Geological Survey 

Research 1962

Short Papers in Geology, Hydrology 

and Topography, Articles 180-239

GEOLOGICAL SURVEY PROFESSIONAL PAPER 450-E
Short Papers in Geology 

Hydrology, and Topography 

Articles 180-239

GEOLOGICAL SURVEY RESEARCH 1962

GEOLOGICAL SURVEY PROFESSIONAL PAPER 450-E

Scientific notes and summaries of investigations 

prepared by members of the Conservation, 

Geologic, Topographic, and Water Resources 

Divisions in the fields of geology, hydrology, 

topography, and related sciences

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON : 1963
UNITED STATES DEPARTMENT OF THE INTERIOR 

STEWART L. UDALL, Secretary

GEOLOGICAL SURVEY 

Thomas B. Nolan, Director

For sale by the Superintendent of Documents, U.S. Government Printing Office

Washington 25, D.C.
FOREWORD

This collection of 60 short papers on subjects in the fields of geology, hydrology, topography, 

and related sciences is the last of a series released as chapters of Professional Paper 450. The 

papers in this chapter report on the scientific and economic results of current work by members 

of the Conservation, Geologic, Topographic, and Water Resources Divisions of the United States 

Geological Survey. Some of the papers announce new discoveries or present observations on 

problems of limited scope; other papers draw conclusions from more extensive or continuing 

investigations that in large part will be discussed in greater detail in reports to be published in 

the future.

Chapter A of this series presents a synopsis of results from a wide range of work done during 

the 1962 fiscal year.

/t

THOMAS B. NOLAN,

Director.

in

CONTENTS

Page 

Foreword. ___-__-_--__-___-_-___________________________________________________-______________--_----------_- Ill

GEOLOGIC STUDIES 

Economic geology

180. Mineralization associated with a magnetic anomaly in part of the Ely quadrangle, Nevada, by A. L. Brokaw,

G. B. Gott, D. R. Mabey, Howard McCarthy, and Uteana Oda_________________________________ El

181. Preliminary report on alkalic intrusive rocks in the northern Wet Mountains, Colorado, by R. L. Parker and F. A.

Hildebrand__________________________________________________________________ 8

Engineering geology

182. Landslides near Gardiner, Montana, by H. A. Waldrop and H. J. Hyden___________________-__----_-_-------- 11

Structural geology

183. The Crooked Creek disturbance, southeast Missouri, by T. H. Kiilsgaard, A. V. Heyl, and M. R. Brock_________ 14

184. Origin of the Middlesboro Basin, Kentucky, by K. J. Englund and J. B. Roen____ _____________ _______________ 20

185. Thrusting developed after folding in the Valley and Ridge province, southwest Virginia, by L. D. Harris________ 22

186. Reversal of throw along a line of low-angle thrust faulting near San Pedro de Atacama, Chile, by R. J. Dingman_ 25

187. Gravity and magnetic anomalies in the northern Oquirrh Mountains, Utah, by D. R. Mabey, E. W. Tooker, and

R. J. Roberts_____________________________________________________-___--______-_-_--------_--------- 28

Stratigraphy

188. Comparison of Oquirrh Formation sections in the northern and central Oquirrh Mountains, Utah, by E. W. Tooker

and R. J. Roberts.________________________________________________________________ 32

189. Windy Gap Volcanic Member of the Middle Park Formation, Middle Park, Colorado, by G. A. Izett, R. B. Taylor,

and D. L. Hoover__________________________________________________________-____-_-_-----_---------- 36

190. Shallow halite deposits in the Flowerpot Shale in southwestern Oklahoma, by P. E. Ward______________________ 40

191. Crinoidal bioherms in the Fort Payne Chert (Mississippian) along the Caney Fork River, Tennessee, by M. V.

Marcher_________-_______________________________________-____-__---_---_-------_---------------- 43

192. Mississippian-Pennsylvanian boundary in northeastern Kentucky, by R. A. Sheppard and Ernest Dobrovolny_ _ _ _ 45

193. Correlation of the Parkwood Formation and the lower members of the Pottsville Formation in Alabama, by W. C.

Culbertson_ __ __ _____________________________________________________--_______--_------------------ 47

194. Pennsylvanian nomenclature in northwest Georgia, by W. C. Culbertson____________________________--__----- 51

195. Pathfinder uplift of Pennsylvanian age in southern Wyoming, by W. W. Mallory___________________,-_----_---- 57

196. Unconformity marking the Jurassic-Cretaceous boundary in the La Ligua area, Aconcagua Province, Chile, by

W. D. Carter___________________________________________________________________-_-----------_------ 61

197. Relations of the Navajo and Carmel Formations in southwest Utah and adjoining Arizona, by J. C. Wright and

D. D. Dickey_________________________________________________________________-----_ 63

Mineralogy, geochemistry, and petrology

198. Bentonite beds of unusual composition in the Carmel Formation, southwest Utah, by L. G. Schultz and J. C. Wright_ 67

199. Sampling a zoned galena crystal for lead isotope study, by R. S. Cannon, Jr., K. L. Buck, and A. P. Pierce______ 73

200. Epidote and related minerals in two deep geothermal drill holes, Reykjavik and Hveragerdi, Iceland, by G. E.

Sigvaldason_ __ _ ___________________________________________________________________________________ 77

201. Epidote in hot-spring systems, and depth of formation of propylitic epidote in epithermal ore deposits, by D. E.

White and G. E. Sigvaldason______________________________________-_______-___------------------------ 80

202. Thorium and uranium in some volcanic rocks from the circum-Pacific province, by David Gottfried, Roosevelt Moore,

and Esma CampbelL________________________________________---________----------------------------- 85

203. The quartz diorite line in northwestern North America, by J. G. Moore. Arthur Grantz, and M. C. Blake, Jr_____ 89

204. Syenite complex older than the Idaho batholith, Big Creek quadrangle, central Idaho, by B. F. Leonard.__ ______ 93

205. Trondhjemite in the Riggins quadrangle, western Idaho, by Warren Hamilton.___________________----_------- 98

206. X-ray determinative curve for olivines of composition Fogo-95 from stratiform and alpine-type peridotites, by P. E.

Hotz and E. D. Jackson_________________________________________________------------_-------------- 101

207. Formation of "salt cups" near San Pedro de Atacama, Chile, by R. J. Dingman____________ ___ ______-------_- 103

Geochronology

208. Lead-alpha ages of zircon in quartz monzonite porphyry, Thiel Mountains, Antarctica a preliminary report, by

A. B. Ford, H. A. Hubbard, and T. W. Stern._______________________ .____.--_._-.--__ - - 105

209. Evaluation of the Pa231 /U- Th230 /U method for dating Pleistocene carbonate rocks, by J. N. Rosholt and P. S.

Antal..____________________________________________________________________ ._____-_--.------------- 108
VI CONTENTS

Page 

Sedimentary petrology

210. Relation of the pore volume of silty sediments to overburden load, particle size, and sorting, by R. H. Meade__ .___ Elll 

Geomorphology and glacial geology

211. Menan Buttes, cones of glassy basalt tuff in the Snake River Plain, Idaho, by Warren Hamilton and W. B. Myers._ 114

212. Comparison of drainage on topograpic maps of the Piedmont province, by E. V. Giusti and W. J. Schneider _._____ 118

213. Correlation of some glacial deposits in New Mexico, by G. M. Richmond.___________________________________ 121

Analytical methods

214. Ion-exchange separation and spectrophotometric determination of cadmium, by Claude Huffman, Jr____________ 126

215. Effect of copper on the precipitation of tellurium with hypophosphorous acid using selenium or gold as a collector,

by H. W. Lakin and C. E. Thompson_____. ___________________________________________ _____________ .. 128

216. Spectrophotometric determination of fluorine with thoron, by B. L. Ingram-__________________________________ 130

TOPOGRAPHIC STUDIES 

Mapping techniques

217. Portable surveying tower, by J. L. Buckmaster and W. D. Murphy------__________________________________ 131

218. Airborne control system, by J. L. Buckmaster, H. B. Loving, and T. O. Dando______________________________ 133

HYDROLOGIC STUDIES 

Engineering hydrology

219. Effects of major water-table changes in Kings and Queens Counties, New York City, by N. M. Perlmutter and

Julian Soren___ _______ _________________ ____________________________ _________ ______________ ________ 136

220. Infiltration rates in weathered crystalline rocks at the Georgia Nuclear Laboratory, Dawson County, Georgia, by

J. W. Stewart..--_____________._..______________-__--.__----__-_-______ ._ ___________________________ 140

221. Evapotranspiration and the relation of ground water to surface water in the Pond Creed basin, Oklahoma, by W. E.

Clark______________________________________________________________________________________________ 142

222. Estimating the specific capacity of a well, by A. N. Turcan, Jr ____________ __________________________________ 145

223. Aerial observation of ice cover to locate areas of ground-water inflow to streams, by S. W. Wiitala and T. G. Newport- _ 148 

Ground water

224. Permeability of glacial till, by S. E. Norris_____ ____- .___ _______.____ _____.___._____.___ _____________ 150

225. Ground water in Cenozoic fill in collapse structures, southeastern Eddy County, New Mexico, by J. B. Cooper-___ 152

226. Relation of fresh water to salt water at Centre Island, Nassau County, New York, by John Isbister____ _______ 154

227. Highly productive aquifers in the Tacoma area, Washington, by K. L. Walters-_ . ____________________________ 157

Surface water

228. Long-term flood frequencies based on extremes of short-term records, by W. J. Conover and M. A. Benson_______ 159

229. Flood peaks related to hydrologic factors in the Southwest, by M. A. Benson________________________________ 161

Quality of water

230. Dissolved-salt contribution to Great Salt Lake, Utah, by A. M. Diaz _________________________________________ 163

Analytical hydrology

231. Orientation of axes for calculating distribution of transmissibility from water-level altitudes, by R. W. Stallman____ 165

232. A method of determining the storage-outflow characteristics of nonlinear reservoirs, by John Shen____________ 167

233. An analog solution of the turbulent-diffusion equation for open-channel flow, by John Shen____________________ 169

234. Measurement of sediment-laden flow by means of a circular orifice, by Nobuhiro Yotsukura_---___- ____________ 172

Experimental hydrology

235. Use of a neutron moisture probe to determine the storage coefficient of an unconfined aquifer, by W. R. Meyer_-_- 174

236. Centrifuge technique for determining time-drainage relations for a natural sand, by R. C. Prill and A. I. Johnson__ 177

237. Behavior of detergents (ABS), bacteria, and dissolved solids in water-saturated soils, by H. G. Page, C. H. Wayman,

and J. B. Robertson_________________________________________________________________________________ 179

238. Adsorption of the surfactant ABS35 on kaolinite, by C. H. Wayman, J. B. Robertson, and H. G. Page___ ________ 181

239. An application of the Gibbs adsorption equation to detergent solutions, by C. H. Wayman__ ______________ .... 184

INDEXES 

Subject____-___--_-______-_--.______________________________________________________________ _____________ 187

Author.____________ _ _______________________________________________________________________________ _____ 188
GEOLOGICAL SURVEY RESEARCH 1962

SHORT PAPERS IN GEOLOGY, HYDROLOGY, AND TOPOGRAPHY, ARTICLES 180-239

GEOLOGICAL STUDIES

ECONOMIC GEOLOGY

180. MINERALIZATION ASSOCIATED WITH A MAGNETIC ANOMALY IN PART OF THE ELY

QUADRANGLE, NEVADA

By ARNOLD L. BKOKAW, GARLAND B. GOTT; DON K. MABEY; HOWARD MCCARTHY, and UTEANA ODA, Denver,

Colo. "f Menlo Park, Calif.; Denver, Colo.

Geologic mapping in the Ely quadrangle, Nevada, 

in 1958-60 revealed structural irregularities and rock 

alteration suggesting that the Ward Mountain area 

of the Egan Range contains concealed intrusive rocks 

and possible accumulations of metals of economic, 

value. An aeromagnetic survey tested this theory, and 

a large magnetic anomaly, similar to one near Ruth, 

Nev., was outlined in the southwest quarter of the 

quadrangle. Subsequently, a geochemical study tested 

the possibility that the magnetic anomaly reflects a 

buried mineralized intrusive body.

The Egan Range is a major north-trending mountain range in east-central Nevada. The central part 

of the range, which is the subject of this article, lies 

south-southwest of Ely between Steptoe Valley on 

the east and the White River valley on the west (fig. 

180.1). The average altitude of the crest of the range 

in this area is about 10,000 feet above sea level, and 

the highest peak is 10,933 feet. Near Ruth the altitude decreases to about 7,000 feet. The floors of the 

adjoining valleys are about 6,000 to 6,500 feet above 

sea level.

The Paleozoic sedimentary rocks are a sequence of 

dolomite, limestone, shale, and sandstone ranging in 

age from Ordovician to Permian (fig. 180.2). These 

rocks are overlain in places along the eastern flank of 

the range by early Tertiary sedimentary rocks and 

by*somewhat younger volcanic rocks. Gravity data 

indicate several thousand feet of Cenozoic sedimentary rocks (valley fill) in Steptoe Valley but generally less in the White River valley.

Small quartz monzonite dikes of early Tertiary or 

Late Cretaceous age intrude the Paleozoic rocks on

the east flank of Ward Mountain, and similar intrusive rocks are exposed near Ruth.

The pre-Tertiary rocks have been complexly folded 

and displaced by thrust and high-angle faults, whereas 

the Tertiary rocks have been only slightly folded and 

tilted.

An aeromagnetic survey of the central part of the 

Egan Range was flown at an elevation of 11,500 feet 

along north-south lines spaced about 2 miles apart. 

The resulting magnetic data are shown by contours on 

figure 180.1.

A positive magnetic anomaly extends over most of 

the range from Ruth south to the edge of the map. 

Two local positive anomalies are superimposed on the 

more extensive anomaly. The northern anomaly is 

centered over the south edge of the Robinson mining 

district, where disseminated copper deposits occur in 

an east-trending zone of metamorphosed sedimentary 

rocks and altered monzonite porphyry. The nearsurface igneous rocks associated with the copper deposits can account for only part of the measured northern anomaly, and a major part of it must be caused 

by a buried igneous body of greater areal extent.

The southern magnetic anomaly is centered over the 

west edge of the Egan Range where the range is composed of nonmagnetic Paleozoic sedimentary rocks. 

The magnetic anomaly is interpreted as due to a large 

mass of igneous rocks underlying the thickest sedimentary rocks of the Egan Range and the eastern 

edge of the White River valley. The zone of high 

magnetic intensity connecting this anomaly with the 

anomaly near Ruth suggests that the igneous bodies

El
E2 GEOLOGICAL SURVEY RESEARCH 1962

115°00'

EXPLANATION

^_ 2650 r""^

^___________y 

Magnetic contours

Hachured where enclosing areas 

of lower intensity. Alined 

dashes are flight lines

Igneous intrusive 

rocks

39°15'

5 MILES

I

39°00'

FIGURE 180.1. Aeromagnetic map of the central Egan Range. Values are total magnetic intensity relative to an 

arbitrary datum. Contour interval is 10 and 50 gammas. Flight level was 11,500 feet above sea level. 

Magnetic survey was flown and compiled under the supervision of J. L. Meuschke. Geology is generalized 

from McJannet (1960).

FIGURE 180.2. Generalized geologic map of the southwest part of the Ely quadrangle, Nevada. Qal, alluvium 

(Quaternary); TV, volcanic rocks, and Tsp, Sheep Pass Formation of Winfrey (1958) (Tertiary); TKm, 

monzonite porphyry (Tertiary or Cretaceous); PP, Permian and Pennsylvanian, undifferentiated; Me, 

Chainman Shale (Mississippian); Mj, Joana Limestone, and Mjj (stipple), Joana jasperoid (Mississippian); D, 

Devonian, undifferentiated; SO, Silurian and Ordovician, undifferentiated; Oep, Eureka Quartzite and Pogonip 

Group (Ordovician). Geology by A. L. Brokaw and D. R. Shawe.
ARTICLE 180 E3

114°55'

R.62 E. R.63 E.

2 MILES

Contact Fault Thrust fault 

Dashed where approximately located Dashed where approximately located. U, Sawteeth on upper plate

upthrown side, D, downthrown side Area of geochemical anomalies
E4 GEOLOGICAL SURVEY RESEARCH 1962

believed to produce the two local anomalies are connected at depth.

The geologic features (fig. 180.2) of the general 

area of the southern magnetic anomaly tend to confirm 

this interpretation. In gross aspect this part of the 

Egan Range appears to be a simple tilted fault block 

similar to many other ranges in Nevada and Utah. 

In detail, however, the structure is more complex. 

The sedimentary rocks have been cut by numerous 

faults that displace the beds from a few tens of feet 

to several thousand feet and divide the range into 

several smaller structural blocks.

Overturned folds and thrust faults mapped within 

and beyond the limits of the area shown in figure 180.2 

indicate that thrust faulting was important in the 

early development of the range. Along Rowe Creek, 

below the outcrop belt of the Chainman Shale, the 

beds are overturned and dip to the west on the limb 

of a recumbent fold that is overturned to the east. 

The attitude of the fold indicates that a thrust plate 

moved to the east over the rocks now exposed. The 

small klippen of Devonian rocks 1 and 3 miles south 

of Rowe Creek may represent remnants of this plate.

Just north of Rowe Creek and near the center of 

the southern magnetic anomaly (shown in detail in 

fig. 180.3), the rocks are cut by the Rowe Creek fault, 

which displaces the rocks several thousand feet in a 

left-lateral sense. The fault is terminated on the east 

by a northwest-trending reverse fault. This reverse 

fault appears to die out to the southeast in the Chainman Shale, although its counterpart or continuation 

may be one of the strong faults that cut the Pennsylvanian and Permian rocks to the southeast. Similarly, 

one of the left-lateral faults that extend across Ward 

Mountain may be the offset segment of the Rowe 

Creek fault. Where the Joana Limestone is cut by 

the reverse fault and related fractures, it is altered to 

jasperoid, and many of the minor fractures are filled 

with secondary calcite and iron-oxide minerals. The 

exposed lower and middle Paleozoic limestones and 

dolomites also are silicified in the areas of structural 

deformation.

Some evidence suggests that the rocks near the magnetic anomaly, particularly in the vicinity of the klippen, have been differentially elevated, forming a structural high that might be due to doming related to a 

buried intrusive. The 1,500-foot difference in elevation between the 2 klippen shown on the map is more 

than can be accounted for by the known faulting.

The similarity of the two magnetic anomalies, the 

existence of major ore bodies in the area of the northern anomaly and of abandoned small mines on Ward 

Mountain, the presence of minor intrusive masses in 

the area of Ward Mountain, and the presence of jasperoid, all suggest that the area of the southern anomaly may also contain significant ore deposits. A geochemical study was therefore made.

Semiquantitative spectrographic analyses for 30 elements, including Cu, Pb, Mo, Ag, Sn, and W, were 

made in the field by U. Oda, M. DeValliere, and W. W. 

Janes, using a truck-mounted spectrograph. As, Sb, 

and Zn were determined by field-type chemical methods of analysis (Ward and others, 1963). Hg analyses were made spectrographically by Paul Barnett using a large sample charge in special electrodes.

The analytical results show that anomalous amounts 

of Ag, Hg, Cu, Pb, Zn, Mo, Sb, As, W, and Sn are 

present in jasperoid and iron-rich fracture fillings in 

the limestones and dolomites along and adjacent to 

major faults directly over the magnetic anomaly (figs. 

180.4 and 180.5). The Rowe Creek fault and the 

northwest-trending fault that is roughly parallel to 

the outcrop belt of the Joana Limestone show the 

highest concentrations of metals and probably were 

the principal conduits for metal-bearing solutions. 

Ba, Cr, Fe, Mn, and Ti also are present in abnormally 

high concentrations in the anomalous area. Unaltered 

limestones and calcite veins are not appreciably enriched in any of the metals except perhaps Hg and As.

The high concentration of metals along the faults 

and over the magnetic anomaly may be a leakage halo 

from metal deposits in and near the postulated concealed intrusive. The quartz monzonite dikes on the 

east side of the area are probably connected at depth 

with this postulated intrusive. The extensive alteration of limestone to jasperoid suggests that similar 

dikes are present at depth in the Rowe Creek area, 

under the geochemical anomalies. The depth to the 

intrusive rocks is not known; however, the low mobility of some of the elements in a limestone environment 

suggests that the intrusive may be at a relatively shallow depth.

Depth analyses of the geophysical data indicate that 

the top of the mass causing the major magnetic feature (fig. 180.1) is about 3,000 feet above sea level or 

3,000 feet below the surface of the White River valley. 

Details of the magnetic field were not well defined by 

this survey, but the gradients associated with the 

southern magnetic high indicate that parts of the mass 

are nearer the surface than others. Its surface is probably irregular and may be at a quite shallow depth in 

some places.
ARTICLE 180 E5

114°55'

ARE OF GEOCHEMICAL,

R.62E. R.63E.

2 MILES

FIGUKE 180.3. Aeromagnetic map, southwest part of Ely quadrangle. Interval 10 gammas.
E6

R.62 E. R.63 E.

GEOLOGICAL SURVEY RESEARCH 1962

R.62E. R.63 E.

T.14N.

SILVER

>50 10-50 5-10 

R.62E. R.63E.

T.14N.

>300

COPPER

30-70 <30

T.14N.

ARSENIC

100-7000 30-100 

R.62E. R.63 E.

<30

T.14N.

MOLYBDENUM

>200 100-200 50-100 <50.

2 MILES

J

FIGURE 180.4. General distribution, in parts per million, of Ag, As, Cu, and Mo in jasperoid and in iron oxide in vugs and

fractures.
ARTICLE 180 E7

R.62 E. R.63 E. R.62 E. R.63E.

T. 14 N

LEAD AND ZINC

R.62 E. R.63 E.

T. 14 N

>0.5

MERCURY

0.2-0.5 <0.2

T.14N

12

ad*4 !

13

J- 24

Rowe..

\ 17

TIN AND TUNGSTEN

R.62 E. R.63 E.

T. 14 N

ANTIMONY

10-100

FIGURE 180.5. General distribution, in parts per million, of Pb, Zn, Sn, W, Hg, and Sb in jasperoid and in iron oxide in vugs

and fractures.
E8 GEOLOGICAL SURVEY RESEARCH 1962

REFERENCES

McJannet, George, compiler, 1960, Geologic map of east central 

Nevada, in Guidebook to the geology of east-central Nevada, Intel-mountain Association of Petroleum Geologists, 

llth Annual Field Conference, 1960, held with Eastern 

Nevada Geological Society.

Ward, F. N., Lakin, H. W., Oanney, F. C., and others, 1963,

Analytical methods used in geochemical exploration by 

the U.S. Geological Survey; U.S. Geol. Survey Bull. 1152, 

100 p.

Winfrey, W. M., Jr., 1958, Stratigraphy, correlation, and oil 

potential of the Sheep Pass formation, east-central Nevada, in Am. Assoc. Petroleum Geologists Rocky Mtn. 

Section, Geol. Rec., 8th Ann. Mtg., p. 77-82.

181. PRELIMINARY REPORT ON ALKALIC INTRUSIVE ROCKS IN THE NORTHERN WET MOUNTAINS, COLORADO 

By RAYMOND L. PARKER and FRED A. HILDEBRAND, Denver, Colo.

Two genetically related alkalic intrusive complexes 

in the northern Wet Mountains of south-central Colorado have been discovered by the U.S. Geological Survey (fig. 181.1). The largest complex is centered 

around McClure Mountain about 13 miles north of 

Westcliffe, and it occupies an area of about 20 square 

miles. It is roughly circular in plan, is cut by radiating faults, and consists principally of mafic and felsic 

alkalic rocks that have concentric distribution. A 

smaller complex at Gem Park, about 5 miles southwest of the McClure Mountain complex, is roughly 

rectangular in plan, is about 2 square miles in area, 

and consists mainly of mafic alkalic rocks. Both complexes intrude gneissic granite and other metamorphic 

rocks of Precambrian age; the Gem Park complex is 

partly overlain by rhyolite of Tertiary age. Both 

complexes are cut by carbonatite dikes, some of which 

contain niobium, rare earths, and thorium. The distribution and extent of niobium, thorium, and rareearth minerals in the McClure Mountain and Gem 

Park complexes are not yet known, but because of the 

similarity to other alkalic complexes in the world that 

contain minable deposits, these two complexes seem 

worthy of exploration to determine their economic 

potential.

Principal rocks comprising the McClure Mountain 

and Gem Park intrusives are: pyroxene-olivine-plagioclase rocks, biotite-hornblende syenite, mafic nepheline-bearing rocks, nepheline syenite, and fenite. The 

pyroxene-olivine-plagioclase rocks are the oldest and 

the nepheline-bearing rocks are the youngest although 

the span in age among the rocks is probably not great. 

They no doubt are related and belong to the same 

comagmatic series.

Dark-gray to dark-brown medium-grained to very 

coarse pyroxene-olivine-plagioclase rocks are prominent along the eastern border of the McClure Mountain complex, and form most of the Gem Park intrusive body. These rocks range from those composed 

mostly of pyroxene to those composed mostly of plagioclase (labradorite). The pyroxene-rich rocks and 

those with about equal proportions of pyroxene, olivine, and plagioclase are the most abundant. Some 

pyroxene-rich rocks contain concentrations of magnetite, and some contain abundant apatite.

Biotite-hornblende syenite intrudes the pyroxeneolivine-plagioclase rocks on the east side of the McClure Mountain complex. The rock is coarsely crystalline, commonly has a trachytoid texture and is 

composed mostly of perthite and albite-oligoclase 

(about 80 percent), and brown biotite and pale-green 

hornblende, occurring singly or together (10 to 20 percent) . In places the alinement of biotite and feldspar 

crystals defines a strong foliation that is attributed to 

flowage during intrusion.

Nepheline syenite, the most abundant rock type in 

the McClure Mountain complex, forms the core of the 

complex and intrudes biotite-hornblende syenite. The 

rock is medium grained to very coarse and is composed of sodic orthoclase, highly variable amounts of 

nepheline, and small to moderate amounts of sodalite 

and biotite or hornblende. Sphene is abundant in 

some varieties. The nepheline is clear and has a 

greasy luster. It weathers readily, leaving recessed 

light-gray or light-brown grain surfaces between intersecting white perthite crystals. The nepheline syenite contains abundant large xenoliths of biotitehornblende syenite near the contact between the nepheline syenite and the biotite-hornblende syenite.

Mafic nepheline-bearing rocks, which are enclosed 

and cut by light-gray nepheline syenite, are prominent 

on McClure and Elkhorn Mountains. These rocks are 

composed of highly variable amounts of nepheline, 

orthoclase, amphibole, brown biotite, and sphene. 

Some of these rocks are amphibole ijolite.
ARTICLE 181 E9

105°30' 105°25'

Gneissic granite 

and related rocks

R. 73 W. R. 72 W.

2 MILES 

J

FIGURE 181.1. Generalized geologic map of the alkalic rocks at Gem Park and McClure Mountain, Fremont and Ouster Counties,

Colo. Unpatterned area not mapped.

Very coarse to medium-grained fenite composed is not present at the border of the pyroxene-olivinemostly of orange feldspar occurs at the border of the plagioclase rocks.

complex where nepheline syenite and biotite-horn- Abundant orange or light-gray syenite porphyry

blende syenite have intruded the gneissic granite and dikes and a few mafic fine-grained dikes (not shown

other associated metamorphic rocks. In general, fenite on fig. 181.1) intrude both complexes.
E10 GEOLOGICAL SURVEY RESEARCH 1962

Carbonatite dikes occur in the McClure Mountain 

complex but are more abundant in the Gem Park 

complex. They are generally light brown to dark 

reddish brown and range in width from less than 1 

foot to more than 10 feet. The larger dikes pinch 

and swell and branch into smaller dikes along their 

strikes. Some of the dikes extend beyond the border 

of the mafic rocks well into the bordering gneissic 

granite. Preliminary mineralogic examination shows 

that the carbonatite dikes are composed principally 

of calcite and dolomite. Some dikes contain rareearth-bearing apatite, nontitaniferous magnetite, niobian (?) ilmenite, ancylite, an unidentified thorium 

mineral, two unidentified cerium-group minerals (one 

of which may be bastnaesite), and a light-blue fibrous 

mineral (possibly riebeckite). Sodium niobate has 

been reported by the authors from vermiculitic mafic 

rock near the center of Gem Park (Parker and others, 

1962).

The alkalic rocks in the McClure Mountain complex 

and at Gem Park are probably of late Precambrian 

age. An albite syenite body, partly mapped by Christman and others (1959) in the northwestern part of

the McKinley Mountain area only 6 miles southeast 

of the McClure Mountain complex, was determined by 

the Larsen zircon method to have an average age of 

595 million years. This syenite and the rocks of the 

McClure Mountain complex all postdate the regional 

metamorphism of Precambrian age, and are probably 

about the same age.

The thorium vein deposits and alkalic alteration in 

the McKinley Mountain area probably are related 

genetically to the McClure Mountain complex because 

most of the veins, syenite dikes, and feldspathized 

zones in that area strike toward the complex; also the 

rock alteration and the mineral assemblage in the 

veins are types that commonly accompany alkalic 

intrusions.

REFERENCES

Christman, R. A., Brock, M. R., Pearson, R. C., and Singewald, 

Q. D., 1959, Geology and thorium deposits of the Wet 

Mountains, Colorado, A progress report: U.S. Geol. Survey Bull. 1072-H.

Parker, R. L., Adams, J. W., and Hildebrand, F. A., 1962, 

A rare sodium niobate mineral from Colorado: Art. 61 

in U.S. Geol. Survey Prof. Paper 450-C, p. C4-C6.
ARTICLE 182 

ENGINEERING GEOLOGY

182. LANDSLIDES NEAR GARDINER, MONTANA 

By H. A. WALDROP and H. J. HYDEN, Denver, Colo.

Ell

Glide blocks, slumps, and earthflows of late Pleistocene and Recent age were noted in tlie course of geologic mapping in the Gardiner quadrangle in 1960 and 

1961. The landslides are in Yellowstone National 

Park, just southwest of Gardiner, Mont. (fig. 182.1).

The north and east slopes of Sepulchre Mountain 

are disintegrating into complex intersecting glide 

blocks and slumps. Large earthflows have begun in 

this area and have run down over alluvial terraces 

above the Gardner and Yellowstone Rivers. The landslides form a crudely transitional sequence; they begin 

as block glides and slumps, then fragment and move 

as earthflows. Three generations of earthflows were 

distinguished.

Weed (1893, p. 16) mapped the earthflows as moraines and ascribed their form to molding under glacier ice. Despite the advent of aerial photography 

and increased knowledge of the nature of landslide 

and glacial deposits, they have been described and 

mapped as glacial drift ever since. Their manner of 

origin is evident on aerial photographs (fig. 182.2), 

which show they have the structure and topographic 

form typical of earthflows. Moreover, except for scattered patches of glacial erratics on the oldest flows, 

the deposits consist entirely of fragments of shale, 

sandstone, basalt, and volcanic breccia in a plastic 

bentonitic clay matrix. All these rocks crop out under 

or upslope from the landslides.

STRATIGRAPHY AND STRUCTURE

Upper Cretaceous shales and conglomeratic sandstones of the Judith River Formation dip northward 

under the landslide area. Wilson (1934) measured 

bentonite beds totaling 26 feet in thickness in this 

sequence.

Sepulchre Mountain is built up of acid and basic 

andesitic breccias of Tertiary age (Iddings and Weed, 

1894) that lie unconformably on the Judith River Formation. The basic breccia overlies the acid breccia 

and makes up the bulk of the mountain. The acid 

breccia that crops out at the base of Sepulchre Mountain includes light-colored quartz-bearing tuffs more 

acid than andesite. Some of these tuffs are bentonitic. 

Other exposures in the Gardiner area reveal bentonite 

on the erosional unconformity at the base of the acid

breccia. This bentonite may have been reworked from 

the Judith River Formation.

Remnants of highly weathered basalt flows of Tertiary age lie on Judith River beds under the bentonite 

at the base of the acid breccia.

BLOCK GLIDE AND SLUMPS

Large straight fissures transect the ridges of Sepulchre Mountain, marking pull-away zones of glide 

blocks. The largest of these fissures (J., fig. 182.2) 

is 2,700 feet long, up to 40 feet deep, and 40 to 60 

feet wide. Slickensides in basic breccia of the fissure 

walls indicate that a few feet of downthrow on the 

north block preceded separation.

Large slump blocks lie transverse to the main ridges. 

Sag ponds, fissures, and scarplets characterize the 

slump topography. The original ground surface has 

been rotated by a slump movement to form steep backslopes on the blocks. .Contacts between the Judith 

River, acid breccia, and basic breccia appear haphazardly on the faces of slump blocks. Slope wash and 

forest cover prevent detailed mapping of the intricate 

glide and slump structures.

EARTHFLOWS

The earthflows consist of a heterogeneous assortment 

of shale, sandstone, conglomerate, volcanic breccia, and 

basalt. Fragments range from 10-foot blocks to clay 

size. Much of the fine matrix in the earthflows weathers with the puffy "popcorn" texture characteristic of 

bentonitic shales. Proportions of these constituents 

vary from one flow to another. In most of the earthflows the dominant rocks are basic breccia, and shale 

and sandstone of the Judith River Formation. In the 

westernmost flow, however (5, fig. 182.2), the coarse 

fraction is almost entirely acid breccia. This flow is 

the best preserved and apparently was the most fluid 

of the postglacial earthflows. It might be termed a 

mudflow.

Interglacial earthflows probably preceded the last 

glaciation of the valley. They consist of the same 

heterogeneous materials as later flows but lack definite flow structure and have less hummocky surfaces. 

Scattered patches of glacial erratics Precambrian

667422 O - 63 - 2
E12 GEOLOGICAL SURVEY RESEARCH 1962

gneiss and schist and Tertiary volcanics appear only 

on the interglacial flows. The greatest concentration 

of erratics is on an earthflow that lies just below a 

remnant of lateral moraine. Clearly the moraine is 

younger than the earthflow, for it lies between the 

flow and its source area. The topography of the 

earthflow was smoothed and the internal structure was

blurred by the glacial action, but in general the plastic, cohesive nature of the flow deposits inhibited 

plucking and offered resistance to glacial erosion.

Postglacial earthflows (fig. 182.3) have concave 

slump-excavated heads, lateral ridges and troughs, and 

spatulate, convex, pressure-ridged toes typical of large 

complex earthflows. Their outlines are sharp, and

00'

1MJLE

.CONTOUR INTERVAL 80 FEET 

DATUM IS MEAN SEA LEVEL

FIGURE 182.1.' Geologic map of Gardiner landslide area. Units of Quaternary age: Qfs, historic earthflows; Qbgs, block glide 

and slump; Qal, alluvium; Qf2j postglacial earthflows; Qg, lateral moraine; Qfj, interglacial earthflows. Tertiary age: Tp, 

andesite porphyry sill; Tb, basalt. Late Cretaceous age: Kjr, Judith River Formation. Hachured areas represent blockglide fissures. Long-dashed lines, approximate contacts; short-dashed lines, inferred contacts.
AETICLE 182 E13

their hummocky topography is well preserved. They 

cut across and incorporate portions of the older interglacial flows, and are in turn cut by small historic 

flows.

Historic flows are recently reactivated parts of older 

flows. They all bear signs of recent movement: fresh 

pull-away scarps, sharp new lateral ridges and troughs,

Horberg, 1940) deepened the valley and oversteepened 

slopes. Retreat of the glacier removed lateral support 

from materials weakened by frost action. Thawing 

of permafrost may have reduced stability of slopes 

and released water for lubrication. Most of the Judith 

River rocks dip northward toward the valley. Block 

gliding and slumping probably started along bentonite

FIGURE 182.2. Aerial photograph of Gardiner landslide area. A, block-glide fissure. B, westernmost earthflow. (Photograph

GS-JA 1-19.)

fresh transverse and radial cracks, tilted trees, exposed beds in this sequence, and along the unconformity at

tree roots, and bulging and splitting of turf. Most the base of the acid breccia. Block glide and slump

are in watercourses or below flow-impounded ponds. fissures provided new access for water, which faciliFACTORS FACILITATING MOVEMENT tated further sliding and led to disruption of glide

Bentonite beds, dip of rocks, glaciation, and earth- -blocks and slumps into earthflows. Gardiner has been

quakes were contributing causes of the Gardiner land- an active seismic area; earthquakes probably triggered

slides. Advances of Yellowstone Glacier (Weed, 1893; many of the landslides. , , , ..
E14 GEOLOGICAL SURVEY RESEARCH 1962

Sepulchre 

Mountain Electric Peak

Mt Everts

FIGURE 182.3. Panorama southwest across the postglacial earthflows. Symbols match those on geologic map.

cut in bedrock of the Judith River Formation predate the flows.

Stream terraces

BEFEBENCES

Horberg, Leland, 1940, Geomorphic problems and glacial geology of the Yellowstone Valley, Park County, Montana: 

Jour. Geology, v. 48, no. 3, p. 275-303.

Iddings, J. P., and Weed, W. H., 1894, Description of the Livingston quadrangle: U.S. Geol. Survey Geol. Atlas, Folio 1.

Weed, W. H., 1893, The glaciation of the Yellowstone Valley 

north of the Park: U.S. Geol. Survey Bull. 104, 41 p., 

4 pis.

Wilson, C. W., 1934, Section of Paleozoic and Mesozoic rocks 

measured at Cinnabar Mountain, Park County, Montana 

and at Mount Everts, Yellowstone National Park, Wyoming: Am. Assoc. Petroleum Geologists Bull., v. 18, no. 3, 

p. 368-369.

STRUCTURAL GEOLOGY

183. THE CROOKED CREEK DISTURBANCE, SOUTHEAST MISSOURI

By THOR H. KIILSGAARD, ALLEN V. HEYL, and MAURICE K. BROCK, Washington, B.C., Beltsville, Md.,

Denver, Colo.

The Crooked Creek disturbance, in southeast Missouri, about 7 miles south of Steelville, is one of a 

number of similar geologic structures in the central 

part of the United States which have many points in 

common. These structures are essentially circular 

areas as much as a few miles in diameter that are 

peripherally bounded by high-angle faults, within 

which are highly deformed and brecciated rocks. The 

central part or core of the deformed area in many 

structures consists of older rocks that have been uplifted above their original positions. The origin of 

the structures has been a subject of conjecture, but 

two concepts have been most commonly advanced: 

subterranean explosion and meteoritic impact.

^ Rolla

Modified from Geologic Map of Missouri 

Division of Geological Survey and Water 

Resources. 1961, and from unpublished 

structural map of the base of the Roubidoux Formation compiled by Mary H. 

McCracken

10 20 30 MILES

FIGURE 183.1. Major faults in the Crooked Creek and Decaturville areas.
ARTICLE 183 E15

The authors visited the Crooked Creek area in November 1961, during the course of studies of lead-zinc 

deposits in the central region of the United States. 

Previously Heyl and Brock had examined several similar structures, including Serpent Mound, Ohio (Heyl 

and Brock, 1962); Jeptha Knob, Ky.; the diatremelike disturbance at the Soward farm, west of Rosiclare, 111.; and several diatremes in the vicinity of 

Avon, Mo. Following the visit to Crooked Creek, 

Heyl and Brock also examined the Decaturville disturbance, Missouri (fig. 183.1), the Kentland disturbance, Ind., and the Holloford crater, Ontario.

GEOLOGIC FEATURES OF THE DISTURBANCE

The Crooked Creek area is underlain by sedimentary rocks of Cambrian and Ordovician age, which dip

a few feet per mile to the northwest. The disturbance 

is at the west end of the major Palmer fault zone and 

about 10 miles south of the known south end of the 

major Cuba fault. It is a complex ovate structure 

about 3 miles wide and 4 miles long, bounded by a 

connected series of peripheral normal faults. Within 

the faulted boundary is an encircling synclinal graben 

of Roubidoux and Jefferson City Formations, inside 

of which is a peripherally faulted, uplifted central 

core of older dolomite and shale (fig. 183.2). Hendriks (1954, p. 53) notes that these older rocks are 

about 1,000 feet above their normal stratigraphic positions. The central circular core is not a simple uplifted dome but rather a partly collapsed dome of 

younger rocks encircled by an anticline of older rocks. 

A north-trending "dog-leg" horst crosses the collapsed

R. 4W. EXPLANATION

I Qal I 

Alluvium

| Ojc~] 

Jefferson City Dolomite

Roubidoux Formation

| Og~| 

Gasconade Dolomite

Ce 

Eminence Dolomite

Potosi Dolomite

I Cdd I 

Derby-Doe Run Dolomite

[ d [ 

Davis Formation

I Cb I

Bonneterre Dolomite 

Br2 $B» #Brl

Missouri Division of 

Geological Survey and 

Water Resources, 1947

1000 0 1000 

I

2000 FEET 

I

Geology by H. E. Hendriks, 

1947 with modifications by 

T. H. Kiilsgaard, A. V. Heyl, 

and M. R. Brock, 1961

Intrusive breccia

ftumber refers to area described in text

Contact 

D ___

Fault

Dashed where approximately located, 

dotted where concealed; U, upthroum 

side; D, downthrown side

*~ Folds 

Anticline Syncline

Strike and dip of beds

Mine or prospect

FIGURE 183.2. Geologic map of the central part of the Crooked Creek disturbance showing intrusive breccias, mines, and prospects.
E16 GEOLOGICAL SURVEY RESEARCH 1962

part of the core, and a number of radial faults extend 

outward from it. Intricately contorted shale of the 

Davis Formation crops out along the crest of the anticline. Alined breccia zones in the core area suggest a 

more complex fault pattern than is shown on the map.

BRECCIAS

Sedimentary rocks in the central uplifted core of 

the disturbance are faulted, intensely shattered, and 

brecciated. Most of the faults are of small displacement but are marked by alinement of brecciated zones, 

or by juxtaposition of tilted or contorted rocks of different lithologies. Two types of breccias are recognized : shatter and intrusive. The shatter breccias are 

far more common and may be seen in most of the central area. Some shatter breccias consist of rocks of 

the same lithology fractured in diverse directions, but 

without rotation of the multisized undisplaced fragments. Equally abundant, however, are shatter breccias in which angular to subangular fragments of the 

same lithology have been dislodged and rotated so 

that all trace of bedding is obliterated. The shattered 

and pulverized fragments are imbedded in a lithified 

matrix of ground rock and microbi sccia, of uniform 

lithology.

Intrusive breccias are pipe-like masses of angular to 

subrounded fragments, imbedded in a pulverized 

groundmass, that have been injected or intruded into 

overlying rocks. The intrusive breccia exposed at localities 1 and 2 (fig. 183.2) may be a single mass. It 

consists largely of coarsely crystalline, nondrusy lightto pinkish-gray dolomite fragments intruded into 

drusy fetid pale-brown Potosi Dolomite, and, in part, 

into more finely crystalline light-gray dolomite 

mapped as Derby-Doe Run Formation. The breccia 

fragments are dolomite characteristic of the Davis 

and Bonneterre Formations. The groundmass of the 

breccia, which forms the bulk of the rock, is gray to 

buff dolomite microbreccia and pulverized argillaceous 

rock flour. Chert, druse, and quartz crystals so typical of the Potosi Dolomite were not seen in the intrusive breccia. At localities 3 and 4 breccia has intruded the Bonneterre and Davis Formations, 

respectively. In these places it consists of fragments 

of coarsely crystalline gray dolomite and sandstone 

in a matrix of cemented sand grains, small angular 

fragments of dolomite, and pulverized argillaceous 

rock flour. The sandstone is lithologically similar to 

the Lamotte Sandstone, which underlies the Bonneterre Dolomite in southeast Missouri, and the fragments probably came from the Lamotte. In places 

the dolomite fragments have been leached out, and, 

at locality 4, barite crystals line some of the dolomite 

cavities.

SHATTER CONES

Shatter cones occur at several localities in the 

strongly brecciated central part of the disturbance. 

They are radially striated blunt pressure cones with 

large apical angles and range from y% to 4 inches in 

height. The cones shown in figure 183.3.4 are from 

the core area and are typical of the larger ones seen, 

both in size and form. Shatter cones of similar size 

and appearance (fig. 183.35) were found in the Decaturville disturbance, except that they were sharper 

and had smaller apical angles than the cones from 

Crooked Creek.

FIGUKE 183.3A, Shatter cones in Potosi Dolomite, from road 

cut in Crooked Creek disturbance. Cone A was inclined 

downward at 60°, cone B inclined upward. Cone A is a rough, 

blunt conical structure with dimple in center. Apex of partly 

obscured cone B is at point of arrow. B, Cluster of sharp 

shatter cones from lead prospect at center of Decaturville 

disturbance, which are similar in appearance to many of those 

seen at Crooked Creek.
ARTICLE 183 E17

The shatter cones found by Hendriks (1954) are 

exposed in a streambed a few hundred feet east of the 

fork in the road along the west line of sec. 17 (fig. 

183.2). Most of the cones in this area point upward, 

as Hendriks described them. Near the west side of 

the locality the axes of the cones trend about N. 50° 

E., and the apexes point upward 10° to 20° above the 

horizontal. About 100 feet farther to the east the 

axes trend about N. 15° W. Countless cones were 

seen in other outcrops in this streambed over a distance of about a thousand feet, and the orientation 

of the cones was found to vary considerably from one 

locality to the next, with respect both to inclination 

to the horizontal plane and to the plane of bedding. 

At most places the more readily apparent cones stand 

out in relief from the fairly flat bed of the stream, 

the apexes pointing upward, but close inspection reveals that almost as many cones point downward. 

Commonly 2 or more cones point in different directions within the same hand sample, and in several 

samples 2 cones were observed base to base, the apexes 

pointing in opposite directions (fig. 183.3^4.).

Numerous shatter cones were found in highly brecciated Potosi Dolomite in fresh roadcuts between the 

fork of the road on the west line of sec. 17 and the 

southwest corner of sec. 17. Here most of the cones 

point downward, but cones having other orientations 

also are common.

MINERALIZED AREAS

Several places in the central part of the Crooked 

Creek disturbance are mineralized by galena, sphalerite, marcasite, barite, and vein quartz. Hendriks 

(1954, p. 77) states that several thousand tons of 

barite was mined from residual red clay, thin stringers, 

and veins and lenses in the Davis Formation in the 

SE1/4SE1/4NE1/4 sec. 18 (fig. 183.2). Barite crystals 

line vugs and fill veinlets in sandstone breccia 100 

yards northwest of these diggings.

A small tonnage of galena was mined about 100 

years ago in the SW^NWi/4 sec. 17 (fig. 183.2) from 

a cluster of shallow shafts called the Metcalf diggings (Winslow, 1894, p. 685). The lead was mined 

from residual clay and dolomite bedrock of the upper 

Bonneterre Dolomite. Some prospecting, and perhaps 

some mining also, is shown by the several shallow 

shafts in the breccia body at the west edge of the 

Metcalf diggings.

Galena, sphalerite, barite, pyrite, and marcasite are 

deposited in and along a north-trending breccia zone 

in the SE corner of the NEi/iSW^, sec. 17 (fig. 

183.2). The minerals are associated with and cemented by veinlets of milky quartz that cut and replace bleached and brecciated Potosi Dolomite. The 

sulfides are well crystallized and coarse grained. Individual crystals and grains of galena, yellow-brown 

sphalerite, and marcasite range from 1 to 3 cm in 

maximum dimension; plates of white barite are as 

much as 6 cm long. In general the mineral characteristics closely resemble those of minerals in the 

barite-lead deposits in the area near Palmer, Mo. 

These deposits are largely in residual clays derived 

from weathering of the Potosi and Eminence Formations, but they are concentrated near the Palmer 

fault zone.

None of the sulfides or the barite in any of the 

mineralized occurrences that were examined were brecciated or broken by faulting. Their unfractured condition, plus the fact that at the deposit in the SWi/4 

of sec. 17 the sulfides are in quartz that has replaced 

and cemented brecciated dolomite, clearly demonstrates that sulfide mineralization occurred after brecciation and fault movement had ceased. Indications 

are that fractured and brecciated areas provided channelways for the mineralizing solutions as well as favorable sites for mineral deposition. The deposits 

are epigenetic and can be dated as post-Crooked Creek 

disturbance in age.

ORIGIN OF THE DISTURBANCE

The disturbance was first studied by Hughes (1911, 

p. 48-52), who described it as an area of sharp folding 

accompanied by faulting. Hendriks (1954, p. 52-73) 

mapped it and discusses four ways in which it could 

have originated: salt-dome intrusion, igneous intrusion, subterranean explosion, and meteoritic impact 

and explosion. He discards the salt-dome concept because salt deposits are unknown in the subsurface of 

southeast Missouri, and discards the igneous-intrusion 

possibility because no intrusive igneous rocks have 

been found in the disturbance, either at the surface 

or in any of the several holes drilled in the area. He 

surmises that either a subterranean explosion or meteoritic impact could have formed the disturbance but 

concludes that orientation of shatter cones and bilateral symmetry caused by maximum and minimum depression of the encircling syncline are best explained 

as having originated by meteoritic impact.

Fox and others (1954, p. 1252-1253) attribute the 

disturbance to igneous activity connected with renewed 

movement along the Palmer and Cuba faults. They 

conjecture that sea water from a covering epicontinental sea descended along the open fractures, was heated 

by the magma, and caused gaseous explosions that 

fractured the overlying rocks. Fragments of igneous
E18 GEOLOGICAL SURVEY RESEARCH 1962

rock were prevented from reaching the surface by 

weight of the sea water and overlying rocks.

The preponderance of evidence reviewed by the authors favors an origin of the disturbance by subterranean explosion at the intersection of the Cuba and 

Palmer faults. The Palmer fault zone has been 

mapped eastward from the disturbance for about 36 

miles, to a point where the fault branches. The north 

branch continues to the northeast and is believed to 

merge with the great southeast-trending Ste. Genevieve fault system, which, in turn, merges with other 

fault systems extending east to the Appalachians 

(Heyl and Brock, 1961, fig. 294.1). The Cuba fault, 

which has been traced for more than 30 miles, would 

intersect the disturbance area if it continues on strike 

from its southernmost exposures (fig. 183.1). Hendriks did not find the fault immediately north of the 

disturbance, but most of this area is coArered by deep 

residuum and dense forest cover. Dake (1930, p. 

182) recognized a southward continuation of the Cuba 

fault and suggested that it might be a continuation 

of the Palmer fault. Hayes (1962), although not referring to the Cuba or the Palmer faults in his map 

of the configuration of the Precambrian surface, recognizes the probability that northwest- and northeasttrending structural lineaments pass through the disturbance area.

Hendriks (1954, p. 72) suggested that the Palmer 

fault formed after deformation of the Crooked Creek 

area, using as evidence the termination of the fault 

at the disturbance. There is, however, a good possibility that the Palmer fault extends west of the 

disturbance, for south of Holla (fig. 183.1) a fault 

zone along the projected strike of the Palmer fault 

could very well be a part of the latter zone. The 

great known extent of the Palmer fault zone and of 

other zones with which it is connected indicate a regional zone of weakness along which fault movement 

has been intermittently active, possibly since Precambrian time. Precambrian faulting also is suggested 

by Hayes' (1962) map of the configuration of the 

Precambrian surface, which shows a pronounced westtrending regional ridge extending more than 40 miles 

east of the Crooked Creek disturbance and more than 

20 miles to the west, whose steeper north side coincides very closely in location to that of the Palmer 

fault zone. Near Ironton, Mo., where Precambrian 

rocks crop out, several alined, elongate ridge fronts 

are considered to be ancient fault-line scarps, and the 

same could hold for the buried west-trending escarpment passing through the Crooked Creek area.

The plausibility of structural control of the Crooked 

Creek disturbance also is influenced by consideration

of the somewhat similar Decaturville disturbance, 

about 70 miles to the west. This latter disturbance 

is an uplifted central core of highly brecciated sedimentary rocks that is peripherally bounded by highangle faults. Brecciated dolomite in the core area 

is mineralized by galena, sphalerite, pyrite, marcasite, 

and reportedly (H. B. Hart, oral communication,

1961) by rare earths, yttrium, niobium, and germanium. In the central part of the core is a body of 

granite pergmatite, probably of Precambrian age, 

surrounded by a mylonitic sericite schist zone. Krishnaswamy and Amstutz (1960) describe the disturbance 

as resulting from tectonic movement and associated 

igneous activity. The dome is along or very near the 

axis of the Proctor anticline (Marbut, 1908, p. 61; 

Mary McCracken, Missouri Division of Geological 

Survey and Water Resources, written communication,

1962). The Decaturville disturbance also is at or very 

near the intersection of the anticlinal axis and the 

southeast projection of a major southeast-trending 

fault shown a few miles west of Camdenton on the 

geologic map of Missouri by the Missouri Division 

of Geological Survey and Water Resources (1962). 

A remote but nevertheless interesting structural possibility is that the dome also is affected by an intersection of the Proctor anticline and the westward 

continuation of the Palmer fault zone.

The continuity of the narrow "dog-leg" horst (fig. 

183.2) extending intact through the collapsed central 

part of the shattered and brecciated core area of the 

Crooked Creek disturbance most certainly would have 

been completely destroyed had the disturbance formed 

by meteoritic impact. On the other hand, the structural pattern displayed in the disturbance has its 

counterpart in many calderas and fractured domes of 

undisputed vocanic origin (Wilbur Burbank, oral 

communication to Heyl and Brock, 1961). The bounding faults of horsts in these areas can best be explained as having formed first as radial tension fractures which develop during upwelling of the whole 

structure. With escape of the pent-up gases or molten 

material, chiefly through the central area of the dome, 

the central part subsided into the resulting void. 

Often an unbroken arch of rock is preserved between 

essentially parallel sets of radial fractures. These 

blocks are arched upward and they tend to contain 

rock of older age than the subsided rock on either 

side. Such is the case of the horst in the Crooked 

Creek disturbance. Its presence favors origin by explosion from below rather than origin from an explosion resulting from meteoritic impact.

Hendriks (1954, p. 54) observed that cones he had 

seen in the Crooked Creek disturbance pointed up-
ARTICLE 183 E19

ward, and from this he postulated that explosion occurred above the level now exposed. He used upward orientation of the cones to support a meteoritic 

origin for the disturbance.

Branca and Fraas (1905), in their work on the 

Steinheim Basin in southern Germany, first recognized conical features as pressure phenomena formed 

by explosion. In describing the Wells Creek and 

Kentland structures, Bucher (1933) reported shatter 

cones and emphasized their significance as high-pressure indicators. More recently shatter cones have 

been found in other peculiar circular structures described by Bucher as cryptovolcanos, including the 

Serpent Mound, Flynn Creek, and Decaturville disturbances. Dietz (3961), in his studies on the genesis 

of shatter cones by impact and explosion, states that 

apexes of the cones point in the direction of the highpressure source.

In the Crooked Creek area the sedimentary rocks 

were at least as flat lying before deformation as they 

are today. If the structure formed from meteoritic 

impact it therefore can be assumed that a relatively 

constant angle should have formed between inclined 

apexes of the shatter cones and the bedding of the 

country rock. Such is not the case, however, for the 

cones are oriented in all directions and display no 

uniformity with respect to bedding.

Diverse orientation of shatter cones suggests multiple centers of impact and explosion within the central part of the Crooked Creek disturbance, rather 

than pressure from a single source as should result 

from meteoritic impact. Multiple orientation of cones 

within relatively small breccia blocks implies that 

pressure acted on the blocks from more than one 

direction. The cones apparently are pressure phenomena, but whether the pressure was generated from 

above or from below, or both (fig. 183.3J.), is debatable.

Origin of the disturbance from an underlying explosion also is indicated by the presence of the intrusive breccia in which the fragments have been 

forced upward through the overlying rocks. Identical 

intrusive breccias are to be seen at other diatremelike disturbances related in origin to tectonic forces 

and igneous activity. A good example is the breccia 

at the diatreme on the O. K. Cash farm, southeast of 

Avon, Mo. This diatreme contains not only intrusive 

breccia but also intruded igneous rocks, barite, and 

fluorspar. The oldest of the several types of breccia 

at the Decaturville disturbance is identical to intrusive 

breccia at Crooked Creek. At the Soward farm, west

of Rosiclare, 111., a diatreme-like disturbance contains 

intrusive breccia and brecciated igneous rocks. Similar sandstone and dolomite breccias, in many of which 

the matrix is composed only of ground and relithified 

sedimentary rocks, occur in the Montreal area, Canada, 

and are described by Clark (1952, p. 100-105) as 

diatreme breccias.

In summary, location of the Crooked Creek disturbance at the intersection of regional faults, presence of intrusive breccia identical to intrusive breccia 

associated with known diatremes, presence of postbreccia sulfide mineralization, and similarity of its 

structure to that of other diatreme-like disturbances 

all suggest that the Crooked disturbance originated 

from a subterranean gaseous explosion.

BEFEBENCES

Branca, W. and Fraas, F., 1905, Das Kryptovulkanische Bechen 

von Steinheim; K. Preuss. Akad. Wiss. Abb., Berlin, 

p. 1-64.

Bucher, W. H., 1933, Cryptovolcanic structures in the United 

States: Internat. Geol. Gong., 16th, London 1933, Kept, 

of the United States v. 2, p. 1055-1084.

Clark, T. H., 1952, Montreal area, Laval and Lachine areas: 

Province of Quebec Dept. of Mines, Geol. Surveys Branch, 

Geol. Kept. 46, 159 p.

Dake, C. L., 1930, Geology of the Potosi and Edgehill quadrangles: Missouri Bureau of Geology and Mines, v. 23, 

2d ser., 230 p.

Dietz, R. S., 1961, Astroblemes: Sci. Am. v. 205, no. 2, p. 50-71.

Fox, J. H., Alien, V. T., and Heinrich, Ross, 1954, Crooked 

Creek "Cryptovolcanic" structure, Steelville, Missouri 

[abs.] : Geol. Soc. America Bull., v. 65, no. 12, pt. 2, 

p. 1252-1253.

Hayes, W. C., 1962, Map showing configuration of the Precambrian surface: Missouri Geol. Survey and Water Resources.

Hendriks, H. E., 1954, The geology of the Steelville quadrangle: Missouri Geol. Survey and Water Resources, v. 36, 

2d ser., 82 p.

Heyl, A. V., and Brock, M. R., 1961, Structural framework of 

the Illinois-Kentucky mining district and its relation to 

mineral deposits: Art. 294 in U.S. Geol. Survey Prof. 

Paper 424-D, p. D3-D6.

Heyl, A. V., and Brock, M. R., 1962, Zinc occurrence in the 

Serpent Mound structure, Ohio: Art. 148 in U.S. Geol. 

Survey Prof. Paper 450-D, p. D95-D97.

Hughes, V. H., 1911, Reconnaissance work: Missouri Geol. 

Survey and Mines Bienn. Report for 1909-10.

Krishnaswamy, D. S., and Anistutz, G. C., 1960, Geology of the 

Decaturville disturbance in Missouri [abs.]: Econ. Geology, v. 55, no. 6, p. 1340.

Marbut, C. F., 1908, The geology of Morgan County: Missouri 

Bureau of Geology and Mines, v. 7, 2d ser., 94 p.

Missouri Division of Geological Survey and Water Resources, 

1962, Geological map of Missouri.

Winslow, Arthur, 1894, Lead and zinc deposits: Missouri Geol. 

Survey, v. 6-7.
E20 GEOLOGICAL SURVEY RESEARCH 1962

184. ORIGIN OF THE MIDDLESBORO BASIN, KENTUCKY 

By KENNETH J. ENGLUND and JOHN B. ROEN, Washington, D.C.

Work done in cooperation with the Kentucky Geological Survey

The Middlesboro Basin is the unique development 

of a circular, alluviated, topographic depression in a 

part of the Appalachian Plateaus that is characterized by rugged hillsides, narrow V-shaped valleys, 

and a minimum of flat bottom land. The relatively 

broad expanse of the basin with its low, rolling hills 

contrasts sharply with the fringing mountains which 

rise 1,000 to 1,900 feet above the valley floor. Nearly 

all the lowland, about 4 miles in diameter, is occupied 

by the city of Middlesboro. The conspicuously different topography and the intense deformation of 

rocks underlying the basin have attracted geologic 

interest since the early investigations of Ashley and 

Glenn (1906) and Eich (1933). They described the 

physiography of the basin and noted the deformed 

condition of the underlying rocks. This deformation 

was attributed by them to collapse or subsidence of an 

arch following relatively recent movement along faults 

associated with the Cumberland overthrust sheet in 

which the basin is situated. Recent geologic mapping 

does not substantiate their theory of tectonic origin, 

but indicates that meteoric impact is the probable 

mode of origin.

The rocks exposed in and adjacent to the Middlesboro Basin are the Lee Formation and part of the 

overlying Breathitt Group, both of Pennsylvanian 

age. The Lee consists mostly of thick-bedded to 

massive quartzose sandstone and quartz-pebble conglomerate in contrast to less competent shale, siltstone, 

sandstone, and coal of the Breathitt Group. Detailed 

mapping and correlation of thin stratigraphic units 

show that around the periphery of the basin the 

Breathitt Group is relatively undisturbed. A shallow 

syncline surrounds the basin and, locally, low-amplitude anticlines separate the syncline from the more 

intensely deformed rocks in the basin. Mining in the 

Hance and Mason coal beds in the syncline surrounding the basin does not reveal unusual or intense 

deformation.

The outer edge of the intensely deformed rocks of 

the Breathitt Group is delineated in most places in 

the Middlesboro Basin by normal faults with nearly 

vertical fault planes. Faults are most conspicuous in 

the northwest part of the basin where brecciated sandstone is abundant in several fault slices. In addition 

to the faults shown in figures 184.1 and 184.2, deformation consists of numerous small unmapped faults, 

fractures in sandstone, and tight folds in shale. Many 

of these minor structures and the attitude of beds 

show circular trends around the center of the basin. 

Toward the center of the basin the strike of the beds 

becomes more variable and the dip increases to vertical or overturned.

At the center of the basin, the core of this circular 

structure consists of conglomeratic sandstone of the 

Lee Formation which appears to be bounded by steeply 

inclined normal faults. Deformation is most intense 

here, with the result that much of the bedding is 

obscured and individual quartz grains are generally 

shattered. Fractures are commonly confined to the 

individual subrounded to rounded quartz grains and 

do not pass from grain to grain. Parallel arrangement of fractures in the grains is common, and in 

some grains different sets of parallel fractures intersect to form a rhombohedral pattern. Many angular 

quartz grains, in the silt to very fine sand size, probably resulted from the brecciation of larger grains.

Discovery of previously unrecognized Lee Formation in the center of the Middlesboro Basin shows 

that rocks underlying the middle of this depression 

are structurally high, a significant factor that rules 

out the earlier interpretations which attributed the 

structure to an adjustment collapse resulting from 

movement along the Pine Mountain overthrust fault 

and Eocky Face fault. A tectonic origin, possibly 

associated with regional structure, is further questionable as shattered quartz grains similar to those 

in the core 9f the basin have not been found in 

petrographic examination of intensely deformed Lee 

sandstone from the principal fault zones of the Cumberland overthrust block.

The character of the rock deformation and the associated structural relations, as disclosed by recent geologic mapping and as interpreted by the authors, 

indicates that the Middlesboro Basin may be an ancient meteor impact scar. Several aspects of the basin, 

such as the nearly circular shape, are particularly 

more typical of impact structures than structures 

commonly associated with Appalachian folding and 

faulting, which tend to be elongate or linear except 

for fensters which are roughly circular in outline. 

However, known fensters in nearby areas (Miller and
Qw
E22 GEOLOGICAL SURVEY RESEARCH 1962

FIGURE 184.2. Generalized section along line A A' of figure 184.1.

Fuller, 1954) contain exposures of younger rocks 

framed b}T older rocks of the thrust sheet, a relation 

not observed in the Middlesboro Basin. The possibility that the disturbance exhibited in the rocks of 

the Middlesboro Basin is caused by the proximity of 

a thrust fault slightly below the basin floor is not 

substantiated by oil and gas test wells drilled within 

4 miles of the basin. These wells show that the Pine 

Mountain overthrust fault is at a depth of about 

2,000 feet below sea level, and they do riot indicate a 

rising trend in the fault plane. Additional evidence 

of meteor impact is indicated by the structurally 

high and intensely shattered core, which may be the 

result of rebound from an impact explosion. Diminishing deformation outward from the center, a marginal syncline, and upturned beds on the periphery 

of the basin are also suggestive of meteor craters. 

Because the basin is completely breached by drainage,

it can be expected that many of its original physiographic features have been modified or obliterated by 

erosion. Silica glass and meteorite fragments are yet 

to be found, but these could have been eroded from 

the area or concealed by extensive lowland alluviation. 

If a meteoric origin, as postulated here, should be 

firmly established the Middlesboro Basin would serve 

as a clue to the structure underlying relatively recent 

depressions of similar origin.

REFERENCES

Ashley, G. H., and Glenn, L. C., 1906, Geology and mineral 

resources of part of the Cumberland Gap coal field, Kentucky : U.S. Geol. Survey Prof. Paper 49, 239 p.

Miller, R. L., and Fuller, J. O., 1954, Geology and oil resources 

of the Rose Hill district the fenster area of the Cumberland overthrust block Lee County, Virginia: Virginia 

Geol. Survey Bull. 71, 383 p.

Rich, ,T. L., 1933, Physiography and structure at Cumberland 

Gap: Geol. Soc. America Bull., v. 44, no. 6, p. 1219-1236.

185. THRUSTING DEVELOPED AFTER FOLDING IN THE VALLEY AND RIDGE PROVINCE,

SOUTHWEST VIRGINIA

By LEONARD D. HARRIS, Knoxville, Tenn.

Eich (1934) concluded from his study of the Cumberland thrust block of southwest Virginia, southeast Kentucky, and north-central Tennessee that major 

thrust faults developed as bedding-plane thrusts in 

incompetent zones early in the erogenic cycle, when 

the rocks were little deformed. These faults tended 

to remain as bedding-plane thrusts until frictional 

resistance impeded their growth at depth. They

then crosscut or ramped diagonally upward to higher 

and younger incompetent zones, where they again became bedding-plane thrusts. Major folds in the allochthonous plates developed as a result of subsequent 

movement. Thus, lie assumed that the major folds 

of the area were rootless structures that had no counterparts below the thrust fault.
ARTICLE 185 E23

About 25 miles northeast of the area Eich discusses, 

detailed mapping in the Duffield (Harris and Miller, 

1958) and Stickleyville quadrangles has confirmed 

that most of the major faults did develop as beddingplane thrusts in incompetent zones. However, an 

analysis of the relation of thrusts to major folds in 

the area does not support the concept that faults 

developed early in the erogenic cycle. Rather, thrusting seems to have been initiated after folding.

Structure in the Duffield and Stickleyville quadrangles is characterized by a series of northeast-trending folds and thrust faults. The major structures, 

from northwest to southeast, are the Cedars syncline, 

Sandy Eidge anticline, Wallen Valley thrust fault, 

Powell Mountain anticline, Red Hill thrust fault, Red 

Hill anticline, Hunter Valley thrust fault, Rye Cove 

syncline, and Clinchport thrust fault (fig. 185.1). The 

Wallen Valley, Hunter Valley, and Clinchport faults 

are regional thrust faults, having displacements on 

the order of 2 miles or more and continuous traces of 

90 miles or more. The Red Hill thrust has a displacement of about three-quarters of a mile and a trace of 

about 20 miles.

Over much of the western part of the area the 

Wallen Valley fault is a bedding-plane thrust dipping 

about 40° SE. In the north-central part of the area 

the fault changes from a bedding-plane thrust to 

one that abruptly crosscuts from Upper Cambrian 

rocks on the south limb of the Powell Mountain anticline through the axial region and into the north limb, 

where it again becomes a bedding-plane thrust in 

Upper Ordovician rocks. In the area of crosscut the 

fault dips uniformly 5° SE. (fig. 185.1, section A-A'). 

This uniform dip seems to rule out the possibility 

that the Powell Mountain anticline resulted from 

folding of the thrust plates in later stages of deformation. Nor does it seem' likely, considering the fact 

that the fault apparently dies out about 2 miles northeast of the Duffield quadrangle, that enough drag occurred during movement to warp the rocks of the 

allochthonous plate into a major anticline. The relation suggests that the Powell Mountain anticline was 

largely developed before the Wallen Valley thrust 

fault and that the fold controlled the position of the 

developing fault. Subsequent movement along the 

fault displaced the anticline northwestward from its 

root structures in the autochthonous plate (section 

A-A'}.

The Red Hill thrust fault and the doubly plunging 

Red Hill anticline exhibit on a smaller scale the relation -just discussed for the Wallen Valley thrust and 

the Powell Mountain anticline (section B-B'). However, the complex geology of the Red Hill thrust and 

the Red Hill anticline cannot be conveniently shown 

at the scale of the map accompanying this article. 

For a more detailed treatment of this small-scale 

structure see Harris and Miller (1958).

The development of the Rye Cove syncline before 

formation of the Hunter Valley thrust fault is strongly 

suggested by the manner in which the fault conforms 

to the syncline. The Hunter Valley fault developed 

in older beds on the south limb of the syncline and 

crosscut to younger beds as it passed through the 

downwarped axial region and into the north limb 

(section A-A'}.

The trace of the Hunter Valley thrust fault in 

general trends diagonally through the area from 

southwest to northeast. Near the center of the area, 

however, the trace swings abruptly southward around 

a closely spaced series of fault slices in the autochthonous plate. A longitudinal section (section C-C'} 

through the area of slices shows that the Hunter Valley fault must have arched up TOO to 1,000 feet higher 

in this locality in order to clear the slices. These 

slices probably were in place early in the development 

of the Hunter Valley thrust plate, forming an obstruction. Subsequent movement across the obstruction warped the Rye Cove syncline enough to cause 

the eastern and western parts of the syncline to plunge 

in opposite directions away from the obstruction. An 

additional effect of the obstruction was to impede 

movement of the allochthonous plate to the degree 

that differential movement was initiated. Differential 

movement resulted in the western part of the syncline being skewed northward relative to the eastern 

part. The projected axial trace of the Rye Cove syncline bends around the obstruction in the autochthonous plate.

These structural relations suggest that thrust faults 

developed later in the erogenic cycle than folding. 

Faults apparently were initiated as bedding-plane 

thrusts in the oldest incompetent zones, and followed 

these zones until they encountered preexisting folds, 

where the beds were downflexed. Downflexing inhibited further development at the same level; however, it facilitated crosscutting of the fault to a higher
E24 GEOLOGICAL SURVEY RESEARCH 1962

82° 52'30"

EXPLANATION

Middle Cambrian 

(lower part) and 

Lower Cambrian

Contact

Fault

U, on upthroum side

Major thrust fault

______ 

Middle Cambrian

(upper part)

Thrust fault 

T, on upper plate

Anticline

Showing plunge

 t "

Sync line

Shburing plunge; dashed 

where approximately 

located

HUNTER VALLEY THRUST ,c'

lOOO'-P

o'-C

\ " 

FIGUBE 185.1. Geologic map and structure sections of the Stickleyville and part of the Duffield quadrangle, Virginia. (Geology

of the Duffield quadrangle and section A A' modified from Harris and Miller, 1958.)
ARTICLE 186 E25

level by bringing younger incompetent stratigraphic 

zones at or near the level of the fault. Under these 

conditions, faults do not simply ramp upward along 

steep diagonal shears from older incompetent rocks to 

younger because of increased frictional resistance; instead, they change stratigraphic position simply by 

cutting across preexisting folds.

REFERENCES

Harris, L. D., and Miller, R. L., 1958, Geology of the Duffield 

quadrangle, Virginia: U.S. Geol. Survey Geol. Quad. Map 

GQ-111.

Rich, J. L., 1934, Mechanics of low-angle overthrust faulting 

as illustrated by Cumberland thrust block, Virginia, Kentucky, and Tennessee : Am. Assoc. Petroleum Geologists 

Bull., v. 18, p. 1584-1596.

186. REVERSAL OF THROW ALONG A LINE OF LOW-ANGLE THRUST FAULTING NEAR SAN PEDRO

DE ATACAMA, CHILE

By ROBERT J. DINGMAN, Washington, D.C.

Work done in cooperation with the Instituto de Investigaciones Geoldgicas, Santiago, Chile, and under the auspices of the U.S.

Agency for International Development

San Pedro de Atacama is located in the northern 

part of the Salar de Atacama in eastern Antofagasta 

Province (fig. 186.1). The Quebrada (canyon) Tarnbores drains southeastward into the Llano de Paciencia, which is an extension of the Salar between 

the Cerros de la Sal and the northern part of the 

Cordillera de Domeyko. This area was mapped geologically as an adjunct to the mapping of the Tulor 

quadrangle, most of which is south of the area of this 

article.

The rocks in the Llano de Paciencia area include 

the San Pedro and Tambores Formations of early 

Tertiary age and ash-flow deposits of late Tertiary 

(Pliocene(?)) age. The San Pedro Formation consists of beds of poorly lithified shale, siltstone, finegrained sandstone, salt, and gypsum at least 2,100 

meters thick which interfinger to the west with conglomerate and sandstone of the Tambores Formation. 

Both formations are overlain by a series of ash-flow 

deposits that are intercalated with alluvial sand and 

gravel. The lower ash-flow deposits are thin, poorly 

lithified, and irregular in extent, apparently having 

been extensively eroded before deposition of the upper 

ash flow, which is 30 to 50 meters thick and in some 

localities is welded to a dense hard rock. The ashflow deposits are overlain in and near the borders of 

the Llano de Paciencia by Quaternary piedmont deposits, largely mud-flow material, and by alluvial deposits of sand, clay, and gravel.

70° 69° 68° 67°

TARAPACA 

PROVINCE

BOLIVIA

TULOR Volcan . XT 

QUADRANGLE_ Licancabur/

Sr-,^\

ANTOFAGASTA

'*' '&'#)] ^e Atacama.1 

Salar de^i^ Toconao /

Atacama p : ./ /

/

Siglia

PROVINCE

25

26 ATACAMA

PROVINCE

Tilomonte 7

-¥

0 V

ular

ARGENTINA

100 200 KILOMETERS

FIGURE 186.1. Index map of part of northern Chile.
E26 GEOLOGICAL SURVEY RESEARCH 1962

The sediments of the San Pedro and Tambores 

Formations were deposited in a basin which adjusted 

downward under the accumulating lodd of sediments. 

As a result these formations were gently tilted toward the center of deposition, and gravitational gliding occurred in a southeasterly direction from the west 

side toward the center of the basin. As described 

by Dingman (1962) the Cerros de la Sal, a few miles 

south, are a series of domes, anticlines, and synclines 

produced as a result of this gliding.

The gliding continued even after the Quaternary 

piedmont deposits were laid down, and it produced 

strong horizontal compressive stress on the ash-flow 

beds. The ash-flow deposits in the Tambores area 

are slightly welded, and they reacted to the compression as a relatively brittle and easily fractured 

unit. The fracturing that occurred in the ash flows 

forms a well-developed system of conjugate joint sets 

intersecting at an angle of 28° to 30°. The joint sets 

strike N. 20° E. to N. 30° E. The ash flows in most 

localities are almost horizontal, and the fracture planes 

dip from 13° to 15° to the east and to the west.

Low-angle thrust faulting took place along planes 

parallel to one of the joint sets, generally along the 

set that dips toward the center of the basin. Major 

zones of thrust faulting developed along the western 

border of the Llano de Paciencia and along both the 

western and eastern slopes of the Cerros de la Sal.

The line of thrust faulting west of the Llano was 

studied in detail near the mouth of the Quebrada 

Tambores (fig. 186.2). Along this line most of the 

displacement occurred on planes dipping 13° to 14° 

E., but in at least one locality the fracture system 

dipping 15° W. became the fault plane. The usual 

displacement with the eastern block upthrown is shown 

in figure 186.3A. Two kilometers to the northeast 

along the same line of faulting, the western block 

is upthrown (fig. 186.35). The eastern block is upthrown in the next exposure to the north. As shown 

in the photographs the fault plane is well exposed 

and the relative displacement is clear. Between the 

areas of the eastward-dipping fault plane and the 

area of the west ward-dipping fault plane, the brittle 

ash flow was probably so crushed that fractures or 

fault planes were obliterated. This crushed zone may 

have facilitated erosion to form the broad valleys cut 

through the ash flow in these areas.

It is not considered likely that the eastward- and 

westward-dipping faults represent a single twisted or 

warped fault plane. The faults are probably nearsurface features and may curve toward the horizontal 

to become bedding plane faults at shallow depths.

If the two faults are not a part of the same fault 

plane, then a near-vertical shear zone oriented normal 

to the line of faulting must separate the eastward- and 

westward-dipping faults; however, no evidence of a 

shear zone was observed.

EXPLANATION

Tt, Tambores Formation 

Tsp, San Pedro Formation

Steeply dipping fault

Showing relative horizontal movement

Low-angle thrust fault

Showing dip; dotted where concealed; 

barbs on upper plate

Contact

Dashed where approximately located

Anticline

Overturned anticline

Northern boundary of 

the Tulor quadrangle

A 

Locality at which photograph taken

FIGURE 186.2. Geologic sketch map of the Quebrada Tambores

area.
ARTICLE 186 E27

FIGURE 186.3. Low-angle thrust faulting in the Quebrada 

Tambores. Photographs were taken southward from points 

A and B on geologic sketch map (see map for explanation of

The geology of the area is remarkable in that here 

a brittle superficial bed was subjected to horizontal 

stress. In this area the tuff bed on the southeast side 

of the fault line remained relatively stationary while 

the northeast side moved over or under, because movement of the opposite sense would have required the 

dowiislope block to have glided against the force of 

gravity. Along most of the line, underthrusting was 

dominant and only in the areas shown in figure 186.35 

did overthrusting occur. Underthrusting, overthrusting, and crushing of the tuff at random intervals would

symbols). A, Fault plane dips 14° E. B, Fault plane dips 

15° W.

have been normal under the conditions that prevailed 

at the time of faulting. The author has no explanation for the predominance of underthrusting other 

than that it may be a vagarity of distribution. It 

might also be an index of the variation of the thickness and competence of the ash-flow beds at the line 

of faulting.

