Articles 

1-59

GEOLOGICAL SURVEY 

RESEARCH 1963

Short Papers in Geology 

and Hydrology

GEOLOGICAL SURVEY PROFESSIONAL PAPER 475-B

Short Papers in Geology 

and Hydrology 

Articles 1-59

GEOLOGICAL SURVEY RESEARCH 1963

GEOLOGICAL SURVEY PROFESSIONAL PAPER 475-B

Scientific notes and summaries of investigations prepared 

by members of the Conservation, Geologic, and Water 

Resources Divisions

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 59 articles is one of a series to be released in 1963 as chapters of Professional Paper 475. The articles report on scientific and economic results of current work 

by members of the Geologic, Water Resources, and Conservation Divisions of the United 

States Geological Survey. Some of the papers present the results of completed parts of continuing investigations; others announce new discoveries or preliminary results of investigations that will be discussed in greater detail in reports to be published in the future. Still 

others are scientific notes of limited scope, and short papers on methods and techniques.

Chapter A of this series will be published later in the year, and will present a synopsis 

of results of work done during the present fiscal year.

XTHOMAS B. NOLAN,

Director.

in

CONTENTS

Page 

Foreword..----____-__-_-___-_-___--____________________._________________________--_-__------_-_--_-___-_---__ in

GEOLOGIC STUDIES 

Geochemistry, mineralogy, and petrology

1. Significance of O18/O16 and C13/C12 ratios in hydrothermally dolomitized limestones and manganese carbonate replacement ores of the Drum Mountains, Juab County, Utah, by T. S. Levering, J. H. McCarthy, and 

Irving Friedman_____________________________________________________________________________________ Bl

2. Thortveitite associated with fluorite, Ravalli County, Mont., by R. L. Parker and R. G. Havens __-____-__-____ 10

3. Beryllium in the tin deposits of Irish Creek, Va., by F. G. Lesure, T. H. Kiilsgaard, C. E. Brown, and M. E. Mrose__ 12

4. Beryllium and fluorine in mineralized tuff, Spor Mountain, Juab County, Utah, by W. R. Griffitts and L. F. Rader, Jr. 16

5. Beryllium and fluorine content of some silicic volcanic glasses from Western United States, by W. R. Griffitts and

H. A. Powers._________._.__.__.______...___.________._.______._.__-________._. _.___ _ 18

6. Some new data on the arsenic content of basalt, by A. J. Bartel, E. J. Fennelly, Claude Huffman, Jr., and L. F.

Rader, Jr. ________________...___.__._.___..__._ ______.__._________._______ .._ 20

7. Preliminary relations in the system Na2B4O7-Ca2B6Ou-H2O, by B. B. Hanshaw. _______________________________ 24

8. Variation in minor-element content of desert varnish, by H. W. Lakin, C. B. Hunt, D. F. Davidson, and Uteana Oda_ 28

9. Relation of ion-exchange capacity to mineral composition and grain size of weathered crystalline rocks at the Georgia

Nuclear Laboratory, Dawson County, Ga., by J. W. Stewart---___-_---_---_----_-------------_----------- 32

10. Cordierite-bearing mineral assemblages in Precambrian rocks, Central City quadrangl -, Colorado, by P. K. Sims

and D. J. Gable...._______..___.____..._____.__________.__._._____..__ ____.. _-_ _ 35

11. Bedded barite deposits of the Shoshone Range, Nev., by K. B. Ketner _______-___-_-________--_---_-----_--- 38

12. Accretionary lapilli in rocks of the Carolina slate belt, Stanly County, N. C., by H. W. Sundelius,______-__-----_ 42

13. Composition and origin of siliceous mudstone in the Carlin and Pine Valley quadrangles, Nevada, by K. B. Ketner

and J. F. Smith, Jr___.___________________________________.___ ___________________ 45

14. Tuffaceous sandstones in the Triassic Chinle Formation, Colorado Plateau, by R. A. Cadigan _________________ 48

15. Petrographic characteristics of some welded tuffs of the Piapi Canyon Formation, Nevada Test Site, Nev., by J. T.

O'Connor.---------______-_----________-----_____-_---_-----__------------- T _----------------------- 52

Geochronology

16. Potassium-argon and lead-alpha ages for stratigraphically bracketed plutonic rocks in the Talkeetna Mountains,

Alaska, by Arthur Grantz, Herman Thomas, T. W. Stern, and N. B. Sheffey--_______-_--_-____--___-__----- 56

17. Radium migration and its effect on the apparent age of uranium deposits at Ambrosia Lake, N. Mex., by H. C.

Granger _ ___-_-___-______________________________-___--______-_____--_--__---_--_-_--------------- 60

Geophysics

18. Gravity survey of the Gold Meadows stock, Nevada Test Site, Nye County, Nev., by D. L. Healey and C. H. Miller. _ 64

19. Depiction of soil-covered structures by infrared aerial photography, by W. A. Fischer __________________--_--- 67

20. Two nomographs for computation of standard equations in earth-resistivity interpretation, by J. H. Scott--------- 71

Stratigraphy and structural geology

21. New Cambrian, Ordovician, and Silurian formations in the Independence quadrangle, Inyo County, Calif., by D. C.

Ross_____-.________.__._____.__.______________________._____.____-_______ 74

22. The Livingston Group of south-central Montana, by A. E. Roberts__________-________---__-_-___-__-_------ 86

23. Age of certain post-Madison rocks in southwestern Montana and western Wyoming, by J. T. Dutro, Jr., and W. J.

Sando_____..____________________________________________ __________________ 93

24. Emendation of the Kelvin Formation and Morrison(?) Formation near Salt Lake City, Utah, by M. D. Crittenden, Jr_ 95

25. Apache Creek Sandstone Member of the Pierre Shale of southeastern Colorado, by G. R. Scott and W. A. Cobban. 99

26. Stratigraphic section at Island Beach State Park, N.J., by P. R. Seaber and John Vecchioli- ______________---_-- 102

27. Aerial reconnaissance of the outer Shumagin Islands, Alaska, by Arthur Grantz_________________________------- 106

28. Structural influence on development of linear topographic features, southern Baranof Island, southeastern Alaska,

by D. A. Brew, R. A. Loney, J. S. Pomeroy, and L. J. P. Muffler________________________________------_- 110

29. Geology of the Gua'nica-Guayanilla Bay area, southwestern Puerto Rico, by I. G. Grossman_______________---_-_ 114

Paleontology and paleoecology

30. Caradocian (Middle Ordovician) fossiliferous rocks near Ashland, Maine, by R. B. Neuman_____________________ 117

31. Miocene vertebrates from Middle Park, Colo., by G. A. Izett and G. E. Lewis-------------_---_---------,---_- 120

32. Paleoecology of the Permian Phosphoria Formation and related rocks, by E. L. Yochelson_______________________ 123
VI CONTENTS

Geomorphology, glacial geology, and glaciology Page

33. Quaternary events along the unglaciated lower Ohio River valley, by L. L. Ray_-------_-_--------------------- B125

34. Karst topography in the Gros Ventre Mountains, northwestern Wyoming, by W. R. Keefer__ ___________________ 129

35. Eruptions of water and sand resulting from an earthquake near Concepci6n, Chile, by Kenneth Segerstrom, Lorenzo

	Casertano, and Carlos Galli O_-__________________-_______------------_-------------------------------- 131

36. Paradise debris flow at Mount Rainier, Wash., by D. R. Crandell_____________----------_--_----------------- 135

37. Preliminary report on glaciology and glacial geology of the Thiel Mountains, Antarctica, by B. G. Andersen_______ 140

38. Influence of snow cover on frost penetration, by D. B. Krinsley________________-___-__--_-___---------_--_--- 144

Astrogeology

39. Origin and nature of the probable skeletal fuzz on the Moon, by C. R. Warren____-______-_-_-_---------_------ 148

Marine geology

40. Pillow structures of submarine basalts east of Hawaii, by J. G. Moore and R. K. Reed__________________________ 153

Economic geology

41. Estimates of world bauxite reserves and potential resources, by S. H. Patterson__-________--_--_--_-________-_- 158

Analytical techniques

42. Lead reference sample for isotopic abundance ratios, by M. H. Delevaux._____________________________________ 160

43. Synthesis of liebigite, by Robert Meyrowitz, D. R. Ross, and A. D. Weeks______-----_-___-_-___-________-_-_- 162

44. Separation of tellurium from iron and gold using tributyl phosphate and ether, by C. E. Thompson and H. W. Lakin__ 164

45. Determination of traces of boron in halite and anhydritic halite rocks, by F. S. Grimaldi and F. O. Simon___-____. 166

46. Determining density and porosity of tuff containing zeolites, by D. D. Dickey and E. F. Monk__________________ 169

HYDROLOGIC STUDIES 

Engineering hydrology

47. Land subsidence in the Arvin-Maricopa area, San Joaquin Valley, Calif., by B. E. Lofgren______________________ 171

48. Waterpower investigations of lakes in Alaska, by Arthur Johnson.___________________________________________ 176

Ground water

49. Ground water in the Nahud outlier of the Nubian Series, Kordofan Province, Sudan, by H. G. Rodis and Wilson

Iskander________________________________________-_________-___-__-__----_-------_-----------___--- 179

50. Contamination of ground water by sea-water intrusion along Puget Sound, Wash., an area having abundant precipitation, by G. E. Kimmel____-_______________________________________________________ 182

51. Influence of land-surface conditions on ground-water temperatures in southwestern Suffolk County, Long Island,

N.Y., by E. J. Pluhowski and I. H. Kantrowitz____________________________________________ 186

52. Seasonal changes in the chemical quality of shallow ground water in northwestern Alaska, by A. J. Feulner and R. G.

Schupp__--_________-____.___________________._________________________. _- 189

Analytical hydrology

53. Changes in ground-water movement and bank storage caused by flood waves in surface streams, by H. H. Cooper,

Jr., and M. I. Rorabaugh_____________________________________________________________________________ 192

54. Preparation of type curves for calculating T/S of a wedge-shaped aquifer, by I. S. Papadopulos._________________ 196

55. Effect of the injection scheme on the spread of tracers in ground-water reservoirs, by Akio Ogata__________--___- 199

Experimental hydrology

56. Effect of base-level changes on bedding development in a laboratory flume, by A. V. Jopling___________________ 203

57. Factors influencing the survival of Escherichia coli in detergent solutions, by C. H. Way man, J. B. Robertson, and

H. G. Page______--____________________________________-__-__-______--_----------__-_____________--_- 205

58. Effect of detergents on the viscosity of water containing bacteria and clay in suspension by C. H. Wayman, H. G.

Page, and J. B. Robertson___________________________________________________________________________ 209

59. Adsorption of the surfactant ABS35 on montmorillonite, by C. H. Wayman, J. B. Robertson, and H. G. Page____ 213

INDEXES 

Subject...-._._...___.___._._.______________-_______ . ____-__-__ _-________--_---_--________________- 217

Author.._.___._._____---_-.-_.______.--_ __-.______-- -----_.--___-_---_____---_.-_.____..___ 219
Article 1

SIGNIFICANCE OF O"/O" AND C13/C" RATIOS IN HYDROTHERMALLY

DOLOMITIZED LIMESTONES AND MANGANESE CARBONATE REPLACEMENT ORES OF

THE DRUM MOUNTAINS, JUAB COUNTY, UTAH l

By T. S. LOVERING, J. HOWARD McCARTHY, and IRVING FRIEDMAN, Denver, Colo.

Abstract. The OW/OW ratio of the carbonate rocks ranges 

widely and decreases as a hydrothermal conduit is approached. 

However, the C^/O13 ratio is almost constant within a carbonate 

unit, even where the unit changes from limestone to dolomite 

or manganese carbonate, but the ratio is significantly different 

from unit to unit. The isotopic relations suggest that ttie 

hydrothermal dolomitizing solution had a high chloride content 

and a low CCV2, SO*"2, and F"1 content.

In the Drum Mountains of central Utah, about 40 

miles northwest of Delta, Cambrian carbonate beds are 

much faulted and locally are mineralized and hydrothermally altered. Limestone has been changed to 

dolomite and in many places dolomitization is obviously 

related to fracture control. Because work elsewhere 

(Engel and others, 1958) had shown a relation between 

the isotopic composition of hydrothermally altered 

carbonate rocks and distance from hydrothermal conduits, further investigation seemed warranted and the 

Drum Mountain area was chosen for study. Some 50 

samples were collecte'd from 6 different beds and most 

of these were analyzed for the isotopes of carbon and 

oxygen.

GEOLOGY

The Drum Mountains form a westward-tilted structural unit, in which about 9,000 feet of Precambrian 

and Cambrian quartzite is overlain by 3,000 feet of 

Cambrian limestone and dolomite. The beds are cut 

by many east- and northeast-trending faults, most of 

which are earlier than quartz monzonite dikes, pebble 

dikes, and vein deposits. Tertiary volcanic rocks overlap the Cambrian sediments. The geology and ore deposits have been described by Crittenden and others 

(1961). These geologists did not correlate the formations with standard sections, but instead designated the

1 Presented at the Vernadsky Centennial Jubilee Celebration, Moscow, 

U.S.S.R., March 1963.

carbonate units by letters and the shale units by numbers, in sequence from the bottom unit. Dolomite A 

and limestone B are replaced by manganese carbonate 

ore near their intersection with some of the mineralization faults. Feeding fissures locally contain 

veins in which rhodochrosite and base metal sulfides 

are present, and near some intrusions the carbonate 

beds are pyrometasomatized. The relation of dolomite 

to limestone in some places clearly indicates fracture 

control of the dolomitized limestone. Increasing recrystallization of dolomite as certain fractures are approached also suggests that some fractures were hydrothermal conduits during dolomitization. Dolomite 

is the earliest and the most extensive product of hydrothermal alteration of the limestone and has been replaced in turn by manganese carbonate in certain beds. 

Later than the dolomite and probably later than the 

manganese carbonate are extensive reefs of jasperoid 

which are primarily controlled by fractures. All the 

fissure-controlled jasperoid tested by the writers has 

the characteristics of productive jasperoid as defined 

by Levering and Hamilton (1962).

The hydrothermal dolomite is coarser grained than 

the limestone from, which it formed, and in places evidence of two or more episodes of recrystallization of 

the hydrothermal dolomite can be observed. Quite 

commonly the hydrothermal dolomite follows the upper 

contact of a limestone bed where, it lies beneath a less 

permeable bed such as shale, argillaceous limestone, or 

dolomite. Dolomitization was accompanied by an appreciable increase in manganese, and the manganese 

carbonate mineralization may represent only the final 

stage of a continuing carbonate alteration in which 

magnesium dominated the early phase and manganese 

the later phase.

ART. 1 IN U.S. GEOL. SURVEY PROP. PAPER 475-B, PAGES B1-B9. 1963.

Bl
B2 GEOCHEMISTRY, MINERALOGY, AND PETROLOGY

VARIATION IN SO18 IN CARBONATE ROCKS AND ORES 

Relation to beds

The range of SO18 values 2 within individual beds 

(tables 1.1 and 1.2, p. B8-B9) tends to be quite similar 

in given are&s and is clearly related to fracture control 

rather than to individual beds (figs. 1.1 and 1.2). 

There is a general rise in the SO18 values in the samples 

from the higher beds but the upper limit of the SO18 

values is very similar in the samples of the least altered 

part of dolomite A, the lowest carbonate bed, and in 

dolomites N and L, the highest beds sampled. Figures 

1.1 and 1.2 show that all samples having delta values 

less than 14°/00 are close to a mineralized fracture. The 

conclusion seems inescapable that the change in delta 

values of the oxygen is related to hydrothermal alteration and not to the primary composition of the rocks.

Relation to ore and mineralizing fractures

As shown in figure 1.1, samples 3 taken along the 

contact of limestone G and dolomite H, between the 

Staats fault and the Last Chance fault, are instructive. The upper part of limestone G has been dolomitized near its contact with dolomite H, and the zone of 

hydrothermal dolomite reaches its maximum thickness 

near two intersecting east-trending faults about midway between the Staats fault and the Last 'Chance fault 

(coordinates D,V, fig. 1.1), reaching deepest into limestone G at a dolomite breccia vein just north of sample 

34. This sample also has the lowest SO18 value of the 

samples collected in limestone G. Samples collected 

northward show a progressive increase in SO18 to 

sample 25, which has the highest delta value of the 

series, and is about midway between the dolomite breccia vein near 34 and similar dolomite breccia just south 

of 20 and 21. Such a change strongly suggests that the 

east-trending zone followed by the breccia was a conduit through which hydrothermal solutions rose and 

spread out along the contact of dolomite H and lime2 The isotopic composition of an element is commonly reported in "delta units," 

units that give the per mil (°/00) difference of the O 18/O« or C 13/C12 ratio from the 

ratio in a standard; thus SO18 is given by

o»

oi1

 i

018

o;«

X 1,000,

where O* is the unknown and O« is the standard.

The standard used for O18/O16 in this study is "mean ocean water," and for C13/C12 

is the "standard belemnite" shell originally chosen as a standard by Urey and others 

(1951).

8 Sample locations are shown on figures 1.1 and 1.2 and are identified 

by map coordinates (as B.III) in table 1.1 and also in the text where 

any specific sample is mentioned.

stone G. South of sample 34 the zone of hydrothermal 

dolomite cuts sharply back towards the contact with 

dolomite H. The dolomite changes abruptly into limestone : sample 35 is from the same bed as 34 and only 3 

inches to the south, but it is limestone. The delta value 

for the dolomite, however, is 13.5 °/00 in contrast to 

16.9 °/00 for the limestone from which it formed. This 

latter value is almost the same as that of hydrothermal 

dolomite sample 25 (16.8°/00), midway between the two 

dolomite breccia veins. The relation of the hydrothermal dolomite to a specific feeding fissure seems clear 

here from spatial relations; and the progressive change 

of the SO18 from 13.5 to 16.8°/00 as the zone of dolomite 

is followed northward away from the breccia vein is 

persuasive evidence for low SO18 values in the vicinity 

of hydrothermal conduits. In harmony with this view 

we find values of only 14.3 to 14.7°/00 in dolomite samples 20 and 21 collected between the jasperoid vein and 

the dolomite breccia seam a few hundred feet south of 

the Staats fault (E, V). Similar changes are shown in 

sample pair 41 and 42 (C,IV; C,III) from north of 

the Last Chance fault in the upper part of dolomite H, 

and in samples from dolomite L south of the Last 

Chance fault (B,III).

The area shown .in figure 1.1 shows evidence of orestage mineralization in few places, whereas the Staats 

mine area (fig. 1.2), 1,400 feet to the east, has produced 

a substantial amount of manganese oxide and manganese carbonate replacement ore from dolomite A and 

limestone B near the Staats No. 2 shaft. Dolomite A 

is normally a sandy argillaceous dolomite but changes 

to a magnetite-rich hydrothermal limestone (sample 

13, D,XI) just south of the southerly Staats fault 

where it is close to shallow intrusive rocks.

The relations of oxygen isotopes to the feeding fissures so clearly demonstrated in the area of figure 1.1 

are also apparent in the more highly mineralized and 

faulted areas shown in figure 1.2. Samples 14 and 15 

(D,IX), of manganoan dolomite (3 to 6 percent Mn) 

from the Staats No. 2 shaft, show the lowest delta 

values (6.3 and 5.0°/00 ) of any specimens collected in 

this study. In general the delta values increased to 

the south and east away from the strongly mineralized 

area. The marked difference of the delta values between sample pairs 2 and 3 (C,XIV) and 4 and 5 

(C,XIII), taken on opposite sides of the Keystone 

fault, suggests that solutions moved southward from 

the Keystone fault. The range of delta values for 

oxygen shown by samples from the dolomitized base of 

limestone B is almost as great as that for dolomite A 

and similarly, the lowest values are found in the 

manganoan dolomite sample 12 (C,XI), which was 

taken close to the Last Chance fault.
LOVERING, MCCARTHY, AND FRIEDMAN 

31 m 12

B3

R.nw.

200 400 600 800 1000 FEET

I____I____\____I I

D 

5.0 to 7.9

Variation in the delta values of isotope ratios

 

14.0 to 16.4

A 

8.0 to 10.9

O 

1.0 to 13.9

X 

16.5 to 19.0

FIGTJBE 1.1. Relation of SO18 to geology, Ibex mine area. All faults are premineralization and are shown 

by heavy lines where mineralized. Numbers are field numbers of samples. Letters indicate carbonate 

units, from youngest to oldest: dolomite N, limestone M, dolomite L, dolomite K, limestone J, limestone I (stippled), dolomite H, limestone G. All formations are of Middle Cambrian age. Arrows 

point in direction of increasing delta values. Geology from Crittenden and others (1961, pi. 20).
B4

vrrr H

GEOCHEMISTRY, MINERALOGY, AND PETROLOGY

I TT "YTT TTTT

R.IIW. R.IOW.

200 400 600

I

800 1000 FEET

Variation in the delta values of isotope ratios

a

5.0 to 7.9

A 

8.0 to 10.9

O 

II.0 to 13.9 14.0 to 16.4 16.5 to 19.0

FIGURE 1.2. Relation of SO18 to geology, Staats mine area. All faults are premineralization and are shown by heavy lines 

where mineralized. Numbers are field numbers of samples. B, limestone B or younger beds; crosshatch, shale 1; A, 

dolomite A (limestone B, shale 1, and dolomite A are Middle Cambrian) ; stipple, Prospect Mountain Quartzite (Lower 

Cambrian). Geology from Crittenden and others (1961, pi. 20).

The relation of samples 6, 7, and 8 (C,XII) taken 

in and near a pebble-dike zone a few hundred feet 

north of the Keystone fault is suggestive. Limestone 

B samples 6 and 7 from 5 feet north and 5 feet south 

of the pebble-dike fracture have similar delta values 

(12.7 and 12.0°/00 ), but the porous friable manganoan 

limestone in the pebble-dike zone itself has a higher 

value (15.1 °/00 ), suggesting that after an early stage 

when solutions of higher temperature affected the wallrocks for several feet on either side, the sheared limestone in the pebble-dike zone equilibrated with later 

solutions of different composition or temperature 

which incidentally supplied appreciably more manganese. All relations shown in figure 1.2 are in harmony 

with the conclusion reached in the Ibex mine area: the 

low delta values of oxygen are related to hydrothermal 

conduits and high temperatures.

VARIATION IN SO3 IN CARBONATE ROCKS AND ORES

Relation to beds

The range of delta values for C18 is very much smaller 

than that for O18, and with a single exception, all values lie between +0.1%0 and 3.7%0 . The precision 

of the 8C13 values is approximately ±0.2°/00, whereas 

the precision for the oxygen values is approximately 

±0.1%0 . As shown in table 1.2, there is no relation 

between the range of SO18 values and 8C18 values. The 

maximum range of 8C1S from hydrothermal dolomite, 

manganoan dolomite beds, and manganese carbonate 

replacement ore, excluding leached ("sanded") dolomite fragments in veins, is 0.9°/00 . This range is exhibited in dolomite A, in the dolomitized base of 

limestone B, and in dolomite H; and the standard deviation for SC18 at a 95-percent-confidence limit ranges
LOVERING, MCCARTHY, AND FRIEDMAN B5

from ±0.36°/00 to ±0.306°/00 only slightly more than 

the precision of the isotopic analysis.

The variation of the carbon isotopes from one bed 

to another, however, may be much greater than the 

variation within the bed. As shown in table 1.2, the 

scatter of the delta values of the carbon isotopes about 

the mean for a given bed is small, as measured by the 

standard deviation (<r), but the means are significantly 

different in different beds; for dolomite A, 8C13 is

 3.2°/00 (±0.36); for limestone B just above, it is

 1.9°/00 (±0.33); and for limestone G it is -0.25%0 

(±0.175).

Values for 8C13 reported by Craig (1953) show a 

range of 5.7°/00 (from +2.4 to 3.3) for marine limestones, and a range of 5.0 (from +2.7 to 2.3) for 

marine dolomites.

The range in 8C13 found in the least altered dolomite 

and limestone of the Drum Mountain samples is from

 3.7 to +0.5°/00, values that seem reasonable for marine carbonate rocks.

Although the lowest values of 8C13 are characteristic 

of the two lowest carbonate beds in the section, there 

is no consistent variation in the SC13 with distance 

above the base. As shown in figure 1.3, the sampling 

adequately indicates the lack of relation of delta values 

to distance above base but falls far short of providing a 

Qi3/Qi2 profiie Of the section.

Relation to ore and mineralizing fractures

As noted in the foregoing, there is no appreciable 

variation in the C13/C12 ratio related to the alteration 

of a limestone to dolomite or manganiferous dolomite, 

nor is there any observable change in the ratio related 

to distance from the hydrothermal conduits that were 

apparently followed by the altering solutions. Nevertheless there is a marked change in the C13/C12 ratio in 

some types of hydrothermal alteration. Samples 62 

and 63 (E,II, fig. 1.1) were obtained from altered fragments of dolomite in a vein of siderite and calcite. A 

dolomite fragment (63) adjacent to the wall had been 

slightly leached or "sanded," and although the outside 

of the fragment was friable, the interior was moderately 

dense. Its 8C13 value is 0.2°/00 , essentially that of 

specimen 57 (F,I), a sample of the unaltered limestone 

country rock ( 0.4°/00 ); in contrast sample 62, which 

had been reduced to a friable dolomite sand, has a 

8C13 value of - 2.1 °/00 , and its lower SO18 value of 6.9 °/00 

indicates activity of hydrothermal rather than supergene solutions. The calcite and siderite matrix of these 

dolomite fragments indicates solutions that carried a 

substantial amount of bicarbonate, and it is therefore 

not surprising that the carbon isotopes of the much 

leached dolomite sand show a marked change in value.

Ool L

Dol K

LSI

Dol K

LsG

1 f D

LS D

Dot A

-4

X

x-»-XX

x t xx

<a us>

JMI-U. .

X X 

X XgXI

x x! x _£>

X 

X Xj(X X

1 1 1 1

kO -3.0 -2.0 -kO 0 +1

SC13

EXPLANATION

10 QO

o 

350 u

0.

S 

300 w

z

o 

365 g

o

30 ~ 

I-"

Ul 

Ul

u. 5 2

co" 

to

15 2

o

X

o -

Dolomite

Leached (sanded) 

dolomite or limestone

Limestone

© ©

Leached (sanded) 

limestone

FIGURE 1.3. Range of SO18 values (solid and dashed lines) 

and average (dots) for each bed sampled.

Sample 13 (C,XI, fig. 1.2) has the lowest 8C13 value 

found (~5.8°/00 ) and represents dolomite A after it 

was pyrometasomatized into a calcite-magnetite rock 

containing minor pyrite. This type of alteration is 

associated with nearby intrusions that are later jthan 

the early stage of hydrothermal dolomitization. The 

exict relations could not be established in the field, but 

the introduction of magnetite and calcite, and the disappearance of dolomite, required different solutions and 

took place at an entirely different time than the hydrothermal dolomitization and formation of manganese 

carbonate ore.

THEORETICAL CONSIDERATIONS

Theoretical isotope-exchange reactions based on thermodynamic data have been substantiated by measurement of isotopic f ractionation in nature, and conversely 

isotopic measurements have shed new light on geologic 

processes. In general, exchange of light isotopes depends on their concentration and on the composition 

and amount of solution that contains them; the temperature of the system and the degree of equilibration 

attained are also of major importance. In a study of
B6 GEOCHEMISTRY, MINERALOGY, AND PETROLOGY

the Leadville Limestone, Engel and others (1958) demonstrated the temperature dependence of oxygen-isotope concentration in carbonates and chert. Assuming 

from geologic evidence an abundance of hydrothermal 

solution of fairly constant isotopic composition, they 

stated, "The highest SO18 values are obtained from 

bedded calcite, dolomite, and chert which have not been 

recrystallized by hydrothermal water (unaltered beds). 

The lowest SO18 values are from the hydrothermal calcite, dolomite, and quartz closest to ore which therefore 

seem to have recrystallized at highest temperature." 

This relation has been further verified by Clayton and 

Epstein (1958) and checked experimentally by McCrea 

(1950), who precipitated calcium carbonate in the laboratory and determined the O18/O16 fractionation factor.

The correlation of low SO18 values with hydrothermal 

conduits in the carbonates of the Drum Mountains thus 

is consistent with previous observations and known 

thermodynamic considerations of isotope-exchange 

reactions.

The isotopes of carbon can be expected to fractionate 

in roughly the same manner as those of oxygen. Although carbon-isotope fractionation in nature has been 

noted chiefly in biological processes, variation in temperature and isotopic concentration of an inorganic 

solution will cause isotopic-exchange reactions with 

ionic carbon compounds as with ionized oxygen compounds. The amount of C13 found in sedimentary and 

shell carbonates is in keeping with the calculated fractionation factors for the system carbonate-bicarbonatecarbon dioxide as worked out by various investigators, 

including Urey and Grieff (1935), McCrea (1950), and 

Craig (1953). The fractionation of carbon isotopes 

characteristic of the metabolic processes of different 

organisms plays an important role in the isotopic composition of precipitated carbonates. C12 tends to concentrate most in the organic carbon of low forms of 

plant life. Landergren (1954) states that the carbon 

isotope ratio is determined by the geochemical conditions prevailing during precipitation of the carbonate 

sediments, but so far as carbon is concerned the geochemical environment is largely the result of biochemical reactions.

As noted by Parker (1963) there is a spread in the 

SC13 between animals of different species of as much 

as 15%0 , although there is a spread of only 1°/00 between individuals of the same species. A given flora and 

fauna contributes organic carbon and carbon dioxide 

to an accumulating sediment, and this sediment should 

have a characteristic 8C13 value throughout the area in 

which the ecological conditions are the same. In the 

area studied by Parker a small bay in the upper

Laguna Madre, Tex. the 8C13 value of the bay water 

was +6.7%0 , whereas that of the sediment in the bay 

averaged about -13.3%0 . The 8C13 values of fish and 

shellfish in general ranged from about 8 to 18% 0 

and the 8C13 values for algae ranged from about 11 

to nearly 40%0 . The different constituents of an organism, showed differences in their ability to fractionate 

the available carbon; in general the fatty acids in an 

organism were 2 to 5 %0 less negative than the lipid 

fractions and both substances were more negative than 

the total organism. According to Parker (1963), 

"Fatty acids isolated from some of the organisms were 

always depleted in C13 relative to the whole organism. 

Organic matter preserved in the sediments should reflect in its C13/C12 ratio not only the species of the organism from which it is derived but the type of 

molecule preserved." The nearly constant value of the 

C13/C12 ratio in a given bed is thus consistent with a 

nearly constant biological and depositional environment while the bed was being deposited, and the isotopic 

variation from bed to bed is caused by changes in the 

isotopic composition of the carbon available during 

deposition, changes that in turn ultimately reflect the 

biochemistry of organic processes.

Weyl (1960), in discussing dolomitization by circulating brines, points out that most subsurface waters 

contain large amounts of calcium and magnesium and 

relatively small amounts of total carbonate, and states 

that "it is possible to show that dolomitization must 

take place approximately by the mole-for-mole replacement of calcite without introduction of external CO2 ." 

This process is represented by the equation Mg+2 

+ 2 aCO3-»CaMg(lCO3 ) 2 +Ca+2 . Such a mechanism of 

dolomitization is not only suggested by the lack of 

carbon-isotope fractionation in the Drum Mountain 

area but would indeed account for that lack.

The marked change in the C13/C12 ratio found in 

some other types of hydrothermal alteration means 

conversely that carbon was introduced in the solutions 

causing such alteration.

COMPOSITION OF HYDROTHERMAL SOLUTIONS

The change from limestone to dolomite is accompanied by a very slight increase in porosity, from about 

iy2 percent to 4 percent, and by a slight increase in 

permeability. Although the carbon isotopes of the 

dolomite and of the limestone from which it formed 

are indentical, the alteration was accompanied by a 

marked change in SO18 such as is appropriate to recrystallization in hydrothermal solutions. There was 

no significant change in either organic or inorganic 

carbon; the chief chemical change was, of course, the
LOVERING, MCCARTHY, AND FRIEDMAN B7

substitution of magnesium for about half of the calcium, but some manganese and iron were also added; 

these elements, however, rarely make up more than 

a small fraction of 1 percent of early formed hydrothermal dolomite. Negative evidence is the lack of 

fluorite or anhydrite either close to the hydrothermal 

channels or in the transition zone from hydrothermal 

dolomite to unaltered limestone, places where such minerals might be expected if the hydrothermal solutions 

contained appreciable fluorine or sulf ate. The relative 

constancy of the C13/C12 ratio in different formations 

where they pass from limestone to dolomite in the 

Drum Mountains indicates that the magmatic solutions 

contained little or no carbonate or bicarbonate in solution to equilibrate with the carbonate of the dolomite. 

Similarly the absence of fluorite or anhydrite indicates 

insufficient F-1 and SO4"2 to saturate the solutions with 

their calcium salts. It seems safe to assume, therefore, 

that the solutions were dominantly chloride brines containing much magnesium and appreciable amounts of 

iron and manganese. It is worthy of note that one of 

the earliest experiments in which calcium carbonate 

was converted to dolomite was carried out by heating 

calcium carbonate with a solution of magnesium 

chloride to 200 °C under a pressure of 15 atmospheres 

(Clark, 1924, p. 566).

Although the late veinlets of manganese carbonate 

both the black manganoan calcite and rhodochrosite 

were probably deposited from solutions containing appreciable bicarbonate, it is probable that the earlier 

manganoan dolomites containing as much as 6 percent 

manganese formed from chloride brines similar to those 

that converted limestone to dolomite but richer in manganese. It is quite in harmony with the field evidence

to suppose that the dolomitizing solutions became progressively richer in manganese chloride with time and 

that the manganoan replacement of the dolomite represents equilibration of the dolomite with later solutions 

that had a higher ratio of manganese to magnesium 

than the earlier solutions.

CONCLUSIONS

The evidence presented above seems to warrant the 

following conclusions:

1. The alteration of limestone to hydrothermal dolomite in the area studied was accomplished by hot solutions containing magnesium chloride with little if any 

bicarbonate, fluoride, or sulf ate ions.

2. The O18/O16 ratios decrease from the outer fringe 

of dolomitic alteration inward toward the feeding fractures, reaching their lowest values where there is the 

greatest evidence of higher temperatures and more intense mineralization.

3. The C13/C12 ratios remain relatively constant 

where limestone is dianged to hydrothermal dolomite 

or manganiferous dolomite but show substantial variation in other types of hydrothermal alteration.

4. C13/C12 ratios'should prove useful in correlating 

carbonate beds, especially where limestone is altered 

to hydrothermal dolomite in local but unpredictable 

areas.

5. Alteration by post-dolomite hydrothermal solutions containing carbon dioxide gave larger negative 

8C13 values than those of the original dolomite. This 

is true both for sanded dolomite in carbonate veins and 

for magnetite-calcite replacement deposits formed from 

dolomite.
B8 GEOCHEMISTRY, MINERALOGY, AND PETROLOGY

TABLE 1.1 Delta values of O18 and C13 of some limestones, 

dolomites, and manganese carbonate ores from the Drum 

Mountains, Utah

[Delta values determined by Irving Friedman]

Field No.

Map coordinates

 ,

Rock type

Distance above 

Cambrian Lower 

quartzite (feet)

Distance from

premineralization 

(feet) faul 

percent) (weight Mn 

o

0 

0

o

Remarks

Dolomite N

75 A.II Dolomite ... 4,510 10 17.4 -0.4 Basal bed.

Dolomite L

44

47 

46

45

BllII 

B.III

Average, bas

Dolomite ...

.....do... ... 

 . .do ...

.... .do... ...

4,360

4,360 

4,360

4,360

150

25 

100

300

<0.02

<.02 

<.02

<.02

18.2

18.5 

19.0

19.1

-0.3

-.2 

-.6

-.6

-.45

Recrystallized, 

light gray. 

Do. 

Recrystallized, 

vuggy. 

Recrystallized.

Do.lomite K

51 

48

A.III 

A.III

Average fonr

Dolomite ... 

.... .do... ...

3,950 

4,300

<50 

>200(?)

15.1 

18.7

lation ___________________

0.0

.5

.25

Just above Limestone J. 

Black, finely 

crystalline, 

syngenetic(?) .

Limestone J

43 B,IV Dolomite. . 3,885 50 17.2 0.1 Fine grained, dark 

gray.

Limestone I

62 

63

57

E.II

Dolomite ... 

 do.......

Limestone2,870 

2,870 

2,580

0 

0

>500

0.30

.75

6.9 

11.5 

14.6

-2.1 

-.2 

-.4

"Sand" in sideritecalcite vein. 

Sanded, 1 ft from 

62. 

Midway between 

jasperoid and Joy 

fault.

Dolomite H

38

39 

21 

59

60 

37

42

41

58

D,V

D,V 

E.V

D'V

C.IV

Dolomite...

Limestone . 

Dolomite ... 

Calcitized 

dolomite. 

Dolomite ... 

 ..do. ..

 -do ...

... -do ... 

- do......

2,255

2,197 

2,197 

2,525

2,525 

2,197

2,545

2,475 

2,560

M

50

<5

<7.5 

»75

>50 

>500

Average, formation, dolomite and limestone 

Average, lower half of formation, dolomite 

stone.. -----

Average, upper half of formation, d

0.10

.20 

.02

.02 

.15

......

13.0

14.4 

14.7 

14.9

15.6 

16.4

16.0

17.1 

17.1

and lime-

-0.6

-.2 

.0 

- .5

- .3

.2

- .7

.1 

- .6

.3

-.15 -.40

Fault, 1-inch 

throw, between 

38 and 39. 

3 ft south of 38. 

10 ft above 20. 

About 5 ft from 

Staats fault. 

2^ ft from 60. 

About 100 ft 

southeast of 39. 

Close to Last 

Chance fault.

15 ft south of 

mineralized 

jasperoid.

TABLE 1.1 Delta values of O18 and C33 of some limestones, 

dolomites, and manganese carbonate ores from the Drum 

Mountains, Utah Continued

[Delta values determined by Irving Friedman]

Field No.

8

1 

§

Rock type

Distance above

Lower Cambrian 

quartzite (feet)

Distance from

premineralization 

fault (fee)

"8 

a

! or

6

0

3 

O

Remarks

Limestone G

34

20 

29 

28 

27 

26 

25 

35

D,V

E,V 

D,V 

D,V 

D,V 

D,V 

D,V 

D,V

Dolomite ...

.....do . 

Limestone. 

Dolomite ... 

.....do .

.....do 

.....do... ... 

Limestone.

2,165

2,195 

2,195 

2,195 

2,195 

2,195 

2,995 

2,165

Average, dolomite and limestone (t(

<100

50 

80 

110 

160 

240 

275 

<100

0.10

.20 

.2 

.10 

.10 

.10 

.10 

.05

13.5

14.3 

16.1 

16.2 

16.4 

16.5 

16.8 

16.9

>p30ft) ... ...

-0.5

-.2 

.0 

-.1

-.3 

-.2 

.0

-.22

Coarse, recrystallized, 3 in. from 

35, same bed. 

50 ft from metallized jasperoid. 

6 in. below 

dolomite. 

75 percent recrystallized. 

40 percent recrystallized. 

25 percent recrystallized. 

10 percent recrystallized. 

Unaltered limestone, 3 in. 

south of hydrothermal dolomite 34.

Limestone B

12

7 

6

9

8

1 

2 

5

C,XI

C.XII 

C.XII

C,XI 

C.XII

B,XIV

c.xrv 

c,xiv

Manganoan 

dolomite. 

Dolomite ..

Calcitized 

dolomite. 

Dolomite.. Manganoan limestone. 

Dolomite . .

.....do .

... -do... ...

Average, dolomite and llm 

Average, dolomite (No. 8 o

40

40 

45

40

45(7)

40 

40 

40

<5

105 

95

50 

100

400 

100 

10

3.0

.2 

.6

.4 

1.0

.75 

.5 

.6

6.7

12.0 

12.7

14.6 

15.1

15.6 

16.5 

16.7

mitted)...... ....... .......

-1.5 

-1.9

-1.9

-2.4 

.0

-1.9 

-2.0 

-2.4

-1.75 -2.0

From manganese 

prospect pit.

5 ft north of 

pebble dike. 

5 ft south of 

pebble dike. 

Very fresh 

Sanded manganoan limestone 

from pebbledike zone. 

Coarsely crystalline. 

Medium-fine 

crystalline. 

Fine- to mediumfine crystalline.

Dolomite A

15

14 

13

3 

11

4 

10

D.IX

D,IX 

C.XI

C,XIV 

C,XI

c,xm C,XI

Average, dol 

Average, dol

Manganoan dolomite.

Manganoan calcite. 

Limy dolomite. 

Dolomite ...

..... do 

 .do... ...

smite and lim 

[>mite only

<20

<20 

5

2

2 

2

<60

<60 

80

20 

3

10 

40

3.0

5.0

.7

.2

.75

.5 

1.4

5.0

6.3 

9.5

9.6 

13.4

18.2 

19.0

-3.2

-2.8

-5.8

-3.1 

-3.3

-3.2 -3.7

-3.6 -3.2

Some dark early 

MnCOs.

White late 

MnC03 . 

Pyrometasomatic.

Altered and 

porous.

Very fresh.
LOVERING, MCCARTHY, AND FRIEDMAN

TABLE 1.2. Summary of _O18 and _C13 values and variations for samples listed in table 1.1

B9

Unit

Dolomite L____-___ ____ ____

Dolomite K__ _ _ _____

Dolomite H______ _ __ _____

Limestone G _ -___ ___,_ ___

Limestone B 2 __ __ ___ ____

Total 

number of

samples

4

3

3

9

8

7

6

Max

19. 1

18.7

14. 6

17.1

16.9

16.7

19.0

W/oo)

Min

18.2

15.1

6.9

13.0

13.5

6.7

5.0

Range

0.9

3.6

7.7

4.0

3.4

10. 0

14.0

Max

-0.3

.5 -.2

.2

.0 -1.5

-2.8

Min

-0.6

.0

-2. 1

-.7

-.5

-2. 4

-3.7

-O

.Range

0.3

.5

1.9

.9

.5

.9

.9

E<Yoo)

IHO»

0.22

.3

1. 12

.31

. 175

. 33

.36

Median

-0.45

. 1

-.4

-.3

-.25

-1.9

-3.2

Avg

-0.45

.2 -.90

-.24

-.22

-2.00 -3.2

i Where a is the standard deviation, computed from range of SC® values and total number of samples, 

a Sanded manganoan limestone (sample 8) from pebble-dike fissure omitted. 

» Magnetite-calcite replacement rock (sample 13) omitted.

REFERENCES

Olark, F. W., 1924, Data of geochemistry: U.S. Geol. Survey 

Bull. 770, p. 566.

Clayton, R. N., and Epstein, S., 1958, The relationship between 

OM/OM ratios in co-existing quartz, carbonate, and iron 

oxides from various geological deposits: Jour. Geology, 

v. 66, p. 352-373.

Oraig, Harmon, 1953, The geochemistry of the stable carbon 

isotopes: Geochim. et Cosmochim. Acta, v. 3, p. 61, 65.

Crittenden, M. D., Jr., Straczek, J. A., and Roberts, R. J., 1961, 

Manganese deposits in the Drum Mountains, Juab and Millard Counties, Utah: U.S. Geol. Survey Bull. 1082-H, 

p. 493-544.

Engel, A. E. J., Clayton, R. N., and Epstein, S., 1958, Variations 

in isotopic composition of oxygen and carbon in Leadville 

limestone in its hydrothermal and metamorphic phases: 

Jour. Geology, v. 66, p. 374-393.

Landergren, Sture, 1954, On the relative abundance of the 

stable carbon isotopes La marine sediments: Deep-Sea Research, v. 1, p. 98-120.

Lovering, T. G., and Hamilton, J. C., 1962, Criteria for the recognition of jasperoid associated with sulflde ore: Art. 63 in 

U.S. Geol. Survey Prof. Paper 450-C, p. C9-C11.

McCrea, J. M., 1950, On the isotopic chemistry of carbonates 

and a paleotemperature scale: Jour. Chem. Physics, v. 18, 

p. 849-857.

Parker, P. L., 1963, Biogeochemistry of the stable isotopes of 

carbon in a marine bay [abs.]: Geol. Soc. America Spec. 

Papers, no. 73, p. 213.

Urey, H. C., and Grieff, L. J., 1935, Isotopic exchange equilibria: 

Am. Chem. Soc. Jour., v. 57, p. 321-327.

Urey, H. C., Lowenstam, K. A. Epstein, S., and McKiuney, C. R., 

1951, Measurement of paleotemperatures and temperatures 

of the Upper Cretaceous of England, Denmark, and the 

southeastern United States: Geol. Soc. America Bull., v. 62, 

no. 4, p. 399-^16.

Weyl, P. K., 1960, Porosity through dolomitization: conservation of mass requirements: Jour. Sed. Petrology, v. 30, 

p. 85-90.
Article 2

THORTVEITITE ASSOCIATED WITH FLUORITE, RAVALLI COUNTY, MONTANA

By RAYMOND L. PARKER and RAYMOND G. HAVENS, Denver, Colo.

Abstract. Thortveitite, (Sc, Y) 2 Si2C>7, has been found associated with fluorite at the Crystal Mountain fluorite deposit 

near Darby, Ravalli County, Mont. This is the first occurrence 

of the mineral to be reported in the Western Hemisphere and 

is a new mode of occurrence for the mineral, previously found 

only in granitic pegmatites.

The rare scandium mineral thortveitite, ((Sc, Y) 2 

Si2O7 ), has been found in the Crystal Mountain fluorite 

deposit 26 miles east of Darby, Ravalli County, Mont. 

This is the first occurrence of the mineral to be reported 

in the Western Hemisphere.

Thortveitite has been reported from only two other 

areas: southern Norway, where it is found in several 

pegmatites in the Iveland-Evje district (Schetelig, 

1922; Neumann, 1961); and Madagascar, where it occurs in the Befanamo pegmatite (Lacroix, 1920) and 

in the pegmatite district of Berere (Behier, 1960, p. 20).

The Crystal Mountain fluorite deposit has been described by Taber (1952, 1953) and by Weis and others 

(1958). The deposit consists of tabular bodies of 

fluorite in coarse-grained biotite granite containing 

xenoliths of biotite-quartz-plagioclase gneiss, hornblende-plagioclase gneiss, and pegmatitic granite. 

The tabular bodies of fluorite, and the foliation in the 

gneissic xenoliths, strike about north and dip 20°-30° E. 

Dikes of granite cut the fluorite bodies and the 

xenoliths.

Thortveitite occurs in the main ore body at Crystal 

Mountain in dark-purple fluorite that contains considerable biotite and subordinate sphene, quartz, 

oligoclase, rare-earth-bearing apatite, and minor green 

amphibole. Brown vitreous grains of metamict fergusonite (Weis and others, 1958) and a green opalescent metamict mineral, probably thorianite, are dispersed in this rock also; both of these minerals are 

strongly radioactive.

The thortveitite crystals within fluorite are embayed 

locally by fluorite, and some crystals contain inclusions 

of fluorite, suggesting that the two minerals were 

formed nearly contemporaneously.

The thortveitite forms elongate subhedral to euhedral 

crystals, many as long as 3 mm, which range from colorless and clear to smoke colored and translucent. Some 

of the smoke-colored crystals contain alternating clear 

and translucent lamellae that appear to be normal to 

the c-axis. The smoke-colored grains possibly reflect 

radiation damage caused by nearby crystals of fergusonite and thorianite. As seen in thin section, some of 

the thortveitite crystals are twinned polysynthetically.

X-ray diffraction powder patterns of the Crystal 

Mountain thortveitite and thortveitite from the Tuftane pegmatite, Iveland, Norway, seem identical. The 

strongest lines (in angstrom units) and estimated intensities (measured by F. A. Hildebrand, U.S. Geological Survey) of the Crystal Mountain mineral are: 

3.12 broad (10), 2.93 (7), 1.647 (5), 2.17 (4), 5.10 (3), 

2.05 (2),2.60 (2).

A semiquantitative spectrographic analysis of the 

Crystal Mountain thortveitite is given in the accompanying table. The thortveitite differs principally 

from that of Norway and Madagascar (Levinson and 

Borup, 1960) in its lower zirconium content and lack

Semiquantitative spectrographic analysis of thortveitite from Crysta'

Mountain, Ravalli County, Mont. 

[Analyst: Raymond G. Havens, U.S. Geological Surveyj

Si__.

Sc___

¥____

Dy_._

Er_._

Gd_ _

Yb_-

Ho 

Tb_._

Percent

______ Major

constituent.

______ Major

constituent.

----- 3

_____ .5

_____ .5

--_-. .3

__-_- .3

--_-. .15

_._.- .07

Percent 

0.07 

.05 

. 3 

. 15 

. 15 

.03 

.015

1

2

15

15

Percent

0. 15 

. 15 

.07 

.03 

.02 

. 015 

.01 

.007 

.0005

Looked for but not detected: Na, K, P, Ag, As, Au, B, Ba, Bi, 

Cd, Co, Cu, Ga, Ge, Hf, Hg, In, Li, Ni, Pb, Pd, Pt, Re, Sb, 

Sr, Ta, Te, Tl, U, W, Zn, Eu. Usual sensitivities for some 

elements do not apply because of high background and interference by rare-earth lines.

BIO

ART. 2 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B10-B11. 1963.
PARKER AND HAVENS Bll

of hafnium. The content of titanium, thorium, calcium, 

niobium, and possibly small amounts of other elements 

in the Crystal Mountain thortveitite may be due to 

minute inclusions of sphene, thorianite, fluorite, and 

fergusonite, although crystals selected for analysis 

appeared free of inclusions under the binocular 

microscope.

The association of thortveitite with fluorite in the 

deposit at Crystal Mountain represents a new type of 

occurrence, which contrasts strongly with the occurrences in granitic pegmatites of Norway and 

Madagascar.

REFERENCES

Behier, Jean, 1960, Contribution a la Mineralogie de Madagascar : Ann. Geol. Madagascar, no. 29, p. 19-20.

Lacroix, Alfred, 1920, Sur I'existence a Madagascar d'un silicate 

de scandium et d'yttrium, la thortveitite: Acad. Sci. (Paris) 

Comptes rendus, v. 171, p. 421-423.

Levinson, A. A., and Borup, R. A., 1960, New data on the hafnium, zirconium, and yttrium content of thortveitite: Am. 

Mineralogist, v. 45, p. 712-715.

Neumann, Henrich, 1961, The scandium content of some Norwegian minerals and the formation of thortveitite, a reconnaissance study: Norsk Geol. Tidsskr., v. 41, p. 197-210.

Schetelig, Jakob, 1922, Thortveitite, a silicate of scandium: 

Norsk Geol. Tidsskr., v. 6, p. 233-244.

Taber, J. W., 1952, Crystal Mountain fluorite deposits, Ravalli 

County, Montana: U.S. Bur. Mines Kept. Inv. 4916, 8 p.

 1953, Montana's Crystal Mountain fluorite deposit is 

big and high grade: Mining World, v. 15, no. 7, p. 43-46.

Weis, P. L., Armstrong, F. C., and Rosenblum, Samuel, 1958, 

Reconnaissance for radioactive minerals in Washington, 

Idaho, and western Montana, 1952-1955: U.S. Geol. Survey 

Bull. 1074-B, p. 19-21.

680231 O 63
Article 3

BERYLLIUM IN THE TIN DEPOSITS OF IRISH CREEK, VIRGINIA

By F. G. LESURE, T. H. KIILSGAARD, C. E. BROWN, and MARY E. MROSE, 

Washington, D.C.

Abstract. Thirty-three chip 'and grajb samples from two adits 

and several caved cuts of the old Irish Creek tin mine, Bockbridge County, Va., range in beryllium content from less than 

0.0006 to as much as 0.67 percent BeO. Although only a few 

samples contained visible beryl, nearly all the samples contained beryllium in amounts greater than average for granodiorite, the host rock of the deposit. Further prospecting for 

beryllium-bearing material may be warranted.

Recent discoveries of beryllium associated with tinr 

deposits on the Seward Peninsula, Alaska (Sainsbury 

and others, 1961), reemphasize a geochemical association worthy of further study. The tin-beryllium occurrence at Irish Creek, Rockbridge County, Va. (fig. 3.1), 

has been known for many years, although ho published 

Be analyses are. known to the authors. The geology 

and mineralogy of this deposit have been described by 

Koschmann and others (1942), and by Glass and others 

(1958).

A reconnaissance trip was made to the Irish Creek 

area on August 28,1962, to obtain samples for beryllium 

determination. 'Sampling was confined to the No. 1 

and the No. 2 workings, as denned by Koschmann and 

others (1942, p. 276).. The two adits at the No. 1 workings were found to be open although partly flooded. 

All underground entries at the No. 2 workings were 

caved, and so also were the banks of the various prospect pits, some to the point where the pits could hardly 

be recognized. Chip samples were taken across veins 

and across adjacent wallrock. Grab samples were taken 

across dumps in those places where banks of the prospect pits were caved and the veins no longer were exposed. Most of the samples from surface workings 

consisted of weathered granodiorite, greisen, and vein 

quartz. The underground samples from the two adits 

in the No. 1 workings were less weathered. Descriptions of the samples and the beryllium content recalculated to BeO are given in the accompanying table.

79° 10'

37° 

52' 

30"

L

Base modified from U.S. Geol. Survey 

Vesuvius quadrangle, Virginia

1 MILE

CONTOUR INTERVAL 250 FEET

FIGURE 3.1. Map showing the location of the No. 1 and No. 2 

workings of the Irish Creek mine, Rockbridge County, Va.

GEOLOGY

The Irish Creek deposits are in a hypersthene granodiorite (Koschmann and others, 1942, p. 227), which 

forms large areas of the Blue Ridge in central Virginia (Bloomer and Werner, 1955, p. 582). In the 

southern part of the mine area the granodiorite is in 

contact with older granitic gneisses and chloritic schists,

ART. 3 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B12-B15. 1963.

B12
LESURE, KIILSGAARD, BROWN, AND MROSE B13

and between the No. 1 and No. 2 workings, granodiorite is cut by younger basic dikes and aplite. The tin 

deposits occur in quartz veins in the granodiorite and 

in greisen zones that border the veins. The veins are 

irregular in trend and range in thickness from an inch 

to several feet. The greisen, which forms a narrow 

zone 1 inch to several feet thick between the quartz 

and unaltered granodiorite, consists of muscovite, 

quartz, and fluorite with smaller amounts of cassiterite, 

beryl, wolframite, and other minerals (Koschmann and 

others, 1942, p. 280). Numerous other greisen-bordered 

quartz veins have been reported in the Irish Creek area 

by Ferguson (1918, p. 13-14).

BERYLLIUM

Beryllium is present as beryl and phenacite in the 

quartz veins and associated greisen (Ferguson, 1918, 

p. 11; Glass and others, 1958, p. 78-79). Beryl, which 

is the more abundant, is generally at the edge of the 

quartz veins (Koschmann and others, 1942, p. 281). 

Minor amounts of phenacite have been found with 

beryl or cutting chlorite and fluorite (Glass and others, 

1958, p. 79). Beryl was the only visible beryllium mineral in the samples that we collected. The mineral is 

pale green, yellow, or white and forms slender hexagonal prisms that range" from 1 to 10 mm in length and 

0.5 to 5 mm across. Longer crystals have been reported by Ferguson (1918, p. 8) and Koschmann and 

others (1942, p. 281).

Only 1 sample from Irish Creek contains as much as 

0.67 percent BeO, but 14 samples range from 0.005 to 

0.044, and 14 from 0.0008 to 0.0047; only 4 samples 

have less than 0.0006 percent BeO (see table). The 

average beryllia content of intermediate igneous rocks 

including granodiorite, the host rock at Irish Creek, 

is 0.0007 percent (Warner and others, 1959, p. 22-23). 

A comparison of the Irish Creek samples with 96 samples from quartz-tungsten, quartz-gold, and other vein 

deposits from 21 mining districts in the United States 

(fig. 3.2) indicates only a slightly greater average BeO 

content for the Irish Creek deposits and in general a 

similar spread of values.

Recently, however, greisen zones around a granitic 

intrusive mass in Colorado were found to contain oregrade beryllium deposits (Sharp and Hawley, 1960). 

The higher than average beryllium content in the 

granodiorite at Irish Creek, the enrichment of the greisen zones, and the numerous quartz-tin veins in an area 

of 12 square miles along the creek (Koschmann and 

others, 1942, p. 272) suggest that additional prospecting 

is warranted.

FIGURE 3.2. A, BeO content of 33 samples of vein quartz, 

greisen, and granodiorite from Irish Creek, Va.; B, BeO content of 92 samples from 21 mining districts in the United 

States, from quartz-gold, quartz-tungsten, and other vein 

deposits that average more than 0.001 percent BeO (data 

from Warner and others, 1959, tables 14-16, 45, 56, 58, 66).
B14 GEOCHEMISTRY, MINERALOGY, AND PETROLOGY 

Beryllia in samples from Irish Creek, Rockbridge County, Va.

[Quantitative analysis by direct-reading spectrometer; analyses by Sol Herman and Harold Westley. Reported as percent Be, recalculated to percent BeO. Sample-location

descriptions refer to Koschmann and others, 1942, fig. 30, pi. 43, p. 276, 286]

No.

1

2

3

4

5

6

7

8

9

10

11

12

13

14

15

16

17

18

19

20

21

22

23

24

25

26

27

28

29

30

31

32

33

Sample

Type

Grab__---__-_-

_____do--_--__.

_..._do________

Chip __. _ ___

_____do____-__

__.__do___-____

_.___do-___-___

_____do_____._.

..._.do_. -_-___

Grab__---__-__

..__.do_-___-__

Chip_. _.----_.

