Mines Library University of Nevada - Reno M|N1imii nm/ ° F N E V * D A. RENO

Reno, Nevada 83557-0044 MINES LIBRARY 3 233 00303 6738 'THESIS

132 7- UNIVERSITY OF NEVADA G . X RENO

GEOLOGY OF THE DESERT PEAK GEOTHERMAL ANOMALY CHURCHILL COUNTY, NEVADA

A THESIS SUBMITTED IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF Master of Science in Geology by Mines Library University of Nevada - Reno Reno, Nevada 89557-0044

John Edward Hiner

M a r c h 1979 The thesis of John Edward Hiner is approved:

> .. . Thesis Advisor

De m T Graduate School

University of Nevada Reno

M a r c h 1979 i s *

11

acknowledgements

I "ish to thank the following persons and institutions;

Phillips Petroleum Company: for finaneia! aid and for access to

temperature, geochemical, and geophysical data, aerial photography,

rn-house reports, drillhole information, cuttings, cores, lithologic logs, and temperature profiles;

Nevada Air National Guard- for inf™ . a • or infra-red imagery of the Hot Springs Mountains and small-scale standard aerial photography of the Desert Peak Geothermal field;

Mackay School of Mines: Professors J. Lintz E R | p-ps uitz, t.K. Larson, and Dean A. Baker III, for encouragement, invaluable advice and f aavice, and time spent with - in the field and office; Dr. M. Campana, for advice and direction

concerning hydrologic aspects of the area and critical review of the groundwater section;

Nevada Bureau of Hines and Geology, Mr. E. Bingler, for advice on

drafting technique, discussions on Miocene stratigraphy, and for thin

sections of cuttings from wells 29-1, B21-1, and B21-2.

«eep appreciation is extended to the Reno staff of Phillips Petro­

leum Company, particularly W.R. Benoit and R.T. Forest, who gave freely

» their time and expertise, and provided useful advice, insight, on-the-

ritique of my efforts, and critical review of the manuscript.

Thanks are extended to Ms. S. Hughes, for aid and advice in drafting

I am especially indebted to my wife, Shirley, who displayed saintly patience, gave me encouragement, and typed the manuscript. Ill

ABSTRACT

Field mapping in the northern Hot Springs Mountains, Churchill

County, Nevada indicates that the intersection of N25E trending Basin and Range faults with an earlier N6QE trending fault zone corresponds with the location of the Desert Peak Geothermal Field. Uplift of the thermally anomalous area occurred twice, first along the N60E trending zone, probably in late Miocene, and again along N25E trending faults in the Pliocene. The two episodes of uplift brought Mesozoic greenstones which contain the geothermal reservoir close to the surface, thereby permitting detection and potential economic exploitation of the blind geothermal resource. Geology, water chemistry and groundwater patterns indicate that the Desert Peak Geothermal Field is structurally and hydrologically unrelated to the Brady's Hot Springs thermal anomaly,

with the possible exception of a shared heat source„ iv

TABLE OF CONTENTS

Introduction and Purpose 1

Procedure 2

Location and Physical Features 4

Previous Work 7

Regional Geology 10

Rock Units 17

Mesozoic Rocks 19

Tertiary Rhyolites 23

Chloropagus Formation 27

Desert Peak Formation 32

Truckee Formation 35

Younger Volcanic Rocks 38

Tertiary Andesite 40

Desert Queen Intrusive 42

Quaternary Deposits 44

Structure of the Northern Hot Springs Mountains 46

Geologic History 55

Ore Deposits 57

Desert Peak Geothermal Field and Geologic Structure 59

Summary 74

Conclusions 76

References Cited 79 V

FIGURES

1. Location Diagram 5

2. Regional Features 9

3. Batholith Distribution 11

4. Mesozoic Stratigraphy 21

5. Tertiary Rhyolite Stratigraphy 26

6. Chloropagus Formation 31

7. Desert Peak Formation 34

8. Truckee Formation 37

9. Simplified Structure Map 45

10. Drape Fold Origin 47

11. Rhombohedral Horst Formation 54

12. Characteristic Temperature Profiles 61

13. Groundwater Movement 64

14. Conceptual Model of Geothermal System 73

PLATES (IN POCKET)

1. Geology of the Northern Hot Springs Mountains

2. Cross-sections of Northern Hot Springs Mountains INTRODUCTION and PURPOSE

In" 1976, Phillips Petroleum Company culminated a five year exploration effort by drilling a successful geothermal well in the northern part of the Hot Springs Mountains, northwestern Churchill

County, Nevada. Exploration techniques consisted of shallow tempera­ ture-gradient holes (200-300 feet; 61-92 meters), deep stratigraphic tests (1000-2000 feet; 305-610 meters), followed by the geothermal deep test. Geochemistry, geophysics, and areal geology were not included in the exploration architecture. This thesis constitutes an effort to establish a relationship between surface rock distribution, geologic structure, and the localization of the geothermal reservoir. 2

PROCEDURE

Field mapping began in November 1977. Forty-five man-days were spent mapping approximately 57 square miles (148 sq. km) in the northern

Hot Springs Mountains. Data were recorded on USGS aerial photography * (series GS-EL, 1947: scale-1:43000). Information was then transferred

by inspection to computer-corrected USGS 7.5" orthophotographic quad­

rangles: Two Tips NE, Soda Lake NW, Soda Lake NE, Fireball Ridge SE,

Desert Peak SW, and Desert Peak SE (scale-1:24000), and then traced

onto a mylar screen of the topography. The mylar of the topography

is a magnification of four USGS 15' quadrangles: Desert Peak (1951),

Fireball Ridge (1957), Soda Lake (1951), and Two Tips (1957).

Geologic cross-sections were constructed utilizing the surface

geology and information from Phillips stratigraphic test holes (here­

after referred to as strat. tests), and deep geothermal wells.

Infra-red imagery of the northern Hot Springs Mountains was tone-

correlated with surface ground temperatures by Dr. J. Lintz and me.

Temperature measurements were made with an 18 inch (46 cm) mercury-in­

glass thermometer, and ground temperatures 12 inches (30 cm) below the

surface were taken every 500 feet (150 m) along two traverses: one

crossed the distinctly anomalous Brady's Thermal Fault Zone, and the

other crossed the Desert Peak Geothermal Field.

Geochemical and groundwater data were incorporated to define the

hydrologic regime and groundwater flow systems. Geochemical data were * taken from in-house reports to Phillips Petroleum Company and consist

of water analyses of thermal waters from drillholes and of available

surface waters in the area. Water-level data were obtained from a 3 report by Olmstead and others (1975) on the Brady's Hot Springs System, and from Phillips Petroleum Company drillholes. Possible recharge rates, water needs, waste fluid reinjection, reservoir characteristics, and legal and institutional aspects of the geothermal discovery are beyond the scope of this paper, and therefore are not treated.

Reconnaissance gravity and magnetic data gathered by Phillips

Petroleum Company personnel, and an electrical resistivity study run by

Geonomics, Inc. for Phillips provided generalized confirmation of major geologic aspects and were used to make a first approximation of basement tectonic trends.

Thermal data from Phillips were utilized in the attempt to establish the relationship between geologic structure and the thermal anomaly.

Fifty-eight rocks were thin sectioned and selected rock samples were stained for feldspar analysis. 4

LOCATION and PHYSICAL FEATURES

The Hot Springs Mountains are located approximately 15 miles

(24 km) northeast of Fernley, Nevada, in the northwestern corner of

Churchill County, Nevada (figure 1). Brady's Hot Springs, a well- known rest stop during the westward migration, is located 4 miles

(6.4 km) west of Desert Peak. The Desert Peak Geothermal Field is situated in the northern Hot Springs Mountains approximately 2 miles

(3.2 km) south of Desert Peak.

The Hot Springs Mountains are a low range of northeast trending hills with generally subdued topography. Elevations vary from 4000 feet (1220 m) in the southwest portions to 5365 feet (1635 m) at

Desert Peak. Present-day climate in the region is arid. Rainfall is estimated at 4 to 6 inches (10-15 cm) per year (Harrill, 1970). Most of the precipitation is received during winter months, although occasional summer thunderstorms furnish small additional amounts., Rainfall records from nearby Fernley show a January average of .80 inches (2.1 cm), while the July average is .22 inches (.56 cm). Average January low temperatures are about 15°F (-9°C), and average July high temperatures are about 90°F

(32°C). Temperatures can be extreme, however, and may vary between winter lows of -25°F (-33°C) and summer highs of 110°F (43°C)„

Vegetation is sparse but variable in response to soil conditions, water availability and quality, and altitude. A Basin Sage community dominates the higher elevations, giving way to Shadscale populations on alluvial fans and on the flats. Interestingly, surface geothermal manifestations are characterized by saltgrass (Distlichus stricta) and samphire (Salicornia rubra). A more complete inventory of present flora may be found in Axelrod (1956, p.93) 5

FIGURE 1. Location diagram There are no springs in the Hot Springs Mountains, and waters around and near the margins of the hills are moderately saline.

This plus the sparse vegetative cover effectively limits animal life to a coyote and rabbit population, and substantial numbers of lizards, snakes, and rodents.

The Hot Springs Mountains are accessible via Interstate Highway

80 on the west and Nevada State Highway 95 on the east. Numerous

unpaved roads and jeep trails of fair to poor quality traverse the

interior reaches of the northern Hot Springs Mountains and provide

access to most of the area with four-wheel-drive vehicles. However,

large portions of the Hot Springs Mountains are veneered by wind-blown

sand, which can impede off-road travel in conventional two-wheel-drive.

Also, mud can be a year-round problem in and near the playas around

the Hot Springs Mountains.

Land use is limited at present to cattle grazing and geothermal

activities. Geothermal Food Processors Inc. recently commenced the

world's first geothermal-powered food processing operation at Brady's

Hot Springs. The Desert Peak geothermal field is in the process of

development. A notable lack of scenic or unique features has minimized

recreational use. 7

PREVIOUS WORK

The earliest published information on the Hot Springs Mountains is contained in the report on the Geology of the Fortieth Parallel, in which

King (1878, vol. 1, p. 423) defined the Truckee Formation from exposures on Fossil Hill in the northeastern part of the Hot Springs Mountains.

Russell (1885) compiled a now-famous monograph on Pleistocene Lake

Lahontan, and therein described Brady's Hot Springs and surrounding area.

Geologists of the Southern Pacific Company conducted reconnaissance field mapping and mineral occurrence inventories in the 1950's and early 1960's

(unpublished reports to Southern Pacific Company). Yen (1950) studied molluscs from the type section of the Truckee Formation. Vanderburg

(1940) examined the mining districts of Churchill County, and gives a brief description of the Desert District, which is situated in the north­ eastern portion of the Hot Springs Mountains. Willden and Speed (1974), as part of a comprehensive study of the geology and mineral deposits of

Churchill County, mapped in the Hot Springs Mountains. Axelrod (1956) mapped a small portion of the northern Hot Springs Mountains as part of

an analysis of the fossil flora of the area. His work represents the most detailed study on the stratigraphy and structure to date. Harrill

(1970) conducted a water resources appraisal of Granite Springs Valley

and surrounding areas, which includes the western part of the Hot Springs

Mountains groundwater regime. Oesterling and Anctil (1962) discussed

the geothermal potential, geology, and economics of Brady's Hot Springs.

Olmstead and others (1975) appraised the hydrogeology of selected geo­

thermal systems in Nevada, including Brady's Hot Springs system.

Geologists of Phillips Petroleum Company, principally R.T. Forest and 8

W.R. Benoit, have compiled substantial amounts of information on strati­ graphy, geochemistry, and temperature distribution during the course of exploration and discovery of the Desert Peak Geothermal Field; this information is contained in in-house reports that were made available to me. 9

FIGURE 2 regiona I features 10

REGIONAL GEOLOGY

The Hot Springs Mountains lie in west-central Nevada near a transi­ tional zone between the Basin and Range province and the Sierra Nevada province (Bonham, 1969). They also form part of the western boundary of the Carson Sink depression (figure 2).

