna Basin; these low values are caused mainly by lowdensity sedimentary fill within the basins. The high IRA values in the Laramie, Bighorn, and Wind River Mountains 

(+75, +65, and +45 mGal, respectively) and other lesser 

mountain ranges are not associated with dense, mafic 

Precambrian basement rocks. This implies that the local 

Airy-Heiskanen compensation model assumed deeper roots 

under the mountain ranges than are actually present. In other 

words, the mountains either do not have local roots or the 

roots are significantly shallower than assumed by the AiryHeiskanen model. 

PREVIOUS GEOPHYSICAL STUDIES 

The first known published gravity study is by Chamberlin (1935), who demonstrated that the isostatic character of 

the region is anomalous. Later, Malahoff and Moberly 

( 1968), in a regional gravity study of the State of Wyoming 

recognized that the mountain blocks within the State were 

rootless. In this section, we briefly discuss the most significant previously published gravity and seismic studies. The 

locations of the studies and appropriate referenceSo'are shown 

on plate 1. 

Wind River Mountains-Green River Basin 

The Wind River Mountains are the most extensively 

studied of any Laramide mountain range, and a careful 

review of key studies there illustrates the evolution of structural concepts during recent decades. Berg (1963) published 

a line-drawing interpretation of a single-fold analog seismic 

profile across the southwest flank of the range beneath which 

he interpreted a thrust fault flattening to the northeast. Berg 

and Romberg (1966) published a gravity model consistent 

with this seismic interpretation. The COCORP deep-crustal 

seismic-reflection line was acquired in 1977 and is the first 

published multifold seismic line across the Wind River 

Mountain front (Smithson and others, 1978). It confirms 

Berg's thrust fault model, and following its publication a 

number of multifold seismic profiles across the Wind River 

Mountains (Lowell, 1983; Basham and Martin, 1985) and 

other ranges (Gries, 1981, 1983; Skeen and Ray, 1983; Gries 

and Dyer, 1985) were published that confirm crustal shortening was dominant over vertical uplift. 

Lowell's (1983) multifold seismic line shows a wedge of 

granite thrust over the Green River Basin with the fault flattening to the northeast and at least 16 km of overhang (fig. 

1 ). Lowell did not give the precise location of this profile, 

but the structure shown is so similar to that shown by Berg 

(1963) that the two are probably very close to each other. 

Smithson and others' (1978) publication of the COCORP 

deep-crustal seismic line across the Wind River Mountains 

is probably the most important single seismic profile in 

advancing evidence for thrusting (30°-35°) in the upper 

crust within the Laramide province. A reprocessed version 

of this COCORP line (Sharry and others, 1986) has significantly improved the section in the middle and lower crust 

and suggests that imbricated thrusts flatten out to the northeast between 20 and 30 km depth with the main thrust at a 

Gravity and Seismic Reflection Data, Wyoming and Southern Montana E3 
SOUTHWEST NORTHEAST 

ol===~==~~~~~~~~~----~-------------------~--~-~~-----------. 

(f) 

0 

6 2 u 

w 

(f) 

z 

4 

0 5 KILOMETERS 

Figure 1. Seismic line across the Wind River Thrust. Reprinted with permission from Lowell (1983, fig. 2, p. 5). Approximate 

location of seismic line shown on plate 1. 

depth of 22 km (fig. 2). This seismic profile is 140 km long 

and shows data to a depth of 40 km. Strong reflectors in the 

southwest were observed to the base of the Paleozoic sedimentary fill in the Green River Basin; the Pinedale Anticline 

is just west of the main thrust fault, and the reflector at 33 km 

depth is interpreted as the Moho. 

A two-dimensional gravity model along the COCORP 

profile (Hurich and Smithson, 1982) suggests that the Wind 

River Thrust flattens out in the lower crust and that the thrust 

uplifted higher density lower crust (2.88 g/cm3) into the middle crust but did not penetrate below the Moho (fig. 3). The 

flattening of the thrust in the middle to lower crust demonstrated in both Sharry and others' (1986) version of the 

COCORP seismic line and the Hurich and Smithson gravity 

model is consistent with the positive isostatic gravity anomalies across the range discussed in the previous section that 

imply the range is rootless. 

Owl Creek Mountains/Casper Arch-Wind 

River Basin-Granite Mountains-Great 

Divide Basin 

Case and Keefer ( 1966) published two gravity models 

that cross the north flank of the Wind River Basin (fig. 4A). 

These show that the fault at this northern boundary is a highangle reverse fault dipping to the north. The eastern of the 

two profiles continues southward across the Granite Mountains and into the Great Divide Basin, where a low-angle 

E4 Application of Structural Geology to Mineral and Energy Resources 

fault dips northward beneath the south flank of the Granite 

Mountains (fig. 4B). Karasa (1976) showed two profiles over 

the Granite Mountains that he also interpreted as being 

underlain by low-angle north-dipping faults. 

Ray and Berg (1985) published a high-quality multifold 

seismic profile over the northeast flank of the Wind River 

Basin and Owl Creek-Casper Arch Thrust, which they interpreted as flattening to the northeast with at least 16 km of 

overhang (fig. 5). 

Although not as spectacular as the Wind River Mountains COCORP deep-crustal seismic line and gravity model, 

gravity models and seismic lines across the Owl Creek 

Mountains, Granite Mountains, and Casper Arch all support 

the low-angle thrust fault model in which faults flatten out 

within the crust. 

Beartooth Mountains-Bighorn Basin 

Bonini and Kinard (1983) showed two alternative gravity 

models over the northeast flank of the Beartooth Mountains 

into the Bighorn Basin (fig. 6). Model 1 is an upthrust model 

similar to that of Prucha and others ( 1965); model 2 is a 30° 

thrust that has approximately 12 km of overhang and is a 

much closer fit to the residual curve than the high-angle 

thrust. Proprietary magneto-telluric and seismic exploration 

methods (Smith, 1987), as well as drilling into the 

Precambrian rocks and an underlying sedimentary section, 

have been employed in this area. 
SOUTHWEST NORTHEAST 

0~----------------~~------------------------------------------------~--------~ 

---

<n10 

1-

a: 

l.U 

~~-~ l.U 

~ 

0 

.....1 

:;2 20 

~ 

:i 

h: 

l.U 

0 30 

VERTICAL EXAGGERATION X 1 

-- - - 40~------~----~-------,------,------,-------,------,------,-------,------,------,--L 

0 10 20 30 40 50 60 70 80 90 100 110 

DISTANCE, IN KILOMETERS 

Figure 2. Interpretation of reprocessed version of COCORP deep-crustal seismic-reflection line across the Wind River Mountains. 

(Sharry and others, 1986) showing imbricate listric thrust fault systems that splay off three decollement zones (01

, 0 2, and 0 3). The 

main thrust appears to flatten onto a decollement (02) between 17 and 26 km depth. The original seismic displays, given in Sharry 

and others (1986), are the best seismic evidence to date for the thrust flattening in the lower crust. Location of line shown on plate 1. 

-120 

(f) -140 .....1 

<{ 

(.9 

::::i -160 ::::! 

~ GREEN RIVER BASIN 

z 

~ -180 0 

.....1 

<{ 

~ 

0 

z 

<{ 

a: 

l.U 

::J 

(.9 

::J 

0 

co 

0 

-200 0 • 

-220 

-240 

-260 

0 20 40 

10 

ffi 0 1====~~~~~·2.~2~5~;~·2~.4~3~~~~~===--­

tu 2.54 ~ s 1 a rf------------==~~:2=;. ===~~~ 

:;2 

z 20 

:r:' 

h: 

~ 30 

0 

0 0 

0 0 

0 0 0 

0 

EXPLANATION 

--Bouguer anomaly 

• Lower crust offset 

0 0 

o o o Lower crust and upper 

mantle offset 

DISTANCE, IN KILOMETERS 

2.73 2.60/ 

~"-::_-------

0 

0 WIND RIVER BASIN 

0 

0 

0 

0 

0 0 

0 0 0 0 

p.35 ;2.65 ;2.43 

2.54 

2.65' 

"" 3.30 ------- 40 ----------MOHO---------~--------------------=-~----

Figure 3. Gravity profile across the Wind River Mountains. Reprinted with permission from Hurich and Smithson (1982, fig. 

4b, p. 1555). Location of gravity profile shown on plate 1. 

Gravity and Seismic Reflection Data, Wyoming and Southern Montana E5 
I'T'I 

0"1 

> ""C 

'2. 

~: iS" 

::::s 

a. 

~ c 

0. 

~ 

2!.. 

~ 0 

~ 

Q 

~ 

~-

i 

Col 

::::s 

Q.. 

I'T'I 

::::s 

~ 

~ 

~ 

"' 

l"'l 

~ 

:1 ~ 0::({) 

::2'-' 

o<:: 

z<.:l ctj 

>-- i--::2' 

~z a:-

<.:l 

-150 

-170 

~ -180 0::(/) 

::2'-' 

o.:: 

z~ 

<::-' -' >-- i--::2' 

>z 

~-

z 

0 

z 

_,.,x---x 

Computed gravity anomaly -...xr ~~~~on d~~~~t~ contrasts ;/ 

,r 

<.:l --=-l~.::.:._"' --X-~f --X--x- OWL 

WIND RIVER I ~_,. ,. Area of anomalous I CREEK 

Gravity anomaly computed for 

density contrasts of -0.3 and 

SOUTHWES~OUNTAINS WIND RIVER .~ t BASIN ~~ a;,sp~,~~~~~d ~~ Mrs 

12.000 pCr Pzr t;:v :::::; grav1ty anomaly 

Boool~~ I TKu 

SEA LEVEL Up=_ 0.3 g/cm 3 l'v\zr 

8000 

16.000 up=-0.15g/cm 3 

24.000 

B 

Position of fault 

approximate; diP unknown 

-0.15 g/cm3 <up1 ) \ )(-x-:;;:. /ry Gravity anomaly computed for 

.--'' / density contrasts of -0.4 x--x--x" / and -0.2 g/cm 3 (up2 ) 

o-..-o-~ ~ 

~ 

~t Subsurface geology is ~.;:: ~ schematic for computa- !"0:: MTS o 

SW ~':"of grav1ty anomal1es Ji ~ 

8000 ~ 

I NORTH 

WIND BASIN Tu OWL CREEK MTS 

Teu l'v\zr TKu 

sr • LfV[l Emigrant Trail y TKu 8 9--~~ 

thrust fault 1 ~~~~;~~~ 4 

fault Zone of 

~ 

24.000 Area of anomalous mass assumed rn computation _ of gravity anomalies 

Zone of 

normal faulting 

pCr 

5 0 5 10 15 MILES 

_______ :.::::-~rusting pCr 
Q 

I» 

::::. 

-< 

I» 

:s 

c. 

)(I 

~· 

3 

;::;· 

~ 

I'D = 

0. 

s· 

:s 

0 

I» 

F 

~ 

~ 

3 

:r 

OQ 

I» 

a. 

~ 

s. 

:r 

I'D 

.. 

:s 

~ Q 

a 

I» 

:s 

I» 

1"1'1 

....... 

1/) 

..J 

< 

~ 

:::i 

= ~ 

z 

~ < 

~ 

0 

z 

< 

~ 

> 

< a: 

~ -230 

-240 

-250 

SEA LEVEL 

[:] 

Post-lower Eocene 

rocks 

Gravity anomaly computed 

for density contrasts of / 

-0.3 and -0.2 g/cm~_," 

b 

_..,-"' 

p; / 

/ 

/ / 

,.a" / 

o...._ _ft' ---o---

Crooks Mtn 

TKu 

A p = -0.2 g/cm3 P.r. 

/ / 

A p = -0.3 g/cm3 Hypothetical thrust f ault 

G 

Lower Eocene and 

younger rocks 

0 5 10 MILES 

EXPLANATION 

G 

Lower Eocene, Paleocene, 

and uppermost Creta· 

ceous (lance Formation) 

rocks 

8 

Mesozoic (pre-Lance) 

rocks 

/ 

Jf 

/ / / 

/ 

/ 

Jj'----

Tu 

pCr 

Tu 

~ _ /Emigrant Trail 

--.:-........._thrust fault 

pCr 

8 

Paleozoic rocks 

0 

Precambrian rocks 

Figure 4. Gravity profiles across the Wind River Basin. Reprinted from Case and Keefer (1966, figs. 4 and 5, p. C125 and C127). Location of gravity profiles 

shown on plate 1. Note that the top of figure is an expanded version of the southern end of profile B. 
ITI 

= 

> "'1:1 

:?.. 

~· 

::::!'. 

Q 

= 

Q 

-

~ '"'I 

c 

0. 

c 

'"'I 

~ 

~ Q 

~ 

Q 

~ ::I 

1'1) 

'"'I 

~ 

Ql 

::I 

Q. 

ITI 

1'1) = 

'"'I 

~ 

~ 

fl 

Q 

c 

'"'I 

t"l 

fl 

A 

8,000 

PROJECTED W.A. MONCRIEF PROJECTED 

18-1 Tepee Flats W.A. MONCRIEF 

SENE 18 3 16-1 Tepee Flats 

PROJECTED . - 7 N·86W NESW 16-37N-86W 

COASTAL OIL & GAS EL. 

5966 Kb~ EL: 5900 Kb 

2-7-36-86 Bullfrog Unit CIG EXPLORATION SENW 7-36-N-86W 1-6-36-86W Bullfrog Unit W.A. MONCRIEF W.A. MONCRIEF 

EL: 6370 Kb ~ SENW 6-36N-86W 27-1 Tepee Flats 14-1 Tepee Flats 

EL: 6203 Kb NWNE 27-37N-86W SWNE 14-37N-86W 

~ -3!t EL: 5665 Kb EL: 5934 Kb 

-3!t -¢- ~ -¢-

A' 

8,000 

WIND RIVER FORMATION 

4,000 4,000 

FRONTIER 

MOWRY-LAKOTA 

JURASSIC-TRIASSIC 

SEA 

LEVEL SEA 

LEVEL 

PALEOZOIC ~~-- ~,_ 1,:_/1-

!.'/'iPREC,:A!\~IaRIAN·_:-~~(--~'~'.:>t~-...';'1~-/, ,._, ___ , 

-4,000 f--

LANCE-MEETEETSE-MESAVERDE -4,000 

L.U f--

L.U 

~ 

z' -8,000 C> 

f= 

~ 

-8,000 6- <{ 

~ -12,000 

f= 

<{ 

~ -12,000 

-16,000 -16,000 

-20,000 -20,000 

-24,000 

0 5 KILOMETERS 

-, - " . . ' - - -24,000 

Figure 5. Interpreted cross section across the Casper Arch Thrust. Kf, Frontier Formation; Kmr-Kml, Mowry Shale to Lakota clastics (informal unit). Reprinted with permission from Ray and Berg (1985, fig. 3, p. 52). Line of cross section shown on plate 1. 
C') 

.. 

Dol 

< 

~· 

Dol 

:s 

Cl. 

~ ;;;· 

3 

;:;· 

tl) 

~ 

0. s· 

:s 

0 

Dol 

F 

~ 

1 

3 

:r 

OQ 

Dol 

8.. 

[ 

r 

:s 

f 

lit 

:s 

Dol 

m 

1.0 

SOUTHWEST 

-160 

-170 

-180 

(/) 

....I 

~ -190 ....I 

....I 

~ 

(/) 

a: 

w 

fw 

:E 

0 

....I 

~ 

z 

z 

0 

~ > w 

....I 

w 

0 

-10 

-20 

0 

+2. 

0 

-2 

-4 

-6 

-8 

-10 

DEGREE 

------- ___ __I!:I~-------~ ~<v<:::> ~<f9Jej 

0 

MODEL 1 

--------------

-·-.- .11_op£L 2 -......:: ......... 

·----..., ~£stouA.L ·,._ 

10 I I I I I I I I I I I I I 

+ ....I 

<( 

:::> 

0 

(i5 

NORTHEAST 

-130 

-140 

w 

a: 

pOL -;No>o\11'1,_ ~~ ~ ~ ~-- ---:..--:..:----

-150 

-160 

MODEL 1 -- - -- _ ..- ----

-- • -·-RE5iou"AL: ·-• -·-• +20 

+10 

7 ..--·--·--

.---.a=-.-30 ~ 0:..:::::::-. ~ ·-~oo~\. 2. 

G-----------------------------------------------------~0 

-20 

30 

I I I 

DISTANCE, IN KILOMETERS 

40 50 I I I I I I I I I I I I 

Roberts I 

D=2.4 to 2.65 glee 

STILLWATER COMPLEX 

D=2.85 glee 

I I I I I I I I 

60 

I I I I 

N 45°30' 

109° 03.2' w ~ 

~ 

Figure 6. Gravity profile across the Beartooth Mountain front. Residual curve is the observed Bouguer gravity minus a third-degree polynomial estimate of the regional 

field. Reprinted with permission from Bonini and Kinard (1983, fig 4, p. 93). Location of gravity profile shown on plate 1. 
Farther to the south, Stone (1985a, b) presented seismic 

profiles (plate 1) on both the east and west flanks of the Bighorn Basin that show a number of west-dipping thrusts 

which offset basement by less than 1 km of displacement on 

individual faults. 

Bighorn Mountains-Powder River 

Basin-Black Hills 

Seismic profiles across the east flank of the Bighorn 

Mountains show that two major west-dipping thrust faults 

are present just west of Buffalo, Wyoming (Grow and others, 

1988). The dominant thrust dips 20°-30° to the west and has 

6-11 km of basement overhang. The lesser fault (Buffalo 

Deep) (Blackstone, 1981) is a few kilometers east of the 

main mountain front and has less than 1 km of offset. A seismic line just west of Buffalo shows both faults clearly (fig. 

7). This line was provided by Chevron USA, Inc., and was 

reprocessed by the U.S. Geological Survey (Miller and others, 1988). This profile shows a fault with 30° west dip and 

6 km of overhang. The Arco Kenney Ranch well located 

near this profile penetrated about 1 km of granite before 

entering sedimentary rocks of the Powder River Basin. A 

second seismic profile from about 20 km north of Buffalo 

shows a fault with 20° west dip and about 11 km of overhang 

(Grow and others, 1988). The Gulf Granite Ridge exploration well, within 1 km of this line, penetrated 1.8 km of granite before entering sedimentary rocks of the Powder River 

Basin. 

A gravity and magnetic profile and models across the 

Powder River Basin (fig. 8) synthesize the structure of the 

Precambrian basement from the Bighorn Mountains to the 

Black Hills (Robbins and Grow, 1990). Constraints on the 

model include the Buffalo seismic line (fig. 7), seismic lines 

in the Powder River Basin by Moore (1985), the structure 

contour map of the basement (plate 2) by Blackstone ( 1990), 

and modeling along the east flank of the Bighorns by Jenkins 

(1986). The normal density of the basement is assumed to be 

2.73 g/cm3, based on measurements ofHurich and Smithson 

(1982). Short-wavelength gravity and magnetic highs and 

lows are observed over the main part of the Powder River 

Basin in contrast to the very gently west dipping basement 

surface. High and low density (2.58-3.01 g/cm3) and magnetic (I=20x10-6-3,200x10-6 emu/cm3) bodies within the 

upper basement must be assumed in order to fit observed 

gravity and magnetic curves over the Powder River Basin. 

This implies that significant lithologic heterogeneity exists 

in the upper part of the basement beneath this basin. 

Densities of sedimentary rocks in the basin range from 

2.23 to 2.58 g/cm3, based on a limited review of well logs 

and bore-hole gravity measurements (Jenkins, 1986). The 

average density is about 2.45 g/cm3, in the deepest part of the 

basin. These density values for the Powder River Basin are 

somewhat lower than density values for sedimentary rocks 

used by Hurich and Smithson (1982) in the Green River 

Basin of from 2.25 to 2.65 g/cm3 and an average of more 

than 2.50 g/cm3. The greater thickness of the entire section 

is to the west, especially for Paleozoic and Mesozoic rocks, 

and the greater proportion of Tertiary rocks is in the east, 

which may account for the lower average density values for 

sedimentary rocks in the Powder River Basin. Although rock 

density studies of more regional extent are needed in the 

Powder River Basin, the values used in our gravity model 

(fig. 8) are the best currently available. 

The excellent fit between the calculated and observed 

gravity curves using a model in which density anomalies are 

confined to the upper 8 km (fig. 8) implies that additional 

density bodies or offsets are not required beneath the Bighom Mountains; that is, the main thrust fault flattens out in 

the middle crust without cutting the Moho or a major density 

boundary within the crust. No deep-crustal reflection data 

are available for the Bighorn Basin or Bighorn Mountains at 

this time, but both our shallow seismic data and gravity 

model for the Bighorn Mountains suggest that the main 

west-dipping thrust fault is analogous to the Wind River 

Thrust except that it has about half as much displacement 

and a detachment zone somewhat shallower in the crust. 

Other Geophysical Studies Supporting 

low-Angle Thrusting 

Three studies in southern Wyoming also support the lowangle thrusting model; however, pre-Laramide basementinvolved deformation in the southern Wyoming mountains 

during the Paleozoic adds complications beyond the scope of 

this study. A COCORP seismic profile over the Laramie 

Range in southeastern Wyoming (plate 1) shows smaller 

scale thrust faults dipping 30°-35° that can be traced to 

depths of 10 km and have offsets of approximately 3 km 

(Brewer and others, 1982; Johnson and Smithson, 1986). A 

gravity study of the Shirley Mountains and Hanna basin also 

concludes that low-angle north-dipping thrusts underlie the 

south flank ofthe Shirley Mountains (Oliver, 1970). An eastwest gravity model over the Rock Springs and Sierra Madre 

Uplifts in southwestern Wyoming by Bankey and 

Merewether ( 1990) shows a smaller scale listric thrust deep 

beneath the basin deposits. 

HORIZONTAL GRAVITY GRADIENT 

In deciding how to interpret the HGA gravity map (plate 

5), we assumed a simple two-body fault-contact model (fig. 

9) for which we computed the response for six different fault 

dip angles. The model assumes a 6-km-deep basin in which 

the sedimentary rocks have a -0.2 g/cm3 density contrast 

with respect to surrounding basement rocks. The six fault dip 

ElO Application of Structural Geology to Mineral and Energy Resources 
C) 

Dl 

< 

· 

AI 

:I 

0.. 

~ 

3 

r;· 

~ ~ 

~ 

!l 

s· 

:I 

0 

e:. 

~ 

~ 

6 

3 

:;-

OQ 

AI 

:I 

0.. 

~ c 

;. 

~ 

~ 

:I 

(./) 

0:: 

IJ.J 

1-

IJ.J 

~ 

0 

_J 

WEST 

;;;25 

z 

:i 

li: 

IJ.J 

0 

15 10 

KENNEY RANCH 

DRILL HOLE 

DISTANCE, IN KILOMETERS 

5 

AMERICAN 

QUASAR 

DRILL HOLE 

EAST 

0 

LEWIS SHALE 

MORRISON FORMATION 

TENSLEEP SANDSTONE 

(PENNSYLVANIAN) 

PRECAMBRIAN ROCKS 

~ Figure 7. Seismic line 4557 across the east flank of the Bighorn Mountains. UJ, Upper Jurassic rocks; UK, Upper Cretaceous rocks. Migrated depth. From Chevron 

g USA, Inc., used with permission; reprocessed by Miller and others (1988). Location of seismic line shown on plate 1. ~ ~ :I 

AI 

~ ... 
~ 

t.J 

> "'CC 

"E.. 

[ 

cs· 

::s 

Q 

..... 

~ c 

r'l 

~ !!. 

C') 

~ 

0 

~ 

6' 

~ :;· 

~ 

!!. 

AI 

::s 

Q. 

m 

::s 

~ 

~ 

~ 

~ c 

Cll 

~ 

WEST SOUTHWEST 

I

50 mGals 

CALCULATE 

4 I Rlr.l-lnRN BIGHORN it1ouNTAINS 

~(/) . a: 2 zw ol-

~~ 0 

>0 

w...J uj ~ -2 

-4 

0 50 

BOUGUER ANOMALY (OBSERVED) 

100 150 

DISTANCE, IN KILOMETERS 

Figure 8. Gravity and magnetic profiles across the Powder River Basin. Location of profile shown on plate 1. 

EAST 

200 250 
angles are 70° normal, vertical, and 20°, 30°, 45°, and 60° 

reverse (thrust faults). Although the amplitude of the anomalies over such faults will vary depending on the actual density contrasts and depth of basin fill, the position of the 

maximum gravity gradient is dependent solely on the fault 

angle. Whereas the position of the maximum gradient is 

directly over a vertical fault, the maximum gradient position 

is displaced toward the mountain range (with respect to the 

surface trace of a fault) for reverse faults and toward the 

basin for normal faults. Although the maximum gradients 

are within 2 km of the surface trace in all six cases, intermediate gradients are displaced 10-20 km toward the range for 

the low-angle thrusts (20° and 30°). 

The surface traces of the major range-bounding faults 

(plate 2) are plotted on the HGA map (plate 5) in order to 

compare gravity gradients with fault-trace positions. Examination of plate 5 reveals that intermediate gradients are displaced toward the ranges by 20 km or more in the Bighorn, 

Wind River, Owl Creek, and Granite Mountains and most of 

the other mountain ranges. Complications may arise where 

shallow density variations exist within uplifted basement 

blocks, where the core of the uplift has later collapsed, or 

where postorogenic debris has onlapped the leading edge of 

the eroded mountain front. Interference patterns between 

multiple faults could also confuse the interpretation. In spite 

of these complications, the overall pattern of asymmetric 

displacement of the maximum and intermediate gradients 

toward the mountain ranges provides additional evidence 

that the Laramide mountain ranges in this region are dominated by low-angle thrust faults. 

If the gravity data are of uniform quality and close 

enough spacing, the HGA map can be a valuable reconnaissance tool for estimating whether a particular range front 

fault is dominantly normal, vertical, or reverse. We have 

already mentioned the significant gaps in station spacing 

within this region, and caution should be exercised in using 

the HGA map in areas where these gaps exist. Two-dimensional modeling of the gravity profiles with constraints from 

seismic reflection and magneto-telluric profiles would be 

desirable to confirm inferences based on the HGA map. As 

more gravity stations with close spacing become available in 

the future, the value of the HGA maps for structural reconnaissance studies in the oil, gas, and mineral industry will 

increase significantly. 

ISOSTATIC COMPENSATION 

The purpose in generating an IRA gravity map (plate 4) 

is to preserve the signatures of anomalous densities within 

the upper crust while removing the effects of mountain roots 

assumed to lie directly beneath the elevated topography in 

the Airy-Heiskanen model. Because IRA values in the study 

area range from a high of + 75 mGal to a low of -77 mGal 

and average near zero, the technique appears to have successfully removed the long-wavelength effects of the 

regional topographic gradient (plate 1) from the CBA values 

(plate 3), which range from -75 to -285 mGal. This nearzero average implies that the area is isostatically compensated at regional scales, but that the Airy-Heiskanen model 

is not the dominant isostatic system involved in the compensation of this area and is not the optimum model for computing IRA maps. With average IRA values near zero and a total 

range of 152 mGal (versus 210 mGal for the CBA map), the 

IRA map is a significant improvement, however, over the 

CBA map in terms of its resolution for emphasizing upper 

crustal structures. 

In order to assess the magnitude of the gravitational 

effect of the mountain roots in the IRA map, we computed a 

two-dimensional Airy-Heiskanen isostatic model from the 

Green River Basin to the northern Powder River Basin (fig. 

I 0). The model crosses the highest peaks in the Wind River 

and Bighorn Mountains (also in Wyoming) and traverses the 

main topographic gradient from southeastern Montana to 

western Wyoming (plate 1). In making an Airy-Heiskanen 

isostatic correction, the weight of the topography above sea 

level is compensated for by an assumed local depression of 

the crust-mantle boundary. The weight of the topography 

above sea level must be balanced by the buoyancy of the 

crustal root beneath the depth of compensation. We assumed 

a density of 2.67 g/cm3 for the topography above sea level, a 

depth of compensation of 30 km, and a density contrast of 

-0.4 g/cm3 at the Moho. Note that the computed depth to 

Moho increases from 38 km in the northeast to 47 km in the 

southwest and has local maximums of about 58 km beneath 

the Wind River and Bighorn Mountains (fig. 10). 

The .calculated effect of the assumed isostatic root varies 

from -160 to -260 mGal from northeast to southwest and is 

a reasonably good first-order fit to the observed CBA profile 

in the high plains and intermontane basins; however, CBA 

values are 75-100 mGal higher than the effect of the roots 

over the Wind River and Bighorn Mountains; that is, the 

two-dimensional IRA values are +75 to +100 mGal over the 

main range, slightly higher values than those on the IRA map 

due to differences between two- and three-dimensional models as well as to slight differences in density contrast and 

depth of compensation. The values used for the IRA map 

result in slightly smaller anomalies, but the main conclusions 

are not affected by these minor differences between the twoand three-dimensional models or the assumed densities and 

depths of compensation. 

The local root beneath the Bighorn Mountains causes a 

local downwarp of more than 40 mGal below the regional 

trend of the effect of the isostatic root (fig. 10). In the previous section, where we reviewed the seismic profiles and twodimensional gravity models for the Wind River and Bighorn 

Mountains, we noted evidence for flattening of the thrust 

faults within the crust; that is, these Laramide mountains 

probably lack local roots. Therefore, the Airy-Heiskanen or 

Gravity and Seismic Reflection Data, Wyoming and Southern Montana E13 
~ 

~ 

> 

"t:) 

~ g· 

cs· 

:I 

0 

... 

~ c 

t"' 

2" 

~ 

C1 

2 

0 

~ 

0 

~ 

=· 

~ 

!!.. 

AI 

:I 

Q. 

"" :I 

~ 

~ ;:IC 

~ 

~ 

~ t"' 

~ 

"' 

-10 (f) 

_J 

<( 

~ _J 

_J 

~ 

~ 

>-' _J 

<( 

-30 I ~ 0 

z 

<( 

~ 

> <( 

a:: 

<.9 

-50 

-3 

(f) 

a:: SEA LIJ t;J LEVEL 

~ 

0 

_J 

Q 

~ 

I , 3 Ia.. 

LIJ 

0 

6 

80 90 

zoo . THRUST/ 

FAULT 

/ /I 

100 

DISTANCE, IN KILOMETERS 

\ 

Figure 9. Gravity anomaly- fault dip comparison model. Asterisks indicate maximum gradient. 

~p=-0 2 g/ cm3 

110 
SOUTHWEST NORTHEAST -100,--------,--------~--------~--------~--------~--------~--------~------~ 

(/) 

_J 

EXPlANATION 

o Observed 

<( 

~ -200 

• Calculated cfrBOUGUER ANOMALY 

ooo oo oOooooo 

o o oo 0 

_J o oooooO 

~ 0 

0 

o Oo o 

-300~------~--------~---------i ________ _L ________ _L ________ _L ________ ~------~ 

0,---------,---------,---------,---------~--------,--------,---------,--------~ 

(/) 

a: 

UJ 

fUJ 

~ 

0 

_J 

~ 30~----------------------------------------------------------------------------~ 

z LOWER CRUST DENSITY= 3.0 g/cm3 

:I 

h: 

UJ 

0 

200 400 

DISTANCE, IN KILOMETERS 

Figure 10. Airy-Heiskanen isostatic compensation model across Wyoming. Line of section shown on plate 1. 

local mechanism for compensation is not the optimum system, at least for shorter wavelength mountain ranges such as 

the Wind River and Bighorn ranges where their widths (or 

wavelengths) are less than 100 km. A modified AiryHeiskanen algorithm that would average the regional topography for wavelengths greater than 100 km would smooth 

out the effect of isostatic root curve (fig. 1 0), and this would 

result in smaller isostatic anomalies. With a smaller range of 

IRA anomalies, the resolution for discriminating anomalies 

caused by upper crustal structure would be enhanced. Such 

maps would be considerably more valuable for oil, gas, and 

mineral exploration than the present CBA or IRA maps and 

would also increase the resolution in the derivative HGA 

maps. Alternatively, Jachens and others (1989) suggested 

that the IRA maps can be filtered to remove the longer 

wavelength effects. This important area for future gravity 

research could yield valuable insight in areas of energy-rich 

basins like the Laramide province of Wyoming. 

The most important conclusion from the IRA map (plate 

4) and two-dimensional gravity model (fig. 10) is that the 

Laramide province of Wyoming is near isostatic equilibrium 

at regional scales; that is, at wavelengths longer than 100 km. 

The large local positive values over the mountain ranges 

imply, however, either that the ranges are underlain by highdensity crust or that the assumption in the IRA model of 

local roots beneath the ranges is, in fact, erroneous. Because 

most outcrops of basement rocks in the region lack evidence 

for large volumes of dense mafic rocks, the IRA map implies 

that the mountain ranges are rootless and not in local isostatic balance. This local imbalance happens because the local 

weight of the ranges on the crust and lithosphere is distributed over wavelengths longer than 100 km. 

ROCK DENSITY 

Hurich and Smithson (1982) could not model the entire 

gravity anomaly over the Wind River Mountains using only 

contributions from contrasts between basin-fill deposits and 

basement blocks in the ranges. They had to include an 

uplifted dense lower crustal layer (2.88 g/cm3) into the upper 

crust (2.73 g/cm3) along the Wind River Fault (fig. 3). 

Smithson and others (1978) and Karasa (1976) encountered 

the same difficulty. The contribution from these uplifted 

lower crustal rocks (fig. 3) is as much as +22 mGal over the 

range and decreases to less than 2 mGal at distances of 100 

km on either side of the mountain range. At least two alternative explanations (or combinations) can provide approximately +20 mGal. 

1. If the average density contrast between basin-fill 

deposits and the basement was slightly larger than that determined by Hurich and Smithson (1982), then a smaller contribution from the lower crust would be necessary. For 

example, a simple model using a basin-fill contrast of-0.06 

g/cm3 relative to the basement also produces a positive of 

Gravity and Seismic Reflection Data, Wyoming and Southern Montana El 5 
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DISTANCE, IN KILOMETERS 

Figure 11. Sediment density contrast model across Wind River Thrust. 

approximately 20 mGal (fig. 11). Therefore, the models are 

sensitive to relatively small changes in average sediment or 

basement density. Our gravity model over the Bighorn 

Mountain front and the Powder River Basin used basin-fill 

density values about 0.05 g/cm3 lower than those used by 

Hurich and Smithson, and we did not need a contribution 

from the lower crust. 

2. Our gravity and magnetic model of the Powder River 

Basin demonstrates that significant short-wavelength (<50 

km) lateral density and magnetization anomalies are present 

within the upper few kilometers of the basement (fig. 8). 

Longer wavelength (>50 km) lateral density variations averaging +0.05 g/cm3 in the upper 10 km of the basement could 

also provide contributions of approximately +20 mGal. 

Although we agree with Hurich and Smithson that the 

gravity models require that the thrust faults flatten before 

reaching the Moho, we are not convinced that the thrusts 

uplift dense lower crust. Very small differences in basement 

and basin-fill densities could alter this conclusion. 

A thorough discussion of rock density measurements is 

beyond the scope of this paper, and the reader is referred to 

the following references: Smithson (1971), Karasa (1976), 

Beyer and Clutsom (1978, 1980), Kososk.i and Robbins 

(1979, 1980a, b), Parks (1979), Hurich and Smithson (1982), 

Wagoner (1985), and Bankey and Merewether (1990). 

Although the density studies ofHurich and Smithson (1982) 

and Bankey and Merewether (1990) are probably the most 

comprehensive to date, additional studies using bore-hole 

gravity and (or) a larger sample of density logs throughout 

Wyoming are needed to confirm those results. 

STRUCTURAL SYNTHESIS 

A 400-km-long cross section from the Green River Basin 

to the Powder River Basin synthesizes the major geological 

and geophysical observations across Wyoming (fig. 12). The 

section is slightly modified from Peterson (1983) and shows 

the basic concept for flattening of the thrusts within the middle crust. Gries (1981, 1983), Lowell (1983), and Stone 

(1988) have proposed similar models. 

The maximum thrust displacement is approximately 20 

km in the Wind River Mountains and about 10 km in the Bighom Mountains. Uplift in the Wind River Mountains started 

in Late Cretaceous time and in the Bighorn Mountains in 

Paleocene time, and major deformation ceased in Eocene 

time in both ranges (Perry and others, this volume). This difference in onset times may account for differences in relative 

amounts of deformation in the two ranges. Smaller thrust 

E16 Application of Structural Geology to Mineral and Energy Resources 
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faults are in the Wind River and Bighorn Basins and an intermediate thrust is along the southwest flank of the Owl Creek 

Mountains. 

The cross section also shows Moho depths and crustal 

and mantle velocities based on seismic refraction studies 

(Braille and others, 1989). In the southwestern part of Wyoming, the Moho is at a depth of about 35 km, in good agreement with the Moho reflector on the Wind River COCORP 

profile (Sharry and others, 1986). In northeastern Wyoming, 

the Moho is at a depth of approximately 50 km. Although no 

deep-crustal refraction measurements are available in the 

central region of the section, we have dashed in a Moho dipping gently to the northeast with no local mountain roots 

beneath the mountain ranges. Note that the inferred regional 

dip of the Moho based on refraction data is opposite to the 

regional dip calculated for the two-dimensional Airy-Heiskanen isostatic profile (fig. 1 0). Therefore, although the 

Airy-Heiskanen model is a good first approximation to the 

long-wavelength components of the Bouguer gravity field, 

the actual mechanism for isostatic compensation must be 

dominated by lateral density variations in the crust and mantle that allow the actual Moho to rise to the southwest. Both 

crustal and upper mantle compressional velocities are lower 

in the southwest than in the northeast and support the inference of lateral density changes in the crust and upper mantle. 

Lateral density variation is a Pratt isostatic mechanism 

(Heiskanen and Vening Meinesz, 1958), and probably is the 

more dominant isostatic mechanism in the Laramide province of Wyoming. 

The CBA, IRA, and HGA gravity maps herein add to the 

body of geophysical and geologic data in the Laramide province that demonstrates the dominance of low-angle thrusts 

that probably flatten within the crust. Understanding the 

crustal structure and mechanism of isostatic compensation 

will ultimately allow improved isostatic algorithms to be 

developed, which will further enhance the resolution of IRA 

and HGA gravity maps for hydrocarbon and mineral exploration. 

CONCLUSIONS 

1. The +50 to+ 75 mGal isostatic residual anomaly (IRA) 

gravity values over Laramide mountain ranges imply that the 

mountains have much smaller roots than assumed in the 

local Airy-Heiskanen model; that is, they are rootless. 

2. The -50 to -77 mGal IRA values over the foreland 

basins in front of the Laramide mountain ranges are caused 

by the low-density basin fill. 

3. The near-zero average IRA values across Wyoming 

and southern Montana indicate that the Airy-Heiskanen 

model is a good first approximation to the complete Bouguer 

anomaly (CBA) gravity values and that the region is near 

isostatic equilibrium for wavelengths longer than 100 km. 

Gravity and Seismic Reflection Data, Wyoming and Southern Montana E17 
4. Modified IRA maps based on topography averaged 

over wavelengths longer than 100 km would correct for the 

erroneous roots assumed in the local Airy-Heiskanen model 

and thus reduce the peak to trough range of values and 

enhance the resolution for smaller geologic features of 

importance in energy and mineral exploration. Further 

research is needed in this area. 

5. Lateral variations in crustal and upper mantle compressional velocities imply that densities also vary laterally. 

Therefore, actual isostatic compensation must involve a 

combination of both Airy-Heiskanen and Pratt systems. 

6. The IRA and horizontal gradient anomaly (HGA) 

gravity maps support previous seismic reflection, twodimensional gravity modeling, and exploration drilling data 

indicating that low-angle (20°-30°) faulting is the dominant 

style of Laramide deformation. Basement overhangs of 

10-20 km are documented in some ranges, such as the Bighorn and Wind River Mountains. 

7. The high gravity gradients are displaced toward the 

mountain ranges on the HGA maps and further support the 

concept of low-angle thrust faulting. 

8 The COCORP Wind River seismic profile, as well as 

two-dimensional gravity models there and in the Bighorn 

Mountains, indicate that the thrust flattens out in the middle 

or lower crust and that the mountains are rootless, as proposed by Peterson (1983). 

9. Energy and mineral exploration can benefit from the 

gravity maps presented in this paper, and when the numerous 

gaps in the field data are filled the value of the maps will 

improve significantly. 

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E18 Application of Structural Geology to Mineral and Energy Resources 
Hurich, C.A., 1981, Gravity interpretation of the southern Wind 

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porosity estimates from borehole gravity data in limestones 

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__ 1980a, In-situ bulk-density estimates from borehole gravity 

data in the Madison Group test well no. 3, Yellowstone 

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80-784, 11 p. 

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65-77. 

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E20 Application of Structural Geology to Mineral and Energy Resources 
Chapter F 

Pre-Late Eocene Structures and their Control of 

Gold Ore at the Drum Mine, West-Central Utah 

By C.j. NUTT and C.H. THORMAN 

U.S. GEOLOGICAL SURVEY BULLETIN 2012 

APPLICATION OF STRUCTURAL GEOLOGY TO MINERAL AND 

ENERGY RESOURCES OF THE CENTRAL AND WESTERN UNITED STATES 
CONTENTS 

Abstract Fl 

Introduction Fl 

Geologic setting Fl 

DrumMine F3 

Low-angle and bedding-parallel faults F4 

Discussion F5 

References cited F7 

FIGURES 

1-3. Maps showing: 

1. Location of Drum Mine, Detroit Mining District and Drum Mountains, westcentral Utah Fl 

2. Geology of Drum Mine area F2 

3. Detailed geology of Drum Mine F4 

4. Chart showing generalized stratigraphy of Cambrian rocks at Drum Mine F6 

5. Photograph showing southwest pit, Drum Mine F6 

TABLE 

1. Age determinations of igneous rocks in the Detroit Mining District, west-central 

Utah F3 

Fll Contents 
Pre-Late Eocene Structures and their Control of 

Gold Ore at the Drum Mine, West-Central Utah 

By C.j. Nutt 1 and C. H. Thorman 1 

Abstract 

The Drum Mine in west-central Utah exploited disseminated gold from two orebodies in Middle Cambrian carbonate 

rocks and shale and Eocene to Oligocene igneous rocks and 

related pebble dikes and sills. The mine is in the central part of 

the Drum Mountains and in the Detroit Mining District, which 

is in part underlain by late Eocene to Oligocene porphyry. Ore 

deposition at the mine was controlled by faults parallel with or 

at low angles to bedding and their associated ramp structures 

and by high-angle fractures and faults with little offset. The 

low-angle structures are primarily younger-over-older faults 

that thin units and are concentrated in or near shale units. Igneous and pebble sills, many of which are ore bearing, follow the 

bedding-parallel faults and ramps. East-west- to northeasttrending faults bound the mine area and are postulated to be 

tear faults associated with low-angle faulting. These high-angle 

faults controlled, and therefore predate, a paleovalley filled 

with flows of 42-Ma Drum Mountains Rhyodacite. A pre-late 

Eocene age of low-angle faulting is based on the post-lowangle-fault age of Eocene to Oligocene sills and pebble sills 

localized along low-angle faults and the pre-42-Ma age of the 

associated high-angle faults that controlled the paleovalley. 

INTRODUCTION 

The Drum Mine in the eastern Great Basin of west-central Utah (fig. 1) exploited two orebodies of disseminated 

gold hosted by sedimentary and igneous rocks. The orebodies were mined in two pits, herein referred to as the northeast 

and southwest pits (fig. 2). At the time of our study from 

1987 to 1989, part of the main orebody in the southwest pit 

and pods of ore in the northeast pit were exposed; the rest of 

the ore had been removed. Our work shows that ore deposition was controlled by faults parallel with or at low angles to 

bedding and their associated ramp structures and by highangle fractures and faults. The low-angle to bedding-parallel 

faults are mostly younger-over-older faults that thin units; 

1

U.S. Geological Survey, Box 25046, MS905, Denver, Colorado 

80225. 

rarely units are thickened. Although thinning of units in the 

Great Basin is commonly interpreted as the result of middle 

Tertiary or younger extension, we interpret the low-angle to 

bedding-parallel faults to be older than late Eocene and probably related to Jurassic or Late Cretaceous to early Tertiary 

(Sevier-age) deformation. 