REFERENCE

Dingman, R. J., 1962, Tertiary salt domes near San Pedro de 

Atacama: Art. 147 in U.S. Geol. Survey Prof. Paper 

450-D, p. D92-D94.

667422 O - 63 - 3
E28 GEOLOGICAL SURVEY RESEARCH 1962

187. GRAVITY AND MAGNETIC ANOMALIES IN THE NORTHERN OQUIRRH MOUNTAINS, UTAH 

By DON E. MABEY, E. W. TOOKEE, and EALPH J. EOBEETS, Menlo Park, Calif.

The northern Oquirrh Mountains (fig. 187.1), about 

20 miles west of Salt Lake City, Utah, extend from 

the vicinity of the Utah Copper mine at Bingham to 

Great Salt Lake. Several peaks in the northern part 

of the range reach 9,000 feet above sea level, about 

4,500 feet above the surrounding valleys.

10 20 MILES

_I

FIGURE 187.1. Index map of part of north-central Utah showing 

area of gravity map (fig. 187.2). A A', generalized location 

of gravity and magnetic profiles shown in figure 187.3.

The northern Oquirrh Mountains are composed 

mostly of sedimentary rocks of Mississippian to Permian age. Small intrusive bodies cut these rocks near 

Bingham; sedimentary and volcanic rocks of Tertiary 

age crop out on the east side of the range.

The rocks at the north end of the range (Tooker 

and Eoberts, 1961) and their correlatives in the Bingham area (Welsh and James, 1961) are now recognized as two different sequences that are juxtaposed 

along the North Oquirrh thrust fault (Art. 188). The 

thrust fault dips north and strikes east across the 

range south of Nelson Peak. The upper plate is composed of rocks assigned to the Eogers Canyon sequence, which is about 10,000 feet thick. Eocks in this 

sequence are folded along northeast-trending axes and 

are complexly faulted. The plate is broken into two 

blocks by a major north-trending tear fault, the Garfield fault. The lower plate of the North Oquirrh 

thrust is composed of rocks assigned to the Bingham 

sequence, which is about 21,000 feet thick. These 

rocks are folded along northwest-trending axes and 

at Bingham are intruded by granitic stocks. This sequence is correlated with rocks in the upper plate of 

the Charleston thrust, which is exposed in the Wasatch 

Eange and assumed to underlie this area.

Antelope Island, about 8 miles north of the Oquirrh 

Mountains and on the projected trend of the range, 

is largely composed of metamorphosed Precambrian 

and Cambrian rocks which Crittenden (1961) believes 

are autochthonous.

GRAVITY DATA

Only a few measured densities are available for 

rocks exposed in the Oquirrh Mountains and on Antelope Island and concealed at depth. The average density of nine representative samples of Paleozoic rock 

in the Oquirrh Mountains is 2.65 g per cm3 No important density difference was measured for the two 

sequences of Mississippian, Pennsylvanian, and Permian rocks. The average density of the Precambrian 

rocks cropping out on Antelope Island and in the 

Wasatch Eange has not been accurately determined 

but is estimated to be about 2.85 g per cm 3, on the 

basis of measurements made on 2 of the more abundant 

rock types and on densities estimated from descriptions of other units.

A gravity map of the valleys west and north of the 

Oquirrh Mountains by Cook and Berg (1961), which 

suggested that a major gravity anomaly exists in the 

range, has been supplemented by additional gravity 

stations within the Oquirrh Mountains and in the valley on the west side of the range. These gravity data 

are reduced to the Bouguer anomaly assuming the 

standard density of 2.67 g per cm 3 for the material 

in the range. Terrain corrections were made through 

zone N on the Hayford charts. These new data and 

those of Cook and Berg are compiled in figure 187.2 

and contoured at a 5-mgal interval.

The Bouguer anomaly values include corrections 

for large elevation differences and terrain effects, but 

uncertainties in these corrections are introduced by 

possible deviations from the assumed rock density
ARTICLE 187 E29

(2.67 g cm 3 ) used to make elevation corrections and 

from errors in making terrain correction. A density 

change of 0.03 g per cm 3 would change the Bouguer 

anomaly difference between the highest and lowest 

station by a factor of about 2 mgals. The largest 

terrain correction is about 45 mgals, but most are 

much smaller. The maximum error for these corrections is about 10 percent, thus the maximum error 

in the Bouguer anomaly values could be about 7 

mgals; however, most of the stations are believed accurate to about 2.5 mgals, or one-half the contour 

interval on figure 187.2.

Bouguer anomaly values show two features of special interest: (1) a marked increase of values north 

of the surface trace of the North Oquirrh thrust, and 

(2) higher values west of the Garfield fault. Values 

are nearly constant south of the North Oquirrh thrust, 

varying over a range of 5 mgals. At the north end 

of the range the values decrease toward a gravity 

low over Quaternary sediments beneath Great Salt 

Lake.

40°45

40°30

The anomaly values on the south end of Antelope 

Island (Cook and Berg, 1961) are about 18 mgals 

higher than the highest values found in the Oquirrh 

Mountains (fig. 187.3) and increase northward to the 

north end of the island. If a correction is made for 

the Quaternary sediments, the anomaly values continue to rise northward over the interval between 

the Oquirrh Mountains and Antelope Island. The 

total Bouguer anomaly difference between the Oquirrh 

Range south of Nelson Peak and the highest values 

on Antelope Island is about 45 mgals (fig. 187.3). 

Anomalies of this amplitude over pre-Tertiary rocks 

in the Basin and Range province are usually associated with regional topographic features and are believed to indicate isostatic compensation for regional 

topography; they generally are not associated with 

individual ranges as narrow as the Oquirrh Mountains 

(Mabey, 1960).

A major discontinuity in regional topography occurs along a west-trending line at the north ends of 

the Oquirrh and Stansbury Mountains. The elevaEXPLANATION

Tertiary sedimentary rocks

Intrusive rocks

Rogers Canyon Bingham sequence 

sequence

Paleozoic sedimentary rocks

Gravity contours

112-00'

FIGURE 187.2. Bouguer gravity anomaly map of the northern part of Oquirrh Mountains, Utah; contour interval 5 mgals. 

Gravity data are partly from Cook and Berg (1961). Geology modified from Cook (1961).
GEOLOGICAL SURVEY RESEARCH 1962

1200

-140

-150

= -160

-170

-180

Bouguer gravity anomaly

FEET , 

10,000

SEA 

LEVEL

-10,000

-20,000

4 P Nelson Peak

Ant£!£PL!?!^d____ GREAT SALT LAKE ^ */^

_ Precambrian basement ***"' "^ ^ / R°eers Ca"y°n sequence / &g 

L *- >. ^ -^'^

Oquirrh

Bingham sequence

Basement profile-** ^ / S^ " ' ^-"'S ^ 

**^/ ** *«..- "" / CHARLESTON THRUST

Alternate basement profiles <T^^ (depth and relation to basement unknown)

v X *\

?,,,,? MILES ^^ *y^_ ^^Basement profile

Mountains

rocks Intrusive 

A

_

-

FIGURE 187.3. Magnetic and gravity profiles from Antelope Island to the central Oquirrh Mountains with diagrammatic section

showing generalized topography and inferred structure.

tion of the ranges and adjoining valley floors increases 

southward, and near the highest parts of the two 

ranges a small east-trending range nearly connects 

the two. The result is a west-trending regional topographic arch about 40 miles wide close to the westward 

projection of the axis of the Uinta arch. This regional topographic arch is superimposed on a more 

general southward rise in the regional elevations.

A part, but not all, of the gravity difference between 

the Oquirrh Mountains and Antelope Island undoubtedly reflects isostatic compensation of the regional 

topography, but the amount of the gravity difference 

that is attributed to the topography is dependent upon 

the type of compensation that is assumed. If a major 

part of the gravity difference is related directly to 

regional topography, it probably indicates adjustment 

for mass anomalies in the upper crust. This would 

require relatively local compensation but would be 

consistent with the indicated isostatic response to

changes of the water depth in Lake Bonneville as 

reported by Gilbert (1890, p. 365-392).

The south edge of the gravity high in the Oquirrh 

Mountains coincides with the trace of the North 

Oquirrh thrust (fig. 187.3). Because no significant 

density contrast exists between the rock exposed at 

the surface along this thrust, the mass anomaly producing the gravity anomaly presumably involves a 

density interface at depth, such as the surface of the 

basement rock. The gravity data suggest that the 

basement rock north of the trace of the North Oquirrh 

thrust is closer to the surface than it is south of the 

thrust. The gravity anomaly increases northward to 

Antelope Island, reflecting a thinning wedge of sedimentary rocks under the north end of the Oquirrh 

Mountains and Great Salt Lake.

Alternate basement profiles that are consistent with 

the available gravity data are illustrated in the diagrammatic section in figure 187.3. One profile (shown
ARTICLE 187 E31

by dashes) assumes that the gravity anomaly in the 

Nelson Peak area is produced by a displacement of 

the basement rock on the North Oquirrh thrust. The 

alternate profile (shown by dots and dashes) assumes 

that the North Oquirrh thrust does not penetrate basement rock and is unrelated, or perhaps indirectly related, to the gravity anomaly and the inferred relief 

on the basement rock. If a density contrast of 0.15 

g per cm 3 , which is a reasonable estimate, exists 

between the Precambrian basement rock and the 

younger sedimentary rocks, the total gravity difference between the south end of Antelope Island and 

the Oquirrh Mountains south of Nelson Peak can be 

produced by about 25,000 feet of sedimentary rocks 

overlying the basement south of Nelson Peak.

The Garfield fault, which has been interpreted by 

Roberts and Tooker (1961, p. 36-45) as a tear fault, 

is also reflected in the gravity contours. Just north 

of the North Oquirrh thrust near Nelson Peak the 

anomaly values are higher on the west side of the fault. 

This gravity expression decreases northward, and is 

not apparent at the north end of the range. The 

gravity difference across the Garfield fault suggests 

upward displacement of the basement rocks, placing 

them several thousand feet higher on the west side 

of the fault. The mechanism of this displacement is 

not clear, but, if the North Oquirrh thrust penetrates 

basement rocks in this area, lateral movement on the 

Garfield tear fault could produce apparent vertical 

displacement of the basement surface.

MAGNETIC DATA

A magnetic profile (fig. 18Y.3) constructed from a 

regional aeromagnetic contour map of the region shows

a magnetic anomaly at the south end of the profile 

produced by some of the post-Paleozoic intrusive rocks 

in the Bingham area. From this anomaly to the 

north end of the range the magnetic intensity increases 

northward at a normal rate. The absence of any 

magnetic anomaly essentially eliminates the possibility that concealed intrusive bodies may account for 

the gravity anomaly underlying the north end of the 

range. The magnetic gradient immediately north of 

the Oquirrh Mountains is probably produced by a 

contrast in magnetic properties in the basement rock.

REFERENCES

Cook, D. R., ed., 1961, Geology of the Bingham mining district 

and northern Oquirrh Mountains: Utah Geol. Soc. Guidebook to the Geology of Utah, no. 16, pi. 1.

Cook, K. L., and Berg, J. W., Jr., 1961, Regional gravity survey along the central and southern Wasatch Front, Utah: 

U.S. Geol. Survey Prof. Paper 316-E, p. 75-89.

Crittenden, M. D., Jr., 1961, Magnitude of thrust faulting in 

northern Utah: Art. 335 in U.S. Geol. Survey Prof. Paper 

424-D, p. D128-D131.

Gilbert, G. K., 1890, Lake Bonneville: U.S. Geol. Survey Mon. 

1, 437 p.

Mabey, D. R., 1960, Regional gravity survey of part of the 

Basin and Range Province: Art. 130 in U.S. Geol. Survey 

Prof. Paper 400-B, p. B283-B285.

Roberts, R. J., and Tooker, B. W., 1961, Structural geology of 

the north end of the Oquirrh Mountains, Utah, in Cook, 

D. R., ed.: Utah Geol. Soc. Guidebook to the Geology of 

Utah, no. 16, p. 36-48.

Tooker, B. W., and Roberts, R. J., 1961, Stratigraphy of the 

north end of the Oquirrh Mountains, Utah, in Cook, D. R., 

ed.: Utah Geol. Soc. Guidebook to the Geology of Utah, 

no. 16, p. 17-35.

Welsh, J. B., and James, A. H., 1961, Pennsylvanian and 

Permian stratigraphy of the central Oquirrh Mountains, 

Utah, in Cook, D. R., ed.: Utah Geol. Soc. Guidebook to 

the Geology of Utah, no. 16, p. 1-16.
E32 GEOLOGICAL SURVEY RESEARCH 1962 

STRATIGRAPHY

188. COMPARISON OF OQUIRRH FORMATION SECTIONS IN THE NORTHERN AND CENTRAL OQUIRRH

MOUNTAINS, UTAH

By E. W. TOOKER and EALPH J. EGBERTS, Menlo Park, Calif.

The Oquirrh Formation is well exposed in the northtrending Oquirrh Mountains in north-central Utah 

(fig. 188.1), about 20 miles west of Salt Lake City. 

The rocks exposed in the northern part of the range 

differ markedly from those of nearly equivalent age 

span in the central part; these two stratigraphic sequences are separated by the North Oquirrh thrust, a 

major east-west fault that crosses the range about 9 

miles south of Great Salt Lake.

The major characteristics of the rock sequence in the 

north end of the range have been described by Tooker 

and Eoberts (1961), and the rocks in the central part 

by Welsh and James (1961). These features and additional data of the writers for the central part of 

the range are summarized in tables 188.1 and 188.2 

and in figure 188.2. More complete data on the faunas 

in these sections will be published later.

The stratigraphic sections are referred to as rock 

sequences, the Eogers Canyon sequence in the northern part and the Bingham sequence in the central part. 

The Oquirrh Formation in both sequences is divided 

into three members on the basis of gross lithology; a 

lower limestone, a middle cyclical limestone, shale, 

and quartzite, and an upper quartzite. The members 

in the two sequences have some differences in age 

spans, and it is therefore difficult to distinguish them 

on the basis of age. For the present these general 

terms are used informally; later, if warranted, the 

members may be given formal names.

The Oquirrh Formation, which includes the 

Oquirrh (?) Formation at its base (fig. 188.2), of the 

Rogers Canyon sequence has an aggregate thickness of 

about 10,000 feet; its base is placed at the contact of an 

interbedded shale and limestone and thick overlying 

limestone. The lower member, 2,800 feet thick, is 

composed of fossiliferous limestone and sandy, cherty, 

and shaly limestone. The bottom half of the member 

contains a brachiopod, coral, and bryozoan assemblage 

of Late Mississippian age (Helen Duncan and Mackenzie Gordon, Jr., written communication, 1962). 

Beds in the top half contain brachiopods and corals 

of Early Pennsylvanian age. The middle member, 

3,500 feet thick, is composed of cyclically repeated 

sandy, cherty, and shaly limestone, shale, and quartzite and sandstone beds that contain fusulinids and 

brachiopods of Atoka and Des Moines (Middle Pennsylvanian) age (Tooker and Eoberts, 1961). The 

upper member, 3,700 feet thick, is composed of quartzite, sandstone, and thin interbedded limestone in the 

lower part and interbedded dolomite and quartzite in

112° in'

41*

EXPLANATION

Steep fault

Thrust fault

Barbs on upper plate

©

Area underlain by thin 

sequences

Location of sections

1. Mount Raymond sequence

2. Rogers Canyon sequence

3. Bingham sequence

4. Timpanogos sequence

FIGURE 188.1. Map of north-central Utah, modified after 

Crittenden (1959), showing general geographic and geologic 

features and the location of stratigraphic sections shown on 

figures 188.2 and 18S.3.
ARTICLE 188 

TABLE 188.1. Summary of stratigraphy, Rogers Canyon sequence, northern Oquirrh Mountains, Utah

E33

System

Permian

\CARBONIFEROUS 

i

Pennsylvanian 1

i

|

Provincial Series

Leonard

Wolf camp

Des Moines

Atoka

Morrow

Chester

Formation

Park City 

Formation

Oquirrh 

Formation

Oquirrh (?) 

Formation

Unnamed 

formation

Member

Grandeur 

Member

Upper

Middle

Lower

Thickness (feet)

750

3,700

3,500

1,400

1,400

10, 000

300+

Character

Interbedded limestone and shale (42 percent), quartzite (53 percent), and dolomite (5 percent).

Quartzite (85 percent), limestone (8 percent), and 

dolomite (7 percent).

Limestone and shale (72 percent), and quartzite (28 

percent) .

Limestone (100 percent), includes sandy, cherty, and 

shaly limestone varieties.

Shale and limestone.

TABLE 188.2. Summary of stratigraphy, Bingham sequence, central Oquirrh Mountains, Utah

System

Permian

CARBONIFEROUS

Pennsylvanian

Mississippian

Provincial Series

Wolf camp

Virgil

Missouri

Des Moines

Atoka

Morrow

Chester

Formation

Diamond Creek Sandstone of 

Welsh and James (1961).

Kirkman Limestone of Welsh 

and James (1961).

Oquirrh Formation

Manning Canyon Shale

Member

Upper

Middle

Lower

Thickness (feet)

400 ±

1 10, 600

8,700

1,900

21, 200

2 1, 100

Character

Sandstone and quartzite.

Limestone.

Quartzite (90 percent) and limestone (10 percent).

Limestone (55 percent) and 

quartzite (45 percent).

Limestone (95 percent) and 

quartzite (5 percent).

Shale and limestone.

1 Upper half estimated from Welsh and James (1961).

the top part that contain fusulinids and gastropods 

of Wolf camp (Early Permian) age (Tooker and Eoberts, 1961). The Oquirrh Formation is capped by a 

unit assigned to the Grandeur Member of the Park 

City Formation. This unit, which is 750 feet thick, 

comprises a fossiliferous basal limestone that contains brachiopods and gastropods of Leonard (Early 

Permian) age (Tooker and Roberts, 1961), a medial 

phosphatic shale, and an upper dolomite.

The Oquirrh Formation of the Bingham sequence 

aggregates about 21,200 feet in thickness and rests on

2 Gilluly (1932).

the Manning Canyon Shale of Mississippian and 

Pennsylvanian age (Gilluly, 1932). The lower member, 1,900 feet thick, consists of sandy, shaly, and 

fine-grained limestone, and some thin interbedded 

quartzite; it contains brachiopods, corals, and bryozoans of Early Pennsylvanian age (Helen Duncan 

and Mackenzie Gordon, Jr., written communication, 

1962). The middle member, 8,700 feet thick, is composed of cyclical limestone, shale, and quartzite beds 

that contain fusulinids, brachiopods, corals, and bryozoans of Atoka, Des Moines, and Missouri (Middle
E34 GEOLOGICAL SURVEY RESEARCH 1962

and Late Pennsylvanian) age (Raymond Douglass, 

written communication, 1961; Helen Duncan and Mackenzie Gordon, Jr., written communication, 1962). The 

upper member, 10,600 feet thick, is composed of quartzite and interbedded limestone that contains fusulinids

3 

BINGHAM

EXPLANATION

Quartzite

\m

Kirkman Limestone 

of Welsh and James 

(1961)

Oquirrh Formation

Park City 

Formation 

Grandeur 

Member

Manning Canyon 

Shale

FIGURE 188.2. Generalized columnar sections of the Rogers 

Canyon and Bingham sequences, northern and central Oquirrh 

Mountains, Utah.

of Missouri, Virgil, and Wolf camp (Late Pennsylvanian and Early Permian) ages (Welsh and James, 

1961). The Oquirrh Formation in this sequence is 

overlain by the Kirkman Limestone of Welsh and 

James (1961) of Wolf camp age.

The Rogers Canyon sequence is similar to the Bingham sequence in gross appearance because in each the 

Oquirrh Formation has three members of generally 

similar lithologies. The underlying and overlying 

rocks are different, however, both in lithology and age; 

and rocks of Late Pennsylvanian age were not identified in the Rogers Canyon sequence but are abundant 

in the Bingham sequence. Beds with similar lithologies are not identical and cannot be mapped across 

the North Oquirrh thrust.

The regional stratigraphic implications of the sequences in the Oquirrh Mountains are not fully 

known; however, some understanding may be gained 

by a comparison of these rocks with well-known sequences of similar rocks (Baker, Huddle, and Kinney, 

1949) in the Wasatch Mountains (fig. 188.3). The 

Timpanogos (thick) sequence, more than 26,000 feet 

thick, resembles the Bingham sequence. The Oquirrh 

Formation in both sequences rests on the Manning 

Canyon Shale of Mississippian and Pennsylvanian 

age and consists of three parts: a lower limestone, a 

middle interbedded limestone and quartzite, and an 

upper quartzite. These units are overlain by the Kirkman Limestone of Early Permian age. The Mount 

Raymond (thin) sequence, about 2,200 feet thick, is 

composed of the Doughnut Formation, a thin basal 

shale and thicker overlying limestone of Mississippian 

age; the Round Valley Limestone of Early Pennsylvanian age; and the Weber Quartzite of probable 

Middle Pennsylvanian age. Overlying the Weber 

Quartzite is the Grandeur Member of the Park City 

Formation, 342 feet thick, composed of a lower limestone of Leonard age (M. D. Crittenden, Jr., written 

communication, 1962), a medial phosphatic zone, and 

an upper dolomite. The thick Bingham and Timpanogos sequences are on the upper plate of the 

Charleston thrust fault. The thin Mount Raymond 

sequence is autochthonous and is presumed to be a 

shelf deposit.

The Rogers Canyon sequence, of intermediate thickness, is on the upper plate of the North Oquirrh 

thrust, which shows apparent south or southeast movement and is interpreted as a remnant of a sequence 

of rocks intermediate between the thin Mount Raymond and thick Bingham and Timpanogos sequences. 

In general, the Rogers Canyon rocks possess aspects 

of both the thin and thick sequences. The rocks above
ARTICLE 188 E35

TIMPANOGOS

EXPLANATION 

Quartette

Sandstone

Sandy limestone

Limestone

Cherty limestone

Shaly limestone

Shale

Locally with phosphorite 

nodules

FEET 

6000

- 4000

- 20OO 

10OO

u 0

VERTICAL 

SCALE

3 

BINGHAM

Kirkman Limestone of 

Welsh and James (1961)

Dolomite

1

MOUNT 

RAYMOND

Park City FormationN 

Grandeur Member

Weber Quartzite 

Round Valley Limestone 

Doughnut Formation/

Park City Formation 

Grandeur Member

FIGURE 188.3. Generalized columnar sections of the Mount Raymond, Rogers Canyon, Bingham, and Timpanogos sequences 

and intervening thrust faults. Location of sections shown on figure 188.1.

and below the Oquirrh Formation in Rogers Canyon 

are comparable with those bounding the Weber 

Quartzite at Mount Raymond. The lithologic character and three-part division of the Oquirrh Formation rocks in the Rogers Canyon sequence are comparable with those of the Bingham sequence. The 

rocks of Late Pennsylvanian and Early Permian age, 

which are thick in the Bingham sequence, thin and 

in part pinch out to the north and east in the Rogers 

Canyon and Mount Raymond sequences. The rocks 

of Early and Middle Pennsylvanian age also thin 

shelfward. Finally, the shale facies common to the

thick sequences at the Mississippian and Pennsylvanian boundary is represented in the thinner sequences by a limestone facies.

In summary, the Mount Raymond sequence is believed to have formed on the shelf, the Timpanogos 

and Bingham sequences in the more rapidly subsiding part of the basin, and the Rogers Canyon sequence in a region intermediate between shelf and 

basin. The Rogers Canyon sequence is the only representative of this intermediate facies that has thus 

far been recognized. Comparable facies may exist in 

north-central Utah. The rocks of the Rogers Canyon
E36 GEOLOGICAL SURVEY RESEARCH 1962

and Bingham sequences, which were originally deposited under different regimens, were juxtaposed in 

the Oquirrh Mountains on the North Oquirrh thrust 

fault.

REFERENCES

Baker, A. A., Huddle, J. W., and Kinney, D. M., 1949, Paleozoic geology of north and west sides of Uinta Basin, Utah: 

Arn. Assoc. Petroleum Geologists Bull., v. 33, no. 7, p. 1161- 

1197.

Crittenden, M. D., Jr., 1959, Mississippian stratigraphy of the 

central Wasatch and western Uinta Mountains, Utah, in 

Williams, N. C., ed., Guidebook to the geology of the 

Wasatch and Uinta Mountains transition area: Intermountain Assoc. Petroleum Geologists Guidebook 10th 

Ann. Field Conf., 1959, p. 63-74.

Gillnly, James, 1932, Geology and ore deposits of the Stockton 

and Fail-field quadrangles, Utah: U.S. Geol. Survey Prof. 

Paper 173, 171 p.

Tooker, E. W., and Roberts, R. J., 1961, Stratigraphy of the 

north end of the Oquirrh Mountains, Utah, in Cook, D. R., 

ed., Geology of the Bingham mining district and northern 

Oquirrh Mountains: Utah Geol. Soc. Guidebook to the 

Geology of Utah, no. 16, p. 17-35.

Welsh, J. E., and James, A. H., 1961, Pennsylvanian and 

Permian stratigraphy of the central Oquirrh Mountains, 

Utah, in Cook, D. R., ed., Geology of the Bingham mining 

district and northern Oquirrh Mountains: Utah Geol. Soc. 

Guidebook to the Geology of Utah, no. 16, p. 1-16.

189. WINDY GAP VOLCANIC MEMBER OF THE MIDDLE PARK FORMATION, MIDDLE PARK, COLORADO

By G. A. IZETT, K. B. TAYLOR, and D. L. HOOVER, Denver, Colo.

A distinctive sequence of andesitic and trachyandesitic volcaniclastic rocks at or near the base of the 

Middle Park Formation of Late Cretaceous and Paleocene age is here named the Windy Gap Volcanic 

Member. The type locality is on the north side of 

U.S. Highway 40, at Windy Gap, 4 miles west of 

Granby, Colo. The measured type section is given 

at the end of this article, and the areal extent of the 

member in central Middle Park, Colo., is shown in 

figure 189.1.

The lowest volcanic member of the Middle Park 

Formation has long been recognized as a mappable 

unit in Middle Park. Marvine (1874) first mapped 

the volcanic unit and applied the name "doleritic 

breccia." Later, Cross (1892) recognized the volcanic 

sequence as the basal unit of a thick complex assemblage of shales, sandstones, grits, conglomerates, and 

volcanic breccias, to which he applied the name Middle Park Formation. He noted that the basal volcanic 

unit was andesitic rather than dolerite as described 

by Marvine. Lovering (1930) described and mapped 

the volcanic unit along the Granby anticline and reported that the volcanic breccia "'* * * is a valuable key 

to the geologic section." Tweto (1947; 1 1957) informally recognized the unit and described its distribution

HOT SULPHUR SPRINGS QUADRANGLE 

(IZETT AND HOOVER)

1 Tweto, Ogden, 1947, Precambrian and Laramide geology of the 

Vasquez Mountains, Colorado : Michigan Univ. thesis, 191 p.

FIGURE 189.1. Index map showing generalized outcrop of the 

Windy Gap Volcanic Member of the Middle Park Formation, 

Middle Park, Colo.
ARTICLE 189 E37

and relation to the geologic history of Middle Park. The 

writers traced the volcanic unit in outcrop in much 

of central Middle Park and mapped it in detail in 

the Hot Sulphur Springs and Bottle Pass quadrangles.

Using the nomenclature proposed by Fisher (1961), 

the Windy Gap Member may be said to consist of 

complexly intercalated multicolored breccia; lapillistone; tuff; epiclastic volcanic conglomerate, sandstone, and siltstone; and local thin fluidal flows and 

a few dike-like bodies. It ranges from 0 to about 

1,150 feet in thickness, but averages about 600 feet 

in thickness in much of the area. Sorting and 

stratification within the member are variable. The 

clastic material of the upper part of the unit is generally well sorted and ranges from thin to thick 

bedded; that of the middle and lower parts is very 

poorly sorted and bedded. The beds are massive, well 

indurated, and form prominent ridges.

The Windy Gap Member in most areas rests unconformably on the Pierre Shale of Late Cretaceous 

age, but locally the unconformity cuts across the 

Pierre, and the Windy Gap rests on the Niobrara Formation of Late Cretaceous age. The contact between 

the Windy Gap and the older rocks is sharp and 

marked by a few feet of reworked shale overlain by 

poorly sorted volcaniclastic rocks. Fossils collected 

by the writers from near the top of the Pierre at 

Windy Gap were identified by W. A. Cobban as 

Inoceramus cf. /. convexus Hall and Meek, and referred to the zone of Baculites scotti, which is correlative to the upper part of the Hygiene Sandstone 

Member of the Pierre Shale (Scott and Cobban, 1959). 

In areas where the Windy Gap Member pinches out 

into the lower part of the Middle Park Formation, 

volcanic conglomerates mterfmger with arkosic conglomerates.

The upper contact of the member is gradational. 

Volcanic breccias and epiclastic volcanic conglomerates, sandstones, and siltstones grade upward into 

grits that generally contain a farily high percentage 

of large Precambrian granitic pebbles plus microcline, quartz, and muscovite granules. Olive-gray and 

yellowish-brown siltstone and claystone as well as thin 

discontinuous limestone and poor-quality coal beds 

overlie the member at Windy Gap. Generally the 

contact was set arbitrarily at the lowest occurrence of 

large amounts of Precambrian granitic rock and microcline fragments.

Volcanic breccias form a large part of the Windy 

Gap Volcanic Member, and they are well exposed in 

the roadcut at Windy Gap. The breccias consist of

variable amounts of angular to subrounded andesite 

and trachyandesite fragments that range in size from 

coarse ash to blocks as much as 5 feet in diameter, 

set in a fine-grained matrix that contains crystals of 

feldspar, pyroxene, hornblende, and biotite. Figure 

189.2 shows a typical sample of the breccia. Rocks 

other than volcanics are rare, but boulders of Precambrian granitic rocks as much as 2 feet in diameter, 

and some granules and sand-sized grains of microcline and quartz, occur near the base of the member 

at Windy Gap and in areas where the member is thin.

FIGURE 189.2. Typical volcanic breccia of the Windy Gap 

Volcanic Member.

The volcanic-rock fragments in the Windy Gap are 

composed of several types of andesite and trachyandesite with porphyritic varieties distinguished by differences in abundance and mineralogy of phenocrysts. 

A few fragments have been tentatively identified as 

porphyritic trachyte. Figure 189.3 shows a photomicrograph of volcanic breccia from near the base of 

the member. The groundmass of many fragments 

consists chiefly of zoned plagioclase (An40-50 ), diopsidic augite, magnetite, and small amounts of hornblende, biotite, and brown glass. Other fragments also 

contain considerable amounts of microcrystalline or 

cryptocrystalline soda-rich alkali feldspar, as identified by X-ray methods. The groundmass may be felty 

or trachytic and ranges from microcrystalline to fine 

grained. Phenocrysts as much as 5 mm in diameter 

may form up to a third of the volume of the porphyritic varieties. The phenocrysts are mainly sub-
E38 GEOLOGICAL SURVEY RESEARCH 1962

-i, s*r - _ - fcjf~."i'v-KJ?M^sr.»?.>- - - * . .--r.*- Y^-^^ON* 

;; -># JB. - -jfr^v 'i%i<1 ' **- f^-.^.^-. £/>i:^" ^'%l^^' ..^ ^- * % :*,\.«...* --. *^T. ^T --.Vi - Hsr*^ t^s^V ^"v*vir': " :f$^^«- ^ -, ^:^^l

^rfT^aL.. . * ". * - - .»?'.! J ! * -J - ^ . .«

"» «-*%

FIGURE 189.3. Photomicrograph of volcanic breccia from near 

base of Windy Gap Member showing clastic trachyte (t), 

trachyandesite (ta), and andesite (a) fragments in matrix of 

plagioclase, pyroxene, hornblende, quartz, and magnetite.

hedral to euhedral zoned plagioclase (An40_65 ) and 

euhedral diopsidic augite (2F 53°-55°); other phenocrysts seen in a few fragments are oxyhornblende, 

common hornblende, quartz, alkali feldspar, and biotite. Hornblende and orthopyroxene (?) ghosts are 

common in some fragments. Other minerals in the 

rock fragments, identified by X-ray methods, include 

zeolites (analcite, heulandite, thompsonite, stilbite, and 

phillipsite?), nontronite, celadonite, chlorite, calcite, 

limonite, quartz, and albite.

Lavas and dikes are rare, but at least one flow and 

several narrow dikes of trachyandesite are exposed 

a short distance above the base of the volcanic unit, 

near the forks of Pole Creek. The flow is about 25 

feet thick, has a sharp base, and has a brecciated top 

consisting of a mixture of coarse debris of the flow 

in a fine-grained epiclastic matrix. The flow rock 

is brown altered porphyritic trachyandesite. The 

groundmass consists of microlites of alkali feldspar, 

andesine, clinopyroxene, and magnetite, set in a matrix 

of a brown mineral, perhaps nontronite, that probably 

formed from glass. The phenocrysts constitute about 

a fourth of the rock and are augite, strongly zoned 

andesine-labradorite, and magnetite.

No fossils diagnostic of age were found in the 

Windy Gap Member; however, pollen identified by 

E. H. Tschudy, U.S. Geological Survey, from a bed 

of carbonaceous debris that overlies the member indicates that the member is of Late Cretaceous age. More 

collections are needed to confirm this determination.

The Windy Gap Member is locally the basal unit 

of a thick complex assemblage of shales, sandstones, 

grits, and conglomerates with subordinate volcaniclastic rocks (Middle Park Formation) that were deposited in the Middle Park Basin in Late Cretaceous and 

early Tertiary time. The source of the Precambrian 

detritus was doubtless the bounding Front and Park 

Ranges. Plugs of andesite and augite syenite near the 

forks of Ute Bill Creek and at the headwaters of 

Eightmile Creek may be the source of some of the 

rocks in the Windy Gap Member. In Late Cretaceous 

time, the surface expression of the plugs may have

Windy Gap Volcanic Member of the Middle Park Formation, 

sec. 27, T. 1 N., R. 77 W., Grand County, Colo.

Middle Park Formation: peet

8. Mudstone, reddish-brown, structureless; contains 

well-rounded andesite granules, pebbles, and cobbles; overlain northward by poor-quality coal beds 

and thin-bedded limestone ; forms slope _________ 50

Windy Gap Volcanic Member:

7. Conglomerate, volcanic; contains well-rounded reddish-brown, greenish-gray, and grayish-red andesite fragments as much as 2 ft in diameter; matrix 

is yellowish-brown volcanic sandy mudstone; 

grades upward into overlying unit; fractures in 

cobbles are coated with heulandite, forms ridge __ 44

6. Covered; float is reddish-brown soft siltstone; forms

slope, _________ _ ________________ _______ 300

5. Sandstone, volcanic, dark-yellowish-brown, conglomeratic; contains rounded andesite fragments 

as much as 6 inches in diameter; laminated to 

thick bedded, well bedded; some units lenticular; 

heulandite in matrix, concentrations of magnetite 

along laminae; topmost unit is locally chertified 

and contains abundant carbonaceous debris; forms 

ridge_ ___________________-_-------_--_- __ _ 124

4. Covered; float is brownish- red soft siltstone; forms

slope _ _________________-_----_--_---------_- 75

3. Breccia, volcanic; greenish-gray matrix; varicolored 

boulders, of angular to subangular trachyte, 

trachyandesite, and andesite fragments, up to 5 ft 

in diameter; a few stratification planes outlined by 

fine-grained volcanic sandstones but generally 

massive and poorly sorted; highly indurated, 

fractures are coated with zeolites (stilbite, analcite, 

thompsonite, and heulandite) and calcite; grades 

downward into underlying unit; forms massive 

cliff______ ________ __ ________ _ __ __ _-__ 500

2. Breccia, same as unit 3 except that basal part contains fragments of granitic rocks, quartz, and 

microcline as much as 2 ft in diameter; a few 

lenticular tuff beds; contact with underlying 

Pierre Shale not well exposed. _ ________________ 100

Total thickness______-__-------_____ ______ 1, 143

Pierre Shale:

1. Shale, medium-olive-gray, silty, structureless; interbedded with light-gray very fine grained sandstone; forms sIope_________-_-_-_______------- 25
ARTICLE 189 E39

been andesite domes similar to those described by Williams (1932). The rapid destruction of these and 

other domes by running water and debris flow could 

result in a tabular mass of volcaniclastic rocks that 

would have characteristics in common with the Windy 

Gap Member. The Tuscan and Mehrten Formations 

of California are similar in many respects to the 

Windy Gap Member, and a similar origin for them 

has been proposed by Anderson (1933) and Curtis 

(1954).

REFERENCES

Anderson, C. A., 1933, The Tuscan Formation of northern California, with a discussion concerning the origin of volcanic 

breccias: California Univ. Dept. Geol. Sci. Bull., v. 23, 

no. 7, p. 215-276.

Cross, Whitman, 1892, The post-Laramie beds of Middle Park, 

Colorado: Colorado Sci. Soc. Proc., v. 4, p. 192-213.

Curtis, G. H., 1954, Mode of origin of pyroclastic debris in the

Mehrten Formation of the Sierra Nevada: California

Univ. Pubs. Geol. Sci. Bull., v. 29, no. 9, p. 453-502. 

Fisher, R. V., 1961, Proposed classification of the volcaniclastic sediments and rocks: Geol. Soc. America Bull., v. 72,

no. 9, p. 1409-1414. 

Levering, T. S., 1930, The Granby anticline, Grand County,

Colorado: U.S. Geol. Survey Bull. 822-B, p. 71-76. 

Marvine, A. R., 1874, Report of the Middle Park division:

U.S. Geol. and Geog. Survey Terr. 7th Ann. Rept (for

1873), p. 83-192. 

Scott, G. R., and Cobban, W. A., 1959, So-called Hygiene Group

of northeastern Colorado, in Rocky Mtn. Assoc. Geol.

Guidebook llth Ann. Field Conf., Washakie-SandwashPiceance Basins, Colorado, 1959: p. 124-131. 

Tweto, Ogden, 1957, Geologic sketch of southern Middle Park,

Colorado, in Rocky Mtn. Assoc. Geol. Guidebook 9th Ann.

Field Conf., North and Middle Park Basins, Colorado,

1957: p. 18-31. 

Williams, Howel, 1932, The history and character of volcanic

domes: California Univ. Dept. Geol. Sci. Bulls., v. 21,

no. 5, p. 51-146.
E40 GEOLOGICAL SURVEY RESEARCH 1962

190. SHALLOW HALITE DEPOSITS IN THE FLOWERPOT SHALE IN SOUTHWESTERN OKLAHOMA

By PORTER E. WARD, San Juan, Puerto Rico

Work done in cooperation with the U.S. Public Health Service

An investigation of the source of salt contributing 

to mineralization of the Arkansas and Red Rivers 

indicated that shallow deposits of halite in the Flowerpot Shale of Permian age were associated with salt 

springs (Ward, 1960, 1961;. This article summarizes 

the results of test drilling in an area of salt springs 

along the Elm Fork of the Red River in Beckham 

and Harmon Counties, Okla.

Subsurface studies are based on a 4-inch core hole 

drilled by the Corps of Engineers, U.S. Army, in 

October 1960 and 7 cable-tool holes drilled in July 

and August 1961 (fig. 190.1). The core hole was 226 

feet deep and the cable-tool holes ranged in depth 

from 37 to 255 feet.

The formations exposed in the area are of Permian 

age and are, in ascending order, the Flowerpot Shale,

R. 26 W.

2 MILES

Area ofreport

OKLAHOMA

EXPLANATION

CX,

Salt spring

0 3

Test hole and 

number

FIGURE 190.1. Map of the Harmon-Beckham County 

salt-spring area, showing test holes, major salt springs, 

and lines of geologic sections. Sections A-A' and 

B-B' are shown on figure 190.3.

Blaine Gypsum, and Dog Creek Shale. The Dog 

Creek Shale crops out south of the salt-spring area, 

but it is unrelated to the problem under investigation 

and is not discussed.

The Blaine Gypsum consists mainly of beds of massive gypsum separated by red and gray gypsiferous 

shale beds. In many places the gypsum beds are underlain by thin impure dolomite. The Blaine was at 

the land surface at the sites of all the test holes, therefore the entire section of the Blaine is not represented 

in any of the holes. The core hole started a few feet 

below the top of the Blaine and penetrated 123 feet 

of Blaine consisting of 75 feet of gypsum, 10 feet of 

dolomite, and 38 feet of gypsiferous shale. Halite 

was not found in the Blaine in the subsurface, but a 

few salt casts were found on the outcrops.

The Flowerpot Shale consists primarily of reddishbrown and gray gypsiferous shale characterized by 

numerous intersecting selenite veins. More than 80 

feet of the Flowerpot is exposed along the Elm Fork 

of the Red River. The thickness of the Flowerpot is 

not known, but test-hole 2 (figs, 190.1, 190.2) started 

about 20 feet below the top of the formation and apparently remained in it to the bottom of the hole, a 

depth of 220 feet.

Salty zones in the Flowerpot Shale are exposed in 

several places a few feet above the bottom of steepwalled canyons in the area. The salt in these zones 

may have been dissolved from subsurface halite beds 

and brought to the surface by ground water; presumably the water evaporated at the surface, leaving a 

salt residue on the rocks. However, the salt in these 

zones may have been deposited with the shale. Small 

voids bearing impressions of salt hoppers are present 

in many of the salty zones (fig. 190.3). Also, hoppershaped casts composed generally of dolomite, and 

rarely of gypsum, were found on the weathered outcrop of the Flowerpot Shale. One such cast consisting of impure gypsum measured 6 inches on a side. 

Both the voids and the casts prove that halite formerly 

occurred in the shale. The time of removal of the 

halite is unknown; but salt is being removed continually from the subsurface through solution as indicated by the local salt springs, all of which issue from 

the Flowerpot.
ARTICLE 190 E41

A'

1950'n

1900'-

1850'-

1800'

1750'-

1700'-

1650'-

1600'-

1950'-i

1900'-

1850'-

1800'

1750'-

1700'-

1650'

EXPLANATION

Shale Dolomite Gypsum Salt

(left, bedded; right 

disseminated in 

shale)

1MILE

FIGURE 190.2. Generalized geologic sections of the Harmon-Beckham County salt-spring area. Datum is sea

level.
E42 GEOLOGICAL SURVEY RESEARCH 1962

FIGURE 190.3. Small void in the Flowerpot Shale 

bearing impressions of hopper-shaped halite 

crystals.

Halite was found in the Flowerpot Shale in the 

subsurface both disseminated in the shale and in lenticular beds. Test-holes 6 and 7 and the core hole penetrated bedded halite, and all the holes penetrated salty 

shale (fig. 190.2). Bedded halite was reached at depths 

of 30, 197, and 149 feet below the land surface (fig. 

190.2), corresponding to altitudes of 1,720, 1,725, and 

1,764 feet, respectively.

The core hole penetrated a zone in the Flowerpot 

Shale between depths of 197 and 200 feet that contained salt crystals and solution voids, many of which 

ranged in diameter from i/2 to 1 inch. The upper few 

inches of the zone contained cubic-shaped voids, formerly occupied by salt crystals. Several inches lower, 

many of the voids contained partly dissolved halite 

crystals. The lower part of the zone contained cubic 

halite crystals but no voids. Veins of bright-redorange halite as much as an inch wide, that fill vertical and near vertical cracks in the shale, were cored 

at a depth of about 208 feet (fig. 190.4). The color 

evidently is imparted by small amounts of iron diffused in the halite.

The test holes were not deep enough to establish 

the full thickness of the halite-bearing beds. TestFIGURE 190.4. Section of a 4-inch core from 

the Flowerp ot Shale from a depth of about 

208 feet in sec. 16, T. 6 N.. R. 26 W. The veins 

are composed of fibrous reddish-orange halite.

hole 6 penetrated salty shale between depths of 90 and 

170 feet, the bottom of the hole. Salty shale was interbedded with halite at depths of 148 to 170 feet. 

Although the areal extent of halite in the Flowerpot 

Shale has not been determined, subsurface studies 

show that it is widespread in western Oklahoma and 

adjacent areas of Kansas and Texas. Individual salt 

beds or zones apparently pinch out within short distances, and sections devoid of halite may be expected 

in some places. In other places, bedded halite and 

halite-bearing shale occur within a few feet of the land 

surface.

REFERENCES

Ward, P. E., I960, Relation of mineral springs to Permian salt 

deposits [ahs.] : Geol. Soc. America Bull., v. 71, no. 12, 

pt. 2, p. 1999.

 1901, Shallow halite deposits in northern Woodward 

and southern Woods Counties, Oklahoma: Oklahoma Geological Survey, Oklahoma Geology Notes, v. 21, no. 10, 

p. 275-277.
ARTICLE 191 E43

191. CRINOIDAL BIOHERMS IN THE FORT PAYNE CHERT (MISSISSIPPIAN) ALONG THE CANEY FORK

RIVER, TENNESSEE

By MELVIN V. MARCHER, Nashville, Term.

WorJc done in cooperation with the Tennessee Division of Geology

Factors that control the development of permeability and porosity in carbonate rocks are of primary interest to hydrologists who are concerned with locating 

and developing ground-water supplies in limestone 

terranes. The well-developed crinoidal bioherms along 

the Caney Fork River in east-central Tennessee provide excellent examples of both primary and secondary permeability. Only the general features and lithologic characteristics of these bioherms are described 

in this article. Additional study now underway will 

define, as closely as possible, the role of rock fabric 

and diagenesis in the development of porosity in the 

biohermal beds. The results of such a study should 

be equally applicable to lithologically similar beds not 

associated with bioherms.

Crinoidal bioherms, which are locally well developed 

along the Caney Fork River in the reach extending 7 

miles downstream from Great Falls Dam, occur at or 

near the base of the Fort Payne Chert of Early Mississippian age. The base of most bioherms is obscured 

by talus and vegetation. At least one small bioherm, 

near Great Falls Dam, is underlain by Fort Payne 

Chert. Two other bioherms rest directly upon the 

Chattanooga Shale, which underlies the Fort Payne; 

at these localities the Chattanooga is bowed slightly 

downward, presumably from the weight of the bioherms. Where the outlines of the bioherms can be 

determined, they are oval or elongate moundlike 

masses. The long axis of most bioherms trends from 

about N. 30° W. to about N. 30° E. (fig. 191.1). The 

dimensions of individual bioherms range from tens 

to hundreds of feet in both width and length and from 

20 to nearly 100 feet in thickness. Preliminary study 

indicates that a completely developed bioherm consists of three major parts: a core, inner flanking beds, 

and outer flanking beds (fig. 191.2). In most of the 

bioherms, the beds making up the core are flat lying 

whereas the flanking beds drape over the core and dip 

outward 5° to 20°. In some incompletely developed 

bioherms, the inner flanking beds are missing and the 

outer flanking beds are in direct contact \vith the core.

The overlying beds of Fort Payne Chert arch over 

the bioherms and consist of uniformly bedded very 

fine grained, calcareous and siliceous siltstone which 

locally contains scattered crinoidal debris. The lithologic break between the overlying beds and the MOEXPLANATION

TENNESSEE \

Area of biohermal 

development in the 

Fort Payne Chert

Trend of

biohermal

axis

Area of 

report

FIGURE 191.1. Distribution of bioherms along the Caney Fork

River.

hermal beds is very distinct, and the contact between 

them represents a submarine unconformity. At the 

margins of the bioherms, however, the enclosing beds 

grade laterally into the biohermal beds. Arching of 

the beds overlying the bioherms is the result of initial 

dip and gradually dies out upward.

The strata that make up the cores of the bioherms 

are thin-bedded greenish-gray argillaceous limestone 

and greenish- or reddish-gray calcareous shale. Crinoid columnals as much as 1 inch in diameter are 

common. In thin sections, it can be seen that the rock 

consists mainly of crinoidal and algal debris in a 

fine-grained dolomitic and clayey matrix. Shreds of 

brown organic(?) material are abundant. Many of 

the crinoid fragments have corroded edges, suggesting

667422 O - 63 -
E44 GEOLOGICAL SURVEY RESEARCH 1962

EXPLANATION

Enclosing beds

Siltstone, very cherty; contains 

scattered crinoidal debris

Core

Crinoidal limestone,fine-grained, very 

argillaceous and shaly, dolomitic; 

contains much algal material

Inner flanking beds

Algal limestone, fine- to mediumgrained, dolomitic, very irregularly 

bedded; contains scattered crinoidal 

debris

Outer flanking beds

Crinoidal and algal limestone, fine- to 

very coarse grained, very massive, 

gypsiferous; contains irregular 

patches of vuggy dolomite and limestone rubble

FIGURE 191.2. Sketch section of a completely developed

bioherm.

that some recrystallization has occurred, and some are 

partly or completely silicified. In places the matrix 

also has been silicified. Algal material forms microlaminated sheaths on fossil fragments or occurs as 

small (0.05 mm or less) spherical bodies, many of 

which are filled with clear crystalline calcite. Parts 

of the core consist mainly of fine- to medium-grained 

silty dolomitic limestone that contains only minor 

amounts of crinoidal and algal debris.

Where present, the inner flanking beds ordinarily 

are thin and irregularly bedded and consist of paleyellowish-brown fine- to medium-grained dolomitic 

limestone that contains various amounts of fossil debris. Under magnification, the inner flanking beds 

are seen to contain much algal material in a slightly 

argillaceous and dolomitic matrix. Most of the crinoid fragments, which are randomly distributed 

throughout the rock, have irregular edges as a result 

of dolomitization. Spherical algal bodies are abundant and some have been filled with calcite. Irregular patches in the rock and some of the crinoid fragments have been replaced by microcrystalline quartz.

The outer flanking beds are massive and lenticular 

and consist mainly of medium- to coarse-grained dolomitic limestone that contains much crinoidal and algal debris. In some thin sections from the outer flanking beds, it could be seen that practically all the 

crinoid fragments had been replaced by cryptocrystalline or crystalline quartz and that part of the matrix

had been irregularly silicified. In at least one thin 

section, parts of the silicified patches had been replaced by dolomite. Large dolomite crystals, possibly 

due to replacement of chert by dolomite (Walker, 

1962, p. 238), are common in the silicified patches.

Parts of the outer flanking beds are a rubble consisting of angular fragments and nodules of finegrained slightly argillaceous limestone in a matrix 

of fine- to medium-grained dolomitic limestone that 

contains various amounts of crinoidal and algal debris. 

Many of the large crinoid stems are surrounded by 

clear crystalline calcite. Most of this calcite is secondary in origin, but some appears to be due to recrystallization of the fine-grained matrix.

Some of the rubble zones in the outer flanking beds 

contain irregular vugs and seams that are partly filled 

with powdery gypsum. Some of the smaller vugs and 

seams are filled with a mixture of crystalline calcite 

and gypsum, and a few are lined with dolomite rhombohedra.

Other parts of the outer flanking beds consist of 

highly porous and permeable dolomite. Dolomitization has destroyed the outline of most of the fossil 

fragments, especially the crinoid columnals, but a few 

spherical algal bodies, some filled with clear crystalline dolomite or calcite, are relatively unaffected. 

Some of the openings or pores in the rock have been 

partly filled with carbonate minerals or are lined with 

dolomite crystals.

The bioherms began to form when the depth of the 

Fort Payne sea was optimum for establishment of 

bioherm-f orming crinoid colonies on the sea floor. For 

a time the growth of the bioherms kept pace with a 

slowly rising sea level, but at times the crests of the 

bioherms reached the level where wave action prevented further growth. At such times the uncemented debris on the crests of the bioherms would be 

swept off and accumulate on the flanks. Further rises 

in sea level submerged the bioherms to a depth such 

that their crests again afforded optimum conditions 

for the growth of crinoids. Fluctuations in sea level 

occasionally allowed waves to sweep debris off the 

crests onto the flanks of the bioherms. The zones of 

rubble in the outer flanking beds may have been formed 

by sliding of material down the sides of the bioherms 

when the angle of repose was exceeded or when storms 

swept large amounts of coarse debris from the crests 

of the bioherms. The angularity of much of the rubble indicates that the fragments were not transported 

for a very great distance. Eventually the rate of sealevel rise outpaced the rate of biohermal growth and
ARTICLE 192 E45

created environmental conditions inimicable to the 

survival of the crinoid colonies.

Preliminary study indicates that the porosity of the 

outer flanking beds is both primary and secondary in 

origin. Primary openings appear to have resulted 

from a failure of fine sediment or cement to be deposited in the original open spaces in the rigid framework of limestone rubble (Murray, 1960, p. 61). Sea 

water trapped in such openings may account for the 

gypsum that lines or fills vugs in the rubble zones. 

Openings of secondary origin are of two types: relatively large, irregularly distributed conduits dissolved

by circulating ground water, and small uniformly distributed pores developed as a result of dolomitization 

(Weyl, 1960, p. 85-90).

REFERENCES

Murray, R. C., 1960, Origin of porosity in carbonate rocks: 

Jour. Sed. Petrology, v. 30, no. 1, p. 59-84.

Walker, T. R., 1962, Reversible nature of chert-carbonate replacement in sedimentary rocks: Geol. Soc. America Bull., 

v. 73, no. 2, p. 237-242.

Weyl, P. K., 1960, Porosity through dolomitization: Conservation-of-mass requirements: Jour. Sed. Petrology, v. 30, 

no. 1, p. 85-90.

192. MISSISSIPPIAN-PENNSYLVANIAN BOUNDARY IN NORTHEASTERN KENTUCKY

By EICHAED A. SHEPPAED and ERNEST DOBEOVOLNY, Denver, Colo.

Work done in cooperation with the Kentucky Geological Survey

A. thick sandstone unit forms the walls of Box Canyon, south of Carter Caves State Park near Olive Hill, 

Ky. (fig. 192.1). The sandstone unit is overlain by 

shale and limestone assigned to the Pennington (?) 

Formation of Late Mississippian age, and the Pennington (?) is overlain by the Lee Formation, which is 

considered to be of Early Pennsylvanian age. The 

sandstone, 43 feet thick, is white to light gray, fine to 

medium grained, quartzose, thick bedded, and conspicuously crossbedded. It is lithologically similar to 

sandstones in the Lee Formation and unlike other 

sandstones in the Pennington Formation.

The stratigraphic relations of the sandstone are 

shown in figure 192.2. A core drilled on the shoulder 

of Box Canyon in 1961 penetrated the Pencil Cave 

beds (drillers' term) below the sandstone. The Pencil Cave beds and a limestone generally found above 

it are correlated respectively with the Hardinsburg 

Sandstone and Glen Dean Limestone of western Kentucky (McFarlan and Walker, 1956, p. 7).

The Pennington (?) Formation consists chiefly of 

shale, but includes a few thin beds of limestone and 

sandstone. The shale generally is light green to greenish gray, fossiliferous, and locally silty. Eed shale

FIGURE 192.1. Index map of measured sections and drill hole 

 in vicinity of Carter Caves State Park, Ky.
E46 GEOLOGICAL SURVEY RESEARCH 1962

A

Roadcut along State 

Highway 182 about 

0.6 mile southeast 

of State Highway 2

c

Drill hole on Box 

Canyon about 100 

feet east of State 

Highway 182

Top of hole 

USGS20209-PC/]

EXPLANATION

LEE FORMATION £t^

D

Roadcut on hill 

south of Tygarts 

Creek

LEE 

FORMATION

Shale

Shale with thin 

limestone lenses

Shale with thin 

sandstone interbeds

Limestone

Sandstone

USGS 20206-PC 

Fossil locality

FEET 

r20

-15

-10 

r 5

-0

PENCIL CAVE

(DRILLERS'

TERM)

EQUIVALENT OF 

GLEN DEAN 

LIMESTONE

PENCIL CAVE

(DRILLERS'

TERM)

FIGURE 192.2. Sections of Mississippian and Pennsylvanian rocks in vicinity of Carter Caves State Park, Ky.

was observed in the upper beds in the core at Box 

Canyon and in nearby outcrops. The most continuous beds are two limestones, one at or near the base 

of the formation and another near the top. Both are 

gray, finely crystalline, argillaceous, and fossiliferous. 

Weathering along argillaceous laminae causes the limestone to separate into plates as much as 1 inch thick. 

Thin, discontinuous yellow-weathering limestones are 

present in the upper part of the formation; fossils 

were not found in these. With the exception of the 

thick sandstone at Box Canyon, the sandstones are 

light tan to light olive, very fine to fine grained, thin 

bedded, and quartzose. Green shale chips occur locally. The sandstones generally are fossiliferous; they 

contain abundant spireferoid brachiopods and, locally, 

fish remains. Perhaps the most diagnostic feature is 

the abundance of casts of trails on the bottoms of 

certain beds.

Marine fossils collected from the Pennington (?) 

Formation in this area (figs. 192.1 and 192.2) were 

identified by J. T. Dutro, Jr., written communication, 

1962) who commented as follows:

All the assemblages indicate a Late Mississippian age * * *. 

There is no indication of an Early Pennsylvanian age for any

of these collections. The occurrence of Diaphragmus cf. 

D. cestriensis (Worthen) * * * points to a late Chester age. 

The other invertebrates, especially brachiopods, support this 

general age assignment.

Diaphragmus cf. D. cestriensis (Worthen) was collected above the sandstone at locality C.

Fossils collected from localities A and C (figs. 192.1 

and 192.2) are:

Locality A:

USGS 20204-PC:

Horn coral, indet.

Echinoderm debris, indet.; abundant

Fenestrate bryozoans, indet.; abundant

Ramose bryozoans, undet.

Productoid brachiopod, indet.

Spirifer cf. S. leidyi Hall

Composite, subquadrata (Hall)

Eumetria cf. E. vera Hall

"Spiriferina" spinosa (Norwood & Pratten) 

USGS 20205-PC:

Spirifer sp.

Pelecypod fragments, indet.; abundant

Tracks and trails, indet. 

USGS 20206-PC:

Echinoderm debris, indet.

Orthotetes cf. 0. kaskaskiensis (McChesney)
ARTICLE 193 E47

Productoid fragment, indet.

Spirifer cf. S. leidyi Hall

Dielasma? sp.

Fish tooth, undet. 

Locality C:

USGS 20207-PC:

Orthotetes cf. 0. kaskaskienses (McChesney)

Allorisma sp.; abundant 

USGS 20209-PC:

Echinoderm debris, indet.; abundant

Fenestrate bryozoans, indet.

Ramose bryozoans, undet.

Orthotetes cf. 0. kaskaskiensis (McChesney)

Diaphragmus cf. D. cestriensis (Worthen)

Ovatia ovata (Hall)

Spirifer cf. S. increbescens Hall

Composita subquadrata (Hall)

Aviculopecten sp.

Bellerophon sp.

Straparollus (Euomphalus) sp.

Murchisonia? sp.

Although the sandstone body is shown as a lens in 

figure 192.2, it may be a tongue of the Lee Formation. 

Englund and Smith (1960) have demonstrated intertonguing of the upper part of the Pennington Formation and the lower part of the Lee Formation in 

southeastern Kentucky. If the sandstone is a tongue 

of the Lee Formation, the Lee includes beds of Late 

Mississippian age at this locality. In either case it is 

clearly indicated that deposition of quartzose sandstones typical of the Lee Formation began during Late 

Mississippian time.

REFERENCES

Englund, K. J., and Smith, H. L., 1960, Intertonguing and 

lateral gradation between the Pennington and Lee Formations in the tri-state area of Kentucky, Tennessee' and 

Virginia [abs.] : Geol. Soc. America Bull., v. 71, p. 2015.

McFarlan, A. C., and Walker, F. H., 1956, Some old Chester 

problems correlations along the eastern belt of outcrop: 

Kentucky Geol. Survey Bull. 20, 36 p.

193. CORRELATION OF THE PARKWOOD FORMATION AND THE LOWER MEMBERS OF THE POTTSVILLE

FORMATION IN ALABAMA

By WILLIAM C. CULBERTSON, Denver, Colo.

In Alabama the Pottsville Formation of Pennsylvanian age, the principal coal-bearing formation, contains one or two resistant conglomeratic quartzose 

sandstone (orthoquartzite) units at its base that contrast with the less resistant subgraywacke sandstone 

beds in the rest of Pottsville. In the Cahaba and 

Coosa coal fields the two quartzose sandstone units 

are named the Shades Sandstone Member and the 

Pine Sandstone Member; in the Plateau coal field the 

two units are the "Lower Conglomerate" and the "Upper Conglomerate" of McCalley (1891), but in the 

Warrior coal field only one unit has been named, the 

Boyles Sandstone Member (fig. 193.1). According to 

the geologic map of Alabama by Butts (1926), the 

Pottsville is underlain by the Mississippian Parkwood 

Formation in the Cahaba and Coosa coal fields and on 

Blount Mountain, but elsewhere it is underlain by the 

Pennington Formation or the Floyd Shale, both of 

Mississippian age. It is the purpose of this article to 

clarify the correlation of lower members of the Pottsville Formation and to redefine the top boundary, and

the extent, of the Parkwood Formation. (See fig. 

193.1.)

The correlation of the basal sandstone members of 

the Pottsville Formation has been hampered by the 

concept that there is only one quartzose sandstone unit 

in the Warrior field, compared to two units in each 

of the other coal fields (for example, Butts, 1910, 

p. 8, and Shotts, 1961). The Boyles Sandstone Member was named and described by Butts as follows:

The basal part of the Pottsville is everywhere in the Birmingham quadrangle a sandstone. In this particular field 

[Warrior] it is called the Boyles sandstone member because 

it is Avell displayed in Boyles Gap. The sandstone is coarse, 

thick bedded, quartzose, and in places conglomeratic at the 

base. Where thickest it changes to a flaggy, finer-grained and 

perhaps argillaceous rock toward the top * * *. It is 100 feet 

thick where exposed in Sayreton and Boyles gaps.

Recently the Boyles Sandstone Member was exposed by the excavation of a railroad cut through the 

ridge that is formed by this member at a locality 

about 500 feet north of Boyles Gap. The strata are 

here overturned, dipping 45° to 60° SE. The ex-
E48 GEOLOGICAL SURVEY RESEARCH 1962

EXPLANATION

FEET 

r-0

-100

-300

~ - 400

INDEX MAP OF NORTHERN ALABAMA 

Shaded areas are coal fields

0 25 50 MILES

PenningtorK?)' 

Formation

FIGURE 193.1. Columnar sections of Parkwood Formation and lower part of Pottsville Formation in Alabama, 

and map showing location of sections: (1), core of the Brilliant hole (Semmes, 1929, p. 155); (2), measured 

section along Mill Creek; (3), measured section at new railroad cut at Boyles Gap; (4), generalized from 

core of Wham well (Semmes, 1929, fig. 24); (5), generalized measured section by Rothrock (1949, fig. 2); 

and (6), unpublished description of core of FP-1 by N. M. Denson and R..K. Hose in 1944.
ARTICLE 193 E49

posed rocks were measured and described by the author as follows:

Section of the Boyles Sandstone Member of the Pottsville Formation 

in new railroad cut through Sand Mountain, NW/* 1 sec. 6, 

T. 17 S., R. 2 W., Jefferson County, Ala.

[Fig. 193.1, col. 3] 

Pennsylvanian:

Pottsville Formation: Feet 

Shale, dark-gray; 30 percent of beds medium-darkgray carbonaceous very fine to medium-grained 

sandstone. Lower 50 ft contains 4 coal beds that 

are 2 to 12 in. thick; sharp basal contact_____.___ 386

Boyles Sandstone Member: 

Upper sandstone unit:

Sandstone, light-gray, fine- to medium-grained, 

quartzose, well-indurated, thick-bedded; crossbedded 30 ft above base. Contains several thin 

beds of medium-dark-gray sandstone with plant 

fossils. Abundant flat pebbles of shale and coal 

in lower 3 to 8 ft. Upper 18 ft is medium-darkgray medium- to coarse-grained sandstone. 

Irregular, sharp basal contact ________________ 86

Middle shale unit:

Shale, dark-gray, fissile; the lower 4 ft is 30 percent 

thin lenticular siltstone and very fine-grained 

sandstone beds. At 4 ft above base is a bed 8 ft 

thick of medium-dark-gray very fine- to finegrained sandstone that is interbedded with 10 

percent shale. Sharp basal contact,___^______ 33

Lower sandstone unit:

Sandstone, light-gray, fine- to medium-grained, 

with a few coarse to very coarse grains; quartzose, well indurated, crossbedded in places; in 

beds 1 to 5 ft thick. Lower 31 ft is conglomeratic, containing sparse rounded pebbles of quartz 

as much as % in. in diameter, which in places are 

concentrated in lenses as much as 1 ft thick. 

Lowest 3 ft contains abundant quartz and shale 

pebbles and fragments of coal ________________ 114

Total, Boyles Sandstone Member___________ 233

Unconformity. 

Mississippian: 

Floyd Shale:

Clay shale, dark-gray, fissile with 5 to 10 percent 

beds of buff-weathering fine-grained sandstone.

A comparison of the exposure at the old railroad 

cut at Boyles Gap with that at the new railroad cut 

described above indicates that Butts (1910, p. 8) was 

referring only to the lower 114 feet of sandstone when 

he stated that the Boyles Sandstone Member was a 

single sandstone bed 100 feet thick. However, the 

upper unit is also a "coarse, thick-bedded quartzose 

sandstone," and the upper unit was included by Butts 

(1910) in the Boyles Sandstone Member in other parts 

of the Birmingham quadrangle. For example, at Mill 

Creek (fig. 193.1, col. 2), the Boyles Sandstone Member, as mapped by Butts (1910), consists of an upper

and a lower ridge-forming quartzose sandstone unit 

separated by a unit 120 feet thick that is mostly shale. 

This is the typical lithology of the Boyles in most of 

the Warrior coal field, although locally the middle 

shale unit is cut out by channeling, or locally the 

sandstone units contain one or two interbeds of shale.

In order to make the boundaries of the Boyles 

Sandstone Member consistent throughout the Warrior 

coal field, the Boyles Sandstone Member is here redefined to include the entire 233-foot section described 

at the new railroad cut at Boyles Gap, which is here 

designated as a reference section.

On the basis of the stratigraphic position, unique 

lithologic sequence, and persistence of these lower 

units of the Pottsville, the lower sandstone unit of 

the Boyles Sandstone Member is correlated with the 

Shades Sandstone Member and the "Lower Conglomerate" of McCalley (1891). The upper sandstone unit 

is correlated with the Pine Sandstone Member and 

the "Upper Conglomerate" (fig. 193.1).

The Parkwood Formation was named from exposures in the vicinity of Parkwood, T. 19 S., E. 3 W., 

in the Cahaba coal field. Butts (1910, p. 8) defined 

it as including the 1,500 to 2,000 feet of gray shale 

and sandstone "lying above the base of the sandstone 

making Little Shades Mountain and Bald Ridge * * * 

and below the Brock coal bed * * *." It is underlain 

in some places by the Pennington Formation, in other 

places by the Floyd Shale.

The Brock coal bed is 6 to 10 inches thick and lies 

about 50 feet below the prominent ridge-forming 

Shades Sandstone Member of the Pottsville Formation. Butts selected the Brock coal because, as he 

stated later (Butts, 1926, p. 206), "* * * the base of 

the Pennsylvanian of Alabama, as located at the base 

of the Brock coal bed, seems to lie at a horizon in the 

lower Pottsville as low as the lowest Pennsylvanian 

of Virginia, West Virginia, or Pennsylvania * * *." 

Subsequent paleontological work, however, showed 

that the base of the Pennsylvanian is an indefinite 

boundary more than 500 feet below the Brock coal 

bed (Butts, 1927, p. 13; Moore, 1944, p. 686). In addition, the thinness and limited extent of the Brock 

coal make its base an impractical formation boundary. 

The top boundary of the Parkwood Formation is here 

changed to the more widespread and conspicuous lithologic boundary, the base of the Shades Sandstone 

Member of the Pottsville Formation (fig. 193.1). This 

boundary was used in mapping the Coosa coal field by 

Butts (1940, p. 11) and Rothrock (1949, p. 18).
E50 GEOLOGICAL SURVEY RESEARCH 1962

According to Butts (1910), the Parkwood Formation is not present in the Warrior and Plateau coal 

fields except on Bloimt Mountain where about 300 

feet of Parkwood underlies the Boyles Sandstone 

Member. However, a sequence of sandstone, shale, 

and one or more coal beds underlies the Boyles Sandstone Member and "Lower Conglomerate," and unconformably overlies the Mississippian Pennington Formation in most of the Warrior and Plateau coal fields. 

This sequence part of McCalley's Lower Measures 

(1891, p. 20) is 0 to 150 feet thick along much of 

the outcrop and thickens to more than 600 feet in the 

subsurface in the western part of the Warrior coal 

field. Mellen (1947, p. 1813) noted that this sequence 

should be called Parkwood Formation in western Alabama, where its fauna has both Mississippian and 

Pennsylvanian aspects. The author agrees, and regards the Parkwood as including all rocks in Alabama 

between the basal conglomeratic sandstone of the Pottsville Formation and the Pennington Formation or 

Floyd Shale. The rocks in Tennessee equivalent to 

the Parkwood Formation constitute the Raccoon 

Mountain Formation.

The deposition of the Parkwood Formation is interpreted as beginning in Mississippian time in the 

geosynclinal trough that included the present Cahaba 

and Coosa fields, and continuing into Pennsylvanian 

time. In early Pennsylvanian time, Parkwood sediments were spread over northern Alabama, but a few 

islands existed where no Parkwood was deposited, 

such as Boyles Gap. Subsequently, the basal conglomeratic members of the Pottsville Formation were 

deposited on the Parkwood, filling channels in places

but elsewhere being deposited conformably on the 

Parkwood Formation. The land surface at Boyles 

Gap rose slowly during the deposition of the Boyles 

Sandstone Member, resulting in a much thinner basal 

unit of the Pottsville than elsewhere in Alabama.

REFERENCES

Butts, Charles, 1910, Description of the Birmingham quadrangle, Alabama: U.S. Geol. Survey Geol. Atlas, Folio 175.

 1926, The Paleozoic rocks, in Adams, G. I., Butts, 

Charles, Stephenson, L. W., and Cooke, C. W., Geology of 

Alabama: Alabama Geol. Survey Spec. Rept. 14, p. 40- 

230.

 1927, Description of the Bessemer and Vandiver quadrangles, Alabama: U.S. Geol. Survey Geol. Atlas, Folio 

221.

 1940, Description of the Montevallo and Columbiana

quadrangles, Alabama: U.S. Geol. Survey Geol. Atlas, 

Folio 226.

McCalley, Henry, 1891, Report of the Coal Measures of the 

Plateau region of Alabama, including a report on the 

Coal Measures of Blount County by A. M. Gibson: Alabama Geol. Survey Spec. -Rept. 3, 238 p.

Mellen, F. F., 1947, Black Warrior basin, Alabama and Mississippi : Am. Assoc. Petroleum Geologists Bull., v. 31, no. 

10, p. 1801-1816.

Moore, R. C., Chairman, 1944, Correlation of Pennsylvanian 

formations of North America: Geol. Soc. America Bull., 

v. 55, p. 657-706.

Rothrock, H. E., 1949, Geology and coal resources of the northeast part of the Coosa coal field, St. Clair County, Alabama : Alabama Geol. Survey Bull. 61, 163 p.

Semmes, D. R., 1929, Oil and gas in Alabama: Alabama Geol. 

Survey Spec. Rept. 15, 400 p.

Shotts, R. Q., 1961, Correlations in the "Coal Measures" of the 

Southeast: Alabama State Mine Expt. Sta. Tech. Rept. 

23, p. 427-446.
ARTICLE 194 

194. PENNSYLVANIAN NOMENCLATURE IN NORTHWEST GEORGIA

By WILLIAM C. CTJLBERTSON, Denver, Colo.

E51

Published reports give four different nomenclatures 

for the rocks of Early Pennsylvania!! age in northwest Georgia. The purpose of this article is to rename the rock units of Georgia according to their 

correlation with the rock units of Tennessee, and to 

show their correlation with the equivalent rocks in 

the adjacent part of Alabama.

In northwest Georgia, rocks of Early Pennsylvanian age are about 1,000 feet thick and are the capping rocks on synclinal Sand, Lookout, and Pigeon 

Mountains, and on several small hills (fig. 194.1). 

The youngest Pennsylvanian rocks are preserved in 

Lookout Mountain near Durham. The rocks consist 

of resistant thick massive sandstone beds, at least one 

of which is conglomeratic, interbedded with weak sequences consisting of beds of shale, fine-grained sandstone, and locally a few beds of coal. The rock units 

vary in thickness from place to place, and locally 

sandstone beds fill channels in underlying rocks.

The Pennsylvanian nomenclature of northwest 

Georgia has been strongly influenced by nomenclatures 

published for adjacent Tennessee (see table). The 

first was that of Safford, who in 1869 divided the 

Coal Measures at the Aetna mines in Tennessee (fig.

194.1. loc. 1), 2 miles from Georgia, into Lower and 

Upper Coal Measures separated by a Lower Conglomerate and an Upper Conglomerate. The Upper Conglomerate is only a sandstone at its type locality (fig.

194.2. col. Ib), but it is equivalent to the bed that 

Safford called the Conglomerate in the rest of Tennessee, and that he later called Sewanee Conglomerate 

(Safford, 1893, p. 89-98).

The first nomenclature applied directly to northwest Georgia was that of Hayes (1892, p. 49), who 

divided the Coal Measures into two formations, the 

Lookout Sandstone below and Walden Sandstone 

above. The contact was placed at the top of "a bed 

of conglomerate 10 to YO feet in thickness which is a 

constant feature throughout the greater part of the 

field." Nowhere in this report on northeast Alabama, 

northwest Georgia, and southernmost Tennessee, nor 

in later reports on this same area, did he mention a 

second conglomerate or indicate how his units correlated with those of Safford. A few years later, 

McCallie in a report on the coal deposits of northwest 

Georgia said that Hayes' Lookout Sandstone corresponded to Safford's Lower Coal Measures "including

EXPLANATION

Core hole Measured section

Composite section

General locality

R.8E. R.10E.

10 MILES

FIGURE 194.1. Index map showing areas of outcropping Pennsylvanian rocks (shaded area) and location of stratigraphic 

sections.
E52 GEOLOGICAL SURVEY RESEARCH 1962

History of vtratigraphic nomenclature of Pennsylvanian rocks in northwest Georgia and South Tennessee

1 CARBONIFEROUS

55 <!

M

55

2 J 

^

CO

W PM

55

d

s

£j

03 

03

03 

03 W

a

Aetna (Etna) 

mines, Tenn. 

(SafEord, 1869, 

p. 383-384; 1893)

(Unit missing)

Upper Coal 

Measures

Upper Conglomerate (Sewanee 

conglomerate)

Unnamed unit

Lower 

Conglomerate 

(Cliff Rock)

Lower 

Coal Measures

Mountain 

limestone

Northwest 

Georgia, 

northeast 

Alabama, and 

south 

Tennessee 

(Hayes, 1892, 

p. 49)

(Unit missing)

Walden 

sandstone

Lookout 

sandstone

Bangor 

limestone

South Tennessee 

<Nelson, 1925, 

p. 39)

Lee group

Rockcastle 

conglomerate

Vandever 

shale

Newton 

sandstone

Eastland 

shale lentil i

Herbert 

conglomerate >

Whitwell 

shale

Sewanee 

conglomerate

Unnamed 

member

a.2 

U 

9 Warren 

 3 Point 

12 sandstone

S 

O Unnamed 

member

Pennington 

formation

Bangor limestone

Geologic map of 

Georgia (Georgia 

Div. of Mines, 

Mining, and 

Geology, 1939)

(Unit missing)

Potts ville 

formation

Pennington shale

Bangor limestone

Northwest Georgia, 

Chart 6 of 

Pennsylvanian 

Subcommittee 

(Moore, 1944)

(Unit missing)

Walden 

formation

Upper 

conglomerate

g ~| Unnamed 

g member

** 

g Lower 

§ conglomerate

Unnamed 

member

Pennington 

formation

Not shown

Northwest Georgia 

(Johnson, 1946)

(Unit missing)

Rockcastle 

sandstone

Vandever 

shale

Bonair 

sandstone

Whitwell 

Shale

a Sewanee 

12 member

1

g Gizzard 

 ^ member

3

Pennington 

shale

Bangor limestone

South Tennessee 

(Wilson and others, 

1956, p. 1)

Crab Orchard Mountains group

a 

o

 a

5

Rockcastle 

conglomerate

Vandever 

formation

Newton 

sandstone

Whitwell 

shale

Sewanee 

conglomerate

Signal Point 

shale

Warren Point 

sandstone

Raccoon 

Mountain 

formation

Pennington 

formation

Not shown

Northwest Georgia 

(this article)

(Unit missing)

Formation Orchard Mountains Crab

Gizzard Formation

Vandever 

Member

Newton 

Sandstone 

Member

Whitwell 

Shale 

Member

Sewanee 

Member

Signal Point 

Shale Member

Warren Point 

Member

Raccoon 

Mountain 

Member

Pennington 

Shale

Bangor Limestone

1 According to Wilson and others (1956, p. 4) the Eastland Shale Lentil and Herbert Conglomerate of Nelson are equivalent to the Whitwell Shale and Newton Sandstone 

of Nelson, so the names Eastland and Herbert are discarded.

the Lower Conglomerate and underlying rocks" 

(MoCallie, 1904, p. 19, 99).