Information concerning pre-Mesozoic stratigraphy and structure west of the Carson Sink is sparse. Available evidence from adjacent regions north and east of the Carson Sink suggest that west-central Nevada was the site for deposition of eugeosynclinal rock types throughout most of the Paleozoic (Silberling and Roberts, 1962: Bonham, 1969). The late

Paleozoic Antler and Sonoma Orogenies disrupted and telescoped the eugeosynclinal facies rocks eastward (Roberts and others, 1958), almost certainly affecting the Hot Springs Mountains region. Unfortunately, documentation of these effects is lacking, largely because the evidence has been covered or obliterated by post-Paleozoic rocks.

Rocks exposed in the Hot Springs Mountains region range in age from

Triassic(?) to Holocene. Mesozoic rocks comprise metamorphosed sedimen­ tary and volcanic rocks, which are intruded by mafic and siliceous plutons.

Mesozoic rocks are extensively exposed in the ranges east of the Carson

Sink, in the West Humboldt Range and in the Trinity Range (figure 2).

In general, well stratified Mesozoic rocks are more prevalent east of the

Carson Sink, whereas to the west relatively nondescript metasedimentary, metavolcanic, and granitic intrusive rocks predominate (figure 3). Where

Mesozoic strata are exposed in west-central Nevada, their age and strati­ graphic relationships are unclear (Bonham, 1969; Johnson, 1977; Willden and Speed, 1974), owing to a lack of readily identifiable marker beds or 11

120 118 H6 ||4

outline of Nevada showing eastern limit of Mesozoic granitic intrusions. West of dashed line more than 50% of exposed pre-Tertiary rock is granitic. 12 lithologically distinct units, discontinuous outcrops, and disruption by intrusions of probable Cretaceous age. In Pershing County, 20 miles

(32 km) northeast, Johnson (1977, plate 1) assigned lithologically homogeneous rocks exposed in the West Humboldt Range to the Juro-Triassic pelitic deposits of the Auld Lang Syne Group. A large hornblende gabbro intrusive, associated differentiates, and comagmatic extrusive units are exposed in ranges along the northwest, north and east borders of the

Carson Sink (Willden and Speed, 1974).

Early Mesozoic structure in the region is inadequately known.

Structural complexities and their continuity are difficult to ascertain, owing to regional metamorphism and the -homogeneity of the pre-intrusive

Mesozoic section.

Middle Jurassic folding and faulting is widespread in west-central

Nevada. In many places, such as the Sand Springs Range (figure 2), folds are refolded, implying at least two phases of diastrophism. Near the

Hot Springs Mountains refolding occurs in the West Humboldt and Trinity

Ranges (Speed, 1968; Johnson, 1977).

In middle Jurassic time, intrusion of a large, dish-shaped lopo- lithic gabbro complex resulted in folding and thrust faulting. The lopolith intruded a small tectonic basin partially exposed at present in the Clan Alpine Mountains 30 miles (80 km) east of the Hot Springs

Mountains, in the Stillwater and West Humboldt Ranqes, and the

Mopung Hills (Speed, 1968). The depositional and structural history of the lopolith is extremely complex. Rocks of the lopolith comprise intrusives of diorite to gabbro composition and comagmatic extrusive

units of basalt. The gabbro complex is floored by pelitic sediments of

the Auld Lang Syne Group and intruded a syntectonic suite of middle 13

Jurassic carbonates and arenites which collected in the small Boyer Ranch

depositional basin (Speed and Jones, 1969). Two major thrust fault zones

are associated with the lopolith: the Fencemaker Thrust and the Wildhorse

Thrust. The Fencemaker Thrust is exposed in the Humboldt and Stillwater

Ranges and in the Clan Alpine Mountains. The Wildhorse Thrust is exposed

in the West Humboldt Range. Pelitic rocks occur in both upper and lower

plates, and the presumably autochthonous rocks have been metamorphosed by

the gabbro intrusive. Since the Wildhorse Thrust also cuts the lopolith,

Speed (1975) infers that intrusion and thrusting were concomitant. The

presence of several other thrusts in the West Humboldt Range (Speed, 1976)

may be indicative of a pervasive imbricate thrust structure along the

western margin of the lopolith.

The Mesozoic rocks were regionally metamorphosed to greenschist facies

prior to the Cretaceous intrusion of granitic plutons (Bonham, 1969).

Locally intense contact metamorphism, fracturing and faulting of country

rock resulted from the granitic invasion.

Cenozoic geology in west-central Nevada reflects a period of intense

volcanism complicated by pervasive Basin and Range faulting. Rocks of

Tertiary age are primarily volcanic, consisting of complexly interfingered

sequences of flows, tuffs, and shallow intrusive rocks. Fluvial and

lacustrine intercalations occur sporadically, and locally are differen

tiated to formation rank. Quaternary rocks are mostly alluvial and

lacustrine sediments. Cenozoic rocks are separated from the pre Tertiary

section by an unconformity of regional extent (Albers, 1964). Cenozoic

deformation comprises high-angle normal faulting and tilting of basin

and range blocks. Locally, folding is significant (Roberts and others,

1967). 14

01igocene(?) volcanics ranging from basalt to rhyolite in composi­ tion commonly overlie pre-Tertiary rocks in northern Washoe County and in Pershing County (Bonham, 1969: Burke, 1973, and figure 2, this report).

These rocks thin out to the south and east, where Miocene silicic vol­ canics form the base of the Tertiary section (Willden and Speed, 1974).

Deposition of Miocene units was accompanied or followed shortly by block faulting which generated small and probably somewhat elongate, enclosed basins, although some basins may have represented no more than undulations on the surface of the volcanic field (Axelrod, 1956). Lacustrine sedi­ mentation in these basins accompanied continued extrusion of intermediate and mafic volcanics in late Miocene and Pliocene time. The sediments are generally thin, discontinuous, highly variable laterally, and amount to

only minor additions to the volcanic pile. A few basins, however, were

sufficiently large and long-lived to accept great amounts of sediment and

volcanic rocks, giving rise to such predominantly sedimentary accumula­

tions as the Esmeralda, Humboldt, and Truckee Formations (Axelrod, 1956).

Source areas for the voluminous volcanic flows and ash-flow tuffs have

been recognized in the form of collapse calderas. Caldera depressions

have been recognized in Churchill County (Riehle and others, 1972), in

the southern Tobin and southern Trinity Ranges in Pershing County (Burke

and McKee, 1973; Willden and Speed, 1974). Other calderas undoubtedly

occur in the region, as well as other types of vents yet unrecognized.

Basin and Range physiography is largely the result of deformation

which began in early Miocene time (Zoback and Thompson, 1978; Johnson,

1977). Widespread and pervasive high-angle faulting raised, and in many

cases tilted, large blocks of the earth's crust, creating the well-known

horst and graben structure of present day Nevada. Major faults in the 13 region trend north to northeast, bounding the ranges on one or both sides and commonly fragmenting the interior parts of the ranges.

The origin of Basin and Range structure has intrigued many geolo­ gists. The consequence of many years of detailed and regional studies is the development of two principal theories on the origin of Basin and

Range structure. One holds that observed tensional features are directly related to deep-seated extension of the substratum. The second theory interprets Basin and Range extension as an adjustment to regional hori­ zontal shear (Bonham, 1969). Horizontal shear forces marginal to and within the Basin and Range are expressed as northwest trending topographic and structural lineations (Sales, 1966). One such oft-cited structural element, the Walker Lane (Locke and others, 1940), lies about 6 miles

(10 km) southwest of the Hot Springs Mountains. Originally defined as a topographic lineament (Bonham, 1969), recent geophysical work indicates that the Walker Lane is a major, deep-seated crustal disruption of the characteristic Basin and Range pattern in western Nevada (Trexler, 1978).

Local crustal shortening, expressed by folds in Tertiary rocks, has occured in many places in west-central Nevada (Willden and Speed, 1974;

Axelrod, 1936). Fold wave lengths vary from a few feet to thousands of feet (a meter to hundreds of meters), and fold axes are generally sub- horizontal. Fold limbs may dip as much as 70°, but more commonly dips

range from 10° to 40° (Willden and Speed, 1974). Large folds trend north

or northeasterly. Smaller folds often trend and plunge southeast. The

folds usually are near and subparallel to the north-northeast trending

Basin and Range faults, and may be closely related. The folds possibly

originated through basement block faulting or from local compressive

stresses caused by faulting. Several ranges, for example the Hot Springs 16

Mountains, probably owe their existence to folding as much as to faulting

(Wallace, 1964).

Rocks of Quaternary age comprise mostly alluvial and lacustrine deposits which cover well over half the region. Isolated pediment gravels mantle many low-lying areas. Quaternary volcanics consist of

basalt and basalt tuff exposed at Upsal Hogback, Rattlesnake Hill, and

at Soda Lake, all in the Carson Desert (Morrison, 1964). Deposits of

Pleistocene Lake Lahontan underlie the Carson Sink, and are approximately

coeval with the basalts there (Willden and Speed, 1974). Lake Lahontan

deposits are also found in most of the valleys in the region. Finally,

unconsolidated alluvium and local sheet and dune sands deposited by wind

and water now largely obscure older rocks at lower elevations and

occasionally at higher elevations within the region. 17

ROCK UNITS

Introduction

Five major rock types are exposed in the northern Hot Springs

Mountains: (1) a Tertiary (?) hornblende-quartz diorite intrusive near the Desert Queen Mine (plate 1); (2) a basal Tertiary sequence of silicic volcanics ranging from rhyolitic to dacitic ash flows and tuff-breccias;

(3) Tertiary intermediate to basic volcanics, principally andesite and basalt; (4) lacustrine and fluviatile Tertiary sedimentary rocks com­ posed of sandstone, siltstone, shale, diatomite, limestone and volcanic ash; (5) Quaternary Lake Lahontan sediments, alluvium, playa deposits, and sheet and dune sands.

The Tertiary and Quaternary rocks in this area rest unconformably upon Mesozoic greenstones, quartzites, phyllites, marbles, and mafic intrusive rocks. The Mesozoic rocks are not exposed in the Hot Springs

Mountains but are important because the Desert Peak geothermal reservoir

occurs in them. A systematic description from well cuttings is included

for later reference.

The Tertiary and Quaternary rock sequence may exceed 4000 feet

(1220 m) in thickness. Volcanic rocks predominate, although locally the

intercalated sediments form important additions to the stratigraphy.

Phillips well 29-1 intersected over 4500 feet (1370 m) of Tertiary vol­

canics with interbedded sediments, but the aggregate thickness may be

exaggerated by tilting. Surface stratigraphic measurements are impre­

cise at best, owing to discontinuous exposures, a monotonous volcanic

section, and a lack of any areally persistent marker beds. 18

The names applied to the rocks in the mapped area are field terms based on hand specimen examination but have generally been confirmed by thin section study. Because the majority of volcanic rocks are porphyritic with aphanitic groundmasses, the assigned names may not correctly reflect the chemical composition of the rocks. 19

MESOZOIC ROCKS

Information concerning the Mesozoic rocks comes mainly from

Phillip's deep geothermal tests, and strat. test 3 (plate 1).

Mesozoic rocks are not exposed in the Hot Springs Mountains, but this description is included because the geothermal reservoir occurs in presumably Mesozoic metavolcanics. Mesozoic rocks are exposed in ranges to the north and northeast. To the northeast rocks exposed in the Mopung

Hills and West Humboldt Range are pelitic sediments with minor carbonate

and arenite intercalations considered by Johnson (1977) to belong to the

Auld Lang Syne Group. Metavolcanic rocks are exposed to the north in

the Trinity Range.