GEOLOGIC SETTING 

The Drum Mine is in the central part of the Drum Mountains, in an area underlain by Upper Proterozoic to Lower 

Cambrian quartzite and subordinate shale and siltstone, Middle Cambrian carbonate rocks and subordinate shale, and 

Tertiary igneous rocks and minor fanglomerate and conglomerate. A late Eocene to Oligocene quartz monzonite 

porphyry system underlies the central part of the Drum 

10 20 MILES I I I 

I I I I 

10 20 30 KILOMETERS 

Figure 1. Location of Drum Mine (solid circle), Detroit Mining 

District (diagonal line pattern), and Drum Mountains (dark 

screen) in west-central Utah. Mountain ranges are shown by 

light shading; approximate outline of Thomas Caldera is shown 

by dashed line. 

Structural Control of Gold Ore, Drum Mine, Utah F1 
N 

QUATERNARY 

t ii i ii I Alluvium 

TERTIARY 

Tuff 

Mt. Laird porphyry dike 

0.5 KILOMETER 

Drum Mountain Rhyodacite 

Fanglomerate 

-

M~~e~~e~~~~eRIAN 

Swasey Limestone 

Whirlwind Formation 

Dome Limestone 

Drum mine 

Southwest pit\:,. ... -'\ 

' \ 

I \, 

\ 

\ 

0.5 MILE 

EXPLANATION 

Chisholm Formation, undifferentiated 

south of Drum Peak Fault 

Upper shale 

Middle thin-bedded limestone 

Lower shale 

Brown-weathering altered Howell 

Limestone and Tatow Member of 

Pioche Formation, undifferentiated 

Howell Limestone 

Marker shale shown by dotted line 

Tatow Member of Pioche Formation 

LOWER CAMBRIAN EB Lower member of Pioche Formation 

· and upper part of Prospect Mountain 

Quartzite, undifferentiated 

Figure 2. Geology of Drum Mine area. Modified from Nutt and others (1991 ). 

F2 Application of Structural Geology to Mineral and Energy Resources 

20 ____L_ 

Contact-Dashed where approximately 

located 

Fault-Dashed where approximately 

located, dotted where concealed 

Low angle to bedding-parallel fault that 

thins and locally thickens beds-Hachures 

on upper plate, dashed where approximately 

located 

Strike and dip of beds 

Pebble dike 

Breccia 

Jasperoid 

Approximate outline of the Drum Mine 

property; includes pits, tailings, and 

buildings 
Table 1. Age determinations of igneous rocks in the Detroit Mining District, west-central Utah 

[nd indicates not determined] 

Age (Mal 

Unit Fission track 40Ar/39Ar Comment 

Felsite dike 

Mt. Laird TufT 

1

35.6 nd 239 336.5 

Minimum age because zircon may have been reset during alteration. 

Extrusive equivalent of porphyry in mining district. Ar-Ar age from 

what is called Mt. Laird TufT in Keg Mountains, Utah. 

Drum Mountains Rhyodacite 242 nd None. 

1 Robert A Zimmermann (U.S. Geological Survey, written commun., 1989). 

2Lindsey (1982). 3Shubat and Snee (this volume). 

Mountains, also known as the Detroit Mining District, and is 

expressed at the surface by dikes and plugs of altered porphyry and numerous pebble dikes. Gold- and copper-bearing 

veins and manganese replacement deposits mined from the 

late 1800's to the mid-1900's are related to this porphyry 

system. Nutt and others (1991) proposed that the Drum gold 

deposit is also porphyry related. 

The Drum Mountains are in the Cretaceous to early Tertiary Sevier Thrust Belt, typified by easterly directed thrusting and folding. In addition, the Drum Mountains are in the 

eastern Great Basin part of the Cenozoic Basin and Range 

Province, a region that has undergone considerable extension. Typically, ranges in this part of the Great Basin are 

tilted blocks of unmetamorphosed rocks that appear little 

deformed; however, Hintze (1978) documented stratigraphic 

thinning that he attributed to Sevier-age bedding-parallel 

faults in the House and Fish Springs Ranges west of the 

Drum Mountains (fig. 1). 

The Drum Mountains are in the midst of Eocene to Oligocene calc-alkaline and Miocene bimodal volcanic fields 

that extend west from the Tintic Mountains to the eastern 

side of the Deep Creek Range (fig. 1 ). In the vicinity of the 

Drum Mountains, volcanism commenced with the eruption 

of flows of the Drum Mountains Rhyodacite (about 42 Ma), 

followed by voluminous eruption of tuffs and caldera formation from about 39 Ma to 35 Ma (Lindsey, 1982). The 

Detroit Mining District is just south of the ring fracture fault 

of the Eocene to Oligocene Thomas caldera (fig. 1 ). Lindsey 

( 1982) proposed that the porphyry in the Drum Mountains is 

an intrusive phase of the caldera-related Mt. Laird Tuff. 

Table 1 shows age determinations of some igneous rocks in 

the Detroit Mining District. 

DRUM MINE 

The Drum Mine is in the lower part of the Cambrian miogeoclinal carbonate section, just above the transition from 

Late Proterozoic and Early Cambrian siliciclastic sedimentation to Cambrian carbonate sedimentation (fig. 2). A 

southwest-dipping section of Cambrian rocks cropping out 

in the mine includes, from oldest to youngest, the lower 

member of the Pioche Formation (quartzite, siltstone, and 

shale), the Tatow Member of the Pioche Formation (dolomitic limestone), and the Howell Limestone (thick-bedded 

limestone) in the northeast pit and the Chisholm Formation 

(shale and thin-bedded limestone) and Dome Limestone 

(thick-bedded limestone) in the southwest pit (figs. 3, 4). The 

best exposed orebody in the mine is in the Chisholm Formation, which consists of a lower shale, middle thin-bedded 

limestone, and an upper shale (fig. 4). Gold-bearing clayaltered rock and jasperoid are concentrated in the shale units, 

primarily near contacts with overlying and underlying limestone. 

At least five types of igneous rocks, as well as numerous 

pebble dikes and sills, crop out near or in the mine (figs. 2, 

3): the Drum Mountains Rhyodacite (about 42 Ma) forms the 

edge of the northeast pit and probably once covered the 

entire mine area; Mt. Laird porphyry sills and dike (39-36.5 

Ma) cut the sedimentary rocks in the mine and surrounding 

area; felsite dikes and sills (35.6 Ma or older) crop out in the 

mine and to the north; a muscovite-bearing tuff (unknown 

age) is exposed in the northeast pit; and a vent-facies tuff 

(unknown age) is present east of the mine. Igneous-related 

breccias and pebble dikes and sills, mostly associated with 

Mt. Laird porphyry and felsite, are common in the mine area. 

Pebble dikes and sills range from almost completely granulated quartzite fragments to scattered quartzite clasts in clay 

matrix. Clay-altered Mt. Laird porphyry is present both as 

fragments and as matrix material. The abundance of igneous 

and pebble dikes and sills suggests that the mine area is 

underlain by porphyry. 

The mine area is in a present-day topographically low 

area, partly filled with Drum Mountains Rhyodacite, that is 

interpreted to be an exhumed paleovalley in Paleozoic rocks. 

Fanglomerate containing Cambrian carbonate clasts crops 

out along the northern margin of the paleovalley, and conglomerate overlies Paleozoic rocks in part of the mine area. 

The paleotopography had a trellis geometry similar to the 

present drainage, as demonstrated by the presence of flow 

rocks in one of the present-day trellis tributaries to the main 

valley. 

Structural Control of Gold Ore, Drum Mine, Utah F3 
/ 

/ 

Northeast pit 

/ / .r-r-"""T ""' - (..;./ 

Approximate outline / 

//~}~~~ 

Southwest pit 

of pit""",// 

I 

I 

I 

I 

I 

200 

I I 

50 

I 

400 FEET I 

100 METERS 

N 

Figure 3 (above and facing page). Detailed geology of Drum Mine. Modified from Nutt and others (1991 ). 

LOW-ANGLE AND 

BEDDING-PARALLEL FAULTS 

Previously undetected faults that are parallel with or at a 

low angle to bedding are throughout the Drum Mountains. 

These faults, herein referred to as bedding-parallel or lowangle faults, are mostly younger-over-older structures and 

F4 Application of Structural Geology to Mineral and Energy Resources 

are primarily in shaly and thin-bedded units that are poorly 

exposed. Faults are most prominent in shales near the contact 

between the Upper Proterozoic to Lower Cambrian quartzite 

section and the overlying Cambrian carbonate section and in 

shale-rich units such as the Chisholm Formation. 

In the Drum Mountains, domains of extremely attenuated 

rocks are bounded by high-angle faults. The mine area is an 
""'1 v .,. "> 

~~"~:~ 

0 0 0 0 

Q Q D Cl Q. 0 

EXPLANATION 

TERTIARY 

Felsite 

Mt. Laird porphyry dike and sill 

Lithic-rich tuff, relative age uncertain 

Drum Mountains Rhyodacite 

Conglomerate and clay-altered igneous 

rock 

CAMBRIAN 

~ Dome Limestone 

>>::3 ?> f:£2~ 

:):/.()\ 

Chisholm Formation 

Upper shale 

Middle thin-bedded limestone 

Lower shale 

Howell Limestone 

Marker shale 

Pioche Formation 

Tatow Member (altered Howell 

Limestone included) 

Lower member 

---- Contact-Dashed where approximately 

located 

---··High-angle fault fracture-Dashed where 

approximately located 

.J....1.J..J.. Low-angle to bedding-parallel fault that 

thins and locally thickens beds-Hachures 

on upper plate 

2 5 

Unit cut by numerous bedding-parallel faults 

that cut out and thicken beds. Offset ranges 

from inches to probably at least hundreds 

of feet 

_j__ Strike and dip of beds 

~ Pebble dike and sill 

D Jasperoid 

'-""' Ore zone exposed in pit; consists of 

altered pebble sill and dike, shale, and 

thin-bedded limestone 

-5,800- Topographic contour-In feet 

example of these highly attenuated domains. The Chisholm 

Formation thins dramatically in the vicinity of the Drum 

Mine (figs. 2, 3), and the upper shale, about 45 ft (14 m) 

thick, is cut out southeast of the mine. The fault block that 

includes the mine area is bounded by two east-west- to northeast-trending faults about three-quarters of a mile (1 km) 

apart (fig. 2). The southern fault, the Drum Peak Fault, is 

well exposed, whereas the northern fault, the PV Fault, is not 

exposed but is interpreted to underlie the northern margin of 

the paleovalley based on small-scale folding along its edge. 

We propose that the Drum Peak and PV Faults are tear faults 

bounding the deformed block that contains the mine and that 

this deformation prepared the area structurally for mineralization. 

Low-angle to bedding-parallel faults are widespread in 

the Drum Mine (fig. 3). Most are in shale units and near contacts separating units of contrasting competency-the lower 

member of the Pioche Formation, at and near the contact 

separating the clastic rocks from the overlying carbonate 

section, and the Chisholm Formation. In many cases, it is 

difficult to identify one main fault; rather, numerous faults 

having offsets ranging from inches to at least hundreds of 

feet are in favorable intervals such as near the upper and 

lower contacts of shale units. 

The best examples of low-angle to bedding-parallel 

faults and associated structures are in the Chisholm Formation in the southwest pit. Both the upper and lower shales are 

folded, broken, and cut by low-angle faults. Breccia is most 

extensive near shale-limestone contacts and in the limestone 

hanging wall. At the bottom of the southwest pit, where the 

shale is cut out, the fault is marked by folded thin-bedded 

limestones. Ramps are associated with the faults and uniformly step up to the northeast. Pebble sills, most of which 

are ore bearing, are concentrated along low-angle faults at 

shale-limestone contacts and particularly above ramps. In 

general, the best ore was mined from ramp areas (Jumbo 

Mining Company, oral commun., 1989). Figure 5, a photograph of the bottom of the southwest pit, shows some of 

these features . 

High-angle structures that show little to no displacement 

were important conduits for gold-bearing fluids. Many of 

these structures are filled with pebble dikes. The most prominent of these is the King Tut Fault, marked at the surface by 

a pebble dike and at depth by ore at its intersection with shale 

units. Flexure zones above ramps typically have arrays of 

radiating fractures, but the fractures generally do not extend 

more than a hundred feet or so. We observed no high-angle 

faults that sole into the low-angle faults nor any rotation of 

blocks above the highly attenuated shale units. 

DISCUSSION 

Gold ore at the Drum Mine was localized both by the 

low-angle to bedding-parallel faults and by high-angle faults 

and fractures. Ore remaining in the southwest pit is concentrated along low-angle and bedding-parallel faults and associated ramps that are best developed in and near shales in the 

Chisholm Formation. 

The timing of the low-angle and bedding-parallel faulting is mainly constrained by crosscutting igneous rocks. 

Pebble sills, primarily associated with the 39-36.5-Ma Mt. 

Laird porphyry and 35.6-Ma or older felsic dikes, were 

localized along low-angle faults and ramps and therefore 

postdate faulting. The sills are undeformed except for minor 

postemplacement shears. Also, pebble dikes are along 

Structural Control of Gold Ore, Drum Mine, Utah F5 
365ft Cliff-forming thin-bedded to massive gray Chisholm Formation 

(111m) limestone. Oncolitic and oolitic locally. 

Upper olive-green to brown shale, red-weathering 

Dome Limestone 200ft oncolitic limestone; middle ledge-forming MIDDLE (61 m) thin-bedded gray limestone; basal olive-green to CAMBRIAN 

tan shale, G/ossop/eura-bearing limestone. 

Howell Limestone 330ft Cliff-forming thin-bedded to massive gray 

(101m) limestone, oncolitic locally. oolitic limestone at 

top. Green shale locally in part. 

Tatow Member 55ft Brown-weathering dolomitic limestone. (17m) 

r·························· I Pioche Black-weathering fine-grained quartzite and 

LOWER Formation 360ft h micaceous quartzite, green to brown micaceous 

CAMBRIAN Lower member (1 10m) ::: 

1<::: :::::::::::::::::::::::::::::::::::\. shale and siltstone. 

l:>~ 

Figure 4. Generalized stratigraphy of Cambrian rocks at the Drum Mine. 

Figure 5. Photograph showing southwest pit, Drum Mine. Bedding-parallel fault cuts out upper shale of Chisholm Formation (c) 

and places Dome Limestone (d) on middle limestone of the Chisholm Formation (c). Limestone beds are folded where shale is 

attenuated. On right side of photograph, gold-bearing pebble sill is localized above ramp. Bench is about 25ft (7 .6 m) above road. 

F6 Application of Structural Geology to Mineral and Energy Resources 
high-angle faults and fractures, many of which are associated with low-angle faults, although some probably formed 

or reopened during porphyry intrusion. 

A second and more tenuous constraint on the formation 

of low-angle faults is based on the age of the paleovalley that 

includes the mine area. The paleovalley must be older than 

the Drum Mountains Rhyodacite that filled it at about 42 Ma. 

If, as we postulate, the paleovalley is controlled by tear faults 

related to the low-angle faults, then low-angle faults predate 

erosion of the paleovalley as well as later 42-Ma flows. Erosion followed a period of uplift, probably forming the paleovalley in the Late Cretaceous to early Tertiary. It should be 

noted that Anderson (1983) believed that throughout this 

part of Utah the basal volcanic rocks are detached and that 

what has been interpreted as the base of the paleo valleys is a 

detachment surface. This interpretation would negate our 

argument on age constraint and imply that all low-angle 

faults are middle Tertiary or younger in age. The presence of 

fanglomerate, which does not contain volcanic clasts and 

underlies the volcanic rocks, is consistent, however, with the 

paleovalley interpretation. Also, volcanic flow rocks that are 

depositional on Cambrian limestones in a paleotributary of 

the trellis paleodrainage strongly support our interpretation 

of a normal, undisturbed basal Tertiary unconformity. We 

believe that little, if any, movement has occurred at the base 

of the Tertiary volcanic sequence in this area. 

These age constraints date the low-angle faults and associated ramps as being older than late Eocene. Although not 

necessarily of common origin, younger-over-older faults of 

probable Mesozoic age have been documented to the north 

by Allmendinger and Jordan (1984) in the Newfoundland 

Mountains, Utah, and by Miller and others (1987) in the Pilot 

Range, Utah and Nevada. Recognition of younger-overolder faults similar to those in the Drum Mountains is, as of 

now, restricted to the hinterland of the Sevier Thrust Belt, a 

distribution that suggests they are related to Sevier deformation. However, Jurassic deformation, variably included in 

the Sevier orogeny or considered a separate event, is increasingly being recognized in the Great Basin (Allmendinger and 

Jordan, 1984; Miller, 1987; Thorman and others, 1990) and 

may have occurred in this part of Utah. Regardless of exact 

timing, these low-angle faults predate middle Tertiary extension and imply significant pre-late Eocene deformation in 

the area. 

REFERENCES CITED 

Allmendinger, R.W., and Jordan, T.E., 1984, Mesozoic structure of 

the Newfoundland Mountains, Utah-Horizontal shortening 

and subsequent extension in the hinterland of the Sevier belt: 

Geological Society of America Bulletin, v. 95, p. 

1280-1292. 

Anderson, R.E., 1983, Cenozoic structural history of selected areas 

in the eastern Great Basin, Nevada-Utah: U.S. Geological 

Survey Open-File Report 83-504, 47 p. 

Hintze, L.F., 1978, Sevier orogenic attenuation faulting in the Fish 

Springs and House ranges, western Utah: Brigham Young 

University Geology Studies, v. 25, pt. 1, p. 11-24. 

Lindsey, D.A., 1982, Tertiary volcanic rocks and uranium in the 

Thomas Range and northern Drum Mountains, Juab County: 

U.S. Geological Survey Professional Paper 1221, 71 p. 

Miller, D.M., Hillhouse, W.C., Zartman, R.E., and Lanphere, M.A., 

1987, Geochronology of intrusive and metamorphic rocks in 

the Pilot Range, Utah and Nevada, and comparison with 

regional patterns: Geological Society of America Bulletin, v. 

99, p. 866-879. 

Nutt, C.J., Thorman, C.H., Zimbelman, D.R., and Gloyn, R.W., 

1991, Geologic setting and trace-element geochemistry of 

the Detroit mining district and Drum gold mine, Drum 

Mountains, west-central Utah, in Raines, G.L., Lisle, R.E., 

Schafer, R.W., and Wilkinson, W.H., eds., Geology and ore 

deposits of the Great Basin, Symposium Proceedings, 1990, 

Reno, Nevada: Geological Society of Nevada and U.S. Geological Survey, v. 1, p. 491-509. 

Thorman, C.H., Ketner, K.B., and Peterson, Fred, 1990, The Elko 

orogeny-Late Jurassic orogenesis in the Cordilleran miogeocline: Geological Society of America Abstracts with Program, v. 22, no. 3, p. 88. 

Structural Control of Gold Ore, Drum Mine, Utah F7 

Chapter G 

High-Precision 40 Ar/39 Ar Geochronology, 

Volcanic Stratigraphy, and Mineral Deposits of 

Keg Mountain, West-Central Utah 

By MICHAEL A. SHUBAT and LAWRENCE W. SNEE 

U.S. GEOLOGICAL SURVEY BULLETIN 2012 

APPLICATION OF STRUCTURAL GEOLOGY TO MINERAL AND 

ENERGY RESOURCES OF THE CENTRAL AND WESTERN UNITED STATES 
CONTENTS 

Abstract G1 

Introduction G 1 

Geologic setting G2 

Geochronology G2 

40 Arf39 Ar dating methods G2 

40 Arf39 Ar data GS 

Volcanic stratigraphy and igneous rocks G6 

Andesite of Keg Spring G6 

Tuff of Dead Ox G6 

Tuff of Keg Mountain GS 

Quartz monzonite porphyry G10 

Mt. Laird Tuff G10 

Dacite porphyry intrusive rocks G10 

Joy Tuff G10 

Rhyolite porphyry G10 

Dell Tuff G12 

Alkali rhyolite G 12 

Structure G12 

Mesozoic compressional structures G 12 

Caldera and cauldron structures G 12 

Flint Spring cauldron G 12 

Keg cauldron G13 

Thomas caldera G 13 

Picture Rock caldera G13 

High-angle normal faults G 13 

Mineral deposits G13 

Volcanic-hosted gold deposits G14 

Dome Hill prospect G 14 

Lead Hill prospect G 14 

Conclusions G14 

References cited G15 

FIGURES 

1. Map showing location of Keg Mountain, west-central Utah G2 

2. Composite age-spectrum diagram for nine samples from Keg Mountain G3 

3. Map showing simplified geology of Keg Mountain G9 

4. Complete Bouguer gravity map of the north-central part of Delta, Utah, 1 °X2° 

quadrangle G11 

TABLES 

1. 40 Arf39 Ar age-spectrum data for Keg Mountain, Utah G4 

2. Summary of 40 Ar/39 Ar age-spectrum dates for Keg Mountain, Utah GS 

Gil Contents 
High-Precision 40 Ar;39 Ar Geochronology, Volcanic 

Stratigraphy, and Mineral Deposits of 

Keg Mountain, West-Central Utah 

By Michael A. Shu bat 1 and Lawrence W. Snee2 

Abstract 

High-precision 40Arf39Ar age-spectrum dates combined 

with results of geologic mapping define the Eocene and Oligocene history of Keg Mountain, which straddles the center of 

the Deep Creek-Tintic mineral belt in west-central Utah. The 

chronology of volcanic events consists of (1) eruption of the tuff 

of Dead Ox and subsidence to form the Flint Spring cauldron 

(undated but presumed to be late Eocene), (2) eruption of the 

dacitic tuff of Keg Mountain (36.77±0.12 Ma) and subsidence 

of the Keg cauldron, followed by resurgent doming, (3) eruption of the Mt. Laird Tuff (36.54±0.06 Ma), collapse of the Thomas caldera, and intrusion of hypabyssal dacite (36.49±0.15 

Ma), and (4) eruption of the Joy Tuff (34.88±0.06 Ma), collapse 

of the Dugway Valley cauldron and possibly coeval Picture 

Rock caldera, and intrusion of rhyolite porphyry (35.14±0.15 

Ma), (5) eruption of the Dell Tuff (approximately 32 Ma) from 

an unknown source, and (6) eruption of late Miocene alkali 

rhyolite. 

These calderas and cauldrons, the margins of which are 

poorly defined and mostly inferred, outline the eastern part of 

a large, late Eocene to early Oligocene igneous complex that 

encompasses the Thomas Range, Keg Mountain, and northern 

Drum Mountains area. Associated with eruptive members of 

the igneous complex are hypabyssal intrusive rocks and mineral deposits. In detail, however, spatial relationships between 

inferred caldera margins, associated hypabyssal intrusions, and 

mineral deposits are not well understood. Mineral deposits in 

the area consist of volcanic-hosted gold prospects, polymetallic veins, and precious-metal-enriched jasperoid bodies that 

are along high-angle, small-displacement normal faults apparently unrelated to inferred caldera margins. 

1

Utah Geological Survey, 2363 South Foothill Drive, Salt Lake City, 

Utah 84109-1491. 

2

U.S. Geological Survey, Box 25046, MS 963, Denver, Colorado 

80225. 

INTRODUCTION 

Keg Mountain in west-central Utah (fig. 1) is underlain 

by a diverse assemblage of volcanic and subvolcanic rocks 

that span the interval from the late Eocene to late Miocene. 

The Eocene and Oligocene history of this area is of particular 

interest because it provides a view of the onset and development of one of western Utah's enigmatic east-trending mineral (and magmatic) belts, the Deep Creek-Tintic mineral 

belt (Shawe and Stewart, 1976). New dates and results of 

geologic mapping presented in this study show that hypabyssal intrusions related to caldera-forming events extend 

many kilometers (> 1 0) from known or inferred caldera margins. Comparison of results from this study with preliminary 

results from the adjacent Simpson Mountains (Y ambrick and 

Snee, 1989) suggests that high-precision 40 Arf39 Ar agespectrum dates can be used to correlate non-ash-flow volcanic and subvolcanic rocks related to caldera-forming events 

between adjacent ranges. 

Previous regional geologic studies by Erickson (1963), 

Shawe (1972), Lindsey and others (1975), Lindsey (1982), 

and Morris (1987) note the complex volcanic geology of 

Keg Mountain and stress the need for detailed mapping. 

Some unresolved problems stemming from these investigations are location of the eastern sides of the Thomas caldera 

and Dugway Valley cauldron (Lindsey, 1982), verification 

of Shawe's (1972) proposed Keg caldera, and location of 

suspected caldera margins related to the Joy and Dell Tuffs 

(Lindsey, 1982). The locations of these caldera and cauldron 

margins are of increasing economic interest because of the 

possible genetic link between hypabyssal intrusions related 

to caldera-forming events and epithermal gold mineralization in the Keg and Drum Mountains. This report presents 

the results of a joint Utah Geological and Mineral 

Survey-U.S. Geological Survey investigation of the geology 

and mineral potential of Keg Mountain as part of the Delta 

1 °X2° quadrangle study of the Conterminous United States 

Mineral Assessment Program (CUSMAP). 

40ArP9

Ar Geochronology, Keg Mountain, Utah Gl 
~UTAH Area of 

map 

JUAB COUNTY 

BERYLLIUM ROAD 

Sevier Desert 

39'15' .___ ________ .a__ ________ __j 

25 KILOMETERS 

Figure 1. Location of Keg Mountain, west-central Utah. 

GEOLOGIC SETTING 

Within the Keg Mountain area is limited exposure of the 

Cambrian, upper part of the Prospect Mountain Quartzite 

and Middle Cambrian carbonate rocks, typical of these units 

exposed elsewhere in the region. These units record an Early 

Cambrian marine transgression and Middle Cambrian deposition of miogeoclinal carbonate strata. During the Cretaceous Sevier orogeny the Prospect Mountain Quartzite was 

thrust over the carbonate rocks; this thrust fault is exposed in 

Keg Mountain. Geologically, however, Keg Mountain is primarily composed of Cenozoic volcanic rocks and subvolcanic intrusive rocks. These Cenozoic rocks lie near the axis of 

the broad, east-trending Deep Creek-Tintic mineral belt 

(Shawe and Stewart, 1976; Stewart and others, 1977). 

Numerous high-angle normal faults that formed during 

regional extension in the late Eocene to late Miocene control 

the location of many of the mineral deposits at Keg Mountain and in surrounding districts. 

Lindsey (1982), Lindsey and others (1975), and Shawe 

(1972) separated volcanism and related mineralization in the 

Thomas Range, Drum, Keg, and Desert Mountains into three 

stages. Using fission-track dates, Lindsey (1982) quantified 

the following stages. The oldest stage of activity (late 

Eocene, 42-39 Ma) consisted of eruptions of calc-alkaline, 

intermediate composition (rhyodacite-quartz latite) volcanic 

rocks and related intrusive rocks. In the Thomas Range and 

Drum Mountains, this oldest stage culminated with the eruption of the Mt. Laird Tuff and the formation of the Thomas 

caldera (Lindsey, 1982). Mineral deposits related to this 

stage include those of copper, manganese, and gold in the 

Drum Mountains and may include polymetallic vein and 

replacement and gold at Keg Mountain. Eruptions of rhyolitic ash-flow tuffs, continued caldera subsidence, and 

intrusion of felsic stocks and plugs characterize the middle 

stage of activity (early to middle Oligocene, 38-32 Ma). In 

the Thomas Range, Lindsey (1982) related subsidence of the 

Dugway Valley cauldron with eruption of the Joy Tuff. The 

middle stage apparently lacks associated mineral deposits. 

The youngest stage of activity (Miocene, 21 to 7~ Ma) consists of bimodal alkali rhyolite-basalt volcanism. Lithophile 

mineral deposits that formed during this stage include the 

world-class beryllium deposit at Spor Mountain (fig. 1) 

(Lindsey, 1977) and uranium (Lindsey, 1982) and fluorspar 

deposits in the Thomas Range. 

GEOCHRONOLOGY 

40Arj39Ar Dating Methods 

Samples from the Keg Mountains were dated using the 

40 Arf39 Ar age-spectrum technique, a variant of the conventional Kl Ar method. To obtain a date using this technique, 

the sample of unknown age and a standard of known age are 

irradiated together in a nuclear reactor to produce 39 Ar from 

39K by fast-neutron bombardment. After irradiation, the 

40 Ar RadiogeniJ39 Ar Potassium ratios of sample and standard 

are determined. An 40 Arf39 Ar isotopic date of a sample is 

calculated according to the relationship 

fu=li'Aln(JF+ 1), 

where tu is the calculated date of the sample, A. is the decay 

constant for radioactive decay of 40K to 40 Ar and 40Ca, J is 

related to neutron flux during irradiation, and F is the ratio 

of 40 Ar R (radiogenic 40 Ar) to 39 ArK (potassium-derived 

39 Ar) of the sample. The decay constants used in this study 

are those recommended by Steiger and Jager (1977): 

/...e=0.58lxl0-10/yr, A.~=4.962xl0-10/yr, and A.=A.E+'A~= 

5.543xlQ-10fyr. The flux parameter, J, is calculated according to the relationship l=(e'A.tm-1 )1(40 Ar 39 Ar K)m where tm is 

the age of the primary flux "monitor" (that is, standard) and 

(40Ar~39ArK)m is the measured and corrected ratio of the 

standard. The standard for this experiment was hornblende 

MMhb--1 with %K=1.555, 40ArR=1.624x1o-9 mollg, and 

K-Ar date=520.4 Ma (Alexander and others, 1978; Samson 

and Alexander, 1987). In order to determine the actual 

40Ar~39 ArK ratio of a sample or standard, however, it is necessary to correct for the presence of atmospheric 40 Ar and 

irradiation-produced interfering isotopes such as 40 Ar (from 

40K), 39 Ar (from 42Ca), and 36 Ar (from 40Ca and 35Cl). 

G2 Application of Structural Geology to Energy and Mineral Resources 
40 I 

f--

"" 

38 Ir--..., 

.....__ 

~ I l 36.6±0.1 Ma 

hL... Ul I I 

~L I 

..... 

rl llr-~ 

::::> 

z 

z 

<{ 

<i: 36 <..9 

UJ 

~ 

~ 

w· I <..9 _j 35.0±0.2 Ma <{ 

1-

z - 34 UJ 

a: 

<{ 

a.. 

a.. 

<{ 

32 -

I 

50 100 

PERCENT 39ArK RELEASED 

Figure 2. Composite age-spectrum diagram for nine mineral separates from six rocks, Keg Mountain, Utah, showing two 

statistically distinct groups of 40Arf39Ar plateau dates, the younger at 35.0±0.2 and the older at 36.6±0.1 Ma. The plateau 

dates of all samples within each group are statistically identical within 2cr analytical precision. Samples of the younger group 

inclu?e ~anidi.ne and biotite from the basal vitrophyre of the Joy Tuff (KMR-1-3) and sanidine and biotite from a hypabyssal 

rhyol1te mtrus1on (KMR-4-6). Samples of the older group include hornblende and biotite from the base of the Mt. Laird Tuff 

(KMR-4-8), biotite from the Mt. Laird Tuff (NKM-21-12), biotite from the basal vitrophyre of the tuff of Keg Mountain 

(KP-6-5), and biotite from a hypabyssal dacite porphyry intrusion (NKM-6-3). 

These corrections are well defined, but, because the mathematical formulation for F is complex, it is not presented here. 

(Our formulation ofF is modified from Dalrymple and others, 1981, to include corrections for all interferences.) 

We selected six samples from Keg Mountain that 

yielded nine mineral separates used for detailed highprecision 40 Arf39 Ar dating. The samples include hornblende and biotite from the basal Mt. Laird Tuff, biotite 

from the middle part of the Mt. Laird Tuff, biotite from 

the basal vitrophyre of the tuff of Keg Mountain, biotite 

from a hypabyssal dacite porphyry, sanidine and biotite 

from a hypabyssal rhyolite porphyry, and sanidine and 

biotite from the basal vitrophyre of the Joy Tuff. Minerals 

were separated using standard techniques and are greater 

than 99.9 percent pure. Detailed 40 Arf39 Ar data are listed 

in table 1, and a summary of the dates is given in table 2. 

A composite age-spectrum diagram for samples analyzed 

in this study is shown in figure 2, and sample localities 

are shown on figure 3 and included in table 1. 

The isotopic composition of argon was measured at the 

U.S. Geological Survey, Denver, Colorado, using a 

MAP215 Series rare-gas mass spectrometer made by Mass 

Analyzer Products Limited. Abundances of five isotopes of 

argon (40 Ar, 39 Ar, 38 Ar, 37 Ar, and 36 Ar) were measured for 

every sample. Argon was released from the samples in 

12-14 temperature steps; abbreviated results are listed in 

table 1. Radiogenic 40 Ar is total 40 Ar derived from natural 

radioactive decay of 40K after all corrections for non-decayderived 40Ar, including atmospheric 40Ar and 40K-derived 

40Ar, have been made. K-derived 39 Ar is total 39 Ar derived 

from the epithermal neutron-induced reaction 39K(n,p)39 Ar 

after corrections for non-39K-derived 39 Ar, including 42Caderived 39 Ar, are made. F is the quantity resulting from the 

division of radiogenic 40 Ar by the amount of K-derived 

39 Ar. Quantities for radiogenic 40 Ar and K-derived 39 Ar are 

given in volts of signal measured on a Faraday detector by a 

digital voltmeter. These quantities can be converted to 

moles, using the mass spectrometer sensitivity at time of 

40Ar/39Ar Geochronology, Keg Mountain, Utah G3 
Table 1. 40Ar/39Ar age-spectrum data for Keg Mountain, Utah 

[Sample localities shown in figure 3. 40 Ar R is radiogenic 40 Ar, 39 ArK is potassium-derived 39 Ar] 

Temperature 40ArR 39ArK Apparent age and error 

(

0

C) 40Ar R 39Ar K F 39Ar/37 Ar (percent) (percent) (Ma at 1a) 

Sample KMR-1-3/23/DDIJ~ Joy Tuff, basal vitrophyre~ sanidine~ 102.1 mg~ measured 40 Ar/36 Ar a =299.5~ 

J-value=0.007431±0.25% (Ia)~ lat 39°50'48.73" N., long 112°52'22.18" W. 

500 0.06512 0.02860 2.277 15.87 17.5 0.2 30.26±2.80 

600 0.21056 0.08308 2.534 30.96 78.7 0.5 33.66±0.82 

700 0.54243 0.20856 2.601 40.54 60.6 1.2 34.53±0.51 

750 1.03892 0.39621 2.622 45.76 93.3 2.2 34.81±0.15 

800 1.63759 0.62452 2.622 51.82 95.5 3.5 34.81±0.14 

850 2.36865 0.89993 2.632 61.46 93.9 5.0 34.94±0.18 

900 2.95837 1.12915 2.620 73.84 97.3 6.2 34.78±0.10 

950 5.24909 1.99358 2.633 93.42 98.2 11.0 34.96±0.10 

1,000 4.80260 1.82517 2.631 115.43 98.1 10.1 34.93±0.10 

1,050 6.93128 2.63075 2.635 131.75 98.2 14.5 34.98±0.10 

1,100 8.64343 3.28430 2.632 149.34 98.3 18.2 34.94±0.10 

1,150 8.99432 3.40018 2.645 156.17 98.1 18.8 35.12±0.10 

1,250 3.89366 1.46577 2.656 129.61 93.2 8.1 35.26±0.10 

1,400 0.33891 0.12319 2.751 69.78 47.3 0.7 36.51±0.38 

Total-gas date 2.635 34.98±0.11 

Plateau date (700-1, 1 00°C) 71.8 34.92±0.16 

Sample KMR-1-3/25/DDIJ~ Joy Tuff, basal vitrophyre; biotite~ 50.2 mg~ measured 40 Ar/36 Ar a =299.5~ 

J-value=0.007426±0.25% (Ia); lat 39°50'48.73"N., long ll2°52'22.18"W. 

500 0.03568 0.03102 1.150 19.73 

600 0.20935 0.14345 1.459 41.29 

700 1.26122 0.52197 2.416 94.19 

750 0.99149 0.38304 2.588 105.05 

800 0.77367 0.29221 2.648 106.61 

850 0.75824 0.28920 2.622 94.91 

900 1.35843 0.51814 2.622 118.87 

950 3.15754 1.20841 2.613 112.43 

1,000 4.25305 1.62684 2.614 84.75 

1,050 4.08619 1.54554 2.644 62.32 

1,150 1.27066 0.48007 2.647 11.74 

1,300 0.06602 0.02544 2.595 7.38 

Total-gas date 2.579 

Preferred date (850-1,150°C) 

measurement of 9. 736x 1 o-13 moles argon per volt of signal. 

The detection limit for argon at the time of this experiment 

was 2x10-17 moles. The measured 40Arf36Ar ratio used for 

mass discrimination correction is 299.5. 

All samples were irradiated in a single irradiation package for 30 hours at 1 megawatt in the U.S. Geological Survey 

TRIGA reactor in Denver. The ]-value for each sample was 

determined from adjacent standards; errors of 0.25 percent 

(68 percent confidence level) in the calculated ]-value were 

determined experimentally by calculating the reproducibility 

of multiple monitors. Corrections for irradiation-produced, 

interfering isotopes of argon were made by measuring 

production ratios for the interfering isotopes of argon produced in pure K2S04 and CaF2 irradiated simultaneously 

with the samples of this study. Those ratios, as determined 

from four measurements of each salt, with errors at 68-

percent confidence level, are 

11.9 0.4 

31.7 2.0 

60.7 7.4 

81.9 5.4 

85.1 4.1 

83.8 4.1 

84.0 7.3 

85.0 17.1 

88.3 23.0 

91.2 21.9 

91.0 6.8 

27.4 0.4 

80.2 

(36Arf37 Ar)ca =2.69±0.01x1~, 

(39 Arf37 Ar)ca =7.08±0.03x1~, 

(38Arf37 Ar)ca =2.90±0.18x1o-s, 

(40Arf39 Ar)K =8.76±0.37x10-3, 

15.34±2.61 

19.44±0.48 

32.08±0.19 

34.35±0.14 

35.13±0.20 

34.79±0.27 

34.78±0.16 

34.67±0.10 

34.69±0.10 

35.08±0.10 

35.11±0.13 

34.43±5.89 

34.22±0.17 

34.84±0.14 

(37 Arf39Ar)K =1.39±0.57x10-4, and 

(38Arf39 Ar)K =1.30±0.01x10-2• 

Corrections were made for additional interfering isotopes of 

argon produced from irradiation of chlorine using the 

method described by Roddick (1983). Measured quantities 

of 37 Ar and 39 Ar were corrected for radioactive decay, and 

in table 1 the 39 Arf37 Ar ratios were corrected for this decay 

as well as for interfering argon isotopes. By multiplying the 

39 Arf37 Ar ratios by approximately 0.5, the relative approximate K/Ca distribution of the samples may be obtained. 

Error estimates for apparent ages of individual temperature 

steps were assigned by using the equations of Dalrymple and 

others (1981); however, the equations were modified to 

G4 Application of Structural Geology to Energy and Mineral Resources 
Table 1. 40Ar/39Ar age-spectrum data for Keg Mountain, Utah-Continued 

[Sample localities shown in figure 3. 40 Ar R is radiogenic 40 Ar, 39 ArK is potassium-derived 39 Ar] 

Temperature 40 Ar R 39 ArK 

(°C) 40Ar 8 

39Ar K F 39Ar/37 Ar (percent) (percent) 

Sample KMR-4-6/22/DD13; Hypabyssal rhyolite porphyry; sanidine; 112.3 mg; measured 40 Ar/36 Ar a =299.5; 

J-value=0.007380±0.25% (lo); lat 39°48'3.40" N., long 112°52'26.95" W. 

600 0.38830 0.14676 2.646 5.29 72.3 0.8 

700 1.18169 0.44145 2.677 6.87 72.1 2.3 

750 1.01737 0.38392 2.650 63.83 95.8 2.0 

800 2.54557 0.95339 2.670 81.86 93.7 4.9 

850 2.75714 1.04152 2.647 105.60 97.3 5.3 

900 3.93424 1.48459 2.650 123.56 98.2 7.6 

950 5.32784 2.00846 2.653 137.12 98.3 10.3 

1,000 5.17896 1.95629 2.647 153.88 98.3 10.0 

1,050 6.54139 2.46716 2.651 156.27 98.1 12.6 

1,100 6.63168 2.49090 2.662 148.85 96.9 12.7 

1,150 14.51355 5.44169 2.667 151.64 91.8 27.8 

1,250 1.88149 0.69523 2.706 120.40 82.9 3.6 

1,400 0.13923 0.05151 2.703 50.40 31.9 0.3 

Total-gas date 2.660 

Plateau date (850-1,150°C) 86.3 

Sample KMR-4-6/27/DDB; Hypabyssal rhyolite porphyry; biotite; 62.1 mg; measured 40Ar/36Ar0 =2995; 

J-value=0.007409±0.25% (lo); lat 39°48'3.40" N., long 112°52'26.95" W. 

500 0.02788 0.02762 1.009 13.80 6.3 0.3 

600 0.21071 0.08909 2.365 15.45 34.1 1.0 

700 0.83796 0.29766 2.815 102.93 63.4 3.4 

750 0.78662 0.28854 2.726 232.15 87.9 3.2 

800 1.21031 0.45221 2.676 202.17 89.7 5.1 

850 0.94621 0.35555 2.661 157.91 88.3 4.0 

900 1.35650 0.50591 2.681 132.03 88.8 5.7 

950 3.42385 1.28393 2.667 162.36 87.6 14.5 

1,000 5.56171 2.09056 2.660 153.11 89.8 23.5 

1,050 5.71675 2.15001 2.659 77.66 90.9 24.2 

1,150 3.28766 1.23540 2.661 28.15 90.6 13.9 

1,300 0.27950 0.10264 2.723 21.83 51.4 1.2 

Total-gas date 2.663 

Plateau date (800-1, 150°C) 90.9 

Apparent age and error 

(Ma at 1o) 

34.89±0.35 

35.29±0.15 

34.94±0.18 

35.20±0.14 

34.90±0.10 

34.94±0.10 

34.98±0.10 

34.91±0.10 

34.96±0.10 

35.10±0.10 

35.16±0.10 

35.67±0.14 

35.63±1.21 

35.07±0.11 

35.04±0.15 

13.44±3.28 

31.34±0.60 

37.24±0.18 

36.08±0.58 

35.42±0.17 

35.22±0.17 

35.49±0.13 

35.30±0.13 

35.21±0.10 

35.19±0.10 

35.22±0.14 

36.04±0.66 

35.25±0.16 

35.25±0.13 

allow the option of choosing the larger of separately derived 

errors in the F-value--either a calculated error from differential equations or an experimental error determined from 

the reproducibility of identical samples. Age plateaus were 

determined by comparing contiguous gas fractions using the 

critical value test of Dalrymple and Lanphere ( 1969), and the 

error was determined using equations of Dalrymple and others (1981). 

40Arf39Ar Data 

Age spectra for the nine dated minerals are relatively 

uncomplicated and are summarized in table 2. The minor 

amounts of excess argon or argon loss displayed by most 

spectra had no effect on apparent age interpretations. Eight 

of nine spectra have apparent age plateaus defined by 70-99 

percent of the released 39 Ar from the sample. The age spectrum for biotite sample KMR-1-3 is slightly disturbed and 

displays minor excess argon but yields a preferred date, 

defined by weight-averaging the 850°C-1,150°C temperature steps, that is statistically indistinguishable from the plateau date for the coexisting sanidine. Two other samples 

from which two minerals were dated are KMR-4-6 and 

KMR-4-8. In both cases the coexisting minerals yielded statistically indistinguishable dates. Without discriminating 

data, there is no reason to select the apparent age of either 

mineral over the other in a pair; therefore, the best estimates 

of crystallization age for samples KMR-1-3, KMR-4-6, 

and KMR-4-8 are the averages of the dates for each pair; 

that is, 34.88±0.06, 35.14±0.15, and 36.58±0.06 Ma, respectively. 