Nelson (1925) published a new detailed nomenclature for southern Tennessee in which he recognized 

the Mississippian Pennington Formation in rocks previously included in the Lower Coal Measures of Safford (1869). In the second change in Georgia nomenclature, the Georgia Division of Mines, Mining, and 

Geology (1939) recognized the Pennington Shale in 

Georgia, but placed all beds above the Pennington 

Shale into an undivided Pottsville Formation. This

usage was followed by Sullivan (1942, p. 32). In a 

third change in the nomenclature for Georgia, the 

Pennsylvanian Subcommittee of the National Research 

Council Committee on Stratigraphy (Moore, 1944, 

chart 6) divided the rocks into the Lookout Formation and Walden Formation, as did Hayes (1892), 

but they included an Upper Conglomerate and a 

Lower Conglomerate within the Lookout Formation. 

Two years later, Wanless (1946, p. 8), in a comprehensive study of the Southern Appalachian coal field, 

used essentially the same nomenclature and also cor-
ARTICLE 194 E53

related the rock units with the named units of Nelson 

in Tennessee.

The fourth nomenclature for northwest Georgia was 

published by Johnson (1946) as a result of a geologic 

investigation of Sand and Lookout Mountains, Ga., 

in connection with a core-drilling program by the 

U.S. Bureau of Mines. His names are derived from 

Nelson's (1925) nomenclature for Tennessee, but also 

include the Lookout Sandstone of Hayes.

It is my conclusion that (1) the Lookout Formation 

as defined by the Pennsylvanian Subcommittee is a 

different unit than that defined by Hayes at the type 

locality, and (2) Johnson's nomenclature is based on 

a miscorrelation with the rock units of Tennessee. In 

the type area (Lookout Mountain, Ga.) the top bed of 

the Lookout Sandstone is the Lower Conglomerate, 

not the Upper Conglomerate as shown by the Pennsylvanian Subcommittee; and the top bed correlates 

with the Warren Point Sandstone, not the Sewanee 

Conglomerate as implied by Johnson's use of the name 

Sewanee Member of the Lookout Sandstone.

These conclusions were reached during a reconnaissance trip into Lookout, Sand, and Raccoon Mountains, using Johnson's geologic map (Johnson, 1946), 

his annotated air photographs, and other published 

reports. On Lookout Mountain, the bed Johnson calls 

Sewanee Member of the Lookout Sandstone is the 

"conglomerate and heavy sandstone" at the top of 

Hayes' Lookout Sandstone at Johnson Crook, and is 

the same bed that Wanless (1946, p. 24) calls Warren 

Point Sandstone Member of the Lookout Sandstone 

(fig. 194.2, cols. 2, 3, and 4). On Sand Mountain at 

Cole City, Ga., the Sewanee Member of Johnson is 

the top bed of the Lookout Sandstone of'Hayes (1894), 

the Lower Conglomerate of McCallie (1904, p. 98) 

and McCalley (1891, p. 44), and the Warren Point 

Sandstone Member of Wanless (1946, pi. 21, col. 2). 

Johnson's map also adjoins the type locality (fig. 

194.1, loc. 8) of the Raccoon Mountain Formation of 

Wilson and others (1956, p. 19), where the Sewanee 

Member of Johnson directly overlies the Raccoon 

Mountain Formation and is thus the same unit as 

their Warren Point Sandstone (see table). At Aetna 

Mines on Raccoon Mountain, a comparison of the measured sections of Hayes, Safford, and Nelson (fig. 194.2, 

cols, la, Ib, and 1) also confirms the identity of the 

top bed of the Lookout Sandstone with the Lower 

Conglomerate, a fact that was noted in 1904 by J. J. 

Stevenson (p. 125), and with the Warren Point Sandstone Lentil of Nelson (1925).

Many authors have stated that the Lookout Sandstone includes the Sewanee at its top, but this is true 

only in Tennessee, exclusive of the southernmost part,

where the Sewanee is largely conglomeratic and the 

older Warren Point is only a sandstone. In the type 

area described above and in southernmost Tennessee, 

Hayes placed the top of the Lookout at the top of 

the conglomeratic Warren Point Member or its equivalent, and included the generally nonconglomeratic 

Sewanee Member in the Walden Sandstone. Although 

Johnson placed the top of the Lookout Sandstone at 

the same horizon as Hayes, the result of his miscorrelation of the top bed is that his named units do 

not correlate with the units of the same name in 

Tennessee; for example, the Vandever Shale in Georgia is a lower bed than the Vandever Shale of Tennessee.

Johnson (1946) has demonstrated that rocks of 

Pennsylvanian age in northwest Georgia contain many 

recognizable units which deserve formal recognition. 

A new nomenclature is here defined for northwest 

Georgia that consists of 2 formations and 7 members. 

The names correspond generally to the names of the 

equivalent rock units in adjacent Tennessee as described in the recent study of Pennsylvanian rocks by 

Wilson and others (1956) or by Nelson (1925).

The table and figure 194.2 (cols. 5, 6, and 7) show 

the new nomenclature and its correlation with the 

generalized sections of Johnson for northwest Georgia, 

a composite section of Lookout Mountain, Ala., and a 

composite section in Tennessee that was adapted from 

Wilson and others (1956), using their description of 

the units on Walden Ridge and Lookout Mountain 

wherever possible. The section on Lookout Mountain, 

Ala., and a description of the rocks were furnished 

by N. M. Denson and R. K. Hose, who, in 1944, made 

a geologic investigation of coal-bearing rocks on Lookout Mountain near Fort Payne. The only published 

result of this investigation is a chapter entitled "Description of deposits" by N. M. Denson in a U.S. 

Bureau of Mines report (Coulter, 1947, p. 4-5).

The new members in northwest Georgia are described in detail below, primarily on the basis of the 

descriptions of Johnson (1946), supplemented by the 

measured section of Wanless (1946, p. 24 and 32) and 

by personal observation. The thicknesses of the units 

are derived from Johnson (1946) because most of 

his data are based on both core drilling and fieldwork. 

Measurement of the rocks in northwest Georgia has 

always been hampered by poor exposures, with the 

result that different authors have reported widely 

varying thickness for the same interval. Published 

reports are in essential agreement as to the thickness 

of the upper 400 to 500 feet of the Pennsylvanian 

section, but the thickness of the interval between the 

base of Lower Conglomerate (Castle Rock coal zone)
E54

FEET 

r 0

-100

-200

-300

-400

L 500

GEOLOGICAL SURVEY RESEARCH 1962 

EXPLANATION

Conglomeratic sandstone 

or conglomerate

Sandstone Shale Sandy shale, siltstone, 

or shaly sandstone

Coal bed Covered

AETNA (ETNA) MINES, TENNESSEE

la

(Hayes, 1894)

lb

Safford (1869, 

p. 383-384)

1C

Nelson (1925, 

p. 149, p. 44)

Correlation line

Dashed where uncertain; 

queried where unknown

LOOKOUT MOUNTAIN, GEORGIA 

3

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(Johnson, 1946)

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FIGURE 194.2. Columnar sections of Pennsylvania strata. See figure 194.1 for locations. Formation names of each author 

shown on left side of column, coal bed names on right except where labeled otherwise. Wanless' correlation of the
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E56 GEOLOGICAL SURVEY RESEARCH 1962

and the No. 4 or Tatum coal bed on Lookout Mountain 

has been variously reported to be 420 feet (Johnson, 

1946), 675 feet (Hayes, 1894), 780 feet (Wanless, 

1946, pi. 25, col. 9), and 1,200 feet (Sullivan, 1942, 

p. 45). As Johnson's measurement is derived from 

a core hole it should be accurate. The larger thicknesses probably were the result of miscalculating the 

thickness of rocks underlying the weathered plateau 

surface of Lookout Mountain.

GIZZARD FORMATION

The Gizzard Formation was named from its exposures on Little Fiery Gizzard Creek in Marion 

County, Tenn., 2 miles south of Tracy City in Grundy 

County (Safford, 1869; Nelson, 1925). In northwest 

Georgia the Gizzard Formation contains three members: the Raccoon Mountain, the Warren Point, and 

the Signal Point Shale.

Raccoon Mountain Member. In northwest Georgia 

the Raccoon Mountain Member is the sequence of 

gray shale, fine-grained sandstone, and lenticular coal 

beds that lie unconformably above the Mississippian 

Pennington Shale and below the cliff-forming conglomeratic sandstone called Warren Point Member. 

Wilson and others (1956) named the Raccoon Mountain Formation from its exposure at locality 8 (fig. 

194.1), about 0.4 mile northwest of Whiteside, Tenn. 

It varies considerably in thickness, ranging from 353 

feet at locality 8 to about 70 feet on the west side of 

Lookout Mountain, Ga. On Lookout Mountain, Ala., 

the equivalent beds average about 50 feet in thickness 

and do not exceed 150 feet. The coal bed that occurs 

at or near the top of this unit is widely known as the 

Etna, Cliff, or Castle Rock coal bed.

Warren Point Member. The Warren Point Member is the massive quartzose (orthoquartzite) sandstone containing scattered to abundant small pebbles 

of quartz or quartzite that forms the main cliff face 

on Lookout and Sand Mountains. Its type locality in 

Tennessee is at Warren Point, i/2 mile north of Monteagle, Grundy County, Tenn. It ranges from about 

50 to 200 feet in thickness. The equivalent bed in 

the Walden Ridge area of Tennessee is reported by 

Wilson and others (1956, p. 4) to be 200 feet thick. 

A comparison of this thickness with the 100-foot thickness shown by Nelson (fig. 194.2, col Ic) indicates 

that the top of the Warren Point is placed at a higher 

horizon by Wilson and others. Using their definition, 

the bed equivalent to the Warren Point Member in 

Lookout Mountain, Ala., is 100 to 150 feet thick in 

most places, ranging up to as much as 250 feet thick,

and contains shale partings in the upper part. The 

Underwood coal bed and associated shale (fig. 194.2, 

col 7) are present only locally.

/Signal Point Shale Member. The Signal Point 

Shale Member in northwest Georgia is 6 to 50 feet 

thick and consists of gray shale with locally a thin 

coal bed and a thin bed of sandstone. The type locality of the Signal Point Shale Formation, as named 

by Wilson and others (1956), is on Tennessee State 

Highway 8 at locality 9 (fig. 194.1). It crops out 

around the margin of Lookout Mountain, Ga. The 

equivalent sequence in Lookout Mountain, Ala., consists of 15 to 50 feet of shale and fine-grained sandstone and 2 coal beds that have been mined intermittently (fig. 194.2, col. 7).

CRAB ORCHARD MOUNTAINS FORMATION

Wilson and others (1956) named the Crab Orchard 

Mountains Group from its exposures in the Crab 

Orchard Mountains, Cumberland County, Tenn. In 

northwest Georgia the Crab Orchard Mountains Formation consists of four members: the Sewanee, Whitwell Shale, Newton Sandstone, and Vandever.

Sewanee Member. The Sewanee Member (Johnson's Bonair Sandstone) is, according to Johnson, a 

sandstone 150 to 200 feet thick, that is similar in 

lithology to the Warren Point Member described 

above except that it does not contain pebbles, and is 

less resistant to weathering. According to Wanless 

(1946, p. 24), it contains a few pebbles on the west 

brink of Lookout Mountain, and according to Denson 

(oral communication, 1962) it contains pebbles on 

Pigeon Mountain. Its type locality is at Sewanee, 

Franklin County, Tenn., about 21 miles west of locality 10 (fig. 194.1). It ranges from 150 to 200 feet 

in thickness in Georgia and is the surface rock on 

most of Look-out Mountain, Ga. On Lookout Mountain, Ala., it is a coarse-grained sandstone as much as 

150 feet thick with a conglomeratic basal part that 

averages 60 feet in thickness.

Whitwell Shale Member. The Whitwell Shale 

Member is a shale and sandy shale sequence, 150 to 

200 feet thick, that underlies about 10 square miles 

in the center of Lookout Mountain, Ga. Its type locality is at Whitwell, Tenn. (fig. 194.1, loc. 10). It 

contains the No. 4 or Tatum coal bed that has been 

mined at several places on Lookout Mountain. On 

Lookout Mountain, Ala., the sequence above the Upper 

Conglomerate includes the thin Sewanee and Tatum 

coals, and is probably equivalent to the Whitwell 

Shale Member.
ARTICLE 195 E57

Newton Sandstone Member. On Lookout Mountain, 

Ga., the coarse crossbedded bench-forming sandstone, 

about 110 feet thick, that underlies the prominent Durham coal bed is named the Newton Sandstone Member. Its type locality is at Newton in southwest 

Cumberland County, Tenn.

Vand&ver Member. The upper 300+ feet of shale 

and sandstone exposed on Lookout Mountain is named 

the Vandever Member from its correlation with the 

Vandever Shale, whose type locality is at Vandever, 

Cumberland County, Tenn. It contains the extensively mined thick Durham coal bed at its base.

REFERENCES

Coulter, D. M., 1947, Coking coal deposits on Lookout Mountain, DeKalb and Cherokee Counties, Alabama: U.S. Bur. 

Mines, Kept. Inv. 4030, 89 p.

Georgia Division of Mines, Mining, and Geology, 1939, Geologic Map of Georgia : scale 1:500,000.

Hayes, C. W., 1892, Report on the geology of northeastern 

Alabama and adjacent portions of Georgia and Tennessee: 

Alabama Geol. Survey Bull. 4, 86 p.

 1894, Ringgold atlas sheet, Georgia and Tennessee: 

U.S. Geol. Survey Geol. Atlas, Folio 2.

Johnson, V. H., 1946, Coal deposits on Sand and Lookout

Mountains, Dade and Walker Counties, Georgia: U.S. 

Geol. Survey Prelim. Map.

McCalley, Henry, 1891, Report on the Coal Measures of the 

plateau region of Alabama: Alabama Geol. Survey Spec. 

Rept. 3, 238 p.

McCallie, S. W., 1904, A preliminary report on the coal deposits of Georgia: Georgia Geol. Survey Bull. 12, 121 p.

Moore, R. C., ch., 1944, Correlation of Pennsylvanian formations of North America: Geol. Soc. America Bull., v. 55, 

no. 6, p. 657-706.

Nelson, W. A., 1925, The Southern Tennessee coal field: Tennessee Div. Geol., Bull. 33-A, 239 p.

Safford, J. M., 1869, Geology of Tennessee : Nashville, Tennessee, 

550 p.

 1893, The topography, geology, and water supply of 

Sewanee, Franklin County, Tennessee: Tennessee State 

Board of Health Bull., v. 8, no. 6,, p. 89-98.

Stevenson, J. J., 1904, Carboniferous of the Appalachian Basin: 

Geol. Soc. America Bull., v. 15, p. 37-210.

Sullivan, J. W., 1942, The geology of the Sand-Lookout Mountain area, northwest Georgia: Georgia Div. Mines, Mining 

and Geology Inf. Circ. 15, 68 p.

Wanless, H. R., 1946, Pennsylvania!! geology of a part of the 

Southe'rn Appalachian coal field: Geol. Soc. America Mem. 

13, 162 p.

Wilson, C. W., Jr., Jewell, J. W., and Luther, E. T., 1956, 

Pennsylvanian geology of the Cumberland Plateau: Tennessee Div. Geol. Folio (Nashville), 21 p.

195. PATHFINDER UPLIFT OF PENNSYLVANIAN AGE IN SOUTHERN WYOMING

By WILLIAM W. MALLORY, Denver, Colo.

The Uncompahgre, Front Range, and Zuni-Defiance 

uplifts have long been recognized as the principal tectonic elements of the ancestral Rocky Mountains during the Pennsylvanian Period in Colorado and adjacent states. Recent compilation and synthesis of Pennsylvanian stratigraphy in Wyoming for the Pennsylvanian folio (in preparation) of the U.S. Geological 

Survey paleotectonic map series indicates that a fourth 

uplift of moderate size in southern Wyoming is related tectonically to the other three. It had its greatest extent and most pronounced topographic expression 

in Atoka (Middle Pennsylvanian) time when it extended from Laramie northward to Casper and westward into the southeast corner of Fremont County

(fig. 195.1). The name Pathfinder uplift,1 hereby 

proposed for this ancient land, is taken from Pathfinder Mountain in T. 29 N., R. 84 W., south-central 

Natrona County, Wyo., which is centrally located on 

the site of the uplift.

The interpretation of the Pathfinder uplift is based 

largely on a series of measured sections and careful 

paleontologic determinations by Love and others

i The name Pathfinder uplift was applied in 1932 by T. S. Levering 

to the structural element in central Wyoming now called the Sweetwater uplift (U.S. Geol. Survey and Amer. Assoc. Petroleum Geologists, 1961) or Granite Mountains uplift (Bndlich, 1879). This tectonic feature is largely a Laramide rejuvenation of the western 

segment of the uplift of Pennsylvanian age described here. Therefore, 

both Levering and the author consider that redefinition here of the 

name Pathfinder uplift is permissible.
E58 GEOLOGICAL SURVEY RESEARCH 1962

108° 107° 106" 105°

20 40 MILES

_J_____I

FIGURE 195.1. Index map showing major features and 

localities. Pathfinder uplift and extreme northern 

end of Front Range uplift of ancestral Rocky Mountains of Atoka time, solid-line pattern; Sweetwater 

(or Granite Mountains) uplift of Laramide age, 

broken-line pattern; outcrop area of Precambrian 

rocks in present ranges of the Rocky Mountains, 

stipple; Laramie Range occupies area between 

Laramie and Casper.

(1953), Thomas and others (1953), and Love (1954) 

in the vicinity of the Laramie Range and the Sweetwater uplift. Additional evidence is provided by 

sample logs of wells in the area, most of which were 

prepared by the American Stratigraphic Co. of Denver. Several other recently published reports give 

information on the fossil content and correlation of 

Pennsylvania!! rocks in southern Wyoming or nearby 

areas (Agatston, 1954; Henbest, 1954, 1956; Mallory, 

1960; and Maughan and Wilson, 1960). The history 

of the uplift from its inception in Morrow (Early 

Pennsylvanian) time to its disappearance in Virgil 

(Late Pennsylvanian) time is traced in the Pennsylvanian folio; only its status in Atoka time is described in this article.

The Amsden Formation of Mississippian and Pennsylvanian age extends throughout much of Wyoming, 

where it is divided into an upper unit 0 to 200 feet 

thick, and a lower red clastic unit 0 to 150 feet thick.

Except for local lenses of clastic sedimentary rock, 

the upper part of the Amsden Formation in Wyoming 

is a carbonate unit. Its age is essentially Atoka although locally it may include strata a little older or 

younger. Its average thickness is about 150 feet. The 

unit is present throughout most of Wyoming except 

on the Pathfinder uplift, an area of approximately

7,300 square miles in the southern part of the State, 

where rocks of Atoka age are absent. Three lines of 

evidence, namely the thickness and the lithology of 

the upper part of the Amsden and the overlap of 

younger beds, support the concept that the Pathfinder 

uplift area was emergent and tectonically positive in 

Atoka time. At 30 wells and surface sections on the 

site of the uplift, rocks of Atoka age are lacking. 

An additional 35 control points in the surrounding 

area show that the upper part of the Amsden Formation thickens uniformly away from the uplift nearly 

everywhere at a rate of 10 feet per mile; locally in 

southern Albany County the rate of thickening is 

much greater, and is 65 feet per mile just south of 

Laramie (fig. 195.2). If the uplift were post-Atoka 

in age, isopachs of Atoka strata would be expected 

to have a pattern discordant at least to some degree 

with the outline of the uplift.

Four distinct divisions of the Pathfinder uplift can 

be designated on the basis of areal geology at the end 

of Atoka time (fig. 195.2): (1) a northeastern segment where Madison Limestone of Mississippian age 

was exposed, (2) a central belt where sandstone and 

shale largely of Morrow age (the lower part of the 

Amsden Formation) were exposed, (3) a southern 

extension where igneous and metamorphic rocks of 

Precambrian age were exposed, and (4) a western 

extension where the Madison Limestone (Mississippian) probably formed all or much of the surface. 

Tertiary rocks lie on Precambrian rocks over much or 

all of the area of the Sweetwater uplift, a Laramide 

tectonic feature comprising most of the western extension of the Pathfinder uplift. On the margins of 

the Sweetwater uplift, however, Pennsylvanian strata 

rest on the Madison Limestone of Mississippian age. 

Hence the Madison probably extended partly or entirely over this part of the Pathfinder uplift at the 

close of Atoka time. Strong Laramide upward movement of the Sweetwater uplift caused removal of all 

or most rocks younger than Precambrian and older 

than Tertiary; therefore, documentable interpretation 

of the paleogeology of the western extension at the 

close of Atoka time is not possible. Figure 195.2 

shows Madison Limestone on the western extension of 

the Pathfinder uplift, but local inliers of older Paleozoic and basement rocks may have been present.

The lithology of the upper part of the Amsden 

Formation surrounding the uplift is directly related 

to the rocks exposed on the uplift in Atoka time and 

to the tectonic activity in each of the four divisions 

designated on figure 195.2. The western extension is 

fringed with local aprons of sandy and shaly strata 

that grade away from the uplift into the regional
ARTICLE 195 E59

108° 107° 106° 105°

EXPLANATION

AREAL GEOLOGY OF PATHFINDER UPLIFT

/ \ x ^ / ' \ V > I - ',- - -: / H \Z C/o.«~ <f* >< >«

.^irii^^^^^S^^i'-SM'-' 1

^.^^-;^^^A^^?:Y

Red shale and sandstone of Morrow age 

(lower part of Amsden Formation)

Y////A

Madison Limestone of Mississippian age

Precambrian granitic and metamorphic rocks

LITHOLOGY OF UPPER PART OF AMSDEN 

FORMATION

Arkosic conglomerate, sandstone, and

siltstone 41

Sandstone

Approximate contact

Shaly sandstone

Calcareous sandstone

i _ _ | 

Shale

Carbonate rock

Isopach

Showing thickness, in feet, of rocks of Atoka 

age. Interval is 50 feet

0125 A 180

Well Surface section

Control points

Number is thickness of upper part of Amsden Formation. P, paleontologic data are available from 

Pennsylvania^ rocks to aid in correlation

10 

I

20 MILES

FIGURE 195.2. Paleogeologic map showing areal geology of Pathfinder uplift at end of Atoka time and thickness and lithology of 

strata of Atoka age (upper part of Amsden Formation) in area surrounding the uplift.

667422 O - 63 -
E60 GEOLOGICAL SURVEY RESEARCH 1962

carbonate facies. The southern extension, where igneous and metamorphic rocks were exposed, is surrounded by arkose of Atoka age. The northeastern 

segment, where the Madison Limestone was exposed, 

also is fringed with sandy and shaly strata which 

grade outward into the regional carbonate facies. 

Along the northern and southern shores of the central 

belt, where Morrow strata (sandstone and shale of the 

lower part of the Amsden) were exposed, the regional 

carbonate facies of the upper part of the Amsden 

flanks the uplift.

The aprons of sand and shale that were deposited 

locally on the fringes of the western extension of 

the Pathfinder uplift indicate that only moderate uplift occurred in that area. Shale and sandstone of 

the lower part of the Amsden Formation, and perhaps 

locally some Paleozoic strata older than Morrow, were 

stripped off and redeposited in the Atoka sea.

The southern extension of the Pathfinder uplift in 

Albany County was strongly uplifted, allowing igneous 

and metamorphic rocks of Precambrian age to be 

actively weathered, eroded, and deposited on the southern shores of the uplift as arkose, which attains a 

thickness of 300 feet south of Laramie. The northeastern segment was moderately uplifted, allowing 

shale and sandstone of the lower part of the Amsden 

to be removed and redeposited in the Atoka sea. The 

Madison was then exposed to some degree of subaerial wasting, but uplift was not great enough to 

result in stripping away this formation. The central 

belt must have been a lowland nearly awash in marine 

waters, because clastic strata of the lower part of the 

Amsden were not removed and adjacent clear marine 

waters accordingly were not contaminated with mud 

and sand.

The Pathfinder uplift was progressively submerged 

during Des Moines, Missouri, and Virgil time, when

its influence was limited to controlling the direction 

of longshore currents and patterns of local sand deposition in the Tensleep, Casper, and Minnelusa Formations.

REFERENCES

Agatston, R. S., 1954, Pennsylvanian and Lower Permian of 

northern and eastern Wyoming: Amer. Assoc. Petroleum 

Geologists Bull., v. 38, no. 4.

Endlich, F. M., 1879, Report on the geology of the Sweetwater 

district: U.S. Geol. Geog. Survey Terr. (Hayden) llth 

Ann. Rept, p. 3-158.

Henbest, L. G., 1954, Pennsylvanian foraminifera in Amsden 

Formation and Tensleep Sandstone, Montana and Wyoming, in Billings Geol. Soc., Guidebook 5th Ann. Field 

Conf.: p. 50-53.

 1956, Foraminifera and correlation of the Tensleep 

Sandstone of Pennsylvanian age in Wyoming, in Wyoming 

Geol. Assoc. Guidebook llth Ann. Field Conf.: p. 58-63.

Love, J. D., 1954, Tentative diagrammatic correlation of Tensleep, Amsden, Casper, and Hartville Formations in Wyoming, in Wyoming Geol. Assoc. Guidebook 9th Ann. Field 

Conf.: chart (in pocket).

Love, J. D., Henbest, L. G., and Denson, N. M., 1953, Stratigraphy and paleontology of Paleozoic rocks, Hartville area, 

eastern Wyoming: U.S. Geol. Survey Oil and Gas Inv. 

Chart OC-44.

Levering, T. S., 1932, Field evidence to distinguish overthrusting from underthrusting: Jour. Geology, v. 15, no. 7, 

p. 656-658.

Mallory, W. W., 1960, Outline of Pennsylvanian stratigraphy 

of Colorado, in Guide to the geology of Colorado: Geol. 

Soc. America, Rocky Mtn. Assoc. Geologists, and Colorado 

Sci. Soc., p. 23-33.

Maughan, E. K., and Wilson, R. F., 1960, Pennsylvanian and 

Permian strata in southern Wyoming and northern Colorado, in Guide to the geology of Colorado: Geol. Soc. 

America, Rocky Mtn. Assoc. Geologists, and Colorado Sci. 

Soc., p. 34-42.

Thomas, H. D., Thompson, M. L., and Harrison, J. W., 1953, 

Stratigraphy of the Casper Formation, in Fusulinids of 

the Casper Formation of Wyoming: Wyoming Geol. Survey Bull. 46, pt. 1, p. 5-14.

U.S. Geological Survey and American Association of Petroleum 

Geologists, 1961, Tectonic map of the United States.
ARTICLE 196 E61

196. UNCONFORMITY MARKING THE JURASSIC-CRETACEOUS BOUNDARY IN THE LA LIGUA AREA,

ACONCAGUA PROVINCE, CHILE

By W. D. CARTER, Washington, B.C.

Work done in cooperation with the Institute de Investiffaciones Geoldfficas, Santiago, Chile

The recent discovery of ammonites of Early Cretaceous (Neocomian) age in the Patagua Formation 

places the Jurassic-Cretaceous boundary at the base 

of that formation (Carter and others, 1961). Subsequent mapping in the La Ligua quadrangle in the 

northwest part of Aconcagua Province, Chile (fig. 

196.1), shows that the strata of Cretaceous age rest 

on rocks of Jurassic age with angular discordance. 

This discordance, marked by successive northward 

overlap of Jurassic strata by the Patagua Formation 

(fig. 196.2) is not readily apparent because in many

70°

100 200 KILOMETERS

FIGURE 196.1. Index map of part of central Chile showing area of study.

places the attitude of the strata above and below the 

discordance is nearly the same.

Rocks of the Jurassic System were divided by 

Thomas (1958) into three formations: The Quebrada 

del Pobre, Ajial, and Melon Formations.

The Quebrada del Pobre Formation (oldest), which 

crops out east of La Ligua and extends northward 

into the valley of Eio Petorca, rests unconformably 

on the La Ligua Formation of Triassic age. It consists of gray arkosic sandstone and quartz-pebble conglomerate interlayered with dark-gray siltstone, calcareous siltstone, and silty limestone. The fine-grained 

units contain a rich faunal assemblage. The presence

of Arietites 'basulcatum, Brug. and Cardinia d&nsestriata, Jaworsky indicate that the Quebrada del Pobre 

Formation is of Early Jurassic (Lias) age. The type 

section is 1,250 meters thick (Thomas, 1958, p. 25). 

Similar thicknesses extend northward as far as the 

Rio Petorca where the sequence is truncated on the 

west by a granodiorite batholith of Jurassic age and 

on the east is cut by the Jurassic-Cretaceous unconformity and overlapped by Lower Cretaceous strata.

The Ajial Formation, consisting largely of volcanic 

lavas, breccias, and tuffs, crops out to the east and 

rests conformably on the Quebrada del Pobre Formation. Thin light-gray to green locally fossiliferous 

sandstone beds are interlayered with the volcanic 

rocks. The fossils indicate that the Ajial Formation 

is also of Early Jurassic age. The formation is as 

much as 1,300 meters thick in the La Ligua quadrangle (Thomas, 1958). At Engorda on the Rio 

Petorca in the northern part of the quadrangle, the 

formation is overlapped by the Patagua Formation.

The Melon Formation (Thomas, 1958), which rests 

conformably on the Ajial Formation, was divided 

into three members: The basal Nbgales Member, the 

Horqueta Member, and the upper Patagua Member. 

The Patagua Member, on the basis of fossil evidence, 

was redefined as the Patagua Formation by Carter 

and others (1961) and placed in the Cretaceous System. The nature of the basal contact of the Patagua 

was not clear then.

The Nogales Member, about 450 meters thick, is 

composed chiefly of marine sandstone and limestone 

at the type area near Nogales, 30 kilometers south of 

La Ligua. Ammonites and pelecypods of Middle 

Jurassic age abound in the strata. Among them, Corvalan identified Emileia (stephanoceras] , which is 

similar to E. singularis Gottsche, and indicative of the 

Bajocian stage of the Middle Jurassic (Thomas, 1958). 

Near the southern limit of the La Ligua quadrangle, 

the sedimentary rocks of the Nogales thin and interfinger with volcanic lavas, breccias, and tuffs very 

similar to those of the Ajial Formation. The Nogales 

is overlapped north of the Rio Petorca by the Patagua 

Formation.

The Horqueta Member is marked by a basal sandstone overlain by a distinctive sequence of volcanic 

breccia consisting of large blocky fragments. The
E62 GEOLOGICAL SURVEY RESEARCH 1962

EXPLANATION

:F- }-.. > 4- + -fv>« A If \. ' ' ' ' V^'SSvSl

XaLiguaX +Jgd -"i ^ v < "I-'.' ." '.". &81It\ +O-K + 4, f +J 7 ^ * f.- ./ tis&m

+ + * i<fVs,i±±? < ^ < v v/."-.\>//-'K«:i^

* * t^y^r/^ ^ V".- ; ->mtti^ * + "* />/ v?>^. > ,- N AX-'-' '^^ Vi^PSy

Granodiorite

32°30'

+ ^ + + ""MlLON QUADRANGLEl.V^iiHC/. 71°00' La Ligua Formation

Contact

Approximately located

Strike and dip 

of beds

10KILOMETERS

FIGURE 196.2. Generalized geologic map showing the unconformable relation between Jurassic and Cretaceous strata in the

La Ligua quadrangle and vicinity, Aconcagua Province, Chile.

member crops out as a narrow band of prominent 

cliffs. It is thickest, 1,700 meters (Thomas, 1958, p. 

36), at the type section on Cerro Horqueta, about 18 

kilometers southeast of La Ligua. Although Thomas 

shows the Horqueta extending at least 30 kilometers 

north of Cerro Horqueta, recent mapping (fig. 196.2) 

indicates that it is overlapped by Cretaceous strata 

near Cerro Iglesia, only about 8 kilometers to the 

north.

The oldest Cretaceous unit is the Patagua Formation, a thick series of marine limestone that rests unconformably on Jurassic rocks. The discovery of 

Thurmanniceras, and other index fossils of Valanginian age in Argentina and Chile, resulted in redefining Thomas' (1958) Patagua Member of the Melon 

Formation of Jurassic age as the Patagua Formation 

of Early Cretaceous age (Carter and others, 1961). 

At Cerro Iglesia the base of the Patagua Formation 

is marked by coarse-pebble to cobble conglomerate 

overlain by and, in places, interfingered with thin 

layers of compressed coquina composed largely of 

poorly preserved ammonites and pelecypods. Where 

exposed, both north and south of this locality, the 

base of the formation is usually marked by lenticular 

yellowish-brown medium- to fine-grained sandstone, 

siltstone-pebble conglomerate, and gray tuffaceous 

sandstone. These basal units are overlain by a series 

of thin-bedded limestones, siltstones, and fine-grained
ARTICLE 197 E63

sandstones intercalated with tuff, mainly in the upper 

part of the sequence. The total average thickness 

of the formation is about 1,100 meters (Carter and 

others, 1961, p. 1895).

Conglomerate lenses that mark the base of the 

Patagua Formation confirm that the contact with the 

underlying strata was an erosional surface, that is an 

unconformity. The strata above and below the contact show agreement of strike and dip in the southern 

half of the La Ligua quadrangle and southward 

through the adjacent Melon quadrangle (fig. 196.2). 

In these places the average strike is due north with 

dips of 25° to 35° E.

North of the Rio Petorca, however, the strike of the 

Cretaceous strata swings westward; the strata of 

Jurassic age are sharply truncated and show minor 

angular unconformability in dip. This suggests that 

the basal Cretaceous contact is a regional disconformity with local minor angular unconformity. Such

angular unconformity may be the result of orogeny 

associated with the intrusion of granitic masses of 

Jurassic age that crop out in the Coastal Cordillera of 

Chile. Part of one such mass is exposed at La Ligua 

and west of the Cretaceous contact. On the other 

hand, Ruiz and others (1961, p. 1558) have suggested 

that there may have been epeirogenic uplift throughout central Chile during Late Jurassic time.

REFERENCES

Carter, W. D., Perez, E., and Aliste, N., 1961, Definition and 

age of the Patagua formation, Province of Aconcagua, 

Chile: Am. Assoc. Petrol. Geol. Bull., v. 45, p. 1892-1896.

Ruiz, C., Aguirre, L., Corvalan, J., Rose, H. J., Jr., Segerstrom, 

K., and Stern, T. W., 1961, Ages of batholithic intrusions 

of northern and central Chile: Geol. Soc. America Bull., 

v. 72, p. 1551-1560.

Thomas, H., 1958, Geologia de la Cordillera de la Costa entre 

el valle de La Ligua y la Cuesta de Barriga: Inst. de 

Invest. Geologicas, Bull. 2, Santiago, Chile.

197. RELATIONS OF THE NAVAJO AND CARMEL FORMATIONS IN SOUTHWEST UTAH AND ADJOINING ARIZONA

By J. C. WEIGHT and D. D. DICKEY, Beltsville, Md., and Denver, Colo.

Work done in cooperation with the U.S. Atomic Energy Commission

The Navajo Sandstone of Late Triassic(?) and 

Jurassic age is overlain by the Carmel Formation of 

Middle and Late Jurassic age throughout most of 

Utah and northern Arizona. The Navajo is a crossbedded eolian sandstone and the Carmel an evenly 

bedded marine limestone, shale, and siltstone unit. 

The contact between the two formations is distinct 

because of the markedly contrasting lithologies. In 

much of central Utah the contact is a smooth plane, 

which truncates the crossbeds of the Navajo and represents a hiatus of unknown duration. But in southwestern Utah the two formations intertongue extensively (fig. 197.1, cross section). The largest tongue 

of Navajo appears to be a marginal marine, rather 

than eolian, deposit reworked from an older part of 

the Navajo Sandstone.

Several thin metabentonites near the base of the

Carmel provide a reliable datum which aids in detailed correlation of the formation.

Location of measured stratigraphic sections shown on figure 197.1

1

?,

3

4

5

fi

7

8

9

10

11

Locality

Pintura

Cave Canyon _ _

Mount Carmel Junction J _ .

Kanab Creek _ _

Lick Wash. __ __ _ _

Hackberry Canvon

Judd Hollow 12 L_-_---_-____

Wahweap Creek __ _ _

Block Mesa 1 .

Section

32

2

29

8

25

4

1

9

31

4

12

Township

40 S.

41 S.

39 S.

41 S.

41 S.

41 S.

40 S.

41 S.

43 S.

44 S.

43 S.

Range

17 W.

15 W.

13 W.

9 W.

8 W.

6 W.

4 W.

1 W.

1 E.

4 E.

8 W.

1 Discussed by Schultz and Wright (Art. 198).

2 Described by Phoenix (1963). 

» Lat 36°33'30" N., long 111°03' W.
EXPLANATION

Outcrop of base of 

Carmel Formation

Measured stratigraphic section

Test well reaching base 

of Carmel Formation

XQ

!*\/ 3 

^X 2

Line of section

IRON

^Antimony Canyon 

Carmel side of contact

GARFIELD

.__I

Cedar City c £f-s 

<0

O r-l

Bryce Canyon rf j 

National Park^J pWASHINGTON 

0 St George

MOHAVE

ARIZONA

5 6

FEET 

1-3000

-2000

Carmel 

Formation

-1000

Navajo 

Sandstone

60 MILES

VERTICAL EXAGGERATION X 132

EXPLANATION 

Crossbedded sandstone 

Red siltstone with some sandstone 

Red claystone with some siltstone

Gray calcareous shale

Gray limestone and dolomite

Gypsum

Bentonite

Circle sriows sample position

Regional angular unconformity

.

__ __ 

COCONINOParia PI

FIGURE 197.1. Map and section showing contact of the Carmel Formation and Navajo Sandstone in southwestern Utah and adjoinin

denotes anticline in Navajo, referred to in text, that was beveled during deposition of Thousand Pockets Tongue of 
ARTICLE 197 E65

NAVAJO SANDSTONE

The Navajo Sandstone extends over most of the 

Colorado Plateau, and in Zion National Park the 

white and red cliffs reach spectacular heights of more 

than 2,000 feet. The sandstone is composed almost 

entirely of quartzose well sorted fine- to mediumgrained sand, and contains wedge sets 30 to 50 feet 

thick of high-angle, sweeping, gently arcuate crossbeds. The character of the sandstone is uniform 

through its whole thickness and extent, and it has 

been consistently interpreted as an eolian deposit.

The Navajo thickens westward across both Utah 

and northern Arizona, from 800 feet southeast of Kaibito to about 2,300 feet in Zion National Park (fig. 

197.1, cross section). In southwestern Utah the formation intertongues along both its lower and upper 

contacts. The Lamb Point Tongue at the base has 

been described by Wilson (Averitt and others, 1955, 

p. 2521).

TEMPLE CAP MEMBER

The Temple Cap Member (Gregory, 1950, p. 89) of 

the Navajo Sandstone is a tonguelike unit about 150 

feet thick which forms a prominent cap on East Temple, West Temple, and many other buttes at the entrance to Zion Canyon. The member consists chiefly 

of sandstone nearly identical in lithology and character of bedding to the rest of the Navajo, but separated from it by a bed of siltstone, 5 to 25 feet thick. 

The bed weathers to form a red tree-covered bench or 

recess at the base of the member. The Temple Cap 

Member cannot be distinguished from the rest of the 

Navajo much farther east than Mount Carmel Junction because the basal siltstone pinches out. The 

western limit is in the gap in outcrops between Zion 

National Park and the Pine Valley Mountains.

THOUSAND POCKETS TONGUE

The Thousand Pockets Tongue (Phoenix, 1963) of 

the Navajo Sandstone intertongues with the Carmel 

throughout the Paria Plateau and Paria Valley. The 

tongue was recognized independently at about the 

same time by D. A. Phoenix in geologic mapping, by 

Kathleen MacQueen (1957) in photogeologic mapping 

of the Paria Plateau, and by the authors of this article in stratigraphic studies at the north end of the 

Paria Valley. The Thousand Pockets Tongue has 

since been traced along the outcrop throughout central 

Kane County. Probable correlatives, to which the 

name cannot yet be confidently applied, are found in 

eastern Kane County and in central Garfield County.

From the area of the Glen Canyon Dam, where the 

tongue is about 200 feet thick, it extends north and 

northwestward more than 50 miles to where it pinches

out. The tongue is composed of well-sorted quartzose 

sandstone, lithologically similar to the Navajo Sandstone but deposited in even beds as well as in crossbeds. A few interbeds of reddish flat-bedded silty 

sandstone extend for thousands of feet within the 

tongue. Slumped and contorted crossbeds occur 

throughout the tongue. Toward the southeast, where 

it joins with the Navajo, large-scale sweeping crossbeds very similar to those in the Navajo are common. 

Toward the distal limits of the tongue, tabular sets 

of even strata and smaller crossbeds predominate. 

The tabular sets are truncated at the top along extensive smooth bedding planes.

The Thousand Pockets Tongue cannot be distinguished easily from the rest of the Navajo southeast 

of the pinchout of the underlying Judd Hollow Tongue 

of the Carmel Formation. However, an unusually 

persistent horizontal truncation plane, underlain by a 

foot or two of sandstone impregnated with black iron 

oxide, extends eastward a few miles from the pinchout on the Paria Plateau. The strata above the black 

band have tabular sets, even beds, and slump structures similar to those described in the Thousand 

Pockets Tongue. Similar features indicate less certainly that the uppermost part of the Navajo Sandstone, along the southeast and eastern fronts of the 

Kaiparowits Plateau, is probably equivalent to the 

Thousand Pockets Tongue. In central Garfield 

County near Escalante and Boulder, prominent interbeds of sandstone similar to beds of the Thousand 

Pockets Tongue are present both on outcrop and in 

test wells which penetrate the lower part of the 

Carmel Formation; these sandstone beds also probably correlate with the Thousand Pockets Tongue.

SOURCE OF THE THOUSAND POCKETS TONGUE

South of the Glen Canyon Dam, local erosion of 

the Navajo over a buried anticline prior to deposition 

of the Carmel Formation has been described by McClymonds (1961). The area from which a minimum 

of 25 feet of sand was eroded, was estimated from 

McClymond's isopach map of the Navajo Sandstone 

and plotted on figure 197.1. Probably more than 300 

feet of sand was eroded from the thickest, or northcentral, part of this area. It is inferred that the sand 

eroded from this anticline was redeposited to form 

the Thousand Pockets Tongue.

This interpretation implies that the material of the 

Thousand Pockets Tongue was transported northward 

for distances as great as 70 miles. The thinning of the 

tongue to the northwestward and the overall relation 

of the predominantly continental Navajo Sandstone 

to the marine Carmel Formation fit this concept. The
E66 GEOLOGICAL SURVEY RESEARCH 1962

flat bedding, the numerous smooth truncation planes, 

and the intertonguing of the Thousand Pockets Tongue with the marine Carmel are criteria that the tongue 

is an aqueous deposit (Baars, 1961, p. 187-189). 

Flame structures (Kelling and Walton, 1947, p. 484- 

485), observed in the tongue near Adairville, are consistently overturned and dragged to the north. This 

northward overturning is also consistent with the 

presumed northward slope of the surface at the time 

of deposition. All the observed features thus indicate 

northward transport of material in the Thousand 

Pockets Tongue, even though the southward dip of 

crossbeds nearly everywhere in the main mass of the 

Navajo indicates eolian transport in that direction 

(Poole, 1962; Wells, 1960, p. 126; Phoenix, oral communication, 1962).

CARMEL FORMATION

The Carmel Formation along the Hurricane Cliffs 

contains a thin lower sequence of gypsiferous red beds, 

a thick middle sequence of gray limestone and shale, 

and an upper sequence of gypsiferous red siltstone 

and sandstone. The lower red beds are stratigraphically equivalent to the Temple Cap Member of the 

Navajo Sandstone (Cook, 1957, p. 34). These red 

beds are only known in Washington County and 

southeastern Iron County.

Thin bentonite clay beds, altered from volcanic-ash 

falls, are widely distributed near the base of the 

Carmel Formation. They are dark purple or light 

yellowish green, and contrast conspicuously with the 

white or reddish-brown enclosing strata. The clay is 

in layers about half an inch to 2 inches thick at most 

outcrops, but locally is as much as several feet thick. 

The thicker pods presumably represent ash washed 

into shallow hollows soon after it fell. Near Gunlock 

and in the Pine Valley Mountains one bentonite is 

near the middle of the several-hundred-foot thickness 

of red beds in the basal part of the Carmel. Several 

slightly later bentonites, extending from Gunlock to 

Judd Hollow, are at the base of the limestone in the 

Carmel, enclosed in 6 inches to 5 feet of clastic strata 

probably reworked from the underlying Navajo. The 

mineralogy and petrology of these bentonites are discussed by L. G. Schultz and J. C. Wright (Art. 198).

The gray limestone and shale in the middle of the 

Carmel, and the gypsiferous red beds at the top, form 

an extensive twofold division of the formation (Gilluly and Reeside, 1928, p. 73-74). The red beds extend beyond the limey beds to the east and southeast.

In southwestern Utah, the Carmel and younger 

Jurassic formations are truncated by a regional angular unconformity and are overlain by Upper Cretaceous strata. Although this unconformity successively 

truncates the Morrison, Summerville, Entrada, and 

Carmel Formations, its angularity is only half a degree or less. It has been clearly recognized only near 

the Glen Canyon Dam.

JUDD HOLLOW TONGUE

The limestone of the lower part of the Carmel 

Formation passes southeastward beneath the Thousand Pockets Tongue of the Navajo to form the Judd 

Hollow Tongue of the Carmel (Phoenix, 1963). On 

the west side of the Paria Valley, the tongue is about 

100 to 200 feet thick; on the east side of the valley, 

about 50 to 100 feet; and near the head of Judd Hollow on the Paria Plateau, it pinches out. From west 

to east the limestone thins, grades from relatively 

pure carbonate to argillaceous limestone, and from 

gray to reddish gray. Interbedded red siltstone and 

sandstone increase eastward. The uppermost bed of 

limestone in the Judd Hollow Tongue, which is ripple 

marked and oolitic near the East Kaibab monocline, 

grades to ripple-marked limey siltstone that may be 

traced a little east of Judd Hollow. Fossils have been 

found in the tongue at the northwest end of the Paria 

Valley. Calcareous strata north of Escalaiite probably grade to very silty beds south of the town.

AGE RELATIONS

In southwest Utah, the limestone of the Carmel contains assemblages of fossil pelycopods and gastropods. 

The Carmel in central Utah, and equivalent units of 

the Twin Creek Limestone in northern Utah, contain 

these same assemblages, as well as critical ammonites, 

thus establishing their Bajocian or early Middle Jurassic age (R. W. Imlay and N. F. Sohl, oral communication, 1961). The main mass of the Navajo Sandstone is considered Late Triassic (?) on the basis of 

scanty paleontologic evidence (Lewis and others, 1961, 

p. 1439). However, the Temple Cap Member, which 

immediately underlies the bentonite at Mount Carmel 

Junction, and the Thousand Pockets Tongue are Jurassic because they intertongue with the Carmel Formation of definite Jurassic age.

REFERENCES

Averitt, Paul, Dettermann, J. S., Harshbarger, J. W., Repenning, C. A., and Wilson, R. F., 1955, Revisions in correlation and nomenclature of Triassic and Jurassic Formations in southwestern Utah and northern Arizona: Am. 

Assoc. Petroleum Geologists Bull., v. 39, p. 2515-2524.

Baars, D. L., 1961, Permian blanket sandstones of Colorado 

Plateau, in Geometry of sandstone bodies: Am. Assoc. 

Petroleum Geologists, p. 179-207.
ARTICLE 198 E67

Cook, E. F., 1957, Geology of the Pine Valley Mountains, Utah: 

Utah Geol. and Mineralog. Survey Bull. 58, 111 p.

Gilluly, James, and Reeside, J. B., Jr., 1928, Sedimentary rocks 

of the San Rafael Swell and some adjacent areas in eastern Utah: U.S. Geol. Survey Prof. Paper 150-D, 110 p.

Gregory, H. E., 1950, Geology and geography of the Zion Park 

region, Utah and Arizona: U.S. Geol. Survey Prof. Paper 

220, 200 p.

Kelling, G., and Walton, E. K., 1957, Load-cast structures: 

their relationship to upper-surface structures and their 

mode of origin: Geol. Mag., v. 94, p. 481-490.

Lewis, G. E., Irwin, J. H., and Wilson, R. F., 1961, Age of the 

Glen Canyon Group on the Colorado Plateau: Geol. Soc. 

America Bull., v. 72, no. 9, p. 1437-1440.

McClymonds, N. E., 1961, Effects of a buried anticline on

ground water in Navajo Sandstone in the Copper MinePreston Mesa area, Coconino County, Arizona: Art. 321 

in U.S. Geol. Survey Prof. Paper 424-D, p. D79-D82.

McQueen, Kathleen, 1957, Photogeologic map of Lees Ferry 

NW, Coconino County, Arizona: U.S. Geol. Survey Misc. 

Geol. Inv. Map 1-196.

Phoenix, D. A., 1963, Geology of the Lees Ferry area, Coconino 

County, Arizona: U.S. Geol. Survey Bull. 1137. [In 

press.]

Poole, F. G., 1962, Wind directions in late Paleozoic to middle 

Mesozoic time on the Colorado Plateau: Art. 163 in U.S. 

Geol. Survey Prof. Paper 450-D, p. D147-D151.

Wells, J. D., 1960, Stratigraphy and structure of the House 

Rock Valley area, Coconino County, Arizona: U.S. Geol. 

Survey Bull. 1081-D, p. 117-158.

MINERALOGY, GEOCHEMISTRY, AND PETROLOGY

198. BENTONITE BEDS OF UNUSUAL COMPOSITION IN THE CARMEL FORMATION, SOUTHWEST UTAH

By L. G. SCHULTZ and J. C. WEIGHT, Denver, Colo., and Beltsville, Md.

Persistent thin beds of bentonite of Jurassic age 

occur at or near the base of the Carmel Formation in 

southwest Utah. Their stratigraphic position and 

geographic distribution are described by Wright and 

Dickey (Art. 197). The bentonites consist of darkpurple or pale-green claystone beds generally an inch 

or two in thickness that appear conspicuously different from the enclosing red beds. Field criteria that 

they are bentonites include their lack of ordinary 

detrital silt and sand, the conspicuous pseudohexagonal books of black lustrous biotite which decrease in 

size and abundance from the bottom to the top of the 

clay bed, and the broad lateral extent of the thin beds.

Thin-section study shows that the bentonites are 

composed of about 1 to 10 percent relatively large 

crystals of feldspar, quartz, and biotite in a very fine 

clay groundmass. Some thin sections show faint remnant structures interpreted as fine-grained shards. 

A few angular fragments apparently of altered volcanic material are outlined because their hematite 

content differs from that of the surrounding groundmass. Clay in the groundmass generally is oriented 

parallel to the bedding, a feature not characteristic 

of most bentonites; wrapping of this bedding around 

the larger, presumably volcanic crystals indicates the 

clay orientation is due at least in part to compaction

during alteration. The larger volcanic crystals appear 

fresh, except that the biotite in a few samples is 

altered partly or wholly to clear, slightly birefringent 

vermiculite. Feldspar phenocrysts are in part euhedral and partly broken into very irregular angular 

particles, indicating aerial rather than stream transportation.

Samples of both the Carmel bentonites and enclosing sediments have been separated into coarse 

(>10/x) and fine fractions. Both size fractions as 

well as the bulk samples were studied by X-ray diffraction; mounts of the coarse fraction were studied 

optically (see table). Note that samples CC-8, MC-T, 

MC-3u, and MC-31, from bentonite beds shown in 

their proper stratigraphic sequence within the Carmel 

Formation in sections at Cottonwood Canyon and 

Mount Carmel Junction, are repeated at the bottom 

of the table under "Bentonite beds". The bentonite 

beds are composed primarily of kaolinite, illite, and 

mixed-layer clays, and contain little or no montmorillonite that is typical of most bentonite beds. Furthermore, except for the apparent absence of IMd muscovite polymorph (listed on the table under "Illite") in 

the bentonites, the clays in the bentonites are similar 

to those in the adjacent sediments. Thus, the clay 

minerals themselves do not indicate that the thin clay
E68 GEOLOGICAL SURVEY RESEARCH 1962

Mineralogy of samples of the Carmel

Sample No. Description

Position in section, or locality (see 

Wright and Dickey, Art. 197)

Abundance of minerals in parts per 10 of whole rock estimated from X-ray 

analysis [Tr., 1-2 percent; ?, doubtful; > greater than; =, approximately 

equal amount; +, detected in coarse fraction only]

Kaolinite

Illite

Amount Type '

Mixed-layer clay

Amount Components 2

Montmorillonite

Quartz

COTTONWOOD

CC-4._...

-6-.___

-2. ....

_i

-3_.._.

 9

-8.....

-7.....

Gray clayey limestone. _ _

Silty biotitic limestone __ _

CARMEL FORMATION

Distance above lowest

prominent limestone

Feet

150

132

125

63

55

30

Distance below lowest

prominent limestone

Feet

60

90

100

?

Tr

Tr

Tr

Tr

Tr

%

Tr

3

%

2

Vz

?

IMd?

IMd?

IMd?

IMd?

1M

?

1M

IMd?

Itf

1

1

1

1

3

2

5/2

1

M, CL__-_.

M>I..__._

M>I.__.__

M>C1.____

M>C1_____

M>I..._._

C1 = M>I__

M>I___-._

Cl, M, !___.

1

2

2

2

3H

2

1

1

1

%

2

4

u

3

MOUNT CARMEL

-16....

-14....

-10....

 Q

-7.....

-6.....

-5.....

-4. ....

-3 1

-2

_i

-12

-11.-..

Gray clayey limestone. _ _

Buff limestone _

.-_.. do.. _.-_.____._.__-___.

CARMEL FORMATION

Distance above base of 

Carmel 

Feet Inches 

50

40

30-40

30-40

71-82

55-70

46-47

45

36-43

14-34

8-12

2

1

0-1

NAVAJO SANDSTONE

Distance below base of 

Carmel 

. Feet 

3

6

u.

Tr

Yzi/

y

1

u

1

1

1

5

1

]/

1

Tr

i/2

1

3/2

5/2

1

3

3

3

Tr

Tr

IMd

IMd?

IMd?

IMd

1M

1M

1M

1M

1M

1M

1M

1M

IMd

1M

?

3

1

Yz

/2

%Yz

4

Tr

1

2

1

Tr

I>M__ __

I>M

T\> jyr

I ^> 1VI

I>M

I>M

1^* 1VI

1^* 1VI

I>M

I>M

I>M

I>M__ .

I^>M__

I>M_._-

4

1^

iy>

6

7

3

6)4

7

4

_j_

4^

9

9

8/2

See footnotes at end of table.
ARTICLE 198 E69

Formation and the Navajo Sandstone

Abundance of minerals in parts per 10 of whole rock estimated from X-ray analysis 

[Tr., 1-2 percent; ?, doubtful; > greater than; =, approximately equal amount; +, 

detected in coarse fraction only] Continued

Potassium feldspar

Amount Type 3

Calcite Calcian 

dolomite

Hematite Biotite

Microscopic examination of coarse grains [A, abundant; c, common; r, rare]

Sanidine (low 

2V)

Microcline or 

orthoclase

Biotite Apatite Hornblende Zircon

CANYON

Tr

Tr

Tr

Tr

Tr

Tr

* 1

Tr

4 1

Mono __

Mono, Tri _

Mono, Tri _

Mono __ __

Tri > Mono _

Tri>Mono_

Tri > Mono.

Mono? _

Mono, TrL.

3

8

8

5

3

Tr

%

2K

?

?

2/2

5%

Tr

1

+

+

+

+

+

c

c

c

A

c

c

A

A

c

A

A

A

c

c

A

c

A

c

r

c

JUNCTION

%

Tr

Tr

Tr

Tr

Tr

Tr

Tr

%

1

%

Tr

Tr

%

Try-2

%

Tri, Mono _

Tri> Mono .

Tri > Mono.

Tri_. ______

Tri__._____

Tri > Mono.

Tri, Mono

Mono _

Tri > Mono.

Tri________

Tri>Mono_

Mono _

Tri > Mono _

Tri. _______

Tri_. ______

Tri_. ______

1

5

Tr

1A

21A

7

7

2

?

?

%

+

+

+ ?

+

+

%

?

c

A

A

A

A

c

A

A

A

A

A

c

A

A

A

A

A

c

c

c

A

r

r

c

r

c

c r
E70 GEOLOGICAL SURVEY RESEARCH 1962

Mineralogy of samples of the Carmel Formation

Sample No. Description

Position in section, or locality (see 

Wright and Dickey, Art. 197)

Abundance of minerals in parts per 10 of whole rock estimated from X-ray 

analysis [Tr., 1-2 percent; ?, doubtful; > greater than; =, approximately 

equal amount; +, detected in coarse fraction only]

'Kaolinite

Illite

Amount Type»

Mixed-layer clay

Amount Components 2

Montmorillonite

Quartz

BENTONITE BEDS IN THE

Western, most

PP-7_____

CC-8._.__ Green, biotitic _____ ______

% 2*4

2

1M

1M

6

5/2

Cl, M, !.___

M>I______

1

2

+

%

Intermediate

pp-1

_o

_3

 4

MC 7

-3 u_. ...

-31.---

Purple slightly biotitic- _ _

Red slightly biotitic ______

Green biotitic ______ _ _

Purple very biotitic _ _ _ _

Purple biotitic_ ______ ____ .

do

do

do

do

Tr

1

2

3^

5

2/2

3/2

3

5

5/2

3

1M

1M

1M

1M

1M

1M

1M

6/2

6

5

3

4

2

1/2

Cl, M______

M>I, V___

Cl, M, Cl-_

I>M__

I>M___

I>M__

2

Tr

1

_[_

_[_

_l_

+ ?

_[_

_j_

Eastern, most

PP-5_____

-6_._._ _____do-_-_---_-_----____.__

Judd Hollow _______ _ _

_-_-_do-_-_---__.-----_-

9J/

1

1 ? 6

8

C1>M, I__.

C1 = I______

? +

+

1 lM=mainly 1M museovite; lMd=mainly IMd muscovite. 

' I=illite; M=-montmorillonite; Cl=chlorite-like.

beds are bentonites. Their designation as such is 

based primarily on their nonclay minerals. The most 

obvious mineralogic feature suggesting that the Carmel bentonites are not normal clastic clay beds is 

their low quartz content; quartz is so rare that it is 

generally detectable only in the coarse fraction. The 

only feldspar recognized in the bentonites is a clear, 

generally untwinned, monoclinic potassium feldspar. 

Judging from the 201 lattice spacing (Bowen and 

Tuttle, 1950) the composition of the feldspar ranges 

from Or80Ab2o to Or65Ab35. The 2F for the 46 feldspar grains measured ranges between 18° and 34°. 

Such material must be sanidine. In addition to sanidine and the previously mentioned euhedral biotite, 

the samples also contain numerous stubby euhedral 

apatite crystals and a few hornblende and zircon 

crystals. The character of this mineral suite, particularly the high-temperature nature of the feldspar, 

indicates that the thin clay beds were derived from 

alteration of a rhyolite ash, and are therefore bentonites. The only evidence observed that any part of 

the clay beds is nonvolcanic in origin is that, of several hundred feldspar grains observed, two from sam3 Mono=monoclinic (sanidine or orthoclase); Tri=triclinic (microeline).

pie PP-6 had grid twinning and therefore are microdine.

The noteworthy feature of the Carmel bentonites 

is the deviation of their clay minerals from normal 

montmorillonitic bentonites. Kaolinite and illite are 

present in most of the Carmel bentonites and are the 

dominant components of some. Presence of the 1M 

muscovite polymorph may be a fairly common result 

of low-temperature diagenetic alteration of rhyolitic 

materials; in contrast, most illites formed by soil or 

metamorphic processes probably are IMd or 2M muscovite polymorphs; the 1M muscovite polymorph has 

also been reported as an alteration product of rhyolite 

from Yellowstone Park (Yoder and Eugster, 1955, 

p. 249). Also noteworthy in the Carmel bentonites 

are the abundance and variability of the mixed-layer 

clays. These clays differ from the mixed-layer illitemontmorillonite common in Paleozoic bentonites 

(Weaver, 1953) by their variability in proportions 

of the different layers, and most notably by the preponderance of chlorite-like layers in several of the 

Carmel samples. A small amount of vermiculite (see 

table) is altered from biotite.
ARTICLE 198 E71

and the Navajo Sandstone Continued

Abundance of minerals in parts per 10 of whole rock estimated from X-ray analysis 

[Tr., 1-2 percent; ?, doubtful; > greater than; =, approximately equal amount; +, 

detected in coarse fraction only] Continued

Potassium feldspar

Amount Type'

Calcite Calcian 

dolomite

Hematite Biotite

Microscopic examination of coarse grains [A, abundant; c, common; r, rare]

Sanidine flow 

2P)

Microcline or 

orthoclase

Biotite Apatite Hornblende Zircon

CARMEL FORMATION 

seaward localities

Tr

Tr

Mono __

Mono?_ _ _

Tr Tr 5 Tr

+

A

A

A c

c

r? r

localities

Tr

Tr

+

Tr

Tr

Tr

Tr

Mono.__ _

Mono _ _

Mono__ - _

Mono _

Mono__

Mono___-__

Mono.____

?

?

5 Tr

Tr

+ ?

+

%

A

A

A

A

A

A

A

c

A

A

c

A

c

c

landward localities

+

Tr

Mono__ -__

Mono___ __

/2

1

+

+

A

A

A

c

r

«About half plagloclase. Biotite partly altered to vermiculite.

Clays in the Carmel sediments above and below the 

bentonites are similar to the clays in the bentonites 

(see table) and therefore presumably also may be derived at least in part from rhyolitic tuff. Other evidence of volcanic components in the sediments are 

some sanidine, biotite, apatite, and rarely hornblende 

and zircon similar to that in the bentonites. However, the enclosing sediments differ from the bentonites by also containing a preponderance of nonsanidine 

feldspar (microcline), the IMd muscovite polymorph, 

and much more quartz; these are probably nonvolcanic detrital constituents. The inferred volcanic origin of much of the clay and the noiivolcanic origin 

of most of the coarser elastics are consistent if much 

of the clay is derived by devitrification of larger tuff 

particles originally deposited with the sediments.

Because 1M muscovite in the Carmel Formation is 

interpreted as being of volcanic origin, and because 

none of the Navajo samples contain other evidence of 

volcanic components, the presence of 1M muscovite in 

sample MC-1 (table) at the top of the Navajo Sandstone seems anomalous. The friable, weathered appearance of sample MC-1, and its location just below

a Carmel bentonite bed, suggest that the clay may 

have infiltrated downward from the overlying bentonite bed.

No textural criteria were recognized to indicate the 

sequence of development of the clay minerals in the 

Carmel bentonites; thus, we can only speculate. The 

presence of sanidine, biotite, and small amounts of 

quartz in all the Carmel bentonites indicates that they 

were altered from a rhyolite tuff of fairly uniform 

composition. Thus, the original composition of the 

ash apparently did not control the course of alteration 

to the variety of clay minerals presently found. Because the phenocrysts generally are unaltered, alteration of the ash was probably no more intensive than 

devitrification of the glass; thus, the kaolinitic bentonites probably are not the products of strong leaching. Experiments with formation of synthetic clays 

(Julian Hemley, oral communication, 1962) indicate 

that a single poorly organized montmorillonoid phase 

forms from a glass much more rapidly than better 

organized phases like the Carmel illite or kaolinite. 

We therefore hypothesize that the rhyolite ash first 

altered to such a poorly organized mass, which, with
E72 GEOLOGICAL SURVEY RESEARCH 1962

time, tended to separate into better organized, more 

stable phases such as the illite. During formation of 

the initial montmorillonoid phase from the ash, considerable SiO2 must have been removed; variable 

amounts of other elements probably also could be 

leached with relative ease from such a poorly crystalline material, thereby explaining, for example, how 

most of the potassium could be leached from some 

samples without alteration of the relatively stable 

feldspar. If little of the potassium originally in the 

rhyolite ash (K2O/A12O3 about ys or 1/4) were removed, the potassium-rich montmorillonoid would 

convert largely to illite (K2O/A12O3 about 1/4); if 

much potassium were removed, abundant mixed-layer 

clay or montmorillonite would result. If alumina in 

excess of that needed for the three-layer clays remains, 

some kaolinite would form. Thus, the highly variable 

composition of the Carmel bentonites would result 

from differential leaching of the chemical components 

of the original ash.

Several factors may have caused variable conditions 

of leaching of the rhyolite tuff. One such factor is 

differential oxidation due to relative degree of marine 

or nonmarine conditions. The Carmel bentonite beds 

probably were deposited on a tidal flat at about the 

time the Carmel sea encroached from the west. Most 

of the samples from intermediate and eastern areas 

farthest from the encroaching sea (see table) are red 

or purple due to included hematite, presumably produced by oxidation during repeated exposure to wetting and drying. Several of these bentonites have

relatively high kaolinite contents. In contrast, bentonite samples from the two more seaward localities 

are green, indicating that they were more continually 

submerged and were exposed to relatively little wetting, drying, and the accompanying oxidation and 

leaching; these samples contain very little kaolinite. 

Physical conditions during deposition, however, seemingly cannot explain all the observed mineralogical 

variations, such as the highly illitic bentonites at Mt. 

Carmel Junction, which contrast with highly chloritic 

mixed-layer bentonites from localities both landward 

and seaward of Mt. Carmel Junction. Also, at Cottonwood Canyon, samples throughout a stratigraphic 

interval of 250 feet contain fairly montmorillonitic 

clays, whereas clays at Mt. Carmel Junction are all 

illitic. The strand line and environment of deposition 

did not remain static during deposition of all these 

rocks. Therefore, the course of alteration of the 

rhyolitic debris in the Carmel must also have been 

influenced by some such factor as the chemistry of 

connate water, which could circulate and uniformly 

affect a considerable thickness of rocks after burial 

of the sediments.

REFERENCES

Bowen, N. L., and Tuttle, O. F., 1950, The system NaAlSi3O8-

KAlSi3O8-H2O: Jour. Geology, v. 58, no. 5, p. 489-511. 

Weaver, C. B., 1953, Mineralogy and petrology of some Ordovician K-bentonites and related limestones [Pa.]: Geol.

Soc. America Bull., v. 64, no. 8, p. 921-943. 

Yoder, H. S., and Bugster, H. P., 1955, Synthetic and natural

muscovites: Geochim. et Cosmochim. Acta, v. 8, nos. 5/6,

p. 225-280.
ARTICLE 199 

199. SAMPLING A ZONED GALENA CRYSTAL FOR LEAD ISOTOPE STUDY

By KALPH S. CANNON, JR., KATHARINE L. BUCK, and ARTHUR P. PIERCE, Denver, Colo.

E73

Work done in cooperation with Division of Research, U.S. Atomic Energy Commission

Nier (1938; Nier and others, 1941) was the first to 

show that the isotopic composition of lead in lead 

minerals of ore deposits varies significantly from one 

deposit or locality to another, and that ore-lead in 

deposits of younger geologic age tends to contain 

higher proportions of the radiogenic isotopes Pb206, 

Pb207 , and Pb208 from decay of uranium and thorium. 

He learned further that lead from Joplin, Mo., compared with lead from other regions, was variable and 

anomalous in isotopic composition. His discoveries 

prompted further investigations in many laboratories, 

including a long-term study of the isotopic variations 

of lead and their geologic significance by the U.S. 

Geological Survey.

The initial goal of the Survey program was to learn 

how to carry out geologic sampling procedures so that 

lead isotope findings would be meaningful. In sampling ore deposits one needs to know what kinds of 

lead isotope variations are expectable within individual mineral grains or crystals, within a hand specimen, 

within a single vein, ore shoot, or deposit, or within 

the broader scope of an entire district, metallogenic 

province, or class of ore deposit. Thus forearmed, 

one can design logical lead isotope experiments to 

solve specific problems in the geology of ore deposits.

A first step is to learn how uniform the isotopic 

composition of lead may be throughout single mineral 

grains or crystals. The first such test was made on 

galena from the Tri-State mining district (MissouriKansas-Oklahoma) because Nier's analyses of three 

different galena samples from "Joplin, Mo." showed 

measurable variations in lead isotope ratios. A large 

crystal was chosen to facilitate selective internal sampling. Individual growth zones of the crystal were 

sampled to test whether the isotopic composition of 

ore-lead was constant as the crystal grew. Isotope 

analyses of lead from these successive growth zones 

are not identical. Instead, they show variations 

greater than analytical error. These variations seemingly are systematic, with radiogenic isotopes progressively more abundant in outer growth layers, excepting the outermost skin of the crystal (sample A). 

These variations appear analogous to those found by 

Nier in ore deposits of different ages; that is, radiogenic isotopes appear to be more abundant in samples 

of younger age. They suggest a possibility that this

crystal, may have grown during a very long span of 

geologic time.

These empirical observations seemed of such fundamental importance to research on the isotope geology 

of lead that only the general nature of the results was 

reported (Cannon, 1951, 1952, 1953; Cannon and others, 1961) while attempts to elucidate details were 

continued. Our sampling procedures, together with 

lead isotope analyses made in 1951 and 1952 (accompanying table), are reported here in detail for the 

first time. These will help in evaluating other studies

Lead isotope ratios of samples cut from zoned galena crystal 

TSO-MK

Calculated

Sample 

No.

A. 

B_ ....

B_ ....

C- ....

D.. ...

E..... ...

F-.-____

H __

I... -

Analysis 

No.

GS/54

GS/55

GS/60

GS/118

GS/56

GS/119

GS/57

GS/58

GS/59

7/3/51..--

7/11-12/51..

7/24-26/51..

1/23/52...-.

7/12-13/51..

1/24/52 _ ..

7/18-20/51..

7/16-18/51..

7/20-24/51..

Instru2B

2B

2B

6

2A

6

2B

2A

2A

MKH/JWR. -

MKH/JEP---

DW/JWR..--

DW/JWR__-

MKH/JWR..

DW/JWR....

MKH/JWR- -

MKH/JWR--

MKH/JWR__

Pbaoe

Pb204

91 73

22.18

22.20

22.42

21.81

21.75

21.36

21.28

21.29

pb207

Pb'M

15.38

15.68

15.62

15.94

15.78

15.79

15.63

15.58

15.55

Pbzos

Pb204

40.04

41.43

40.69

40.17

40.02

39.70

Measured :

Lab. No.

GS/54..... 

GS/55....------ 

GS/60 -------- -

GS/118-. -- 

GS/56.. ---------

G8/119....-..-.- 

GS/57 - ----------

G8/58 ..... ... . 

GS/59. _ ------

pb204

pb206

0. 0461

±. 0002 

.0452

±.0004 

.0450

±.0006 

.0446

±. 0003 

.0459

±. 0002 

.0460

±.0002 

.0469

±.0003 

.0469

±.0007 

.0470

±.0006

Pb204

pb207

0.0650

±.0001 

.0640

±. 0005 

.0640

±.0008 

.0627

±.0003 

.0635

±.0002 

.0634

±.0002 

.0639

±. 0010 

.0642

±. 0010 

.0645

±. 0010

pb204

Pb208

0. 0250

±.0001 

.0244

±.0002 

.0245

±.0004 

.0242

±.0001 

.0248

±. 0002 

.0246

±.0002 

.0249

±.0009 

.0250

±.0005 

.0252

±.0005

Pb207

pb206

0. 7076

±.0017 

.7065

±. 0012 

.7040

±.0009 

.7110

±. 0019 

.7233

±.0011 

.7264

±.0019 

.7319

±.0045 

.7325

±. 0016 

.7303

±.0026

pb206

pb208

0. 5427

±.0008 

. 5434

±. 0012 

.5439

±.0010 

.5412

±. 0009 

.5414

±.0007 

.5345

±. 0011 

.5316

±.0045 

.5319

±.0045 

.5363

±. 0024

Pb207

Pb20B

0.3841

±. 0003 

.3832

±. 0002 

.3324

±.0019 

.3851

±.0009 

.3913

±.0006 

.3886

±. 0005 

.3887

±. 0029 

.3897

±.0005 

.3913

±.0016

1 Initials of the assayer, M. K. Hill or Dora Whaley, followed by initials of the supervisor, J. W. Redniond or J. E. Parham, of the Mass Assay Laboratory, Y-12 Plant, 

Carbide and Carbon Chemicals Co., Oak Ridge, Tenn.

2 "The errors quoted are an expression of the precision of this measurement only. 

The accuracy is estimated at less than 1 percent from known sources of systematic 

errors." (Statement of laboratory.)
E74 GEOLOGICAL SURVEY RESEARCH 1962

of variations within galena crystals which were sampled by other methods (Kulp and others, 1953, p. 8; 

Austin and Slawson, 1961), as well as more general 

studies in Mississippi Valley mining districts (Farquhar and Gumming, 1954; Eckelmann and Kulp, 

1959). The uncertainty of the mass-spectrometric 

measurements tends to be large, however, as compared 

with the natural variations. Therefore we are continuing to try to make new measurements of the precision needed. We hope to interpret these phenomena 

quantitatively as soon as analyses of appropriate quality are at hand.

Our galena crystal is from the Netta mine in the 

Picher-Miami area of the Tri-State mining district. 

This mine is in Ottawa County, Okla., one-third mile 

west of Picher, and about 20 miles south-southwest 

of Joplin, Mo. The specimen (Field No. TSO-MK) 

was collected on October 17, 1949, by Mr. Boodle Lane 

of Galena, Kans., and was obtained from him the 

same day. Information given by Mr. Lane implies 

that the crystal was collected in abandoned workings 

at a depth of about 250 feet, from a mine pillar 400 

feet north and 600 feet east of the southwest corner 

of the Netta tract. This pillar, according to geologic 

mapping by Carl Addison (McKnight and others, 

1944), contains high-grade ore in "M" bed, the most 

important ore horizon of the district.

The specimen was an imperfect crystal measuring 

about 7.0 by 7.5 by 9.0 cm and weighing about 2.5 kg. 

It was attached along one cube-edge to jasperoid, and 

its external form implies that it grew from this wall 

of jasperoid into an open space. Growth and lineage 

structures seen in cross section (fig. 199.1) likewise 

indicate that the crystal grew asymmetrically outward 

from the jasperoid wall. Internal discontinuities imply that crystal growth was periodic, possibly interrupted at times by partial solution. At some stages 

of growth the crystal evidently had fairly regular 

cube faces, but the present crystal faces are made up 

of a complex and extremely uneven assemblage of 

cubic and octahedral forms. A few younger, small 

crystals of dolomite, sphalerite, and marcasite have 

grown on these rough surfaces. In the last major episode of growth history the two cube faces farthest 

from the wall enlarged by outward growth and acquired a sprinkling of marcasite crystals, whereas a 

face adjacent to the wall was etched to depths of a 

centimeter or more and then dusted with sphalerite 

crystals.

The crystal was sawed in half, nearly parallel to 

one cubic cleavage and normal to the wall and the 

longer axis of the crystal. The sawed faces were 

polished and then the various etch methods and reagents recommended by Ramdohr (1955, p. 496) and 

by Frondel and others (1942) for revealing internal 

growth structures of galena were tried, without success. After many failures a modification of the hot 

HC1 etch of cleavage surfaces described by Becke 

(1884) successfully revealed the growth structures in 

this crystal. The sawed surfaces were rough ground 

to develop a matte or pitted surface that is essentially 

a mosaic of minute cleavage faces. On this matte 

surface a radial pattern of lineage structures was well 

exhibited, and the center of this pattern was taken as 

the nucleus of crystal growth for this particular cross 

section (fig. 199.15, sample I). This surface was first 

etched by immersion in cold 1:7 (v/v) HNO3 for several minutes until it began to darken or show iridescent tarnish, and then by immersion in hot fuming 

1:2 (v/v) HC1 until light and dark growth structures 

appeared and developed maximum contrast.

The etched surface is predominantly gray, marked 

by alternating lighter and darker lines indicating the 

concentric growth layers, and by relatively large white 

triangular octahedral-face loci (fig. 199.1). The galena 

can be expected to contain different concentrations of 

Ag atoms under the lighter and darker areas (Frondel and others, 1942), and the patterns, therefore, suggest fluctuations in the ratio of Ag and Pb atoms 

available to the growing crystal. Several of these 

fluctuations evidently reflect sudden variation, others 

more gradual change, in the Ag/Pb ratio. Certain 

growth zones are marked by inclusions of other sulfide and carbonate minerals which originated as overgrowths on former exterior faces of the crystal similar 

to those on present exterior faces. The principal zones 

of inclusions seem to coincide with discontinuities 

within the crystal that are revealed by etch ^patterns. 

These several lines of evidence taken together indicate 

that the crystal's growth was interrupted a number of 

times. Such evidence that galena has had a complex 

history of repetitive deposition and dissolution has 

been observed throughout the Miami-Picher area 

(E. T. McKnight, written communication, 1961).

Samples weighing approximately 50 mg were cut 

from successive layers on this etched surface (fig. 

199.15). Sample I was drilled from the focus of 

radiating lineages, believed to represent the nearest 

approach to the growth nucleus in this cross section. 

Samples F, G, and H were sawed from slots 1 mm wide 

on three sides of the innermost recognizable growth 

layer: F was cut at one cube-face position, G at an 

adjacent cube-face position, and H at an octahedronface position. Samples E, D, C, and B were similarly 

cut at cube-face positions from successively younger 

growth layers outside of sample-position G. Finally,
ARTICLE 199 E75

3cm

r

FIGUEE 199.1. A, Photograph of Tri-State galena crystal TSO-MK as sectioned, etched, 

and sampled. B, Photograph retouched to emphasize light and dark etch patterns, 

growth zones, arid crystallographic spacing of samples A through I (described in text).