__--_do-__-_-

_.___do___._-_.

-__._do____-__

Grab__-___--

_.-_.do_._.

Chip___.___ _

Grab-._-_- chiP _ __:___

.._-_do____

-_._do_____.

_____do________

 __do_____.

-_-.do_____.__

-_...do__.._

_ -_do__._

_____do____

-_-__do__-____

-___do-____.__

---.do________

 __do______._

Grab________

Length 

(feet)

15

q

1

o

4

10

1 O

7

7

6

q

5

o

14

1 P;

14 ^

4

4

1 C

4

5

o

5

3

2

Description

A^o. 2 workings 

Sheared granodiorite, relatively fresh, from dump of southwesternmost trench near

the 2,775-ft contour.

Gneiss, partly weathered, 115 ft S. 48° E. of southwesternmost trench, near the 2,800-

ft contour.

of southwestern shaft, near 2,750-ft contour.

Sheared granodiorite, weathered, on southeast side of sample 4_ ______ _ _ _ _ _____

24 and 34 ft from northwest end of the long trench near the 2,725-ft contour, westcentral part of workings.

Greisen, weathered, sample 9 area, 39 ft from west end of trench. _ _ _______ __

Granodiorite, mylonite, and greisen, from dump 36 ft northeast of old shaft in gulch,

near the 2,675-ft contour, west-central part of "workings. 

Highly altered granodiorite, small quartz veins and greisen, weathered, along south

wall of caved adit, where it crosses the 2,700-ft contour, near center of workings.

north side of northernmost shaft, elevation about 2,685 ft.

caved shaft, elevation about 2,715 ft, eastern part of workings.

No. 1 workings 

Weathered granodiorite from dump of 3 easternmost pits of workings, on the 2,750-ft

contour.

2,650-ft contour.

near east end of long cut, near the 2,575-ft contour. 

Greisen, 2 small quartz veins, 1 to 2 inches thick, partly weathered, at eastern end

of long cut, near the 2,525-ft contour. 

2 quartz veins, 2 to 3 inches thick, and greisen, on the south side and about 50 ft

from the eastern end of the long cut.

back of the adit, 8 ft from face of north fork of west adit. 

Across back of adit; greisen, small quartz veinlets }i to}£ inch wide, 39 ft west of fork

in west adit.

Silicified granodiorite in f ootwall of a shear zone at point where the west fork of west

adit turns northeast.

Greisen along 2-inch quartz vein, greisen zone }£ to 1^ inches on each side, visible

beryl in greisen, 2 ft from face of east adit.

Partly weathered greisen, vein quartz, and sheared granodiorite from dump of shaft

of workings.

BeO

(percent)

0. 0008

<. 0006

<. 0006

.0028

.0014

.0039

.0028

.0017

.0050

. 0033

. 0044

. 0047

. 0236

. 0033

.0022

<. 0006

.0011

.0008

<. 0006

.0141

.0050

.0067

.0111

. 0444

.0133

.0011

.0083

.0158

.67

. 0089

.0050

.033

. 0144
LESURE, KIILSGAARD, BROWN, AND MROSE B15

REFERENCES

Bloomer, R. O., and Werner, H. J., 1955, Geology of the Blue 

Ridge region in central Virginia: Geol. Soc. America Bull., 

v. 66, no. 5, p. 57&-606.

Ferguson, H. G., 1918, Tin deposits near Irish Creek, Virginia: 

Virginia Geol. Survey Bull. 15-A, 19 p.

Glass, J. J., Koschmann, A. H., and Vhay, J. S., 1958, Minerals 

of the cassiterite-bearing veins at Irish Creek, Virginia, and 

their paragenetic relations: Econ. Geology, v. 53, p. 65-84.

Koschmann, A. H., Glass, J. J., and Vhay, J. S., 1942, Tin deposits of Irish Creek, Virginia: U.S. Geol. Survey Bull. 

936-K, p. 271-296.

Sainsbury, C. L., Helz, A. W., Annell, C. S., and Westley, Harold, 

1961, Beryllium in stream sediments from the tin-tungsten 

provinces of the. Seward Peninsula, Alaska: Art. 151 w 

U.S. Geol. Survey Prof. Paper 424-C, p. C16-C17.

Sharp, W. N., and Hawley, C. C., 1960, Bertrandite-bearing 

greisen, a new beryllium ore, in the Lake George District, 

Colorado: Art. 35 in U.S. Geol. Survey Prof. Paper 400-B, 

p. B73-B74.

Warner, L. A., and others, 1959, Occurrence of nonpegmatite 

beryllium in the United States: U.S. Geol. Survey Prof. 

Paper 318, 198 p.
Article 4

BERYLLIUM AND FLUORINE IN MINERALIZED TUFF, 

SPOR MOUNTAIN, JUAB COUNTY, UTAH

By WALLACE R. GRIFFITTS and L. F. RADER, JR., 

Denver, Colo.

Abstract. The content of beryllium in a mineralized tuff 

bed varies directly with the fluorine content. Beryllium concentrations show a bimodal distribution with a large mode in 

the range 20 to 100 ppm and a smaller mode, representing-the 

ore zone, in the range 1,000 to 10,000 ppm.

Beryllium minerals are associated with minor 

amounts of fluorite in nonpegmatitic deposits of many 

districts (Norton and others, 1959). However, no published study of the quantitative relationship between 

beryllium and fluorine in mineralized rocks has been 

made. The Spor Mountain, Utah, area was found suitable for such a study because the west side of Spor 

Mountain has zones of beryllium-rich rhyolitic tuff that 

also contain fluorite (Staatz and Griffitts, 1961). The 

Vitro Minerals Co. kindly supplied drill-hole samples 

to supplement this study; adequate representation of 

the lower parts of the tuff bed would not have been 

possible without the samples.

The distribution of beryllium and fluorine in the 

mineralized tuff bed at Spor Mountain is shown in 

figure 4.1. The upper part of the tuff bed contains the 

greatest concentrations not only of beryllium and 

fluorine but also of iron, manganese, zinc, lead, tin, and 

rare earths. However, the zones of greatest concentration of these other elements coincide only approximately with the beryllium-rich zones. The abundance 

of most of the above elements is greater in the lower 

part of hole 1 than in the lower part of hole 2, the 

reverse of the trend noted for the abundance of beryllium and fluorine. The rank-correlation coefficient 

(Dixon and Massey, 1951, p. 261) calculated for beryllium and fluorine in 30 samples from each drill hole 

gives values of 0.636 and 0.656, respectively. The critical value for significance at the 1-percent level for 30

samples is listed as 0.432. Thus the correlation between 

beryllium and fluorine is significant to a very high 

degree.

The plot of the beryllium concentration against the 

fluorine content of the samples (fig. 4.2) shows that the 

concentrations of the two elements vary directly. For 

convenience the concentrations are grouped into ranges 

with midpoints in the series 1, 1.5, 2, 3, 5, and 7, and 

these midpoints are plotted in the diagram. Figure 4.2 

shows that most of the samples are grouped in 2 concentration ranges, namely, 1,000 ppm of Be or more, 

corresponding to the ore zone, and 100 ppm of Be or 

less, delineating a submarginal mineralization of widespread extent. The concentration range between 100 

and 1,000 ppm of Be contains relatively few samples, 

indicating a sharp boundary of the ore zone. Similar 

sharp boundaries between ore and nonore rocks have 

been observed in other beryllium districts.

The concentrations of the two elements must reach a 

maximum before either approaches 100 percent. Study 

of selected specimens suggests that the maximum beryllium content might be between 7 and 10 percent, associated with a fluorine content of about 30 percent.

REFERENCES

Dixon, W. J., and Massey, F. J., Jr., 1951, Introduction to statistical analysis: New York, MeGraw-Hill Book Co., Inc., 

370 p.

Norton, J. J., Griffitts, W. R., and Wilmarth, V. R., 1959, Geology and resources of beryllium in the United States: United 

Nations Internat. Oonf. Peaceful Uses Atomic Energy Proc., 

2d, Geneva 1955, p. 21-34.

Staatz, M. H., and Griffitts, W. R., 1961, Beryllium-bearing tuff 

in the Thomas Range, Juab County, Utah: Econ. Geology, 

v. 56, p. 941-950.

ART. 4 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B16-B17. 1963.

B16
GRIFFITTS AND RADER B17

100 200 FEET

BERYLLIUM, IN PARTS PER MILLION FLUORINE, IN PERCENT

>1000 100-1000 50-100 <50 >1.5 1-1.5 0.5-1 

FIGUBE 4.1. Distribution of beryllium and fluorine in the mineralized tuff bed at Spor Mountain.

<0.5

1-

z

UJ

o

£

UJ

0.

z

UJ

z

cr

O

u.

3

3

2

1.5

1

0.7

0.5

0.3,

0.2

0.15

n i

i

-

_

-

-

- 1

- 7

-8

- 1

-

i

1

2

1

8

6

1

1

i

1

2

3

2

7

3

i

l 1 1 I 1 1 I A l l i

,, -

2 11]

73 1

6

91 1

21

-

-

iii i i iii

15 20 30 50 70 100 150 300 500 1000 1500 2000 5000 

BERYLLIUM, IN PARTS PER MILLION

FIGURE 4.2. Relation between 'beryllium and fluorine in mineralized tuff at Spor Mountain. Numbers indicate number 

of analyses falling at each coordinate intersection.
Article 5

BERYLLIUM AND FLUORINE CONTENT OF SOME SILICIC VOLCANIC 

GLASSES FROM WESTERN UNITED STATES

By WALLACE R. GRIFFITTS and HOWARD A. POWERS, Denver, Colo.

Abstract. Analyses of more than 170 samples of silicic volcanic glass show that the Pearlette Ash Member of the Sappa 

Formation of Pleistocene age from the Great Plains contains 

the most beryllium and fluorine, that glass of Pliocene age from 

southern Idaho contains less than the Pearlette, and that glass 

of Pleistocene age from the Cascade Range contains the least 

beryllium and fluorine.

Beryllium and fluorine are commonly associated in 

nonpegmatitic mineralized rocks, and present data suggest a possible primary association in granitic rocks. 

Inasmuch as glassy volcanic rocks have been less affected by postmagmatic changes than granitic rocks, 

they may yield more information on the berylliumfluorine relations in magmas. An examination of more 

than 170 analyses of silicic volcanic glasses from Western United States shows that beryllium-rich glasses 

tend also to be rich in fluorine (fig. 5.1).

A marked difference in the composition of glasses of 

different provinces also shows very clearly in figures 

5.1, 5.2, and 5.3. Most of the samples that contain more 

than 0.10 percent fluorine are of Pliocene glass from 

southern Idaho or the Pearlette Ash Member of the 

Sappa Formation, which is found in Pleistocene strata 

over much of the Great Plains. The fluorine content 

of samples of the Pearlette averages 0.14 percent. Most 

of the samples that contain 0.10 percent or less of

LZ 

11

§10

_J

d 9

5

oc 8

LU

a.

w 7

oc ,.

< 6 

a.

? 5

J4

> 3

m 2

1

n

_

 x

 A X X

4 -

 A X X

2 7

 A A X

42 2 5

X X 

 » A AAX AX

2 3 2 46

 A A A X

2 32

x 

25547 23

 A» A A A A

565

92**

-

1 1 1 I 1 1 1 1 1 1 1 1 1

x

x

X A X

3

X X

4 2

X X

2

X A

3

X X

A i

1 1 1 1 1 1 1 1

0 .01 .02 .03 .04 .05 .06 .07 .08 .09 .10 .11 .12 .13 .14 .15 .16 .17 .18 .19 .20 .21 .22

FLUORINE, IN PERCENT

FIGURE 5.1. Relation of beryllium content to fluorine content of silicic glasses. Number indicates number of analyses represented by symbol, when more than one. Triangles indicate samples from 

southern Idaho (Pliocene) ; dots, from Cascade volcanoes or an unknown source (mostly Pleistocene) ; and crosses, Pearlette Ash Member of the Sappa Formation (Pleistocene).

B18

ART. 5 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B18-B19. 1963.
GRIFFITTS AND POWERS B19

30 

V) 

Ul

_l 

Q.

I 20

"0 .01 .02 .03 .04 .05 .06 .07 .08 .09 .10 .11 .12 .13 .14 .15 .16 .17 .18 .19 .20 .21 .22

FLUORINE, IN PERCENT

FIGURE 5.2. Frequency diagram for fluorine. Solid line, all samples; short-dashed line, samples from 

southern Idaho (Pliocene) ; and long-dashed line, Pearlette Ash Member of the Sappa Formation 

(Pleistocene).

fluorine are from the Cascade Eange, although some are 

from southern Idaho and one was of the Pearlette Ash 

Member. These samples average 0.06 percent fluorine.

401 

123456789 10 11 

BERYLLIUM, IN PARTS PER MILLION

FIGUBB 5.3 Frequency diagram for beryllium. Solid 

line, all samples; short-dashed line, samples from 

southern Idaho (Pliocene) ; and long-dashed line, Pearlette Ash Member of the Sappa Formation (Pleistocene).

The frequency diagram for fluorine (fig. 5.2) shows 

a markedly bimodal distribution of fluorine concentrations with a low mode corresponding to the samples from 

the Cascade Eange and a high mode corresponding to 

the Pearlette samples. The frequency plot for beryllium (fig. 5.3) shows a trimodal distribution. This, 

too, results from the combining of 2 statistical populations as the Cascade Range samples have a simple 

unimodal distribution and a mean of 4.4 parts per million of beryllium, and the Pearlette samples have a 

bimodal distribution and a mean of 7.4 ppm of 

beryllium.

The unimodal log-normal distribution for beryllium 

reported by Coats and others (1962) probably results 

from the inclusion of samples from many populations 

with different mean beryllium concentrations.

REFERENCE

Coats, R. R., Barnett, P. R., and Conklin, N. M., 1962, Distribution of beryllium in unaltered silicic volcanic rocks of 

the western conterminous United States: Econ. Geology, 

v. 57, p. 963-968.
Article 6

SOME NEW DATA ON THE ARSENIC CONTENT OF BASALT

By A. J. BARTEL, E. J. FENNELLY,

CLAUDE HUFFMAN, JR., and L. F. RADER, JR., Denver, Colo.

Abstract. Data are presented on the arsenic content of 48 

samples of basalt from 6 different areas. On the basis of this 

work and 66 determinations previously published by others, a 

median value of 1 ppm is indicated for the arsenic content of 

basalt and diabase.

Published data on the arsenic content of rocks are 

limited. Tremearne and Jacob (1941), in a paper on 

the arsenic content of phosphates, summarized most 

of the data published prior to 1940 on arsenic in rocks, 

minerals, coals, soils, sediments, and waters. More recently, Onishi and Sandell (1955) reported on the 

arsenic content of igneous rocks, minerals, chondrites, 

and shales.

Most common igneous rocks contain only trace 

amounts of arsenic. Values reported by Onishi and 

Sandell range from 1.5 ppm for granite to about 3.5 

ppm for silicic volcanic rocks, with 2.0 ppm as an 

average for basalt and diabase.

New data to supplement those available on the arsenic 

content of basalt have been obtained recently by the 

analysis of 48 basalts selected as representative of those 

studied by Eader and others (1963) for zinc content. 

The* analytical method was described by Eader and 

Grimaldi (1961) and has been used in this laboratory 

not only to determine the arsenic content of shales 

(Tourtelot, 1962) but also of many other samples. The 

standard deviation of the determinations on 12 basalts 

from Idaho analyzed by 2 different chemists is 0.35 

ppm 'of As for the range 2.0 ppm of As or less. A 

reagent-blank correction equivalent to 0.5 ppm of As 

has been made on all determinations.

The arsenic content of basalt based on 48 samples 

ranges from 0.1 to 6.0 ppm, with a weighted average 

of 1.0 ppm of As (table 6.1). Only 1 sample was found 

to contain more than 2 ppm of As. The preponderance 

of samples from Idaho and Oregon resulted from an 

effort to determine the range of arsenic in basalts

within a single province as compared to the total range 

found in basalt.

TABLE 6.1. New data on the arsenic content of basalt 

[Analysts, Ardith Bartel, E. J. Fennelly, and C. Huffman, Jr.]

Location

Connecticut, central trap rocks.. ___ _

Idaho, Snake River Plain. __ _______ _

New Mexico, Jemez Mountains. ______

Number of samples

4 

4 

5 

15 

4 

16

48

Arsenic content 

(ppm)

Range

0 1-0. 8 

8-1.4 

1-6.0 

3-1.0 

3-2.0 

1-2.0

. 1-6. 0

Average

0.3 

1. 1 

1.7 

.9 

1.0 

.8

1 1.0

i Weighted.

Table 6.2 (p. B22) presents complete standard rock 

analyses and determinations of 14 minor elements on 4 

samples from each of the 6 areas. These samples were 

selected to represent the total range of Variation among 

the samples within eac'h area in their content of SiO2 , 

MgO, CaO, K2O, and Na2O. Despite this careful selection based on composition, geologic age, and location, no 

correlation of arsenic with other elements, either major 

or minor, has been discovered. The greatest amounts 

of sulfur and zinc are present in the Idaho samples 

(table 6.2), and the most copper in the Hawaii samples. 

However, rank correlations of these elements to arsenic 

have shown no significance. The spectrographic determinations (Ga through Sr) have a coefficient of 

variation ranging from 5.4 percent for nickel to 20.3 

percent for strontium (Bastron and others, 1960, p. 

180).

The Weighted average value of 1.0 ppm of As in 

48 basalts from 6 areas of the United States included 

in our study differs slightly from that found by Onishi 

and Sandell (1955, tables 5, 10). Their data in the

B20

ART. 6 IN U.S. GBJOL. SURVEY PROF. PAPER 475-B, PAGES B20-D23. 1963.
BARTEL, FENNELLY, HUFFMAN, AND RADER B21

0 5 10 15 20 

NUMBER OF SAMPLES

FIGUBE 6.1. Frequency distribution of the arsenic content of 

basalt and diabase.

range 0 to 6.6 ppm of As (66 samples) have been plotted 

with our data td show the frequency distribution of 

arsenic (fig. 6.1). The samples are all from basalt 

areas of the United States except for 4 from Japan 

and 1 from Sicily. The plot of data from 114 samples 

gives a median value of 1.0 ppm of As as the content 

for basalt and diabase.

REFERENCES

Bastron, Harry, Barnett, P. R., and Murata, K. J., 1960, Method 

for the quantitative spectrochemical analysis of rocks, minerals, ores, and other materials by a powder D-C arc technique: U.S. Geol. Survey Bull. 1084-G.

Onishi, H., and Sandell, E. B., 1955, Geochemistry of arsenic: 

Geochim. et Cosmochim. Acta, v. 7, p. 1-33.

Rader, L. F., and Grimaldi, F. S., 1961, Chemical analyses for 

selected minor elements in Pierre shale: U.S. Geol. Survey 

Prof. Paper 391-A, 45 p.

Rader, L. F., Swadley, W. C., Huffman, Claude, and Lipp, H. H., 

1963, New chemical determinations of zinc in basalts and 

rocks of similar composition: Geochim. et Cosmochim. Acta. 

[In press]

Tourtelot, H. A., 1962, Preliminary investigation of the geologic 

setting and chemical composition of the Pierre shale, Great 

Plains region: U.S. Geol. Survey Prof. Paper 390, 74 p.

Tremearne, T. H., and Jacob, K. D., 1941, Arsenic in natural 

phosphates and phosphate fertilizers: U.S. Dept. Agriculture Tech. Bull. 781, 40 p.
B22 GEOCHEMISTRY, MINERALOGY, AND PETROLOGY

TABLE 6.2. Composition of basalts representative of sfa> diverse areas examined for arsenic content 

[Arsenic, sulfur, and zinc determined cbemlcally by A. Bartel, I. Frost, and C. Huflman; other minor elements determined spectrographlcally]

California, Mount Lassen area

1 2 3 4

Connecticut, central trap rocks

6 6 7 8

Hawaii, Island of Hawaii

9 10 11 12

Minor elements (parts per million)

As.. ..._._.___

S .

Zn__ ..._._.___

Ga

Sc---__._---.

Y_. ._...__._.

Co...... ._-_..

Cu...... ......

Ni....._......

Zr... .........

V__ __.__._--.

Cr__.__ .......

Ba... ........

Sr. ...____.__.

0. 1

<20

70

10

50

40

50

100

200

90

300

400

100

300

O Q

^9fj

75

18

41

36

34 inn

60 i«m

270

250

360

510

0.1

/OH

64

10

on

QfJ

30

70

100

200

200

300

700

500

0. 1

*">n

79

10

QfJ

OfJ

4.0

70

200

100

200

4.00

300

1,000

1.4

8O

78

1 3

40

9O

4.9

140

110

110

260

260

110

160

0

0

/9n

7Q

16

4.O

20

4.9

89

110

14.0

260

250

150

99O

1.0

QQfJ

96

1 5

70

Of)

30

69

QO

70

300

41

110

100

1.0

4.O

84

16

40

20

42

130

110

130

260

260

280

14.0

0.3

^f)C\

110

16

40

63

46

210

130

230

370

170

220

700

0. 1

60

100

20

40

60

50

200

100

200

300

400

130

700

1.7

<20

100

18

42

40

40

190

300

150

280

880

71

260

0.4

40

120

17

40

68

45

200

110

270

400

140

250

740

Major oxide composition (percent)

SiO2 -. _ ...

A12O3 . --._-.---

Fe2 O3 -------_-

FeO__ _.._....

MgO__ _...._.

CaO. ..__...._

Na20__ .......

K20._......._

H20+_________

H20-_..._.___

C02 ._. .......

Ti02 .. ..._....

P205 . _ -

MnO. ........

Total.......

48.93

17.63

.98

8.08

9. 18

10.67

2.81

. 23

.07

.02

.01

.99

.10

. 16

99.86

49.98

17.77

2.79

5.93

5.84

10.09

2.93

.64

1. 12

1.07

.08

1. 16

.36

. 14

on on

52.53

17.66

1 fiQ

5 8ft

6.33

8.16

3

CQ

1 4.8

94.

.48

.01

i 35

. 41

. 14

QQ Q3

53.76

1fi Q4.

2 89

3 77

o f\9

9 no

3 07

5Q

. 11

OQ

.01

87

90

. 11

QQ QO.

4.Q 88

14. 97

2 Q1

6

QQ

7.41

10.14

9 77

.48 1 90

2 O9

1 1Q

 10

.17

100. 07

50.47

14. 14

9 8°.

7 on

7.38

8 9.8

424

.38

9 7Q

4.4.

.09

1. 10

.15

. 17

99.85

51.67

13.88

3.60

8 Q1

5.61

9 81

9 4.Q

.52

.92

1. 14

.05

1. 11

19

.23

100. 06

51.78

14 44

2.01

8.21

7.63

10.52

1.88

.50

1.02

4.8

.06

. 17

10

. 19

100. 02

50. 55

13.84

2.98

9.32

6. 19

10. 10

2.61

.67

.01

.02

.01

3.31

.30

.18

100. 09

50.73

13. 74

1. 83

9.45

7.23

11.22

2.20

.46

.02

. 14

.01

2.62

.22

.17

100. 04

50.87

13. 12

5.93

5. 53

9.46

9.81

2.26

.30

.17

.07

.01

2. 17

.23

-. 17

100. 10

50.99

13.73

3.39

9. 36

5.42

9.38

2.75

.80

.01

.03

.01

3.57

.39

.18

100. 01

1. 2-D-1543. Basalt, Pliocene, Pleistocene. Near Hole in 

Ground Crater, junction of dirt road and State Highway 

44, Calif. Collected by G. A. Macdonald. Dorothy 

Powers and P.R. Barnett, analysts.

2. 4-D-1530. Basalt, Pliocene. Near Jenny Creek, north of 

Klamath River, Calif. Collected by G. A. Macdonald. 

Margaret Seerveld and P. R. Barnett, analysts.

3. 10-D-1246. Basalt, Pliocene or Pleistocene. Near Ash 

Pan Butte, State Highway 89, Manzanita Lake quadrangle, 

California. Collected by G. A. Macdonald. Dorothy 

Powers and P. R. Barnett, analysts.

4. ll-D-1236. Basalt, Pliocene or Pleistocene. Near Manzanita 

Lake, 1 mile north of Divide, State Highway 89, Calif. 

Collected by G. A. Macdonald. Dorothy Powers and P. R. 

Barnett, analysts.

5. ll-E-2277. Basalt, Late Triassic, Talcott Basalt. Near 

Sunset Rock Park, Conn. Collected by H. E. Simpson. 

June Goldsmith and Nancy Conklin, analysts.

6. 9-E-2271. Basalt, Late Triassic, Talcott Basalt. Near

Farmington, Conn. Collected by H. E. Simpson. June 

Goldsmith .and Nancy Conklin, analysts.

7. 23-E-2065. Basalt, Late Triassic, Holyoke Basalt. Near 

Berlin, Hartford County, Conn. Collected by C. E. 

Fritts. Dorothy Powers and J. C. Hamilton, analysts.

8. 17-E-2270. Basalt, Late Triassic. Dike near River Glen, 

Conn. Collected by H. E. Simpson. June Goldsmith 

and Nancy Conklin, analysts.

9. ll-B-409. Olivine basalt, eruption 3/25/55. Kilauea, Hawaii. Collected by G. A. Macdonald. Marjorie Balazs 

and P. R. Barnett, analysts.

10. 14-C-815. Basalt, Recent (Prehistoric). Kilauea, Hawaii. 

Collected by H. A. Powers. Dorothy Powers and P. R. 

Barnett, analysts.

11. 20-E-2146. Olivine basalt, Recent ''Prehistoric). Kahuku 

Pali, Hawaii. Collected by G. D. Fraser. June Goldsmith and J. C. Hamilton, analysts.

12. 22-B-410. Basalt, eruption 3/3/55. Kilauea, Hawaii. 

Collected by G. A. Macdonald. Marjorie Balazs and 

P. R. Barnett, analysts.
BARTEL, FENNELLY, HUFFMAN, AND RADER B23 

TABLE 6.2 Composition of basalts representative of six diverse areas examined for arsenic content Continued

Idaho, Snake Biver Plain

13 14 15 16

New Mexico, Jemez Mountains area

17 18 19 20

Oregon, Mount Hood area

21 22 23 24

Minor elements (parts per million)

As ---------

S__ .. -----._--

Zn._... ...... .

Ga.. ...... ...

Sc______-.___-

¥____---_._-- Co..........:.

Cu_. .. .......

Ni____ ....... .

ZT.... ........

V. ............

Cr_.._ --------

Ba__ ----..-

Sr

1.0

500

150

20

40

100

50

40

90

400

250

200

700

500

0.6

80

120

16

40

50

43

53

170

230

240

400

390

160

0.3

250

130

18

40

80

39

67

94

370

300

250

580

300

0.8

50

140

16

39

60

37

61

61

230

240

190

460

230

0.3

260

82

10

30

50

60

40

100

80

200

200

400

500

0.5

<20

86

10

30

70

40

100

100

200

200

200

400

500

1.2

<20

68

10

30

70

40

100

200

200

200

200

1,000

2 000

2.0

<20

100

30

70

40

50

80

100

200

200

400

800

900

0.7

<20

84

16

56

46

42

22

87

110

360

260

250

170

0.4

<20

110

16

58

52

41

26

81

100

360

270

260

190

0. 1

<20

71

13

36

36

42

46

200 -150

170

900

220

280

0.3

<20

70

14

30

36

28

56

50

190

200

170

290

310

Major oxide composition (percent)

SiO2

AhOa

Fe203 --_.----

FeO. .__......

MgO _ __----

CaO______----

Na20_____----

K2 0. ...__.. -

H20+______--

H20-___-_--

C02_ __----

Ti02 ._-______-

P205 __- ......

MnO_-_------

Total. ------

45.37

14. 11

2.98

11.98

6.74

10.05

2.24

.47

.77

.63

.02

3. 54

1. 11

.22

100. 23

45.57

15. 07

1.60

12.06

9.48

9.50

2. 57

.62

. 02

. 06

.00

2. 63

.56

. 17

99.91

45. 67

14.96

1.62

12.71

7.26

9.49

2.77

.89

. 02

. 18

. 01

3. 23

.99

.23

100. 03

47.06

14. 19

3. 18

11.36

6.59

9. 57

2. 54

. 64

.48

.38

. .02

3. 12

.60

. 21

99.94

46.48

16.00

3.79

6. 50

8.63

8. 17

2.55

. 65

2. 99

2. 16

.03

1. 36

. 30

. 17

99.78

48.48

16.37

3.91

6.62

6. 71

9.59

3.75

.78

1.06

. 18

.06

1.84

. 51

. 16

100. 02

49.98

16.45

5.88

5.22

6. 06

9. 52

3.32

.91

. 19

.22

.00

1. 64

.37

. 17

99.93

51.40

16. 33

7.06

1. 98

5.34

8.28

3.82

2.02

.57

.65

.01

1. 52

.68

. 15

99.81

47.47

16.01

4. 62

6.26

6. 67

10.30

2. 53

.48

1. 92

1. 72

.06

1.63

.27

. 16

100. 10

49.38

15. 39

4. 87

6.24

6.49

10. 55

2. 71

.49

1.48

. 59

.02

1.44

.22

. -18

100. 05

49.92

15.27

2.29

5.99

9.81

8. 59

3.09

.81

.86

1.30

.03

1.45

.30

. 14

99.85

50.70

16.96

3.77

4. 21

6. 16

8. 64

3. 52

.98

1.32

1.41

.04

1.67

.36

. 14

99.88

13. 9-A-196. Basalt, Pliocene. Tuana Gulch, Idaho. Col- 19. 

lected by H. A. Powers. Lois Trumbull and P. R. Barnett, 

analysts.

14. 6-D-1649. Olivine basalt, Recent. Near Wendell, Idaho.

Collected by H. A. Powers. Ruth Kittrell and Nancy 20. 

Conklin, analysts.

15. 8-D-1655. Olivine basalt, late Pleistocene. Near Dietrich, 

Idaho. Collected by H. A. Powers. Ruth Kittrell and 

Nancy Conklin, analysts. 21.

16. 23-D-1684. Basalt, middle Pleistocene. Clover Creek, 

Idaho. Collected by H. E. Malde. Paula Montalto and 

Nancy Conklin, analysts. 22.

17. 1-53-1758. Basalt, Pliocene. Rio del Oso Wash, Black 

Mesa quadrangle, New Mexico. Collected by C. S. Ross 

and R. L. Smith. Lucille Kehl and P. R. Barnett, analysts. 23.

18. 6-54-175. Basalt, Pliocene. Santa Clara Canyon, west of 

Ranger Station, Santa Clara quadrangle, New Mexico. 

Collected by C. S. Ross and R. L. Smith. Lois Trumbull 24. 

and P. R. Barnett, analysts.

8-A-64. Olivine basalt, Pliocene. Near Warm Springs, San 

Antonio Valley, Jemez Springs quadrangle, New Mexico. 

Collected by C. S. Ross and R. L. Smith. Jean Theobald 

and P. R. Barnett, analysts.

2-52-1567. Olivine basalt, Miocene or Pliocene. Near 

Boyd Ranch, Peralta Canyon, Jemez quadrangle, New 

Mexico. Collected by C. S. Ross and R. L. Smith. L. C. 

Peck and R. S. Harner, analysts.

2-D-1781. Basalt, Miocene. Columbia River Basalt, bottom flow, Johnnycake Mountain, Oreg. Collected by R. E. 

Wilcox. Ruth Kittrell and P. R. Barnett, analysts.

11-D-1781. Basalt, Miocene. Columbia River Basalt, 

middle flow, Johnnycake Mountain, Oreg. Collected by 

R. E. Wilcox. Ruth Kittrell and P. R. Barnett, analysts.

12-D-1776. Olivine basalt, Eocene, Clarno Formation.

Bottom flow, Cottonwood Creek, Oreg. Collected by R. E.

Wilcox. Ruth Kittrell and P. R. Barnett, analysts. 

16-E-1777. Olivine basalt, Eocene, Clarno Formation.

Middle flow, Cottonwood Creek, Oreg. Collected by R. E.

Wilcox. Ruth Kittrell and P. R. Barnett, analysts.
Article 7

PRELIMINARY RELATIONS IN THE SYSTEM Na2B4<VC«2B6On-H2O

By BRUCE B. HANSHAW, Washington, D.C.

Abstract. Few thermochemical data are available, but the 

symmetry of the equilibrium constants provides an approach 

for studying the system Na2B4O7-Ca2B8Oii-H2O. A semiquantitative phase diagram constructed by using this knowledge plus 

the record phases coexisting naturally in presumed equilibrium is presented. New relations among the sodium borates 

are suggested, and experimental methods for studying the three 

component system are outlined.

Few thermochemical data are available for the system Na2O B2O3 H2O and virtually none are available 

for the system CaO-B2O3-H2O. However, the composition of most of the minerals found in the borate 

deposits of Western United States falls in the ternary 

plane Na2B4 07-Ca2B6Oii-H2O (fig. 7.1) of the quaternary tetrahedron Na2O-CaO-B2O3-H2O. By making use of the similarity of the equilibrium-constant 

expressions for reactions between phases in this ternary 

system and by utilizing published information on the 

coexistence of naturally occurring phases, it is possible 

to construct a schematic three-dimensional phase diagram for the system.

In general, when the activity of water is decreased 

without temperature change the higher hydrates give 

up water and form lower hydrates. Increased temperature at constant activity of water has the same effect.

Christ and Garrels (1959) determined phase relations 

among the sodium borates. At 20 °C the dehydration of 

borax (Na2B4O7 -10H2O) to tincalconite (Na2B4O7 

 5H2O) occurs at a water-vapor pressure of 10 mm, 

which corresponds to a water activity of 0.58. Kernite 

(Na2B4O7 '4H2O) is not stable at 20°C but becomes 

stable at higher temperature. Tincalconite forms from 

borax in dry desert air; the reaction is readily reversed 

by increasing the activity of water. At the surface, tincalconite forms from kernite only after kernite has 

reverted to borax. Borax and tincalconite have similar 

crystal-chemical structures but the structure of kernite 

differs (Christ and Garrets, 1959). This structural difference helps to explain the apparently anomalous behavior of the hydrates.

Some of Menzel's data (in Christ and Garrels, 1959) 

for the sodium borates were used in this study. Using 

methods outlined by Latimer (1952, p. 359 and following pages), C. V. Guidotti, Harvard University (oral 

communication, 1960), calculated new entropy data for 

the soditim borates and from these he recalculated the 

free energy of reactions. On a graph of temperature 

versus activity of water, Guidotti's borax-tincalconite 

boundary is essentially in agreement with Menzel's. 

Guidotti's equilibrium curve for the pair borax-kernite, 

however, does not agree with Menzel's. According to 

Menzel, kernite should not exist as a stable phase over 

the range of temperature and water activity at or near 

the earth's surface. Because kernite apparently is a

H,O

Na2 B 4 O7 '10H2 Oj 

Borax

Ca 2 B6 Ou -13H20 

Inyoite *

7H2O 

Meyerhofferite

Na2 B4O7'5H2 O 

Tincalconite

Na 2 B4 O7 4H2O C

Kernite Colemarfite

FIGUEE 7.1. Naturally occurring phases in the system

ART. 7 IN U.S. GBOL. SURVEY PROP. PAPER 475-B, PAGES B24-B27. 1963.

B24
HANSHAW B25

stable phase at surface conditions as judged by its widespread occurrence, Guidotti's borax-kernite boundary 

was used in this study. Using his free-energy values, 

borax inverts, directly to kernite above 51 °C. As observed from natural occurrences (Christ and Garrels, 

1959) the conversion of borax to kernite occurs slowly, 

but, once formed, kernite persists metastably at lower 

temperatures where it may slowly revert to either borax 

or tincalconite.

Kurnakova and Nikolaev (1948) and later Garrels 

and Christ (written communication, April 29, 1957) 

pointed out the following relations between borax, 

ulexite, and colemanite.

Writing the reaction for equilibrium between borax 

and ulexite:

5Na2B407 .10H20+4Ca+2=4NaCaB609.8H2O

+6Na+1 +2H+1 +17H2O,

the equilibrium constant is

+18+12

[Ca+2]4

17 (1)

where the brackets denote activity. 

For equilibrium between ulexite and colemanite:

6NaCaB6CV8H20+4Ca+2=5Ca_B6011 -5H20

+6Na+1 +2H+1 +22H2O,

and the equilibrium constant is

rNa+1l 6 rH+1l2 If _ irNd, j [ji j

 (2)

The expressions for the two equilibrium constants 

have the same form and differ only in the power to 

which the activity of water is raised. If all the usually 

considered equilibrium reactions between the various 

minerals in the three-component system Na2B4Or 

Ca2B6On-H2O are written, the general form for all the 

equilibrium constants is

[Na+1] 6 [H+1] 2 

= [Ca+2] 4 (3)

where n is the number of moles of water for each 

reaction. The moles of water for each reaction are 

listed in the accompanying table. Thus, it is possible 

to plot phase relations in this system as functions of log 

[Na+1] 6 lH+1] 2/[Ca+2]4, log [H2O], and temperature. Two 

of the axes on the phase diagrams are in log units because 

such plots tend to give straight lines. Inverse temperature is used as the third axis in order to take advantage 

of a form of the van't Hoff equation which states that

/d log K\ _

V d(l/T) }P 2.3035

When this function is plotted over a limited temperature 

range, the resulting curve generally has a large radius of 

curvature. On a schematic diagram, using these axes, 

equilibrium surfaces may be approximated by planes. 

Before attempting to construct such a three-dimensional 

diagram, it is useful to consider the reported field 

occurrences of phases coexisting in presumed equilibrium.

Schaller (1916) reported colemanite, meyerhofferite, 

and inyoite from veins which crosscut lacustrine shales 

and sandstones at Mount Blanco, Calif. Ulexite and 

borax occur at the land surface. Inyoite probably resulted from the action of ground water on borax and 

ulexite, and was deposited in fissures as the solutions 

percolated downward. Inyoite is a primary mineral 

in some deposits (Nikolaev and others, 1938) and inverts to one of the lower hydrates upon exposure to 

desert air. Meyerhofferite is also stable underground 

(Schaller, 1916); colemanite forms when meyerhofferite is exposed to drier air. Christ and Garrels (1959) 

report that after a drift is opened in a borax deposit, 

tincalconite forms from borax. Meyerhofferite probably requires a higher activity of water than tincalconite because meyerhofferite is commonly observed 

in newly opened drifts prior to dehydration caused 

by exposure to air. Colemanite appears to be stable at 

the humidity and temperature at the land surface.

Moles of water in the stoichiometric reactions indicated

For stoichiometric reaction between indicated phases 

Ulexite-meyerhofferite. . _--_-------_--___

Ulexite-iny oite _______-_-_-_--_-__-----_-_-

Ulexite-colemanite__ _ __-_----------------_-

Tincalconite-ulexite__ _ _____________________

Tincalconite-probertite- .___-_----_--------_

Borax-ulexite. _______--_-____-__-------.---

Borax-probertite ___________________ ____-

Probertite-colemanite. _ ___--__-__---------_

Kernite-probertite. -_____--_____-----------

Number of 

molet of 

water, n

12

-18 

22

-8

4

17

29

4

-1

Schaller (1929) reported that in the Kramer district, 

California, crystals of kramerite (probertite) crosscut 

clay, kernite, and borax, and that tincalconite is always 

secondary, formed by dehydration of borax and by 

hydration of kernite. Christ and Garrels (1959) correlate the occurrence of kernite in the subsurface with 

a temperature of 58 0 ±:50C, depending on the depth. 

Therefore, by association, probertite is presumed to be 

the stable sodium-calcium borate hydrate near this 

temperature. Neither ulexite nor colemanite was reported to occur with kernite. However, the pair kernite-colemanite may be stable at high temperature and 

(or) low activity of water. In the mines at Kramer, 

tincalconite occurs as a thin film on kernite crystals.

Foshag (1931) reported probertite in veins crosscutting kernite and borax near Kyan, Calif., which indi-
B26 GEOCHEMISTRY, MINERALOGY, AND PETROLOGY

FIGURE 7.2. Schematic diagram of phase relations in the system Na2B4O7-Ca2B6Ou-H2O at 1 atmosphere pressure.

cates a probable secondary origin for the probertite. 

Probertite also occurs there with colemanite. Foshag 

suggests that probertite forms after ulexite as a result 

of burial; presumably this change is due to increased 

pressure and temperature.

By using a combination of physicochemical data and 

reported coexistence of natural phases, figure 7.2 was 

constructed to explain relations in the ternary system. 

Values on the vertical axis represent little more than 

a guess at the real numbers.

The boundary between kernite and tincalconite (fig. 

7.2) is in doubt. Eeported occurrences and the geometric requirements of the diagram suggest that kernite 

and tincalconite coexist in equilibrium under certain 

conditions, but the slope of the two-phase boundary 

curve is not known.

Experimental work is necessary to fix the positions 

of the phase boundaries. Two major problems to be 

overcome are: Measurement of the activity of Ca+2 and 

accurate determination of the activity of water. The 

latter is most critical because of very high exponents 

for the activity of water in the expressions for the 

equilibrium constants. C. L. Christ and A. H. Truesdell, U.S. Geological Survey (oral communication), 

have begun experimental work on this system; activity 

of the calcium ion may be determined by Ca-sensitive 

glass electrodes, and activity of water by an appropriate 

manometer, or an electrode sensitive to vapor pressure 

of water.

Additional data no doubt will alter the position of 

some of the phase boundaries on figure 7.2, but the general configuration of the diagram should not be radi-
HANSHAW B27

cally changed. With additional experimental data, 

minerals in this system will be unusually useful 

indicators of conditions in the geologic past.

REFERENCES

Christ, C. L., and Garrels, R. M., 1959, Relations among sodium 

borate hydrates at the Kramer deposit, Boron, California: 

Am. Jour. Sci., v. 257, p. 516-528.

Foshag, W. F., 1931, Probertite from Ryan, Inyo County, 

California: Am. Mineralogist, v. 16, p. 338-341.

Kurnakova, A. G., and Nikolaev, A. V., 1948, Isothermal

solubility of the system Na2O-CaO-B2O3-H2O at 25° C:

Akad. Nauk Izv., SSSR, Otdel. Khim. Nauk, no. 4, p.

377-382. 

Latimer, W. M., 1952, The oxidation states of the elements and

their potentials in aqueous solutions: Englewood Cliffs,

N.J., Prentice-Hall, Inc., 392 p. 

Nikolaev, A. V., and others, 1938, The synthesis of inyoite:

Doklady Akad. Nauk, U.R.S.S., v. 18, p. 431-432. 

Schaller, W. T., 1916 Inyoite and meyerhofferite, two new

calcium borates : U.S. Geol. Survey Bull. 610. 

 1929, Borate minerals from the Kramer district, Mohave

Desert, California: U.S. Geol. Survey Prof. Paper 158-1.

680231
Article 8

VARIATION IN MINOR-ELEMENT CONTENT OF DESERT VARNISH

By H. W. LAKH C. B. HUNT, D. F. DAVIDSON, and UTEANA ODA, 

Denver, Colo.

Abstract. The cobalt content of desert varnish from Death 

Valley and northeastern Nevada shows a close correlation with 

the manganese content of the desert varnish, and barium, 

lanthanum, molybdenum, nickel, lead, and yttrium show a general correlation. Thus the content of these elements in desert 

varnish may not be meaningful as an indicator of mineralized 

zones. Correlation with ore provinces may be represented by 

the high boron content of varnish from Death' Valley, Calif., 

by the high arsenic and antimony content of that from the 

Edna Mountains, Nev., and by the high copper content of that 

from the Antler Peak-Copper Basin area, Nevada.

Data on the trace-element content of desert varnish 

published by Engel and Sharp (1958) aroused interest 

in the possibility of using desert varnish as a geochemical regional-prospecting tool.

Samples from 6 sites in Death Valley, Calif., and 

from 6 sites in the northeastern quarter of Nevada were 

used to study the variation of the minor-element content of the varnish. Semiquantitative spectrographic 

analyses were made on three samples from each site: 

(1) the varnish, (2) the rock chips or pebbles remaining 

after removal of the varnish, and (3) the soil adjacent 

to the pebbles. The soil samples are of the 3- to 4-inchthick light-gray fine-grained very porous surface layer 

in which the varnished pebbles are embedded.

The varnish was removed from the pebbles by dissolution by the reducing action of a dilute solution of 

ammonium oxalate and oxalic acid (pH 4.0), activated 

by ultraviolet light (A. S. DeEndredy, Ghana Soil 

Survey, oral communication to H. W. Lakin, August 

1960). About 1 kilogram of varnished pebbles was 

placed in a 600-milliliter beaker, covered with the buffered oxalate solution, and the whole placed under an 

ultraviolet light for 24 hours. The oxalate solution 

was filtered, evaporated to dryness, and ignited in a 

furnace at 500°C. The weight of the oxides thus obtained ranged from 0.2 to 1.4 grams. Except for wet 

chemical determination of iron, manganese, and calcium

all other values were obtained by Semiquantitative 

spectrographic analysis.

In the accompanying table (p. B30-B31), the data are 

assembled in order of decreasing manganese content in 

the varnish. Varnish on the pebbles from Death Valley 

generally has a much higher manganese content than 

varnish found in northeastern Nevada; the iron content 

is fairly high in all the samples.

Engel and Sharp (1958, p. 489) in their study of 

desert varnish in the Mojave Desert and Death Valley 

state that "Field observations and some chemical data 

show that varnish is layered. Where the uppermost 

layer of black material has been removed, the underlying part is brown. Chemical analyses of successive 

leaches suggest a relative decrease in Mn inward from 

the surface." The low manganese content of all but 

one of the samples of varnish from northeastern Nevada 

may indicate that the high-manganese outer layer of 

varnish from Death Valley and the Mojave Desert is 

absent; possibly it was not formed or has been removed.

There is a striking positive correlation between the 

cobalt content and the manganese content of all the 

varnish samples. The cobalt content of the varnish 

ranges from 2,000 ppm in the sample with the highest 

manganese content to 30 ppm in that with the lowest 

manganese content; the pebbles and soil have a quite 

uniformly low cobalt content (<5-20 ppm). The 

amount of barium, lanthanum, molybdenum, nickel, 

lead, and yttrium in the varnish also generally decreases with decreasing amounts of manganese. Varnish with a low manganese content is enriched in 

beryllium. Because the variation in content of these 

elements appears to be dependent upon the amount of 

manganese in the varnish, the content of these elements 

in desert varnish is not likely to be a meaningful indicator of mineralized zones.

ART. 8 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B28-B31. 1963.

B28
LAKIN, HUNT, DAVIDSON, AND ODA B29

Nevertheless, some correlation with ore provinces 

may be represented by the data. Varnish (N-l, N-2) 

collected in a region whose rocks are presumed to have 

a high arsenic, antimony, and vanadium content contains anomalously large amounts of these elements. 

But varnish (N-3) collected from pebbles in an alluvial 

fan 3 miles from a similarly mineralized area does not 

contain anomalously large amounts of these elements. 

Varnish from site N-10, west of Eureka, Nev., contains

much silver, lead, zinc, and little copper; the ore mined 

upstream contains all four metals. Varnish (N-5, 

N-6) from a copper mining district has a relatively 

high copper content. The high boron content in the 

Death Valley varnish .correlates with high boron in the 

Death Valley area.

REFERENCE

Engel, C. G., and Sharp, R. P., 1958, Chemical data on desert 

varnish: Geol. Soc. America Bull., v. 69, no. 5, p. 487-518.
Partial analyses of desert varnish, varnish-free pebbles, and soil from 12 sites in California and Nevad

[Chemical analyses for iron, manganese, and calcium, by J. B. McHugh, H. L. Neiman, and J. H. Turner; semiquantitative spectrographic analyses for all other elem

Sample

Mn Fe Mg Ca Ti

Percent

Ag As B Ba Be Cd Co Cr Cu Ga La Mo Ni Pb Sb Sc Sn 

Parts per million

CALIFORNIA 

Death Valley; Death Valley Canyon fan

DV-2:

Soil...... 

28

.05

.07

11

1

3

3

.2

1.5

6.2

1.5

0.5

.2

.7

<1

<1

<1

< 1,000

<500

<500

1,000

20

30

> 10, 000

300

500

1

<1

2

<50

<20

<20

2,000

5

15

70

15

50

1,500

5

20

<20

5

15

500

<50

70

100

<2

2

1,000

7

30

1,000

<10

15

<200

<100

<100

50

<5

15

<1

<1

<1

Death Valley; Hell Gate fan

DV-6:

Soil......... -

28

.07

.07

22

.7

2

5

.3

2

1.5

10

0.2

.7

.3

<1

<1

<1

<1,000

<500

<500

700

30

50

> 10, 000

300

500

<1

1

2

<50

<20

<20

1,500

<5

10

100

20

50

1,000

7

20

<20

5

20

500

50

70

100

<2

<2

500

7

20

300

20

20

<200

<100

<100

100

<5

10

<1

<1

<1

Death Valley; Trail Canyon fan

DV-3:

Soil-.-.... ....

16

.05

.05

31

3

1.5

5

1

2

0.7

10

0.15

1

.5

<1

<1

<1

<1,000

<500

<500

1,500

30

30

> 10, 000

1,000

500

<1

2

2

<50

<20

<20

1,500

15

10

70

70

50

1,500

20

20

<20

20

15

200

100

70

100

<2

<2

700

30

20

500

15

20

<200

<100

<100

30

10

10

<1

<1

<1

Death Valley; Mustard Canyon hills

DV-5:

Soil. ..... ... ...

14

.03

.05

29

1.5

1.5

5

.2

1.5

0.7

10

0.3

1

.5

<1

<1

<1

<1,000

<500

<500

1,000

15

150

>10,000

500

700

1

<1

2

^fJC\

<20

<20

1,500

5

7

150

30

50

1,000

10

15

<20

10

15

Knn

<50

50

500

<2

2

500

7

20

500

10

15

<200

<100

<100

150

5

10

<1

<1

<i

Death Valley

DV-1:

Soil . ..

12

.15

.1

35

5

2

>5

.3

1.5

1.5

2

0.15

.5

.5

<1

<1

<1

<1,000

<500

<500

1,000

10

15

10, 000

2,000

2,000

<1

1

2

<50

<20

<20

1,000

5

7

50

<5

30

200

5

10

<20

20

20

<50

100

70

100

<2

<2

300

5

10

150

20

20

<200

<100

<100

20

<5

10

<1

<i

<1

NEVADA 

Humboldt County; fan on north side of Antler Peak

N-5:

Soil . .... 

6.4

.05

.07

30

3

5

3

1

1.5

0.5

1

3

0.15

.3

1

<1

<1

<1

< 1,000

<500

<500

150

20

70

7,000

1,000

1,000

1

1

2

<100

<50

<50

700

5

15

150

70

150

1,000

30

100

30

10

15

<50

50

100

100

<2

5

200

20

30

150

15

30

<200

<200

<200

<10

10

20

<2

<1

<1

CALIFORNIA 

Death Valley; Artists Drive area

DV-4:

Soil..... 

6

.07

.07

32

3

3

>5

1.5

2

3

7

.5

.7

<1

<1

<1

<1,000

<1,000

<1,000

700

30

50

1,000

1,000

<1

1

^Kft

<20

<20

700

10

15

Kfi

50

70

500

20

30

<20

15

20

i in

so

70

5

500

20

30

150

20

20

<200

<100

<100

50

10

15

<1

<1

<1

NEVADA 

Humboldt County; east flank of Edna Mountains; near quartz gold prospect

N-l:

Soil . ..... . -

3

.01

.07

18

1

5

1.5

1.5

1.5

1.5

3

0.3

.07

.5

<1

1

<1

7,000

<500

<500

100

20

70

7,000

9 fifiA

1,500

5

<1

3

<100

^ EA

<50

300

<5

15

200

30

70

700

30

50

20

5

20

<50

<50

70

20

<2

<2

300

20

30

150

<10

30

3,000

500

700

20

<5

15

<20

<10

<10
Partial analyses of desert varnish, varnish-free pebbles, and soil from 12 sites in California and Nevada Continu

[Chemical analyses for iron, manganese, and calcium, by J. B. McHugh, H. L. Neiman, and J. H. Turner; semiquantitative spectrographic analyses for all other eleme

Sample

Mn Fe Mg Ca Ti

Percent

Ag As B Ba Be Cd Co Or Cu Ga La Mo Ni Pb Sb Sc Sn 

Parts per million

Eureka County; wash west of Eureka

N-10:

Soil 

1.5

.7

.1

29

5

3

7

2

2

2.4

10

5

0.5

.3

.5

2

1

3

<1,000

<500

<500

50

20

30

1,000

1,000

1,000

2

2

2

<100

<50

<50

50

20

10

150

100

50

100

50

50

20

15

20

<50

<50

50

10

30

3

70

15

20

1,000

70

500

<200

<200

<200

30

15

10

50

<10

20

Humboldt County; small fan below outcrop of shale of Comus Formation, Edna Mountains

N-2:

T*pTVKlpQ

Soil 

0.96

.007

.07

24

1

3

2

.15

1.5

2.0

2

2

0.3

.15

.5

<1

2

<1

3,000

<500

<500

200

20

50

1,500

2,000

1,000

5

<1

3

100

<50

<50

100

<5

10

500

150

100

1,000

30

50

30

5

15

<50

<50

50

30

<2

2

500

30

70

70

<10

20

1,500

300

<200

20

5

15

<20

<10

<10

Humboldt County; fan from Osgood Mountains, 2.8 miles east of Getchell camp

N-3:

Soil .

0.52

.03

.07

23

5

5

2

1

1.5

0.7

3

3

0.3

.7

.5

<1

<1

<1

<1,000

<500

<500

50

70

70

1,000

3,000

1,000

 3

1

2

<100

<50

<50

50

I"

15

150

70

70

300

30

30

30

10

15

<50

50

50

30

5

2

150

30

30

50

10

15

<200

<200

<200

20

15

15

<20

<10

<10

Lander County; 50 yards downslope from Turquoise mine in Copper Basin

N-6:

Soil 

0.4

.01

.05

20

7

3

1.5

.3

1

0.6

.15

2

0.2

.5

.7

<1

1

<1

<1,000

<500

<500

150

15

50

700

700

1,000

2

3

2

<100

<50

<50

30

<5

10

200

70

70

1,500

300

50

30

15

20

<50

70

50

15

7

2

150

15

20

20

15

20

<200

<200

<200

10

10

15

<20

<10

<10
Article 9

RELATION OF ION-EXCHANGE CAPACITY

TO MINERAL COMPOSITION AND GRAIN SIZE OF WEATHERED CRYSTALLINE ROCKS

AT THE GEORGIA NUCLEAR LABORATORY, DAWSON COUNTY, GEORGIA

By J. W. STEWART, Lakeland, Fla.

Work done in cooperation with U.S. Atomic Energy Commission and the U.S. Air Force

Abstract. Samples of weathered crystalline rocks at the 

Georgia Nuclear Laboratory were found to have ion-exchange 

capacities ranging from 1.0 to 13.7 milliequivalents per 100 

grams. In general, the higher ion-exchange capacities characterized the samples having the higher clay content. The 

principal minerals were kaolinite, muscovite, quartz, vermiculite-biotite, gibbsite, and vermiculite.

The relation of ion-exchange capacity to the mineral 

composition and grain size of weathered crystalline 

rocks at the Georgia Nuclear Laboratory was determined from samples collected at various depths below 

land surface from the side of a borrow pit and during 

the drilling of 33 monitoring wells. The borrow pit 

and 3 of the wells are near the REL (Radiation Effects 

Laboratory) waste-disposal pit, and the 30 other wells 

are near REF (Radiation Effects Facility) wastedisposal pits 1 and 2. The samples from the borrow 

pit were collected at depths of 1 to 27 feet below land 

surface and those from the wells at depths of 0 to 

18 feet. Laboratory tests of the samples were made 

by the U.S. Geological Survey, Denver, Colo.

The REL waste-disposal pit is in the southeastern 

corner of the Georgia Nuclear Laboratory area and is 

1.2 miles southeast of REF pit 1. The altitude of the 

land surface at the REL pit is about 1,112 feet. 

Quartz-mica schist, garnet-biotite schist, and amphibolite schist underlie the pit, and a small pegmatite vein 

composed largely of quartz and muscovite is exposed in 

its northern corner. The foliation at the pit strikes 

northeast and dips about 73° SE.

The saprolite at the REL site is thicker and more 

deeply weathered than at the REF site, largely because 

the fine-grained rocks are more intensely sheared than 

at the REF site. In addition, erosion probably has

removed more saprolite at the REF site where the topographic relief is about 200 feet, than it has at the REL 

site where the topographic relief is about 100 feet. 

Test drilling near the REL pit revealed that the saprolite is about 137 feet thick, or about 11 times thicker 

than that at REF pit 1 and about 2% times thicker 

than that at REF pit 2. At both the REL and REF 

sites the schistose structure is preserved in the highly 

weathered saprolite beneath the soil zone.

The samples of saprolite were dispersed in distilled 

water with sodium metaphosphate added as a dispersing 

agent and fractionated by repeated centrifuging and 

decanting to separate the silt (2-02 microns) and the 

clay (<2 microns) fractions. The principal minerals 

in the silt and clay were identified from the X-ray diffraction patterns of each fraction, and estimates of their 

relative abundance were expressed as parts in 10.

In order of relative abundance, the principal minerals 

in the 11 samples of saprolite from the borrow pit were 

kaolinite, gibbsite, hematite, quartz, muscovite, goethite, 

and vermiculite. By far the most abundant mineral 

was kaolinite, which averaged about seven-tenths of the 

silt fraction and about four-tenths of the clay fraction. 