Dates for the nine minerals fall into two statistically distinct groups (fig. 2), 36.6±0.1 and 35.0±0.2 Ma, that define 

two distinct periods of magmatism. The tuff of Keg Mountain, Mt. Laird Tuff, and a hypabyssal dacite porphyry were 

40ArP9Ar Geochronology, Keg Mountain, Utah GS 
Table 1. 40Ar/39Ar age-spectrum data for Keg Mountain, Utah-Continued 

[Sample localities shown in figure 3. 40 Ar R is radiogenic 40 Ar, 39 ArK is potassium-derived 3

9 Ar] 

Temperature 40 Ar 

(°C) 40Ar R 39Ar K F 39Ar/37 Ar (perce~t) Sample NKM--6-3/29/DD13~ Hypabyssal dacite porphyry, biotite~ 67.4 mg~ measured 40Ar;36Ar0 =299.5~ 

J-value=0.007395±0.25% (Ia)~ lat 39°49'24.53" N., long 112°53'34.35" W. 

500 0.06138 0.06406 0.958 15.16 9.4 

600 0.17083 0.10871 1.571 4.59 35.9 

700 1.50395 0.56855 2.645 78.93 54.6 

750 2.42808 0.86649 2.802 185.50 88.3 

800 5.52870 1.99433 2.772 320.71 96.1 

850 3.74547 1.34822 2.778 277.33 96.0 

900 1.82843 0.66352 2.756 100.62 87.6 

950 1.98268 0.72116 2.749 6.45 86.8 

1,000 2.05554 0.74390 2.763 9.71 89.0 

1,050 2.68882 0.97950 2.745 66.45 89.6 

1 '150 0.49706 0.18255 2. 723 14.64 82.7 

1,300 0.00887 0.00594 1.492 3.24 3.2 

Total-gas date 2.728 

Plateau date (800-1, 150°C) 

39ArK 

(percent) 

0.8 

1.3 

6.9 

10.5 

24.2 

16.3 

8.0 

8.7 

9.0 

11.9 

2.2 

0.1 

80.4 

Sample KP--{)-5/28/DD13~ Tuff of Keg Mountain, basal vitrophyre~ biotite~ 55.0 mg~ measured 40 Ar/36 Ar a =299.5~ 

J-value=0.007420±0.25% (la)~ lat 39°47'46.00" N., long 112°53'29.96" W. 

500 0.06530 0.03487 

600 0.11228 0.04253 

700 0.30849 0.11201 

750 0.86382 0.30429 

800 1.47215 0.52130 

850 1.49688 0.53668 

900 1.78396 0.64114 

950 2.62071 0.94540 

1,000 3.97124 1.43109 

1,050 5.74289 2.07489 

1,150 2.52579 0.90807 

1,300 0.02514 0.00969 

Total-gas date 

Plateau date (850-1,150°C) 

1.873 

2.640 

2.754 

2.839 

2.824 

2.789 

2.782 

2.772 

2.775 

2.768 

2.782 

2.595 

2.776 

12.80 

19.66 

43.96 

85.27 

136.63 

159.41 

230.13 

265.82 

88.47 

235.51 

140.48 

16.05 

11.8 

45.8 

41.9 

75.8 

90.3 

93.1 

94.1 

92.1 

90.8 

90.2 

89.0 

13.0 

0.5 

0.6 

1.5 

4.0 

6.9 

7.1 

8.5 

12.5 

18.9 

27.4 

12.0 

0.1 

86.4 

Apparent age and error 

(Ma at 1a) 

12.74±0.87 

20.84±0.81 

34.95±0.11 

37.00±0.12 

36.61±0.10 

36.69±0.10 

36.39±0.13 

36.31±0.14 

36.49±0.11 

36.25±0.17 

35.96±0.31 

19.79±8.48 

36.03±0.15 

36.49±0.15 

24.89±1.97 

35.00±1.70 

36.49±0.61 

37.60±0.13 

37.41±0.13 

36.95±0.16 

36.87±0.12 

36.73±0.13 

36.77±0.11 

36.67±0.10 

36.85±0.12 

34.40±7.45 

36.78±0.14 

36.77±0.12 

emplaced during the first episode. Within the resolution of 

the data for these three units, the differences in emplacement 

history as recognized in the field (see later section) are not 

distinguishable in the argon data. The Joy Tuff and a 

hypabyssal rhyolite porphyry plug were emplaced during the 

later magmatic episode. Again, within the resolution of the 

argon data, an isotopic age difference for these two units 

cannot be resolved. 

named the unit the Keg Spring andesite and latite, and Shawe 

(1972) and Staub (1975) referred to it as the Keg Spring 

andesite. The andesite of Keg Spring probably is a constructional volcanic high that was formed from multiple vents. 

Apparent interfingering of units indicates that eruption of the 

andesite of Keg Spring spanned the interval from pre-tuff of 

Dead Ox (late Eocene) to post-Mt. Laird Tuff (<36.54±0.06 

Ma) eruption. 

VOLCANIC STRATIGRAPHY AND 

IGNEOUS ROCKS 

Andesite of Keg Spring 

The oldest volcanic unit at Keg Mountain, based on field 

relations, is the andesite of Keg Spring (M.A. Shubat and 

G.E. Christenson, 1989, unpublished mapping). It consists 

of a heterogeneous sequence of dark-colored andesite flows 

and andesitic mudflow breccia. Erickson (1963) originally 

Tuff of Dead Ox 

The tuff of Dead Ox (M.A. Shu bat and G .E. Christenson, 

1989, unpublished mapping) consists of three intercalated 

lithostratigraphic units: (1) a stratified tuff member, (2) a 

megabreccia member, and (3) a lithic-crystal ash-flow tuff 

member. The stratified tuff consists of thin-bedded to laminated volcanic sandstone, siltstone, and ash-rich tuff. Lithiccrystal tuff is present as distinct beds and also as the matrix 

of megabreccia. Coarsely lithic portions (2~0 em diameter 

G6 Application of Structural Geology to Energy and Mineral Resources 
Table 1. 40Ar/39Ar age-spectrum data for Keg Mountain, Utah-Continued 

[Sample localities shown in figure 3. 40 Ar R is radiogenic 40 Ar; 39 ArK is potassium-derived 39 Ar) 

Temperature 40 Ar R 

(

0

C) 40Ar R 39Ar K F 39Art37 Ar (percent) 

Sample NKM-21-12/26/D013; Mt. Laird Tuff; biotite; 57.3 mg; measured 40Ar;36Ar0 =299.5; 

J-value=0.007430±0.25% (Ia); lat 39°48'55.99" N., long 112°55'51.93" W. 

500 0.02223 0.01927 1.154 10.58 5.8 

600 0.10302 0.04913 2.097 12.66 28.2 

700 1.19991 0.43847 2.737 46.37 50.4 

750 2.93219 1.06674 2.749 112.40 84.1 

800 3.96776 1.44129 2.753 190.38 89.8 

850 2.18086 0.78986 2.761 188.57 83.1 

900 0.99403 0.36113 2.753 68.38 68.1 

950 1.32051 0.48162 2.742 3.31 69.1 

1,000 1.68199 0.61257 2.746 1.11 73.3 

1,050 2.59559 0.94584 2.744 8.77 82.3 

1,150 1.32329 0.48045 2.754 3.54 77.0 

1,300 0.01956 0.00480 4.071 0.22 4.7 

Total-gas date 2.741 

Plateau date (700-1,150°C) 

Sample KMR-4-8/21/0013; Mt. Laird Tuff, base; hornblende; 310 mg; measured 40Ar;36Ar0 =299.5; 

J-value=0.007357±0.25% ( Ia); lat 39°51 '00.09" N., long 112°52'41.43" W. 

500 0.02351 0.01087 2.163 0.39 3.5 

700 0.21900 0.06423 3.409 1.29 24.3 

800 0.19170 0.06208 3.088 1.44 63.0 

850 0.10858 0.03586 3.028 1.47 63.7 

900 0.10127 0.03260 3.106 0.96 57.5 

950 0.38293 0.13478 2.841 0.24 65.6 

1,000 3.09179 1.11026 2.785 0.17 84.2 

1,050 5.34234 1.91651 2.788 0.16 90.2 

1,100 1.52856 0.54926 2.783 0.14 88.1 

1,150 0.68760 0.24725 2.781 0.13 85.1 

1,200 0.19654 0.07082 2.775 0.13 70.3 

1,400 0.17336 0.05950 2.914 0.14 13.3 

Total-gas date 2.806 

Plateau date (1,000-1,200°C) 

39ArK 

(percent) 

0.3 

0.7 

6.6 

15.9 

21.5 

11.8 

5.4 

7.2 

9.2 

14.1 

7.2 

0.1 

98.9 

0.3 

1.5 

1.4 

0.8 

0.8 

3.1 

25.9 

44.6 

12.8 

5.8 

1.6 

1.4 

90.7 

Apparent age and error 

(Ma at 1a) 

15.40±5.62 

27.89±1.26 

36.31±0.17 

36.47±0.10 

36.53±0.10 

36.63±0.13 

36.52±0.27 

36.38±0.11 

36.43±0.12 

36.41±0.10 

36.55±0.18 

53.75±9.61 

36.37±0.16 

36.48±0.14 

28.48±3.37 

44.69±1.11 

40.53±0.92 

39.74±2.90 

40.76±1.95 

37.32±0.56 

36.59±0.14 

36.62±0.10 

36.56±0.16 

36.54±0.30 

36.46±0.75 

38.26±0.79 

36.86±0.31 

36.59±0.29 

clasts) of the lithic-crystal tuff resemble "mesobreccia" 

described by Lipman (1976) and weather to form cobbleand boulder-strewn slopes with poorly exposed matrix. 

The megabreccia member of the tuff of Dead Ox consists 

of clasts of Paleozoic rocks, conglomerate, andesitic mudflow breccia, and andesite set in poorly welded lithic tuff 

matrix. Lithic-crystal tuff matrix is poorly exposed because 

of mantling of megabreccia by Quaternary Lake Bonneville 

lacustrine deposits but is seen as a thin rind of tuff coating 

megabreccia clasts. Clast sizes are from less than 0.2 to 240 

min diameter, but most clasts are between 3 and 60 m. Most 

quartzite and some limestone clasts in the megabreccia are 

pervasively brecciated. In places clasts are monolithologic, 

angular, matrix-supported breccia with a fine-grained matrix 

of comminuted material, whereas in other places clasts are 

intact but highly fractured and sheared. The former textures 

could have been produced in a subvolcanic environment by 

"gas-explosion" and hydrothermal activity, as reviewed by 

Shawe and Snyder (1988). The latter textures could represent blocks deformed during thrusting and later incorporated 

into the megabreccia. In contrast to the Paleozoic clasts, 

andesite clasts are only locally brecciated. The juxtaposition 

of surficially derived andesite (debris-flow facies) clasts and 

shattered Paleozoic clasts argues for an origin involving the 

mixing of eruptive megabreccia (Shawe and Snyder, 1988) 

and landslide megabreccia (Lipman, 1976). Such mixing 

could occur during an Aira-type eruption (Aramaki, 1984), 

during which a volcanic depression is filled with megabreccia and tuff and lacks distinct ring-faults. 

Field relations suggest that the tuff of Dead Ox underlies 

the tuff of Keg Mountain and predominantly overlies, but in 

part interfingers with, the andesite of Keg Spring. 

40Ar/39Ar Geochronology, Keg Mountain, Utah G7 
Table 1. 40Ar/39Ar age-spectrum data for Keg Mountain, Utah-Continued 

[Sample localities shown in figure 3. 40 Ar R is radiogenic 40 Ar, 39 ArK is potassium-derived 39 Ar) 

Temperature 40Ar R 39Ar K 

(

0

C) 40Ar R 39Ar K F 39ArP7 Ar (percent) (percent) 

Apparent age and error 

(Ma at 1a) 

Sample KMR-4-8/24/DD13~ Mt. Laird TufT, base~ biotite~ 78.2 mg~ measured 40Ar!36Ar =299.5~ 

J-value=0.7424±0.25% (lot lat 39°51'00.09" N., long 112°52'41.43" W. 

500 0.06137 0.05137 1.195 11.53 9.0 0.5 

600 0.12916 0.05911 2.185 19.41 42.0 0.6 

700 0.36972 0.13581 2.722 27.57 48.1 1.3 

750 1.06817 0.38646 2.764 55.93 76.4 3.7 

800 2.01958 0.72495 2.786 103.09 89.4 7.0 

850 2.59186 0.93159 2.782 139.57 92.8 9.0 

900 3.62307 1.30943 2.767 162.10 93.8 12.6 

950 5.48573 1.98479 2.764 45.77 92.1 19.2 

1,000 5.12590 1.86121 2.754 4.41 88.5 18.0 

1,050 4.66806 1.69934 2.747 18.34 87.3 16.4 

1,150 3.27674 1.18838 2.757 12.98 86.6 11.5 

1,300 0.08719 0.03137 2.780 1.93 37.6 0.3 

Total-gas date 2. 751 

Plateau date (900-1,150°C) 77.6 

Table 2. Summary of 40Ar/39Ar age-spectrum dates for Keg Mountain, Utah 

15.93±0.97 

29.03±1.54 

36.10±0.49 

36.64±0.14 

36.93±0.13 

36.88±0.12 

36.68±0.10 

36.64±0.10 

36.51±0.10 

36.42±0.10 

36.56±0.12 

36.85±2.69 

36.47±0.13 

36.56±0.11 

Sample no. Unit Mineral Date (Ma±1 a) Character of age spectrum 

KMR-1-3 Joy TufT, basal vitrophyre Sanidine 34.92±0.16 Plateau~ minor inherited argon. 

Biotite 34.84±0.14 Preferred date~ disturbed with minor argon loss. 

KMR-4-6 Hypabyssal rhyolite porphyry Sanidine 35.04±0.15 Plateau; minor excess argon. 

Biotite 35.25±0.13 Plateau; minor excess argon. 

NKM-6-3 

KP-6-5 

NKM-21-12 

KMR-4-8 

Hypabyssal dacite porphyry Biotite 36.49±0.15 Plateau; minor excess argon. 

TufTofKeg Mountain, basal vitrophyre Biotite 36.77±0.12 Plateau; minor excess argon. 

Mt. Laird TufT Biotite 36.48±0.14 Plateau. 

Mt. Laird TufT, base Hornblende 36.59±0.29 Plateau; minor excess argon. 

Biotite 36.56±0.11 Plateau; minor excess argon. 

Constraints on the minimum age of the unit are the observations that (1) the tuff of Dead Ox is cut by a dacite porphyry 

plug dated at 36.49±0.15 Ma (sample NKM-6-3) and (2) the 

megabreccia member contains no clasts of tuff of Keg 

Mountain or Mt. Laird Tuff. A constraint on the maximum 

age of the unit is the observation that clasts of the andesite of 

Keg Spring are common in the megabreccia member. 

Tuff of Keg Mountain 

The tuff of Keg Mountain (M.A. Shubat and G.E. 

Christenson, unpublished mapping, 1989) consists of darkcolored, densely welded, moderately crystal-rich ash-flow 

tuff of dacitic composition that has a local black vitrophyric 

base. Abundant bronze-weathering biotite, prominent on 

surfaces parallel with layering, characterizes the unit. As discussed above, our dates show a slightly older age for the tuff 

of Keg Mountain compared to the Mt. Laird Tuff but do not 

provide statistically unambiguous eruption ages. Field relations, discussed below, suggest that the tuff of Keg Mountain 

is older than the Mt. Laird Tuff. First, several dacite porphyry plugs, one of which we dated at 36.49±0.15 Ma (sample NKM-6-3, table 2), that are compositionally similar to 

the Mt. Laird Tuff cut the tuff of Keg Mountain. Second, 

there is a consistent stratigraphic succession from the tuff of 

Keg Mountain to Mt. Laird to Dell Tuff with increasing elevation in the north-central part of Keg Mountain; this stratigraphic superposition supports the suggested age sequence. 

Third, the tuff of Keg Mountain is absent south of an arcuate 

fault in the Picture Rock Hills Quadrangle (the northern edge 

of the Picture Rock caldera(?), fig. 3). Therefore, if the tuff 

of Keg Mountain is younger than the Mt. Laird Tuff, then its 

absence south of the fault requires confinement of the tuff of 

Keg Mountain to the Keg cauldron or stratigraphic thinning 

of the tuff of Keg Mountain from almost 200 m to zero over 

a distance of 2 km. If the tuff of Keg Mountain is older, then 

the Mt. Laird Tuff covers the tuff of Keg Mountain south of 

this fault. 

G8 Application of Structural Geology to Energy and Mineral Resources 
113°00' 

39°52'30" 

FLINT SPRINGS 

CAULDRON 

PICTURE ROCK HillS 

QUADRANGlE 

5 KILOMETERS 

39°37'30" L.__ _____________________ _j 

112°45' 

MINERAL OCCURRENCES 

* 

0 

A 

2 

3 

Volcanic-hosted gold 

Polymetallic vein 

Jasperoid-related 

Dome Hill 

Copper Hill 

Lead Hill 

\J 

KEG MOUNTAIN 

RANCH 

QUADRANGlE 

Alkali rhyolite, 6.7-10 Ma 

(Plavidal, 1987) 

Dell Tuff, 32.0±0.6 Ma 

(Lindsey, 1982) 

Joy Tuff (left) and rhyolite 

porphyry intrusions (right) 

Mt. Laird Tuff (left) and 

dacite porphyry intrusions (right) 

Quartz monzonite porphyry (left), 

36.6± 1.6 Ma (Lindsey, 1982), and 

other intrusions (right) 