667422 O - 63 - 6
E76 GEOLOGICAL SURVEY RESEARCH 1962

sample A was scraped from the gray oxidized surface 

of the exterior cube face adjacent to sample B.

The samples of lead sulfide were converted to lead 

iodide by R. G. Milkey, U.S. Geological Survey Washington laboratory. Lead isotope analyses of the iodide samples were made on several 6-inch Nier-type 

60° mass spectrometers in the Mass Assay Laboratory, 

Y-12 Plant, Carbide and Carbon Chemicals Co., Oak 

Ridge, Tenn., under the general supervision of Roger 

F. Hibbs. In July 1951, samples A, B, D ,F, H, and 

I were analyzed on instruments 2A and 2B; sample B 

lead iodide was split and submitted as two independent samples. To check the systematic variations indicated by these analyses, intermediate samples C and E 

were analyzed in January 1952 on instrument 6. The 

last 2 analyses are imperfectly consistent with the first 

7: ratios among Pb206 , Pb207, and Pb208 are consistent, 

but the reported abundance of Pb204 tends to be lower 

than in the earlier analyses and is more nearly in 

accord with Nier's analyses of Joplin samples. However, each batch of analyses is judged to be internal^ 

consistent. The duplicate analyses of sample B are 

well within experimental error, and anatyses of samples F and H from the same growth layer are practically identical. Precision of each batch of anatyses 

is judged to be about 1 percent for ratios involving 

Pb204, and 0.2 percent for ratios involving only the 

major isotopes.

For each sample the ratios actually measured, and 

their calculated log errors, are reported in the second 

part of the table. For each analysis there is about a 

90-percent probability that, under the same operating 

conditions, the given instrument would find ratios 

within these limits. The ratios to Pb204 given in the 

first part of the table, however, are calculated values 

obtained by adjusting the six measured ratios. Figure 199.2 is a graphic plot of reciprocals of the Pb207/ 

Pb206 ratios, the most orderly of the measured ratios.

The general nature of the results is evident. Within 

the visibly zoned part of the crystal, between inner 

growth z'one F-H and outer growth zone B, the data 

suggest systematic outward increases in the Pb206/ 

Pb207 ratio and in total radiogenic lead. As referred 

to the number of atoms of Pb204, Pb206 increases outward by about 4 percent, Pb208 by about 2 percent, 

and Pb207 increases very slightty or remains nearty 

constant. Only the largest of these variations exceed 

the uncertainty of the measurements, yet the variations 

seem to be systematic and consistent except for innermost sample I and outermost sample A. The incongruous decrease in radiogenic isotopes in outermost 

sample A .seems to be a "skin effect," similarly noted 

in galena crystals from Hansonburg, N. Mex., by 

Austin and Slawson (1961).

Interpreting the data at face value, this galena crystal had a growth history whose details are recorded by 

individual growth laj^ers and by the isotopic variations 

from layer to layer. The lead deposited from oreforming solutions was progressively enriched in radiogenic isotopes as the crystal grew. These radiogenic 

increments appear to account for about 1 percent of 

the atoms in the entire crystal. One interpretation of 

the recorded variations would suggest that the crystal 

grew for a hundred million years or more, but that the 

episode of crystal growth had ended before the crystal 

was collected. Because of the analytical uncertainty 

of the data, new lead isotope analyses of greater precision are now being made, to permit better interpretation of the complex history of this crystal and its 

bearing on the genesis of Mississippi Valley ores.

H-58 

1-59 F-57 r~r E-119 I D-56 I 

SAMPLES ANALYZED c_ 118

B-55 

B_6Q A_54

1.42 

1.41

§ 1.40 

f 1.39 s x

 o 1.38

CL

1.37

1.36

200

i

300

i

400

I 500

I

600

HYPOTHETICAL VOLUME OF GROWING CRYSTAL, IN CUBIC CENTIMETERS

FIGURE 199.2. Variations in the isotope ratio Pb206/Pb207, based on analyses of samples cut from zoned galena crystal TSO-MK.

Hypothetical volume=(2 X linear distance) 3 .
ARTICLE 200 E77

REFERENCES

Austin, C. F., and Slawson, W. F., 1961, Isotopic analyses of 

single galena crystals: A clue to history of deposition: 

Am. Mineralogist, v. 46, p. 1132-1140.

Becke, Friedrich, 1884, Aetzversuche am Bleiglanz: Tschermaks mineralog. petrog. Mitt., v. 6, p. 237-280.

Cannon, R. S., Jr., 1951, Isotope geology of lead, in Rabbitt, 

J. C., Summary of the research work of the Trace Elements Section, Geochemistry and Petrology Branch, for 

the period Oct. 1-Dec. 31, 1951: U.S. Geol. Survey TEI218, p. 25-27, issued by U.S. Atomic Energy Comm. Tech. 

Inf. Service, Oak Ridge, Tenn.

 1952, Geological Survey's work on isotope geology of 

uranium, thorium, and their decay products: U.S. Geol. 

Survey TEI-209, 12 p., issued by U.S. Atomic Energy 

Comm. Tech. Inf. Service, Oak Ridge, Tenn.

1953, Isotope geology of lead, in Trace Elements research Quarterly progress report, Jan. 1-March 31, 1953: 

U.S. Geol. Survey TEI-270, p. 54-56, issued by U.S. Atomic 

Energy Comm. Tech. Inf. Service, Oak Ridge, Tenn.

Cannon, R. S., Jr., Pierce, A. P., Antweiler, J. C., and Buck, 

K. L., 1961, The data of lead isotope geology related to 

problems of ore genesis: Econ. Geology, v. 56, p. 1-38.

Eckelmann, F. D., and Kulp, J. L., 1959, Lead isotopes and ore 

deposition in the Southeast Missouri lead district [abs.]: 

Geol. Soc. America Bull., v. 70, p. 1595.

Farquhar, R. M., and Cumming, G. L., 1954, Isotopic analyses 

of anomalous lead ores: Royal Soc. Canada Trans., v. 48, 

ser. 3, p. 9-16.

Frondel, Clifford, Newhouse, W. H., and Jarrell, R. F., 1942, 

Spatial distribution of minor elements in single-crystals: 

Am. Mineralogist, v. 27, p. 726-745.

Kulp, J. L., Owen, H. R., Eckelmann, W. R., and Bate, G. L., 

1953, Isotopic composition of some samples of common 

lead: U.S. Atomic Energy Comm., NYO-6197, 13 p., issued 

by U.S. Atomic Energy Comm. Tech. Inf. Service, Oak 

Ridge, Tenn.

McKnight, E. T., Fischer, R. P., and others, 1944, Zinc-lead 

deposits in part of the Picher field, Ottawa County, Okla., 

and Cherokee County, Kans.: U.S. Geol. Survey Mineral 

Inv. Prelim. Map, 6 sheets.

Nier, A. O., 1938, Variations in the relative abundances of the

isotopes of common lead from various sources: Am. 

Chem. Soc. Jour., v. 60, p. 1571-1576.

Nier, A. O., Thompson, R. W., and Murphey, B. F., 1941, The 

isotopic constitution of lead and the measurement of geological time. Ill: Phys. Rev., 2d ser., v. 60, p. 112-116.

Ramdohr, Paul, 1955, Die Erzmineralien und ihre Verwachsungen : Berlin, Akademie-Verlag, p. 496.

200. EPIDOTE AND RELATED MINERALS IN TWO DEEP GEOTHERMAL DRILL HOLES, REYKJAVIK AND

HVERAGERDI, ICELAND

By GUDMUNDTJR E. SlGVALDASON i

An extensive review of the volcanic geology and 

thermal activity of Iceland is given by Barth (1950). 

More recent work on surface alteration by hydrothermal solutions has been done by Sigvaldason (1959) 

and on subsurface temperature relations by Bodvarsson and Palmason (1962). The present article 

contains results of a preliminary study of hydrothermal alteration at depth in two Icelandic thermal 

areas, with special emphasis on the distribution of 

epidote and related minerals.

Thermal activity in Iceland has been classified into 

two main categories: high-temperature areas more or 

less confined to centers of postglacial volcanic activity, and low-temperature areas with no direct relation to Pleistocene or Recent volcanism. The city of 

Reykjavik lies within one of the low-temperature 

areas. Surface indications of thermal activity are 2 

small hot springs with temperatures of 88°C. The 

area is covered by as much as 100 m of olivine dolerite 

of Pleistocene age. Below the lavas is a thin series 

(10 to 20 m) of interglacial deposits, which in turn 

overlies Tertiary plateau basalts. The thickness of

1 Present address, University Research Institute, Reykjavik, Iceland.

the Tertiary basalts in this area has been estimated 

at 2,000 m by seismic studies (Tryggvason and Bath, 

1961), but the character of the underlying rocks is 

unknown except that higher densities are indicated.

Drilling for hot water for domestic heating has been 

carried out over a period of 3 decades. Most of the 

drill holes range in depth from 100 to 600 m, but 

recently several holes have been put down to 1,000 m 

and 1 hole is 2,200 m deep. The deep holes were 

drilled with a large rotary drill, which delivers rock 

fragments as cuttings with an average size of 0.1 

to 0.5 cm.

Samples were marked as being from the depth of the 

drill at the time of sampling. For the present report 

no correction has been made for the error involved, 

as it is believed to be of minor importance, especially 

in the upper parts of the hole. More serious errors 

could be introduced by mixing of rock fragments from 

different depths by abrasive action of the drilling 

rods and mud on bore walls. This can result in misinterpretation of the downward extension of a particular mineral phase, but correct information is obtained on the first appearance of a mineral.
E78 GEOLOGICAL SURVEY RESEARCH 1962

The original composition of the Tertiary lavas can 

only be discussed on a broad basis. Walker (1959, 

1960) describes sections 5,000 and 10,000 m thick in 

the Tertiary plateau of eastern Iceland, OH vine 

basalts make up about 23 percent of the section, and 

48 percent are lavas of the tholeiitic type. No regularity in the sequence of different lava types was 

detected. Similar compositions can be expected of 

lavas in the Reykjavik area, but definite data on this 

point were not obtained because of the hydrothermal 

alteration. Most of the minerals of the basalt have 

undergone complete alteration, but some less altered 

fragments reveal the general basic character of the 

original rock. A single sample of cuttings often 

contains fragments with rather different crystal sizes, 

presumably resulting from mixing of grains from the 

marginal fine-grained parts of lava flows and the more 

coarsely crystallized central parts. As a rule the finegrained material is more altered because of higher 

susceptibility to hydrothermal solutions and closer 

contact to flow channels.

The olivine dolerite that mantles the Reykjavik 

area is generally quite fresh. Below this cover, temperatures increase at the rate of 0.5°C per meter (fig. 

200.1) ; alteration begins with montmorillonite, which 

extends a few tens of meters down into the Tertiary 

lavas. Montmorillonite is accompanied by abundant 

calcite and zeolites. Mordenite has been identified 

from depths of 176 and 185 m. Other zeolites not yet 

identified with certainty occur at shallower depths.

TEMPERATURE, IN DEGREES 

CENTIGRADE

80 100 120 140

FIGURE 200.1. Temperature-depth relations in a drill hole 

at Reykjavik, Iceland, with data on mineral distribution.

Chlorite is the most characteristic alteration product below the zone of montmorillonite and down to 

the bottom of the hole. The chlorite is usually present as a replacement mineral, but it occurs also as 

cavity linings and amygdule fillings. X-ray diffraction patterns of the material show that the second 

order of the basal reflection is consistently stronger 

than the first order, indicating an iron-rich type. 

Minerals associated with chlorite and occurring more 

or less continuously from below the montmorillonite 

zone down to the bottom of the hole are albite, quartz, 

calcite, and laumontite. Albite replaces plagioclase 

more or less completely, especially in strongly altered 

fragments. Partial albitization of plagioclase is found 

in coarser grained fragments. Quartz has not been 

found as a groimdmass constituent, but well-developed 

quartz prisms are commonly found in the cuttings, 

indicating deposition in amygdules. Calcite is sporadic throughout the section as amygdule fillings and 

replacement of other minerals. Laumontite first appears at 160 m and is by far the commonest zeolite 

at greater depths. Nontronite is found between 820 

and 1,650 m as brown-colored fine-grained aggregates 

dispersed in the groundmass. Coarser grained crystals 

were observed at 1,000 m. Pumpellyite was found just 

below the olivine dolerite in the sedimentary section 

and continues in the Tertiary basalts down to a depth 

of 200 m. The index of refraction (n^, = 1.70) indicates an iron-rich variety (Seki, 1961). Clinozoisite 

is present as a minor constituent between 200 and 660 

m, with indices of refraction na, = 1.710 and 

ny, = 1.715. This mineral has anomalous interference colors but no pleochroism. The optical sign is 

positive, which distinguishes it from epidote. Epidote was first noted as needles in a vesicle filling with 

calcite at 900 m. From here down to the bottom 

of the hole epidote is found in increasing amount. 

In the upper part of this section epidote appears to 

be confined to veins and vesicles, where it occurs with 

calcite and zeolite. At greater depths epidote replaces 

the groimdmass as well, and crystals up to a considerable size' (1-2 mm) are developed.

The high-temperature thermal area near Hveragerdi, 

50 km east of Reykjavik, has been described by Earth 

(1950), Einarsson (1951), and Bodvarsson (1961). 

The area lies within the zone of Recent volcanism, 

with the last eruption in this general area occurring 

in A.D. 1000. The rocks consist of basaltic tuffs and 

lavas with minor intrusive rhyolite. Several holes 

ranging in depth from 300 to 1,200 m have been 

drilled in this area. Rocks penetrated at depth are 

Tertiary basaltic lavas. Depth to the contact is un-
ARTICLE 200 E79

certain but is very probably not deeper than 100 to 

200 m.

Temperatures in the drill holes show distinct maximums, indicating horizontal stratification of thermal 

water with a source of heat outside the drilled area.

After discovery of epidote at and below 900 m in 

the Reykjavik drill hole just described, a specific 

search for epidote and related minerals was made in 

cuttings from a drill hole 655 m deep at Hveragerdi. 

Pumpellyite, clinozoisite and epidote were found in 

this hole at much shallower depths than at Reykjavik 

and at higher temperatures. Figure 200.2 shows the 

distribution of these minerals. The sequence of alteration products with depth in Hveragerdi is not yet 

known in detail, but the distribution of a few minerals 

is indicated in figure 200.2. In a thin zone near the 

surface, acid solutions have converted the rock into 

kaolinitic alteration products. Montmorillonite is the 

dominant alteration product from depths of a few 

meters to about 100 m, where chlorite becomes dominant. Abundant calcite and zeolites are found throughout the drilled section below the acid zone. In the 

upper part zeolites are abundant but their identities 

are still uncertain. Laumontite first appears at 225 m 

and is the dominant zeolite below about 250 m.

Qualitatively, alteration follows similar paths in the 

two areas, the main differences being more complete 

replacement by hydrothermal minerals in the hightemperature area and appearance of particular minerals or assemblages at different depths, commonly 

shallower at Hveragerdi.

In order to interpret these differences, especially in 

regard to the occurrence of epidote, two possibilities 

are presented below.

The epidote and perhaps some of the zeolite minerals 

are not related to present thermal activity but are 

relict from an earlier regional metamorphism. Toward 

the end of the Tertiary both areas consisted of a pile 

of basalt flows, considerably thicker than at the present time. Metamorphism took place at depth, similar 

to that described in eastern Iceland by Walker (1959, 

1960); temperatures were consistent with a normal 

geothermal gradient. Glacial erosion during the 

Pleistocene resulted in the present relief, and the epidote should be considered as a relict, metastable phase 

in the present hot-spring systems.

The second and favored possibility is the formation of epidote, and perhaps the zeolites, essentially 

under present conditions. Factors influencing differences in mineralogy between the two thermal areas 

are pressure, temperature, and activity of dissolved 

components in hydrothermal solutions. Occurrence of

TEMPERATURE, IN DEGREES 

CENTIGRADE

160 180 200 220

£ <D ID £ 2 O

I *1I iff I

240 S to J3 OO5 5

200

400

600

FIGURE 200.2. Temperature-depth relations in a drill hole at 

Hveragerdi, Iceland, with data on mineral distribution.

epidote at considerably greater depth in the low-temperature area favors this latter possibility. In addition, epidote is very rare in the regional zeolitic metamorphism of basaltic lavas of Iceland (Walker, 1960, 

p. 516) but is commonly associated with zeolites in 

the propylitic type of hydrothermal alteration of 

basaltic and andesitic lavas, as reviewed in the following paper (Art. 201).

REFERENCES

Barth, Tom F. W., 1950, Volcanic geology, hot springs and 

geysers of Iceland: Carnegie Inst. Washington Pub. 587, 

174 p.

Bodvarsson, Guunar, 1961, Physical characteristic of natural 

heat resources in Iceland: United Nations Conference on 

New Sources of Energy, preprint. Rome. 1961.

Bodvarsson, Gunnar, and Palniason, G., 1961, Exploration of 

subsurface temperatures in Iceland: United Nations Conference on Xew Sources of Energy, preprint, Rome, 1961.

Einarsson, Trausti, 1951, Geology of the Hengill area: Verkfraedingafelag Islands, Timarit v. 36, p. 49-82 (in Icelandic with English summary).

Seki, Y., 1961, Pumpellyite in low grade metamorphism: Jour. 

Petrology, v. 2, p. 407-423.

Sigvaldason, Gudmundur E., 1959, Mineralogische Untersuchungen uber Gesteinszersetzung durch postvulkanische 

AktivitJit in Island; Beitrage zur Mineralogie mid Petrographie, v. 6, p. 405-426.

Tryggvason, Eysteinn, and Bath, M., 1961, Upper crustal structure of Iceland: Jour. Geophys. Research, v. 66, p. 1913- 

1925.

Walker, George P. L., 1959, The geology of the Reydarfjordur 

area, eastern Iceland: Geol. Soc. London Quart. Jour., 

v. 114, p. 367-393.

 19(50, Zeolite zones and dike distribution in relation to 

the structure of the basalts of eastern Iceland: Jour. 

Geology, v. 58, p. 515-528.
E80 GEOLOGICAL SURVEY RESEARCH 1962

201. EPIDOTE IN HOT-SPRING SYSTEMS, AND DEPTH OF FORMATION OF PROPYLITIC EPIDOTE IN

EPITHERMAL ORE DEPOSITS

By DONALD E. WHITE and GTJDMUNDUR E. SIGVALDASON/ Menlo Park, Calif.

Before 1961, epidote was one of the very few alteration or gangue minerals of epithermal ore deposits 

that had not been identified in hot-spring systems. 

Within a period of 1 year, however, epidote was found 

at considerable depth in 4 different hot-spring systems. 

Naboko and Piip (1961) first noted epidote in drill 

core from Pauzhetsk Hot Springs, Kamchatka, where 

a few crystals were found near 750 m in depth. Steiner 

(1963, and written communication, 1961) then found 

epidote at Wairakei, New Zealand, near and below 

425 m in depth. Sigvaldason, without knowledge of 

the previous discoveries, made a specific search for 

and found epidote in cuttings from a deep drill hole 

in Reykjavik, Iceland, and later from a Hveragerdi 

drill hole, as described in the preceding article (Art. 

200).

EPIDOTE IN METAMORPHIC ENVIRONMENTS

Much interest has recently been focused on epidote 

in the relatively low grade green schist and zeolitic 

facies of metamorphic rocks. Epidote is a common 

constituent of rocks of the green-schist facies (Fyfe, 

Turner, and Verhoogen, 1958, p. 166-167), which 

typically contains quartz, albite, epidote, chlorite (especially the iron-rich varieties), muscovite, paragonite, 

talc, actinolite, pumpellyite, calcite, and dolomite.

A zeolite facies and several subfacies have been 

proposed (Coombs, 1954, 1960; Fyfe and others, 1958; 

Coombs and others, 1959) for rocks that bridge the 

gap between diagenesis and metamorphism. In southern New Zealand, tuffs and volcanic graywackes have 

been converted in part to rocks that include, in approximate order of appearance downward, heulandite 

(and the high-sodium variety clinoptilolite), analcime, celadonite; laumontite, albite, adularia, chlorite, 

calcite; and pumpellyite, prehnite, epidote, actinolite, 

muscovite, and stilpnomelane. The section is now 28,000 

feet thick and may have had as much as 10,000 feet 

of cover removed by erosion. A little epidote occurs 

in the middle and upper parts of the section and is 

probably clastic in origin (Coombs and Fyfe, oral 

communication, 1962). The part of the section with 

unquestioned epigenetic epidote was buried to depths 

on the order of 25,000 to 30,000 feet; load pressures 

were at least 2,500 bars and temperatures were probably near 250°C.

1 Present address, University Research Institute, Reykjavik, Iceland.

Zen (1961) has suggested that the zeolite facies 

can form when the chemical potential of CO2 is low 

relative to H2O, but it does not develop in many sedimentary environments characterized by calcite, quartz, 

and kaolinite or pyrophyllite. Eeactive components 

such as volcanic glass, unstable minerals of intermediate and basaltic volcanic rocks, and a low organicmatter content all favor a low chemical potential of 

CO2 relative to H2O in sediments after burial. Orderdisorder relations may influence formation of epidote 

from plagioclase (Christie, 1962; Eutland, 1962), and 

possibly also from zeolites.

EXPERIMENTAL SYNTHESIS OF EPIDOTE

Numerous attempts have been made to synthesize 

epidote and its related minerals clinozoisite and zoisite. 

Efforts at less than 5,000 bars of water pressure and 

500°C have not been successful (Coombs and others, 

1959, p. 90; Fyfe and others, 1958, p. 169). Clinozoisite has been observed at 600°C and about 12,000 

bars of pressure (Fyfe and others, 1958, p. 170). 

Merrin (1960) has reported synthesis of epidote at 

525°C and 3,000 bars of pressure; the slope of the 

boundary of the epidote field is positive, and passes 

through 650°C and 4,000 bars of pressure. Pistorius 

and others (1962) have reported zoisite and clinozoisite 

at temperatures of 400°C to 800°C and pressures of 

10 to 15 kilobars; starting materials influence the 

phases found, and equilibrium may not have been 

attained in many of the experiments. Pellizer (1957) 

induced alteration of a diabase to epidote-albite-calcite-quartz-sericite-clay by watei; and CC>2; conditions 

ranged from 375°C to 540°C, about 220 to 340 bars 

of water pressure, and pH of 3.8 to 9. Finally, Fyfe 

(1960) has synthesized epidote from its components, 

but seeded with epidote, at 2,000 bars water pressure 

and 465°C to 630°C. Zeolites are believed to be the 

stable phases below 300°C.

EPIDOTE IN PROPYLITIC ALTERATION

In many thick piles of volcanic rocks, hydrothermal 

alteration has partly or completely changed the original rocks to varied assemblages of minerals that include chlorite, albite, and epidote, as well as adularia, 

various zeolites, celadonite, actinolite, pumpellyite, 

prehnite, sericite, pyrite, calcite and other carbonates, 

quartz, and some other minerals. There is not complete agreement in defining the term "propylitic"
ARTICLE 201 E81

alteration, in specifying which of the above minerals 

are essential, or in designating the kinds of rocks involved (Lindgren, 1933, p. 457-458; Coats, 1940; 

Calkins, 1944; Nakovnik, 1954; Wilshire, 1957; and 

G. A. Thompson and White, written communication). 

The rocks are usually andesites or basalts, and the 

altered assemblages are green; chemical analyses typically show significant gains in water, CC>2, and commonly sulfur, but proportions of other components 

are nearly unchanged. Epidote-bearing rocks are 

associated in some localities with green zeolitic rocks 

containing little or no epidote; the latter have commonly been considered as a separate zeolitic alteration, 

distinct from propylitic alteration. In at least some 

localities, but perhaps not everywhere, epidote-bearing 

rocks grade upward into zeolitic rocks containing 

little or no epidote.

The propylitic and zeolitized rocks commonly occur 

near centers of volcanic activity; epithermal gold and 

silver deposits are very common in the vicinity, but 

generally the ore deposits are more intimately associated with different hydrothermal alterations that involve marked chemical changes such as notable additions of quartz, sericite, K-feldspar, alunite, or other 

minerals (Nolan, 1933). The depth of formation of 

the propylitic and zeolitic alterations has seldom if 

ever been determined with confidence, but erosion is 

commonly considered to have removed a few hundreds 

to several thousands of feet of cover. Original depths 

of formation are very unlikely to have been as great 

as the 20 to 30 km estimated for the epidote-rich zeolitic facies, as previously mentioned, in spite of the 

strikingly similar mineralogy of the two assemblages, 

one hydrothermal and the other metamorphic. Schistose mineral orientation, for example, is typically lacking in the hydrothermal environment, and is thus consit ent with a much shallower depth.

EPIDOTE IN HOT-SPRING SYSTEMS

The depths and corresponding temperatures of the 

epidote found in hot-spring systems are shown in 

figure 201.1, with some data on temperatures elsewhere 

in the drill holes.

The Wairakei thermal system is perhaps the most 

intensively explored and studied hot-spring system 

in the world (Steiner, 1953; Grange, 1955; Ban well 

and others, 1957; Coombs and others, 1959, p. 70-72). 

Steiner (1963; written communication, 1961) has recently found epidote at depths as shallow as 420 in, 

although it is more abundant at deeper levels.

Naboko and Piip (1961) have described alteration 

assemblages at Pauzhetsk, Kamchatka. Dacitic and 

andesitic tuffs and tuff-breccias, stratified nearly horizontally, occur to 800 m, the greatest drilled depth; 

andesite is abundant or dominant from 263 to 383 m 

and below 570 m. The rocks are considered to be 

Quaternary in age to 668 m and Tertiary at greater 

depths. Three major alteration zones are recognized:

(a) Kaolinite (perhaps with montmorillonite) 0 to 5 m 

depth; formed by near-surface attack by sulfuric acid.

(b) Zeolites and feldspar, 30 to 240 m; andularia and 

laumontite are abundant in the upper part, and laumontite and an unidentified but similar zeolite in the 

lower part with minor adularia, calcite, chlorite, or 

pyrite. (c) Propylitized green rocks from 250 to 

800 m, with chlorite, calcite, and pyrite as the characteristic minerals, and locally with hydrous mica, 

gypsum, anhydrite, an unidentified zeolite, and adularia; in the lowest of 4 propylitic subzones, epidote 

in individual grains with chlorite and zeolite was 

identified near 750 m, and albitized plagioclase near 

670m.

Temperatures attain a maximum of 180° C in the 

interval 125 to 400 m; below 400 m, temperatures decrease somewhat (Naboko and Piip, 1961, p. 110). 

Elsewhere, Piip and others (1961, p. 4) report that 

". . . in the central part of the thermal field tapped 

by the deep rotary well (west of the main heat liberation zone), a rise of temperature to 180 to 185°C is 

noted only down to depth 400 m. Still farther down, 

the temperature declines to 130 to 140° C."

Sigvaldason (Art. 200) described briefly the hydrothermal alterations at Hveragerdi and Reykjavik, 

Iceland, with emphasis on the distribution of epidote 

(see fig. 201.1, this article).

In each of these four hot-spring areas, epidote occurs in volcanic rocks of Tertiary or early Quaternary 

age. Conceivably each could have been buried deeply 

and altered, with the present distribution of minerals 

at least in part unrelated to the present hydrothermal 

environment. Except for Pauzhetsk, which has been 

drilled less than 50 m below the recognized epidote, 

this mineral becomes increasingly abundant at greater 

depth. This suggests that it is either forming now in 

the lower parts of the system, or that former epidote 

has been destroyed in the upper parts but is stable or 

metastable below.

Chemical analyses of the waters of the four spring 

systems are shown in the accompanying table. Of major significance is the fact that all 4 are from 8 to 9.4 in 

pH. At such a moderately high pH, all or most available CO2 is combined as carbonate and bicarbonate, and 

CO2 activity is low relative to H2O. These are precisely the conditions that favor formation of zeolites 

according to Zen (1961), and this is especially true 

where reactive volcanic rocks are available.
E82

100

GEOLOGICAL SURVEY RESEARCH 1962

TEMPERATURE, IN DEGREES CENTIGRADE

150 200 250

200

400

600

800

Boiling-point curve; upper limit of 

temperature in water-saturated 

environment in fractured, competent rocks

EXPLANATION

Pauzhetsk, I 

Kamchatka |

Shallowest occurrence 

of epidote in each drill 

hole

Temperature near 140 C(?) 

(see text)

To 2,200 meters and 146°C 

(nearly constant) below 

about 1,300 meters

FIGURE 201.1. Temperature-depth relations in drill holes containing epidote in four geothermal areas. Dashed curve for 

Wairakei is generalized for many drillholes. Dashed and queried curve for Pauzhetsk is inferred from published reports.

Some negative evidence is provided by the fact that 

epidote is absent or extremely rare at shallow depths 

in all studied hot-spring systems (White, 1955, p. 

103-122, and unpublished data). Possibly favorable 

altered andesitic rocks at Steamboat Springs, Nev. 

(Sigvaldason and White, 1961, 1962), and lathe in 

the Upper Basin of Yellowstone Park have been specifically searched for epidote, but it is virtually absent 

at depths ranging up to 685 feet. At Steamboat 

Springs a few probable grains of clastic epidote were 

identified from 11 feet in depth in drill hole GS-5 

(temperature 50°C), along with fresh clastic grains 

of hornblende.

The common occurrence of epidote in the propylitic 

altered rocks of epithermal ore deposits at depths that 

could originally have been a few thousands of feet, 

but almost certainly were not as deep as the green 

schist or the deeper part of the zeolite facies of metamorphism, is additional supporting evidence that the 

epidote of the four hot-spring systems is forming now, 

either stably or met ast ably.

Coombs and others (1959, p. 81-83) have carefully 

distinguished between load pressure and water pressure. Most hydrothermal syntheses are carried on 

under conditions of PH2o = Ptotai. The reaction
ARTICLE 201 

Chemical analyses, in parts per million, of waters from the hot spring systems known to contain epidote at depth

E83

Temperature (degrees C.) __ __

pH__ ___-_ _______________

SiO2 ________________

Al_. _________________ _______

Fe._ 

Ca ___. -_____-_______-_-_.

Mg-- 

Na _ ___ ___ ___ _ _____

K____. ______________________

Li___ _ ____________________

NH4 _ _______________________

As 

HCOa----.----------.---.-

CO3 -___-----_---___-__-_

SO4 

d. _-

F __ _______________________

Br _ _______________________

I______________.____________

PO4 -_-__--_--_ _________

B __________________

H2S___--__--__ ___ _

CO2 .... .....................

Total..-. ________

Pauzhetsk, Kamchatka

Spring »

100.6 

8.4

190 

None found 

None found 

64 

10 

1,010 

88

1. 1 

37 

9 

83 

1,680 

.8 

3.2 

None found

31

3,207

Deep well ' 

(120-132 m)

8.1

106

53 

4.9

| 977

.7

18 

24 

55 

1,540

0

2,779

Deep well 2 

(300-400 m)

2(9?)

131

119 

7.0

940

. 7

12 

71 

164 

1,400

2,845

Deep well 2 

(300-800 m)

2(8?)

113

158 

4.6

833

.7

13 

20 

363 

1,280

2,785

Walrakel, 

New Zealand

Well 4 3

>100 

8.6

386

26 

<. 1

f 1, 130 

\ 146 

12 

.9

35

35 

1,930 

6.2

26 

1. 1 

11

3,745

Hveragerdi, 

Iceland

Well*

100 

9.4

283 

.37 

. 12 

2.8 

.0 

174 

10 

. 1 

. 1 

.08 

24 

57 

72 

152 

2.1 

.5 

.0 

.08 

.85 

5.2

784

Reykjavik, 

Iceland

Well «

100 

9.2

172 

.18 

.00 

2. 1 

.4 

59 

2.3 

.0 

.0 

.00 

0 

35 

20 

36 

1.2 

. 1 

. 1 

. 14 

.04 

.2

333

»Ivanov (1958, p. 202-207).

' Naboko and Piip (1961, p. Ill); no pH given for deep samples from the drill hole 

containing epidote but proportions of COs to HCOs suggest 8 to 9.

' Well in eastern part of field, typical NaCl type according to Wilson (in Grange, 

1955, p. 37).

analcite + quartz >- albite + H2O, for example, 

occurs at about 280°C and 1,000 bars (Pn2o = P total}- 

This is probably similar to conditions of metamorphism of fine-grained water-saturated sedimentary 

rocks, where compaction and escape of water is still 

occurring but residual interstitial water and water 

released from hydrous phases cannot escape rapidly. 

The situation is very different, however, in the hotspring and probably the epithermal ore-deposit environments, where most or possibly all of the water is 

meteoric in origin (Craig and others, 1956; White, 

1955). White (1957) has suggested that meteoric water commonly circulates in fractured and permeable 

rocks to depths of as much as 10,000 feet. In systems 

involving downward penetration of meteoric water, 

pressure on the water cannot normally exceed hydrostatic pressure, so that PSi0 «=* Vs PIOOA] Coombs and 

his associates (1959, p. 81-82) believe that in such circumstances, and considering a series of related minerals of differing degrees of hydration, equilibrium will

t Collected 8/31/58 by Q. Bodvarsson, analysis by C. E. Roberson, U.S. QeolSurvey; also determined Mn, Cu, Pb, Zn, NOs, and NOs, 0.00 ppm.

s Eeykjavik II, Pvottalaugar; collected 10/2/58 by Q. Bodvarsson, analysis by 

H. C. Whitehead, U.S. Qeol. Survey; also determined Mn, Cu, Pb, Zn, NOs, 

NOs, 0.00 ppm; OH, 4 ppm.

favor a lower stage of hydration than where Ps2 o 

Plead* In the analcite-albite reaction considered above, 

for example, calculations suggest that albite stability 

will be lowered from about 280° C where /V2 o = P total 

= 1,000 bars to only 100° C at PH20 = Vs P total = 

1,000 bars. The calculated curve, incidentally, has a 

flat negative slope similar to that of the upper limit 

of recognized epidote (fig. 201.1). The same reasoning applied to the series heulandite-pumpellyite-prehnite-epidote-anorthite suggests that epidote (and each 

hydrous mineral) is stable at much lower temperatures 

and H2O pressures in hot-spring-epithermal environments than in normal metamorphic environments.

If epidote is actually stable only below the line indicated on figure 201.1, we should still expect to find 

relict metastable epidote in the high parts of some 

spring systems, as in the cited occurrence at Steamboat Springs. [Relatively unstable minerals such as 

pyroxene and hornblende are commonly preserved 

metastably in the upper hundred feet or so even in
E84 GEOLOGICAL SURVEY RESEARCH 1962

the hottest of spring systems. Epidote antedating the 

activity may also be preserved metastably at shallow 

depths but be decomposed at intermediate depths where 

reaction rates are higher.

SIGNIFICANCE OF EPIDOTE, IN THE EPITHEKMAL 

ENVIRONMENT

If the flatly inclined line of figure 201.1 does mark 

an upper limit for formation of epidote in hydrothermal environments, it will be useful in defining limits 

for depth of original cover in areas of altered volcanic 

rocks and epithermal ore deposits. If epidote occurs 

at the present land surface, for example, erosion of 

at least 400 m (1,300 ft) of cover is indicated. Some 

of the zeolites may likewise prove useful when their 

distribution in hot-springs systems is better understood.

E,EFEREN"CES

Banwell, C. J., Cooper, B. R., Thompson, G. E. K., and McCree, 

K. J., 1957, Physics of the New Zealand Thermal Area: 

New Zealand Dept. Sci. Indus. Research Bull. 123, 109 p.

Calkins, F. C., 1944, Outline of the geology of the Comstock 

Lode district, Nevada: U.S. Geol. Survey Prelim. Rept., 

35 p.

Christie, O. H. S., 1962, Feldspar structure and the equilibrium 

between plagioclase and epidote: Am. Jour. Sci., v. 260, 

p. 149-152.

Coats, R. R., 1940, Propylitization and related types of alteration and the Comstock Lode: Econ. Geology, v. 35, p. 1-16.

Coombs, D. S., 1954, The nature and alteration of some Triassic sediments from Southland, New Zealand: Roy. Soc. 

New Zealand Trans., v. 82, p. 65-109.

 1960, Lower grade mineral facies in New Zealand: 

Internat. Geol. Cong., 21st, Copenhagen 1960, Proc., pt. 13, 

p. 339-351.

Coombs, D. S., Ellis, A. J., Fyfe, W. S., and Taylor, A. M., 

1959, The zeolite facies, with comments on the interpretation of hydrothermal syntheses: Geochim. et Cosmochim. Acta, v. 17, p. 53-107.

Craig, H., Boato, G., and White, D. B., 1956, The isotopic geochemistry of thermal waters: Nuclear Sci. Ser., Rept. 19, 

Nuclear processes in geologic settings, p. 29-44.

Fyfe, W. S., 1960, Stability of epidote minerals: Nature, v. 

187, p. 497-498.

Fyfe, W. S., Turner, F. J., and Verhoogen, J., 1958, Metamorphic reactions and metamorphic facies: Geol. Soc. America Mem. 73, 259 p.

Grange, L. L., eel., 1955, Geothermal steam for power in New 

Zealand: New Zealand Dept. Sci. Indus. Research Bull. 

117, 102 p.

Ivanov, V. V., 1958, The principal regularities of the formation 

and distribution of the thermal waters of Kamchatka:

Akad. Nauk SSSR., Trudy Lab. Vulkanol., v. 13, p. 186- 

211. (In Russian)

Lindgren, Waldemar, 1933, Mineral deposits, 4th ed.: New 

York, McGraw-Hill Book Co., Inc., 930 p.

Merrin, Seymour, 1960, Synthesis of epidote and its apparent 

P-T stability curve [abs.]: Geol. Soc. America Bull., v. 

71, p. 1929.

Naboko, S. I., and Piip, B. I., 1961, Recent metamorphism of 

volcanic rocks in the region of the Pauzhetsk Hot Springs 

(Kamchatka) : Akad. Nauk SSSR, Trudy Lab. Vulkanol., 

v. 19, p. 99-114. (In Russian)

Nakovnik, N. I., 1954, Mineral facies, genesis, and practical 

importance of propylitized rocks: Zapiskl Vsessoyuz, 

mineralog. Obslichestra., v. 83, p. 85-94. (In Russian)

Nolan, T. B., 1933, Epithermal precious-metal deposits, in Ore 

deposits of the Western States: Am. Inst. Min. Bng., 

Lindgren vol., p. 623-640.

Pellizer, R., 1957, Transformazioni spirimentali alle condizione 

pneumatolitiche e idrotermali di rocce oflolitiche Appenniniche: Rendiconti Soc. Mineral. Ital., v. 13, p. 294-324. 

(In Italian)

Piip, B. I., Ivanov, U. V., and Averyev, V. V., 1961, The hydrothermal waters of Pauzhetsk, Kamchatka, as a source of 

geothermal energy: United Nations Conference on New 

Sources of Energy, preprint, p. 1-12.

Pistorius, C. W. F. T., Kennedy, G. C., and Sourirajan, S., 

1962, Some relation between the phases anorthite, zoisite, 

and lawsonite at high temperatures and pressures: Am. 

Jour. Sci., v. 260, p. 44-56.

Rutland, R. W. R., 1962, Feldspar structure and the equilibrium between plagioclase and epidote: A reply: Am. 

Jour. Sci., v. 260, p. 153-157.

Sigvaldason, G. E., and White, D. E., 1961, Hydrothermal 

alteration of rocks in two drill holes at Steamboat Springs, 

Washoe County, Nevada: Art. 331 in U.S. Geol. Survey 

Prof. Paper 424-D, p. D116-D122.

 1962, Hydrothermal alteration in drill holes GS-5 and 

GS-7, Steamboat Springs, Nevada: Art. 153 in U.S. Geol. 

Survey Prof. Paper 450-D, p. D113-D117.

Steiner, Alfred, 1953, Hydrothermal rock alteration at Wairakei, New Zealand: Econ. Geology, v. 48, p. 1-13.

 1963, Hydrothermal epidote: Mineralog. Mag. [In 

press]

White, D. E., 1955, Thermal springs and epithermal ore deposits, in Pt. 1 of Bateman, A. M., ed., Economic geology, 

50th anniversary volume, 1905-55: Urbana, 111., Econ. 

Geology Pub. Co., p. 99-154.

 1957, Thermal waters of volcanic origin: Geol. Soc.

America Bull., v. 68, p. 1637-1658. 

Wilshire, H. G., 1957, Propylitization of Tertiary volcanic

rocks near Ebbetts Pass, Alpine Co., Calif.: Calif. Univ.

Pub. in Geological Sciences, v. 32, p. 243-272. 

Zen, E-an, 1961, The zeolite facies: An interpretation: Am.

Jour. Sci., v. 259, p. 401-409.
202.

ARTICLE 202 E85 

THORIUM AND URANIUM IN SOME VOLCANIC ROCKS FROM THE CIRCUM-PACIFIC PROVINCE

By DAVID GOTTFRIED, EOOSEVELT MOORE, and ESMA CAMPBELL, Washington, D.C.

New data are available on the thorium and uranium 

contents of suites of volcanic rocks from the Mariana 

Islands in the western Pacific Ocean, eastern Oregon, 

and southwestern British Columbia. The suites are 

fairly representative of petrographic provinces of 

calcic-magma types that are found in the circumPacific province, peripheral to the margins of the 

Pacific Basin. The relatively narrow zone between 

the circum-Pacific province and the Pacific Basin province, the andesite line, is significant to problems bearing on the petrogenetic evolution of contrasting magma 

types and has been discussed by Schmidt (1957) with 

regard to the origin of the volcanic rocks of the Mariana Islands. In brief, the characteristic volcanic 

suites in the circum-Pacific province consist of basalt, 

andesite, dacite, and rhyolite. Their origin and composition are believed to be related to orogeny and to 

the presence of a sialic layer. Within the Pacific 

Basin, characterized by crustal stability and the absence of a sialic layer, the volcanic rocks are predominantly tholeiitic olivine basalts with relatively 

insignificant amounts of andesites and late alkalic 

differentiates.

The precision and accuracy of the methods for the 

determination of thorium and uranium have been discussed by Larsen and Gottfried (1960).

MARIANA ISLANDS

The volcanic rocks of Saipan consist mainly of hypersthene andesites of late Eocene to late Oligocene 

age and dacites of late Eocene age (Cloud and others, 

1956; Schmidt, 1957). Chemically the rocks are characterized by high SiOg, Al2Os, and CaO contents and 

a very low K2O content as compared with ordinary 

calc-alkalic suites. On the basis of Peacock's (1931) 

classification of igneous-rock series, the volcanic suite 

of Saipan is highly calcic with an alkali-lime index 

of approximately 65.

Basaltic rocks are abundant on Guam according to 

Schmidt (1957), and corresponding rock types on 

Guam are similar to those of Saipan. The volcanic 

rocks from the northern Mariana Islands (Pagan, 

Agrihan, and Alamagan) were derived from Quaternary volcanoes and are chiefly basaltic and andesitic flows. The andesites are similar to those of 

Saipan and to those of the Izu Peninsula region of 

Japan.

The thorium and uranium contents of 26 samples 

of chemically analyzed rocks that were supplied by 

E. G. Schmidt, J. I. Tracey, Jr., and Gilbert Corwin 

of the Geological Survey were determined. These 

data and the Th/U ratios are listed in table 202.1, 

and those of Saipan, Pagan, and Guam are plotted 

on Larsen type (1938) variation diagrams on figure 

202.1. Hence these data can be compared directly 

with suites from other localities (Larsen and Gootfried, 

1960, Gottfried and others, 1962).

The dacites of Saipan are the most siliceous rocks 

in the Mariana Islands that were analyzed for thorium 

and uranium. They average only a little over 1 ppm 

(part per million) of thorium and uranium. The 

andesites average 0.7 ppm of uranium and 1.1 ppm 

of thorium. Except for one sample with a Th/U 

ratio of 2.9, the Th/U ratios of the rocks of Saipan 

fall between 0.4 and 1.8, and average 1.2.

The basalts and andesites of Pagan average about 

0.5 and 0.7 ppm of uranium and 0.9 and 1.2 ppm 

of thorium respectively. The Th/U ratios of the rocks 

average 2.1 for the basalts and 2.2 for the andesites.

o

The two basalts from Agrihan and Alamagan contain 

about 1 ppm each of thorium and uranium.

The basaltic rocks of Guam average about 0.4 ppm 

of uranium and in this respect are similar to those of 

Pagan. However their average thorium content 

1.4 ppm is greater, and hence the basalts of Guam 

have a larger Th/U ratio. The Th/U ratios have a 

wide scatter; omitting 1 sample which has a Th/U 

ratio of 14, the average Th/U ratio is 2.9. One sample 'of andesite from Guam contains 0.35 ppm of uranium and 0.72 ppm t)f thorium.

Five samples of andesite and 1 sample of gabbro 

from the Palau islands were analyzed for uranium 

only; the andesites average 0.26 ppm and the gabbro 

0.14 ppm. These data indicate that these andesites 

have a lower uranium content than those of the Mariana Islands.
E86 GEOLOGICAL SURVEY RESEARCH 1962

TABLE 202.1. Thorium and uranium contents and Th/U ratios 

of volcanic rocks from Saipan, Pagan, Guam, Alamagan, and 

Agrihan, Mariana Islands

Sample No.

S 135_____.

S 636 A__._

S 107-.. _.

S 15_______

S 67 A_____

S 139____ __

S 448-_____

S 317_-_-_-

PC-67_____

P-8-______

P-5________

P-6________

PC-75.. ...

PC-55_. __.

PC-89... _.

PC-61_____

Go 7-ld__._

Ee 4-2c. ...

Go 7-lc_ 

Dh 6-lb____

Dh6-lc_...

Fd2-lb____

G-S48-85A.

MSS 37.. 

AG-4______

AL-7__.___

Rock types

Saipan

Augite-hypersthene

andesite. 

Augite andesite _ 

Augite-hypersthene

andesite. 

do- _ - . _

do. __._._

Hornblende-bearing dacite vitrophyre.

Pagan

Augite basalt ______

augite basalt. 

_ _do_ _

. __do _ . _

Andesite

Basalt. ___ _ _

Guam

Basalt. _________ _

_ _do_

. _do._ _ _

_ _do_ _ __ ....

_do_ _-__ _

do-_ __ ...

Augite basalt

Andesite _ _ _ __

Agrihan

Augite basalt ______

Alamaga

augite basalt.

Position 

% SiO2 

+K20 

 CaOFeO-MgO

-1. 9

-.9

+ 1.2

+ 2.3

+ 5.5

+ 24.4

+ 26. 0

+ 26.7

-11.9

-11.0

-9.3

-8.7

-4.7

-4. 5

-1. 1

+ 8.5

-12.0

-8. 5

-7. 6

C Q

-4.8

-3.5

+ 3.2

+ 4.8

-9.9

n

-6.7

Thorium 

(ppm)

1. 2

1.2 

. 4

1.4

.6

1.6

1. 1

.88

0.81

.79

. 53

. 77

1. 1

.72

1.4

1.9

2.4

1. 1

. 6

1.5

2.0

1.3

.81

. 72

1.0

0. 98

Uranium 

(ppm)

0.70

1. 1 

.90

.48

.53

1. 5

.82

1. 1

0. 21

.48

.44

. 53

. 54

.23

.58

1.3

0. 17

.41

.28

.41

.43

.41

. 90

.35

1. 1

0. 95

Th/U

1.7

1.8 

.4

2.9

1. 1

1. 1

1.3

.8

3.9

1.7

1.2

1. 5

2.0

3. 1

2.4

1. 5

14. 1

2.7

2.2

3.7

4.7

3 2

. 9

2. 1

1.0

1.0

MOUNT GARIBALDI AREA. SOUTHWESTERN BRITISH 

COLUMBIA

The volcanic rocks from southwestern British Columbia are Quaternary in age and range in composition from basalt to rhyodacite. A report on the 

petrology of the suite of rocks has been published by 

Mathews (1957), who furnished the samples for this 

study. Chemically the rock series is characterized by 

relatively large amounts of A12C>3 and CaO and low 

K2O. The suite is strongly calcic as indicated by the 

Peacock (1931) alkali-lime index of 64.0 and is similar in many respects to the volcanic rocks of the High 

Cascades (Mathews, 1957). Analyses for thorium and 

uranium have been made on 13 samples (table 202.2 

and figure 202.2). Thorium and uranium are present

TABLE 202.2. Thorium and uranium contents and Th/U ratios 

of volcanic rocks from the Mount Garibaldi area, southwestern 

British Columbia

Sample 

No.

526-._-_

381_-___

532__-__

451_____

554_____

522_____

511___-_

527_.__.

420__-_-

547_--_-

421_____

535____.

469_____

Rock type

Gabbro. _______ __

Basalt. _______ ____

do

Andesite breccia ___

_____do-__------

_____do _

_.___do __

____.do _-

do _

_____do ___ ____

_____do__---_-___--_

Position M SiOz-fK2OCaOFeOMgO

-18.0 

-11.3 

-9.3 

+ 2.0

+ 4.4 

+ 6.6 

+ 7.0 

+ 12.0 

+ 12. 5 

+ 13. 0 

+ 15. 0 

+ 19.0 

+ 24. 1

Thorium 

(ppm)

1.4 

1. 1 

1. 1 

1.6 

1.3 

1.0 

. 99 

2.0 

2.6 

3. 1 

2. 5 

2.7 

4.0

Uranium 

(ppm)

1. 2 

1. 1

.47 

1. 1 

.90 

. 90 

1. 1 

1. 1 

1.6 

1. 2 

1.3 

1.9 

1. 6

Th/U

1. 2 

1.0 

2.3

1. 5 

1.4 

1. 1 

. 9 

1.8 

1.6 

2. 6 

1.9 

1.4 

2. 5

in only slightly greater amounts in the dacitic rocks 

than in the mafic rocks. The increase is less than 

twofold. The average Th/U ratio of the mafic rocks 

and the dacites is about 1.5. The one sample of 

rhyodacite has a Th/TJ ratio of 2.5. The suite as a 

whole has a Th/U ratio similar to that of the andesites and dacites of Saipan.

STRAWBERRY MOUNTAIN, EASTERN OREGON

The upper sequence of volcanic rocks of Strawberry 

Mountain ranges from basalt to rhyolite but consists 

chiefly or basaltic and hypersthene andesites. The
ARTICLE 202 E87

 

I

-iuz

50-0

"* 5 n

Z> 0

 

LJ

,3.9 -14.1

-10 

POSITION

10 " -10 

POSITION

10

SAIPAN PAGAN GUAM

FIGURE 202.1. Variation diagrams showing the distribution of thorium, uranium, and Th/U ratios in volcanic rocks from Saipan

Pagan, and Guam.

structural and petrologic relations of these volcanic 

rocks of Miocene age to those of the Columbia Eiver 

Plateau has been discussed by Thayer (1957), who 

points out that these rocks are remarkably similar 

to the lavas of the High Cascades. Chemical analyses 

of these rocks were made by S. M. Berthold, E. A. 

Nygaard and E. N. Eccher of the U.S. Geological 

Survey. The alkali-lime index of these volcanic rocks 

is about 62, thereby placing the suite in the calcic 

series.

The thorium and uranium contents of eight samples 

are listed in table 202.3 and plotted on figure 202.3.

tnt- 4

(T

o^o

D (T 2

I- 1

TABLE 202.3. Thorium and uranium contents and Th/U ratios 

of volcanic rocks of Strawberry Mountain, Oreg.

< Z

0-0

Sample No.

TPT 47-60. . 

TPT 48-125.

TPT 48-122-

TPT 48-21 __ 

TPT 48-23..

TPT 48-132_

TPT 48-131_ 

TPT 48-133.

Eock types

Olivine basalt _

Basaltic andesite__

Andesite (plug)____

Pyroxene andesite. 

-__-_do-_-_-___--_

Pyroxene dacite___

Hornblende dacite_ 

Rhyolite

Position }i

S1O2+ K2OCaOFeOMgO

-7.6 

-1.0

+ 4. 5

+ 9. 0 

+ 8. 9 

+ 19. 2

+ 20. 2 

+ 22. 6

Thorium 

(ppm)

1. 3 

1. 5

2. 2

3. 2 

2. 8 

5. 5

5. 8 

6. 4

Uranium 

(ppm)

0. 46 

. 61

.75

1. 1 

1. 1 

2. 2

2. 2 

2. 5

Th/U

2. 8 

2. 4

2. 9

2. 9 

2. 5 

2. 5

2. 6 

2.6

-20 -10 0 10 20 

POSITION y3 SiO2 + K2O-CaO-FeO-MgO

30

r . , , n ,, , -. . ,, There is approximately a fivefold increase in the

thorium and uranium contents of the most siliceous

FIGURE 202.2. Variation diagram showing the distribution of 

thoriunl) uranium, and Th/U ratios in volcanic rocks from the

Mount Garibaldi area, southwestern British Columbia.
E88 GEOLOGICAL SURVEY RESEARCH 1962

7 r

oe 

< z

QDOC 3

f 2

oe

<z 3 Q-O

-20 -10 0 10 20 

POSITION V^SiOj + KzO-CaO-FeO-MgO

30

FIGURE 202.3. Variation diagram showing the distribution of 

thorium, uranium, and Th/U ratios in volcanic rocks from 

Strawberry Mountain, eastern Oregon.

rocks over the amounts contained by the most mafic 

rocks.

The Th/U ratio of all the rock types is strikingly 

similar and averages 2.6. The Th/U ratio of the 

andesites is about V/2 times greater than the Th/U 

ratio of the andesites of Saipan and also about 1% 

greater than the Th/U ratio of the andesites and dacites 

of the Mount Garibaldi area.

DISCUSSION

The thorium and uranium contents and especially 

the Th/U ratios of many of the volcanic rocks of 

the circum-Pacific province are in marked contrast to

those of volcanic rocks from the Pacific Basin province 

and of batholithic rocks from the Western United 

States (Larsen and Gottfried, 1960). Most of the 

andesites and dacitic rocks discussed here have Th/U 

ratios ranging from about 1% to about 2^. The 

bulk of the lavas from Hawaiian volcanoes are tholeiitic basalts which have an average Th/U ratio of 

3.2, and the average Th/U ratio of the southern California, Sierra Nevada, and Idaho batholiths falls between 3i/£ and 5y2.

Some petrologists (Tilley, 1950; Waters, 1955; 

Byers, 1961) believe that andesitic rocks of the orogenic belts were formed by the assimilation of geosynclinal sediments by basaltic magma. The low 

Th/U ratios may reflect this proposed origin for the 

andesitic rocks. It is interesting to note that the 

curves for thorium and uranium on the variation diagrams (figs. 202.2 and 202.3) show a somewhat linear 

variation as compared to the curvilinear trends shown 

by the rocks of the southern California batholith 

(Larsen and Gottfried, 1960). But to adequately discuss the geochemical cycle of thorium and uranium 

in the erogenic belts, data are needed on the thorium 

and uranium contents of graywackes, argillites, and 

cherts which characterize the sediments of the eugeosynclinal belts.

The data presented show that strong regional variations of thorium and uranium in igneous rocks can be 

closely correlated with the chemical characteristics 

of the rock series and their tectonic setting. Similar 

conclusions were reached by Chao and Fleischer (1960) 

in their study of the abundance of zirconium in igneous rocks.

REFERENCES

Byers, F. M., Jr., 1961, Petrology of three volcanic suites,

Umnak and Bogoslof Islands, Aleutian Islands, Alaska:

Geol. Soc. America Bull., v. 72, p. 93-128. 

Chao, E. C. T., and Fleischer, Michael, 1960, Abundance of

zirconium in igneous rocks: Internat. Geol. Gong., 21st,

Copenhagen 1960, Proc., pt. 1, p. 106-131. 

Cloud, P. E., Schmidt, R. G., and Burke, H. W., 1956, General

geology [of Saipan, Mariana Islands]; U.S. Geol. Survey

Prof. Paper 280-A, p. 1-126. 

Gottfried, David, Moore, Roosevelt, and Caemmerer, Alice,

1962, Thorium and uranium in some alkalic igneous rocks

from Virginia and Texas: Art. 27 in U.S. Geol. Survey

Prof. Paper 450-B, p. B70-B72. 

Larsen, E. S., Jr., 1938, Some new variation diagrams for

groups of igneous rocks: Jour. Geology, v. 46, p. 505-520. 

Larsen, E. S., 3d, and Gottfried, David, I960, Uranium and

thorium in selected suites of igneous rocks: Am. Jour.

Sci., v. 258-A, p. 151-169.
ARTICLE 203 E89

Mathews, W. H., 1957, Petrology of Quaternary volcanics of 

the Mount Garibaldi map-area, southwestern British Columbia: Am. Jour. Sci., v. 255, p. 400-415.

Peacock, M. A., 1931, Classification of igneous rock series: 

Jour. Geol., v. 39, p. 54-67.

Schmidt, R. G., 1957, Petrology of the volcanic rocks [of 

Saipan, Mariana Islands]: U.S. Geol. Survey Prof. Paper 

280-B, p. 127-175.

Thayer, T. P., 1957, Some relations of later Tertiary volcanology and structure in eastern Oregon: Internat. Geol.

Cong., 20th, Mexico City 1956, Section 1 Cenozoic volcanology, v. 1, p. 231-245. 

Tilley, C. E., 1950, Some aspects of magmatic evolution-: Geol.

Soc. London Quart. Jour., v. 106, pt. 1, p. 37-61. 

Waters, A. C., 1955, Volcanic rocks and the tectonic cycle:

Geol. Soc. America Spec. Paper 62, p. 703-722.

203. THE QUARTZ DIORITE LINE IN NORTHWESTERN NORTH AMERICA 1

By JAMES G. MOORE, ARTHUR GRANTZ, and M. C. BLAKE, JR., Menlo Park, Calif.

Buddington (1927) demonstrated that the granitic 

rocks of southeastern Alaska show a regional zonation 

in composition: they pass progressively from quartz 

diorite on the west, through granodiorite, to quartz 

monzonite on the east. A similar relation has been 

described in the western conterminous United States 

(Moore, 1959) where it is possible to delineate a line 

(called the quartz diorite line) roughly parallel to 

the Pacific Coast, which separates a western zone in 

which quartz diorite is the dominant granitic rock 

from an eastern zone in which granodiorite or quartz 

monzonite is the dominant granitic rock. The granitic 

rocks considered are chiefly of Jurassic and Cretaceous 

age, but they include some Tertiary intrusives and possibly some Paleozoic intrusives. In the present paper 

the quartz diorite line is extended across Alaska and 

the significance of the regional chemical differences 

that it helps to outline is explored.

Data regarding the dominant granitic rock type in 

each of approximately 100 mapped areas in Alaska 

and Canada (British Columbia and Yukon) are shown 

on figure 203.1. In order to compare the data, the 

same system is used in classifying the granitic rocks 

of all the mapped areas, even if the author of a published report used a different system. The classification system is the same system used in the earlier 

compilation of data on granitic rocks of the western 

conterminous United States (Moore, 1959). On figure 

203.1 symbols are used for three main groups of 

granitic rocks: (a) quartz diorite, (b) granodiorite, 

and (c) quartz monzonite or granite.

iThis article -was first printed as Art. 183 in U.S. Geol. Survey 

Prof. Paper 424-C. It is reprinted here to include the list of references on fig. 203.1 that -was inadvertently omitted from the first 

printing.

The granitic rocks of Alaska show an increase in 

the proportion of K-feldspar relative to plagioclase 

inland from the southern continental margin; this 

change is very similar to that found in the granitic 

rocks of California, Oregon, and Washington. In 

each area the granitic rocks become progressively 

richer in K and Si, and poorer in Na, Fe, and Mg 

inward from the continental margin.

The quartz diorite line drawn across Alaska (fig. 

203.1) separates a zone adjacent to the southern coast 

in which quartz diorite is the dominant granitic rock 

from a zone to the north where granodiorite, quartz 

monzonite, or granite is the dominant granitic rock. 

The quartz diorite line does not mark an abrupt 

change in composition, and its position is dependent 

on the classification system employed.

From the latitude of Ketchikan to that of Skagway, 

the line is close to the international boundary between 

Alaska and Canada. It trends northwestward from 

Skagway on the north side of the St. Elias and 

Wrangell Mountains. The line reaches its farthest 

north point about 70 miles south of Fairbanks from 

where it curves to the southwest and ultimately passes 

out to sea at Bristol Bay. The line remains north of 

the Aleutian Islands, many of which are underlain by 

quartz diorite, but is probably south of the Pribilof 

Islands. The line is subparallel to the arcuate pattern 

of the regional structural elements of southern Alaska.

By combining all the data on^ie-positfon of the 

quartz diorite line in northwestern North America, 

the line can be traced for more than 3,500 miles from 

the Gulf of California to Bristol Bay in the Bering 

Sea (fig. 203.2). The quartz diorite line is approximately parallel to the west coast of North America,
E90 GEOLOGICAL SURVEY RESEARCH 1962

.«.»- *.^

^Tairbanks ik 76 I 

 61 76 * * (

*"1 

 " 1 "& *'" *",;T^' \

-i ./*" Xv

EXPLANATION

Dominant granitic rock types:

A30 

Quartz diorite

18 

Granodiorite

22 

Quartz monzonite or granite

Areas with significant amounts of quartz diorite 

present though not dominant

Quartz diorite line (northern limit of quartz diorite 

as the dominant granitic rock)

c^11

pA.CIFIC

SL, o \^i" "u^ "r /U./IOA/I v., 4l>> 

* / ? ^N£ «/ r/ 3ZA!

/ 3Sf tt~L.

'^Ifhk rK ,^\y^4|5>

tffCb

4

JDIAK

"ANDS 50 0 50 100 150 MILES i i 1 1 1 1

WO °

->^TM" *** *" *-^.» .»* 

FIGURE 203.1. Dominant granitic rock types in some areas of Alaska and adjoining parts of Canada. Numbers refer to references

listed on opposite page.
ARTICLE 203

REFERENCES USED IN COMPILING FIG. 203.1

E91

8.

9.

10.

11.

12.

13.

14.

15.

16.

17.

18.

19.

20.

21.

22.

23.

24.

25.

27.

28.

29.

30.

31.

32.

33.

35.

36.

37.

38.

39.

40.

41.

42.

43.

44.

45.

Buddington, A. F., 1927, Jour. Geology, v. 35, p. 224. 

Becker, G. F., 1898, U.S. Geol. Survey 18th Ann. Kept.,

Pt. Ill-a, p. 1-86. 

Byers, F. M., Jr., 1959, U.S. Geol. Survey Bull. 1028-L,

p. 267-369. 

Becker, G. F., 1898, U.S. Geol. Survey 18th Ann. Kept.,

Pt. Ill-a, p. 1-86; and Byers, F. M., Jr., 1959, U.S.

Geol. Survey Bull. 1028-L, p. 267-369. 

Byers, F. M., Jr., 1957, U.S. Geol. Survey Bull. 1024-1,

p. 179-213. 

Fraser, G. D., and Snyder, G. L., 1959, U.S. Geol. Survey

Bull. 1028-M, p. 371-408. 

Grantz, Arthur, 1956, U.S. Geol. Survey Bull. 1024-D,

p. 95-106. 

Kennedy, G. C., and Waldron, H. H., 1955, U. S. Geol.

Survey Bull. 1028-A, p. 1-18. 

Capps, S. R., 1937, U.S. Geol. Survey Bull. 880-C, p. 111-

184. 

Powers, H. A., Coats, R. R., and Nelson, W. H., 1960, U.S.

Geol. Survey Bull. 1028-P, p. 521-554. 

Rossman, D. L., 1959, U.S. Geol. Survey Bull. 1058-B,

p. 33-59.

 1959, U.S. Geol. Survey Bull. 1058-E, p. 139-216. 

Sainsbury, C. L., 1957, U.S. Geol. Survey Bull. 1024-G,

p. 141-161. 

Seitz, James, 1959, U.S. Geol. Survey Bull. 1058-C, p. 61-

120.

Moffit, F. H., 1938, U.S. Geol. Survey Bull. 894, 137 p. 

Mertie, J. B., Jr., 1938, U.S. Geol. Survey Bull. 903, 96 p. 

Cady, W. M., Wallace, B. E., Hoare, J. M., and Webber,

E. J., 1955, U.S. Geol. Survey Prof. Paper 268, 132 p. 

Sainsbury, C. L., written communication, 1960. 

Mertie, J. B., Jr., 1938, U.S. Geol. Survey Bull. 903, 96 p. 

Coats, R. R., 1956, U.S. Geol. Survey Bull. 1028-E, p. 83-

100. 

Juhle, Werner, 1955, U.S. Geol. Survey open-file report,

74 p.

Buddington, A. F., 1929, U.S. Geol. Survey Bull. 807, 124 p. 

Grant, U. S., and Higgins, D. F., 1910, U.S. Geol. Survey

Bull. 443, 89 p. 

Kennedy, G. C., 1953, U.S. Geol. Survey Prof. Paper 251,

46 p.

26. Hoare, J. M., written communication, 1960. 

Robinson, G. D., and Twenhofel, W. S., 1953, U.S. Geol.

Survey Bull. 998-C, p. 59-84. 

Gault, H. R., Rossman, D. L., and Flint, G. M., Jr., 1953,

U.S. Geol. Survey Bull. 998-D, p. 15-58. 

Gault, H. R., and Fellows, R. E., 1953, U.S. Geol. Survey

Bull. 998-A, p. 1-13. 

Ray, R. G., 1954, U.S. Geol. Survey Bull. 1004, 86 p.; and

Ray, J. C., 1933, U.S. Geol. Survey Bull. 849-C, p. 165-

228. 

Bressler, C. T., 1951, U.S. Geol. Survey Bull. 963-C,

p. 81-92. 

Kennedy, G. C., and Walton, M. S., Jr., 1947, U.S. Geol.

Survey Bull., 947-C, p. 39-64. 

34. Twenhofel, W. S., Robinson, G. D., and Gault, H. R.,

1947, U.S. Geol. Survey Bull. 947-B, p. 7-38. 

Moffit, F. H., 1943, U.S. Geol. Survey Bull. 933-B, p. 103-

174.

 1954, U.S. Geol. Survey Bull. 989-D, p. 65-218.

Guild, P. W., and Balsley, J. R., Jr., 1942, U.S. Geol. Survey Bull. 936-G, p. 171-187.

Brabb, E. E., and Miller, D. J., 1960, U.S. Geol. Survey 

open-file report, 29 p.

Keller, A. S., and Reiser, H. N., 1959, U.S. Geol. Survey 

Bull. 1058-G, p. 261-298.

Mertie, J. B., Jr., 1925, U.S. Geol. Survey Bull. 773, p. 215- 

267.

Martin, G. C., and Katz, F. J., 1912, U.S. Geol. Survey 

Bull. 485, 138 p.

Watson, K. DeP., 1948, British Columbia Dept. Mines 

Bull., 25, 74 p.

Watson, K. DeP. and Mathews, W. H., 1944, British Columbia Dept. Mines Bull. 19, 52 p.

Aitken, J. D., 1959, Canada Geol. Survey Mem. 307, 89 p.

Kerr, F. A., 1948, Canada Geol. Survey Mem. 246, 94 p.

46. Kindle, E. D., 1952, Canada Geol. Survey Mem. 268, 68 p.

47. Kerr, F. A., 1948, Canada Geol. Survey Mem. 248, 84 p.

48. Muller, J. E., 1958, Canada Geol. Survey Prof. Paper 58-9, 

9 p.

49. Cockfield, W. E., and Bell, A. H., 1926, Canada Geol. Survey Mem. 150, 63 p.

50. Brown, A. S., and Jeffery, W. G., 1960, British Columbia 

Dept. Mines Prelim. Map (Southern Queen Charlotte 

Island).

51. Canada Geological Survey, 1957, Canada Geol. Survey 

Prelim. Ser. Map 9 1957.

52. Barker, Fred, 1957, U.S. Geol. Survey Geol. Quad. Map 

GQ-100.

53. Canada Geological Survey, 1957, Canada Geol. Survey 

Prelim. Ser., Map 9 1957.

54. Hopkins, D. M., written communication, 1960.

55. Park, C. F., Jr., 1933, U.S. Geol. Survey Bull. 849-G, p. 381- 

424.

56. Ross, C. P., 1933, U.S. Geol. Survey Bull. 849-H, p. 425- 

468.

57. Reed, J. C., 1933, U.S. Geol. Survey Bull. 849-D, p. 231- 

287.

58. Ross, C. P., 1933, U.S. Geol. Survey Bull. 849-E, p. 289- 

333

59. Moffit, F. H., and Mertie, J. B., Jr., 1923, U.S. Geol. Survey 

Bull. 745, 149 p.

60. Capps, S. R., 1927, U.S. Geol. Survey Bull. 791, 92 p.; 

Waring, G. A., 1936, U.S. Geol. Survey Bull. 861, 57 p.; 

and this report.

61. Prindle, L. M., 1913, U.S. Geol. Survey Bull. 525, 220 p.

62. Paige, Sidney, and Knopf, Adolph, 1907, U.S. Geol. Survey 

Bull. 327, 71 p.

63. Capps, S. R., 1916, U.S. Geol. Survey Bull. 630, 130 p.

64. Hummel, C. L., written communication, 1961.

65. Wahrhaftig, Clyde, 1958, U.S. Geol. Survey Prof. Paper 

293-A, p. 1-68; and Moffit, F. H., 1915, U.S. Geol. 

Survey Bull. 608, 80 p.

66. Sable, E. G., written communication, 1961.

67. MacKevett, E. M., written communication, 1961.

68. Muller, E. H., and Alien, H. B., written communication, 

1959.

69. Reed, John C., Jr., written communication, 1961.

70. Wahrhaftig, C., written communication, 1961; and this

report. 

71-75. This report.

76. Mertie, J. B., Jr., 1937, U.S. Geol. Survey Bull. 872, 

276 p.

77. Tailleur, I. L., oral communication, 1961; and Smith, P. S., 

and Mertie, J. B., Jr., 1930, US. Geol. Survey Bull. 

815, 351 p.

78. Moffit, F. H., 1938, U.S. Geol. Survey Bull. 904, 54 p.

79. 1935, U.S. Geol. Survey Bull. 886, 38 p.

80. Chapin, Theodore, 1918, U.S. Geol. Survey Bull. 668, 67 p.

81. Mertie, J. B., Jr., 1937, U.S. Geol. Survey Bull. 868-D, 

p. 145-178.

82. Moffit, F. H., 1918, U.S. Geol. Survey Bull. 675, 82 p.

83. Eakin, H. M., 1916, U.S. Geol. Survey Bull. 631, 88 p.

84. Smith, P. S., 1917, U.S. Geol. Survey Bull. 655, 162 p.

85. Capps, S. R., 1935, U.S. Geol. Survey Bull. 862, 101 p.

86. Spurr, J. E., 1899, U.S. Geol. Survey 20th Ann. Rept., 

Pt. 7-B, p. 31-264.

87. Matzko, J. J., Jaffe, H. W., and Waring, C. L., 1958, Am. 

Jour. Sci., v. 256, p. 529-539.

88. Brooks, A. H., 1911, U.S. Geol Survey Prof. Paper 70, 

234 p.

89. Hoare, J. M., and Coonrad, W. L., 1959, U.S. Geol. Survey 

Misc. Geol. Inv. Map 1-292.

90. 1959, U.S. Geol. Survey Misc. Geol. Inv. Map 1-285.

91. Brown, J. S., 1926, U.S. Geol. Survey Bull. 783-D, p. 97- 

150.

92. White, M. G., 1952, U.S. Geol. Survey Circ. 185, 13 p.

93. Brosge", W. P., written communication, 1961.

94. Miller, D. J., oral communication, 1961.

95. Sharp, R. P., and Rigsby, G. P., 1956, Am. Jour. Sci., 

v. 254, p. 110-122.

96. Mertie, J. B., Jr., 1940, U.S. Geol. Survey Bull. 918, 97 p.

667422 O - 63 - 7
E92 GEOLOGICAL SURVEY RESEARCH 1962

Greatest extent of early 

Tertiary marine transgression

Continental slope 

Dashed line outer edge continental shelf

FIGURE 203.2. The position of the quartz diorite line in 

northwestern North America and its relation to other 

features of the continental margin.

and especially to the inner edge of the continental 

slope. On the average, it is about 150 miles inland 

from the inner edge of the continental slope.

Two prominent landward embayments are present 

in the trend of the quartz diorite line. In Alaska the 

line curves northward 300 miles from the continental 

slope, and in Oregon and Washington, the line curves 

eastward more than 400 miles. Both embayments are 

reflected in more subdued landward curves in the 

trend of the edge of the continental shelf.

Although the embayment in Oregon and Washington is only faintly reflected by the configuration of the 

coast line or of the inner edge of the continental slope, 

it is paralleled more closely by coast lines of the geologic past. The line representing the greatest transgression of seas of early Tertiary age (dotted line, fig. 

203.2) is deflected markedly into the embayment in 

Washington and Oregon. Moreover, some evidence 

suggests that the continental slope within this embayment may have been father east than at present during early Tertiary time. Eocene rocks of the Olympic 

Peninsula in northwestern Washington are marine 

throughout: chiefly pillow basalt, thin-bedded graywacke, argillite, and red foraminiferal limestone with 

associated manganese deposits. Such an assemblage 

of rocks, and the absence of shallow-water sedimentary structures in them, suggest deposition in deep 

water, possibly on or beyond the continental slope.

The quartz diorite line is approximately parallel to 

major Mesozoic and Cenozoic tectonic elements, such 

as elongation of granitic intrusives, foliation of metamorphic rocks, axes of folds, and strike of large 

faults. Major Quaternary volcanoes are alined parallel to and slightly on the seaward side of the line (see 

fig. 203.2).

It is evident from the relations described above and 

shown on figure 203.2 that the quartz diorite line is a 

fundamental feature of the western continental margin of North America. It trends approximately parallel to the edge of the continent as indicated today 

by its relation to the shore line and to the continental 

slope, and as indicated in the past by its relation to 

early Tertiary shore lines. The difference in chemical 

composition of granitic rocks on each side of the line 

is presumably inherited from chemical differences in 

the crust, which existed before the granitic rocks were 

emplaced (Moore, 1959). A possible explanation of 

this difference is that west of the quartz diorite line 

the crust was thinner than in the area east of it and 

composed of a larger percentage of eugeosynclinal 

sediments and volcanic rocks that are relatively rich 

in sodium. The character of the Precambrian rocks 

west of the line is little known because exposures are
ARTICLE 204 E93

extremely rare. However, the authors believe that 

Precambrian granitic rocks are practically nonexistent 

west of the line. If this is true, then the Mesozoic 

and Cenozoic granitic rocks there are largely in their 

first cycle, that is, they were emplaced in a basement 

containing no older granitic rock.

East of the quartz diorite line, before emplacement 

of the Mesozoic granitic rocks, the crust was probably 

thicker and was composed of a larger percentage of 

platform-type sediments as well as great volumes of 

granitic rock. Thus, east of the line the crust was 

relatively rich in potassium relative to sodium. The 

Mesozoic and Cenozoic granitic rocks generated in 

this crust therefore contain a larger proportion of 

potassium feldspar relative to sodium-calcium feldspar than the granitic rocks of the continental border.

Quartz diorite appears to be the dominant granitic 

rock encircling the entire Pacific basin. A detailed 

search has not been made of the literature to establish 

the position of the line in the southern or western

Pacific. However, evidence at hand suggests that the 

line passes west of Kamchatka and between Japan 

and Korea. A similar change in composition of granitic rocks seems to occur slightly north of New 

Guinea, and in northern New Zealand. The quartz 

diorite line probably lies from 200 to 600 miles on 

the outside (landward side) of the andesite line (Marshall, 1911, and Macdonald, 1949), and consequently 

the island arcs of the Pacific border are between the 

andesite line and the quartz diorite line.

REFERENCES

Buddington, A. F., 1927, Coast Range intrusives of southeastern Alaska: Jour. Geology, v. 35, p. 224.

Macdonald, G. A., 1949, Hawaiian petrographic province: Geol. 

Soc. America Bull., v. 60, p. 1541-1596.

Marshall, Patrick, 1911, Oceania: Handb. der Regionalen 

Geologie, v. 7, no. 2, 36 p.

Moore, J. G., 1959, The quartz diorite boundary line in the 

western United States: Jour. Geology, v. 67, no. 2, p. 

198-210.

204. SYENITE COMPLEX OLDER THAN THE IDAHO BATHOLITH, BIG CREEK QUADRANGLE, CENTRAL IDAHO

By B. F. LEONARD, Denver, Colo.

An arcuate complex of syenite and quartz syenite 

involved with amphibolite is exposed in the Ramey 

Ridge area, northeast corner of the Big Creek quadrangle, Idaho and Valley Counties, Idaho. (See fig. 

204.1.) The name Ramey Ridge Complex is here 

given to the units of syenitic rock shown on that figure. 

The syenitic rocks of the Ramey Ridge Complex are 

the only outcropping felsic plutonic rocks demonstrably older than the Idaho batholith that have been 

reported from this region. These rocks were earlier 

interpreted as slightly older facies of the Idaho batholith itself (Leonard in Larsen and others, 1958, p. 51; 

Ross m Shenon and Ross, 1936, fig. 2 and p. 8-9). 

The complex has not been conclusively dated, but the 

evidence from geologic relations and lead-alpha determinations on zircon suggests a Paleozoic age.