Hematite, absent in the silt fraction, was the second 

most abundant mineral in the clay fraction; it probably 

resulted from oxidation and hydration of iron-bearing 

minerals such as muscovite and goethite. Vermiculite 

was absent in the silt fraction but was present in most 

of the clay fraction; presumably it resulted from alteration of some muscovite by hydration and leaching.

The ion-exchange capacity of the samples of saprolite 

from the borrow pit ranged from 1 to 5 meq per 100 g 

(milliequivalents per 100 grams). The larger ionexchange capacities characterized the samples collected

ART. 9 IN U.S. GEOL. SURVEY PROP. PAPER 475-B, PAGES B32-B34. 1963.

B32
STEWART B33

from the side of the pit at depths of 9.4 to 15.0 feet 

below the land surface.

Because the values for these samples were so much 

smaller than those for the samples collected during the 

drilling of the nearby wells, they are not considered 

truly representative. It is believed that the 4-year 

exposure of the saprolite to heavy rains has resulted 

in the removal of much of the clay and in the consequent reduction of the ion-exchange capacity of the 

material still in place. Washing out of clay under such 

conditions has been observed in more recently excavated 

pits in the Georgia Nuclear Laboratory area.

Nineteen samples of saprolite from depths of 0 to 

16 feet were collected during the drilling of the 8 wells 

near the EEL waste-disposal pit. In samples collected 

at depths of less than 4 feet below the land surface, the 

average silt content was about 29 percent by weight 

and average clay content about 40 percent (fig. 9.1). 

In samples from the depth interval 4 to 16 feet, the average silt content was 40 percent by weight and the clay 

content 10 percent. The silt fraction from a depth of 

0 to 4 feet contained, in order of abundance, kaolinite, 

quartz, muscovite, goethite, and gibbsite, and from a 

depth of 4 to 16 feet, kaolinite, muscovite, and quartz. 

Similarly, the clay fraction from a depth of 0 to 4 feet 

contained kaolinite, vermiculite, goethite, gibbsite, muscovite, and quartz, and from a depth of 4 to 16 feet, 

kaolinite, muscovite, and goethite.

Kaolinite, the predominant mineral in the saprolite, 

averaged 6 to 7 parts in 10 of the silt fraction and 7 

parts in 10 of the clay fraction. Muscovite, the second 

most abundant mineral, averaged 1 to 2 parts in 10 of 

both the clay and silt fractions. Vermiculite, absent in 

the silt samples, was present in several of the clay fractions; on the other hand, feldspar was present in several silt fractions but was absent in the clay fractions. 

Quartz constituted about 1 to 2 parts in 10 of the silt 

fractions but occurred only as a minor constituent of the 

clay fractions.

The ion-exchange capacity of the samples collected 

during the drilling of the REL wells ranged from 2.8 

to 8.1 meq per 100 g. The samples from depths of less 

than 4 feet had an ion-exchange capacity greater than 

4.5, those from depths of 5 to 10 feet from 4.1 to 6.5, 

and those from 11 to 16 feet less than 4.9. In general, 

the larger values of ion-exchange capacity were obtained 

for the most highly weathered saprolite that had the 

greatest percentage of clay by weight.

Waste-disposal pits REF-1 and REF-2 are in the 

north-central part of the Georgia Nuclear Laboratory 

area. REF pit 1 is near the crest of a steep hill about 

600 feet northeast of REF pit 2, which is on top of the 

highest hill in the REF site. The altitude of the land

Gi 0

Silt Clay

(29 percent) (40 percent)

Depth 0-4 feet

Silt Clay

(40 percent) (10 percent)

Depth 4-16 feet

REL SITE "~ ' ~ ~ "

Clay

(18 percent) 

I8 feet

REF SITE

FIGURE 9.1. Average mineral composition of saprolite from wells 

at the REL and REF sites. Gi, gibbsite; G, goethite; K, 

kaolinite ; M, muscovite ; Q, quartz ; V, vermiculite ; VB, 

vermiculite-biotite ; O, other minerals.

surface at REF pit 1 is 1,172 feet and at REF pit 2 

is 1,245 feet.

Quartz-biotite schist, biotite-garnet schist, and 

garnet-biotite-plagioclase schist underlie both these 

pits. Quartz veins ranging from y8 of an inch to 18 

inches thick are exposed on the west side of pit 1. The 

foliation in the area strikes northeast and dips about 

84° SE. at REF pit 1 and about 68° SE. at REF pit 2. 

The saprolite is 55 feet thick in the vicinity of REF 

pit 2, from which point it thins in all directions; at 

REF pit 1 it is about 12 feet thick.

In the vicinity of the REF pits, 97 samples of saprolite from depths of 0 to 18 feet were collected during 

the drilling of the 30 wells. Samples from depths of 

less than 5 feet below the land surface averaged about 

26 percent silt and 30 percent clay by weight and from 

depths of 5 to 18 feet averaged 35 percent silt and 18 

percent clay. In order of abundance, the silt fraction 

from a depth of 0 to 5 feet contained quartz, vermiculite-biotite (hydrobiotite), muscovite, and kaolinite, 

and from a depth of 5 to 18 feet contained vermiculitebiotite, muscovite, kaolinite, and quartz. Similarly,
B34 GEOCHEMISTRY, MINERALOGY, AND PETROLOGY

the clay fraction from a depth of 0 to 5 feet contained 

kaolinite, gibbsite, goethite, muscovite, vermiculite, 

and vermiculite-biotite, and from a depth of 5 to 18 

feet contained kaolinite, muscovite, gibbsite, goethite, 

and vermiculite-biotite. (See fig. 9.1.)

Vermiculite-biotite, which was not reported in 

samples from the BEL site, averaged about 3 parts 

in 10 of the silt fraction in the samples from the KEF 

site. Kaolinite, muscovite, and quartz each averaged 

about 2 parts in 10 of the silt fractions. The average 

contents of principal minerals in the clay fractions were 

kaolinite, 6 parts in 10; gibbsite, 1 part in 10; muscovite, 

1 part in 10; goethite, 1 part in 10; and vermiculitebiotite, less than 1 part in 10. Gibbsite was present 

only as a trace hi the silt fractions, and quartz was 

present only as a trace in the clay fractions. Several 

of the silt and clay fractions showed traces of chlorite 

and lepidocrocite.

The ion-exchange capacity of the samples from the 

REF site ranged from 2.5 to 13.7 meq per 100 g. Most 

of the ion-exchange values greater than 4.6 were for 

samples from depths of 5 feet or less. With few exceptions these samples contained the highest percentage of 

clay by weight, indicating again that the ion-exchange 

capacity tends to be greater if the average particle size 

is smaller. The ion-exchange capacity of the samples 

from depths of 6 to 10 feet ranged from 2.7 to 8.9 meq 

per 100 g, and from depths of 11 to 18 feet ranged 

from 2.5 to 8.0 meq per 100 g. In both these depth

ranges, the larger ion-exchange capacities characterized 

those samples having the greater clay content.

The principal minerals in the samples of saprolite 

collected during the drilling of all 33 wells consisted of 

the following, in order of relative abundance: kaolinite, 

muscovite, quartz, vermiculite-biotite, gibbsite, vermiculite, goethite, and feldspar. At both the REL and 

KEF sites, kaolinite was the most abundant mineral, 

constituting as much as 5 to 8 parts in 10 of some 

samples. The range in ion-exchange capacity at the 

REL site (2.8* to 8.1 meq per 100 g) was slightly less 

than that at the REF site (2.5 to 13.7 meq per 100 g), 

largely because the saprolite at the REF site contained 

more clay and vermiculite-biotite, vermiculite, and 

gibbsite. At both sites the largest values of ionexchange capacity characterized the samples from 

depths of less than 5 feet. The shallow samples generally contained more clay than the samples from 

greater depths. Although the ion-exchange capacity 

tended to decrease with depth, the decrease was not uniform nor smoothly progressive.

The generally low ion-exchange capacity of samples 

from both sites indicates that the saprolite has a relatively low capacity for base exchange. Nevertheless, 

the overall efficiency of the saprolite for retention of 

radionuclides probably is fairly high because tlie saprolite is thick, the water in the saprolite dilutes the infiltrating radioactive wastes, and the rate of lateral 

movement of water within the saprolite is very slow.
Article 10

CORDIERITE-BEARING MINERAL ASSEMBLAGES IN PRECAMBRIAN ROCKS, 

CENTRAL CITY QUADRANGLE, COLORADO

By P. K. SIMS and D. J. GABLE, Minneapolis, Minn., and Denver, Colo.

Abstract. Cordierite occurs with anthophyllite in small 

lenses within a microcline gneiss unit and in 'biotite-cordieritegarnet-sillimanite gneiss within a biotitic gneiss unit of the 

Idaho Springs Formation. The cordierite of both occurrences 

evidently originated by dynamothermal metamorphism of sedimentary rocks, with little addition of material except possibly 

magnesium.

Cordierite is common in the Precambrian metamor-" 

phic rocks of the Idaho Springs Formation in the central part of the Front Range. It occurs with anthophyllite in small bodies within a microcline-bearing 

paragneiss of granitic composition near Central City, 

and with sillimanite and garnet in biotite gneiss in adjacent areas. The rocks are part of the succession of 

high-grade metasedimentary gneisses that predate the 

Precambrian intrusive igneous rocks of the Front 

Range (Moench and others, 1962, p. 37-39; Lovering 

and Goddard, 1950, p. 19-29).

TABLE 10.1. Representative modes (volume percent) of 

cordierite-anthophyllite rocks and associated rock types

Biotite.........:..............

Garnet. _____ . ___ ......

Sillimanite.. _ .... .... ...

Apatite.......... .............

Chlorite.. ....__. .............

Epidote. __ ..............

Total.. 

Average grain diameter, in

Cordierite-anthophyllite 

rocks

1

3.8 

90.0

6. i

5.4

0.4 

0.3

100.0

2 

44.4

29.6 

2.2

0.7 

8.8 

13.9

Tr.

Tr.

0.1

0.3

100.0

0.5

3

51.5

27.1 

10.4

0.1 

8.0

Tr.

2.7

0.2

100.0

0.5

4

.12.7 

14.5 

29.6 

1.3

Tr. 

1.9

Tr.

Tr.

Tr.

100.0

Anze

Hornblendetnd cummingtonite-bearing 

rocks

5

20.9 

16.7

47.2

2.8

1.7

0.1

4.3

1.2

5.1

100.0

AJ178-80

0.3

6

8.7 

.H.I

30.2

2.0

0.1

0.9

2.6

0.4

1.0

100.0

Anei-M 

0.3

Cordicritebiotite 

rocks

7

59.4 

1.8 

21.2

17.1 

0.3

0.2

100.0

0.5

Cordierite-anthophyllite rocks occur in scattered 

lenses within the microcline gneiss. The lenses are a 

maximum of 300 feet long, elongate in the plane of the 

foliation, and sharply delineated from the enclosing 

gneiss. At places they are adjacent to thin layers of 

amphibolite or a calcic hornblende gneiss. In most 

bodies the cordierite-anthophyllite rocks are interlayered with hornblende and hornblende-cummingtonite 

rocks, generally on a scale of a few feet. One lens is 

associated with cordierite-biotite rocks. Cordierite is 

not found with either hornblende or cummingtonite. 

Modes that are representative of the rock types, and indicate their extreme variation in composition, are given 

in table 10.1. The texture of all the rocks is granoblastic. Garnet and cordierite tend to form poikiloblasts in a mosaic of slightly smaller mineral grains.

The cordierite-bearing rocks contain the following 

mineral assemblages:

Cordierite-anthophyllite rocks

Anthophyllite-biotite-cordierite - magnetite - quartz 

Anthophyllite-biotite-cordierite-garnet - magnetitequartz 

Anthophyllite-biotite-cordierite-magnetite - plagioclase-quartz 

Anthophyllite-biotite - cordierite - garnet - magnetite-plagioclase-spinel 

Anthophyllite-biotite-cordierite-magnetite-spinel

Cordierite-biotite rocks

Biotite-cordierite-quartz 

Biotite-cordierite-garnet-plagioclase-quartz

Green spinel and less commonly corundum are present 

locally in some thin sections of cordierite-anthophyllite rock that contain quartz.

The cordierite-bearing biotite gneiss is a distinct unit 

within the Idaho Springs Formation. It forms disART. 10 IN U.S. GBOL. SURVEY PROP. PAPER 475-B, PAGES B35-B37. 1963.

B35
B36 GEOCHEMISTRY, MINERALOGY, AND PETROLOGY

continuous layers that conform generally to the gross 

layering in the biotite gneiss units. One body, a maximum of about 1,000 feet thick, can be traced continuously along strike for 2i£ miles; others are more 

lenticular. Most bodies lie within larger masses of sillimanitic biotite-quartz gneiss, and grade sharply into 

it. Microcline-rich pegmatites are associated with the 

bodies commonly as lenses or pods lying parallel to the 

foliation in the gneiss. Lenses of amphibolite, calcsilicate gneiss, granodiorite, and hypersthene-bearing 

quartz diorite are found near many bodies of cordieritebearing gneisses.

The cordierite-bearing biotite gneiss consists of the 

following general mineral assemblages:

Biotite-cordierite-potassic feldspar-sillimanite

Biotite-cordierite-garnet

Biotite-cordierite-garnet-sillimanite

Biotite-cordierite-potassic feldspar-garnet-sillimanite

Each of the assemblages contains magnetite, quartz, 

and plagioclase; spinel is a local mineral (table 10.2). 

The rocks have a granoblastic texture and except for 

the development of porphyroblasts of garnet, cordierite 

and, rarely, magnetite, the minerals are intergrown in 

a mosaic pattern. In a few sections, cordierite was observed to embay garnet and to surround it in welldefined coronas free of sillimanite and biotite. This 

suggests that the cordierite developed at the expense of 

garnet, probably by the reaction

garnet+biotite+sillimanite+quartz ^cordierite 

+magnetite+potassic feldspar+H2O 

suggested previously by Schreyer and Yoder (1961, p. 

150). Magnetite and potassic feldspar developed in this 

way tend to cluster adjacent to the cordierite. Mineral 

phases that were not consumed in the reaction were enclosed poikilitically in the cordierite.

TABLE 10.2. Representative modes (volume percent) of 

cordierite-bearing sillimanite-'biotite gneiss

Cordierite __________________

Sillimanite ............... ............

Garnet. ___________________

Biotite ________ .... _ . __ ------

Muscovite- ________________

Clinozoisite ________ . __ .... ....

Spinel- ______ ....................

Chlorite . .............................

Total..... ........ ............

Composition of plagioclase- _ ..........

1

3.9

2.2

30.4

26.7

8.2

3.9

22.0

2.7

Tr.

100.0

0.7

2

14.2

2.2

24.5

19.2

7.3

3.6

25.0

3.9

Tr.

0.1

Tr.

Tr.

100.0

0.3

3

Tr.

12.1

27.3

5.9

0.1

33.1

15.2

6.3

Tr.

Tr.

Tr.

100.0

0.3

4

7.7

3.9

26.7

11.7

3.2

17.4

27.7

1.4

Tr.

Tr.

0.2

0.1

100.0

5

15.3

37.2

31.7

1.5

Tr.

8.3

6.0

Tr.

Tr.

Tr.

100.0

0.3

The cordierite-bearing rocks were produced by dynamothermal metamorphism, accompanying the plastic 

deformation of the region (Moench and others, 1962, p. 

39-45). The mineral assemblages are characteristic of 

rocks that span the boundary of the amphibolitegranulite metamorphic facies of Eskola (1952, p. 162- 

169).

The assemblages of the cordierite-bearing biotite 

gneisses developed by recrystallization of original 

argillaceous sediments, probably with little addition or 

subtraction of materials. The cordierite apparently 

marks layers that contained more magnesia that those 

that yielded the associated sillimanitic biotite-quartz 

gneiss. Rocks containing similar assemblages are found 

in many other regions, as noted by Schreyer and Yoder 

(1961) and Barker (1962).

The origin of the cordierite-anthophyllite and associated rocks is less certain. .These rocks are included 

within a rock unit of presumed metasedimentary origin, 

and have a fine-scale lithologic layering that resembles 

the layering in other metasedimentary rocks of the area. 

The cordierite-anthophyllite rocks have abnormally 

high percentages of both ferrous iron and magnesia 

and abnormally low amounts of lime, soda, and potassium compared to known sedimentary or igneous rocks. 

The associated hornblende- and cummingtonite-bearing 

rocks also have a high content of ferrous iron and magnesia, but contain substantial lime (6.53 percent CaO 

in one sample) in contrast to the cordierite-anthophyllite rocks. Judging from the -modes, the cordieriteanthophyllite gneiss is similar in composition to the 

cordierite-biotite gneiss except for having less potassium. Possibly the rocks were derived from original 

iron-rich sediments by the metasomatic addition of 

magnesia; primary sediments of this type contain 

negligible alkalies. A metasomatic origin has been 

proposed for similar rocks from other regions by Eskola 

(1914, p. 254), Bugge (1943, p. 103), Tilley (1937), 

and others.

REFERENCES

Barker, Fred, 1962, Cordierite-garnet gneiss and associated

microcline-rich pegmatite at Sturbridge, Massachusetts,

and Union, Connecticut: Am. Mineralogist, v. 47, p. 907-918. 

Bugge. J. A. W., 1943, Geological and petrological investigations

in the Konggberg-Bamble formation : Norges Geol. Und. no.

160, 150 p. 

Eskola, Pentti, 1914, On the petrology of the Orijarvi region in

southwestern Finland: Comm. Geol. Finlande Bull., v. 40,

277 p. 

 1952, On the granulites of Lapland : Am Jour. Sci., Bowen

vol., p. 133-171.
SIMS AND GABLE B37

Lovering, T. S., and Goddard, E. N., 1950, Geology and ore 

deposits of the Front Range, Colorado: U.S. Geol. Survey 

Prof. Paper 223, 319 p.

Moench, R. H., Harrison, J. E., and Sims, P. K., 1962, Precambrian folding in the Idaho Springs-Central City area, 

Front Range, Colorado: Geol. Soc. America Bull., v. 73, 

p. 35-58.

Schreyer, W., and Yoder, H. S., Jr., 1961, Tetrographic guides 

to the experimental petrology of cordierite, in Annual 

report of the Director of the Geophysical Laboratory, 1960- 

1961: Carnegie Inst. Washington Year Book 60, p. 147-152.

Tilley, C. E., 1937, Anthophyllite-cordierite granulites of the 

Lizard: Geol. Mag., v. 74, p. 300-309.
Article 11

BEDDED BARITE DEPOSITS OF THE SHOSHONE RANGE, NEVADA

By KEITH B. KETNER, Denver, Colo.

Work done in cooperation with the Nevada Bureau of Mines

Abstract. Barite is interbedded with chert and minor 

amounts of limestone in Paleozoic bedded chert units. Evidence 

of replacement is widespread. Neighboring sulflde deposits are 

normal mesothermal deposits in altered wallrocks, are related 

to intrusive rocks, and are discordant with bedding. Barite deposits are monomineralic, occur in unaltered host rock, are related to bedded chert rather than to intrusive rocks, and tend 

to be concordant. They are probably epigenetic replacements 

of chert and limestone but are genetically unrelated to neighbor-,, 

ing sulflde deposits.

In the United States, barite is obtained mainly from 

three distinct types of deposits: (1) fracture fillings 

and replacements in carbonate rocks, (2) residual mantles on barite-rich carbonate rocks, and (3) beds in siliceous sedimentary rocks. Vein and mantle deposits are 

well known. This study, based on five deposits in the 

Shoshone Range, Nov., records some of the important 

features of the less well known bedded deposits. The 

deposits were briefly described previously by Gianella 

(1941, p. 294) and Brobst (1958, p. 96).

The Greystone, Bateman Canyon, and Slaven Canyon 

deposits (fig. 11.1) are in the Slaven Chert of Devonian 

age. The Mountain Springs deposit is in a chert unit 

of a formation mapped originally as the Pumpernickel 

Formation of Carboniferous age (Ferguson and others, 

1951); however, recent discoveries of fossils show that 

the host rocks there are of Devonian age. The Argenta 

deposit and others farther northeast are in chert formations similar to the Slaven in lithology but of unproved 

age. Small deposits are also found in bedded chert of 

Ordovician age (James Gilluly, oral communication, 

1959).

The host rock consists almost entirely of bedded 

chert, but includes subordinate limestone and argillite. 

Shaly laminae at 1- to 3-inch intervals cause the chert 

to break up into platy fragments. In many beds concentrations of organic matter, clastic particles, or radiolaria form nonfissile layers about a millimeter thick.

R. 44 E. R. 46 E. R. 47 E.

6 MILES

EXPLANATION

X 

Bedded barite deposit Metallic sulfide deposit

FIGTJBE 11.1. Outline of northern part of Shoshone Range 

showing principal bedded barite deposits and metallic sulfide deposits.

Stylolites nearly parallel to bedding are common. 

Typically the chert is dark gray to black, less commonly 

it is greenish gray.

The chert beds are composed largely of quartz and 

chalcedony with minor amounts of pyrite in various 

stages of oxidation, carbonaceous matter, illite ( ?), and 

traces of other minerals too small to be identified. Most

ART. 11 IN U.S. GEOL. SURVEY PROP. PAPER 475-B, PAGES B38-B41. 1963.

B38
KETNER B39

grains of quartz and chalcedony large enough to be seen 

clearly with the microscope have interpenetrating contacts. The grain size ranges from about 0.002 to 0.5 

mm in diameter. Some of the large grains are clearly 

clastic particles, distinguishable from the rest by their 

regular outlines and concentration in beds; others have 

resulted from secondary growth along fractures. 

Radiolaria are abundant in many beds.

Limestone lenses in the host rock are poorly bedded. 

They weather brown but are dark gray on freshly 

broken surfaces. The limestone is composed almost 

entirely of calcite but contains minor amounts of quartz 

and chalcedony. The limestone lenses are much more 

coarsely crystalline than enclosing chert beds; the 

grains are commonly 1 to 4 mm in diameter. The 

coarse grain size seems to be due to recrystallization 

because many radiolaria replaced by calcite are enclosed 

within single calcite crystals. Quartz and chalcedony 

are in the form of remnants of beds bounded by stylolites and of small secondary deposits which have replaced calcite. Radiolaria are less abundant in the 

limestone than in the chert. Brachiopods of the genus 

Halorella have been found in limestone lenses near the 

Greystone mine.

No mineralogic, chemical, or textural differences were 

detected between the host rock closely surrounding and 

interbedded with barite and the host rock at various 

distances up to 2 miles from known barite deposits.

Barite deposits consist of interbedded baritev chert, 

and minor amounts of limestone (fig. 11.2). They 

range in size from a few feet to about 2,000 feet in longest exposed dimension. Outcrops are generally somewhat elongate parallel to host-rock bedding. Deposits 

commonly consist of a large irregular ore body associated with smaller pods at distances of a few hundred 

feet or less.

Like the host-rock chert, the barite contains thin 

shaly beds at intervals of one to a few inches and internal nonfissile laminae about 1 mm thick. Some separations between laminae resemble stylolites in their 

jagged cross sections and dark color. The microscopic 

texture of barite closely resembles that of chert. Barite 

is commonly dark gray of about the same shade as much 

of the associated chert and limestone.

An outstanding feature of the barite deposits is their 

mineralogic simplicity. Minor amounts of quartz and 

calcite are intimately mixed with barite in some beds. 

These may be vestigial remains of chert and limestone 

beds or, possibly epigenetic additions contemporaneous 

with introduction of the barite. Minute cubes of iron 

oxide scattered irregularly throughout the barite are 

apparently altered pyrite. Microscopic veinlets of 

quartz, calcite, and barite transect some of the barite

FIGURE 11.2. Interbedded barite and chert, 

Argenta deposit. A white stripe has been 

painted across all barite beds to distinguish 

them from the chert beds.

beds. At most, only small amounts of quartz, calcite, 

and pyrite have been introduced with the barite. The 

widespread occurrence of these minerals as veinlets in 

the host rock, distant from any known barite mineralization, suggests that only barite has been introduced in 

unusual concentrations.

Microf ossils resembling those present in the host rock 

are rather common in some barite beds. Brachiopods 

of the genus Halorella are sparsely distributed in some 

barite beds of the Greystone and Mountain Springs deposits. All the fossils are baritized and only the gross 

features are preserved.

The trace-element assemblages of associated barite, 

limestone, and chert beds are very similar; however, 

there are certain consistent differences as shown in the 

accompanying table. In the Slaven Canyon deposit 

the barite contains less manganese and magnesium than
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KETNEE B41

the limestone. Generally the barite contains less magnesium and boron but more strontium than the chert.

Replacement of both chert and limestone are indicated by the following somewhat ambiguous evidence. 

Although individual barite beds tend to be remarkably persistent, mapping of entire deposits indicates 

some discordance between the barite and host-rock bedding trends that are not easily explained by faults or 

unconformities. Baritized radiolaria and brachiopods 

prove replacement, at least locally. Because brachiopods do not occur in chert, replacement of limestone 

is specifically indicated. Microscopic textural evidence 

such as bedding textures of chert invaded discordantly 

by blades of barite also indicate local replacement. 

The similarity of texture and bedding in chert and 

barite indicates possible replacement, but the dissimilarity between texture and bedding of limestone lenses 

and barite indicates that no replacement took place. 

Perhaps original limestone textures have been altered 

except where they have been preserved by baritization.

Bedded barite deposits differ from neighboring metallic sulfide deposits in several important respects. 

Whereas the sulfide deposits are associated with a northwest-trending band of intrusive rocks, the barite deposits apparently are unrelated to concentrations of 

intrusives but rather are associated with bedded chert, 

especially Devonian chert. Metallic sufide deposits are 

controlled by shear zones and are clearly discordant 

with bedding, whereas the barite deposits are controlled 

to a large extent by bedding and tend to be concordant. 

The sulfide deposits contain the common mesothermal

mineral assemblage of pyrite, arsenopyrite, chalcopyrite, galena, sphalerite, and tetrahedrite, whereas few, 

if any, minerals but barite have been introduced to form 

the barite deposits. The sulfide deposits are commonly 

associated with recrystallized, bleached, and sericitized 

rock but barite deposits are remarkably free from host 

rock alteration.

In spite of these contrasts with the obviously epigenetic metallic sulfide deposits, the extensive textural 

evidence of replacement, great stratigraphic thickness 

of all deposits in relation to lateral extent, and the 

abrupt discordance revealed in some places by mapping 

and unexplained by visible faults or unconformities, 

indicate an epigenetic origin for barite too. It seems 

likely that the bedded barite deposits are replacements 

of both chert and limestone but are not obviously related to igneous intrusives or to other epigenetic deposits of the area.

The invariable association of bedded barite deposits 

with bedded chert formations and the apparent association with Devonian chert can be used in prospecting. 

Deposits under shallow cover might be discovered by 

detailed gravity surveys of areas underlain by bedded 

chert.

REFERENCES

Brobst, D. A., 1958, Barite resources of the United States:

U.S. Geol. Survey Bull. 1072-B. 

Ferguson, H. G., Muller, S. W., and Roberts, R. J., 1951, Geology

of the Mount Moses quadrangle, Nevada: U.S. Geol. Survey

Geol. Quad. Map GQ-12. 

Gianella, V. P., 1941, Barite deposits of northern Nevada:

Am. Inst. Mining Metall. Engineers Trans., v. 144, p. 294-299.
Article 12

ACCRETIONARY LAPILLI IN ROCKS OF THE CAROLINA 

SLATE BELT, STANLY COUNTY, NORTH CAROLINA

By HAROLD W. SUNDELIUS, Washington, D.C.

Abstract. Four occurrences of accretionary lapilli have been 

recognized in rocks of the Carolina slate belt in the Mount 

Pleasant quadrangle, North Carolina. The lapilli occur in lithic 

tuffs of intermediate composition. The presence of these lapilli 

suggests subaerial deposition of the enclosing rocks and volcanism from a vent located within several miles of the depositional site of the lapilli.

Four occurrences of accretionary lapilli have been 

found in the east-central part of the Mount Pleasant 

quadrangle, Stanly County, N.C. Accretionary lapilli 

are pellets of volcanic ash that show a concretionary 

structure. These lapilli also suggest subaerial deposition of the rocks in which they occur (Moore and Peck, 

1962, p. 191; Kittmann, 1962, p. 75,144). As far as the 

author is aware, this is the first reported occurrence 

of such lapilli in the southeastern part of the Piedmont 

province.

The Mount Pleasant quadrangle is located in the 

eastern part of the Piedmont of south-central North 

Carolina (fig. 12.1). The eastern two-thirds of the 

quadrangle is in the Carolina slate belt; the western 

one-third of the quadrangle is in the Charlotte belt 

(King, 1955, p. 346-350, map in pocket).

The rocks of the slate belt in this area are interbedded and intermixed tuffaceous argillite, sandstone, 

and siltstone, crystal and lithic tuffs, and minor flows, 

probably of Ordovician -age. Most of the volcanic material is felsic to intermediate in composition. Diabase 

dikes and gabbroic intrusive bodies are also present. 

The rocks occur in a southwest-plunging syncline and 

anticline and display both bedding- and axial-plane 

cleavage.

The rocks containing the lapilli are intermediate 

lithic tuffs that are mildly metamorphosed. Most of 

the tuff occurs in units 10 to 100 feet thick and is interbedded with and grades into tuffaceous sandstone and 

siltstone. Individual tuff units can be traced for distances ranging from several hundred feet to 2y2 miles.

FIGTTRE 12.1. Index map, Mount Pleasant quadrangle, North 

Carolina, showing occurrences of accretionary lapilli (X) 

The tuffs are fine grained, gray to greenish gray, and 

range from massive to well bedded.

The tuffs consist of a very fine grained matrix with 

an average grain size of about 0.1 mm, in which occur 

lithic fragments, crystals, crystal fragments, and accretionary lapilli (fig. 12.2). With high magnification, 

the fine-grained matrix is seen to be brown to yellowish brown and to consist chiefly of altered ash. Angular lithic fragments 0.1 to 5 mm in diameter constitute 

as much as 20 percent of the rock and are commonly

ART. 12 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B42-B44. 1963.

B42
SUNDELIUS B43

darker than the matrix. Crystals and crystal fragments are 1 mm or less in length.

The tuffs are composed mainly of plagioclase, potassium feldspar, and quartz, accompanied by a suite of 

secondary minerals that includes chlorite, clinozoisiteepidote, quartz, sericite, actinolite, calcite, hematite, 

and limonite. Plagioclase is generally in the albiteoligoclase range, is polysynthetically twinned, and in 

places, exhibits compositional zoning. It occurs as 

partially sericitized laths and as aggregates altered to 

sericite, chlorite, epidote, and calcite. Potassium feldspar is less abundant than plagioclase and is generally 

cloudy and partially sericitized. Both plagioclase and 

potassium feldspar occur as subhedral to euhedral crystals and crystal fragments in the matrix, in the lithic 

fragments, and in the lapilli. Original ferromagnesian minerals are altered to chlorite, clinozoisiteepidote, calcite, and actinolite and are no longer 

recognizable. Chlorite occurs as fibrous aggregates, 

some of which are oriented in a radial pattern, and in 

anhedral masses with rims of epidote. Clinozoisiteepidote generally occurs as anhedral grains and aggreFIGTJRE 12.2. Photomicrograph of large accretionary lapillus in altered intermediate lithic tuff. Note ellipsoidal 

shape of lapillus and concentric structure with fine-grained 

periphery. Note that core material is finer grained than 

that of host rock. Plane-polarized light. 

680231 O 63 4

gates. Quartz occurs both in anhedral grains and as 

stubby, euhedral hexagonal grains. A chemical analysis of tuff containing accretionary lapilli shows it to 

be intermediate in composition between andesite and 

dacite.

The accretionary lapilli are concentrated in 0.5- to 

5-cm beds of fine sand- to silt-sized tuff just above layers of medium to coarse sand-sized tuff. Lapilli beds 

are generally overlain by layers of fine silt and claysized tuff in which few lapilli occur. Lapilli have been 

found at several places in the lithic tuffs, but individual 

beds have not been traced beyond the outcrops in which 

they occur.

The lapilli are generally oblate spheroids or ellipsoids 3 by 1.5 by 2 mm in size and are -commonly flattened parallel to the bedding. Flattening of the lapilli 

is probably due to compaction of the beds in which 

they occur. Lapilli elongated parallel to prominently 

developed axial-plane cleavage were noted in a few 

places. The longest dimension of the lapilli examined 

ranges from 0.9 to 7.1 mm, and their median apparent 

diameter is 3.7 mm. The ratio of the major to minor 

axes of the lapilli in sections perpendicular to the bedding ranges from 1.0 to 2.0, and averages 1.3.' These 

figures are of the same magnitude as those reported by 

Moore and Peck (1962, p. 186) for lapilli from the 

western United States.

The lapilli are the same greenish gray as the tuff in 

which they occur. Because of the similarity in color 

and texture between lapilli and tuff, they are commonly 

quite difficult to detect and may easily be overlooked. 

On some weathered rock surfaces the lapilli stand out 

in relief and appear very much like small concretions; 

on others, lapilli weather out and leave small ellipsoidal 

or spherical voids.

Each lapillus commonly consists of a core surrounded 

by a finer grained peripheral zone. The core constitutes 

10 to 80 percent of the individual, lapillus and is commonly greater than 50 percent. It consists of ash 

similar to that of the enclosing tuff; however, the grain 

size of the tuff in the host rock is somewhat larger than 

in the lapilli.

Most of the lapilli observed have a single distinct concentric shell around the core. However, some of the 

lapilli have as many as three concentric shells around 

their cores (fig. 12.3). In such lapilli, the grain size 

decreases outward in each shell. Generally, the shell is 

continuous around a lapillus core, but there are cases 

where the shell is interrupted or pinches out at one or 

more places. Also, some shells are breached by or 

molded around angular lithic fragments. The outermost rim of the lapilli shells contains perceptible fibrous 

micaceous material that is commonly oriented tangen-
B44 GEOCHEMISTRY, MINERALOGY, AND PETROLOGY

FIGTJKE 12.3. Photomicrograph of accretionary lapillus in 

altered intermediate lithic tuff. Note lapillus in center of 

picture with fine-grained multiple shells around core. Left 

side of lapillus rim is flattened and broken. Also, note 

lapilli fragments to right of large lapillus and at upper 

right. Lapilli and host rock are composed of the same 

material. Plane-polarized light.

tially to the curvature of the lapilli. A few lapilli 

have formed by the coalescing of two separate cores.

Broken lapilli of a variety of sizes are mixed with 

unbroken lapilli (fig. 12.3). Although lapilli and lapilli

fragments are generally separated by tuffaceous material, in places they accumulate as closely packed 

spheres.

In a recent paper dealing with accretionary lapilli, 

Moore and Peck (1962, p. 189-191) summarized much 

of what is known about their formation. They reviewed older ideas and discussed several possible mechanisms by which accretionary lapilli could form. Of 

the mechanisms described, the one most plausible for 

the lapilli discussed here involves the accretion of moist 

volcanic ash around agglutinated cores of ash in an 

eruptive cloud of ash and water vapor. Cooling of 

water vapor in an eruptive cloud and subsequent condensation apparently resulted in the aggregating of ash 

particles to form lapilli cores. As the cores fell through 

the ash cloud, concentric layers formed around the 

cores by accretion of ash on the moist core surfaces.

The presence of accretionary lapilli is suggestive of 

subaerial deposition of the enclosing rocks, for long 

immersion in water would probably result in the dispersion of the accreted ash particles (Moore and Peck, 

1962, p. 191, 192; Rittmann, 1962, p. 75, 144). The 

presence of accretionary lapilli further suggests that 

eruptive vents from which the volcanic ash was ejected 

were above water at the time of eruption and were 

located within a few miles of the lapilli occurrences 

(Moore and Peck, 1962, p. 191).

REFERENCES

King, P. B., 1955, A geologic section across the southern Appalachians :an outline of the geology in the segment in 

Tennessee, North Carolina, and South Carolina, in Russell, 

R. J., ed., Guides to southeastern geology: Geol. Soc. 

America, p. 346-350.

Moore, J. G., and Peck, D. L., 1962, Accretionary lapilli in volcanic rocks of the western continental United States: Jour. 

Geology, v. 70, no. 2, p. 182-193.

Rittmann, Alfred, 1962, Volcanoes and their activity: New York, 

Interscience Publishers, 305 p.
Article 13

COMPOSITION AND ORIGIN OF SILICEOUS MUDSTONES

OF THE CARLIN AND PINE VALLEY QUADRANGLES, NEVADA

By KEITH B. KETNER and J. FRED SMITH, JR., Denver, Colo.

Work done in cooperation with the Nevada Bureau of Mines

Abstract. Ordovician, Silurian, and Devonian rocks in the 

upper plate of the Roberts Mountains thrust fault and postthrust Lower Mississippian rocks are largely siliceous mudstones. Analyzed representative samples contain 63 to 94 percent SiO2 and 1 to 13 percent A12O3 . Minerals are mainly quartz, 

chalcedony, illite, and dolomite. Grains are clastic quartz silt, 

clastic dolomite silt, fossil tests, and illite all in a mosaic 

matrix of apparently nonclastic quartz. The matrix quartz 

seems to be a chemical precipitate unrelated to local vulcanism.

In northeast Nevada, rocks of Ordovician through 

Devonian age include eastern (carbonatic), western 

(siliceous), and transitional (mixed) rock assemblages 

telescoped by the Roberts Mountains thrust of latest 

Devonian or earliest Mississippian age (Roberts and 

others, 1958). Mississippian through Permian rocks, 

which overlap the thrust fault, include limestone and 

a variety of clastic rocks whose siliceous particles were 

derived from uplifted western and transitional assemblages. The dominant rock types in the western, transitional, and basal-overlap assemblages can be described 

by the term "siliceous mudstone." As used here the 

term embraces all siliceous sedimentary rocks made up 

of grains of silt size and smaller (less than 0.06 mm). 

Included are such rocks as chert, siliceous shale, siliceous argillite, cherty siltstone, and intermediate types 

difficult to classify.

Detailed stratigraphy of western- and transitionalassemblage rocks is unknown owing to poor exposures, 

complex structure, and scarcity of fossils. Data are 

necessarily derived from studies of isolated and widely 

scattered exposures. The samples sectioned and analyzed in this investigation, however, are firmly dated 

by fossils as to system. Ordovician rocks of the western and transitional rock assemblages in the Carlin and 

Pine Valley quadrangles consist mostly of siliceous 

mudstone aggregating hundreds of feet in thickness,

and lesser amounts of limestone, quartzite, and greenstone. Fossils found include graptolites, chitinous 

brachiopods, hystrichospheres (?), sparse radiolaria (?), and unidentified spheroids about 5 microns 

in diameter. Silurian rocks of the western assemblage 

are represented by a single dated sample. However, 

we believe it to be representative of the several hundred 

feet of siliceous mudstone beds that we assign to the 

Silurian western asemblage. Graptolites were the only 

fossils found. Devonian rocks of the western and transitional facies include siliceous mudstone strata aggregating several hundred feet in thickness, and less abundant limestone and impure clastic dolomite. Fossils 

found include eurypterids (appendages only), conodonts, goniatites, and abundant radiolaria (?).

The stratigraphy of the overlap assemblage is better 

known; Mississippian rocks of this assemblage include 

basal units consisting partly of siliceous mudstone lithically similar to some rocks of the western and transitional assemblages. These basal units are divided into 

a very siliceous lower unit A of Kinderhook age and an 

overlying more argillaceous unit B of Kinderhook to 

Osage age (see table). The lower unit bears certain 

lithic similarities to the Pilot Shale but as dated it is 

somewhat younger. Fossils found in it include conodonts, spicules, and radiolaria. The upper unit is similar in many respects to the Chainman Shale. Fossils 

found include goniatites, brachipods, and corals. The 

Joana Limestone is not present.

The siliceous mudstones have certain features in common regardless of age. They are generally dark brown 

or dark gray to black when fresh but commonly 

weather light gray. On many weathered surfaces, beds 

an inch to a few inches thick are visible. Microscopic 

examination reveals internal nonfissile laminations a

ART. 13 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B45-B47. 1963.

B45
B46 GEOCHEMISTRY, MINERALOGY, AND PETROLOGY

fraction of a millimeter thick in most specimens. 

Microfossils are sparse. Thin section, X-ray, and 

chemical analyses indicate that quartz, including chalcedony, is the dominant mineral; silica composes 63 

to 94 percent of the analyzed samples as shown in the 

accompanying table. Angular detrital quartz grains 

0.01 to 0.03 mm in diameter compose on the average 

less than 10 percent of the rock; the groundmass material is composed mainly of interlocking grains of quartz 

and chalcedony, which range widely in size but average about 0.005 mm in diameter. Aluminous minerals 

including illite, potash feldspar, and kaolinite compose an estimated 3 to 20 percent of most samples. 

X-ray diffraction patterns indicate that illite is the 

dominant aluminous mineral. The few grains of potash feldspar seen in thin section are angular and appear 

to have grown in place. It is not known whether the 

illite and kaolinitp are detrital or are alteration products of other detrital minerals. Dolomite is a common 

apparently clastic component in some samples, for example 553-C and 553-E, and small amounts of brown 

organic debris, amorphous carbon, iron oxides, and 

water are present in most specimens. Pyrite, marcasite, barite, jarosite, alunite, and gypsum are present 

in some. The sulfides (pyrite, marcasite) have replaced quartz, whereas the sulf ates tend to occur in veinlets. Aluminous sulfates may have formed during 

weathering by reaction of sulfuric acid, derived from 

oxidation of iron sulfides, with aluminous silicates. 

For many samples, the carbon content indicated in the 

table may be abnormally low because weathered material was analyzed. The mineralogic occurrence of 

many of the elements listed is unknown. The quantities of iron, sodium, potassium, titanium, magnesium 

(in excess of magnesium in dolomite), and zirconium, 

however, correlate well enough with aluminum to sugAnalyses of samples of Paleozoic siliceous mudstone, Carlin and Pine Valley quadrangles, Nevada

[Chemical analyses by Paul Elmore, Samuel Botts, Gillison Chloe, and I. C. Frost, 1962; spectrographic analyses by Nancy Conklin, 1962]

Chemical analyses (weight percent)

Age and unit

Mississippian: 

Unit B. ...

Unit A. __ .

Silurian.. ..........

Sample No. 1

580... .......

SR-132 ...

161 

218 

872 

823-B. 

577-.--.--..

822-B .

822 A.......

825-.-- ..

818 ..

508- 

553-D.......

553-C 

553-E. ......

SiO2

71.7

67.6

89.4

89.3

86.4

84.2

84.0

QA K

84.4

80.4

CO 0

QA A

QO 9

87.8

79 7

77.8

63.0

A12 03

9 Q

10 ft

4 Q

3.2

4.8

6.6

4.0

1.1

1.5

5.7

c A

4.2

6.1

e 0

5.1

Fe2 O3

2 A

4.1

.16

89

.36

79

00

.41

99

1 S

1.5

1 C

.36

.63

1.0

sn

1.0

FeO

2

0

1.4

9fi

07

99

.15

.56

19

oo

.40

.17

9fi

.11

.14

.27

.86

MgO

1 Q

1 Q

.35

.47

.40

.43

.74

no

.15

QQ

°.Q

39

49

1.1

3.6

CaO

1 9

CO

.10

9Q

.02

.04

.67

.14

.14

.12

.17

97

.10

.12

.47

1 ft

5.6

Na2 O

0.24

7°.

.12

.12

.12

.14

.12

.06

.07

.10

.08

.06

.11

.13

.08

.08

.08

K2 O

9 fi

3 9

1.3

.63

1.2

1.6

8Q

1 Q

.26

1 ft

2

0

2.0

1.1

1.3

1.1

1 9

1.4

TiO2

0.69

.82

.23

.16

.29

.33

.22

.04

.06

.32

.34

°.n

.06

.22

.25

.28

.25

.24

P2 05

0.27

.30

.10

.17

.08

.12

.13

.06

1.2

.16

.12

.08

.10

.07

.08

1 9

2.0

MnO

0.04

.02

.02

.02

.00

.00

.01

.01

.01

.02

.00

.01

.01

.00

.00

.01

.02

.04

Loss 

on 

ignition

5.4

6.1

1.7

1.7

3.5

4.7

7.0

2.0

3.3

3.8

7.1

2.0

1.9

2.2

3.1

8.2

8.1

14.9

Total 

C

0.73

.51

.39

.70

.43

.16

3.64

.69

1.37

.36

1.81

.20

.73

.15

.23

2.46

3.42

5.12

CO2

0.92

.07

<.05

.06

<.05

<.05

.06

<.05

<.05

<.05

<.05

<.05

<.05

.06

<.05

<.05

1.8

7.6

Acid 

insoluble'

;

2.1

2.7

2.4

Semiquantitative spectrographic analyses (weight percent)

Age and unit

Mississippian: 

Unit B...

Unit A.--

Devonian

Silurian... .....

Ordoviciai).--

Sample No. 1

" 80

SR-132 

161 

218 

872

823-B 

525 

577 

822-B . 

822-A 

506 

825

818 

564 

508 

553-D553-C, .

553-E. ......

B

.007 

.015

.005

.01

.01

.007

.003

.003

.015

.01

ftfl^

.002

.007

.007

.007

.007

.005

Ba

0.2

.15 

.3

1.5

1 ^

.3

.2

.15

.15

.3

.2

.07

.07

.1

.15

1.

Cr

0.01

.01

.005

.005

ftfl^

ftO^

nn°.

.0015

.007

.005

003

.0015

.01

.007

.007

Cu

0.007

.007 

ft A 3

.01

.003

fifteL

ftfW

ftfW

.01

001

ftft7

on*;

.007

rwi7

.015

Ga

O ftO9

.003

.0005

nnn?

.0015

0

0

0

.001

.0015

001

0

.0007

.0007

.0015

.001

.0015

La

0

.005 

0

0

0

0

0

0

0

.005

.005

0

0

0

0

0

0

0

Mo

.001 

0

0

0

.0015

0

.001

.015

.003

0

.0007

0

0

.0007

0

.0005

Ni

.007 

.001

.003

.005

.001

.007

.002

.003

.007

.0015

.002

.0007

.0007

.003

.005

.005

Pb

.0015 

.0015

0

0

0

0

0

0

.0015

.0015

.0015

0

0

0

. 0015

0

.0015

Sc

0.0015

.002 

.0007

.001

.001

.0015

0

0

0

.001

.001

0

0

.0005

.0005

0

0

.001

Sr

0.015

.015 

.005

.007

.015

.015

.005

.003

.01

.02

.015

.003

.002

.005

.005

.05

.05

.05

V

0.02

.02 

.015

.015

.015

.015

.07

.005

.015

.07

.007

.015

.02

.01

.015

.02

.015

.02

Y

0.003

.003 

.0015

.0015

.001

.0015

0

0

.001

.002

.001

.0015

.0015

.001

.002

.001

.002

.003

Yb

0. 0003

.0003 

.0002

. 00015

. 00015

.00015

0

0

.0001

.0003

.0001

.0002

. 00015

.00015

.0002

.00015

.0002

.0003

Zr

0.02

.015 

.007

.003

.007

.01

.002

0

0

.007

.007

.02

.001

.005

.005

.005

.015

.007

Samples arranged in order of silica content within stratigraphic units. 2 Chiefly barite.
KETNER AND SMITH B47

gest that these elements occur mainly in illite and the 

other aluminous minerals. The phosphorous content 

seems to correlate with the quantity of chitinous (?) 

debris observed in thin section.

The siliceous mudstones appear to be little-altered 

primary deposits, rather than replacements of preexisting beds of another lithic type such as volcanic 

ash or limestone. Relict volcanic and carbonatic textures are not visible in thin section, and the good preservation of some delicate organic structures and fine 

bedding precludes the possibility that volcanic or other 

rocks have been replaced with complete loss of original 

texture.

Although siliceous tests compose as much as 40 percent of a few beds, particularly those of Devonian age, 

on the whole they cannot account for more than a small 

fraction of the free silica present. Recognizably clastic 

quartz composes only a small fraction of the free silica.

The remaining silica of uncertain origin forms a 

mosaic of grains a few microns in diameter. The silica, 

must be either recrystallized very fine detrital or organic silica, or else a chemical precipitate. Because 

the mosaic or interlocking texture of the fine-grained 

groundmass could hardly have formed by recrystallization of original detrital or organic silica without 

obliterating observed radiolaria, we tentatively conclude that the groundmass is a chemical precipitate.

Whether the groundmass quartz was chemically precipitated or organically derived, it would seem to necessitate unusually high concentrations of dissolved silica. 

Marine vulcanism, which is commonly thought to supply such high concentrations, prevailed both locally and 

regionally in Ordovician time, but local exposures of 

Devonian, Silurian, and Mississippian rocks contain 

no volcanic deposits. Apparently local contemporaneous vulcanism was not required for the formation of 

siliceous mudstone.

REFERENCE

Roberts, R. J., Hotz, P. E., Gilluly, James, and Ferguson, H. G.. 

1958, Paleozoic rocks of north-central Nevada: Am. Assoc. 

Petroleum Geologists Bull., v. 42, no. 12.
Article 14

TUFFACEOUS SANDSTONES IN THE TRIASSIC CHINLE FORMATION, COLORADO PLATEAU

By ROBERT A. CADIGAN, Denver, Colo.

Abstract. Petrographic and chemical analyses of six sandstone samples from southern outcrops of the Petrified Forest 

Member of the Chinle Formation show the sandstones to be 

sodic arkoses and feldspathic tuffs with textures definitive of 

sediment derived from a moderately active tectonic environment. Source rocks probably included latitic to rhyolitic extrusive and intrusive rocks.

Detrital volcanic debris was recognized in the Chinle 

Formation (Late Triassic) of northeastern Arizona by 

Merrill (1911, p. 22) and by Alien (1930, p. 285).- 

Later investigations showed volcanic debris to be 

characteristic of the Chinle throughout most of the 

Colorado Plateau region. Petrologic studies of the 

Chinle Formation are in progress, and preliminary results reported below show the general character of 

tuffaceous sandstone typical of the formation in three 

widely separated localities in the southern part of the 

Colorado Plateau region (fig. 14.1).

The most tuffaceous sandstones of the Chinle Formation are in the Petrified Forest Member, which is comLORADO

ingate

NEW 

MEXICO i

250 MILES

FIGURE 14.1 Extent of Triassic rocks in the Colorado 

Plateau region, and location of sampled areas.

posed of interbedded mudstone and sandstone in thick 

beds (table 14.1). The mudstone is purple, red, orange, 

brown, pale green, and gray, and the sandstone is 

characteristically gray. In northeastern Arizona and 

northwestern New Mexico, some of the sandstone strata 

form a thick continuous unit referred to as the Sonsela 

Sandstone Bed (Stewart and others, 1957), which contains the petrified trees and logs of the Petrified Forest 

National Monument. The variegated mudstones of the 

Petrified Forest Member color the Painted Desert of 

northeastern Arizona.

TABLE 14.1. Stratigraphic vhart of the Chinle Formation in 

northeastern Arizona

Formation

Wingate Formation

Chinle Formation

Moenkopi Formation

Member

Rock Point i

Owl Rock

Petrified Forest

Sonsela Sandstone 

Bed

Lower

Sbinarump

Thickness 

(feet)

800

400

100

900

150

0-60

Description

Structureless reddish-brown 

horizontally laminated siltstone and crossbedded sandstone.

Thin limestone strata interbedded with structureless 

red siltstone.

Variegated bentonitic claystone and clayey sandstone 

with conspicuous sandstone 

and conglomerate strata. 

The latter are dominant in 

the Sonsela Sandstone Bed 

and in the Shinarump 

Member.

i The same lithologic unit as the Church Rock Member of the Chinle Formation 

in Utah (Stewart, 1957, p. 460).

Six samples, all from the Petrified Forest Member, 

were studied: 2 samples from near the southern border 

of the Petrified Forest National Monument, Ariz., the 

type area of the member; 2 from the Fort Wingate area, 

McKinley County, N. Mex.; and 2 from the Valley of 

Fire area, Clark County, Nev. (fig. 14.1). Results of 

modal analyses made by the point-count method are

B48

ART. 14 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B48-B51. 1963.
CADIGAN B49

given in table 14.2. The samples from the Petrified 

Forest National Monument and the Fort Wingate area 

are impure tuffs, technically feldspathic (sedimentary) 

tuffs, and those from the Valley of Fire are arkoses 

with a minor tuff content.

Mineral components were generally identified by their 

optical properties. Indeterminate clay mixtures, which 

present.an anomalous optical appearance compared to 

the better crystallized, relatively pure clay minerals, 

were identified by X-ray analysis. "Corrensite" was

TABLE 14.2. Modes of sandstone from the Petrified Forest 

Member of the Chinle Formation

[In percent by volume]

Mineral components

Petrified 

Forest 

National 

Monument l

2L2067 «L2069

Fort

Wingate '

2 L 2366 L2367

Valley of 

Fire'

L2356 L2358

Mean 

modal 

composition, 

by volume

I. Cementing components (except silica)

a. Carbonates and sulfates 

(calcite) ______ ..... 

b. Red interstitial iron oxides

Total 

0.8

.8

0.2 

.2

.4

0.6

.6

2.0 

.2

2.2

0.2 

13.2

13.4

2.4

2.4

0.6 

2.7

3.3

II. Siliceous components

a. Quartz grains and overb. Quartzite fragments ........

c. Chert, detrital- -- ..........

d. Chert, authigenic _____

Total--.. _ ... ... 

29.0

1.8

.2

31.0

32.8

.8

33.6

44.0 

.2 

.2

44.4

40.0 

1.4 

1.6

43.0

24.2 

.8 

.2

25.2

41.6 

3.0

44.6

35.3 

.9

.7

36.9

III. Feldspathic components

a. Potassic feldspar. - .........

c. Kaolinitic clays (alteration 

products). .................

Total . 

4.4

12.0

16.4

3.8

12.4

16.2

17.0

17.0

12.0

12.0

0.2 

46.6

5.2

52.0

12.4 

328.8

.8

42.0

3.5 

21.5

1.0

26.0

IV. Dark-mineral and mica components

a. Mica flakes and books (biotite, chlorite) ................

b. Chlorite and mica clays ....

c. Micaceous and mafic rock

d. Heavy minerals (sp gr 

2.90+) 

e. Miscellaneous (opaque and 

unidentified grains) ........

Total 

0.4 

.4

0.8

0.4 

.6

*.6

1.6

1.0 

7.0

1.4

1.4

10.8

0.2

.4

.2 

1.8

2.6

0.2 

.8

.2

«.6 

.6

2.4

0.2 

.8

.2

1.2

0.4 

1.6

.3 

.3 

.6

3.2

V. Volcanic components

a. Tuff and felsite fragments 

(silicified) 

b. Montmorillonitic clays (alteration products) ...........

c. Altered ash (chiefly shards,

Total 

Grand total _ _ -. .....

37.2

13.8

51.0

100

34.2

14.0

48.2

100

16.6

10.6

27.2

100

12.2 

1.2

26.8

40.2

100 .

6.8 

.2

7.0

100

9.8

9.8

100

19.5

.2 

10.9

30.6

100

1 See table 14.3 for identification of samples.

2 From the Sonsela Sandstone Bed. 

* Plagioclase partly altered to chlorite. 

4 Includes a little glauconite. 

6 Includes authigenic epidote.

identified (sample L2367) by J. C. Hathaway (1959, 

written communication). Aluminous montmorillonite 

mixed with chlorite was identified in samples L2358, 

L206T, and L2069 by L. G. Schultz (1962, oral communication) . X-ray and optical data suggest that the 

clays in all the samples are generally high-magnesium, 

low-iron, and high-aluminum varieties of chlorite and 

montmorillonite.

Cementing components of the sandstones (table 14.2) 

are chiefly interstitial calcite and nearly opaque iron 

oxide-impregnated clay. Siliceous components consist 

of quartz grains and minor overgrowths, quartzite fragments, and detrital chert. Feldspathic components are 

primarily orthoclase and albite; minor interstitial 

kaolinite present in some samples is included with the 

feldspathic components because it is presumed to be an 

alteration product of them. Dark-mineral and mica 

components are principally chlorite, altered biotite, and 

finely crystalline mixtures of chlorite and mica clays. 

Nonopaque heavy minerals included with this group 

are principally apatite, authigenic epidote, glauconite, 

and zircon; opaque heavy minerals are leucoxene, magnetite, and ilmenite. The volcanic components consist 

of devitrified, silicified, and chloritized fragments of 

the following: volcanic glass or vitrophyre; vitric, 

crystal, and lithic tuff; and felsite. Montmorillonitic 

clays are included in this group because they are intergrown with chlorite and amorphous and microcrystalline silica in interstitial aggregates that probably represent devitrified hydrolyzed volcanic ash.

Tuff and vitrophyre fragments include both sodic 

and potassic varieties. The sodic variety predominates 

and in some samples is the only one observed. Some 

sodic tuff contains albite in euhedral phenocrysts or in 

poorly crystallized anhedral forms. Potassic fragments were identified by the yellow stain produced by 

sodium cobaltinitrite after etching with hydrofluoric 

acid. All original glass fragments are divitrified and 

show contorted flow or perlitic structures. The altered 

glass which forms the matrix of the crystal tuff fragments has a birefringence of 0.001 or less, but has 

indices of refraction approximately the same as quartz, 

which suggests that it has been altered to a poorly 

crystallized form of quartz.

The chemical composition of the six samples as determined by standard rock analysis is shown in table 14.3, 

which also includes, for comparison, average composition of some common clastic sedimentary rocks. With 

certain exceptions, the samples in general resemble 

average arkose. The sodium oxide content is near that 

of average graywacke, but the potassium oxide content 

is lower than in any of the compared rocks. Magnesia 

is comparable to that of average sandstone.
B50 GEOCHEMISTRY, MINERALOGY, AND PETROLOGY

TABLE 14.4. Calculated distribution of chemical elements among components of tuffaceous sandstones in the Petrified Forest

Member of the Chinle Formation

[Analyst: Margaret C. Lemon]

Sioj . _______ ____ ......