Keg Tuff 

~~~~~CJ~a~on~" ~ t~ ~o~f'o. ~v<l Dead Ox Tuff; lithic-crystal 

tuff member (left) and megabreccia member (right) 

• • • • • • • 

Andesite of Key Spring 

Paleozoic rocks; upper plate 

(left) and lower plate 

Thrust fault 

High-angle fault 

.,.. .,.. .,.. .,.. Caldera/ cauldron boundary 

* Mt. Laird Tuff vent area? 

Figure 3. Simplified geology of Keg Mountain, west-central Utah. Solid numbered circles are locations of dated samples listed 

in table 2: 1, KMR-1-3; 2, KMR-4-6; 3, NKM-6-3; 4, KP-6-5; 5, NKM-21-12; 6, KMR-4-8. 

40Ar/39Ar Geochronology, Keg Mountain, Utah G9 
Quartz Monzonite Porphyry 

The quartz monzonite porphyry is a holocrystalline intrusive rock containing coarse (2-12 mm) phenocrysts of plagioclase, quartz, biotite, hornblende, pyroxene, and 

magnetite set in a fine-grained hypidiomorphic-granular 

matrix of quartz, plagioclase, and potassium feldspar. The 

porphyry forms a stock about 4.8 km long by 1.6 km wide in 

the south-central part of the Keg Pass Quadrangle (fig. 3). 

Lindsey and others (1975) dated the porphyry by the fissiontrack method at 36.6±1.6 Ma. 

Mt. Laird Tuff 

The Mt. Laird Tuff (Lindsey, 1979, 1982) is a regionally 

extensive unit consisting of moderately welded ash-flow 

tuff, tuff breccia, lapilli tuff, accretionary lapilli-block tuff, 

probable flow rocks, and hypabyssal intrusive rocks of rhyodacitic to quartz latitic composition. Its most distinctive 

characteristic is the presence of abundant, coarse (2-12 mm) 

phenocrysts of white plagioclase in all facies. Our dates indicate an extrusion age of 36.54±0.06 Ma for the Mt. Laird 

Tuff (samples KMR-4-8 and NKM-21-12, table 2), which 

is statistically identical to a 36.4±1.6 Ma date reported by 

Lindsey (1982). Because the fission-track date of the overlying Joy Tuff is about 38 Ma, Lindsey considered the true age 

of the Mt. Laird Tuff to be about 39 Ma. Over most of Keg 

Mountain, the Mt. Laird Tuff appears to be outflow from the 

Thomas caldera; however, mapping indicates an apparent 

Mt. Laird vent in the north-central part of Keg Mountain (fig. 

3). Rocks in this apparent vent consist of an accretionary 

lapilli-block tuff facies that grades into a hypabyssal porphyry facies; these are cut by pebble dikes containing fragmented and unfragmented Mt. Laird clasts. In addition, 

within the vent the tuff has steeply dipping layering. 

Dacite Porphyry Intrusive Rocks 

The hypabyssal dacite porphyry consists of abundant, 

coarse (2-10 mm) phenocrysts of plagioclase, quartz, 

biotite, hornblende, and magnetite set in a matrix of plagioclase, quartz, and potassium feldspar microphenocrysts and 

aphanitic material. Dacite porphyry forms several small 

plugs, each less than 900 min diameter, in the center of the 

Keg Pass Quadrangle (fig. 3). Our date of 36.49±0.15 Ma 

(sample NKM-6-3, table 2), as well as mineralogic and petrologic similarities, indicates that these rocks are comagmatic with the Mt. Laird Tuff. 

Regional aeromagnetic data (Kucks, 1991) show a circular, high-amplitude magnetic high (with a paired dipole low 

to the north) that straddles the northern part of the eastcentral edge of the Keg Pass Quadrangle (fig. 4). We interpret this anomaly to represent a buried dacite porphyry stock 

because of its proximity to several mapped dacite porphyry 

plugs and the possible Mt. Laird vent area described above 

(fig. 3). Magnetotelluric resistivity data (D. Campbell, oral 

common., 1990) support this interpretation. This interpretation has significance for exploration because anomalous 

amounts of gold are commonly detected in argillic-pyritic 

alteration halos around this type of intrusion. 

joy Tuff 

The Joy Tuff (Lindsey, 1979) is a regionally extensive 

unit that consists of moderately to densely welded, moderately crystal-rich, rhyolitic ash-flow tuff that locally has a 

black basal vitrophyre and overlying black fiamme-rich 

zone. Several exposures of the basal vitrophyre of the Joy 

Tuff resting on the Mt. Laird Tuff show that it clearly overlies the Mt. Laird Tuff; therefore, field relations are consistent with our date of 34.88±0.06 Ma (sample KMR-1-3, 

table 2). Our date shows a marked discordance with nine 

fission-track ages reported by Lindsey (1982) that average 

38.0±0.7 Ma. Although the improved accuracy of the 

40 Arf39 Ar age-spectrum dating method may be solely 

responsible for this discrepancy, it is possible that the unit 

mapped as Joy Tuff at Keg Mountain is lithologically similar 

to but younger than the Joy Tuff at the type locality. The 

exposed thickness of the Joy Tuff is thin ( <24 m) in most of 

the Keg Mountain area, but south of an arcuate fault (figs. 2, 

3) it abruptly thickens to more than 180m. Lindsey (1982) 

believed the source of the Joy Tuff to be within the Dugway 

Valley cauldron (fig. 4). 

Rhyolite Porphyry 

The rhyolite porphyry consists of coarse (as large as 1 

em) phenocrysts of orthoclase, quartz, plagioclase, and 

biotite set in an aphanitic matrix. Rhyolite porphyry forms 

plugs that lie within two linear, subparallel, northeast-trending belts, one of which is about 14 km long, that cut through 

the west-central part of Keg Mountain. Lithology and our 

date of35.14±0.15 Ma (sample KMR-4-6, table 2) for one 

of these plugs are similar to those of the Joy Tuff and suggest 

that the rhyolite porphyry unit is comagmatic with and a 

hypabyssal equivalent of the Joy Tuff. A fission-track age of 

30.8±1.8 Ma reported by Lindsey and others (1975) for 

another rhyolite porphyry plug raises the possibility that 

more than one age of rhyolite intrusions may be present. 

Figure 4 (facing page). Complete Bouguer gravity map of the 

north-central part of the Delta 1 °X2° quadrangle, Utah (from 

Bankey and Cook, 1989). Stippled areas are magnetic highs 

(from Kucks, 1991 ). Thick lines are known or inferred caldera 

margins as discussed in text. Contour interval 2 mGal. 

Gl 0 Application of Structural Geology to Energy and Mineral Resources 
10 KILOMETERS 

40 Ar j39 Ar Geochronology, Keg Mountain, Utah G 11 
Several lines of evidence suggest that a buried pluton, 

which we interpret to be a granitic body related to the rhyolite porphyry plugs exposed at the surface, underlies the Keg 

Pass area (figs. 2, 3). Deep resistivity (AMT) soundings collected along a north-south profile across Keg Mountain 

(Campbell and Visnyei, 1989) show a resistive body at depth 

beneath Keg Pass that has a resistivity signature (>200 ohmmeters) typical of many igneous rocks in the region (D. 

Campbell, oral commun., 1990). Regional aeromagnetic 

data (Kucks, 1991) show a high-amplitude magnetic ridge 

extending from Keg Pass to a point about 8 km to the south 

(fig. 4). This magnetic signature is similar to the signatures 

of many known intrusions in the Deep Creek-Tin tic mineral 

belt. 

Dell Tuff 

The Dell Tuff (Lindsey, 1979) is a pink to tan, poorly to 

moderately welded, crystal-rich, rhyolitic ash-flow tuff containing 2-8-mm phenocrysts of quartz, sanidine, plagioclase, 

and biotite. Staatz and Carr (1964) first described the tuff, 

and Lindsey (1979) formally named it for exposures in the 

Thomas Range. The minimum thickness of the Dell Tuff at 

Keg Mountain is 150 m. Recent paleomagnetic work by J.L. 

Hanna (University of Vermont, oral commun., 1989) shows 

that the Dell Tuff at the type locality ("The Dell," in the 

Thomas Range) has a normal magnetic polarity and that the 

Dell Tuff at Keg Pass has a reverse polarity; the Dell Tuff 

therefore consists of at least two cooling units. Lindsey and 

others (1975) dated three samples from Keg Pass, ranging in 

age from 32.5±1.6 to 33.8±1.3 Ma, and Lindsey (1982) 

established an average age of32.0±0.6 Ma for the formation. 

The source of the Dell Tuff has not been found; however, 

Lindsey (1982) speculated that the source should lie in or 

near Keg Mountain or the Thomas Range. 

Alkali Rhyolite 

Alkali rhyolite at Keg Mountain is present in two units, 

the lower unit is an unnamed rhyolite (Plavidal, 1987) and 

the upper unit is the Topaz Mountain Rhyolite (Lindsey, 

1979). Work by Plavidal (1987) shows that the two units are 

compositionally gradational; the Topaz Mountain Rhyolite 

was probably derived from the older unnamed rhyolite parent through crystal fractionation. Both units consist of flows, 

domes, and dikelike shallow intrusions of white, gray, and 

purplish alkali rhyolite. Phenocrysts in the unnamed rhyolite 

consist of coarse ( 1 em), abundant (20-40 percent) plagioclase and sanidine and lesser quartz and biotite. The Topaz 

Mountain Rhyolite contains sparse (10-15 percent), small (2 

mm) phenocrysts of quartz, sanidine, plagioclase, and 

biotite. Black vitrophyre is at the base of some flows and 

domes in both formations. Flows and domes probably were 

erupted from local vents, with dikelike bodies serving as 

feeders. Stratified tuff horizons are in both formations as discontinuous, wedge-shaped (less than 60 m thick) intercalations between flows. Stratified tuff is pale tan to orange and 

massive to thin bedded and consists of rhyolitic tuff and volcaniclastic sandstone. Stratified tuff was deposited as airfall, ash-flow, and ground-surge eruptions and as waterlain 

deposits (Lindsey, 1979; Bikun, 1980; Plavidal, 1987). The 

exposed thickness of alkali rhyolite at Keg Mountain is 450 

m. Plavidal (1987) reported K-Ar dates of 6.7±0.3 and 

6.9±0.3 Ma for the unnamed rhyolite. Fission track ages of 

7.8±0.6 and 8.2±0.5 Ma for the Topaz Mountain Rhyolite at 

Keg Mountain reported by Lindsey and others (1975) indicate a discrepancy in geochronology. 

STRUCTURE 

Mesozoic Compressional Structures 

Morris (1983, 1987) first identified in the northern Keg 

Mountain area and Slow Elk Hills a Sevier-age thrust that 

locally consists of several imbricate slivers. Exposed lower 

plate rocks of this thrust are probable Howell Limestone 

(Middle Cambrian) (Hintze and Robison, 1975), and upper 

plate rocks include the Prospect Mountain Quartzite, Pioche 

Formation, and Howell Limestone. The thrust may represent 

as little as 350 m or as much as 2,000 m of stratigraphic 

separation. Upper plate quartzite is pervasively brecciated 

throughout the area. Lower plate carbonate rocks have been 

locally altered to jasperoid immediately beneath the thrust. 

Ubiquitous white calcite veinlets in lower plate carbonate 

rocks may be related to the thrusting event. Axes of smallamplitude ( <60 m), upright, map-scale folds and minor folds 

(east vergent) in lower plate carbonate rocks plunge gently 

toward the north-northeast. The thrust surface is sharp and 

well exposed in many locations, consisting of a thin ( <4 em) 

cataclasite zone and no ductile deformation on the footwall. 

Caldera and Cauldron Structures 

Flint Spring Cauldron 

The source of the tuff of Dead Ox remains unidentified, 

but our interpretation of an eruptive origin (or at least an 

eruptive component) for the megabreccia member of the tuff 

of Dead Ox suggests a nearby vent. We tentatively propose 

that the tuff of Dead Ox erupted from the Flint Spring cauldron (M.A. Shubat and G.E. Christenson, unpublished data, 

1989) in the northwest part of the Keg Pass Quadrangle (figs. 

3, 4). Two lines of evidence support this proposition: (1) the 

presence of a subtle gravity low in this area and (2) the 

absence of a mappable caldera margin (structural or topographic wall) bounding the megabreccia terrain. The absence 

G12 Application of Structural Geology to Energy and Mineral Resources 
of a thick section of the tuff of Dead Ox, a pronounced gravity low near the megabreccia terrain, or recognizable ring 

faults argues against a large local source caldera for the tuff 

of Dead Ox. 

Keg Cauldron 

To date, the tuff of Keg Mountain has only been recognized at Keg Mountain and therefore is probably a locally 

derived unit. We propose that it was erupted from a cauldron, 

termed the Keg cauldron (M.A. Shubat and G.E. Christenson, unpublished data, 1989), centered on the quartz monzonite porphyry stock in the south-central part of the Keg Pass 

Quadrangle (figs. 3, 4). As defined here, the Keg cauldron 

differs substantially from the Keg caldera defined by Shawe 

(1972). The term "caldera" does not apply because of the 

absence of clearly defined margins. Evidence for the existence of the Keg cauldron is the thickening of the tuff of Keg 

Mountain from zero in the center of the Keg Pass Quadrangle to more than 165 m only 3 km to the south. The distribution of, and dips measured in, the tuff of Keg Mountain 

outlines a broad, low-relief dome centered on the quartz 

monzonite porphyry stock. Because of similarities in modal 

composition, age, and spatial distribution, we consider this 

stock to be comagmatic with the tuff of Keg Mountain and 

view it as a resurgent intrusion that caused doming. 

Thomas Caldera 

We place the eastern edge of the Thomas caldera along 

the western margin of Keg Mountain (figs. 3, 4), as originally drawn by Shawe (1972) and Lindsey (1982), for stratigraphic and geophysical reasons. We interpret the Mt. Laird 

Tuff at Keg Mountain to be a thin outflow facies of the formation; thus the caldera margin must lie west of Keg Mountain. Detailed mapping shows that where top and bottom 

contacts are exposed only thin sections of Mt.Laird Tuff are 

present between the tuff of Keg Mountain and the Joy Tuff 

and that over the center of the proposed Keg cauldron the Mt. 

Laird Tuff is missing. We found no stratigraphic or structural 

evidence indicating that the margin of the Thomas caldera 

passes through or east of Keg Mountain. Based on geophysical evidence, a likely place to draw the eastern margin of the 

Thomas caldera is a persistent gravity gradient along the 

western margin of Keg Mountain (fig. 4); in this case the 

gravity low of Dugway Valley is interpreted to represent, in 

part, less dense intracaldera fill. The absence of late Tertiary 

to Quaternary fault scarps along the western margin of Keg 

Mountain argues against this gradient being only the result 

of basin-range faulting. This gradient forms the eastern margin of an oval-shaped, north-elongate gravity low (fig. 4) 

that we associate with the Thomas caldera. The closure of 

this gravity anomaly at its northern and southern ends and 

the presence of bedrock near the axis of the low suggest control by rocks other than basin-fill sediments. The gravity low 

may represent subvolcanic intrusion of Miocene Topaz 

Mountain Rhyolite in addition to fill of the Thomas caldera. 

Picture Rock Caldera 

ARCO geologists identified and named the Picture Rock 

caldera while exploring for minerals in the Thomas Range 

and Keg Mountain area during the late 1970's and early 

1980's (Anaconda Geological Document Collection, University of Wyoming, Laramie). The caldera is primarily 

defined by an oval-shaped gravity low (fig. 4) that is coincident with low-resistivity material (Campbell and Visnyei, 

1989) that could be intracaldera fill. An arcuate fault, concave to the south, mapped in the north-central part of the Picture Rock Hills Quadrangle (fig. 3), separates a thick section 

(> 180 m) of Joy Tuff to the south from thin sections (<24m) 

to the north. We interpret this fault as the structural wall of 

the Picture Rock caldera and associate the caldera with the 

eruption of the Joy Tuff. Insufficient data exist to determine 

the relationship between the Picture Rock caldera and the 

Dugway Valley cauldron. As discussed above, it is possible 

that rocks mapped as Joy Tuff in the Thomas Range and at 

Keg Mountain may be distinct units, in which case the Picture Rock caldera would be the source for the Joy-like ashflow tuff exposed in the Picture Rock Hills. 

High-Angle Normal Faults 

Many small-displacement, high-angle normal faults at 

Keg Mountain cut units as young as the 6-Ma Topaz Mountain Rhyolite (Plavidal, 1987) but do not cut Quaternary 

units. Well-exposed field relations and geochronology 

tightly constrain the onset of at least some of the high-angle 

faulting. One north-striking fault in the center of the Keg 

Pass Quadrangle drops the base of the tuff of Keg Mountain 

15 m down to the west and is filled by a dacite porphyry 

plug. Our dates for the tuff of Keg Mountain and one of the 

dacite porphyry intrusions indicate that the age of this fault 

is between 36.77±0.12 and 36.49±0.15 Ma. High-angle 

faults at Keg Mountain are in all orientations, and the 

amounts of stratigraphic separation across these faults, 

where demonstrable, are small ( <60 m). Apparently, these 

faults formed throughout the period of volcanism from the 

early Oligocene to Pliocene. Examination of slickensides 

surfaces of some faults yielded a normal sense of slip. 

Although small in displacement and sparsely distributed, 

high-angle faults exert a strong control on the localization of 

mineral deposits at Keg Mountain. 

MINERAL DEPOSITS 

Morris (1978) first described pyritic alteration in the Keg 

Mountain area, and since then lead, silver, copper, and gold 

40ArP9Ar Geochronology, Keg Mountain, Utah G13 
have been discovered (fig. 3). These deposits form three 

main categories: volcanic-hosted gold, sediment-hosted 

polymetallic (lead-zinc-silver dominated) vein and replacement, and carbonate-hosted jasperoid prospects. 

Volcanic-Hosted Gold Deposits 

The most explored volcanic-hosted gold deposit at Keg 

Mountain is about 1 km west of the Copper Hill prospect 

(fig. 3). Industry geologists discovered the prospect in 1987 

and explored it with 14 drill holes in 1988 and 1989. Gold is 

concentrated along a high-angle, northwest-striking, southwest-dipping normal fault that juxtaposes the Mt. Laird Tuff 

(hanging wall) against the lithic-crystal tuff member of the 

tuff of Dead Ox. Within a few meters of the fault the tuff of 

Dead Ox is intensely argillized and pyritized. The highest 

values obtained to date average 1.02 ppm Au over a 6-m 

interval. Other areas in which gold is anomalous (fig. 3) are 

adjacent to pebble dikes cutting the Mt. Laird Tuff, near contacts of dacite porphyry plugs, and along high-angle faults. 

Dome Hill Prospect 

The Prospect Mountain Quartzite hosts a polymetallic 

vein and replacement deposit (Cox and Singer, 1986) at 

Dome Hill (fig. 3). Dome Hill constitutes the southwestern 

half of a dome that is cut by several high-angle faults, the 

most prominent of which consists of a silicified rib along the 

elongate northwest-trending axis of the hill. On the surface, 

host rocks are quartzite and lesser interbedded shale; in core, 

carbonate rocks also form the host. The mineral deposit at 

Dome Hill consists of highly oxidized stockwork and 

replacement with anomalous concentrations of silver, lead, 

copper, gold, and bismuth. Primary sulfide minerals are 

pyrite, galena, tetrahedrite-tennantite, and rare chalcopyrite. 

Analyses of mineralized samples show high concentrations 

of silver (as much as 27 oz/ton in bulk samples) and bismuth 

(as much as 3.7 percent) and locally high values of gold (6.5 

ppm), base metals, arsenic, and antimony. 

Faults and fractures control the exposed distribution of 

mineralized rock at Dome Hill, and highly anomalous samples are along the dominant northwest-trending fault along 

the crest of the hill. Mineralized rock apparently extends for 

more than 900 m to the northwest of the hill to an area of buried mineralized rock discovered by drilling (fig. 3). Drill 

holes in this area tested a northeast-trending self-potential 

responder coincident with northeast-striking air-photo lineaments. Drill core shows the presence of quartz veinlets in 

carbonate host rock containing pyrite, galena, and sphalerite 

and several intercepts of disseminated sulfides in silicified 

breccia zones. One stockwork zone contains 0.70 oz/ton Ag, 

0.55 percent Pb, and 0.54 percent Zn over a 3-m interval. 

Lead Hill Prospect 

Two types of mineralized rock are present at Lead Hill: 

(1) carbonate-hosted gold-, silver-, zinc-, and molybdenumbearing jasperoid and (2) small, massive pods of galena 

replacing limestone. Jasperoid is along high-angle faults and 

fracture zones and forms resistant, dark-brown masses about 

2-9m long and 1-3m wide. Sparse ore minerals injasperoid 

are fine-grained pyrite and galena. Gangue minerals are 

almost entirely microcrystalline, chalcedonic silica with a 

ubiquitous iron oxide stain and trace of sericite. Geochemical analyses show that the lead, silver, molybdenum, gold, 

and zinc are contained in the jasperoid. The highest values 

obtained from jasperoid samples are 1.89 percent Pb, 7.88 

oz/ton Ag, 450 ppm Mo, 0.59 ppm Au, and 280 ppm Zn. The 

second type of mineralized rock at Lead Hill is small ( <30 

em diameter) pods of almost massive galena that replaced 

limestone adjacent to a granodiorite dike exposed in a 

prospect trench at the north end of Lead Hill. Analyses of the 

galena-rich pods also show the presence of silver (0.66 

oz/ton) and trace amounts of zinc. 

CONCLUSIONS 

The combination of detailed geologic mapping and highprecision 40 Arf39 Ar geochronology allowed us to identify 

the following sequence of Eocene and Oligocene calderaand cauldron-forming events at Keg Mountain (in order of 

decreasing age): eruption of the tuff of Dead Ox from the 

Flint Spring cauldron, the tuff of Keg Mountain from the 

Keg cauldron, the Mt. Laird Tuff from the Thomas caldera, 

and the Joy Tuff (or a Joy-like tuff) from the Picture Rock 

caldera. Geochronology provides a temporal link between 

the petrographically similar Mt. Laird Tuff and hypabyssal 

dacite porphyry intrusive rocks, which we take as evidence 

for a comagmatic origin. Correlation of hypabyssal dacite 

intrusive rocks with the Mt. Laird Tuff implies that the area 

affected by this magmatic event is much larger than the areal 

extent of the Thomas caldera (about 300 km2). The distance 

spanned from the dacite intrusions at Keg Mountain to similar intrusions in the Drum Mountains, on the opposite side of 

the Thomas caldera, is more than 35 km. 

Work in the Simpson Mountains (fig. 1) by Yambrick 

and Snee (1989) resulted in discovery of rocks of similar 

lithology and age (36.5±0.2 Ma) to the Mt. Laird-related 

rocks described here. Y ambrick and Snee also dated a rhyodacitic ash-flow tuff at 35.1±0.2 Ma, indistinguishable in age 

from Joy-related rocks dated in this study at 35.0±0.2 Ma. 

(In addition, R.A. Y ambrick and L.W. Snee dated the rhyolite of Judd Creek from the Simpson Mountains and determined that it was extruded at 35.9±0.2 Ma, a time when 

volcanism apparently did not occur at Keg Mountain.) These 

similarities suggest that high-precision 40 Arf39 Ar age-spectrum dating can be used in conjunction with petrochemistry 

G14 Application of Structural Geology to Energy and Mineral Resources 
and petrography to correlate volcanic and subvolcanic rocks 

between adjacent ranges. For example, at 36.5±0.2 Ma intermediate rocks similar in composition to the Mt. Laird Tuff 

were erupted and intruded in the area from th~~ Drum to 

Simpson Mountains and thus affected an area as much as 60 

km in one dimension. Assuming the other dimension of this 

tract to be at least 10 km yields a minimum area of 600 km2, 

twice that of the Thomas caldera. Perhaps of equal or more 

significance, Keg Mountain and the Simpson Mountains 

were affected by volcanic activity that produced rocks of 

similar composition (rhyolite, rhyodacite and quartz latite) 

during the same time interval (36.5-35.0 Ma). 

In a gross sense, mineral deposits at Keg Mountain probably are associated with the eastern part of a large, late 

Eocene to early Oligocene igneous complex thalt spans the 

Thomas Range, Keg Mountain, and northern Drum Mountains area. These mineral deposits consist of volcanic-hosted 

gold prospects, polymetallic veins, and precious-metalsenriched jasperoid bodies that are along high-angle, smalldisplacement normal faults. In detail, however, there are no 

clear spatial relationships between inferred caldera margins, 

associated hypabyssal intrusions, and mineral deposits (fig. 

3). The widespread distribution of hypabyssal intrusions 

related to the Mt. Laird Tuff and Joy Tuff magmatic systems 

suggests the possibility of concealed, larger intrusions 

related to these systems. Mineral deposits at Keg Mountain 

may have been formed by hydrothermal systems induced by 

possible concealed intrusions. 

REFERENCES CITED 

Alexander, E.C., Michelson, G.M., and Lanphere, M .. A., 1978, A 

new 40 Art39 Ar dating standard: U.S. Geological Survey 

Open-File Report 78-701, p. 6-8. 

Aramaki, Shigeo, 1984, Formation of the Aira caldera, southern 

Kyushu, -22,000 years ago: Journal of Geophysical 

Research, v. 89, no. BlO, p. 8485-8501. 

Bankey, Viki, and Cook, K.L., 1989, Complete Bouguer gravity 

map and related geophysical maps of the Delta 1 °X2° 

quadrangle, Utah: U.S. Geological Survey Miscellaneous 

Field Studies Map MF-2081-A, scale 1:250,000. 

Bikun, J. V., 1980. Fluorine and lithophile element mineralization at 

Spor Mountain, Utah: Tempe, Arizona State University, 

M.S. thesis, 195 p. 

Campbell, D.L., and Visnyei, Andrew, 1989, Audiomagnetotelluric soundings in the Delta Utah 1 °X2° 

quadrangle: U.S. Geological Survey Open-File Report 

89-538, 71 p. 

Cox, D.P., and Singer, D.A., 1986, eds., Mineral deposit models: 

U.S. Geological Survey Bulletin 1693, 379 p. 

Dalrymple, G.B., Alexander, E.C., Lanphere, M.A., and Kraker, 

G.P., 1981, Irradiation of samples for 40 Art39 Ardating using 

the Geological Survey TRIG A reactor: U.S. Geological 

Survey Professional Paper 1176, 56 p. 

Dalrymple, G.B., and Lanphere, M.A., 1969, Potassium-argon 

dating: San Francisco, W.H. Freeman, 251 p. 

Erickson, M.P., 1963, Volcanic geology of western Juab County, 

Utah, in Sharp, B.J., and Williams, N.C., eds., Beryllium and 

uranium mineralization in western Juab County, Utah: Utah 

Geological Society, Guidebook to the Geology of Utah, no. 

17, p. 23-35. 

Hintze, L.F., and Robison, R.A., 1975, Middle Cambrian 

stratigraphy of the House, Wah Wah, and adjacent ranges in 

western Utah: Geological Society of America Bulletin, v. 86, 

p. 881-891. 

Kucks, R.P., 1991, Aeromagnetic and related maps of the Delta 

1 °X2° CUSMAP quadrangle, Utah: U.S. Geological Survey 

Miscellaneous Field Studies Map MF-2081-B, 7 p., scale 

1:250,000. 

Lindsey, D.A., 1977, Epithermal beryllium deposits in waterlaid 

tuff, western Utah: Economic Geology, v. 72, no. 2, p. 

219-232. 

___ 1979, Geologic map and cross sections of Tertiary rocks in 

the Thomas Range and northern Drum Mountains, Juab 

County, Utah: U.S. Geological Survey Miscellaneous 

Investigations Series Map 1-1176, scale 1:62,500. 

___ 1982, Tertiary volcanic rocks and uranium in the Thomas 

Range and northern Drum Mountains, Juab County, Utah: 

U.S. Geological Survey Professional Paper 1221, 71 p. 

Lindsey, D.A., Naeser, C.W., and Shawe, D.R., 1975, Age of 

volcanism, intrusion, and mineralization in the Thomas 

Range, Keg Mountain, and Desert Mountain, western Utah: 

U.S. Geological Survey Journal of Research, v. 3, no. 5, p. 

597-604. 

Lipman, P.W., 1976, Caldera-collapse breccias in the western San 

Juan Mountains, Colorado: Geological Society of America 

Bulletin, v. 87, p. 1397-1410. 

Morris, H.T., 1978. Unprospected zone of pyritic alteration in westcentral Utah: U.S. Geological Survey Open-File Report 

78-791, 8 p. 

___ 1983, Interrelations of thrust and transcurrent faults in the 

central Sevier orogenic belt near Leamington, Utah, in 

Miller, D.M., Todd, V.R., and Howard, K.A., eds., Tectonic 

and stratigraphic studies in the Great Basin: Geological 

Society of America Memoir 157, p. 75-81. 

___ 1987, Generalized geologic map of the Delta 2° 

quadrangle, Tooele, Juab, Millard, and Utah Counties, Utah: 

U.S. Geological Survey Open-File Report 87-185, scale 

1:250,000. 

Plavidal, K.R., 1987, The geology and petrology of the eastern Keg 

Mountains, Juab County, Utah: Salt Lake City, University of 

Utah, M.S. thesis, 137 p. 

Roddick, J.S., 1983, High precision intercalibration of 40Art39 Ar 

standards: Geochimica et Cosmochimica Acta, v. 47, p. 

887-898. 

Samson, S.D., and Alexander. C.E., 1987, Calibration of the 

interlaboratory 40 Art39 Ar dating standard, MMhb-1: 

Isotope Geoscience, v. 66, p. 27-34. 

Shawe, D.R., 1972, Reconnaissance geology and mineral potential 

of the Thomas, Keg, and Desert calderas, central Juab 

County, Utah: U.S. Geological Survey Professional Paper 

800-B, p. B67-B77. 

Shawe, D.R., and Snyder, D.B., 1988, Ash-flow eruptive 

megabreccias of the Manhattan and Mount Jefferson 

Calderas, Nye County, Nevada: U.S. Geological Survey 

Professional Paper 1471, 28 p. 

Shawe, D.R., and Stewart, J.H., 1976, Ore deposits as related to 

tectonics and magmatism, Nevada and Utah: American 

Institute of Mining Engineers, Annual Meeting, Las Vegas, 

Nevada, 1977, Transactions, v. 260, p. 225-232. 

Staatz, M.H., and Carr, W.J., 1964, Geology and mineral deposits 

of the Thomas and Dugway Ranges, Juab and Tooele 

40ArP9Ar Geochronology, Keg Mountain, Utah G15 
Counties, Utah: U.S. Geological Survey Professional Paper 

415, 188 p. 

Staub, A.M., 1975, Geology of the Picture Rock Hills quadrangle, 

Juab County, Utah: Salt Lake City, University of Utah, M.S. 

thesis, 87 p. 

Steiger, R.H., and Jager, E., 1977, Subcommission on 

geochronology; convention on the use of decay constants in 

geo- and cosmochronology: Earth and Planetary Science 

Letters, v. 36, p. 359-354. 

Stewart, J.H., Moore, W.J., and Zietz, Isidore, 1977, East-west 

patterns of Cenozoic igneous rocks, aeromagnetic 

anomalies, and mineral deposits, Nevada and Utah: 

Geological Society of America Bulletin, v. 88, no. 1, p. 

67-77. 

Yambrick, R.A., and Snee, L.W., 1989, Early Oligocene calcalkaline and middle Miocene bimodal volcanism, Simpson 

Mountains, west-central Utah: Geological Society of 

America Abstracts with Program, v. 21, no. 7, p. A245. 

G16 Application of Structural Geology to Energy and Mineral Resources 
Chapter H 

Structural Setting of the Chief Mining District, 

Eastern Chief Range, Lincoln County, Nevada 

By PETER D. ROWLEY, LAWRENCE W. SNEE, 

HARALD H. MEHNERT, R. ERNEST ANDERSON, 

GARY J. AXEN, KELLY J. BURKE, F. WILLIAM SIMONDS, 

RALPH R. SHROBA, and STEPHEN D. OLMORE 

U.S. GEOLOGICAL SURVEY BLILLETIN 2012 

APPLICATION OF STRUCTURAL GE<JLOGY TO MINERAL AND 

ENERGY RESOURCES OF THE CENTRAL AND WESTERN UNITED STATES 
CONTENTS 

Abstract Hl 

Introduction Hl 

History and production H3 

Geology H6 

Stratigraphy H6 

Intrusive rocks H8 

Caliente caldera complex H9 

Structure HlO 

Mineral belts HlO 

Pre-upper Oligocene detachment faults HlO 

Miocene faults Hll 

Upper Miocene and Pliocene basin-range faults H12 

Ore deposits H13 

Conclusions H14 

References cited H15 

FIGURES 

1. Photograph showing mines in upper Cobalt Canyon, Chief Mining District, 

Nevada H3 

2, 3. Maps showing: 

2. Tectonic setting of Chief Mining District H6 

3. Generalized geology of Chief Mining District and vicinity H7 

TABLE 

1. Chemical analyses of grab samples of hydrothermally altered and mineralized rocks 

from the Chief Mining District, Lincoln County, Nevada H4 

H II Contents 
Structural Setting of th1e Chief Mining District, 

Eastern Chief Range, Lincoln County, Nevada 

By Peter D. Rowley1

, Lawrence W. Snee2, Harald H. Mehnert2, R. Ernest Anderson3, 

Gary j. Axen4, Kelly j. Burke5, F. William Simonds 1, Ralph R. Shroba 1, and Stephen D. Olmore6 

Abstract 

The Chief Mining District produced at least 2,000 ounces 

of gold, 11,000 ounces of silver, and minor amounts of lead 

and copper between 1870 and the 1960's. The production was 

from fissure-vein precious-metal deposits containing highly 

oxidized mineral aggregates within breccia zones of mostly 

low-angle faults. The main veins consist of scorodite and 

related minerals that replaced arsenopyrite and of barite, 

galena, quartz, and iron- and manganese-oxide minerals. The 

deposits probably represent the lower part of an epithermal 

gold-silver system that was subsequently deeply eroded. The 

likelihood of finding additional precious-metal economic orebodies is not good, but the district has potential for a porphyrycopper deposit. 

Most veins are in Proterozoic and Cambrian quartzite 

(Stirling Quartzite, Wood Canyon Formation, Zabriskie 

Quartzite) and Cambrian limestone (Highland Peak Formation). A large shallow intrusive body of quartz monzonite porphyry, the Cobalt Canyon stock, underlies the district and was 

the source of the mineralizing fluids, perhaps from convective 

cells of heated meteoric water. 40ArJ39Ar dates for minerals in 

the stock average 24.8 Ma; the stock may represent an early 

phase of igneous activity in the Caliente caldera complex, 

whose boundary is about 4 km south of the mining district. 

The Chief Mining District is near the western end of the 

east-trending Delamar-lron Springs mineral belt. This belt may 

have resulted from igneous activity along early Tertiary eaststriking faults, which also may have defined othe1' east-striking 

1

U.S. Geological Survey, Box 25046, MS 913, Denver, Colorado 

80225. 2

U.S. Geological Survey, Box 25046, MS 963, Denv.;:r, Colorado 

80225, 3

U .S. Geological Survey, Box 25046, MS 966, Denv,er, Colorado 

80225. 4

Department of Earth and Planetary Science, Harvard University, 

Cambridge, Massachusetts. 5

Department of Geology, Northern Arizona University, Flagstaff, 

Arizona. 

lV.s. Geological Survey, Ciudad Bolivar, Venezuela. 

structural features in the area including the Timpahute Lineament, which intersects the mining district. Other epithermal 

gold districts in and near the Caliente caldera complex also 

contain east-striking features that suggest that younger (reactivated) faulting, igneous activity, and mineralization were all 

locally controlled by east-striking faults. 

Recent detailed geologic mapping demonstrates at least 

three episodes of Tertiary extensional faulting in the district. 

The oldest episode of faulting was along the low-angle Stampede Detachment Fault Zone, a major east-dipping structure of 

pre-late Oligocene age (pre-32 Ma), with upper plate movement to the east, that separates Zabriskie Quartzite in the lower 

plate from Highland Peak Formation in the upper plate. Faults 

of this episode provided sites for fissure-vein orebodies. The 

second episode of extensional faulting took place in the 

Miocene (by at least 19 Ma and continuing to less than 12 Ma) 

and is characterized by high-angle strike-slip, oblique-slip, and 

normal faults and, outside the district, by detachment faults. An 

early phase of this episode may have controlled emplacement 

of the Cobalt Canyon stock and thus epithermal mineralization. The youngest episode, basin-range faulting, formed much 

of the present topography; the fau Its strike north, postdate 

basalts of 12 Ma, and also cut basin-fill sedimentary rocks of 

the Panaca Formation (as young as 3 Ma) and Quaternary 

deposits. 

INTRODUCTION 

The Chief Mining District in Lincoln County, Nevada, 

also known as the Caliente Mining District, produced at least 

2,000 ounces of gold, 11,000 ounces of silver, and minor 

amounts of lead and copper between 1870 and the 1960's. 

The metals came from a cluster of mines (fig. 1) in a 6-km2 

area on the eastern side of the Chief Range, about 8 km north 

of Caliente, Nevada. Ore bodies are in quartzite of Late Proterozoic and Early Cambrian age and limestone and dolomite 

of Early and Middle Cambrian age; the ore resulted from 

emplacement of a quartz monzonite porphyry stock of 

Structural Setting of Chief Mining District, Nevada Hl 
Oligocene age. This report summarizes the structural control on the orebodies based on data collected during detailed 

bedrock mapping. It also includes preliminary 40 Arf39 Ar 

dates that constrain the timing of faulting and mineralization. 

The Chief Mining District is in the eastern Great Basin 

(fig. 2) at the western end of the east-trending Delamar-Iron 

Springs mineral belt ofShawe and Stewart (1976). The Chief 

Range is a basin range bounded on the east by the deep 

Panaca fault-block basin. The main locus of igneous activity 

in the area is the large Tertiary Caliente caldera complex of 

Nevada and Utah; the northern rim of this complex is as 

close as 4 km to the mining district. 

Prior to our studies, the only geologic information published on the Chief Mining District was a report based on a 

reconnaissance study of the ore bodies by a master geologist, 

Eugene Callaghan, that was written in 1936 while production in the district was declining. As far as we know, production ceased in the 1960's and the mining district was 

subsequently abandoned; few buildings remain standing 

(fig. 1 ), and, because most shafts, adits, and tunnels have collapsed, few orebodies are accessible. Even where freshly 

exposed, most orebodies are extensively oxidized and the 

ore minerals are complex hydrous oxides that are difficult to 

identify without extensive microscopic and geochemical 

study, which we have not done. Thus we cannot contribute 

much additional information on the orebodies, except for 

chemical analyses of grab samples (table 1). Much later than 

Callaghan's study, the geologic setting of the area was established by milestone reconnaissance 1 :250,000-scale geologic maps of Lincoln County, Nevada, by Tschanz and 

Pampeyan (1970) and Ekren and others (1977). 

The new data in this report result from two 1:24,000-

scale geologic mapping projects. The first, beginning in 

1985, involved structural studies by G.J. Axen (University of 

Northern Arizona and Harvard University) and his students 

in the Highland Range and nearby areas about 10-15 km 

north of the Chief Mining District and their collaboration 

with J.M. Bartley (University of Utah), W.J. Taylor (University of Utah), and D.R. Lux (University of Maine at Orono) 

who worked in the North Pahroc Range, southern Schell 

Creek Range, and southern Fairview Range, about 30 km 

west and north-northwest of the mining district. The purpose of this work was to define geometry, timing, and mechanisms of Tertiary extension in the area, the products of 

which include two major detachment fault zones, the older 

Stampede Detachment and the younger Highland Detachment (Axen and others, 1988). In order to understand these 

fault zones, the Paleozoic stratigraphy of the area was 

refined. Mapping progressed southward, the most recent 

being that of Burke ( 1991) in the northern Chief Range. The 

second project, beginning in 1986, involved quadrangle 

mapping by the U.S. Geological Survey (USGS) near Caliente in an effort to understand the evolution of the Caliente 

caldera complex. This project was later incorporated into a 

major USGS mapping project that encompasses the Basin 

and Range-Colorado Plateau (BARCO) transition zone in 

southeastern Nevada and southwestern Utah. Mapping of the 

Caliente caldera complex started along its northwestern margin in the Indian Cove quadrangle (Rowley and Shroba, 

1990, 1991 ), just east of the mining district, and in the Eccles 

(R.E. Anderson, unpub. data, 1988) and Caliente quadrangles (P.D. Rowley, unpub. data, 1988), southeast and south 

of the mining district. Later mapping of the Chief Mountain 

quadrangle (Rowley and others, 1991), which embraces the 

mining district and most of the Chief Range, as well as mapping of the Pahroc Spring SE quadrangle (Swadley and 

Rowley, 1992) and mapping in progress ofthe Caliente NW, 

Delamar, and Dow Mountain quadrangles, provides additional geologic framework for this report. 

Six new 40 Ar/39 Ar mineral dates from three samples and 

one K-Ar mineral date from one sample help to constrain 

temporal relationships in the Chief Mining District. These 

data are part of a larger K-Ar and 40Arf39 Ar geochronologic 

study being done within the context of the BARCO project. 

The 40 Arf39 Ar isotopic dating method is a variant of the conventional K-Ar method; for 40 Arf39 Ar analysis, samples are 

irradiated in a nuclear reactor to produce 39 Ar from 39K. In 

this study, 50-100 mg of sanidine and biotite and about 300 

mg of hornblende were irradiated in the USGS Denver 

TRIG A reactor for 30 hours at 1 megawatt. Each sample was 

irradiated adjacent to an aliquant of standard MMhb--1 hornblende for which K=1.555 percent, 40ArK=l.624x1o-9 

mol/g, and K-Ar date=520.4 Ma (Samson and Alexander, 

1987). Samples were experimentally degassed in a doublevacuum resistance furnace via step heating in 12-15 steps 

under ultra-high vacuum for about 20 minutes at each temperature step. The mass 40, 39, 38, 37, and 36 isotopes of 

argon were analyzed using a Mass Analyser Products series 

215 rare-gas mass spectrometer. Raw isotopic data were corrected for volume, mass discrimination, trap current, radioactive decay of 37 Ar and 39 Ar, and interfering isotopes of 

argon. Decay constants of Steiger and Jager (1977) were 

used to recalculate dates of any cited samples that were analyzed before 1977. Our preliminary ages were first reported 

in Rowley, Snee, and others (1990) and Snee and others 

(1990). Isotopic data are available from the authors. 

Classifications of volcanic and plutonic rocks are those 

of the International Union of Geological Sciences (Le Bas 

and others, 1986, and Streckeisen, 1976, respectively). Relative offset on some faults that contain slickensides was 

determined using the methods of Angelier and others (1985) 

and Petit (1987). 

Acknowledgments.-We gratefully acknowledge numerous helpful discussions with R.B. Scott concerning the structural setting and regional stratigraphy, with G.B. Allen and 

S.P. Bingham (both of Alma American Mining Corporation), J.E. Elliott, and E.B. Ekren regarding epithermal mineral deposits, and with W C Swadley pertaining to surficial 

deposits. We thank R.B. Blakestad and R.R. Kern (both of 

Homestake Mining Company) for access to companyH2 Application of Structural Geology to Mineral and Energy Resources 
Figure 1. Photograph, looking west, showing mines in upper Cobalt Canyon, Chief Mining District, Lincoln County, Nevada. 

Building to right of car was a former boarding house, and building farther up the canyon and to left is at entrance to tunnel of the 

Contact Mine; both are underlain by Highland Peak Formation, beyond which are Stirling Quartzite, Wood Canyon Formation, 

and Zabriskie Quartzite. A dump at the Gold Stake Mine is farther up canyon. The Advance Mine is in gulch left of central hill 

on horizon, and the Old Democrat Mine is at bottom of gulch to right of central hill. 

confidential data on the Chief Mining District, and we are 

grateful to these geologists and to Homes take for permission 

to publish some of those data, as identified in the text. We 

thank J.L. Askey and J.B. Harlan (both of FMC Gold Company), Charley Martin (Worldwide Resources), and R.R. 

Kern for information on the Taylor (Easter) Mine. Inclusion 

of chemical analysis of mineralized and altered rock samples 

(table 1) would not have been possible without the kind cooperation and rapid work of B.M. Adrian, R. Hill, D.E. Detra, 

R.M. O'Leary, R.W. Leinz, and C.M. McDougal of the 

USGS, and dating could not have been done without mineral 

separations by D.M. Cheney. Mapping in the northern Chief 

Range and adjacent areas by Axen and Burke was partly 

funded by grants to them from Texaco, Marathon, and Mobil 

oil companies, by grants to Burke from Sigma Xi and Northern Arizona University, and by National Science Foundation 

Grant EAR-14455-AC2 to J.M. Bartley (University of 

Utah). Editorial assistance was provided by C.H. Thorman 

and K.S. Kellogg. We thank C.H. Thorman for technical 

review of the manuscript and Ron, Betty, Chad, and Brenda 

Young of Caliente for help and many kindnesses during our 

field work. 

HISTORY AND PRODUCTION 

The date when the first ore deposits in the Chief Mining 

District were found is unknown, but it probably was shortly 

after the initial discoveries in 1868 in the Pioche Mining District, 27 km to the north (Callaghan, 1936). The Chief district 

was organized in 1870 and had an estimated production of 

$25,000 during the next few years. Production records for 

the district are incomplete. The first recorded shipment of 

ore was in 1896, followed by shipments of unknown value in 

subsequent years (Callaghan, 1936). An estimated 2,000 

ounces of gold, 11,000 ounces of silver, 45,000 pounds of 

lead, and 2,000 pounds of copper, with a total value of about 

$70,000, were tabulated for the period 1907-1953 (Callaghan, 1936, table 1; Tschanz and Pampeyan, 1970, table 

28). There is little information on production from individual 

Structural Setting of Chief Mining District, Nevada HJ 
J: 

""' 

> 

"'C 

~ 

r:;· 

~ Table 1. Chemical analyses of grab samples of hydrothermally altered and mineralized rocks from the Chief Mining District, Lincoln County, Nevada 

s· 

::I [Flame atomic absorption analyses ofTe, Tl, and Au by R. Hill, R. O'Leary, and B.H. Roushey; DC-ARC AES analyses of other elements by B.M. Adrian, D. Detra, and R.T. Hopkins. Ca, Fe, Mg, Na, and 

s.. Ti in percent, all other values in parts per million. Lower limit of determination of each element given in parentheses. N, not detected at limit of detection; H, interference; <,detected, but below limit of 

~ detection;>, greater than. Pat or below limit of detection except in samples 777 (0.5 percent) and 812 (0.2 percent); Bi not detected except in sample 781 (15 ppm); Ge, Sn, and Th not detected; Nb at or c: 

Q. below limit of detection except in sample 812 (20 ppm); W only detected in sample 781 (20 ppm)] c: 

~ 

C) Sample ~ number 378a 469b 751 765 777 778 781 793 794 802 812 866 1121c 1122a 0 

0 

O"Q Chemical analyses of samples -< 

0 Te (0.05) N N N 0.4 1.4 3.6 0.9 1.7 0.2 < N 0.1 N N 

~ Tl (0.05) < 0.35 0.15 0.35 0.95 0.2 0.3 0.2 2.2 3.1 0.5 2.8 0.55 0.45 

:;· 

~ Au (0.05) N N N N N N 1.1 N 1.1 N N 0.05 N N ~ Ca (0.05) 0.2 >20 0.2 0.15 0.5 0.05 < 15 10 7 0.2 1 1.5 5 Q.l 

::I Fe (0.05) 2 1 0.7 2 3 20 7 7 2 7 0.3 7 5 7 0.. 

ITI 

::I 

~ Mg (0.02) 0.3 1.5 0.1 0.7 0.1 0.03 0.1 0.3 2 2 0.3 0.3 1 3 O"Q 

-< Na (0.2) N < < 2 N < 0.2 N < < 0.2 0.5 >5 5 ~ 

~ Ti (0.002) 0.3 0.07 0.02 0.5 0.07 0.02 0.2 0.03 0.007 0.1 1 0.02 0.5 0.7 ~ 

~ Ag (0.5) N N N < 1 5 20 1 10 < < 100 N N n 

~ As (200) N N N N 500 10,000 10,000 300 500 200 N 300 N N 

B (0.10) N 10 10 70 50 10 2,000 15 15 50 150 70 < < 

Ba (20) 70 30 70 300 1,000 700 100 200 5,000 2,000 150 500 1,000 1,000 

Be (1) 2 < 2 < 5 < 2 <1 <1 < 2 1.5 <1 1 

Cd (20) N N N N 70 N N N N N N 70 N N 

Co (10) 10 N < < 15 N 15 10 < N N 10 15 20 

Cr (10) 100 10 10 < < 10 150 < < 15 70 20 50 50 

Cu (5) 20 5 5 5 500 50 100 15 30 7 5 150 30 30 

Ga (5) 5 15 N 70 N H 20 N 7 15 50 7 50 50 

La (50) < N N 70 100 N 50 N N N 100 N 100 100 

Mn (10) 150 700 200 100 >5,000 50 2,000 5,000 5,000 5,000 200 5,000 300 1,000 

Mo (5) N N N N 15 10 50 < 10 15 N N N N 

Ni (5) 20 5 5 < 70 < 7 10 7 5 < 5 15 20 

Pb (10) 20 10 N 30 20 10,000 20,000 500 10,000 100 50 15,000 50 50 

Sb (100) N N N N < 200 700 < < N N 150 N N 

Sc (5) 7 10 < 10 N N 10 7 < 5 10 5 7 15 
!a 

2 

0. 

c 

~ 

~ :r 

OQ 

s. 

n 

:r 

;:s· 

-

~ :;· 

:;· 

OQ 

c 

~-

~· 

z 

I'T> 

< 11.1 

0.. 

11.1 

:I 

l.n 

Table 1. Chemical analyses of grab samples of hydrothermally altered and mineralized rocks from the Chief mining district, Lincoln County, 

Nevada-Continued 

Sample 

number 378a 469b 751 765 777 778 781 793 794 802 812 866 

Sr (100) < 200 < 200 500 150 1,000 100 1,000 200 N 100 

v (10) 100 < 20 100 100 100 100 20 15 50 150 500 

y (10) 15 15 < 20 30 N 10 20 15 15 20 15 

Zn (200) N N N N 700 300 15,000 500 3,000 500 N 10,000 

Zr (10) 

378a 

469b 

751 

765 

777 

778 

781 

793 

794 

802 

812 

866 

112lc 

1122a 

150 50 10 150 100 10 200 10 

Lithology and location of samples 

Fault zone between Zabriskie Quartzite and Cobalt Canyon stock, lat 37°40'09" N., long 114°31'31" W. 

Metamorphosed limestone at contact with Cobalt Canyon stock, lat 37°39'33" N., long 114°30'51" W. 

Hydrothermally altered and quartz veined Cobalt Canyon stock, lat 37°40'36" N., long 114°30'14" W. 

< 50 

Hydrothermally altered local lava flows and tuffs near contact with Cobalt Canyon stock, lat 37°41'11" N., long 114°30'58" W. 

300 70 

Quartz veins, mineralized rocks in Zabriskie Quartzite from prospect pit 250m east of Gold Stake vein, lat 37°41'19" N., long 114°31'38" W. 

Quartz veins and mineralized rocks in Wood Canyon Formation from a shaft on Gold Stake vein, lat 37°41'16" N., long 114°31'46" W. 

Mineralized quartz veins on north side of a shaft of Republic Mine, 1at 37°41'17" N., long 114°32'03" W. 

Mineralized rocks in Highland Peak Formation brecciated by Stampede Detachment, from dump ofSOA Mine, lat 37°41'56" N., long 114°31'17" W. 

Mineralized vein in Highland Peak Formation at entrance to Gold ChiefMine, lat 37°41'50" N., long 114°31'20" W. 

Mineralized vein in fault zone in Highland Peak Formation at Lucky Chief No. 3 Pit, lat 37°41 '51" N., long 114°32'01" W. 

Hydrothermally altered Cobalt Canyon stock, lat 37°42'17" N., long 114°32'05" W. 

Mineralized rocks in fault zone of Lucky Hobo Mine, lat 37°41'42" N., long 114°31'51" W. 

Hydrothermally altered Cobalt Canyon stock, lat 37°40'29" N., long 114°30'04" W. 

Sulfide minerals in hydrothermally altered chilled contact of Cobalt Canyon stock, lat 37°40'34" N., long 114°30'16" W. 

1121c 1122a 

300 700 

100 150 

20 30 

N N 

200 200 
40 80 KILOMETERS 

EXPLANATION 

Caliente caldera complex 

Boundary of aeromagnetic 

ridge; from Zietz and others 

(1976, 1978) 

Axes of Pioche and Delamar-lron 

Springs Mineral Belts of 

Shawe and Stewart (1976) 

Timpahute Lineament of 

Ekren and others (1976) 

Axis of Intermountain Seismic 

Belt (Smith and Sbar, 1974) 

Colorado Plateau-Great Basin 

boundary 

Figure 2. Tectonic setting of Chief Mining District, Lincoln 

County, Nevada. B, Beaver; C, Chief Mining District; Ca, Caliente; Ce, Cedar City; D, Delamar; F, Fillmore; Fr, Frisco; H, 

Hur.rican.e;. M, M.i lf~rd; Ma, Marysvale; P, Pioche; Pe, Pennsylvania M1~1ng D1stnct; S, Stateline Mining District; T, Taylor 

(Easter) Mme; Te, Tempiute Mining District. 

mines other than some records that show that the Gold Chief 

Mine was by far the principal producer and the largest mine 

in the district. By 1911, the Gold Chief had an inclined shaft 

about 130 m long and drifts totalling about 300 m in length 

(Callaghan, 1936). Production was insignificant after 1941, 

and mining probably ceased in the 1960's. As a result of 

renewed interest in heap-leach processing of disseminated 

gold ore, claims recently were staked throughout the district 

as well as in the area adjacent to the main mass of the Cobalt 

Canyon stock (described below), which is centered about 1.5 

km south of the district. An area north of the Lucky Hobo 

Mine was apparently drilled by Texas Gulf International in 

1982 (Tingley, 1984). Homestake Mining Company did a 

geologic and geochemical study, authored by S.D. Olmore, 

of the district in 1985, including 12 reverse-circulation drill 

holes. 

Prospects on the northern town limits of Caliente are considered to be part of the Chief Mining District (Tingley, 

1984 ), but, because no production records are known for the 

area near these prospects and no ore bodies are exposed there, 

they are not discussed in this report. 

GEOLOGY 