The chief components of the complex are alkali 

syenites. They lack feldspathoids and sodic pyriboles, 

usually contain a little quartz, and are very slightly 

enriched in cerium and niobium.

The syenite complex, interpreted as an assemblage 

of slightly to highly contaminated magmatic rocks, 

was emplaced mostly within old metavolcanic rocks 

correlated with those recently described from an area 

5 miles to the southwest (Leonard, 1962). A narrow 

belt of the old metavolcanics is preserved along Beaver 

Creek as the footwall of the syenite complex, and as a 

septum separating the syenite complex from younger 

intrusive rocks to the west. The exposed metavolcanics comprise dark-green amphibole-rich hornfels 

and amphibolite; green, fine-grained, biotite-rich hornfels; gray to green, very fine grained, micaceous 

"metasiltstone"; and some mylonitized felsic rocks. 

White quartzite on Ramey Ridge and near the county 

line is interpreted as a remnant of the hanging wall of 

the syenite complex.

The syenite complex is divisible into a highly contaminated outer part and a less contaminated inner 

part. In each part, the main components are syenite 

and quartz syenite of two distinct facies, A and B.
E94 GEOLOGICAL SURVEY RESEARCH 1962

2 MILES

FIGURE 204.1. Geologic map of part of Big Creek quadrangle, 

Idaho, showing Ramey Ridge Complex. Map units, from 

oldest to youngest: ms, metasedimentary rocks; mv, metavolcanic rocks; m, metasedimentary and metavolcanic rocks; 

q, white quartzite; Ramey Ridge Complex, Paleozoic(?) 

l^rsa, syenite highly contaminated with amphibolite, and 

I^rs, syenite slightly contaminated with amphibolite; PC, 

Casto Volcanics, Permian(?); Ki, Idaho batholith rocks, 

Cretaceous; Q, surficial deposits, Quaternary. Conventional 

symbols denote contacts, bedding, foliation, and lineation. 

Faults, Tertiary dikes, and prospects are omitted. The local 

border of the Idaho batholith is just west of the map edge.

The letters are used provisionally because geographic, 

mineralogic, color, and fabric designations are unsatisfactory, owing to the range of properties shown by 

the two facies. The distinctions are that the mafic 

minerals in facies A are usually in discrete small 

clots, while those in B seldom are; and that where 

one facies cuts the other, A is older and slightly more

mafic, is generally darker and coarser, and generally (?) has a higher ratio of ferromagnesian minerals to magnetite.

The outer part of the syenite complex (fig. 204.1) 

is a mixture of syenite, amphibolite, and hybrid rocks. 

Where exposures are good, five major rock types are 

distinguishable. Arranged according to decreasing 

content of dark minerals and roughly in order of 

emplacement, these are: amphibolite and mafic hornfels, mafic syenite, syenite pegmatite with coarse mafic 

aggregates, syenite and quartz syenite of facies A, 

and syenite and quartz syenite of facies B. Each 

type of rock locally grades into one or more of its 

fellows, though elsewhere contacts are sharp and a 

definite sequence of emplacement is observable. Darkgreen to dark-gray fine-grained amphibolite and mafic 

hornfels are the oldest components of the complex. 

They occur as angular inclusions, a few inches to a 

few feet across, in a matrix of syenite or mafic syenite. 

In places, the amphibolite is slightly migmatitic, but 

most of it is outwardly little modified. The inclusions have evidently been derived from the old metavolcanic wallrocks. Typically, the inclusions are found 

in white to gray medium- to coarse-grained indistinctly foliated syenite (facies A) in which the mafics 

are not quite uniformly distributed. In some varieties, 

the planar structure given by small mafic clusters is 

emphasized by subparallel tablets of alkali feldspar. 

The rather dark color of much of this syenite is due 

more to the gray feldspars than to the quantity of 

mafic minerals, which seldom exceeds 15 percent. Locally, syenite of facies A grades into syenite of facies 

B, a pale-gray, pinkish, or yellowish rock in which the 

mafics are rather uniformly distributed. In many 

places, syenite B cuts sharply across an igneous breccia 

of generally darker syenite A and amphibolite, forming dikes a few millimeters to several centimeters 

wide. The proportions of amphibolite and syenite are 

difficult to estimate, owing to poor exposure and inherent variation; probably amphibolite and other dark 

rocks form more than 10 but less than 40 percent of 

the mapped unit.

The inner part of the syenite complex (fig. 204.1) 

is distinguished from the outer part by relative scarcity 

of amphibolite or hornfels inclusions. Many outcrops 

show no dark inclusions; others show only a few small 

fragments in an area of many square feet. Sporadic 

layers of mafic syenite are present, but syenite pegmatite with knots of hornblende, magnetite, and 

sphene has not been found. Gray syenite of facies A 

predominates. Locally, this is cut by narrow dikes 

of light syenite of facies B. The only significant 

difference in these two main facies of syenite in the
ARTICLE 204 E95

inner and outer units (fig. 204.1) is that hornblende 

as a varietal mafic is effectively restricted to the outer 

unit, the one in which amphibolite inclusions are relatively abundant. In the outer unit, hornblende is 

present locally in syenites of both facies A and B, 

whereas hornblende is absent from nearly all syenites 

of both facies in the inner unit. This distribution 

strongly suggests that hornblende in the outer unit 

is largely or wholly xenocrystic, or is derived from 

the reconstitution of mafic xenoliths.

In mineral composition, the two main facies (A 

and B) are simple. Essential minerals are alkali feldspar, magnetite, and biotite or hornblende. Mafic minerals seldom exceed 15 percent. Quartz is a varietal 

mineral: a few syenites have none, many have 1 to 3 

percent, and the quartz syenites have about 10 percent. 

Feldspathoids have not been found. Most of the 

rocks conform to the definition of part of the classic 

nordmarkites, and the photomicrograph figured by 

Johannsen (1937, p. 6) would be illustrative of many 

specimens from the Ramey Ridge Complex. Microantiperthite, generally carlsbad twinned, is the predominant feldspar. However, some microperthite is 

present locally. The determination of the host feldspar is so difficult in many of the intergrowths that 

the term "perthite" will be used in this report. Patchtwinned albite is the principal feldspar in a few specimens. Cryptoperthite, anorthoclase, and other alkali 

feldspars may be present; they have not yet been 

studied. Biotite is more common than hornblende, 

though both may be present. The biotite is dark 

green; the hornblende is blue green. Magnetite is 

commonly clustered with the ferromagnesian mineral. 

The fine-grained, more felsic rocks (generally quartz 

syenites) commonly have 2 or 3 percent magnetite and 

less biotite or hornblende. Apatite and zircon are 

ubiquitous. The zircons range from minute to relatively large (2 or 3 mm). Some of the large ones are 

pink, even in thin section, and are commonly the 

coarsest mineral present, except for the perthite. 

Sphene, either discrete or as rims on magnetite, is 

present in nearly half of the 36 specimens examined. 

Small granules of fluorite are found in more than 

half the specimens. Allanite is present in a few. 

Chevkinite, a complex Ce-Ti silicate, has been positively identified in one sample and is suspected in 

others. Some chevkinite is rimmed by very fine 

grained allanite(?), according to X-ray identification 

by Fred A. Hildebrand. Pyrite is rare, but pseudomorphs of reddish goethite after accessory pyrite are 

rather common, and many syenites show crack filling 

and local replacement by iron oxides, hisingerite (?), 

and nontronite( ?). Two-thirds of the specimens show

slight alteration to epidote, pumpellyite, chlorite, and 

leucoxene.

The fabric of facies A and B ranges from undeformed to cataclastic and granoblastic. Undeformed 

varieties (represented by less than a third of the 

specimens) have medium to coarse feldspars, generally 

carlsbad twinned, either crudely equigranular or distinctly elongate. The mafic minerals are usually much 

finer, occurring as patches between the feldspars. 

Quartz, if present, is also interstitial to the feldspars. 

Contacts between adjacent perthites are straight sided, 

interlocking, or implicated. Alternation of equant 

and elongate feldspars gives a thin-layered appearance 

to some otherwise homogeneous, undeformed syenites. 

Incipient deformation of the primary fabric has caused 

peripheral granulation of a few perthites, yielding a 

little mortar of sodic plagioclase. With increasing 

deformation, more plagioclase mortar has formed. 

Advanced deformation has caused granulation of part 

or all of the body of the perthites, yielding a mixture 

of fine sodic plagioclase and microcline but preserving a few perthite porphyroclasts. Some syenite has 

new biotite growing athwart feldspar contacts, or narrow veinlets of very fine hornblende cutting across 

primary planar structures. The distribution of cataclastic and granoblastic fabrics has no evident relation to contacts or primary flowage features. The 

changes in fabric and proportions of feldspars are very 

similar to those in some rocks from the Adirondack 

Mountains described by Buddington (1939, p. 251- 

303). Changes in the syenites in Idaho are likewise 

taken as evidence of dynamothermal metamorphism.

Chemically analyzed rocks of the two main syenite 

facies have 62.6 to 67.3 percent SiC>2. Data for a 

syenite of intermediate SiC>2 content are given in the 

table below. This syenite belongs to facies B. Judged 

from many thin sections, the bulk composition of the 

complex is slightly less silicic.

The presence of 0.3 percent chevkinite was unexpected. This strongly absorbing, dark-brown to opaque 

accessory mineral was identified by X-ray by Fred A. 

Hildebrand (film no. D-6260). Young and Powers 

(1960) described chevkinite as minute crystals and 

primary inclusions in rhyolitic ash from the Cenozoic 

of Idaho and other Western States. Dwight L. 

Schmidt (oral communication, 1962) states that he 

has not found or heard of chevkinite in any of the 

major placer deposits of central and south-central 

Idaho. So far as I know, the mineral has not previously been identified in plutonic rocks of the region. 

Chevkinite accounts for the Ce and some of the Ti 

in the syenites. The source of the Nb has not yet 

been determined. The presence of Ce, other rare-
E96 GEOLOGICAL SURVEY RESEARCH 1962

Chemical and modal data for a syenite from the Ramey Ridge

Complex

[Field No. L/54/1051; lab. No. B-232. Collected from talus, west side of Hand Creek 

trail, 2,200 feet northeast of Joe Creek (not shown on map) or about 1 mile upstream 

frommouth of Hand Creek, northeast ninth, Big Creek quadrangle, Idaho County, 

Idaho.

Fine-grained, subporphj ritie; faint planar structure given by crude laminae of 

elongate subparallel feldspars alternating with those of stubby, randomly oriented 

feldspars. Cataclasis moderate: peripheral granulation common; discontinuous 

lenticles of crushed material sparse. Represents the youngest faeies of rock forming the syenite complex]

CHEMICAL ANALYSIS 

[Lois Trumbull, analyst, 1955]

Si02 ... ... ...

TiO2 ... ... ... ....

AljOj. ...............

FeaOa-- ~

FeO

MgO.._ _ .

CaO._. .......... ....

NajO--..... ......

KjO.... ._ ......

HjO+_ 

H20-_-_ 

PjOj...... ...........

MnO. ...............

Weight 

percent 

............ 64.95

_- . .46

 . .. 16.52

.- - 3.37

 1.60

-. .- .21

 .83

... . 6.30

..... . 5.02

. .- .14

04

 .-. .09

. - -- .19

BaO_-_- .... ... .... 

C02 - --

Cl-.... ...... ... ... ... .

F__... ...................

S-.--. ..............

Total.......... ........

Total (corrected) ___ .

Spgr... . 

Weight 

percent 

........ 0.03

........ .02

..... .03

........ .16

........ .00

........ 99.96

........ .08

....-.- 99.88

........ 12.71

MODE 2

Perthite... ----- ___

Chevkinite- _____

Volume 

percent 

............ 4.9

. _- 3 1.6

............ 76.1

....... ___ - 411.5

1.1

 1.0

 2.7

 .3

Fluorite. __ ..... ___ .

Total..----.-.........

Volume 

percent 

........ 0.2

2

........ .2

........ .2

....... Tr.

--- 100.0

CIPW NORM

Q

ab.... ... ... ... ....

an ...... .........

hi-.... ...............

Weight 

percent 

............ 8.05

 ... 29.67

 53.06

. _ . 2.12

............ .10

._ _- . .52

. ______ . 4.43

11 .-- ..............

fr _-_ - .-- . .

Total. ...... ...........

Weight 

percent 

........ 0.30

 - .88

._ .20

..... .18

....... .05

.... . 99.56

SPECTROCHEMICAL ANALYSIS (SEMIQUANTITATIVE) 

[Paul R. Barnett, analyst, 1955]

Ba.... ................. O.OX

Be-........ . . .OOOXCe_.......... - .OX

Or-......................... .OOOXCu......... ............ . .OOOX

Ga.......................... .OOX

La.............. .OXNb ....................... O.OXNd- ..................... .OXSc-. .OOOX

Sr .- - ---. .OOX+

Y........................... .OOX+

Yb.. . - - .OOOX+

ZT........................... .OXLooked for but not detected: Ag, As, Au, B, Bi, Cd, Co, Dy, Er, Qd, Qe, Hf, Hg, 

In, Ir, Li, Mo, Ni, Os, Pb, Pd, Pt, Re, Rh, Ru, Sb, Sm, Sn, Ta, Te, Th, Tl, U, V, 

W, Zn.

1 Determined by M. Balazs.

2 Average of 3 thin sections; 4,441 points counted.

3 May include some cryptoperthite and anorthoclase.

4 Sodic oligoclase, recrystallized. Some albite, only slightly perthitic, is also 

present.

earth elements, and Nb in slightly greater amounts 

in the syenite complex than in other rocks of the ]8ig 

Creek area may offer a clue to concentrations of such 

elements in the bedrock or placers of the area.

The gross structure of the syenite complex is still 

to be determined. Geologic mapping has not been extended north and east of the area shown on figure 

204.1, so that an interpretation based on only part of 

the complex is given here. For this part, the key 

structure is the Hand Creek syncline, whose nose lies 

between Boulder and Hand Creeks. Here, the remnant of old metavolcanics outlines the southwest limb 

and part of the nose of the southeast-plunging syncline whose northeast limb is outlined by inclusions 

and foliation in the syenite complex. This sector of 

the complex is grossly concordant with the old metavolcanics, though syenite dikes cross the contact locally. 

The syncline is flanked on the southwest by an anticline of metavolcanics, but the syenite complex cuts 

across the limb and nose of this structure along an 

irregular, steeply-dipping surface that is exposed in 

cliffs along Beaver Creek. Whether the patch of 

white quartzite on Ramey Ridge originally lay in the 

Hand Creek syncline or in a subparallel syncline 

cannot be determined, owing to the scarcity of outcrops in the critical area.

The highly contaminated outer unit of the syenite 

complex is separated from the cylindrical (?), slightly 

contaminated central mass along a surface or narrow 

zone that, where well exposed, is highly irregular. 

In places, the southwest contact of the inner unit is 

crudely conformable to the structure of the old country rock, but mostly it is markedly discordant. On 

figure 204.1, this discordance is shown only by the 

perpendicular relation between the internal structure 

of the inner unit and the contact between the two 

units. On a topographic base, the discordant relations are more obvious, and the field evidence indicates that some syenites of the inner unit intrude 

those of the outer unit.

The northwest contact of the inner unit cannot be 

accurately determined, owing to disintegration of rocks 

in the adjoining units.

Some linear magmatic flow in the complex is suggested by the subparallel, elongate perthites in undeforined syenite, as well as by scarce elliptical clots 

of amphibolite.

Superimposed on the earlier tectonic and flow structures are patches and zones of mylonite. The largest 

of these strikes north-northeast, roughly parallel to 

Hand Creek and 0.3-0.7 mile east of the creek. The 

mylonite zone is a few hundred to a thousand feet 

wide. Dikes of Tertiary age, mostly granophyric,
ARTICLE 204 E97

have been intruded here and there along the zone but 

are not restricted to it. Elsewhere in the Big Creek 

quadrangle, a few mylonite zones cut the Casto Volcanics and the rocks of the Idaho batholith. Therefore, it is reasonable to suppose that some of the 

mylonite zones in the syenite complex are younger 

than the batholith. However, the presence of mortar 

fabric and granoblastic patches in syenite pebbles 

taken from conglomerate in hornfelsed, only slightly 

sheared, prebatholith volcanics (Casto) clearly indicates that considerable deformation of the syenite occurred before the Idaho batholith was emplaced. Since 

some specimens of bedrock from the syenite complex 

show mylonite veins cutting mortar gneiss and granoblastic syenite, parts of the complex must have been 

deformed at least twice. It is reasonable to interpret 

one episode of deformation as prebatholith and one as 

postbatholith, though in any one specimen or outcrop it would be rash indeed to attempt assignment of 

a mylonite to a particular episode of deformation. If 

the syenite complex was appreciably deformed during 

emplacement of the Idaho batholith, evidence of that 

effect has not yet been recognized.

In spite of some uncertainty in interpreting the 

gross structure and evolution of the syenite, these 

points are established: (1) the syenite complex was 

emplaced after major deformation of the old metavolcanics had developed a system of rather tight folds 

that now plunge gently to moderately southeast; (2) 

syenite intrudes the white quartzite exposed on Ramey 

Ridge; (3) the white quartzite is unconformably overlain by Casto Volcanics; (4) aplites and alaskites of 

the Idaho batholith intrude the syenite complex and 

the Casto Volcanics; (5) only a mile southwest of 

the syenite complex, rounded pebbles and cobbles of 

syenite are found in greenstone conglomerate referred 

to the Casto Volcanics. Until syenite is found cutting across a tight fold in the white quartzite, the 

relation between syenite and the principal folding of 

the quartzite remains indeterminate. However, all the 

evidence from other areas in the quadrangle indicates 

that old metavolcanics and overlying white quartzite 

have undergone the same major episodes of folding.

Thus it is reasonable to assume that the syenite is 

probably younger than the major folding of the 

white quartzite. The white quartzite may be late 

Precambrian, Cambrian, or Ordovician (Leonard, 

1962). The Geological Survey's official age assignment of the Casto is Permian(?). So long as the 

question mark, instead of the name Permian, is borne 

in mind, the assignment is appropriate enough for 

this nonfossiliferous formation. An age of Triassic 

or even Jurassic is as reasonable as Permian, since 

the volcanics and associated sedimentary rocks in all 

three systems are rather similar in the Cordilleran 

region. (See, for example, Bostwick and Koch, 1962; 

Nolf and Taubeneck, 1962).

The broadest time span into which the syenite 

complex can be fitted is therefore late Precambrian 

to perhaps Cretaceous; the narrowest is perhaps Ordovician to Permian. Provisionally, the complex is regarded as Paleozoic. A closer approximation of its 

age awaits the results of isotopic studies of the component rocks and minerals.

REFERENCES

Bostwick, D. A., and Koch, G. S., Jr., 1962, Permian and

Triassic rocks of northeastern Oregon: Geol. Soc. America

Bull., v. 73, p. 419-422. 

Buddington, A. F., 1939, Adirondack igneous rocks and their

metamorphism: Geol. Soc. America Mem. 7, 354 p. 

Johannsen, Albert, 1937, A descriptive petrography of the

igneous rocks, v. 3, The intermediate rocks: Chicago,

Univ. Chicago Press, 360 p. 

Larsen, E. S., Jr., Gottfried, David, Jaffe, H. W., and Waring,

C. L., 1958, Lead-alpha ages of the Mesozoic batholiths of

western North America: U.S. Geol. Survey Bull. 1070-B,

27 p. 

Leonard, B. F., 1962, Old metavolcanic rocks of the Big Creek

area, central Idaho: Art. 5 in U.S. Geol. Survey Prof.

Paper 450-B, p. B11-B15. 

Nolf, Bruce, and Taubeneck, W. H., 1962, Permo-Triassic

stratigraphic revisions, Wallowa Mountains, northeastern

Oregon [abs.] : Geol. Soc. America Rocky Mountain Section, Provo Mtg. Program, p. 30-31. 

Shenon, P. J., and Ross, C. P., 1936, Geology and ore deposits

near Edwardsburg and Thunder Mountain, Idaho: Idaho

Bur. Mines and Geology Pamph. 44, 45 p. 

Young, E. J., and Powers, H. A., 1960, Chevkinite in volcanic

ash: Am. Mineralogist, v. 45, p. 875-881.
E98 GEOLOGICAL SURVEY RESEARCH 1962

205. TRONDHJEMITE IN THE RIGGINS QUADRANGLE, WESTERN IDAHO 

By WARREN HAMILTON, Denver, Colo.

Trondhjemite is widespread in the 30-minute Riggins quadrangle of west-central Idaho, forming two 

large plutons and several smaller ones as well as 

abundant dikes and sheets in gneissic complexes. The 

name "trondhjemite" was introduced by Goldschmidt 

(1916, p. 77) for light-colored granitic rocks containing essential oligoclase and quartz but having little or 

no K-feldspar, and having biotite as the dominant 

mafic mineral. Such rocks are strikingly different 

from common quartz diorite on the one hand and 

granite on the other, and the distinguishing name 

"trondhjemite" is desirable to emphasize the unusual 

composition.

The interior of the Idaho batholith, which is probably of middle Cretaceous age (Jaffe and others, 

1959), consists largely of massive uniform quartz monzonite and granodiorite. West of this is a gneissic 

and schistose border zone about 30 miles wide within 

which the proportion of intercalated granitic material 

and the grade of regional metamorphism decrease 

westward (Hamilton, 1960). The trondhjemite occurs 

largely in the western half of the gneissic part of 

this border zone.

A pluton of trondhjemite crops out beneath Columbia River Basalt and surficial deposits in the southwest part of the Riggins quadrangle and has an area 

probably greater than 50 square miles. The trondhjemite is white and is flaked by plane-parallel aggregates 

of large biotite crystals which enclose granular green 

epidote. Texture is xenomorphic granular and faintly 

gneissic. Plagioclase is anhedral calcic oligoclase that 

typically shows slight normal-oscillatory zoning; lamellar acline or pericline twinning is common and 

many grains are finely grid twinned. Orthoclase is 

lacking or rare. Quartz anhedra are elongate parallel 

to the foliation. Biotite forms 5 to 10 percent of the 

rock and where most abundant is accompanied by a 

little hornblende. A little muscovite in the more silicic 

phases is mostly in large extremely irregular crystals 

of replacement origin.

Epidote is present in practically all of this trondhjemite pluton in amounts up to 2 percent. The epidote 

is green and hence conspicuous in hand specimen, but 

it is only weakly pleochroic in thin section. Much 

of it is in euhedral crystals within or against aggregates of biotite but some is extremely poikilitic, enclosing chiefly biotite. Both types are of replacement 

origin: single thin sections contain grains of epidote 

showing all degrees of replacement of biotite, from

highly irregular and poikilitic epidote to compact euhedral crystals. Oscillatory zoning of the plagioclase 

shows that the rock crystallized from magma, and 

the epidote is regarded as a late-magniatic reaction 

product.

The pluton was intruded semiconcordantly into older 

gneisses to the east and north, and produced contact 

miffmatites ~ in a zone as wide as two-thirds of a mile.

Along Little Salmon River to the north of the pluton 

varied quartz dioritic gneisses are cut intricately by 

large and small dikes of relatively dark trondhjemite 

that is gneissic parallel to the walls of the dikes. These 

dikes are cut in turn by massive, uniform, lightcolored trondhjemite, and these are cut by sodic pegmatites. The various dikes bulk larger than does the 

host gneiss near the pluton. Locally phyllonitized 

quartz diorite is cut by uncrushed dikes of trondhejemite. Near Sixmile Creek east of the main trondhejemite pluton, broadly uniform mafic quartz diorite 

gneiss is cut by an intricate network of aplitic trondhjemite, partly as discrete dikes, both concordant and 

discordant, but largely as a pervasive matrix enclosing 

blocks of quartz diorite. Contacts are extremely irregular and crosscutting in detail (fig. 205.1A).

Smaller plutons stocks and semiconcordant sheets 

 of trondhjemite similar to that of the plutoa described are present in the center of the quadrangle. 

One such pluton is broken by a west-directed thrust 

fault (Hamilton, 1960) and is sheared to phyllonite 

near the fault.

Trondhjemite gneiss forms abundant intercalations 

in granodioritic and quartz monzonitic gneisses in the 

southeastern part of the quadrangle. Some of these 

layered complexes are much sheared.

A long pluton of trondhjemite gneiss crops out over 

an area about 5 miles long and several miles wide 

within the north-central part of the Riggins quadrangle, and is exposed for an additional 10 miles of 

length north of the quadrangle. The northern part 

of the west contact dips eastward, and is a semiconcordant intrusive contact in which the trondhjemite 

lies above marble and metavolcanic rocks; the marble 

is contact metamorphosed and metasomatized. The 

metamorphic rocks are the footwall of the same thrust 

fault that displaces the small pluton of trondhjemite 

noted above. The southern part of the west contact 

of the trondhjemite gneiss dips westward beneath 

hanging-wall metamorphic rocks. The gneiss is interpreted to have been intruded as a thick wedge along
ARTICLE 205 E99

FIGUBE 205.1. A, Trondhjemite-quartz diorite migmatite. The dark gneissio quartz diorite is injected intricately by white 

trondhjemite, and both are cut by a dike of sodic pegmatite. B, Shear-folded trondhjemite layer in mafic gneiss. Both 

trondhjemite and host gneiss are foliated and lineated cataclastically. Scale is 6 inches long.
E100

the thrust fault and to have been folded subsequently. 

The southern and eastern contact of the trondhjemite 

gneiss pluton is a zone several miles wide of trondhjemite intercalated in varied gneisses. Tight to open t 

fold shapes defined by foliation in the trondhjemite 

and by distribution of gneiss layers in the contact 

zone appear to be folds imposed on the trondhjemite 

rather than products of its intrusion into a pre-existing fold system, or folds produced by quasimagmatic 

flow. Figure 205.15 shows shear-folded gneiss in the 

eastern border zone of the pluton.

The trondhjemite gneiss is streaked by discontinuous folia and trains of biotite, and has a distinctive 

sugary appearance. Plagioclase is near An2o in most 

specimens and is as sodic as An1B in some; zoning 

is evenly gradational without discontinuities, and is 

normal in some places but reverse in others. Most specimens carry a few percent of untwinned potassic 

feldspar. Mafic minerals total 5 to 12 percent, locally 

more. Intergrown with the biotite of most specimens 

is conspicuous green epidote and generally a little 

hornblende, but many specimens have muscovite instead accompanying the biotite. Tiny crystals of pink 

garnet are present locally. Quartz is abundant and 

in much of the pluton is markedly elongate, with the 

long dimension and the c crystallographic direction 

subparallel to foliation. This orientation, the crystalloblastic fabric of all minerals including orthoclase, 

and the presence of poikilitic and reverse-zoned oligoclase indicate metamorphic recrystallization of the 

rock at high grade high enough to recrystallize orthoclase. The gneiss is monotonously uniform over 

large areas, and layered rock is uncommon within it.

The trondhjemite and allied rocks are highly distinctive mineralogically because of their low content 

of mafic minerals and their sodic plagioclase. They 

are comparably distinctive chemically, differing from 

common types of granitic rocks of similar SiOz content in all major oxides (fig. 205.2). Relative to granitic rocks of the normal types, in which the proportion 

of potassic feldspar increases and that of plagioclase 

decreases with increasing silica content, the trondhjemites contain 2 percent more A12OS , 1 to 2 percent 

more Na2O, and a little more CaO, but only about 

% as much iron and MgO, and ys as much K2O.

FIGUEE 205.2. Silica-variation diagram of chemical analyses 

of trondhjemites and allied dike rocks of Riggins quadrangle, 

shown by points. Curves show general trends of variations 

in rocks of the common series quartz diorite-granodioritequartz monzonite-granite in the Idaho, Sierra Nevada, and 

southern California batholiths.

20

15

,10

O

rt 

O

rt 

fc

60 65 70 75 

SiO2 , IN WEIGHT PERCENT

80
ARTICLE 206 E101

The trondhjemite and allied rocks of the Riggins 

quadrangle also have relatively high ratios of ferric 

to ferrous iron. This is expressed mineralogically by 

the almost ubiquitous epidote. Were there less ferric 

iron, then presumably the calcium which entered the 

epidote would instead have crystallized in plagioclase, and the plagioclase would be slightly more calcic 

than it is.

The trondhjemite of the Riggins quadrangle varies 

widely in occurrence. One trondhjemite pluton is 

broken by a thrust fault which is intruded by another 

pluton, and the later of these plutons is itself highly 

deformed and metamorphosed. The large pluton of 

the southwest part of the quadrangle is unmetamorphosed, and dikes from it cut phyllonite perhaps 

formed at the time of thrusting. Trondhjemite east 

of this pluton is sheared cataclastically and reconstituted at low grade. As all of the trondhjemite is in 

the western border zone of the Idaho batholith, within 

which the proportion of granitoid material decreases 

systematically westward, it is considered to be related genetically to the batholith and to have formed,

despite its complex relations, within a part of Cretaceous time before the emplacement of the massive 

interior of the batholith. Thrusting, metamorphism, 

and intrusion within the border zone are viewed as 

related phases of continuing activity.

Although the metamorphism is of "regional" type, 

the heat for it was probably partly introduced by 

granitic magmas and partly generated by extreme deformation of the rocks, and was largely unrelated to 

geothermal gradients controlled by depth of burial.

REFERENCES

Goldschmidt, M., 1916, Ubersicht der Eruptivgesteine Im 

kaledonischen Gebirge zwischen Stavanger und Trondhjem: Kristiania Videnskaps. Skr., I. Math.-Naturv. 

Klasse, no. 2, 140 p.

Hamilton, Warren, 1960, Metamorphism and thrust faulting 

in the Riggins quadrangle, Idaho: Art. 103 in U.S. Geol. 

Survey Prof. Paper 400-B, p. B230-B231.

Jaffe, H. W., Gottfried, David, Waring, C. L., and Worthing, 

H. W., 1959, Lead-alpha age determinations of accessory 

minerals of igneous rocks (1953-1957) : U.S. Geol. Survey Bull. 1097-B, p. 65-148.

206. X-RAY DETERMINATIVE CURVE FOR OLIVINES OF COMPOSITION

ALPINE-TYPE PERIDOTITES

FROM STRATIFORM AND

By PRESTON E. HOTZ and EVERETT D. JACKSON, Menlo Park, Calif.

The X-ray determinative method developed by Jackson (1960), using five natural olivines in the range 

Fogo-go from the Peridotite member of the Stillwater 

Complex, Montana, was applied to five chemically analyzed olivines from alpine- type peridotites from 

northwestern California and southwestern Oregon. 

The alpine-type olivines ranged in composition from 

Fooo-95 and appeared to fall on the extrapolation of 

the curve for stratiform types. The data for olivines 

from both provinces was therefore combined to construct a curve for the range Fo80_95.

Following the procedures used by Jackson (1960),

olivine concentrates of five specimens of enstatite peridotite and dunite from alpine-type peridotites at four 

localities in northwestern California and southwestern 

Oregon were prepared and split to obtain samples 

for chemical analysis and X-ray measurements. The 

Fo content was calculated from the chemical analyses 

as the atomic ratio of Mg to total octahedral cations. 

X-ray measurements were made of the difference in 

degrees 20 (CuK«i) between the olivine (062J peak, 

and the lithium fluoride (220) peak for subsamples 

of the analyzed olivines, again following Jackson's 

method.
E102 GEOLOGICAL SURVEY RESEARCH 1962

The pertinent data for both the stratiform and 

alpine-type olivines are summarized in the table below.

Fo content and X-ray measurements of olivine samples

Field No. Foi

Mean of 12

readings 2 20062

(olivine) -26220

(LiF) (degrees)

Standard

deviation

of 12 readings 

(degrees)

Stillwater Complex 3

52MV-9-______ _

55MV-29____..____

55BE-37 4 _ _______

55MV-26_______

55BE-44 «______. ___________

55MV-40- _______________

55MV-49___-._____ _________

80. 5

84. 0

84. 7

84. 9

85. 8

85. 8

89. 7

3. 019

2. 958

2. 945

2. 947

2. 928

2. 926

2. 857

0. 001 £

. 001 £

. 00 K

. 002£

. 0012

. 001C

. 0014

California and Oregon

H45-57_______.____________

NM1-61_________. _________

D9-61___ __________________

KN2-61 «.______

C7-58._ ________-__

NM4-61_______ _

90. 3

90. 9

91. 2

91. 2

91. 4

94. 2

2. 842

2. 834

2. 830

2. 830

2. 833

2. 769

. 0017

. 0021

. 0014

. 0014

. 0021

. 0032

1 Atomic ratio of Mg to total octahedral cations.

2 CuKai.

s From Jackson (1960, table 197.1). 

* Hidden duplicate of 55MV-26. 

«Hidden duplicate of 55MV-40. 

» Hidden duplicate of D9-61.

A regression analysis of the data in the table was 

made, considering each of the 13 measured sets to be 

individual determinations, and taking the Fo value 

computed from the chemical analysis as the independent variable. The equation for the straight-line 

curve (fig. 206.1) is: 20062 (olivine) -20220 (LiF) = 

4.4722-0.018015 Fo. The regression constant for the 

combined data differs less than 0.4 percent and the 

regression coefficient differs less than 1.0 percent from 

the equation computed for the Stillwater olivines 

alone (Jackson, 1960, p. B433). The 95 percent confidence intervals for the combined regression equation are ±0.008° 20 at 88 mol percent Fo and 

±0.009° 20 at 80 and 95 mol percent Fo. These 

limits correspond to an uncertainty of approximately 

±0.44 mol percent Fo at Fo88 , and ±0.49 mol percent 

Fo at Fo80 and Fo95 .

We conclude that the relation between the (062) 

spacing and the Fo content is the same for olivines 

from both Stillwater and California-Oregon peridotites, and that the determinative curve (fig. 206.1) is

3.02 -

3.00

O Olivines from the Stillwater Complex /

Olivines from northwestern California/// 

and southwestern Oregon '

At Fooo the 95 percent confidence

interval is ±0.008

2.80 -

2.78

2.76

95 94 93 92 91 90 89 88 87 86 85 84 83 82 81 80

Fo (ATOMIC RATIO Mg/TOTAL OCTAHEDRAL CATIONS)

FIGURE 200.1. X-ray determinative curve for stratiform and 

alpine-type olivines of composition Foso-95-

probably valid for olivines from both stratiform and 

alpine-type environments. Preliminary data for olivines from the Hawaiian volcanic province (K. J. 

Murata, oral communication, 1962) indicate, however, 

that use of the determinative curve leads to errors of 

greater than 2 mol percent Fo, suggesting that the 

curve is not applicable to olivines from such volcanic 

rocks.

REFERENCE

Jackson, E; D., 1960, X-ray determinative curve for natural 

olivine of composition Fo8(M)0 : Art. 197 in U.S. Geol. 

Survey Prof. Paper 400-B, p. B432-B434.
ARTICLE 207

207. FORMATION OF "SALT CUPS" NEAR SAN PEDRO DE ATACAMA, CHILE 

By ROBERT J. DINGMAN, Washington, D.C.

E103

Work done in cooperation icith the Institute de Geologicas Investigaciones, Santiago Chile, and under the auspices of the

Agency for International Development

During the geologic mapping of the Tulor quadrangle near San Pedro de Atacama, Chile (fig. 207.1), 

fragile pure-white structures having the approximate 

size and shape of an inverted tea cup (fig. 207.2) 

were observed on outcrops of salt beds in the San 

Pedro Formation. The occurrence of these "salt cups," 

as the author proposes to call them, is restricted to 

smooth horizontal surfaces in the Cerros de la Sal, 

in the southeast corner of the Tulor quadrangle. The 

cups are composed of almost pure sodium chloride as 

are the underlying salt beds of the San Pedro Formation, which are 97+percent NaCl. In horizontal section the salt cups are almost perfectly round and are 

2 to 4 inches in diameter at the land surface. In vertical section they are smoothly rounded domes of 

which the ratio of the base to the altitude is approxi20'

21°

22'

23'

24'

25'

26'

71° 70 c 69° 68° 67 C

TARAPACA 

PROVINCE

c

BOLIVIA

TULOR i Volcan 

QUADRANGLE_ Licancabuiv I A^_..-^%

Pedro

Solar

Atacama )':";:':':/., / '::' ' '. ''

ANTOFAGASTA

PROVINCE

ATACAMA

PROVINCE

/.f.-'. : >: (-, i de Atacama.1

Toconao j 

i

I I 

Tilomonte S

< ;

w V

-!__X

Pular

ARGENTINA

100 200 KILOMETERS

FIGURE 207.1. Index map of part of northern Chile 

showing location of the Tulor quadrangle.

mately 2:1. The thickness of the wall is ys to %6 inch 

at the base and decreases upward to approximately 

y16 inch at the top of the cup. The salt cups may be 

open or completely closed at the top. They are open 

at the top during the youthful stage, closed in the 

mature stage (fig. 207.3), and open again during the 

olcl-age stage (fig. 207.4). Beneath the salt cup a 

round tubular passage extends downward into the salt. 

The irregularity of the passage prevents determination of its depth, although one tube was probed to a 

depth of 32 inches.

The cups are formed as a result of solution and redeposition of salt from the underlying salt beds. The 

tubular passages under the cups probably extend 

downward into either an intersecting system of fractures or a cave or tunnel that forms part of the internal drainage system of the salt domes. Under 

changing atmospheric conditions, the flow of air 

through such a connection would be similar to that 

through a sucking and blowing well. The upward 

moving air during the blowing stage would be moist 

and would carry a certain amount of salt in solution. 

Deposition by crystallization would occur at the land 

surface owing to the lower vapor pressure of the dry 

outside air. Because fog or periods of high humidity 

are, very uncommon in this area, the air drawn into

FIGTJKE 207.2. Salt cup in late stage of construction 

almost closed.

Vent is
E104 GEOLOGICAL SURVEY RESEARCH 1962

the cups is dry and so does not take the salt back into 

solution. The gradual sealing of the vent at the top

FIGURE 207.3. Mature salt cup. Vent is completely closed.

f ~ - - » ; - -

of a cup indicates that the cup was formed by a very 

gentle current of air. Actual movement of air out 

of or into the cups could not be observed because 

the wind blows continuously in this area.

As shown in figure 207.5, hairlike growths of crystals 

extend inward from the inner surface of the cups. 

Except for these growths, the inner surface of the 

cups is very smooth, whereas the outer surface is 

relatively rough. Because the wall of a cup is very 

light in weight and highly porous, air movement 

through the wall probably occurs even after the vent 

has been sealed. If the wall were not porous, a 

closed cup would rupture during changes of air pressure.

The cups are extremely delicate structures and probably are destroyed by raindrops. In the San Pedro 

area, however, the average rainfall is less than 4 

inches per year and several years may pass without 

any precipitation. The occasional moderate to heavy 

rains probably provide the moisture that eventually 

results in the formation of new cups. Like desert 

flowers, it is likely that salt cups "bloom" after a 

rain.

FIGURE 207.4. Salt cups in the old-age stage. Erosion results 

from abrasion by windblown sand or from solution by raindrops. The large size of the mound indicates that many generations of salt cups have formed at this site. FIGURE 207.5. Salt cup broken open to show internal structure.
ARTICLE 208 El05 

GEOCHRONOLOGY

208. LEAD-ALPHA AGES OF ZIRCON IN QUARTZ MONZONITE PORPHYRY, THIEL MOUNTAINS, ANTARCTICA 

A PRELIMINARY REPORT

By A. B. FORD, H. A. HTJBBARD, and T. W. STERN, Washington, D.C.

Work supported ~by the National Science Foundation

The Thiel Mountains in the interior of West Antarctica (fig. 208.1) consist mainly of cordierite-bearing hypersthene-quartz monzonite porphyry that has 

been intruded by plutons of coarse-grained biotite 

granite (Ford and Aaron, 1962). The porphyry is one 

of the oldest rock units in the area, but its geologic

age is unknown because of the absence of fossiliferous 

rocks.

Little is known regarding the geologic age of plutonic rocks in West Antarctica. Quartz monzonite 

basement rocks in the Horlick Mountains, approximately 150 miles west of the Thiel Mountains, are

93°W. 92 °W 90°W. 89°W.

NSample 2

Aaron Glacier

-Sample 1

McMurdo Sound I (f> 

Ross

90 °E.

5 MILES

85° 

GO'S.

85° 

15'S

FIGURE 208.1. Index and geologic sketch maps, Thiel Mountains, Antarctica. Hypersthene-quartz monzonite porphyry is

shown by small "v"s, younger intrusive granitic rocks by " + "s.
E106 GEOLOGICAL SURVEY RESEARCH 1962

overlain unconformably by fossiliferous Lower Devonian to Permian clastic rocks (Long, 1962). Granitic intrusions that may be Tertiary in age occur in 

the southern part of the Palmer Peninsula (Knowles. 

1945), approximately 800 miles northeast of the Thiel 

Mountains. Granitic rocks occur in many isolated 

nunataks between the Palmer Peninsula and the 

Thiel Mountains, but their age is unknown.

The occurrence of hypersthene-bearing granitic 

rocks in West Antarctica has not been previously reported, although charnockites and other hypersthenebearing rocks are common among crystalline rocks exposed along the coast of East Antarctica. Their presence in the Thiel Mountains therefore suggests an extension of the East Antarctic shield, presumably of 

Precambrian age, at least locally into West Antarctica.

The Thiel Mountains are part of the Transantarctic 

Mountains that cross the continent from the Ross 

Sea to the Weddell Sea. According to Hamilton 

(1962), the present Transantarctic Mountains mark 

the approximate site of a late Precambrian-Cambrian 

geosynclinal belt that underwent orogeny during Cambrian to Ordovician time. On the basis of a potassium-argon analysis of biotite from paragneiss of the 

McMurdo Sound area, Goldich and others (1958) dated 

an orogenic event at about 500 m.y. approximately 

the end of Cambrian time (Kulp, 1961).

Two bulk samples of hypersthene-quartz monzonite porphyry were collected in the Thiel Mountains 

by a U.S. Geological Survey party in January 1961. 

Lead-alpha ages of zircon concentrates from these 

samples (table 208.1) are the first radiometric ages 

to be determined for rocks from the interior of Antarctica. Additional large rock samples collected in 

the 1961-62 season will give age data for the younger

TABLE 208.1. Lead-alpha ages of zircon from hypersthene 

quartz monzonite porphyry, Thiel Mountains, Antarctica

Sample

1_

2: Less magnetic split.

More magnetic split.

Location

lat 85°10'S., 

long 90°30'W. 

Green Valley; 

lat 85°05'S., 

long 90°15'W. 

..-.-do ...._.._

Alpha 

counts 

per milligram 

per 

hour

100

124 

161

Pb 

(ppm)i

26.4

32. 9 

36. 7

Calculated 

age 2 (m.y.)

con i *7n

630 ±70 

530 ± 60

intrusive rocks as well as providing a check on the ages 

reported here.

Zircon from Green Valley (sample 2) was concentrated into a less magnetic fraction and a more 

magnetic fraction. Two determinations from sample 

2 give an average age of about 580 ±60 m.y. The 

average age of samples 1 and 2 is about 600±70 m.y., 

and suggests an orogenic event in late Precambrian or 

early Cambrian time.

The dated porphyry samples are quartz monzonitic 

in composition (table 208.2), but elsewhere the composition of the porphyry varies locally to granodiorite. 

The porphyry is medium to dark gray and is fresh 

to only slightly weathered. Iron oxides thinly coat 

some rock surfaces. A thin coating of alteration 

products on feldspars on exposed rock surfaces makes 

the highly porphyritic texture of the rock very obvious in some places. Unweathered surfaces appear 

almost even grained.

Microscopically, angular and blocky plagioclase 

phenocrysts show oscillatory zoned andesine centers 

that are commonly surrounded by sodic oligoclase 

rims. Many crystals show truncated zones which indicate that the crystals have been broken. Potassium-feldspar phenocrysts commonly have outer parts 

crowded with very small, rounded quartz inclusions. 

The distribution of the quartz inclusions suggests that 

late porphyroblastic growth has occurred around earlier phenocrysts. Quartz phenocrysts are generally 

rounded and deeply embayed. Hypersthene, the principal mafic mineral, is commonly surrounded by rims

TABLE 208.2. Modes of dated samples of hypersthene-quartz 

monzonite porphyry, Thiel Mountains, Antarctica l

[X, present in very small amounts; figure in parentheses is number of zircon grains 

in approximately 600 mm2 area]

1 Lead determinations by Harold Westley and Charles Annell, U.S. Geological 

Survey. All values are averages of duplicate determinations.

2 Constants: C=2,485; ft=1.56X10-«. Age rounded to nearest 10 m.-y.

Chlorite- _ _-.-________-_-__-

Rutile . .____....._.____.__._

Sericite ______ -..

i

23

31

32

2

3

7

X

X(76)

X

X

X

X

X

(

Light-colored 

fraction

34

28

32

1

1

3

X

X(74)

X

X

X

X

X

2

Dark-colored 

fraction

30

25

31

7

3

1

X

X(172)

X

X

X

X

Each mode is based on about 1,800 points counted by the Chayes method.
ARTICLE 208 E107

of brown and green biotite and locally of chlorite. 

Biotite is almost entirely associated with hypersthene, 

from which it appears to be an alteration product. 

Biotite in porphyry samples from near contacts with 

the younger granite plutons is porphyroblastic and 

highly sieve textured. Cordierite occurs both as large 

euhedral phenocrysts and as anhedral crystals associated with plagioclase and potassium feldspar. Cordierite is commonly altered to pale-green chlorite and 

sericite.

Most zircon occurs as separate grains in a dense 

xenomorphic-granular groundmass of quartz and potassium feldspar. Dark-brown pleochroic halos are 

common around very small zircon grains in biotite, 

and yellow pleochroic halos around zircon in cordierite. Small zircon grains are also associated with 

opaque minerals and with hypersthene. Most zircon 

is clear and colorless. The grains range in diameter 

from less than 0.01 to 0.45 mm and average about 

0.05 mm.

Quartz and feldspar phenocrysts range from 1 to 

5 mm in diameter. Groundmass crystals average 0.05 

mm. The groundmass is coarser near contacts with 

the younger intrusions. The porphyry appears homogeneous in most outcrops, as at sample 1 locality, 

but locally, as at sample 2 locality, poorly defined 

layered and lens-shaped bodies of very-dark-gray porphyry are present. The darker color of the layers 

and lenses is apparently due to the greater amount 

of dense groundmass. Groundmass constitutes 60 percent of the very-dark-gray porphyry in contrast to 45 

percent of the medium-gray porphyry.

Several features indicate that the porphyry has been 

recrystallized. These are (1) a hornfelsic or granulitic groundmass texture, (2) cordierite that appears 

to have replaced feldspar, (3) poikiloblastic rims on 

potassium-feldspar phenocrysts, (4) poikiloblastic biotite crystals, and (5) alteration rims of biotite and 

chlorite around hypersthene grains. Because these 

features are much more common in the vicinity of contacts with the biotite granite, they are probably effects of thermal metamorphism related to the intrusions. Therefore, zircon ages reported here may reflect 

the age of the younger intrusions rather than crystallization of the porphyry.

REFERENCES

Ford, A. B., and Aaron, J. M., 1962, Bedrock geology of the

Thiel Mountains, Antarctica: Science, v. 137, no. 3532,

p. 751-752. 

Goldich, S. S., Nier, A. O., and Washburn, A. L., 1958, A<H)/K40

age of gneiss from McMurdo Sound, Antarctica: Am.

Geophys. Union Trans., v. 39, no. 5, p. 956-958. 

Hamilton, Warren, 1962, Petrochemistry of probable Paleozoic

granitic rocks from the Ross Sea region, Antarctica:

Art. 225 in U.S. Geol. Survey Prof. Paper 424-C, p. 0209-

C212. 

Knowles, P. H., 1945, Geology of southern Palmer Peninsula,

Antarctica: Am. Philos. Soc. Proc., v. 89, no. 1, p. 132-145. 

Kulp, J. L., 1961, Geologic time scale: Science, v. 133, no. 3459,

p. 1105-1114. 

Long, W. E., 1962, Sedimentary rocks of the Buckeye Range,

Horlick Mountains, Antarctica : Science, v. 136, no. 3513,

p. 319-321.
El08 GEOLOGICAL SURVEY RESEARCH 1962

209. EVALUATION OF THE Pa231/U-Th230/U METHOD FOR DATING PLEISTOCENE CARBONATE ROCKS 1

By J. N. ROSHOLT and P. S. ANTAL, Denver, Colo., and The Marine Laboratory, Miami, Fla.

Work done in cooperation ivith The Marine Laboratory, University of Miami

One of the potential radioactive-dating methods is 

that of measuring the Th230/U and the Pa2ai/U ratios 

in stratigraphically controlled samples. The Th230/U 

ratio provides the most direct approach for determining age in the time range of 40,000 to 250,000 years, 

whereas the use of both ratios provides a more stringent test of the validity of age determinations.

A small amount of previous work has been done 

on variations of this uranium-decay method. Barnes 

and others (1956), in their study of coral-limestone 

cuttings from a drill core taken on a Pacific Ocean 

atoll, were the first to obtain data on the Th230/!! 

ratio in Pleistocene rocks. Sackett 2 made more extensive analyses on this type of material and estimated 

ages from the Th230/U ratio in Pleistocene rocks. 

Tatsumoto and Goldberg (1959) extended the study 

of the ionium-uranium dating technique to other carbonates of the marine environment, primarily oolites 

and coral from Atlantic Ocean islands. An evaluation of this dating method, using Pa231 and Th230, on 

archeologically correlated bone specimens has been 

given by Rosholt (1958, p. 235).

If the calculated age is to represent the true age 

of the material analyzed, three assumptions must be 

valid: (1) the uranium was introduced into the material over a short interval compared to the age of 

the daughter products; (2) no measurable amounts 

of daughter products were introduced directly from 

the surroundings; (3) measurable quantities of uranium or daughter products were not subsequently 

leached from the sample. In most naturally occurring 

material there is bound to be some deviation from all 

three of these assumptions. The purpose of this investigation is to evaluate the overall effect of the deviations in carbonate-rock material for which good 

Pleistocene stratigraphic correlations are available. 

The carbonates include marine limestones, calcareous 

tufa, and cave stalagmites.

RESULTS

Uranium was determined by the polarographic 

method (Antal, 1961) and it is expressed as parts per

1 Contribution No. 418, The Marine Laboratory, University of 

Miami. Research was supported in part by USABC Contract at 40-1 

' (2411) and NSF Grant G-15744 to the Marine Laboratory.

2 Sackett, W. M., 1958, Ionium-uranium ratios in marine deposited 

calcium carbonates and related materials: Washington Univ., St. 

Louis, Ph.D. thesis.

million of uranium. Radiochemical methods (Rosholt, 1957; Rosholt and Dooley, 1960) were used for the 

determination of all the radioactive daughter products, 

which are expressed as equivalents to the parent nuclides in terms of equivalent parts per million. The 

description of the type of material, location, and 

method of Pleistocene stratigraphic correlation for 

marine limestones and terrestrial carbonate rocks are 

listed in tables 209.1 and 209.3, respectively. The 

analytical data, isotope ratios used for calculation, 

apparent ages obtained from these ratios, and estimated ages based on Pleistocene correlation, where 

available, are shown in tables 209.2 and 209.4, respectively.

TABLE 209.1. Description of marine limestones

[Collected by J. Hoflmeister]

Sample Description and Location 

Hoff-G___-__ Miami Oolite, calcarenite just below crust. 

Overlies contact between oolite and Key Largo 

Limestone, Big Pine Key, Fla.

-F______ Breccia containing considerable organic matter

from hole in Key Largo Limestone.

-C___-__ Key Largo Limestone. Calcarenite under top 

crust, Grassy Key, Fla., 30 miles from Big 

Pine Key.

-A______ Key Largo Limestone. Coral rock just under

crust, top sample of core 2, Grassy Key, Fla.

-B_ _____ Key Largo Limestone. Coral rock and weathered

crust at 16-20 foot depth in core 2.

-D______ Key Largo Limestone. Coral rock at 106-113

foot depth in core 2.

-E______ Key Largo Limestone. Coral rock in form of

large pebbles at 156-165 foot depth in core 2, 

just above quartz sand in underlying deposit.

The suite of Florida marine limestones (samples 

Hoff-A to Hoff-G) was collected in 1960 by J. E. 

Hoffmeister, of the University of Rochester. The 

samples contain a mixture of aragonite and calcite. 

The uranium content in all the samples is lower than 

the equivalent Pa231 or Th230, therefore no direct ratio of daughter product to uranium can be calculated. 

To provide a comparison between samples where no 

direct daughter to parent ratio is applicable, apparent minimum and maximum ages are calculated as 

described by Rosholt .(1961, p. 1396). If the assumptions made to obtain the apparent minimum and 

maximum ages approach validity, the actual age of
ARTICLE 209 

TABLE 209.2. Analytical data and apparent ages of marine limestones

E109

Sample

Hoff-G___.__ -....__..

-F. __________ _

-C_______________

-A.________....._

-B___.___________

-D_____. _________

-E___-__..____.__

Eesidue 

(percent)

0. 16 

1. 1 

.58 

. 59 

.40 

.41 

. 31

u

Ppm

1. 02 

.20 

.35 

.46 

.50 

.29 

. 61

Pa23i Th23» Ra2M Thzss

Equivalent ppm

1. 3 

. 54 

2. 9 

1. 5 

4. 4 

4. 3 

4.2

1. 2

. 33 

1. 5 

.86 

4. 1 

3. 5 

3. 7

1. 2 

.32 

2.0 

1.0 

3.7 

2.8 

3.2

0.05 

. 16 

.06 

.05

.04

Th23«/Pa«i

0. 92 

.61 

. 52 

. 56 

. 93 

. 81 

.88

Apparent age (years)

Minimum

150, 000 

45, 000 

23, 000 

33, 000 

95, 000 

77, 000 

88, 000

Maximum

>250, 000 

70, 000 

40, 000 

50, 000 

>250, 000 

180, 000 

230, 000

the sample should be between the two apparent age 

values.

Samples Hoff-A, -B, -D, and -E were selected 

from drill core 2, on Grassy Key, Fla. The core extended to approximately 170 feet, through the Key

Largo Limestone of Sanford (1909, p. 215). Sample 

Hoff-B contained a large amount of oolite (G. Rusnak, oral communication, June 1961) intermixed with 

a weathered crust. A similarity in apparent ages is 

evident between sample Hoff-B and Miami Oolite

TABLE 209.3. Description of terrestrial carbonate rocks

Sample Carbonate type Location Stratigraphic description

61 ER-11. 

60 ER-5_.

-7-

-11. 

61 ER-10.

-I..

-7.

-6.

-4.

-9.

60 ER-16.

-1.

Limestone sinter. 

Calcareous tufa_

Lake marl.

Calcareous tufa.

.doStalagmite, calcite___

.do.

.doLa Vache Cave near Tarascon (Aliege),

French Pyrenees. Collected by L. Zotz. 

Thoming, 4 miles south of Buxton, South

Africa. Collected by K. P. Oakley and

F. G. Orffer, 1953. 

Exposed at gravel quarry near Zeifen, Kreislaufen on Salzach River, southeast

Bavaria, 427 meters above sea level.

Submitted by H. Gross. 

Hungarian site Tata. Collected by L.

Vertes. 

Ehringsdorf, Germany. Collected by G.

Behm-Blanke, Weimar. 

Grotte-abri Lachaud Terrasson, Dordogne,

France, 1942. Submitted by Madame

Leroi-Gourhan. 

"Neanderthal Footprint Cave," Tana della

BASUA, Toirano, Liguria, North Italy.

Collected by Andrea Nersi and K. P.

Oakley, March 1961. 

.._..do _________.____..-__.-._-._-_.-..

.do. .do.

.do.

Stalagmite calcite 

prisms.

Stalagmite, porous___

Stalagmite, calcite 

and aragonite.

Grotto Romanelli, south Italy, from Blanc 

collection. Submitted by K. P. Oakley.

Santa Croce Bisceglie, Italy. Collected by 

L. Cordini, submitted by K. P. Oakley.

Grottedel'Hygene, Arcy-sur-Cure (Yonne). 

Collected by Madame Leroi-Gourham.

Swartkrans, Transvaal, South Africa. 

lected by K. P. Oakley, 1953.

ColOverlies rich Magdalenian deposit of culture

dated at 9700 ±200 years by C14 . 

Lower Gamblian age.

Interstadial lake. Bed of 1-meter thickness 

intercalated in Lower Wurm gravel, below 

Main Wurm ground moraine at depth of 

8 meters.

Age is probably greater than 50,000 years.

Eemian age.

Stalagmite fragment containing considerable 

noncalcareous matter. Proto-Magdalenian I and proto-Madgalenian II.

Outermost stalagmite blocking entrance to 

cave. Contains traces of charcoal. Late 

Wiirm age probable and Main Wiirm age 

possible.

Upper stalagmite forming inner part of mass 

blocking entrance to cave. Main or Late 

Wiirm age probable.

Lower stalagmite blocking entrance to cave. 

Probable age is Main or Late Wiirm.

Stalagmite layer NIV.X (-H). Age is probable Early Wurm.

Stalagmite containing small calcite prisms 

intermixed with charcoal underlying middle Mousterian. Age is probable Early 

Wiirm.

Small conical stalagmite reworked by early 

Mousterians and placed by them on early 

Mousterian soil.

Underlying breccia with Paranthropus.
E110 GEOLOGICAL SURVEY RESEARCH 1962

TABLE 209.4. Analytical data and apparent dates of terrestrial carbonate rocks 

[Uranium analyses by P. Antal]

Sample

61 ER-11__________

60 ER-5-___-______

-7__---_-_---

-11_. -._-____

61 ER-10----------

-1 ______ ___

-8__--_______

 7

-6__--_-__-__

-4-_______-_-

-9____-_-____

60 ER-16_____-____

-1

Residue 

(percent)

2.3 

. 15 

8. 85 

2. 5 

1. 8 

2.0 

. 13 

. 033 

. 034 

. 002 

. 68 

1. 1 

. 06

u

Ppm

0.028 

. 045 

. 049 

.028 

. 021 

. 026 

. 020 

. 110 

. 083 

. 012 

. 10 

. 056 

. 8

pa231 Th230 Ra226 Th228

Equivalent ppm

0. 046 

. 076 

. 36 

. 85 

. 44 

. 43 

. 048 

. 39 

. 34 

. 42 

. 71 

. 13 

1. 06

0. 024 

. 045 

. 31 

. 40 

. 31 

. 36 

. 024 

. 25 

.23 

. 25 

.36 

. 09 

1. 03

0. 12 

. 038

.31 

. 27 

.36 

. 017

. 16

. 77 

. 14 

. 97

0. 013 

. 004 

. 24 

. 13 

. 006 

. 14 

. 012 

<. 005 

<.005 

<. 005 

. 50 

. 081 

<. 01

Th23o/pa23i

0. 52 

. 59 

. 86 

. 47 

. 70 

.84 

. 50 

. 64 

. 67 

. 60 

. 51 

. 69 

. 97

Apparent date of uranium 

introduction (years)

Minimum

23, 000 

40, 000 

90, 000 

10, 000 

60, 000 

80, 000 

18, 000 

50, 000 

56, 000 

40, 000 

21, 000 

60, 000 

160, 000

Maximum

40, 000 

65, 000 

220, 000 

16, 000 

120, 000 

200, 000 

28, 000 

90, 000 

100, 000 

70, 000 

33, 000 

110,000 

>250, 000

Estimated age (years) 

based on Pleistocene 

correlation

< 9, 700 (O date) 

25,000 (approx.) 

40,000-50,000 

> 50, 000 

65,000-95,000 

11,000-25,000 

11,000-40,000 

30,000-65,000 

30,000-65,000 

50,000-65,000 

50,000-65,000 

> 60, 000 

> 200,000

sample Hoff-G, reflecting the influence of the large 

oolite content of Hoff-B. For core 2, exclusive of 

Hoff-B, a gradual increase of apparent age with depth 

is indicated. The apparent maximum ages of samples 

Hoff-D and Hoff-E, near the bottom of the core, and 

the oolite samples are in fair agreement with the 

Th230/U age, probably maximum ages, determined 

by Tatsumoto and Goldberg (1959, p. 205) on similar 

samples. Thus only very approximate estimations of 

chronology can be obtained from marine limestones 

such as these which have largely recrystallized to calcite.

All the terrestrial carbonate samples shown in tables 

209.3 and 209.4 contain equivalent Pa231 in excess of 

the uranium content, therefore minimum and maximum apparent ages had to be calculated for these 

samples. The first five samples represent forms of 

calcareous tufa, which, in general, show a greater 

discrepancy between calculated apparent age and estimated stratigraphic age than do the following stalagmite samples. Stalagmite samples 61 ER-6, -7, -8, 

and 60 ER-1 are dense microcrystalline carbonates 

from which it was anticipated that accurate ages 

might be obtained on the three stratigraphically 

younger specimens 61 ER-6, -7, and -8. Significant 

amounts of uranium apparently were lost from the 

samples during their history, however, and accurate 

ages were not obtained. The apparent minimum ages 

of the three samples appear to be close to their estimated ages. Approximately the same age as the 

minimum age would be obtained for the three samples 

if a continual, steady loss of uranium during this time 

interval were assumed.

Although accurate ages cannot be obtained from 

the stalagmites, some interesting geochemical processes are indicated. The high Pa231 and Th230 concentrations show a previous existence of uranium in 

the 1 to 5 part per million range about the amount of 

uranium that would be expected if the stalagmites 

were initially formed as aragonita. Tatsumoto and 

Goldberg (1959, p. 205) have shown that inorganic 

calcium carbonate precipitates from sea water as aragonite that contains 3 to 5 parts per million uranium, 

a value somewhat higher than that in the stalagmites 

because of the higher uranium content in sea water 

than in ground water.

CONCLUSION

Most uranium is incorporated in carbonates at the 

time of their formation; subsequently, a significant 

amount of uranium is lost by leaching, and the major 

part of the original uranium is lost if the carbonate 

rock recrystallizes from aragonite to calcite. On the 

other hand, only a small amount of the total observed 

uranium is present in bone and wood at the time of 

burial; the major fraction of the uranium is incorporated after burial.

At this stage in the investigation of Pleistocene 

dating by uranium decay, no really suitable type of 

material has been found that indicates an accurate 

dating potential. None of the material covered by 

this study shows promise. Tatsumoto and Goldberg 

(1959) analyzed two selected pieces of aragonitic Key 

Largo Limestone, collected by R. Ginsburg from a 

different locality than the samples described in table
ARTICLE 210 Elll

209.1. The large difference in ThE30/U ages, 130,000 

and 230,000 years, of the two aragonitic limestones, 

which should be approximately the same age (R. Ginsburg, oral communication, Aug. 1961), does not indicate reliable age determinations.

It is doubtful that calcitic samples will be found 

that can be analyzed readily for Pa231 in addition to 

Th230, and that have good dating potential. Primary 

calcite may have dating potential, but because of the 

very low uranium content the Pas31 content is considerably below the present limits of detection. No 

aragonitic stalagmites or stalactites have been analyzed for uranium and daughter products, although 

Murray (1954, p. 486) has found that nearly pure 

aragonitic calcium carbonate exists in some caves. Further investigation of this type of material should be 

made to determine its dating potential.

REFERENCES

Antal, P., 1961, Fehlerquellen bei der polarographischen Bestimmung von Mikrogramraengen Uran (nach der Anreicherung aus festen Proben und Wassern) : Mikrochimica 

Acta, v. 1961, p. 235-244.

Barnes, J. W., Lang, E. J., and Potratz, H. A., 1956, Ratio of 

ionium to' uranium in coral limestone: Sci., v. 124, p. 175-

176.

Murray, J. W., 1954, The deposition of calcite and aragonite 

in caves: Jour. Geology, v. 62, p. 481-492.

Rosholt, J. N., Jr., 1957, Quantitative radiochemical methods 

for the determination of the sources of natural radioactivity: Anal. Chemistry, v. 29, p. 1398-1408.

 1958, Radioactive disequilibrium studies as an aid in 

understanding the natural migration of uranium and its 

decay products: United Nations Internat. Conf. on the 

Peaceful Uses of Atomic Energy, 2d, Geneva, 1958, v. 2, 

p. 230-236.

 1961, Uranium migration and geochemistry of uranium

deposits in sandstone above, at, and below the water table, 

Part I. Calculation of apparent dates of uranium migration in deposits above and at the water table: Econ. 

Geology, v. 56, p. 1392-1403.

Rosholt, J. N., Jr., and Dooley, J. R., Jr., 1960, Automatic 

measurements and computations for radiochemical analyses : Anal. Chemistry, v. 32, p. 1093-1098.

Sanford, Samuel, 1909, The topography and geology of Southern Florida: Florida Geol. Survey, 2d Ann. Rept, p. 177- 

231.

Tatsumoto, M., and Goldberg, E. D., 1959, Some aspects of the 

marine geochemistry of uranium: Geochim. et Cosmochim. 

Acta, v. 17, p. 201-208.

SEDIMENTARY PETROLOGY

210. RELATION OF THE PORE VOLUME OF SILTY SEDIMENTS TO OVERBURDEN LOAD, PARTICLE SIZE,

AND SORTING

By ROBERT H. MEADE, Menlo Park, Calif.

Studies of saturated unconsolidated to semiconsolidated silty sediments in western Fresno County, Calif., 

show that a progressive reduction in pore volume results from an increasing overburden load. The intergranular processes involved in the compaction of some 

of these sediments were discussed in an earlier paper 

(Meade, 1961), to which the reader is referred for 

information on the geographic location and a short 

descriptive introduction to the studies. This article 

summarizes some of the pore-volume relations observed 

in the sediments. The results reported here supersede 

the porosity-depth relations shown in figure 324.2 of 

the earlier article. In the earlier study, the use of 

depth as a variable was misleading because pumping 

had reduced the artesian pressures in the deeper sediments far below their natural levels, thereby increasing the effective overburden (grain-to-grain) load in 

these zones and causing the relation between depth 

and effective load to become irregular. In this article, 

therefore, pore volume is plotted against effective load. 

Furthermore, the samples for the present study were 

selected more carefully and given more thorough statistical treatment.

Samples for this study were taken from flood-plain 

and alluvial-fan deposits cored in three holes in western Fresno County: 14/13-11, 16/15-34, and 19/17-22 

(locations shown in Meade, 1961, fig. 324.1). The 

numerical data were taken mainly from porosity and 

particle-size measurements that were made in the Geological Survey Hydrologic Laboratory (Johnson and
E112 GEOLOGICAL SURVEY RESEARCH 1962

Morris, 1962). Effective overburden loads were computed by E. E. Miller by subtracting the fluid pressures in the sediments as measured at the time the 

sediments were cored from the total load (bulk 

weight) of the overlying sediments.

The statistics that describe the observed pore-volume relations were computed on a digital computer 

and are listed in the table below. For explanations 

of the statistical terms and assumptions involved in 

the computations, see Ezekiel (1941). The sediments 

were divided into four groups by size: silty sands, 

sandy silts, clayey silts, and silty clays. Void ratio, 

rather than porosity, is used as a measure of pore 

volume because its correlation with effective overburden load seems to be approximately linear (Engelhardt, 1960, p. 39^1). The median diameter (Md<f>)

Details of correlation and regression between void ratio and other

factors

[e=void ratio (ratio of pore volume to volume of solids), log L=logarithm (base 10) 

of effective overburden load (kg per cm2), Md«=median particle diameter (<t> units, 

see Inman, 1952, p. 133), QZ>$ = quartile deviation (particle sorting, <t> units, see 

Krumbein, 1936, p. 102-103)]

Group

Letter

designation

in figure

210.1

Number of

samples

Coefficient of

linear

correlation

Probability

of significance

of correlation'

Regression equation

Standard

error of

estimate

of void

ratio

Simple linear regressions of void ratio on effective overburden load

Clayey silt-.. __

Silty clay 

A

B

C

D

39

28

40

28

-0.40

-.68

-.62

-.29

>0.99

>> 00

>.99

.92

e=1.02-0.23 log L

e-1.36- .41 log L

e=1.35- .40 log L

e=1.01- .16 log L

0.12

.07

.08

.12

Significant simple linear regressions of void ratio on particle-size measures

Silty clay.... _ ..

A

B

D

39

28

28

-0.44

.29

.35

--» QQ

.93

.96

e-0.81-0.099 QD«

(2)

e=.30+.054 Afd*

0.11

.12

Multiple linear regressions of void ratio on effective overburden load, particle size,

and sorting

Silty sand.. _ ...

Sandy silt. .......

Clayey silt... ....

Silty clay-. .......

Silts..............

A

B

C

D

B+C

39

28

40

98

68

0.56

.80

.67

do

.67

>0.99

>.99

>.99

QQ

]> 99

C=0.94-0.181ogL

+0.044 Md« 

-0.10Q.D0 

e-l.10-.41 log L

+.083 Md* 

-.096 QD* 

«=1.19-.401ogL

+.038 Md* 

-.056 QD«

+.049 Md* 

«-1.29-.401ogL

+. 019 Mdt 

-.031 QD*

0.11

.06

.08

.11

.08

i Based on "Student" 's J-test (see Ezekiel, 1941, p. 318-324). 

* Correlation between e and Mrf«; equation not computed.

and quartile deviation (QD<f>] are used as measures of 

particle size and sorting, respectively. For explanations of these measures, see Inman (1952, p. 133) and 

Krumbein (1936, p. 102-103). Whether the relations 

of Mdcf) and QD<f> to void ratio are linear is not certain, 

but they are assumed to be so in this study.

Simple regressions of void ratio on effective overburden load are shown in figure 210.1. The sandy 

and clayey silts of groups B and C show the most 

unequivocal decrease in void ratio with increasing 

overburden load. The slopes of the regression lines 

for both groups are virtually identical (about -0.40), 

regardless of whether the load is assessed alone or in 

conjunction with the particle-size factors: compare the 

regression coefficients of the load terms in the simpleand multiple-regression equations in the table below.

The relation of void ratio to load in the silty sands 

in group A is less clear, because it is obscured by a 

significant correlation between void ratio and sorting 

(fig. 210.2, left). The sorting deteriorates (that is, 

QD<f) increases) systematically with depth and therefore influences the apparent slope of the void ratioload line. The multiple-regression equation for group 

A indicates that the slope of this line should be less 

steep than illustrated in figure 210.1 A

The slope of the void ratio-load line for the silty 

clays in group D is anomalously steep. A significant correlation between void ratio and Md$ (fig. 

210.2, right) partly obscures the relation between void 

ratio and load, but not enough to explain the anomalously steep slope of the void ratio-load line.

Figure 210.3 is a comparison of the results observed 

in Fresno County and those summarized by Skempton 

(1953, p. 52-55) from observations made elsewhere 

in the United States and in Britain. Skempton's 

curves (fig. 210.3, left) show that, with decreasing 

particle size, the void ratio at a given load increases, 

and the void ratio decreases more rapidly under increasing load. In figure 210.3, right, the regression 

lines from figure 210.1 (solid) are superimposed on 

Skempton's curves (dashed). Skempton's nomenclature based on Atterberg limits does not coincide 

with that of Shepard (1954), which is used in figure 

210.1. The silty sands of group A correspond to 

Skempton's "silts," the sandy and clayey silts of 

groups B and C correspond to his "silty clays," and 

the silty clays of group D correspond to his "clays." 

Even when differences of nomenclature are taken into 

account, the Fresno County sediments are more porous 

than the group of sediments whose void ratios were 

summarized by Skempton. Perhaps this is due to a 

higher proportion of montmorillonite in the Fresno
to K O- | gi o g> 

' Bil O a o, o < g- 0 2 <o S o 3- r° S» 

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EFFECTIVE OVERBURDEN LOAD , IN 

KILOGRAMS PER SQUARE CENTIMETER

f o.

EFFECTIVE OVERBURDEN LOAD , IN 

KILOGRAMS PER SQUARE CENTIMETER
E114 GEOLOGICAL SURVEY RESEARCH 1962

County sediments (Meade, 1961, p. D89), but this supposition would have to be checked against the claymineral compositions of the other sediments.

Whereas the slopes of the regression lines for 

groups A, B, and C nearly correspond to the appropriate slopes of Skempton's curves for "silts" and 

"silty clays," the regression line for group D is anomalous. Were the analogy with Skempton's curves carried through, line D would be well to the right of 

line B-C and would have a less steep slope. Physical 

explanations for this anomaly do not come to mind 

readily. As mentioned above, it is not explained by 

variations in median size. If the cause of the anomaly 

were elastic rebound of the clays between the time of 

coring and the time the pore volumes were measured, 

rgression line D would appear well to the right of 

the other lines. Perhaps the anomaly is due to variations in some other factor or factors that were not 

assessed statistically (sorting or fabric, for example). 

Or perhaps and I favor this explanation the errors 

inherent in sampling ("errors" in the statistical sense, 

that is) were so great that they obscured the true 

relation. In any case, the results of the statistical 

study of pore volume in the silty clays of group D are 

inconclusive.

The above evidence indicates that, under effective 

overburden loads in the range from 10 to 70 kg per 

cm2 , the rate of change of void ratio with effective 

overburden load (expressed as Iog10 kg per cm2 ) in 

the sandy and clayey silts seems to be about -0.40, 

and that the rate of change in the silty sands may be 

near -0.20.

REFERENCES

Engelhardt, Wolf von, 1960, Der Porenraum der Sedimente: 

Berlin, Springer, 207 p.

Ezekiel, Mordecai, 1941, Methods of correlation analysis (2d 

ed.) : New York, John Wiley & Sons, 531 p.

Inman, D. L., 1952, Measures for describing the size distribution of sediments: Jour. Sed. Petrology, v. 22, p. 125-145.

Johnson, A. I., and Morris, D. A., 1962, Physical' and hydrologic properties of water-bearing deposits in the Los Bantos-Kettleman City area, California: U.S. Geol. Survey 

open-file rept., 179 p.

Krumbein, W. C., 1936, The use of quartile measures in describing and comparing sediments: Am. Jour. Sci., v. 32, 

p. 98-111.

Meade, R. H., 1961, Compaction of montmorillonite-rich sediments in western Fresno County, California: Art. 324 in 

U.S. Geol. Survey Prof. Paper 424-D, p. D89-D92.

Shepard, F. P., 1954, Nomenclature based on sand-silt-clay 

ratios: Jour. Sed. Petrology, v. 24, p. 151-158.

Skempton, A. W., 1953, Soil mechanics in relation to geology: 

Yorkshire Geol. Soc. Proc., v. 29, p. 33-62.

GEOMORPHOLOGY AND GLACIAL GEOLOGY

211. MENAN BUTTES, CONES OF GLASSY BASALT TUFF IN THE SNAKE RIVER PLAIN, IDAHO

By WAEREN HAMILTON and W. BRADLEY MYERS, Denver, Colo.

The Menan Buttes are two well-preserved cones of 

glassy olivine-basalt tuff that rise from the Snake 

River Plain about 20 miles north of Idaho Falls in 

southeastern Idaho. The cones were erupted through 

the water-saturated flood plain of the Snake River, 

and to this is due their very unusual character. Each 

cone is about 2 miles long in a northeast direction 

(fig. 211.1), but only about two-thirds as wide, and 

each has a comparably oblate, large summit crater 

about 1/2 mile long and 200 to 400 feet deep. The 

crater rim of the northern cone stands as high as 

800 feet above the surrounding plain (fig. 211.2), and 

the southern one 550 feet.

The northeast-trending Snake River Plain is about 

50 miles wide at the longitude of Menan Buttes. It 

has a broad medial upland surface underlain largely 

by basalt flows of late Quaternary age, and broad 

marginal alluvial lowlands within which flow most of 

the streams. The Menan Buttes are at the northwest 

edge of the southern lowland. The edge of the medial 

upland is here the river-eroded front of an aa flow 

of olivine basalt (fig. 211.1). The flow surface is 

ragged and nearly free of loess; thus, the flow is probably of Recent age, although it probably is older than 

the Menan Buttes. Henrys Fork of the Snake River 

was diverted southward from its general southwest
43°47'

U.S. Geological Survey photographs 

Coordinates are approximate 1 MILE

FIGURE 211.1. Vertical aerial photographs of the Menan Buttes, Idaho, mounted as a stereotriplet.
E116 GEOLOGICAL SURVEY RESEARCH 1962 

Prevailing wind > 

SOUTHWEST

6000' -|

50OO' -

4000'

NORTHEAST

1 MILE

FIGURE 211.2. Section of North Menan Butte.

course by the building of the buttes and joins the main 

fork of the Snake at the east edge of the south cone. 

Most of the basalt vents of the Snake River Plain 

are marked by broad, low lava domes. Cinder cones 

are uncommon, and tuff cones like the Menan Buttes 

are rare.

The asymmetry of the Menan Buttes cones and their 

craters clearly demonstrates northeastward drift of 

pyroclastic ejecta parallel to the prevailing winds. The 

cones are markedly asymmetric elongate northeastward and their craters lie toward the southwest ends 

of their midlines. The craters are similarly asymmetric, their northeast walls being the gentlest and 

the farthest from the vents and standing 100 to 200 

feet higher than the rest of the rims.

The cones were built by small fragments of solid 

glass that were erupted explosively within the craters 

and then fell to form the inner walls of the craters 

as well as the outer slopes of the cones. Tuff beds 

of the outer slopes steepen gradually upward and are 

near the angle of repose (35°) only in places near 

the rim. The crater walls are dip slopes on beds dipping inward at about the angle of repose, and many 

beds roll over sharply at the rim to dip outward at 

angles of 20° to 35° on the upper slopes of the cones. 