A1 2 03 ... .......................

Fe2 O3- _ ........... _ ......

FeO... ..

MgO.... .........................

CaO. ..__ ._____.

NajO..... ........................

K2 0........ ......................

HjO+... .........................

HaO-._. ..................

TiOj.. .......................

P»08 ...... ................ ........

MnO.. - ..

CO3 .. ........ .

Cl........... ...... ....

F. ...

S_. ....

Total.... ...................

1 2 3 4

Tuff

75.93 

10.46 

.87 

.56 

1.74 

1.05 

2.31 

1.09 

2.30 

2.83 

.17 

.06 

.07 

.33 

.03 

.03 

.03

99.86 

.03

99.83

76.92 

10.39 

1.02 

.36 

1.58 

.80 

2.29 

.95 

2.24 

2.65 

.29 

.05 

.04 

.11 

.01 

.03 

.01

99.74 

.02

99.72

77.96 

10.23 

.75 

.70 

1.92 

.56 

2.55 

.24 

2.33 

1.92 

.40 

.09 

.03 

.03 

.02 

.03 

.03

99.79 

.03

99.76

80.09 

8.95 

1.30 

.49 

2.00 

.97 

2.31 

.10 

2.10 

.67 

.25 

.08 

.04 

.58 

.00 

.02 

.01

99.96 

.01

99.95

5 6

Arkose

71.67 

14.20 

3.30 

.12 

.27 

.86 

5.55 

.53 

1.30 

.26 

.48 

.02 

.02 

.79 

.01 

.01 

.16

99.55 

.08

99.47

81.38 

9-85 

.92 

.02 

.27 

.24 

3.59 

1.83 

.69 

.17 

.14 

.02 

.01 

.01 

.02 

.01 

.08

99.30 

.04

99.26

Average 

of 

tuffs

77.73 

10.01 

.99 

.53 

1.81 

.85 

2.37 

.60 

2.24 

2.02 

.28 

.07 

.04 

.26 

.02 

.03 

.02

Average 

of 

arkoses

76.53 

12.03 

2.11 

.10 

.27 

.55 

4.57 

1.18 

1.00 

.22 

.31 

.02 

.02 

.40 

.02 

.01 

.12

Average 

of the 6 

samples

77.32 

10.68 

1.36 

.38 

1.30 

.74 

3.10 

.79 

1.83 

1.42 

.29 

.05 

.04 

.31 

.02 

.02 

.05

Average sandstone i

78.66 

4.78 

1.08 

.30 

1.17 

5.52 

.45 

1.32 

1.33 

.31 

.25 

.08 

Tr. 

5.04 

Tr.

Average orthoquartzite 2

93.16 

1.28 } «

.07 

3.12

} .39

} .65 

.03

2.01

Average 

graywacke 3

64.7 

14.8 

f 1.5 

I 3.9 

2.2 

3.1 

f 3.1 

\ 1.9 

I 2.4 

I -7 

.5 

.2 

.1 

1.3

.2

Average 

arkose 4

76.37 

10.63 

2.12 

1.22 

.23 

 1.30 

1.84 

4.99

} .83

.41 

.21 

.25

.54

1. Feldspathie orthoquartzitie tuff; Sonsela Sandstone, 3 ft above base; 600 ft notth of intersection of U.S. Highway 260 with the southern boundary of Petrified Forest National Monument, Navajo County, Ariz. Sample No. L2067, USQS Serial No. F2788.

2. Feldspathie orthoquartzitie tuff; Sonsela Sandstone, 22 ft above base; same location as 1. Sample No. L2069, USQS Serial No. F2789.

3. Feldspathie orthoquartzitie tuff; Sonsela Sandstone, 140 ft above base; Fort Wingate Ordnance Depot, MeKinley County, N. Mex. (see. 30, T. 15 N., E. 16 W). Sample 

No. L2366, USOS Serial No. F2790.

4. Feldspathie orthoquartzitie tuff; Petrified Forest Member, 170 ft above base of Sonsela Sandstone; same location as 3. Sample No. L2367, USQS SerialNo. F2791. 

- krkose; Petrified Forest Member, 134 ft above base of Chinle Formation; Valley of Fire, Clark County, Nev. (sec. 26, T. 17 S.,~ " *" " ' " - T 

No. F2786.

5. Arkose; Petrified Forest Member, 134 ft above base of Chime Formation; Valley of Fire, Clark County, Nev. (sec. 26, T. 17 S., R. 67 E). Sample No. L2356, USQS Serial 

No. F2786.

6. Tuffaeeous arkose; Petrified Forest Member, 269 ft above base of Chinle Formation; same location as 5. Sample No. L2358, USQS Serial No. F2787.

1 Averages of 253 undifferentiated sandstones (Clarke, 1924, p. 547).

2 Averages of 8 typical orthoquartzites (Pettijohn, 1957, p. 298).

3 Averages of 23 typical graywackes (Pettijohn, 1957, p. 307).

4 Averages of 7 typical arkoses (Pettijohn, 1957, p. 324).

In table 14.4, by combining modes and chemical analyses, an attempt is made to show distribution of the 

chemical elements among the mineral components of the 

rocks. Such calculations require many assumptions as 

to composition of mineral and rock fragments, and the 

results must be regarded as only approximate.

Textural measurements of the samples are summarized in table 14.5. The averages of the statistical measures, with the exception of those of L2356, indicate that 

the sandstones are medium grained and moderately to 

poorly sorted, with moderately skewed and moderately 

peaked grain-size distributions. An analysis of such 

textural features has been made previously (Cadigan, 

1961). The average sorting (standard deviation) and 

average mean grain size suggest that the sediments were 

derived from source areas undergoing uplift at a moderate rate, and that they were deposited in an area 

undergoing subsidence at a moderate rate. Average 

skewness and kurtosis indicate low to moderate amounts 

of reworking or transportation.

Stewart and others (1957) have interpreted the 

Petrified Forest Member of the Chinle Formation as a 

stream deposit made up of debris from volcanic and

sedimentary rocks derived from a source area in southern Arizona and vicinity. The present study in general 

confirms this interpretation. The source area must have 

contained large bodies of volcanic and possibly intrusive rocks, probably latitic to rhyolitic in composition.

REFERENCES

Alien, V. T., 1930, Triassic bentonite of the Painted Desert: 

Am. Jour. Sci. ser. 5, v. 19, no. 112, p. 283-288.

Cadigan, B. A., 1961, Geologic interpretation of grain-size distribution measurements of Colorado Plateau sedimentary 

rocks: Jour. Geol., v. 69, no. 2, p. 121-142.

Clarke, F. W., 1924, The data of geochemistry: TJ.'S. Geol. 

(Survey Bull. 770.

Merrill, G. P., 1911, The fossil forests of Arizona: Washington, 

B.C., U.S. Govt. Printing Office, 23 p.

Pettijohn, F. J., 1957 Sedimentary rocks: 2d ed., New York, 

Harper and Bros., 718 p.

Stewart, J. H., 1957, Proposed nomenclature of part of Upper 

Triassic strata in southeastern Utah: Am. Assoc. Petroleum 

Geologists Bull., v. 41, no. '3, p. 441-465.

Stewart, J. H., Poole, F. 'G., and Wilson, R. F., 1957, Triassic 

studies, in Geologic investigations of radioactive deposits, 

Semiannual progress report, June 1 to Nov. 30, 1957: U.S. 

Geol. 'Survey TEI 700, p. 115-122, issued by U.S. Atomic 

Energy Comm. Tech. Inf. Service, Oak Ridge, Tenn.
CADIGAN

TABLE 14.4. Calculated distribution of chemical elements among .components of tuffaceous sandstones in

Member of the Chinle Formation

B51

the Petrified Forest

SiOj A1O3 ~ FeO MgO CaO Najb KaO H,0+ HZ O- TiOs PjOs MnO CO2 Cl F S

Computed 

mean 

modal 

composition (percent 

by 

weight)

Cementing components (except silica)

Total-... - 

1.07

1.07

0.86

.86

0.57

.57

0.34

.04

.38

0.18

.18

0.08

.08

0.27

.27

0.61 

2.80

Siliceous components

hert, detritaL- _____ ____________

Chert, authigenic.. -. ... . ...

Total-. 

35.35

.90

.70

36.95

35.35 

.90 

.70

Feldspathic components

Potassie feldspar.. ... _____ ... _ ..........

Total... ..............................

2.15

14.40

.45

17.00

0.62

4.06

.38

5.06 ......

0.01

.01 ......

0.14

.14

2.68

2.68

0.59

.01

.60

0.14

.14

0.02

. .02 ...... ...... ......=...... ..... . ......

3.37

21.29 

.99

Dark-mineral and mica components

Mica flakes and books _______ ............

Heavy minerals (sp gr2.90+)__ ................

0.18 

.40

.14

\

0.11 

.37

.12

 .

0.37

~6.~37~ 0.09 

.28

 

0.03

 0.05

.04

0.01 

.19

.02

.12

"O.'OJT

.23

 

0.29

......

0.05 0.04

-...--

0.04 0.02 0.02 0.05

0.44 

1.64 

.32

1.26

Volcanic components

Tuft and felsite fragments (silieified) . . _____ 

Montmorillonitie clays (alteration products) . . .

Altered ash (chiefly shards and clay) ..........

Total.. ..................................

Grand total- ______ . _____ _ _.

17.60 

.11

4.10

21.81

77.55

0.92 

.06

3.21

4.19

10.71

 

0.42

.42

1.36

..- 

0.38

 

0.94

.94

1.31

0.04 

.01

.14

.19

.74

0.43

------

.43

3.11

0.10

------

.10

7Q

0.17

1.01

1.18

1.84

0.17 

.01

.89

1.07

1.43

......

 

------

0.29

......

------

0.05

... 

......

0.04

......

......

 

0.31

......

......

......

0.02

.... ..

......

0.02

......

......

0.05

19.43 

.19

10.71

.......

100. 00

TABLE 14.5. Grain-size distribution measures of sandstone samples from the Petrified Forest Member of the Chinle Formation

Sample No.i

L2067---.

L2069.-- _

L2366- ........... .. ..

L2367. ................ -.

L2356 ».-_--

L2358. ............ ..............

Statistical measures of the grain-size distribution

Mean

Millimeters

0.173 

.217 

.097 

.131 

.243 

.159 

.155

*

2.54 

2.21 

3.38 

2.95 

2.04 

2.66 

2.75

Median

Millimeters

0.397 

.376 

.177 

.218 

.259 

.214 

.276

Standard 

deviation

2.907 

2.306 

2.112 

1.925 

.842 

1.590 

1.968

Skewness

0.860 

1.309 

1.111 

1.195 

1.200 

1.056 

1.106

Kurtosis

1.825 

6.229 

4.165 

5.516 

8.942 

6.082 

4.763

Grain-size distribution by percentiles (phi units)

Pa

0.00 

.61 

1.61 

1.39 

1.42 

.86 

.89

P^

0.33 

.74 

1.85 

1.58 

1.52 

1.09 

1.12

Pie

0.70 

.96 

2.11 

1.84 

1.68 

1.48 

1.42

Pso

1.34 

1.42 

2.51 

2.21 

1.95 

2.23 

1.94

Pw

5.05 

2.92 

4.36 

3.81 

2.70 

3.76 

3.98

P«5

11.04 

9.46 

9.26 

7.82 

3.60 

5.69 

8.65

P98

13.31 

14.00 

11.38 

10.07 

4.50 

7.01 

11.16

i See table 14.3 for identification of samples.

8 Grain-size distribution measures from thin section; not used in computation of mean values.
Article 15

PETROGRAPHIC CHARACTERISTICS OF SOME WELDED TUFFS

OF THE PIAPI CANYON FORMATION, NEVADA TEST SITE, NEVADA

By J. T. O'CONNOR, Denver, Colo.

Work done in cooperation with the U.S. Atomic Energy Commission

Abstract. Significant petrographic characteristics of four 

welded-tuff cooling units include the quartz-alkali feldsparplagioclase feldspar ratios of the phenocrysts and the K2O-Na2O 

ratios of the alkali feldspar phenocrysts. Some welded tuff, 

previously thought to belong to the cooling unit of the Tiva 

Canyon Member, is actually part of the Rainier Mesa.

The Piapi Canyon Formation of Miocene or younger 

age (Poole and McKeown, 1962) at the Nevada Test 

Site, southern Nye County, Nev., is a sequence of rhyolitic ash-flow tuffs with minor interlayers of air-fall 

tuffs and water-laid tuffaceous rocks. A brief description and the distinguishing features of some of these 

tuffs is given by Hinrichs and Orkild (1961). In order 

to supplement the previous work by establishing petrographic criteria for correlation of the tuff units, a systematic petrographic study of the welded tuffs was 

undertaken. This article is based on preliminary results of the study of four units of the Piapi Canyon 

Formation. These are, in ascending order:

1. The Topopah Spring Member, a multiple-flow 

cooling unit, which may show compound cooling in the 

western part of the Nevada Test Site.

2. The tuffs of Spring Canyon, a multiple-flow cooling unit which occurs mainly in the western part of 

the Nevada Test Site.

3. The Tiva Canyon Member, a compound cooling 

unit.

4. The Rainier Mesa Member, a composite sheet. 

This member contains several multiple-flow cooling 

units.

Each of these cooling units generally includes a 

bedded air-fall tuff at its base. The units are separated by bedded tuffaceous sedimentary rocks. The 

terminology followed is that of Smith (1960).

Modal analyses of thin sections of specimens from 

nine measured stratigraphic sections (fig. 15.1) and

various selected localities at the Nevada Test Site were 

made with a Chayes counter. The measured sections 

include sections 11 through 15 of Hinrichs and Orkild. 

The phenocrysts were tabulated separately from the 

groundmass, and mineral amounts were calculated as 

percentage of total phenocrysts. Counting was continued until the variation of mineral amounts was less 

than a few percent per hundred phenocryst points. A 

few of the thin sections from the Tiva Canyon and 

Topopah Spring members (cooling units) contained 

too few phenocrysts for this small variation. The 

quartz-alkali feldspar-plagioclase feldspar ratios from 

the modal analyses are plotted on a triangular diagram 

in figure 15.2.

The composition and optical properties of the feldspar of the Piapi Canyon Formation were determined 

from mineral separates from the same specimens from 

which the thin sections were cut. The compositions 

of the alkali feldspars were determined with an X-ray 

diffractometer by the position of the (201) reflection 

(Tuttle and Bowen, 1958). The samples were homogenized and a quartz internal standard was used. The 

optical properties of the alkali feldspars were determined on a spindle stage (Wilcox, 1959).

The anorthite content of the plagioclase feldspar of 

samples from the lower part of the Rainier Mesa Member (composite sheet) was determined by fusing a bead 

of plagioclase and measuring its refractive index 

(Schairer and others, 1956). Other plagioclase compositions were estimated from albite-twin extinction 

angles.

Phenocrysts in the Topopah Spring Member (cooling 

unit) characteristically consist of two feldspars; quartz 

is uncommon. One feldspar is sanidine having an optic 

angle ranging between 35° and 45°, a composition 

near Ab5oOr5 o. The other feldspar is a high-temperaART. 15 IN U.S. GBOL. SURVEY PROF. PAPER 475-B, PAGES B52-B55. 1963.

B52
0'CONNOR B53

116°45' 116-30' 116°00'

Mine Mountain 

SHOSHONE MOUNTAIN South

E f": ;"*% "" ® : JACKASS ®15 

Tdt>opaf> DIVIDE

Spring *'° '

FIGURE 15.1. Index map of the Nevada Test Site and vicinity showing the location of sections (solid circles) measured in 

the Piapi Canyon Formation in this study. The numbered sections refer to Hinrichs and Orkild (1961).
B54 GEOCHEMISTRY, MINERALOGY, AND PETROLOGY

EXPLANATION

Rainier Mesa Member composite sheet

Tiva Canyon Member cooling unit 

a

Tuffs of Spring Canyoncooling unit

Topopah Spring Membercooling unit

Quartz

Alkali 

feldspar

Plagioclase 

feldspar

FIGTTRE 15.2. Modal analyses of the Piapi Canyon Formation. Mineral contents are in volume percent of total 

phenocrysts calculated to 100 percent.

ture modification of plagioclase having an estimated 

composition near Ab7oAn3 o. Quartz makes up less than 

5 percent and generally less than 2 percent of the 

phenocrysts.

Mafic-minera] phenocrysts occur only sparingly in 

the main part of the Topopah Spring Member. Biotite 

is the most abundant mafic mineral but magnetite, 

pyroxene, amphibole, sphene, and tscheffkinite (chevkinite) also are found. The mafic minerals increase in 

abundance in the upper flow of the cooling unit (fig. 

15.3).

The Spring Canyon cooling unit is mineralogically 

similar to the upper part of the Topopah Spring Member. The Spring Canyon cooling unit contains biotite 

throughout, however, and is further characterized by 

abundant rock-fragment inclusions.

The Tiva Canyon Member contains no quartz phenocrysts in most of the thin sections studied. Quartz 

occurs in this cooling unit, as in the other welded tuffs 

studied, as granophyric growths within pumice cavities. 

Anorthoclase (Ab65Or35 ) makes up nearly all of the 

phenocrysts in the main part of this cooling unit. 

Some variation in the composition of the alkali feldspar 

phenocrysts, particularly toward NaAlSi3O8 -rich feldspars, is noted. This variation is strongest in the 

uppermost flows of the Tiva Canyon Member in which 

plagioclase feldspar is relatively more abundant in the

UNIT LITHOLOGY MINERALOGICAL VARIATION

^ «^ ¥

= o Q..E

CD p"

; . ^Plagioclase 

\ ' < feldspar

/Biotite and\ 

pyroxene L

^Quartz

50

Volume percent 

of total phenocrysts

100

EXPLANATION FEET 

-0

Densely Vitrophyre 

welded tuff

FIGURE 15:3. Measured section of the Piapi Canyon Formation in Spring Canyon, Nevada Test 'Site, showing the mineralogical variation of the tuffs. Solid lines bound 

mappable units, short dashed lines bound ash flows within 

the cooling units.
0'CONNOR B55

modal analyses. The optic angle of the alkali feldspar 

in the Tiva Canyon Member is generally near 45°. 

The plagioclase feldspar phenocrysts are strongly resorbed and replaced by anorthoclase. Eeplacement 

phenomena are not as abundant in the plagioclase 

phenocrysts of the Topopah Spring Member and tuffs 

of Spring Canyon.

The Eainier Mesa Member has the most variable 

mineralogy of the tuffs of the Piapi Canyon Formation. 

This composite sheet generally contains abundant 

quartz, alkali feldspar, and plagioclase feldspar phenocrysts. However, near the margin of the Timber Mountain caldera, thin quartz-poor densely welded ash flows 

occur within the Eainier Mesa (fig. 15.3).

The lowest cooling unit of the Rainier Mesa Member 

has the largest percentage of quartz phenocrysts (35 

to 50 percent of the total phenocrysts). The alkali 

feldspars from this cooling unit have a small optic 

angle (2F*=28°) and a relatively high KAlSi3O8 

content (Or61Ab39 ). Alkali feldspars from younger 

cooling units in this member have larger optic angles 

=(as large as 45°) and lower KAlSi3O8 contents. The 

quartz-poor ash flows of this member contain an alkali 

feldspar with a composition near Ab61Or39 . Plagioclase 

in specimens of welded tuff from the lower part of the 

Eainier Mesa Member have a composition between 

Ab78An22 and Ab85An15 . The KAlSi3O8 content of the 

plagioclase was ignored in determining these values. 

Mantling and replacement of plagioclase by alkali 

feldspar occur in the welded tuffs of the Eainier Mesa, 

particularly the quartz-poor flows.

Mafic minerals are concentrated in at least two of the 

upper cooling units of the Eainier Mesa Member 

(fig. 15.3).

CONCLUSIONS

The quartz-alkali feldspar-plagioclase feldspar ratios of the phenocrysts are the most useful petrographic characteristics found for correlating the units 

in the welded tuffs of the Piapi Canyon Formation. 

The precise mineral percentages are more characteristic 

of individual ash flows than of the entire cooling units, 

but a general correlation can usually be observed within 

the cooling units.

The percentage of the rock consisting of phenocrysts 

varies widely with the degree of welding and with the 

ash flow involved. For this reason the modal analyses 

are plotted as percentage of total phenocrysts rather 

than percentage of total rock.

The percentage of mafic-mineral phenocrysts in the 

tuffs varies widely within the cooling units. Generally 

there is an increase in abundance of mafic minerals in 

the upper part of the cooling units.

Petrographic examination reveals that Hinrichs' and 

Orkild's (1961) identification of the Tiva Canyon 

Member at Jackass Divide (loc. 13, fig. 15.1), Four 

Corners Butte (14), Yucca Pass (16), Massachusetts 

Mountain (17, 18), French Peak (19), and Scarp Canyon (20) was in error. Specimens of this unit contain 

abundant quartz, sanidine, and plagioclase phenocrysts. 

The sanidine has a small optic angle (2Fa.=28°) and a 

relatively high KAlSi3O8 content (Or6 iAb39 ). These 

features are characteristic of the lowest cooling unit of 

the Eainier Mesa Member, not the Tiva Canyon Member. The Tiva Canyon Member is now known to extend 

approximately 4 miles east and 1 mile south of the Jackass Divide area; it was not recognized in earlier work 

in this area because of complex faulting.

REFERENCES

Hinrichs, E. N., and Orkild, P. P., 1961, Eight members of the 

Oak Springs formation, Nevada Test Site and vicinity, Nye 

County, Nevada: Art. 327 in U.S. Geol. Survey Prof. Paper 

424-D, p. D96-D103.

Poole, F. G., and McKeown, F. A., 1962, Oak Spring Group of 

the Nevada Test Site and vicinity, Nevada: Art 80 in U.'S. 

Geol. Survey Prof. Paper 450-C, p. C60-C62.

Schairer, J. F., Smith, R. L., and Chayes, Felix, 1956, Refractive indices of plagioclase glasses : Carnegie Inst. Washington Year Book 55, p. 195-197.

Smith, R. L., 1960, Zones and zonal variations in welded ash 

flows: U.S. Geol. 'Survey Prof. Paper 354-F, p. 149-159.

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

light of experimental istudies in the system NaAlSiaOgKAlSisOs-SiOa-HsO: Geol. Soc. America Mem. 74, p. 153,

Wilcox, R. E., 1959, Use of spindle stage for determining refractive indices of crystal fragments: Am. Mineralogist, v. 

44, p. 1272-1293.

Williams, P. L., 1960, A stained, slice method for rapid determination of phenocryst composition of volcanic rocks: Am. 

Jour. ScL, v. 258, p. 148-152.
Article 16

POTASSIUM-ARGON AND LEAD-ALPHA AGES

FOR STRATIGRAPHICALLY BRACKETED PLUTONIC ROCKS

IN THE TALKEETNA MOUNTAINS, ALASKA

By ARTHUR GRANTZ; HERMAN THOMAS, T. W. STERN, and NOLA B. SHEFFEY, 

Menlo Park, Calif./ Washington, D.C.

Abstract. A tie between radioactivity and stratigraphic 

time scales resulted from dating the Kosina 'batholith. Two 

radiogenic methods indicate the batholith is 160-165 million 

years old; geologic mapping suggests emplacement between 

Toarcian (Early) and Oxfordian (Late Jurassic) time, and 

perhaps close to the Early-Middle Jurassic boundary.

Six potassium-argon ages of biotite and four leadalpha ages of zircon, determined for a stratigraphically bracketed pluton in the eastern Talkeetna 

Mountains of south-central Alaska, indicate that the 

pluton is about 160-165 million years old. The pluton intrudes rocks of Sinemurian to Toarcian (Early 

Jurassic) age and was the source of many boulders in 

conglomerate of Oxfordian (Late Jurassic) age. The 

geologic record further indicates that emplacement may 

have begun very early in Middle Jurassic time. The 

data thus suggest a tie between the Toarcian-Oxfordian 

interval (and perhaps the Early-Middle Jurassic 

boundary) on the stratigraphic time scale and about 

160-165 million years on the radioactivity time scale.

GEOLOGIC ENVIRONMENT

The dated rocks are from the Oshetna Kiver drainage of the eastern Talkeetna Mountains. Half of them 

are from the plutonic rocks shown in the northwest 

part of figure 16.1 rocks which occupy a large area in 

the drainage of Kosina Creek to the north of the area 

of figure 16.1 and for convenience are called the Kosina 

batholith. The other half are from boulders in the 

Naknek Formation near the Little Oshetna Kiver (see 

table 16.1 and fig. 16.1).

The Talkeetna Formation, which is intruded by the 

Kosina batholith, is widespread in the southern Talkeetna Mountains and consists of a thickness of 

perhaps 2 miles or more of predominantly marine sedimentary and volcanic rocks. It contains ammonites 

(identified by K. W. Imlay, written communications, 

1961 and 1962) which are representative of parts of 

the Sinemurian, Pliensbachian, and Toarcian (including upper Toarcian) Stages of the Early Jurassic. 

Pectens of the genus Weyla from the Talkeetna (identified by S. W. Muller, oral communications, 1961 and 

1962) belong to species which range through these same 

stages. Three of the fossil localities are between 1 

and 2 miles from the batholith (fig. 16.1). Locality A 

(USGS Mes. loc. 28659), in hornfelsed beds intruded 

by an apophysis of the batholith, contains Weyla and 

is therefore of Early Jurassic age; locality B (USGS 

Mes. Iocs. 25938, 28660, 28661, 28662, and 28663), in 

beds within three-quarters of a mile of another apophysis, contains the ammonites CrucUobiceras (two species), Acanthopleuroceras, and Radstockiceras, and the 

pecten Weyla dufrenoyi d'Orbigny an association 

which is of early Pliensbachian age; locality C (USGS 

Mes. loc. 26722) contains Crucilobiceras of late Sinemurian to earliest Pliensbachian age.

The Naknek Formation, deposited in the Matanuska 

geosyncline from early Oxfordian to late Kimmeridgian or early Portlandian time, contains abundant 

fresh-appearing plutonic clasts which we think originated in the Kosina batholith. These clasts occur in 

conglomerate beds at the base of the formation and at 

several higher levels. The beds contain cobbles and 

boulders, and some attain a thickness of 1,000 to 1,500 

feet; yet, within a few miles south of their northernmost outcrops, they lens out into siltstone and shale. 

The boulders of plutonic rock that were dated came 

from about 400 feet above the base of such a conglomerate at the bottom of the Naknek Formation (fig. 

16.1). This conglomerate lenses out southward into

ART. 16 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B56-B59. 1968.

B56
GRANTZ, THOMAS, STERN, AND SHEFFEY B57

148°00' 147°30'

EXPLANATION

5 MILES

Geology in north third after Capps 

(1940, pi. 2), and 0. J. Ferrians and 

J. R. Williams (written communication, 1954); in south part after 

Grantz (1960)

Naknek Formation of 

Late Jurassic age

Jns, siltstone and shale 

Jnc, conglomerate above

base of formation 

Jnbc, conglomerate at

base of formation

Kosina batholith of 

Jurassic age

Talkeetna Formation

Volcanic and sedimentary rocks 

of Early Jurassic age. Includes satellitic masses of 

plutonic rocks

Contact

Dashed where approximately 

located

Fault

Dashed where approximately 

located

la 

X

Sample locality

Early Jurassic fossil locality

FIGURE 16.1. Generalized geologic map of Oshetna River area, Alaska, showing location of samples dated by K-Ar and

Pb-a methods, and fossil localities.

TABLE 16.1. Source of dated samples

No. on fig. 16.1 and 

tables 16.2 and 16.3

2.. ...................

S.. ...................

i. ..-.-...-...__..-.

5

6... 

Field No.

59AGzM26 

TO A fl<7~M"iV7

TO A n<7~M~'yi

59AGzM25-I -

59AGzM25-II -----

59AGzM25-III. --

Rock type Source

Kosina batholith, 4,000 ft from its southeast contact. 

Kosina batholith, 2,500 ft from its southeast contact. 

Kosina batholith, 3,000 ft from its southeast contact. 

Rounded boulders from conglomerate unit

(Jnbc on fig. 16.1) at base of Naknek Formation (Late Jurassic). Boulders collected

about 400 ft above base of the conglomerate at 

place where it is about 1,500 ft thick. Size 

of boulders in outcrop face: I, 2X4}^ ft; II, 

1X2 ft; III, 1X2 ft.

Locality

Talkeetna Mountains (A-3) quadrangle, lat 62°12'50" N., long 148°06'- 

35" W. 

Talkeetna Mountains (B-2) quadrangle, lat 62°21'17" N., long 147'49°- 

12" W. 

Talkeetna Mountains (B-2) quadrangle, lat 62°21'22" N., long 147°49'- 

18" W. 

Talkpetna Mountains (A-2) quadrangle, lat 62°09'52" N., long 147°44'-

40" W.

siltstone containing Cardioceras martini Reeside of 

early Oxfordian age and Buchia concentrica (Sowerby) of late Oxfordian and early Kimmeridgian age 

(K. W. Imlay, written communication, 1953).

A reconnaissance study of the Kosina batholith was 

made by Chapin (1918, p. 42-43 and pi. 2), who reported that the intrusive body is dominantly quartz 

diorite. However, G. D. Eberlein (oral communication, 1962), who also made a reconnaissance study of 

the area, has stated that in its central part the batholith 

is.a composite body consisting of granodiorite (with 

biotite the predominant ferromagnesian mineral) as 

well as of rocks that are more mafic and more salic than 

granodiorite. These include quartz monzonite containing both muscovite and biotite, amphibolite paragneiss, 

and tactite. In the recent study only the southeast
B58 GEOCHRONOLOGY

margin of the batholith was visited. This area is richer 

in hornblende than the central part and is composed 

mainly of hornblende-biotite granodiorite, much of 

which borders on quartz diorite, some on quartz monzonite. (The modal classification of granitic rocks 

adopted is that of Moore, 1959, p. 198, which is modified from Johannsen, 1931.) Quartz diorite and diorite are also present. The approximate composition 

of the granodiorite is quartz, 25-30 percent; plagioclase 

(mainly sodic andesine), 35-50 percent; orthoclase 

(largely interstitial or intergrown with quartz), 5-15 

percent; biotite (partly altered to chlorite), 10-15 percent; and hornblende, 10 percent. The modes cited 

here and below are based on point counts mainly by 

W. L. Griffin and M. C. Blake, of the U.S. Geological 

Survey. Minor accessories include opaque minerals, 

apatite, zircon, and sphene.

The three boulders of plutonic rock from the Naknek 

Formation that were dated are hornblende-biotite 

granodiorite and quartz diorite. Their approximate 

composition is quartz, 15-30 percent; sodic andesine, 

35-55 percent; orthoclase, <5-15 percent; and biotite 

and hornblende, each about 10 percent. Opaque minerals, apatite, sphene, and zircon are minor accessories. 

The boulders are thus similar to the dominant rocks 

of the south margin of the Kosina batholith. The similarity suggests that the boulders came from the adjacent batholith; this idea is strengthened by the similarity in radioactivity ages of the boulders and the south 

margin of the batholith. However, on the basis of our 

present knowledge of the Kosina batholith and the 

plutonic rocks in south-central Alaska generally, their 

identity cannot be unequivocally established.

GEOLOGIC AGE OF THE KOSINA BATHOLITH

The Kosina batholith appears to have been emplaced 

between Toarcian and Oxfordian time, but the field 

data do not pinpoint the time of intrusion. However, 

possible times of emplacement are suggested by unconformities within the sedimentary section of the nearby 

Matanuska geosyncline (see fig. 16.2). The rocks of 

the geosyncline, which are dominantly marine, crop out 

along the southern margin of the Talkeetna Mountains 

and were derived from the area immediately to the 

north. They rest imconformably upon the Talkeetna 

Formation and range in age from early Bajocian (Middle Jurassic) to Maestrichtian (Late Cretaceous).

Chapin (1918, p. 43) thought that emplacement of 

the pint on occurred between deposition of the Chinitna 

Formation (Callovian) and the imconformably overlying Naknek Formation (Oxfordian), and perhaps 

accompanied the uplift that ended Chinitna sedimentation. Unconformities also occur beneath beds of late

Bathonian age and beneath beds of early Callovian age. 

However, more significant than these is the unconformity between the Talkeetna Formation and the lower 

Bajocian rocks at the base of the section deposited in 

the Matanuska geosyncline. The tectonic event marked 

by this unconformity ended widespread Early Jurassic 

(Talkeetna Formation) deposition which had been 

characterized by volcanism in the Talkeetna Mountains 

area. Within the domain of this deposition, a provenance area was created at this time in the Talkeetna 

Mountains, and a nonvolcanic geosyncline (Matanuska) 

was created in the area to the south. The Kosina batholith may have been intruded at about the ToarcianBajocian time boundary (Early-Middle Jurassic), because the unconformity at this boundary represents the 

major tectonic event of the region in the ToarcianOxfordian interval.

RADIOGENIC AGE DETERMINATIONS

The mean of 3 K-Ar dates (table 16.2) on 2 samples 

of the Kosina batholith is 165 m.y., and the mean of 3 

K-Ar dates on 3 boulders of plutonic rock from the 

Naknek Formation is 158 m.y. The mean age of the 

5 samples dated is 162 m.y. The mean of 3 Pb-a. dates 

(table 16.3) of samples for which K-Ar dates are available is 165 m.y. Each is compatible, within the limits 

of analytical error, with the corresponding K-Ar date. 

One sample, for which no K-Ar date could be obtained, 

yielded a Pb-a date of 125 ±15 m.y. The dates

Stratigraphic limits 

for emplacement of 

southeast part of 

Kosina batholith

Series

a

1a 

n

T3 $

5 _5

Jurassic Lower 

Stage

Purbeckian

Portlandian

Kimmeridgian

Oxfordian

Callovian

Bathonian

Bajocian

Toarcian

Pliensbachian

Sinemurian

Hettangian

.Deposits of the Matanuska 

geosyncline

Naknek Formation

/..Conglomerate (Jnbc) at base of 

1 Naknek Formationfrom which the 

1 Jdated boulders were collected 

 < First common plutonic detritus 

f Chinitna Formation

1 Sandstone 

>Tuxedni Formation

Matanuska geosyncline

> Talkeetna Formation in 

the Talkeetna Mountains

'

FIGURE 16.2. Stages of the Jurassic showing stratigraphic position of the Naknek Formation and the plutonic boulders that 

were dated, and the inferred time-stratigraphic limits of the 

southeast margin of the Kosina batholith.
GRANTZ, THOMAS, STERN, AND SHEFFEY B59

by the Geological Survey are reported here for the first 

time.

TABLE 16.2 Potassium-argon ages of Motile from southeast margin of Kosina batholith and from boulders of similar lithology 

in the Naknek Formation

[Age of sample la from Evernden and others (1961, p. 88, sample KA.431); all other 

determinations by U.S. Geological Survey]

No. on fig. 

16.1

KI

(percent)

K« (ppm) Radiogenic 

Ar<° (ppm)

Radiogenic

Ar" 

K*»

Calculated 

age (millions of 

years)

Samples from Kosina batholith

la_-----._

lb_-----_-

3. --------

3.63

3.39 

5.60

4. 10 

6. 76

0. 0394 

. 0710

0. 00961 

.0105

169 

155 

170

Samples of boulders from Naknek Formation

4___ -----

5---.--.__

6... ------

3.31

4. 77 

2. 00

4. 00 

5.76 

2. 42

0. 0393 

. 0551 

. 0236

0. 00982 

. 00956 

. 00975

160 

155 

160

Decay constants:

Ae = 0.589XlO- 10 per yr.

X/S =4.76X10-10 per yr. 

Abundance ratio: K40/K=0.0118 atomic percent.

i Determinations by P. Elmore and I. Barlow.

TABLE 16.3. Lead-alpha ages of zircon from southeast margin of 

Kosina batholith and from boulders of similar lithology in the 

Naknek Formation

[Alpha-activity measurements by T. W. Stern; spectrographic analyses of lead by 

NolaB. Sheffey]

No. on fig. 16.1

Alpha counts 

per milligram 

per hour

Pb (ppm)

Calculated age 

(millions of 

years)i

Samples from Kosina batholith

2... ____...__...-__

3.--. ---_-___-_-__-

300 

219

15 

(14, 16) 

14. 5 

(13. 5, 15. 5)

125±15 

165 ±20

Samples of boulders from Naknek Formation

4-.____. _...______.

6__--___-.-__-__-__

100 

130

7.2 

7.9

(8. 2, 7. 6)

180 ±20 

150±15

i Pb-« ages (rounded to nearest 5 million years) were calculated from the equation: fi pi.

t = > where i is the calculated age, in millions of years; Cis a constant based upon 

a

the ratio Th/U, which was assumed to be 1 in the zircon samples dated, and has a 

value of 2,485; Pb is the lead content, in parts per million; and a is the alpha counts 

per milligram per hour.

The scatter in the K-Ar dates and the difference between dates la and Ib may be related to the comparatively low K content of the biotites dated. This is due 

to the partial alteration of biotite to chlorite in the 

samples dated; however, the K-Ar dates obtained are 

thought to be approximately correct in spite of the 

alteration.

The difference between the mean of the batholith 

samples (165 m.y.) and the mean of the boulder samples 

(158 m.y.) is also within the limits of accuracy and may 

have no special significance. However, the difference is 

possibly due in some way to the cycle of erosion, transportation, and weathering to which the boulders were 

subjected in Late Jurassic time, and to the long period 

during which they remained buried in marine sedimentary rocks. Alternatively, it might be due to differences 

in the cooling history between the part of the batholith 

sampled in outcrop and the part represented by the 

boulder samples. Nevertheless, it seems best at present 

to consider all of the samples as one group with a mean 

age of 162 (160-165) m.y. This age for the emplacement of the southern margin of the Kosina batholith 

establishes a tie between the radioactivity and the stratigraphic time scales in the Toarcian-Oxfordian stratigraphic interval. The age of the emplacement may 

even approximately date the Toarcian-Bajocian boundary, and thus the boundary between Early and Middle 

Jurassic.

REFERENCES

Capps, S. R., 1940, Geology of the Alaska Railroad region: U.S.

Geol. 'Survey Bull. 907, 201 p. 

Chapin, Theodore, 1918, The Nelchina-'Susitna region, Alaska:

U.S. Geol. 'Survey Bull. 068, 67 p. 

Evernden, J. F., Curtis, G. H., Obradovich, J., and Kistler, R.,

1961, On the evaluation of glauconite and illite for dating

sedimentary rocks "by the potassium-argon method:

iGeochim. et Cosmochim. Acta, v. 23, p. 78-99. 

Grantz, Arthur, I960, Geologic map of Talkeetna Mountains

(A-2) quadrangle, Alaska and the contiguous area to the

north and northwest: U.S. Geol. 'Survey Misc. Geol. Inv.

Map 1-313, scale 1: 48,000. 

Johannsen, Albert, 1931, A descriptive petrography of the

igneous rocks, v. 1: Chicago, 111., Univ. Chicago Press,

342 p. 

Moore, J. G., 1959, The quartz diorite 'boundary line in the

western United States: Jour. Geology, v. 67, p. 198-210.

680231 O 68
Article 17

RADIUM MIGRATION AND ITS EFFECT ON THE APPARENT AGE 

OF URANIUM DEPOSITS AT AMBROSIA LAKE, NEW MEXICO

By HARRY C. GRANGER, Denver, Colo.

Abstract. Considerable radium (Ra226 ) has migrated out of 

the uranium ores and has been partly reconcentrated in barite, 

cryptomelane, and along the surfaces of mudstone bodies. The 

loss of Ra226 , which ultimately decays to Pb208, indicates that 

the apparent Pb^/Pb208 age of the ores would be too great and 

that the Pb^/U288 age would be too small.

Radiochemical analyses of uranium from deposits at 

Ambrosia Lake, N. Mex., indicate migration of various 

isotopes of the uranium-decay series. Discordant ages 

of these and other uranium deposits probably can be 

attributed partly to the migration of radium.

The possible loss of radium from uranium, ore has 

generally been considered less likely as a cause of 

anomalous age determinations than other processes such 

as loss of radon or addition of older radiogenic lead. 

A few investigators (for example, Kennedy, 1961, p. 

121-122) have suggested that uranium ores may be deficient in lead because of radium migration, but few 

data have been presented in support of this idea.

Migration of members of the U238 and U235 radioactive-decay series during the last few hundred thousand years can be evaluated by comparing the equilibrium relations among these members (Rosholt, 

1958). Because the ratio of U238 to U235 (about 138:1) 

is almost constant in natural samples, the radioactive 

equilibrium (secular equilibrium) of members of one 

series can be compared directly with members of the 

other series. The two series are shown below:

Half lives of major isotopes of the U238 and U235 series 

U238 series U236 series

Uranium (U238 ) 4.5 Uranium (U235 ) 7.1

X109 yrs X108 yrs 

Uranium (U234 ) 2.5 X Protactinium (Pa231 )

105 yrs 3.4X104 yrs 

Ionium (Th230 ) 8X Radium (Ra223 ) 11.2 days

104 yrs Lead (Pb207 ) stable 

Radium (Ra226 ) 1,620

yrs 

Lead (Pb206 ) stable

In theory, when a uranium deposit is somewhat more 

than 1 million years old, all members of each decay 

series should be approximately in radioactive equilibrium. In deposits less than 1 million years old the 

concentration of daughter products in the U238 series 

will be less than that required for radioactive equilibrium, and in deposits less than about 200,000 years old 

the concentration of daughter products in the U235 

series will be less than that required for radioactive 

equilibrium. The concentration of a daughter product is expressed as the amount of uranium (U234, U235, 

and U238 ), in percent, in radioactive equilibrium (secular equilibrium) required to produce the measured 

amount of daughter product. These equivalent concentrations of the various radioisotopes are denoted by 

e (for example, ePa231, called equivalent protactinium) 

to distinguish them from eU, which is the amount of 

uranium and its daughter products in radioactive equilibrium that would emit the measured amount of 

radioactivity.

Migration of uranium or its daughter products can 

cause addition or subtraction of certain members of the 

decay series so recently that members have had insufficient time to attain radioactive equilibrium. To calculate the magnitude of migration of members of the 

decay series, a reliable index is needed for comparison 

of the concentrations of individual members. Although uranium is the best index for theoretical or 

laboratory work 'because it supports the radioactive 

equilibrium, it is not necessarily a satisfactory index 

for natural samples because of its relative mobility.

Protactinium (Pa231 ) probably is the radioisotope 

least likely to migrate from its site of origin under 

natural conditions, and it is more abundant than most 

other daughter products because it has a relatively 

long half life. The movement of Pa231 is limited by 

hydrolysis or by reaction with ions in the ground water 

which tend to form precipitates or colloids, and, perART. 17 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B60-B63. 1963.

B60
GRANGER B61

haps, by ion exchange with organic substances (Elston, 

1954). For these reasons, Pa231 is the best index for 

the evaluation of excesses or deficiencies of the other 

members of the U235 series.

Eosholt (1960, 1961) proposed that both Pa231 and 

Th230 (ionium) can be used as indices, rather than 

uranium, because neither migrates in measureable quantities from where they are produced by decay of the 

parent isotopes. Pa231 and Th230, therefore, should be 

present in constant proportions for a series of samples 

cut from a single ore body in a uniform simple environment, This proportion is not constant at Ambrosia 

Lake even for adjacent samples; thus, either Pa231 or 

Th230 must have migrated or one of the parent isotopes 

must have migrated.

Within the prefault ores (Granger and others, 1961) 

the ePa231 generally exceeds the total U concentration, 

which indicates either that Pa231 has been added or 

uranium subtracted. Probably uranium has been removed, for protactinium is much less soluble than 

uranium under most natural conditions.

The correlation between U and eTh230 in the orebearing rocks at Ambrosia Lake (fig. 17.1) is excellent, 

suggesting that U238 and Th230 have had time to reach 

radioactive equilibrium. If uranium has moved, as 

suggested above by its radioactive-equilibrium relation 

to Pa231 , the good correlation between the concentrations of uranium and eTh230 suggests that Th230 is also 

mobile in this environment, to about the same degree 

as uranium.

10.0

i.o

o.i

o.oi

o.ooi

U

0.001 0.01 0.1 1.0 10.0 

EQUIVALENT IONIUM (eTh230 ), IN PERCENT

FIGURE 17.1. Correlation between uranium and equivalent 

ionium (eTh230 ) in M samples from Ambrosia Lake district, 

New Mexico.

Chemical and radiochemical analyses were made of 

94 samples of ore-bearing sandstone from the Ambrosia 

Lake district. Eadiochemical analyses were made by 

John N. Eosholt, Jr.; elJ analyses were made by Lorraine M. Lee; U analyses by the volumetric method 

were made by Henry H. Lipp; and U analyses by the 

fluorometric method were made by Dorothy L. Ferguson. Much of the interpretation is based on discussions with John N. Eosholt, Jr., and Ealph S. 

Cannon, Jr.

The samples range in uranium content from 0.001 to 

15.2 percent, Most samples are from suites that extend 

across ore layers and a foot or more into the enclosing, 

relatively barren sandstone on either side. Both prefault and postfault ores (Granger and others, 1961) 

contain coffinite, and most of the samples are essentially 

unoxidized. The distribution of radioactive daughter 

products is similar in both prefault and postfault ore 

bodies, although the postfault ore bodies cannot be 

sampled with the same geometrical precision as the prefault ore bodies because of greater irregularities in 

shape. The following interpretations of radium migration are based largely on data from prefault ore, but 

they should apply also to the postfault ore.

Data from 38 of the samples, which constitute 5 

suites cut across prefault ore layers, indicate that the 

ore is about 20 percent deficient in uranium relative to

3.0 

2.0 

1.0 

0.5-1

0.5-

ORE

1.0-

0.5-

TJ ePa^tTh2 U

SANDSTONE ADJACENT TO ORE

FIGURE 17.2. Abundance of various radioisotopes in uranium 

ore and rock adjacent to ore bodies at Airibrosia Lake, 

N. Mex., relative to the equivalent protactinium, which is 

arbitrarily set at 1.0. U, uranium; ePa231, equivalent 

protactinium; eTh230, equivalent ionium; eR226, equivalent 

radium.
B62 GEOCHRONOLOGY

Pa231 (fig. 17.2). The first few feet of relatively barren 

sandstone on either side of the ore is about 47 percent 

deficient in uranium. The ore and adjacent rock taken 

together are about 21 percent deficient in uranium.

In the same group of samples, Ra226 is nearly 41 percent deficient in the ore and is about 220 percent in 

excess relative to Pa231 (fig. 17.2) in the sandstone adjacent to the ore. The deficiency in the ore represents 

a much greater amount of Ra226, because of the high 

uranium content of the ore, than the excess in the relatively barren sandstone adjacent to the ore; the average 

deficiency of Ra226 in the ore and adjacent rock is about 

32 percent.

Some of the Ra226 lost from the ores probably can 

be traced to samples that have an anomalously high 

radioactivity or radium or lead content. For example, 

Granger and others (1961) describe high concentrations 

of Ra226 in radioactive barite in and near some of the 

bodies. The presence of such concentrations of Ra226, 

not associated with parent isotopes, is strong evidence 

of relatively recent migration of radium. The most 

likely sources of the anomalous radium in radioactive 

barite are the nearby or enclosing ore bodies.

In like manner, strongly to weakly radioactive cryptomelane that partly replaces mudstone galls has been 

found in several mines. Because of its low uranium 

content and high barium content, the radioactivity 

probably is attributable largely to radium substituting 

chemically for barium. Furthermore, the strongly 

radioactive cryptomelane is relatively enriched in lead, 

suggesting that the lead is mostly radiogenic and is derived from radioactive decay of part of the radium.

Most of the radium concentrated outside the ore 

bodies is probably contained in mudstone. Gammaray logs of drill holes throughout the Ambrosia Lake 

area show anomalous radioactivity associated with both 

the upper and lower surfaces of most mudstone lenses. 

Analyses of the outer 1 or 2 inches of mudstone layers 

near ore commonly disclose anomalously high radioactivity coupled with an abnormally high lead content. 

Within mudstone layers, however, the radioactivity is 

essentially in balance with the uranium content, and 

the lead content is low.

Migration of uranium or daughter products, whether 

modern, sporadic, or continuous throughout the history 

of a rock, can affect significantly the apparent age of 

individual rock samples as determined by uraniumand lead-isotope ratios. If the relative deficiency of 

uranium in the Ambrosia Lake ores is due entirely to 

modern loss, and not due to continuous or sporadic loss 

since the original deposition, the apparent Pb206/U238 

and Pb *oyU885 ages should be too great. The Pb207/

Pb206 ratio, however, should represent a reasonably accurate absolute age if the decay series maintained radioactive equilibrium until the time of uranium loss.

Loss of radium does not affect the relative uranium 

content of the ore but does cause an eventual deficiency 

of lead. Because of its much longer half life, Ra226 

is more likely to escape from the ore and migrate for 

some distance than Ra223 ; thus, proportionately less 

Pb206 than Pb207 will be produced in the ore body. Accordingly, the Pb206/U238 age will appear to be too small 

and the Pb207/Pb206 age too great, but the Pb207/!!235 

age should be reasonably accurate.

The Pb206 deficiency in Ambrosia Lake ores resulting 

from long-continued loss of Ra226 is, by inference, relatively greater than the uranium deficiency resulting 

from modern loss. Therefore, within the ore, Pb207/ 

Pb206 ages should appear to be much too great, Pb207/ 

U235 ages somewhat too great, and Pb206/U238 ages too 

small.

In the rock adjacent to the ore, however, uranium apparently has been lost, but the excess of Ra226 predicts 

an excess of Pb206. Hence, the Pb206/U238 age should 

appear to be much too great and the Pb207/U235 age 

somewhat too great, but the Pb^/Pb206 age should appear to be much too small.

Until the relative migration of uranium and its 

daughter products can be better evaluated, isotope age 

determinations on samples from the Ambrosia Lake 

district cannot be considered accurate. Because of the 

apparent paucity of nonradiogenic lead,1 a simple ratio 

between total lead and total uranium may indicate the 

true age of mineralization just as accurately as the 

apparent isotope age. The Pb/U ratio computed from 

semiquantitative spectrographic analyses for lead and 

chemical analyses for uranium of 27 high-grade (> 1.0 

percent U) prefault ore samples from various mines 

at Ambrosia Lake suggests an age of about 55 million 

years.

Somewhat greater ages are indicated if low-grade ore 

and rock adjacent to ore bodies are included in the age 

calculations. The total lead and uranium contents of 

75 samples in 9 suites extending across prefault ore layers and the adjacent wallrocks suggest an average age 

of about 100 million years. Even this apparent age, 

however, probably is less than the true age as only a 

part of the radiogenic lead lost from the ore because 

of radium migration is contained in the adjacent rock.

iMore than 100 rock samples analyzed by semiquantitative spectrographic methods contained too little lead to report «0.0007 percent 

Pb). Many of these were postfault ore samples containing up to 4.86 

percent uranium, indicating that postfault ore is of very recent age and 

that the host rock was essentially devoid of lead before postfault ore 

deposition.
GRANGER B63

REFERENCES

Elston, R. E., 1954, The chemistry of protactinium, in Seaborg,

G. T., and Katz, J. J., eds., Actinide elements: New York,

McGraw-Hill Book Co., p. 117. 

Granger, H. O., Santos, E. S., Dean, B. G., and Moore, F. B.,

1961, Sandstone-type uranium deposits at Ambrosia Lake,

New Mexico an interim report: Econ. Geology, v. 56, no. 7,

p. 1179-1209. 

Kennedy, V. C., 1961, Geochemical studies of mineral deposits

in the Lisbon Valley area, San Juan County, Utah: U.S.

Geol. Survey open-file rept., 150 p.

Bosholt, J. N., Jr., 1958, Radioactive disequilibrium studies as 

an aid in understanding the natural migration of uranium 

and its decay products, in United Nations Internat. Conf. 

Peaceful Uses Atomic Energy, 2d, Geneva, 1955, Proc., p. 

230.

 1960, -A study of uranium migration in sandstone-type 

ore deposits: Art. 21 in U.S. Geol. Survey Prof. Paper 400-B, 

p. B41-B42.

 1961, Uranium migration and geochemistry of uranium 

deposits in sandstone above, at, and below the water table, 

pt. 1, Calculation of apparent dates of uranium migration 

in deposits above and at the water table: Econ. Geology, 

v. 56, no. 8, p. 1392-1403.
Article 18

GRAVITY SURVEY OF THE GOLD MEADOWS STOCK, 

NEVADA TEST SITE, NYE COUNTY, NEVADA

By DON L. HEALEY and CARTER H. MILLER, Denver, Colo.

Work done in cooperation with the U.S. Atomic Energy Commission

Abstract. The Gold Meadows stock intrudes clastic Paleozoic 

rocks and is overlain by Tertiary volcanic rocks. The interpreted configuration of the intrusive was determined by threedimensional and two-dimensional analysis of gravity data. The 

stock is elongated to the northeast with steep slopes on the 

southwest, east, and northeast sides.

The Gold Meadows stock, which lies about 45 miles 

north of Mercury, Nev., is included in a regional 

gravity survey of the Nevada Test Site and adjacent 

areas that was begun in 1958. The geology of the area 

was mapped by Gibbons and others (1959) of the U.S. 

Geological Survey. Superposition of the gravity data 

on the geologic map offers a means to estimate the size 

and shape of the stock at depth. An aeromagnetic 

survey was flown in 1961, but the interpretation of that 

data is not yet completed.

The stock is composed of quartz monzonite of 

Cretaceous (?) to Tertiary (?) age that was intruded 

into older sedimentary rocks. It crops out over an 

area 2*4 miles long in a northeasterly direction and 

about % of a mile wide. Remnants of the sedimentary 

rocks are exposed on the east and south sides of the 

stock and dip into it. The remnants are quartzite of 

the Wood Canyon Formation of Early Cambrian and 

Early Cambrian (?) age, or of the Sterling Quartzite 

of Precambrian age. These sedimentary and intrusive 

rocks are overlain unconformably by welded and nonwelded tuff and rhyolite of Tertiary age and surface 

deposits of Quaternary age (fig. 18.1A). The tuff and 

rhyolite dip away from the stock at angles of 20° to 35°, 

suggesting that the quartz monzonite was a topographic 

high on the prevolcanic erosion surface (Gibbons and 

others, 1959, p. 16).

Approximately 70 gravity measurements were made 

to define the anomaly associated with the stock. The

data were reduced to sea level using an elevation factor 

which corresponds to a density of 2.67 g per cc, so 

that, as a first approximation, variations in the Bouguer anomaly values are proportional to the thickness 

of the low-density Tertiary and Quaternary deposits. 

The survey is referenced to an observed absolute gravity value at McCarran Field, Las Vegas, Nev. (Woollard, 1958). Terrain corrections were applied to a 

radial distance of 9 miles. A regional gradient of 

approximately 1 mgal per mile to the north was removed from the data prior to interpretation.

Density determinations indicate that the tuff and 

rhyolite have an average bulk density of 2.03 g per cc, 

the quartz monzonite a density of 2.61 g per -cc, and the 

quartzite an average of 2.62 g per cc. Since the bulk 

densities of the quartzite and quartz monzonite are 

very similar, the older rocks are considered as part of 

the Gold Meadows stock for the purposes of this article. 

An approximate density contrast of 0.6 g per cc between the stock and overlying rocks was used in the 

interpretation of the gravity data.

The complete Bouguer gravity data (fig 18.1 B] indicate that the anomaly over the Gold Meadows stock 

is roughly symmetrical and is slightly elongated to the 

northeast. A secondary nose trends northward off the 

map. The stock is nearly surrounded by a gravity low 

that slopes off regularly on three sides but forms a 

fairly well defined trough on the southeast.

The three-dimensional method described by Talwani 

and Ewing (1960) was used to compute the size and 

shape of the Gold Meadows stock at depth. As this 

method is not sensitive to small changes around the 

perimeter of the body, several two-dimensional profiles (after Dobrin, 1952, p. 96) were constructed to

ART. 18 IN U.S. GEOL.. SURVEY PROF. PAPER 475-B, PAGES B64-B66. 1963.

B64
HEALEY AND MILLER B65

EXPLANATION

(Cretaceous? to Tertiary?)

Quartzite 

(Cambrian? and Precambrian)

 U12b.07-2

 4958

Drill hole

Showing number, and elevation, 

in feet, of top of pre- Tertiary 

rocks

B

i

Structure contour

Drawn on top of undivided 

quartz monzonite and older 

rocks. Contour interval 500 

feet. Datum mean sea level

A _____________ A' 

Line of section

Gravity station

Complete Bouguer gravity 

Contour interval 5 milligals. 

All numbers are negative. Hachures on lowest co-ntours

NEVADA

\_Gold Meadows _.

\ . stock Nevada '

\ \ Test Site I

\ I JMercury

\

2 MILES

Las Vegas

°-Y~J36° 

\

\\

FIGURE 18.1. A, Generalized geologic and paleotopographic map of the Gold Meadows area; geology generalized 

from Gibbons and others (1960) ; B, Complete Bouguer anomaly map of the Gold Meadows area.
B66 GEOPHYSICS

Quaternary and 

Tertiary deposits

Quaternary and 

Tertiary deposits

2 MILES

FIGURE 18.2. Generalized geologic section showing top of undivided quartz monzonite and older 

rocks, as determined by analysis of gravity data. Line of section shown on figure 18.14.

supplement the three-dimensional interpretation and 

refine the computed configuration of the stock and its 

adjacent quartzite bodies.