Stratigraphy 

Most rocks in the Chief Mining District (fig. 3) consist of 

resistant, reddish-brown, pink, gray, and white, mostly thinbedded, planar and crossbedded, dense, vitreous orthoquartzite correlated by Callaghan (1936) with the Prospect 

Mountain Quartzite of Late Proterozoic and Early Cambrian 

age (details on regional stratigraphy of this and other Paleozoic units in the area are in Merriam, 1964, and Tschanz and 

Pampeyan, 1970). Stewart (1974, 1984) measured a stratigraphic section through this continental and near-shore 

marine quartzite sequence in Antelope Canyon, 6 km south 

of the district, and correlated it with Stirling Quartzite of 

Late Proterozoic age, Wood Canyon Formation of Late Proterozoic and Early Cambrian age, and Zabriskie Quartzite of 

Early Cambrian age. We followed Stewart's 'Stratigraphy in 

our mapping. The Stirling Quartzite is composed of lightcolored, unfossiliferous, generally pure quartzite that is typically coarser grained than the two younger formations and 

is at least 600 m thick (base not exposed); a prominent basalt 

lava flow is in its upper part. The Wood Canyon Formation 

is a dark-colored sequence of interbedded quartzite, siltstone, and shale that is about 320m thick and that contains 

animal tracks, burrows, and a thin basalt flow in its upper 

part. The Zabriskie Quartzite is mostly white, pure quartzite 

that is as thick as 140m and that contains local animal tracks 

and burrows. 

The Lower and Middle Cambrian Pioche Shale conformably overlies the Zabriskie Quartzite and is well exposed in 

an incomplete section in and north of Antelope Canyon. It 

consists of a soft, green, locally fossiliferous, micaceous 

marine shale and subordinate siltstone, sandstone, and limestone and is as thick as 250m in the Pioche area. Conformably overlying the Pioche in the Antelope Canyon area are 

two thin Middle Cambrian marine units, the Lyndon Limestone, a resistant medium- and dark-gray limestone about 20 

m thick, and the Chisholm Shale, a soft brown nonmicaceous 

shale about 30 m thick. 

Lower parts of Middle and Upper Cambrian Highland 

Peak Formation are exposed in the Antelope Canyon area. 

This formation consists of resistant, light- to dark-gray, 

black, and white, marine limestone and dolomite that has a 

maximum thickness of about 1,500 m in the Pioche area 

(Merriam, 1964 ). Numerous members have been mapped 

north and west of the Chief Range, where they are well 

Figure 3 (facing page). Generalized geology of Chief Mining 

District and vicinity, Lincoln County, Nevada. A, Advance 

Mine; C, Contact Mine; G, Gold Chief Mine; H, Lucky Hobo 

Mine; L, Lucky Chief Mine; 0, Old Democrat Mine; R, Republic Mine; S, Gold Stake Mine; So, SOA Mine. 

H6 Application of Structural Geology to Mineral and Energy Resources 
1W35' 

QT 

QT 

Cenozoic 

114"30' 

PANACA 

( 

s) or 

'§.: 

f 

5 KILOMETERS 

I 

..c;: .' 

!/ 

I 

) BASIN 

I 

II 

I 

I I 

I I 

I I 

I I I I 

,' I 

I 

EXPLANATION 

Surficial and basin-fill sediments 

Tuffs and sedimentary rocks of 14-12 Ma 

Hiko Tuff 

Lava flows of Indian Cove 

Quichapa Group 

Cobalt Canyon stock 

Oligocene tuffs and lava flows 

Cambrian limestone and shale 

Cambrian and Proterozoic quartzite 

-"--~~ High-angle fault-Dashed where approximately 

located, dotted where concealed 

_, __ _.,._ Newman Canyon Detachment-Teeth on 

upper plate 

(\ D A Stampede Detachment-Teeth on 

upper plate 

.....,. ....... .._...._ Delamar caldera-Dotted where concealed 

.l ... l ••• ~ •• Clover Creek caldera-Concealed 

Structural Setting of Chief Mining District, Nevada H7 
exposed (Axen, 1986; Lewis, 1987; Axen and others, 1988; 

Sleeper, 1989; Burke, 1991), but in the Antelope Canyon 

area only the basal Peasley Member and overlying Burrows 

Member (Merriam, 1964}are exposed. In the mining district, 

light-gray limestone only about 50 m thick is exposed; it is 

contact metamorphosed by the underlying Cobalt Canyon 

stock (see below) and can only be assigned tentatively to the 

Highland Peak Formation. 

Tertiary volcanic rocks unconformably overlie Proterozoic and Cambrian units in an area north of Antelope Canyon 

that is within 3 km of the southern margin of the mining district. Most of the volcanic rocks here are complexly faulted 

by a series of oblique-, normal-, and strike-slip faults; these 

same rocks are better exposed west of the Chief Range. The 

lowest volcanic unit (Oligocene) in the faulted area north of 

Antelope Canyon consists of hydrothermally altered local 

lava flows and tuffs intertongued and overlain by the Isom 

Formation, which consists of a series of mostly brown, 

crystal-poor, trachytic, regional ash-flow tuffs with a probable caldera source on the northern side of Escalante Valley 

north of Modena, Utah, and has a 40 Arf39 Ar date of about 27 

Ma (Best, Christiansen, and Blank, 1989). 

The overlying Quichapa Group contains, in ascending 

order, regional ash-flow sheets of the Leach Canyon Formation, Swett and Bauers Tuff Members of the Condor Canyon 

Formation, and Harmony Hills Tuff (Anderson and Rowley, 

1975). The Leach Canyon Formation is a soft, pink, poorly 

to moderately welded, crystal-poor, rhyolite ash-flow tuff; it 

has an average age of 24.0 Ma based on three K-Ar dates on 

biotite, sanidine, and plagioclase by Armstrong (1970) and 

two fission-track dates on zircon by Kowallis and Best 

(1990), but there is significant variation in these dates 

(26.7-21.6 Ma). The Swett and Bauers Tuff Members are 

resistant, reddish-brown and light-purple, densely welded, 

crystal-poor, rhyolitic ash-flow tuff; the Bauers has two 

40 Arf39 Ar dates averaging 22.8 Ma (Best, Christiansen, and 

others, 1989, table R3 ). The Harmony Hills Tuff is a resistant, greenish-gray, moderately welded, crystal-rich, andesitic ash-flow tuff that has a probable age of about 22.5-22 

Ma based on isotopic dates on overlying (Rowley and others, 

1989) and underlying rocks. 

A sequence of resistant, fresh, mostly gray, red, and 

brown lava flow rocks and subordinate flow breccia and volcanic mudflow breccia and minor fluvial sandstone rests 

unconformably on the Harmony Hills or older rocks 2.5 km 

southeast of the mining district. This sequence was named 

the lava flows of Indian Cove (Rowley and Shroba, 1991) for 

the area where it attains its greatest thickness (about 500 m) 

5 km southeast of the mining district. It probably formed as 

isolated or coalesced volcanic domes composed of highalkali andesite and trachyandesite and subordinate highalkali dacite and trachydacite. A lava flow from the unit 

yielded a Miocene plateau 40 Arf39 Ar date on sanidine of 

21.9±0.1 Ma (L.W. Snee, preliminary unpub. data, 1990; 

Snee and others, 1990). Pinching out against this dome 

complex, but locally exposed on its northern flank about 2.5 

km southeast of the mining district (Rowley and Shroba, 

1991), is resistant, tan, poorly to moderately welded, crystalpoor, rhyolite ash-flow tuff of the Hiko Tuff (Dol goff, 1963 ), 

which has an average 40 Arf39 Ar date of 18.6 Ma (Taylor and 

others, 1989). Locally derived sedimentary rocks and interbedded ash-flow tuff of the Kane Wash Tuff (14 Ma; Novak, 

1984) overlie the Hiko Tuff south and west of Caliente and 

are in turn overlain by the tuff ofEtna (which we earlier miscorrelated with the Ox Valley Tuff of Cook, 1960); these 

units are included in unit Ts of figure 3. 

The youngest rocks in the area (unit QT of fig. 3) are fluvial and lacustrine sedimentary rocks exposed in areas surrounding the Chief Range. Most of these sedimentary rocks 

are basin-fill deposits associated with the youngest (basinrange) episode of extensional faulting, but sandstone, conglomerate, and air-fall tuff at least 200m thick (sedimentary 

rocks of Newman Canyon of Row ley and others, 1991) that 

predate basin-fill sedimentary rocks south of the Chief 

Range and overlie the tuff of Etna have yielded a K-Ar date 

on sanidine of 13.8±0.9 Ma (H.H. Mehnert, preliminary 

unpub. data, 1990; Snee and others, 1990). Among the basinfill deposits, the most extensive unit fills the Panaca Basin 

east of the mining district. This unit consists of soft, tan, 

salmon, and light-green sandstone, limestone, siltstone, 

mudstone, and claystone of the upper Miocene and Pliocene 

Panaca Formation. Only the upper part, about 400 m thick, 

of the formation is exposed (Rowley and Shroba, 1991) and 

based on vertebrate fossils it is as young as the Blancan III 

mammalian age of 3.7-3.2 Ma (Repenning, 1987). Overlying the Panaca Formation are mostly coarse grained surficial 

deposits ranging in age from Pliocene to Holocene. Chief 

among these deposits is a series of alluvial-fan deposits of 

different ages that underlie five or more well-expressed geomorphic surfaces. In contrast to the fine-grained sediments 

of the Panaca Formation that were deposited in a closed 

basin, these coarse sands, gravels, and fanglomerates, as well 

as rounded fluvial gravels, are interpreted to represent deposition following the development of through-flowing drainage in the Panaca Basin via Meadow Valley Wash to the 

Colorado River, which is about 170 km south of the mining 

district. 

Intrusive Rocks 

A large, irregular, discordant intrusive mass of quartz 

monzonite porphyry, herein called the Cobalt Canyon stock, 

underlies most of the Chief Mining District at shallow depth 

and is exposed as plutons, plugs, dikes, sills, and faulted slivers within and near the district. Hydrothermal solutions from 

the stock and (or) circulating meteoric groundwater heated 

by the stock are believed to be the source of most metals in 

the district. Callaghan ( 1936) was the first to map parts of the 

stock and to recognize its association with the orebodies. 

H8 Application of Structural Geology to Mineral and Energy Resources 
Ekren and others (1977) mapped the main mass of the stock, 

which is centered about 1.5 km south of the district. Our 

mapping reveals separate plutons, plugs, dikes, and sills (fig. 

3) and interprets these bodies to be apophyses of one stock at 

depth. In addition to the plutonic rocks shown in figure 3, 

dikes, small plugs, and faulted slivers of the same lithology 

have been noted in several mines and prospects. From north 

to south, these bodies include ( 1) a plug 250 m northeast of 

the entrance to the Gold Chief Mine (Callaghan, 1936, fig. 2) 

and drill-hole cuttings (Homestake Mining Company) of 

quartz monzonite porphyry on the road (from a drill hole 

sited on the road) 100m southeast of that mine, which indicate the presence of plutonic rock at depth; (2) a silicified 

dike in "the fault west of the Gold Chief Mine" (presumably 

the Stampede Detachment, which is exposed here) and a sill 

in limestone just south of the mine (Callaghan, 1936, p. 10); 

(3) drill-hole cuttings of pyrite-bearing porphyry on the road 

450 m southeast of the entrance to the Gold Chief Mine; (4) 

several plugs near the Lucky Hobo Mine (Callaghan, 1936, 

fig. 2, p. 29) and an altered dike along the Stampede Detachment exposed in a prospect trench about 200 m west of the 

mine; (5) a "lens" exposed in the main shaft and samples 

from the dump of the Old Democrat Mine (Callaghan, 1936, 

p. 24); (6) chips of porphyry next to a vertical shaft along the 

road in the bottom of Cobalt Canyon less than 100 m east 

down the canyon from the Old Democrat Mine and drill-hole 

cuttings of fresh porphyry along the same road another 100 

m down the canyon; (7) several plugs and dikes between the 

Advance and Contact Mines (Callaghan, 1936, fig. 2); (8) a 

plug on the ridge just south of the Contact Mine and an 

altered dike in the mine itself (Callaghan, 1936, p. 10, 

21-23); and (9) a 1-m-wide dike in the Stampede Detachment exposed in a trench on top of the ridge just south of the 

Contact Mine. 

Rock of the Cobalt Canyon stock is gray, tan, green, and 

khaki, locally grusy, quartz monzonite porphyry that consists of 35-55 percent phenocrysts of plagioclase, subordinate clinopyroxene, and minor biotite and iron-titanium 

oxide minerals in a fine-grained, holocrystalline groundmass 

of sanidine, subordinate to minor quartz, and minor clinopyroxene, iron-titanium oxide minerals, and biotite. The intrusive mass has fine-grained holocrystalline chilled margins 

and, in northwestern exposures, includes a large body of 

black, coarse-grained pyroxenite. Locally the intrusion is 

deuterically and hydrothermally altered at and near intrusive 

and fault contacts. The Cobalt Canyon stock is of Oligocene 

age, based on a 24.8-Ma average of two 40Arf39Ar plateau 

dates on mineral separates, 24.9±0.1 Ma on sanidine and 

24.7±0.1 Ma on biotite (L.W. Snee, preliminary unpub. data, 

1990; Snee and others, 1990), from a sample collected from 

the chilled margin of the main mass about 1 km east-southeast of the district (Rowley and others, 1991). 

A second variety of intrusive rock is named the porphyry 

of Meadow Valley Wash (Rowley and Shroba, 1990, 1991). 

The unit consists of widely scattered dikes and plugs, too 

small to be shown in figure 3, of high-alkali dacite that was 

intruded along major oblique-slip and strike-slip faults in the 

area. The light-gray and pink, locally grusy rock is characterized by large (as long as 2 em) feldspar phenocrysts. The 

rock consists of about 35 percent phenocrysts of plagioclase, 

subordinate biotite, subordinate to minor sanidine, quartz, 

and hornblende, minor iron-titanium-oxide minerals, and 

traces of clinopyroxene, all in a glassy groundmass. It is 

named for exposures in the canyon of Meadow Valley Wash 

about 5 km south-southeast of the mining district, but the 

same rock has been mapped along a fault about 3 km southsouthwest of the district, where it crops out along the 

Meadow Valley Wash and Chief Canyon Fault Zones. It also 

is present along the Gravel Pit Fault Zone and at several 

other locations in the northern Caliente caldera complex. 

Plugs of similar lithology were noted along a northweststriking fault zone just east of Panaca Summit, about 30 km 

northeast of the mining district (unit Tp of Best and others, 

1991), and in a major north-northwest-striking fault zone 

(Swadley and Rowley, 1992) on the northern side of the 

Delamar Mining District, about 30 km southwest of the 

Chief Mining District (P.D. Rowley, unpub. mapping, 

1990). A sample from a dike of the porphyry of Meadow 

Valley Wash intruded along the Chief Canyon Fault Zone in 

the canyon of Meadow Valley Wash is Miocene in age, 

based on a 19.4-Ma average of two 40 Ar/39 Ar plateau dates 

on mineral separates, 19.3±0.1 Ma for a well-defined spectrum on sanidine and 19 .5±0.1 Ma for a slightly less defined 

spectrum on hornblende; 22.5 Ma for a disturbed spectrum 

on biotite from the same sample was discounted (L.W. Snee, 

preliminary unpub. data, 1990; Snee and others, 1990). 

Caliente Caldera Complex 

The Caliente caldera complex contains two small epithermal gold districts and may control the large nearby 

Delamar epithermal gold district (see later discussion). The 

caldera complex is also the source of several ash-flow tuffs 

exposed near the Chief Mining District. Williams ( 1967) 

was the first to suggest, on the basis of regional stratigraphy 

and isopach data, that the Leach Canyon and Condor Canyon 

Formations were derived from centers in the CalientePanaca area. Noble and others (1968) and Noble and McKee 

(1972) did the first geologic reconnaissance study in this 

area and proposed that the Caliente depression, which they 

defined as a circular volcano-tectonic basin at least 32 km in 

diameter that extends from about 7 km east of Caliente eastward almost to the Utah State line, was the source ofregional 

ash-flow sheets of the Harmony Hills, Racer Canyon, Hiko, 

and Ox Valley Tuffs. Reconnaissance mapping by Ekren and 

others (1977), combined with Bouguer gravity data, much 

refined the extent and location of the feature, which they 

called the Caliente caldron complex. Instead of being circular, they determined that it is elongated east-west (64 km by 

Structural Setting of Chief Mining District, Nevada H9 
29 km), extending from 25 km west of Caliente to almost the 

Utah State line, and that it is centered farther south than previously inferred; they recognized that the Hiko Tuff was 

derived from the western part of the complex. 

Our detailed mapping in the north-central and northwestern parts of the complex and our reconnaissance studies in 

the northeastern part, combined with aeromagnetic and Bouguer gravity data (Blank and Kucks, 1989), indicate that the 

Caliente caldera complex extends into Utah and thus is 

somewhat larger (at least 80 km east-west by 35 km northsouth) than was previously thought. The complex consists of 

several inset calderas of widely different ages, four of which 

we have named (Rowley and Siders, 1988). The oldest is the 

Clover Creek caldera, the source of the Bauers Tuff Member 

(22.8 Ma; Best, Christiansen, and others, 1989) and perhaps 

the source of the petrologically similar Swett Tuff Member, 

both of the Condor Canyon Formation. The Clover Creek 

caldera, although complexly faulted and mostly covered, is 

4 km south-southeast of the mining district and underlies the 

north-central part of the Caliente caldera complex. The 

Delamar caldera, the source of the Hiko Tuff (18.6 Ma; Taylor and others, 1989), is 8 km south of the mining district and 

makes up the western part of the caldera complex. The other 

two named calderas produced unnamed ash-flow tuffs that 

are not known to crop out outside the caldera complex; these 

tuffs are as young as about 15 Ma (Mehnert and others, 

1989). The tuff of Etna and at least one other ash-flow 

sequence probably also were derived from the Caliente 

caldera complex. The Delamar and younger calderas clearly 

formed during an episode of major faulting (described later) 

and thus are characterized by a complex combination of 

faults and structural and topographic caldera walls (Rowley, 

Anderson, and others, 1990). The Clover Creek caldera may 

have formed during the same episode of faulting. 

Although the southern exposed margin of the Cobalt 

Canyon stock (24.8 Ma) is almost adjacent to the Clover 

Creek caldera (22.8 Ma), the difference between the two 

ages of these bodies argues against a genetic tie. Nonetheless, the stock may have been synchronous with an early, as 

yet undocumented, phase of magmatism in this long-lived 

caldera complex. In fact, isopach data led Williams ( 1967) to 

suggest that the source of the Leach Canyon Formation 

(24.0? Ma) was the northern part of the Caliente caldera 

complex, but this hypothesis has not been verified because 

this part of the complex is only partly mapped and is mostly 

buried by much younger rocks. 

STRUCTURE 

Mineral Belts 

Mineralization in the Chief Mining District probably was 

controlled by both regional tectonic features and local structures. Numerous mineral belts have been proposed for many 

years to describe alignments of mining districts in the Great 

Basin. Bagby (1989) summarized some of this literature and 

showed two major provinces of mineral belts, one in central 

and western Nevada in which precious metals are the most 

significant products and in which districts are aligned along 

northwest-trending zones and one in eastern Nevada and 

western Utah in which base metals are the most significant 

products and in which districts are aligned along generally 

east-trending zones. 

The Chief Mining District is within the Delamar-Iron 

Springs mineral belt of Shawe and Stewart (1976), which 

also includes the Tempiute, Delamar, and Iron Springs mining districts and is south of the Pioche mineral belt (fig. 2). 

It might be more appropriate to call these alignments igneous 

belts because they contain most of the volcanic and intrusive 

rocks in southwestern Utah and southeastern Nevada and are 

well delineated by aeromagnetic anomalies (fig. 2) (Zietz 

and others, 1976, 1978). Cenozoic and perhaps older easttrending alignments of structures and igneous centers also 

are well known in this part of the Great Basin (Ekren and 

others, 1976; Rowley and others, 1978). One of these, the 

Timpahute Lineament (fig. 2) of Ekren and others (1976), 

intersects the Chief district; Ekren and others (1976, 1977) 

suggested that it controlled emplacement of the Cobalt Canyon stock, the northern side of the Caliente caldera complex, 

and numerous other features. Stewart and others (1977) also 

called attention to the east-trending grain of the geology of 

the Great Basin, especially patterns of Tertiary igneous 

activity. Other, mostly older (Oligocene to early Miocene), 

east-trending topographic alignments, dikes, and other features were noted by Best (1988) in southeastern Nevada and 

southwestern Utah. Also noteworthy is the fact that most 

faults strike east in the Bull Valley Mountains of Utah 

(Hintze, 1980) and Clover Mountains of Nevada, which are 

east and southeast of the Chief district, and that the Intermountain Seismic Belt (Smith and Sbar, 1974) swings westsouthwest through southern Utah and Nevada (fig. 2), centered at about the latitude of the Caliente caldera complex. 

The probable control of the belts, lineaments, and other 

aligned features is the same, namely old, deep-seated, eaststriking faults or basement flaws, some of which have been 

reactivated during the Tertiary (Rowley and others, 1978). In 

stark contrast, almost all major younger (post-middle 

Miocene) faults strike north, northeast, and northwest. 

Pre-Upper Oligocene Detachment Faults 

Recent geologic mapping demonstrates at least three episodes of Tertiary extensional faulting in the Chief district. 

These episodes were recognized also by Bartley and others 

(1988), Taylor and others (1989), and Taylor (1990) in an 

east-west transect north of the district. The oldest episode 

involves a north-striking, east-dipping, low-angle fault, first 

recognized by Callaghan (1936), that separates Zabriskie 

Hl 0 Application of Structural Geology to Mineral and Energy Resources 
Quartzite on the west from Highland Peak Formation on the 

east; Callaghan also recognized that the fault is a major 

structure because the entire thickness of the Pioche Shale is 

faulted out. He called it a "probable" thrust fault and noted 

that it is better exposed in Antelope Canyon than in the mining district and that many orebodies are in the wide breccia 

zones developed along the structure. Tschanz and Pampeyan 

(1970, plate 3) traced this structure northward along the eastem side of the Chief Range and correlated it with a west-dipping low-angle fault on the western side of the Highland 

Range, which is north of the Chief Range; they named the 

structure the "Highland thrust fault" and considered it to be 

of "Laramide" age. In a brief discussion of thrust faults in 

Lincoln County, however, Tschanz and Pampeyan (1970, p. 

82-83) cited field relations, especially in the northern Bristol 

Range about 50 km north of the Chief district, that suggest 

that this fault may be Tertiary in age and may not be a thrust. 

The name "Stampede detachment'' was proposed by Axen 

and others (1988) for outcrops near Stampede Gap in the 

northwestern Highland Range that Tschanz and Pampeyan 

correlated with their Highland thrust; Axen and colleagues 

mapped the east-dipping Stampede Detachment along the 

eastern Chief Range and southwestern Highland and Bristol 

Ranges and in the Pioche area (Bartley and others, 1988, fig. 

2). Axen and others (1988) recognized that the low-angle 

normal fault is related to extension that predates the oldest 

(late Oligocene) volcanism in the area; movement of the 

upper plate of the Stampede Detachment was to the east. The 

detachment is primarily a bedding-parallel fault in the 

Pioche Shale or, less commonly, in the Chisholm Shale or 

lower Highland Peak Formation, and the rocks in this part of 

the section are generally attenuated and locally cut out. The 

age of the Stampede Detachment is poorly constrained, but 

Axen and others (1988) and we consider it to be Tertiary; the 

geometry of the detachment resembles faults mapped by 

Nutt and Thorman (this volume) in the Drum Mountains of 

western Utah that they consider to be of Sevier(?) age. 

Throw on the Stampede Detachment decreases southward, 

and the fault has not been recognized at the latitude of the 

Caliente caldera complex. The fault is well exposed in the 

Chief District despite having been contact metamorphosed. 

Gouge and fault breccia as thick as 20 m along the fault are 

favorable sites for mineralized rocks. 

Miocene Faults 

The major episode of faulting in and near the Chief Mining District is Miocene in age (at least as old as 19-12 Ma). 

Faults of this episode consist of low-angle normal (detachment) faults and high-angle strike-slip, oblique-slip, and normal faults (fig. 3). Those faults that strike east or northwest 

tend to have right slip or oblique (right) slip, whereas those 

that strike north or northeast tend to have left slip or oblique 

(left) slip. An example of a major high-angle oblique-slipfault zone is the Chief Canyon Fault Zone (fig. 3). Where 

exposed about 2 km southeast of the district, the fault zone 

strikes north and is almost vertical; we consider it to be an 

oblique-slip fault based both on strike-slip and dip-slip slickensides and on reidel shears (see Angelier and others, 1985; 

Petit, 1987) that indicate its eastern side was offset down and 

to the right. Farther south, it veers south-southeast and 

requires the same sense of oblique slip to explain slickensides and geologic relations. The amount of movement on 

this part of the fault is constrained by outflow Bauers Member to the north and intracaldera Bauers Member to the 

south; on the basis of this poorly constrained margin of the 

Clover Creek caldera, we estimate right slip of 1 km or more. 

Where the Chief Canyon Fault Zone cuts the Delamar 

caldera, west of lower Barnes Canyon about 10 km southsoutheast of the district, the zone is expressed as a narrow, 

complex horst of propylitically and argillically altered rocks 

in which the predominant offset is normal (R.E. Anderson, 

unpub. data, 1989). The northern part of the Chief Canyon 

Fault Zone, east and northeast of the district, is not exposed 

but was inferred on the basis of a steep Bouguer gravity gradient (Blank and Kucks, 1989); considering that the fault 

here strikes north and separates the 700-m-high Chief Range 

from the Panaca Basin, it is likely that this segment of the 

fault zone is younger than the southern part of the fault zone 

and that it is a basin-range fault (see next section). 

Other faults west of the Chief Canyon Fault Zone and 

4-6 km south ofthe mining district confirm that major deformation occurred during this Miocene episode (fig. 3). The 

Meadow Valley Wash Fault Zone strikes north along the 

western side of the canyon of Meadow Valley Wash; a 2-kmlong wedge of the Clover Creek caldera on the east against 

Bauers outflow tuff and Oligocene volcanic rocks on the 

west indicates at least 2 km of left slip and probable displacement down to the east. The fault zone is joined on the east, 

just north of Antelope Canyon, by the north-northeast-striking Gravel Pit Fault Zone that separates another 2-km-long 

wedge of the Clover Creek caldera on the east from Stirling 

Quartzite, Wood Canyon Formation, and Zabriskie Quartzite on the west. Field evidence argues against the Clover 

Creek caldera having been west of the fault, so we interpret 

the Gravel Pit Fault Zone as an oblique-slip fault of at least 

2 km of left slip and probable down throw to the east; limited 

slickenside data, including reidel shears, support this interpretation. Dikes and plugs of the porphyry of Meadow Valley Wash, as much as 300 m across, have been intruded 

along the Meadow Valley Wash, Gravel Pit, and Chief Canyon fault zones. Preliminary 40 Arf39 Ar dates on the porphyry indicate that most deformation on these faults had 

taken place by 19 Ma. Other nearby faults having the same 

attitude and throw and interpreted to belong to the same 

Miocene episode are younger, whereas still others may be 

older. 

High-angle faults are abundant in the Chief Range, but, 

because most rocks are Proterozoic and Cambrian quartzite, 

the age of faulting and amounts of slip are not well 

Structural Setting of Chief Mining District, Nevada H11 
constrained. Abundant slickensides in the brittle quartzite 

indicate mostly oblique, normal, right, and left slips on these 

faults. Probably most date to the Miocene episode of faulting, but we cannot rule out the possibility that some faults in 

the Chief Range predate the Miocene episode and thus may 

be of the same age as the Stampede Detachment or perhaps 

older; because of their unusual strike, a series of east -striking 

faults, some shown on figure 3, are particularly suspect as 

older features. In addition, some faults in the interior of the 

Chief Range, like those that bound it, must be younger basinrange faults (see next section). The age of faults is somewhat 

better constrained in the mining district because most faults 

there offset Tertiary volcanic rocks and the Cobalt Canyon 

stock. Unfortunately, few shears were found in the grus of 

the stock, but mapping indicates that the southwestern and 

northern sides of the stock are bounded by faults. The Old 

Democrat Fault Zone (fig. 3) is typical of the deformation 

along faults in the Chief Range. It is a significant fault zone 

that has several splays, local fault-breccia zones as wide as a 

few meters, and opposing dips to strata on opposite sides of 

the zone. The fault plane locally dips steeply east and in 

other places dips steeply west; quartzite east of the fault zone 

is younger than quartzite to the west. Slickensides on several 

splays at the Old Democrat Mine indicate oblique slip on 

east-dipping fault planes and right slip and reverse slip on 

west-dipping fault planes. The Old Democrat, Advance, and 

perhaps Lucky Hobo mines, as well as several unnamed 

prospects, are along the Old Democrat Fault Zone and its 

splays; apparently these mines were developed in old (24.8 

Ma) orebodies offset by the fault zone, although metals are 

present along some faults and may indicate remobilization of 

metals by groundwater during or after deformation. The earliest phase of high-angle faulting of the Miocene episode 

probably is in fact Oligocene in age and synchronous with 

emplacement of the Cobalt Canyon stock and epithermal 

mineralization, as in the Walker lane of western Nevada 

(see, for example, John and others, 1989; Hardyman and 

Oldow, 1991). 

A major low-angle normal fault, the Highland Detachment of Axen ( 1986), is exposed at the northwestern end of 

the Chief Range (Burke and Axen; 1990; Burke, 1991). 

From here it can be traced northward as a fairly continuous 

feature along the western sides of the Highland and northern 

Bristol Ranges. Parts of it were mapped as the "Laramide" 

Highland Thrust Fault and other parts were mapped as the 

"post-Laramide" Bristol Thrust Fault (Tschanz and Pampeyan, 1970). Axen (1986), Lewis (1987), Axen and others 

(1988), Bartley and others (1988), Sleeper (1989), and Burke 

and Axen (1990) interpreted exposures of the Highland 

Detachment along the Chief, Highland, and Bristol Ranges 

to be the breakaway of the detachment. In contrast to the 

much older Stampede Detachment, the Highland Detachment dips west and movement of its upper plate was to the 

west. These geologists have not identified the Highland 

Detachment to the east, and they interpreted that it continues 

in the subsurface under ranges to the west. 

The hanging wall of the Highland Detachment, on the 

western sides of the Chief and Highland Ranges, contains a 

complex sequence of intertongued, mostly soft, coarsegrained landslide deposits and mudflow breccia, minor alluvial deposits, and rhyolitic air-fall and possibly ash-flow tuff 

that comprise the McCullough Formation of Axen and others 

(1988). These rocks are syntectonic, having been deposited 

contemporaneously with movement on the detachment 

(Axen and others, 1988; Sleeper, 1989; Burke and Axen, 

1990; Burke, 1991). One of the air-fall tuff beds in the 

McCullough Formation, perhaps from a source in the Caliente caldera complex, has a 40 Arf39 Ar date of 15.3 Ma (Axen 

and others, 1988; Taylor and others, 1989), thereby giving a 

minimum age for detachment faulting. Andesitic volcanic 

mudflow breccia, andesitic flow breccia, and small andesitic 

lava flows, called the volcanic rocks of Klondike Spring 

(Rowley and others, 1991), that underlie the McCullough 

Formation may also be syntectonic with the Highland 

Detachment. 

In the Caliente caldera complex south of the mining district, north-northwest-striking high-angle right-slip, obliqueslip, and normal faults are abundant and one low-angle 

detachment fault (Newman Canyon Detachment) is exposed 

(Bowman, 1985; R.E. Anderson and P.D. Rowley, unpub. 

mapping, 1986-1990; Angelier and others, 1987; Mehnert 

and others, 1989; Michel-Noel and others, 1990; Rowley and 

others, 1990). In places caldera margins are a complex combination of faults and topographic and structural margins, 

arguing that these faults were active during formation of the 

Delamar caldera (18.6 Ma; Taylor and others, 1989) and 

younger calderas (as young as 15 Ma; Mehnert and others, 

1989). Fault offsets that decrease upward in the stratigraphic 

section ("growth faults" of Bowman, 1985) indicate that the 

faults were also active during deposition of postcaldera tuffs 

of the Kane Wash Tuff (14 Ma; Novak, 1984) and the overlying tuff of Etna (also 14 Ma). These tuffs are unconformably overlain by the conglomerate of Ash Canyon (Rowley 

and Shroba, 1990, 1991) and intertongued basalt lava flows 

that have K-Ar dates of about 12 Ma (Mehnert and others, 

1989) in exposures about 12 km south-southeast of the Chief 

Mining District. Although our early mapping indicated that 

the conglomerate and basalt also unconformably overlie 

faults of the Miocene episode (Mehnert and others, 1989), 

later mapping indicates that some of these faults· cut rocks at 

least as young as the conglomerate and basalt (Snee and others, 1990). The similar ages of many of the faults, whether 

detachment or high angle, in the Caliente caldera complex, 

Chief Range, and Chief district indicate that the faults 

formed in a major Miocene eposide, from at least 19 to 12 

Ma, the main extensional episode in this part of Nevada. 

Upper Miocene and Pliocene 

Basin-Range Faults 

Displacements on Neogene faults have strongly influenced development of much of the present-day large-scale 

H12 Application of Structural Geology to Mineral and Energy Resources 
topography in the eastern Great Basin, especially imprinting 

this topography with a prominent north-south grain. These 

Neogene faults are known as basin-range faults and are generally younger than about 10 Ma (Anderson and others, 

1983). They may have formed under an extension direction 

(with least principal stress oriented west-northwest) that had 

been rotated clockwise about 60° from the Miocene extension direction (with least principal stress oriented westsouthwest) (Anderson and Ekren, 1977; Zoback and others, 

1981; Anderson and others, 1983; Anderson, 1989). Near the 

Chief Mining District, most basin-range faults strike north, 

such as those in the Panaca Basin (fig. 3). The oldest exposed 

basin-fill deposits, the conglomerate of Ash Canyon and 

intertongued 12-million-year-old basalts, are cut by mostly 

north-striking basin-range faults. Younger Tertiary basin-fill 

sedimentary rocks, including the Panaca Formation, are also 

cut by such faults. Some basin-range faults in the area are 

Quaternary in age, such as north-striking faults in Delamar 

Valley about 20 km west of the Chief Mining District 

(Tschanz and Pampeyan, 1970; Ekren and others, 1977). 

Similar faults control the position of some north-trending 

streams about 15 km south-southeast of the mining district 

(R.E. Anderson, unpub. mapping, 1989). 

Old basin-range faults are not as easy to identify as 

younger ones, but many old faults have greater displacement 

than younger ones. The north-trending Chief Range is probably a basin range, bounded by basin-range faults whose displacements must be at least that of the 700-m height of the 

range above the adjacent Panaca Basin. Faults on the western 

and southern sides of the range (upper left part of fig. 3) are 

locally well exposed and consist of high-angle faults containing slickensides that indicate normal slip. The buried 

northern part of the Chief Canyon Fault Zone, extending 

north from just east of the mining district, is interpreted to be 

the range-bounding fault that separates the Chief Range from 

the Panaca Basin. If so, it is much younger than the southern 

part and perhaps is a reactivated part of the Miocene obliqueslip fault zone that is farther to the south. 

ORE DEPOSITS 

Epithermal fissure-vein precious-metal deposits are common in and adjacent to calderas (see, for example, Hayba and 

others, 1986; Heald ano others, 1987; Lipman, this volume). 

Most such deposits are related to the plumbing systems of 

calderas, in which late-forming intrusions produce convection cells of meteoric water. Orebodies form in the upper 

parts of these plumbing systems, and thus younger, less 

eroded calderas are generally better sites for exploration than 

are older buried ones. The Caliente caldera complex is 

poorly known geologically and has not been adequately 

explored for epithermal gold; however, it contains relatively 

young calderas, suggestive of good exploration potential. 

The Chief Mining District, near the northern edge of the 

complex, and the major Delamar Mining District (Callaghan, 1937; Tchanz and Pampeyan, 1970; Tingley, 1984), 

near the southwestern edge of the complex (fig. 2), are both 

deeply eroded epithermal gold systems. In both districts, 

most orebodies are in brittle quartzite of the Stirling Quartzite, Wood Canyon Formation, and Zabriskie Quartzite, 

which must have shattered like glass during faulting, preparing the rocks for mineralization, which took place during 

intrusive activity that apparently was synchronous with 

faulting. Both districts probably are of the adularia-sericite 

type (Heald and others, 1987) of epithermal deposits. Analyses of grab samples (Tingley, 1984, appendix B, samples 

17 40-1755) from the Delamar district show gold values of 

as much as 22 ppm, silver values of as much as 300 ppm, and 

anomalous values of copper, lead, and zinc. East-striking 

structures, including rhyolite dikes associated with some 

ore bodies, are common in the district (Callaghan, 1937). The 

Delamar district has recently (1990-91) been explored and 

studied by Alma American Mining Company and FischerWatt Gold Company (S.P. Bingham and G.B. Allen, written 

commun., 1990). 

Another known adularia-sericite type of epithermal system underlies the Taylor (or Easter) Mine, located within the 

caldera complex 13 km east-northeast of Delamar and 15 km 

southwest of Caliente (fig. 2). Tingley (1984, appendix B, 

sample 579) reported values of gold of 18 ppm and of silver 

of 70 ppm from a grab sample from the mine. Sparse slickenside data suggest that the structure is simple, namely a normal fault zone in Hiko intracaldera tuff (Delamar caldera) 

that strikes N. 80° E. and dips 50° to the north; this fault may 

be a dilational response to right slip along a nearby concealed east-striking fault (R.R. Kern, oral commun., 1990). 

Banded quartz veins that contain comb structure and follow 

the fault have low base-metal contents but relatively high 

gold contents, and locally the tuff is intensely hydrothermally altered within a zone that is at least 100 m wide on 

either side of the fault. The ore bodies are in a high-level epithermal system (G.B. Allen and S.P. Bingham, oral commun., 1990) that has been previously drilled and explored by 

Homestake Mining Company and other firms (Charley Martin, oral commun., 1990) and in 1990-91 by FMC Gold 

Company (J.L. Askey and J.B. Harlan, oral common., 1990). 

In addition, the small, poorly known Pennsylvania mining district on the southern margin of the caldera complex 

(fig. 2) is probably an epithermal gold system (Tchanz and 

Pampeyan, 1970; Tingley, 1984); values of gold of as much 

as 180 ppm and of silver of as much as 1,500 ppm, as well as 

anomalous values of copper, lead and zinc, were reported 

from grab samples (Tingley, 1984, appendix B, samples 

1701-1707). 

Ore bodies in the Chief Mining District are deeply eroded 

epithermal fissure-vein deposits that occupy breccia and 

low-angle faults of the Stampede Detachment Fault and 

high-angle faults and joints within quartzite. Most of the 

Structural Setting of Chief Mining District, Nevada H13 
following descriptions of the ore deposits are those of Callaghan (1936). Callaghan noted evidence for replacement of 

some country rocks by ore minerals-for example, quartzite 

in the Old Democrat Mine (fig. 3) and limestone or dolomite 

in the Lucky Chief Mine. He found that all ore is oxidized, 

except for some sulfide minerals (galena) in the Republic 

and Gold Chief mines. He recognized three types of vein 

deposits. The first, which has the highest amounts of gold 

and silver in the district, is the arsenian type, which is found 

in the Advance, Old Democrat, Republic, and Lucky Hobo 

mines. Typically, the ore in these veins is a green to brown, 

vuggy aggregate of supergene scorodite that replaced arsenopyrite and is locally replaced in tum by a brown micaceous mineral (arsenosiderite?) and by iron-oxide minerals; 

quartz, jarosite, cerussite, descloizite, mimetite, beudantite, 

and manganese-oxide minerals also are present. The second 

type of vein deposit, which is the most economically productive in the district although it contains relatively low proportions of gold and silver, is characterized by a network of 

tabular barite crystals and scattered grains of galena; chalcedonic quartz is present as veins in barite, and cavities 

between barite crystals locally contain quartz, calamine, 

mimetite, and manganese- and iron-oxide minerals. This 

type of deposit is present along the Stampede Detachment in 

the Gold Chief and Lucky Chief mines. The third type of 

vein deposit, from which most of the lead was produced, 

includes irregular and pipelike masses of cerussite in the 

Republic and Lucky Chief mines; in addition to cerussite, 

ores contain fine-grained quartz, opal, plumbojarosite, and 

manganese- and iron-oxide minerals. A fourth type of vein, 

which is barren or contains low amounts of gold and silver 

but is exposed in many prospect pits, consists of lenses and 

small bodies of red iron oxide in quartzite resulting from 

fault breccia cemented by iron oxide and comb quartz. In 

addition, Tingley ( 1984) noted fine-grained disseminated 

sulfide minerals and silicified quartzite wallrock in the 

Advance Mine. Furthermore, in 1985, Homestake Mining 

Company discovered, by reverse-circulation drilling, a large 

body of disseminated pyrite (2-1 0 percent by volume) and 

phyllically altered rock of the Cobalt Canyon stock at about 

110 m depth below, east, and south of the Gold Chief Mine. 

Although chemical analyses reveal anomalous amounts of 

arsenic and several other elements, the body does not contain 

appreciable amounts of gold or silver, but it is suggestive of 

a porphyry copper deposit (data from a 1986 report by S.D. 

Olmore of Homestake Mining Company that was made 

available for publication courtesy of Homestake Mining 

Company in 1990). 

The only published chemical analyses of orebodies in the 

Chief District are by the USGS and include atomic absorption analyses of As, Au, Sb, and Zn and semiquantitative 

spectrographic analyses of other elements for 12 grab samples (Tingley, 1984, appendix B) and flame atomic absorption analyses of Te, Tl, and Au and direct-current arc and 

alternating-current spark emission spectrographic field 

analyses of other elements for 14 grab samples that we collected (table 1). The analyses suggest that the various veins 

are not much different geochemically from each other except 

for those in mines in the Stampede Detachment, which contain somewhat less arsenic. There is no evidence from these 

meager geochemical data that the veins are of more than one 

age. The different vein types of Callaghan ( 1936) cannot be 

discerned using the geochemical data, although those veins 

that are high in arsenic (arsenian type) are in the Advance, 

Republic, Lucky Hobo, and Gold Stake mines, as he noted; 

these veins, however, contain some of the highest values of 

lead and other base metals, in contrast to his descriptions of 

the minerals in these veins. Overall, the analyses show that 

veins contain gold and silver as well as relatively high values 

of other indicator elements (As, B, Sb, Te, Tl; samples were 

not analyzed for Hg or U) typical of epithermal preciousmetal deposits (Silberman and Berger, 1986). Anomalous 

values of gold and silver are in veins in the Gotd Chief (highest overall gold value, at 3.9 ppm; sample 120 of Tingley, 

1984), Lucky Hobo (highest overall silver value, at 2,000 

ppm; sample 123 of Tingley, 1984), Republic, Advance, 

Lucky Chief, SOA, and Gold Stake mines and in a prospect 

(sample 126 of Tingley) southwest of the Contact Mine. 

Even more anomalous values are those of lead and zinc and, 

to a much lesser extent, of copper; high values of iron, magnesium, and manganese also are characteristic. These higher 

values of base metals and silver, in fact, provide evidence of 

vertical geochemical zoning in the Chief district (Silberman 

and Berger, 1986) and indicate, along with the field evidence, that the district has been deeply eroded. It is likely 

that the veins containing the highest amounts of gold and silver were removed by erosion long ago. 

CONCLUSIONS 

The Chief Mining District is a minor mining district, but 

the lessons it provides may help in finding ore deposits elsewhere in and near the Caliente caldera complex. The district 

contains deposits of a deeply eroded epithermal fissure-vein 

system. The heat source that led to the transportation of most 

metals was provided by the Cobalt Canyon stock (24.8 Ma), 

which underlies the district. Metals most likely were leached 

from country rocks, then precipitated during convective 

overturn of groundwater heated by the stock; some metals, 

however, may have been derived from hydrothermal solutions from the intrusion itself. The stock may represent an 

early magmatic event in the adjacent Caliente caldera complex, although it predates the known main episode of magmatism (23-14 Ma) in the complex, which has controlled 

mineralization of three other epithermal gold districts. The 

tectonic control on the location of the stock and the caldera 

complex is not well constrained but may result partly from 

an east-trending grain in underlying rocks, perhaps the result 

of pre-middle Tertiary east-striking faults, some of which 

H 14 Application of Structural Geology to Mineral and Energy Resources 
were reactivated, and partly from the Miocene episode of 

faulting. 

Most orebodies in the district are in brecciated fault zones 

of the Tertiary Stampede Detachment system, an east-dipping bedding-parallel normal fault that separates quartzite 

below the detachment from limestone and dolomite above it; 

the detachment was active prior to late Oligocene volcanism, 

well before mineralization. As in the major Delamar Mining 

District to the southwest, brittle quartzite was extremely 

shattered in and near faults, providing sites for mineralization. At some of these sites, ore was localized along highangle faults and joints in quartzite. Some of these structures 

probably formed during emplacement of the Cobalt Canyon 

stock, which may also have localized a porphyry copper ore 

deposit. Structures of the Miocene episode of high-angle 

oblique-slip faulting that began before 19 Ma may have 

localized some metals. Early phases of the Miocene episode 

of faulting in fact may have been synchronous with, and provided control on, emplacement of the Cobalt Canyon stock. 

Some ore could have resulted from younger igneous activity 

such as the porphyry of Meadow Valley Wash or from remobilization of existing orebodies by groundwater heated during deformation. The history and structure of the Chief 

District have many similarities to those of epithermal gold 

districts in theW alker lane of western Nevada (John and others, 1989; Hardyman and Oldow, 1991). Evidence for 

geochemical zoning in which amounts of gold decrease with 

depth and evidence for deep erosion that removed most of 

the roof over the Cobalt Canyon stock indicate that the likelihood is low for discovering new gold orebodies in the district. The possibility of finding other gold deposits in and 

near the Caliente caldera complex, however, is probably 

much better. 

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469,55 p. 

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volcanic units of the Caliente volcanic field, Lincoln County, 

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5, p. 17-24. 

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115, p. 435-436. 

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H16 Application of Structural Geology to Mineral and Energy Resources 
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A300, p. 407-434. 

Structural Setting of Chief Mining District, Nevada H17 

Chapter I 

Extensional Geometry in the Northern Grant 

Range, East-Central Nevada-Implications for 

Oil Deposits in Railroad Valley 

By KAREN LUND and L. SUE BEARD 

U.S. GEOLOGICAL SURVEY BULLETIN 2012 

APPLICATION OF STRUCTURAL GEOLOGY TO MINERAL AND 

ENERGY RESOURCES OF THE CENTRAL AND WESTERN UNITED STATES 
CONTENTS 

Abstract 11 

Introduction 11 

Geologic setting 12 

Attenuation structures 13 

Styles of attenuation 13 

Geometry of attenuation structures 15 

Age of extension 17 

Implications for oil exploration 17 

Conclusions 18 

References cited 18 

FIGURES 

1, 2. Maps showing: 

1. Location of study area and nearby oil fields, east-central Nevada 12 

2. Tectonic features and geology of northern Grant Range 14 

3. Cross sections showing tectonic relations in northern Grant Range 16 

Ill Contents 
Extensional Geometry in the Northern Grant Range, 

East-Central Nevada-Implications for 

Oil Deposits in Railroad Valley 

By Karen Lund1 and L. Sue Beard2 

Abstract 

Heterogeneous extension in the northern Grant Range is 

manifested by a stacked set of curviplanar low-angle attenuation faults that formed concurrent with arching of the range 

about a north-northwest axis. The style and amount of attenuation was controlled by the I ithologic character and structural 

depth of the units and, locally, by the geometry of the arch. On 

the east side of the range, a stacked array of low-angle attenuation faults are subparallel with bedding, attenuation is distributed across the many stacked fault zones, and except at highest 

crustal levels these faults do not cut to the surface. On the west 

side of the range (steeper side of the arch), the curviplanar lowangle attenuation faults converge into a single west-dipping 