Layering is concordant to the gentle slopes around 

the base of the cones. This is the general case on the 

steep upper slopes also, where there are however many 

aberrant dips, both inward and outward, indicating 

a complex history of overlapping falls. A number of 

irregular bedding features in the cones can be seen by 

stereoscopic study of figure 211.1. No fluidal lava features were found on the cones.

Conspicuous layering throughout the tuff is marked 

by irregular variation in size of glass fragments. Individual layers are a fraction of an inch to several 

feet thick and are generally thickest high on the cones. 

The tuff is firmly coherent but porous, and most is 

soft enough so that it thuds and crumbles when struck 

by a hammer. The tuff is uniform drab, dark olive

gray. It consists largely of sand-size to granule-size 

fragments of clear, unaltered basaltic glass that is 

dark in hand specimen but light brown in thin section. There is much pore space, and tiny fragments 

of similar glass only partly fill the interstices. Palagonite and other amorphous and cryptocrystalline material discontinuously coat most fragments and give 

them a dull, drab appearance in hand specimen.

Glass fragments were solid and too cold to flow 

when they were erupted. Fragments are subround to 

angular, are undeformed, and consist of equant granules rather than shards or droplets. Many fragments 

are microvesicular; a few vesicles are filled, mostly by 

calcite. Elongate vesicles are common within granules but orientation is random, and neither welding 

nor secondary flowage was recognized.

The glass is of olivine-basalt composition that is 

probably typical of the region despite the unusual occurrence. The refractive index of the glass in one 

specimen ranges from 1.610 to 1.616 for yellow light,1 

indicating a silica content of about 4T percent according to the curves presented by Williams and others 

(1954, fig. 7). The glass contains scattered euhedral 

crystals of forsteritic olivine ( 2F = 85°-90°) and 

euhedral laths of labradorite, and sparse octahedra of 

magnetite. No pyroxene or ilmenite was seen. The 

Quaternary basalt of the Snake River Plain is olivine 

bearing and slightly alkaline, typically contains 45 to 

48 percent SiO2, and has late-crystallizing augite and 

ilmenite (Howard A. Powers, written communication, 

1961) and is thus like the Menan Buttes material insofar as they can be compared.

The cones were erupted upon the flood plain of the 

Snake River and contain abundant xenolithic cobbles, 

pebbles, and sand grains from the river alluvium. 

Xenoliths are largest and most numerous near the 

crater rims on both inner and outer sides, but are 

plentiful in many places about the lower slopes also.

1 The fast, accurate focal-masking method of Cherkasov (1957), as 

improved by Ray E. Wilcox (oral communication, 1962), was used.
ARTICLE 211 E117

Rounded river cobbles have a common maximum diameter of 6 inches and some are as large as 1 foot. 

Cobbles and pebbles consist partly of locally derived 

basalt and rhyolite, and partly of aplite, granite, 

quartzite, and other resistant rocks from more distant 

sources. Sand grains of quartz and of rhyolite obsidian and welded tuff, all also derived from river 

alluvium, are conspicuous throughout the tuff of the 

buttes. Xenolithic material totals perhaps 1 or 2 

percent of the volume of the tuff.

The unusual character of the eruptions that formed 

the Menan Buttes was obviously due to the presence 

of water-saturated flood-plain deposits immediately beneath the vents. The cones were erupted across the 

courses of the Snake River and Henrys Fork in the 

wettest part of the flood plain. Upon contact with 

abundant water at very shallow depth, the lava was 

chilled, slightly palagonitized, and erupted explosively. Glass fragments were solid when they hit the 

ground and probably were chilled to solidity during 

extrusion. We picture the eruption as a steam-charged 

spray of solid fragments.

Interbedded with the underlying gravels are basalt 

flows, for angular blocks of basaltic lava also form 

abundant inclusions in the tuff. Basalt blocks as 

large as 3 to 5 feet in diameter are numerous near the 

crater rims and attest to the explosive power of the 

eruptions, but blocks are smaller and fewer on the 

lower slopes of the cones.

North and South Menan Buttes are of similar very 

late Pleistocene or, more probably, early Recent age. 

Both show comparably small amounts of dissection 

by shallow radial gullies, the deepest gullies draining 

the long lee slopes of the cones. Both cones could 

have formed within as short a period as a few months. 

Before eruption of the cones, Henrys Fork flowed 

where North Butte now stands and had a course from 

upper left to lower right of the left picture of figure 

211.1; building of the buttes diverted the river southward, and it now joins the Snake River at the east 

edge of South Menan Butte. The rivers have cut 

bluffs along the eastern and southern sides of the 

southern butte, but the cones must have been erupted 

upon the flood plain at virtually its present level.

Between North and South Menan Buttes is a low, 

arcuate ridge forming a part of the rim of a third 

tuff cone (fig. 211.1). This cone has been breached 

by erosion along its southwest edge, and overlapped 

by South Menan Butte along its southeast edge. Three 

small narrow ridges trending northeastward from the 

north side may represent fissure eruptions. The crater 

is longer and wider than the craters forming the 

buttes, but its rim stands only 200 feet above the

plain and its slopes are gentler than those of the 

main cones. The base of the middle cone is buried 

by flood-plain alluvium and the cone remnants are 

more dissected than are comparable parts of the main 

cones.

The Menan Buttes and intervening cone remnant 

were apparently erupted along a north-northwesttrending fissure zone of recurrent basalt eruptions. A 

small, much eroded remnant of an old cone lies on 

the same trend !/2 mile south of the southeast corner 

of figure 211.1 (Stearns and others, 1938, pi. 4), and 

another small cone lies just beyond the northwest 

corner of figure 211.1.

Tuff cones similar to the Menan Buttes probably 

have similar origins in eruptions through water-saturated materials. (Basalt eruptions in the far more 

common relatively dry environments produce flows, 

domes, fountains, and cinder cones.) The several tuff 

cones which Howard A. Powers (oral communication, 

1962) has seen in other parts of the Snake River Plain 

were erupted, like the Menan Buttes, through watersaturated surficial material. Tuff cones in the Hawaiian Islands occur only near the coast or as offshore 

islets, thus also in areas where abundant water was 

encountered by the erupting basalt (Wentworth and 

Macdonald, 1953, p. 68). A number of cratefed tuff 

cones in the Pinacate volcanic field of northwestern 

Sonora (Jahns, 1959) are similar to the Menan Buttes, 

and at least one (Cerro Colorado) appears identical 

in most respects. Although Jahns considered the 

craters to have formed by collapse, the marked elongation of the Cerro Colorado cone and crater parallel 

to the prevailing wind, and the downwind asymmetry 

of the crater, indicate that crater at least to have 

formed by explosive rather than collapse processes.

The Menan Buttes are shown with a 10-foot contour interval on the 1:24,000-scale Menan Buttes topographic quadrangle map of the Geological Survey. 

The shaded-relief edition of this map is included in 

the Survey's sets of 25 and 100 topographic maps 

intended for use in the teaching of physiography and 

map interpretation (Upton, 1955; the buttes were 

termed cinder cones by Upton). The three vertical 

aerial photographs of the buttes (fig. 211.1) have an 

original scale of about 1:27,600, and are available 

from the Map Information Office, U.S. Geological 

Survey, Washington 25, B.C. (photographs 2-25, 2-26, 

2-27, project GS-LV, Oct. 8, 1950).

REFERENCES

Cherkasov, Yu. A., 1957, O primenenii "fokalnogo ekranirovaniya" pri izmereniyakh pokazateley prelomleniya immersionnym metodom [Application of "focal screening" 

to measurement of indices of refraction by the immersion
E118 GEOLOGICAL SURVEY RESEARCH 1962

method], in Sovremennye metody ruineralogicheskogo, issledovaniya gornykh porod, rud, i mineralov : Gosudar. 

Nauch.-tekh. Izdat. Lit. Geol. Okh., Moscow, p. 184-207. 

(English translation in Internat. Geol. Rev., 1960, v. 2, 

p. 218-235)

Jahns, R. H., 1959, Collapse depressions of the Pinacate volcanic field, Sonora, Mexico, in Heindl, L. A., ed., Southern Arizona Guidebook II: Ariz. Geol. Soc., p. 165-184.

Stearns, H. T., Crandall, Lynn, and Seward, W. G., 1938, 

Geology and ground-water resources of the Snake River

Plain in southeastern Idaho: U.S. Geol. Survey WaterSupply Paper 774, 268 p.

Upton, W. B., Jr., assembler, 1955, A set of one hundred 

topographic maps illustrating specific physiographic features [index map]: U.S. Geological Survey.

Wentworth, C. K., and Macdonald, G. A., 1953, Structures and 

forms of basaltic rocks in Hawaii: U.S. Geol. Survey 

Bull. 994, 98 p.

Williams, Howel, Turner, F. J., and Gilbert, C. M., 1954, 

Petrography: San Francisco, Freeman & Co., 406 p.

212. COMPARISON OF DRAINAGE ON TOPOGRAPHIC MAPS OF THE PIEDMONT PROVINCE

By ENNIO V. GIUSTI and WILLIAM J. SCHNEIDER, Washington, D.C.

Because topographic maps portray only a selected 

part of the total drainage of an area, they may be 

inadequate for certain detailed hydrologic studies. 

Actual differences in drainage between basins are due 

to differences in geology, topography, and climate. 

However, the proportions of the streams shown on 

topographic maps differ also because of differences in 

map scale, date of publication, and publishing organization; these differences are analyzed in this article 

for maps of the Piedmont province.

For the purposes of this study, the Piedmont province may be considered reasonably homogeneous in 

geology, topography, and climate. Thus, a random 

selection of maps of the Piedmont from Virginia 

southward into Alabama practically eliminates differences between maps that are attributable to real variations in the number of streams per unit area.

The variables used in this study are drainage area 

(A), stream order (^), and number of streams of a 

given order (Nu). Indices A and u are based on a 

stream classification proposed by Horton (1945) and 

modified by Strahler (195Y). Although stream frequency and drainage density are related, stream frequency was chosen as a variable for this study because it is probably a more reliable measure of the 

morphogenesis of the area. Use of the number of 

streams rather than stream length eliminates the additional variability inherent in current map-making procedures caused by the process of editing stream lengths. 

The three variables were measured on 108 topographic 

maps compiled by various organizations at various 

map scales, as shown in table 212.1. Maps of the

Geological Survey were subdivided into two groups 

according to year of publication because a major 

change in standards for drainage portrayal was made 

in 1954.

TABLE 212.1. Stream frequency portrayed in various map series

Group

No.

1

2

3----.

5

6

7--__-

Number 

of maps

13

10

18

15

12

6

15

19

Publishing 

agency '

AMS, SCSUSGS .-

USGS .__

USGS _-

USGS __

USGS .

Period of 

publication

1943-57---.

Before 1954_

Afterl954__

Before 1 954_

After 1954_ _

1943-57_-_-

1891-1907- _

1952-57. __.

Scale

1 : 24, 000

1:24,000

1:24,000

1:62, 500

1 : 62, 500

1:62, 500

1:125,000

1 : 250, 000

Number 

of firstorder 

streams

per square 

mile

3. 2

3. 2

1. 4

1. 25

. 43

1. 9

2. 1

. 14

i AMS, Army Map Service; SCS, Soil Conservation Service; USGS, Geological 

Survey.

Iii addition to the 108 topographic maps analysed, 

the complete drainage network for the Yellow River 

basin in Georgia was studied photogrammetrically 

from aerial photographs taken at 7,200 feet. The 

relation between the number of first-order streams and 

drainage area for subareas of the Yellow River basin 

is shown in figure 212.1. The relation is of the form

Nuj. = k Ab.

In this relation for the Yellow River basin, as well 

as for those for the standard topographic maps, the 

exponent b did not differ statistically from unity at
ARTICLE 212 E119

500 -

DRAINAGE AREA (A), IN SQUARE MILES

FIGURE 212.1. Relation between number of first-order streams and drainage areas for subareas of the Yellow River basin,

Georgia.

the 95-percent confidence level. Accordingly, the basic 

equation was modified to

Nui = k A,

in which the constant k is an index of the number 

of first-order streams per unit area. This relation is 

shown in figure 212.2 for each of the map groups 

listed in table 212.1. Table 212.1 also lists the number of streams per square mile for each map group. 

The ratio of these quantities as listed in table 212.2, 

can be used to convert counts of first-order streams 

determined for any map group to a comparable value

for a different map group. Thus roughly comparable 

measures of stream frequency can be computed from 

maps of different types.

BEFEBENCES

Horton, R. E., 1945, Erosional development of streams and 

their drainage basins hydrophysical approach to quantitative morphology: Geol. Soc. Arner. Bull., v. 56, no. 3, 

p. 275-370.

Strahler, A. N., 1957, Quantitative analysis of watershed geomorphology; Am. Geophys. Union Trans., v. 38, no. 6, 

p. 913-920.
GEOLOGICAL SURVEY RESEARCH 1962

EXPLANATION 

1:250,000 scale

x 1:125,000 scale

 1:62,500 scale

* 1:24,000 scale

0.05 0.1 0.5 5 10 50 100 

DRAINAGE AREA^J.IN SQUARE MILES

500 1000 5000

FIGURE 212.2. Relation between number of streams and area for maps of different scale.

TABLE 212.2. Factors to convert stream frequency from a map at a given scale to the comparable number on another type of map.

Multiply if entering from the side; divide if entering from the top.

[Publishing organizations abbreviated as in Table 212.1]

Type of map

Infrared photogrammetry

1:24,000

AMS, SCS

1 : 24,000

USGS

Before 1954

1:24,000

USGS

After 1954

1 : 62,500

AMS

1:62,500

USGS

Before 1954

1:62,500

USGS

After 1954

1:125,000

USGS

1 : 250,000

AMS

Infrared 

photogrammetry

1

22

22

51

37

57

165

34

507

1:24,000 

AMS, SCS

0.045

1

1

2.3

1.7

2. 6

7.5

1. 5

23

1:24,000 

USGS before 

1954

0. 045

1

1

2. 3

1. 7

2. 6

7. 5

1. 5

23

1:24,000 

USGS after 

1954

0.02

. 44

.44

1

.74

1. 12

3.25

.67

10

1:62,600

AMS

0.027

.59

.59

1. 35

1

1. 52

4. 4

.90

13. 5

1:62,500 

USGS before 

1954

0.018

. 39

. 39

.90

. 66

1

2.9

.60

8.90

1:62,500 

USGS after 

1954

0.006

. 135

. 135

. 31

.23

. 34

1

. 20

3.07

1:125,000 

USGS

0.03

. 66

.66

1. 5

1. 1

1. 68

4. 90

1

15

1:250,000 

AMS

0.002

.044

.044

. 10

.074

. 112

.325

.067

1
ARTICLE 213 E121

213. CORRELATION OF SOME GLACIAL DEPOSITS IN NEW MEXICO 

By GERALD M. EICHMOND, Denver, Colo.

Clear evidence of Pleistocene glaciation in New 

Mexico is mostly confined to the higher parts of the 

Sangre de Cristo Mountains, and to the Sierra Blanca 

in the south-central part of the State (see index map, 

fig. 213.1). An ancient till, correlated with the Cerro 

Till of the San Juan Mountains, has also been reported from the Canjillon Divide area in the northcentral part of the State, west of the Eio Grande 

(Ellis, 1935; Smith, 1936).

Glaciation in the Sangre de Cristo Mountains has 

been reported by Salisbury (1901) and Stone (1901), 

and subsequently by Ellis (1931). Eay (1940) was 

the first to recognize multiple glaciation in the region. 

He described moraines of five "substages" of the Wisconsin "glacial stage" in the vicinity of Wheeler 

Peak, northeast of Taos, in the valleys draining Jicarita Peak, southeast of Taos, and in the vicinity of 

Lake Peak, northeast of Santa Fe. He also showed 

that deposits described as moraines in the Moreno 

Valley (Ellis, 1931), are not of glacial origin. Smith 

and Eay (1941) further reported a complex of Wisconsin moraines on the northeast side of Sierra Blanca 

Peak below a cirque mentioned briefly by Antevs 

(1935).

In order to correlate Eay's "substages" with the 

glaciations of the Wind Eiver Mountains of Wyoming 

(Blackwelder, 1915; Eichmond, 1948, 1957a, 1962; 

Moss, 1951b; Holmes and Moss, 1955), the writer 

briefly re-examined the succession of glacial deposits 

along Eio Hondo on the slopes of Wheeler Peak (fig. 

213.1), along Eio Nambe on the slopes of Lake Peak 

(fig. 213.2) and on Sierra Blanca Peak (fig. 213.3). 

New topographic maps in two of these areas have 

made possible a more accurate means of plotting end 

moraines and of determining their altitudes than was 

available to Eay. Deposits of glaciations recognized 

by the writer are correlated with those described by 

Eay (1940) as shown in the table on page E122.

CRITERIA OF CORRELATION

Deposits of three glaciations from oldest to youngest, the Buffalo, Bull Lake, and Pinedale were described in the Wind Eiver Mountains of Wyoming 

by Blackwelder in 1915. Subsequent work has demonstrated 3 pre-Bull Lake glaciations (Cedar Eidge, 

Sacagawea Eidge, and Dinwoody Lake) (Eichmond, 

1962) included in deposits formerly called Buffalo, 2

stades of Bnll Lake Glaciation (Eichmond, 1948; 

Moss, 1951a; Holmes and Moss, 1955), 3 stades of 

Pinedale Glaciation (Eichmond, 1948, 1962), and 2 

stades (Temple Lake and Gannett Peak) of Neoglaciation (Hack, 1943; Moss 1951a, 1951b; Eichmond, 

1957b). Criteria for recognition of these glaciations 

and their subdivisions in the Wind Eiver Mountains 

and elsewhere have been described in many papers 

(Blackwelder, 1915, 1931; Fryxell, 1951; Moss, 1951a; 

Holmes and Moss, 1955; Eichmond, 195Ya, 1962; Nelson, 1954).

The following characteristics of the deposits in the 

Wind Eiver Mountains were used as the basis for 

correlation with the deposits in New Mexico described here.

Cedar Ridge and Sacagatwea Ridge Glaciation. 

deeply weathered; red clayey soils; lack glacial topography; confined to interstream divides.

Dinwoody Lake Glaciation. Weathering as above, 

but forming poorly preserved broad, mature moraines 

on canyon walls or floors.

.Locally deposits of all three glaciations are superposed and separated by thick red clayey soils.

Bull Lake Glaciation. Commonly two broad, mature moraines (early and late stades) in canyons; dissected by axial and tributary streams; mature zonal 

soils 3 to 4 feet thick. The two tills grade into separate outwash deposits, which, where superposed, are 

separated and overlain by mature zonal soils similar 

to those on the moraine deposits.

Pinedale Glaciation. Commonly 3 well-preserved 

rough, bouldery moraines (early, middle, and late 

stades) in canyons upstream from moraines of the 

Bull Lake Glaciation. Immature zonal soils 1 to 2 

feet thick. Outwash gravels underlie separate terraces or form compound fills. Locally thin azonal 

soil or dark alluvial silty-clay layer separates outwash 

of early and middle stades.

Neoglaciation. Temple Lake Stade small moraines 

or rock glaciers in cirques. Thin azonal soil, tundra 

or scrub spruce vegetation. Gannett Peak Stade 

fresh moraines or rock glaciers above those of Temple 

Lake Stade; no soil; commonly barren, but locally 

with lichen on boulders and sparse pioneer vegetation; 

moraines lie in front of existing glaciers or in recently evacuated cirques.
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105°25'

3 MILES

J

FIGURE 213.1. Moraines in the valley of Rio Hondo, Wheeler 

Peak, Sangre de Cristo Mountains, N. Mex. Be and Bl, 

early and late stades of Bull Lake Glaciation; Pe, Pm, and 

P I early, middle, and late stades of Pinedale Glaciation; 

T and G, Temple Lake Stade and Gannett Peak Stade of 

Neoglaciation. Index map: 1, Wheeler Peak; 2, Lake Peak; 

3, Sierra Blanca.

are extremely blocky, in part owing to removal of the 

interstitial fine material by melt water. The blocky 

material extends along the lower southwest slopes of 

the basin, where it was mapped as Wisconsin IV "substage" by Eay. The soil on both the outer and inner 

moraines has an immature Brown Podzolic profile like 

that on the end moraine at 10,200 feet. Because of 

the hummocky surface, the abundance of boulders, and 

the immature but zonal character of the soil profile, 

these moraines are correlated with the early, middle, 

and late stades of the Pinedale Glaciation.

Within and overlapping the lips of the 11 small 

cirques that encircle the main basin are small blocky 

moraines and rock glaciers that were referred by Eay

to his Wisconsin V "substage." Their average altitude is 11,850 feet, though one extends down across 

the uppermost Pinedale moraine. The blocks are covered by lichen and, where fine material is exposed at 

the surface, the moraine bears a thin azonal soil. 

These features are characteristic of the Temple Lake 

Stade of Neoglaciation.

Four of the cirques contain one or two still younger 

moraines or rock glaciers that have no soil and only 

a few sparse plants. These deposits are fresh, but do 

not appear active. They are correlated with the 

Gannett Peak Stade of Neoglaciation.

LAKE PEAK

Along the Eio Nambe, on the northeast slope of 

Lake Peak (altitude 12,409 ft) broad, evenly sloping 

lateral moraines, 60 to 80 feet high, lie on rock shoulders within the relatively open canyon at altitudes 

of between 10,000 and 10,600 feet near the confluence 

of the north and south forks and about 100 feet above 

the stream (fig. 213.2) The end moraine was probably

R.10E 11 E.

3 MILES

J

FIGURE 213.2. Moraines in the valley of Rio Nambe on 

Lake Peak, Sangre de Cristo Mountains, N. Mex. 

Pre-B, pre-Bull Lake glaciation; Be and Bl, early and 

late stades of Bull Lake Glaciation; Pe, Pm, PI, early, 

middle, and late stades of Pinedale Glaciation; T, 

Temple Lake Stade of Neoglaciation.

667422 O - 63 - 9
E124 GEOLOGICAL SURVEY RESEARCH 1962

at an altitude of just below 10,000 feet and a little over 

2 miles from the cirque heads. The surface of the 

moraines is very smooth, and very few boulders are exposed. The till contains many deeply disintegrated 

stones, mostly of granite, gneiss, and schist, but some 

are fresh and a few preserve soled and striated surfaces. 

The material is mantled by a compact reddish (7.5T7?) 

clayey sandy soil more than 6 feet thick. The thickness, color, and clay content of the soil suggest that 

the till is of pre-Bull Lake age. The location of the 

deposits within the canyon and the preservation of 

morainal topography suggest correlation with the Dinwoody Lake Glaciation (Richmond, 1962) the youngest 

of the three pre-Bull Lake glaciations, in the Wind 

River Mountains.

Higher in the canyon of the South Fork of the Rio 

Nambe, at 10,100 and 10,500 feet altitude, are two 

other broad moraines on the surface of which boulders 

are fairly numerous. The till contains disintegrated 

and fresh stones in roughly equivalent proportions and 

bears a mature Brown Podzolic soil 3 to 4 feet thick. 

A typical soil profile consists of 6 inches of dark-brown 

(1.5YR 4/2) loose micaceous silty sand, 14 inches 

of dark-brown (7.517? 4/4) slightly plastic micaceous 

stony silty sand, and 20 inches of strong-brown (7.5 

TR 5/6) compact plastic micaceous silty sand having 

a fine angular blocky structure, beneath which is lightbrown (7.5J7? 6/3), compact stony sandy till. The 

lower moraine is the Wisconsin II of Ray (1940). 

The outer slopes of the upper moraine extend over a 

rock lip, but the inner slopes are clearly of depositional origin. Along the north fork, a pair of large 

lateral moraines of similar character on the north 

side of the valley indicate glacial termini at about 

10,400 and 10,500 feet, and at a distance of about iy2 

miles from the cirque heads. The mature topography, 

abundance of surface boulders, and soil profile on 

these two lateral moraines, and the fact that their 

terminal portions have been removed by erosion, suggest their correlation with the early and late stades 

of Bull Lake Glaciation.

Along the South Fork of the Rio Nambe, at altitudes of 10,750 and 11,100 feet, and on the north 

fork at 10,600 and 10,850 feet, are broad, high, steepfronted bouldery, hummocky end moraines that are 

but little dissected. The deposits are forested and 

have an immature Brown Podzolic soil that consists 

of 1 inch of black humus, iy2 inches of gray silty 

sand (A2 horizon), and 20 inches of brown (10T7? 

4/3), stony silty sand (B horizon), underlain by 

fresh pale-gray-brown till. The outer moraine on the 

south fork impounds a peat bog; the inner overlies a 

rock lip that retains a tarn. The surface character,

soil, and relative position of these deposits suggest 

their correlation with the early and middle stades of 

Pinedale Glaciation. The ice at the maximum of 

these advances extended about 1 mile from the cirque 

heads.

A single low forested moraine with an immature 

Brown Podzolic soil lies on the lip of the cirque 

basin of the North Fork of the Rio Nambe at an 

altitude of 11,250 feet. Within the cirque of the 

south fork at 11,400 feet is a forested, blocky deposit 

that has a similar soil and is probably a rock glacier. 

The character of these deposits suggests that they 

represent the late stade of Pinedale Glaciation.

The rock-glacierlike deposit on the south fork is 

overlapped by a distinctly younger lobate, blocky 

rock glacier at an altitude of 11,500 feet. The front 

of the deposit is 100 feet high, and covered with 

spruce. It has a dark-brown (IOTR 4/3) silty sandy 

azonal soil about 8 inches thick, and probably represents the Temple Lake Stade of Neoglaciation. Fresh 

talus presumably of Gannett Peak age lies at the foot 

of the cirque headwall. Along the northwest-facing 

sector of the cirque on the north fork, a well-developed 

protalus rampart that has a soil and vegetation characteristic of the Temple Lake Stade lies at the foot of 

active talus.

SIERRA BLANCA

From the shallow cirque on the northeast side of 

Sierra Blanca Peak (altitude 12,003 feet), a bouldery 

lateral moraine extends along the north side of the 

canyon of the North Fork of the Rio Ruidoso to a 

terminal moraine 50 feet high at an altitude of 9,850 

feet about a mile from the cirque headwall (fig. 213.3). 

The deposit is trenched about 30 to 40 feet by the creek, 

and bears a mature Brown Podzolic soil about 3 feet 

thick. The B horizon of the soil is strong-brown 

(7.5T7i? 5/6) and contains noticeable illuviated clay, 

though it is not plastic. The boulders are mostly of 

granite porphyry, and many, both on and in the till, 

are deeply disintegrated. A second broad terminal 

moraine of similar character lies at an altitude of 

10,400 feet. It also is about 50 feet high and is littered with boulders, some 10 to 12 feet in diameter. 

These two moraines are believed to represent the early 

and late stades of Bull Lake Glaciation. A small 

lake basin back of the upper moraine is filled with 

arkosic sand and gravel which extends upslope as an 

outwash plain into still another bouldery morainal 

ridge, about 30 feet high, which forms an arcuate loop 

on the steep slope leading into the cirque at an altitude of 10,500 feet. The moraine is grass covered
ARTICLE 213 E125

. Sierra Blanca

Peak

12,003 ft

03 )

VT. f > -

Z

' -. 

'

-

R. 11 E.

3 MILES

J

FIGURE 213.3. Moraines on Sierra Blanca Peak, Sierra 

Blanca, N. Mex. Be and Bl, early and late stades of Bull 

Lake Glaciation; Pe, Pm, and PI, early, middle, and late 

stades of Pinedale Glaciation.

except for a few spruce along its outer slope and 

lateral extensions. It is but little dissected, bears an 

immature Brown Podzolic soil about 10 inches thick, 

and is yellowish brown (10T7? 5/4). A similar moraine, nearly 100 feet high, lies across the slope at an 

altitude of 10,900 feet, and a third, about 60 feet high, 

at 11,300 feet. These three moraines are believed to 

represent the early, middle, and late stades of Pinedale 

Glaciation.

The cirque floor is between 11,400 and 11,500 feet 

in altitude. Talus along the cliffed headwall is stable 

and covered by soil and grass or spruce, and spruce 

extends as scattered clumps to the top of the peak. 

No deposits having the characteristics of the Temple 

Lake or Gannett Peak Stades of Neoglaciation were 

seen, and no snowbanks were present on July 10, 

195T.

A steep basinlike slope on the northwest side of the 

peak may have been occupied by ice during the Bull 

Lake Glaciation, but not in Pinedale time. The slope 

is steep and no moraines have been preserved.

REFERENCES

Antevs, Ernst, 1935, The occurrence of flint and extinct animals in pluvial deposits near Clovis, New Mexico, Pt. 2 

of Age of Clovis Lake clays: Philadelphia Acad. Nat. 

Sci. Proc., p. 304-312.

Blackwelder, Eliot, 1915, Post-Cretaceous history of the mountains of central western Wyoming: Jour. Geology, v. 23, 

p. 97-117, 103-217, 307-340.

 1931, Pleistocene glaciation in the Sierra Nevada and 

Basin Ranges: Geol. Soc. America Bull., v. 42, p. 865-922.

Ellis, R. W., 1931, The Red River lobe of the Moreno glacier: 

New Mexico Univ. Bull., v. 4, p. 1-26.

 - 1935, Glaciation in New Mexico: New Mexico Univ. 

Bull., Geol. Ser., v. 5, p. 1-31.

Fryxell, F. M., 1951, Grand Teton National Park: U.S. Geol. 

Survey topographic map. (Text on back)

Hack, J. T., 1943, Antiquity of the Finley Site: Am. Antiquity, 

v. 8, no. 3, p. 235-241.

Holmes, G. W. and Moss, J. H., 1955, Pleistocene geology of 

the southwestern Wind River Mountains, Wyoming: Geol. 

Soc. America Bull., v. 66, p. 629-654.

Moss, J. H., 1951a, Early man in the Eden Valley: Pennsylvania Univ. Mus. Mon., v. 6, 124 p.

 1951b, Late glacial advances in the southern Wind

River Mountains, Wyoming: Am. Jour. Sci., v. 249, no. 12,

p. 865-883. 

Nelson, R. L., 1954, Glacial geology of the Frying Pan River

drainage, Colorado: Jour. Geology, v. 62, p. 325-343. 

Ray, L. L., 1940, Glacial chronology of the southern Rocky

Mountains: Geol. Soc. America Bull., v. 51, p. 1851-1918. 

Richmond, G. M., 1948, Modification of Blackwelder's sequence

of Pleistocene chronology in the Wind River Mountains,

Wyoming [abs.] : Geol. Soc. America Bull., v. 59, p. 1400-

1401.

 1957a, Three pre-Wisconsin glacial stages in the Rocky 

Mountain region: Geol. Soc. America Bull., v. 68, p. 239- 

262.

 1957b, Correlation of Quaternary deposits in the Rocky 

Mountain region, U.S.A.: Internat. Congress of INQUA, 

5th, Madrid-Barcelona, 1957, Resumes des communications, 

p. 157.

1962, Three pre-Bull Lake tills in the Wind River

Mountains, Wyoming: Art. 159 in U.S. Geol. Survey

Prof. Paper 450-D, p. D132-D136. 

Salisbury, R- D-, 1901, Glacial work in the western Mountains

in 1901: Jour. Geology, v. 9, p. 718-731. 

Smith, H. T. U., 1936, Periglacial landslide topography of

Canjillon Divide, Rio Arriba County, New Mexico: Jour.

Geology, v. 44, p. 836-860. 

Smith, H. T. U., and Ray, L. L., 1941, Southernmost glaciated

peak in the United States: Science, v. 93, p. 209. 

Stone, G. H., 1901, Note on the extinct glaciers of Arizona and

New Mexico : Science, v. 14, p. 798.

667422 O - 63 - 10
E126 GEOLOGICAL SURVEY RESEARCH 1962 

ANALYTICAL METHODS

214. ION-EXCHANGE SEPARATION AND SPECTROPHOTOMETRIC DETERMINATION OF CADMIUM

By CLAUDE HUFFMAN, JR., Denver, Colo.

An anion-exchange resin procedure has been adapted 

for the separation of microgram amounts of cadmium 

from milligram amounts of zinc, lead, thallium, cobalt, nickel, bismuth, copper, ferric iron, and manganese before its spectrophotometric determination 

with dithizone. This separation is based in part on 

the anion-exchange resin work of Kraus and Nelson 

(1955) describing the absorption behavior of many 

metal chlorocomplexes, and also on the zinc procedure 

of Huffman, Lipp, and Rader (1963). Cadmium and 

zinc are absorbed on an anion-exchange resin (Dowex 

1, X-8, chloride form) from 1.2M hydrochloric acid. 

Zinc is eluted with O.OLfl/ hydrochloric acid while 

cadmium remains on the resin. In this investigation, 

it was determined that cadmium is quantitatively 

eluted with l.OM nitric acid. The cadmium content 

is then determined by processing an aliquot of the 

eluate through the last dithizone-extraction step described by Saltzman (1953).

The method is applied to the determination of a 

few tenths of a part per million or more of cadmium 

in mineralized silicate rock samples. Cadmium analyses are shown for a National Bureau of Standards 

standard zinc ore sample no. 113, standard diabase 

W-l, standard granite G-l and a rock sample containing sphalerite (S-4-A).

APPARATUS AND REAGENTS

Resin column. Column is 24 cm long, with an inside diameter 

of 0.8 cm. Prepare the resin column for use by adding 

a water slurry of anion-exchange resin (Dowex 1, X-8, 

50- to 100-mesh, chloride form) until the slurry is 10 cm 

in depth. Condition the resin before use by passing hydrochloric acid through the column as described in "Procedure." Regenerate the used resin by passing 100 ml 

water through the column, followed by 50 ml l.OM hydrochloric acid.

Hydrochloric acid, 1.2M and 0.01M.

Nitric acid, l.OJf.

Hydroxylamine hydrochloride, 20 g in 100 ml water.

Sodium hydroxide, 40-percent, potassium cyanide, 0.05-percent, 

400 g and 0.5 g respectively in 1 liter water.

Dithizone solution. 10 mg dithizone in 1 liter redistilled 

chloroform. Keep in refrigerator.

Tartaric acid, 20 g in 100 ml water.

Standard cadmium solution, 2 jig Cd per ml in 0.1M hydrochloric acid. To obtain this solution, dilute a stronger 

stock solution prepared from cadmium metal.

Spectrophotometer and paired 1-cm cells.

PROCEDURE

Transfer a 1.0-g sample to a platinum dish. Process a reagent blank with the samples. Add 10 ml demineralized water, 10 ml nitric acid, 10 ml hydrofluoric acid, and let stand overnight. Add 10 ml 

perchloric acid, and 5 ml sulfuric acid. Place dish 

on a hotplate and fume off the acids until about 3 ml 

remains. Cool. Wash down the sides of the dish 

with water, add 5 ml perchloric acid and repeat the 

fuming until all the acid is evaporated. Do not bake 

the salts. Cool. Add 10 ml hydrochloric acid and 

25 ml water to the dish, cover, and digest on a steam 

bath for 30 minutes. Transfer the solution to a 

100-ml volumetric flask, cool, dilute to volume with 

water, and mix.

Condition the resin column for use by passing 20 

ml 0.01J/ hydrochloric acid through it, followed by 

20 ml 1.2M hydrochloric acid. Kepeat this operation 

two or three times and leave the 1.2J/ hydrochloric 

acid in the column with about 2 ml of solution above 

the resin.

Quantitatively transfer an aliquot of the sample 

solution containing less than 50 ^g Cd to the resin 

column after diluting the aliquot volume to 50 ml 

with 1.2J/ hydrochloric acid. Regulate the flow rate 

of the solution through the column to about 1 ml per 

minute by adjusting the stopcock at the bottom of the 

column. When the flow stops, discard the solution 

that has passed through, wash the column with 50 ml 

1.2J/ hydrochloric acid, and discard the wash solution. Elute the zinc by passing 65 ml 0.01J/" hydrochloric acid through the resin at a flow rate of about 

1 ml per minute. Save this solution only if zinc analysis is required. Wash the resin with 10 ml water, 

then place a clean 100-ml volumetric flask under the 

resin column and elute the cadmium by passing 60 ml 

l.QM nitric acid through the column at a flow rate of 

1-ml per minute. Add 2 g tartaric acid to the flask 

and dilute to volume with water.

Transfer a 25-ml portion of the eluate solution containing less than 14 fig Cd to a 125-ml separatory funnel. Add 0.50 ml hydroxylamine hydrochloride 

solution, 5 ml sodium hydroxide-potassium cyanide 

solution, and 15 ml dithizone solution to the separatory 

funnel and immediately shake for 1 minute. Allow
ARTICLE 214 E127

the phases to separate and drain the chloroform into 

a 1-cm cell. Measure the absorbance at 518 nut, using 

the reagent blank as a reference solution.

To establish the standard spectrophotometric curve, 

add 2, 4, 8, and 12 /^g Cd to 25 ml 2-percent w/v 

tartaric acid solution and proceed, starting with the 

dithizone extraction step described in the preceding 

paragraph.

REMARKS

The amount of 1M nitric acid required to elute 50 

/Ag Cd from the resin column was determined. Elution tests with 15, 30, 45, and 60 ml 1M nitric acid 

showed recoveries of 48, 90, 99, and 99 percent, respectively. A plot of these data shows that cadmium is 

quantitatively eluted with about 43 ml of acid. Elution tests with equivalent amounts of 0.001 M hydrochloric acid gave poor recoveries.

Possible interference in the procedure by other 

metals that react with dithizone was investigated by 

processing milligram amounts of zinc, lead, thallium, 

cobalt, bismuth, copper, ferric iron, and manganese 

through the procedure. With the exception of zinc, 

these data (table 214.1) show little or no interference 

from the metals listed. The test with zinc, equivalent 

to 0.8 /Ag Cd, was found by X-ray fluorescence measurements to be due to cadmium impurity in the zinc 

sulfate salt used in the test. Analysis of NBS standard sample 113 (Tri-State concentrate zinc ore) conTABLE 214.1. Interference of milligram amounts of diverse 

metals on the cadmium determination after separation by union 

exchange. (No cadmium added in each test)

Metal added

Amount of interMilligrams of ference as micrometal added grams of cadmium

Fe(III).

3

5

2

4

5

4

5

150

5

»0. 8

< 1

. 2

< 1

< 1

. 1

. 3

< 1

< 1

taining 61.1 percent zinc was straightforward; no interference with the method was noted.

This method is compared in table 214.2 with other 

methods for the determination of cadmium in 2 zinc 

sulfide samples and 2 standard silicate rock samples, 

W-l and G-l. The agreement between methods is 

good for the zinc sulfide samples. Eecent values reported by Fleischer and Stevens (1962) for W-l and 

G-l are in good agreement with those obtained by this 

method. Earlier values cited by Ahrens and Fleischer 

(1960) for W-l and G-l are clearly in error. The 

cadmium values for W-l and G-l were obtained by 

processing duplicate 2.0-g portions, one spiked with 

20 /Ag Cd, through the procedure. The recovery of 

cadmium added to W-l and G-l was found to be 

better than 97 percent.

TABLE 214.2. Comparison of cadmium values determined by the 

method described and by other methods

Sample Rock type

Cadmium in percent

Method described Other methods

1 This interference was determined to be cadmium impurity in the zinc sulfate 

salt solution used in the test.

NBS 113_ Zincore____- 0.410.400.440.40 1 0. 41

S-4-A--__ Sphalerite,._ .16 .14 .16 .15 2. 14

W-l__-_- Diabase.-.__ <. 00003 3. 00003

G-l------ Granite-__-_ <. 00003 3. 000006

1 X-ray fluorescence, oral communication from J. J. Tregoning.

2 Controlled-potential electroanalysis, Charles Kinser, analyst. 

' Recommended magnitude values cited by Fleischer and Stevens (1962) based 

on polarographic, colorimetric, neutron-activation, and spectrogi aphic analysis.

REFERENCES

Ahrens, L. H., and Fleischer, M., 1960, Report on trace constituents in granite G-l and diabase W-l, in, Stevens, 

R. E., and others. Second report on a cooperative investigation of the composition of two silicate rocks: U.S. Geol. 

Survey Bull. 1113, p. 83-111.

Fleischer, Michael, and Stevens, R. E., 1962, Summary of new 

data on rock samples G-l and W-l: Geochimica et Cosmochimica Acta, v. 26, p. 525-543.

Huffman, C., Lipp, H. H., and Rader, L. F., 1963, Spectrophotometric determination of micro quantities of zinc in 

rocks: Geochimica et Cosmochimica Acta. (In press)

Kraus, K. A., and Nelson, F., 1955, Anion exchange studies of 

the fission products, in International Conference on the 

Peaceful Uses of Atomic Energy, v. 7, A conf. 8/P/837, 

U.S.A., 23 June. [In English]

Saltzman, B. E., 1953, Colorimetric microdetermination of cadmium with dithizone: Anal. Chem., v. 25, p. 493.
E128

215.

GEOLOGICAL SURVEY RESEARCH 1962

EFFECT OF COPPER ON THE PRECIPITATION OF TELLURIUM WITH HYPOPHOSPHORUS ACID USING

SELENIUM OR GOLD AS A COLLECTOR

By H. W. LAKIN and C. E. THOMPSON, Denver, Colo.

As a part of a more comprehensive study of analytical methods for the determination of traces of tellurium, tests were made of the completeness of reduction and precipitation of tellurium with hyphophosphorous acid from very dilute solutions of tellurium 

(IV) using selenium or gold as a collector.

A number of shale samples were digested with oxidizing acids and hydrofluoric acid; the solutions evaporated to dryness and the residues dissolved in 4N or 

6N hydrochloric acid. Tellurium-129 was added to 

these solutions together with selenium or gold. These 

solutions were reduced with either stannous chloride, 

sulfur dioxide, or hypophosphorous acid to determine 

the effectiveness of the reducing agents. Large excesses of the reductants failed in some cases to give 

any precipitate and in others gave incomplete precipitation of the tellurium and gold or selenium. Stannous chloride was the most satisfactory, but the precipitate was contaminated with much tin. In solutions 

of chalcopyrite, however, tellurium and gold or selenium were quite readily reduced to the element by 

hypophosphorous acid. In QN hydrochloric acid, ferric iron is not readily reduced to ferrous iron by sulfur dioxide or hypophosphorous acid. Gold (III) is 

not reduced to the element by hypophosphorous acid 

in &N hydrochloric acid containing ferric chloride; 

tellurium and selenium are incompletely reduced. 

However, cupric chloride is rapidly reduced by hypophosphorous acid in 4N to QN hydrochloric acid, and 

the cuprous ion readily reduces ferric iron to the ferrous state.

These observations led to the series of experiments 

outlined below to test the effectiveness of hypophosphorous acid as a reductant and selenium and gold 

as carriers for tellurium with and without added 

copper.

To 400 ml of QN hydrochloric acid were added the 

amounts of tellurium-129, copper as cupric chloride, 

and selenium noted in table 215.1; 20 ml of 50-percent 

hypophosphorous acid was added, the solutions were 

heated on the steam bath for 1 hour, and allowed to 

stand overnight. The solutions were filtered with 

suction through very fine sintered-glass filters. The 

filtrate was evaporated to near dryness and transferred to test tubes. The radioactivity of these solutions was measured in a scintillation counter using as 

a detector a well-type 1.5-inch thallium-activated sodium iodide crystal. Similar data are given for the

TABLE 215.1 Effect of copper on the precipitation of elemental 

tellurium from 6N hydrochloric acid using selenium as a carrier

Te129

added

(MB)

0. 5

. 5

1.0

1.0

Activity 1 of 

added Te 129 in

same volume

as filtrate

(counts per

10 min)

2,652

2,500

5,304

5,000

Se added

(mg)

2

4

2

4

Activity of filtrate 

(counts per 10 min)

NoCu

added

2, 135

870

4,747

3,276

2mg Cu

added

381

110

680

2 0

Te precipitated G*g)

NoCu

added

0. 10

.32

. 10

.65

2mg Cu

added

0.45

.48

.87

1.00

i Background of 3,576 counts subtracted from all values given in the table. 

1 Counts less than average background.

precipitation of tellurium from 47V hydrochloric acid 

in table 215.2.

TABLE 215.2 Effect of copper on the precipitation of elemental 

tellurium from 4N hydrochloric acid using selenium as a carrier

Tei»

added

(Mg)

0. 5

. 5

. 5

1.0

1.0

-1. 0

Activity ' of

added Te12* in

same volume

as filtrate

(counts per

10 min)

2,473

2,208

2,208

4,946

4,416

4,416

Se added

(mg)

4

10

10

4

10

10

Activity of filtrate

(counts per 10 min)

NoCu

added

502

943

2mg Cu

added

90

109

71

225

173

207

Te precipitated G*g)

NoCu

added

0.4

.81

2mg Cu

added

0.49

.47

.49

.96

Q8

.98

i Background of 3,576 counts subtracted from all values given in the table.

The data in tables 215.1 and 215.2 show that 90 to 

100 percent of the tellurium is precipitated from 4A7 

and 67V hydrochloric acid solutions containing only 1.2 

to 2.5 parts per billion (109 ) of tellurium. Both the 

concentration of selenium and the presence of copper 

in the solution affect the recovery of tellurium.

The effect of foreign ions on the precipitation of 

tellurium from a IN hydrochloric acid solution containing 25 ppb tellurium is illustrated in table 215.3. 

These data indicate that about 95 percent of the tellurium in complex solutions can be precipitated with 

selenium as a carrier in the presence of copper. In a 

similar series using 67V hydrochloric acid, the recovery ranged from 95.6 to 100 percent; the average loss 

in 13 experiments was 2.1 percent. In the absence of 

copper, no tellurium was precipitated from a 67V hydrochloric acid solution containing aluminum, iron, 

potassium, calcium, magnesium, and sodium.
ARTICLE 215 E129

TABLE 215.3. Completeness of precipitation of tellurium in the 

presence of foreign ions using selenium as a carrier

[Base solution: Each 400 ml of 4AT HC1 contains 0.7 g Al, 1 g K, 0.9 g Fe, 0.3 g Mg, 

0.7 g Ca, 0.4 g Na, 10 mg Se, 10 mg Cu, 10 Mg Te 129]

Elements added

Base solution+0.8 g Fe, 0.7 g 

Pb, 0.4 g Mn. ___________

Base solution + 0.4 g Cu, 0.2 g 

Zn

Base solution + 0.2 g Ni, 0.5 g 

Co_______________________

Base solution +0.7 g Hg, 0.4 g 

Sb,0.3g As_----_---------

Base solution +1 mg Au___

Activity ' of 

10 Mg Te128 in 

same volume as 

filtrate (counts 

per 10 min)

14, 763

14, 763

18, 453

14, 763

12, 303

18, 453

Activity 

of filtrate 

(counts 

per 10 

min)

639

679

6Q4

643

958

928

Te precipitated 

(«)

9.57

9. 54

9. 62

9. 57

9.21

9. 50

i Background of 3,587 counts substracted from all values given in the table.

By reduction with hypophosphorous acid, 99.6 percent of the tellurium was removed from solution by 

2 precipitations with 2 mg of gold from 300 ml of 

either 2/V or 4A" hydrochloric acid, each solution containing 1 g of iron, 0.4 g of copper and 5 //,g of tellurium-129. One precipitation removed 99.7 percent 

of the tellurium from QN hydrochloric acid.

TABLE 215.4. Precipitation of tellurium from 0N hydrochloric 

acid using gold as a carrier

Te'»

added

fog)

0.5

.5

1.0

1.0

5.0

Activity ' of

added Te»« hi

same volume

as filtrate

(counts per

10 min)

2,652

2,500

5,304

5,000

26, 519

Au added

(mg)

2

8

2

8

2

Activity of filtrate

(counts per 10 min.)

NoCu

added

440

128

738

122

7,525

2mg Cu

added

251

2 0

743

2 0

652

Te precipitated (Mg)

NoCu

added

0.42

.48

.82

. 98

3.6

2mgCu

added

0.45

.50

. 82

1.00

4.9

The data in table 215.4 show the efficiency of gold 

as a carrier for tellurium. The tellurium-129 and gold 

were precipitated from 400 ml of 6A" hydrochloric 

acid by addition of 20 ml of 50-percent hypophosphorous acid*and heating at 80°-90°C. for 1 hour. 

The filtration was made the following day. The filtrates were evaporated to a small volume and the 

radioactivity of the tracer tellurium-129 remaining 

in the filtrate was compared to reference solutions of 

known concentration.

Precipitation of tellurium was incomplete from solutions containing 5 parts per million of gold (2 mg 

gold dissolved in 400 ml solution) but was complete 

for those containing 20 ppm of gold. Even in these 

solutions to which no iron was added, copper appears 

to enhance the precipitation of tellurium.

The effect of foreign ions on this precipitation is 

illustrated by the data in table 215.5. These data indicate that gold is a somewhat better carrier for tellurium than selenium is.

TABLE 215.5. Completeness of precipitation of tellurium in the 

presence of foreign ions using gold as a carrier

[Base solution: Each 400 ml of 6ATHC1 contains 0.7 g Al, 1 g K, 0.9 g Fe, 0.3 g Mg, 

0.7 g Ca, 0.4 g Na, 8 mg Au, 10 mg Cu, 10 Mg Te 126]

Elements added

Base solution + 1 g Fe+1 g 

Pb+1 g Mn. _____________

Base solution+ 1 g Cu+ 1 g Zn_ 

Base solution-!- 1 g Ni+ 1 g Co_ 

Base solution + 1 g Sb+1.6 g 

Bi-_-__________ __________

Base solution +1 g Hg + 0.9 g

As --------------------

Activity » of 

10 ME Te'29 in 

same volume as 

filtrate (counts 

per 10 min)

14, 880

6,200

12, 400 

12, 400

12, 400

10, 630

Activity 

of filtrate 

(counts 

per 10 

min)

196

2 00

241 

138

248

192

Te precipitated 

G«)

9.87

10.00

9.81 

9. 89

9.80

9. 82

1 Background of 3,687 counts subtracted from all values given in the table. 

* The activity in counts per 10 min was less than the average background.

' Background of 3,368 counts subtracted from all values given in the table. 

2 The activity in counts per 10 min was less than the average background.
E130 GEOLOGICAL SURVEY RESEARCH 1962

216. SPECTROPHOTOMETRIC DETERMINATION OF FLUORINE WITH THORON 

By BLANCHE L. INGRAM, Washington, D.C.

The method of Grimaldi and others (1955) for the 

spectrophotometric determination of fluorine with 

thoron has been modified to permit absorbance measurements with 5-cm cells, giving increased reliability 

in the 0-20 jug range. Also discussed is a problem 

which may arise with blank determinations in the use 

of unacidified thorium nitrate reagent.'

REAGENTS AND APPARATUS

Standard fluoride solution. 1 ml contains 1 fig fluoride. Prepare from pure sodium fluoride.

Thorium nitrate solution. 1 ml contains the equivalent of 

25 fig thorium dioxide and 0.1 ml perchloric acid.

Thoron solution. o-(2-Hydroxy-3,6-disulfo-l-naphthyazo) benzenearsonic acid disodium salt, a 0.10 percent aqueous 

solution.

Perchloric acid, 70 percent.

Sodium hydroxide solution, 2 percent aqueous solution.

Phenolphthalein, 0.1 percent alcoholic solution.

Spectrophotometer, Beckman DU, supplied with 5-cm cells.

PROCEDURE

Prepare a distillate of the sample as directed in 

Grimaldi and others (1955). Transfer a 15-20 ml 

aliquot of the distillate to a 25-ml volumetric flask. 

Pipet 1-, 3-, 5-, 10-, and 20-ml aliquots of the fluoride 

standard into 25-ml volumetric flasks. Designate a 

25-ml flask as 0 /xg F and another as a reference solution. Add sufficient distilled water to all flasks to give 

a volume of about 15 to 20 ml. Add 1 drop of phenolphthalein to each and then sodium hydroxide dropwise until the indicator changes to pink. Then add, 

by means of pipets or burets, to all flasks except the 

reference, 1 ml hydroxylamine hydrochloride, 2.00 ml 

thorium nitrate, and 1.00 ml thoron reagent. To the 

reference flask, add 0.2 ml HCKX, 1 ml hydroxylamine hydrochloride, and 1.00 ml thoron. Adjust all 

volumes to 25 ml with distilled water and mix. Measure the absorbance of the solutions at 545 niju, after 15 

minutes using a Beckman DU spectrophotometer and 

5-cm cells. A straight-line relation should be obtained 

between absorbance and fluoride concentration. One 

microgram of fluoride decreases the absorbance of the 

thorium-thoron color by about 0.008.

PROBLEM WITH BLANK DETERMINATIONS

The spectrophotometric method of Grimaldi and 

others gave absorbance readings for 0 /*g standards 

which were often too low. Zero standards and blank

distillates seemed to contain fluorine that could not be 

explained. These difficulties were traced to the use 

of old thorium nitrate reagent. For best results, the 

reagent referred to as "thorium nitrate solution for 

spectrophotometry, 1 ml contains 50 jug of thorium 

dioxide," should be prepared fresh when required. 

Alternatively, it can be prepared containing 0.1 ml 

HCICX per milliliter, in which case the addition of 

0.2 ml HCKX is omitted in the procedure. The acidified solution is stable. Absorbance readings at several 

fluorine concentrations, using thorium nitrate reagent 

of varied age with and without acid are given in the 

table below.

Accurate pipetting can not be overemphasized in 

this method. Poor pipetting techniques can be a 

source of error that is very apparent in blank determinations. As examples, tests using 5-cm cells showed 

that the absorbance of a 0 jug standard with 0.05 ml 

thorium nitrate reagent in excess was 0.015 higher, 

and one with 0.05 ml thoron in excess was 0.019 

higher than one without excess of either reagent. 

Pipetting errors may be kept to a minimum by mixing the reagents in appropriate ratios just before addition and adding all as one reagent with one pipetting operation.

Absorbance readings using thorium nitrate reagent with and without

acid 1

F in 25 ml 

G*g)

0.. ............

20.. .. .........

40.. .......... .

60-- 

80. 

100.... - 

Blank distilThorium nitrate reagent without 

acid

Fresh

0.475 

.440 

.403 

.364 

.326

.475

1 

week 

old

0.463 

.437 

.393 

.354 

.314 

.278

.465

2 

weeks 

old

0.462 

.437 

.398 

.357 

.311 

.281

.464

4 

weeks 

old

0.462 

.435 

.398 

.358 

.318 

.284

.468

1 year 

old a

0.383 

.416 

.382 

.346 

.307

.418

Thorium nitrate reagent 

containing 0.1 ml HClOi 

per ml

Fresh

0.479 

.447 

.406 

.362 

.326 

.289

.479

1

week 

old

0.479 

.443 

.404 

.364 

.325 

.291

.478

2

weeks 

old

0.479 

.445 

.407 

.364 

.324 

.289

.479

4 

weeks 

old

0.479 

.444 

.408 

.366 

.326 

.288

.479

i2-cm cells used for all measurements. 

2 Not same Th(NOs)4 reagent.

REFERENCE

Grimaldi, F. S., Ingram, B. L., and Cuttitta, Frank, 1955, 

Determination of small and large amounts of fluorine in 

rocks: Anal. Chemistry, v. 27, p. 918-921.
ARTICLE 217

TOPOGRAPHIC STUDIES

E131

MAPPING TECHNIQUES

217. PORTABLE SURVEYING TOWER 

By JAMES L. BUCKMASTER and WILBTJRN D. MTJRPHY, Washington, D.C.

With the introduction of electronic distance-measuring instruments into control surveys in 1956, water 

bodies, vegetation, and rough terrain were no longer 

a handicap to accurate distance measurements provided stations were intervisible. Older reconnaissance 

methods such as observing the terrain from the top 

of a truck, from fire towers, or by climbing trees, were 

unsatisfactory. Commercial platforms that operated 

hydraulically were available, but these were costly and 

could serve only for reconnaissance. Standard triangulation towers were heavy, cumbersome, and required a trained crew to transport and erect them. In 

order to plan a control project and select suitable 

station sites, a portable tower was needed that could 

be erected quickly.

A lightweight aluminum tower designed by the 

U.S. Geological Survey combines the qualities of 

portability and stability, and it can be easily and 

quickly erected (fig. 217.1). The number of parts in 

the tower is kept to a minimum by welding each horizontal triangular brace assembly into a one-piece unit, 

and by making the tower sections similar. A flared 

base section is used for better stability at heights 

greater than 40 feet.

The triangular brace assemblies are vertically 

spaced 6 feet apart while the corner tubes are either 

6 or 12 feet long, permitting separate assembly of 

6- and 12-foot tower sections. Each 6-foot section or 

panel of the outer tower, defined by two corner tubes 

and two horizontal members, is diagonally crossbraced 

by %-inch aircraft cable, giving a maximum strengthweight ratio. An aluminum ladder leads to a working platform which has a waist-high railing and a 

nylon netting between the railing and the floor.

For reconnaissance work, only the outer tower is 

used. For measuring distances and angles, however, 

an inner tower is used to provide greater stability and 

rigidity. The construction of the inner tower is similar to that of the outer tower, except that each 6-foot 

section is braced with a diagonal tube that serves as 

both a tension and a compression member.

For transporting the tower, a special trailer unit, 

witji accessory box, can be easily attached between 

any two adjacent horizontal brace members (fig.

217.1). The trailer-mounted tower can be towed by 

field truck or station wagon. For most ground transportation, tower height is limited to 50 feet. It is expected that a helicopter can transport a fully assembled tower, in heights to 75 feet, where stations are 

inaccessible to ground vehicles.

A 50-foot tower can be erected by two experienced 

men in about 2 hours, using a hand winch and raising 

boom. Normally it is raised, lowered, and transported 

fully assembled. In the raising operation, the tower 

pivots about two lower feet. An imaginary line passing through these two points is called the hinge axis. 

Two lateral cables are attached to the top of the 

tower and fastened to anchors on opposite sides of the 

tower along the hinge axis. The raising boom is positioned vertically over the base of the tower. The 

boom is guyed to the same two anchors that hold the 

tower in lateral position, to restrain the tower and 

boom from side movement, but allow them to rotate 

about the hinge axis. The raising cable is attached 

to the top of the tower and passed over the raising 

boom to a winch anchored approximately 50 feet beyond the tower base. As the cable is winched in, the 

tower rises in an arc about its hinge axis. The process is reversed in lowering the tower.

When in position over a station, the tower is leveled 

with adjustable foot screws which rest on foot plates. 

The outer structure is guyed with aircraft cables fastened to screw-in type earth anchors or natural objects 

such as trees. The inner unit is held to the ground 

with earth anchors and a cable sling.

A tower may be mounted on a truck for reconnaissance (fig. 217.2). A 50-foot section of tower rests 

horizontally above the cab and bed and is erected at 

the rear of the truck by a hand winch and cable system. Lateral outriggers are extended to stabilize the 

truck and tower before erection. With this system, 

one man can erect the tower in 15 minutes. Height 

is limited more by the truck's ability to traverse woods 

roads than by strength of the tower. Because of the 

ease and rapidity with which it can be erected, truckmounted towers are being designed to be erected in 

heights up to 70 feet.
E132 GEOLOGICAL SURVEY RESEARCH 1962

FIGUEE 217.2. Fifty-foot tower mounted on a truck.

FIGTJEE 217.1. Assembled tower, showing inner and outer 

tower; trailer unit is attached.
ARTICLE 218 

218. AIRBORNE CONTROL SYSTEM

By JAMES L. BUCKMASTER, HUGH B. LOVING, and THOMAS O. DANDO, Washington, D.C.

E133

Investigations are being conducted to devise methods of executing control surveys from an airborne 

platform, a radical departure from conventional surveying techniques. In the airborne control system 

(ABC) a helicopter serves as an aerial platform above 

points for which geodetic positions are to be obtained. 

In position, the craft serves as a target for angle 

measurements from one or more ground stations and 

as the "remote" station for the electronic distancemeasuring equipment.

In the system, as sketched in figure 218.1, the two 

existing ground-control stations are occupied by a 

theodolite observer, an operator for the electronic distance-measuring instrument, and a recorder. If additional height is required at the ground stations, 

portable surveying towers can be brought in by helicopter, fully assembled, or can be towed in by vehicle.

Since angle and distance measurements from one of 

the ground stations determine the position of the 

station sought, and a like set of measurements from 

the other station provides a check, it is not mandatory 

that the ground control be positioned to produce 

strong intersection angles as in conventional triangulation. The ground stations therefore can be selected 

with considerable freedom.

In addition to the pilot, the airborne team includes 

an engineer to operate the distance-measuring instrument to determine helicopter height and to verify the 

identification of the ground points. The pilot must 

maintain a hovering attitude while the angle and distance measurements are being made.

The helicopter serves as an aerial platform over the 

station on which a geodetic position is desired. The 

installation in the helicopter consists of one unit of 

the electronic distance-measuring instrument, a hoversight, a height indicator, and a rotating beacon.

The hoversight (fig. 218.2) consists of a siliconedampened pendulum which has a self-contained light 

source for projecting a beam of collimated light 

through a semitransparent mirror. With the aid of 

another mirror the observed views an erect image of 

the terrain below and the image of the light source. 

The image of the filament projected against the terrain image indicates the nadir, thus defining the vertical. The hoversight enables the pilot to position the 

craft above the point on which geodetic control is 

desired.

The height indicator is a device that enables the 

pilot to maintain a specified height above the ground. 

It contains a reel of line with a weighted end that is 

lowered to the ground. As the line is payed out it 

passes around a calibrated drum which is geared to a 

dial counter graduated in feet. Contact with the 

ground is maintained by a motor-driven fluid slip

FIGURE 218.1. The airborne control system.
E134 GEOLOGICAL SURVEY RESEARCH 1962

FRONT-SURFACE 

MIRROR

TO NADIR

COLLIMATOR

FIGURE 218.2. Diagram of hoversight.

clutch attached to the reel. Static tests show the unit 

to be accurate within 0.1 foot in 35 feet.

The rotating beacon is mounted on the exterior of 

the helicopter cabin on the vertical axis of the hoversight. It is the target on which angles are observed 

from the ground stations. Theodolites are used at 

both ground stations for the observation of both vertical and horizontal angles.

Measurement procedures are started and stopped 

by the engineer in the helicopter. Gusty winds may 

cause the craft to wander from the point, making it 

necessary to disrupt the measurements until the pilot 

regains his position.

Current practice is to turn 6 direct and 6 reverse 

horizontal angles, and 1 to 4 sets of vertical angles. 

Generally, as much time is required for the distance 

determination as is required for turning angles.

The number of vertical angles required depends on 

their intended use. If the vertical angles are to be 

used for slope reductions alone, one set will suffice 

(no special effort is made to hold the craft at a fixed 

height) and these can be measured at the same time 

the horizontal angles are turned. Vertical angles for 

supplemental control elevations must be more precise; 

consequently, they are repeated more times and read 

as an independent function while the craft is being 

held at a known height.

For each distance determination the operator performs a full measuring sequence consisting of an 

initial coarse reading, repeated fine readings, and a 

final coarse reading. The mean fine readings are used 

for refining the distance. Meteorological data are 

taken at the ground and air stations; on the basis of 

these data a correction for the index of refraction is 

made, to the distance.

Limitations of the present installation prevent making simultaneous distance determinations to the ground 

stations; consequently, one ground crew makes both 

distance and angle determinations while the other 

measures only angles. Upon completion of this set 

of measurements, the procedure is reversed.

Different combinations of angles and distances afford alternate methods of position computations such 

as triangulation, trilateration, or a combination of the 

two. The choice of computing methods and the number and kinds of observations made depend on the desired accuracy and the check requirements. Where 

an unchecked position will suffice, angles and distances 

need be measured from only one station, and a hovering 

period of about 6 minutes over the point is ample 

time to produce results equivalent to those from a spur 

traverse-line location or a strong intersected point.

Three separate tests have been conducted wherein 

existing second-order triangulation was repositioned 

using the ABC system. A comparison of the results 

of the last ABC measurements with the triangulated 

control is shown in the table below. As indicated in

Comparison of ABC tests with triangulation 

[Tests made Sept. 16-18,1961]

Remote helicopter

station, using

hoversight

Hover

height

(feet)

Triangulated

geodetic

distance

(meters)

Uncertainty of

control l

(meters)

Geodetic

distance

based on

ABC

(meters)

Difference in

geodetic

distance

(meters)

Proportional

difference 2

Ground station Krech

Gibbs 

Culvers Gap 

(RM2)..._ 

Normanock LOT .

70

180

75

72

30

122 

50

9922. 38

5637.94

6714. 94

14704. 39

20615. 93

23009. 74 

27659. 01

±0.50

±.28

±.34

±.74

±1.03

±1.15 

±1.38

9922. 63

5638. 06

6714. 45

14703. 60

20615. 32

23008. 55 

27658. 76

+0.25

+ 19

-.49

-.79

-.61

-1.19 

-.25

1:39,689

1:46,983

1:13,704

1:18,613

1:33,795

1:19,336 

1:110,636

Ground station Andover eccentric tower

Gibbs 

Culvers Gap 

(RM2). .... .

Normanock LOT...

42

62

55

62

70

82 

50

7705. 51

12101. 76

7048. 45

15212. 96

21048.87

22418. 68 

25803. 44

±0.38

±.60

±.35

±.76

±1.05

±1.12 

±1.29

7705. 36

12101. 86

.7048. 31

15212. 23

21048. 52

22417. 82 

25801. 72

-0.15

+.10

-.14

-.73

-.35

-.86 

-1.72

1:51,370

1:121,017

1:50,346

1:20,840

1:60,140

1:26,068 

1:15,002

1 Based on 1:20,000 accuracy.

2 Based on triangulated geodetic distance.

the last column, only four ABC measurements differ 

from the triangulated distances by more than 1:20,000. 

Three of these lines were measured in earlier tests of 

the ABC system. The distances obtained in the earlier
ARTICLE 218 E135

tests agree with the distances obtained in the last test to 

at least 1:50,000; accordingly there is no reason to 

doubt that second-order triangulation is sufficiently 

accurate to be used as a comparative standard. An 

additional test is scheduled using a more accurate 

control net.

In the ABC system, accuracy is largely a function 

of hover height and the number of observations. 

Accuracy falls off almost linearly with increasing

hover height. At present, a 200-foot hover height 

seems to be a practical limit, but tests are being conducted to develop a device which will greatly increase 

the practical hover height. The system shows much 

promise as a practical method of extending basic 

horizontal and mapping control. In its present state 

it should be competitive with conventional methods 

and should afford substantial savings in areas of 

limited access.
E136 GEOLOGICAL SURVEY RESEARCH 1962

HYDROLOGIC STUDIES

ENGINEERING HYDROLOGY

219. EFFECTS OF MAJOR WATER-TABLE CHANGES IN KINGS AND QUEENS COUNTIES, NEW YORK CITY

By N. M. PERLMUTTER and JULIAN SOREN, Mineola, N.Y.

Work done in cooperation with New York State Water Resources Commission

Kings and Queens Counties, 2 of New York City's 

5 political subdivisions, occupy about 189 square miles 

of western Long Island (fig. 219.1). Consumption of 

water in 1961 by the 4.5 million residents and by industry totalled about 640 mgd (million gallons per 

day), of which 93 mgd was pumped from local wells 

(see table) and the remainder was imported from 

surface reservoirs owned by the city of New York. 

Public supplies in Kings County and in northern and 

southern Queens County are entirely imported from 

surface sources. However, two private companies supply ground water to about 740,000 people in central 

Queens, and about 800 industrial and commercial wells 

provide water for cooling and air conditioning in 

both counties.

The ground-water reservoir ranges in thickness from 

about 50 to 800 feet and consists of a shallow unconfined aquifer and two deep confined (artesian) aquifers. The shallow aquifer is composed of outwasli deposits of Pleistocene age, and the confined aquifers are

Ground-water pumpage, in millions of gallons per day, in 1961 '

County

Kings ______

Queens _ ___ _

Total__ __ __________

Public 

supply

0

53. 9

53. 9

Industrial 

and commercial

22. 5

16. 9

39.4

Gross

22. 5

70. 8

93. 3

Nets

9. 1

61.0

70. 1

1 Data from New York State Water Resources Commission.

2 Net pumpage equals gross pumpage minus amount returned through diffusion 

(recharge) wells.

partly deposits of Pleistocene age and partly deposits 

of Late Cretaceous age. Crystalline bedrock of Precambrian (?) age forms the lower boundary of the 

zone of saturation. The aquifers are separated by 

leaky and discontinuous beds of silt, sandy clay, and 

clay; hence, they are poorly to well connected hydraulically. Salt water, which surrounds the area on

three sides, has invaded all the beds at the shoreline, 

and in some places has moved inland as much as 3 

miles (Lusczynski, 1952, fig. 2; Perlmutter and others, 

1959, fig. 4).

This article is concerned with fluctuations of the 

water table in the unconfined aquifer in the upper 50 

to 150 feet of the zone of saturation. Maps of the 

water table have been published by Veatch and others 

(1906), Suter (1937), Jacob (3945), Lusczynski and 

Johnson (1951), and Lusczynski (1952). The latest 

map (as of 1961) is shown in figure IB.

In 1903, Kings and Queens Counties were supplied 

with ground water by about a dozen private water 

companies, numerous wells and ponds owned'by the 

city of New York, and some domestic and industrial 

wells. Most of the water used was discharged to 

tidewater through public sewers, but in less urbanized 

sections much of the water was returned to the shallow 

aquifer through cesspools and septic tanks. Figures 

219.L4. and C, and D show contours and profiles of 

the water table in 1903 under conditions which may 

reflect some pumping but which are as close to natural 

as can be depicted from early data. The highest 

altitudes of the water table were about 20 feet in 

northeastern Kings County and about 60 feet in northeastern Queens County. Most of the ground water 

discharged north and northwest into the East River 

and south and southeast into Jamaica Bay and the 

Atlantic Ocean. The aquifers contained mostly fresh 

water in 1903, except in some places near the shore 

where salt water was present naturally or had moved 

toward wells near the shore.

By 1933, nearly all the ground water used for public supply and a considerable amount of self-supplied 

industrial pumpage was wasted to sewers; a deep 

extensive cone of depression had developed on the 

water table; and some public-supply wells in Flatbush, 

in central Kings County, had been contaminated by 

sea-water encroachment. These conditions, and the 

increasing use of ground water for air-conditioning,
74° 00' 73"45' 74° 00'

40° 

37' 

30"

Woodhaven , Land surface Jam

^__ 1903 _

FIGURE 219.1. Contours of the water table in Kings and Queens Counties, Long Island, N.Y., in 1903 (A) and 1961 (B), and profile

County (C) and Queens County (D). Altitude of land surface and water levels in feet referred to mean sea level; contour inte

in 1903 after Wiggin in Suter (1937). Stippled area is salty water.
E138 GEOLOGICAL SURVEY RESEARCH 1962

resulted in the passage of legislation that required 

the approval of the New York State Water Resources 

Commission for the drilling of all wells, except agricultural, yielding more than 70 gpm (gallons per 

minute). Under this legislation and its amendments, 

the commission generally ruled that water pumped at 

rates in excess of 70 gpm (45 gpm since 1954) for 

air-conditioning and cooling must be returned to the 

source aquifer.

By 1936, despite conservation measures, the water 

table had declined to a record low and was below sea 

level in a large cone of depression which included 

most of Kings County and the adjoining part of 

southwestern Queens County (Suter, 1937, fig. 26). 

The deepest point in the cone was about 35 feet below 

sea level in the Williamsburg section of northern Kings 

County (fig. 219.W). The depression of the water 

table in Kings resulted in a decrease in the rate of 

downward leakage to the deep confined aquifers and 

in an increase in ground-water underflow from Queens. 

From computations based on the volume of the cone 

of depression and an estimated specific yield of 25 

percent for the glacial-outwash deposits, it is calculated that at least 49 billion gallons of fresh water 

was removed from storage in the shallow aquifer in 

Kings County between 1903 and 1936. Probably two or 

three times that amount had been "mined," as nearly 

all the fresh water in the shallow aquifer had been 

replaced or contaminated by sea water. The decline 

of the water table was due to excessive pumping from 

both shallow and deep aquifers, to wasting to sewers of 

all water from public-supply and some industrial wells, 

and to a substantial decrease in natural recharge owing 

to extensive paving of streets and building construction.

In June 1947 the wells of a private water company, 

which had supplied as much as 27 mgd in central 

Kings County, were shut down as a result of condemnation proceedings by the city of New York, and 

a sharp rise of the water table began. Between 1947 

and 1950, the ground water in storage in the watertable aquifer increased by about 20 billion gallons 

(Lusczynski, 1952, p. 10). However, a significant part 

of this increment probably consisted of salt water that 

continued to move toward the center of the cone of 

depression after the cessation of pumping. By December 1961 the water table in northern Kings County 

had recovered as much as 40 feet, and it was above 

sea level in most of the county except for a small area 

in the northern part (fig. 219.15). It is estimated

that a net gain in storage of about 18 billion gallons 

was made between 1950 and 1961.

The recovery of the water table created seepage 

problems in parts of the subway system and in inadequately waterproofed basements in buildings constructed during the period of low water table. Seepage at several subway stations in Flatbush increased 

from less than 20 gpm in 1947 to as much as 1,000 

gpm in 1961, according to records of the New York 

City Transit Authority. The use of diffusion wells 

to return used water from air-conditioning systems has 

been discontinued in some places where local recharge 

mounds on the rising water table have reached basement levels. The principal remedy used to date to 

overcome local seepage problems has been the operation of sump pumps that discharge the water into 

public sewers. Other remedies such as tile fields and 

rings of dewatering wells generally have been regarded as too costly to install and maintain.

Hydraulic gradients have been reversed in most of 

Brooklyn from landward to the normal seaward direction, and a slow flushing of salt water from the 

aquifers is in progress. However, salty water containing several hundred to as much as 14,000 ppm 

(parts per million) of chloride still occupies parts of 

the unconfined aquifer in shoreline areas, and water 

containing as much as several thousand parts per 

million of chloride occurs in the confined aquifers beneath Flatbush despite the 15-year shutdown of the 

public-supply wells and the increased rate of recharge 

resulting from the rise of the water table.

In contrast, the water table in most of Queens 

County declined during the period 1950-61, and the 

cone of depression noted as early as 1933 was greatly 

enlarged. In 1961, water levels in the center of the 

cone in the Woodhaven-Jamaica area were about 15 

feet below sea level. Comparison of figures 219.L4 

and B suggests that the water table has declined as 

much as 20 feet since 1903 along the central part of 

the Queens-Nassau boundary. It is estimated that 

underflow from Nassau to Queens increased from a 

negligible amount in 1903 to about 10 to 13 mgd in 

1961. Underflow from Queens to Kings was reduced 

from a substantial amount in 1935 to a negligible 

quantity in 1961. Some ground water now flows from 

Kings County toward the Woodhaven area in Queens 

County (fig. 219.15). Calculations based on the net 

decline of the water table and a specific yield of 25 

percent indicate that about 51 billion gallons of fresh 

water was removed from storage in the water-table 

aquifer in Queens County between 1903 and 1961.
ARTICLE 219 E139

A continued slow rise of the water table in Kings 

County may result in the eventual filling of the small 

residual cone of depression in the northern part of 

the county. An increase of about 11 billion gallons of 

ground water in storage in the shallow aquifer and 

an additional 2 to 10 feet of recovery would be required to restore the water table to its 1903 level. 

Because of changes in hydrologic conditions, the 1903 

level may never be restored, and full restoration would 

not be desirable in any case.

The original average rate of natural recharge of 1 

mgd per square mile probably has been reduced to 

about 0.25 to 0.5 mgd per square mile because of 

building construction, paving, and sewering. Eecharge with fresh water and rising water levels have 

resulted in some flushing of salt water from the 

aquifers, but another 30 to 40 years may be required 

to flush out the remaining salt water that intruded 

the aquifers. Some deepening of the water-table depression in Queens County may be anticipated. Sea 

water will continue to move landward so long as the 

depression is below sea level.

The aquifers in Kings and Queens Counties constitute a valuable emergency source of water for civil 

defense. Probably 20 mgd of fresh water could be 

pumped in Kings and 60 mgd in Queens for many 

years. Larger amounts could be pumped temporarily 

in both counties for periods of several months to a 

year without serious damage. However, continued

pumping at much higher rates would result in eventual deterioration of the quality of most of the water 

owing to sea-water encroachment.

The case histories of overdevelopment in Kings and 

Queens Counties and the cost in terms of declining 

water levels and deterioration in water quality are 

lessons that should be heeded in planning large-scale 

development of water supplies in Nassau and Suffolk 

Counties on Long Island and in other coastal areas 

where similar hydrologic conditions prevail.

REFERENCES

Jacob, C. E., 1945, The water table in the western and central 

parts of Long Island, New York: N.Y. State Water Power 

and Control Comm. Bull. GW-12.

Lusczynski, N. J., 1952, The recovery of ground-water levels 

in Brooklyn, New York, from 1947 to 1950: U.S. Geol. 

Survey Circ. 167.

Lusczynski, N. J., and Johnson, A. H., 1951, The water table 

in Long Island, N.Y., in January 1951: New York State 

Water Power and Control Comm. Bull. GW-27.

Perlmutter, N. M., Geraghty, J. J., and Upson, J. E., 1959, The 

relation between fresh and salty ground water in southern 

Nassau and southeastern Queens Counties, Long Island, 

New York: Econ. Geology, v. 54, p. 416-435.

Suter, Russell, 1937, Engineering report on the water supplies 

of Long Island: New York State Water Power and Control Comm. Bull. GW-2.

Veatch, A. C., Slichter, C. S., Bowman, Isaiah, Crosby, W. O., 

and Horton, R. E., 1906, Underground water resources 

of Long Island, New York: U.S. Geol. Survey Prof. 

Paper-44.
E140

220.