Mesozoic rocks penetrated by the geothermal wells include altered

basalts, quartzites, marble, phyllite and hornfels, and rhyolitic(?)

tuffs (figure 4). Altered basalts, or greenstones, predominate in the

two successful wells, B21-1 and B21-2, whereas well 29-1 and strat. test

3 intersected mostly metasedimentary rocks. The reason for the variance

is not presently understood.

Greenstones in B21-1 comprise a thick sequence of flows with only

minor intercalations of metasedimentary rocks. The greenstones are char­

acteristically dull, dark green, aphanitic rocks which have been altered

to calcite, chlorite and other secondary minerals. Relict porphyntic

texture can be found, but generally the rocks appear to have been aphanitic

and microcrystalline.

Metasedimentary rocks in well 29-1 consist of quartzite, hornfels

or phyllite, and marble. Greenstones are present, but constitute only

30% to 40% of the section. Quartzite beds are generally light gray to 20 gray-brown, dense and fine-grained. Only rarely is a trace of clastic texture visible. A few quartzite beds are pyritic and occasionally spotted with chlorite flakes. Hornfelsic rocks are gray to black, amorphous, and may be silicified siltstones. Cuttings of these rocks occasionally show a foliation and sheen characteristic of phyllite. The marbles are generally light colored, ranging from milky white to light brown, and are usually finely crystalline although coarsely crystalline

intervals were seen. A section of intermediate to acid tuffs occurs in

29-1 which are light green, highly silicified, and very fine-grained. A

few intervals may have relict phenocrystic texture, with remnant quartz

and feldspar phenocrysts. As with the greenstones, minerals in the tuffs

have been altered to secondary minerals.

Fractures in the greenstones are generally free of calcite or

silica coatings. Rarely, older(?) fractures cemented with silica can

be inferred from cuttings samples, but this is tenuous. In general, to

judge from the drilling rate and electric logs, fractures appear to be

confined to fault zones, and are not an overall pervasive feature of the

metavolcanic rocks. Well 29-1 intersected the thickest Mesozoic section, 2755 feet

(840 m) of mostly metasedimentary rocks with metavolcanic intervals

(figure 4). The well bottomed in a^hornblendite intrusive which may be

genetically related to the Humboldt Lopolith (Speed, 1976). It is possible

that the greenstones are comagmatic extrusives, but evidence is inconclu­

sive Age of the sequence is assumed to be Mesozoic, and possibly Juro-

riassic, on the basis of similarity between well cuttings and exposures

Hills, and Trinity Range. Some of the in the West Humboldt Range, Mopung 21-2 29-1 21-1

basalt flows basalt flows homfels & phyllite quartzite'' basalt v quartzite & \ basalt interbeds hornfels basalt flows basalt flows scale homfels rhyolitic tuffs tuffs? 500ft. 0 / marble & tuff known fractures basalt

marble — SEA LEVEL meta-sandstone basalt flows

r) interbedded FIGURE A. Stratigraphy of Mesozoic rocks basalt & homfels from Phillips Pet. Co. deep geothermal wells

quartzite beds

fa.

Lang Syne Group, but incomplete understanding of Mesozoic stratigraphy in the region, and particularly in the wells, precludes any attempt at correlation at present. The hornblendite gave discordant K-Ar dates from plagioclase and hornblende; plagioclase at 9 m.y. and hornblende at

25 m.y. The significance of these dates is unknown, since high tempera­ tures at depth may have adjusted the dated minerals to younger ages. 23

Tr - TERTIARY RHYOLITES

The oldest Tertiary rocks in the mapped area consist of silicic volcanic units ranging from lithic tuff-breccias to crystal-vitric tuffs.

Rock composition varies from andesite to rhyolite, but rhyodacite and dacite predominate (figure 3). Surface exposures are poor and appear discontinuously from Rhyolite Ridge (informal name, plate 1) northeast­ erly to the vicinity of Cinnabar Hill. Correlation of individual units from outcrop to outcrop or with occurrences in the wells is difficult, and only a few units can be correlated with confidence from well to well.

Areal distribution of the rhyolites is unknown, although in the wells the sequence is thinning southward. Thickness in the northernmost well B21-2 is 2130 feet (655 m) and in well 29-1 the sequence slims to 1250 feet

(380 m).

The sequence is roughly divisible into three parts: a lower rhyolite to rhyodacite series, a medial dacite series with minor andesite, and an upper rhyolite to rhyodacite series. The rhyolites range from white to lavender and are commonly vitrophyric. Most units have quartz, k-spar, and biotite when phenocrysts are visible. Rhyodacites are generally gray, gray-green, and varying shades of pink to purple. Quartz and biotite are almost universally present, and k-spar and plagioclase appear in varying amounts. Lithic fragments are visible in thin section but difficult to identify in cuttings. Lithic-rich rocks are most prevalent in the upper portion of the sequence. The dacite and andesite

sequence is characterized by flows with distinct plagioclase phenocrysts

in a generally brown, gray, and rust-colored aphanitic matrix. Where

fresh, plagioclase is clear and euhedral, but it weathers to a character 24 istic chalky-white with indistinct phenocryst borders. Biotite is the prevalent mafic mineral, but pyroxene was observed. The upper sequence of rhyolites and rhyodacites consists of several units of gray to violet quartz-biotite-k-spar-plagioclase crystal-lithic tuffs and white to green-white, pink and brown k-spar-biotite-quartz vitric crystal tuffs.

Quartz is less frequent than in lower rocks, and biotite occurs as obvious, large euhedral books. Lithic fragments, usually of more mafic rocks, but also as finer-grained equivalents of the parent rock, occur with varying frequency in cuttings.

Two units at or near the base of the sequence are correlative in all three geothermal deep tests: a white quartz vitrophyre and a lavender or violet rhyodacite. The white vitrophyre is glassy, and quartz-crystal- lined vugs are visible in cuttings. In thin section small anhedral quartz and rare k-spar phenocrysts are visible. The matrix is fluidally- banded with an appearance best described as graphic. The lavender rhyodacite is a vitric crystal tuff which characteristically has small quartz and sparse biotite phenocrysts. Subhedral plagioclase and k-spar are common.

Individual unit thicknesses vary from 15 feet (4.5 m) to as much as 200 feet (61 m). In B21-2 the dacite series has the same appearance for 580 feet (176 m), but there are probably several flows of similar composition which cannot be distinguished in cuttings.

Examination of well cuttings reveals the presence of small gravel interbeds at the top of the rhyolite sequence. The gravels consist of subangular to subround fragments of rhyolitic rocks and frosted quartz grains. Fragment size ranges from very fine to at least pebble size, judging from curvatures seen on some cuttings fragments. The gravels are 25 considered to be evidence for some amount of relief and erosion, and therefore, unconformity between the rhyolites and overlying Chloropagus

Formation.

No attempt was made to correlate these rhyolites with other ash-flow tuffs in the region. Some parts of the rhyolite sequence may be correla­ tive with units in the Old Gregory Formation (Axelrod, 1956), rocks described by McJannet (1957) or the ash-flow tuff subdivision of the now- abandoned Hartford Hill Rhyolite (Bingler, 1978).

The rhyolite sequence underlies the Mio-Pliocene Chloropagus

Formation and unconformably rests on Mesozoic rocks. The rhyolites are therefore assumed to be Miocene, but lower units may be as old as

Oligocene. Since regional correlation is tenuous, radiometric ages of other possibly comparable formations are not cited. gray-whlt* qtz-blot-plag rhyodaelt*

B2I-I pink qtz-biot-k »par oryotol tuff

qtz vltrophyr*

andiolt* a doelt* flow*

rod-brown to gray blot-plag-ktpor with minor qtz rhyodaoito

qray blot-ktpor-plog rhyodaoito

brown matrix blot-kopor rhyollt* whit* & pink qtz-k opor-blot-plag rhyodaoito

gray a brown daolto

& t> ~k t> - <-b » -

whit* aphanlttc qtz vltrophyr* ■4

^ 7 M lavtndtr k spar-plag-qtx rhyodflcltt

hj CTn 27

Tb - CHLOROPAGUS FORMATION

A series of andesitic to basaltic flows, agglomerates, and tuff breccias, with minor intercalations of water-lain tuffs and shales was

named the Chloropagus Formation (Axelrod, 1956, p. 95) for exposures at

Green Hill in the northern Hot Springs Mountains. Outcrops of this for­

mation are excellent in the western foothills of the mountains, and also

in isolated areas southwest of Cinnabar Hill.

The Chloropagus Formation consists chiefly of dark brown, black and

reddish vesicular basalt and andesite flows. Water-lain tuffs, tuff

breccias, tuffaceous sandstone and siliceous shale form prominent inter­

beds of extremely local character in the vicinity of Green Hill

(figure 6). Most of the basalt and andesite flows are highly vesicular, locally

scoriaceous, and commonly amygdaloidal with quartz and calcite fillings.

Flow thicknesses vary from five feet (1.5 m) to 25 feet (8m).

Olivine basalt, the most common rock in the formation, is typically

dark brown or red-brown on weathered surfaces and dark grey to black on

fresh fractures. In thin section these rocks are seen to contain

plagioclase phenocrysts (labradorite?), olivine commonly altered to

iddingsite, and pyroxene in a finely crystalline groundmass of feldspar

and pyroxene. Some rocks have a vesicular matrix of black glass.

The andesites usually weather to a medium or dark brown. Fresh

surfaces are gray to brown-gray. Basaltic pyroxene andesites are most

common, although hornblende andesites were seen. In thin section, the

basaltic andesite is seen to contain plagioclase and pyroxene phenocrysts, 28 commonly in cumulophyric arrangement (clusters of phenocrysts), within a felty groundmass of plagioclase (oliogoclase?) microlites and dark red-brown glass. K-spar phenocrysts occur rarely in the hornblende andesites.

Mud-flow breccias and agglomerates occur locally throughout the formation. These are vari-colored green and gray, consisting of pebble to boulder-sized clasts of andesite and basalt in a tuffaceous silt to sand-sized clastic matrix. According to Axelrod (1956, p. 95), the agglomerates commonly grade laterally to finer-grained, better-sorted lapilli tuffs.

The intercalated sediments almost always include a tuffaceous com­ ponent in make-up. The dominant lithology is tuffaceous siltstone or shale, with thin interbeds of tuffaceous sandstone. The sediments grade laterally both north and south to water-lain tuffs and tuff breccias of andesitic composition. The shales are locally silicified, and from one such shale Axelrod (1956, p. 97) recovered the Chloropagus flora. A more complete description of both the sedimentary units and the flora

can be found in Axelrod (1956).

Formation thickness varies from approximately 400 feet (122 m) to

a maximum of 2600 feet (792 m) in geothermal well 29-1. Axelrod

(1956, p. 95) managed to piece together 1920 feet (585 m) from surface

exposures which, as with other formations in the area, are highly

faulted. Also, the preponderance of volcanic rocks of similar character

make stratigraphic positioning difficult, since the intercalated lake

beds cannot be trusted to occur with any lateral continuity. Several

thin sedimentary intervals were encountered in strat. test 1 and the

deep tests, but cannot be related to surface exposures with any certainty 29

Stratigraphic relationships in the three deep tests and the strat. tests indicate that the Chloropagus Formation is thickening southward, which may indicate a source in that direction. Interestingly, Axelrod

(1956) implied a volcanic source to the south on the basis of current

directions in the intercalated sediments. However, the rapid change in thickness of 400 to 2600 feet (122 to 792 m) over a distance of slightly more than 2 miles (3 km) has undoubtedly been enhanced by faulting and erosion.

The base of the Chloropagus Formation is not exposed in the Hot

Springs Mountains. At the north end of Rhyolite Ridge (informal name, plate 1) the contact between the unnamed rhyolite unit (Tr) and

Chloropagus is veneered by alluvium, but rock attitudes imply a con­ formable relationship. Elsewhere, varying rock types near the inferred contact suggest that some amount of unconformity exists between the

Chloropagus and the underlying rhyolites. The extreme thickness varia­ tions in the wells appear to substantiate an unconformable relationship.