fault zone along which the total amount of extension is probably greater than in other parts of the range. Here, the stratigraphic separation (or juxtaposition) of Mississippian units over 

Middle Cambrian units results in an extensional culmination. 

Windows into the culmination expose a distinct geometry in 

the curviplanar low-angle fault array which indicates that lowangle attenuation faulting was synchronous with arching of the 

range. 

Cross sections incorporating seismic and drill-hole data 

suggest that the low-angle attenuation faults in the range 

extend into Railroad Valley to the west of the Grant Range with 

no significant offset by high-angle normal faults. Therefore, the 

topographic expression of the Grant Range and Railroad Valley 

may be due to synchronous arching and low-angle faulting. 

These new data and interpretations provide insight into the 

enigmatic structures buried in Railroad Valley. Thus, both 

petroleum source and reservoir rocks in Railroad Valley oil 

fields are in relatively immature but extensively fractured rocks 

of the upper plate of the extensional culmination. 

1

U.S. Geological Survey, Box 25046, MS905, Denver, Colorado 

80225. 2

U.S. Geological Survey, Flagstaff, Arizona 86001. 

INTRODUCTION 

The northern Grant Range and adjoining Railroad Valley 

are in the middle of the northern Basin and Range province 

in east-central Nevada (fig. 1). The rocks in the Grant Range 

show effects of complex deformation but are nonetheless 

dominantly in proper stratigraphic order and upright. Railroad Valley, to the west, is one of the widest and deepest valleys in Nevada and contains the first oil well and most 

productive oil fields in the State. 

Previous mapping in the Grant, Horse, and White Pine 

Ranges (fig. 1) shows that the Paleozoic rocks are cut by 

many low-angle faults, but two different conclusions have 

been made about the age and origin of the faulting. Mesozoic 

to early Tertiary compressional tectonism was thought to 

have caused the younger-on-older low-angle faults observed 

in Paleozoic rocks of the northern Grant Range (Kirkpatrick, 

1960; Huttrer, 1963; Hyde, 1963; Hyde and Huttrer, 1970), 

as well as both the younger-on-older and older-on-younger 

low-angle faults in the southern Grant Range (Cebull, 1967, 

1970). Alternatively, mapping of both Paleozoic and Tertiary rocks in the northern Grant, Horse, and White Pine 

Ranges by Moores and others (1968) resulted in the conclusion that, because the structural geometry of Tertiary rocks 

is essentially the same as that of underlying Paleozoic rocks, 

folding and low-angle younger-on-older faults were formed 

by late Tertiary block uplift, basement extension, and compression. 

Recent work demonstrates that both of these deformational events and structural processes (late Mesozoic compression and late Tertiary extension) occurred in the Grant 

Range. In the central Grant Range, a pre-Late Cretaceous 

macroscopic east-vergent overturned fold in metamorphosed 

Cambrian rocks (Cebull, 1967), which probably formed in 

the Mesozoic (Fryxell, 1988), was cut by two sets of Tertiary 

low-angle attenuation faults-early east-dipping faults cut 

by later west-dipping faults (Fryxell, 1984, 1988). In the 

northern Grant Range, Cambrian and Lower Ordovician 

Extensional Geometry in the Northern Grant Range, East-Central Nevada 11 
10 20 KILOMETERS 

Figure 1. Index map showing location of study area and nearby oil fields (enclosed by dotted lines), east-central Nevada. 

rocks were also metamorphosed and ductilely deformed at 

moderate to deep structural levels before being sliced and 

attenuated by a series of younger-on-older faults that deform 

rocks of all ages exposed in the range. These younger-onolder faults are part of a system of low-angle attenuation 

faults (faults that thin the section) (Lund and Beard, 1986, 

1987; Camilleri and others, 1987; Lund and others, 1987, 

1988; Camilleri, 1989). 

The most recent structural event has traditionally been 

thought to be formation of the Grant Range and Railroad 

Valley by a down-to-the-west, steep normal fault(s) along 

the west flank of the range, a model that is in accordance 

with classic Basin and Range styles (Stewart, 1978). A steep 

range-front fault and related minor faults are commonly 

depicted on maps and subsurface structural and seismic 

interpretations (Moores and others, 1968; Bortz and Murray, 

1979; Guion and Pearson, 1979; Vreeland and Berrong, 

1979; Effimoff and Pinezich, 1981; Anderson and others, 

1983; Poole and Claypool, 1984; Kleinhampl and Ziony, 

1985; Flanigan, 1988; Read and Zogg, 1988; Veal and others, 1988). Minor offset of Holocene alluvial fans and lake 

terraces, as well as the presence of a prominent line of 

springs and a sharp gravity gradient (Lund and others, 1987, 

12 Application of Structural Geology to Mineral and Energy Resources 

1988; Blank, 1991), indicates that some faulting has 

occurred along the range front. However, a sign-bit 

Vibroseis seismic line across the Grant Canyon oil field 

(Potter and others, 1991, fig. 1) also published by Veal and 

others (1988), indicates that there is no significant offset of 

the low-angle fault system by a major range-front fault at 

that latitude. 

The recently mapped structural geometry of the northern 

Grant Range and the interest in the geology of Railroad Valley because of oil exploration have provided reasons to try to 

integrate the new field data from the range with published 

drill-hole and seismic data from the valley. Our purpose is 

to provide a preliminary model for understanding the geometry and timing of formation of the valley relative to structures exposed in the range and for evaluating the structural 

setting of the oil occurrences in Railroad Valley. 

Acknowledgments.-Geologic mapping was accomplished during evaluation of the Blue Eagle and Riordans 

Well wilderness study areas by the U.S. Geological Survey. 

We are indebted to the numerous project members whose 

mapping contributed to our understanding of the geology of 

the northern Grant Range: B. Nevins, R. Demetrion, A.R. 

Pyke, D. Schneidereit, J.L. Snee, M. Elrick, J.E. Fryxell, and 

P.A. Camilleri. Discussions and (or) field trips with F.G. 

Poole, F.J. Kleinhampl, P.M. Sheehan (University of Wisconsin at Madison), R.J. Ross, Jr., J.G. Johnson (Oregon 

State University), and M.E. Taylor helped us with stratigraphic problems. We thank R.B. Scott, D.M. Miller, C.H. 

Thorman, H.R. Blank, Jr., W.J. Perry, Jr., C.J. Potter, J.A. 

Grow, and S.E. Boyer (University of Washington) for helpful discussions about aspects of the structure. Reviews by 

W.J. Perry, Jr., and A.R. Wallace improved the manuscript. 

GEOLOGIC SETTING 

During the Paleozoic, this area was on the continental 

shelf of western North America and was the site of deposition of more than 6,000 m (Kellogg, 1963; Moores and others, 1968) of predominantly carbonate rocks. In the northern 

Grant Range, Upper Cambrian to Middle Ordovician carbonate rocks are variably metamorphosed to greenschist 

facies; they include phyllite, phyllitic marble, and calcareous 

and dolomitic marble (Lund and others, 1987, 1988; Camilleri, 1989). Rocks higher in the stratigraphic section show no 

visible effects of metamorphism. Upper Ordovician to Middle Devonian units form a thick dolostone section (Ely 

Springs, Laketown, Sevy, and Simonson Dolostones); the 

Middle and Upper Devonian Guilmette Formation is a thick 

unit of interbedded limestone and dolostone. The Upper 

Devonian and Lower Mississippian Pilot Shale is absent 

because of nondeposition (Poole and others, 1977). The 

Lower Mississippian Joana Limestone contains both massive encrinitic limestone and dark limestone with argillaceous partings. The Mississippian Chainman Shale is the 
older of two oil source beds in the section (Poole and 

Claypool, 1984 ). The thick Upper Mississippian to 

Pennsylvanian Ely Limestone is the youngest Paleozoic unit. 

Rocks of the northern Grant Range were not deformed during the Paleozoic orogenies; the area is east of the regions 

most affected by deformation associated with the Antler 

(Devonian and Mississippian) and Sonoma (Permian and 

Triassic) orogenies (fig. 1). 

Mesozoic terrigenous deposits are absent because of 

either nondeposition or erosion. Lacustrine deposits of the 

Sheep Pass Formation are Paleocene and Eocene in age in 

the northern Grant Range (Fouch, 1979); the Sheep Pass Formation is the second principle oil source rock (Poole and 

Claypool, 1984). Beds of the Sheep Pass Formation disconformably overlie Paleozoic strata (Moores and others, 1968). 

These beds accumulated in local basins (Fouch, 1979) that 

may have been formed by upper crustal structural extension 

(Vandervoort and Schmitt, 1990) in the hinterland during 

Sevier (late Mesozoic to early Tertiary) thrust faulting. 

As much as 1,700 m of Oligocene to Miocene rhyolitic 

welded and nonwelded tuffs were deposited disconformably 

on the older rocks (Moores and others, 1968). They have an 

uneven distribution and both local and distant sources 

(Cook, 1960; Scott, 1965; Moores and others, 1968; Best and 

others, 1989). The tuffs were deposited prior to or during the 

early phases of crustal extension. Syntectonic sedimentary 

rocks of the Miocene and Pliocene Horse Camp Formation 

both gradationally and unconformably overlie the volcanic 

rocks (Moores, 1968). As much as 3,300 m of the Horse 

Camp Formation were deposited in the northern Grant 

Range and in Railroad Valley; this thickness is a composite 

figure and less than this can be expected in the Horse Camp 

sedimentary basins. In actual thickness, Oligocene to 

Holocene deposits probably do not exceed 3,500 min Railroad Valley or the Horse Camp Basin in the northern Grant 

Range (Moores, 1968; Bortz and Murray, 1979; Flanigan, 

1988; Read and Zogg, 1988; Veal and others, 1988). 

ATTENUATION STRUCTURES 

The structural geometry of the northern Grant Range is 

dominated by a stacked set of younger-on-older low-angle 

attenuation faults that omit strata; most of these faults are at 

low angles to bedding and separate the stratigraphic units 

into structural plates. Both the strata and low-angle attenuation faults are asymmetrically arched about a north-northwest-trending axis. Along the range crest, beds are almost 

horizontal, whereas those on the east side dip east and those 

on the west flank dip west (fig. 2); the west limb of the arch 

is steeper than the east. The style and amount of attenuation 

on the low-angle faults and in the intervening plates was controlled by lithologic character and structural position (depth) 

and is also related to location on the arch. 

Styles of Attenuation 

On the east side of the range at highest structural levels, 

attenuation was accomplished by large-scale block rotation 

above low-angle attenuation faults. Large blocks of Mississippian to Tertiary rocks, which generally amount to a kilometer or more of stratigraphic section, form east-dipping 

ridges on the east side of the range crest. Bedding attitudes 

in these blocks are from about 20° to 80° to the east. The 

blocks were tilted during attenuation, mostly along lowangle attenuation faults in the Chainman Shale and Sheep 

Pass Formation but also in the Joana Limestone, at various 

levels in the volcanic section, and in the syntectonic Horse 

Camp Formation. These faults acted as sole faults or decollement zones. Thus, Mississippian to Pennsylvanian and 

younger rocks are structurally decoupled from older rocks. 

At the decoupling zones, stratigraphic section is attenuated 

or eliminated; silicification (jasperoid formation) is particularly common along faults in the Chainman Shale and Joana 

Limestone. The rotated blocks are remarkably intact, and 

there is little internal brecciation. 

Rocks of the middle and lower structural levels are separated into multiple structural plates by a prominent series of 

stacked low-angle attenuation faults and are not involved in 

the large-scale rotated blocks that characterize structures of 

the higher level. Faults from middle and lower levels do not 

cut to the surface in the northern Grant Range. 

At middle structural levels, the massive dolostones that 

dominate the Ordovician to Devonian rocks are separated 

into several structural plates or domains by bounding lowangle attenuation faults that are subparallel with bedding. 

Thus, individual structural plates are commonly lithologic 

packages, and, conversely, stratigraphic units are generally 

attenuated and bounded by low-angle attenuation faults. 

Important faults of this type are present at the top and bottom 

of the Guilmette Formation, near the base of the Simonson 

Dolostone, near the Silurian-Devonian boundary, and at the 

top of the Eureka Quartzite. The low-angle attenuation faults 

are, in part, products of (and result from) many high- to 

moderate-angle, small- to medium-scale faults that 

coalesced into the specific stratigraphic horizons. Thus, the 

low-angle attenuation faults at these stratigraphic horizons 

acted as zones of decoupling between separately extending 

structural plates. 

Attenuation internal to structural plates at middle levels 

occurred by many different types of extensional structures. 

Minor low- to moderate-angle attenuation faults, which are 

difficult to map, caused bedding discordance throughout the 

section. On the west side of the range, major sets of downto-the-east, medium-scale attenuation faults are isolated 

within several different structural plates. These minor fault 

sets are antithetic to movement on the main system and 

caused bedding to rotate to the west, thereby steepening the 

dip of rocks westward into Railroad Valley. Elsewhere, 10-

m-scale domino-style attenuation faults are also present. 

Paleozoic carbonate rocks at middle structural levels are 

Extensional Geometry in the Northern Grant Range, East-Central Nevada 13 
14 Application of Structural Geology to Mineral and Energy Resources 
CJ 

u 

-

D 

EXPlANATION 

Alluvial and colluvial deposits (Quaternary) 

Horse Camp Formation (Pliocene and Miocene), undivided 

volcanic rocks (Oligocene), and parts of Sheep Pass 

Formation (Eocene and Paleocene) 

Granite (Late Cretaceous) 

Ely Limestone (Pennsylvanian and Upper Mississippian), 

Chainman Shale (Mississippian), and parts of Joana 

Limestone (Lower Mississippian) 

Guilmette Formation (Upper and Middle Devonian) and 

parts of Joana Limestone (Lower Mississippian) 

Simonson Dolostone (Middle Devonian), Sevy Dolostone 

(Lower Devonian and Upper Silurian), Laketown Dolostone 

(Silurian), and parts of Ely Springs Dolostone (Lower 

Silurian and Upper Ordovician) 

Eureka Quartzite (Middle Ordovician), parts of Ely Springs 

Dolostone (Upper Ordovician), and parts of Pogonip 

Group (Middle Ordovician) 

Pogonip Group (Middle and Lower Ordovician), variably 

metamorphosed 

Metasedimentary rocks (Upper and Middle Cambrian), 

undivided 

---Contact 

-....,....- High·angle normal fault-Dip greater than approximately 

30°. Dashed where approximately located; dotted where 

concealed; bar and ball on downthrown side 

... 

. . . 

--+--

---q40 

_!}_ 

Low-angle fault-Dip less than approximately 30° 

Fault that omits strata-May occur in metamorphosed 

or unmetamorphosed rocks; dotted where concealed. 

Decoration on upper plate 

Fault that repeats strata-Saw teeth on upper plate 

Strike and dip of bedding 

Strike of vertical bedding 

Strike and dip of overturned bedding 

Strike and dip of foliation 

Plunge and trend of minor fold, showing sense of asymmetry 

---1.{''--l+r--- Overturned syncline 

-+G+9+-- Inverted synformal anticline, verging north 

Trace of layering and axial planar cleavage in meta· 

morphosed rocks 

Figure 2 (above and facing page). Tectonic map of northern 

Grant Range, east-central Nevada. Lines of sections A-A', 

8-8' -8", and C-C' (fig. 3) are also shown. Location of subsurface data used in cross sections indicated by names of drill 

holes: Shell 1, GC1 (Grant Canyon 1 ), GC2 (Grant Canyon 2), 

and Thorn 1. Geology mapped by K. Lund and L.S. Beard, 

1984-1987; B. Nevins, 1984; R. Demetrion, A.R. Pyke, and 

D. Schneidereit, 1985; P.A. Camilleri, M . Elrick, and j.E. Fryxell, 1986; J.L. Snee, 1987; geology of Tertiary volcanic rocks 

simplified from Scott (1965). 

commonly brittlely attenuated and form spectacular breccias; millimeter-scale fragments dominate the dolostone 

breccias, whereas centimeter- to meter-scale fragments characterize breccias in the limestone beds. 

Below the Eureka Quartzite at lower structural levels, 

Middle Ordovician and older rocks are metamorphosed and 

ductilely deformed (presumably during Mesozoic shortening). Brittle attenuation structures are superimposed on earlier ductile structures (Lund and others, 1987, 1988; 

Camilleri, 1989). Macroscopically, the styles of attenuation 

at this lowest level are similar to those of the middle level 

except that brittle deformation is less pervasive. Low-angle 

attenuation faults that are subparallel with bedding and with 

cleavage also attenuate strata at this level, although associated minor attenuation faults are less common. Breccia 

zones generally only a few meters thick perhaps reflect the 

deeper structural level or the pre-existing metamorphic 

fabric. 

Geometry of Attenuation Structures 

The stacked low-angle attenuation faults display a peculiar geometry. On the east side of the range, the low-angle 

attenuation faults are subparallel with bedding and with each 

other. On the west limb of the arch, the stacked array of lowangle attenuation faults from all structural levels converges 

into a single 18°-30° west-dipping low-angle attenuation 

fault zone that extends westward toward Railroad Valley 

(fig. 3). Where the stacked array of low-angle attenuation 

faults converges, an extensional culmination is formed. This 

culmination is well exposed in extensional eyelid windows 

(Boyer and Elliot, 1982) in several canyons (or valleys) on 

the west side of the range. In the windows, a stack of at least 

five major low-angle attenuation faults, which are seen on 

the east side of the windows, converges to form a single 

west-dipping fault on the west side of the windows. The convergence forms a consistent geometric pattern: deeper lowangle attenuation faults are more arched than the higher level 

faults and appear to be truncated where they merge with the 

higher level faults (Lund and Beard, unpublished mapping, 

1984-1987; Lund and Beard, 1987; Camilleri, 1989, 1990; 

Lund and others, 1991). This pattern in a stacked fault 

system, where older faults are more tightly folded than 

younger, indicates that folding and faulting were contemporaneous (Gilluly, 1960). 

The single west-dipping fault zone cuts down section in 

the footwall to the west and includes lenses and slivers (as 

thick as 200 m) of the intervening cut-out plates; these slivers are similar to features described elsewhere as chaos 

structures (Wright and Troxel, 1969), nappes (Wernicke, 

1981; Lister and Davis, 1989), and extensional duplexes 

(Gibbs, 1984; Root, 1990) and have been noted in and east 

of the White Pine Range to the north (Walker and others, 

1986). 

Additionally, several examples of macroscopic to mesoscopic overturned folds and older-on-younger faults (faults 

that omit section, fig. 2) are present at specific locations with 

respect to the low-angle attenuation fault system (Lund and 

others, 1991). These enigmatic structures are in the upper 

Extensional Geometry in the Northern Grant Range, East-Central Nevada 15 
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~ c: 

@ 

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B B' B" 

zl;; -c 

Dj: 

~~ <DO 

GC 1 GC 2 Thorn 1 

c C' 

5 KILOMETERS 

Figure 3. Cross sections showing tectonic relations in northern Grant Range, east-central Nevada. Lines of sections shown in figure 2; units same as in figure 2; no vertical 

exaggeration. Subsurface information generalized from drill hole data: Shell1 (Bortz and Murray, 1979); GC1 (Grant Canyon 1 ), GC2 (Grant Canyon 2), and Thorn 1 (Veal 

and others, 1988). Surface data from Lund and others (1987, 1988). 
plate to the west-dipping fault, and in areas where the 

stacked low-angle attenuation faults converge. They are also 

in or near east-trending down-to-the-north fault zones that 

are transverse to the main north-northwest-trending arch of 

the range (fig. 2). Where asymmetric, structures near the 

major down-to-the-west fault are west vergent. Those on the 

down-to-the-north zones show north vergence. These structures are local, isolated from each other by extensional 

faults, and cannot be reconstructed as continuous pre-extensional features. The local shortening structures, although 

possibly the result of Mesozoic compression, were alternatively possibly caused by complex extensional duplexing 

and (or) by local shortening associated with the overall 

extension. 

The sense of motion along the low-angle attenuation 

faults is inferred from the geometry of the fault system. No 

asymmetric ductile fabrics or striated surfaces were found 

associated with the fault zones (Lund and Beard, unpublished mapping, 1984-1987; Camilleri, 1989), and, probably 

because the fault zones are subparallel with bedding along 

most of their length, no offset elements were identified along 

the low-angle faults. Observation of the faults in the fanned 

array indicates that relatively more stratigraphic section is 

omitted and apparent displacement increases westward 

along each fault. The increase in apparent displacement may 

be accounted for in part by mesoscopic extensional deformation within the structural plates between each of the attenuation faults; this internal strain may accumulate westward as 

increased slip on individual attenuation faults. In addition, 

the entire fault system cuts down section to the west and fanning in the fault system converges westward (Lund and 

Beard, unpublished mapping, 1984-1987; Camilleri, 1989; 

Lund and others, 1991 ). This geometry suggests that the 

extensional motion was to the west; this is opposite to the 

sense that would be inferred from an imbricate fan in a thrust 

setting in which the thrust faults branch in the direction of 

motion (Gilluly, 1960; Boyer and Elliot, 1982). In summary, 

the major west-dipping low-angle attenuation fault has the 

geometry of a west-dipping, west-directed ramp that defines 

a discrete zone of localized extension to the west, compared 

with the more diffuse extension in the east part of the range 

where the fan of attenuation faults is found. 

Age of Extension 

Deposition of the Oligocene volcanic sequence probably 

mostly predates extension in the northern Grant Range 

because there is little angular discordance at the base of or 

within the volcanic sequence (Scott, 1965; Moores and others, 1968); however, deposition of the Miocene and Pliocene 

Horse Camp Formation (which remains poorly dated) probably spanned a period of major extension because units display the geometry of growth-faulted sedimentary deposits 

and include syntectonic deposits (Moores, 1968; Moores and 

others, 1968). Lower strata ofthe Horse Camp Formation are 

tilted as much as 80°, and tilting decreases upward such that 

younger strata exhibit only minor tilting (Moores and others, 

1968). In the central Grant Range, extensional faults cut a 

tuff dated at 27.8 Ma (Fryxell, 1988). Taylor and others 

(1989) show that, in the Quinn Canyon Range at the south 

end of the Grant Range, early extension occurred between 32 

and 27 Ma and that peak extension occurred about 14 Ma. 

IMPLICATIONS FOR OIL EXPLORATION 

Oil from source rocks in the Mississippian Chainman 

Shale and Paleocene and Eocene Sheep Pass Formation has 

been produced from reservoirs in Oligocene volcanic rocks, 

Sheep Pass Formation, Upper Mississippian to Pennsylvanian Ely Limestone, Middle and Upper Devonian Guilmette 

Formation, Upper Silurian and Lower Devonian Sevy and 

Middle Devonian Simonson Dolostones, and even from 

Middle Ordovician Eureka Quartzite (Poole and Claypool, 

1984; Flanigan, 1988; Read and Zogg, 1988). The oil traps 

have long been recognized as structural (Bortz and Murray, 

1979; Foster, 1979), but the low-angle attenuation faults in 

the northern Grant Range have only recently been modelled 

in cross sections of Railroad Valley (Flanigan, 1988; Read 

and Zogg, 1988). Although these new models incorporate 

low-angle attenuation faults into the cross sections, neither 

the overall structural geometry nor individual faults have 

been previously linked or integrated from the Grant Range to 

Railroad Valley. 

Published seismic lines for the east side of Railroad Valley show prominent low-angle reflectors, but these reflectors 

are unexplained on interpretive cross sections (Bortz and 

Murray, 1979; Anderson and others, 1983; Veal and others, 

1988). Many drill holes in the Eagle Springs, Kate Spring, 

and Grant Canyon fields have been drilled through upper 

Paleozoic (or younger) rocks directly into early Paleozoic 

metamorphic rocks or into plutonic rocks (Bortz and Murray, 1979; Flanigan, 1988; Veal and others, 1988); the plutonic rocks are thought to be equivalent to the Late 

Cretaceous Troy granite (Fryxell, 1988). These metamorphic and plutonic rocks form the lower plate culmination. 

Thus, the low-angle fault system that dominates the structural geometry of the range can also be extended beneath the 

valley. For the Grant Canyon oil fields (fig. 3), the depth 

information from drill holes and the attitudes of reflectors on 

seismic reflection profiles show that the major west-dipping 

low-angle fault on the west side of the northern Grant Range 

continues into Railroad Valley at the same attitude as in the 

range and with no significant offset (Potter and others, 

1991). For Kate Spring and Eagle Springs oil fields to the 

north, modification of the controlling low-angle attenuation 

fault system by minor steeper normal faults is more plausible. Significantly, the Railroad Valley oil reservoirs are all in 

complexly faulted and brecciated rocks (Bortz and Murray, 

1979; Poole and Claypool, 1984; Flanigan, 1988; Read and 

Extensional Geometry in the Northern Grant Range, East-Central Nevada 17 
Zogg, 1988; Veal and others, 1988) that are in the upper 

plate to this major west-dipping fault. 

Previously published petroleum exploration models indicate that oil was found in areas where seismic studies show 

topographic highs buried beneath valley fill (Bortz and Murray, 1979; Foster, 1979; Flanigan, 1988). These highs were 

interpreted to be horst blocks formed by minor high-angle 

normal faults that were part of a high-angle range-front fault 

system. Alternatively, our study indicates that these buried 

highs may reflect the irregular topography of the pre-Quaternary (pre-burial) erosional surface that developed on the 

complexly faulted upper plate of the low-angle attenuation 

fault system. 

CONCLUSIONS 

The major structural event preserved in the northern 

Grant Range is middle to late Tertiary extension that was 

manifested by synchronous low-angle attenuation faulting 

and arching. Cross sections incorporating seismic and drillhole data show that the low-angle attenuation fault system in 

the range extends into Railroad Valley. Although high-angle 

faults do occur along the range front based on surface geologic and geophysical data, the low-angle attenuation fault 

system in the area of Grant Canyon oil field is probably not 

offset (fig. 3); more modification of the low-angle fault system could be expected to the north. Thus, the relief between 

the range and valley may have been largely caused by arching synchronous with formation of the low-angle attenuation 

fault system, which dips only 18°-30° west, rather than by a 

high-angle range-front fault. 

This structural geometry suggests the need for a new 

model for petroleum exploration in the region. In Railroad 

Valley, both petroleum source and reservoir rocks are in the 

upper plate to the structural culmination that was formed by 

the low-angle attenuation fault system; rocks in the upper 

plate are relatively immature and extensively fractured. The 

intense fracturing of upper plate rocks and the juxtapositioning of immature source rocks over less fractured, metamorphosed rocks were instrumental in localizing petroleum 

reservoirs in Railroad Valley. 

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evolution of some basins in the northern Basin and Range 

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Geology of White River Valley, the Grant Range. eastern 

18 Application of Structural Geology to Mineral and Energy Resources 

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Boyer, S.E., and Elliot, David, 1982, Thrust systems: American 

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1196-1230. 

Camilleri, P.A., 1989, Superposed compressional and extensional 

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Camilleri, P.A., Lund, Karen, and Beard, L.S., 1987, Superposed 

compressional and extensional strain in metamorphosed 

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Cebull, S.E., 1967, Bedrock geology of the southern Grant Range, 

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___ 1970, Bedrock geology and orogenic succession in southern 

Grant Range, Nye County, Nevada: American Association 

of Petroleum Geologists Bulletin, v. 54, p. 1828-1842. 

Cook, E.F., 1960, Great Basin ignimbrites, in Boettcher, J.W., and 

Sloan, W.W., Jr., eds., Geology of east-central Nevada: 

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Effimoff, Igor. and Pinezich, A.R., 1981, Tertiary structural development of selected valleys based on seismic data. Basin and 

Range province, northeastern Nevada, in Extensional tectonics associated with convergent plate boundaries: Royal Society of London Philosophical Transactions, ser. A, v. 300, no. 

1454, p. 435-442. 

Flanigan, D.M.H., 1988, Kate Spring field discovery, Nevada 

Basin and Range: The Mountain Geologist, v. 25, p. 

159-169. 

Foster, N.H., 1979, Geomorphic exploration used in the discovery 

of Trap Spring oil field, Nye County, Nevada, in Newman, 

G.W., and others, eds., Basin and Range symposium and 

Great Basin field conference: Denver, Rocky Mountain 

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Fouch, T.D., 1979, Character and paleogeographic distribution of 

Upper Cretaceous(?) and Paleogene nonmarine sedimentary 

rocks in east-central Nevada, in Armentrout, J.M., Cole, 

M.R.. and TerBest. H., Jr., eds., Cenozoic paleogeography of 

the western United States: Society of Economic Paleontologists and Mineralogists, Pacific Coast Paleogeography Symposium 3, p. 97-112. 

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of North Carolina, Ph.D. thesis, 139 p. 

___ 1988, Geologic map and descriptions of stratigraphy and 

structure of the west-central Grant Range, Nye County, 

Nevada: Geological Society of America Map and Chart 

Series MCH064, 15 p., scale 1:24,000. 

Gibbs, A.D., 1984, Structural evolution of extensional basin margins: Geological Society of London Journal, v. 141, p. 

609-620. 

Gilluly, James, 1960, A folded thrust in Nevada-Inferences as to 

time relations between folding and faulting: American Journal of Science, Bradley Volume. v. 258-A, p. 68-79. 
Guion, D.J., and Pearson, W.C., 1979, Gravity exploration for 

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and others, eds., Basin and Range symposium and Great 

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Huttrer, G.W., 1963, Structure and stratigraphy of the central Grant 

Range, Nevada: Seattle, University of Washington, M.S. thesis, 59 p. 

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Grant Range, Nevada: Seattle, University of Washington, 

M.S. thesis, 63 p. 

Hyde, J.H., and Huttrer, G.W .. 1970, Geology of central Grant 

Range, Nevada: American Association of Petroleum Geologists Bulletin, v. 54, p. 503-521. 

Kellogg, H.E., 1963, Paleozoic stratigraphy of the southern Egan 

Range, Nevada: Geological Society of America Bulletin, v. 

74, p. 685-708. 

Kirkpatrick, D.H., 1960, Structure and stratigraphy of the northern 

portion of the Grant Range, east-central Nevada: Seattle, 

University of Washington, M.S. thesis, 79 p. 

Kleinhampl, F.J., and Ziony, J.I., 1984, Geology of northern Nye 

County, Nevada: Nevada Bureau of Mines and Geology Bulletin 99A, 172 p. 

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core complexes and detachment faults formed during Tertiary continental extension in the northern Colorado River 

region, U.S.A.: Journal of Structural Geology, v. 11, p. 

65-94. 

Lund, Karen, and Beard, L.S., 1986, Structural history of the northern Grant Range, east-central Nevada-Overprinting of 

structural styles: Geological Society of America Abstracts 

with Programs, v. 18, p. 392. 

___ 1987, Heterogeneous extensional geometries in the northern Grant Range, east-central Nevada: Geological Society of 

America Abstracts with Programs, v. 19, p. 751. 

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Hamilton, M.M., 1988, Mineral resources of the Riordans 

Well Wildemess Study Area, Nye County, Nevada: U.S. 

Geological Survey Bulletin 1731-H, 16 p. 

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the northern Granty Range and Railroad Valley, Nye county, 

Nevada: Implications for oil occurrences, in Flanigan, 

D.M.H .. Hansen, Mike, and Flanigan, T.E., eds .. Geology of 

White River Valley, the Grant Range, eastern Railroad Valley and western Egan Range, Nevada: Nevada Petroleum 

Society, 1991 Fieldtrip Guidebook, p. 1-6. 

Lund, Karen, Nash, J.T., Beard, L.S., Blank, H.R., Jr., and Tuftin, 

S.E., 1987. Mineral resources of the Blue Eagle Wilderness 

Study Area, Nye County, Nevada: U.S. Geological Survey 

Bulletin 1731-D, 19 p. 

Moores. E.M., 1968, Mio-Pliocene sediments, gravity slides, and 

their tectonic significance, east-central Nevada: Journal of 

Geology, v. 76. p. 88-98. 

Moores, E.M., Scott, R.B.. and Lumsden, W.W., 1968, Tertiary tectonics of the White Pine-Grant Range region, east-central 

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Hydrocarbon source rocks of the greater Rocky Mountain 

region: Denver, Rocky Mountain Association of Geologists, 

p. 179-229. 

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580-581. 

Extensional Geometry in the Northern Grant Range, East-Central Nevada 19 

Chapter J 

Late Paleozoic Structure of the 

Southern Part of the Uinta Basin, 

Utah, from Seismic Reflection Data 

By CHRISTOPHER J. POTTER, REX TANG, and 

TIMOTHY J. HAINSWORTH 

U.S. GEOLOGICAL SURVEY BULLETIN 2012 

APPLICATION OF STRUCTURAL GEOLOGY TO MINERAL AND 

ENERGY RESOURCES OF THE CENTRAL AND WESTERN UNITED STATES 
CONTENTS 

Abstract Jl 

Introduction Jl 

Regional setting J2 

Regional late Paleozoic stratigraphy-Correlation with reflection data J3 

Data J4 

Structural patterns and thickness variations J6 

Discussion J6 

References cited J8 

PLATE 

[Plate is in pocket] 

1. Seismic reflection lines in southern part of Uinta Basin, Utah. 

FIGURES 

1, 2. Maps showing: 

1. Regional geologic setting of study area, northeastern Utah J2 

2. Late Paleozoic geologic setting of study area in southern part of Uinta 

Basin J2 

3. Chart showing upper Paleozoic stratigraphy of study area J3 

4, 5. Maps showing: 

4. Location of seismic reflection lines examined in study J3 

5. Thickness variations for Doughnut Formation to Kaibab Limestone 

interval J4 

6. Sketch showing seismic reflection patterns observed for late Paleozoic faulting in 

study area JS 

7, 8. Cross sections showing: 

7. Variations in thickness of Doughnut Formation to Kaibab Limestone 

interval J6 

8. Imbricate thrust structure east of Price J6 

9. Map showing Paleozoic geologic setting of study area and major faults identified in 

study. J7 

J II Contents 
Late Paleozoic Structure of the Southern Part of the 

Uinta Basin, Utah, from Seismic Reflection Data 

By Christopher J, Potter 1, Rex Tang2, and Timothy J. Hainsworth3 

Abstract 

Seismic reflection data from the southern part of the 

Uinta Basin near Price, Utah, reveal a network of late Paleozoic faults that produced abrupt variations in stratigraphic 

thicknesses in a structurally complex 30-mi (48 km)-wide 

northwest-southeast-trending trough. This zone is west of the 

ancestral Uncompahgre Uplift, north of the Emery Uplift, and 

east of the Oquirrh Basin. In this zone, Pennsylvanian and Permian clastic rocks locally are several thousand feet thicker than 

in directly adjacent areas to the north and south. 

Faults that displaced strata below the Kaibab Limestone 

but did not displace the Kaibab were active between Chesterian (top of Doughnut Formation) and Leonardian (Kaibab 

Limestone) time. High-angle reverse and normal faults occur in 

a block-faulting pattern similar to that observed in upper Paleozoic rocks in the Piceance basin in northwestern Colorado; 

thrust faults occur in a local north-south trending belt. In localized zones of strong Mesozoic and (or) Cenozoic reactivation 

of the late Paleozoic faults, original fault styles are not clear. 

West of Price, along the northern edge of the Emery 

Uplift, the principal Paleozoic faults are northwest-striking, 

northeast-directed reverse faults that have as much as 5,000 ft 

(1,525 m) of throw. This fault set is almost colinear with the 

Uncompahgre Fault (which also underwent late Paleozoic 

reverse movement) but dips in the opposite direction. Pennsylvanian rocks are locally absent on reverse-fault-bounded structural highs. Steep northwest-striking normal faults are also 

present. 

East and northeast of Price, west-directed thrust faults 

(and several east-directed backthrusts) controlled late Paleozoic deformation. High-angle reverse and normal faults are 

also present. 

The major northwest-striking reverse faults and the westdirected thrusts accommodated displacements that are compatible with late Paleozoic oblique sinistral-reverse slip on the 

Uncompahgre Fault. The major reverse faults record a shorten1

U.S. Geological Survey, Box 25046, MS939, Denver, Colorado 

80225. 2

The Louisiana Land and Exploration Company, 2950 North Loop 

West, Suite 1200, Houston, Texas 77092. 3

P.O. Box 461841, Aurora, Colorado 80046. 

ing direction parallel to that associated with the Uncompahgre 

Fault. The west-directed thrust faults probably formed in 

response to the strike-slip component of movement of the 

Uncompahgre block as it moved west against the late Paleozoic basin. The probable genetic tie between the Uncompahgre fault and faults in the study area implies that the latter faults 

were most active during Desmoinesian time. Northwest-trending normal faults may record a relaxation of regional compressive stresses following Ancestral Rockies orogenesis. This 

interpretation relates all of these late Paleozoic faults to Ancestral Rockies tectonism. Most of the late Paleozoic faults in the 

study area have undergone only minor post-Paleozoic reactivation or deformation, although major Mesozoic and Cenozoic reactivation locally has occurred. 

INTRODUCTION 

Regional unconformities and thick sedimentary basins in 

the central Rocky Mountains, Colorado Plateau, and northeastern part of the Great Basin preserve an impressive record 

of late Paleozoic ("Ancestral Rockies") orogenesis, but 

many regional structures of this age have been obscured by 

younger tectonism, basin development, and volcanism. With 

few exceptions, major late Paleozoic faults at the margins of 

Precambrian exposures have been reactivated or overprinted 

by Laramide-age (Late Cretaceous to Eocene) or younger 

faults. In order to understand late Paleozoic faulting, it is 

necessary to separate younger movement from Paleozoic 

movement or to identify unreacti vated Paleozoic structures. 

Identification of late Paleozoic faults that have not been 

reactivated (or have undergone only minor reactivation) is 

best accomplished using subsurface (well and seismic) data 

(for example, Frahme and Vaughn, 1983 ). In this study, seismic reflection data from the southern part of the Uinta Basin, 

near Price, Utah, are used to illustrate the style of late Paleozoic deformation northwest of the ancestral Uncompahgre 

Uplift. 

Acknowledgment.-We thank Grant Norpac, Inc., for 

permission to publish seismic reflection sections EU-1-85, 

EU-5, EU-6, and EU-9. 

Late Paleozoic Structure from Seismic Reflection Data, Uinta Basin, Utah Jl 
REGIONAL SETTING 

The study area (fig. 1) is on the south flank of the Laramide-age Uinta Basin, directly north of the San Rafael 

Swell. Tertiary faults (for example, the Clear Creek and Fish 

Creek Grabens; Osmond, 1965), as well as numerous late 

Paleozoic fault blocks and small basins that can be defined 

on seismic reflection data (this report), are present in the 

area. The Paleozoic basins and faults accommodated displacements in a zone bounded by four major late Paleozoic 

geologic features: the ancestral Uncompahgre Uplift, the 

Emery Uplift, the Paradox Basin, and the Oquirrh Basin (fig. 

2). Buried late Paleozoic faults have also been identified in 

seismic data from the Piceance Basin of northwestern Colorado, northeast of the Uncompahgre Uplift (Waechter and 

Johnson, 1985), and the San Juan Basin of northwestern 

New Mexico (Huffman and Taylor, 1989). The style and 

regional significance of buried late Paleozoic structures in 

the study area are examined in this report. 

The late Paleozoic faults that bound the ancestral 

Uncompahgre Uplift (fig. 2) can be traced into the study area 

in the subsurface (Stone, 1977). The Uncompahgre Fault, on 

the southwest side of the uplift, underwent southwestdirected reverse motion in late Paleozoic time (Frahme and 

40' 

39' 

UINTA BASIN 

-.../ 

[if 

ct 

iii 

i! 

rl I I Jurassic outcrop limit 1 

0 50 MILES 

61-----r-----'~.,..__.,-+5-o -KI~LO_M_.J~TERS 

Figure 1. Regional geologic setting of study area (screened 

pattern), northeastern Utah. CCG, Clear Creek Graben; FCG, 

Fish Creek Graben. 

12 Application of Structural Geology to Mineral and Energy Resources 

112° 111 ° 110'J 109° 41".-------.l---------,-----------,----~.------,-!11 ,---- ~---, 

OQUIRRH ~ 

BASIN WYOMING I SHELF 

I 

40°-

EMERY 

39°- UPLIFT 

0 50 100 MILES 

I 

0 50 100 KILOMETERS 

Figure 2. Late Paleozoic geologic setting of study area in 

southern part of Uinta Basin, Utah. 

Vaughn, 1983; White and Jacobsen, 1983). Stone (1977) 

inferred that there was also a significant left-lateral component to late Paleozoic displacement on the Uncompahgre 

Fault, whereas Stevenson and Baars (1986) argued for rightlateral motion. The Garmesa Fault forms the northeast 

boundary of the uplift. It underwent much less displacement 

than the Uncompahgre Fault; movement was oblique, dominated by left-lateral slip (Stone, 1977). Late Paleozoic displac;ements on these faults diminished to zero within, or just 

east of, the study area, where displacements were accommodated along a system of intersecting shorter faults. Laramide 

reactivation occurred along the Garmesa Fault, east of the 

study area (Osmond, 1965), but apparently did not occur 

along the northwestern part of the Uncompahgre Fault (Heyman, 1983). 

Szabo and Wengerd (1975) documented a subsurface late 

Paleozoic fold and fault belt striking N. 60° W. beneath the 

Paradox Basin adjacent to the south flank of Uncompahgre 

Uplift. The northwestern end of this belt merges into the 

"Price graben" of Szabo and Wengerd (1975) between Green 

River and Price. The "Price graben" formed in late Paleozoic 

time and is bounded on the northeast by the Uncompahgre 

Fault and on the southwest by the Emery Fault (fig. 2) 

(Szabo and Wengerd, 1975); part of it occupies the southeastern part of the study area. The present study indicates 

that it is not a true graben and that the late Paleozoic structure 

of the Price area is more complex than that portrayed by 

Szabo and Wengerd (1975). 

Hite (1975) documented a northeast-striking fault set in 

the Paradox Basin and presented evidence for left-lateral 
displacement along these faults. Some of this displacement 

is clearly Jurassic or younger, but the earlier history of this 

fault set is poorly understood (Hite, 1975). 

Southwest of Price, the Emery Fault, oriented approximately N. 30° W., forms the northeastern boundary of the 

late Paleozoic Emery Uplift (Szabo and Wengerd, 1975) 

(fig. 2). Several N. 60° W.-trending faults developed within 

the Emery Uplift, oblique to the Emery Fault (Szabo and 

Wengerd, 1975). 

The San Rafael Swell (fig. 1) developed during the Laramide across the grain of the Emery Uplift. Laramide deformation is responsible for the present northerly regional dip in 

the study area, which is at the north-plunging end of the San 

Rafael Swell and on the north-dipping south flank of the 

Uinta Basin. Younger fault zones are prominent in the westem part of the study area on the Wasatch Plateau. These 

include the north-trending Clear Creek Graben and the 

northwest-trending Fish Creek Graben (Osmond, 1965) (fig. 

1). The seismic data examined in this study demonstrate that 

the Fish Creek Graben reactivated late Paleozoic faults. 

Regional late Paleozoic StratigraphyCorrelation With Reflection Data 

A stratigraphic column for the upper Paleozoic rocks of 

the study area is shown in figure 3. These strata are penetrated by deep wells at several locations, and key reflections 

from this interval can be traced on seismic reflection lines 

that can be tied to these wells (fig. 4). Identifications of key 

stratigraphic horizons were calibrated using sonic logs and 

velocity surveys in the wells. 

z 

<( 

~ a: 

LU 

c... 

z 

<( 

z 

<( 

> _J 

>-

(/) 

z 

z 

LU 

c... 

z 

<( 

0::: 

c... 

U5 

(/) 

U5 

(/) 

~ 

OCHOAN 

GUADALUPIAN 

LEONARDIAN 

WOLFCAMPIAN 

VIRGILIAN 

MISSOURIAN 

DESMOINESIAN 

~ ATOKAN '\.._ 

MORROWAN 

CHESTERIAN 

MERAMECIAN 

OSAGEAN 

KINDERHOOKIAN 

Weber 

and 

Morgan 

Formations 

Limestone 

Figure 3. Upper Paleozoic stratigraphy of study area in 

southern part of Uinta Basin, Utah. 

UTAH COUNTY 

DUCHESNE COUNTY --~---------~------

CARBON COUNTY 

BOOK CLIFFS 

I 

Sunnyside 0 

Mt. Fuel+ 

10 MILES 

10 KILOMETERS 

Figure 4. Location of seismic reflection lines examined in 

this study (heavy lines) and deep wells that penetrate upper 

Paleozoic section. Numbered lines (1, 5, 6, 9) refer to GrantNorpac seismic lines EU-1-85, EU-5, EU-6, and EU-9, 

which are shown on plate 1. Wells: Arco, Arco Hiawatha 1; 

Arcadia, Arcadia Telonis 1; American Quasar, American 

Quasar 311 Drunkards Wash; Hunt, Hunt 116 State; Mt. 

Fuel, Mountain Fuel 1 Sunnyside. 

The deepest Paleozoic reflector generally recognized in 

the area is the top of the Mississippian carbonate rocks correlative with the Redwall and Leadville Limestones. Franczyk ( 1991) uses the term Madison Limestone in the area 

north of the Emery Uplift, and the name Deseret Limestone 

is commonly used in industry logs of well and seismic data 

in the vicinity of Price. These rocks will be called Red wall in 

this report, following Hintze (1988). On the seismic lines 

examined in this study, the Redwall reflection (lines EU-5 

and EU-9, plate 1) is a doublet that cannot be traced entirely 

across any of the seismic lines, probably because of signal 

penetration problems. The relatively unreflective interval 

above the Redwall represents the Humbug Formation and 

the Doughnut Formation (commonly referred to as the Manning Canyon Shale on well logs in the study area). A welldefined doublet at the base of an overlying set of banded 

reflections (lines EU-5 and EU-9, plate 1) marks the top of 

the Doughnut. The banded reflection sequence corresponds 

to mixed clastic and carbonate rocks of the Pennsylvanian 

and Permian "undifferentiated Weber and Morgan Formations" (Franczyk, 1991 ). A prominent reflection from the 

Kaibab Limestone (of Leonardian and locally Guadalupian 

age; Hintze, 1988) lies above the Weber and Morgan reflections. 

In this study, local thickness variations are documented 

in the interval between the Doughnut and Kaibab reflections 

(fig. 5). This interval is occupied mainly by the Weber and 

Morgan Formations. In many cases, the thickness of the 

Weber and Morgan abruptly changes across Pennsylvanian 

or Permian faults. On seismic sections, these faults displace 

Mississippian reflections but do not displace the Kaibab 

reflection (lines EU-5 and EU-9, plate 1, fig. 6). They are 

Late Paleozoic Structure from Seismic Reflection Data, Uinta Basin, Utah JJ 
the primary evidence used to interpret the late Paleozoic 

structural evolution of the region and are mapped in the subsurface in this study (fig. 5). On some seismic sections, a 

"PI" (for lower Permian) reflection is identified within the 

Pennsylvanian and Permian section (fig. 6). This reflection 

is tentatively correlated with the top of an interbedded 

marine limestone, sandstone, and shale unit of Virgilian and 

Wolfcampian age (lower Cutler beds of Loope and others, 

1990; includes rocks assigned to Elephant Canyon Formation by Baars, 1962). Similar to the Kaibab, the PI reflection 

is not typically cut by late Paleozoic faults. 

DATA 

An extensive network of intersecting seismic lines was 

examined in this study; locations of selected seismic lines 

are shown in figure 4. Sonic logs and velocity surveys in 

\ ~'\ 

' ( X' 

' J 

I I 

l, 

wells along the seismic lines were used to verify the correlation between reflections and the stratigraphic section. Twoway travel times for the Doughnut to Kaibab interval were 

measured along these lines, and these travel times were converted to the thicknesses shown on figure 5. This conversion 

was calibrated using several deep wells (Hunt 1-16 State, 

American Quasar 31-1 Drunkards Wash, Arcadia Te1onis 1) 

that penetrate upper Paleozoic rocks near the reflection lines 

(fig. 4 ). Because the average seismic velocity for this stratigraphic interval differs in each of these wells and would certainly vary throughout the study area, thickness values are 

approximate. 

Fault geometry and sense of offset along late Paleozoic 

faults were determined based on identification of hangingwall and footwall cutoffs of Mississippian reflections. 

Faults illustrated on figure 5 were identified on more than 

one seismic line in almost every case. Identification of several of the faults on three or more seismic nnes permitted 

EXPLANATION 

Md-Pk interval, >4,500 ft 

c=J Md-Pk interval, 3,500-4,500 ft 

c=J Md-Pk interval, 2,000-3,500 ft 

c=J Md-Pk interval, <2,000 ft 

Thrust fault 

Reverse fault 

Steeply dipping normal fault 

DUCHESNE COUNTY 

---CARBON COtJNTY '-------------

0 

Sunnyside 

0 5 10 15 20 MILES 

0 10 20 30 KILOMETERS 

Figure 5. Thickness variations for Doughnut Formation (Md) to Kaibab Limestone (Pk) interval; faults identified or inferred based on reflection data are shown. Lines of sections X-X' and Y-Y' (fig. 7) are also shown. A, 8, C, D, and E are 

faults shown on plate 1 and discussed in text. 

J4 Application of Structural Geology to Mineral and Energy Resources 
TYPICAL SEISMIC REFLECTION PATTERNS Pk----] 

-

--

M~ 