GEOLOGICAL SURVEY RESEARCH 1962

INFILTRATION RATES IN WEATHERED CRYSTALLINE ROCKS AT THE GEORGIA NUCLEAR LABORATORY,

DAWSON COUNTY, GEORGIA

By J. W. STEWART, Atlanta, Ga.

Work done in cooperation with the U.S. Atomic Energy Commission and the U.S. Air Force

The rate at which water infiltrates the weathered 

crystalline rocks at the Georgia Nuclear Laboratory 

was determined at 2 waste-disposal pits by means of 

18- and 24-inch single-ring infiltrometers. The rings 

were driven into the saprolite to a depth of 6 inches, 

and a constant water depth of 6 inches was maintained 

in each ring. A total of 8 tests were made at 6 locations in the REL (Radiation Effects Laboratory) 

disposal pit, and a test was made at each of 3 locations in REF (Radiation Effects Facility) disposal 

pit 2. Of the 11 tests, 6 were made with 24-inch 

rings and 5 were made with 18-inch rings. With the 

exception of 1 test, the results obtained with the 24- 

inch ring were consistently higher than those obtained 

with the 18-inch ring.

Both disposal pits were constructed in saprolite 

and average about 4 to 5 feet in depth. The pits are 

trapezoidal in profile and the slope of the sides is 3:5. 

The bottom dimensions of the REL pit are 60 by 25 

feet and those of REF pit 2 are 50 by 15 feet.

The REL pit is near the top of a small ridge in 

the southeastern corner of the Georgia Nuclear Laboratory about 1.1 miles southeast of REF pit 2. The 

altitude of the land surface at the pit is about 1,112 

feet above mean sea level. Quartz-mica schist, garnetbiotite schist, and amphibolite schist underlie the pit, 

and a small pegmatite dike is exposed in the north 

corner of the pit. The rocks are deeply weathered; 

in a test well near the edge of the pit the saprolite is 

about 135 feet thick, or about 2y2 times thicker than 

that at REF pit 2. The foliation strikes northeast 

and dips about 73° SE.

REF pit 2 is on the top of the highest hill in the 

north-central part of the Georgia Nuclear Laboratory. 

The topographic relief in the vicinity of this pit is 

about 200 to 300 feet, and the altitude of the land 

surface at the pit is 1,245 feet, or about 133 feet 

higher than at the REL pit. REF pit 2 is underlain 

by quartz-biotite schist, biotite-garnet schist, and garnet-biotite-plagioclase schist. The saprolite at this 

pit is about 55 feet thick but thins rapidly in all 

directions; at distances of 50 feet both northwest and 

northeast of the pit, it is about 12 to 15 feet thick. 

The foliation strikes northeast and dips about 68° SE.

In the REL pit the material tested was largely 

clay and silt, mica-quartz-garnet schist saprolite, and

kaolinized feldspar; in REF pit 2 the material was 

biotite-garnet schist saprolite and biotite schist saprolite. (See fig. 220.1.) Although jointing was evident 

at all but one of the test locations, the joints were 

filled and apparently did not affect the infiltration rate. 

The infiltration rate for the 8 tests in the REL pit 

ranged from 0.06 to 1.24 in. per hr. (inches per hour), 

and for the 3 tests in REF pit 2 ranged from 0.96 to 

1.94 in. per hr. The greatest infiltration rate was obtained for biotite-garnet schist saprolite in REF pit 2 

(location 18-4), and the smallest infiltration rate was 

obtained for mica-quartz-garnet schist saprolite and 

kaolinized feldspar in the REL pit (locations 18-1 and 

18-2).

Throughout the REL pit the saprolite is mantled 

by a layer of silt and clay that ranges in thickness 

from 11/2 to 4 inches. However, the permeability of 

the silt and clay proved not to be the major factor 

controlling infiltration into the saprolite inasmuch as 

the greatest infiltration rate was determined where 

the layer of silt and clay is thickest. The greatest 

infiltration rate (1.24 in. per hr.) was obtained at 

location 24-3. Although the material tested at this 

location consists partly of mica-quartz-garnet schist 

saprolite similar to that at the location of all the other 

tests in the same pit, it also consists partly of granitegneiss saprolite, a material not present at the other 

locations. In general, saprolite derived from granitegneiss contains less clay and is more permeable than 

saprolite derived from biotite schist or mica-quartzgarnet schist. Wells drilled in granite-gneiss saprolite 

commonly yield ample supplies of water for domestic 

and small industrial uses.

In REF pit 2, the greatest infiltration rate (1.94 in. 

per hr.) was obtained at location 18-4, where the 

tested material consisted entirely of biotite-ganiet 

schist saprolite. The saprolite at location 24^t was 

similar to that at location 18^t except that the saprolite contains zones of secondary limonite cementation, micaceous fracture coatings, and weathered biotite schist. The weathered biotite schist 'comprises 

about 15 percent of the total material tested at location 24^t. The infiltration rate there was 1.09 in. per 

hr., or considerably less than the rate obtained for 

the biotite-garnet schist saprolite at location 18^L 

At location 24-5, where about 50 percent of the sap-
ARTICLE 220 E141

18-1

18-2

18-3

24-3 

REL DISPOSAL PIT

7//Z

18-4

24-4

24-5 

REF DISPOSAL PIT 2

EXPLANATION

Kaolmized feldspar concentration

Dark micaceous fracture coatings

Biotite-garnet schist saprohte Quartz veins

Secondary limonite cementation

Geology by C. W. Sever, Jr., 1958

FIGURE 220.1. Geologic sections in infiltrometer rings, Georgia Nuclear Laboratory.

and 6 inches deep.

The rings are 18 and 24 inches in diameter

667422 O - 63 - 11
E142 GEOLOGICAL SURVEY RESEARCH 1962

rolite consists of weathered biotite schist, the infiltration rate was 0.96 in. per hr. the smallest rate obtained in BEF pit 2. It appears, therefore, that the 

percentage of weathered biotite schist in the saprolite 

was the principal factor controlling infiltration of 

water at the different locations.

Four of the eight tests in the EEL pit were made 

when the soil was dry, no rain having fallen for 4 to 6 

days preceding the tests or during the tests. The other 

4 tests were made when the soil was wet to moderately wet owing to rain for a 1- to 11-day period 

before the tests were begun and also during the tests.

Tests were made under two different conditions of 

soil moisture at locations 24-2 and 18-2 in the KEL 

pit. One of the tests at location 24-2 was made when 

the soil was dry and the weather was hot and humid; 

the average infiltration rate was 0.49 in. per hr. During the 23-hour period of the test, the air temperature 

ranged from 68°F to 90°F and the water temperature 

ranged from 70°F to 82°F. Although the changes in 

air and water temperature seemed to produce erratic 

changes in the infiltration rate, the effect was minor. 

Evaporation loss was minimized by covering the 24- 

inch ring with an aluminum plate and canvas. The 

other test at location 24 2 lasted 59 hours and was 

made when the saprolite was wet owing to 11 consecutive days of rain (total rainfall 7.15 in.) preceding 

the test. The average infiltration rate was 0.46 in. 

per hr., or nearly the same as the rate obtained for 

the earlier test when the saprolite was initially dry. 

The two tests at location 18-2 were made with an 

18-inch ring under conditions similar to- those made 

at location 24-2. The average infiltration rate was 

0.06 in. per hr. during each of the tests at location 18-2.

All the tests in REF pit 2 were made when the 

saprolite was wet owing to heavy rains before and 

during the tests. As a result, the effect of rain on the 

infiltration rate could not be determined.

The widely differing infiltration rates were due 

largely to the differences in mineral composition of 

the saprolite, to the condition of the saprolite (dry 

or saturated) before and during the test periods, and 

possibly to the use of rings of different size. Other 

factors which probably affected the infiltration rates 

to a lesser degree were the presence or absence of 

quartz veins, differences in the thickness of the clay 

and silt cover, and changes in the temperature of the 

air and water.

Because a large volume of material was tested, the 

results obtained with the 24-inch ring infiltrometer 

probably are more representative of the actual infiltration rate of the saprolite.

221. EVAPOTRANSPIRATION AND THE RELATION OF GROUND WATER TO SURFACE WATER IN THE POND

CREEK BASIN, OKLAHOMA

By WILLIAM E. CLARK, Gainesville, Fla.

If a water-table aquifer discharges into a stream, 

the flow of that stream during periods of no overland 

runoff is related to the stage of the water level in the 

aquifer. However, the relation between dry-weather 

streamflow and water-level fluctuations in wells is not 

direct because variations in the rate of evapotranspiration cause differences in the amount of water measured as streamflow. To establish the relation between 

the dry-weather streamflow and ground-water levels 

in the Pond Creek basin, Oklahoma, an analysis was 

made of pertinent data for the period 1951-57. A 

preliminary analysis of hydrologic data for this basin 

was made by the author (Clark, 1956).

The Pond Creek basin is in west-central Oklahoma. 

The drainage from an area of 319 square miles is 

measured at the Geological Survey gaging station 

on Pond Creek near Fort Cobb, Caddo County. About 

500 feet from the creek at a point about 3.5 miles upstream from the gaging station is an observation well, 

referred to in this article as well 46. The position of 

the water level in this well is considered to be a reliable index to the general ground-water level in the 

part of the aquifer adjoining Pond Creek because it 

varies linearly with the average water-level position 

in 16 other wells near the creek. To take into account the varying effect of evapotranspiration on the
ARTICLE 221 E143

A/\/V V \/\

FIGURE 221.1. Diagram showing the relation of the dry-weather flow of Pond Creek to the weather index and depth

to water in well 46.
E144 GEOLOGICAL SURVEY RESEARCH 1962

dry-weather flow of Pond Creek, a weather index was 

computed for given dates in the period 1951-57. The 

weather index for each given date was derived by 

averaging the maximum air temperature, in degrees 

Fahrenheit above freezing, for that date and the preceding 9 days, and multiplying the average thus obtained by the number of daylight hours on that date. 

The results of a correlation of the dry-weather flow 

of Pond Creek with the depth to water in well 46 

and the weather index are shown by the diagram in 

figure 221.1. As would be expected, the relation between dry-weather flow and the depth to water is 

linear when the weather index is constant. From the 

diagram, the dry-weather flow of Pond Creek can be 

determined for various combinations of the weather 

index and the depth to water in well 46.

The weather index indicates chiefly the effects of 

weather on the evapotranspiration of ground water, 

but also includes the effects of weather on the evaporation of water from the surface of the creek.' When 

the weather index is less than 280 degree-hours, there 

is little or no decrease in streamflow due to evapotranspiration. A weather index of about 280 ordinarily 

occurs when the average daily high temperature is 

about 60°F. When the weather index exceeds about 

280, the dry-weather flow, for a constant stage of the 

ground-water level, decreases with an increase in the

weather index. The tendency of the curve to flatten 

for larger values of the weather index indicates that 

there is a maximum rate at which phreatophytes transpire ground water.

The effect of evapotranspiration on the dry-weather 

flow of Pond Creek can be derived from the diagram 

in figure 221.1 as follows: First, determine the dryweather flow corresponding to the stage of the water 

level in well 46 and to a weather index of less than 

280 degree-hours. Then, determine the dry-weather 

flow corresponding to the stage of the water level 

and to the actual weather index. The difference between these dry-weather flows is the decrease in 

streamflow due to evapotranspiration for the particular combination of water level and weather.

The amount of the decrease in streamflow due to 

evapotranspiration varies not only with the weather 

index but also with the stage of the ground-water 

level. As may be determined from the diagram, if 

the weather index is 1,000 degree-hours, the decrease 

in streamflow due to evapotranspiration of ground 

water is 20 cfs (cubic feet per second) when the 

depth to water in well 46 is 13 feet and is 12 cfs 

when the depth to water is 17 feet. Thus, a 4-foot 

decline of the water level in well 46 reduces the effect of evapotranspiration on the streamflow by about 

40 percent.

1000 pr

100

1953 1954

FIGURE 221.2.- -Comparison of measured streamflow of Pond Creek a'nd the dry-weather flow as determined by the weather index

and depth to water in well 46.
ARTICLE 222 E145

Evapotranspiration accounts for a large percentage 

of the variation in the dry-weather flow of Pond 

Creek. For example, when the depth to water in well 

46 is 15 feet and the weather index is 1,000 degreehours, the dry-weather flow is about 11 cfs, a reduction of about 16 cfs, or 60 percent, due to evapotranspiration.

The dry-weather flow as determined from the diagram and the measured streamflow for the 1954 water 

year are shown graphically in figure 221.2. During 

periods of no overland runoff, the dry-weather flow 

as determined from the diagram and the measured 

streamflow generally agree closely. However, when

the measured streamflow consisted partly of overland 

runoff, the dry-weather flow as determined from the 

diagram represents the potential dry-weather flow, or 

the flow that the stream would have had if there had 

been no overland runoff. Some periods, such as July 

and August, when actual streamflow was slightly less 

than dry-weather flow determined from the diagram, 

indicate that other factors may at times be significant.

REFERENCE

Clark, W. E., 1956, Forecasting the dry-weather flow of Pond 

Creek, Oklahoma; a progress report: Am. Geophys. Union 

Trans., v. 37, no. 4, p. 442-450.

222. ESTIMATING THE SPECIFIC CAPACITY OF A WELL

By A. N. TURCAN, JR., Baton Rouge, La.

Work done in cooperation with Louisiana Department of Public Works and, Department of Conservation,

Louisiana Geological Survey

Although equations and theories regarding the hydraulic properties of an aquifer and the water-producing ability of wells have been derived and discussed 

byTheis (1935), Muskat (1937), Jacob (1945), Ahrens 

(1958), and Todd (1959), the performance of a well 

cannot be predicted readily by the direct application 

of any one formula. The intent of this article is to 

present some of the equations, together with simple 

examples, which aid in predicting the specific capacity of a 100-percent-efficient well that taps an isotropic, homogeneous artesian aquifer of large areal 

extent. Considered, in order, are the computations of 

the theoretical specific capacity of a well that is 

screened opposite the full thickness of the aquifer and 

of a well that is screened opposite only part of the 

aquifer. Because all aquifers differ in some respects

from the postulated "ideal" conditions and because 

few, if any, wells are 100-percent efficient, predictions 

of well performance can, at best, be only approximate.

As used in this article, the terms describing hydraulic properties of an aquifer and the performance 

of a well are defined as follows: The coefficient of 

transmissibility (T) of an aquifer is the rate of flow 

of water, in gallons per day, through a vertical strip 

of the aquifer 1 foot wide extending the full saturated height of the aquifer under a unit hydraulic 

gradient. The coefficient of storage (S) of an aquifer 

is the volume of water released from or taken into 

storage per unit surface area of the aquifer per unit 

change in the component of head normal to that sur-
E146 GEOLOGICAL SURVEY RESEARCH 1962

face. The specific capacity of a well is its yield in 

gallons per minute per foot of drawdown.

If a well is screened opposite the full thickness of an 

"ideal" aquifer and if, when the well is pumped, the 

water level inside the well is the same as that immediately outside the well, the specific capacity (I7) of this 

100-percent efficient well may be expressed in terms of 

the Theis nonequilibrium formula (Theis, 1935, p. 520), 

as follows:

 O U 1 T (1)

in which

<?=rate of discharge from the well, in gallons per

minute ;

Sw=water-level drawdown in the well, in feet; 

T is as previously defined; and 

W(u) is known as "the well function of u."

Values of W(u) corresponding to values of u from 10~15 

to 9.9 may be obtained from a table presented originally 

by Wenzel (1942, p. 89) and reproduced by Ferris and 

others (1962, p. 96-97). The value of u is obtained 

from

u=-

Tt

in which

rw =radius of the pumped well, in feet; 

£=time since pumping began, in days; and 

S and T are as previously defined.

The relation of the specific capacity of a 100-percentefficient well to well diameter and to the hydraulic 

characteristics of the aquifer is shown by a graph that is 

based on the Theis nonequilibrium formula and was 

prepared by Hex K. Meyer (Theis and others, 1954, 

fig. 2). If the coefficients of transmissibility and storage for an "ideal" aquifer are known, this graph provides a quick and easy method for estimating the 

specific capacity of the well that has been pumped 

continuously for 1 day.

The following example illustrates the use of the Theis 

nonequilibrium formula to estimate the specific capacity 

of a 100-percent-efficient well that is screened opposite 

the full aquifer thickness.

Example 1. A water user requires 1,000 gpm 

(gallons per minute) from a well that has a radius 

(rw) of 0.5 foot and is screened opposite the full thickness of an artesian aquifer. The aquifer is 100 feet 

thick, has a coefficient of transmissibility of 100,000 

gpd (gallons per day) per foot, and has a coefficient of 

storage of 0.001. No nearby pumping will affect the 

water level in the well. The user desires to know what

the theoretical specific capacity of his well will be at the 

end of 1 day of continuous pumping.

The value for u is determined, as follows, by substituting appropriate values for the unknowns in 

equation 2 :

^1.87(0.5) 20.001^ 

U~ 100,000X1 ~

The corresponding value of W(u) is 18.6055, and the 

theoretical specific capacity of the well at the end of 

1 day is computed from equation 1, as follows:

Y= 100,000

114.6X18.6055=47 gpm per foot of drawdown.

In most wells the screen is set opposite only part of 

the full thickness of the aquifer. Thus, even in an 

"ideal" aquifer, the flow lines to and into the well are 

modified and the well's theoretical specific capacity is 

decreased by a certain percentage. The problem of 

partial penetration (that is, of wells screened opposite 

only part of the aquifer) is discussed in detail by Muskat (1937), Jacob (1945), and Ahrens (1958). To 

predict the yield of a well that is screened opposite 

either the upper or the lower part of an "ideal" aquifer, 

Jacob (1945) presented the following empirical expression of a formula derived by Kozeny (1933):

(3)

in which

Q discharge of a well screened opposite only part

of the aquifer; 

Q0=discharge of a well screened opposite the full

thickness of the aquifer; 

 p=ratio of length of screen to full thickness of the

aquifer, expressed as a percentage; 

rM,=radius of the pumped well, in feet; and 

ra=thickness of the aquifer, in feet.

Because equation 3 was derived for steady-state 

conditions, it is not applicable unless the rate of drawdown has become so slow that for all practical purposes 

steady-state conditions can be assumed to exist. 

Generally, pumping for 1 day produces such conditions. 

Even if the well is screened opposite the middle part 

of the aquifer or is screened opposite two or more 

different intervals, equation 3 provides results that 

are satisfactory for most general engineering and design 

purposes.

The specific capacity of a well that is screened opposite only part of an aquifer (YA) can be expressed in 

terms of the specific capacity of a well that is screened
ARTICLE 228 E147

opposite the full thickness of the same aquifer (F) by 

dividing both sides of equation 3 by the drawdown 

(sw), as follows:

cos 2

1+7 2pm cos V

For ease of discussion, the quantity in brackets in 

both equation 3 and equation 4 is referred to as the 

productivity ratio (PR), a term used by Craft and 

Hawkins (1959, p. 300) in similar calculations to determine the production of oil wells. Thus, equation 4 

may be expressed

YA=Y[PR]. (5)

A graph presented by Muskat (1937, fig. 84) has been 

modified by extension of the curves and changes in 

nomenclature to provide an easy means for determining 

the productivity ratio for wells screened opposite 

different percentages of the full thickness of the aquifer (fig. 222.1). The productivity factor is a dimensionless term which expresses the relative yield of wells 

that are screened opposite different percentages of 

aquifer thickness. The curves in figure 222.1 can be 

both extrapolated and interpolated with considerable 

accuracy because, for aquifers more than 50 feet thick, 

the relation of the productivity factor to aquifer thickness is almost exactly linear (Muskat, 1937, p. 275).

The successive steps in computing the specific capacity of a well that has a radius of 0.25 or 0.5 foot 

and is screened opposite only part of an isotropic and 

homogeneous aquifer of known thickness are as follows:

1. Using figure 222.1, determine the productivity 

factor for the well whose specific capacity is to be 

computed.

2. Using the same illustration, determine the productivity factor for a well that has the same radius, 

that taps an aquifer of the same thickness, but that 

is screened opposite the full thickness of the aquifer.

3. Divide the productivity factor determined in 

step 1 by the productivity factor determined in step 2 

to obtain the productivity ratio (PR}.

4. Multiply the productivity ratio by the specific 

capacity (7) of a 100-percent-efficient well that taps 

the full thickness of the aquifer. (See equation 5.)

The following example illustrates use of the graph 

(fig. 222.1).

Example 2. A water user requires 1,000 gpm from 

a well that has a radius of 0.5 foot and that is

p=ratio of length of 

screen to full thickness of aquifer, 

expressed as a 

percentage

40 80 120 160 200 240 280 320 360 400 

THICKNESS OF AQUIFER, IN FEET

FIGURE 222.1. Relation of aquifer thickness and productivity 

factor to well radius and percent of aquifer screened (after 

Muskat, 1937, fig. 84).

screened opposite 50 percent of the full thickness of 

an artesian aquifer. The aquifer is 100 feet thick, 

has a coefficient of transmissibility of 100,000 gpd per 

foot, and has a coefficient of storage of 0.001. No 

nearby pumping will affect the water level in the well. 

The user desires to know the specific capacity of his 

well at the end of 1 day of continuous pumping.

From figure 222.1, the productivity factor for a 

well that has a radius of 0.5 foot and that is screened 

opposite 50 percent of the full thickness of an aquifer 

is 0.95. Also from figure 222.1, the productivity factor for a well that has a radius of 0.5 foot and that 

is screened opposite the full thickness of an aquifer 

is 1.5. Thus, the productivity ratio (0.95 -=- 1.5) is 

computed to be 0.63. This value, multiplied by 47 

gpm per foot of drawdown (the specific capacity of 

a well screened opposite the full thickness of the 

same aquifer), gives 30 gpm per foot of drawdown 

for the specific capacity of the well screened opposite 

50 percent of the aquifer thickness.
E148 GEOLOGICAL SURVEY RESEARCH 1962

REFERENCES

Ahrens, T. P., 1958, Water well engineering: Water Well 

Jour., v. 12, no. 12, p. 18, 19, 28, 30-32.

Craft, B. C., and Hawking, M. F., 1959, Applied petroleum 

reservoir engineering: Englewood Cliffs, N.J., PrenticeHall, Inc., 437 p.

Ferris, J. G., Knowles, D. B., Brown, R. H, and Stallman, 

R. W., 1962, Theory of aquifer tests: U.S. Geol. Survey 

Water-Supply Paper 1536-E, p. 69-174.

Jacob, C. E., 1945, Partial penetration of pumping well, adjustments for: U.S. Geol. Survey open-file rept, 7 p., 

4 figs.

Kozeny, J., 1933, Theorie und Berechnung der Brunnen: 

Wasserkraft und Wasserwirtschaft, v. 28, p. 101.

Muskat, Morris, 1937, The flow of homogeneous fluids through

porous media: New York, McGraw-Hill Book Co., Inc.,

763 p. 

Theis, C. V., 1935, The relation between the lowering of the

piezometric surface and rate and duration of discharge

of a well using ground-water storage: Am. Geophys.

Union Trans., pt. 2, p. 519-524. 

Theis, C. V., and others, 1954, Estimating transmissibility

from specific capacity: U.S. Geol. Survey open-file rept.,

lip. 

Todd, D. K., 1959, Ground water hydrology: New York, John

Wiley & Sons, Inc., 336 p. 

Wenzel, L. K., 1942, Methods of determining permeability of

water-bearing materials, with special reference to discharging-well methods: U.S. Geol. Survey Water-Supply

Paper 887, 192 p.

223. AERIAL OBSERVATION OF ICE COVER TO LOCATE AREAS OF GROUND-WATER INFLOW TO STREAMS

By S. W. WHTALA and T. G. NEWPORT, Lansing and Escanaba, Mich.

Correlation of open-water reaches and measurement 

of the base flow of streams in the Upper Peninsula of 

Michigan indicate that the inflow of relatively warm 

ground water is sufficient to prevent many small 

streams from freezing over in subfreezing weather. 

Aerial photography and direct aerial observation can 

thus be used under favorable conditions as a reconnaissance method of mapping areas of ground-water 

discharge.

The possibility of using aerial observations to determine areas of ground-water seepage to streams was 

suggested when open reaches of streams were observed on a flight over the Upper Peninsula of Michigan in late January 1962. Streams were at base flow 

following an extended period of relatively cold weather 

and were completely ice covered in most areas. Some 

relatively short reaches were open because of the turbulence induced by rapids and small waterfalls. Open 

Avater in other reaches, however, could not be explained 

by turbulence or by the presence of sewage or other 

wastes in the streams. These observations led to the 

hypothesis that the open water in some reaches was 

caused by the discharge of relatively large quantities 

of warmer (in winter) ground water into the streams.

To test this hypothesis further, investigations were 

made in an area near Marquette, Mich. Two lowaltitude flights, on February 20 and on March 15, 

1962, were made over the area in a light plane furnished by the Michigan Department of Conservation. 

On these flights, observed open reaches of streams 

were marked on maps and were photographed with a 

35-mm camera. On March 17, aerial photographs of 

the area were taken by the Michigan Air National 

Guard from an altitude of 13,300 feet. By using a 

stereoscope, the reaches of open water, even on very 

small streams, could be defined quite easily and accurately on the aerial photographs. Later, the observations and the photographs obtained on the three flights 

were correlated with streamflow measurements made 

on selected streams in the area during a period of base 

flow in August 1962.

A comparison of the ice conditions observed in February and March 1962 on Goose Lake Outlet with 

base flow on August 16, 1962, is presented on figure 

223.1. This stream leaves Goose Lake and flows across 

an outwash plain composed of sand, gravel, and 

boulders. Scattered knobs of Precambrian rock protrude through this sheet of unconsolidated material.

All of the ice cover on Goose Lake Outlet was observed in a reach where the stream is actually losing 

water. The open water just downstream from Goose 

Lake can be attributed to the outflow of lake water,
ARTICLE 223 E149

7 -

i 6

o 

o

LJ 

CO

o: 5

LJ 

Q_

Area of 

report

- '/"%

MICHIGAN/^ j

n

876543210 

DISTANCE ABOVE MOUTH, IN MILES

FIGURE 223.1. Ice-cover observations February and March 

1962, in relation to discharge, August 16, 1962, during a period 

of base flow, Goose Lake Outlet.

which in winter would be somewhat warmer than 

stream water. The open water between the two frozen reaches may be caused by rapids and perhaps by 

slight inflow from a small tributary. The discharge 

graph is dashed between miles 3.8 and 4.95 because 

it represents an estimate of the inflow from the small 

tributary at mile 3.85. Downstream from this tributary, the flow increases rapidly in a reach that was 

open at the time of the winter observations. No other 

tributaries join Goose Lake Outlet in this reach, but 

many seeps along both banks noted when the streamflow measurements were made indicate substantial 

ground-water discharge.

Similar methods were used to investigate several 

other streams near Marquette. The results, including 

those for Goose Lake Outlet, are summarized in the 

table.

The measured base flows of all streams listed in the 

table were less than 8 cubic feet per second. The 

results indicate that in these streams the gain in disStream

Goose Lake Outlet

(below mile 3.8) . 

Bear Creek. _ _______

River near Humboldt.

Ice conditions

Open__ ____

_____do_________

_____do________-

Lengthof 

reach 

(miles)

2.95

3.90

3.55

2. 85

2. 55

Gain in 

discharge 

per mile of 

channel 

(cubic feet 

per second)

1. 56

1.57

1.48

.49

.42

charge in unfrozen reaches as measured the following 

summer was about three times greater than that for the 

reaches that were mostly frozen over.

The investigations described suggest that reaches of 

open water in small streams during prolonged cold 

periods may indicate substantial ground-water discharge to the streams. Extraneous factors such as 

the turbulence due to steep gradient, the presence of 

sewage and other wastes, the decomposition of vegetal matter, and perhaps others may also prevent the 

formation of a complete ice cover on streams. In any 

particular investigation, these factors must be considered.

Winter aerial observations appear to provide a 

quick, simple, and inexpensive means of identifying 

areas that should be considered for more intensive investigation. The method should be particularly useful 

in remote and inaccessible areas where other data are 

not available. It is, of course, adaptable only to relatively small streams in regions where most of the 

streams are completely ice covered for long periods. 

The effect of ground-water inflow upon ice conditions 

would hardly be noticeable on the larger streams 

where such inflow is small in relation to the volume 

of water carried in the stream. In applying the 

method, flights should be carefully scheduled to coincide with periods of cold weather and base flow, 

and should not be made soon after a heavy snowfall 

that may bridge over open areas on small streams.
El 50 GEOLOGICAL SURVEY RESEARCH 1962 

GROUND WATER

224. PERMEABILITY OF GLACIAL TILL

By STANLEY E. NORRIS, Columbus, Ohio

Work done in cooperation with the Division of Water of the Ohio Department of Natural Resources and with the

Miami Conservancy District

Glacial till controls in varying degree the availability of ground water in thousands of square miles in 

the northern United States. Where it mantles broad 

areas of upland, till retards the infiltration of precipitation to underlying aquifers. Also, in valleys where 

the water table is hydraulically continuous with a 

river, a layer of till within the zone of saturation may 

confine the aquifer below it and greatly limit recharge, 

by induced infiltration of streamflow, to the confined 

aquifer.

Despite its hydrologic importance, very few data 

on the permeability of till have been published. One 

reason for this lack of permeability determinations is 

that till is a poor aquifer. No doubt another reason 

is that many hydrogeologists believe that a permeability value, or even a range of such values, for till 

at one place would be meaningless if applied to till 

in other places. The latter idea may stem from a 

widespread misconception that till is a more or less 

random accumulation of drift, ranging widely in its 

physical properties from place to place and lacking 

continuity in many of the common lithofacies characteristics upon which correlation of sediments is 

based. On the contrary, evidence shows that till in 

widely separated areas actually is reasonably uniform 

in permeability and, by inference, in other related 

properties. Most values of the permeability coefficient of till in Ohio are similar to those for till in 

Illinois and South Dakota (fig. 224.1). This is true 

for determinations made both in the laboratory and 

by aquifer tests.

Laboratory determinations of permeability of till 

are reported for 10 samples from Spink and Hand 

Counties, east-central South Dakota; 3 from Montgomery County, southwest Ohio; and 13 from Holmes, 

Wayne, Medina, and Ashland Counties, northeast 

Ohio. Field computations of till permeability based 

on aquifer tests are reported for 5 sites in the Mad 

River valley, Montgomery County, Ohio, and for 7 

sites in southern Illinois. The laboratory tests were 

made in the Hydrologic Laboratory of the Geological 

Survey.

i.o

o.i

0.001

0.01

0.0001

Montgomery 

County, Ohio

FIGURE 224.1. Chart showing range of till permeability in 

Ohio, Illinois, and South Dakota.
ARTICLE 224 E151

The samples from South Dakota were obtained with 

a diamond core barrel and hydraulic drill, either from 

surface exposures or at depths ranging from about 18 

to 110 feet. Presumably they represent till of Wisconsin age, as do also the Ohio samples.

Four samples of surficial till from Montgomery 

County, Ohio, were collected by driving thin-walled 

aluminum cylinders 6 inches in diameter and 6 inches 

in length into the till, thus cutting an "undisturbed" 

core. The hardest and most resistant till sampled 

had a coefficient of permeability of 65 gallons per day 

per square foot. Because this is an excessively high 

value, the sample is believed to have cracked before 

it reached the laboratory. Therefore, the value is not 

shown on figure 224.1.

Sixteen samples of surficial till were collected in 

northeastern Ohio also by driving cylinders into the 

till, but, as the cylinders were smaller 2 inches in 

diameter and about 2 inches in length sampling was 

much easier and the resulting cores are believed to 

have been disturbed less than those collected in the 

larger cylinders. Nevertheless, 3 of the 16 samples 

had either shrunk or cracked en route to the laboratory and so were not tested for permeability.

The permeability of till of early Wisconsin age in 

the Mad River valley at Rohrers Island, near Dayton, Montgomery County, Ohio, was determined by 

analysis of an aquifer test (Norris, 1959). The till 

ranges in thickness from about 11 to 50 feet and lies 

between 2 deposits of water-bearing sand and gravel, 

of which the upper averages about 65 feet in thickness 

and the lower about 50 feet. The permeability of the 

till was determined from its effect on water-level fluctuations in 5 observation wells whose distance from 

the pumped well ranged from 510 to about 4,200 feet.

Walton (1960) determined the permeability of till 

at seven sites in southern Illinois from aquifer tests. 

At 6 of the" sites the till, which ranges in thickness 

from 7 to 16 feet, is presumed to be of Illinoian age; 

at the other site the till is 12 feet thick and is of Wisconsin age. The permeability of the till at one site 

was computed to be 1.6 gallons per day per square 

foot. As this value is nearly three times as great as 

the maximum value for the till at the other sites, it is 

thought not to be representative and so is not shown 

on figure 224.1.

The permeability values determined by analysis of 

the aquifer test at Rohrers Island compare closely 

with those determined in the laboratory for surficial 

tills from elsewhere in Montgomery County, Ohio. 

Moreover, the permeability values for till in Illinois 

are in the same range as those for till in Montgomery 

County, and although the range of permeability values 

of the samples collected in northeast Ohio and in 

South Dakota is relatively large, most values are in 

the same range as that of the till in southwest Ohio 

and in Illinois.

The close grouping of most of the permeability values suggests that, with judgment, they can be extrapolated over large distances and applied with reasonable confidence in estimating the rate of percolation 

through glacial till to underlying aquifers.

REFERENCES

Norris, S. B., 1959, Vertical leakage through till as a source 

of recharge to a buried-valley aquifer at Dayton, Ohio: 

Ohio Dept. Nat. Resources, Div. of Water Tech. Rept. 2.

Walton, W. C., 1960, Leaky artesian aquifer conditions in 

Illinois: Illinois State Water Survey, Dept. Registration 

and Education Rept. Inv. 39.
E152 GEOLOGICAL SURVEY RESEARCH 1962

225. GROUND WATER IN CENOZOIC FILL IN COLLAPSE STRUCTURES, SOUTHEASTERN EDDY COUNTY,

NEW MEXICO

By JAMES B. COOPER, Albuquerque, N. Mex.

Work done in cooperation with the U.S. Atomic Energy Commission

Thick Cenozoic deposits filling a deep depression 

formed by solution of evaporites from rocks of Permian age yield large quantities of potable water to 

wells in southeastern Eddy County, N. Mex. The 

deposits underlie an area of 150 square miles east of 

the Pecos River in New Mexico and extend southward 

into the western part of Texas (fig. 225.1).

Maley and Huffington (1953) described three extensive accumulations of Cenozoic fill in the Delaware 

basin of southeastern New Mexico and western Texas

R.29E. R.30E. TEXAS R.31E.

01234 SMILES 

I I I I I I

EXPLANATION

NEW MEXICO

Area of 

report.

Triassic rocks

Permian rocks

Water well

Numeral indicates 

number of wells 

at location

FIGURE 225.1. Map showing subsurface extent of Cenozoic fill 

and older rocks in southeastern Eddy County, N. Mex. 

Hachures indicate closed depressions at land surface.

and attributed the formation of the depressions that 

contain the fill to the removal, by surface and ground 

water, of great amounts of salt and other evaporites 

from rocks of Permian age. The depressions presumably were formed in late Tertiary time when the 

evaporite rocks were elevated and partly exposed as 

a result of an eastward tilting of the Delaware basin. 

The depression in southeastern Eddy County described here is connected by thin deposits of fill with 

a larger depression to the south in Texas. Nicholson 

and Clebsch (1961, p. 46-47) describe similar collapse 

structures in the adjoining part of southern Lea 

County to the east.

The Cenozoic fill consists of unconsolidated and 

semiconsolidated clay, silt, and sand. Probably most 

of the deposits are of Quaternary age, but in part 

they may be of late Tertiary age. Maley and Huffington (1953, pi. 1) indicate that the fill is more than 1,400 

feet thick at its center near the New Mexico-Texas 

State line, but none of the logs of holes drilled in the 

New Mexico part of the depression show the fill to be 

more than 1,000 feet thick. As there are no surface 

indications of the margins of the collapse depression, 

the extent of the Cenozoic deposits is shown as approximate in figure 225.1. Information from drillers' 

logs of oils tests and water wells and data on the 

chemical quality of the ground water were used to 

determine the approximate extent of the fill. The 

logs were studied and the water samples were analyzed during a detailed investigation of the groundwater resources of the Project Gnome area made by 

Cooper (1962) on behalf of the U.S. Atomic Energy 

Commission.

The Cenozoic fill is the only material in southeastern Eddy County from which large quantities of 

potable water can be obtained. Although its existence 

had been known as early as 1952, its potential as an 

aquifer has been recognized only within the past few 

years. The bedrock of Permian and Triassic age 

yields little or no potable water.

The maximum thickness of saturated Cenozoic deposits in the New Mexico part of the depression is 

about 800 feet. Water is yielded to wells by beds of 

sand, each as much as 50 feet thick and separated by
ARTICLE 225 E153

somewhat thicker beds of silt of clay. For example, 

water well 1 that supplies the Pecos turbine station 

of the El Paso Natural Gas Co. taps four different 

sand zones. The specific capacity of this well, as

FIGURE 225.2. Chemical quality of ground water in southeastern Eddy County, N. Mex. Sum of calcium plus 

magnesium in equivalents per million times 50 equals hardness 

as CaCO3 in parts per million.

determined by pumping the well for 20 hours at about 

200 gpm (gallons per minute), was 10 gpm per foot 

of drawdown. The potential yield of the well is estimated to be more than 1,000 gpm.

At present, 14 wells ranging in depth from about 

300 feet to about 800 feet tap the Cenozoic fill in southeastern Eddy County. Two of the wells, in sec. 5, 

T. 26 S., E. 30 E., are pumped at more than 100 gpm 

each to supply water to the Pecos turbine station; 

all the others are for stock or domestic supplies and 

are pumped at rates of only a few gallons per minute.

Ground water in the area moves generally southwestward. Although some water percolates into the 

Cenozoic fill from the Triassic rocks on the east and 

northeast, the primary source of recharge to the fill 

probably is local precipitation that infiltrates through 

large sinkholes and other surface depressions in the 

central part of the northeast quadrant of the fill area 

(fig. 225.1). Ground water moves out of the fill into 

rocks of Permian age where its chemical quality rapidly deteriorates owing to contact with gypsum and 

saline evaporites.

The chemical quality of the water in the Cenozoic 

fill is far superior for most uses to the highly saline 

water normally present in the Permian and Triassic 

rocks. Figure 225.2 presents chemical analyses of 

water samples from typical wells tapping each of the 

indicated aquifers.

The large quantity of water of relatively good 

chemical quality contained in the Cenozoic fill is a 

valuable natural resource which has not been utilized 

fully because the economic development of the region 

is still in an early stage.

REFERENCES

Cooper, J. B., 1962, Ground-water investigations of the Project 

Gnome area, Eddy and Lea Counties, N. Mex.: U.S. Geol. 

Survey TEI 802, open-file rept, 67 p.

Maley, V. C., and Huffington, R. M., 1953, Cenozoic fill and 

evaporite solution in the Delaware Basin, Texas and New 

Mexico: Geol. Soc. America Bull., v. 64, no. 5, p. 539-545.

Nicholson, Alexander, Jr., and Clebsch, Alfred, Jr., 1961, Geology and ground-water conditions in southern Lea County, 

N. Mex.: New Mexico Bur. Mines and Mineral Resources 

Ground-Water Rept. 6.
El54 GEOLOGICAL SURVEY RESEARCH 1962

226. RELATION OF FRESH WATER TO SALT WATER AT CENTRE ISLAND, NASSAU COUNTY, NEW YORK

By JOHN ISBISTER, Mineola, N.Y.

WorJc done in cooperation with the Nassau County Department of Public WorJcs

Although surrounded on all sides by sea water, 

Centre Island has ample supplies of fresh water owing 

to a fortunate combination of geologic and hydrologic features. Wells yield water of ^excellent quality 

from a shallow unconfined aquifer and from a deeper 

confined aquifer. The fresh-water lens in the shallow 

aquifer floats on salt water and is replenished by infiltration of precipitation on the island. The deep 

aquifer is effectively sealed off from the overlying salt 

water by thick, extensive clay strata and is replenished by infiltration on the main body of Long Island.

Centre Island, in northeastern Nassau County, is 

slightly more than 2 miles long, ranges in width from 

about 0.2 to 0.8 mile, and comprises an area of about 

a square mile (fig. 226.1). Except for a narrow spit 

which connects it to the mainland, Centre Island is 

completely surrounded by salt water in Oyster Bay 

Harbor and Long Island Sound. The long axis of 

the island is formed by a north-trending ridge of hills, 

which rises to a maximum altitude of about 100 feet 

in the southern extremity. Centre Island is nearly 

breached at its greatest width by a low marshy area 

that is partly inundated by tidal waters. Glacial till 

containing large boulders covers most of the irregular 

land surface. Rocky Point at the north end of the 

island derives its name from the abundance of boulders and cobbles that litter the beach.

Centre Island is an incorporated village composed 

almost entirely of homes and large estates; the population in 1960 was 270 persons (Nassau County Planning Commission, 1961, p. 9). Each of the 175 homes 

on the island has its own well, and wastes are disposed of through cesspools.

Water levels mentioned in this article are referred 

to mean sea level at Sandy Hook, N.J., datum. The 

mean tidal range at Oyster Bay Harbor is about 7.3 

feet and the normal range of spring tides is as much 

as 8.6 feet (U.S. Dept. of Commerce, 1962, p. 214).

Well data from nearby parts of Nassau and Suffolk 

Counties indicate that the bedrock beneath Centre 

Island is composed of crystalline rocks of Precambrian to early Paleozoic age. The bedrock surface 

slopes southeast and ranges from about 450 feet below 

sea level at Rocky Point at the north end of the island 

to 550 feet below sea level at Moses Point at the south 

end (fig. 226.2). Unconsolidated sediments of Late

Cretaceous and Pleistocene age, ranging in thickness 

from about 450 to 600 feet, lie on the bedrock. The 

Upper Cretaceous deposits are divided into two formations, the Raritan and the Magothy(?). The older, 

the Raritan Formation, is subdivided into the Lloyd 

Sand Member and an unnamed clay member. The 

younger beds of Cretaceous age are tentatively assigned to the Magothy (?) Formation. The Cretaceous deposits are mantled by beds of glacial till and 

outwash of Wisconsin age referred to informally as 

upper Pleistocene deposits. A deep buried channel 

just south of Centre Island indicates that the Cretaceous sediments were eroded to a depth of nearly 550 

feet below mean sea level (fig. 226.2). Well data from 

nearby Cove Neck suggest that the channel filling consists of beds of the Jameco Gravel, a lower Pleistocene 

outwash deposit; the Gardiiiers Clay, an interglacial 

marine deposit; and upper Pleistocene deposits. Recent deposits form narrow beaches along much of the 

periphery of the island.

The ground-water reservoir consists of two discrete 

aquifers: a shallow aquifer, which consists partly of 

beds of the Magothy (?) Formation and partly of 

Pleistocene deposits, and a deep confined aquifer, 

which consists mostly of the Lloyd Sand Member of 

the Raritan Formation and partly of the Jameco 

Gravel (fig. 226.2).

The shallow aquifer is composed of beds of sand 

and gravel and lenses of silt and clay having a total 

thickness of about 100 to 200 feet. The aquifer contains water under unconfined or semiconfined conditions. The water table ranges in altitude from about 

sea level near the shore to about 4 feet above in the 

southern part of the island. Recharge is by infiltration of local precipitation, which averages about 44 

inches annually. Recharge probably ranges from 11 

inches (0.5 million gallons per day per square mile) 

to 22 inches (1.0 mgd per square mile). Discharge 

from the shallow aquifer is chiefly by evapotranspiration and by submarine outflow; the net withdrawal 

from wells is negligible.

The Lloyd Sand Member, which forms most of the 

deep confined aquifer, is about 200 feet thick. It contains water under artesian conditions owing to tight 

confinement by the overlying clay member and the 

underlying bedrock. Continuity with the deep con-
ARTICLE 226 E155

CENTRE ISLAND SOUND

,OUG

ATLANTIC

25 MILES

FIGURE 226.1. Index map showing location of Centre Island.

fined aquifer of Long Island is maintained beneath 

Oyster Bay Harbor by hydraulic interconnection 

through deposits of the Jameco Gravel, which are 

overlain and confined by the Gardiners Clay (fig. 

226.2). The deep aquifer at Centre Island is recharged chiefly by underflow from the south. Discharge is chiefly by submarine outflow to Long Island 

Sound. A few wells along the shore flow to waste 

most of the time, and water pumped for domestic use 

is returned to the shallow aquifer by cesspools. A 

survey of the flowing wells in the Centre Island area 

was made between 1944 and 1949 by H. L. Frauenthal 

of the Nassau County Department of Public Works.

The shallow aquifer contains a relatively thin lens 

of fresh water underlain by a larger body of salt 

water. Wells screened in the shallow aquifer are 

concentrated in the southern part of the island, where 

the fresh-water lens is thickest. The fresh-water lens 

thins to the north to the point where it is inadequate 

for domestic supply.

The normal range of the tide in Oyster Bay Harbor 

is from about 4 feet above to 4 feet below sea level. 

Onshore winds and changes in barometric pressure 

extend the normal range. Unusually high tides during hurricanes may inundate low-lying areas. Fluctuations of head in the aquifer and of the tide levels 

in the surrounding bodies of salt water causes intermittent movement of salt water into and out of the 

shallow aquifer, resulting in a wide zone of diffusion. 

Because of scant data, the contact of fresh water and 

salt water shown in figure 226.2 is approximate. The 

chloride content of the water in the shallow aquifer 

ranges from 8 parts per million in the fresh zone to 

at least 2,196 ppm in the zone of diffusion, whereas 

the source of the salty water, Long Island Sound, 

contains about 15,700 ppm. Slightly higher than 

normal amounts of chloride in wells tapping the 

fresh ground water may be caused by water drawn 

from the zone of diffusion, by a small amount of 

chloride in the precipitation, by contamination from 

cesspools, or by a combination of the three. Saltwater flooding of the marshy areas also accounts for 

the presence locally of salty water in the shallow 

aquifer.

The deep confined aquifer yields fresh water containing about 3 ppm of chloride. A subbottom depthrecorder survey made in Long Island Sound by Beckmann and others (1959, p. 749-60) suggests that the 

buried terminus of the Lloyd is about 1 mile north 

of Eocky Point. The interface of the fresh and 

salt water in the deep confined aquifer is therefore 

probably less than a mile north of the shoreline.
E156 GEOLOGICAL SURVEY RESEARCH 1962

SEA 

LEVEL

Principal movement 

of ground water

600

FIGURE 226.2. Generalized geologic section of Centre Island, Nassau County, showing source and occurrence of salty ground

water.

Observed water levels in the deep aquifer have not 

changed appreciably in the past 10 years, so it is 

reasonable to assume that the position of the interface 

has not changed significantly during that period. Two 

observation wells located just west of Centre Island 

at Oak Neck Beach and Mill Neck are adequate to 

monitor any future permanent decline in head that 

might signify landward movement of the interface.

REFERENCES

Beckmann, Walter C., Roberts, A. C., and Luskin, B., 1959,

Sub-bottom depth recorder: Geophysics, v. 24, no. 4,

Oct. 1959, p. 749-60. 

Nassau County Planning Commission, 1961, Population sixty:

Mineola, N.Y., Nassau County Planning Conim., 26 p. 

U.S. Dept. of Commerce, Coast and Geodetic Survey, 1962,

Tide tables, high and low water predictions: Washington,

U.S. Govt. Printing Office, 283 p.
ARTICLE 227 E157

227. HIGHLY PRODUCTIVE AQUIFERS IN THE TACOMA AREA, WASHINGTON

By KENNETH L. WALTERS, Tacoma, Wash.

Work done in cooperation with the State of Washington Department of Conservation, Division of Water Resources

The well field of the city of Tacoma (fig. 227.1) 

is in an area underlain by glacial aquifers whose permeability may be the greatest of any in the Puget 

Sound lowland. The dominant physiographic feature 

of the well field is a channel formed by a stream 

that discharged from a glacial lake during the latest 

deglaciation of the Puget Sound lowland. This southR. 2E. R. 3 E.

FIGURE 227.1. Map of Tacoma area, showing location of city 

wells. Contours are in feet above mean sea level.

trending channel, in a part of the city known as South 

Tacoma, is partly filled with gravel; its surface is at 

an altitude of about 225 to 300 feet above sea level, 

about 70 feet lower than the till plain on both sides.

The principal aquifers underlying the well field 

consist of gravel deposits related to glaciations earlier 

than the one responsible for the channel. Probably 

the most important effect that the channel has on the 

productivity of the well field is that it localizes 

recharge in an area where till is thin or absent. These 

aquifers probably underlie a large part of the Tacoma 

area, but they are known to be of unusually high 

permeability only in an elongate area occupied in part 

by the Tacoma well field. The cause of the high 

permeability at this location is not known, but is 

probably due to the deposition of coarser or better 

sorted materials there than were being deposited elsewhere.

The city of Tacoma pumped water from a system 

of comparatively inefficiently constructed wells in the 

southern part of the present well field from about 

1903 until 1933. The selection of this site was influenced by the presence of a water pipeline through 

that area.

More intensive development of the well field was 

started in 1929 with the drilling of a 310-foot well 

of modern construction. This well, city well 1-A, was 

tested at a rate of 3,080 gpm (gallons per minute) 

with a drawdown of 60 feet, indicating a specific 

capacity of more than 50 gpm per foot of drawdown. 

Well 1-A was so productive that the city continued 

to drill more wells along the course of the channel 

as the demand for ground water increased. By the 

end of 1961 the city had 13 public-supply wells in 

operation along a 3-mile reach of the channel. In 

addition to these public-supply wells, the city drilled 

two wells in the north part of the channel for the 

operation of a heat-pump installation at its utilities 

building. The 13 public-supply wells have a combined 

short-term capacity of about 64 million gallons per 

day. Under conditions of continuous pumping, however, the combined daily production is about 51 million gallons; at this pumping rate the drawdown is 

not considered to be excessive.

Data describing the wells and their performance 

are presented in table 227.1.
E158 GEOLOGICAL SURVEY RESEARCH 1962

TABLE 227.1. Data on city-owned wells in the Tacoma well field

Well No. (USGS)

20/3-19Pl__-____.__

-18D1 ___

-18D2 _ _

-30N1_____ _____

-13H1__ _____

-19F1 __________

-13Jl-_____ _

-30L5 -___.__

-30C4____ ______

-18D3 . _ _

-30L6_ ____ .__

20/2-13J2____._ __

20/3-18C2 _ ___.

20/2-13 Al _______

-13A2__ ___ ___

Well No. (city)

1-A ______

2-A ______

2-B_ __ _ _

3-A_ ________

4-A___ __ _______

5-A____________

6-A ______

7-A____________

8-A _____

9-A ______

10-B ______

11-A_____ _ __ _

12-A ________

Utilities Bldg_ _ _

_____do_________

Depth

(feet)

306

172

70

312

204

352

179

298

297

113

83

113

167

260

85

Altitude 

(feet above 

mean sea 

level)

261

243

245

271

246

266

267

266

267

280

257

269

309

246

244

Year 

drilled

1929

1930

1949

1931

1930

1930

1939

1939

1939

1Q4Q

1956

1948

1957

1953

1953

Pumping 

rate 

(gpm) i

3,080

1, 150

(2, 025)

3,400

(3, 600)

3,650

(3, 920)

1,450

4,500

(6,240)

3,770

(3, 200)

920

(3,940)

3,890

(4, 250)

6,000

5,500

1,000

(930)

9,350

5,000

620

2,500

Drawdown 

(feet)

60

60

4. 5

63

92

68

74

67

60

31

25

122

31

2.2

Coefficient of 

transmissibility 

(gpd per ft)

2, 900, 000

270, 800

145, 000

184, 500

1, 250, 000

58, 000

114, 000

753, 000

Specific 

capacity

51

34

800

62

16

92

51

59

71

30

875

41

20

1,135

Static water 

level (feet 

below landsurface 

datum)

39

36

25

41

33

42

27

28

41

59

28

57

84

39

Date

11-60

11-60

11-60

11-60

11-60

11-60

11-60

11-60

11-60

11-60

11-60

11-60

11-60

6-53

1 Pumping rates in parentheses are for period of testing only.

The relation between specific capacity and coefficient 

of transmissibility for individual wells varies considerably. Although most of the wells tap more than 

one water-bearing zone, the wide range in specific 

capacity probably is due more to differences in construction of the individual wells than to variations in 

aquifer characteristics.

The well field is used to supplement surface-water 

supplies, and usually is not pumped continuously for 

long periods of time. The annual pumpage from the 

well field increased from 158 million gallons (485 

acre-feet) in 1930, to 6,754.7 million gallons (20,740 

acre-feet) in 1961 (table 227.2). Since 1930, a total 

of 39,584 million gallons (121,520 acre-feet) has been 

pumped from the well field.

Even though the withdrawal has been substantial, 

periodic measurements of static water levels in the 

field indicate that the depth to water has not changed 

appreciably in 30 years. For example, the earliest 

recorded water-level measurement in well 1-A, on 

February 11, 1930, shows that the depth to water was 

then about 42 feet below the land surface. On December 21, 1960, the depth to water in this well was 

about 38 feet.

The records indicate that ground water can serve 

as a reliable emergency supply to Tacoma in case of 

failure of the surface supply, and that the aquifers 

could yield even more than the present production 

without serious depletion.

TABLE 227.2 Annual pumpage, Tacoma well field

Year 

1930 _____

1931 ____

1932_.______

1933 _____

1934 _____

1935 --

1936 ____

1937__-_____

1938 _____

1939 ____

1940 ____-

1941 _ _

1942 _____

1943 _____

1944 _____

1945 _____

Pumpage from 

well field 

{million gallons)

_______ 158.0

_______ 168.5

____ 381.9

_______ 606.9

_______ 421.0

_______ 278.0

_______ 258.5

_______ 524.0

412.7

----___ 428.7

_______ 880.9

_______ 619.9

_______ 611.3

_______ 208.3

___ __ 822.2

.___._. 1.517.2

Year 

1946 _____

1947 ____

1948 _____

1949 _____

1950 ____

1951 -___

1952_. ______

1953 ____

1954__ ______

1955 -___

1956 _.__

1957 --_-_

1958 _-__

1959 __._

I960 _

1961 _

Pumpage from 

well field 

(million gallons) 

_______ 891.6

_______ 1,262.6

_______ 1,062. 8

_______ 1,059. 5

___.__. 1,289.0

_______ 1, 140.5

_______ 826.7

_______ 1,391. 1

_____-- 992.5

____-._ 2, 143. 5

_______ 3, 194. 5

_______ 1, 141. 8

_______ 1,225.7

_______ 2,722. 7

.___-_. 4,186. 9

_______ 6.754.7
ARTICLE 228 

SURFACE WATER

E159

228. LONG-TERM FLOOD FREQUENCIES BASED ON EXTREMES OF SHORT-TERM RECORDS 

By W. J. CONOVER and M. A. BENSON, Washington, D.C.

The usual length of flood-discharge records does not 

permit the assignment of a reliable probability to 

occasional extraordinary floods. For example, we know 

that 1,000-year or 10,000-year floods may occur at any 

time, and that such floods are actually being experienced, but the period of record is so short that 

we cannot assess the recurrence interval of these rare 

events.

A solution to this problem has been planned that 

consists of three parts. The first part is the reduction of all flood peaks to a comparable basis, so that 

records at all sites may be combined. The second 

part is a study of the degree of statistical dependence 

between flood records and the reduction of all records 

to an equivalent number of independent records. The 

third part is the combining of all the independent 

records and the assigning of recurrence intervals or 

probabilities to events in the combined group.

The first two parts of this solution are believed to 

be within reach, requiring only a study of several alternative procedures and a great deal of work in 

abstracting the hydrologic characteristics of drainage 

basins. The subject of this article is a solution to the 

third part of the problem.

The combining of a set of records that are independent, as in the "station-year" method familiar to 

hydrologists, consists of adding all records end-to-end 

and considering them as a single record in assigning 

probabilities. However, continuous records do not 

exist at all sites where flood data are available, so 

that a modification of the station-year method is necessary.

The solution presented here will enable the evaluation of recurrence intervals for the peak flood events 

at all sites for which the maximum peak within a 

common period of time is known. This solution permits information at long-term gaging stations to be 

combined with that at ungaged sites or short-term 

stations where there is no continuous record but where 

the highest peak is known. The recurrence intervals,

or probabilities, would be defined for the combined 

period of record. For example, by using only the 

highest annual discharges in 50 years at 20 sites where 

the flow at each site is independent of the flow at the 

other sites, it would be possible to rank the extremes 

and to determine their probabilities within a 1,000- 

year period. The recurrence interval of the highest 

extreme would be close to 1,000 years; the assignment 

of recurrence intervals to the other ranked extremes 

is not as obvious. This paper shows how to determine 

recurrence intervals for all of the ranked extremes of 

independent records.

Let nk independent values from any continuous 

distribution be divided at random into n groups of k 

elements each. Order the elements of each group, 

starting from the highest to the lowest, and then order 

the groups according to the value of the highest 

element of each group. Denote the first element of 

the ith group by an. Thus a-n is the largest value of 

all nk values. Also a2 i is greater than &22, but not 

necessarily greater than «i2 . The generalized form of 

the array follows:

Elements 

within each

group

1

2

3

A

i

an

«12

a\k

2

«21

«2fc

Group

3

032

03Jfc

n

a n\

«n2

Cf n3

a n k

i-1 

m=0

Now suppose we consider an additional independent 

value, x, from the same continuous distribution. What 

is the probability that x will exceed a« (^=1, 2, . . ., 

n)?

By use of a basic theorem in order statistics, the 

solution to this problem can be shown to be

n\

667422 O - 63 - 12
E160 GEOLOGICAL SURVEY RESEARCH 1962

where p(att) represents the probability of an additional 

independent value, x, exceeding an (i=l, 2, . . ., n). 

This may be called the exceedance probability. By 

using the above formula, n=20, and k=5Q, the table 

below was computed. To compare the theoretical 

probabilities with the results of sampling from known 

distributions, three trials were made, using an extremevalues distribution for the first trial and the normal 

distribution for the other two trials. In each trial, a 

table of random numbers was used in selecting the 

samples. Twenty groups of 50 random numbers each 

were used in each trial, and the numbers in the groups 

were ranked in the manner previously described. Then 

the actual rank among the 1,000 random numbers was

recorded for each of the numbers that headed a group, 

and the corresponding exceedance probabilities were 

plotted in figure 228.1. Theoretical probabilities given 

by the values shown in the table below were used to 

define a smooth curve of theoretical values. Figure 

228.1 demonstrates the close correspondence between 

the sampled and the theoretical values.

Probabilities and recurrence intervals of extremes, n=20, k=50

0.001

Exceedance 

probability

0. 000999

. 002050

. 003157

. 004329

. 005573

. 006896

. 008313

. 009835

. 01148

. 01328

Recurrence 

interval

P(au)

1001

488

317

231

179

145

120

102

87. 1

75.3

Exceedance 

probability

11_._.__ 0.01524 

12___.__ .01743 

13_.____ .01988 

. 02268 

. 02592 

_ . 02980 

. 03463 

_ . 04102 

_ . 05051 

. . 06913

Recurrence 

interval

T=-

14.

65.6

57.4

50.3

44. 1

38.6

33. 6

28.9

24.4

19.8

14. 5

0.01-

6 8 10 12 14 16 

RANK WITHIN GROUP OF EXTREMES

18

FIGURE 228.1. Theoretical curve compared with random trials, 

for n = 20 and fc = 50. Dots indicate values from trial 1, open 

circles from trial 2, and triangles from trial 3.

The formula derived for this article has made it 

possible to base the probabilities of rare floods on 

only the extreme events experienced at all sites where 

flood information is available, provided, first, that the 

flood data are comparable and, second, that they are 

independent. Although records of peak discharge, in 

their usual form, are not comparable and are not independent, it is expected that future studies will develop methods of expressing flood records on a comparable basis and of reducing them to an equivalent 

set of independent records. The formula developed 

in this paper can then be applied to assign probabilities to the extreme flood events that have occurred.
229.

ARTICLE 229 

FLOOD PEAKS RELATED TO HYDROLOGIC FACTORS IN THE SOUTHWEST

By MANUEL A. BENSON, Washington, D.C.

E161

In a previous investigation, Benson (1962) examined 

the relations between annual peak discharge and many 

hydrologic factors in New England, a humid region. 

A similar investigation has now been completed for 

the western part of the Gulf of Mexico basin, a region comprising most of Texas and New Mexico and 

small parts of Louisiana and Colorado. This basin 

is semiarid to arid in its larger part west of Austin, 

Tex., and humid in its eastern smaller part. Figure 

229.1 shows the extent of the region studied.

Both physiography and climate vary widely in the 

area studied. The altitude ranges from sea level at 

the gulf to 14,000 feet in the headwaters of the Eio 

Grande in Colorado. Mean annual rainfall ranges 

from 58 inches in Louisiana to 8 inches in New Mexico. 

The part of the mean annual rainfall that appears as 

stream runoff ranges from 40 percent in Louisiana to 

1 percent in western Texas and New Mexico and is 

less than 5 percent for two-thirds of the region. The 

annual peak floods (highest momentary peak discharges each water year) tend to occur in different 

seasons in different parts of the region. In Louisiana 

and eastern Texas, annual floods generally are caused 

by winter and spring rains, in central Texas by spring 

storms and fall cyclonic storms, in west Texas and

103° 94°

\COLORADO

I ^' \ i"j\i - Southern boundary

' / \\ °f snowmelt area i

' | ( ^%__^ I I OKLAHOMA

/ ' \ ' T^ L

f NEW MEXICO ) I '-^ ^-^x

^ ) / ) ' X ._. ^\_.aC--'- / f /i \ v «-T>.. <s' S^«-_ J.

FIGURE 229.1. Extent of western part of the Gulf of Mexico 

basin studied, and boundary between snowmelt flood and rain 

flood areas.

most of New Mexico by summer thunderstorms, and in 

Colorado and northern New Mexico by spring snowmelt.

A field and library search for historical information provided data concerning floods that occurred 

before the period of record and helped in determining 

recurrence intervals of the rare floods in recent years. 

This information, much of which was not previously 

available for analysis, provided the principal basis 

for establishing the values of 50- and 100-year floods 

that were used in the analyses. Similar values for 

shorter recurrence intervals were based on annual 

peaks for the period of record.

In the humid-region study in New England, flood 

peaks of recurrence intervals varying from 1.2 to 

300 years were related to 6 statistically significant 

variables. In order of decreasing importance, the 

variables are drainage-area size, main-channel slope, 

an orographic factor, area of lakes and ponds, winter 

temperature, and rainfall intensity and duration.

In the study of the largely semiarid region, these 

and many other physiographic and climatic variables 

were considered in relation to peak discharge. Included were basin altitude, length, rise, shape, and 

orientation; stream-order number; prevailing wind 

direction; total annual snow depth; water content of 

snow on April 1st; spring temperature; mean annual 

precipitation and runoff; runoff-precipitation ratios; 

average annual number of thunderstorm days; and 

the percentage of forested area. In addition, an attempt was made to develop indices based on channel 

geometry that could represent the effect of channel 

storage. A hydrologic soil-cover complex developed 

by the Soil Conservation Service for representing infiltration capacity was used to represent the effect of 

soils on peak flow.

The study showed that better results could be obtained by distinguishing the area where annual peaks 

are caused by snowmelt from the area where annual 

peaks are caused by rain. The boundary between the 

small part of the area affected by snowmelt and the 

larger part, called the rain flood area, is shown in 

figure 229.1 as a dotted line roughly through 'Santa 

Fe, N. Mex. By use of multiple-regression techniques, the relation of peak discharge to hydrologic 

variables for both the rain flood and snowmelt flood 

areas was determined. These relations are linear 

when the logarithms of all quantities are used.
E162 GEOLOGICAL SURVEY RESEARCH 1962

Within the rain flood area, the annual peak discharge was found to be related to seven statistically 

significant variables, which are, in approximate order 

of decreasing importance:

A, drainage-area size, in square miles;

/, rainfall intensity, the 10-year 24-hour rainfall, in inches.

S, maiii-channel slope, in feet per mile;

L, basin length, in miles ;

St, area of lakes and ponds, expressed as percentage of the 

total drainage area and increased by 1 percent;

R, median runoff-precipitation ratio, for months: when annual peaks occur; and

N, average annual number of thunderstorm days.

The form of the relation is

QT=aA»ScStdLeI'NgRh,

where QT is the T-year annual peak discharge in cubic 

feet per second and a, b, c, . . .h are the regression coefficients. Values of the coefficients for selected values 

of T are shown in table 229.1.

TABLE 229.1. Regression coefficients, rain flood area

T (years)

1.2............

2.33 _ ........

5.... _ .......

10.............

25. . .-

50.............

a

375

QQQ

539

9fiQ

182

3,670

b

1.6

1.3

1.1

1.0

1.0

c

0

0 9

.2

.3

.5

.4

d

0 9

-1.3

-1.5

-1.6

-1.0

-1.1

e

-1.4

-1.1

 S

-.7

-.5

-.5

/

0

0

1 1

1.5

i a

1 9

1.4

ff

-0.5

-.7

__ Q

-.7

-.7

-1.0

h

0.7

.4

.3

.3

.3

The results of this study may be compared with 

those of the study for New England. Importance of 

the variables is judged by their relative effect in explaining the original variation in the size of flood 

peaks. The size of drainage area is the most important variable, as in New England. Rainfall intensity 

is second in importance, because of the extreme variability in rainfall characteristics in the rain flood 

area of the Southwest in contrast to New England, 

where rainfall intensity was the least effective variable. Slope is next in importance in the Southwest; 

in New England its effect was second to that of drainage-area size. Basin length is next in importance, 

whereas it was not an effective variable in New England. Large losses to channel infiltration and reductions in peak discharge due to channel storage are 

common in the Southwest where floods often flow 

through channels that were initially dry. Thus, in 

the rain flood area, peak discharge is related inversely 

to channel length as shown by the negative exponent 

of L. Storage in lakes and ponds is an important 

variable in both New England and the Southwest.

Within the snowmelt flood area, peak discharge was 

found related to six statistically significant factors, 

which are:

A, drainage-area size, in square miles;

S, main-channel slope, in feet per mile;

St, area of lakes and ponds, expressed as percentage of the

total drainage area, and increased by 1 percent; 

P, mean annual precipitation, in inches; 

E, altitude, in feet; and 

N, average annual number of thunderstorm days.

Of these, drainage area is by far the most important, 

and the rest are of smaller but nearly equal importance. 

The form of the relation is :

Table 229.2 shows the variation of the coefficients for 

selected values of T.

TABLE 229.2. Regression coefficients, snowmelt flood area

T (years)

1.2. 

2.33.. ....... . ...

5 

10 

25 

log 10 a

21. 0156

15.1943

16. 3258

18. 2182

21. 9216

b

0.2

.3

.4

.3

.2

c

-0.9

-.7

-.6

-.6

-.7

d

-0.9

-.9

-.8

-.9

-1.3

e

5.0

3.6

3.3

3.5

3.4

/

-4.2

-2.8

-3.1

-3.5

-4.5

ff

-4.2

-3.1

-3.0

-3.2

-2.5

Within the snowmelt flood area of the Southwest, 

slope has a negative coefficient, contrary to what was 

found in the rain flood area and in New England. 

The negative coefficient can be explained by the snowmelt characteristics within this area. In the humid 

climate of New England, the normal 3 to 4 inches of 

rain per month in the spring, coupled with rising 

temperatures, results in rapid runoff of combined rain 

and snowmelt water. In the Southwest, where rain 

is scant in most of area, the melting of snow in the 

spring is caused mainly by rising temperatures and 

occurs over a longer period of time. In addition, the 

melting starts at low altitudes and progresses to higher 

altitudes. In a basin with a large rise and hence a 

steep slope, only a small portion of the basin is contributing melt water at one time. Therefore, other 

things being equal, a steep slope is associated with a 

small peak discharge. The negative relation of the 

rate of snowmelt to slope overbalances the effect of 

slope in increasing the rate of travel of the water; 

the net result is a negative relation of peak discharge 

to slope.

The standard error of estimate is a measure of success in explaining the variability in a dependent variable, such as peak discharge, by means of statistically 

significant independent variables, and is also a meas-
ARTICLE 230 E163

ure of the error in predictions made by using the 

developed relation. In the New England study the 

standard error of estimate varied with T and ranged 

from 23 to 37 percent. In the Southwest, the standard 

error for floods with recurrence intervals of 2.33 years 

and more ranged from 42 to 62 percent.

The fact that standard errors are larger in the 

Southwest than in New England is attributed to the 

following:

1. Lack of efficient quantitative indices for the factors of geology and soil influences that are important in this region.

2. Difficulty in defining the higher flood peaks because of shifting sand-bottom streams and short duration of thunderstorm flood peaks.

3. Lack of adequate definition of the noncontributing parts of some drainage areas.

4. Erratic chance occurrence of storms within a 

largely semiarid region.

The error attributable to the first three of these 

items may be reduced as measurement techniques improve and hydrologic knowledge increases. The erratic geographic occurrence of storms, typical of semiarid conditions, is believed to account for a large part, 

perhaps the major part, of the unexplainable deviations present after all known physical and climatic 

factors are taken into account. On this assumption, 

the average flood experience within the region is a 

better basis for establishing hydrologic relations than 

is the average flood experience at any one point. The 

part of the residual error that is attributable to the 

variability of storm occurrence cannot be decreased 

except by a long period of time sampling of runoff 

experience.

REFERENCE

Benson, M. A., 1962, Factors influencing the occurrence of 

floods in a humid region of diverse terrain: U.S. Geol. 

Survey Water-Supply Paper 1580-B. (In press)

QUALITY OF WATER

230. DISSOLVED-SALT CONTRIBUTION TO GREAT SALT LAKE, UTAH 

By A. M. DIAZ, Topeka, Kans.

Work done in cooperation with University of Utah

Estimates based on weighted-average dissolvedsolids content of surface discharge to Great Salt Lake 

indicate that the dissolved-mineral contribution to 

Great Salt Lake by surface streams was about 2 million tons for the 1960 water year. This estimate exceeds previous estimates (Eardley and others, 1957) 

of approximately 1 million tons annually and is based 

on systematic measurements during the water year. 

The estimated contribution of 2 million tons was by 

streams, drains, canals, and springs; it does not include mineral matter contributed by ground-water 

inflow.

Great Salt Lake is underlain by lake deposits of 

Pleistocene Age that on the east side of the lake exi University of Utah Research Fund and Uniform School Fund.

tend to the foot of the Wasatch Mountains. Large 

volumes of ground water move toward the lake 

through these materials and probably contribute a 

sizeable dissolved-mineral increment directly to the 

lake each year. However, the extent to which groundwater inflow affects the total salt contribution cannot 

be determined from available data.

The table below summarizes the computations of 

dissolved-solids contributions for the 1960 water year. 

Fifty-seven percent of the estimated tonnage was contributed by the three major streams in the basin 

the Bear, Weber, and Jordan Rivers; 25 percent was 

from drains discharging industrial and municipal 

wastes; 17 percent was from springs that discharge 

along the lakeshore; and 1 percent was from small 

creeks and drains along the east shore between the
E164 GEOLOGICAL SURVEY RESEARCH 1962

Weber and Jordan River basins. All surface discharge, summarized in the table, enters the lake along 

its north, east, and south shores. The lake is bordered 

on the west by the Great Salt Lake Desert and receives no significant inflow from that area.

Estimated contribution of dissolved solids to Great Salt Lake, 1960

water year

Dissolved

Source salts, in tons 

Bear River and Blue Springs Creek _______________ 700, 000

Weber River and miscellaneous inflow to Ogden Bay_ 140, 000 

Jordan River_____________________________ 320, 000

Kennecott drain, Garfield drain, Salt Lake City sewage canal, and sewage contributed by other 

communities in the east-front area_____________ 510,000

Springs discharging to lake along north and south

shores_____________________________ 340, 000

Small streams along the east shore, not included

above.-____________________________ 24, 000

Total_________________________ 2, 034, 000

Runoff to Great Salt Lake is derived principally 

from precipitation on the Wasatch and Uinta Mountains. Although streams draining these mountains 

are highly regulated, they are the major source of 

inflow to the lake. Rocks in the drainage areas range 

in age from Precambrian to Quaternary. The Precambrian and Paleozoic rocks generally crop out in 

the higher parts of the mountains, whereas Mesozoic 

and Cenozoic rocks crop out at lower elevations. The 

water of streams draining the Precambrian and Paleozoic rocks in the higher parts of the basin is of the 

calcium bicarbonate type and is of excellent quality, 

having a dissolved-solids content of less than 100 parts 

per million. However, the discharge to the lake itself 

is generally of poorer quality, having dissolved mineral matter from the rocks over and through which it 

passed in the lower parts of the basin and having been 

concentrated by addition of domestic sewage and by 

extensive agricultural and industrial use. Conseqently, the quality of water at the mouths of the individual streams varies considerably. The wide variation in chemical quality between the three major 

streams is shown by the annual weighted-average analyses shown in figure 230.1.

The environmental differences between each of the 

major drainage basins is reflected by the predominance of sodium and chloride ions in the Bear River, 

calcium and bicarbonate ions in the Weber River, and 

sodium and sulfate ions in the Jordan River. The 

difference in chemical composition of water in the 

lower Bear and Weber Rivers may be attributed 

mainly to differences in geologic environment, whereas 

the chemical composition of water in the lower Jor51 :Q 55 10 

EQUIVALENTS PER MILHION

Figurelabove; p^ttfern is dissotved^ 

solids cqntent, in parts per njilHon

FIGURE 230.1. Average chemical composition of water in Great 

Salt Lake and its major tributaries, 1960 water year.

dan River is affected significantly by intensive agricultural and industrial use.

The initial phase of this investigation was made 

during a relatively dry period, during which the annual runoff was slightly less than 60 percent of normal. The greatly reduced inflow to Great Salt Lako 

was reflected by a decreasing stage of the lake. At 

the end of the 1960 water year the surface of Great, 

Salt Lake was 4,193.40 feet above mean sea level, tho 

lowest stage in 100 years of record. The previous 

low stage was recorded in 1940 when the lake reached 

a level of 4,193.75 feet. At the 4,193.40-foot level the 

lake contains about 10 million acre-feet of brine, and 

the lake is about 64 miles long and a maximum of 30 

miles wide. The average depth of the lake at this 

stage is less than 10 feet.
ARTICLE 231 E165

During the 1960 water year, 20 samples of brine 

were obtained from different parts of the lake. Chemical analyses of the samples indicate that the brine 

was nearly saturated with sodium chloride, and that 

the average dissolved-solids content of the lake brine 

was about 27 percent. At this concentration, 1 acrefoot of brine contains about 450 tons of salt; therefore the dissolved-salt content of the lake was 4.5 X 

109 tons at the end of the 1960 water year.

REFEBENCE

Eardley, A. J., Gvosdetsky, Vasyl, and Marsell, R. E., 1957, 

Hydrology of Lake Bonneville and sediments and soils of 

its basin: Geol. Soc. America Bull., v. 08, p. 1141-1202.

ANALYTICAL HYDROLOGY

231. ORIENTATION OF AXES FOR CALCULATING DISTRIBUTION OF TRANSMISSIBILITY FROM

WATER-LEVEL ALTITUDES

By R. W. STALLMAN, Denver, Colo.

Subsequent testing by E. A. Sammel (oral communication, July 1962) of a procedure described by Stallman 

(1956), shows that the sensitivity of calculations for 

defining the coefficient of transmissibility, T, from 

water-level altitudes is dependent on the orientation of 

the x and y axes in the field of flow. This article sets 

forth the conditions under which sensitivity is a function 

of axis orientation.

Consider the following differential equation for 

steady ground-water flow in a confined nonhomogeneous aquifer:

where T is the coefficient of transmissibility of the 

aquifer and x and y refer to distances in a cartesian 

coordinate system in the plane of flow in the aquifer. 

It has been proposed (Stallman, 1956) to substitute 

finite-difference approximations of the head differentials 

in equation 1 at two or more points in the field of flow 

and to solve the resulting set of equations for the rate 

and direction of change in T within the aquifer. 

Unfortunately, water-level altitudes applicable to 

equation 1 are seldom known to a high degree of 

accuracy at field sites. Therefore errors occur in the 

estimates of head differentials made from field measurements, and these errors are reflected in the calculated 

rates of change in T derived from finite-difference 

approximations of equation 1.

All the gradients contained in equation 1 and the 

angular displacement of each from the x axis of a 

cartesian coordinate system are shown on figure 231.1. 

Angles to the vectors are taken from a common reference 

line which is at an angle 7 from the x axis. The- effect

FIGURE 231.1. Sketch showing definitions of vectors in field of 

ground-water flow.
E166 GEOLOGICAL SURVEY RESEARCH 1962

on the error, E, of rotating the x axis with respect to 

the position of the vectors is determined by reconstructing equation 1 to account for error and differentiating to find those values of 7 for which dE/dy is 

zero that is, when E is a minimum.

If the error terms due to inaccuracies in the data 

are included, equation 1 may be written as follows:

=E, (2)

where E is the total error in equation 2 due to finitedifference estimates of head differentials, H is the 

gradient of the first derivative of head, f(H) is the 

error in the estimate of (d2/*,/b?/2+d2Vd?/2), / is the 

gradient of head, jx (I) is the error in the estimate of 

dh/bx, and/j,(7) is the error in the estimate of bhfoy.