According to Axelrod (1956, p. 115) floristic and climatic compar­ isons with nearby flora indicate that the Chloropagus Formation is Early

Clarendonian, or Mio-Pliocene in age. Lower portions of the formation may be as old as Barstovian, or late middle to late Miocene. This is supported by a radiometric date for an andesitic tuff in the upper part of the type Chloropagus section, which yielded a late Miocene K-Ar age of 13.9 million years (Evernden and James, 1964, p. 70). Additionally,

a whole-rock K-Ar date of 14.5 million years (early late Miocene) was obtained from a basal basalt of the Chloropagus Formation in the Pah

Rah Range (Bonham, 1969, p. 30). 30

The stratigraphic section described by Axelrod (1956, p. 95) and the sections intersected by the geothermal wells and strat. tests differ greatly, and in fact are noteworthy for their diversity. In general, the formation has a greater proportion of truly volcanic rocks in the southern part of the mapped area and in the region of the geothermal field, whereas volcanic sediments (water-lain tuffs, breccias, tuffa- ceous shales and sandstones) show a proportionate increase to the north

(figure 6). In the geothermal anomaly wells B21-1 and B21-2 cut a

Chloropagus section comprised almost entirely of volcanic rocks. The interbedded tuffs and other sediments reappear to the north and south.

This strongly suggests a pre-existent high in the region of the geo­ thermal anomaly which affected depositional patterns during Chloropagus

time. 31

Brown, black and reddish amygdaloidal basalt

Local interbeds of sand, lime, and tuff Brown, red-brown and greenish basalt and andesite flows

Local thin sedimentary intercalations

Dark brown, black and red vesicular and fine-grained basalt and andesite flows

Fine-grained tuff, tuff-breccia, and siliceous shale

Black and brown vesicular and amygdaloidal basalt and andesite flows

Siliceous shale, silty tuff and tuff-breccia

Green, yellow-green and gray andesitic tuff and tuff-breccia, with massive gray-green andesitic agglomerates, occasional tuffaceous sandstone and silty tuff

Basaltic agglomerate and andesitic tuff-breccia

Black, brown, red and greenish basalt flows with minor andesite flows

300 feet

FIGURE 6

generalized stratigraphy of the Chloropagus Formation 32

Tdp - DESERT PEAK FORMATION

The Desert Peak Formation was named by Axelrod (1956, p. 97) for exposures in the central part of the northern Hot Springs Mountains.

It crops out north and east of Desert Peak, and on the northwest flank of the Hot Springs Mountains, where it overlies the Chloropagus Forma­ tion and is in turn overlain by the Truckee Formation. In this report the formation is divided into two members (figure 7).

The lower member of the Desert Peak Formation consists of inter- bedded siliceous shale, basaltic tuff, and thin basalt flows. A dense olive-gray basalt is the basal unit in the foothills west of Desert

Peak, but evidently pinches out in other directions as it is not seen elsewhere. The basalt is overlain by siliceous shales which form the base of the unit in the other parts of the area. The shales are usually yellow and yellow-brown, appear porcelaneous, are invariably silicified, and vary from .5 to 25 feet (.15 to 7.5 m) in thickness. Individual beds are 5 to 10 inches thick (12 to 20 cm). Regularly interbedded with the shales are fine-grained basaltic tuffs and thin basalt flows. The tuffs are characteristically mustard or green-yellow, 1 to 10 feet

(.3 to 3 m) thick, and commonly weather to a pebbly texture. The basalts are olive, green-gray, densely aphanitic, highly weathered flows between

5 and 15 feet (1.5 and 4.5 m) in thickness. In thin section, the basalts are typically microporphyritic with cumulophyric clusters of plagioclase and pyroxene(?) altered to calcite in an intergranular matrix of plagio­ clase microlites, small crystals of pyroxene(?), and cryptocrystalline material. The lower member of the formation was apparently deposited only locally, because it has thinned out and is not present in outcrops 33 on the margins of the northern Hot Springs Mountains. Thickness there­ fore varies from 0 to approximately 300 feet (91 m).

The upper member consists of thinly bedded diatomite, thin silicified shales, and minor amounts of basaltic tuff. The upper member can be differentiated from the lower member by its consistently thin bedding, white to cream coloration, and a distinctive weathering style referred to by Axelrod (1936, p. 98) as "poker-chip weathering." Beds are commonly % to 2 inches (1 to 5 cm) thick. The diatomite is perva­ sively silicified and exhibits a pearly or resinous luster on fresh sur­ faces. The shales are fissile, with \ to 1 inch (% to 2.5 cm) partings, and are usually diatomaceous. Minor basaltic tuffs associated with the sediments are 1 to 3 feet (.3 to 1 m) thick, although Axelrod noted two thicker accumulations northwest of Desert Peak. The upper member of the formation is about 600 feet (183 m) thick.

The Desert Peak Formation lies conformably on the Chloropagus

Formation, although the lower member was deposited in more restricted waters than was the upper member. The lower member pinches out to the north, south and east within the mapped area, where the upper member laps onto the Chloropagus. The upper member grades upward into the

Truckee Formation.

Fossils have not been recovered from the Desert Peak Formation, but its age is assumed to be early Pliocene on the basis of its stratigraphic position between the Mio-Pliocene Chloropagus and middle Pliocene

Truckee Formation. 34

Upper member: tuffaceous siltstone, diatomaceous siltstone, opalized diatomite, basalt tuff

basaltic tuff

opalized diatomite and thin silicified shale

silicified shale, thin opalized diatomite

Lower member: massive silicified shale, thin interbedded basalt flows basaltic tuff

siliceous shale, basaltic tuff

olive-gray dense aphanitic basalt

FIGURE 7. Generalized Stratigraphy Desert Peak Formation

200ft. 35 Jtf - TRUCKEE FORMATION

The northern Hot Springs Mountains are the type locality of the

Truckee Formation, which was described in the report on the Geology of

the Fortieth Parallel (King, 1878, vol. 1, p. 415). The Truckee

Formation is composed of fluvial and lacustrine sediments and associated

volcanic rocks (figure 8). It is well exposed on the northern and

western margins of the Hot Springs Mountains, but is absent in the inte­

rior portions of the mountains due to erosion. The formation is also

preserved beneath younger volcanic rocks in the southeastern part of the

northern Hot Springs Mountains, where all members of the formation appear

to be thinning out. The formation is divisible into three members, as

described below.

The lower member of the Truckee Formation consists of yellow-brown

palagonite, basaltic lapilli tuff, gray mollusc-rich sandy coquina, gray-

white limestone, basaltic pebble breccia, water-lain ash, tuff breccias,

tuffaceous sandstone, diatomaceous shale, and basalt. The basalt is a

black, dense, fine-grained to glassy rock. Fine-grained phases have

plagioclase microlites and common phenocrysts of olivine altered to

iddingsite. Glassy phases generally occur at the tops and bottoms of

flows. Total thickness of the lower member is 95 to 100 feet

(29 to 31 m).

The middle member is predominantly thin-bedded to massive white

diatomite interbedded with minor amounts of light-gray vitric tuff and

thin discontinuous beds of gray sandstone. Measured thickness of the middle member is about 850 feet (259 m). 36

Small outcrops of the upper member were mapped, but the best expo­ sures occur northwest of the mapped area, in sections 17, 19, and 20,

T23N-R27E. The upper member consists mostly of slabby, gray-brown lime­ stone with interbedded gray sandstone, very thin diatomite beds, and basalt pebble conglomerate. The limestone is commonly sandy and usually exhibits good bedding varying from 4 to 3 inches (1.5 to 2 cm) to

10 to 20 feet (3 to 6 m) in thickness. Thickness of the upper member is approximately 850 feet (259 m) (Axelrod, 1956, p. 102).

The top of the Truckee Formation was not seen. The lower contact is gradational into upper Desert Peak siltstones and tuffs within a ver­ tical 20 feet (6 m). According to Axelrod (1956, p. 101) the lower mem­ ber of the Truckee Formation pinches out to the northwest. All three members appear to be thinning to the southeast, and it seems reasonable to assume that the basin in which the "type" Truckee Formation accumulated was of limited extent, and may have been only 10 to 15 miles (16 to 25 km) in diameter.

Age of the Truckee Formation is almost certainly Pliocene. The lower member has yielded fossils considered by Yen (1950) as Pliocene.

Fossils found in the middle member outside the mapped area confirm Yen's conclusion. Axelrod (1950, 1956) considers the lower member to be Late

Clarendonian (early Pliocene), and that the middle member to be as young as

Hemphillian (.middle Pliocene). Fossils have not yet been recovered from the upper memeber, but the younger portions of the upper member.may be as young as Blancan (late Pliocene), (Axelrod, 1956, p. 102). 37

Upper member; gray, gray-brown, and tan sandy limestone, paper-thin diatomite, basaltic pebble conglomerate

Middle member; thin bedded to massive soft white diatomite, light gray vitric ash, thin gray discontinuous sandstone

Lower member: basaltic pebble breccia, grav-white limestone & coquina, water-lain ash, tuffaceous silt, sandstone, diatomaceous shale, basalt

Scale FIGURE 8. Generalized stratigraphy of the 0 300ft. Truckee Formation 38

Tyv - YOUNGER VOLCANICS

The younger volcanic rocks are predominantly olivine basalts. The basalts are dark gray to black rocks which usually weather to a mottled

dark brown and black. The basalts are porphyritic, locally vesicular, and usually scoriaceous near flow tops. In thin section the rocks typi­ cally contain phenocrysts of subhedral plagioclase and olivine.- The groundmass contains plagioclase microlites and minor granular olivine in a cryptocrystalline matrix of deuteric minerals(?) and glass.

Individual flows vary from about 15 feet to 50 feet (4.5 to 15 m).

Judging from exposures in the southern Hot Springs Mountains, total thickness may be several hundred feet (several tens of meters). Due to faulting and erosion, maximum thickness in the mapped area is about

100 feet (30 m).

Rocks of this unit occur in the western part of the northern Hot

Springs Mountains. Erosional remnants of the younger volcanics often cap ridges and small mesas thereby preserving the easily eroded sedi­ ments underneath. The basalts are not seen east of the Telephone Road

(informal name, plate 1).

The younger volcanics unconformably overlie the Truckee

Formation and the Desert Peak Formation in the western portion of the northern Hot Springs Mountains. Since the lower contact is unconform- able with the middle Pliocene Truckee Formation, age of the basalt sequence is presumably late middle to late Pliocene. An upper age limit is harder to ascertain. Within the mapped area the basalt sequence had been extensively faulted and eroded prior to the deposition of Pleisto­ cene Lake Lahontan sediments. Further, field relationships imply that the basalts are older than the andesitic ash flow sequence described next, tentatively establishing the age of the basalt sequence as wholly

Pliocene, and probably late middle to early late Pliocene (early

Blancan?). 40

Ta - TERTIARY ANDESTTF

A sequence of andesitic tuffs, flow breccias, and flows occurs

east of the Telephone Road (informal name, plate 1) in the northern Hot

Springs Mountains. The little-studied sequence consists primarily of

hornblende-plagioclase ash-flow tuffs. The rocks are light brown to

purplish-gray on fresh surfaces, and weather in a platy fashion to a

red-brown or lavender hue. In thin section the rocks are seen to con­

tain small, subparallel, broken plagioclase laths, subhedral hornblende

and minor potassium feldspar phenocrysts in a matrix of ash, pumice,

glass and cryptocrystalline material. Eutaxitic texture and fluidal

banding are common. Cumulophyric clusters of plagioclase, potassium

feldspar and pyroxene occur sparingly. Chemical analysis may show these

rocks to be more properly classified as trachy-andesites. Lithic frag­

ments, usually of more mafic volcanic rock types or as finer-grained

equivalents of the parent rock, are common in the basal units but

decrease upward in abundance. Central portions of individual tuff units

are generally vitrophyric, grading upward and downward into less welded

parts, with concomitant increase of ash content. The andesite unit lies

with angular unconformity on sediments interpreted as Truckee Formation,

Desert Peak Formation, and Chloropagus Formation. Folding and faulting

preceded, and in part controlled deposition of the andesite unit. In

general, the thickest accumulations occur in the troughs of synclines

and alongside scarps created by downdropped blocks prior to extrusion

of the andesite sequence. Extent of erosion is unknown, but the unwelded upper portion of at least one cooling unit has been removed, exposing

the dense, vitrophyric and resistant welded center, creating a stripped 41 structural surface (Bingler, pers. commun., 1977). Original thickness of the unit is therefore unknown. Geologists of the Southern Pacific

Co. measured about 500 feet (152 m) of andesite east of the mapped area, but thicknesses measured by me range from 50 to 100 feet (15 to 30 m).