Md ~-/ Md ~ '-'-...------Md-~Md Mr ~ -~ ~Mr '-.., 't--MrFigure 6. Seismic reflection patterns observed for late Paleozoic faulting in study area. Subhorizontal solid and dashed 

lines are schematic representations of reflections. In this idealized diagram, Mississippian reflectors are offset by faults and 

Permian reflectors are not. Examples of actual data (plate 1) 

demonstrate that in some cases Permian and Mesozoic reflectors have been gently folded above Paleozoic faults by minor 

Laramide reactivation of faults. Pk, Kaibab Limestone; PI, top 

of Upper Pennsylvanian-Lower Permian interbedded limestone, shale, and sandstone sequence; Md, Doughnut Formation; Mr, Redwall Limestone. 

detailed observation of along-strike variations in offset and 

fault geometry. 

Seismic data that illustrate the structural styles are shown 

on plate 1. These Vibroseis data (lines EU-1-85, EU-5, 

EU-6, and EU-9) were recorded by Grant Norpac, Inc., in 

July 1983 (EU-5), June 1984 (EU-6, EU-9), and June 1985 

(EU-1). Lines EU-5, EU-6, and EU-9 were acquired using 

a 16-second, 14-80-Hz source sweep, a shot point interval of 

440ft (134m), 96 geophone groups with 220-ft (67 m) spacing, and a nominal stacking fold of 24. Line EU-1 was 

acquired using a 16-second, 10-56-Hz source sweep, a shot 

point interval of 330 ft (101 m), 120 geophone groups with 

165-ft (50 m) spacing, and a nominal stacking fold of 30. 

Seismic lines EU-5 and EU-9 form a 17-mi (27 km)-

long northeast-southwest transect in the vicinity of Helper 

(plate 1 ). Along these seismic lines, the Doughnut to Kaibab 

interval gradually thickens from southwest to northeast from 

approximately 1,500 to 3,600 ft (457-1,097 m), and three 

northeast-directed reverse and thrust faults (A, B, C) intersect the lines. The geometry of each of these faults is constrained by hanging-wall and footwall cutoffs of 

Mississippian reflectors (Redwall and Doughnut). Based on 

these relationships, faults A, B, and C dip approximately 

30°, 50°, and 10° SE., respectively. Updip from fault A, a 

local increase in dip in the Mesozoic reflectors indicates 

minor Laramide (?) reactivation of this fault. 

Local discordances (truncation or onlap) within the 

Pennsylvanian and Permian banded zone attest to uplift of 

individual fault blocks. Near the Arcadia Telonis 1 well (line 

EU-5, plate 1), inclined reflections in the banded Pennsylvanian and Permian sequence in the hanging wall of fault A are 

locally truncated. The PI reflection appears to be folded 

above, and perhaps offset by, fault A. A dipping reflection in 

the interval between the PI and Pk reflections (fig. 5) is truncated and overlain by subhorizontal reflections. To the north, 

the PI reflection continues undeformed above faults B and C. 

Just south of fault C, there is a southward-onlapping pattern 

at a lower stratigraphic level within the banded Pennsylvanian and Permian reflections. Thus, within these two seismic 

lines, several local unconformities provide evidence for 

localized erosion and (or) nondeposition at slightly different 

stratigraphic positions. 

Line EU-6 (plate 1 ), which extends to the northwest from 

the intersection oflines EU-5 and EU-9 near Helper (fig. 4), 

illustrates gradual northwesterly thinning of the Doughnut to 

Kaibab interval on the upthrown (southeast) side of a minor 

steeply dipping fault (D). The fault geometry is not well constrained by the seismic data here, and the dip of Doughnut 

reflections away from the fault in both directions suggests 

that an arch is locally developed. Southeast of the fault, a 

local angular unconformity is defined by the onlapping of 

Pennsylvanian and Permian strata above the gently southeast 

dipping Doughnut reflector. It is possible that fault-induced 

arching produced a submarine, or barely emergent, topographic high across which younger layers onlapped. Arching 

of reflectors above fault Din the Kaibab through Dakota section demonstrates that post-Paleozoic deformation persisted 

in this zone and probably included Laramide reactivation of 

the fault. 

Line EU-1 (plate 1) is a 26-mi (42 km)-long, north-south 

seismic line in the western part of the study area. Data quality on line EU-1 is inferior to that of the other seismic lines 

discussed here, especially in an 8-mi (13 km) interval 

(denoted by shading along the top of the line drawing in plate 

1) near Schofield Reservoir (fig. 4 ). The poorer data quality 

may be related to raypath complexities produced by young 

structures; line EU-1 follows the Clear Creek Graben, and 

the Fish Creek Graben crosses the seismic line near 

Schofield Reservoir (Osmond, 1965). The data quality deteriorates significantly in the Fish Creek Graben, and the graben itself is not well imaged. The Mesozoic or younger faults 

that are apparent on plate 1 may be related to the Clear Creek 

Graben; the seismic line is highly oblique to the strike of this 

north-trending fault zone. 

Although the Paleozoic section is not well imaged on line 

EU-1, some significant trends are apparent. The southern 

part of the line is on the north flank of the Emery Uplift, as 

evidenced by the absence of the Doughnut Formation along 

the southern one-third of this line and in the nearby Arco 

Hiawatha 1 well. North of fault E, the Redwall to Kaibab 

interval abruptly thickens and discontinuous Doughnut 

reflections may be present (plate 1). In the northern part of 

the line EU-1, there are obvious Doughnut reflections. The 

abrupt variation in thickness of the late Paleozoic section 

across fault E indicates that the fault was active during late 

Paleozoic time. Other faults on this seismic line may have 

undergone Paleozoic movement, but all of these faults also 

underwent younger movement, as shown by the offset of the 

Kaibab and younger reflectors. Thus, it is difficult to obtain 

a clear picture of late Paleozoic movement for these faults. 

Late Paleozoic Structure from Seismic Reflection Data, Uinta Basin, Utah J5 
STRUCTURAL PATTERNS AND 

THICKNESS VARIATIONS 

The overall pattern of faults and thickness variations is 

shown on figure 5. In most of the study area, the thickness of 

the Doughnut to Kaibab interval varies from about 1 ,500 to 

about 4,500 ft (457-1,372 m); the variation is principally 

controlled by local syndepositional faulting. The interval is 

thinnest (1) west and southwest of Helper, corresponding to 

the north flank of the Emery Uplift, and (2) around Sunnyside, at the northwest end of the ancestral Uncompahgre 

Uplift (fig. 5). 

West of Price, the principal Paleozoic faults are northwest-striking, northeast-directed reverse faults that step off 

of the northern and northeastern margins of the Emery Uplift 

and have as much as 5,000 ft (1 ,525 m) of throw (figs. 5, 7). 

Across one of these faults, the preserved thickness of the 

Doughnut to Kaibab interval varies from zero (where the 

entire interval has been locally eroded) to more than 5,000 ft 

{1,525 m). Steep northwest-striking normal faults are also 

present. 

North of Helper, a thickened late Paleozoic section in an 

east-west striking subbasin occupies a footwall position with 

respect to two bounding reverse faults that have opposing 

displacement directions (figs. 5, 7). The more northerly of 

the two reverse faults merges to the southeast with a northerly striking, west-directed thrust system that controlled late 

Paleozoic deformation east and northeast of Price. 

The structural style in this thrust system resembles that of 

a foreland thrust belt, in contrast to the block-faulting style 

that is evident to the west. Figure 8 illustrates imbricate 

thrust structure observed on a seismic line about 6 mi (1 0 

km) east of Price. A few backthrusts, as well as high-angle 

reverse and normal faults, are also present (fig. 5). 

The reflection patterns indicate that many of the late 

Paleozoic faults in this region are essentially growth faults, 

regardless of geometry (normal, thrust, or reverse). 

Although local truncation and onlapping relationships 

within the Morgan and Weber Formations are indicative of 

sw NE 

FEET 1' (k X' 

I 

5,000 I I 1/; d {v \Md 

y 

J Y' 

FEET I I 0 10 MILES 5,0001 1A t} \ ~ 0 10 KILOMETERS 

Md 

Figure 7. Cross sections X-X' and Y-Y' showing variations 

in thickness in Doughnut Formation (Md) to Kaibab Limestone (Pk) interval, and traces of major faults. Datum is top of 

Kaibab Limestone; no vertical exaggeration. Shaded area is 

Md-Pk interval. Lines of sections are shown in figure 5. 

J6 Application of Structural Geology to Mineral and Energy Resources 

FEET 

SEA_ 

LEVEL 

SW NE 

-------Md 

2 MILES 

f-----,----'----r-~ 

-----Mr 

3 KILOMETERS 

Figure 8. Cross section interpreted from migrated seismic 

section about 6 mi (1 0 km) east of Price, illustrating imbricate 

thrust structure that is part of small north-south-striking thrust 

system. Thrust faults shown here have produced about 1.2 mi 

(2.0 km) of shortening. Solid lines are traced from reflections, 

dashed lines are interpreted continuations of stratigraphic horizons, and stippled lines are inferred faults. Vertical exaggeration approximately 0.85. Stratigraphic symbols denote top of 

unit as inferred from seismic data. Pk, Kaibab Limestone; PI, 

top of Upper Pennsylvanian-Lower Permian interbedded 

limestone, shale, and sandstone sequence; Md, Doughnut Formation; Mr, Redwall Limestone. 

local emergence of fault blocks, most of the faulting probably occurred in a relatively shallow submarine setting during 

Pennsylvanian and Permian deposition and did not result in 

widespread subaerial exposures. 

Mesozoic and Tertiary reactivation and overprinting are 

most apparent near the Clear Creek and Fish Creek Grabens 

and near Sunnyside (fig. 1). Young faulting occurred in other 

parts of the study area, but in most cases the late Paleozoic 

structures have not been greatly modified by younger faulting. 

DISCUSSION 

Fault patterns in the study area present a complex picture 

dominated by high-angle (block) faulting in the western part 

of the study area and by thrust faulting in the eastern part. 

Reverse and thrust faulting are predominant, but normal 

faults are locally important. 

A distinctive style of faulting is associated with each of 

the two Ancestral Rockies uplifts that bound the study area. 

Northwest-striking, steeply southwest dipping reverse faults 

bound the Emery Uplift on its northern and northeastern 

sides, whereas a north-striking, west-directed local thrust 

system bounds the ancestral Uncompahgre Uplift at its 

northwest end. The other faults may comprise an interference pattern between these two principal zones, as will be 

discussed later in this section. 

The northwesterly strike of reverse faults that step off the 

north end of the Emery Uplift is almost parallel to the strike 

of the Uncompahgre Fault, known to be a major late Paleozoic reverse fault (fig. 9). The faults are also parallel to the 

N. 60° W.-striking faults documented by Szabo and 

Wengerd (1975) within the Emery Uplift. Although the 
Uncompahgre Fault dips northeast and the faults north of the 

Emery Uplift dip southwest, the orientation of all of these 

faults is compatible with regional northeast-southwest shortening. 

West-directed thrusting at the northwest end of the 

Uncompahgre Uplift requires that the Uncompahgre block 

moved west with respect to the late Paleozoic sedimentary 

basin in the study area. Such movement agrees with Stone's 

(1977) conclusion that the Uncompahgre Fault experienced 

significant strike-slip motion in addition to its well-documented reverse slip (Frahme and Vaughn, 1983). Several 

miles of left-lateral motion on the Uncompahgre Fault could 

have been transferred to the west-directed thrust system, in 

which shortening occurred in front of the west-driven 

Uncompahgre Uplift. 

The curved fault pattern north and northwest of Helper is 

part of a complex zone of accommodation between the 

Uncompahgre and Emery Uplifts. The two uplifts are 

bounded by en echelon reverse fault zones (fig. 9) that 

require either a strike-slip zone or a complex curved fault 

zone between the two uplifts. Northwest-southeast motion 

of the Uncompahgre or Emery blocks would further complicate the structure in the intervening region. Along the 

curved faults north and northwest of Helper, high-angle 

reverse displacement (and local low-angle thrusting) 

occurred on east to northwesterly striking segments. Individual reverse faults pass into low-angle thrusts at their 

southeast ends; to the south, this fault geometry is replaced 

by the north-south thrust system described above. The 

curved faults in this accommodation zone appear to have 

been flattened and rotated toward the orientation of the 

north-south thrust system; such rotation is a further expression of shortening associated with northwesterly movement 

of the ancestral Uncompahgre Uplift. 

The fault pattern in and near the study area is consistent 

with a general northeast-southwest orientation for regional 

compression. Slight variations in the regional stress field 

during Pennsylvanian and Permian time could have produced alternating reverse and strike-slip motion on major 

pre-existing zones of weakness that trend about N. 60° W. 

(For example, regional compression oriented between N. 

45° E. and N. 15° E. could produce reverse displacement on 

the Uncompahgre Fault and the northwesterly trending faults 

north of the Emery Uplift; easterly excursions of regional 

compression direction would produce sinistral strike slip on 

the same faults.) The reflection data provide information 

only on dip-slip motion; there may have been a considerable 

strike-slip component of motion on some of the northwesterly striking subsurface faults in the study area. 

Normal faults are also present, both parallel and orthogonal to the dominantly northwest trends of the reverse faults. 

Northeasterly striking normal faults are compatible with the 

northeast-southwest regional compression direction proposed above. The northwesterly striking normal faults may 

have developed in response to "post-Ancestral Rockies" 

39° 

00' ' 

EMERY UPLIFT 

0 

0 

EXPLANATION 

~ Thrust fault 

~ Reverse fault 

-~ Steeply dipping normal 

fault 

Green 

River 0 

30 MILES 

40 KILOMETERS 

Figure 9. Paleozoic geologic setting of study area including 

major faults identified in this study. 

(and pre-Kaibab) relaxation of the northeast-southwest compression. It is also possible that movement on some of these 

faults included a strike-slip component. 

It is difficult to closely constrain the timing of movement 

on the late Paleozoic faults discussed here. The faults displace reflectors correlated with the Meramecian and Chesterian Doughnut Formation and do not offset reflectors 

correlated with the Leonardian Kaibab Limestone. In examples cited above, local unconformities (surfaces of truncation and onlap) were identified within the banded reflection 

zone in the Weber and Morgan Formations; these are surfaces of erosion and (or) nondeposition above fault blocks 

and are overlain by reflectors that are not affected by the late 

Paleozoic faulting. The PI reflection is stratigraphically 

higher than almost all ofthese local unconformities and postdates most of the faults. If the PI reflector marks the top of 

the Virgilian and Wolfcampian interbedded carbonate and 

clastic sequence, as suggested above, almost all of the late 

Paleozoic faulting would have to be pre-Wolfcampian. Relationships near fault A on line EU-5, illustrated in plate 1, are 

the only examples of slightly younger Permian deformation 

cited here. Thus, the late Paleozoic faulting in the study area 

was certainly post-Chesterian and pre-Leonardian and probably pre-Wolfcampian in most cases. If, as suggested above, 

the faults are genetically tied to the displacement on the 

Uncompahgre Fault, they were most active in Desmoinesian 

time because Desmoinesian sedimentation patterns strongly 

suggest that the Uncompahgre Uplift was rising most rapidly 

at this time (Szabo and Wengerd, 1975; Johnson and others, 

in press). 

late Paleozoic Structure from Seismic Reflection Data, Uinta Basin, Utah J7 
Kluth and Coney (1981) interpreted Ancestral Rockies 

orogenesis as the remote foreland expression of the 

Ouachita-Marathon orogeny, which was produced by collision between South America-Africa and a peninsular projection of the North American craton. Budnick (1986) 

proposed that the Ancestral Rockies were formed along the 

Wichita megashear in response to the collision of eastern 

North America with Africa (Alleghanian orogeny). These 

two models have different kinematic implications for the 

structures in and near the study area (Budnick, 1986). A 

right-lateral component of displacement on northwesterly 

striking faults would be consistent with the Kluth and Coney 

model, whereas a left-lateral component of displacement 

would be consistent with the Budnick model. Because the 

study area is small, the buried fault patterns cannot rule out 

either model; however, the fault patterns are most compatible with the model of Budnick (1986), which requires leftlateral shear along northwesterly trending faults and east- to 

northeast-directed compression. The relevant observations 

in the study area include the reverse component of movement on northwesterly striking faults and the local northsouth thrust belt that provides evidence for westward motion 

of the Uncompahgre block relative to the Paleozoic sedimentary rocks in the Price area. This study produced no evidence, however, that the Uncompahgre Fault and the study 

area are part of a megashear along which 72-93 mi (120-150 

km) of left-lateral displacement occurred, as suggested by 

Budnick (1986). Instead, our analysis implies that strike-slip 

displacement on the Uncompahgre Fault was probably less 

than 6 mi (10 km). 

REFERENCES CITED 

Baars, D.L., 1962, Permian system of the Colorado Plateau: American Association of Petroleum Geologists Bulletin, v. 46, p. 

149-218. 

Budnick, R.T., 1986, Left-lateral intraplate deformation along the 

Ancestral Rocky Mountains; implications for late Paleozoic 

plate motions: Tectonophysics, v. 132, p. 195-214. 

Frahme, C.W., and Vaughn, E.B., 1983, Paleozoic geology and 

seismic stratigraphy of the northern Uncompahgre front, 

Grand County, Utah, in Lowell, J.D., ed., Rocky Mountain 

foreland basins and uplifts: Rocky Mountain Association of 

Geologists Guidebook, p. 201-211. 

Franczyk, K.J., 1991, Phanerozoic stratigraphic and time-stratigraphic cross sections along line C-C', Uinta and Piceance 

J8 Application of Structural Geology to Mineral and Energy Resources 

basin area, southern Uinta Mountains to northern Henry 

Mountains: U.S. Geological Survey Miscellaneous Investigations Series Map 1-2184-C. 

Heyman, O.G., 1983, Distribution and structural geometry offaults 

and folds along the northwestern Uncompahgre uplift, westem Colorado and eastern Utah, in Averett, W.R., ed., Northem Paradox Basin-Uncompahgre uplift: Grand Junction 

Geological Society Guidebook, p. 45-57. 

Hintze, L.F., 1988, Geologic history of Utah: Brigham Young University Geology Studies Special Publication 7, 202 p. 

Hite, R.J., 1975, An unusual northeast-trending fracture zone and 

its relations to basement wrench faulting in northern Paradox 

Basin, Utah and Colorado: Four Comers Geological Society 

Field Conference, 8th, Canyonlands Guidebook, p. 217-223. 

Huffman, A.C., Jr., and Taylor, D.J., 1989, San Juan Basin faulting-More than meets the eye [abs.]: American Association 

of Petroleum Geologists Bulletin, v. 73, p. 1161. 

Johnson, S.Y., Chan, M.A., and Konopka, E.H., in press, Pennsylvanian and Early Permian paleogeography of the UintaPiceance basin region, northwest Colorado~ and northeast 

Utah: U.S. Geological Survey Bulletin 1787-CC. 

Loope, D.B., Sanderson, G.A., and Verville, G.J., 1990, Abandonment of the name "Elephant Canyon Formation" in southeastern Utah; physical and temporal implications: The 

Mountain Geologist, v. 27, p. 119-130. 

Kluth, C.F. and Coney, P.J., 1981, Plate tectonics of the Ancestral 

Rocky Mountains: Geology, v. 9, p. 10-15. 

Osmond, J.C., 1965, Geologic history of site of Uinta Basin, Utah: 

American Association of Petroleum Geologists Bulletin, v. 

49, p. 1957-1973. 

Stevenson, G.M., and Baars, D.L., 1986, The Paradox-A pullapart basin of Pennsylvanian age, in Peterson, J.A., ed., 

Paleotectonics and sedimentation in the Rocky Mountain 

region, U.S.: American Association of Petroleum Geologists 

Memoir 41, p. 513-539. 

Stone, D.S., 1977, Tectonic history of the Uncompahgre uplift, in 

Veal, H.K., ed., Exploration frontiers of the central and 

southern Rockies: Rocky Mountain Association of Geologists Field Conference, 1977, Guidebook, p. 23-30. 

Szabo, E., and Wengerd, S.A., 1975, Stratigraphy and tectogenesis 

of the Paradox Basin: Four Comers Geological Society Field 

Conference, 8th, Canyonlands Guidebook, p. 193-211. 

Waechter, N.B., and Johnson, W.B., 1985, Seismic interpretation in 

the Piceance Basin, northwest Colorado, in Gries, R.R., and 

Dyer, R.C., eds., Seismic exploration of the Rocky Mountain 

region: Rocky Mountain Association of Geologistli and Denver Geophysical Society, p. 247-258. 

White, M.A., and Jacobson, M.I., 1983, Structures associated with 

the southwest margin of the ancestral Uncompahgre uplift, in 

Averett, W.R., ed., Northern Paradox Basin-Uncompahgre 

uplift: Grand Junction Geological Society Guidebook, p. 

33-39. 
Chapter K 

Audio-Magneto-Telluric Investigation at 

Turkey Creek Caldera, Chiricahua Mountains, 

Southeastern Arizona 

By ROBERT M. SENTERFIT and DOUGLAS P. KLEIN 

U.S. GEOLOGICAL SURVEY BULLETIN 2012 

APPLICATION OF STRUCTURAL GEOLOGY TO MINERAL AND 

ENERGY RESOURCES OF THE CENTRAL AND WESTERN UNITED STATES 
CONTENTS 

Abstract Kl 

Introduction Kl 

Data acquisition and processing K2 

Resistivity patterns interpreted from one-dimensional modeling K3 

Profile A-A'-The eastern ring structure K4 

Profile B-B'-The northern moat and caldera interior K4 

Profiles C-C' and D-D' -Cretaceous rocks northeast of the caldera K5 

Profiles E-E' and F-F-West flank of Turkey Creek caldera K6 

Conclusions K8 

References cited K9 

FIGURES 

1. Map showing location of audio-magneto-telluric soundings in Turkey Creek caldera 

area, Chiricahua Mountains, northeastern Arizona K2 

2-7. Contoured resistivity profiles, Turkey Creek caldera area: 

2. A-A' K3 

3. B-B' KS 

4. C-C' K5 

5. D-D' K6 

6. E-E' K7 

7. F-F' K8 

Kll Contents 
Audio-Magneto-Telluric Investigation at Turkey Creek 

Caldera, Chiricahua Mountains, Southeastern Arizona 

By Robert M. Senterfit 1 and Douglas P. Klein 1 

Abstract 

Electromagnetic induction data obtained using distant 

field sources, mostly of natural origins in the audio-magnetotelluric frequency range of 4.5-27,000 Hz, were analyzed to 

depict the geoelectric structure of the Turkey Creek volcanic 

center, an inferred mid-Tertiary rhyolitic caldera about 20 km 

in diameter and located about 10 km south of the Chiricahua 

National Monument in the Chiricahua Mountains of southeastern Arizona. The data for each station consist of scalar electromagnetic measurements at discrete frequencies for two 

orthogonal magnetic and electric field pairs. Thirty-two stations were occupied, spaced about 0.5 km apart. Observations 

made along profiles across the caldera ring structure and on 

either side of the structure indicate that a unit having a resistivity of about 2,500 ohm-m probably represents the monzonite 

within the mapped caldera ring structure. At a depth of about 

2,000 m a unit of similar resistivity in the Cretaceous-Precambrian sedimentary rocks surrounding the caldera is inferred to 

be overlain by a unit of lower resistivity that probably represents Cretaceous sedimentary rocks and weathered near-surface material. 

High resistivity values are noted for the deeper core intrusive rocks in contrast with the lower resistivity values for the 

tuffs and flow rocks. High resistivity units associated with 

inferred ring structures are present on the east side of the structure and absent on the north and west flanks of the volcanic 

center. 

INTRODUCTION 

Data from 32 audio-magneto-telluric (AMT) soundings 

(4.5-27,000 Hz) were collected in the summer of 1989 in the 

area of the Turkey Creek caldera in the Chiricahua Mountains of southeastern Arizona (fig. 1). The objective of this 

work was to aid in the geologic mapping of the structural and 

lithologic relationships in the area by providing information 

1

U.S. Geological Survey, Box 25046, MS960, Denver, Colorado 

80225. 

on the electrical structure of the subsurface. It seemed likely 

that there would be distinguishable resistivity contrasts 

between Tertiary tuff units, pre-Tertiary sedimentary units, 

and Tertiary intrusive rocks of this area. The sounding data 

were collected along six profiles (fig. 1) located to provide a 

reconnaissance picture of the geoelectrical structure to interpretative depths of about 2-5 km. Soundings were placed 

across the west and east edges of an inferred caldera structure (du Bray and Pallister, 1991) at locations mostly dictated by available access. 

The AMT method consists of recording natural electromagnetic field variations and their corresponding induced 

electric fields in the Earth at discrete frequencies. The principles of the AMT method correspond to those of the 

magneto-telluric (MT) method (Cagniard, 1953; Vozoff and 

others, 1963; Vozoff, 1972), except that the signals measured are at higher frequencies and originate mainly from 

atmospheric electrical disturbances (spherics) rather than 

from ionospheric or magnetospheric phenomena. Strangway 

and others (1973) discussed the potential and limitations of 

the AMT method in mineral exploration. Case histories of 

the AMT method in mineral and geothermal environments 

have been described by Strangway and Kozier (1979), 

Hoover and others (1976), and Hoover and Long (1976). 

Previous AMT or MT studies investigating the structural and 

lithologic relations in volcanic areas include those by 

Hoover and others (1976) and Hermance and others (1984) 

(Long Valley, California); Leary and Phinney (1974), Stanley and others (1977), and Stanley (1982) (Yellowstone, 

Wyoming and Idaho); Long (1985) (Questa, New Mexico); 

and Fitterman and others (1988) (Newberry volcano, Oregon). 

Much of the geology of the Turkey Creek caldera is 

exposed in cross section as a result of extensional Basin and 

Range faulting and associated erosion (Marjaniemi, 1969; 

Latta, 1983). The caldera formed about 25 Ma, during the 

Oligocene, in response to eruption of silicic ash-flow tuff of 

the Rhyolite Canyon Tuff (fig. 1). Precaldera volcanic rocks 

of the area include rhyolitic tuffs to intermediate-composition flow rocks. A large central monzonite porphyry unit is 

AMT Investigation, Turkey Creek Caldera, Chiricahua Mountains, Arizona Kl 
109°30' 

32°00',-----------------------,--------r-------------r--.,.:--=-,,-------------------, 

/ 

/ 

10 KILOMETERS 

~ 

~ 

G 

EXPLANATION 

Quaternary slump breccia 

Tertiary monzonite porphyry 

(core and ring) 

Oligocene Rhyolite Canyon TuffWelded tuff 

Tertiary moat lavas and tuff of 

Fife Canyon 

Cretaceous Bisbee Formation 

---Contact-Dashed where inferred 

-...,.-- Fault 

---Structural margin of caldera 

.._25 AMT station location 

A--A' AMT profile 

31°40' '----------------------------------'-------------------' 

Figure 1. Location of audio-magneto-telluric soundings in Turkey Creek caldera area, Chiricahua Mountains, southeastern 

Arizona. Profiles A-A' through F-F' are presented as contoured, composite one-dimensional cross sections in figures 2-7. 

Generalized geology from du Bray and Pal lister (1991 ). 

intrusive into the domed intracaldera tuff (Lipman, 1984). 

The caldera is bordered by Cretaceous sedimentary rocks of 

the Bisbee Formation to the east where the ring structure is 

best exposed. The caldera borders are less certain to the 

north where they disappear under rhyolite flow rocks and 

tuff, to the west where the caldera is bordered by alluvium, 

and to the south where adjacent calderas may be present. 

DATA ACQUISITION AND PROCESSING 

The scalar AMT system used by the U.S. Geological Survey (Hoover and others, 1976, 1978; Hoover and Long, 

1976) records in analog form two sets of orthogonal 

magnetic (H) and electric (E) field amplitudes for each of 

several frequency bands. The system has the capability for 

16 narrow-band frequencies distributed logarithmically in 

the range of 4.5-27,000 Hz. Typically, measurements at all 

bands are attempted, but atmospheric signal conditions can 

preclude acquisition of reliable data for some frequencies. 

The 32 stations in this report were observed in May 1989. 

The signals sought were all natural in origin except for the 

14,000-Hz and the 27,000-Hz bands, which partly utilize 

navigation and communication signals. 

Simultaneous peaks from each pair of orthogonal E-H 

records were scaled and combined to compute apparent 

resistivity (in ohm-meters, proportional to [E/H]) for the 

two observed E and H field orientations. Typically, 8-16 

K2 Application of Structural Geology to Mineral and Energy Resources 
samples from each analog record form the basis for computing estimates of resistivity/north-south (Rns) and 

resistivity/east-west (Rew). The reduced data were edited to 

reject data considered inconsistent with the trend of the 

sounding curve (the bilogarithmic plot of apparent resistivity versus frequency). Selection of data for rejection was 

performed independently on the Rns and Rew curve for 

each station. A smooth approximation to the data is 

required by the electromagnetic response of the Earth 

(Weidelt, 1972). Furthermore, shifts between the two 

curves can have interpretable geologic significance with 

regard to heterogeneity or anisotropy in the Earth (Vozoff, 

1972; Strangway and others, 1973). 

The observation sites were to some extent limited by 

accessibility by road; however, most stations were sited at 

the most optimum locations for determination of structures 

and lithologic contrasts within, across, and outside the ring 

structure. At sites where the ring structure was assumed to be 

near, theE-lines were oriented parallel with and perpendicular to the inferred structure such that the apparent resistivity 

for both E-parallel and E-perpendicular polarizations was 

SOUTH SOUTHWEST NORTH NORTHEAST 

C/) 

a: 

f-

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:;;;:; 

~ 

A A' 

STATION 

16 15 13 

- SEA -++--\--\-----\----+--+-----,t..-.+-1---+-~.:___j_ 

~ LEVEL 

0 

5 

~ -1 > 

-2 

-3 

3 

HORIZONTAL DISTANCE, IN KILOMETERS 

Figure 2. Contoured resistivity-depth section for 

profile A-A' across eastern ring structure of Turkey 

Creek caldera. Resistivity contours are based on onedimensional modeling (Bostick, 1977) at each AMT 

station. The modeling fits an approximation to the 

geometric mean of observed Rxy and Ryx curves at 

each station. Contours are in units of ohm-meters at 

roughly equal log intervals (multiples of 10, 16, 25, 

40, 63, and 1 00). 

directly measured. Apparent resistivity values thus obtained 

can be compared with those from theoretical two-dimensional type-curves corresponding to the appropriate polarization (Strangway and others, 1973). At sites where the 

location of the ring structure was not known or could not be 

inferred, the E-lines were oriented north-south and eastwest, respectively. 

An estimate of the resistivity distribution with depth was 

obtained by transforming the smoothed data to a resistivitydepth function using Bostick's method (Bostick, 1977; 

Bostick and others, 1977, Goldberg and Rotstein, 1982). The 

results of this transformation provide a one-dimensional 

approximation of the depth-resistivity distribution at each 

station. The response of the resulting layered model for each 

station was computed using the forward-modeling algorithm 

described by Schmucker ( 1971) and is presented along with 

the observed data for each profile. The resulting one-dimensional interpretations based on the above computations are 

presented as contoured composite sections for each of the 

profiles. 

Limitations of one-dimensional models are that inaccuracies and poor resolution may result because of scatter and 

bias in the measurements and because of the limited range 

and number of frequencies observed (Oldenburg, 1979). 

More important in regard to the data in our study, however, 

is the fact that almost all the data show a response to lateral 

resistivity variations as indicated by a shift between the Rns 

andRew parts of the sounding curve. This shift is referred to 

as apparent anisotropy. In our one-dimensional analysis, the 

curve transformed using Bostick's method (1977) was a 

hand-smoothed average of the Rns and Rew data. The data 

points were subjectively weighted based on their quality. If 

the data are considered equally, the resulting curve is equivalent to using the geometric mean sounding curve, which for 

two-dimensional structures is an invariant curve even for 

tensor (phase-amplitude measurements) data (Berdichevsky 

and Dimitriev, 1976; Eggers, 1982; Park and Livelybrooks, 

1989). Nevertheless, direct information about the dimensionality of structures is lost in this approach. 

RESISTIVITY PATTERNS INTERPRETED FROM 

ONE-DIMENSIONAL MODELING 

To illustrate the resistivity structure, six profiles were 

constructed. Each profile illustrates the distribution of resistivity with depth as interpreted by contouring the results of 

one-dimensional Bostick transformations for the soundings 

along the profile. Because the Bostick method initially provides a smoothly varying depth-resistivity curve, the depths 

for a logarithmic sequence of resistivity values (multiples of 

10, 16, 25, 40, 63, and 100 ohm-m) were selected directly 

from the transformation for each station. The sounding 

curves and the calculated response of the one-dimensional 

layered model for the Bostick transformation are shown for 

each station on the profile. 

AMT Investigation, Turkey Creek Caldera, Chiricahua Mountains, Arizona KJ 
Profile A-A'-The Eastern Ring Structure 

Profile A-A' (fig. 2) extends from south-southeast (station 17) to north-northeast (station 8) across a part of the 

eastern ring structure of Turkey Creek caldera. The ring 

structure proper, or the combined zone of ring dikes and 

breakaway faults of the caldera collapse, is believed to be 

roughly centered in a zone marked by a quartz monzonite 

intrusive body (fig. 1). The quartz monzonite body is bordered on the southwest by welded tuff of the Rhyolite Canyon Tuff, partly covered by slump breccia, and bordered on 

the northeast by sedimentary rocks of the Cretaceous Bisbee 

Formation. 

Station 13 is at or near the inferred contact of the porphyry of the core structure with the ring. Station 14 is at an 

exposed contact between the annular ring of quartz monzonite and monzonite porphyry (both combined into the Tmp 

unit of fig. 1) and the Bisbee Formation composed of arkose, 

graywacke, thin limestone, basalt, and diabase. 

The resistivity of the near-surface rocks along profile 

A-A' ranges from 63 to about 160 ohm-m or more. These 

resistivity values are typical for fractured and weathered surface rocks or alluvium of this survey, although lower values 

commonly are measured in basin environments. The thickness of the surface layer (not explicit in the contours of fig. 

2 because of the scale and contour interval) varies from 

20-40 m within the ring structure to 400 m in the outlying 

rocks. 

The 160 to 400 or 630 ohm-m resistivity zone is similar 

in configuration to the deeper and higher resistivity zone 

(>630 ohm-m). The contact between the lower resistivity 

tuffs or sedimentary rocks (<630 ohm-m) and the higher 

resistivity (>630 ohm-m) core or ring intrusive rocks is 

believed to be roughly marked by the transition resistivity 

contours of 630-1,000 ohm-m. Note particularly the zone 

between stations 13 and 9, where a sharp lateral change in 

resistivity and the high resistivity of the near-surface material correspond in part to the mapped quartz monzonite and 

other intrusive rocks (unit Tmp) near the ring structure. The 

high resistivity values (greater than 630 or 1,000 ohm-m) 

extend from a depth of about 50 m to about 2-2.5 km, where 

the contours flatten to form a ubiquitous high-resistivity, 

semihorizontal layer that is indicated in the lower frequencies of the data at all stations. In the vicinity of the ring structure at stations 4 and 14, the high resistivity values are 

inferred to represent unaltered parts of the intrusive rock 

unit. At depth, both below the ring structure and elsewhere, 

the high resistivity values may be the result of either the 

intrusive rocks of the monzonite porphyry unit and related 

intrusive rocks or other basement rocks. 

Although station 13 is also in the area of mapped intrusive rocks, resistivity values here are significantly lower than 

those measured between stations 4 and 14. These lower values, less than 250 ohm-m, suggest either local alteration or a 

higher fault and fracture density with included pore fluids. 

The distinct separation of the two apparent resistivities (Rns 

andRew) at stations 4, 13, and 14 is indicative of major lateral contrasts in resistivity. 

The high resistivity values of the ring structure are clearly 

bounded by lower resistivity values of less than 250 ohm-m, 

typically less than 160 ohm-m. In the area to the east, below 

stations 9 and 8, the lower resistivity values are associated 

with sedimentary rocks of the Bisbee Formation. To the 

west, the similar low resistivity values probably are associated with slump breccia at stations 15, 16, and 17. The transition to higher resistivity values is rather sharp at a depth of 

about 2.2 km below stations 9 and 8, whereas below stations 

15, 16, and 17 the transition from 400 to 1,000 ohm-m is 

more gradual and of considerable thickness. These data most 

likely indicate units consisting of altered intercaldera tuff or 

altered and fractured intrusive or basement rocks above the 

higher resistive basement. This zone of intermediate resistivity (400-1,000 ohm-m) is at a depth of about 1.8 km below 

station 15 and at decreasing depths and is thicker toward station 16. 

Profile 8-8'-The Northern Moat and 

Caldera Interior 

Profile B-B'(fig. 3) begins at station 23, above the Tertiary moat lavas of Fife Canyon and trends south-southeast 

across the contact with the welded tuffs of the Rhyolite Canyon Tuff at station 21. The profile crosses the contact with 

the ring monzonite porphyry at stations 18 and 17, where 

profile B-B' intersects profile A-A'. For the most part, profile B-B' is parallel with and inside of the inferred ring structure; station 22 is on the inferred ring structure and station 23 

is outside of the inferred caldera boundary. The inferred 

caldera boundary at station 22 is uncertain because this area 

is covered by rhyolite flow rocks and tuffs. 

The contoured section for profile B-B' (fig. 3) shows a 

low-resistivity surface layer of 40-160 ohm-m between stations 23 and 21. The low-resistivity layer is about 700 m 

thick at station 23; it thins southward to station 22, then thins 

rapidly to 30m at stations 19 and 21, and pinches out by station 20. This layer is interpreted to represent the configuration of the rhyolite tuffs along profile. These tuffs are not 

significantly altered at the surface but must be sufficiently 

porous and have sustained sufficient pore fluid ionic activity 

to cause this relatively low resistivity. We have no petrophysical data to verify this, but the near-surface data are 

quite conclusive of the low resistivity values. The base ofthe 

zone is, however, rather uncertain because there are no sharp 

gradients in the contour intervals to at least the 1,000-ohmm contour. Although we assume that resistivity values 

greater than 1,000 ohm-m are representative of basement 

rocks (probably core intrusive rocks of unit Tmp), it would 

be highly arbitrary to assign interfaces without further oneand two-dimensional modeling trials. 

Bearing in mind the above, we propose that the generalized zones of 40-160 ohm-m, 250-400 ohm-m, and 1,000 

K4 Application of Structural Geology to Mineral and Energy Resources 
NORTHWEST 

B 

-2 

-3 

22 

STATION 

SOUTHEAST 

B' 

19 21 20 18 

Resistivity 

10,967 ohm-m~ 

at 14,500 m 

HORIZONTAL DISTANCE, IN KILOMETERS 

Figure 3. Contoured resistivity-depth section for profile 8-B' 

in northern moat area and interior of Turkey Creek caldera. Resistivity contours are based on one-dimensional modeling (Bostick, 1977) at each AMT station. The modeling fits an 

approximation to the geometric mean of observed Rxy and Ryx 

curves at each station. Contours are in units of ohm-meters at 

roughly equal log intervals (multiples of 10, 16, 25, 40, 63, and 

100). 

ohm-m and greater represent, respectively, shallow and 

"wet" tuffs, intermediate-depth tuffs of probably lower 

porosity ("drier"), and intrusive basement rocks. The upper 

tuffaceous rocks thin to the southeast, and the high-resistivity basement rocks significantly shallow to the southeast. 

The correspondence between variations in the thickness of 

the 250-400-ohm-m zone, as based on steps in the 400-ohmm contour, and the location of mapped faults near stations 19 

and 22 (du Bray and Pallister, 1991) suggests that zone 

thickness is in part fault controlled. 

Station 22 is at the inferred ring structure boundary. We 

consider the lack of an anomaly below stations 22 and 23 

comparable to that on profile A-A' at stations 4 and 14 to be 

significant. The data show that near station 22 unaltered 

intrusive rocks are no shallower than about 1.5 km. It is possible that intrusive rocks or faults associated with a ring 

structure may have been altered to the extent that there is little resistivity contrast with the tuffs; however, based on comparison with profile A-A', it is more probable either that the 

caldera boundary is north of station 23 or south of station 22 

or that faults and intrusive rocks associated with a ring structure are buried beneath tuffs more than 1-1.5 km deep. 

Profiles C-C' and D-O' -Cretaceous Rocks 

Northeast of the Caldera 

Profile C-C' (fig. 4) is entirely along the Cretaceous Bisbee Formation in an area of high relief. The short ( 1.6 km) 

north-south profile begins at station 5 and merges to the 

south with profile A-A' at station 8. Station 3 is at the top of 

Onion Gap Pass, a topographic saddle along profile. Most 

observations were taken along a relatively narrow ridge, and 

the largest topographic variations are lateral to the profile. 

The data along profile show the lack of a significant layer 

of low resistivity (not explicit on figure 4 because of scale 

and contour interval). The relatively high resistivity values 

(greater than 1,000 ohm-m) near the surface indicate the 

absence of surface alluvium and weathered and conductive 

pore-water rocks near the surface and are considered typical 

(/) 

a:: 

UJ 

f2.5 

:;g 1.5 

g 

;;.:2 

~ 

~-

:z 

~ 0.5 

0 SEA 

;:3 § LEVEL 

> -0.5 

-1 

NORTH SOUTH 

c 

5 

.6. 

STATION 

6 7 3 

A 

1 

HORIZONTAL DISTANCE, 

IN KILOMETERS 

C' 

8 

Figure 4. Contoured resistivitydepth section for profile C-C' in 

area of Cretaceous rocks northeast 

of Turkey Creek caldera and 

merging with profile A-A' across 

the caldera boundary. Resistivity 

contours are based on one-dimensional modeling (Bostick, 1977) at 

each AMT station. The modeling 

fits an approximation to the geometric mean of observed Rxy and 

Ryx curves at each station. Contours are in units of ohm-meters at 

roughly equal log intervals (multiples of 10, 16, 25, 63, 40, and 

100). 

AMT Investigation, Turkey Creek Caldera, Chiricahua Mountains, Arizona K5 
C/) 

= 

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=: ~ SEA 

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WEST 

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STATION 

11 

EAST 

D' 

10 

... 

HORIZONTAL DISTANCE, IN KILOMETERS 

Figure 5. Contoured resistivity-depth 

section for profile 0-0' in area of Cretaceous rocks east of Turkey Creek caldera, 

along East Turkey Creek drainage. Resistivity contours are based on one-dimensional modeling (Bostick, 1977) at each 

AMT station. The modeling fits an approximation to the geometric mean of observed Rxy and Ryx curves at each station. 

Contours are in units of ohm-meters at 

roughly equal log intervals (multiples of 

1 0, 16, 25, 40, 63, and 1 00). 

of topographically high, well-drained terrane. The lowest 

resistivity values are at depths where fluids probably have 

become stabilized. 

Considerable variation in the one-dimensional modeled 

resistivity values for the upper 1-1.5 km may suggest significant heterogeneity in the Bisbee Formation; however, some 

of the variation may be an artifact of severe topographic 

effects that are unaccounted for in one-dimensional modeling. A small isolated low (less than 100 ohm-m) is at a depth 

of 0.1-0.5 km at station 7 and, with smaller dimension, at 

station 5. Resistivity values greater than 250 ohm-m are 

between these lows. The wedge of 1,600-ohm-m material 

between stations 3 and 8 is poorly constrained; it results from 

the high resistivity in the one-dimensional model for station 

3. The resistivity structure below station 8 is best seen in profile A-A '(fig. 2). 

Combined contour models for the east end of profile 

A-A'and the south end of profile C-C' suggest that the configuration of basement beneath the Bisbee Formation is 

probably represented roughly by the 630-ohm-m contour at 

a depth of 1-5 km. Because of the uncertainty associated 

with the combined effects of resistivity variations in the Bisbee Formation, topography, and limitations of one-dimensional modeling, it is not considered proper to assess the 

detailed configuration of the slope or structure on the basement rock. 

Profile D-D' (fig. 5) consists of three stations, 10, 11, 

and 12, along the East Turkey Creek drainage in the Cretaceous Bisbee Formation. These data were acquired in an area 

of much more favorable topography for AMT soundings 

than that for profile C-C', and the apparent heterogeneity in 

the resistivity of the Bisbee as contoured for profile C-C' is 

not reflected in profile D-D'. The data taken along profile 

D-D' show significant shifts (anisotropy) between the Rns 

and Rew curves; however, the one-dimensional modeling 

produces relatively smoothly varying resistivity variations 

that are believed to reflect the major aspects of the subsurface basement structure. 

The large shifts in the lower frequencies at stations 12 

and 11 may well be related to the sharply thickening Bisbee 

Formation east of station 11, as inferred from the 250-630-

ohm-m contours. Part of the higher resistivity west of station 

11 is inferred to reflect the edge of the quartz monzonite 

intrusive rocks of the Turkey Creek caldera ring structure. 

East of station 11, the 1,000-ohm-m contour, or perhaps the 

630-1 ,000-ohm-m zone, is inferred to reflect the general 

configuration of igneous-metamorphic basement rocks. 

Profiles f-f' and f-f'-West Flank of 

Turkey Creek Caldera 

The location of the west flank of the Turkey Creek 

caldera is problematic and highly inferential from the standpoint of geologic mapping because the caldera boundary 

area is not as well exposed as it is on the east flank. Two 

roughly parallel AMT profiles were established across the 

inferred boundary to explore the geoelectric structure in an 

attempt to provide some subsurface information on this 

questionable boundary. 

Profile E-E'(fig. 6) extends from west to east up Rock 

Creek Canyon, from station 32 on the west to station 31 on 

the east. Stations 32 and 29, which have resistivity values of 

100 ohm-m or less from the surface to a depth of about 1,000 

m, indicate the presence of a significant thickness of alluvium or perhaps alluvium overlaying relatively low resistivity moat lavas of Fife Canyon (compare with profile B-B'). 

Higher resistivity values of 400 ohm-mare encountered at 

about 2,000 m depth, but resistivity values typical of intrusive or basement rocks (as on profile A-A') were not measured along the eastern part of this profile. 

At station 29, at the entrance of Rock Creek Canyon, the 

welded tuffs and rhyolite lavas of the Rhyolite Canyon Tuff 

are shown in the AMT data as the shift between the Rns and 

K6 Application of Structural Geology to Mineral and Energy Resources 
1.5 

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a: 0.5 1-

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~ LEVEL 

~-

z -0.5 

t;; 

0 

~ 

-1 

~ -1.5 ~ 

-2 

-2.5 

-3 

WEST 

E 

29 

-- / '-..... ----------/ ------------

STATION 

30 

EAST 

E' 

HORIZONTAL DISTANCE, IN KILOMETERS 

Figure 6. Contoured resistivity-depth section for profile E-E' on west flank of 

Turkey Creek caldera, along Rock Creek Canyon. Resistivity contours are 

based on one-dimensional modeling (Bostick, 1977) at each AMT station. The 

modeling fits an approximation to the geometric mean of observed Rxy and 

Ryx curves at each station. Contours are in units of ohm-meters at roughly 

equal log intervals (multiples of 10, 16, 25, 40, 63, and 1 00). 

Rew curves (fig. 6). Farther east at station 30, the low-resistivity layer thins to less than 100m and disappears between 

stations 30 and 31. At the east end of the profile, station 31 

on the monzonite porphyry of the core intrusion shows resistivity values greater than 1,600 ohm-m from the surface 

downward to the maximum depth sensed. 

Profile F-F' (fig. 7) extends from west (station 1) to east 

(station 2) along Turkey Creek Canyon road. In comparison 

with profile E-E', the low-resistivity layer, less than 400 

ohm-m, is limited to the upper 200 mat station 1, thins to the 

east toward station 27, and disappears east of station 27. 

Because profiles E-E' and F-F' are essentially parallel with 

each other, the higher resistivity rocks(> 630 ohm-m) may 

be assumed to be the same unit beneath both profiles. These 

high-resistivity rocks, assumed to represent the intrusive 

unit, apparently are at depths of less than 1-1.5 km along the 

western section of profile F-F'. This suggests a north-trending deepening of the intrusive unit in the area of the west 

ends of profiles F-F' and E-E'. 

High resistivity values of 1,000 ohm-m or greater are 

within 30 m of the surface in the area of stations 26, 25, and 

14. In this area, anomalous lower resistivity values (less than 

1,000 ohm-mat 0.5 km) are beneath stations 26 and 24. To 

a lesser magnitude, the buried low-resistivity anomaly is also 

sensed at station 25. Outcrops of intracaldera welded tuff of 

the Rhyolite Canyon Tuff between stations 26 and 25 may be 

responsible for this anomaly. The Rhyolite Canyon outcrop 

widens south of station 26 (fig. 1) and is also present within 

about 1 km of stations 25, 24, and 2. It is likely that the contoured anomaly on profile F-F' (fig. 7) results from lateral 

sensing of this unit, which we assume has lower resistivity 

than the intrusive unit. Thus, the contours are not believed to 

represent a low-resistivity anomaly at depth. Similarly, we 

do not place great significance on the low resistivity values 