Depth control was provided by three drill holes (fig. 

18.L4). The Tertiary rocks are 3,670 feet thick in 

Exploratory 1, 2,420 feet thick in Ul2b.07-2, and 2,110 

feet thick in U12.E.

The configuration computed for the stock is represented by paleotopographic contours superimposed on 

the geologic map in figure 18.1 A.

It is elongate parallel to the long axis of the quartz 

monzonite outcrop and has a steep southeast slope and 

a moderately steep northwest slope. A generalized 

geologic section (fig. 18.2) shows the computed contact between the Tertiary and Quaternary deposits and 

the undivided quartz monzonite and quartzite. No 

separation of the stock and the quartzite was attempted

because their similar densities render such a contact 

meaningless.

REFERENCES

Dobrin, M. B., 1952, Introduction to geophysical prospecting: 

New York, McGraw-Hill Book Co., Inc., 435 p.

Gibbons, A. B., Hinrichs, E. N., Dickey, D. D., McKeown, F. A., 

Poole, F. G., and Houser, F. N., 1959, Engineering geology 

of test sites in granite and dolomite at Gold Meadows, 

Climax, and Dolomite Hill, Nevada Test Site, Nye County, 

Nevada Preliminary report: U.S. Geol. Survey TEMR-884.

Gibbons, A. B., Hinrichs, E. N., Hansen, W. R., and Lemke, W., 

1960, Preliminary geologic map of the Tippipah Spring NW 

quadrangle, Nye County, Nevada: U.S. Geol. Survey TEI754, open-file report.

Talwani, Manik, and Ewing, Maurice, 1960, Rapid computation 

of gravitational attraction of three-dimensional bodies of 

arbitrary shape: Geophysics, v. 25, no. 1, p. 203-225.

Woollard, C. P., 1958, Results for a gravity control network at 

airports in the United States: Geophysics, v. 23, no. 3, 

p. 520-535.
Article 19 

DEPICTION OF SOIL-COVERED STRUCTURES BY INFRARED AERIAL PHOTOGRAPHY

By WILLIAM A. FISCHER, Washington, D.C.

Abstract. A series of anticlines and synclines beneath residual soils in east-central New Mexico is depicted by infrared 

photography. The photographs show a series of light and 

dark streaks parallel to structures mapped in adjacent areas. 

The positions of the areas of brightest tone correspond closely 

to the positions of anticlinal axes.

In a semiarid lowland area of little relief in eastcentral New Mexico (fig. 19.1), much of the bedrock 

is obscured by residual soils. Field geologic mapping 

of this area by K. J. Hackman and W. A. Fischer revealed that the dip of most exposed bedrock is related 

to local subsidence associated with the solution of limestone and gypsum, and not to regional tectonic structures. Thus, because of erratic dips and residual-soil

FIGURE 19.1. Location of the area in east-central New Mexico 

included in conventional aerial photograph (fig. 19.2) and in 

infrared aerial photograph (fig. 19.3), and distribution of 

major rock units..

cover, tectonic structures of probable regional significance are masked and cannot be mapped in parts of 

the area by conventional field mapping methods or 

from conventional aerial photography (fig. 19.2).

Infrared aerial photographs of the area were taken 

especially for this study by the U.S. Coast and Geodetic 

Survey on April 3,1960 (fig. 19.3). These photographs, 

which record energy reflected from the earth's surface 

having wavelengths of approximately 0.72 to 0.86 

microns, show a series of alternating light and dark 

streaks parallel to the regional structural trend. The 

bright streaks are visible on the exposures preceding and 

succeeding the photograph shown in figure 19.3 in the 

same flight line and also in part on 1 exposure taken at 

90° to the flight line and overlapping approximately the 

eastern third of the area shown in figure 19.3. Contrast 

between the bright streaks and the surrounding terrain 

diminishes toward the margins of all the photographs; 

this may relate to characteristics of the taking lens.

Figure 19.4 shows three curvilinear densitometer records representing film-density measurements along 

traverses A-A'9 B-B', and C-C' (fig. 19.3); the positions of -anticlinal axes as mapped in the field are also 

shown. Similar film-density measurements were made 

across a diapositive copy of the photograph shown in 

figure 19.3 at spacings of 1 centimeter. An aperture 

diameter of 1 centimeter was used in the densitometer to 

integrate density measurements over relatively large 

areas and thus reduce the effect of small shadows or 

cultural features. Density measurements were assigned 

arbitrary numerical values and plotted on the photograph. Lines were drawn connecting areas having 

equal tone values to produce figure 19.5, an isotonal map 

depicting the areas of brightest photographic tone and 

the positions of anticlinal axes and synclinal troughs as 

mapped in the field. The areas of brightest tone appear 

to be shifted to the right with respect to the anticlinal 

axes. This asymmetry in part relates to the time lag of 

the recorder and direction of scan of the densitometer

ART. 19 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B67-B70. 1963.

B67
B68 GEOPHYSICS

FIGURE 19.2. Conventional aerial photograph of the area. Residual soils obscure the bedrock in the lower left part of the 

area. Approximate scale 1:51,000. (Photograph by U.S. 

Department of Agriculture.)

FIGUEE 19.3. Infrared aerial photograph of approximately the 

same area shown in figure 19.2. A-A', B-B', and C-C' indicate location of densitometer traverses shown in figure 19.4. 

Approximate scale 1:77,000. (Photograph by U.S. Coast 

and Geodetic Survey.)

EXPLANATION

Axis of known anticline

FIGURE 19.4. Curvilinear densitometer records showing relative 

light transmission through a positive-film copy of photograph 

shown in figure 19.3. Location of densitometer traverses 

shown in figure 19.3. Positions of anticlinal axes plotted 

from field structural measurements.

Areas of brightest 

photographic tone

EXPLANATION

Axis of anticline

A'

B'

C'

Axis of syncline

FIGURE 19.5. Map showing areas of brightest tone on figure 

19.3, and positions of anticlinal and synclinal axes as mapped 

in the field.
FISCHER B69

(from left to right). The areas of brightest photographic tone, however, are elongate in a direction parallel to the regional structural trend and appear to correspond to anticlinal axes.

Interpretation of conventional aerial photographs 

suggests that in this area soils are thinner over the axes 

of anticlines than over the troughs of synclines. Limited field observation of soil thickness by Fischer and 

Hackman supports this interpretation and suggests that 

soils near the crests of anticlines are commonly less than 

3 inches thick. The distribution of areas believed overlain by thin soils corresponds closely with the brighter 

images on the infrared photography.

Measurements of the infrared reflectance of soils 

and grass from the test area are not available. However, samples of the limestone which underlies much of 

the area and relatively dry grasses collected from other 

localities were illuminated with a strong tungsten source 

and observed with a 'Snooperscope (Model M-2). A 

Snooperscope is an image-converting device that presents a visible image of infrared radiation of approximately the same wavelengths as those recorded on the 

infrared aerial photographs. Figure 19.6 is a photograph taken through the Snooperscope of two clumps 

of grass set in white paper sample boxes and a sample 

of the limestone oriented so that a weathered surface 

faces the camera. This photograph shows that the 

limestone reflects more infrared radiation of these 

wavelengths than the grass. This observation leads the 

author to believe that the relatively bright images appearing on the infrared aerial photographs relate to a 

relative sparsity of grass cover (greater percentage of 

soil and rock fragments exposed to vertical view) as 

contrasted to a relatively denser cover of grass in areas 

having thicker soils. Measurements by Krinov (1947, 

reflectance curves 62, 97, and 100) of similar terrain in 

Russia also show that some dry soils and rocks have a 

higher infrared reflectivity than some grasses.

Figure 19.7 is a photograph recording visible light 

reflected from the same objects shown in figure 19.6. 

The limestone reflects more visible light than the grass, 

but the contrast is not as great as that shown on the 

infrared photograph (fig. 19.6). This difference in 

contrast can best be seen by comparing the brightness 

of the limestone sample to that of the white sample 

boxes.

Variations in growth stage or moisture content of 

vegetation are probably not responsible for the variations in infrared radiation recorded on the infrared 

photographs. Measurements by Krinov (1947, reflectance curve 64) and observations by Ives (1939, p. 434) 

show that mature vegetation in lush growth reflects 

more infrared light than dry vegetation. Thus, if variFIGTJKE 19.6. Photograph taken through Snooperscope showing infrared radiation from a, semidry grass, and 6, limestone. Size of s~ample trays 6 inches by 6 inches.

FIGURE 19.7. Conventional photograph ©f objects shown in 

figure 1&.6.

ations in plant moisture or development were a significant cause of the observed tone differences, the tone 

distribution on the infrared photography of this area 

would probably have been reversed. That is, vegetation * 

covering areas underlain by the thinner soils would 

likely have been drier and (or) growth retarded and 

have reflected less infrared radiation (photographed 

with a darker tone) than vegetation growing on areas 

underlain by relatively thicker soils.

Field investigations by Fischer and Hackman reveal 

that the composition of the surface of the lowlands 

area shown in figure 19.3 is relatively uniform, and 

consists largely of residual soils derived from, underlying limestone. Fragments of sandstone and monzonite are present locally in the alluvium, and beds of 

caliche, and fragments of sandstone are present locally 

on the surface. The distribution of these materials,
B70 GEOPHYSICS

however, does not correspond to the distribution of 

tones on the infrared photographs and it is unlikely 

that tone distribution relates to composition of surface 

materials.

It is also unlikely that differences in moisture content of soils at the surface are a significant cause of the 

tone differences, because according to records of precipitation (U.S. Weather Bureau, 1960) the area was 

quite dry at the time of photography (April 3, 1960). 

No measurable precipitation was recorded at any of the 

5 most proximate weather stations for the 8 days preceding the time of photography. These same stations 

recorded totals of less than 0.2 inch of precipitation for 

the entire month of March (U.S. Weather Bureau, 

1960).

Study of these infrared photographs suggests that 

infrared photography may have application in structural studies of some semiarid areas. It will be necessary, however, to take additional infrared photographs 

of the area to ascertain whether any of the special conditions that prevailed at the time these photographs 

were taken (for example, dry surface, time of year, 

overexposure) are significant to the postulated relation 

between tone on infrared aerial photographs and 

structural features.

REFERENCES

Ives, R. L., 1939, Infrared photography as an aid in ecological

surveys: Ecology, v. 20, no. 3, p. 433-^39. 

Krinov, E. L., 1947, Spectral reflectance of natural formations,

chap. 2: Laboratoriia Aerometodov, Akad. Nauk SSSR,

Moscow, 271 p. [Natl. Research Council of Canada Tech.

Translation TT 439, translated by E. Belkov, 1953]. 

U.S. Weather Bureau, 1960, Climatological data New Mexico:

v. 66, no. 4, April 1960, p. 56-67, 70.
Article 20

TWO NOMOGRAPHS FOR COMPUTATION OF STANDARD EQUATIONS 

IN EARTH-RESISTIVITY INTERPRETATION

By JAMES H. SCOTT, Denver, Colo.

Abstract. Two nomographs have been developed to aid in 

making interpretations of resistivity field data. One represents 

the equation

Q= P2 P1

and the other

P2 + P1

\ . % __

Pi P2

Electrical earth-resistivity measurements under favorable conditions may be interpreted by fitting 

theoretical curves to field data. Field values of apparent resistivity are plotted against electrode separation 

on log-log graph paper, and 2- or 3-layer theoretical 

curves (Roman, 1934; Wetzel and McMurry, 1937) 

are selected, combined, and adjusted to obtain a 

modified theoretical curve which fits the plotted field 

values. During the interpretation process two equations are used repeatedly.

The first is

 i P2~P1 '

P2 + P1

(1)

where Q is the reflecting factor, pi is the effective 

resistivity of the overburden, and p2 is the resistivity 

of the lower layer, which is assumed to extend downward to infinity. Q is a dimensionless quantity, and 

Pi and p2 are measured in units of ohm-length (for 

example ohm-meters, ohm-feet, etc.). It matters not 

what particular units of resistivity are used as long as 

both pl and p2 are measured in the same units. Equation 1 is usually solved for p2 but occasionally for Q 

or Pl.

The second equation,

*1 . C2 ti~T~t 2

 I = > 

Pl P2 P

(2)

is used to obtain the effective composite resistivity, p', 

for two layers having thicknesses t\ and tz and respective resistivities of pi and p2 . Occasionally equation

2 is solved for one of the parameters other than p'. 

Resistivities pi, p2 , and p' are measured in units of 

ohm-length, and thicknesses ti and t2 are measured in 

units of length. Any units may be used as long as the 

same resistivity units are used for pb p2 , and p', and 

the same units of length are used for t\ and t2.

Nomographs representing equations 1 and 2 are 

presented in figures 20.1 and 20.2. Design equations 

for the nomographs were derived using methods 

described by Epstein (1958). The p scales and the t 

scales on the nomographs may be multiplied by any 

factor to extend the range of usefulness of the nomographs. However, if one of the p or t scales is multiplied by a factor, its companion scale (or scales) must 

also be multiplied by the same factor. For example, 

in figure 20.2, if the ti scale is multiplied by 102, then 

the tz scale must also be multiplied by 102, or if the pi 

scale is multiplied by 10"1 then the p2 and p' scales 

must also be multiplied by 10"1 .

The nomograph of figure 20.1 is used in the usual 

way; the value of the unknown parameter is found by 

drawing a straight line through values of the known 

parameters and observing where the line intersects the 

scale of the unknown parameter. The nomograph of 

figure 20.2 is used in a different manner. First, pi is 

plotted against tz on the left half of the nomograph. 

Then p2 is plotted against £1 on the right half of the 

nomograph. The desired value of p' is found by drawing a straight line between the two plotted points and 

observing where it intersects the central vertical scale. 

With a little experimenting one can see how to determine parameters other than p' by drawing a straight 

line through the known value of p' and the known p, t 

pair (pi, tz or p2 , £1).

The main advantage of using the nomographs is the 

speed and ease with which results can be obtained. Although numerical solutions of equations 1 and 2 are

ART. 20 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B71-73. 1963.

B71
B72 GEOPHYSICS

/c

1000

900-

800-

700-

600 -

500 -

400:

300 :

200 :

100. 

90^

80 i

70 {

60:

50:

40^

30 -_

20-

10-

'i C 

r - 0.98 n

L- -.95

; -.90 :

: -.80 :

: -.70

- -.60-

" -.50-

'- -.40-

-.30 :

-.20-

-.10 :

o .

\j

3 /

LQ :

-.10

-.20

-.30

-.40 ;

-.50 :

-.60 ;

-.70 :

7. BO :

-.90 :

-.95 :

.98 ^

P2 pi FIGURE 20.1. Nomograph representing the equation Q=. P2~rpi

ho

: 20

:30

^40

:50

:60

^70

-80

r9Q

L 100

-200

-300

-400

-500

-600

-700

800

-900

1000

more accurate than nomograph solutions, experience 

has indicated that nomograph results are adequate for 

most resistivity-interpretation problems. However, if 

resistivity or thickness values cover a very wide range it 

is advisable to check the nomograph results with hand 

calculations if high accuracy is required.

REFERENCES

Epstein, L. T., 1958, Nomography: New York, Interscienee 

Publishers, Inc., 144 p.

Roman, I., 1934, Some interpretations of earth-resistivity data: 

Am. Inst. Mining Metall. Engineers Trans., v. 110, p. 183-200.

Wetzel, W. W., and McMurry, H. V., 1937, A set of curves to 

assist in the interpretation of three layer resistivity problem : Geophysics, v. 2, no. 4, p. 329*-341.
SCOTT B73

I 

1 00

1000 

500

200

, m

70 

60

50 

40

30 

25

20 

19

18 

17

16 

15

14 

13

12 

11

10

\ P' L 

00 90 80 70 60 50 40 30 20 10 0 10 20 30 40 50 60 70 80 90 1<

"* 

2

DO00

1000 

500

200

70 

60

50 

40

30 

25

20 

19 

18 

17

16 

15

14 

13

12 

11

10

100 90 80 70 60 50 40 30 20 10 0 10 20 30 40 50 60 70 80 90 100

FIGURE 20.2. Nomograph representing the equation + = , * 

Pl P2 P
Article 21

NEW CAMBRIAN, ORDOVICIAN, AND SILURIAN FORMATIONS

IN THE INDEPENDENCE QUADRANGLE, INYO COUNTY, CALIFORNIA

By DONALD C. ROSS, Menlo Park, Calif.

Work done in cooperation with the California Division of Mines and Geology

Abstract. The Bonanza King Formation is conformably overlain by limestone and siltstone of the Lead Gulch Formation 

(Upper Cambrian), and by gray dolomite of the Tamarack 

Canyon Dolomite (Upper Cambrian). The overlying Mazourka 

Group consists of two formations: siltstone and limestone of the 

Al Rose Formation (Lower Ordovician), and fossiliferous sdlty 

blue-gray limestone, siltstone, and chert of the Badger Flat 

Limestone (Middle Ordovician). A conformable sequence continues through the Middle and Upper Ordovician, and is capped 

by argillaceous, bioclastic limestone rich in coral debris of the 

Vaughn Gulch Limestone (Silurian), which grades laterally 

northward to graptolite-bearing calcareous shale, siltstone, and 

argillaceous limestone of the Sunday Canyon Formation 

(Silurian).

Six new formations are named in a conformable sequence of Cambrian, Ordovician, and Silurian strata in 

the Independence quadrangle (fig. 21.1), to call attention to lithologic contrasts of these units with their 

stratigraphic equivalents in nearby areas, particularly 

to the south and east.

CAMBRIAN FORMATIONS

Conformably overlying the Bonanza King Formation of Middle and Late Cambrian age is the Lead 

Gulch Formation, which comprises about 300 feet of 

thin-bedded limestone, siltstone, dolomite, chert, and 

shale. It is in turn conformably overlain by about 900 

feet of thin- to thick-bedded gray dolomite with some 

chert that is named the Tamarack Canyon Dolomite. 

Details of distribution of the Lead Gulch Formation 

and the Tamarack Canyon Dolomite can be obtained 

from a preliminary geologic map of the Independence 

quadrangle, a black and white map at a scale of 1:48,000 

(Boss, 1962). On this map the term Nopah Formation 

(Upper Cambrian), which included both formations, 

was used. The two formations were differentiated in 

the field, however, and both will be shown in color on

a subsequent geologic map. Late Cambrian trilobites 

in the Lead Gulch Formation and the gross lithology 

of the two formations suggest affinities with the Nopah 

Formation and the Dunderberg Shale. Local names 

are proposed for use in the Inyo Mountains, however, to 

point up lithologic variations from the type Nopah and 

Dunderberg, and to express some degree of doubt as to 

exact correlation with These units at this time.

Lead Gulch Formation

The Lead Gulch Formation is here named for exposures in Lead Gulch, a short tributary of Mazourka 

Canyon (fig. 21.1). The formation extends as a thin, 

relatively continuous, though faulted and folded, 'belt 

from about the latitude of Independence north to the 

edge of the quadrangle. The belt of outcrop continues 

north into the adjoining Waucoba Mountain quadrangle 

for about 1 mile, where it is cut out by Mesozoic granitic rocks-and overlapped by Cenozoic basalt. To the 

south, rocks similar to the Lead Gulch Formation were 

seen at several points along the front of the Inyo Mountains east of Lone Pine (15 miles south of Independence 

on U.S. Highways 6 and 395).

The type section of the formation is here designated 

as being on a spur along the east wall of Mazourka Canyon, 7,500 feet S. 85° E. from the SE. cor. sec. 36, T. 11 

S., K. 35 E. Many other good exposures of at least partial sections are found along the belt of outcrop, but one 

of the most striking and easily accessible is along the 

ridge about 750 feet northeast of the Whiteside mine 

east of the Mazourka Canyon road.

The thickness of the Lead Gulch Formation at the 

type section is 280 feet, but structural complications at 

almost all outcrops of this relatively incompetent unit 

preclude accurate measured sections. Probably a maximum of 300 feet would be a fair estimate for the thickART. 21 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B74-B85. 1963.

B74
ROSS B75

118°00'

Mesozoic granitic 

rocks

Mesozoic granitic rocks

;Ty >e;seci ori iif Tarriarac 

Canyon E'pjom t&

EXPLANATION Type;;sectioniof A

Younger Paleozoic rocks

Silurian rocks

Svg, Vaughn Gulch Limestone; coral bearing 

Ssc, Sunday Canyon Formation; graptolite 

bearing

Type section of Sunday

Canyon Formation

Undirrerentiated Middle and

Upper Ordovician- Formations

Mazourka Group

Includes Al Rose Formation (Lower Ordovician)

and Badger Flat Limestone (Middle Ordovician)

Tamarack Canyon Dolomite and Lead

Gulch Formation (Upper Cambrian)

Mesozoic granitic rpcks

Older Paleozoic rocks

Kearsarge (Citrus) 

(Abandoned)

Type section of 

Vaughn Gulch 

Limestone

Independence 

Cenozoic alluvial deposits quadrangle

36°45

FIGUKE 21.1. Generalized geologic map of the Independence quadrangle, Inyo County, Calif.

680231 O 65
B76 STRATIGRAPHY AND STRUCTURAL GEOLOGY

ness of the formation. At a great many places it is 

thinned considerably or is entirely squeezed out between 

the bounding thick dolomites.

The contact with the underlying Bonanza King For- 

 mation is probably conformable judging by the similarity of attitudes near the contact, though the actual 

contact is almost invariably structurally disturbed. 

The contact with the overlying Tamarack Canyon Dolomite is definitely conformable, and can be seen best at 

the type locality.

The Lead Gulch Formation comprises limestone, siltstone, dolomite, chert, and shale interlayered in a regular-bedded sequence in which beds are most commonly 

1/2 to 2 inches thick, but locally as thick as 5 inches. 

Thinly laminated siltstone that weathers out in relief in 

bright orange and red tints, and blue-gray to mediumgray limestone are the dominant lithologies. A distinctive fissile olive-brown to dark-green shale, or its 

metamorphic equivalent, occupies as much as 20 feet at 

the very base of some sections. This unit, though not 

noted at the type section, was probably originally present but has been squeezed out. Much of the unit, and 

particularly the silty and marly material of the orangeweathering beds, has undergone thermal metamorphism. Minerals, such as diopside, wollastonite, tremolite, scapolite, biotite, and antigorite are present in varying amounts.

The physical appearance of outcrops of the Lead 

Gulch Formation is the most striking of any unit in the 

Inyo Mountains. The interlayered regular-bedded sequence of bluish-gray-weathering limestone and orangeweathering siltstone, which is particularly evident 

where the section has been metamorphosed, is an excellent stratigraphic and structural marker. The basal 

dark-green shale beds are particularly susceptible to 

structural squeezing and were not noted in many outcrops. They serve locally to indicate the base of the 

formation, particularly north of Lead Gulch where 

overturning of the formation could be deduced despite 

the metamorphism of the usually distinctive bounding 

units to coarse buff massive indistinguishable dolomites.

Superficially, outcrops of the Lead Gulch Formation 

might be confused with parts of the overlying Al Rose 

Formation. In most outcrops where a reasonable 

thickness is exposed, however, it becomes apparent that 

the Lead Gulch Formation is characterized by regular 

thin interbedding of the silty and limy units in relatively equal proportions, whereas the Al Rose Formation typically has irregular bedding and wide variations in the relative proportions of limy and silty 

material. Commonly silty material greatly exceeds 

limy material, which occurs in less resistant lenses or

"eyes." The limestone "eyes," so typical in the Al Rose, 

are rare in Lead Gulch outcrops.

The type section (see below) has perhaps a greater 

proportion of thin-bedded dolomite than most other 

outcrops, but interbedded units of thin-bedded dolomite 

and thin-bedded limestone are common, particularly 

near the top of the formation. The top of the formation is the uppermost thin-bedded limestone. The base 

of the formation is sharp and is the lowest fine-grained 

clastic rock or limestone that overlies the dolomite of 

the Bonanza King Formation.

Type section of the Lead Gulch Formation in the Independence 

15-minute quadrangle, Inyo County, Calif.

[location of section by California grid, zone 4 : 2,273,300 ; 586,400]

Tamarack Canyon Dolomite. 

Conformable contact. 

Lead Gulch Formation (280 ft) :

Limestone, medium-dark-gray ; interbedded with 

orange-weathering silty layers in very thin beds ^ to 

1 inch thick _______________ 52 

Covered _ 15 

Dolomite, medium-dark-gray; very thin bedded in irregular layers __ 44

Covered _____ 35

Dolomite, medium- to dark-gray ; thin platy layers 5

Covered (scattered platy dolomite float; fragments of

Lingula) 50

Limestone, medium-bluish-gray ; laminated to very thin 

bedded. Interlayered with regular layers of siltstone 

to y2 inch thick, which weathers grayish to orange. 

Some black chert in nodular beds to 4 inches thick. 

Lingula present. Beds folded near base and contact 

contorted- 79 

Covered contact, presumably fault cutting out originally

conformable shale. 

Bonanza King Formation.

Careful search for fossils yielded only acrotretid 

brachiopods (black chitinous shells) and fragments 

possibly of trilobites, until the third field season when 

Fred K. Miller found a fossiliferous bed on the first 

spur south of the type section of the Tamarack Canyon 

Dolomite that yielded small agnostid trilobites, acrotretid brachiopods, and echinoderm parts. This collection was identified and studied by A. R. Palmer, who 

reports (written communication, 1961) :

The dominant identifiable trilobites are agnostids representing the genera Homagnostus and Pseudagnostus. Homagnostus is a characteristic agnostid in the lower Nopah faunas * * *. 

It is certainly from1 beds no older than the Late Cambrian Aphelaspis zone of Dresbach age. This is just about the westernmost 

dated Upper Cambrian locality in the U.S. (exclusive of Alaska) .

The finding of this fossil locality lead to the discovery 

of other trilobites about 21/4 miles farther south, which 

were identified and reported on by A. R. Palmer as follows (written communication, 1962) :
ROSS B77

The specimens all represent a species of what is presently being called Loganellus, and this sort of trilobite characterizes the 

beds just above the Dunderberg interval over much of Nevada.

The age of the Lead Gulch Formation is therefore 

considered to be Late Cambrian.

Beds lithologically similar to the Lead Gulch Formation and in the correct stratigraphic position to be correlative are also found along the south edge of the 

Waucoba Mountain quadrangle, in the Waucoba Wash 

quadrangle on the west side of Saline Valley, and in 

the New York Butte quadrangle east of Lone Pine. 

These areas respectively north, east, and south of the 

Independence area indicate that the Lead Gulch 

lithology is a persistent mappable feature through much 

of the Inyo Mountains. A similar lithology in the 

Quartz Spring area, which contains fossils indicating 

a Late Cambrian age, constitutes the lower 250 feet 

of unit 1 of the Nopah Formation (McAllister, 1952, 

p. 9). Another occurrence of this same lithology, shown 

to me by H. G. Stephens, is at the base of the Nopah 

Formation along the west front of the Panamint Eange 

north of California Highway 190. The Lead Gulch 

Formation is considered to be approximately equivalent 

to the unnamed basal unit of the Nopah Formation 

in Hazzard's (1937, p. 320) type area.

The term Dunderberg might be applied to these beds, 

as it has been used informally for the beds lithologically similar to the Lead Gulch at Quartz Spring and 

in the Panamints. However, the lithology and fauna 

of the Catlin Member of the Windfall Formation 

(Upper Cambrian) at the Nevada Test Site (Barnes 

and Byers, 1961) also appear to be similar to the Lead 

Gulch Formation. It thus appears that the Lead Gulch 

Formation probably includes the Dunderberg interval 

plus some of the Windfall interval.

Tamarack Canyon Dolomite

The Tamarack Canyon Dolomite is here named for 

exposures near the mouth of Tamarack Canyon (fig. 

21.1). The type section is here designated to be along 

the bottom of a subsidiary canyon of Mazourka Canyon 

about 5,700 feet N. 59° E. of the SE. cor. sec, 36, T. 11 

S., E. 35 E. Its distribution parallels that of the Lead 

Gulch Formation. It is a relatively competent unit 

and is generally well exposed throughout the belt of 

outcrop.

At the type section the formation is 910 feet thick. 

Where variations in thickness exist they are thought to 

be the result chiefly of faulting and folding. Probably 

900 feet, is a reasonable average thickness for this unit. 

Both the overlying and underlying contacts are conformable.

The formation (see section below) is dominantly 

laminated to thick-bedded very-light-gray to mediumgray dolomite that weathers normally to a monotonous 

dull-gray surface. Commonly outcrops that look 

grossly thick bedded to massive prove to be thin bedded 

on close observation; the bedding is shown chiefly by 

a fluted weathered surface that accentuates minor differences in the thin dolomite layers. Black chert forms 

nodules and nodular beds as much as a few inches thick. 

It is scattered to abundant in the dolomite, but though 

widespread, it is absent in many outcrops. Chert 

nodules are particularly abundant near the mouth of 

Lead Gulch and to the west across the Mazourka Canyon road. Hints of crossbedding in some outcrops suggest clastic dolomite, but'most dolomite is sufficiently 

recrystallized to mask any original clastic texture. 

Near granitic contacts, particularly in the Lead Gulch 

area, the dolomite becomes quite coarse grained and 

dazzling white on fresh surfaces.

Type section of the Tamarack Canyon Dolomite in the Independence 15^mimtte quadrangle, Inyo County, C'a-lif.

[Location of section by California grid, zone 4: Base of section, 

2,271,500; 590,000; top of section, 2,270,500; 590,0001

Al Rose Formation. Feet

Conformable contact.

Tamarack Canyon Dolomite (910ft) : 

Dolomite, medium-dark-gray; weathers medium light 

gray; generally massive, but in part very thin 

bedded. Rilled weathered surface accentuates the 

thin bedding_______ ___ 272 

Dolomite, medium- to dark-gray; weathers medium light 

gray to light gray. Very thin bedded to laminated 

with rilled weathered surfaces in part. Flaggy to 

slabby parting. Minor black chert near the base. 123 

Dolomite, same as overlying unit, but black chert 

nodules and lenses as thick as 4 inches and as long 

as 2 ft are more abundant-_ 515

Conformable contact.

Lead Gulch Formation.

The base of the formation is taken as the base of 

the dolomite above the highest limestone in the underlying Lead Gulch Formation. It is quite probable that 

the contact between these two formations is an interfingering one and that different horizons are taken as 

the contact in different areas. The top of the formation 

is marked by the highest dolomite overlain by the 

limestone, siltstone, and shale of the Al Eose Formation. Both the upper and lower contacts are 

pronounced color breaks; the gray dolomite contrasts 

strongly with the brown weathering bounding formations.

No fossils have been found in this unit, but on the 

basis of the fossil content of the bounding formations, 

it could be Late Cambrian or Early Ordovician or both. 

A Late Cambrian age is tentatively assigned to the 

Tamarack Canyon because the absence of clastic quartz 

and silty interlayers in this unit suggests affinity with
B78 STRATIGRAPHY AND STRUCTURAL GEOLOGY

the Nopah Formation (Upper Cambrian) rather than 

with the lower part of the Pogoni^ Group (Lower and 

Middle Ordovician), in which clastic material is widespread. The lithologic break between the Al Rose 

Formation with its abundant clastic material and the 

essentially clastic-free Tamarack Canyon Dolomite is 

thus tentatively equated with the contact between the 

Pogonip Group and the Nopah Formation and the 

boundary between the Ordovician and the Cambrian.

Though both the Nopah and Tamarack Canyon 

Formations are predominantly dolomite, a possibly 

significant difference exists. The Nopah in its type area 

in the Nopah Range (Hazzard, 1937, p. 320), in the 

Quartz Spring area (McAllister, 1952, p. 9), and at 

many other localities including the west flank of the 

Last Chance Range only 25 miles northeast of the Independence quadrangle, has characteristic banding in a 

wide range of shades of gray. This banding, similar to 

that which is so diagnostic of the Bonanza King Formation, is absent in the Tamarack Canyon Dolomite. The 

seemingly attractive hypothesis that Tamarack Canyon 

is Nopah that has lost its banding from incipient 

metamorphism is not valid because the underlying 

Bonanza King Formation, with the same opportunity 

for metamorphism, preserves its diagnostic banding 

quite well. The lack of this diagnostic and very widespread color banding of the Nopah is the chief reason 

for proposing a local name for presumably correlative 

rocks in the Inyos.

The Tamarack Canyon Dolomite is tentatively correlated with the Nopah Formation of the Quartz Spring 

area (McAllister, 1952, p. 9) minus its lowest member, 

which may be correlative with the Lead Gulch Formation. The possibility of correlation of the Tamarack 

Canyon with, lower Pogonip alone or combined with 

some part of the Nopah cannot be excluded at this time. 

This dolomite immediately below the Mazourka Formation of Phleger (1933) is regarded by Langenheim 

and others (1956, p. 2088, fig. 3) as Pogonip. Dull-gray 

dolomites in the correct stratigraphic position to be correlative with the Tamarack Canyon were seen in the 

Inyo Mountains east of Lone Pine, and possibly unit "a" 

of the Pogonip Group of the Darwin area (Hall and 

MacKevitt, 1962, p. 8) is also a correlative of the upper 

part of the Tamarack Canyon Dolomite.

ORDOVICIAN FORMATIONS

New data for the rocks immediately overlying the 

Upper Cambrian beds suggest the need to modify the 

nomenclature of that part of the Ordovician section 

named the Mazourka Formation by Phleger (1933, p. 

2-3). He briefly described "a succession of argillaceous 

shales and limestones of lower Middle Ordovician age,

675 feet in thickness," in which he recognized 2 members, a lower calcareous shale 125 feet thick and an 

upper argillaceous limestone 550 feet thick. Somewhat 

earlier Kirk (in Knopf, 1918, p. 35), in the first description of Ordovician rocks from this area, referred to 

about 500 feet of argillaceous limestone containing 

Chazyan fossils, which is almost certainly the same interval as the upper member of Phleger's Mazourka 

Formation. Langenheim and others (1956, p. 2085, 

2091) augmented previous fossil collections and measured stratigraphic sections that included Phleger's Mazourka Formation.

Detailed study of the Independence quadrangle has 

confirmed Phleger's Mazourka Formation as a mappable unit. I propose that the Mazourka Formation 

be raised to group status and that the two members 

Phleger described be given formational rank. The 

newly proposed Mazourka Group thus contains the Al 

Rose Formation (Lower Ordovician), comprising 

400(?) feet of brown-weathering, in part graptolitebearing siltstone and limestone, which is overlain by 

the Badger Flat Limestone (Middle Ordovician), comprising 600 feet of fossiliferous silty blue-gray limestone, siltstone, and chert. The continuity and distinctive nature of these 2 formations along a strike length 

of about 15 miles plus the recognition of similar lithologies north and southeast of the Independence quadrangle make them comparable to the other mapped formations in the area. In addition, group status for the 

Mazourka would give it a somewhat parallel position 

with the Pogonip Group, with which it is correlative, 

certainly in part, and possibly in its entirety.

The name Pogonip Group is not used in the Independence area because of differing opinions as to how 

much of the Inyo Mountains section is equivalent to the 

Pogonip. Langenheim and others (1956, p. 2087, 2091) 

considered the Mazourka Formation of Phleger the 

uppermost formation of the Pogonip Group. The underlying dolomite, which I consider to be Upper Cambrian and correlative with part of the Nopah Formation, is part of the Pogonip Group by their usage. 

Farther south in the Inyo Mountains, in the New York 

Butte quadrangle, C. W. Merriam (written communication, 1959) also includes units in the Pogonip Group 

that are probably correlative with rocks referred to the 

Upper Cambrian in the Independence quadrangle. 

These differences of opinion as to what constitutes 

Pogonip in the Inyos make it advisable to retain the 

local term Mazourka in the Independence area for the 

present.

The type sections of the Al Rose Formation and 

the Badger Flat Limestone, which together make up 

the Mazourka Group (see stratigraphic section), are
ROSS B79

here designated as being on a spur along the east wall 

of Mazourka Canyon. The top of the composite section is 6,000 feet S. 70° E. from the SE. cor. sec. 36, T. 

11 S., R. 35 E.; it was measured up the crest of the 

northeast-trending spur. To reach the section, turn 

east off U.S. Highway 6 and 395 at the Mazourka Canyon road at the south edge of Independence. About 

15 miles from, the highway a branch road to the right 

goes to talc prospects near the base of the section. 

Phleger's (1953, p. 2) designation of the type section 

of his Mazourka Formation was "an unnamed canyon 

which is the first tributary canyon entering Mazourka 

Canyon on the east below the Elbow in Mazourka 

Canyon." The precise location of this "elbow" is not 

known, but probably the ridge on which I measured 

the Mazourka Group is about 500 feet southeast of the 

canyon that Phleger referred to.

Type section, of the Mazourka Group in the Independence 15- 

minute quadrangle, Inyo County, Calif.

[Location of section by California grid, zone 4>: Base of section 

2,272,800; 586,000; top of section, 2,271,400; 585,000]

Barrel Spring Formation of Phleger (1933). 

Conformable contact. 

Mazourka Group (about 1,000ft) :

Badger Flat Limestone (586ft) : Feet

Limestone, medium-gray, very thin and irregular bedding; interbedded lenses and nodular beds of 

brown-weathering silty material. Rare chert layers as much as % inch thick. Abundant fossil 

fragments__________________________ 41

Quartzite, dark-gray to light-gray; beds 6 inches to 

1 foot thick, weather to brownish knobby surface. 

Some limy cement____________________ 5

Limestone, silty, medium-gray; weathers distinctive 

dark yellowish orange to light brown. Pelmatozoan 

fragments__________________________ 12

Limestone, medium-gray, very thin, irregular bedding; abundant nodules and lenses of yellowishweathering silty material. Pelmatozoan fragments____________________________ 23

Quartzite, medium-gray; weathers brownish and 

knobby. Some limy cement and lenticular limestone layers-_______________________ 8

Limestone, medium-gray to dark-gray, irregular 

lenses of yellow- to brown-weathering silty material give distinctive irregular nodular appearance 

to outcrops. Fossils locally abundant_______ 489

Limestone, silty(?), gray, very thin bedded, minor

black chert_________________________ 8 

Al Rose Formation (about 400 ft) :

Shale and mudgtone, medium-d'ark-gray to olive-gray; 

weathers light brown to moderate reddish brown in 

layers to 2 inches thick ; interbedded gray limestone 

in beds to % inch thick. Makes distinctive dark 

outcrop. Graptolites near top____________ 52

Mazourka Group (about 1,000 ft) Continued

Al Rose Formation (about 400 ft) Continued feet 

Siltstone and mudstone, medium-gray ; weathers dark 

yellowish orange to light brown, very irregularly 

interbedded with medium-gray limestone in very 

thin beds. Where silty beds are dominant, limestone lenses weather out leaving diagnostic "eyes." 

Outcrops have an overall brown appearance. Fossil 

fragments locally. Structural contortions in this 

relatively incompetent unit preclude accurate measurement of thickness________ ___ 300-400 

Limestone, medium-gray, very thin bedded to laminated ; black chert near base. A 6-inch bed of edgewise conglomerate associated with sandy bioclastic 

layers about 3 feet below top 13

Conformable contact.

Tamarack Canyon Dolomite.

Details of the distribution of both formations of the 

Mazourka Group can be obtained from a preliminary 

black and white geologic map of the Independence 

quadrangle by Ross (1962). On this map the lower and 

upper units of the Pogonip Group equate respectively 

to the newly named Al Rose Formation and Badger 

Flat Limestone.

Al Rose Formation

The Al Rose Formation is here named for exposures 

east of Al Rose Canyon, a tributary to Mazourka Canyon (fig. 21.1). The formation crops out as a relatively 

continuous but faulted belt along almost the entire 

length of the Independence quadrangle. To the north 

the belt continues about a mile into the adjacent 

Waucoba Mountain quadrangle, where it is cut out by 

Mesozoic granitic rocks and overlapped by Cenozoic 

deposits of Owens Valley. Southward, also, the belt of 

outcrop is interrupted by Mesozoic granitic plutons, but 

at several places along the front of the Inyo Mountains 

east of Lone Pine in the New York Butte quadrangle, 

the distinctive lithology of the Al Rose Formation has 

been recognized. The minimum length of the discontinuous belt of exposure is about 30 miles.

The thickness of the Al Rose Formation is difficult to 

determine accurately because the unit is relatively incompetent and consequently folded and faulted. At the 

type section, I would estimate the thickness to be about 

400 feet. Phleger (1933, p. 2) reports a thickness of 

125 feet for his lower member, which is presumably the 

Al Rose Formation of this report, but no map accompanies his report and I was unable from his description 

to locate the place where his section was measured. 

Langenheim and others (1956, p. 2087) give a thickness 

of 106 feet to beds lithologically equivalent to the Al 

Rose Formation in their principal measured section, but 

strike faulting has thinned the section.
B80 STRATIGRAPHY AND STRUCTURAL GEOLOGY

The base of the formation is commonly structurally 

disturbed or faulted out, but locally undisturbed parts 

of the contact with similar attitudes above and below 

indicate that the Al Rose Formation conformably overlies the Tamarack Canyon Dolomite. The contact with 

the overlying Badger Flat Limestone is also conformable.

Outcrops of the Al Rose Formation typically have 

orange- to red-brown surfaces and are readily distinguishable from the overlying and underlying grayweathering carbonate units. Siltstone, mudstone, shale, 

and less commonly chert, all of which are commonly 

hornfelsed, have very thin irregular bedding and are 

dominant in the formation. Medium-gray to bluishgray generally silty limestone is subordinate to the 

reddish-weathering fine-grained clastic rocks. Commonly the limestone forms elongate lenses which, being 

less resistant, weather back as holes or "eyes" in the 

outcrop. This gives rise to a very diagnostic weathered 

surface, particularly in hornfelsed outcrops.

The uppermost unit of the formation, which is 52 feet 

thick at the type locality, is a much more regularly 

interbedded sequence of 1- to 2-inch beds of gray limestone and light-brown-weathering shale and siltstone. 

In some areas this unit has a contrastingly darker brown 

outcrop than the rest of the formation. This unit, which 

bears an important graptolite fauna, is probably equivalent to unit 4 of the section of Langenheim and others 

(1956, p.2087).

The base of the formation is defined as the lowest 

occurrence of fine-grained elastics and limestone, below 

which is thin- to thick-bedded dolomite. The top of the 

formation is defined as the uppermost occurrence of 

brown-weathering beds, which are in contact with the 

overlying blue-gray layers of the Badger Flat Limestone. The contact between the dominantly brownweathering Al Rose Formation and the overlying bluegray-weathering Badger Flat Limestone is one of the 

most distinctive contacts in the area.

Near the top of the Al Rose Formation, in the regularly bedded unit, graptolites collected from several 

localities were identified by R. J. Ross, Jr. (written 

communication, 1961), as Didymograptus protobifidus 

Elles and species of Phyllograptus, the first Early Ordovician fossils that have been reported from the Inyo 

Mountains. They are considered by R. J. Ross, Jr., to be 

of late Arenig age and indicate a horizon high in the 

Ninemile Formation of the Pogonip Group. Near the 

base of the formation, trilobites and phosphatic brachiopods are also present, which suggests possible correlation with the Goodwin Limestone of the Pogonip 

Group (R. J. Ross, Jr., written communication, 1961).

On the basis of its fossils the Al Rose Formation is 

assigned an Early Ordovician age.

Examination of the section studied by McAllister in 

the Quartz Spring area, 30 miles to the east, suggests 

that units 4 through 7 of the Pogonip (1952, p. 11) are 

equivalent to the Al Rose Formation. Similar comparison suggests equivalence of units "b" and "c" of the 

Pogonip Group in the Darwin area, 30 miles to the 

southeast (Hall and MacKevett, 1962, p. 8).

Badger Flat Limestone

The Badger Flat Limestone is here named for exposures in and near Badger Flat (fig. 21.1). The formation crops out in a faulted belt parallel to the Al 

Rose Formation and with almost identical distribution.

At the type section the Badger Flat Limestone is 

586 feet thick. The formation has a relatively constant thickness, though variations exist due to structural complications and possibly to variations in 

original thickness. The common range of thickness is 

500 to 600 feet in relatively undisturbed sections, which 

compares favorably with Phleger's (1933, p. 2) measurement of 550 feet at his type section. Kirk's original 

description of about 500 feet of fossiliferous argillaceous limestone (in Knopf, 1918, p. 35) was probably 

also of the Badger Flat Limestone. Thickness measurements by Langenheim and others (1956, p. 2085, 

2088, 2090) are somewhat lower, reflecting structural 

thinning in areas where faulting was not readily apparent until larger areas were mapped in detail.

The Badger Flat Limestone rests conformably on the 

Al Rose Formation and is in turn conformably overlain 

by the Barrel Spring Formation of Phleger (1933, p. 

5) of Middle Ordovician age.

Outcrops of the formation generally weather shades 

of gray and blue gray. The blue color of outcrops 

is diagnostic, as is nodular irregular bedding. Darkto medium-gray limestone is the dominant rock type 

(see stratigraphic section). Almost invariably, specimens studied in thin section reveal an abundance of 

silty quartz grains scattered in a matrix of clastic calcite. Much of the limestone is thus calcarenite or calcilutite. Irregular lenses and beds of light-gray, 

orange, and reddish-brown-weathering silty and marly 

material occur in varying amounts throughout the 

formation. These more resistant layers weather out in 

relief as spines and ridges or in exotic irregular patterns to give a distinctive appearance to weathered 

outcrops. Black chert, both as nodules and nodular 

beds, is locally abundant, particularly in the lower part 

of the formation.

The base of the formation is defined as the lowest 

occurrence of blue-gray limestone, and this contact
ROSS B81

with the upper unit of the brown-weathering Al Rose 

Formation is readily apparent. The top of the formation is placed at the top of the highest blue-gray limestone in contact with the impure quartzite or sandy 

limestone at the base of the Barrel Spring Formation 

of Phleger (1933). Thin quartzite layers near the top 

of the Badger Flat Limestone in some sections, including the type section, can cause confusion, but the basal 

unit of the Barrel Spring is several tens of feet thick, 

whereas the quartzite units in the Badger Flat are less 

than 10 feet thick.

Fossils are particularly abundant in the Badger Flat 

Limestone. In the interval from 100 to 200 feet below 

the top of the formation, gastropods as large as 3 inches 

across and belonging at least in part to the genera 

Palliseria and Maclurites are locally abundant enough 

to be a mappable feature. In this zone, Palliseria 

robusta Wilson, a guide to the Whiterock Stage of 

Cooper (1956) of lowest Middle Ordovician, has been 

identified by E. L. Yochelson (written communication, 

1961). This species is also present in the Antelope 

Valley Limestone of the Pogonip Group (Yochelson, 

written communication, 1961).

Brachiopods and trilobites have also been collected 

from several different beds in this formation. Most 

abundant are the brachiopods Plectorthis mazourkaensis Phleger and Plectarthis patula, Phleger (R. J. Ross, 

Jr., written communication, 1961), which are most common about midway in the formation. Pelmatozoan 

fragments are also widespread and abundant, and conceritric structures as large as half an inch across and 

resembling the algal form G-irvanella are present. On 

the basis of its fossils the Badger Flat Limestone is 

assigned a Middle Ordovician age.

Comparison with the Quartz Spring area suggests 

correlation of unit 8 of the Pogonip (McAllister, 1952, 

p. 11) with the Badger Flat Limestone. In the Darwin area, unit "d" of the Pogonip Group (Hall and 

MacKevett, 1962, p. 8) is probably also a correlative. 

It is interesting to note that at Darwin and the Quartz 

Spring area the upper part of the Pogonip, in common 

with the Badger Flat Limestone, is medium gray and 

somewhat bluish, irregularly bedded, contains abundant silty material, and bears a distinctive gastropod 

zone. At both Darwin and Quartz Spring, however, 

the carbonate is almost entirely dolomite, whereas in 

the Badger Flat Limestone it is entirely calcite.

Other Middle and Upper Ordovician rocks

Conformably overlying the Badger Flat Limestone 

of the Mazourka Group is a succession of three formations of Middle and Late Ordovician age. These units, 

named from base to top respectively, the Barrel Spring

Formation, the Johnson Spring Formation, and the 

Ely Spring Formation, have all been used previously 

in the Independence quadrangle, and have been retained 

in the current mapping as previously defined. Brief 

descriptions of their lithology and their distribution can 

be obtained from the preliminary geologic map of the 

quadrangle by Ross (1962).

SILURIAN FORMATIONS

The uppermost Ordovician formation is conformably 

overlain by Silurian strata that are lithologically quite 

distinct from the dolomite that represents the Silurian 

most commonly elsewhere in the southwestern part of 

the Great Basin. In the southeast part of the Independence quadrangle the name Vaughn Gulch Limestone is proposed for a thin-bedded sequence of argillaceous limestone and bioclastic limestone rich in coral 

debris. This f acies grades laterally northward to graptolite-bearing calcareous shale, siltstone, and argillaceous limestone for which the name Sunday Canyon Formation is proposed.

Rocks now known to be of Silurian age in the Independence quadrangle were first reported by Waring 

and Huguenin (1917, p. 54), but probably the first field 

identification of these same rocks was by Kirk in 1912 

(in Knopf, 1918, p. 36-37). A section of 1,430+ feet 

was measured east of Kearsarge and fossils were collected by Kirk, who assigned these beds to the Devonian. 

Kirk later recognized that these beds may be Silurian 

(Nolan, 1943, p. 153).

Stauffer later (1930, p. 86) examined this same section. His measured section probably includes some of 

the overlying Mississippian beds and underlying Ordovician beds, but he tabulated an impressive array of fossil forms from beds totaling about 1,300 feet, which he 

referred to the Devonian.

Paleontologic study by Waite (1953, p. 1521) led him 

to assign a Silurian age to the beds in the Kearsarge 

area. Waite also made the first note of the fact that 

the coral-rich limestone facies of the Kearsarge section 

is largely replaced by a graptolite-bearing shaly facies 

to the north. This recognition by Waite of two Silurian 

facies within the Independence quadrangle is a significant contribution to the geology of the region.

Vaughn Gulch Limestone

The Vaughn Gulch Limestone is here named for exposures along Vaughn Gulch and the ridge to the north 

(fig. 21.1). Vaughn Gulch is a small canyon tributary 

to Owens Valley near the mouth of Mazourka Canyon. 

The formation is exposed as a discontinuous faulted 

and folded belt extending from a point about 2 miles 

southeast of Vaughn Gulch north for about 7 miles to
B82 STRATIGRAPHY AND STRUCTURAL GEOLOGY

Water Canyon. The outcrop is terminated at the south 

end by a granitic intrusive body and a fault. The nearest Silurian rocks on strike to the south appear about 

12 miles to the southeast, in the New York Butte quadrangle, but these are dolomite and quartzite of the Hidden Valley Dolomite (C. W. Merriam, written communication, 1962). The north limit of the Vaughn Gulch 

Limestone is placed in Water Canyon. Kocks of 

Vaughn Gulch lithology extend as tongues for several 

miles north of Water Canyon, but the canyon makes a 

practical mapping break.

The type section is along the ridge northwest of 

Vaughn Gulch in the NE^4 sec. 8, T. 13 S., R. 36 E. 

The section here is very well exposed and copiously f ossiliferous, and only minor folds and faults disturb the 

homoclinal section. Elsewhere in the belt of outcrop, 

folds, faults, and, in places, rather intense contact metamorphism affect the formation.

The thickness at the type section is about 1,500 feet. 

The contact with the underlying Ely Springs Dolomite 

of Late Ordovician age is conformable; the base is 

placed above the highest dolomite or massive chert of 

the Ely Springs. The contact with the overlying Perdido Formation of Mississippian age is one of erosional 

unconformity. The top of the formation is generally 

an obvious contact of coarse noncalcareous clastic rocks 

overlying fossiliferous limestone.

The Vaughn Gulch Limestone is dominantly thinbedded argillaceous and silty limestone (see section 

below). The limestone is generally medium to dark 

gray and much has a blue cast. Bioclastic limestone, 

crowded with fragments of corals, sponges, bryozoans, 

and pelmatozoans, is the most diagnostic lithology of 

this formation. In general the impure layers weather 

in shades of red, orange, and yellow, and are less fossiliferous. Black chert in nodules and beds as thick 

as 3 inches is comon in unit 7 near the top of the formation, somewhat less common near the base, and scattered elsewhere. There are no obvious mappable units 

within the formation, and the units of the type section 

(see below) are distinguished on the basis of gross color 

differences and varying ratios of bioclastic limestone 

to argillaceous limestone.

Type section of the Vaughn Gulch Limestone in the Independence 15-minute Quadrangle, Inyo County, Calif.

[Location of section by California grid, zone 4: Base of section, 

2,270,900; 544,200; top of section, 2,270,200; 542,900]

Perdido Formation.

Erosional unconformity.

Vaughn Gulch Limestone (1,518 ft) : Feet 

7. Limestone, dark-gray; weathers generally medium 

gray in 1- to 6-ineh laminated beds interlayered 

with l^ 0 - to 3-inch black chert layers. 'Some limestone 'beds weather grayish orange to moderate

Vaughn Gulch Limestone (1,518 ft) Continued.

7. Limestone, dark-gray Continued. Feet 

reddish brown, suggesting argillaceous impurities. 

Chert decreases near top of unit and limestone 

becomes 'bioclastic; coral fragments abundant 'at 

upper contact____ __ 96

6. Limestone, medium- to dark-bluish-gray, dense, laminated. Lesser argillaceous limestone _ 70

5. Argillaceous limestone and siltstone, dark-gray; 

weathers light gray to grayish orange; platy to 

shaly fracture. Few fossils; generally poor 

exposure.______ ..____ 105

4. Argillaceous limestone with red-weathering surfaces, 

laminated, shaly parting. Interbedded bioclastic 

limestone and dense bluish limestone abundant, 

but not as common as in unit 3 484

3. Limestone, medium-dark-gray to medium-bluishgray ; 6-inch to 1-foot bioclastic beds are dominant, 

in which coral material is abundant; orangeweathering argillaceous layers subordinate. Minor 

chert_______________,_____________ 215

2. Limestone, like unit 1, but greater proportion of 

argillaceous limestone; bioclastic layers are subordinate. Slope has overall yellowish cast. Some 

black chert in lenticular layers and nodules______ 363

1. Limestone, medium-dark-gray to medium-bluish-gray, 

laminated to thin-bedded; alternating sequence of 

bioclastic limestone and dense blue limestone. 

Thin-bedded orange-weathering argillaceous limestone subordinate. Thin black chert 'beds and 

nodules also present 185 

Conformable contact. 

Ely Springs Dolomite.

Systematic changes in lithology were not evident 

from the type section either south toward the dolomite 

f acies or north toward the calcareous shale f acies. The 

northernmost exposures of the Vaughn Gulch Limestone that are unfaulted are on a north-northeasttrending ridge about half a mile east of the NE. cor. 

sec. 36, T. 12 S., K. 35 E. Here bioclastic layers are 

less frequent and argillaceous material appears to be 

more abundant, but less than 300 feet of beds survived 

pre-Perdido erosion, so a comparison with the type 

section is not necessarily meaningful.

The abundance of fossils at the type section was 

first noted by Kirk (in Knopf, 1918, p. 36, 37), who 

made only a modest collection because of time limitations on his study. Stauffer, who examined the section first noted by Kirk, tabulated a prodigious list 

of fossils (1930, p. 86-89) keyed to units in his measured section, which he referred to the Middle Devonian 

(1930, p. 91). Waite .(1953, p. 1521), concluded that 

f aunal elements in the same section preclude a Devonian 

age. He specifically mentioned the forms Conchidium, 

Pycnostylus guelphensis Whiteaves, and Atrypina cf. 

A. disparilis (Hall). Waite also noted that new and 

better material shows that the Calceola of Stauffer is 

Rhizophyllwri) a characteristic Silurian genus, and con-
ROSS B83

eluded that the fossils indicate a late Niagara or early 

Cayuga age for the section.

Collections from the type section of the Vaughn 

Gulch Limestone have also been made by C. W. Merriam, who stated (written communication, 1963) :

The faunas consist very largely of corals, only Atrypa and 

rhynchonellids (Eatonia Mcostata Stauffer) being at all common among the brachiopods. The large dasycladacean algae 

(Verticillopora annulata Rezak) are most prolific here and 

provide a tie with the Hidden Valley of the type area as well 

as with the Roberts Mountains formation of central Nevada 

and the Laketown dolomite of western Utah.

Among corals of the limestone facies at Mazourka Canyon 

are many conforming to the general features of Strombodcs. 

Others are assigned to Chonophyllum, Rhizophyllum, Heliolites, Alveolites and Clad-opora. Also present are large cyathophyllids and bushy forms resembling Disphyllum.

The Vaughn Gulch Limestone is considered to be of 

Silurian age.

Sunday Canyon Formation

The Sunday Canyon Formation is here named for 

a small tributary canyon of Mazourka Canyon about 

1 mile west of the belt of Silurian outcrop (fig. 21.1). 

The formation extends in an almost continuous band 

from Water Canyon (about iy2 miles southeast of 

Sunday Canyon) north for about 6 miles to Badger 

Flat. North of Badger Flat the formation appears to 

have been completely removed before the deposition 

of the overlying Mississippian beds. The southern 

limit of the formation is placed at Water Canyon, 

where the two facies can most conveniently be differentiated.

The type section is along Bonanza Gulch, the first 

east-trending tributary to Mazourka Canyon north of 

Water Canyon. The base of the section is 9,000 feet 

N. 58° E. from the NE. cor. sec. 25, T. 12 S., E. 35 E.

Type section of the Sunday Canyon Formation in the Independence 15-minute quadrangle, Inyo Courity, Calif.

[Location of section by California grid, zone 4: base of section, 

2,273,200; 570,500; top of section, 2,272,100; 570,000]

Perdido Formation.