In finite-difference methods, <¥h/dx2 and &h/(ty2 generally are evaluated separately. Furthermore, the error 

in each estimate is approximately inversely proportional to the value of the second differential. If the 

second differentials are expressed in trigonometric form, 

it is seen from the definition of H and from figure 1 that

and

_H cos (T+T) 

where A is a constant.

[_ +H sin

(3)

(4)

Similarly, errors in the estimates of the first differential are approximately inversely proportional 

to the respective values of the first differentials. Thus 

the second and third terms in equation 2 can be written 

as follows:

(a+7)[/cos (0+7) 

, C

COS (0+7)] (5)

and 

bT p"N» I "1 Lfy+^(/) \=B sin sn

sin (0+7)], (6)

where B is the gradient of transmissibility and C is 

a constant.

The total error, E, is given by the sum of all the 

terms on the right-hand side of equations 3 to 6. The 

derivative of the total error is found to be

ay

(r+7)-sin

TA ["sin3 (T+7)-cos3 (T+T)"| 

H L sin2 (r+7) cos2 (r+7) J 

sin (0-a)

/ sin2 (0+7) cos2 (0+7)

1

J

(7)

If dEfdy is set equal to zero, inspection of equation 

7 shows that the error is minimal when 0 a=0 and 

T+7=ir/4 radians. In a physical sense this means 

that E"is least if the slope of the water surface is parallel 

to the direction of maximum spatial rate of change of 

T and if the spatial rate of change in the first differential of head is greatest along a line which is at an angle 

of ?r/4 radians from the x axis. However, such an 

orientation of vectors is found in few regions of natural 

flow.

Integration of equation 7 yields

(r+7)+COS 

TA. (T+7)+csc 

. T2

T>ri 

-2 ^ sin (0-a)[cot 2(0+7)]?i. (8)

If 7i 72 is assumed to be very small, the change in 

error produced by a small angular rotation of the 

reference axis of figure 231.1 can be approximated. 

Then if 2(0+71 )=ir/2 or 3?r/2 radians, the last term 

in equation 8 vanishes. Therefore, if 71 is taken to 

be zero, the error in the last term can be minimized 

by orienting the x axis 7r/4 radians from the gradient 

of head. In regions where T=0+ir/2 or T=0+ 

3ir/2 and 0=7r/4, all terms in equation 8 are zero.

Any wide deviation from the ideal angular relations 

in the flow field which produces zero or negligible error 

may preclude the application of equation 1 for accurate 

estimation of the rate of change in T from water-level 

altitudes. Error cannot be defined accurately from 

equation 8 in all cases because the constants A and C 

generally are unknown. In the general flow problem, 

the investigator should select T and 0, with 7=0, 

so that all the bracketed terms in equation 8 are small. 

For most problems such a selection will suffice for producing relatively accurate estimates of the distribution 

of T.

REFERENCE

Stallman, R. W., 1956, Numerical analysis of regional water 

levels to define aquifer hydrology : Am. Geophys. Union 

Trans., v. 37, no. 4, p. 451-460, August.
ARTICLE 232 El67

232. A METHOD OF DETERMINING THE STORAGE-OUTFLOW CHARACTERISTICS OF NONLINEAR RESERVOIRS

By JOHN SHEN, Phoenix, Ariz.

In the analysis of small-basin flood characteristics, 

it is often desirable to define the storage-outflow relations, from certain observed outflow hydrographs. Such 

storage effects may be those due to manmade reservoirs 

or those due to natural storage in ponds, channels, or 

minor depressions.

Whereas a linear storage-outflow relation is generally assumed to be applicable to many natural basins, 

experience by Mitchell (1962) with streams in Illinois 

indicates that nonlinear storage is a condition that 

occurs with sufficient frequency to warrant careful 

consideration.

Methods for determining the linear storage-outflow 

characteristics from the recession outflow hydrographs 

have been suggested by Linsley and others (1958, 

p. 151-155) and Mitchell (1962). A more generalized 

method which is applicable to nonlinear as well as 

linear reservoirs is described below.

Fundamentally, the storage-outflow relation of a 

nonlinear reservoir may be expressed (Mitchell, 1962)

(1)

in which S is the storage, Q is the rate of outflow, 

K is the storage coefficient, and x is an exponent. 

Values of K and x may be variable, but for many sites 

they appear to be nearly constant.

The basic relation for inflow, outflow, and storage 

is expressed by the equation of continuity:

(2)

where 7 is'the rate of inflow and t is the time.

Upon differentiating equation 1 and substituting it 

into equation 2, we have

Thus,

I-Q=KxQ*- 1

dQ_ I-Q 

dt KxQx~ 1 '

dt (3)

(4)

For the recession limb of an outflow hydrograph, the 

inflow ceases (7=0), hence

dQ_ Q dt~ *- 1 '

or

dQ

dt (5)

The curvature of the recession hydrograph is represented by its second derivative

d*Q^ x-2 

dt2

or

___J^ x-2 

dt2 ~K' x r-l (6)

The significance of equation 5 may be realized if we 

rewrite it in finite-difference form and take the logarithm 

of both sides. Hence we have

log -=-(*-2) log Q-log (Kx). (7)

It may readily be seen that equation 7 is a linear logarithmic equation and may thus be represented by a 

straight line on log-log graph paper. If the value of

 ¥ is plotted as the ordinate and the value of Q as

the abscissa, then the slope of the line would be represented by (x 2) with the point of intercept (log 

Q=0) at Kx. Hence, the values of K and x may be 

determined separately.

From the slope of the line, (x 2), the value of x 

may be derived. Furthermore, at Q=l, log Q=0, 

and equation 7 becomes

or

A*

_ 

Kx (8)

Hence, the value of Km&y be determined once the value 

of x is known.

Two special cases arise which are of particular interest:

When x=2, equation 5 becomes

^2__L_. (Q) 

dt~~ 2K ^ }

Thus, the recession curve is a straight line on a linear 

scale having a constant slope of ^-

When x=l, equation 5 becomes

dQ__ Q dt~ K' (10)

and the reservoir is linear. The value of K can, of 

course, be readily determined from the recession hydrograph or by any other conventional methods.
El68 GEOLOGICAL SURVEY RESEARCH 1962

For illustration, figure 232.1 shows a typical outflow

10 11 12

FIGURE 232.1. A typical outflow hydrograph of a nonlinear 

reservoir. K=2, x = %.

hydrograph of a nonlinear reservoir, for which K= 2 

time units and x=%. The hydrograph of inflow to 

the reservoir is assumed to have the form of an isosceles 

triangle.

Figure 232.2 shows a plotting of versus Q on

log-log scales for the recession limb of the outflow hydrograph shown in figure 232.1. It may be seen that the 

plotted points define a straight line having a slope of 

1.5, hence x=%. Moreover, the line intercepts the

point (Q=1.0, -T?=I-°)- Tnus> the val^e of K can

be computed from equation 8 as 2.0.

The method described above furnishes a simple 

means of analyzing the storage-outflow characteristics

1.0 

0.8

0.6

0.4 

0.3

I

0.2

0.10

0.08

0.06

0.04

0.03

0.1 0.2 0.3 0.4 0.6 0.8 1.0 2

Q 

FIGURE 232.2. Plotting of -^ versus Q.

of small drainage basins. Further evaluation of this 

method as applied to field data is being undertaken in 

Illinois under the direction of William D. Mitchell.

REFERENCES

Linsley, R. K., Jr., Kohler, M. A., and Paulhus, J. L. H., 1958, 

Hydrology for engineers: New York, McGraw-Hill Book 

Co., Inc., 340 p.

Mitchell, W. D., 1962, Effect of reservoir storage on peak 

flow: U.S. Geol. Survey Water-Supply Paper 1580-G.
ARTICLE 233 E169

233. AN ANALOG SOLUTION OF THE TURBULENT-DIFFUSION EQUATION FOR OPEN-CHANNEL FLOW

By JOHN SHEN, Phoenix, Ariz.

The basic equation for two-dimensional turbulent 

jvo. , , -, . -, ,TVT i i diffusion in an open channel may be taken (Nobuhiro

Yotsukura, written communication. May 10, 1962) as ' ' J ' '

b.D

be b£=_b_ 

ba; by

bc (1)

where c is the concentration of diffused material, t is 

the time, x is the distance in downstream direction, y 

is the vertical distance from the channel bottom, u is 

the velocity in x direction, and D is the diffusion coefficient assumed to be equal to the coefficient of momentum exchange. Thus, equation 1 represents the 

turbulent-diffusion process in spatially two-dimensional 

steady flow, for which u as well as D is a function of 

y only.

Electrically, equation 1 may be simulated by a twodimensional resistance-capacitance network. A generalized scheme for such a net is shown in figure 

233.1. The upper boundary for the net is the water

For a particular node under steady-flow condition, ^ 

is of constant value. Thus letting -^ =j(y) K and 

substituting it into equation 2, we have

(3)

by V by/ by2 ^ by

And, at a node point (/, m) (see fig. 1), we may write 

equation 3 in finite-difference form as

*y( b /n D Wt.~ bC\ D7-, ~ / Ci ,m+l\Cl ,m 1 ^Cl, m\ ( W )

Ci.m-i Ci.

or

^ / ^,\\ /D j£ \ 

* ~^~' 2Ay /

' 

««

- - -1

V

^, (.

' ;

i

p. . .., ,c 

^ <

, ,» ,:-,,

; i

J, 4»+

) ^,.1

h ^

1 l

-*. ..

'.

).. ..- '*-. J-f-liV

i

m-

~ -« ~~~3 As

@ 1 and

i

equ

byA

F

dev

figu

"»-l^ A2/ A2/

equation 4 may be rewritten as

1 ) (ci. m-i ei, m). (5) 2Ay2

From Kirchhoff's law, an analogous equation may be 

developed for the electric-resistance network shown in 

figure 233.2. The electrical equation is

i)+ (Em-i Em). (6)

Equation 6 is seen to be analogous to equation 5.

FIGURE 233.1. A schematic two-dimensional net for turbulent 

diffusion.

surface and the lower boundary is the channel bottom. Thus, iA is equivalent to ^- ( D ^ V and Em+i, Em and 

The solution to equation 1 is developed term by term ^ ^ ^ 

in the following discussion. E^ are equivalent to c,, m+,, c,. m and c,. m_! respectively. The necessary conditions for the two equations 

SOLUTION FOB THE TEBM (D ^ to be completely identical are

The term ^- [D^ ] may be rewritten as TO+1/2 i 

oy \ by/ and

_b /D *c\=D WCtdD be 

by V by/ by2 "1" by * by (2) Rm-V2==~i

2Ay2

(7)
E170 GEOLOGICAL SURVEY RESEARCH 1962

FIGURE 233.2. A resistance network for solving the term

*YD fc\

Oy\ OyJ

Hence the values of Rm+y2 and Rm-x may be determined 

once the values of Dm+1 , Dm, and Dm-\ are known.

<V SOLUTION FOB THE TERM u ^Lda;

dc Similarly, if we write the term u ^ in finite-difference J> da;

form, we have

dc

U

dc

2teL

(Ci+i, m Cl

\ 2Ax

or

~U . 

(8)

Again, an electrically analogous network may be 

developed as shown in figure 233.3. From figure 233.3, 

we may write the electrical equation,

(9)

1+1/2 1-1/2

Comparing equations 8 and 9, we have the following 

equivalents:

dc

1+1 = C i+i,

and

2Ax

(10)

Hence, the values of Rt+y2 and Ri-y2 may be determined 

from the known value of um. It is seen that in order

FIGURE 233.3. A resistance network for solving the term u

to achieve the complete analogy, Ri+y2 must be a negative resistor; that is, the current flows in the opposite 

direction with respect to the voltage potential.

Whereas a negative resistor can be manipulated by 

means of operational amplifiers, an alternate approach 

would be considerably simpler.

If we rewrite equation 8 as

u

dc 

ox

Urn '2Ax (U)

we may accomplish its electrical counterpart by means 

of one operational amplifier at each node. Figure 233.4

FIGURE 233.4. An operational amplifier circuit for solving the

dc

term u^-. 

da;

shows such a circuit, in which the operational amplifier 

is used as an adder-subtractor.

From the circuit in figure 233.4 we may derive the 

following relation:

IB=- (12)
Again, comparing equations 11 and 12, we have

ARTICLE 233 E171

iB= u dc

I, m

and

2Ax

(13)

Thus, equation 11 is satisfied.

SOLUTION FOB THE TERM

dc

The term ^ can be easily simulated by connecting

a capacitor at each node point, as shown in figure 

233.5. As the current-voltage relation for a capacitor is

(14)

dc FIGURE 233.5. A capacitor circuit for solving the term j-'

dc 

we have the equivalents ic ?r> when C l.

Thus, by interconnecting the circuits in figures 233.2,

233.4, and 233.5 at a common node (l,m) (fig. 233.6), 

we have

^e *A + £B- (15)

And, upon substituting equations 6, 9 and 14 into 

equation 15, we get

(16)

FIGURE 233.6. Schematic circuit diagram at a node for the 

diffusion equation. Values of R and R' are 1 megohm or 

higher.

which satisfies the complete diffusion equation

dc d /^ dc\ dc 3-7=5 ( D ^ I u ^-'

di d?/\ dy/ dx

The foregoing analysis indicates that the resistance 

values of Rm+y2 and Rm-y2 are functions of D. Likewise, the values of Ri+y2 and Ri-y2 are functions of u. 

Thus, for each flow condition (at a certain Reynolds 

number) these resistors are to be set according to the 

predetermined values of D and u at a particular node.

dc Boundary conditions are such that ^-=0 at the water

surface and at the channel bottom. Initial condition, 

c0, may be simulated by applying a voltage source at 

one or a series of nodes. Hence, it may be a varying 

function of time as well as space. The corresponding 

concentrations of the diffused material may then be 

measured at any node as a continuous voltage output.
E172 GEOLOGICAL SURVEY RESEARCH 1962

234. MEASUREMENT OF SEDIMENT-LADEN FLOW BY MEANS OF A CIRCULAR ORIFICE *

By NOBUHIRO YOTSUKURA, Washington, D.C.

When a sand-water mixture undergoes a rapid acceleration, as through an orifice constriction, the sand and 

water behave differently because inertial forces predominate the motion. An approximate analysis of such 

flow through a circular orifice was worked out on the 

basis of the one-dimensional energy equation.

The equivalent-flow area required for the sediment 

particles may be defined by the equation

QS=KAUS , (1)

where Qs is the volumetric rate of flow of particles per 

unit time, A is the cross-sectional area of the pipe, Us is 

the mean particle velocity, and K is the coefficient to 

give the equivalent-flow area of the particles when 

applied to the total area A. Then, the rate of flow of 

water Qw is given by

where Uu is the velocity of the water. From the 

definition of mean volumetric concentration Cv, Qs and 

Qw are also given by the relations

and

Qs=CvQ, (3)

(4)

where Q is the rate of flow of the mixture.

Equations 1, 2, 3 and 4 are combined into an equation

l-K Cv = U 

K 'l-C, U

(5)

When there is no slippage of the particles relative to 

the water, Us is equal to Uw. Then equation 5 shows 

that K=CV . If Ug is not equal to Uw, however, K is 

not equal to Cv .

Consider a circular orifice constriction in a horizontal 

pipe as illustrated in figure 234 . 1 . As a first approximation, it is assumed that the energy loss between sections 

1 and 2 is negligible. Since the flow of the mixture is 

continuous, the energy per unit volume of mixture at 

sections 1 and 2 can be considered equal. Writing the 

energy equation per unit volume of mixture,

-\_, 

FIGURE 234.1. Sketch of flow through an orifice.

where p denotes the pressure and p the density. If it is 

assumed that 17^=17,1, then K±=CV and Ui can be 

used to replace Usi and Uwi, and equation 6 becomes

* .

'

(7)

To compute pi pz as a function of Ui, Cv, pw, and ps, it 

is necessary to eliminate U^, Us2 , and K2 ; thus three 

more equations containing these variables must be 

derived.

The first and second of these equations are obtained 

by applying continuity for the water flow and the 

sediment flow. Thus,

and

(8)

(9)

1 Based in part on dissertation for the degree of Doctor of Philosophy, Colorado State University, October 1961.

The third equation can be obtained by considering the 

acceleration effects. It is assumed that the accelerative 

force on the sediment mass is largely a result of pressure 

gradient. The drag as well as the virtual mass resist-
ARTICLE 234 El 73

ance are neglected. Applying the momentum principle 

between sections 1 and 2 separately on the water flow 

and the sediment flow,

and

Cvp1A1 K2p2A2=PsQs (Us2

(lOa)

(10b)

If each side of equation lOa is divided by the corresponding side of equation lOb, and a term (Cv K2 )A2p2 

is assumed to be negligibly small compared with both 

C^Aip^ A2p2 ) and (1 Cv )(Alpl A2p2 ), it can be 

shown that

 i ps

Us2 Ui

(11)

where 7 refers to specific weight.

Substituting equations 8, 9, and 11 into equation 7, 

it is seen that

where

o= > O__ IS

a=w n^l

S-l

o "' a-\-aCv ,

and

S-\ n

a= - auv .

i, (12)

(13)

(14)

(15)

(16)

This is the equation of flow through an orifice, and it 

takes into account the differential acceleration of 

sediment and water. It should be noticed that A2 is 

the area of the vena contracta rather than the orifice 

area; thus, strictly speaking, A2 is not constant.

The comparable equation assuming a homogeneous 

mixture is given by

(17)

where the mixture density p is given by 

p=(l (18)

Equations 12 and 17 are compared numerically below. 

For simplicity, equation 12 can be written as

U2

(12')

By comparing \}/pw and (a2 l)p, or 2 and > the vaCL 1 pit

lidity of the assumption of the homogeneous mixture 

can be examined. In preparing the following table, S 

was assumed to be equal to 2.65.

\

\ a

c, \

(percent) \

5. __ . _ . _

10..,. . .

15 

20 . . .

25 

30.... . .

*

02-1

1.062

1.114

1.180

1.240

1.300

1.288

1.67

p

pw

1.082

1.165

1.248

1.330

1.413

1.495

Ratio

1.019

1.046

1.058

1.073

1.087

1.161

*

02-1

1.061

1.122

1.250

1.314

1.388

2

P

Pw

1.082

1.165

1.248

1.330

1.413

1.495

Ratio

1.020

1.038

1.057

1.064

1.075

*

02-1

1.065

1.128

1.188

1.260

1.332

1.410

3

P

pw

1.082

1.165

1.248

1.330

1.413

1.495

Ratio

1.016

1.033

1.051

1.056

1.060

1.060

To obtain the mixture velocity corrected for the 

nonhomogeneous effect, the velocity determined by the 

assumption of homogeneous mixture is divided by the 

square root of the ratio in the table. It is seen that, 

except for a case of C,=30 percent and a=1.67, the 

correction is less than 4 percent.

Because of the assumption employed in the analysis, the result is not applicable if deposition of sediment takes place near the orifice. Even though the 

assumptions of negligible drag and virtual mass resistance seem reasonable, they certainly will have some 

effect if more elaborate calculations are made. Also 

a small energy loss due^to the impact of sediment on 

the upstream face of the orifice as well as the wall 

friction loss are to be expected. Since each of these 

small effects, except that of virtual mass resistance, 

will lead to a higher value of P\-P<i than that given 

by equation 12, the true correction is somewhat larger 

than that given by the approximate theory used in 

this article.
E174 GEOLOGICAL SURVEY RESEARCH 1962 

EXPERIMENTAL HYDROLOGY

235. USE OF A NEUTRON MOISTURE PROBE TO DETERMINE THE STORAGE COEFFICIENT OF AN

UNCONFINED AQUIFER

By WALTER R. METER, Garden City, Kans.

WorTc done in cooperation with the State Geological Survey of Kansas, the Kansas State Board of Health, and the Kansas

State Board of Agriculture

The coefficient of storage of an unconfined aquifer 

is approximately equal to the specific yield x of the 

material that is drained when water is pumped from 

the aquifer. Under field conditions the storage coefficient can be determined by several methods. One of 

these, the aquifer-test method, consists of measuring 

the rate of water-level decline (or recovery) in one 

or more observation wells near a well from which 

water is being pumped (or has been pumped) at a 

steady rate. Another method makes use of a neutron 

moisture probe to determine the difference between the 

moisture content of saturated material and the moisture content of the same material after it has been 

drained. These two distinctly different methods were 

used simultaneously to determine the storage coefficient of alluvium in the Arkansas River valley near 

Garden City, Kans.

In the vicinity of the test site the aquifer is very 

uniform both vertically and horizontally. Its thickness is 54 feet and, although it contains a few lenses 

of clay, the aquifer consists largely of medium to 

coarse sand and fine to coarse gravel. Well 24 34 

Iddb, which was pumped for the test, is 70 feet deep, 

and has a 16-inch casing perforated from 16 feet below land surface to the bottom of the well. The hole 

drilled for the well was 30 inches in diameter; the 

annular space between the casing and the wall of the 

hole was packed with gravel before the well was 

developed. Observation wells and moisture-probe access tubes were installed at distances of 10, 50, 100, 

200, and 450 feet from the pumped well. The observation wells, each about 40 feet deep, were cased with 

li/4-inch iron pipe tipped with a 3-foot sand point. 

The moisture-probe access tubes were cased with

i As applied to a rock or soil; the ratio of (1) the volume of water 

which, after being saturated, it will yield by gravity to (2) its own 

volume.

1.625-inch OD (1.555-inch ID) cold-drawn steel tubing and were plugged at the bottom to prevent entrance of water. The access tube nearest the pumped 

well was 30 feet deep; the others were 25 feet deep. 

The water-level measurements and moisture-content 

determinations presented here are for observation 

well 1 and access tube 2, both of which were 10 feet 

from the pumped well.

The moisture probe and sealer unit used in the test 

were loaned by the Kansas State Agricultural Experiment Station at Garden City. The probe contains 

a BF3 (boron trifluoride) slow-neutron detector tube, 

a 5-millicurie fast-neutron source (radium-266 and 

finely ground beryllium), and a transistor amplifier 

circuit. The number of moderately slow or slow 

neutrons detected per unit of time is a measure of the 

concentration of hydrogen atoms in the material being sampled. The hydrogen content of inorganic materials is largely contained in the molecules of free 

water; hence, the slow-neutron count is a measure of 

the moisture content of the material. The count is 

registered visibly on the sealer's five glow-tube decade 

counters. The moisture-calibration curves supplied by 

the manufacturer for specific types of access tubing 

are used to convert counts per minute to percent 

moisture by volume.

Before the aquifer test was begun, the percent moisture by volume was determined for each foot of saturated material at each access tube. When pumping 

began, the moisture probe was in access tube 2; its 

center was 1.5 feet below the water table, at a depth 

of 18.1 feet below the land surface. Although, during the test, most of the determinations of moisture 

content were made with the center of the moisture 

probe at that depth, some also were made with the 

center of the moisture probe at depths of 17.1 and 

19.1 feet below land surface. The determinations
ARTICLE 235 E175

made during the test, covering about 25 hours, are 

given in the table below.

Water-level measurements and moisture-content determinations 

during aquifer test near Garden City, Kans., November 21-28, 

1960

Water-level measurements and moisture-content determinations 

during aquifer test near Garden City, Kans., November 21 28, 

1960 Continued

Time since 

pumping started 

(minutes)

Depth to water 

below land surface 

(feet) .

Moisture content 

(percent by volume)

Amount of water 

drained (percent 

by volume)

Center of moisture probe 19.1 feet below land surface

pumping started

(minutes)

below land surface

(feet)

(percent by volume) drained (percent

by volume)

72

Center of moisture probe 17.1 feet below land surface 73

0

82

84 

94

480 

481 

1,484

16.65

21. 77

21. 78 

21.83

22.82 

22. 82 

23. 72

29.3

14. 2

13.7 

13.8

11.5 

11.3 

10. 3

y&

15. 1 158

15. 6

15' 5 491 17- 8 i S! 18-° '492

19. 0 L ' 4y2

21. 72

21. 72

21.84 

21. 87

22.09

22.09 

22.84

22.84 

23. 75 

23. 75

31.5

31.0

Jb. 4 

23.8

13. 1

12. 9 

10.9

10. 7 

9.6 

9.5

0

.5

5. 1

7.7

18.4

18. 6 

20. 6

20. 8 

21. 9 

22.0

Center of moisture probe 18.1 feet below land surface

0

2

4

6

8

10

15

20

25

30

43

45

47

48

51

59

60

62

64

66

68

70

80

90

92

109

180

182

184

484

485

1,489

16. 65

20. 72

20. 87

20. 96

21.04

21. 11

21. 20

21. 27

21.30

31. 36

21. 50

21. 51

21. 53

21. 54

21. 55

21.61

21. 62

21. 64

21. 66

21. 67

21. 69

21. 71

21. 76

21. 81

21. 82

21. 91

22. 18

22. 19

22. 19

22. 83

22. 83

23. 74

30. 5

30.5

30. 2

29. 7

31. 1

30. 5

30. 5

31. 1

30. 4

29. 9

29.6

29. 1

28.0

28.0

28. 1

25.5

24.2

23.4

23. 1

22.3

21. 1

20. 9

17.4

15.0

15. 2

14. 3

13. 9

13.8

13. 7

11.3

11.3

10. 3

0

0

. 3

.8

0

0

. 1

. 6

. 9

1. 4

2. 5

2. 5

2. 4

5.0

6. 3

7. 1

7.4

8. 2

9. 4

9. 6

13. 1

15. 5

15. 3

16. 2

16. 6

16. 7

16. 8

19. 2

19. 2

20. 2

Well 24-35-lddb was pumped at a rate of 915 gallons per minute for 9,080 minutes (6 days, 7 hours, 

and 20 minutes). Measurement of the depth to water 

in all the observation wells was made at 1-minute 

intervals for the first 10 minutes, at 2-minute intervals 

for the next 10 minutes, at 3-minute intervals for 9 

minutes, at 5-minute intervals for the next hour,, at 

30-miimte intervals for the next 2% hours, at 6-hour 

intervals for the next 24 hours, and at approximately 

12-hour intervals for the rest of the test. Personnel 

of the Division of Water Resources of the Kansas 

State Board of Agriculture assisted in making the 

water-level measurements. The .depth to water in observation well 1 at the times the determinations of 

percent moisture by volume were made with the moisture probe also are given in the table.

The Thiem equation (Wenzel, 1942, p. 79-82) and a 

modified form of the Theis nonequilibrium formula 

(Cooper and Jacob, 1946) were used in analyzing the 

drawdown data. As determined by the Thiem equation, the coefficient of transmissibility is 127,000 gallons per day per foot and the coefficient of storage 

is 0.19; as determined from the Theis nonequilibrium 

formula, the coefficient of transmissibility is 115,000 

gallons per day per foot and the coefficient of storage 

is 0.21.

Selected measurements of the depth to water and 

of the amounts of water drained from materials at 

depths of 17.1, 18.1, and 19.1 feet, as computed from

667422 O - 63 - 13
E176 GEOLOGICAL SURVEY RESEARCH 1962

the determinations of percent moisture by volume, 

are plotted in figure 235.1 in relation to time since 

pumping started.

Although the data for a depth of 18.1 feet appear 

to indicate that drainage began between 25 and 30 

minutes after pumping started, it cannot be stated 

with certainty that drainage began until about 43 

minutes after pumping started inasmuch as the error 

of accuracy of the determinations is about 1 percent 

moisture by volume. After 1,489 minutes of pumping, 

the percent moisture by volume was 10.3 and drainage 

was virtually complete. This value, therefore, approximates the specific retention of the material at 

that depth. Inasmuch as the percent moisture by 

volume was about 30.5 before drainage commenced, the 

specific yield of that material is about 20.2 percent 

by volume. The curve in figure 235.1 for a depth 

of 18.1 feet shows that once drainage began it proceeded rapidly and was about 80 percent complete 

within 65 minutes after drainage began.

=> 18

tn 22 """/""T"---.^M^ter; 

/ , ^^SL^,

vJk^\ I ^\ I I

10 100 1000 

TIME SINCE PUMPING STARTED, IN MINUTES

FIGURE 235.1. Depth to water in observation well 1 and 

amount of water drained at depths of 17.1, 18.1, and 19.1 feet, 

as determined by neutron moisture probe in access tube 2.

The curve in figure 235.1 for a depth of 19.1 feet 

below land surface shows that drainage began about 

75 minutes after pumping started. At this depth the 

moisture content of the fully saturated material was 

about 31.5 percent by volume and the specific retention 

after 1,492 minutes of pumping was 9.5 percent. If 

drainage at that time is regarded as complete, the 

specific yield of the material at that depth is about 

22 percent by volume.

As the first determination of the moisture content 

at 17.1 feet was not made until 82 minutes after the 

test began, the time that drainage commenced at that 

depth is not known. Inasmuch as the moisture content by volume before the test began was 29.3 and 

the specific retention was 10.3 percent by volume, the 

specific yield was about 19.0 percent by volume when 

drainage approached completion.

Because the storage coefficient of a water-table aquifer is approximately equal to the specific yield of the 

aquifer, the storage coefficient at the test site may be 

regarded as equal to the average of the specific-yield 

values for the depths of 17.1, 18.1, and 19.1 feet. 

That average is 0.205, which is between the values of 

0.19 and 0.21 determined by analysis of the drawdown data. The closeness of the value determined 

from measurements made with the neutron moisture 

probe to those determined by the aquifer-test method 

indicates that the two methods are equally reliable as 

a means of determining the storage coefficient of homogeneous isotropic unconfined aquifers.

REFERENCES

Cooper, H. H., Jr., and Jacob, G. E., 1946, A graphical method 

of evaluating formation constants and summarizing wellfield history: Am. Geophys. Union Trans., v. 27, no. 4, 

p. 526-534.

Wenzel, L. K., 1942, Methods for determining permeability of 

water-bearing materials, with special reference to discharging-well methods: U.S. Geol. Survey Water-Supply 

Paper 887, 192 p.
236.

ARTICLE 236 El77

CENTRIFUGE TECHNIQUE FOR DETERMINING TIME-DRAINAGE RELATIONS FOR A NATURAL SAND 

By K. C. PRILL and A. I. JOHNSON, Denver, Colo.

Work done in cooperation with the California Department of Water Resources

The solution of many ground-water problems is 

aided by a knowledge not only of specific yield, or the 

ultimate drainage of sediments, but also of the amount 

and distribution of moisture in the sediments with 

respect to time since drainage started. Centrifuging 

has been proposed as a promising method of estimating moisture distribution in sediments after gravity 

drainage is complete (Meinzer, 1942, p. 352-356; Marx, 

1956; Miller and Miller, 1955; Prill, 1961). As one 

phase of specific-yield research by the Geological Survey Hydrologic Laboratory, Denver, Colo., the centrifuge method of estimating moisture distribution 

in sediments after various periods of gravity drainage is undergoing evaluation. This method has been 

found to yield, for medium sand, time-drainage data 

that agree closely with those obtained by reading 

tensiometers inserted into a column of the sand that 

is allowed to drain naturally.

Fresno sand, a natural medium sand (fig. 236.1) 

obtained at a test site near Fresno, Calif., was packed 

in duplicate plastic columns about 163 cm (centimeters) long and 2.8 cm in diameter. Tensiometers

100

were located along the columns at 19 positions above 

the base. The sand was saturated and then allowed to 

drain by gravity. Readings of moisture tension within the columns were made at regular intervals. The 

data thus collected were converted to moisture content, in percent of dry weight, by using a calibration 

curve relating moisture content to moisture tension 

for Fresno sand. The moisture content of the sand at 

the different heights above the base of the columns 

then was plotted for 1.5, 5, 24, and 148 hours after 

drainage commenced (fig, 236.2). Because some of

160

5 10 15 20 

MOISTURE CONTENT, IN PERCENT OF DRY WEIGHT

PARTICLE-SIZE DIAMETER, IN MILLIMETERS

FIGURE 236.2. Moisture distribution in columns of Fresno sand 

FIGURE 236.1. Particle-size distribution of Fresno sand sample. at intervals after the start of drainage by gravity.
E178 GEOLOGICAL SURVEY RESEARCH 1962

the tensiometers were slow in responding during the 

early periods of drainage, the moisture values obtained from them are not plotted. The weighted average moisture content (based on the moisture content 

at the tensiometer locations) of 3 different segments 

of each column as well as for the whole column was 

determined for each of the 4 periods of drainage. The 

values are given in the table below.

Moisture content of Fresno sand after drainage by gravity for 

selected periods and after centrifuging at 36.2 g's for calculated 

equivalent time

163-cm column drained by gravity

Segment of 

column 

(centimeters 

above base)

100-163

70-100

0-70

0-163

Length of 

drainage 

period 

(hours)

1. 5

5

24

148

1. 5

5

24

148

1. 5

5

24

148

1. 5

5

24

148

Moisture ' 

content 

(percent of 

dry weight)

9. 5

7.4

5. 8

4.3

10. 4

8.6

6.9

4. 8

15. 1

13. 6

13.3

12. 5

12. 1

10. 4

9. 1

7. 9

4.5-cm column drained by centrifuging

Segment of 

column 

(centimeters 

above base)

2. 8-4. 5

1. 9-2. 8

0 -1. 9

0 -4. 5

Calculated 

drainage 

period 

(seconds)

4. 1

13. 7

66

407

4. 1

13. 7

66

407

4. 1

13. 7

66

407

4. 1

13. 7

66

407

Moisture 

content 

(percent of 

dry weight)

10.2

7.3

5. 8

4. 0

12. 5

8. 5

6.8

4.9

17. 6

13. 5

14.0

13. 8

13. 8

10.3

9.4

8.4

1 Computed from moisture-tension measurements.

Fresno sand then was packed into duplicate columns 

4.5 cm long for drainage by centrifuge. The centrifuge 

acceleration required to equate the small sample in 

the centrifuge to the column drained by gravity was 

obtained from the formula

a=T2 9> (i)

in which LI is the length of the drainage column to 

which the centrifuged sample is being equated, L2 is 

the length of the centrifuged sample, a is the centrifuge 

acceleration, and g is the acceleration due to gravity.

From appropriate substitutions in the formula it was 

determined that 36.2 g's would be required to equate 

the sample 4.5 cm long to the sample 163 cm long.

The relation between the times for equivalent 

drainage by the two methods is given by

(2)

in which ti is the time for gravity drainage, t2 is the time 

for drainage in the centrifuge, and a and g are as 

defined previously. According to equation 2, drainage 

periods of 4.1, 13.7, 66, and 407 seconds in the centrifuge 

under an induced acceleration of gravity of 36.2 g's 

would be equivalent to 1%, 5, 24, and 148 hours, respectively, of natural drainage in the column. The 

actual centrifuge drainage periods were maintained 

within ± 1 second of the calculated drainage periods.

For this study, 4 pairs of samples were centrifuged 

for lengths of time equivalent to column-drainage 

periods of 1.5, 5, 24, and 148 hours. After each 

sample was centrifuged it was divided into fractions 

that were proportionally equivalent to the three different segments of the drainage column, and the moisture content of each fraction was determined by oven 

drying. The values thus determined, together with 

the weighted average of the entire sample, also are 

given in the accompanying table.

The values for moisture content obtained by the 

two methods show a striking similarity, thus indicating the applicability of this centrifuge technique 

for predicting gravity-drainage relations of sandy materials. Present work shows that promising results 

also may be obtained by this centrifuge technique when 

applied to a wide range of materials.

REFERENCES

Marx, J. W., 1956, Determining gravity drainage characteristics on the centrifuge: Am. Inst. Mining Metall. Engineers 

Trans., v. 207, p. 88-91.

Meinzer, O. E., ed., 1942, Hydrology: New York, McGrawHill, Inc., 712 p.

Miller, E. E., and Miller, R. D., 1955, Theory of capillary 

flow pt. l, Practical implications: Soil Sci. Soc. America 

Proc., v. 19, no. 3, p. 267-271.

Prill, R. C., 1961, Comparison of drainage data obtained by 

the centrifuge and column-drainage methods: Art. 434 in 

U.S. Geol. Survey Prof. Paper 424-D, p. D399-D401.
ARTICLE 237 El79 

237. BEHAVIOR OF DETERGENTS (ABS), BACTERIA, AND DISSOLVED SOLIDS IN WATER-SATURATED SOILS

By H. G. PAGE, C. H. WAYMAN, and J. B. EOBERTSON, Denver, Colo.

Work done in cooperation with Federal Housing Administration

Many investigators (including Flynn and others, 

1958; Deluty, 1960; and Walton, 1960) have described 

water pollution by the anionic detergent alkylbenzenesulfonate (ABS). This most commonly used type of 

household detergent is characterized by its resistance 

to biochemical degradation. When septic-tank and 

sewage-plant effluents containing ABS and bacteria 

are infiltrated into unsaturated materials, both of these 

constituents and dissolved solids are potential pollutants of ground-water supplies. Robeck and others 

(1962) determined that the ABS content does not 

significantly influence the rate of growth of coliform 

bacteria (EscTierichia coli) or their movement through 

water-saturated sands. In our study, the comparative 

effectiveness of a coarse-grained, a fine-grained, and a 

colloid-coated soil in removing ABS, dissolved solids, 

and bacteria from sewage effluent was determined under saturated-flow conditions.

Three filtration test columns approximately 4 feet 

(column 1), 2 feet (column 2), and 1 foot (column 

3) long and of 3.75 inches inside diameter were constructed of Lucite. Column 1 was packed with "20-30 

mesh Ottawa sand," column 2 with a sandy loam, 

and column 3 with a colloidal alumina-coated 20-30 

mesh Ottawa sand. The material in the columns was 

packed with a mechanical vibrator (Morris and Kulp, 

1961). Flow rates ranged from 2.2 to 3.6 feet per 

day and were maintained at approximately 2 milliliters per minute by peristaltic pumps. The sewage 

effluent was introduced at the base of the columns and 

filtered upward to minimize the effect of air entrapment and channeling. The apparatus for one of the 

columns is shown in figure 237.1.

Selected results of daily analyses of the sewage 

effluent, before and after nitration through test column 

1, are given in figure 237.2. Coliform (Escherichia 

coli) and total bacteria were affected similarly by 

nitration, and their concentrations differed only in 

magnitude before and after filtration. Determination 

of selected constituents and properties of the sewage 

effluent (before filtration) indicated the following 

ranges of concentration: bacteria, 160,000 to 2,150,000 

per milliliter; detergent (ABS), 3.3 to 5.5 ppm (parts 

per million); dissolved solids, 440 to 1,000 ppm; dissolved oxygen, 0.0 to 5.0 ppm; biochemical oxygen 

demand, 125 to 225 ppm; pH, 7.4 to 8.1; and surface 

tension, 50 to 55 dynes per centimeter. ABS was

3t75'' 

(O

Tygon tubing*

Lucite - 

cylinder

-Tygon tubing

Tygon tubing

FIGURE 237.1. Sketch showing one column of the filtration 

test-column apparatus.

determined by the methylene-blue method, which is 

subject to a variety of limitations (Wayman, 1962). 

Determinations of coliform and total bacteria were 

made by the pour-plate technique, using MacConkey 

and Nutrient agars, respectively, at an incubation 

temperature of 37°C for 48 hours. Other determinations were made by standard laboratory methods.

The 20-30 mesh Ottawa sand in column 1 consisted 

principally of well-sorted rounded frosted quartz and 

a minor amount of mica. The parameters of this 

material are listed in the table (p. E180). Column 1 

(Ottawa sand) was operated continuously for 71 days. 

During the first 10 days of operation, bacteria traveled through the column virtually unaffected. During the next 61 days, the removal of bacteria by nitration ranged from 50 to 99.8 percent and averaged 

about 88 percent. Filtration, however, had little influence on the ABS and dissolved-solids contents (fig. 

237.2). After 71 days of operation, a 1-percent solution of a positively charged colloidal alumina was 

added to the sewage reservoir. The colloidal alumina
E180 GEOLOGICAL SURVEY RESEARCH 1962

20 30 ^ 70 

TIME, IN DAYS

FIGURE 237.2. Typical effects of soil filtering of sewage effluent.

immediately caused flocculation of about 70 percent of 

the negatively charged ABS and reduced the bacteria 

count in the reservoir to zero. No additional ABS 

was removed by percolating this alumina-containing 

solution through the soil column. Addition of a 0.1- 

percent alumina solution removed most of the bacteria 

but was only about half as effective as the 1-percent 

solution in removing ABS (see fig. 237.2). Although 

effective, use of an alumina solution to treat sewage 

effluent on a large scale would be economically prohibitive at the present cost of 16 cents per gallon.

Physical and chemical properties of column material 

[N.D., not determined, but probably insignificant]

Median diameter _ _ _/u

Coefficient of sorting

Surface area (calculated from sieve analPorosity, packed percent

Total soluble salts. _ ____________ppm__ 

Organic matter _________________ _ppm_ _ 

CaCOa ___ ppm 

P2O5 _ _---__ _ ppm _

N as NO3 ______________________ _ppm_ _ 

Available K_ ___________________ _ppm_ _

Column 1, 

20-30 mesh 

Ottawa sand

760

1.02

28

39

N.D. 

N.D. 

N.D. 

N.D.

N.D. 

N.D. 

N.D.

Column 2, 

sandy 

loam

50

>2. 35

534

53

1,400 

20, 000 

2,000 

280

12 

130 

7.3

The sandy loam in column 2 consisted mainly of 

quartz, illite, feldspar, and mica (see table for additional properties). Approximately 57 percent of this 

material had a diameter of less than 61j«,. Column 2 

was operated for 72 days. Chemical analyses of effluent sampled during initial operation of the column 

indicated that the ABS and dissolved-solids contents 

of the sewage had increased markedly in passing 

through the column. This increase presumably resulted from flushing of the high content of organic 

matter in the soil (see table) and could not be attributed entirely to the ABS content. (The analytical 

procedure used was not specific for ABS.) After the 

first day of operation, the amount of bacteria removed 

by the sandy loam ranged from 60 to 96 percent and 

averaged 93 percent. As in the tests of column 1, 

little ABS or dissolved solids was removed by filtration even after 72 days of operation.

The filter medium in column 3 consisted of 20-30 

mesh Ottawa sand coated with colloidal alumina and 

had 34 percent porosity. The column was operated 

for only 3 days. Initial operation removed 99 percent of the bacteria from the sewage and 30 percent 

of the ABS. After only V/2 days of operation, the 

capacity of the column to remove bacteria and ABS 

decreased markedly, and the test was terminated 

quickly.

Bacterial-slime growth, which resulted in pore clogging, was indicated in all three columns by darkening 

of the filter media, presumably largely by iron sulfide (FeS) (Winneberger and others, 1960, p. 32). 

In coarse sand, pore clogging increases uniformly in 

the direction of water flow. In fine-grained soil, pore 

clogging apparently develops in an erratic pattern 

along the more permeable channels. After 71 days 

of operation, pore clogging was minor in coarse sand, 

but after only 15 days it was serious in fine-grained 

materials. In the fine-grained soil of column 2, clogging was so severe that the soil column began to break 

up after 68 days of operation.

In summary, coarse sand and sandy loam removed 

about 90 percent of the bacteria from sewage within 

a few feet of travel, but additional travel does not 

necessarily remove all the remaining bacteria. Dissolved solids and detergent content (ABS) virtually 

were unaffected by filtration of saturated flow. Because bacterial clogging occurs quickly in fine-grained 

soils, coarser sand is preferred as a pollution filter 

for the removal of bacteria. The use of colloidal 

alumina to remove bacteria or ABS seems to be economically unfeasible at the present time.
ARTICLE 238 E181

REFERENCES

Deluty, Jerome, 1960, Synthetic detergents in well water:

Public Health Repts., v. 75, p. 75-77. 

Flynn, J. M., Andreoli, A., and Guerrera, A. A., 1958, Study

of synthetic detergents in ground water: Am. Water

Works Assoc. Jour., v. 50, no. 12, p. 1551-1562. 

Morris, D. A., and Kulp, W. K., 1961, Mechanical uniform

packing of porous media: Art. 305 in U.S. Geol. Survey

Prof. Paper 424-D, p. D31-D32. 

Robeck, G. G., Bryant, A. R., and Woodward, R. L., 1962,

Influence of ABS on the movement of coliform through

water-saturated sandy soils: Am. Water Works Assoc.

Jour., v. 54, no. 1, p. 75-82.

Walton, Graham, 1960, Effects of pollutants in water supplies 

 ABS contamination: Am. Water Works Assoc. Jour., 

v. 52, no. 11, p. 1354-1362.

Wayman, C. H., 1962, Limitations of the methylene blue 

method for ABS determinations: Art. 49 in U.S. Geol. 

Survey Prof. Paper 450-B, p. B117-B120.

Winneberger, J. H., Francis, L., Klein, S. A., and McGauhey, 

P. H., 1960, Biological aspects of failure of septic-tank 

percolation systems final report: Univ. California, Berkeley, College of Eng. and School of Public Health, Sanitary Eng. Research Lab., 125 p.

238. ADSORPTION OF THE SURFACTANT ABS35 ON KAOLINITE 

By C. H. WAYMAN, J. B. EOBERTSON, and H. G. PAGE, Denver, Colo.

Work done in cooperation with Federal Housing Administration

Nondegradable surfactants injected into the ground 

in household waste water may pollute ground-water 

supplies. One of the natural means of removing these 

surfactants from waste water is by their adsorption 

on soil minerals. Because the clay mineral kaolinite 

is widely distributed, its ability to adsorb alkylbenzenesulfonate (ABS), an essential nondegradable surfactant in detergents, was studied.

A radiotracer technique employing ABS tagged 

with S35 was used in the study. The effect of such 

variables as alkyl chain length, pH, concentration of 

ABS, the addition of phosphate and carboxymethylcellulose, and salt content on adsorption of ABS by 

kaolinite was observed.

A sample of kaolinite, presized at 100 percent less 

than 5/x diameter and 96 percent less than 2/x, was 

obtained from Dry Branch, Ga. Additional sizing 

reduced the diameter of the kaolinite to the range from 

0.2/x to 2.0/x and resulted in a product having a surface 

area of 20.5 m2 per g (square meters per gram). Surface 

area was determined by a method described by Emmett 

(1942, p. 3). A typical chemical analysis of kaolinite 

(in percent) is: SiO2 , 45; A12O3 , 38; TiO2 , 0.5 to 2.0; 

loss on ignition, 14; and CaO, Na2O, K2O, and MgO, 

trace amounts. The two forms of S35-tagged ABS, 

dodecyl ABS (C12 ) and pentadecyl ABS (C15 ), differed 

only in the number of carbon atoms in their alkyl

chains; the typical ABS ion can be represented in 

solution by

where R represents the alkyl chain. The specific 

activity of either form was 1 millicurie per millimole.

Experiments were designed to study the adsorption 

of ABS on kaolinite. Sealed sausage casings containing a suspension of 200 mg of kaolinite in 10 ml 

of deionized water were placed in several small polyethylene bottles. A series of solutions of varying concentrations of ABS tagged with S35 were prepared. 

Additional anions and cations were added to some of 

these solutions. Then a 10-ml volume of a different 

ABS solution tagged with 'S35 was added to each of 

the small polyethylene bottles -containing the sealed 

sausage casing. The bottles were capped and were 

placed in a constant-temperature (25°C ± 1°C) agitated water bath. The sausage casing functioned as a 

molecular sieve, being permeable to ABS ions but 

imperable to clay particles. In time, various amounts 

of ABS equilibrated with the kaolinite surface.

Periodically, duplicate 1-ml samples were removed 

from the solutions surrounding the sausage casing and 

were evaporated to dryness on planchets. The dried 

samples were counted in a conventional beta counter 

utilizing a shielded end-window Geiger counter and 

sealer. Then the counting rates of the test samples
E182 GEOLOGICAL SURVEY RESEARCH 1962

were compared to standards containing known amounts 

of S35-tagged ABS. The total number of counts for 

both the standards and the test samples was corrected 

for background, and the amount of ABS adsorbed per 

gram of kaolinite was calculated by the following 

equation:

A (Cs -Ct)VKs

cswk

(1)

where

^lfc =micrograms of ABS adsorbed per gram of

kaolinite;

Cs =the counting rate (counts per second) of 1 ml 

of evaporated ABS35 solution of known 

starting concentration (Ks), corrected for 

background;

Ct=the counting rate (counts per second) of 1 ml 

of evaporated test-sample solution, corrected for background;

KS =ABS35 concentration (parts per million) at the 

beginning of an experiment ;

Wk=weight of kaolinite, in grams; and 

V volume of water used.

Figure 238.1 shows that ABS is more readily adsorbed at low pH than at high pH, and that the 

amount of pentadecyl ABS adsorbed is significantly 

greater than the amount of dodecyl ABS. The kaolinite surface probably contains more positively 

charged sites at low pH than at high pH; consequently, anions like ABS are more readily adsorbed 

at low pH. Other investigators have found that anionic adsorption on clay minerals is significantly dependent on pH (Schofield, 1949; Schofield and Sam300

-200

12

TIME, IN DAYS

son, 1953; Street, 1956; and Holtzman, 1962). At pH 

values greater than 4, ABS is concentrated in the solution for short equilibration periods. This effect, 

indicated by negative values for adsorption of ABS, 

can be attributed to adsorption of water by kaolinite, 

thereby effectively increasing the concentration of 

ABS beyond that of the original solution. A pseudoequilibrium adsorption generally was attained in about 

1 or 2 days, but equilibrium adsorption was not attained, even after 12 days.

Figure 238.2 shows that the adsorption capacity of 

kaolinite for ABS increases as the concentration of 

ABS increases. Adsorption is greatest at low pH.

The amount of ABS adsorbed by kaolinite seems to 

increase in the presence of phosphate and carboxymethylcellulose (fig. 238.3). In a previous study, 

Header and Fries (1952) found that phosphate ion 

in ABS solution was not adsorbed on cotton, a highly 

adsorptive substance for ABS.

The table on page E183 shows the influence of different types of salts on adsorption of ABS by kaolinite. 

Results are presented for solutions of dodecyl ABS 

(4 parts per million) containing approximately 0.01 

molar concentrations of aluminum, calcium, or sodium 

chlorides. Monovalent and divalent chloride salts 

seem to increase ABS adsorption equally well for the 

pH range from 4 to 10. Trivalent salts (for example, A1C13 ) cause greater ABS adsorption, especially 

in the pH range from 4 to 7. This greater adsorption 

might be explained in part by the combined effects of 

critical micelle concentration (CMC) and the SchulzeHardy rule. The CMC is the concentration of surfactant (ABS) necessary to promote formation of

300

200

m 

o: uj O iw ^ 

m ? 

< -i 100

o

g -100

o

-200

10 12

TIME. IN DAYS

FIGURE 238.1. Influence of pH, length of alkyl chain, and time FIGURE 238.2. Influence of.pH, ABS concentration, and time 

on adsorption of ABS by kaolinite. on adsorption of ABS by kaolinite.
ARTICLE 238 E183

Effect of various salts on ABS adsorption on 

kaolinite

Percentage increase in A RS 

adsorbed

NaCl, 584 ppm

10.

CaCl2 , 952 pprn

A1C13, 1,333 ppm

10.

40

30

50

35

50

60

175

200

75

1 Adjusted by adding hydrochloric acid or 

sodium hydroxide.

ionic aggregates (micelles) from single ions. The 

micelles have a much higher degree of surface activity 

than do single ions. The CMC can be reduced if salts 

are in solution (Corrin and Harkins, 1947). For the

200

m

ce uj 

OH

m ?

<-i 100

ce w 

o

S ->

o

-200

EXPLANATION 

10 ppm dodecyl ABS

10 ppm dodecyl ABS + 

10 ppm phosphate

 - 10 ppm dodecyl ABS + 10 ppm 

phosphate+ 1 ppm carboxymethylcellulose

O pH=4 

A pH = 7 - 

El pH = 10

6 8 

TIME. IN DAYS

10 12

FIGURE 238.3. Influence of phosphate, carboxymethylcellulose, 

pH, and time on ABS adsorption by kaolinite.

Schulze-Hardy rule, the higher the valency, the less 

of the cation required to reduce the CMC (Lange, 

1951; Powney and Addison, 1937). If the surface 

activity is increased by the formation of detergent 

micelles, then solutions containing high concentrations 

of salt and highly charged ions (for example, A1+3 ) 

might enhance ABS adsorption on clay minerals.

This study indicates that ABS adsorption on kaolinite is influenced markedly by the length of the 

alkyl chain of the surfactant (ABS), concentration 

of ABS, pH of the solution, and the type and amount 

of salt in solution.

Studies similar to the above are in progress for the 

clay minerals illite and montmorillonite and' for mixed 

clay minerals.

REFERENCES

Corrin, M. L., and Harkins, W. D., 1947, The effect of salt on 

critical concentration for the formation of micelles In 

colloidal electrolytes: Am. Chem. Soc. Jour., v. 69, p. 683- 

688.

Emmett, P. H., 1942, The measurement of the surface areas 

of finely divided or porous solids by low temperature adsorption isotherms: Advances in Colloid Sci., v. 1, p. 1-38.

Holtzman, W., 1962, The application of the Verwey and Overbeek theory to the stability of kaolinite-water systems: 

Colloid Sci. Jour., v. 17, p. 363-382.

Lange, H., 1951, Tiber dem Einfluss gewohnlicher Electrolyte 

auf die Mizellenbildung in Kolloidelectrolytelosung: Kolloid-Zeitschrift, v. 121, p. 66-71.

Meader, A. L., and Pries, B. A., 1952, Adsorption In the detergent process: Indus, and Eng. Chemistry, v. 44, p. 1636- 

1648.

Powney, J., and Addison, C. C., 1937, The properties of detergent solutions. II. The surface and interfacial tensions 

of aqueous solutions of alkali sodium sulfates: Faraday 

Soc. Trans., v. 33, p. 1243.

Schofield, R. K., 1949, Effect of pH on electric charges carried by clay particles: Soil Sci. Jour., v. 1, p. 1-8.

Schofield, R. K., and Samson, H. R., 1953, The deflocculation 

of kaolinite suspensions and accompanying change-over 

from positive to negative chloride adsorption: Clay Minerals Bull., v. 2, p. 45-50.

Street, N., 1956, The rheology of kaolinite suspensions: Australian Chemistry Jour., v. 9, p. 467.
E184 GEOLOGICAL SURVEY RESEARCH 1962

239. AN APPLICATION OF THE GIBBS ADSORPTION EQUATION TO DETERGENT SOLUTIONS

By.COOPER H. WAYMAN, Denver, Colo.

Work done in cooperation with Federal Housing Administration

Surfactants (surface-active agents) or detergents 

cause changes in the surface properties of solutions. 

Changes in surface properties of detergent solutions 

at the air-solution interface can be evaluated by measurement of the solutions' surface tension. The importance of surface tension to water-pollution studies 

has been indicated in adsorption studies by Wayman 

and Eobertson (1962).

In adsorption studies of detergents, the area occupied by a detergent molecule at the air-solution interface is essential to the calculation of amounts of detergents required to form a monomolecular layer on 

solid mineral surfaces (Wayman, 1962). The discussion that follows describes how this area can be determined from an application of surface-tension measurements to the Gibbs adsorption equation.

The original Gibbs adsorption equation showing 

the behavior of surfactants at the air-solution interface with respect to concentration was derived for 

ideal, dilute solutions (Gibbs, 1928, p. 219). Later 

workers who have used surface-tension measurements 

as a mean of studying the behavior of surfactants at 

air-solution interfaces have disagreed as to the validity 

of the equation. For nonionic surfactants, the Gibbs 

adsorption equation seems to be satisfactory (Hutchinson, 1948). For ionizable compounds, some investigators found that the surface excess (the amount of 

surfactant at the air-solution interface compared to 

the amount in the bulk solution) calculated from surface-tension measurements amounts to only half the 

value measured directly (Brady, 1949; Salley and 

others, 1950). Pethica (1954) showed that the equation is valid only in an excess of a neutral salt. Davies 

(1952) achieved satisfactory results in calculating the 

surface excess of cetyltrimethylammonium ions at the 

oil-water interface, and Matuura and others (1958) 

found that the surface excess of alkyl sulfates could 

be calculated reasonably well for concentrations below 

the critical micelle concentration (that concentration 

at which single ions begin to form into a cluster of 

ions). To the writer's knowledge, the Gibbs adsorption equation has not been previously applied in studies of solutions containing the surfactant alkylbenzenesulfonate (ABS).

When surfactants are added to water, their ions or 

molecules tend to concentrate at the surface in preference to the bulk of solution. The surfactant solute can 

be represented in solution as a hydrophilic (waterattracting) group and a hydrophobic (water-repelling) 

group. For the widely used anionic surfactant alkylbenzenesulfonate, the two groups are located as

[M C6H4 

Hydrophobic group Hydrophilic group

where M is the length of the alkyl chain.

For a dilute solution of sodium alkylbenzenesulfonate 

(Na+1X-1), the Gibbs adsorption equation can be expressed as

a« R TTx-idlnCx-i

where

(1)

dy=change in surface tension, 

R=gas constant, 

T= absolute temperature, 

rNa+i=adsorption of sodium ion at the surface, 

IVi=adsorption of the anionic detergent ion at

the surface,

rH+i=adsorption of the hydrogen ion at the surface, and CNan, Cx-i, CH«=various ionic 

concentrations.

If the pH is kept constant and sodium ion is not 

adsorbed at the surface, then rNa«=0 and FH+i^r^-i 

= rHjr with CNa,+i = Cx-i=CN&x, and equation 1 reduces 

to

J- "FT !V /^~"A/xrt T~»rrT " 7 T _ /~y " \ /

Hence, the surface excess, THST (adsorption at the surface), can be calculated from equation 2 if the surface 

tension is measured at various ionic concentrations. 

Figure 239.1 shows the change in surface tension 

at various concentrations of detergent at five different 

temperatures ranging from 2°C to 35°C. At tempera-
ARTICLE 239 E185

o.oooi

ABS CONCENTRATION, IN MOLES PER LITER

0.001 _______0.01 1 I I I I

70

60

 I I 

EXPLANATION

n 2°c

A io°c

O 20"C

V 27°C

O 35°C

50

30

20. I I I

10 100 

ABS CONCENTRATION, IN PARTS PER MILLION

1000 4000

FIGURE 239.1. Relation of surface tension to concentration of ABS and to temperature.

tures of more than 2°C, small changes in concentration produce large changes in surface tension for ABS 

concentrations of as much as the critical micelle concentration (approximately 1,000 ppm). At this concentration, the detergent has less tendency to accumulate in the surface than at lower concentrations. 

After computing the changes in surface tension from 

the slopes of figure 239.1, the surface excesses were 

calculated and then were plotted on figure 239.2. The 

data indicate that detergent anions tend to accumulate 

at the surface and to saturate the air-solution interface at concentrations of about 1,000 ppm, the critical 

micelle concentration; at greater concentrations, desorption seems to occur. Maximum concentration of 

ABS at the air-solution interface is indicated at the 

lowest temperature (2°C).

The data from figure 239.2 can be used to calculate the 

area occupied per detergent molecule at the air-solution 

interface as follows:

S= 1 (3) TN

where

S area occupied per molecule; 

7Vr=Avogadro's number=6.02 X1023 molecules

per mole; 

r=surf ace excess, in moles per cm2 .

If the area occupied per molecule is known, monolayer 

or multilayer adsorption on solid mineral surfaces can 

be more readily determined. Equation 3 was used to 

calculate the area occupied (S) per molecule of a detergent solution containing 50 ppm of sodium alkylbenzenesulfonate at 20°C. Thus,

S=

3.35X10- 10 moles per cm2 X10- 18 cm2 per A2 X6.02X1023 molecules permole 

=49 A2 per molecule,
E186 GEOLOGICAL SURVEY RESEARCH 1962

0.001 0.002

ABS CONCENTRATION, IN MOLES PER LITER 

0.003 0.004 0.005 0.006 0.007 0.008 0.009

-1

-O EXPLANATION 

D 2°C 

A 10°C 

o 20°C 

V 27°C 

O 35°C

1000 2000

ABS CONCENTRATION, IN PARTS PER MILLION

FIGURE 239.2. Relation of surface excess to concentration of ABS and to temperature.

3000

where A2 (1 square angstrom unit) is equivalent to 

10~16 cm2 . For similar detergent solutions, the location 

of the hydrophilic group on the hydrocarbon chain has 

a pronounced influence on S (Meader and Griddle, 

1953). These investigators found the following values 

for S in dilute solutions of detergents: sodium 1-dodecylbenzenesulfonate ($=38 A2 per molecule), sodium 

2-dodecylbenzenesulfonate (49.5), sodium 6-dodecylbenzenesulfonate (120), and hygroscopic dodecylbenzenesulfonate (56). These data suggest that the 

detergent used in our work ($=49 A2 per molecule) 

consisted mainly of sodium 2-dodecylbenzenesulfonate.

REFERENCES

Brady, A. P., 1949, Surface tension of aqueous solutions of 

two foam-fractionated detergents: Jour. Phys. and Colloid Chem., v. 53, p. 56-66.

Davies, J. T., 1952, The application of the Gibbs equation to 

charged monolayers, and their desorption from the oilwater interface: Faraday Soc. Trans., v. 48, p. 1052-1061.

Gibbs, J. W., 1928, Collected works of J. W. Gibbs, v. 1: 

London, Longmans, p. 219.

Hutchinson, E., 1948, Mixed monolayers. I. Adsorbed films 

at air-water interfaces: Jour. Colloid Sci., v. 3, p. 413-424.

Matuura, R., Kimizuka, H., Miyamoto, S., and Shimozawa, E., 

1958, The study of the adsorption of detergents at a solution-air interface by radiotracer method. I. Adsorption 

isotherm for the solution of sodium alkyl sulfates: Chem. 

Soc. Japan Bull., v. 31, p. 532-538.

Meader, A. L., and Griddle, D. W., 1953, Force-area curves of 

surface films of soluble surface-active agents: Jour. Colloid Sci., v. 8, p. 170-178.

Pethica, B. A., 1954, The adsorption of surface-active electrolytes at the air-water interface: Faraday Soc. Trans., 

v. 50, p. 413-421.

Salley, D. J., Weith, A. J., Argyle, A. A., and Dixon, J. K., 

1950, Measurement of the adsorption of surface-active 

agents at a solution-air interface by a radiotracer method: 

Royal Soc. London Proc., series A203, p. 42-55.

Wayman, C. H., 1962, Adsorption of anionic detergent on solid 

mineral surfaces: Art. 117 in U.S. Geol. Survey Prof. 

Paper 450-C, p. C137-C139.

Wayman, C. H., Robertson, J. B., and Page, H. G., 1962, 

Surface tension of detergent solutions: Art. 179 in U.S. 

Geol. Survey Prof. Paper 450-D, p. D190-D192.
SUBJECT INDEX

[For major topic headings such as "Economic geology," "Engineering geology," "Stratigraphy," see under State names or refer to table of contents]

Article

ABS. See Surfactants.

Aeromagnetic surveys, Nevada, Egan Range__ 180

Utah, Oquirrh Mountains..__ 187 

Airborne control system, description 218 

Alabama, stratigraphy, coal fields..__-__ 193 

Alaska, quartz diorite line______ 203 

Alluvium, storage coefficient_.___ ___- 235 

Alteration, hydrothermal, hot springs.. ___ 200

propylitic, hot springs and ore deposits_ 201 

Amsden Formation, Wyoming, geologic

history._______________ 195 

Antarctica, geochronology, Thiel Mountains.. 208 

Aquifers, glacial deposits. - _________ 224,227

unconflned, determination of storage

coefficient__-..._________ 235 

Arid land, hydrology.______________ 229 

Arizona, stratigraphy, northern_____.__ 197

B

Bacteria, removal from sewage by soils.... _ 237

Bentonite beds, composition__________ 198

Bioherms, permeability...___________ 191

C

Cadmium, ion-exchange separation and spectrophotometric determination.... 214 

California, engineering geology, Fresno

County_______________ 210 

sedimentary petrology, Fresno County... 210 

Canada, quartz diorite line. __________ 203 

Carbonate rocks, Pa25i/U-Th"0/U dating

method________________ 209 

Carboniferous. See Mississippian, Pennsylvanian. 

Carmel Formation, Arizona and Utah, relation

to Navajo Formation_______ 197

Utah, bentonite beds___________ 198

Centrifuges, in measurement of soil moisture_ 236

Cerium, in syenite.._______________ 204

Chile, geomorphology, San Pedro de Atacama. 207

stratigraphy, Aconcagua Province.... _ _ 196

structural geology, Salar de Atacama-.._. 186

Clay minerals, association with hot springs_ 200

Collapse structures, ground water in_.____ 225

Colorado, economic geology, Wet Mountains. 181

stratigraphy, Middle Park.....__ ___ 189

Compaction, silty sediments.._ ___ _ __ 210

Copper, effect on precipitation of tellurium. _ 215

Crab Orchard Mountains Formation, Georgia,

definition of new members. _______ 194

Craters, volcanic.._______________ 211 

Cretaceous, Colorado, stratigraphy.____ 189 

Cretaceous-Jurassic boundary, Chile.-..___ 196 

Cryptovolcanoes, Missouri, Crooked Creek... 183

D

Detergents, removal from sewage by soils_____ 237

See also Surfactants. 

Dissolved solids, removal from sewage by

soils._________.__.____. 237

surface water, source. ___________ 230 

Drainage, comparison on topographic maps 212 

Drainage basins, characteristics________ 229

E

Article

Epidote, hydrothermal_____________ 201 

Evapotranspiration, effect of streamflow___ 221

Faults and faulting, Chile, Salar de Atacama. 186 

Floods, frequency based on extremes of shortterm records____________ 228 

frequency in the Southwest________ 229 

Flowerpot Shale, southwestern Oklahoma,

halite deposits..__________ 190 

Fluorine, spectrophotometric determination

with thoron.. ___________ 216 

Fort Payne Chert, Tennessee, bioherms___ 191 

Fresh water, relation to salt water_____ 219,226 

Fluorine, spectrophotometric determination

with thoron..___________ 216

Q

Galena, lead isotope studies. _________ 199 

Georgia, engineering hydrology, Georgia Nuclear Laboratory__________ 220 

stratigraphy, northwest- _________ 194 

Gibbs adsorption equation, application to detergent solutions__________ 239 

Glacial till, permeability-..__________ 224 

Gravity surveys, Utah, Oquirrh Mountains 187 

Great Salt Lake, salt content. _________ 230 

Ground water, coefficient of storage in unconfined aquifers____________ 235 

inflow to streams._____________ 223

H

Halite deposits, relation to salt springs____ 190

Hot springs, clay minerals___________ 200

epidote and zeolites___________ 200,201

Ice cover, relation to ground-water inflow to

streams.._______________ 223 

Iceland, geochemistry, Hveragerdi and Reykjavik-____________-. 200, 201 

Idaho, geomorphology, Menan Buttes____ 211 

metamorphism and petrology, Big Creek

quadrangle.____________ 204

Riggins quadrangle_______.__ 205

Idaho batholith, associated older rocks____ 204

trondhjemite...___________.___ 205

Illinois, ground water, southern. ______ 224

Ion-exchange separations, cadmium_ ___ 214

Ionium dating, carbonate rocks________ 209

Isotopes, lead, growth-zoned crystal..--___ 199

Jurassic, Arizona and Utah, stratigraphy. __ 197

Utah, bentonite beds. ___________ 198

Jurassic-Cretaceous boundary, Chile____. 196

Kaolinite, adsorption of surfactants on___ 238

Kentucky, stratigraphy, northeastern____ 192

structural geology, Middlesboro Basin__ 184

Note. Numbers refer to articles

Article

Landslides, causes of..... 182 

Lead, isotopes, variation with growth zoning. 199 

Lead-alpha age determinations, zircon 208 

Lee Formation, Kentucky, northeastern 192

M

M applng, airborne control system. 218

portable surveying tower 217

Metamorphism, zeolitic fades... -- 201

Meteoritic impact structures, Kentucky 184

Missouri-.. 183

Michigan, engineering hydrology, Upper

Peninsula,.-.. 223 

Middle Park Formation, Windy Gap Volcanic

Member, Colorado... 189 

Mississippian, Alabama, correlations 193 

Tennessee, bioherms---- 191 

Utah, correlations... 188 

Mississippian-Pennsylvanian boundary, Kentucky... .._.... . - 192 

Missouri, structural geology, Crooked Creek

area. 183

Montana, engineering geology, Gardlner 182

geomorphology, Gardlner 182

N

Navajo Formation, Arizona and Utah, relation

to Carmel Formation 197 

Neutron moisture probe, in determination of

aquifer storage coefficient 235 

Nevada, economic geology, Egan Range 180

geochemistry, Egan Range 180

geophysics, Egan Range 180 

New Mexico, flood frequency 229

glacial geology, Sangre de Cristo Mountains and Sierra Blanca... 213

ground water, Eddy County.. 225 

New York, ground water, Centre Island 226

ground water, Kings and Queens Counties. 219 

Niobium, Colorado, Wet Mountains 181

Idaho, central.......... 204

O

Ohio, ground water, Montgomery County 224 

ground water, northeastern 224 

Oklahoma, engineering hydrology, Pond Creek

basin. 221

stratigraphy, Beckham and Harmon

Counties..___ 190 

Olivine, X-ray determinative curve 206 

Open-channel flow, turbulent-diffusion equation-.. - 233

Oquirrh Formation, Utah, correlations 188 

Ore deposits, propylitic alteration.... 201

Orifices, circular, measurement of sedimentladen flow.. 234

Pa23i/U-Th23l)/U dating method, Pleistocene

carbonate rocks__ ._ 209 

Pacific province, thorium and uranium 202 

Parkwood Formation, Alabama, correlation.. 193

E187
E188 INDEX

Article

Pathfinder uplift, Wyoming, definition and

history.______________ 195 

Pennington Formation, Kentucky, northeastern.. 192

Pennsylvanlan, Georgia, nomenclature..__ 194

Utah, Oquirrh Formation_________ 188

Wyoming, Pathfinder uplift_______ 195 

Perldotites, olivlne from alpine and stratiform

types_____--.--._........ 206

Permian, southwestern Oklahoma, halite__ 190

Utah, Oquirrh Formation_________ 188 

Piedmont province, drainage_________ 212 

Pleistocene, New Mexico, correlations..------ 213

Montana, landslides________.___ 182 

Pleistocene continental rocks, dating,.__... 209 

Pollutants, in water-saturated soil______ 237 

Porosity, relation to overburden, particle size,

and sorting...._________. 210 

Portable surveying tower, description..__ 217 

Pottsville Formation, Alabama, correlation_ 193 

Proactlnium dating, carbonate rocks__. __ 209

Q

Quartz dlorite line, northwestern North

America_______________ 203 

Quaternary, Idaho, Snake River Plain... __ 211 

See also Pleistocene.

R

Ramey Ridge Complex, Idaho, definition-... 204 

Rare earths, Colorado, Wet Mountains_.... 181

See oZso specific name.

Recession curves, streamflow._________ 232 

Reservoirs, determination of storage-outflow

characteristics__________. 232

Salt, In surface water______________ 230

Salt cups, formation of.__---------__... 207

Salt springs, source of salt..__________ 190

Article

Salt-water encroachment, New York, Long

Island . 219

Saprollte, water infiltration rates.. 220 

Sediments, silty, compaction. __.. ._ 210 

Shatter cones, occurrence and mode of formation......... ... .. 183

Short-term records, as basis for long-term flood

frequency.. 228 

Soils, measurement of moisture content 236 

removal of detergents, bacteria, and dissolved solids....__ ......- 237

South Dakota, ground water, east-central . 224

Specific capacity, equations for estimating 222

Spectrophotometric determinations, cadmium 214

fluorine ....____ . . 216

Stalagmites, dating..._....__._ .. 209

Storage coefficient, determination for unconflned aquifers_ 235

Storage-outflow characteristics, reservoirs ... 232

Streamflow, effect of evapotranspiration ... 221

measurement of sediment-laden flow._.. 234

relation of accretion to Ice cover . 223

Surfactants, adsorption on kaollnlte.. 238

effect on surface properties of solutions... 239

Surveying, airborne control system 218

portable towers___._-_ 217

Syenite, cesium and niobium. 204

T

Tellurium, precipitation with hypophosphorous acid......_ 215

Tennessee, petrography, Caney Fork River

area._ 191 

stratigraphy, Caney Fork River area. 191

Tertiary, Colorado, stratigraphy 189

Texas, flood frequency . 229

Thorium, clrcum-Paclflc province. 202 

Colorado, Wet Mountains 181

Thoron, In Spectrophotometric determination

of fluorine...-. -.-._------_ 216

Topographic maps, comparison of drainage on. 212

Article

Transmisslbillty, calculating distribution of 235 

Tri-State mining district, lead Isotope studies. 191 

Trondhjemite, occurrence in Idaho batholith.. 209 

Tuff, glassy basalt. . . _____ 211 

Turbulent flow, open-channel... ______ 233

U

Uranium, circum-Paciflc province.______ 202

Utah, geophysics, Oquirrh Mountains.____ 187

mineralogy, southwest___________ 198

quality of water, Great Salt Lake area_ 230

stratigraphy, Oquirrh Mountains____ 188

southwestern__ ___.__ 197

structural geology, Oquirrh Mountains_ 187

V

Virginia, structural geology, Valley and Ridge

province______..._______ 185 

Volcanic rocks, thorium and uranium content. 202 

Volcanlsm, cones of glassy basalt tuft_____ 211

W

Washington, ground water, Tacoma area...- 227

Water infiltration, rate in weathered crystalline rocks_____ ______ 220 

Water stage, analysis.__------ ______ 231

Weathering, effect on water infiltration rates 220 

Well yield, formula for estimating_______ 222 

Wyoming, stratigraphy, southern.. ______ 195

X

X-ray determinations, olivine from peridotltes .... ... .....___ 206

Zeolites, hydrothermal... __________ 200,201 

Zircon, lead-alpha ages______________ 208

AUTHOR INDEX

Antal, P. SArticle

.. 209

Benson, M. A__________________ 228,229 

Blake, M. C., Jr.--------------_.. . 203

Brock, M. R____________________ 183 

Brokaw, A. L. . __. - 180

Buck, K. L- 199

Buckmaster, J. L_____________.__ 217,218

Campbell, Esma...________...._ 202

Cannon, R. S., Jr. 199

Carter, W. D ... .... 196

Clark, W. E.......... ... ... 221

Conover, W. J___________________ 228

Cooper, J. B...... . .- .. ._.. 225

Culbertson, W. C.......__... - ....... 193,194

Dando, T. O.-...-._............_...... 218

Diaz, A. M........_........................ 230

Dickey, D. D............................... 197

Dlngman, R. J............................. 186,207

Dobrovolny, Ernest_______________ 192

Englund, K. J....____.._____......... 184

Ford, A. B.... . .... . . 208

Glustl, E. V .. - 212

Gott, G. B........................ ..-. .- 180

Gottfried, David... - 202

Grantz, Arthur. --_. . 203

Hamilton, Warren... 205,211

Harris, L. D..... .................... 135

Heyl, A. V ......... .... iH3

Hildebrand, F. A. 181

Hoover, D. L-. 189

Hotz, P. E.. .. - 206

Hubbard, H. A.. 208

Huffman, Claude, Jr_. ... 214

Hyden, H. J.... ...... ... .... 182

Ingram, B. L ..._... . ___ 216

Isbister, John-....__._ . ......__ 226

Izett, G. A -._......_ . 189

Note. Numbers refer to articles.

Jackson, E. D. 

Johnson, A. I..

206

236

Kiilsgaard, T. H. ........... ............ 183

Lakin, H. W .... ............ 215

Leonard, B. F ... ... 204

Loving, H. B. _.............__._ 218

Mabey, D. H.............................. 180,187

McCarthy, Howard. . 180

Mallory. W. W_......... ........... __ 195

Marcher, M. V .... .... . 191

Meade, R. H. 210

Meyer, W. R ......... ...._ __ 235

Moore, J. G___________.________ 203

Moore, Roosevelt.. 202

Murphy, W. D.............................. 217

Myers, W. B................................. 211

Newport, T. G ... .......... - 223

Norris, S. E .............................. 224
INDEX E189

Article Article Article 

Oda.Uteana. . - 180 Rosholt, J. N _ ___________.___.._____.____ 209 Thompson, C. E. ........... ................. 215

Tooker, E. W.. ............................ 187,188

J---------- Turcan, A. N, Jr............................ 222 Page, H. G ......... 237,238

 Schultz, L. G__ __________________ 198 

i^arKer, K. 1, ........................ isi John................................. 232,233 Waldrop, H. A. _.___. .__._. 182

Perlmutter, N. M ________________ 219 ,_ , . , , ,. -^ T _ 

. . ' Sheppard, E. A._...................... ... 192 Walters, K. L.. . ......^........ .......... 227

P 11 1 -R r r ^ Sigvaldason.G.E....... ................... 200,201 Ward, P. E._ .......... ......... ..... 190

rrm, ±i.u_._ .-......................._... Zdb Soren, Julian................................. 219 Wayman, C. H. ........... ........... 237,238,239

Stallman, E. W........ ...................... 231 White, D. E__... ..............-..... .. 201

Richmond, G. M...... ....................... 213 Stern, T. W.................................. 208 Wiitala, S. W.. ........... ... .. . .. 223

Roberts, R. J _________________ 187,188 Stewart, J. W. .............. _ ... __ ... . _ 220 Wright, J. C. _ .... __ ... _____ ...... 197,198

Robertson, J. B. __ . _ ....... _ ........ 237,238

Roen, J. B. ______ . ________ ... __ 184 Taylor, R. B ________ .' __ . ______ 189 Yotsukura, Nobuhlro... . ___ . 234

Note Numbers refer to articles.

o