Three discordant radiometric K-Ar age determinations have been obtained from mineral separates from the andesite sequence. Geochron

Laboratories dated plagioclase at 11.2 m.y. The U.S. Geological Survey dated hornblende at 4.3 m.y., and plagioclase at 2.3 m.y. According to

Voegtli (pers. commun. ,• 1978) the radiometric age determinations by the

U.S. Geological Survey are more reliable. The discordance of the radio- metric determinations raises a potentially severe question: if the sediments beneath the andesite unit are correctly correlated with the

Truckee, Desert Peak, and Chloropagus Formations, and if the 11.2 m.y. date is correct, then the age index for Nevada's Tertiary non-marine fossil record needs revision. However, a 2 to 4 m.y. age for the an­ desite sequence is compatible with the fossil evidence. I believe the latter interpretation is correct. 42

Tdi - DESERT QUEEN INTRUSIVF

A small pluton of hornblende-quartz diorite crops out in the vicin­

ity of the Desert Queen Mine (plate 1). In hand specimen the rock is a

dull gray-green on both fresh and weathered surfaces. It is fine­

grained, holocrystalline, and equigranular. Altered sodic plagioclase

(60-70%) partly altered hornblende (15-20%), quartz (5-10%), and altered

potash(?) feldspar (0-5%) are the major minerals seen in thin section.

Iron oxides and other opaque minerals comprise about 5% of the rock.

The diorite is mildly propylitized. Alteration consists of sericite

and epidote after feldspar, and chlorite after hornblende.

The intrusive mass is bounded on the east by a normal fault. The

north, west, and south contacts are inferred to be intrusive into

slightly older or nearly coeval volcanic rocks of the unnamed rhyolite

unit (Tr), and the Chloropagus Formation (Tb). The diorite outcrops

over approximately one third of a square mile (0.8 km).

Willden and Speed (1974, plate 1) assigned a Jurassic age to the

diorite. However, field relationships suggest that the intrusive may

be as young as Tertiary (Miocene?). Volcanic tuffs and flows along the

north margin have been intensely altered by the intrusion of the diorite,

both thermally and dynamically. The volcanic rocks bordering the intru­

sion are lithologically more similar to Tertiary volcanic units than to

any of the Mesozoic rocks exposed in the surrounding region, and may be metamorphosed equivalents of rhyolitic tuffs in the unnamed rhyolite unit (Tr). Andesites of the lower Chloropagus Formation on the west margin of the intrusive have been propylitized, as have dacitic and rhyodacitic rocks of the unnamed rhyolite unit (Tr) on the south margin. 43 The only area where propylitic alteration of the Chloropagus Formation

and the rhyolite unit was observed is in the vicinity of the diorite

intrusive. Presumably the alteration was caused by emplacement of the pluton.

Other Tertiary intrusives, some of quartz diorite composition, have

been mapped m neighboring areas by Johnson (1977), Bonham (1969),

Gianella (1936), and Thompson (1956). In the Olinghouse District about

20 miles (32 km) to the west, Bonham (1969, p. 31) describes composi-

tionally similar rocks which intrude the Hartford Hill Rhyolite and over-

lying Chloropagus Formation. The Davidson Granodiorite near Virginia

City intrudes the Miocene Alta Formation (Thompson, 1956), which is at

least partially a temporal equivalent of the Chloropagus Formation

(Bonham, 1969, p. 32, figure 11).

No radiometric ages from these Tertiary intrusions have been

obtained, but the evidence cited above suggests a late Miocene or earliest

Pliocene age of intrusion. Field relations in the mapped area, although speculative, imply that the Desert Queen diorite intruded the unnamed

rhyolite unit and the lower part of the Mio-Pliocene Chloropagus Forma­ tion. Therefore, it seems reasonable to assume a late Miocene age for emplacement of the Desert Queen quartz diorite pluton (informal name, plate 1). 44 Qal ■Z-JjjjATERNARY n rp n ^ y js

Quaternary deposits is a ge„eral map unit which includes Lake

Lahontan sediments, alluvial fan deposits, pedim8nt ^ ^

deposits, wind-blown sanrl^ „nj sands, and minor amounts of siliceous sinter aroung Brady's Hot Springs.

^ tahontan sediments occur around the margins of the Hot Springs

Hountarns, and comprise subhorirontal thin-bedded, fine-grained silts

and clays. Small wedge-shaped gravel deposits are intercalated locally with finer grained sediments qnif k ’ Spit and bar deposits of locally derived alluvial material define former beachlines.

Alluvial fan deposits cover much of the low lands in the map area.

Fan deposits consist of poorly sorted and poorly bedded silts, sands,

and gravels. The fan deposits overlie Lake tahontan deposits, and in

places have been terraced by Lake Lahontan waters, indicating that the

fan deposits are both coeval and younger than Lake Lahontan sediments.

Gullies cut into Windmill Flat and Ashflow Flat (informal names,

plate 1) reveal that Tertiary sediments are obscured by a thin veneer of

pediment gravels. The gravels are poorly sorted, angular to subround

accumulations derived from nearby alluvial fan deposits.

Several small playas are present in the mapped area. The piaya

iments consist of clay, silt, and minor sand mixed with crystalline

salt deposited as matrix and as thin crusts resulting from evaporation of seasonal waters.

Eolian sands mantle large portions of the mapped area, particularly

along the Telephone Road and in Windmill Flat (informal names, plate 1).

The sand is composed of coarse to very fine grains of quartz, feldspar, magnetite, chert, volcanic rock fragments, and other minor constituents. V. tj *

^ BRADYS THERMAL FAULT BRADY'S THERMAL ANOMALY__

SHIRLEY FAULT-

DESERT PEAK FAULT _

''ES 29-1 i THERMAL ANOMALY- ■ -v*w ■*!?'' & ' 'm m j m ov SPRinrs ...... c

The northern Hot Springe Mountains ere cheraoterireH by „ north_

northeasterly trending etructurel pattern. The formations strike

north-northeasterly, and the folds and faults generally are parallel to

them. The dominant north-northeast structural trend is disturbed in the

south-central portion of the mapped area by a northeast trending rone of

disruption where the rocks have been structurally elevated (figure , and

Plate 1, this report). The older rocks exposed in the central regions

are flanked by increasingly younger formations exposed to the east and

»est. Small-scale folding in Tertiary rocks reflects basement block- faulting.

Several small folds were mapped. The folds involve rocks as young

as Tyv (younger volcanic unit, plate 1). Fold axes trend north-north-

east, roughly parallel to the dominant fault trend in the area, and

usually plunge gently to the south 10° to 20°. Most of the folding is

gentle but asymmetric. Dips of fold limbs are generally 20° to 30°,

although locally dips as high as 63° were measured. Folds are best

preserved m surface exposures of lacustrine sediments of the Truckee and Desert Peak Formations, but involvement of thick unruptured volcanic sequences of the Chloropagus Formation in the cores of the folds implies a deeper-seated origin. Folds mapped in Tertiary rocks near Pyramid

Lake appear to be the result of differential block movement along base­ ment faults (E.R. Larson, pers. commun., 1978). This explanation, termed drape folding by Stearns (1971), requires basement faults which do not appear at the surface, but are proximal to the folds. Rock relationships in the mapped area and in the geothermal wells suggest 47

r

FIGURE 10. Depicting basement block faulting, drape fold formation, and fold disruption by further faulting along same fault traces 48

that numerous faults are present in the deeper Tertiary volcanics (Tr,

for instance) which did not dislocate the overlying sediments. Drape

folding best fits the available evidence, although local compressive

stress associated with normal faulting is an alternate explanation

(Bonham, 1961). Unfortunately, delineation of the entire fold pattern

is severely hampered by discontinuous exposures of folded rocks, and

later intense faulting disrupted and undoubtedly obscured or obliterated

many earlier—formed folds. It is likely that several fold—disruptive

faults now seen at the surface were initially basement faults which

contributed to fold origination. Initial movement occurred in basement

rocks and was absorbed in the overlying sediments by drape folding.

Later movement on the faults was sufficiently large to break the sedi­

ments, too (figure 10). Folding is late Pliocene to early Pleistocene

in age since no rocks younger than Tyv (younger volcanic unit) are

involved.

Two distinct fault sets were observed in the northern Hot Springs

Mountains. The dominant fault pattern trends about N25E and appears

related to Basin and Range tectonic stresses. A subordinate, less well-

defined fault pattern trends about N55E to N70E. Other faults were

recorded of variable orientation which apparently represent local adjust­ ment to stress. A few of these faults trend N10W to N40W, subparallel

to the Walker Lane, and represent the only stress relief with possible

genetic ties to the Walker Lane shear.

Basin and Range north-northeast normal faulting is well-represented in the northern Hot Springs Mountains. Faults trending approximately

N25E are pervasive throughout most of the mapped area, decreasing in

abundance only where alluvial cover conceals the older rocks. The faults 49 are all high-angle (greater than 50°). One fault mapped by Axelrod

(1956) is correctly identified as a high-angle reverse fault (^ mile,

0.8 km, west of Green Hill, plate 1), but all the other faults of this category are normal faults. Fault displacements are difficult to ascer­ tain, owing to a lack of readily identifiable marker beds and discon­ tinuous outcrops, but apparently range from 50 to 100 feet (15 to 30 m) to at least 1000 feet (300 m).

Several major N25E trending faults with significance to geothermal heat localization were mapped. The most obvious is Brady's Thermal

Fault which traverses sections 1, 12, and 13, T22N-R26E. The trace of this fault is revealed by hot springs, fumaroles, sinter deposition, and discontinuous scarp features. Younger volcanic rocks (Tyv) and lower Truckee Formation (Ttl) on the west side of the fault have been dropped down against lowermost Truckee Formation (Ttlm). Total dip- slip on the fault is about 500 feet (150 m).

Another major N25E trending fault bounds the west side of Brady's

Hill (informal name, plate 1). It cuts across sections 5, 6, and 7,

T22N-R27E, and is inferred to continue southward into section 13,

T22N-R26E. Middle Truckee diatomite beds on the west are displaced

downward against upper Chloropagus rocks. Total dip-slip on this fault

isnat least 1000 feet (300 m).

The Shirley fault (figure ?), traceable for at least 6 miles

(10 km), is exposed in the southeastern corner of section 18, T22N-R27E,

and continues north-northeast into section 23, T23N-R27E. At the south

ern exposure, middle Truckee diatomites are dropped down against

Chloropagus Formation on the east. In the north, Desert Peak Formation

on the east is dropped down against Chloropagus basalts on the west. 50

The central portion of the fault has had only minor movement, and oppo­ site senses of motion at either end of the fault provide compelling evidence for a scissors-type movement. Total dip-slip movement on the distal portions of the fault may be as much as 500 feet (150 m).