Erosional unconformity.

Sunday Canyon Formation (683 feet) : Feet 

5. Argillaceous limestone and lesser mudstone, darkgray, dense; weathers grayish orange to light 

brown. Massive to poorly bedded in 1- to 2-foot 

layers, which are interbedded with black chert 

layers as thick as 8 inches. Limestone studded 

with metamorphic minerals, most commonly tremolite(?). About 60 feet below top a 1-foot blue 

bioclastic limestone layer contains favositoid, alveolitoid, and cyathophylloid corals (W. A. Oliver, 

written communication, 1962)____________ 166 

4. Mudstone and chert, dark-gray to black; weathers 

light gray to moderate reddish brown, thin-bedded. 

Minor dark-gray limestone______________ 27

Sunday Canyon Formation (683 feet) Continued. Feet 

3. Limestone, argillaceous, dark-gray; weathers light 

gray or shades of yellow, orange, or red; very thin 

bedded, flaggy parting. Minor bioclastic layers as 

thick as 1 foot, particularly coral-rich bed 114 feet 

below top of unit____________________ 217 

2. Argillaceous limestone, dark-gray; weathers light 

gray in beds generally 1 to 2 feet thick alternating 

with blue-gray limestone layers of about the same 

thickness that are rich in fragments of pelmatozoans, corals, and bryozoans. Argillaceous limestone dominant in upper 100 feet of unit_____ 140 

1. Limestone, argillaceous, and calcareous shale, medium-gray ; weathers light gray to yellowish-gray ; 

very thin bedded to laminated, platy splitting. 

Clastic fossil fragments in lower 55 feet of unit. 

This unit contains Honograptus at several localities, not along the measured section but only a 

short distance to the north. At other localities 

Monograptus occurs from near the base of the formation to as high as 300 feet above the base___ 133 

Conformable contact, poorly exposed. 

Ely Springs Dolomite.

The maximum exposed thickness of the Sunday 

Canyon is 683 feet at the type section. The base of the 

formation appears to be conformable on the Ely 

Springs Dolomite, and the contact is placed at the top 

of the highest dolomite or massive black chert. On 

the basis of fossil evidence, an erosional unconformity 

separates the Sunday Canyon Formation from the 

overlying Perdido Formation of Mississippian age. 

This contact, however, is one of apparent conformity, 

in contrast to the upper contact of the Vaughn Gulch 

Formation, where erosion is evident. Near Badger 

Flat the Sunday Canyon thins and vanishes, presumably due to erosion although faulting may be responsible. Placement of the upper contact of the Sunday 

Canyon Formation is further complicated where shale 

in the Perdido Formation is in contact with shale in 

the Sunday Canyon. However a decided color break 

generally separates light-gray to yellowish calcareous 

shale here assigned to the Silurian from dark-gray, 

red-weathering noncalcareous shale assigned to the 

Perdido Formation of Mississippian age. Thin stringers of bioclastic limestone, which are present in the 

light-colored calcareous shales, are absent in the dark 

noncalcareous shales. These factors have led me to 

place the upper contact of the Sunday Canyon Formation at the break between the calcareous and noncalcareous shales.

The Sunday Canyon Formation contains a mixed 

sequence of calcareous siltstone, calcareous shale, and 

argillaceous limestone (see section, p. B83). The formation is typically thin bedded, and shaly to flaggy 

fragments make diagnostic weathered slopes. Though 

mostly dark to medium gray on fresh surfaces, these
B84 STRATIGRAPHY AND STRUCTURAL GEOLOGY

rocks weather to distinctive light gray, commonly with 

yellow and orange tints. Bioclastic blue-gray limestone beds and lenses (tongues of the Vaughn Gulch 

facies) are more common near Water Canyon, but some 

can be traced almost to Badger Flat. Black chert beds 

and nodules, and subordinate lenses of quartzite and 

calcareous quartz sandstone, are associated with some 

bioclastic layers, particularly in the central part of the 

formation.

Graptolites have been collected in the shaly to flaggy 

argillaceous limestone, calcareous shale, and siltstone, 

generally near the base of the formation, from at least 

10 localities from Water Canyon to the south edge of 

Badger Flat. The graptolites have been identified by 

E. J. Eoss, Jr., who reports the following forms (written communications, 1960, 1961) :

Monograptus vulgaris Wood

Monograptus vomerinus Nicholson

Monograptus aff. M. twnescens Wood 

These forms are considered by him to indicate a late 

Wenlock or early Ludlow (mid-Silurian) age.

The bioclastic limestone tongues of the Vaughn Gulch 

facies in the Sunday Canyon Formation have yielded 

corals which are generally poorly preserved and commonly fragmented. The following coral forms have 

been identified by W. A. Oliver, Jr. (written communications, 1960,1962) :

Alveolites sp.

Favosites sp.

Thamnopora sp.

"Cystiphyllum" sp.

cladoporoid coral

cerioid rugose coral

cyathophylloid(?) coral

cylindrical rugose coral

Oliver indicates that the age assignment based on these 

forms can be no closer than Silurian or Devonian.

Tentaculites cf. T. bellulus Hall has been identified by 

E. J. Eoss, Jr. (written communication, 1961), from a 

locality about 300 feet above the base of the formation 

11/2 miles north of the type section, where it occurs along 

with Monograptus. Ostracodes, sponge spicules, and 

brachiopods that are probably rhynchonellids have also 

been collected from the formation. The Sunday Canyon 

Formation is considered to be of Silurian age.

Facies relations

Throughout a large area of southeastern California 

and parts of southern Nevada the Silurian is represented primarily by cherty dolomite, to which the name 

Hidden Valley Dolomite is generally applied. In the 

Quartz Spring area, where it was originally defined 

(McAllister, 1952, p. 15), this formation is largely

Silurian, although fossils of Early Devonian age are 

found near the top.

The widespread Hidden Valley Dolomite is separated from the Vaughn Gulch Limestone by a severalmile span of granitic rocks, but a paleontologic tie between the two formations is furnished by the large 

dasycladacean algae Verticittopora annulata Eezak, 

which occurs in the Vaughn Gulch Limestone as well as 

in the Hidden Valley Dolomite in the southern part of 

the Inyo Mountains (C. W. Merriam, written communication, 1962). The contact of the Vaughn Gulch with 

the underlying Ely Springs has been placed at an 

obvious mappable point the top of the uppermost dolomite but fossils are found only in the lower part of 

the Ely Springs and it is possible that the upper part 

of the Ely Springs in the Independence quadrangle 

could be equivalent to part of the Hidden Valley. Just 

below the type section of the Vaughn Gulch Limestone 

in the upper Ely Springs, dark-gray limestone and 

bedded chert are present. This unique occurrence of 

limestone in the Ely Springs in the quadrangle could 

represent an intertonguing of the Hidden Valley dolomite facies and the Vaughn Gulch coral-limestone facies 

in the upper part of what has been mapped as Ely 

Springs.

As already discussed, the Vaughn Gulch coral-limestone facies grades northward into the Sunday Canyon 

graptolite facies, which is terminated by erosion or 

faulting north of Badger Flat. Graptolites at the south 

edge of Badger Flat mark the northwesternmost occurrence of fossiliferous Silurian rocks in the Great 

Basin and, in fact, fossiliferous Silurian rocks do not 

reappear to the north or west for several hundred miles. 

The thick stratigraphic section in the Mount Morrison 

roof pendant in the Sierra Nevada, however, may well 

include Silurian rocks in the several thousand feet of 

unfossiliferous beds that conformably overlie graptolite-bearing Ordovician rocks (Einehart and others, 

1959).

REFERENCES

Barnes, Harley, and Byers, F. M., Jr., 1961, Windfall formation (Upper Cambrian) of Nevada Test site, and vicinity, 

Nevada:' Art. 188 in U.S. Geol. Survey Prof. Paper 424-C, 

p. C103-C106.

Cooper, G. A., 1956, Chazyan and related brachiopods: Smithsonian Misc. Colln., v. 127, pt. 1, p. 1025-1245.

Hall, W. E., and MacKevett, E. M., Jr., 1962, Geology and ore 

deposits of the Darwin quadrangle, Inyo County, California : V.S. Geol. Survey Prof. Paper 368, 87 p.

Hazzard, J. C., 1937, Paleozoic section in the Nopah and Resting Springs Mountains, Inyo County California: California Jour. Mines and Geology, Rept. 33, State Mineralogist, p. 273-339.
ROSS B85

Knopf, Adolph, 1918, A geologic reconnaissance of the Inyo 

Range and the eastern slope of the southern Sierra Nevada, 

California, with a, section on Stratigraphy of the Inyo 

Range, by Edwin Kirk: U.S. Geol. Survey Prof. Paper 110, 

p. 19-^8.

Langenheim, R. L., Jr., and others, 1956, Middle and Upper (?) 

Ordovician rocks of Independence quadrangle, California: 

Am. Assoc. Petroleum Geologists Bull., v. 40, no. 9, p. 

2081-2097.

McAllister, J. F., 1952, Rocks and structure of the Quartz 

Spring area, northern Panamint Range, California: California Div. Mines Spec. Rept. 25, p. 1-38.

Nolan, T. B., 1943, The Basin and Range province in Utah, 

Nevada, and California: U.S. Geol. Survey Prof. Paper 

197-D, p. 141-196.

Phleger, F. B., Jr., 1933, Notes on certain Ordovician faunas of 

the Inyo Mountains, California: Southern California Acad. 

Sci. Bull., v. 32, pt. 1, p. 1-21.

Rinehart, C. D., Ross, D. C., and Huber, N. K., 1959, Paleozoic 

and Mesozoic fossils in a thick stratigraphic section in the 

eastern Sierra Nevada, California : Geol. Soc. America Bull., 

v. 70, p. 941-946.

Ross, D. C., 1962, Preliminary geologic map of the Independence 

quadrangle, Inyo County, California: U.S. Geol. Survey 

Mineral Inv. Map MF-254.

Stauffer, C. R., 1930, The Devonian of California: California 

Univ. Publications, Geol. Sci., v. 19, no. 4, 81-118.

Waite, R. H., 1953, Age of the "Devonian" of the Kearsarge 

area, California [abs.] : Geol. Soc. America Bull., v. 64, no. 

12, pt. 2, p. 1521.

Waring, C. A., and Huguenin, Emile, 1917, Inyo County, Fifteenth report of the State Mineralogist, California State 

Mining Bureau, p. 29-134.
Article 22

THE LIVINGSTON GROUP OF SOUTH-CENTRAL MONTANA

By ALBERT E. ROBERTS, Denver, Colo.

Abstract. The western part of the Crazy Mountains basin 

in south-central Montana contains more than 13,000 feet of 

continental deposits of latest Cretaceous and Paleooene age, 

customarily assigned to the Livingston Formation. In the type 

area, the lower half of this sequence is redefined as the Livingston Group, and,is divided into four new units: the Cokedale, 

Miner Creek, Dillman Greek, and Hoppers Formations; the 

upper half is assigned to the Fort Union Formation.

INTRODUCTION

Epeirogenic arching in the vicinity of the Elkhorn 

Mountains, Mont., began late in the Santonian Stage 

of Late Cretaceous time with truncation of the Eagle 

Sandstone and Colorado Shale along parts of the east 

side of the uplift. This uplift and erosion was accompanied by the volcanism which created the thick Elkhorn Mountains Volcanics (Klepper and others, 1957). 

Following withdrawal of the Upper Cretaceous Eagle 

seas in western Montana, the area east of the Bridger 

Range and north of the Beartooth Range gradually 

subsided to form the Crazy Mountains basin. This 

basin, elongated northwest, is 40 to 70 miles wide and 

130 miles long (fig. 22.1). Subsidence and deposition 

were greater in the western part of the basin and were 

accelerated in late Campanian and Maestrichtian time. 

The stratigraphic sequence there indicates continuous 

deposition during latest Cretaceous and Paleocene time. 

More than 13,000 feet of sediments, derived predominantly from volcanic rock, was deposited. These rocks 

have been generally assigned to the Livingston Formation. In this article the Livingston is restricted to include only the lower half of the continental sequence, 

and is raised to group rank and subdivided into four 

new formations.

The thick sequence of sedimentary rocks, consisting 

chiefly of debris of andesitic lava and other volcanic 

rocks "typically developed in the vicinity of Livingston, Mont.," was named the Livingston Formation by 

Weed (1893, p. 21). He divided it into three units: the 

leaf beds (lowest), the volcanic agglomerates, and the

110° 109° 108°

FIGURE 22.1. Index map of Montana showing localities 

relative to area discussed. Hachured line shows approximate limits of Crazy Mountains basin.

conglomerates. The volcanic agglomerate unit occurs 

only locally 35 miles east of Livingston, and its stratigraphic position is uncertain. Peale (1896) extended 

the name Livingston westward to include predominantly pyroclastic deposits in Jefferson Canyon, southwest of Three Forks, Mont.; near Sphinx Mountain in 

the Madison Range; and in the Maudlow area at the 

north end of the Bridger Range (Peale, in Weed, 1893, 

p. 21). Iddings and Weed (1894) and Weed (1899) 

assigned the strata younger than the Eagle Sandstone 

in most of the Crazy Mountains basin to the Livingston 

Formation. Coal-bearing rocks in the northern part 

of the basin were first correlated with the Eagle Sandstone by Stone (1908, p. 78-79); however, he placed the 

Livingston in the Tertiary. Two years later Stone and 

Calvert (1910, p. 551) demonstrated that the Livingston Formation grades laterally northeastward into 

marine and nonmarine beds of the Claggett, Judith 

River, Bearpaw, Lennep, and Lance (Hell Creek) Formations of Late Cretaceous age and the Lebo Member

ART. 22 IN U.S. GBOL. SURVEY PROF. PAPER 475-B, PAGES B86-B92. 1963

BS6
ROBERTS B87

of the Fort Union Formation of Paleocene age. No 

further stratigraphic studies of the Livingston were 

made in the southwestern part of the basin until McMannis (1955, p. 1407) and Richards (1957, p. 420) divided the formation into five generalized lithologic 

units including the Fort Union Formation.

The writer is deeply indebted to James R. Gill, William A. Cobban, and William J. McMannis for geologic 

information and discussions of stratigraphic problems.

STRATIGRAPHY

The stratigaphic sequence overlying the Eagle Sandstone at Cokedale, Mont., is a thick alternating series of 

coarse- and fine-grained continental deposits of the Livingston Group and Fort Union Formation. The coarsegrained units in the lower part of the Livingston Group 

(Cokedale and Miner Creek Formations) are more continuous and sheetlike than the typical stream-channel 

and piedmont alluvial deposits that are common in the 

upper part of the Livingston (Billman Creek and Hoppers Formations) and in the overlying Fort Union 

Formation. Many of the coarse-grained units are 

conglomeratic.

The Livingston Group thins progressively eastward; 

the sequence is 6,445 feet thick at Cokedale but is represented by only 2,700 feet of marine and nonmarine beds 

at Columbus, Mont. It also thins northward to Lennep, 

Mont., where it is represented by 4,260 feet of marine 

and nonmarine beds. The overlying Fort Union Formation likewise thins markedly to the east and north. 

In both these directions, rocks of the Livingston and 

Fort Union become finer grained, better sorted, thinner bedded, lighter in color, and less andesitic in composition owing to an increase in the proportions of quartz 

and orthoclase. Interfingering of the nonmarine Livingston with marine units likewise increases to the east 

and north.

The petrology of the volcanic material in the Livingston Group and Fort Union Formation indicates 

that it was derived principally from the Elkhorn 

Mountains, although several other source areas near the 

southern, western, and northwestern margins of the 

Crazy Mountains basin have been suggested. A local 

source near Flathead Pass in the Bridger Range was 

suggested by Weed (1893, p. 29) and another near the 

jDastle Mountains was suggested by Tanner.1 Another 

local source south of Big Timber, Mont., was first described by Weed (1893, p. 26-29) for his "volcanic 

agglomerates" unit and later by Parsons (1942, p. 1177) 

for the "Livingston igneous series." Billingsley (1916, 

p. 35) stated that the andesite conglomerate at Maudlow

1 Tanner, J. J., 1949, Geology of the Caatle Mountain area, Montana: 

Prlnceton Univ. Ph.D. thesis. [Available on microfilm from TJniv. 

Microfilms, Inc., Ann Arbor, Mich. ]

(equivalent to part of the Livingston) could have no 

other source than the lavas of the Elkhorn Mountains. 

McMannis (1955, p. 1412) and Klepper and others 

(1957, p. 40) also concluded that the thick deposits of 

volcanic material in the Elkhorn Mountains were most 

likely the source of much of the Livingston Formation.

An eastward increase in abundance of quartz and 

other minerals of nonvolcanic origin and a decrease in 

grain size across the Crazy Mountains basin, and the 

presence of volcanic rock known only from the Elkhorn Mountains Volcanics in conglomerate beds of the 

Livingston Group, corroborates the conclusion of Billingsley, McMannis, and Klepper that the predominant 

source was to the west and northwest of the basin.

The volcanic material in the Livingston Group includes fragments of several types and textures of andesitic and possibly some latitic rocks as well as beds of 

bentonite and andesitic ash. Also present in smaller 

amounts are lithic fragments of welded tuff, spherulitic 

and devitrified rocks that may be dacitic in composition, 

crystal fragments of volcanic and nonvolcanic origin, 

and clay minerals. Plagioclase is the most abundant 

mineral of volcanic origin, followed by pyroxene, 

amphibole, biotite, and magnetite. The plagioclase 

ranges from oligoclase to labradorite; the most common 

composition is andesine (An40 ). Some of the finer 

grained clayey rocks contain approximately 45 percent montmorillonite and 55 percent mixed-layer 

mpntmorillonite-illite-vermiculite; some also have small 

amounts of illite and chlorite (L. G. Schultz, oral communication, 1962). Quartz, probably derived from 

granitic, metamorphic, and sedimentary rocks, is the 

principal nonvolcanic constituent. The ratio of quartz 

to plagioclase in the Livingston is approximately 2 to 

1. Orthoclase occurs locally, generally in very small 

quantities, and ranges from a trace in the lower part 

of the Livingston Group to as much as 15 percent in 

the upper part.

Zeolite minerals are common in interstices. Reddishbrown heulandite is the most conspicuous zeolite and 

is present throughout the Livingston, but is most abundant in the Miner Creek Formation. Analcite is present 

in some units as cement, producing a mottled texture 

owing to crystallization of the analcite in clusters. 

Traces of laumontite and clinoptilolite are also present 

in some units. Other cementing agents are silica and 

calcite.

The Livingston Group includes the following formations, in ascending order: Cokedale, Miner Creek, Billman Creek, and Hoppers. The general correlation and 

stratigraphic relations of the type Livingston Group 

near Livingston, Mont., with other areas in Montana 

and Wyoming are shown in figure 22.2. The Cokedale 

Formation is the westward nonmarine equivalent of the
AHdva)Dii,vai,s QJSEV Tvaruonai-s 889
ROBERTS B89

Claggett Shale, Judith River Formation, and Bearpaw 

Shale. The Cokedale is the lower part of the leaf-beds 

member of Weed's (1893, p. 22) Livingston Formation. 

The Miner Creek Formation correlates eastward with 

the lower part of the Hell Creek Formation, and the 

basal sandstone unit, the Sulphur Flats Sandstone 

Member, is the nonmarine f acies of the marine Lennep 

Sandstone. The Miner Creek is the upper part of 

Weed's leaf beds. The Billman Creek Formation correlates approximately with the middle part of the Hell 

Creek Formation. The Billman Creek is in the lower 

part of the conglomerates member of Weed's (1893, 

p. 30) Livingston Formation. The Hoppers Formation 

correlates approximately with the upper part of the 

Hell Creek Formation. The Hoppers is also in the 

lower part of Weed's conglomerates member.

The Livingston Group at Cokedale is overlain by a 

nonmarine sequence 6,615 feet thick that is assigned to 

the Fort Union Formation on the basis of stratigraphic 

position and tectonic implication from rock types. The 

Fort Union includes three lithologic units: a lower conglomeratic sandstone member 975 feet thick, a middle 

member of sandstone and claystone 3,840 feet thick, and 

an upper conglomeratic sandstone member 1,800 feet 

thick, with an erosion surface at the top. The lower 

member is of Late Cretaceous age and the middle member is of Paleocene age, based on fossil evidence. Conglomerate in the Livingston Group consists chiefly of 

Cretaceous volcanic-rock fragments, whereas the conglomerate in the overlying Fort Union contains igneous, metamorphic, and sedimentary fragments derived 

from Precambrian, Paleozoic, and Mesozoic rocks. The 

Fort Union Formation is the upper part of Weed's 

conglomerates member.

The type sections of the formations in the Livingston 

Group were measured from the abandoned town of 

Cokedale, in sec. 26, T. 2 S., R. 8 E., northeastward to 

sec. 7, T. 2 S., R. 9 E., as shown in figure 22.3.

Cokedale Formation. The Cokedale Formation is 

here named for the nonmarine section exposed in the 

Sy2 sec. 23, and NEi/i sec. 26, T. 2 S., R. 8 E., immediately north of the former coal-mining town of Cokedale, Mont. The type section conformably overlies the 

Eagle Sandstone and is 1,550 feet thick. The Cokedale 

consists of siltstone and sandstone with claystone, tuff, 

bentonite, and coal in the lower part.

Siltstone forms about 50 percent of the formation. 

It is usually massive bedded, tuffaceous, and commonly 

olive gray, weathering to yellowish gray. Carbonaceous siltstone containing abundant leaf impressions, 

pollen, and spores in the lower part indicates transition from brackish-water marine to continental deposi11

Cokedale 

(abandoned)

26

12

13

24

0)

25 3O

R. 8 E.

17

2O

29

FIGURE 22.3. Location of measured units in the Livingston Group near Livingston, Mont. 1 and 2, 

traverse of type section of Cokedale Formation ; 3 and

4. traverse of type section of Miner Creek Formation 

(3, type section of Sulphur Flats Sandstone Member) ;

5. 6, and 7, traverse of type section of Billman Creek 

Formation; 8 and 9, traverse of type section of 

Hoppers Formation. (Base from U.S. Geological 

Survey Hoppers 7^-minute quadrangle, Montana.)

tion. At a few places fresh-water mollusks are 

preserved.

Hard ridge-forming sandstone forms about 35 percent of the formation. The beds are generally dusky 

yellow green, weathering light olive; massive to thin 

bedded; crossbedded; and fine to coarse grained, in 

places conglomeratic. The sandstone contains angular 

to subrounded grains of volcanic rock, plagioclase 

(andesine), augite, hornblende, magnetite, biotite, and 

quartz. The cement is mainly silica but includes lesser 

amounts of clay, zeolites, and calcite. Pebbles of volcanic rock include welded tuff that is found only in 

the Elkhorn Mountains Volcamcs west of this area. 

These pebbles were first noted 600 feet above the base 

of the Cokedale Formation. The sandstone contains 

much petrified wood and, rarely, dinosaur bones.

Prominent beds of light-gray calcareous quartzose 

sandstone crop out at 90, 340, and 510 feet above the 

base of the Cokedale Formation; the uppermost bed is 

16 feet thick and probably is the western tongue of 

the Parkman Sandstone. These beds are better sorted
B90 STRATIGRAPHY AND STRUCTURAL GEOLOGY

than other sandstone in the Cokedale and contain a 

distinctive heavy-mineral suite similar to that in the 

Eagle Sandstone at Cokedale (Roberts, 1957, p. 42-45). 

The heavy minerals, in order of decreasing abundance, 

are magnetite, zircon, tourmaline, corundum, garnet, 

hornblende, staurolite, rutile, apatite, hematite, limonite, and leucoxene (R.F. Gantnier, written communication, 1962).

Claystone and sandstone from 660 to 680 feet above 

the base of the Cokedale Formation in the SWl/4 sec. 

23, T. 3 S., R. 8 E., contain an abundant pollen and 

spore flora that according to R. H. Tschudy (written 

communication, 1962) is definitely older than the flora 

in the Mitten Black Shale Member of the Pierre Shale 

and most nearly resembles the flora from the Blair 

Formation of the Rock Springs region, Wyoming. 

Fragments of \Monoclonius sp. that commonly occur 

in the Judith River Formation in Montana and in 

the "Pale Beds" at the top of the Belly River Formation in Alberta were identified by G. E. Lewis (written 

communication, 1963) from a sandstone 710 feet above 

the base of the Cokedale.

The basal contact of the Cokedale Formation and 

the Livingston Group is not distinct. Weed (1893, 

p. 11) distinguished the volcanic sediments of the Livingston Formation from the underlying nonvolcanic 

coal-bearing formation (Eagle Sandstone) on the basis 

of lithology and believed the two were separated by an 

unconformity. Stone and Calvert (1910, p. 761) correctly described the Livingston Formation as conformably overlying the coal measures. Roberts (1957, p. 

47) arbitrarily designated the top of the arkosic sandstone that overlies the uppermost minable coal bed 

(Cokedale No. 5) as the contact between the Livingston 

and the underlying Eagle Sandstone, and this is taken 

as the base of the Cokedale. There is some andesitic 

sandstone in the upper part of the Eagle Sandstone and 

there is some arkosic sandstone and coal in the lower 

part of the Cokedale Formation; nevertheless, the 

boundary used is the best mappable contact in the 

Livingston area. The difference between physical and 

chemical properties of the coals in the upper part of 

the Eagle (bituminous-coking) and the lower part of 

the Cokedale (lignite-noncoking) also supports this 

boundary assignment. The thin black chert-pebble 

conglomerate that marks the top of the Eagle or base 

of the Claggett throughout much of Montana was not 

found in the section at Cokedale.

Miner Greek Formation, The Miner Creek Formation is here named for the nonmarine section exposed 

along Miner Creek in the E^ sec. 19 and NW% sec. 

20, T. 2 S., R. 9 E. (fig. 22.3)". The type section conformably overlies the Cokedale Formation and is 1,350

feet thick. The Miner Creek consists largely of alternating beds of siltstone and sandstone. At the base is 

a prominent ridge-forming sandstone and tuff unit, 160 

feet thick, that is here named the Sulphur Flats Sandstone Member for the section exposed along Miner Creek 

near Sulphur Flats in the SEi/4 sec. 19, T. 2 S., R. 9 E.

Massive-bedded tuffaceous siltstone forms 75 percent 

of the Miner Creek Formation. The siltstone in the 

upper part of the formation is olive gray and weathers 

light olive gray. In the middle of the formation it is 

dusky yellow green, weathering grayish yellow green, 

and in the lower part it is generally grayish olive green, 

weathering grayish yellow green. Red specks and veinlets of heulandite characterize the formation. Analcite 

and laumontite are present in traces.

Sandstone forms about 20 percent of the formation. 

It is generally well bedded and fine grained except for 

that in the Sulphur Flats Sandstone Member, which is 

massive, crossbedded, and poorly sorted. The sandstone is grayish green, weathering light olive gray, and 

is composed of andesitic volcanic-rock fragments, 

plagioclase (andesine), augite, hornblende, and magnetite; the Sulphur Flats Sandstone Member in addition contains much biotite and diopside. A distinctive 

fine-grained calcareous mottled yellowish-gray quartzose sandstone, 6 to 20 feet thick, forms an excellent 

marker bed at the base of the Sulphur Flats Sandstone 

Member. In the lower part of the formation, particularly in the Sulphur Flats Sandstone Member, tuff and 

bentonite are interbedded with the sandstone. Most of 

these fine-grained pyroclastic rocks have been altered 

to indurated siliceous claystone. Petrified wood, pollen, 

spores, and fragments of dinosaur bones are present in 

the Sulphur Flats Sandstone Member.

Stanton (in Stone and Calvert, 1910, p. 655) identified a marine fauna from the Livingston Formation 

near Wilsall, Mont., in the NEi/4 sec. 25, T. 4 N., R. 8 

E., as Late Cretaceous, probably occupying a position 

near the top of the Bearpaw Shale. James R. Gill and 

the writer visited this locality and obtained fossils identified by W. A. Cobban (written communication, 1961) 

as representing the Baculites compressus zone or possibly the slightly younger Baculites cuneatus zone in the 

upper Campanian of Late Cretaceous age (fig. 22.2). 

The marine rocks from which these fossils were obtained 

can be traced from Wilsall northward into the type section of the Lennep Sandstone at the north end of the 

Crazy Mountains (J. R. Gill, oral communications, 

1962). To the south, in the Livingston area, these beds 

are represented by the nonmarine Sulphur Flats Sandstone Member of the Miner Creek Formation.

A carbonaceous claystone 420 feet above the base of 

the Miner Creek Formation in sec. 19, T. 2 S., R. 9 E.,
ROBERTS B91

and a similar clay stone 1,330 feet above the base of the 

Miner Creek in sec. 20, T. 2 S., R. 9 E., contain a diagnostic pollen and spore flora, according to E. B. Leopold 

(written communication, 1962). Of these samples the 

genus of greatest stratigraphic interest is Aqmlapollenites, of which five species occur. The pollen assemblage is closely similar to that of a zone rich in Aquilapollenites in the upper part of the Pierre Shale.

Billman Creek Formation. The Billman Creek Formation is here named for the nonmarine section exposed 

near Billman Creek in the Sy2 sec. 13, T. 2 S., R. 8 E., 

and the Wy2 sec. 18, T. 2 S., R. 9 E. (fig. 22.3). The 

type section conformably overlies the Miner Creek 

Formation and is 2,590 feet thick. The Billman Creek 

is mostly claystone but includes lesser amounts of sandstone, siltstone, and conglomerate. It is less resistant 

to weathering than adjacent formations and generally 

forms valleys. Claystone comprises about 65 percent 

of the formation and is usually massive bedded. Three 

samples average 40 percent montmorillonite, 50 percent 

mixed-layer montmorillonite-illite-venmculite, 5 percent illite, and 5 percent chlorite (L. G. Schultz, oral 

communication, 1962). In the upper part of the formation the claystone is olive black to olive gray, weathering light olive gray; in the middle part it is mostly 

grayish brown, weathering pale yellowish brown; and 

in the lower part it is mainly dusky red, weathering 

grayish red, but partly grayish green, weathering grayish yellow green.

Channel-fill sandstone and conglomerate form about 

25 percent of the formation. They are generally dusky 

yellow green, weathering grayish yellow green, and are 

composed of volcanic-rock fragments, plagioclase 

(andesine), augite, diopside, hornblende, magnetite, 

biotite, quartz, and orthoclase cemented by silica, calcite, clay, or zeolites. A 20-foot continuous sandstone 

unit 96 feet above the base of the Billman Creek Formation contains abundant spherical calcareous sandstone concretions and is an excellent marker bed.

Dinosaur bones of Lance age were described from 

several localities in the upper part of the Billman Creek 

Formation 13 miles east of Bozeman, Mont. (McMannis, 1955, p. 1408). Fresh-water mollusks were 

collected from the middle part of the Billman Creek 

and identified by D. W. Taylor (written communication, 1962). The species are known from much of the 

Cretaceous and could not be precisely dated, but were 

in accord with the age assigned on the basis of fossil 

vertebrates.

Hoppers Formation. The Hoppers Formation is 

here named for the nonmarine section exposed near 

Hoppers Siding on the Northern Pacific Railway in 

the SW% sec. 7 and NW% sec. 18, T. 2 S., R. 9 E. (fig.

680231 O 63 7

22.3). The type section conformably overlies the Billman Creek Formation and is 965 feet thick. The 

Hoppers consists largely of sandstone with interbedded claystone and siltstone.

Ridge-forming sandstone makes up 60 percent of the 

Hoppers Formation and is generally massive to thin 

bedded, crossbedded, and poorly sorted. It is dusky 

yellow green and generally weathers olive gray; however, the basal 140 feet weathers a conspicuous yellowish gray in massive spheroidal shapes. The sandstone 

consists of angular to subrounded grains of volcanic 

rock, plagioclase (andesine), augite, hornblende, magnetite, biotite, quartz, and orthoclase. Silica and calcite 

are the common cement. Many of the sandstone layers 

are channel-fill deposits containing pebbles of clay and 

volcanic rock, and fragments of wood and other plants. 

The basal sandstone beds of the Hoppers are a cliffforming unit composed of massive, crossbedded, calcareous, generally conglomeratic sandstone with interbedded claystone and siltstone. These sandstone beds 

are generally banded by layers of ferromagnesian minerals, particularly magnetite.

About 25 percent of the formation is massive-bedded 

claystone. Iti s olive gray, weathering light olive gray. 

A few beds contain calcareous claystone concretions and 

at least one contains fresh-water mollusks.

Siltstone, forming about 15 percent of the formation, 

is generally gradational with sandstone and similar in 

color and mineral composition.

REFERENCES

Billingsley, Paul, 1916, The Boulder batholith of Montana: Am.

Inst. Mining Engineers Trans, v. 51, p. 31-57. 

Cobban, W. A., 1958, Late Cretaceous fossil zones of the Powder

River basin, Wyoming and Montana, in Wyoming Geol.

Assoc. Guidebook 13th Ann. Field Conf., Powder River

basin: p. 114-119.

 1962a, New Baculites from the Bearpaw Shale and 

equivalent rocks of the western interior: Jour. Paleontology, v. 36, no. 1, p. 126-135.

 1962b, Baculites from the lower part of the Pierre Shale

and equivalent rocks in the western interior: Jour. Paleontology, v. 36. no. 4, p. 704-718.

Cobban, W. A., Erdmann, C. E., Lemke, R. W., and Maughan, 

E. K., 1959, Revision of Colorado Group on Sweetgrass 

Arch, Montana: Am. Assoc. Petroleum Geologists Bull., 

v. 43, no. 12, p. 2786-2796.

Gill, J. R., and Cobban, W. A., 1961, Stratigraphy of- lower and 

middle parts of the Pierre Shale, northern Great Plains: 

Art. 352 in U.S. -Geol. Survey Prof. Paper 424-D, p. D185- 

D191.

 1962, Red Bird Silty Member of the Pierre Shale, a new 

stratigraphic unit: Art. 8 in U.S. Geol. Survey Prof. Paper 

450-B, p. B21-B24.

Hancock, E. T., 1918, Geology and oil and gas prospects of the 

Lake Basin field, Montana: U.S. Geol. Survey Bull. 691-D, 

p. 101-147.
B92 STRATIGRAPHY AND STRUCTURAL GEOLOGY

Iddings, J. P., and Weed, W. H., 1894, Description of the Livinston sheet [Montana] : U.S. Geol. Survey Geol. Atlas, 

Folio 1, 5 p., 4 maps.

Klepper, M. R., Weeks, R. A., and Ruppel, B. T., 1957, Geology of the southern Blkhorn Mountains, Jefferson and 

Broadwater Counties, Montana: U.S. Geol. Survey Prof. 

Paper 292, p. 82.

Lyons, J. B., 1944, Igneous rocks of the northern Big Belt 

Range, Mont.: Geol. Soc. America Bull., v. 55, no. 4, p. 

445-472.

McMannis, W. J., 1955, Geology of the Bridger Range, Montana : 

Geol. Soc. America Bull., v. 66, no. 11, p. 1385-1430.

Parsons, W. H., 1942, Origin and structure of the Livingston 

igneous rocks, Montana: Geol. Soc. America Bull., v. 53, 

no. 8, p. 1175-1186.

Peale, A. C., 1896, Description of the Three Forks sheet [Montana] : U.S. Geol. Survey Geol. Atlas, Folio 24, 7 p., 4 maps.

Richards, P. W., 1955, Geology of the Bighorn Canyon-Hardin 

area, Montana and Wyoming: U.S. Geol. Survey Bull. 1026, 

93 p.

 1957, Geology of the area east and southeast of Livingston, Park County, Montana: U.S. Geol. Survey Bull. 

1021-L, p. 385-438.

Roberts, A. E., 1957, Coal-bearing rocks and mines at Cokedale, 

Park County, Mont., in Billings Geol. Soc. Guidebook 8th 

Ann. Field Conf., Crazy Mountain Basin, Montana, 1957: 

p. 39-47.

Robinson, C. S., Mapel, W. J., and Cobban, W. A., 1959, Pierre 

shale along western and northern flanks of Black Hills, 

Wyoming and Montana: Am. Assoc. Petroleum Geologists 

Bull., v. 43, no. 1, p. 101-123.

Schmidt, R. G., and Zubovic, Peter, 1961, Cobern Mountain overthrust, Lewis and Clark County, Montana: Art. 211 in U.S. 

Geol. Survey Prof. Paper 424-C, p. C175-C177.

Stone, R. W., 1908, Coal near the Crazy Mountains, Montana, 

in coal fields of North* Dakota and Montana: U.S. Geol. 

Stirvey Bull. 341-A, 120 p.

Stone, R. W., and Calvert, W. R., 1910, Stratigraphic relations of 

the Livingston Formation of Montana: Bcon. Geology, v. 5, 

no. 6, p. 551-557; no. 7, p. 652-669; and no. 8, p. 741-764.

Weed, W. H., 1893, The Laramie and the overlying Livingston 

Formation in Montana: U.S. Geol. Survey Bull. 105, 68 p.

 1899, Description of the Little Belt Mountains quadrangle [Montana] : U.S. Geol. Survey Geol. Atlas, Folio 56, 

11 p., 4 maps.

Zapp, A. D., and Cobban, W. A., 1960, Some Late Cretaceous 

strand lines in northwestern Colorado and northeastern 

Utah: Art. 112 in U.S. Geol. Survey Prof. Paper 400-B, 

p. B246-B249.
Article 23

AGE OF CERTAIN POST-MADISON ROCKS IN 

SOUTHWESTERN MONTANA AND WESTERN WYOMING

By J. T. DUTRO, JR., and WILLIAM J. SANDO, Washinston, D.C.

Abstract. Three Late Mississippian faunal zones are reported from the Big Snowy-Amsden interval. Beds with 

Caninia and Spirifer braseriamis are of probable middle Chester 

age. Striatif era-bearing strata are early Chester or latest Meramec equivalents. These zones are correlatives of zones present 

in the Upper Mississippian sequence of the Chesterfield Range, 

southeastern Idaho.

The Late Mississippian age of a part of the strata assigned to the Big Snowy-Amsden interval in southwestern Montana and western Wyoming has been recognized for some time. The history of the development 

of the Mississippian-Pennsylvanian boundary problem, 

as it involves the understanding of rocks assigned at 

various times to the "Amsden," "Big Snowy," or "Quadrant" Formations, has been completely treated by J. 

Steele Williams (1948, p. 327-351). More detailed 

stratJgraphic studies and a more critical analysis of 

some of the fossils were called for by Williams. The 

partial results of some of these studies, undertaken by 

the authors in the past 5 years or so, are summarized 

below.

In the Bridger Range, Mont. (fig. 23.1, loc. 2), McMannis (1955, p. 1401-1404, and table 1) divided this 

interval into five lithologic units. The lower three units 

were assigned to the Upper Mississippian; the upper 

two units were called Pennsylvanian. In the northern 

Gravelly Range, Mont. (fig. 23.1, loc. 3), Hadley (1960) 

recognized both the Big Snowy Group and the Amsden 

Formation; he assigned the Big Snowy Group to the 

Upper Mississippian and the Amsden to the Pennsylvanian and Mississippian, undivided. In the Elkhorn 

Mountains, Mont. (fig. 23.1, loc. 1), Klepper (Klepper 

and others, 1957) assigned beds called Amsden to the 

Mississippian and Pennsylvanian Systems. Rubey's 

(1958) map of the Bedford quadrangle, Wyoming (fig. 

23.1, loc. 4), also shows the Amsden as Mississippian 

and Pennsylvanian.

Recent field studies by the authors in these areas have 

revealed the presence of Upper Mississippian fossils

115° 113°

47'

45'

43°

, Helena

>., ._ MONTANA l ' 'I - '

1 '.WYOVnWfi

25 50 75 100 MILES

FIGURE 23.1. Index map showing areas discussed in 

text. (1) Indian Creek section, Elkhorn Mountains, Mont. (2) South Angler section, Bridger 

Range, Mont. (3) Northern Gravelly Range, 

Mont. (4) Bedford quadrangle, Wyoming. (5) 

Chesterfield Range, Idaho.

and faunal zones hitherto unreported; 2 brachiopod 

zones and 1 coral zone have been identified. These 

zones were first recognized in the Mississippian sequence 

exposed in the Chesterfield Range, southeastern Idaho, 

where the section has been under study by the writers 

since the'summer of 1959 (fig. 23.1, loc. 5).

ART. 23 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B93-B94, 1963.

B93
B94 STRATIGRAPHY AND STRUCTURAL GEOLOGY

The two upper brachiopod zones of the Idaho sequence 

are characterized by Spirifer brazerianus Girty at the 

top and Striatifera brazeriana (Girty) below. Beds 

containing Spirifer brazerianus are approximately the 

same as those containing Caninia, the most abundant 

coral genus of the uppermost Mississippian coral zone 

in the Chesterfield Eange sequence (fig. 23.2).

Caninia has been identified from the uppermost beds 

assigned to the Big Snowy Group in the Gravelly Eange 

by Hadley. The genus was collected also from near 

the top of unit 3 of McMannis' Big Snowy in the South 

Angler section in the Bridger Range. The Indian Creek 

section in the Elkhorn Mountains was measured by the 

writers during the summer of 1962, and Caninia was 

collected from the uppermost beds of the medial limestone unit in the Amsden Formation of Klepper. 

Caninia had been identified earlier by Sando from the 

Amsden Formation of Rubey at the Haystack Peak 

and Covey Cutoff Trail sections in the Bedford quadrangle, Wyoming.

At Indian Creek, fossils indicative of both brachiopod zones have been recognized. Spirifer brazerianus 

occurs together with Caninia in the upper beds. About 

40 feet lower, near the base of the limestone unit, 

specimens of Striatifera are present in large numbers. 

In the Bridger Range, Diapfiragmus was collected from 

the same strata that yielded Caninia. Although Diaphragmus is not restricted to the Spirifer brazerianus 

Zone, it is most common there. Its presence in the 

Bridger Range suggests strongly that the upper brachiopod zone is represented in unit 3 of McMannis. 

Spi.rifer brazerianus and other brachiopods characteristic of the upper brachiopod zone occur together with 

Caninia in the upper beds assigned to the Big Snowy 

Group by Hadley in the northern Gravelly Range. 

Finally, brachiopods characteristic of the Spirifer brazerianus Zone occur with Caninia in the Amsden of 

Rubey from the Covey Cutoff Trail section in the Bedford quadrangle.

The Caninia and Spirifer brazerianus Zones are 

considered Chester, but not latest Chester, in age. 

Younger Mississippian beds are known to be present

in central and western Utah (M. Gordon, Jr., oral 

communication). The Striatifera-bz&ring beds are of 

either early Chester or latest Meramec age.

FIGURE 23.2. Late Mississippian faunal zones in the northern 

Rocky Mountains with suggested correlations to the midcontinent region. Diagonal lines indicate gaps in faunal control; vertical lines indicate probable hiatus.

REFERENCES

Hadley, J. B., 1960, Geology of the northern part of the Gravelly 

Range, Madison County, Montana; in Billings Geol. Soc. 

Guidebook, llth Ann. Field Conf., West Yellowstone earthquake area, Montana, 1960: p. 149-153.

Klepper, M. R., Weeks, R. A., and Ruppel, E. T., 1957, Geology 

of the southern Elkhorn Mountains, Jefferson and Broadwater Counties, Montana: U.S. Geol. Survey Prof. Paper 

292, 82 p.

McMannis, W. J., 1955, Geology of the Bridger Range, Montana : Geol. Soc. America Bull., v. 66, no. 11, p. 1385-1430.

Rubey, W. W., 1958, Geology of the Bedford quadrangle, Wyoming: U.S. Geol. Survey Quad. Map GQ-109.

Williams, J. Steele, 1948, Mississippian-Pennsylvanian boundary 

problems in the Rocky Mountain region: Jour. Geology, v. 

mr no. 4, p. 327-351.
Article 24

EMENDATION OF THE KELVIN FORMATION AND 

MORRISON(?) FORMATION NEAR SALT LAKE CITY, UTAH

By MAX D. CRITTENDEN, JR., Menlo Park, Calif.

Abstract. The Kelvin Formation (Kelvin Conglomerate of 

Mathews, 1931, Cretaceous) is emended to include (a) the 

Parleys Member (new name), characterized by white modular 

limestone and called Morrison(?) Formation by Granger and 

Sharp (1952), and (b) an upper member consisting of conglomerate and siltstone. The Kelvin is correlated with the 

Cedar Mountain Formation.

Stratigraphic units previously assigned a Late Jurassic and Early Cretaceous age in the area immediately 

east of Salt Lake City, Utah, include (a) a thin white 

unit tentatively designated Morrison(?) by Granger 

and Sharp (1952, p. 14) and presumed to be Jurassic, 

and (b) a thick overlying unit consisting of reddishbrown conglomerate that they called the Kelvin Formation and assigned to the Cretaceous. Study of these 

units over a larger area indicates that both are probably Cretaceous. As a result, they are assigned to the 

Kelvin Formation, which is here emended to include 

two members.

KELVIN FORMATION

Name and history. The name Kelvin conglomerate 

was first used by Mathews (1931, p. 48) in describing 

the area of Parleys and Emigration Canyons.

The Kelvin conglomerate is named after Kelvin's Grove, a 

rather prominent and well established locality in Emigration 

Canyon. The formation is composed of very coarse, well 

rounded, well polished boulders cemented with a rather resistant 

cement. The boulders are from a few inches to two feet in 

diameter, the average ranging about 6 to 8 inches. The formation is somewhat variable in color, locally red, a reddish gray, 

and in some localities a pure gray. The formation is roughly 

stratified indicating a still water origin. The thickness of 

the formation appears to be somewhat variable.

Mathews did not include either a map or a measured 

section, and thus it is impossible now to be sure exactly 

what rocks he was talking about. The location of the 

original Kelvins Grove resort is uncertain, and the 

site determined by Granger on the basis of local inquiry (written communication, 1962) is in the unit now

designated as the Preuss Sandstone. Moreover, except 

for the reference to conglomerates described as "pure 

gray," there is nothing in Mathews' description to indicate that he saw the white nodular limestones that 

characterize the unit called the Morrison (?) Formation 

by Granger.

The term Kelvin (?) Formation was applied by Eardley (1944, p. 838) to a section near Peoa on the Weber 

Eiver that is equivalent to the Kelvin in the type area 

but is less conglomeratic.

The term Kelvin Formation was later applied by 

Granger and Sharp (1952, p. 15) and by Granger (1953, 

p. 4) in the type area to about 1,500 feet of pale-red 

siltstone, sandstone, and conglomerate that occupies the 

interval between the top of the white marker bed containing nodular limestone, and the base of the predominantly tan or yellowish sandstones designated the 

Frontier Formation. A measured section of this unit 

was publishe'd by Granger (1953, p. 11-12), but the 

formation was not formally redefined.

The name Morrison Formation was also used by 

Mathews in this area (1931, p. 47), but was applied to 

the beds immediately above the Twin Creek Limestone 

that are designated here and by Granger (1953, p. 4) 

as the Preuss Sandstone. The first use of the term 

Morrison (?) for the thin white marker beds was apparently by Granger and Sharp (1952, p. 14), though 

this use was attributed partly to unpublished work of 

Stokes and others. More recently, Stokes (1959, p. 

113) has suggested that this unit may be Cretaceous.

Definition. The Kelvin Formation as here defined 

consists of two members: a lower one, given the new 

name the Parleys Member, consisting of the distinctive 

light-colored unit called the Morrison(?) by Granger, 

and an upper member, unnamed, comprising the rocks 

to which the name Kelvin Formation was applied by 

Granger. The type section of the Parleys Member is 

in readouts on Highway 40 in the N%SWM sec. 8, T.

ART. 24 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B95-B98. 1963.

B95
B96 STRATIGRAPHY AND STRUCTURAL GEOLOGY

1 S., R. 3 E., and located 1.1 miles west of Parleys Summit. The section measured by Granger (1953, p. 1-12) 

on the north slope of Parleys Canyon about 4 miles 

east of the Mountain Dell Reservoir is designated the 

type section of the upper member. These two sections 

taken together constitute the type section of the formation as a whole.

Parleys Member

The distinctive unit here named the Parleys Member 

forms a prominent white marker about 100 feet thick 

that is plainly visible both on the ground and on aerial 

photographs, in Parleys and Emigration Canyons and 

as far east as the Weber River. It owes its striking 

appearance to beds or scattered nodules of fine-grained 

limestone that weathers pale gray to pure white, though 

such rocks make only about a third of the total thickness in the exposures west of Parleys Summit; the remainder is lavender-gray siltstone (40 percent), sandstone (20 percent), and conglomerate (10 percent). All 

the units are lenticular and vary rapidly in proportion along the strike.

The lower contact is drawn at the base of the first 

pale-gray, lavender-gray, or white beds above the continuous deep-red, chocolate-red, or brick-red sandstone 

of the 'underlying Preuss Sandstone. The lowest beds 

are only conglomeratic locally; in other places they contain nodules of limestone like those above. There is no 

evidence of erosional unconformity. The upper contact is drawn on a similar basis at the highest very light 

colored beds. These generally grade abruptly to the 

dark-red or reddish-brown sandstone and coarse conglomerate that are included in the upper member. The 

type section measured in exposures on U.S. Highway 

40 follows:

Type stratigraphic section of the Parleys Member of the 

Kelvin Formation along U.S. Route 40, in the Ny2 SWV^ sec. 

8, T. 1 8., R. 8 E., Salt Lake County, Utah.

[Section measured by M. D. Crittenden, 1959]

Upper member of the Kelvin Formation:

Sandstone, reddish-brown, coarse-grained, massive; includes

conglomerate at top and along strike 

Parleys Member of the Kelvin Formation: Feet

Siltstone, pale-lavender-gray; contains 30 to 40 percent 

white fine-grained limestone as nodules and concretions_______________________________ 20

Limestone (60 percent) and siltstone (40 percent) ; 

concretions and nodular masses of white to pale-gray 

limestone enclosed in lavender-gray siltstone_____ 11

Siltstone, reddish-brown, friable; contains scattered 

nodules of pale-brown fine-grained limestone (20 

percent) ____________________________ 9

Sandstone, pale-brown, fine-grained, well-cemented__ 2

Siltstone, pale-brown, friable_________________ 3. 5

Limestone, sandy, pale-gray, massive, nodular______ 3

Parleys Member of the Kelvin Formation Continued Feet

Siltstone, lavender-gray; contains nodules of white limestone (30 percent)____________ _ _ 8

Limestone, white -to very-pale-gray, locally sandy, massive; weathers in large nodular forms_________ 36

Siltstone (40 to 70 percent) and limestone (30 to 60 percent) ; medium-lavender-gray siltstone; nodules, concretions, and beds of pale-gray limestone; grades into 

overlying unit_________ 14

Sandstone, lavender-gray to brick-red, fine-grained; 

locally contains lenses of conglomerate; irregularly 

bleached white. _ 12

Claystone, dark-reddish-brown, friable _ 4.5

Siltstone and limestone; lavender-gray siltstone containing nodules and concretions of pale-buff to pinkish 

fine-grained limestone (up to 50 percent) 4

Siltstone, lavender-gray to reddish-brown; contains 

structureless sandstone as lenses % to 1 inch thick 9

Sandstone, pale-tan to pale-reddish-brown, fine-grained 

well-sorted; beds 1 to 3 ft thick _____ 6

Conglomerate composed of pebbles, some cobbles, fragments of friable sandstone; strongly crossbedded, locally well cemented; grades into overlying unit 6

Sandstone, pale-gray to lavender-brown, coarse-grained, 

poorly sorted, crossbedded; contains pebble conglomerate as irregular lenses; small areas of the unit are 

bleached pure white and recemented 9

Total thickness of the Parleys Member of the Kelvin

Formation 157 

Preuss Sandstone:

Sandstone and siltstone, micaceous, brick-red to pale-orangered, locally friable; weathers platy to massive; sandstone 

is ePQissbeidded in 1- to 2-inch beds.

Upper member

The upper member of the ^Kelvin Formation consists 

of pale-red to dark-grayish-red sandstone, siltstone, 

and conglomerate whose color contrasts strongly with 

that of the nearly white Parleys Member below. The 

bulk of the member consists of pale-red to reddishbrown sandstone and siltstone in alternating beds a few 

feet to 20 feet thick. A 90-foot bed of cobble and 

boulder conglomerate occurs at the base of the member, 

but pebbly zones and beds of pebble conglomerate occur sporadically higher in the section and account for 

about 10 percent of the thickness. The thick basal conglomerate is well exposed at the switchback in Emigration Canyon where the road turns southwest leading to 

Parleys Canyon.

Granger and Sharp (Granger, 1953, p. 11-12) 

record a thickness of 1,565 feet for the upper member 

on the north slope of Parleys Canyon opposite Lambs 

Canyon. It thickens and becomes somewhat less conglomeratic eastward, inasmuch as Morris 1 reports 

2,700 feet near Peoa on the Weber River, with pebble 

conglomerates in only the lower part.

» Morris, E. C., 1953, Geology of the Big Piney area : Utah Univ., 

Salt Lake City, M.S. thesis, Dept. Geology, p. 22.
CRITTENDEN B97

Correlation

The identity of the white beds here designated the 

Parleys Member of the Kelvin Formation and their 

Correlation with known Jurassic and Cretaceous rocks 

in adjoining areas have long been a matter of speculation. The historical development of the problem has 

been reviewed recently by Stokes (1959, p. 112), who 

points out that the crux of the matter is to determine 

how much, if any, of the rock in this area is equivalent 

to the Morrison.

The nearest outcrops that are unquestionably Morrison are li/4 miles north of Peoa on the Weber River, 

Some 15 miles east of Parleys Summit. At that place, 

Morris,2 with Stokes' concurrence (1959, p. 112), de- 

$cribed a 285-foot section of sandstone and siltstone 

^ith very minor amounts of limestone, and assigned 

these beds to the Morrison Formation. The colors are 

variable, but the siltstones are predominantly greenish 

j^ray or brownish red, and the sandstones are mainly 

pale pinkish gray. About 200 feet above the base is 

a 40-foot unit of scarlet siltstone that N. C. Williams 

and students have found to be a persistent and useful 

marker.

The base of the Morrison at Peoa is placed at the top 

Of greenish-gray fine-grained glauconitic sandstones appropriately assigned to the Stump Sandstone. The top 

is somewhat arbitrary, according to Stokes (1959, p. 

112), but was placed at the base of the first continuous 

>ed of conglomerate. The only limestones in the Morrison of this section are brownish weathering, impure, 

and dark gray, and occur in two beds 1% and 5 feet 

thick within 10 feet of the base. Several zones of limy 

Sandstone are present higher in the section, but none 

is persistent.

The part of Morris's section between the base and 

the scarlet marker bed has the aspect of typical Morri- 

$on of the Colorado Plateau, according to L. C. Craig 

(written communication, 1959). This conclusion is 

Supported by R. E. Peck (written communication, 1959) 

jwho reported to Craig the recovery of sparse charbphytes of Morrison aspect from the beds surrounding 

phe limestones. Above the scarlet bed, however, according to Craig". . . the sequence becomes dominantly 

lavender or purplish, gastroliths are common, limestone 

nodules form a rubble, and a few thin conglomeratic 

Sandstones are present. These features are more suggestive of the Cedar Mountain of the Plateau." Both 

ptokes and Craig therefore include in the Cretaceous 

till of the beds containing conglomerate in this section. 

! The question of where the white limestones of the 

Parleys Member fit in the Peoa section is now clear.

a Op. cit., p. 23.

They do not in the least resemble the thin dark limestones near the base of the Morrison section; on the 

other hand they are essentially identical with the white 

nodular limestones intercalated with the lavender-gray 

siltstones and conglomerates beginning some 300 feet 

above the scarlet marker. They are thus within the 

Kelvin Formation, and are about 340 feet above the 

top of the Morrison as placed by Craig, and 200 feet 

above it as placed by Stokes.

On the basis of this comparison with the Peoa section, which contains the westernmost recognizable and 

fossiliferous rocks of Morrison aspect, it is concluded 

that the white marker unit of Parleys Canyon is not 

Morrison. The decision to assign it member status depends on the fact that it is intercalated to the east with 

rocks like those above and below, and there becomes 

part of a unit that is difficult to divide on lithologic 

grounds.

As now defined, the Kelvin Formation is correlated 

with the Cloverly Formation of Wyoming, and with 

the Cedar Mountain Formation of southeastern Utah. 

This conclusion was reached tentatively by Stokes 

(1957, chart opposite p. 94) who showed the Morrison 

to be absent in Parleys Canyon, though he did not specifically mention the white beds in question, and represents an extension of the guidelines that Stokes laid 

down earlier (1952, p. 1770) for the division of the 

Jurassic and Cretaceous rocks in the western part of 

the Uinta Mountains: "(1) The Cretaceous beds contain most, and locally all, of the so-called 'gastroliths'; 

(2) the Cretaceous is marked by numerous limy nodules 

which are densely crowded together in places; and (3) 

the prevailing color of the Cretaceous is lavender, or 

purplish, and the color bands are not sharply marked."

The absence of the Stump and Morrison Formations 

.in Parleys Canyon indicates that there is a slight regional unconformity between the Kelvin Formation and 

the underlying Preuss sandstone, though they appear 

conformable in individual outcrops.

Age

The Kelvin Formation as here redefined is regarded 

as of medial Early Cretaceous age on the basis of correlation with the Peoa section, which establishes that 

it overlies the westernmost beds identifiable as Morrison Formation, and underlies the fish-scale-bearing beds 

of the Aspen Shale. No new faunal evidence has been 

obtained, though it is probably significant that R. E. 