(160-630 ohm-m) in the surface layer at station 2. The data 

may be sensing alteration or weathering in the intrusive unit 

at station 2, but one-dimensional modeling here is inadequate for confident inferences. 

The main conclusion we draw from profiles E-E' and 

F-F' is that, within a depth of about 2 km, there is no intrusive structure on the west ends of these lines that corresponds 

to that of the ring boundary to the east (profile A-A', fig. 2). 

If such a ring structure exists in this area, it is buried beneath 

a combination of alluvium and relatively low resistivity moat 

lavas at a depth of at least 2 km, probably 3 km. The data are 

otherwise in accord with geologic mapping. 

AMT Investigation, Turkey Creek Caldera, Chiricahua Mountains, Arizona K7 
22 SEA 

WEST 

F 

STATION 

27 26 25 24 

EAST 

F' 

~ LEVEL ~~--~~------+------~L----f:__HffiH-+-hh-+-+--+++-l---W-\--\--\-t-----l-

;;;:; 

-1 

---- -2 

-3 

-4 

10 

HORIZONTAL DISTANCE, IN KILOMETERS 

Figure 7. Contoured resistivity-depth section for profile F-F' on west flank of Turkey Creek caldera, along Turkey 

Creek Canyon road. Resistivity contours are based on one-dimensional modeling (Bostick, 1977) at each AMT station. The modeling fits an approximation to the geometric mean of observed Rxy and Ryx curves at each station. 

Contours are in units of ohm-meters at roughly equal log intervals (multiples of 10, 16, 25, 40, 63, and 1 00). 

CONCLUSIONS 

The AMT experiment reported here provides (1) a test of 

the electrical signature of an exposed and well-mapped area 

of volcanic intrusive rocks, inferred from geological evidence (Marjanieme, 1969; Latta, 1983; Lipman, 1984; du 

Bray and Pallister, 1991) to be a ring structure of the Turkey 

Creek caldera, and (2) several applied exploratory references 

that may be of value in mapping the volcanic structure at 

depth. Because the interpretations are based on one-dimensional modeling, there may be some distortion in the resistivity sections of the profiles (figs. 2-7), but some of the major 

features shown in the profiles provide a first-order indication 

of subsurface conditions. 

Profiles A-A', B-B', C-C', and D-D' (figs. 2-5) are 

mostly within the area of mapped border structures of the 

rhyolitic volcanic center (fig. 1) (du Bray and Pallister, 

1991). The intrusive rocks here are generally of high resistivity (1 ,000 ohm-m or more), in contrast to lower resistivity 

(typically 160-400 ohm-m) exterior (east)-bordering Cretaceous sedimentary rocks (Bisbee Formation), interior (west) 

bordering slump units, and moat lavas and tuffs (Rhyolite 

Canyon Tuff). The moat tuffs and lavas show particularly 

low resistivity values (63-100 ohm-m) in the upper 1 km 

(profile B-B', fig. 3). These resistivity contrasts make it 

quite straightforward to identify intrusive rocks associated 

with the edge of the volcanic center and locally to identify 

the core intrusive or high-resistivity basement rocks at depth. 

The AMT data provide information to depths of about 

1.5-3.5 km, depending on the resistivity of the nearer surface 

material (resolution at depth decreases when shallower 

material is less resistive). 

At least one area mapped as intrusive rocks (near station 

13, profile A-A', fig. 2) shows anomalously low resistivity 

values (about 250 ohm-m) that may suggest local alteration 

or intense fracture density. Such an area may be a candidate 

for exploration for possible mineralization if geologic evidence for alteration is found. Profile B-B' (stations 23 and 

22, fig. 3), below the inferred northern border of the volcanic 

center, indicates no high-resistivity structures above a depth 

of about 2.5 km. In this area, deeper moat lavas are of higher 

resistivity (greater than 160 ohm-m to as much as 630 or 

1,000 ohm-m) and may thicken toward the north at least 

partly in response to structural (fault) control. The latter conclusion is based on correspondence of the thickening with 

mapped faults (du Bray and Pallister, 1991). 

K8 Application of Structural Geology to Mineral and Energy Resources 
Profiles E-E' and F-F' (figs. 6, 7) were established to 

map the western border of the Turkey Creek volcanic center. 

In this area there are no surface manifestations of border 

intrusive rocks similar to those mapped on the east. The 

AMT data identify the interior (core) intrusion as a high 

resistivity unit (> 1,000 ohm-m) and a possible unit of alluvium and (or) moat lavas and tuffs having resistivity values 

of 60-400 ohm-m. These results are in agreement with geologic mapping except that no evidence of high-resistivity 

intrusive rocks has been found in the area of the inferred border of the volcanic center. The depth of penetration here (at 

the western ends of profiles E-E' and F-F') is about 4 km. 

REFERENCES CITED 

Berdichevsky, M.N., and Dimitriev, V.I., 1976, Basic principles of 

interpretation of magnetotelluric sounding curves, in Adam, 

A., ed., Geoelectric and geothermal studies (east-central 

Europe, Soviet Asia): Budapest, Akademiai Kiado, KAPG 

Geophysical Monograph, p. 1650221. 

Bostick, F.X., Jr., 1977, A simple almost exact method of MT analysis: Workshop on Evaluation of Electrical Methods in the 

Geothermal Environment, Snowbird, Utah, November 1976, 

University of Utah Report, p. 175-177. 

Bostick, F.X., Jr., Smith, H.W., and Boehl, J.E., 1977, A simplified 

one-dimensional magnetotelluric inversion method, in Magnetotelluric and DC dipole-dipole soundings on northern 

Wisconsin: Austin, University of Texas, Final technical 

report prepared under contract N00014-76-C, 0484, Office 

of Naval Research, and Grant GA-38827, National Science 

Foundation, appendix I. 

Cagniard, L., 1953, Basic theory of the magnetotelluric method: 

Geophysics, v. 18, no. 3, p. 605-635. 

du Bray, E.W., and Pallister, J.S., 1991, An ash-flow caldera in 

cross section-Ongoing field and geochemical studies of the 

Turkey Creek caldera, Chiricahua Mountains, SE Arizona: 

Journal of Geophysical Research, v. 96, p. 13435-13457. 

Eggers, D.E., 1982, An eigenstate formulation of the magnetotelluric impedance tensor: Geophysics, v 47, p. 1204-1214. 

Fitterman, D.V., Stanley, W.D., and Bisdorf, R.J., 1988, Electric 

structure of Newberry volcano, Oregon: Journal of Geophysical Research, v. 93, no. B9, p. 10119-10134. 

Goldberg, S., and Rotstein, Y., 1982, A simple form of presentation 

of magnetotelluric data using the Bostick transform: Geophysical Prospecting, v. 30, p. 211-216. 

Hermance, J.F., Slocum, W.M., and Neumann, G.A., 1984, The 

Long Valley/Mono Basin volcanic complex-A preliminary 

magnetotelluric and magnetic variation interpretation: Journal of Geophysical Research, v. 89, p. 8325-8337. 

Hoover, D.B., Frischknecht, F.C., and Tippens, C.L., 1976, Audiomagnetotelluric sounding as a reconnaissance exploration 

technique in Long Valley, California: Journal of Geophysical Research, v. 81, no. 5, p. 801-809. 

Hoover, D.B., and Long, C.L., 1976, Audiomagnetotelluric methods in reconnaissance geothermal exploration: Second United Nations Symposium Development and Utilization 

Geothermal Resources, Proceedings, v. 2, p. 1059-1064. 

Hoover, D.B., Long, C.L., and Senterfit, R.M., 1978, Some results 

from audiomagnetotelluric investigations in geothermal 

areas: Geophysics, v. 43, no. 7, p. 1501-1514. 

Latta, J., 1983, Geochemistry and petrology of the ash flows of 

Chiricahua National Monument, Arizona, and their relation 

to the Turkey Creek Caldera: Tucson, University of Arizona, 

M.S. thesis, 194 p. 

Leary, P., and Phinney, R.A., 1974, A magnetotelluric traverse 

across the Yellowstone region: Geophysical Research Letters, v. 1, no. 6, p. 265-268. 

Lipman, P.W., 1984, The roots of ash-flow calderas in western 

North America-Windows into the tops of granitic 

batholiths: Journal ofGeophysica1 Research, v. 89, no. BlO, 

p. 8801-8841. 

Long, C.L., 1985, Regional audio-magnetotelluric study of the 

Questa caldera, New Mexico: Journal of Geophysical 

Research., v. 90, p. 11270-11274. 

Matjaniemi, D.K .. 1969, Geologic history of an ash-flow sequence 

and its source area in the Basin and Range province of southeastern Arizona: Tucson, University of Arizona, Ph.D. thesis, 176 p. 

Oldenburg, D.W., 1979, One-dimensional inversion of natural 

source magnetotelluric observations: Geophysics, v. 44, no. 

7, p. 1218-1244. 

Park, S.K., and Livelybrooks, D.W., 1989, Quantitative interpretation of rotationally invariant parameters in magnetotellurics: 

Geophysics, v. 54, p. 1483-1490. 

Schmucker, Ulrich, 1971, Interpretation of induction anomalies 

above non-uniform surface layers: Geophysics, v. 36, no. 1. 

Stanley, W.D .. 1982, Magnetotelluric soundings on the Idaho 

National Engineering Facility, Idaho: Journal of Geophysical 

Research, v. 87, no. B4, p. 2683-2691. 

Stanley, W.D., Boehl, J.E., Bostick, F.X., Jr., and Smith, H.W., 

1977, Geothermal significance of magnetotelluric sounding 

in the eastern Snake River Plain-Yellowstone region: Journal of Geophysical Research, v. 82, p. 2501-2514. 

Strangway, D.W., and Kozier, A., 1979, Audio-frequency magnetotelluric sounding-A case history at the Cavendish geophysical test range: Geophysics, v. 44, no. 8, p. 1429-1446. 

Strangway, D.W., Swift, C.M., Jr., and Holmer, R.C., 1973, The 

application of audio-frequency magnetotellurics (AMT) to 

mineral exploration: Geophysics, v. 38, no. 6, p. 1159-1175. 

Vozoff, K., 1972, The magnetotelluric method in the exploration of 

sedimentary basins: Geophysics, v. 37, no. 1, p. 98-141. 

Vozoff, K., Hasegawa, H., and Ellis, R.M., 1963, Results and limitations of magnetotelluric surveys in simple geologic situations: Geophysics, v. 28, no. 5, part 1, p. 778-792. 

Weidelt, P., 1972, The inverse problem of geomagnetic induction: 

Journal of Geophysics, v. 38, p. 257-289. 

AMT Investigation, Turkey Creek Caldera, Chiricahua Mountains, Arizona K9 

Chapter L 

Ash-Flow Calderas as Structural Controls of 

Ore Deposits-Recent Work and 

Future Problems 

By PETER W. LIPMAN 

U.S. GEOLOGICAL SURVEY BULLETIN 2012 

APPLICATION OF STRUCTURAL GEOLOGY TO MINERAL AND 

ENERGY RESOURCES OF THE CENTRAL AND WESTERN UNITED STATES 
CONTENTS 

Abstract Ll 

Introduction Ll 

General caldera model: Relation to mineralization Ll 

Questa magmatic system, New Mexico L3 

Central San Juan caldera cluster, Colorado LS 

Mineralized Mesozoic calderas of southeastern Arizona L 7 

Discussion LlO 

References cited Lll 

FIGURES 

1, 2. Diagrammatic sections showing: 

1. Generalized caldera model L2 

2. Crustal model illustrating evolution of Latir volcanic field and co genetic 

intrusions L4 

3. Map showing generalized geology of central San Juan caldera cluster L6 

4. Diagrammatic north-south section longitudinally along Creede Graben L8 

5. Maps showing generalized geology of four mineralized areas in southeastern 

Arizona interpreted as containing fragments of Late Cretaceous calderas L9 

Lll Contents 
Ash-Flow Calderas as Structural Controls of Ore 

Deposits-Recent Work and Future Problems 

By Peter W. Lipman 1 

Abstract 

Field, petrologic, geochronologic, and geophysical studies of Oligocene ash-flow calderas and associated igneous 

rocks in the southern Rocky Mountains and Mesozoic volcanic 

systems in southeastern Arizona illustrate the applications of 

volcanic structural geology to problems of localization of ore 

deposits and associated hydrothermal systems. Examples 

include (1) molybdenum deposits along the most deeply subsided and recurrently intruded southern margin of the 26-Ma 

Questa caldera; (2) 25-Ma mineralization along the recurrently 

faulted Creede Graben in the central San Juan Mountains; (3) 

little studied hydrothermal alteration south of Creede that 

defines a coherent pattern related to ring intrusions along the 

southern margin of the newly recognized 27.1-Ma South River 

caldera; and (4) localization of porphyry copper mineralization 

along intrusions related to several Late Cretaceous-early Tertiary calderas in southern Arizona. Caldera structures are common loci for mineralization because they are the near-surface 

structural expression of crustal magmatic processes involving 

major petrologic modification of crustal compositions. Misinterpretation of caldera features in complex mineralized terranes commonly has led to inappropriate exploration 

strategies. 

INTRODUCTION 

Caldera structures, ranging in scale from ring faults and 

graben systems to smaller scale faults and joints, can develop 

in response to the rise, differentiation, and solidification of 

upper crustal magmatic systems, commonly of batholithic 

scale (Lipman, 1984, 1988). Hydrothermal systems and 

associated ore deposits form where volcanic structures intersect upper parts of the magma bodies. Ash-flow calderas 

provide important structural control for subsequent mineralizing events (Steven and others, 1974; McKee, 1979; 

Rytuba, 1981 ), but the depths, transport rates, and source of 

1

U.S. Geological Survey, MS910, 345 Middlefield Road, Menlo 

Park, California 94025. 

metals in caldera-related deposits remain poorly understood. 

The stratigraphically complex sequences that record volcanic evolution are typically imperfectly preserved due to subsequent tectonism and erosion, as well as commonly being 

enveloped in a haze of hydrothermal alteration. Both magma 

bodies and volcanic structures evolve rapidly relative to geochronologic resolution, and timing of volcanic structural 

evolution, magmatism, and mineralization is difficult to constrain with adequate precision. A major challenge, both for 

mineral explorationists and volcanologists studying the 

altered and incompletely preserved volcanic rocks associated with ore deposits, is to apply and utilize volcanologic 

concepts and processes identified through study of younger 

and simpler volcanic systems. Recent studies on caldera systems of varying ages and structural levels by myself and others in the U.S. Geological Survey, academia, and industry 

are reviewed to provide insights on interactions between 

caldera-related processes and mineralization. 

GENERAL CALDERA MODEL: 

RELATION TO MINERALIZATION 

Recent general models for structural and volcanic evolution of ash-flow calderas (fig. 1), briefly summarized here, 

relate major structural features to sequential stages of premonitory volcanism, eruption and associated caldera collapse, resurgent magmatism and accompanying uplift of the 

caldera floor, and late hydrothermal activity and mineralization (Smith and Bailey, 1968; Lipman, 1984). 

Most ash flows have erupted from sites of closely preceding volcanism that records initial shallow accumulation of 

caldera-related magma; some radial and circumferential 

structures may be initiated at this stage. Most large calderas 

collapse during ash-flow eruptions as indicated by thick 

ponding of late eruptive products within the concurrently 

subsiding depression. Structural boundaries of calderas are 

single ring faults or composite ring-fault zones that mostly 

dip vertically to steeply inward. Some relatively small 

Ash-Flow Calderas as Structural Controls of Ore Deposition L 1 
-2 

-4 

-6 A 

en 

a: 

1- -2 ~ g 

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~ 

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~ Cl 

-6 

B 

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-6 c 

Figure 1. Generalized caldera model. Modified from Lipman (1984). A, Precollapse volcanism. Clustered intermediate-composition stratovolcanoes grow over isolated small high-level plutons that mark beginning of accumulation of batholithic-size 

silicic magma body that will feed ash-flow eruptions. Uplift related to emplacement of plutons may lead to development of 

arcuate ring fractures. Site of subsequent caldera collapse indicated by dotted lines; heavy arrows indicate upward movement 

of magma. B, Caldera geometry just after ash-flow eruptions and concurrent caldera collapse. Central area of clustered earlier 

volcanoes caves into collapsed area. lntracaldera tuff ponds during subsidence and is an order of magnitude thicker than cogenetic outflow ash-flow sheet. Initial collapse along ring faults is followed by slumping of oversteepened caldera walls and 

accumulation of voluminous collapse breccias that interfinger with ash-flow tuff in the caldera fill. Caldera floor subsides 

asymmetrically and is tilted to left side of diagram. Main magma body underlies entire caldera area and is compositionally 

zoned (or was prior to eruptions), becoming more mafic downward. C, Resurgence and postcaldera deposition. Resurgence 

here is asymmetrical with greatest uplift in area of greatest prior collapse. Extensional graben faults form over crest of dome. 

Some resurgent uplift is accommodated by movement along ring faults in sense opposite that during caldera subsidence. Magma body has risen into volcanic pile and intruded cogenetic intracaldera welded tuff. Original caldera floor has been almost 

entirely obliterated by rise of magma chamber to near level of prevolcanic land surface. Caldera moat is partly filled by lava 

domes and volcaniclastic sediments. Hydrothermal activity and mineralization are dominant late in cycle. 

volume pyroclastic eruptions are associated with incomplete 

hinged caldera subsidences or structural sags; larger systems 

are completely bounded by ring faults. Scalloped topographic walls beyond the structural boundaries of most 

calderas are due to secondary gravitational slumping during 

subsidence, and associated caldera-collapse slide breccias 

L2 Application of Structural Geology to Mineral and Energy Resources 

interleaved with thick tuff are a signature of intracaldera 

accumulations. Most exposed caldera floors are a structurally coherent plate or cylinder bounded by a ring fault or 

dike, indicating pistonlike caldera collapse. Deviations from 

circular shape commonly reflect the influence of regional 

structures. 
Postcollapse volcanism may occur from varied vent 

geometries within ash-flow calderas; ring-vent eruptions are 

common in resurgent calderas and reflect renewed magmatic 

supply. Large intrusions related to resurgence are exposed 

centrally within some calderas; ring dikes and other intrusions along bounding ring fractures are common in regional 

extensional environments. Resurgence within calderas may 

yield a symmetrical dome, accompanied by keystone-graben 

faulting, or more geometrically complex uplift structures; 

resurgence is most common in large calderas (greater than 

10 km diameter) in cratonic crust and is associated with large 

silicic intrusions. In addition to resurgence within single 

calderas, broader magmatic uplift within many silicic volcanic fields reflects isostatic adjustments following the 

emplacement of associated subvolcanic batholiths. 

Hydrothermal activity and mineralization accompany all 

stages of ash-flow magmatism and become dominant late 

during caldera evolution. Some mineralization is closely tied 

to late stages of the caldera magmatic cycle, but much rich 

mineralization is millions of years younger than caldera collapse, where the caldera served primarily as a structural control for genetically unrelated intrusions and their associated 

hydrothermal systems. 

Examples of the diverse relations between caldera evolution and associated mineralization for several representative 

regions in the western United States are summarized below. 

QUESTA MAGMATIC SYSTEM, 

NEW MEXICO 

The 26-18-Ma Questa magmatic system (fig. 2), 

expressed at the surface as the Latir volcanic field, contains 

a relatively simple example of an intracontinental caldera 

related to major mineralization. The volcanic field, caldera, 

and associated intrusions are remarkably exposed in cross 

section due to structural and topographic relief along the 

eastern margin of the Rio Grande Rift in northern New Mexico (Lipman, 1988; Lipman and Reed, 1989). Diverse geologic, petrologic, geochronologic, geophysical, and mineral 

resource studies provide a framework for evaluating interactions between magmatic processes, caldera evolution, and 

mineralization in a extensional cratonic environment. 

Exposed levels increase in depth from the mid-Tertiary 

depositional surface in northern parts of the igneous complex 

to plutonic rocks originally at depths of 3-5 km in the south. 

Erosional remnants of an ash-flow sheet of weakly peralkaline rhyolite (Amalia Tuff) and andesitic to dacitic precursor 

lavas, disrupted by rift-related faults, are preserved as far as 

45 km beyond their sources at the Questa caldera. Broadly 

comagmatic batholithic granitic rocks, exposed over an area 

of 20 by 35 km, range from mesozonal granodiorite to epizonal porphyritic granite and aplite (Lipman and Reed, 

1989). Shallower silicic phases are exposed primarily within 

the caldera. Compositionally and texturally distinct granites 

form resurgent intrusions within the caldera and discontinuous mineralized ring intrusions along its margins; a 

batholithic mass of granodiorite extends for 20 km south of 

the caldera and locally grades upward to granite below its 

flat-lying roof. A negative Bouguer gravity anomaly (15-20 

mGal; Cordell and others, 1986), which encloses exposed 

granitic rocks and coincides with boundaries of the Questa 

caldera, defines boundaries of the shallow batholith that was 

emplaced into lower parts of the volcanic sequence and in 

underlying Precambrian rocks. Paleomagnetic pole positions 

indicate that successively crystallized granitic plutons 

cooled through Curie temperatures during the interval of 

caldera formation, initial regional extension, and rotational 

tilting of the volcanic rocks (Hagstrom and Lipman, 1986). 

Isotopic ages for the intrusions range from 26 Ma within the 

caldera, indistinguishable from the ash-flow sh~et, to as 

young as 18 Main the southern deeper parts of the composite 

batholith (Lipman and others, 1986; K. Foland and G. Czamanske, unpub. 40/39 Ar ages); younger radiometric ages to 

the south probably reflect both continued magmatism and 

slower cooling in deeper parts of the magmatic system. 

These relations indicate that the batholithic complex broadly 

represents the source magma for the volcanic rocks, into 

which the Questa caldera collapsed, and that the magma was 

mostly liquid during regional tectonic disruption. 

Petrologic and isotopic studies indicate that the volcanic 

rocks constitute intermittently quenched samples of upper 

parts of the Questa magma bodies at early stages of crystallization; in contrast, comagmatic granitic rocks preserve an 

integrated record of protracted crystallization of the magmatic residue as volcanism waned (Johnson and others, 

1989). Multiple differentiation processes were active during 

evolution of the Questa magmatic system: crystal fractionation, replenishment by mantle and lower crustal melts of 

varying chemical and isotopic character, mixing of evolved 

and primitive magmas, upper crustal assimilation, and perhaps volatile transfer (Johnson and others, 1990). As a result, 

an evolving cluster of coalesced magma chambers generated 

diverse assemblages of broadly cogenetic rocks within a few 

million years. Modelling of the fractionation processes indicates that large volumes of mantle-derived basalt interacted 

with the lower crust; return of ultramafic residua to the mantle (fig. 2) is required to have kept crustal thickness and surface elevations from having reached impossibly high values. 

Major molybdenum stockwork deposits are associated 

with the most fractionated granitic and aplitic intrusions 

along the southern caldera margin (Leonardson and others, 

1983); similar but lower grade deposits also are present near 

the roof of a late intrusion south of the caldera. Molybdenum 

is concentrated along the roofs of intrusive cupolas that are 

localized within fault-controlled horsts related to regional 

extensional faulting. The highest ore grades are also associated with low-angle fractures influenced by the southern 

Ash-Flow Calderas as Structural Controls of Ore Deposition L3 
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Upper crust 

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Clustered central volcanics 

60~----------------------------------------------------------------------------~ 

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20 

30 

40 

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Upper crust 

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(to depth of several hundred kilometers) 60~----------J-_______________________________ ~ ___ __L 

L4 Application of Structural Geology to Mineral and Energy Resources 
Figure 2 (facing page). Crustal model illustrating generalized 

evolution of Latir volcanic field and cogenetic intrusions. Schematic sections are oriented north-south to minimize complicating effects of developing Rio Grande Rift. Details of magmatic 

features not to scale. Seismic P-wave velocities are regional 

present-day values for southern Rocky Mountains. Modified 

from Lipman (1988). A, Growth of central volcanoes during 

early (waxing) magmatism and initial accumulation of magmas 

in upper crust, before caldera formation (28-26.5 Ma). Primitive basaltic magma (phase 1 diapirs) rises from below lithosphere (subducted slab interacting with asthenosphere?), partly 

equilibrates with lithospheric mantle as it rises, ponds near 

base of crust, causes partial melting in lower crust (xxx), and 

crystallizes (heavy stipple) and fractionates as heat is transferred. Magmas segregating and rising within crust (phase 2 diapirs) are mixtures of fractionated mantle basalt and crustal 

melts. These magmas feed central volcanoes (stippled) and 

small subvolcanic plutons (phase 3 bodies). uplift of igneous 

area (from horizontal short-dashed line; not to scale) is due to 

buoyant ponding of basalt at base of crust, to rise of lower 

crustal melts into upper crust, and to thermal expansion of 

lithosphere associated with magmatism. 8, Later (waning) 

magmatism, during ash-flow eruption, caldera subsidence and 

resurgence, and consolidation of shallow crustal batholith 

(26.5 Ma). Compositionally varied cupolas develop in upper 

crust, along top of broadly laccolithic batholith. Southern roof 

of batholith solidified at crustal level several kilometers deeper 

than subcaldera northern part, perhaps because of more mafic 

mean composition. Compositional gradients near roof, becoming more mafic downward, are indicated by dotted lines. Crystallized and fractionated basaltic magma underplates and 

mixes with lower crustal material, replacing more evolved 

melts that have migrated upward; basal underplated material 

(heavy stipple) consists of ultramafic cumulates that merge in 

physical and chemical properties with lithospheric mantle, or 

may be sufficiently dense to sink deeper into mantle. Thermally 

weakened crust wi II accommodate extension along magmatic 

loci where plate stresses are conducive, permitting later mantle 

basalt to reach surface with less residence time and chemical 

interaction with lower crust (inception of Rio Grande Rift). 

Slow cooling of crustal section, during waning of magmatism 

and later, will cause subsidence over core of batholith, but 

broader regional uplift related to lower density of crust, and 

perhaps thinning of lithospheric mantle, will remain. 

wall of the caldera (Leonardson and others, 1983). The 

worldwide association of molybdenum deposits with highly 

evolved granites indicates that molybdenum is concentrated 

by processes of igneous differentiation rather than mobilized 

from upper crustal country rocks by magma-related hydrothermal circulation. Such magmatic concentration processes 

may be favored by the open-system behavior between the 

lower crust and upper mantle indicated for the Questa magmatic system, in which mantle-derived heat fuels protracted 

crystal fractionation of silicic differentiates. Thus, formation 

of the Questa deposits probably required an especially favorable structural setting that involved concurrently intersecting 

caldera-related and regional rift faults in association with an 

evolved magmatic system driven by open-system interactions between mantle and lower crust. 

CENTRAL SAN JUAN CALDERA 

CLUSTER, COLORADO 

The San Juan volcanic field consists mainly of intermediate composition lavas and breccias, erupted at about 35-30 

Ma from scattered central volcanoes, overlain by at least 15 

widespread voluminous ash-flow sheets erupted at 30-26 

Ma from caldera sources. It thus represents a slightly earlier 

(pre-Rio Grande Rift), larger, and more volcanologically 

complex analogue of the Latir-Questa system. 

The much studied central caldera cluster, active at 28-26 

Ma, contains six calderas (fig. 3), which host the major epithermal base- and precious-metal veins of the Creede District (Steven and Ratte, 1965; Steven and Lipman, 1976; 

Lipman and Sawyer, 1988; Lipman and others, 1989). The 

caldera-related ash-flow sheets of the central cluster represent some of the largest volume pyroclastic eruptions known 

in Earth history (at least 3,000 km3 for the 27.8-Ma Fish 

Canyon Tuft) and vary in composition from dacite to lowsilica rhyolite, in contrast to the high-silica rhyolite tuff 

erupted from the Questa caldera. Erosion has been insufficient to expose any granitic rocks associated with the central 

caldera cluster, but the calderas lie axially within a large negative gravity anomaly interpreted to reflect the presence of a 

composite batholith at shallow depth (Plouff and Pakiser, 

1972). Detailed resolution of the rapid sequence of volcanic 

and structural events within the central cluster is possible 

because of high-resolution 40/39 Ar dates for all major ashflow sheets (Lanphere, 1988). 

Exposed levels of the central calderas are dominated by 

topographic walls, resurgent domes containing keystone graben faults, and intervening moat depressions that localized 

thick postcollapse volcanic and sedimentary deposits (fig. 

3). All caldera structural boundaries (inferred to be ring 

faults) are covered, but combined geologic and geophysical 

data suggest that the aligned north-trending keystone graben 

faults of three western calderas (27.3-Ma Bachelor, 26.6-Ma 

Creede, and 26.1-Ma San Luis) were localized along the 

western segment of the concealed ring fault bounding the 

large earlier (27 .8 Ma) La Garita caldera (Bethke and Lipman, 1987). Creede vein mineralization (25.1 Ma) is associated with late movement along these graben faults. The 

faults thus have had complex movement histories, first as a 

segment of the early La Garita ring fault, then as three separate keystone grabens related to resurgent doming of discrete 

calderas at 0.5-0.6-m.y intervals, and again about 1 m.y. 

later during vein mineralization. Additional periods of fault 

movement are possible, and an important lesson from the 

central cluster is the complex history of mineralized structures that could only be resolved under exceptionally favorable circumstances, provided at Creede by the recurrent 

datable volcanic events within a relatively brief time span. 

Restudy of the key junction area for three of the calderas in 

the northern part of the mining district (Lipman and Sawyer, 

Ash-Flow Calderas as Structural Controls of Ore Deposition LS 
~ 38°00' ~ 

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X \ Baldy\..-

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Mineral \ Stewart 

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10 

L6 Application of Structural Geology to Mineral and Energy Resources 

South Fork0 

MH 

20 KILOMETERS 
Figure 3 (facing page). Generalized geology of central San 

Juan caldera cluster. Caldera margins are indicated by hachured lines (dashed where approximate, dotted where concealed); heavy stipple pattern indicates intracaldera resurgent 

uplifts. Key to calderas: 8, Bachelor; C, Creede; CP, Cochetopa 

Park; LG, La Garita; MH, Mount Hope; SL, San Luis; SR, South 

River. Random hatch pattern indicates early intermediate-composition volcanic rocks (precaldera); unpatterned areas are 

caldera-related ash-flow tuffs and lavas; solid black areas are 

andesite-monzodiorite intrusions along south margin of South 

River caldera; screnned pattern indicates approximate area of 

pyritic and argillic hydrothermal alteration. Generalized faults 

(bar and ball on downthrown side) shown only for the Creede 

Graben area. Modified from Lipman and Sawyer (1988). 

1988) shows how misinterpretation of some caldera-wall 

unconformities as faults and some caldera-fill landslide 

deposits as brecciated intrusions has been confusing for mineral exploration strategies. 

The Creede vein mineralization and associated hydrothermal system cut across the three western calderas, apparently driven by a heat source in the northern part of the 

district (fig. 4). Fluid flow was asymmetrically down the 

paleo- (and present-day) topographic gradient (Bethke and 

Lipman, 1987). Geologic, isotopic, and fluid-inclusion data 

suggest that ore deposition was controlled by special circumstances involving mixing of deep-circulating brines from 

playa sediments of the Creede caldera moat with dilute meteoric fluids. Mining production is dominated by silver and 

base-metal sulfides, and lead isotopic compositions of 

galena indicate that the bulk of the lead, and presumably 

most other metals, in the Creede ores has been leached from 

Precambrian rocks at depths of 3-5 km or more (Doe and 

others, 1979). The ore lead is notably more radiogenic than 

the associated Tertiary volcanic rocks. Thus, special circumstances of volcanic structural control, compositions of ore 

fluids, and sources of metals are all involved in genesis ofthe 

Creede deposits; apparent lack of ore deposits at some other 

San Juan calderas may reflect absence of such overlapping 

favorable conditions. 

Essentially all ores produced from the San Juan region 

are now known to have been localized by caldera structures, 

except for local minor production from the cores of precaldera central volcanoes. Even some previously seemingly 

isolated loci for intense pyritic mineralization and supergene 

alteration along the southern margin of the central cluster 

have recently been found to occur in proximity to ring intrusions along the southern margin of the recently recognized 

South River caldera (fig. 3) (Lipman and others, 1989). The 

presence of this 27.1-Ma caldera was missed during earlier 

studies because, unlike previously recognized calderas of the 

central cluster, the South River caldera did not resurge after 

ash-flow eruption ceased but rather was filled to overflowing 

with postcollapse lavas. The presently recognized population of calderas in the western United States is probably 

heavily biased toward such more readily identifiable resurgent calderas. Accordingly, additional unrecognized mineralized areas may be associated with obscure nonresurgent 

calderas. 

MINERALIZED MESOZOIC CALDERAS OF 

SOUTHEASTERN ARIZONA 

Southeastern Arizona and adjacent areas constitute one 

of the largest porphyry copper provinces in the world (Titley, 1982). Most of the deposits are associated with Late 

Cretaceous-early Tertiary (Laramide-age) intrusions; 

related volcanic rocks, ranging in composition from 

andesite to silicic rhyolite record the development of a 

continental-margin volcanic arc. Porphyry copper mineralization has commonly been regarded as associated with the 

roots of andesitic volcanoes (Sillitoe, 1973; Branch, 1976), 

or with intrusions that did not vent, and with antithetic to 

caldera structures (Sillitoe, 1980). Although silicic volcanism has commonly been considered unrelated to such copper mineralization-in contrast to its generally recognized 

association with porphyry molybdenum systems such as 

Questa-reevaluation of the magmatic setting of the Arizona province suggests a widespread association of silicic 

caldera-related volcanism and porphyry copper mineralization (Lipman and Sawyer, 1985, 1986). 

Recognition of large-scale volcanic features in the mineralized Arizona rocks is complicated by hypogene alteration and multiple structural complexities including 

Laramide compressive deformation, mid-Tertiary low-angle 

extensional faulting, and late Cenozoic block faulting. More 

than half of the pre-Cenozoic bedrock is covered, and only 

fragments of calderas are exposed. The igneous rocks in 

southeastern Arizona also contain striking structural and 

stratigraphic features for which unusual hypotheses have 

been proposed: widespread intrusive pyroclastic phenomena" (Watson, 1968), regional "exotic-block" breccias 

(Simons and others, 1966), and the concept of "volcanic 

orogeny" (Mayo, 1971). Reinterpretation of these features, 

using recent models of caldera formation and silicic volcanism gained through study of active or little eroded terranes, 

suggests the presence of 10-15large Mesozoic caldera fragments (Lipman and Sawyer, 1985). Definitive characteristics of the caldera fragments include (1) ash-flow deposits 

more than 1 km thick, having features of intracaldera ponding; (2) interleaved "exotic block" breccias in tuffaceous 

matrix, interpreted as megabreccias resulting from calderawalllandsliding; and (3) local granitic bodies, interpreted as 

caldera ring intrusions, along arcuate structural boundaries 

of the thick volcanic sequences. Three caldera fragments that 

host porphyry copper mineralization associated with the late 

granitic intrusions have been restudied in detail: Silver Bell, 

Sierrita, and Tucson. 

Ash-Flow Calderas as Structural Controls of Ore Deposition L7 
NORTH 

5 

-1 

/" __.-' " '- '-...__ ...__ _____________ _LPaleosurface 

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

SOUTH 

CREEDE 

CALDERA 

Figure 4. Diagrammatic north-south section longitudinally along Creede Graben along crest of resurgent dome within Bachelor caldera (fig. 3) showing inferred geometry of fossil hydrothermal system in relation to mineralization. Graben faults and 

mineralization extend across three calderas. Volcanic units (ages in Ma): Tab, precaldera andesite flows and breccias; Taf, intracaldera ash-flow tuffs. Inferred caldera-related intrusions: Br, resurgent intrusion of Bachelor caldera; Cr, resurgent intrusion 

of Creede caldera; Ci, ring intrusion of Creede caldera; Si, ring intrusion of San Luis caldera: Mi, mineralization-related intrusion. Solid arrows show inferred circulation of saline fluids derived from Creede Formation; dashed arrows show inferred circulation of dilute meteoric fluids derived from high northern area. No vertical exaggeration. Modified from Bethke and Lipman 

(1987). . 

Upper Cretaceous volcanic rocks (73-69 Ma) are closely 

associated with porphyry copper deposits in the Silver Bell 

Mountains (fig. 5A), where caldera subsidence is indicated 

by thick intracaldera tuff, interlayered collapse breccia, arcuate bounding faults, and ring intrusions (Sawyer, 1989). In 

the Sierrita Mountains (Lipman and Fridrich, 1990) (fig. 

5B), Tertiary tilting is inferred to provide an oblique section 

through an upper crustal magmatic system, including a 

Laramide-age caldera containing thick welded tuffs and 

slide breccias, subcaldera batholith (Ruby Star Granodiorite), and major associated copper deposits (Sierrita-Esperanza, Twin Buttes, Mission-Pima, and San Xavier). Detailed 

mapping and petrologic studies in the Tucson Mountains, 

between the Silver Bell and Sierrita areas, indicate that thick 

ash-flow tuffs and multiple associated lenses of megabreccia 

(Cat Mountain Tuff, Tucson Mountain Chaos) accumulated 

within a Late Cretaceous trap-door caldera that was resurgently -deformed by a large intrusion (Amole pluton); 

porphyry-type alteration and subeconomic copper mineralization occurred locally adjacent to the intrusion (Lipman 

and Fridrich, 1990). Broadly similar mineralization-related 

Mesozoic calderas have also been interpreted for the Santa 

Rita and Tombstone areas (fig. 5), based on field reconnaissance and reinterpretation of published data. 

In the northern Dos Cabezas Mountains, thick varicolored pyroclastic breccias, exposed in oblique cross section 

due to Tertiary rotational faulting, were interpreted as intrusive by Erickson (1986); recent mineral resource studies by 

L8 Application of Structural Geology to Mineral and Energy Resources 

the U.S. Geological Survey report discrete breccia lithologies and intrusive breccia pipes considered to be potentially 

significant structural controls of mineralization (Drewes, 

1985; Drewes and others, 1988). Alternatively, the entire 

breccia sequence, including under- and over-lying welded 

tuffs, has been interpreted as a single thick depositional ashflow cooling unit, of variable lithic content, that filled a 

Laramide-age caldera (Lipman and Sawyer, 1985, 1986). 

The "breccia pipes" mapped by Drewes, rather than being 

intrusive bodies, are marked by unmappable irregular gradations within variably altered massive lithic tuff of the caldera 

fill; some areas of dark rock are propylitically altered, others 

are reddish due to oxidation or bleached by supergene leaching of pyritic rock. Such variable alteration, which is common in caldera-fill sequences, has different (and more 

limited) significance for mineral exploration than brecciapipe structures. 

Other mineralized areas associated with Mesozoic intrusions could represent the roots of calderas that are so deeply 

eroded that only precaldera rocks are preserved. For example, as suggested to me by David Sawyer ofthe U.S. Geological Survey, the major San Manuel-Kalamazoo porphyry 

copper deposit may be a western structurally dismembered 

part, perhaps representing a deeper structural level, of the 

mineralized caldera system that is fragmentarily preserved 

15 km to the east at Copper Canyon (Lipman and Sawyer, 

1985). Both mineralized areas are associated with -69-Ma 

monzogranite porphyry that intrudes upper crustal rocks 
:································· 

AD 

B 

1W30' 

\ 

\ 

~ j 

I 

;/ 

;/ 

// 

y 5 KILOMETERS 

31045 , .--_:_11-=,0o--=-o~o· _________ ---r'"~~~,..,...-, 

0 

pit 

ea.~7'• floo,, 

D 

.. · ••••••••• / 5 KILOMETERS 

EXPLANATION 

Cenozoic rocks 

Granitic rocks 

Postcollapse volcanic rocks 

Collapse breccia 

.··, K;,' ,, . 

· Ka ·· 

··. PI< 

ClJ}J 

Ash-flow tuffs 

Precaldera volcanic rocks 

Pre-Late Cretaceous rocks 

Porphyry copper mine 

Figure 5. Generalized geology of four mineralized areas in southeastern Arizona interpreted as containing fragments of Late 

Cretaceous calderas. Caldera margins shown by hachured lines, dashed where approximately located. See text for discussion 

of individual areas. Modified from Lipman and Sawyer (1985). 

Ash-Flow Calderas as Structural Controls of Ore Deposition L9 
northeast of the Santa Catalina core complex, and they may 

have been disrupted by extensional faults in a manner similar 

to that demonstrated on smaller scale for the San ManuelKalamazoo deposit (Lowell, 1968). Possibly analogous offsets of Laramide-age porphyry deposits by detachment 

faulting have been documented on the east flank of the Sierrita Mountains (summarized in Titley, 1982). 

Relations between mineralization and older Mesozoic 

volcanism in southeastern Arizona, apparently mainly Jurassic in age, are even less understood than those for Laramideage deposits. Fragments of several calderas and associated 

outflow tuffs are preserved in the Canelo Hills-Huachuca 

Mountains, Patagonia Mountains, and Pajarito Mountains 

(Lipman and Sawyer, 1985; Riggs and others, 1990). Local 

mineralization of these volcanic rocks, and association of the 

porphyry copper deposit at Bisbee with a Jurassic intrusion, 

suggests additional mineral potential associated with the 

Jurassic caldera-related rocks. 

An instructive association of probable Jurassic caldera 

volcanism and copper porphyry mineralization is the Courtland-Gleeson District. There, lithologically intermixed 

Paleozoic rocks, previously mapped in detail as a complex 

Mesozoic thrust fault zone (Gilluly, 1956; Drewes, 1981), 

consist of megablocks depositionally enclosed in rhyolitic 

welded tuff that is spatially associated with a large Jurassic 

granitic intrusion. Less altered Cretaceous welded tuff (Sugarloaf Quartz Latite) unconformably overlies (rather than in 

thrust contact as mapped by Gilluly and Drewes) both the 

tuffaceous megabreccia and Jurassic intrusive rocks, suggesting that these rocks are parts of a caldera complex that 

was mineralized in the Jurassic. My reinterpretation of these 

lithologic complexities as due to primary volcanic depositional processes, rather than to Mesozoic thrust tectonics, 

provides a new structural framework for mineral exploration 

in the district. 

The clustering of Mesozoic porphyry copper deposits in 

Arizona may be partly due to level of erosion, intermediate 

between that of the base- and precious-metal veins of a San 

Juan volcanic field and deeper plutonic mineralization such 

as the tungsten skams in the Sierra Nevada. Widespread 

preservation of Mesozoic caldera fragments in southern Arizona, in contrast to limited preservation of outflow volcanic 

sequences, probably reflects both the lower structural levels 

represented by the subsided caldera fills and the greater 

resistance to erosion characteristic of the thick wellindurated caldera-fill material. The spacing of identified 

caldera fragments in southeastern Arizona is comparable to 

that in well-preserved Tertiary volcanic fields such as the 

San Juan region, and the original caldera density may have 

been even greater in Arizona, especially if effects of Tertiary 

extension and cover by basin alluvium are considered. Recognition that many of the most important Arizona ore deposits are structurally localized along caldera margins and 

related subvolcanic structures and intrusions provides a new 

conceptual framework for further exploration for concealed 

deposits. 

DISCUSSION 

The structural association between caldera magmatism 

and varied ore-deposit types, including epithermal base- and 

precious-metal veins, both copper and molybdenum porphyries, and deeper level replacement ores, is now well established (Steven and others, 1976; McKee, 1979; Rytuba, 

1981). Important topics in need of further study include the 

life span of caldera-related hydrothermal systems and associated mineralizing processes. The time constraint depends 

on the longevity of the magmatic heat source, which in tum 

involves interplay between convective-conductive cooling 

and replenishment by addition of new magma to the upper 

crustal system. Especially valuable' will be multidisciplinary 

studies of exceptionally young mineralized systems exposed 

by rapid tectonic uplift; most such systems are outside the 

United States. For example, at a recently recognized 2.5-Ma 

ash-flow caldera in the north-central Caucasus Mountains 

(Russia), associated Pleistocene granites and major molybdenum-tungsten deposits (broadly analogous to Questa, New 

Mexico) have been exposed by 3-5 km of Quaternary uplift 

(present rates as much as 12 mm/yr) due to the impingement 

of Africa and Arabia against the Eurasian craton (Lipman 

and others, 1990). In this system, apparent cooling ages indicated by available K-Ar determinations on the granites vary 

from essentially contemporaneous with the caldera-related 

tuff to as much as 1 m.y. younger; ongoing detailed chronologie studies utilizing isotopic and fission-track methods 

should provide a high-resolution record of magmatic cooling, fluid flow, and temporal evolution in the roots of this 

major mineralized geothermal system. 

Another long-standing question concerns the sources of 

the metals in caldera-hosted and other igneous-related mineral deposits. The relative roles of metal transport from great 

depths by magmas versus concentration of ores from upper 

crustal country rocks has been controversial since the 19th 

century. Oxygen isotope studies show that most igneousrelated ores were deposited from hydrothermal fluids in 

which deeply circulating meteoric water was a major or 

dominant component (Taylor, 1974) and that associated 

igneous bodies served mainly as the thermal engines to drive 

a convecting hydrothermal system (Norton, 1982). Calderas, 

as the surface expressions of large shallow silicic magmatic 

systems, are particularly favorable sites for such processes. 

For example, lead in epithermal veins in the San Juan volcanic field is more radiogenic than in associated igneous rocks 

and mainly must have been leached by hydrothermal fluids 

from Precambrian rocks that underlie the volcanic field (Doe 

L 1 0 Application of Structural Geology to Mineral and Energy Resources 
and others, 1979). Some would argue that most Phanerozoic 

ore deposits represent recycling of metals originally concentrated in the Precambrian, especially as massive sulfide 

deposits generated during processes of crustal formation 

(Hutchinson, 1980). Other metals, such as in molybdenum, 

uranium, and beryllium deposits, are consistently associated 

with distinctively evolved silicic magmas commonly 

emplaced during regional extension; these metals probably 

were concentrated as incompatible components during magmatic differentiation (Christiansen and others, 1983; Stein 

and Hannah, 1985). Even these metals, along with their host 

magmas, probably were mostly recycled from lower cratonic 

crust that has been variably melted and mobilized by mantlegenerated mafic magmas. 

Other priority targets for more detailed studies include 

young or even presently forming caldera-related mineralization encountered in geothermal exploration drilling, as in 

New Zealand and at the Valles caldera in New Mexico 

(Ewers and Keays, 1977; Hulin and others, 1987). Such situations permit direct observation of the deep geometry and 

physical conditions in mineralizing geothermal systems. 

Another attractive target for future research involves 

research drilling through well-documented caldera-related 

ore deposits into deep parts of the fossil geothermal environment and possibly into the intrusive heat source, as proposed 

for Creede, Colorado (Bethke and Lipman, 1987). Crustal 

sections providing natural exposures from the near-surface 

environment down to the fossil magma chamber, as represented by a solidified batholith, occur naturally only due to 

complex structures, as at Questa, New Mexico, and Sierrita, 

Arizona, that leave significant ambiguities concerning the 

original geometry. Research drilling into structurally 

straightforward systems provides the only available method 

to overcome such ambiguities in the third dimension. 

Acknowledgments.-The areal studies and concepts discussed here mainly represent a decade of products from my 

U.S. Geological Survey Roots of Ash-Flow Calderas 

Project, which has been supported mainly by the Geothermal 

Research and National Geologic Mapping Programs. Significant parts of this work were carried out in close collaboration with Philip Bethke, Jon Hagstrom, Clark Johnson, 

David Sawyer, John C. Reed, Jr., and others in the Survey, 

academia, and industry too numerous to acknowledge individually. 

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L 12 Application of Structural Geology to Mineral and Energy Resources 

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