Approximately \ mile (0.4 km) west of Desert Peak, extending

N30E-S30W, lies the apparently high-angle Desert Peak Fault (figure 9) which drops lower Desert Peak shales on the west down against upper

Chloropagus volcanics on the east, a displacement of about 1000 feet

(300 m), based on strat. test 4. This fault is significant in that it may be part of the boundary system for the Desert Peak Geothermal

Field. Age of this fault is uncertain, but it postdates the younger

volcanic unit (Tyv).

The Desert Queen Fault (figure 9) trends about N25E, dips 50° east,

and drops andesite ash-flows (Ta) on the east down against lower

Truckee limestone on the west. Total displacement at the southern

portion of the fault is approximately 100 feet (30 m), but displacement

increases northward to the vicinity of the Desert Queen Mine, where

andesite ash-flows (Ta) on the east are dropped down against the

unnamed rhyolite (Tr). Erosion prior to faulting and faulting prior

to deposition of the andesite account for some apparent displacement,

but dip-slip is still at least 500 feet (150 m). Since the Desert Queen

Fault cuts the andesite sequence, its latest movement must be younger

than 4 m.y. Cage of the andesite). Faults in the southeastern portion

of the mapped area have displaced Truckee and older rocks more than the

overlying andesite ash-flows (Ta). The andesite sequence characteris­

tically has been stepped down on the east, usually in increments of

50 to 100 feet (15 to 30 m). The underlying Truckee, Desert Peak, and 51

Chloropagus Formations have been displaced similarly, but dip-slip ranges from 100 to 500 feet (30 to 150 m). The implication is clear that fault­ ing was initiated prior to deposition of the andesite sequence (Ta).

Continued similar stress during and after andesite deposition was accom­ modated along essentially the same fault planes.

The distribution of rocks in the northern Hot Springs Mountains suggests that the area is a highly-faulted and tilted fault-block trend­ ing north-northeast. Older rocks of the unnamed rhyolite unit (Tr) are exposed in the central portions of the region, and increasingly younger rocks are exposed outward, both east and west (see Willden and Speed,

1974, plate 1). Differential movement of blocks has left the area be­ tween the Desert Peak Fault and the Desert Queen Fault structurally elevated in comparison to surrounding areas. Blocks are incrementally stepped down to the west and to the east.

The second fault set occurs as a vaguely defined zone which is first seen in section 29, T22N-R27E, and trends approximately N60E to the vicinity of the Desert Queen Mine. Faults within the zone vary in trend from N50E to N76E. Most fault traces can be followed or inferred for only short distances, usually about 1 mile (1.6 km), and appear to pre­ date the Basin and Range faults, although the apparent termination of a few of the Basin and Range faults by N60E faults implies renewed or concurrent movement on N60E faults. Slickensides observed at the north end of Rhyolite Ridge (plate 1) rake 52°SE, indicating that at least the last motion of the N60E trending faults had an oblique component. Thus, some amount of horizontal shear contributed to the formation of the zone.

Unfortunately, traces of the N60E faults are mostly concealed beneath alluvium, and actual relationships between N25E and N60E fault sets are inferential at best. 52

Surface evidence suggests that the area in the N60E zone is a horst.

Faults at the southeastern edge of the zone have moved down on the south, whereas faults along the northwestern margin of the zone have moved down to the north. Well relationships suggest that central portions of the horst (sections 14, 15, 21, and 22, T22N-R28E) may be structurally ele­ vated over surrounding terrain as much as 2000 feet (610 m), although movement on individual faults is probably only 200 to 500 feet (61 to

150 m). For instance, the base of the Chloropagus Formation in well

29-1 (plate 1) is nearly 1500 feet (450 m) lower than the formation base in well B21-1, and approximately 2700 feet (823 m) lower than in well

B21-2, although some of the apparent uplift is probably taken up by differential erosion, dipping rock sequences, and later movement along

Basin and Range N25E faults. Nevertheless, the N60E fault zone crudely defines a horst system within the zone boundaries.

The area within the N60E trending fault zone was elevated during the latter stages of deposition of the unnamed rhyolite (Tr), as indi­ cated by sporadic occurrences in the wells of gravels in the upper part of the rhyolite section. This is confirmed by a decrease in thickness of the overlying Chloropagus Formation and an uncharacteristic lack of sedimentary intercalations which are prevalent in the Chloropagus elsewhere in the area. This suggests that uplift along N60E trending faults occurred in late Miocene time. A second uplift occurred later, probably in late Pliocene, along N25E trending faults, principally the

Desert Peak Fault and Desert Queen Fault. A late Pliocene age is implied by the unconformable relationship of the andesite unit (Ta) to underlying rocks, which were folded, faulted and eroded prior to ande­ site deposition. 53

The result of these two distinct faulting episodes is the creation of an en echelon, rhombohedral series of horst blocks defined by the

N25E and N60E faults (figure 11). The en echelon horsts occupy the region from section 29, T22N-R27E to Cinnabar Hill, section 6, T22N-R28E.

The east and west boundaries of the horst complex are the Desert Queen

Fault and the Desert Peak Fault. The north and south boundaries are the outlying faults of the N60E fault zone.

The Desert Queen intrusive is the locus of the Desert District, a small gold-silver camp in the northeastern portion of the Hot Springs

Mountains. Epithermal gold-silver mineralization and alteration has affected both the diorite intrusive and intruded volcanics, and there is definite spatial association of veins and diorite. Whether or not there is an association in time between intrusion and mineralization is unknown. The Desert District will be discussed in the section on ore deposits. c. block showing N.25E. fractures showing rhombohedral horst block on horsted block bordered by N.60E. and N.25E. faults

FIGURE II.

block diagram dep cting formation of rhombohedral horst blocks 55

GEOLOGIC HISTORY

The sequence of known or inferred geologic events in the northern

Hot Springs Mountains is as follows:

(1) late Triassic and early Jurassic deposition of the Auld Lang Syne

Group, a sequence of predominantly pelitic rocks, tentatively

including carbonate and arenite strata and basalt flows in the

Hot Springs Mountains area;

(2) regional metamorphism of early Mesozoic rocks to greenschist facies

(pelitic rocks to phyllite and hornfels, carbonates to marble,

arenite to quartzite, and volcanics to greenstones);

(3) middle Jurassic intrusion of the Humbolt gabbroic lopolith and

deposition of comagmatic basalt units. Intrusion of the lopolith

caused local thrusting at distal portions, notably in the West

Humboldt Range, and in places thermal aureoles were superimposed

on regional greenschist fabric;

(4) late Jurassic history unknown; Cretaceous intrusion of granitic

plutons to the west;

(5) late Cretaceous history unknown; probable depositional hiatus late

Cretaceous to early Tertiary; regional unconformity;

(6) inception of volcanism in 01igocene(?); intermediate to mafic flows

north of area; predominantly silicic volcanics in Hot Springs

Mountains area and south;

(7) continued extrusion of silicic volcanic rocks in Miocene time;

Creation of N60E structural elements in late middle Miocene(?) —

horsting of Telephone Road area (informal name, plate 1), some

topographic relief and erosion; 56

(8) late Miocene-early Pliocene inception of intermediate to mafic

volcanism (Chloropagus Formation); inception of N25E trending

extensional faulting(7);

(9) late Miocene and Pliocene basins formed; continued volcanism

with localized lacustrine and fluvial sedimentation (Chloropagus,

Desert Peak, and Truckee Formations) — topographic relief and

erosion;

(10) mafic volcanism (Tyv); continued block faulting — drape folding

prior to 4 m.y. of Tyv and older rocks; erosion of high areas;

(11) eruption and deposition of andesitic ash flow sequence (Ta);

(12) erosion; creation at higher elevations of modern topographic

elements;

(13) Lake Lahontan sedimentation; inception of geothermal systems(?) —

continued N25E faulting — basaltic volcanism at Soda Lake and

Upsal Hogback (Carson Sink); pedimentation, sinter deposition at

Brady's Hot Springs;

(14) Pleistocene to present alluvial deposition; sheet and dune sands,

alluvial fans; inception of hot aquifers — leakage from geothermal

reservoir, hot springs at Brady's. 57

ORE DEPOSITS

Potentially economic deposits of metallic minerals and non-metallic materials occur in the northern Hot Springs Mountains. Gold-silver deposits of the Desert District occur in veins which cut the Desert

Queen quartz diorite. Non-metallic deposits include diatomite, lime­ stone, volcanic ash, and sand and gravel aggregates.

The Desert District, in the northeastern Hot Springs Mountains, has two mines, the Desert Queen and the Fallon Eagle, and numerous prospect pits. The principal mine, the Desert Queen, may have been discovered as early as 1849 (Vanderburg, 1940). The Desert District has a combined production of less than $50,000. Available figures sug­ gest that the most active years were 1883, 1884, 1938, and 1939. Between those years activity was sporadic, and production never exceeded $1,000 per year (Willden and Speed, 1974).

Production was from quartz veins which generally parallel prominent jointing in the diorite host rock. The veins carry sparse, highly oxi­ dized sulfides, and with the possible exception of small high-grade pockets, are usually low-grade (Cunningham, 1957). Vein exposures and samples taken by Willden and Speed (1974) appear to be representative, and show no promise of enrichment with depth, although Willden and Speed state that "material of economic grade might be developed in the district, particularly if the price of silver is favorable."

Quarries developed in middle Truckee diatomites occur along the northwestern margin of the Hot Springs Mountains. Elsewhere, numerous bulldozer prospect-cuts reveal additional amounts of diatomite. The

Aquafil Company produced excellent quality diatomite from deposits in 58

T23N R27E (Bonham, 1961). Substantial quantities of apparently commer­

cial quality remain available.

Both the upper and lower members of the Truckee Formation contain

limestone of good quality (Cunningham, 1957). Similar deposits mined

near Fernley, Nevada, have been utilized in the manufacture of cement.

Conceivably, the Truckee limestones in the northern Hot Spring

Mountains could be used for the same purpose.

Gray vitric ash beds averaging five feet (1.5 m) in thickness, but

often as thick as 20 feet (6 m), crop out in numerous places in the

area, though not as abundantly as diatomite. Similar material from

nearby referred to as pumice by Tischler (1961) has been mined and used

as lighweight aggregate in the manufacture of concrete structural forms.

The gray vitric ash in the mapped area may constitute an economic

resource.

The Hot Springs Mountains area abounds in Quaternary sand and

gravel. Through the years, small quarries were developed for mainten­

ance of old Highway 40, and later expanded during construction of Inter­ state 80 (Bonham, 1961). The long distance to other potential markets precludes all but local usage. Extensive local usage is not imminent. Prior to the exploration program begun by Phillips Petroleum Com­ pany in 1973, geothermal exploration in the Hot Springs Mountains area was concentrated along the Brady's Thermal Fault (plate 1). In the early 1960's, drilling by Magma Power Co., Earth Energy, Inc., and

Union Oil Co. confirmed the existence of economic temperatures, but unfortunately also proved the fault-related thermal system to be too small to support long-term electric power generation. Thermal activity is restricted to Brady's Thermal Fault which dips steeply westward

(Oesterling and Anctil, 1962). Further exploration has been discouraged because of low flow, inadequate recharge, and severe calcium-carbonate scaling problems in the wellbores. Exploration has been concentrated along the fault and surrounding areas ignored, evidently because they lacked obvious surface thermal characteristics.