Peck obtained charophytes typical of the Morrison from 

the rock immediately above the Stump Sandstone at 

Peoa, but failed to find any in the Parleys Member of 

the Kelvin in Emigration Canyon (letter to Craig, 

1959).
B98 STRATIGRAPHY AND STRUCTURAL GEOLOGY

REFERENCES

Eardley, A. J., 1944, Geology of the north-central Wasatch Mountains, Utah: Geol. Soc. America Bull., v. 55, p. 819-894.

Granger, A. E., 1953, Stratigraphy of the Wasatch Range near 

Salt Lake City, Utah: U.S. Geol. Survey Circ. 296, 14, p.

Granger, A. E., and Sharp, B. J., 1952, Geology of the Wasatch 

Mountains east of Salt Lake City, City Creek to Parleys 

Canyon: Utah Geol. Soc. Guidebook, no. 8, p. 1-37.

Mathews, A. A. L., 1931, Mesozoic stratigraphy of the Central 

Wasatch Mountains: Oberlin College Lab. Bull., new. ser., 

no. 1, 50 p.

Stokes, W. L., 1952, Lower Cretaceous in Colorado Plateau: 

Am. Assoc. Petroleum Geologists Bull., v. 36, No. 9, p. 

1766-1776.

 1957, Jurassic system of the southern flank of the Uinta 

Mountains, in Intermountain Assoc. Petroleum Geologists 

Guidebook, 8th Ann. Field Conf.: p. 92-96.

 1959 Jurassic Rocks of the Wasatch Range and vicinity, 

in Intermountain Assoc. Petroleum Geologists Guidebook, 

10th Ann. Field Conf.: p. 109-114.
Article 25

APACHE CREEK SANDSTONE MEMBER OF THE PIERRE SHALE 

OF SOUTHEASTERN COLORADO

By GLENN R. SCOTT and WILLIAM A. COBBAN, Denver, Colo.

Abstract. An 82-foot-thick sandstone bed that crops out on 

North Apache Creek, Huerf ano County, Colo., is here adopted as 

the Apache Creek Sandstone Member of the Pierre Shale. The 

member, shaly in the lower part and sandy above, overlies the 

transition member and underlies the Sharon Springs Member of 

the Pierre Shale.

The lower part of the Pierre Shale between Pueblo 

,and the Huerfano River contains a persistent sandy 

unit whose upper part was originally called the 

"Apache sandstone," later the Apache Creek Sandstone 

Member of the Pierre Shale. The original description 

jof the Apache Creek Sandstone Member did not designate a type locality or type section. The purpose of this 

[article is to redefine the member, establish its stratigraphic, boundaries, designate a reference locality, describe the reference section, and discuss its fossils, age, 

and correlation.

The first published reference to the Apache Creek 

was by Lavington (1933, p. 399) who stated, "In the 

Walsenburg district this zone [the barren zone] contains a 20-foot sandstone member which H.W.C. Prommel in an unpublished report has called the Apache 

sandstone." Later, in a catalog of stratigraphic names 

used in the Raton basin and vicinity, Mitchell and others 

(1956, p. 131) referred to this unit as the Apache 

Creek Sandstone Member of the Pierre Shale. Johnson and Stephens (1954) reported a 10-foot bed of sandstone approximately 500 feet above the base of the 

Pierre Shale on the western flank of the Greenhorn 

anticline about a mile south of the Huerfano River, 

and stated that this sandstone may be what Prommel 

called the "Apache sandstone." However, Johnson 

(1958, p. 563) stated that this sandstone does not 

crop out in the Walsenburg area.

The reference locality of the Apache Creek Sandstone Member is designated as the west-facing valley 

wall on the east side of a valley tributary to North 

Apache Creek in the SE}4 sec. 30, T. 25 S., R. 67 W.,

Huerfano County, Colo. At the reference locality the 

member overlies nonsandy shale of the transition member of the Pierre Shale and is overlain by hard carbonaceous siltstone or silty shale here assigned to the 

Sharon Springs Member of the Pierre Shale. Both the 

Apache Creek Sandstone and the Sharon Springs 

Members are completely exposed.

At its reference locality the member as here redefined 

is 82 feet thick. A lower sandy shale is 27 feet thick 

and contains small ironstone concretions; a middle concretionary sandstone is 25 feet thick and contains sandstone concretions more than 15 feet in diameter; an 

upper 30-foot-thick sandstone ("Apache sandstone" of 

Prommel?) forms a prominent cliff. The following 

stratigraphic section was measured at the reference 

section.

Reference section measured north of North Apache Creek on 

west face of Mil in SEy^ sec. 30, T. 25 8., R. 67 W., Huerfano 

County, Colo., at reference locality of Apache Creek Sandstone Member of Pierre Shale

[Measured with Jacob staff],

Pierre Shale (part) :

Sharon Springs Member (part) : Ft in 

Shale, olive-gray, sandy in lower 2 ft, silty in upper

9 ft 6 in__________ 11 6 

Apache Creek Sandstone Member: 

19. Sandstone, yellowish-gray, massive; forms 

cliffs; contains 1-ft gray shaly bed in 

middle _____ 15 0 

18. Sandstone, yellowish-gray, platy; softer than

bed 19; locally forms cliff _ 15 0 

17. Sandstone concretion, yellowish-gray to yellowish-orange, 15-ft diameter; separated laterally from similar concretions by soft shaly 

sandstone; contains large smooth baculites 7 0 

16. Sandstone, gray, shaly 1 8 

15. Sandstone, dark-yellowish-orange, concretionary 8 

14. Shale, gray, sandy 14 0 

13. Sandstone, yellowish-gray 2 0

ART. 25 IN U.S. GEOL.. SURVEY PROF. PAPER 475-B, PAGES B99-B101. 1963.

B99
B100 STRATIGRAPHY AND STRUCTURAL GEOLOGY

Pierre Shale (part) Continued

Apache Creek Sandstone Member Continued Ft in 

12. Shale, olive-gray, sandy______________ 4 2 

11. Ironstone concretions, silty, platy________ 2 

10. Shale, olive-gray, sandy___ __________ 6 2 

9. Ironstone concretions and sandy iron-stained 

shale, moderate-yellowish-brown, platy; contain baculites that have weakly ribbed flanks 

(USGS Mes. loc. D3959) ____________ 0 6 

8. Shale, olive-gray, sandy______________ 2 6 

7. Siltstone, moderate-yellowish-brown, platy,

iron-stained ___________________ 4 

6. Shale, olive-gray, sandy_____________ 1 3 

5. Ironstone concretions, moderate-yellowishbrown, platy, silty; contain Inoceramus aff. 

I. cycloides Wegner (USGS Mes. loc. 

D3960) _______________________ 2 

4. Shale, olive-gray, sandy_____________ 8 4 

3. Ironstone concretionary layer, light brown,

platy __ _________________________ 4

2. Shale, olive-gray, silty; contains thin sandstone

lenses _____________________ 3 0

1. Ironstone concretions, olive-gray; weather

light brown____________________ 4

Total ______________________ 82 7 

Transition member (part) :

Shale, olive-gray, silty, noncalcareous.

At the reference locality the member contains sparse 

fossils in sandstone and ironstone concretions. Fineribbed Inoceramus aff. /. cycloides Wegner was found 

in ironstone concretions 12 feet above the base of the 

Apache Creek Sandstone Member. A weakly ribbed 

species of Baculites that is earlier than Baculites obtusus Meek was found 16 feet above the base in ironstone concretions, and a large smooth baculite was 

found 45 feet above the base in a large sandstone concretion. The overlying Sharon Springs Member is 135 

feet thick and contains several faunal zones. Baculites 

obtusus probably ranges through the lower part of the 

Sharon Springs Member. At 25 feet above the base of 

the Sharon Springs Member noded baculites replaced 

by selenite may represent Baculites mcleami Landes.

At Pueblo the Apache Creek is 200 feet thick and 

contains the weakly ribbed baculite throughout. The 

overlying Sharon Springs Member is 113 feet thick and 

contains several fossil zones. In the basal 12 feet of the 

Sharon Springs Scaphites aff. S. spiniger Schliiter and 

an early weakly ribbed form of Baculites obtusus were 

found in beds of layered flat lime'stone concretions. The 

more strongly ribbed form (the typical form) of Baculites obtusus was found in large septarian limestone 

concretions 18 feet above the base of the Sharon

Springs. Flattened gypsiferous baculites that may 

be Baculites mclearni were found about 30 feet above 

the base. Baculites aff. B. asperiformis Meek was 

found 57 feet above the base and Baculites aff. B. perplexus Cobban was found 91 feet above the base in 

large septarian limestone concretions.

The age of the Apache Creek, as based on its relation to Baculites obtusus and Scaphites spiniger, is 

very early late Campanian (Late Cretaceous). The age 

in terms of the Gulf Series of northeast Texas is probably middle Taylor, or about the age of the Wolf e City 

Sand Member of the Taylor Marl.

Fossil zone

Baculites 

perpleacus

Baculites sp. 

(smooth)

Baculites 

asperiformts

Baculites

mclearni

Baculites 

obtusus

Baculites

Late 

form

Early 

form

sp.

(weakly ribbed)

Baculites sp. 

(smooth)

Scaphites 

hippocrepii

Durango region

Lewis 

Shale 

(part)

Cliff 

House 

Sandstone

Menefee Formation

Point 

Lookout 

Sandstone

Mancos 

Shale 

(part)

Pueblo

1 

£

Rusty member 

(part)

Sharon 

Springs 

Member

Apache Creek 

Sandstone Member

Transition 

member

Niobrara 

Formation 

(part)

Black Hills

£

Mitten 

Black 

Shale 

Member 

(part)

Sharon 

Springs 

Member

Gammon 

Ferruginous

< Groat 

<f Sandstone 

£. Bed

Member

Salt Creek, Wyo.

 3

£

>,

3

Unnamed ? 

sandstone C

Sussex Sandstone 

Member

Shannon 

Sandstone 

Member

"K" sandstone

FIGUBE 25.1. Correlation of Apache Creek Sandstone Member 

of Pierre Shale and adjacent units.

As shown on figure 25.1, the Apache Creek Sandstone Member is correlated with the Menefee Formation, near Durango, Colo., on the basis of the transgressive and regressive relations of the sea during Point 

Lookout, Menefee, and Cliff House time (Cobban, 1956, 

p. 27). Along the southwest flank of the Black Hills 

the upper part of the Gammon Ferruginous Member of 

the Pierre Shale contains the weakly ribbed species of 

Baculites which indicates its equivalence to the Apache 

Creek. In the Salt Creek oil field, Wyoming, the 

Sussex Sandstone Member of the Cody Shale contains 

the same baculites and is equivalent to the Apache 

Creek.
SCOTT AND COBBAN B101

REFERENCES

Cobban, W. A., 1956, The Pierre shale and older Cretaceous rocks 

in southeastern Colorado, in Rocky Mtn. Assoc. Geologists 

Guidebook, Geology of the Raton basin, Colorado 1956: p. 

25-27.

Johnson, R. B., 1958, Geology and coal resources of the Walsenburg area, Huerfano County, Colorado: U.S. Geol. Survey 

Bull. 1042-O, p. 557-583.

Johnson, R. B., and Stephens, J. G., 1954, Geology of the La 

Areta area, Huerfano County, Colorado: U.S. Geol. Survey 

Oil and Gas Inv. Map OM-146.

Lavington, C. S., 1933, Montana group in eastern Colorado: Ani. 

Assoc. Petroleum Geologists Bull., v. 17, no. 4, p. 397-410.

Mitchell, J. G., Greene, John, and Gould, D. B., 1956, Catalog of 

stratigraphic names used in Raton basin and vicinity [Colo.- 

N. Mex.], in Rocky Mtn. Assoc. Geologists Guidebook, Geology of the Raton basin, Colorado, 1956: p. 131-135.
Article 26

STRATIGRAPHIC SECTION AT ISLAND BEACH STATE PARK, NEW JERSEY

By PAUL R. SEABER and JOHN VECCHIOLI, Trenton, NJ.

Work done in cooperation with the New Jersey Division of Water Policy and Supply

Abstract. Quaternary, Tertiary, and Cretaceous sediments 

totaling 3,798 feet and 93 feet of Paleozoic or Precambrian biotite gneiss were penetrated in drilling a test well on the Atlantic 

shore of central New Jersey. The sedimentary sequence reflects 

a deeper water, more continuous environment of deposition than 

the outcrop section, but it is similarly characterized by continental and transitional deposits underlying and overlying 

glauconitic marine beds.

In a test hole drilled at Island Beach State Park, on 

the Atlantic shore of New Jersey (fig. 26.1), a thickness 

of 3,798 feet of sedimentary rocks of the Coastal Plain

FIGURE 26.1 Map of New Jersey showing 

location of Island Beach test well.

ranging in age from Late Cretaceous to Eecent was 

penetrated above biotite gneiss that is correlative with 

crystalline rocks of the Piedmont province. The hole, 

designed to test the lithology and water-bearing possibilities of the entire sedimentary section in the central 

Coastal Plain, is the first test in this part of the Coastal 

Plain to reach crystalline basement rocks.

A depth of 3,891 feet was reached by a 9-inch exploratory borehole drilled by the conventional rotary method. 

A total of 226 sidewall cores were taken at irregular 

intervals based on significant lithologic changes indicated by the drill cuttings and the electric and radioactivity logs. The generalized lithologic descriptions 

in the accompanying table are based primarily on detailed lithologic descriptions of the cores supplemented 

by descriptions of drill cuttings and interpretations 

from electric and radioactivity logs.

Correlation of the stratigraphic units in the drill hole 

at Island Beach with those of the outcrop area is based 

mainly on stratigraphic position and lithologic similarity. The descriptions of outcrops were adopted mainly 

from those given by Owens and Minard (1962) for the 

Columbus, N.J., quadrangle and by Minard and Owens 

(1962) for the New Egypt, N.J., quadrangle (table 

below). Precise time-stratigraphic and biostratigraphic correlations have not been attempted as yet.

Probably the most diagnostic mineral in the marine 

formations of the New Jersey coastal plain, as compared 

with other marine formations, is glauconite, which occurs abundantly as discrete grains or interlayered with 

other minerals (Owens and others, 1961, p. 319. Glauconite occurs throughout the section at Island Beach, 

and it is the major constituent or major accessory mineral from the upper part of the Raritan Formation to 

the top of the Eancocas Group a thickness of about 

2,400 feet. In the outcrop sections, glauconite has not 

been reported from the Karitan Formation; it is reported to be a minor constituent in the Magothy ForART. 26 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B102-B105. 1963.

B102
SEABER AND VECCHIOLI BIOS

and thickness of formations exposed from Roosevelt Jo 

Woodstown on the New Jersey coastal plain

[Modified after Minard and Owens, 1960]

Age

Quaternary

"Tertiary

i 

Cretaceous

Series

Pleistocene

Pliocene (?) and 

Miocene (?)

Miocene

Eocene

Paleocene

Late Cretaceous

Group

Rancocas

Monmouth

Matawan

Formation

Cape May

Cohansey Sand

Kirkwood

Manasquan

Vin cento wn 

Hornergtown Sand

Red Bank Sand 

Navesink 

Mount Laurel Sand

Wenonah 

Marshalltown 

Englishtown 

Woodbury Clay 

Merchantville

Magothy 

Raritan

Exposed thickness 

(feet)

(?)

0-160

30-90

0-50

0-60 5-30

0-80 

5-30 

20-70

20-60 

10-22 

40-70 

0-55 

50-70

20-50 

0-150

mation; and it is the major constituent or a major 

accessory from the base of the Matawan Group to the 

top of the Kancocas Group.

Although a general similarity exists between the 

rocks at Island Beach and their outcrop equivalents, 

several striking differences, other than the occurrence 

of glauconite, are readily apparent.

The basement complex of Paleozoic or Precambrian 

iage, which underlies the entire sedimentary sequence, 

is represented by pegmatitic biotite gneiss at a depth of 

3,798 feet. The upper 64 feet is highly weathered and 

has the appearance of saprolite. The gneiss is probably a correlative of rocks of the Piedmont province 

that crop out west of the Fall Line (base of the Cretaceous, fig. 26.1). The Raritan Formation of Late Creitaceous age, which overlies the basement complex, 

! extends from 2,070 to 3,798 feet and is similar only in 

jits lower 1,004 feet to the continental Raritan exposed 

!in outcrop. The upper 724 feet appears to have been 

deposited in a transitional or marine environment and 

to have no exposed lithologic equivalent. The overlying 

Magothy Formation has a deeper water aspect than 

the Magothy in outcrop, the former being thicker, finer 

grained, and more glauconitic and fossiliferous.

The Matawan and Monmouth Groups, which were 

penetrated between 1,162 and 1,810 feet, also reflect a 

deeper water environment of deposition. These deposits are thicker, less quartzose, and more glauconitic, 

calcareous, and fossiliferous than their outcrop equivalents. At Island Beach, the dominantly quartzose sand

beds that are present in the outcrop of the Matawan 

and Monmouth Groups are either finer grained, such 

as the Red Bank Sand and Wenonah Formation, thinner, such as the Mount Laurel Sand, or entirely absent, 

such as the Englishtown Formation (see accompanying log). The Englishtown Formation is considered to 

be absent because no lithologic units comparable to those 

in the outcrop section were found in the appropriate 

stratigraphic interval. The Navesink and Marshalltown Formations are lithologically similar to their 

outcrop equivalents, although the Marshalltown is 

considerably thicker at Island Beach. Lithologic units 

similar to the Merchantville Formation and Woodbury 

Clay occur between 1,598 and 1,810 feet but as alternating interbedded layers rather than distinct adjacent 

formations as in the outcrop.

The overlying Rancocas Group is divided into three 

formations* in the outcrop but could not be subdivided 

at Island Beach. The unit occurs between 400 and 

1,162 feet and grades from a basal glauconitic silt and 

clay unit not present in the outcrop to a glauconite and 

quartz sand at the top. The overlying Kirkwood Formation is similar to its occurrence elsewhere in the 

Coastal Plain. The Cohansey Sand is represented here 

only by a clay unit similar to clays comprising a small 

part of the formation elsewhere in the Coastal Plain. 

The Cape May Formation (46 feet thick) and the 

Beach complex (74 feet) at the top of the section are 

similar to units described elsewhere along the coast.

The thickening of the entire section between the outcrop and Island Beach is the result of thickening of 

individual beds and the addition of lithologic units 

not present in the outcrop section. In the Island Beach 

well, the thickest additional lithologic units occur where 

distinct unconformities are present in the outcrop. 

These units are the marine or transitional deposits (724 

feet thick) at the top of the Raritan and the glauconitic 

clay and silt (370 feet) at the base of the Rancocas 

Group. The presence of these units in the Island Beach 

well indicates that the shoreline lay between Island 

Beach and the outcrop near Trenton during their deposition. The thicker section and additional units also suggest more continuous deposition at Island Beach 

throughout the Cretaceous and Tertiary than at the 

outcrop section.

In summary, the sedimentary sequence penetrated in 

the drilling of the test well at Island Beach reflects 

deeper water and more continuous deposition than does 

the outcrop section, but the sequence in both places is 

characterized by continental and transitional deposits 

underlying and overlying glauconitic marine beds.
B104 STRATIGRAPHY AND STRUCTURAL GEOLOGY

Generalised log of test well at Island Beach State Park, Ocean 

County, N.J.

[Land surface 10 feet above mean sea level]

Quaternary: 

Beach complex:

Quartz sand, pale-grayish-orange, medium- to 

coarse-grained, well-sorted; fossiliferous, 

micaceous _____________________________

Cape May Formation( ?) :

Quartz sand, olive- to light-gray, very fine 

to fine-grained, well-sorted; very lignitic, 

micaceous; interbedded with clay, medium- to dark-gray, calcareous, fossiliferous, 

sparingly glauconitic, lignitic, micaceous, 

containing quartz sand and pebbles______

Thickness Depth 

(feet) (feet)

74 74

46 120

Tertiary:

Cohansey Sand(?) : 

Clay, light-olive-gray, silty, lignitic_____ 30 150

Kirkwood Formation:

Quartz sand, light- to yellowish-gray, very 

fine to fine-grained, well-sorted, silty, 

sparingly glauconitic at base, lignitic, 

micaceous; interbedded with clay, olivegray to brownish-black, silty, lignitic, 

micaceous __________________ 250 400

Rancocas Group:

Sand (glauconite and quartz), olive-gray 

to greenish-black, very fine to mediumgrained, silty and clayey, calcareous, 

fossiliferous, lignitic, micaceous, pyritic; 

interbedded with sandy clay and silt, 

olive- to dark-greenish-gray, fossiliferous, 

lignitic, micaceous, pyritic, containing 

quartz and glauconite____________ 392 792 

Clay, light- to dark-greenish-gray, very calcareous, very fossiliferous, sparingly 

glauconitic, sparingly lignitic, micaceous; 

interbedded with sandy silt, olive- to 

greenish-gray, very calcareous, very fossiliferous, lignitic, micaceous, pyritic, 

containing glauconite and quartz_____ 370 1,162

Cretaceous: 

Monmouth Group: 

Red Bank Sand:

Quartz sand, olive-gray to olive-black, 

very fine to fine-grained, well-sorted, 

clayey, sparingly glauconitic, lignitic, 

micaceous _________________ 16 1,178

Silt, dark-greenish-gray, calcareous, fossiliferous, glauconitic, micaceous____ 12 1,190 

Navesink Formation: 

Clay, olive-gray to greenish-black, calcareous, fossiliferous, micaceous, containing glauconitic sand_________ 36 1,226 

Mount Laurel Sand:

Quartz sand, olive-brown, medium-grained, 

well-sorted, calcareous, sparingly glauconitic; contains laminae of clay, 

grayish-black, lignitic, micaceous____ 12 1, 238

Depth 

(feet)

Cretaceous Continued

Matawan Group: Thickness 

Wenonah Formation: (feet) 

Silt, greenish-black, clayey, calcareous,

sparingly glauconitic, micaceous _ 22 1,260 

Quartz sand, olive-gray to greenish-black, 

very fine to fine-grained, clayey, calcareous, fossiliferous, sparingly glauconitic, lignitic, micaceous_______ 94 1,354 

Marshalltown Formation:

Sand (glauconite and quartz), dark-gray 

to greenish-black, very fine to mediumgrained, silty and clayey, calcareous, 

fossiliferous, lignitic, micaceous, pyritic; 

interbedded with sandy clay and silt, 

greenish-gray to olive-black, calcareous, 

fossiliferous, lignitic, micaceous, pyritic, 

containing quartz and glauconite; contains 4-foot layer of glauconitic limestone _____________________ 244 1,598 

Woodbury Clay and Merchantville Formation : 

Clay, sandy, medium- to dark-greenishgray, very calcareous, very fossiliferous, micaceous, pyritic, containing glauconite and quartz; interbedded with 

silt, dark-gray to dusky-green, calcareous, fossiliferous, micaceous, pyritic, 

sideritic, containing glauconitic sand 212 1,810 

Magothy Formation:

Quartz sand, light- to yellowish-gray, very 

fine to fine-grained, well-sorted, lignitic, 

micaceous, pyritic; sandy silt, olive- to 

medium-dark-gray, clayey, calcareous, 

fossiliferous, lignitic, micaceous, pyritic, 

containing quartz and glauconite; clay, 

dark-gray, silty, micaceous, interbedded 

with quartz sand, yellowish-gray, very 

fine to fine-grained, well-sorted, sparingly glauconitic, lignitic, sideritic; 

clay, dark-greenish-gray, calcareous, fossiliferous, micaceous, interbedded with 

glauconitic sand, dark-greenish-gray, 

very fine to medium-grained, poorly 

sorted, clayey, calcareous, fossiliferous, 

micaceous, containing layers of limestone, light-greenish-gray, fossiliferous, 

micaceous ___________-__ 260 2,070 

Raritan Formation (?) : 

Quartz sand, grayish-orange, mediumgrained, well-sorted, calcareous, sparingly 

glaueonitic, containing laminae of white 

clay ______________________ 22 2,092 

Silt, sandy, olive- to greenish-gray, very 

calcareous, very fossiliferous, micaceous, 

containing glauconite and quartz, interbedded with quartz and glauconite sand, 

olive- to greenish-gray, very fine to coarsegrained, poorly sorted, clayey, very calcareous, very fossiliferous, micaceous, pyritic, sideritic; contains 4-foot layer of 

limestone, light-greenish-gray, fossiliferous, micaceous,.-____ ____ _ 413 2, 505
SEABER AND VECCHIOLI B105

Cretaceous Continued Thickness Depth 

Raritan Formation (?) Continued (feet) (feet) 

Clay, light- to greenish-gray, calcareous, fossiliferous, glauconitic, micaceous; interbedded with quartz sand, light- to yellowish-gray, fine- to coarse-grained, lignitic, 

sparingly micaceous, pyritic, sideritic__ 289 2, 794 

Quartz sand, light- to yellowish-gray, very 

fine to coarse-grained, lignitic, micaceous, 

sideritic; interbedded with clay, light- to 

dark-gray and red, lignitic, limonitic, 

micaceous, pyritic; contains two 3-foot 

layers of shale, black, carbonaceous, micaceous, pyritic__________________ 1,004 3,798 

Paleozoic or Precambrian:

Gneiss, biotite, weathered (saprolite?) ____ 64 3,862 

Gneiss, biotite, with pegmatite veins______ 29 3, 891

REFERENCES

Minard, J. P., and Owens, J. P., 1960, Differential subsidence of 

the southern part of the New Jersey coastal plain since 

early Late Cretaceous time: Art. 82 in U. S. Geol. Survey 

Prof. Paper 400-B, p. B184-B186.

 1962, Pre-Quaternary geology of the New Egypt quadrangle, New Jersey: U.S. Geol. Survey Geol. Quad. Map 

GQ-161.

Owens, J. P., and Minard, J. P., 1962, Pre-Quaternary geology 

of the Columbus quadrangle, New Jersey: U.S. Geol. Survey Geol. Quad. Map GQ-160.

Owens, J. P., Minard, J. P., and Blackmon, P. D., 1961, Distribution of clay-sized sediments in the coastal plain formations near Trenton, New Jersey: Art. 263 in U.S. Geol. 

Survey Prof. Paper 424-C, p. C317-C319.
Article 27

AERIAL RECONNAISSANCE OF THE OUTER SHUMAGIN ISLANDS, ALASKA

By ARTHUR GRANTZ, Menlo Park, Calif.

Abstract. The outer Shumagin Islands, which jut 65 miles 

into the North Pacific from the western part of the Alaska 

Peninsula, consist of slaty argillite and graywacke of late 

Mesozoic age, and biotite granodiorite which has intruded them.

The outer islands of the Shumagin Islands group 

form a geologically distinct island subgroup, composed 

of slaty argillite, graywacke, and granodiorite, which 

projects from 25 to 65 miles into the North Pacific from 

the western part of the Alaska Peninsula (fig. 27.1). 

This subgroup comprises the islands from Nagai seaward (fig. 27.2), and extends from a point 45 miles 

southeast of the Aleutian arc to a point 18 miles from 

the outer edge of the continental shelf (arbitrarily 

placed at the 100-fathom depth curve) and 75 miles 

from the axis of the Aleutian Trench.

It is noteworthy that each major rock type in the 

outer Shumagins was recorded in 1741 by Georg Wilhelm Steller while serving as physician and naturalist 

on Bering's second voyage, the exploration which discovered Alaska.1 These and a few other geologic observations entered in Steller's journal of this voyage establish that this noted botanist and zoologist made the 

first recorded geologic observations in Alaska. Grewingk (1850, p. 173 and pi. 2) noted that much clay slate 

("Thonschiefer") crops out on Nagai, and his map 

showed this island to be underlain by metamorphic 

rocks. Dall (1882, repeated in Dall and Harris, 1892, 

p. 233, and Dall, 1896, p. 807-809) noted that the outer 

Shumagins were composed of granitic and metamorphic 

rocks, but he presented no geologic map. Atwood (1911, 

pi. 6) mapped the outer Shumagins as "volcanic(?)-

1 Steller's journal (in the translation by Stejneger in Colder, 1925, 

p. 79) states of Nagai and the outer Shumagins . . . "This island, as 

well as all the others, consists only of high solid rocks covered with 

vegetation. The rock is mainly a coarse, gray and yellowish graywacke (sic), in some places a gray sandstone; a black, thick slate 

occurs also." Steller's observations were probably made near the low 

valley which transects Nagai Island 9 miles from its southern tip. 

His gray sandstone and black slate are the medium gray graywacke 

and dark gray slaty argillite of Nagai Island. His graywacke as 

rendered in Stejneger's translation was apparently "Felsen Stein" 

rather than the equivalent "Graufels" in the original journal (Golder, 

1925, p. 79, footnote 164) and the phrase "coarse, gray and yellowish 

graywacke" is thought to record the light-gray granitic rock which 

crops out in the sea cliffs near Steller's probable landing place on 

Nagai.

160° 00'

FIGURE 27.1 Index map of soutnwest Alaska.

unsurveyed" although he had noted Ball's observations 

in his text, and Atwood's map was used by Dutro and 

Payne (1954) in compiling their geologic map of 

Alaska. The present article is based upon aerial reconnaissance in the outer Shumagins from a light 

wheel-plane on June 12-14, 1962, supplemented by 

observations at three landing places and study of aerial 

photographs. E. V. Alien, of the U.S. Geological Survey contributed samples and (or) data on bedrock at 

five gravity stations which he established in the outer 

Shumagins in 1961.

The oldest rocks in the outer Shumagins are the slaty 

argillite and graywacke of Nagai, western Big Koniuji, 

and the smaller islands which lie between them (fig. 

27.2). These rocks also form large xenoliths and probably roof pendants in the plutonic rocks on Nagai and 

Big Koniuji. In outcrops at Saddlers Mistake and 

Eagle Harbor, and in hand specimens from Pirate 

Shake, the graywacke is medium or medium dark gray, 

predominantly fine and medium grained, and thin to 

very thick bedded with argillite interbeds. Graded 

bedding, from medium sand at the base to silt at the top, 

is common but not conspicuous. Argillite chips of

ART. 27 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B1O6-B10Q. 1963.

B106
GRANTZ B107

I60°00'

55 C

I59°30'

I

Mzg(?)

Spectacle I.

Mzag 

/^ Bendel I.

Mzag Mzg

Turner I.

T&.Y-:

Near I.

Mzg(?)

10 Miles

BIRD I..

SIMEONOF I.

K\\\J CHERNABURA I.

55 (

I60°00'

EXPLANATION

Volcanic rocks of late

Tertiary or Quaternary age

Biotite granodionte

of late Mesozoic or younger age

Pattern broken where interpreted

from distant views or aerial photoSAND POINT

Big Koniuji I.

Argillite and graywacke 

of late Mesozoic age

Little Koniuji I.

55°

SIMEONOF I.

160° CHERNABURA I. 

SHUMAGIN ISLANDS

27.2. Reconnaissance geologic map of the outer Shumagin Islands.

680231 O 6i3 8
BIOS STRATIGRAPHY AND STRUCTURAL GEOLOGY

sand to cobble size are widespread in the graywacke 

and locally form conglomeratic accumulations. Sole 

markings are not abundant but groove casts, and possible grazing marks, flute casts, and worm-tube fillings, 

are found. The fine-grained tops of some graded beds 

contain small-scale current ripple marks. The argillite is dark and very dark gray, occurs in thin to very 

thick beds, and in places contains thin layers of graded 

silt with small-scale crossbedding. Limestone concretions, small iron sulfide concretions, and thin accumulations of shells were found in a few beds.

The slaty argillite and graywacke are hard, dense, 

and considerably deformed. The density of six samples 

ranged from 2.69 to 2.74 and averaged 2.71. Diagenetic 

and low-grade metamorphic recrystallization have almost or entirely eliminated porosity from these rocks, 

and fracturing and veining are locally common, especially in the graywacke. Dips are moderate to very 

steep, and over most of Nagai are directed toward the 

northwest. Dips are equally steep but more variable 

in direction east of Nagai. Faults are numerous, and 

both open and chevron folds were noted. As seen from 

the air and on two ground traverses, strong axial-plane 

cleavage appears to be absent, but slaty cleavage was 

observed, especially near some faults and in argillite 

interbedded with thick graywacke beds.

Fossils are not abundant in the slaty argillite and 

graywacke, but four collections were obtained by the 

writer at Eagle Harbor, Nagai Island. These contain 

fragments of the late Mesozoic mollusk Inoceramus 

(identified by D. L. Jones) and a few rushlike plants. 

The indicated age and lithologic character of these 

rocks suggest that they are part of the similar sequence 

of late Mesozoic age which underlies Kodiak Island 

and the Kenai-Chugach-St. Elias Mountain system. 

In the latter areas the group names Orca, Sunrise, 

Valdez, and Yakutat have been variously applied to 

these rocks, but until their stratigraphic classification 

is finally established in these larger and better known 

areas, it seems best neither to apply the old nor to coin 

new names for the argillite and graywacke of the outer 

Shumagins.

Plutonic rocks form part of Nagai and most of Big 

Koinuji and the more easterly Shumagins. Their 

presence on Chernabura and Bird Islands is known only 

from one gravity station and from photogeology.- As 

seen from the air, the plutonic rocks intrude the argillite and graywacke on Nagai and Big Koniuji Islands, 

and their age is therefore no older than late Mesozoic.

Samples of the pluton collected at five localities 

on Nagai, the Koniuji, and Simeonof Islands were 

examined with a binocular microscope. Four are lightgray medium-grained biotite granodiorite (classification of Johannsen, 1939) with hypidiomorphic-granular 

texture. The fifth sample, from the southern part of 

Nagai, is biotite adamellite with similar color and 

texture except for coarse grains of late potash feldspar. 

The dark minerals in all samples are chiefly reddishbrown biotite, and minor magnetite and chlorite. Four 

density determinations ranged from 2.63 to 2.67 and 

averaged 2.65. Modes of the four granodiorite samples, in volume percent, determined from stained surfaces follow:

Average 

(percent) (percent)

Quartz _______________________ 24-31 28 

Plagioclase ___ _ 38-18 42 

Potash feldspar_______ __ 10-19 15 

Dark minerals________ 13-17 15

100

The adamellite sample contains, by volume, 32 percent 

quartz, 34 percent plagioclase, 26 percent potash feldspar, and 8 percent dark minerals. The uniform character and the map distribution of these rocks suggest 

that they are part of a batholith which extends over 

half the area of the outer Shumagins. This batholith is at least 20 miles in diameter and could be much 

larger.

Tertiary sedimentary rocks also exist, according to 

Dall (1896, p. 808-809), "On the western edge of 

Nagai... above the metamorphic schists and quartzites, 

but they are greatly altered and consolidated and constitute a small area in comparison with underlying 

strata." These rocks were not recognized in the present aerial reconnaissance.

The outer Shumagins differ markedly in geologic 

character from the inner SHumagins, which expose 

unmetamorphosed and in part poorly consolidated sedimentary rocks of Tertiary age overlain by pyroclastic 

rocks and lavas of late Cenozoic age. The latter constitute Andronica Island and The Haystacks and lie 

within 4 miles of the northwest part of Nagai. The 

inner Shumagins are thus structurally much lower than 

the outer Shumagins, and a flexure and (or) fault which 

strikes northeast and is structurally down on the northwest must lie between Nagai and The Haystacks. If 

Tertiary rocks crop out on Nagai, as reported by Dall, 

then part of this structure would cross the western part 

of this island.

Pleistocene glaciers covered the outer Shumagins and 

extended into areas now beneath the sea. The glaciers 

which originated on the 1,000- to 1,900-foot mountains 

of these islands carved cirques, tarns, and deep Ushaped valleys whose lower ends formed fiords. Postglacial marine erosion has produced wave-cut platforms around these islands and, in places, hanging
GRANTZ B109

cirque valleys and cliffed headlands. An elevated plat- 

;form of low relief, also interpreted to be wave cut, 

icovers three-fourths of Simeonof Island and small areas 

on Chernabura. Hasty reconnaissance suggests that 

ithe elevated surface was possibly formed by more than 

one stand of the sea. In one area on Simeonof Island 

Ithe planed-off bedrock surface stands 15 to 25 feet above 

Sea level at distances of approximately three-fourths 

mile from the ancient strandline. The cutting of this 

elevated platform may be older than the last period of 

glaciation in this area, for on Simeonof Island features 

identified on aerial photographs as glacial deposits 

(D. M. Hopkins and D. S. McCulloch, oral communication, 1962) appear to lie upon it. Its elevation indicates that the area of the platform that is 15 to 25 feet 

above sea level could be of Sangamon (last interglacial) 

kge because Second Beach at Nome, thought to be of 

this age, is 35 to 40 feet above present sea level (HopIkins and others, 1960, p. 53). However, because the 

platform lies in a tectonically active region, its elevation is not sufficient to establish its age. The elevated 

[platform was not recognized on the Koniujis or Nagai; 

if its absence is not due to subsequent glacial or marine 

erosion, then these islands have been depressed rela- 

|tive to Simeonof and Chernabura in the time since the 

platform was cut.

REFERENCES

Atwood, W. W., 1911, Geology and mineral resources of parts of 

the Alaska Peninsula : U.S. Geol. Survey Bull. 467, 137 p.

Ball, W. H., 1882, Note on Alaska Tertiary deposits: Am. Jour. 

Sci., ser. 3, v. 24, p. 67-68.

 1896, Report on coal and lignite of Alaska: U.S. Geol. 

Survey 17th Ann. Rept, pt. 1, p. 763-908.

Ball, W. H., and Harris, G. B., 1892, Correlation papers, Neocene : U.S. Geol. Survey Bull. 84, 349 p.

Butro, J. T., Jr., and Payne, T. G., 1954, Geologic map of Alaska : 

U.S. Geol. Survey, scale 1: 2,500,000.

Golder, F. A., 1925, Bering's voyages, v. 2, Steller's journal of

. the sea voyage from Kamchatka to America and return on

the second expedition, 1741-1742, translated and in part

annotated by Leonhard Stejneger: Am. Geog. Soc. Research

Ser. 2, 290 p.

Grewingk, Constantin, 1850, Beitrag zur Kenntnis der orographischen und geognostischen Beschaffenheit der nordwestKiiste Amerikas mid den anliegenden inseln: Russ. K. 

Mineralog. Gesell., Verhand., 1848-49, p. 76-424, pi. 1-3, 

tables 4-7. [Separately issued by Karl Kray, St. Petersburg, 1850]

Hopkins, B. M., MacNeil, F. S., and Leopold, E. B., I960, The 

coastal plain at Nome, Alaska A late Cenozoic type section for the Bering Strait region: Internat. Geol. Cong., 

21st, Copenhagen 1960, Rcpt, pt. 4, p. 46-57.

Johannsen, Albert, 1939, A descriptive petrography of the 

igneous rocks, 2d ed., v. 1: Chicago, 111., Univ. Chicago 

Press.
Article 28

STRUCTURAL INFLUENCE ON DEVELOPMENT OF LINEAR TOPOGRAPHIC FEATURES, 

SOUTHERN BARANOF ISLAND, SOUTHEASTERN ALASKA

By DAVID A. BREW, ROBERT A. LONEY, JOHN S. POMEROY, and L. J. PATRICK MUFFLER, 

Mcnlo Park, Calif.

Abstract. Joints, foliation, and faults influenced development of the ice-eroded deep fiords and valleys. Northwesttrending linear topographic features parallel the generalized 

strike of the foliation, but are locally controlled by faults. 

Northeast-trending linear features parallel the generalized 

strike of several joint sets, but are locally influenced by faults.

The rarely visited southern part of Baranof Island 

displays some spectacularly rugged glacial topography, 

including many small fiords and deep lake-filled valleys 

that are strikingly linear (fig. 28.1).

These dominantly ice-eroded linear features are compared to the attitudes of joint sets, foliation, and faults, 

which it is assumed influence the erodibility of the bedrock by providing surfaces along which the rock will 

break or otherwise erode more easily.

In a largely hypothetical study Twenhofel and Sainsbury (1958) inferred fault control for almost all prominent linear topographic features in southeastern Alaska, but did not exclude the possibility that some of 

the features could be joint controlled. Peacock (1935) 

concluded that the pattern of fiords in British Columbia corresponds to the underlying pattern of folds, 

fractures, and faults in the bedrock, but did not present detailed evidence. He emphasized (1935, p. 658) 

the parallelism of joint sets and physiographic features 

and noted also that faulting was locally important. 

The present study differs from these previous ones in 

that it is a detailed quantitative study of a much 

smaller area.

Southern Baranof Island consists of metamorphic 

rocks intruded by igneous masses believed to be satellitic to the Coast Range batholith on the mainland.

The northwestern part of the area shown on figure 

28.1 is underlain by consistently northwest-striking 

graywacke and argillite of the Sitka Graywacke of 

Jurassic and Cretaceous age (Loney and others, 1963a). 

To the southeast this unit is intruded and thermally 

metamorphosed by two major granitic complexes. The 

northeastern part of the area shown on the illustration

consists of four major igneous masses intruded into 

amphibolite, schist, and gneiss derived from several 

lithostratigraphic units of probable Mesozoic age. The 

distribution of these several units is shown on a forthcoming preliminary map by Loney and others (1963b). 

The geology of the island will be described in detail in 

a later report.

This study is based primarily on structural observations made from shoreline exposures during reconnaissance mapping. The pattern of fiords provides an 

areally even distribution of observations. The data 

from the intervening ridges are unevenly distributed 

and were not incorporated in the study, although they 

confirm the conclusions drawn from the shoreline data.

The field and analytical procedures used were as follows : Structural observations were taken at about 1-mile 

intervals along all of the coastline shown on figure 28.1. 

Separate pole diagrams of structures in the igneous 

and in.the metamorphic rocks were prepared for each 

subarea, but because of the essential similarity of the 

attitudes in both types of rock the pole diagrams were 

combined to make the contour diagrams. Only the 

best developed joint set observed at each station was 

used in preparation of the joint diagrams. Joints dipping less than 35° were excluded from the analysis 

because they probably have relatively little effect on 

the orientation of the linear topographic features. 

Foliations measured in the metamorphic rocks include 

schistosity and closely spaced shear surfaces; a few 

relict bedding attitudes from unfoliated less metamorphosed rocks are also included. Foliations measured 

in the igneous rocks include schlieren and other mineral 

layering.

The data presented on figure 28.1 permit visual comparison of the structural observations with the orientation of the fiords, larger lakes, and principal streams 

as taken from the Port Alexander 1:250,000 Alaska 

Topographic Series map. In comparing the structural

ART. 28 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B110-B113. 1963.

B110
BREW, LONEY, POMEROY, AND MUFFLER Bill

FIGURE 28.1. Map of southern Baranof Island, southeastern Alaska, showing: fiords, lakes, streams, faults (dashed 

where probable) ; and foliation and joint attitudes for each subarea indicated by Roman numeral. Poles to foliation (F) and joints (J) shown in lower hemispheric equal-area projection; number of poles in parentheses. Solid 

contours indicate 5 (omitted in diagrams of 30 poles or less), 10, 20, and 40 percent per 1 percent of area. Dashed 

contours indicate the limit of the 1-percent circle drawn about single points.
B112 STRATIGRAPHY AND STRUCTURAL GEOLOGY

observations and the topographic orientation, the scatter of both should be kept in mind. Rose diagrams 

showing the azimuth and length of fiords and valleys 

in each subarea were prepared for comparative purposes but are not included here because of space limitations ; most of these diagrams show the expected bimodal 

distribution. Divergencies and apparent inhomogeneities result from the large size of the individual subareas. Although the similarity between the distribution of linear topographic features and structural elements is only approximate, the essential correspondence 

is within the limits imposed by the large size of the 

subareas and the scatter of the topographic and structural data.

JOINTS

Well-developed and persistent, generally northeaststriking joint sets in both the igneous and metamorphic 

rocks appear to exercise the primary influence on the 

northeast-trending linear topographic features. The 

influence of these sets is in places augmented or surpassed by near-parallel major faults, but (as shown on 

fig. 28.1) only a few large northeast-striking fault zones 

were mapped on southern Baranof Island.

The joint diagrams for subareas I-VIII on the west 

side and southern tip of the island show a complex system of steep northeast-striking joint sets; a few sets 

of different strike are important in only 1 or 2 subareas. 

The more diffuse diagrams generally are from subareas 

with greater areas of igneous rock and migmatite outcrop, in which several prominent joint sets are present 

and no one set is dominant. The difference in the position of the maximum of subarea VIII as compared with 

the maximums of the subareas to the north may reflect 

an areal change in lineation attitudes, as most of the 

joints are of the ac type (perpendicular to the fold 

axes).

The diagrams for subareas IX-XIII on the eastern 

side of the island demonstrate that the joint sets measured on the western side of the island extend through 

to the eastern side as far north as subarea X; but the 

pattern shown by these diagrams is disrupted to the 

north in the widespread intrusive and migmatite terrane 

of subareas XI to XIII.

The general correspondence between the strike of the 

dominant joint sets and the linear northeast-trending 

topographic features is shown on figure 28.1. This 

tendency toward parallelism indicates that the joints 

have very probably controlled the trend of these linear 

features. Generally northeast-trending topographic

features which diverge somewhat from the strike of 

the joints may be caused either by the modifying influence of foliation, by faults, by local joint sets, or by 

combinations with joint sets of lesser development that 

are not represented in the diagrams.

FOLIATION

The northwest-trending linear topographic features 

are generally shorter than the northeast-trending features and tend to be parallel to the generalized strike 

of the steeply dipping foliation which ranges from 

N. 24° W. to N. 44° W. Foliation in the isoclinally 

folded Sitka Graywacke and its metamorphosed equivalents consists in many places of pre-metamorphism 

shear surfaces that locally form significant shear zones. 

Because of the isoclinal folding, the foliation and the 

boundaries of the original lithostratigraphic units are 

almost everywhere nearly parallel; it therefore is unlikely that a topographic feature controlled by an original unit would diverge appreciably from the strike of 

the foliation.

The diagrams of subareas I, II, and III on the west 

side of Baranof Island show diffuse maximums representing steeply dipping foliation that strikes between 

N. 65° W. and N. 30° W. The foliation diagram for 

subarea II is not shown because of extreme scatter, 

probably due to the rotation of the foliated xenoliths 

within the igneous mass. Diagrams V to IX show 

simpler, more marked preferred orientations that are 

due to the lesser amounts of igneous rock involved and 

to the more uniform deformation recognized in these 

subareas. The rocks of subarea VIII are strongly 

lineated, and the lineation may have influenced the 

credibility of the bedrock through its modifying effect 

on the foliation. In the diagrams for subareas X to 

XIII the northwest- and west-northwest-striking pattern is continued, but is generally more diffuse because 

of large areas of igneous and mixed igneous and metamorphic rocks in those subareas.

FAULTS

A few major and several minor faults have been 

mapped on southern Baranof Island (fig. 28.1). More 

detailed mapping would undoubtedly reveal many more 

minor faults and possibly a few more major faults. 

Almost all of the faults shown on the illustration have 

associated shear zones that influenced the erodibility of 

the bedrock locally, in many places extensively enough
BREW, LONEY, POMEROY, AND MUFFLER B113

to be expressed on a topographic map at a scale of 

1:63,360 or even smaller. At Big Branch Bay the existence of a fault is inferred from the offset of geologic 

contacts across the bay. This is the only northeasttrending fiord where such offset is present.

The pattern of the demonstrated faults indicates that 

the number of faults striking in the northeast quadrant 

,is about equal to that in the northwest and that the 

I strike of these structures is reflected topographically in 

their immediate vicinity. It is apparent, however, that 

most of the fiords and deep valleys do not occur on the 

:major faults or their extensions.

REFERENCES

Loney, R. A., Berg, H. C., Pomeroy, J. S., and Brew, D. A., 1963a, 

Reconnaissance geologic map of Chichagof Island and northwestern Baranof Island, Alaska: U.S. Geol. Survey Misc. 

Geol. Inv. Map 1-388. [In press]

Loney, R. A., Pomeroy, J. A., Brew, D. A., and Muffler, L. J. P., 

1963b, Reconnaissance geologic map of Baranof and Kruzof 

Islands, Alaska: U.S. Geol. Survey Misc. Geol. Inv. Map. 

[In press]

Peacock, M. A., 1935, Fiord-land of British Columbia: Geol. Soc. 

America Bull., v. 46, p. 633-696.

Twenhofel, W. S., and Sainsbury, C. S., 1958, Fault patterns in 

southeastern Alaska: Geol. Soc. America Bull., v. 69, p. 

1431-1442.
Article 29

GEOLOGY OF THE GUANICA-GUAYANILLA BAY AREA, SOUTHWESTERN PUERTO RICO

By I. G. GROSSMAN, Middletown, Conn.

Work done in cooperation with the Commonwealth of Puerto Rico

Abstract. Previously unmapped serpentinite, chert, and 

associated basic igneous rocks of Cretaceous age (the Benneja 

Complex of Mattson, 1958) and overlying clastic rocks of Oligocene age (the Juana Diaz Formation) crop out near Central 

San Francisco (a sugar mill) south of an east-trending fault 

that is concealed by the alluvium of the Rio Yauco. The south 

side of the fault has been uplifted relative to the north side.

Puerto Rico, the easternmost and smallest of the 

Greater Antilles, is about 110 miles long and about 35 

to 40 miles wide. A central upland, which extends the 

full length of the island, is composed largely of 

igneous, metamorphic, and sedimentary rocks of Cretaceous age and in small part of sedimentary rocks of 

early Tertiary age. The coastal plains flanking this 

upland on the north and south are made up largely 

of sedimentary rocks of early to late Tertiary age, 

and near the coast these beds are mantled by alluvium 

and beach deposits of Quaternary age. All the principal streams flow seaward from the central upland, but 

those on the south coast are much shorter than those 

in the north. (See fig. 29.1) The Rio Yauco, a major 

stream in the southwestern part of the island, makes a 

sharp turn near the sea; no other large river on the 

island exhibits so abrupt a change in direction of flow 

in the lower, more mature part of its course.

Hydrologic studies in southwestern Puerto Rico in 

1962 included the lower part of the drainage basin 

of the Rio Yauco. Geologic mapping in support of 

those studies revealed a previously unrecognized inlier 

that extends about three quarters of a mile along the 

south wall of the Rio Yauco valley west of Central a 

San Francisco. Postulation of a concealed fault 

accounts for not only the inlier but also the atypical 

drainage pattern in the vicinity of Central San 

Francisco.

Of the two rock units exposed in the inlier, the older 

consists of serpentinite, chert, and basic igneous rocks

1 The Puerto Rican term for "sugar mill."

(Kb on map). It is believed that these correlate with 

similar rocks of the Bermeja Complex of Mattson 

(1958, 1960) of Late Cretaceous age, the type area of 

which is in the vicinity of Mayagiiez on the west coast. 

The Bermeja Complex includes the oldest rocks exposed 

in Puerto Rico, and the newly discovered serpentinite 

that here is tentatively assigned to the Bermeja is the 

southeasternmost known outcrop of this metamorphosed igneous rock.

The younger of the two rock units in the inlier consists of conglomerate and sandstone (Tj on map), 

which unconformably overlies the Bermeja Complex. 

It also crops out east of Guanica for a distance of about 

iy2 miles along the south side of a northeast-trending 

valley, where it consists of conglomerate and sandstone 

overlain by shale. Both exposures of this clastic unit 

are regarded by the author to correlate with the lowest member of the Juana Diaz Formation of Oligocene age, as described by Zapp and others (1948). In 

the inlier west of Central San Francisco, the clastic 

unit ranges in thickness from less than 100 feet near 

the east end to more than 250 feet near the west end, 

and dips 17°-35°S. The beds in the outcrop east of 

Guanica also dip southward at an.angle of 12° near 

the east end of the exposure. Fossils found at the 

latter locality in 1955 by C. Wythe Cooke and subsequently identified by him (written communication, 

1956) include Lepidocyclina gigas Cushman, Heterostegina sp., Pecten meseticus Maury, Ampullinopsis 

spenceri (Cooke), and Ostreet, haitensis Sowerby. 

Both the inlier west of Central San Francisco and the 

clastic rocks east of Guanica are within the large area 

mapped by Zapp and others (1948) as the Ponce Limestone of Oligocene and Miocene age. The conglomerate and shale exposed about % mile southeast of 

Guanica and the shale exposed 2 miles west-southwest 

of Central San Francisco (X's, fig. 29.1) have not

ART. 29 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B114-B116, 1963.

B114
GROSSMAN B115

 £_>' ' (; : ^D^-^'vr^rg ..'. '1 *"^TVf' -^f' »

r-j^&gjfcS' * ~;£SAN ~F R A N cis'co" FA" C

v'.'- .Central Saii 

n-:*VFrancisco.'

2 MILES

CARIBBEAN EXPLANATION

66 "OO1

.Qd 

Surficial deposits

Ponce Limestone

10 20 30 40 MILES

Juana Di'az Formation

Upper Cretaceous rocks 

(after Slodowski, 1956)

Bermeja Complex of

Mattson (1958) 

Serpentinite, chert, and assoqiated basic igneous rocks

Contact

p_____ 

u 

Probable fault

i 

Possible fault

Postulated concealed 

fault

Strike and dip of beds

x

Exposures of clastic rocks 

in the Ponce Limestone 

(after Zapp and others, 

1948)

FIGURE 29.1. Geologic map of the Guanica-Guayanilla Bay area, southwestern Puerto Rico.

t>een examined by the writer. They may be beds 

within the Ponce Limestone.

Igneous rocks of Cretaceous age (Ku, fig. 29.1), 

which crop out about 2 miles west of San Francisco, 

yere mapped by Slodowski (1956) as the andesite member of his so-called Ensenada Formation. This unit has 

[not been studied in detail by the writer but appears to 

fbe younger than the rocks he has correlated with the 

iBermeja Complex and older than those correlated with 

the Juana Diaz Formation. Evidently the surface on 

which the Tertiary sediments were deposited was one 

of considerable topographic relief, and deposition of 

the Ponce Limestone began before the highest hills of 

eroded Cretaceous rocks had been completely buried by

elastics of the Juana Diaz.

The Ponce Limestone, which conformably overlies 

the Juana Diaz in the area, is the youngest bedrock unit 

of Tertiary age in southern Puerto Rico. Throughout 

its extent in the Guanica-Guayanilla Bay area, it dips 

5°-20°S. It is shown in figure 29.1 to extend farther 

south than indicated on earlier geologic maps of this 

area; previously, the rugged coastal area between 

Guanica and Punta Ventana was thought to consist 

largely of beach deposits of Quaternary age.

Unconsolidated sediments consisting of alluvium, 

colluvium, beach sand, dune sand, and marsh deposits 

mantle the bedrock at low altitudes. In the GuanicaGuayanilla Bay area, these deposits of Quaternary age
B116 STRATIGRAPHY AND STRUCTURAL GEOLOGY

are in direct contact with, and unconformably overlie, 

the Bermeja Complex, the Juana Diaz Formation, and 

the Ponce Limestone; the upper surface of the surficial 

deposits ranges from about 45 feet above sea level to a 

few feet below sea level.

The serpentinite of the Bermeja Complex northwest 

of Yauco is overlain successively at Yauco by the Juana 

Diaz Formation and the Ponce Limestone, both of which 

crop out in broad east-trending belts. There, as in the 

Guanica-Guayanilla Bay area, the Tertiary rocks commonly dip southward (Grossman, 1962, p. 63). Therefore, a fault is postulated to account for the duplication 

of the stratigraphic sequence in the inlier. Because 

the south-dipping beds of the Ponce Limestone north 

of the inlier would, if projected southward, extend beneath the older Bermeja Complex and Juana Diaz 

Formation, the fault is assumed to underlie the alluvium 

of the lower Rio Yauco and of its tributary at the bend. 

If extended westward, the trace of this concealed fault 

would connect with the fault mapped by Zapp and 

others (1948) east of Guanica (Z-Zf, fig. 29.1). No 

search for evidence of faulting has been made in the 

mile-long intervening upland (indicated by question 

marks on figure 29.1) which is outside the area of this 

study.

The hypothesized east-trending fault would explain 

the present eastward course of the Rio Yauco downstream from the sharp bend and also the eastward course 

of the small tributary that enters the Rio Yauco at the 

bend. That the Rio Yauco formerly followed a more 

direct course to the Carribbean is indicated by the broad 

valley that extends from the present valley of the river 

southeastward past Central San Francisco to the Caribbean Sea just east of Punta Ventana. Blocked at its 

mouth by a sandbar and surmounting dunes, this "abandoned" valley is drained by an undersized ephemeral 

stream that flows southeastward to the blocked exit to 

the sea and then doubles back and empties into 

Guayanilla Bay.

The postulated fault is here named the San Francisco

fault after the sugar mill through whose land it passes. 

Because the fault offsets the Ponce Limestone and is 

masked by deposits of undisturbed alluvium, it postdates consolidation of the Oligocene and Miocene limestone and antedates deposition of the Quaternary alluvium in the east-trending part of the Rio Yauco valley. 

If it connects with the fault mapped by Zapp and others 

(1948) east of Guanica and extends eastward as far as 

Guayanilla Bay, the fault is at least 6 miles long. Possibly it is much longer, extending westward beneath 

the alluvium in the vicinity of Guanica and eastward 

across Guayanilla Bay.

As the trace of the San Francisco fault is approximately parallel to the strike of the beds on both sides, 

it is classified as a strike fault. However, some doubtful 

bedding in the more highly deformed Cretaceous unit 

of the inlier is not parallel to the regional strike. The 

serpentinite in the inlier is directly in line with the 

regional trend of a serpentinite body that extends southeastward from the Mayagiiez area, suggesting that 

strike slip on the fault, if any, was negligible. The 

amount of dip slip is at least several hundred feet and 

could be as much as a few thousand feet. The beds 

south of the fault obviously have moved upward relative to those in the north. Classification of the fault 

as normal or reverse awaits determination of the direction of dip of the fault surface.