Phillips Petroleum Company initiated shallow temperature-gradient hole drilling in the Hot Springs Mountains area in 1973 and culminated the exploration program in 1976 by drilling the first of two successful deep tests, well B21-1 (plate 1). This discovery, the Desert Peak Geo­ thermal Field, is revealed by a thermal anomaly that has temperature gradients of 6°F/100 ft. and higher which underlie an area of approxi­ mately 73 square miles (190 sq. km). The anomaly has a notable lack of surface thermal manifestations, and as such is the first totally blind geothermal discovery in Nevada to date (1978). Reservoir waters are classified as sodium-chloride waters, and are contained in fractured

Mesozoic greenstones. The observed reservoir temperature is as high as

406°F (207°C) which substantiates the chemical geothermometers. The 60 geothermal field was discovered and delineated by drilling a total of 54

shallow temperature-gradient holes (300-500 ft; 91-150 m), 8 strat. tests

(1000-2000 ft; 300-610 m), and confirmed by two deep geothermal producing wells (3000-4000 ft; 900-1280 m).

Exploration was hampered initially by the existence of near-surface,

saline thermal aquifers leaking from the geothermal reservoir. Tempera­

ture gradient patterns inferred from shallow holes and geophysical data

falsely portrayed the position of the geothermal reservoir. Deep test

29-1 (plate 1) was drilled in 1974 on the basis of the shallow tempera­

ture data and an electrical resistivity study, and though it did not

penetrate the geothermal reservoir it did reveal the presence of the

shallow thermal aquifer system. Strat. tests were designed to penetrate

below the aquifer system in order to obtain temperature data untainted

by shallow thermal water flow, and sucessfully delineated the true posi­

tion of the geothermal reservoir.

Temperature gradients above the shallow aquifers and the reservoir

are high. In holes influenced by the shallow thermal flow, temperature gradients range as high as 64 F/100 ft. Drillholes above the reservoir and unaffected by shallow thermal flow characteristically vary from

6 F/100 ft. to 15°F/100 ft., although temperature gradients over the most intense part of the anomaly may reach 40°F/100 ft. Holes drilled outside the reservoir which intersect the thermal aquifers commonly show temperature reversals with depth, or become isothermal (figure 12).

At the time, the thermal aquifers were considered a hindrance, since they were the basis for placement of the expensive, commercially unsuccessful deep test (29-1). In retrospect,however, the geothermal discovery would have been substantially delayed without the presence of DEPTH INCREMENTS u ifune b thermal by aquiferbut influenced rfl outside thermalprofileanomaly T EMPERATURE INCREMENTS oml' rfl outside profilenormal ' rud h Dsr Pa thermal Desert Peak anomaly thearound xm ls f eprtr profilesexamples in temperature and of hra anomalythermal T

IU E 12 FIGURE

qie ad insideaquifer thermal and anomaly profile influenced not bythermal nlecd by thermalinfluenced aquifer rfl vr eevi and reservoir overprofile

I

62 abnormal temperature gradients caused by the aquifers (Benoit, pers. commun., 1978). Identification and delineation of thermal aquifers are now part of the exploration technique. The effects of the thermal aquifers are considered analogous to the alteration halo around an ore deposit. As with the mineral alteration, thermal aquifers constitute not only a larger target, but a means for identification of continued exploration direction. Both the temperature and the hydrologic poten­ tial of the thermal aquifer decrease away from the source and allow the geologist to "chase the aquifer upstream" to its source, hopefully a viable geothermal reservoir.

Geology, geophysics, and geochemistry played a very minimal role in the discovery of the Desert Peak Geothermal Field. The commodity souaht was heat, and only methods to detect heat were used. The dis- 63 produced some interesting connotations. The Hot Springs Mountains apparently act as a hydrologic boundary for at least two major ground- water flow systems: (1 ) the Brady's Hot Springs-Fireball Valley-Fernley system, and (2) the Carson Sink-Desert Peak system (Mifflin, 1968).

Hydrogeologic evaluation is complicated by the superposition of higher-

than-normal geothermal gradients in the area and a paucity of data in

areas away from Brady's Hot Springs.

Groundwater near Brady's Hot Springs occurs in alluvial and

lacustrine deposits at shallow to moderate depths but east of Brady's

Thermal Fault groundwater is contained in joints, fissures, along con­

tacts, and in sedimentary intercalations in the Tertiary volcanic

sequence, as well as in the underlying Mesozoic basement rocks.

Generally depth to water increases eastward from Hot Springs Flat,

except near Brady's Thermal Fault where an elongate mound of water is

found at shallow depths. Mounding of the water-table is also apparent

in an area centered above the Desert Peak Geothermal Field. Brady's

Thermal Fault apparently acts as a steeply-dipping conduit in concert

with the gently inclined lacustrine and alluvial aquifers.

The general direction of water movement in the region is suggested

hy the configuration of the water-table (.figure 13). Near Brady's Hot

Springs two flow patterns are indicated. One is the movement of ground-

water southerly from Hot Springs Flat towards Fernley Sink (Harrill,

1970). The other pattern is upward movement of thermal water along

Brady's Thermal Fault as is implied by the water-table mound near the

fault trace (Olmstead and others, 1973). Similar conditions are implied

in the vicinity of the Desert Peak Geothermal Field where the water-

table configuration reveals an anomalous mound above the geothermal II

A\00 K/‘

— -USGS-4 O 4 0 5 8 ' (USGS-5

USGS-3 H 40 6 1 ' BRADY-S -TYBE-W ATER

2 . 1 B y 5 \ ,T‘ H g s -2o , USG 17qT s / V , \ ; :

4 0 5 s \ r Y/USGS-J& AUhilON DEEP/TEST / f “ •• ,G_— )---- 1-----/- 7 p * [0A'JKi/----— H—- -- r-----... V— J.— 4-&--— s sGS-7 / / // • ' f />[r| \ -c/ '1 Is r- J iiiRR.q- I USGS-I47 / :S./°'o'7«"w ell, T s s s - e ' 1’ I USGS

40650 USGS-16 1 r r ' p y k USGS-H > 409li_ . jU S G S 4 0 5 3 3 W k / t i SGS-I20 Q l ° 6! STRAT.

B21-2, A 4160

STRAT. T E S T 8 4 0 3 8 v A 4 0 7 8 zG USGS \ B2I-I , . A 4310 \ • A s TR.

0USX3S-2O \ V : STRAT. TEST 1.^

25 / ‘Vs /° '28 / i „„*** a DESERT PEAK 29-1 / USG SflSV/'Y*' A * - ~ ~ Rhittrpi water well / 7 dv's water well m R R R R q "i

1^74* •• y*"* J

; J : ,; ;::5: J . : "I

65 reservoir. ' As at Brady's Hot Springs this suggests two flow patterns.

One is groundwater movement towards the Carson Sink and the other is upward movement of thermal water leaking from the geothermal reservoir

(figure 13).

Water chemistry data collected and analyzed by W.R. Benoit (1974) indicate that Brady's thermal waters and Carson Sink-Desert Peak thermal waters are chemically distinct and hydrologically separate waters.

Brady's thermal water is a sodium-chloride water characterized by

SO4/CI ratios between .12 and .37 (average = .29) and total dissolved solids contents ranging from about 2500ppm to 3380ppm. Carson Sink-

Desert Peak water is also a sodium-chloride water, but SO^/Cl ratios are usually about .03 and total dissolved solids contents vary from about 5500ppm to approximately 8500ppm. The only lateral variations in water chemistry of both types appear to be related to temperature change. For instance, boron and silica decrease and magnesium increases as temperature drops (Benoit, 1974). Otherwise the chemistry of each water type is remarkably uniform. Carson Sink-Desert Peak waters under­ lie the northwest part of the Carson Sink, the eastern slopes of the

Hot Springs Mountains, and are inferred to underlie portions of the western slopes as well. Brady's thermal waters underlie Hot Springs

Flat and possibly the far western slopes of the Hot Springs Mountains

(figure 13). In the northern Hot Springs Mountains the hydrologic divide and the topographic divide may not coincide. Available evidence suggests an absence of mixing of Carson Sink-Desert Peak water and

Brady's thermal water which implies that the hydrologic divide is sharp and effective. Even considering the abnormally high temperature gradients, thermal water at Brady's must circulate to at least 5000 feet H B m m n n ^ m m m m m m m m m "i

66

(1525 m) in -order to attain observed temperatures, Therefore, the hydro-

logic divide apparently extends downward at least 5000 feet.

The possibility that Brady's thermal waters and Carson Sink-Desert

Peak thermal waters share a common origin is remote for several reasons.

S0^ contents of Brady's thermal water are more than twice the contents

of Desert Peak thermal waters, and Cl contents are less than half, hence

the wide variance of S0^/C1 ratios. Total dissolved solids contents of

thermal waters collected at flow lines during deep geothermal well tests

vary from an average 3260ppm at Brady's Hot Springs to about 6700ppm at

Desert Peak, a difference of at least 3000ppm. Chemical constituent

ratios show little or no change away from the source. In fact, Carson

Sink-Desert Peak waters are easily recognizable six miles from the

thermal reservoir where cooled samples from the shallow thermal aquifer

system have been sampled and analyzed.

If the Brady's thermal water and Carson Sink-Desert Peak waters

share a common origin, then either Brady's water has been diluted or

Desert Peak water has been highly enriched. The uniformity of constit­

uent ratios implies that dilution has not occurred, as does the agree­

ment of chemical geothermometers with observed temperatures. If Brady's

water were diluted,constituent ratios would be expected to vary, and the

chemical geothermometers would give erroneously low predicted tempera­

tures. There is no evidence presently available which suggests that

dilution has occurred. Conversely, enrichment of a Brady's thermal

water to attain Desert Peak thermal water chemical concentrations is

unlikely. The two geothermal reservoirs occur at approximately the .... 1LM11 i ii ill i! "i

67 and this has not been observed. Chemical changes at depth require

Brady's water and Desert Peak water to separate during upward migration with Desert Peak waters moving through a substantially different geo­ logic environment such as an evaporite sequence. Thick evaporite sequences have not been encountered in any of the deep geothermal tests, and Desert Peak thermal waters have very low 50^ concentrations, sug­ gesting that deep waters have not encountered substantial amounts of evaporites.

The occurrence of a boundary for two groundwater regimes between

Brady's Hot Springs and Desert Peak is probably not coincidental, but rather is indicative of two separate recharge areas for the two thermal systems. Research has shown that most thermal waters were originally meteoric (White, 1970). Isotope data from thermal waters in the northern

Hot Springs Mountains agree with this interpretation (Phillips Petrol­ eum Company, proprietary data), and suggest that initial chemical dif­ ferences of two different groundwaters easily and best account for thermal water differences at Brady's Hot Springs and Desert Peak.

The proximity of Brady's Hot Springs and the Desert Peak geothermal field, and the substantial difference in water chemistry with no evidence

for mixing argue strongly that the two thermal systems are hydrologically

unconnected and have separate recharge systems.

Two flow networks are inferred (figure 13), principally on the

basis of water chemistry, but available water-level data are consistent

with the chemical inference. The most obvious feature of the flownet

is the distinct mounding of the water-table above the thermal anomalies

It seems reasonable to assume that the mounds are related to the thermal

anomalies and two related explanations are invoked to account for their w s m a m m m m ...... „ .'■■M^IN,

68

existence. First, the presence and continued flow of shallow thermal

aquifers shows that thermal waters are moving upward out of the geo­

thermal reservoir into a more orthodox gravitational flow network.

This upward movement of water in thermally anomalous areas is due to

vertical forces that are not well understood. Olmstead and others (1975)

believe the water-table mound at Brady's Hot Springs results from up­

ward-directed forces, termed vertical potential gradients. Vertical

potential gradients may account for the water-table mound above the

Desert Peak Geothermal Field as well. Second, water densities are

lower in the upflowing heated portion of a thermal convection cell.

Decreased water densities result in higher water columns and could thus

be expressed as a temperature-related groundwater mound. Both condi­

tions probably contribute to the groundwater mounding above thermal

anomalies in the Hot Springs Mountains.

The Desert Peak Geothermal Field is still in the early stages of

development. Consequently, knowledge of reservoir characteristics is

inadequate. The areal extent of the reservoir and its thickness are

unknown because its limits are masked somewhat by the "temperature

aureole" emanating upward and outward from it, and a total thickness of

the reservoir has not yet been drilled. Reservoir fluids are contained

in fractures in Mesozoic greenstones. Flow tests indicate more than

adequate recharge is available to sustain long-term electric power gen­

eration. Since the size of the reservoir is unknown, calculation of

heat content which is analogous to oil reserves measured in barrels is

only guesswork. However, available evidence suggests that the highly

fractured greenstone reservoir may be thick and could contain signifi­

cant, as yet untapped fracture horizons equalling or exceeding present pi M M lJUUWW— ■ ■I

69

production (Benoit, pers.common,, 1978). Reservoir pressures are con­ stant and high enough to maintain flow to surface equipment once the

well has been stimulated. In short, preliminary tests indicate that

the Desert Peak geothermal reservoir may be the largest yet discovered in Nevada.

The region underlain by the Desert Peak Geothermal Field is

structurally higher than other parts of the northern Hot Springs Moun­

tains. The field is roughly centered at the mutual corner of sections

14, 15, 22, and 23, T22N-R27E. Most of the surface area above the geo­

thermal field is covered by alluvium but where bedrock is exposed it

consists mostly of rocks of the lower Chloropagus Formation or units

of the unnamed rhyolite sequence (Tr). Temperature distribution deduced

from drillhole data is areally non-linear and shows no obvious associa­

tion with faults observed at the surface. In fact, the lack of linear

temperature distribution strongly suggests that some mechanism other

than high-angle faulting serves to concentrate the heat.

Although high-angle faults do not appear to localize the thermal

anomaly, the outline of the anomaly coincides well with Tertiary faults

trending N25E and N60E, which suggests that high-angle faults may act

as boundaries for the Desert Peak Geothermal Field (figure 9). This is trending fault near the Desert Queen Mine. Both faults are Basin and

Range normal faults and have contributed to the structural elevation

of the field area. Possible north and south boundaries are formed by

the older NbOE trending faults which are probably normal faults but

data are presently inconclusive.

Field mapping, temperature data from drillholes, and detailed

gravity work suggest that the Desert Peak Geothermal Field is contained

in the en echelon, rhombohedral horst blocks described on page 53

(figure 9). Two stages of uplift along two distinct fault sets has

resulted in an area of structural elevation at the intersection of the

two structural trends. Structural elevation of fault blocks in this

area is significant to the geothermal anomaly in at least two ways:

(1 ) the heat aureole surrounding the geothermal reservoir was brought

up to depths shallow enough to be detected by shallow drillhole explor­

ation methods, and (2) Mesozoic rocks which contain the reservoir are

closer to the surface and within range of economic drilling. Also,

reservoir fluids will flow naturally from present discovery depths

without continuous pumping, an expensive and probably prohibitive pro­

cedure at present. Were the reservoir too deep to flow naturally, the

Desert Peak discovery would have been an exploration success but almost certainly an economic failure.

Although the boundary faults define approximate areal limits of the geothermal anomaly they do not in themselves constitute the geo­ thermal reservoir. Tertiary high-angle faults probably act only as boundaries and as a hydrologic divide, although at least one Tertiary fault taps the geothermal reservoir. It functions as the conduit for the small amount of thermal water which leaks upward from the reservoir a

71 into the shallow thermal aquifer system. Other than this minor leakage

Tertiary faults do not hold significant amounts of thermal fluids, even though they extend to depths of at least 4000 to 5000 feet (1220 to 1525 m).

The geothermal reservoir waters are contained in fractures in

Mesozoic greenstones. Tertiary volcanic rocks form an effective cap over the Mesozoic reservoir. The Tertiary rocks are not appreciably hydrothermally altered, indicating that thermal water circulation in the younger rocks has been held to a minimum in spite of the pervasive

Tertiary high-angle faulting. Reservoir fractures are apparently confined to a Mesozoic stratigraphy and structure that is not well known, and correlation of units between the three deep tests is still speculative, thus hampering interpretation of the reservoir fracture system. However, available evidence has some interesting implications.

There is a possible repetition of Mesozoic strata in well 29-1 which bottomed in a hornblendite intrusive possibly related to the middle

Jurassic gabbro lopolith. Greenstone reservoir rocks in wells 21-1 and 21-2 may be comagmatically related to the gabbro intrusive. If so, the fractures in the greenstones and the possible bedding repetition could be related to the thrusting associated with lopolith emplacement.

Bedding repetitions, comagmatic greenstones, gabbro intrusive, and lopolith-related thrusting are present in the West Humboldt Range and

Mopung Hills 12 miles (19 km) to the north. The inference of similar relationships in the Mesozoic section beneath the Hot Springs Mountains is admittedly conjectural but suggests that if thermal fluids circulate in thrust zones in Mesozoic rocks, the area of recharge for the Desert

Peak Geothermal Field may be regional in extent. Flow test data from m

72

well 21-1 indicate that thermal water recharge is greater than would

be expected from a localized vertical fracture system, which lends cre­

dence to the above inference. Alternatively, however, Tertiary fault

zones might intersect pre-existent Mesozoic fracture systems. Shattered

intersections could well provide sufficient porosity and permeability

for deep, high-volume circulation of thermal waters, as well as a basis

for localization. Why some Tertiary faults would act as borders and

hydrologic divides and others would remain permeable is highly specula­

tive .

The source of heat is presently unknown. Available geophysical

I evidence does not support the inference of a shallow magma chamber m below or tangential to the Hot Springs Mountains area. Heat flow out­

side Brady's and Desert Peak thermal anomalies is about average for the

Basin and Range (Olmstead and others, 1975: Phillips Petroleum Co.,

proprietary data). Unless extensive lateral thermal water circulation

exists, as yet undetected, the implication is clear that thermal waters

must circulate to depths of several thousand feet (several hundred

meters) in order to attain the observed temperatures. Geologically, the

relationship between Brady's Hot Springs and the Desert Peak geothermal

anomaly is uncertain. Their proximity suggests a common heat source,

but water chemistry indicates the two systems are unconnected at

shallower depths and since the high observed temperatures suggest cir­

culation as deep as 5000 feet (1525 m) the thermal systems are probably

unconnected at depth as well. It is likely therefore that both thermal

anomalies reflect deep, unconnected thermal water convection systems,

and the proximity of the two anomalies is geologically coincidental.

Thus the heat source itself may be a common factor, but the thermal

convection systems are certainly separate entities (figure 14). 10,000-12,000 feet ?

FIGURE 14 schematic cutaway of conceptual model of Desert Peak- Brady's Hot Springs area. Note separate recharge, water-table mounds, and separate convection systems: Brady's Hot Springs along the fault, and Desert Peak in fractured Mesozoic rocks. 74

SUMMARY

The Desert Peak Geothermal Field is situated in a complexly faulted en echelon series of rhombohedral horst blocks located at the intersec­

tion of two presumably Tertiary fault systems trending N25E and N60E.

Horst-bounding faults may define the borders of the geothermal system.

Structural elevation by horsting has brought the geothermal reservoir, contained in Mesozoic basement rocks, to depths shallow enough to per­ mit detection and potential economic exploitation of the thermal anomaly.

Reservoir waters up to 406°F (207°C) are contained in fractures in

Mesozoic greenstones. The greenstones and the fractures may be geneti­ cally related to emplacement of the middle Jurassic Humboldt Lopolith.

Available evidence suggests that reservoir permeability and porosity are related to Mesozoic tectonic events, although shattered intersec­ tions of Tertiary fault zones and Mesozoic fracture systems cannot be excluded. Nevertheless, localization of thermal waters appears to be a consequence of basement tectonic elements. The reservoir is trapped beneath and maintained by a thick Tertiary volcanic sequence. In spite of pervasive Basin and Range faulting, the volcanic cap is effective, allowing only a small amount of thermal water leakage from the reservoir into a shallow thermal aquifer system.

Hydrologic and geochemical data indicate that the Desert Peak Geo­ thermal Field is distinct and separate from the nearby Brady's Hot Springs thermal system. Reservoir waters and associated thermal convection systems of the two thermal anomalies are unconnected at least to 5000 feet (1220 m), although the proximity of the Desert Peak anomaly to the Brady's anomaly argues strongly for a common heat source.

Water-table data and resultant flownet show not only that the two anomalies are hydrologically separate, but reveal the existence of dis­ tinct groundwater mounding above the thermal anomalies. The mounding is considered to be the effect of vertical potential flow gradients and decreased densities of heated waters in the upward flowing portions of thermal convection cells.

Source of heat for the Brady's Hot Springs system and Desert Peak geothermal system is unknown. Geophysical data presently available appear to preclude a shallow crustal heat source. The association of both Desert Peak and Brady's thermal anomalies with faults suggests deep circulation of thermal waters to at least 5000 feet (1220 m) and upward thermal convection. Flow tests indicate that Brady's thermal system is hampered by inadequate recharge, whereas thermal water recharge of the Desert Peak Geothermal Field is sufficiently high to imply re­ gional interconnection. 76

CONCLUSIONS

The future of geothermal energy, and the geothermal explorationist, depends to a large degree on the geologist's ability to identify and develop viable new geothermal prospects. Previous geothermal explora­ tion has concentrated on areas of surface thermal activity, but the discovery of the Desert Peak Geothermal Field has opened a vital second chapter in geothermal exploration: identification of blind geothermal resources. Blind geothermal resource identification will require the application of every available exploration tool, including increased attention to the relationship of geology to the localization of geo­ thermal resources.

The Desert Peak Geothermal Field was discovered without the use of geology, although post-discovery surface mapping (this thesis) shows a distinct, positive structural relationship between geology and geo­ thermal occurrence. Whether or not this relationship could have been recognized prior to discovery is doubtful. Delineation of structural patterns and recognition of their possible implications might have sig­ nificantly reduced target size and consequently exploration time.

Knowledge of the structural pattern can aid field development in two ways. Future deep tests will be preceded by strat. tests. Strat. tests located on the basis of the known geology will help clarify structural and geothermal elements which are only approximately known now (i.e., fault position, displacement, and attitude, field and hydrologic bound­ aries, and possibly even reservoir characteristics). Secondly, deter­ mination of the stratigraphic succession and comparison of lithology with temperature profiles shows that the unnamed rhyolite unit (Tr) 77 can be trusted to give a true temperature-gradient. Temperature profiles in the overlying Chloropagus Formation, occasionally the host to the shallow thermal aquifer system, can be misleading. In the future, strat. tests can be designed to drill into the unnamed rhyolite until sufficient thickness is penetrated to give reliable temperature profiles. Signi­ ficant reductions of drilling costs may result.

Delineation of structural patterns on a regional basis could help geothermal energy exploration. Trexler and others (1978) investigated lineament analysis as a geothermal exploration technique in Nevada, and concluded that judicial use of the method could be beneficial. The

results of this thesis are similar to Trexler's, though on a smaller scale: intersections or disruptions of major trends, such as the N25E

and N60E fault zones described herein, are good places to begin explor­

ation even though heat localization may be a response to other, as yet

unrecognized, tectonic elements. Also, the recognition of groundwater mounding over thermal anomalies has significance in that it may be a

direct response to thermal water upwelling. Incorporation of accessible hydrologic information into the exploration architecture, both on a

regional and on a detailed basis, represents a potentially useful and

pragmatic exploration tool, if such mounding is present over other

thermal anomalies. Methods could be developed to discriminate between

thermally-caused groundwater mounds and non-thermal phenomena.

The results of this thesis indicate that geologic structure can

provide information significant to geothermal heat localization on a

local scale, and potentially on a regional scale. At Desert Peak, knowl­

edge of geologic structure may simplify field development. Regional

analysis of structural patterns may identify analogous or complementary 78

structural elements which might similarly affect geothermal heat locali­

zation. A beneficial by-product of structural studies is the recogni-

tion of possible groundwater mounding over thermal anomalies.

I

J 79

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