REFERENCES

Grossman, I. G:, 1962, Stratigraphy and hydrology of the Juana

Diaz Formation in the Yauco area, Puerto Rico: Art. 137

in U.S. Geol. Survey Prof. Paper 450-D, p. D62-D63. 

Mattson, P. H., 1958, Geology of the Mayagiiez area, Puerto

Rico: Dissert. Abs., v. 18, no. 1. 

 1960, Geology of the Mayagiiez area, Puerto Rico: Geol.

Soc. America Bull., v. 71, no. 3, p. 319-361. 

Slodowski, T. R., 1956, Geology of the Yauco area, Puerto Rico:

Princeton Univ. and Industrial Lab., Puerto Rico Admin.

de Fomento, 130 p. 

Zapp, A. D., Bergquist, H. R., and Thomas, C. R., 1948, Tertiary

geology of the coastal plains of Puerto Rico: U.S. Geol.

Survey Oil and Gas Inv. Prelim. Map 85 (2 sheets).
Article 30 

CARADOCIAN (MIDDLE ORDOVICIAN) FOSSILIFEROUS ROCKS NEAR ASHLAND, MAINE

By ROBERT B. NEUMAN, Washington, D.C.

Abstract. Basalt conglomerate, tuffaceous mudstone, and 

Other rocks about 150 feet thick contain fossils dominated by 

jWelsh-Baltic brachiopods such as Nicolella cf. N. actoniae 

(Sowerby) and Sampo cf. 8. indentata Spjeldnaes. These rocks, 

ind other both younger and older Ordovician rocks of volcanic 

origin in Maine, have faunas with strong European affinities, 

isuggesting volcanic islands as possible migration routes.

Richly fossiliferous basalt conglomerate, tuffaceous 

Jnudstone, and other rocks of a yet unnamed unit are 

exposed on the floor of a gravel pit 5.6 miles east of 

Ashland, Maine (fig. 30.1). This sequence contains 

brachiopods that have not been found before in North

R. 4 W.

FIGURE 30.1. Index map showing location of gravel pit east 

of Ashland.

America although several are familiar Welsh and 

Baltic forms. This article is submitted in advance of 

completing the study of the fossils in order to call attention to these representatives of the Old World.

The exposures were first discovered by W. H. Forbes 

of Washburn, Maine, to whom the professional geologists working in northeastern Maine are increasingly 

grateful for his continuing discoveries of important 

fossil localities (see Berry, 1960a; Pavlides and others, 

1961). Forbes submitted his first collections from this 

place to A. J. Boucot in 1959, and Boucot referred 

them to me when he found that they contained Ordovipian fossils. Later, in 1961 and 1962, I collected additional material.

, Professor Alwyn Williams of Queen's University, 

[Belfast, Northern Ireland, examined the collection during his visit here in 1962 and suggested the identification of some of the brachiopods, and I am grateful for 

his help. S. D. Berger assisted me in some of the collecting, and in making the sketch of the pit (fig. 30.2).

Geologic mapping in the Presque Isle quadrangle 

(Boucot and others, 1963) and the eastern part of the 

Ashland quadrangle (E. S. Nay lor, California Institute 

of Technology, written communication, 1962) indicates 

that the rocks in the gravel pit lie near the core of an 

anticline. There are few other exposures in the vicinity, however, and none others are like these. Nearby 

exposures -are "greenstone." Younger Ordovician 

argillite and a variety of Silurian rocks occur on the 

flanks of the anticline, according to Naylor.

Excavation of about 10 feet of unconsolidated drift 

in the gravel pit uncovered bedrock in several places 

over an area of about 1 acre, and boulders like those 

of the exposed bedrock are heaped in several places 

in the excavation. Only a small part of the floor of the 

pit is bedrock.

Bedding is visible only at the eastern tip of the pit 

(fig. 30.2) where alternating beds of more and less limy 

mudstone containing well-oriented brachiopod shells 

strike north and dip 30° E. On the basis of this attitude, a thickness of 150 feet is estimated from the outcrop width of 350 feet exposed in the pit.

The well-stratified interval, about 3 feet thick, contains dark-gray mudstone that is largely noncalcareous except for the brachiopod shells; one bed 6 inches 

thick is a muddy limestone. Fine-grained volcanic 

debris that includes small fragments of ash and feldspar 

crystals is concentrated along several bedding surfaces. 

Brachiopods are common, especially a dalmanellid, 

probably Onniella.

Freshly broken mudstone from this interval has a 

strong asphaltic odor, and it contains scattered small 

blebs of a vitreous asphaltic substance. Analysis of this 

substance by Breger (written communication, 1962) 

shows that it is sapropelitic organic residue that lacks 

the free hydrocarbons of crude petroleum.

Massive mudstone without a trace of stratification 

lies above and below the well-bedded mudstone, and

ART. 30 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B117-B119. 1963.

B117
B118 PALEONTOLOGY AND PALEOECOLOGY

Mudstone with fragments 

of volcanic rock, bedding obscure

FIGURE 30.2. Sketch map of the gravel pit 5.4 miles east of Ashland, Maine, showing location of fossil-bearing bedrock 

exposures.

crops out to about the center of the excavation. It contains abundant volcanic debris, ranging from microscopic fragments to angular pieces of basaltic rock as 

large as 6 inches across. The abundance of large fragments is variable, ranging from an occasional isolated 

piece to a breccia of fragments enclosed in a matrix of 

fossiliferous mudstone. Several genera of brachiopods 

and a few rugose corals occur in this rock, some nearly 

filling the space between larger rock fragments.

Boulder conglomerate and interbedded green mudstone are the westernmost and presumably the oldest 

beds exposed. Most of the boulders in the conglomerate 

are basaltic rock, but there are many fragments of finegrained limestone. The matrix of the conglomerate 

consists of poorly sorted angular fragments of basalt, 

commonly of granule size, and includes scattered, somewhat fragmentary brachiopods. One limestone boulder 

contains only slightly compressed graptolites that were 

etched free in the laboratory and identified by Berry 

(written communication, 1962) as a new species of 

Amplexograptus. In addition, there are fragments of 

favositid coral colonies.

Throughout the sequence as a whole, brachiopods are 

by far the most common fossils, followed by rugose and 

tabulate corals. There are also a few trilobites, gastropods, bryozoans, and graptolites, but they are rare.

Preservation of the brachiopods is excellent. Most

of the shells are undeformed, and only those in the 

conglomerate are fragmentary. The nature of the rock 

requires that in order to study the brachiopods the 

shells must be dissolved away, leaving internal and external molds that preserve features required for identification. The molds are so perfect that punctate shells 

are revealed by hairlike threads of matrix that bridge 

the cavities between their interior and exterior surfaces. 

The following brachiopods have been identified to 

date. Essentially the same forms occur throughout the 

exposure, and differ from place to place only in relative 

abundance. For most of them specific affinities are not 

given, as comparable forms are not available for comparison in this country. 

Orthoidea:

Cyrtonotella sp.

Glyptorthis sp.

Nicolella cf. N. actoniae (Sowerby)

Platystrophia sp. 

Clitambonoidea:

Triplecia cf. T. insularis (Eichwald) 

Strophomenoidea:

Chonetoidea sp.

Christiania sp.

Diambonia sp.

Leptaena sp.

Ptychoglyptus sp.

Sampo cf. S. indentata Spjeldnaes 

Dalmanelloidea:

Onniettal sp.

Several of these fossils have not been seen before in 

the United States, although some of them have long 

been known in Great Britain and Europe. The Nicolella is very similar to a species of Caradocian age from 

Estonia (Cooper, 1956, pi. 39E), the Triplecia to one 

from the Caradoc Sandstone in Wales (Hall and 

Clarke, 1892, pi. 11C, fig. 21), and the strophomenoid 

suite to that of the 4b stage in the Oslo region, Norway 

(Spjeldnaes, 1957). Species similar to the Maine 

forms of Christiania, Diambonia, and Glyptorthis also 

occur in the lower Ardmillian Series in the Girvan District, Scotland (Williams, 1962). Williams (1962, p. 

61) correlated the lower Ardmillian with the Wilderness Stage of Cooper (1956), and Berry (1960b, p. 

100-101) considered the 4b stage of the Norwegian succession equivalent to the Wilderness and Trenton Stages 

of Cooper (1956). The rocks of the gravel pit, therefore, were probably deposited during this interval.

The rocks that contain these fossils were obviously 

deposited near the site of contemporaneous volcanism, 

as were some older and younger Ordovician units elsewhere in Maine whose fossils are currently under study.
NEUMAN B119

As at Ashland, most of these faunas seem more European than American. Perhaps these areas of volcanism 

were in the form of a chain of islands that defined 

[paths of migration between Europe and North America 

throughout Ordovician time.

REFERENCES

[Berry, W. B. N., 1960a, Early Ludlow graptolites from the Ashland area, Maine: Jour. Paleontology, v. 34, p. 1158-1163.

 1960b, Correlation of Ordovician graptolite-bearing sequences : Internat. Geol. Cong., 21st, Copenhagen 1960, 

Kept., pt. 7, p. 97-108.

Boucot, A. J., Field, M. T., Fletcher, R. C., Forbes, W. H., 

Naylor, R. S., and Pavlides, Louis, 1963, Reconnaissance 

bedrock geology of the Presque Isle quadrangle, Maine: 

Maine Geol. Survey Bull. [In press]

Cooper, G. A., 1956, Chazyan and related brachiopods: Smithsonian Misc. Colln., v. 127, pt. 1, 1024 p., pt. 2, 269 pis.

Hall, James, and Clarke, J. M., 1892, An introduction to the 

study of the genera of Paleozoic Brachiopoda, pt. 1 of 

Paleontology of New York, v. 8: New York Geol. Survey, 

367 p.

Pavlides, Louis, Neuman, R. B., and Berry, W. B. N., 1961, Age 

of the "ribbon rock" of Aroostook County, Maine: Art. 30 

in U.S. Geol. Survey Prof. Paper 424-B, p. B65-B67.

Spjeldnaes, Nils, 1957, The Middle Ordovician of the Oslo region, Norway, 8. Brachiopods of the suborder Strophomenida: Norsk, geol. tidsskr., v. 37, p. 1-214.

Williams, Alwyn, 1962, The Barr and lower Ardmillian series 

(Caradoc) of the Girvan district, south-west Ayrshire, with 

descriptions of the Brachiopoda: Geol. Soc. London Mem. 3, 

267 p., 25 pis.
Article 31

MIOCENE VERTEBRATES FROM MIDDLE PARK, COLORADO

By GLEN A. IZETT and G. EDWARD LEWIS, Denver, Colo.

Abstract. Fossil vertebrate faunal elements indicate a probable middle Miocene age for tuffaceous clayey siltstones and 

conglomerates that are exposed over large areas and mantle 

older rocks near Kremmling in Middle Park, Colo.

Recent geologic mapping in the Hot Sulphur Springs 

and Kremmling 15-minute quadrangles, Grand County, 

Colo., led to the discovery of well-preserved fossil vertebrates. Study of the fossil vertebrate faunal elements 

found by the writers and re-evaluation of the literature 

concerning these faunal elements indicate that they are 

probably of middle Miocene age.

The fossils were found in a sequence of soft tuffaceous 

siltstones and conglomerates that are exposed over large 

areas near Kremmling, Colo., and that mantle the older 

rocks in the area. The thickness and the exact stratigraphic relations among the rocks here dated as probable middle Miocene and adjacent rock units in the area 

have not been fully determined and await completion 

of stratigraphic studies now in progress.

The siltstones and conglomerates are shown on the 

geologic map of Colorado by Burbank and others 

(1935) as the North Park Formation of Miocene(?) 

age. These rocks have also been called the Troublesome Formation, by Richards.1 He applied the name 

to good exposures along Troublesome Creek near 

Kremmling, Colo., and believed that the rocks might 

be of Oligocene and Miocene age. Later, Levering and 

Goddard (1950, p. 41) used the name Troublesome 

Formation and they suggested that most of the formation might be of Oligocene age.

The Troublesome Formation consists of several distinct rock types. The most common types are lightgrayish-orange, light-greenish-gray, or reddish-brown 

tuffaceous clayey siltstones. The siltstones are structureless and locally limy. Less common rock types that 

are intercalated with the clayey siltstone are thin beds 

of ash, impure limestone beds and nodules, mud-pebble 

conglomerates, chert beds as much as 3 feet thick, and

1 Richards, Arthur, 1941, Geology of the Krernrnling area, Grand 

County, Colorado : Michigan Univ. Ph. D. thesis. (Available from University Microfilms, Inc., Ann Arbor, Mich.)

montmorillonitic claystone beds. Pipy concretions or 

root casts up to 2 inches in diameter are also common 

in the siltstone. In the lower part of the sequence 

coarse conglomerates consisting of Precambrian granitic 

rocks or Tertiary porphyritic volcanic rocks are locally 

present.

Vertebrate fossils were found at several localities in 

Middle Park. An incomplete skull and mandible of a 

horse were found with other bones in limy clayey siltstones that probably rest unconformably on Precambrian gneiss at locality D526, about 5 miles east of 

Kremmling, Colo. An incomplete skull and mandible 

of a rhinoceros were found in grayish-orange limy siltstone at locality D520, about 4J/2 miles east of Kremmling, Colo. An incomplete skull of a merycoidodont 

was found at the base of a road that is cut in grayishorange limy claystones that rest unconformably on the 

Niobrara Formation of Late Cretaceous age at locality 

D522, about 4 miles west of Parshall, Colo. Other 

fossils collected from the area are not yet identified. 

The specimens studied were identified as follows:

Merychippus aff. M. sejunctus (Cope); USGS fossil 

vertebrate locality D526, NW^SE^ sec. 18, T. 1 N., 

R. 79 W., Grand County, Colo.; incomplete skull and 

mandible with parts of skeleton.

l.Aphelops megalodus Cope or lAphelops profectus 

(Matthew); USGS fossil vertebrate locality D520 

NW^SEi/4 sec. 24, T. 1 N., R. 80 W., Grand County, 

Colo.; incomplete skull and mandible.

Merycochoerus aff. M. matthewi Loomis; USGS fossil vertebrate locality D522, NE^SW% sec. 9, T. 1 N., 

R. 79 W., Grand County, Colo.; incomplete skull with 

right and left M2-3.

Cockerell (1908, p. 683) was the first to report middle 

or upper Miocene rocks from near Kremmling in Middle Park, Colo., on the basis of a jaw fragment with 

two teeth of a horse identified by J. W. Gidley as 

Parahippus. Levering (1930, p. 74, pi. 6) mapped over 

20 square miles of Miocene rocks near Granby in Grand 

County, Colo., from which he reported the collection 

of Blastomeryx, a camel comparable to Procamelus, a

ART. 31 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B120-B122, 1963.

B120
IZETT AND LEWIS B121

large Parahippus< and Moropus cf. M. elatus Marsh 

identified by H. J. Cook as "an Arikaree fauna * * * 

Characteristic of the lower Miocene * * *"; but we 

tvould point out that Procamelus suggests upper Miocene, and that Moropus ranges into the middle Miocene, 

ParaMppus into the upper Miocene, and Blastomeryx 

into the middle Pliocene. Eichards 2 reported the collection of bones identified as "Procamelus" and "titanOthere" in the Troublesome Formation. This led 

Lovering and Goddard (1950, p. 41) to suggest that 

'''Oligocene and late Miocene or early Pliocene fossils" 

Occur in Middle Park. We have found no fossils of 

Oligocene or Pliocene age near Kremmling in Middle 

Park and we suspect that the fragmentary character 

of Kichards' fossils may have led to erroneous identi- 

:ications. The "titanothere" might be a large rhinoceros such as Aphelops, and the "Procamelus" might 

>e an earlier genus of camel; both would be of Miocene 

age.

The stratigraphic association of the three genera and 

$pecies found by us in the Troublesome Formation is 

puzzling because the same genera and species were 

reported from rocks supposedly of different ages in 

northeastern Colorado. The problem of association in 

^he Miocene rocks of Middle Park is inseparable from 

the same kind of problem in the Miocene rocks of northeastern Colorado, where discrepant views on correlation

liave been reached chiefly from f aunal studies. Figure

Matthew (1899, Schultz and Galbreath Izett and Lewis 

1901) Falkenbach (1953) (this article) 

Osborn (1909) (1940, 1949)

\ MIOCENE

Upper

Middle

Lower

II 1

1 2

I

4

I I

1 3

I I

2 4

T T T T

7 7 7 

i i i 1

1234

1. Merychippus sejunctus and M. aff. M. sejunctus

2. Aphelops profectus and ?A. profectus

3. Aphelops megalodus and fA megaloflus

4. Merycochoerus

FIGURE 31.1. Stratigraphic distribution of four taxa found 

[ in Middle Park and northeastern Colorado, according to several writers.

2 See footnote, p. B120.

31.1 shows graphically the stratigraphic position of 

genera from Middle Park and from northeastern Colorado, according to several writers who have made 

f aunal studies.

Cope (1874, p. 9-11, 15) found Aphelops megalodus 

and Merychippus sejunctus somewhere along "the line 

of bluffs * * * from * * * the valley of Crow Creek 

* * * eastward for at least sixty miles" in Weld and 

Logan Counties, Colo., in rocks "of the Loup Fork 

epoch." Matthew (1899, p. 65, 71; 1901, p. 358) found 

Merychippus sejunctus and Aphelops profectus in the 

same area together with Merycochoerus proprius, and 

coined the name "Pawnee Creek Beds" for the formation in which they occur. He added (1901, p. 359) that 

"a number of species described by Prof. Cope" (including Aphelops megalodus} "appear from their matrix 

to have come from" the next higher "horizon" in the 

"Pawnee Creek Beds."

Schultz and Falkenbach (1940, p. 279, 288; 1949, p. 

80, 83), who have monographed the merycoidodonts, 

consider that the genus Merycochoerus occurs only in 

middle Miocene rocks equivalent to the Marsland 

Formation of the Hemingford Group of Lugn (1939). 

More recently Galbreath (1953, p. 18-20, 27, 31-37, 

91-110) reinvestigated the "Pawnee Creek" faunas of 

northeastern Colorado. He records his interpretation 

of the age relations of the genera and species we have 

identified from Middle Park as follows': 

Late Miocene, "Eubanks local fauna"

Merychippus sejunctus, "occurrence certain" 

Aphelops megalodus, "occurrence probable" 

Middle Miocene, "Martin Canyon local fauna" 

Aphelops profectus, "known to occur" 

Merycochoerus, "known to occur" 

In summary, we believe that these fauna! elements 

probably represent a single stratigraphic unit in the 

Kremmling area, and we have reached the following 

tentative conclusions:

1. The putative species identified by us from the 

Miocene sediments of Middle Park are probably of 

middle Miocene age, equivalent to that of the Marsland fauna.

2. The Miocene fauna identified by Cook (Lovering, 

1930) and that collected by Kichards 3 are probably 

from the middle Miocene and correlate with the Marsland Formation of Nebraska.

3 See footnote, p. B120.
B122 PALEONTOLOGY AND PALEOECOLOGY

REFERENCES

Burbank, W. S., and others, 1935, Geologic map of Colorado:

U.S. Geol. Survey. 

Cockerell, T. D. A., 1908, A new locality for Miocene mammals:

Science, new ser., v. 28, p. 683. 

Cope, E. D., 1874, Report on the stratigraphy and Pliocene

vertebrate paleontology of northern Colorado: U.S. Geol.

and Geog. Survey Terr. Bull. 1, p. 9-28. 

Galbreath, E. C., 1953, A contribution to the Tertiary geology

and paleontology of northeastern Colorado: Kansas Univ.

Paleont. Contr., art. 4, p. 1-120. 

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

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

Levering, T. S., and Goddard, E. N., 1950, Geology and ore

deposits of the Front Range, Colorado: U.S. Geol. Survey

Prof. Paper 223, 318 p.

Lugn, A. L., 1939, Classification of the Tertiary system in Nebraska : Geol. Soc. America Bull., v. 50, p. 1245-1276.

Matthew, W. D., 1899, A provisional classification of the freshwater Tertiary of the West: Am. Mus. Nat. History Bull., 

v. 12, p. 19-75.

 1901, Fossil mammals of the Tertiary of northeastern

Colorado: Am. Mus. Nat. History Mem., v. 1, pt. 7, p.

355-447. 

Osborn, H. F., 1909, Cenozoic mammal horizons of western

North America: U.S. Geol. Survey Bull. 361, p. 1-90. 

Schultz, C. B., and Falkenbach, C. H., 1940, Merycochoerinae, a

new subfamily of oreodonts: Am. Mus. Nat. History Bull.,

v. 77, p. 213-306.

 1949, Promerycochoerinae, a new subfamily of oreodonts : 

Am. Mus. Nat. History Bull., v. 93, p. 69-198.
Article 32 

PALEOECOLOGY OF THE PERMIAN PHOSPHORIA FORMATION AND RELATED ROCKS

By ELLIS L. YOCHELSON, Washington, D.C.

Abstract. Various associations of fossils suggest that plausi- 

^)le paleoecological interpretations are: Park City Formation, 

limestones deposited in shelf areas and mudflats; Shedhorn 

Sandstone, sandstones on or near beaches; and Phosphoria Forjnation, black shale and chert in basins. Each environment has 

its characteristic fauna.

Between 1945 and 1955 slightly more than 1,500 fossil collections were obtained by Geological Survey 

parties assigned to investigations of the western phosphate field (McKelvey and others, 1959). The collections are from more than 100 measured sections in 

Idaho, Montana, Utah, and Wyoming, and are from 4 

dissimilar rock types. In general, these represent 

deposition of limestone (Park City Formation) and 

sandstone (Shedhorn Sandstone) in shelf areas, and 

black shale and chert (Phosphoria Formation) in basin 

iareas (McKelvey and others, 1959). Each rock facies 

lhas its own characteristic fossil assemblage. The 

nearly complete limitation of each assemblage to a 

particular facies is striking.

Although more than 150 fossil taxa were identified, 

only 25 to 30 taxa comprise 95 percent of the fauna. 

Many fossils are poorly preserved. To insure objectivity, several degrees of identification reliability were 

employed for what is probably the same taxon.

Examination of the various combinations of occurrence and relative abundance of the fossils suggested 

that there might be some consistent associations. Sixty 

such associations were defined; many, but not all, of 

these are thought to represent former life assemblages. 

Each of the 1,500 collections was assigned to one of these 

associations, though some subjective judgments were 

required to smooth out variations introduced by quality 

of preservation and local collecting problems. Some 

stratigraphic sections, sampled in detail, show a consistent vertical sequential change in the fossil associations. Similar consistent changes with reference to the 

basin margin also occur.

The most satisfactory interpretations can be made 

in the shelfward facies, particularly within the middle 

limestone of the Park City Formation, outcropping between the Wyoming-Idaho State line and the eastern 

edge of the Wind River Mountains. Sheldon (1963) 

has interpreted a line of maximum regression within 

the body of the Franson Member. The first fossils above 

the line of regression are abundant Orb'iculoidea (Roemeretta) , a large inarticulate brachiopod. Small nuculoid pelecypods, thought to have been shallow 

burrowing forms, are associated locally. Both fossils 

commonly are found in a light-gray claystone.

The claystone is succeeded by a dense moderately 

bioclastic limestone with abundant masses of fairly 

large ramose bryozoan colonies. Slightly higher in the 

sequence the bryozoans are associated with the brachiopods Derbyia and Spinferina. Derbyia apparently required a hard substrate for cementation; Spiriferina 

could have been attached to the bryozoan colonies by a 

pedicle. Eare pectenoids and mytiloids, both nonburrowing pelecypods, also occur in this association. 

Higher in the sequence the two brachiopods are much 

more abundant while the bryozoan colonies coordinately 

decrease not only in abundance but also, possibly, in 

relative size.

The next higher beds are shaly limestones or calcareous shales characterized by the brachiopods Neospirifer, 

Sphenosteges, Sphenolosia, and "Liosotella" These 

four genera are almost mutually exclusive and there is 

some evidence that they occur in sequence as listed. 

Neospirifer may have been attached by a pedicle in 

youth, but probably was free living in later life. The 

productoids Sphenosteges and Sphenolosia cemented 

themselves to a hard substrate. "Liosotella" was a 

small free-living procluctoid supported on the bottom by 

spines. Mollusks are almost never associated with any 

of these brachiopods. Nonproductoid brachiopods 

occur in small numbers, seemingly at random, with each 

of the genera.

These associations are succeeded in turn by associations characterized by EcJiinauris and, finally, by 

Bathymyonia and (or) Muirwoodia. These three 

brachiopods are spinose moderately large to large proART. 32 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B123-B124. 1963.

680231 0 63- B123
B124 PALEONTOLOGY AND PALEOECOLOGY

ductoids. Almost all occurrences of these associations 

are in calcareous shales. The few collections obtained 

below the line of maximum regression seem to show a 

similar, but reversed, sequence.

The lithologic sequence of the Franson Member is 

best considered as an upward change from claystone to 

limestone and then, gradually, to calcareous shale. 

Other physical evidence also indicates a transgressing 

sea. This is interpreted as representing an environmental change from a harder to a softer bottom. The 

inferred functional morphology of the fossils indicates 

a corresponding change from forms requiring a hard 

substrate to those increasingly able to tolerate and 

thrive on a softer substrate.

Fossils from the Ervay Carbonate Rock Member or 

upper limestone member are like those from the Franson Member and show the same general sequence of 

occurrence. Slight differences suggest that the Ervay 

sea may have had less intense wave and current action. 

Fossils from the Grandeur Member or lower limestone 

are dissimilar. Too few collections are available to suggest any valid paleoecologic interpretations.

Eastward and southward, the Ervay and Franson 

faunas change from bryozoan-brachiopod assemblages 

to molluscan assemblages. Both burrowing and surface-dwelling pelecypods occur. Plagioglypta, a 

scaphopod which seems to have lived partially covered 

by sediments, is common. In the Uinta Mountains of 

Utah, the molluscan limestone of the Franson Member 

is underlain by claystone containing Orbiculoidea 

(Roemerella) and nuculoid pelecypods. These beds 

may be interpreted as widespread shallow-water deposits similar to those of modern mudflats.

Northwestward, the bryozoan-brachiopod-bearing 

shelf limestones and shales are replaced by the Shedhorn Sandstone, which has yielded few fossils. In the 

main, fossils consist of resistant fish remains, scaphopods, and large pelecypods, almost exclusively burrowing forms. The fossils support earlier interpretations 

of this unit as a beach or near-beach sand. This beach 

environment is in partial contrast to the mudflat environment noted previously in central Wyoming and 

northeastern Utah.

The two basin facies are more difficult to interpret.

The Rex Chert Member not unexpectedly contains few 

fossils other than sponge spicules which, however, occur in astronomical numbers. While these do not contribute definitive ecologic information, they are not incompatible with the notion that the chert represents the 

deepest water facies of the Phosphoria. A few limestone lenses in the midst of chert sequences are anomalous. They contain a prolific fauna of bottom-dwelling 

organisms and have many of the physical characteristics 

of reef masses. The fauna in each lens is a limited one 

and differs from that of other nearby lenses. These 

differences may be explained as fortuitous colonization 

by pelagic larvae that floated basin ward.

In the Meade Peak Phosphatic Shale Member of the 

Phosphoria Formation, nine genera of brachiopods and 

mollusks form the bulk of the fauna. More than 30 

associations have been formulated. In contrast to the 

shelf facies, most of these associations seem to occur 

at random and no systematic change of the fauna can be 

demonstrated. Most of the fossils seem to have been 

surface dwellers rather than members of the infauna; 

shelled cephalopods and squids, both presumably pelagic carnivores, occur sporadically. The fossils suggest 

that the bottom sediments may have been somewhat 

harder than one would judge from their diagenetic 

product. The profusion of shelled invertebrates implies a depth range with a maximum depth of several 

hundred feet for the basin.

The fauna and associations within the younger Retort 

Phosphortic Shale Member of the Phosphoria Formation show some differences from those of the Meade 

Peak Member. The distinctive features of the Retort 

may not be significant. Ecologic differences between 

the two phosphatic shales probably are of the same 

order of magnitude as the differences between the upper 

two shelf limestones.

REFERENCES

McKelvey, V. E., Williams, J. Steele, Sheldon, R. P., Cressmann, 

E. R., Cheny, T. M., and Swanson, R. W., 1959, The Phosr 

phoria, Park City, and Shedhorn Formations in the western 

phosphate field: U.S. Geol. Survey Prof. Paper 313-A, p. 

1-47.

Sheldon, R. P., 1963, Physical stratigraphy and mineral resources 

of Permian rocks in western Wyoming: U.S. Geol. Survey 

Prof. Paper 313-B. [In press]
Article 33 

QUATERNARY EVENTS ALONG THE UNGLACIATED LOWER OHIO RIVER VALLEY

By LOUIS L. RAY, Washington, D.C.

Abstract. The alluvial history of the unglaciated Ohio River 

valley, below Louisville, Ky., was intimately affected by glaciers 

invading its drainage basin. Alluviation of a deep bedrock vallejy and the lower courses of tributary streams is interpreted 

partly from the stratigraphy of loess deposits genetically related 

to valley trains.

Below Louisville, Ky., the valley of the Ohio River 

lies beyond the limits of continental glaciation, yet it 

has been intimately affected by the waxing and waning 

of the great Quaternary ice sheets that invaded its 

drainage basin. Serving as a drainage for debris-laden 

melt waters during at least 3 glaciations, and possibly 

4^ the valley has been deeply alluviated by valley-train 

deposits. Tributary valleys, except for that of the 

Wabash River, have been choked in their lower courses 

by lacustrine, backwater clayey silts to the level of the 

valley trains of the Ohio that dammed their mouths. 

The Wabash valley, which also served as a drainageway 

fij-om the ice sheets to the north, likewise contains valley-train deposits, so that tributaries to its lower course 

afe also broadly alluviated by lacustrine, backwater 

clayey silts. Sand dunes and a stratigraphic succession 

of loess deposits along the Ohio valley walls and mantling the adjacent countryside are genetically related to 

tl)ie fluvioglacial deposits within, the valley. Because 

pre-Wisconsin alluvium in the Ohio valley is buried by 

Wisconsin and later deposits, the pre-Wisconsin history 

of Quaternary events is in part interpreted from the 

l<j)ess deposits.

Well records indicate that the bedrock channel of the 

Ohio River below Louisville was deeply alluviated duri^ig the Quaternary at one point to a depth of 200 

ffeet. From Louisville to a point near the confluence of 

tie ancient Ohio (Cache Valley) with the Mississippi

^iver, the bedrock valley floor descends from an elevaon of 309 feet (MacCary, 1955) to slightly less than 

180 feet (Fisk, 1944). Various ages have been assigned 

to the cutting of this deeply buried bedrock channel 

(Wayne, 1952). Regional studies have led to the inference that the channel cutting was essentially, if not 

wholly, completed before the advance of the. first glacial 

i£e sheet, the Nebraskan, and that the channel is, therefore, of preglacial, presumably pre-Quaternary origin. 

Significant bedrock incision by the river since the first 

glacial invasion of its drainage basin is held to be 

unlikely.

Although events during the Nebraskan Glaciation are 

obscure, it appears that this first great ice sheet buried 

and permanently disarranged the ancient MahometTeays river system to the north and east of the preglacial Ohio River. At that time the drainage basin of 

the Ohio was vastly enlarged by diversion of the headwaters of the Mahomet-Teays drainage into the Ohio. 

Presumably valley-train deposits were developed along 

the Ohio, for glacial invasion of a drainage basin results in aggradation along the drainageways. Conversely, ice-free interglaciations result in degradation. 

It appears unlikely that the inferred fluvioglacial deposits were wholly removed by degradation of the Ohio 

during the Aftonian (post-Nebraskan-pre-Kansan) Interglaciation or that the river again eroded its bedrock 

valley bottom. The lack of loess deposits along the 

Ohio valley that might be genetically related to a Nebraskan valley train is not surprising because of 

weathering and erosion during the long Aftonian Interglaciation. Loess deposits, if formerly present, must 

have been limited areally and in thickness, for a Nebraskan valley train would have had a restricted surface 

to serve as source area for deflation of silt sediments. 

Also the valley train would have been confined within 

the high protective walls of the deep bedrock valley so 

that deflation from its surface would have been limited.

Field observations show that ice of the Kansan Glaciation crossed the Ohio River valley. Tills that are 

far more deeply weathered than those of the third or 

Illinoian Glaciation are exposed within the valley of 

the Ohio and south of the river above Louisville (Ray, 

1957; Flint and others, 1959). No outwash debris assignable to a Kansan age has been identified, however, 

within the Ohio valley below the point where it was 

crossed by the Kansan glacier, for presumably the outwash is buried by younger deposits and has not been 

recognized in well records. The presence of a valley

ART. 33 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B125-B128. 1963.

B125
B126 GEOMORPHOLOGY, GLACIAL GEOLOGY, AND GLACIOLOGY

train is indicated, however, by a well-exposed section of 

loess near Yankeetown, Ind., that has a well-defined 

stratigraphic succession of four silt deposits. The lowest, deeply weathered silt deposit overlies weathered 

bedrock and is in turn overlain and protected by a compact colluvium of local rock fragments and silt that is 

in part cemented by iron oxides. Above the colluvium 

is the characteristic sequence of three loess deposits related to the third (Illmoian) and fourth (Wisconsin) 

glaciation. The lowest silt, of pre-Illinoian age, is interpreted as a loess of Kansan age that is composed of 

eolian sediments swept from the surface of a valley 

train of Kansan age. It is unlikely that deposits of 

Kansan loess have been preserved elsewhere unless they 

have been protected from erosion by fortuitous circumstances during the very long post-Kansan, Yarmouth Interglaciation. Normally the younger loesses 

of Illinoian and Wisconsin age rest directly on bedrock. 

During the Yarmouth Interglaciation, earlier loess 

was subject to deep weathering and erosion, and the 

Ohio Eiver degraded its valley-train deposits. Following the Yarmouth, ice sheets of the third glaciation, 

the Illinoian, advanced across the Ohio valley in a few 

places above the Louisville area. Again there is no 

direct evidence of a valley train of Illinoian age, but 

there is ample evidence of its existence because of the 

deposits of Loveland Loess along the lower unglaciated 

valley that were derived from eolian sediments deflated 

from an Illinoian valley-train surface. The Loveland 

Loess commonly overlies weathered bedrock; rarely 

weathered rubble or colluvium (Ray, 1957). Deeply 

weathered and leached of its original carbonates, the 

Loveland Loess along the Ohio bears a remarkable similarity in physical character to, and is in the same stratigraphic position as, those deposits along the Mississippi 

above and below its confluence with the Ohio (Leighton 

and Willman, 1950; Horberg, 1956). The distinct welldrained profile of weathering on the Loveland Loess, 

developed during the following Sangamon Interglaciation, is characterized by a deep-crimson-red crumbly 

silty clay at the top that grades downward to a darkbrownish-yellow compact clayey silt. Erosion during 

the Sangamon Interglaciation completely removed the 

Loveland Loess in many places, so that the succeeding 

younger loess deposits of Wisconsin age commonly rest 

directly on bedrock.

The oldest deposit along the Ohio River valley correlative with the fourth or Wisconsin Glaciation is the 

Farmdale Loess (Ray, 1957, I960).1 Lying directly on 

weathered Loveland Loess or on bedrock, the Farmdale

1 Farmdale Loess at several localities along the Ohio River in northern Kentucky and southern Indiana recently has been labeled "Roxana silt" by Frye and others (1962).

Loess is similar in its distinctive physical characteristics and stratigraphic position to loess along the Mississippi River valley (Wascher and others, 1948; Leighton and Willman, 1950; Ray, 1963). Although normally 

compact and leached, its appearance is quite youthful 

in comparison to the earlier, deeply weathered Loveland 

Loess. The widespread distribution of Farmdale Loess 

along the Ohio leads to the conclusion that a valley train 

of Farmdale age marked the earliest Wisconsin alluviation of the valley, although no valley-train deposits 

have been identified. A unique exposure, however, of 

sandy and wood-bearing humic clayey silts at Owensboro, Ky., has a carbon-14 date of 23,150±500 years 

B.P. (Rubin and Suess, 1956, sample W-270). These 

sandy and clayey silt beds underlie 30 feet of fluvioglacial sand and gravel of the younger valley train of 

Tazewell age, and are believed to represent alluviation 

following waning of ice of Farmdale age in the drain-- 

age basin upstream. Woody materials and pollen from 

this deposit indicate a local flora similar to that of today, with some pollen, presumably from afar, indicating flora of a cooler climate.

No loess of lowan age has been identified along the 

Ohio River above the mouth of the Wabash. With the 

present state of knowledge, it appears that a glacial 

invasion of the upper Ohio drainage basin did not occur 

at that time. A thin loess between the Farmdale and 

the overlying loess of Tazewell age along the lower 

Wabash valley, and the Ohio valley below the mouth 

of the Wabash River suggests, however, that a valley 

train of lowan age, now buried by later outwash of 

Tazewell age, was its source. The suggested presence 

of the valley train implies that the Wabash drainage 

basin as well as that of the Mississippi was invaded by 

an lowan ice sheet.2

Maximum alluviation of the Ohio River valley occurred in Tazewell time when ice sheets of Wisconsin 

age were most widespread in the Ohio drainage basin. 

The valley was choked by outwash debris that formed 

a broad valley train whose surface is now the highest 

terrace within the valley. Notable changes in the 

course of the river resulted from burial of low bedrock 

divides by glacial outwash, for example, near Owensboro, Ky.; also upstream from Paducah, Ky., where a 

low divide between the lower Cumberland and Tennessee Rivers was submerged by backwater alluviation 

of the Cumberland valley, a tributary to the Ohio. 

Burial of this divide permitted the Ohio to abandon 

its ancient course through the Cache Valley of southern

2 The lowan Loess, not recognized as such but referred to in places 

as Peoria (transitional) by Frye and others (1962), has, according 

to their data, distinctive D.I. (diffraction intensity) ratios, in marked 

contrast to those of the Farmdale Loess below and the Tazewell Loess 

above (Roxana silt and Peoria Loess of Frye and others).
RAY B127

Illinois and utilize the lower Cumberland and Tenness|ee valleys as a lower route to the head of the Mississippi 

I£mbayment.

Valley trains along the Ohio and lower Wabash 

Kivers dammed the mouths of their tributaries and 

ponded their lower courses to produce shallow basins 

filled with lacustrine, calcareous backwater clayey silt, 

in places highly fossiliferous (Thornbury, 1950, and 

dithers). The resultant lacustrine flats are today prominent features of the landscape. At one artificial exposure through a loess-mantled bedrock hill and the 

adjacent lacustrine sediments that overlap the lower 

hill slope, the Loveland and Farmdale Loesses on the 

hjill slope continue beneath the lacustrine sediments, but 

the thick overlying Tazewell Loess on the hill slope 

merges imperceptibly with the lacustrine beds. Two 

carbon-14 dates (Rubin and Alexander, 1960, samples 

W-520 and W-645) of wood from lacustrine beds along 

the Ohio indicate a Tazewell age of 18,530±500 and 

19,940±300 years B.P. for the deposits. Postdepositional leaching of carbonates of the lacustrine sediments averages about 3% inches in depth for each 

1,000 years.

All post-Farmdale loess along the Ohio River, believed to be largely of Tazewell age above and lowanTazewell age below the mouth of the Wabash River, 

ig called Peorian for convenience. Deposits are thickest nearest the valley-train source areas and grow thinrier with increasing distance from the source. Where 

thick and well drained, a characteristic postdepositional 

profile of weathering has been developed (Ray, 1963). 

Sand dunes on the Tazewell valley-train surface and on 

tihe adjacent valley walls are loess mantled; dunes along 

valley walls may contain intercalated lenses of loess.

The post-Tazewell sequence of events along the un- 

£;laciated Ohio River valley is complex. Differing terlace sequences in various sectors reflect differing 

sequences of events. Above the mouth of the Tennessee 

Ipver, remnants of two post-Tazewell terraces are 

^ound between the level of the flood plain and the terrjace of Tazewell age. The higher of these two terraces 

is the better represented, and is tentatively assigned a 

Cary age. The age and origin of the lower terrace,

poorly represented by small scattered remnants, is not 

certain.

The terrace level assigned a Cary age is believed to 

be contemporaneous with the surface described in the 

Wabash River valley by Fidlar (1948) and attributed 

by him to cutting throughout the course of the Wabash 

by torrents released from ancient Lake Maumee during retreat of ice of Cary age. Recent field studies, 

however, indicate that the Wabash eroded only in its 

upper reaches. In its lower course it was an. aggrading 

stream, for terrace remnants are underlain by wellbedded deposits of sand and gravel, and the terrace 

surface is overlain by sand dunes that bank against the 

valley wall and that are not loess mantled. A deltalike 

mass of debris that was poured from the Wabash into 

the Ohio River valley appears to have partly ponded 

the Ohio and modified its terrace sequence for many 

miles both upstream and downstream.

Below the mouth of the Cumberland River, the Ohio 

River, flowing in its new course after abandonment of 

the Cache Valley, has been largely influenced by events 

on the Mississippi River. Two low, level terraces above 

the flood plain are believed to represent backwater alluviation contemporaneous with periods of aggradation along the Mississippi. Of these, the higher, on 

which the city of Paducah is situated, may be of Mankato age, and the lower may be contemporaneous with 

alluviation once assigned a late Mankato age in the 

Mississippi valley (Leighton and Willman, 1949) and 

more recently a Valders age (Leighton, 1960). A 

younger, great alluvial fan in the Mississippi valley, on 

which the city of Charleston, Mo., is situated, seemingly 

did not affect the regimen of the Ohio.

The flood plain of the Ohio River represents the last 

event in its history. WTiere the valley is wide and 

broadly alluviated, meanders have developed in the alluvial valley fill, and extensive point bars and alluvial 

islands have formed with surfaces as much as 40 feet 

above the mean low water level of the river before 

canalization. All indications suggest that today the 

artificially canalized river is neither actively aggrading 

nor degrading. In many places, however, it is shifting 

laterally so that its course has been somewhat modified 

in historic time.
B128 GEOMORPHOLOGY, GLACIAL GEOLOGY, AND GLACIOLOGY

REFERENCES

Fidlar, M. M., 1948, Physiography of the lower Wabash Valley 

[Ind.]: Indiana Dept. Conserv., Div. Geology Bull. 2, 112 p.

Fisk, H. N., 1944, Geological investigation of the alluvial valley 

of the lower Mississippi River: Vicksburg, Miss., Mississippi 

River Comm., 78 p.

Flint, R. F., and others, 1959, Glacial map of the United States 

east of the Rocky Mountains: Geol. Soc. America, 2 sheets, 

scale 1:1,750,000.

Frye, J. C., Glass, H. D., and Willman, H. B., 1962, Stratigraphy and mineralogy of the Wisconsinan loesses of Illinois : Illinois State Geol. Survey Circ. 334, 55 p.

Horberg, C. L., 1956, Pleistocene deposits along the Mississippi 

Valley in central-western Illinois: Illinois State Geol. Survey Rept. Inv. 192, 39 p.

Leighton, M. M., I960, The classification of the Wisconsin glacial 

stage of north central United States: Jour. Geology, v. 

68, p. 529-552.

Leighton, M. M., and Willman, H. B., 1949, Itinerary of [2nd 

Pleistocene] field conference: Urbana, 111., auspices of the 

State Geologists, 86 p.

 1950, Loess formations of the Mississippi Valley: Jour. 

Geology, v. 58, p. 599-623.

MacCary, L. M., compiler, 1955, Map of the Louisville area,

Kentucky, showing contours on the bedrock surface: U.S.

Geol. Survey Hydrol. Inv. Atlas HA-5. 

Ray, L. L., 1957, Two significant new exposures of Pleistocene

deposits along the Ohio River valley in Kentucky: Jour.

Geology, v. 65, p. 542-545.

 1960, Significance of loess deposits along the Ohio River 

valley: Art 92 in U.S. Geol. Survey Prof. Paper 400-B, 

p. B211.

 1963, Silt-clay ratios of weathering profiles of Peorian

loess along the Ohio valley: Jour. Geology, v. 71, p. 38-47. 

Rubin, Meyer, and Alexander, Corrinne, 1960, U.S. Geological

Survey radio-carbon dates V: Am. Jour. Sci., Radiocarbon

Supp., v. 2, p. 129-185. 

Rubin Meyer, and Suess, H. E., 1956, U.S. Geological Survey

radio-carbon dates III: Science, v. 123, p. 442-448. 

Thornbury, W. D., 1950, Glacial sluiceways and lacustrine plains

of southern Indiana: Indiana Dept. Conserv., Div. Geology

Bull. 4, 21 p. 

Waseher, H. L., Humbert, R. P., and Cady, J. G., 1948, Loess

in the southern Mississippi Valley; Identification and distribution of the loess sheets: Soil Sci. Soc. America Proc.

1947, v. 12, p. 389-399. 

Wayne, W. J., 1952, Pleistocene evolution of the Ohio and

Wabash valleys [Ind.]: Jour. Geology, v. 60, p. 575-585.
Article 34 

KARST TOPOGRAPHY IN THE GROS VENTRE MOUNTAINS, NORTHWESTERN WYOMING

By WILLIAM R. KEEPER, Laramie, Wyo.

Abstract. Karst features are well developed in gently dipping Mississippian Madison Limestone in Tosi Creek Basin. 

EThe fissures and sinkholes formed partly during ancient periods 

.of karst development, and partly during late Pleistocene and 

(Recent times from glacial melt water flowing across, and perIcolating through, extensively ice-scoured and fractured 

[limestone.

, An unusually large and well-^exposed karst plain, 

iformed partly under alpine climatic conditions during late Pleistocene and Kecent times, is present in 

Tosi Creek Basin in the southeastern part of the Gros 

jVentre Mountains, northwestern Wyoming (fig. 34.1). 

The solution-cavity features, chiefly fissures and sinkholes (fig. 34.2), are in a gently dipping surface cut 

on the Mississippian Madison Limestone. Karst terrain is rare in the alpine regions of the Kocky Mountains, even though the Madison Limestone, as well as 

other thick carbonate sequences, is broadly exposed in 

them; hence, a unique combination of favorable geologic, physiographic, and geochemical conditions must 

have existed in Tosi Creek Basin for the development 

and preservation of such terrain.

The karst features are mostly in thick- to thin-bedded 

dark-gray to black coarsely crystalline limestone 

approximately 200 feet below the top of the Madison 

Limestone. The sinkholes are circular to semicircular 

in plan and, range from a few tens of feet to 1,000 

feet across (fig. 34.2). In some sinks, limestone beds 

dip uniformly toward the centers, with little or no colEXPLANATION

R.112W. R. 111 W.

2 MILES

j FIGURE 34.1. Map showing karst area (hachured) in Tosi Creek Basin. (Geologic mapping by W. R. Keefer and D. B. 

! Andretta, 1957; base from U.S. Geol. Survey topographic map of the Gros Ventre quadrangle, 1910).

ART. 34 IN U.S. GEOL,. SURVEY PROF. PAPER 475-B, PAGES B129-B130. 1963.

B129
B130 GEOMORPHOLOGY, GLACIAL GEOLOGY, AND GLACIOLOGY

lapse along the edges; in others, the beds are completely collapsed inward, leaving encircling walls as 

much as 15 feet high. The floors of most sinks are 

relatively flat. Long narrow depressions, probably 

representing collapse along fissures, are also common 

(fig. 34.2). Visible fissures are generally only a foot or 

two wide, but many are as much as 10 feet deep. Most 

of the karst features show little evidence of mechanical 

weathering. The basin area is virtually devoid of soil 

and vegetation.

Throughout northwestern Wyoming and adjacent regions, the upper part of the Madison Limestone generally contains several zones of caverns and sinkholes 

representing various periods of ancient karst development. Such zones can be readily observed in the Madison in the vicinity of Tosi Creek Basin, and part of the 

present floor of the basin is an exhumed fossil karst. 

Irregular masses of red and tan very coarse sedimentary 

breccia, occupying small depressions along the east side 

of the basin, probably represent the sedimentary filling of some of the ancient caverns and sinkholes. The 

fresh appearance of many of the karst features in the 

Tosi Creek Basin, however, indicates that extensive 

chemical weathering of the carbonate rocks flooring the 

basin has also taken place in relatively recent times, 

probably since the late Pleistocene. This has resulted 

in considerable modification of preexisting caverns and 

sinkholes, as well as in the formation of new ones.

During the Pleistocene glaciations Tosi Creek Basin 

was filled by an icefield that fed large glaciers in the 

valleys of Tosi and Rock Creeks (fig. 34.1). The upper 

surface of the icefield may have reached the top of 

Tosi Peak. By the end of Pinedale Glaciation (latest 

glaciation of the Pleistocene generally considered to

be about 7,500 years ago) the basin floor had been 

scoured into the Madison Limestone. During the recession of these youngest glaciers, large amounts of 

melt water flowed across and percolated through the 

deeply grooved, ice-scoured, and fractured limestone 

surfaces, and into the fossil karst beneath. Bedding 

planes and vertical joints in the limestone, widened by 

frost action, probably served as readymade avenues 

for subterranean streamflow. The ancient sinks and 

caverns likewise facilitated the entrance of glacial melt 

water to underground passages; these openings were 

enlarged and became an integral part of the younger 

system of karst features. The preexisting caverns may 

also have been the sites of extensive accumulation of 

glacial ice which, being buried, did not melt as quickly 

as the surface ice. When melting did occur, the support was removed from some cavern roofs, which collapsed and formed sinkholes.

Much snow now accumulates annually in Tosi Creek 

Basin. Despite the seasonal melting of the snow, there 

is little surface runoff, so most drainage must be underground. Because of the northeast dip (average 7°) 

of the carbonate strata, the flow of underground water 

is directed eastward into Rock Creek and northeastward 

into Tosi Creek. It seems likely, therefore, that karst 

development has been virtually continuous since the 

melting of Pinedale ice. The solvent power of the 

near-freezing waters is apparently adequate to cause 

extensive chemical weathering, even though the initial 

carbon dioxide content is probably very low and there 

is little opportunity for the water to dissolve additional 

carbon dioxide before reaching the limestone. Repeated 

freezing and thawing nearly the year around undoubtedly facilitates breakup and collapse of some features.

FIGURE 34.2. Aerial photographs of a part of Tosi Creek Basin oriented for stereoscopic 

viewing. North is to the left. (Photographs from U.S. Forest Service).
Article 35

ERUPTIONS OF WATER AND SAND RESULTING FROM AN EARTHQUAKE 

NEAR CONCEPCION, CHILE

By KENNETH SEGERSTROM, LORENZO CASERTANO,1 and CARLOS. GALLI O.2

Santiago, Chile

Work done in cooperation with the Universidad de Chile and the 

Institute de Investigaciones Geologicas, Chile, and under the auspices 

of the Agency for International Development

Abstract. The earthquake of May 21, 1960, near Concepcion, 

Chile, was accompanied locally by eruption of water and sand, 

! forming mounds resembling volcanoes in miniature. A sudden 

'increase in transient artesian pressures caused by compressioiial earthquake waves ruptured a surficial clay layer and 

forced water from underlying deposits to rise in fountains.

! INTRODUCTION

At 6:10 a.m. on May 21, 1960, the region of ConcepIcion, in southern Chile, was shaken by an earthquake

of intensity IX (Mercalli scale, as modified in 1931) and

I magnitude 7.5 (Blchter). A news account, which was

suppressed to prevent alarm, stated that eruptions of

warm water and sulfurous gases took place in the Boca

Sur area, on the Pacific coast about 8 km west of Con1 Volcanologist, Universidad de Chile, Santiago, Chile.

2 Geologist, Institute de Investigaciones Geologicas, Santiago, Chile.

cepcion, and immediately south of the mouth of the 

Rio Bio Bio (fig. 35.1). The area is about 150 km west 

of the nearest volcano of Cenozoic age.

On June 4, Segerstrom and Casertano visited the area 

in the company of Carlos Ruiz, Director of the Institute de Investigaciones Geologicas; interviewed people 

who had been at or near the site on May 21; and spent 

several hours observing the eruptive vents and erupted 

material. Observations made in the area by Galli and 

Joaquin Sanchez about 10 days later confirmed the 

earlier observations. The "mud volcanoes" at the mouth 

of the Rio Bio Bio have been described briefly by Segerstrom (1960) and in more detail by Casertano (1962, p. 

21-25), who advanced a theory, later modified by the 

present authors, regarding their mechanism. They were 

mentioned by Galli and Sanchez (1960) and by

EXPLANATION

Undifferentiated

Quaternary

deposits

Beach deposits Tidal mudflat deposits

Dune sand

Bedrock

Strand lines

FIGUEE 35.1. Surficial geologic map of Boca Sur area, near Concepci&n, Chile. 

ART. 35 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B131-B134. 1963.

Bl-31
B132 GEOMORPHOLOGY, GLACIAL GEOLOGY, AND GLACIOLOGY

Watanabe and Karzulovic (1960). Similar phenomena 

in other parts of the world have been described as 

"earthquake fountains" (Richter, 1958, p. 106-109) and 

"sand boils" (Ericksen and others, 1954, p. 109).

Grateful acknowledgment is made for the use of a 

helicopter furnished by officials of the 2nd Naval Zone 

of Chile. Thanks are due R. J. Dingman and G. H. 

Davis, of the U.S. Geological Survey, Carlos Ruiz, and 

Pierre St. Amand, formerly professor of geology and 

physics at the Universidad de Chile, who critically reviewed the manuscript and aided in interpreting the 

cause of the eruptions.

The Rio Bio Bio empties into the Pacific Ocean in a 

region of bedrock hills and coastal plain. The mouth 

of the river is bordered by a tombolo which joins the 

Promontorio de Pompon with the Peninsula de Tetas 

to the north, and a spit which extends southward to the 

Boca Sur area. The hills are made up principally of 

metamorphic rocks, largely mica schist, phyllite, and 

slate, probably of Precambrian age. Granitic rocks 

of unknown age locally intrude the metamorphic complex. Near the mouth of the river, the hills of metamorphic rock are partly covered by fossiliferous beds 

of Late Cretaceous and early Tertiary age (Galli and 

Sanchez, 1960; Veyl, 1961; Galli, 1961). North of the 

Bio Bio these hills form the Tetas Peninsula and south 

of the river they are named Cordillera de Nahuelbuta.

Flanking the hills is a coastal plain, probably of 

Pleistocene age. North of the Bio Bio the plain is 

developed on fluvial and deltaic deposits with four 

terrace levels deposited and cut by the river in Pleistocene and Recent time (Galli, 1962). The fluvial and 

deltaic deposits are made up mostly of dark-gray 

medium- to fine-grained sand. Coarse sand, gravel, and 

large boulders are sparingly distributed. South of the 

Bio Bio the coastal plain is developed on marine deposits of fine to coarse sand. The plain has emerged 

from a submarine platform in Pleistocene and Recent 

time and it features a nearly straight, north-trending 

beach, abandoned strand lines, and subsequent drainage, characteristic of the very youthful stage of the 

erosion cycle.

The observations described below were made in and 

near a low-lying cultivated field about 200 m wide which 

lies between a coastal sandbar about 130 m wide, and 

the slightly higher coastal plain to the east (fig. 35.1). 

The field, which is 0 to 2 m above the high-tide level, 

has brown clayey topsoil 20 to 25 cm thick which overlies dark-gray sand and mud extending to unknown 

depth. Some of the mud contains abundant carbonized 

vegetal matter and approaches peat in composition. 

The dark color of the sand is due to its high content of 

reworked volcanic ash brought down from the volcanoes

of the Cordillera de los Andes by the Rio Bio Bio and 

deposited at and near the coast.

VENTS AND PRODUCTS OF THE ERUPTIONS

When the field was visited on June 4, its surface of 

corn stubble and sparse grass was broken by irregularly spaced fissures and circular to elliptical holes with 

a predominant east-northeast orientation and alinement, roughly parallel to the nearby riverbank. Less 

frequently, the orientation of the cracks and holes was 

about N. 30° W., roughly parallel to the nearest segment of coastline. Most of the fissures were a few 

meters to a few tens of meters long and 2 to 10- cm 

wide, and were filled with sand to a depth of 5 to 10 

cm below the surface. Spacing between the subparallel 

east-northeast-striking fissures was 1.5 m to 10 m or 

more.

Accumulations of gray sand, commonly 1 to 5 cm 

thick but in a few places as much as 10 cm thick, lay 

atop the brown soil along the cracks and around the 

holes, forming ridges and mounds. The mounds were 

circular to slightly elliptical and ranged in diameter 

from 5 to TO cm. Depressions, or craters, in the centers 

of the mounds made most of them resemble volcanoes

B

50 CM

FIGURE 35.2. Plans and profiles of representative sand mounds 

and craters in a cultivated field of the Boca Sur area.
SEGERSTROM, CASERTANO, AND GALLI O. B133

in miniature. The diameter of the depressions in the 

circular mounds ranged from %0 to % the base diameter, and the depth was commonly less than the 

height of the mounds. Where the crater was small 

as compared to the outer diameter of the mound, the 

crest of the mound was flat (fig. 35.2J.); where the 

crater was relatively large, the crest was knife edged 

(fig. 35.2Z?). Several mounds were double or joined 

and resembled miniature twin volcanoes. One circular mound had two holes in the middle, one northeast 

of the other; the holes were close together and only 

about one-tenth as wide as the mound (fig. 35.2#).

Many of the craters formed in depressions that 

[were bordered by fissures striking east-northeast. The 

depressions ranged from roughly rectangular to nearly 

jsemicircular in plan, depending on whether they were 

small grabens bordered by fault fissures on the long 

sides, or fault blocks bordered by a fissure on only one 

jlong side. The depressions were commonly 1 to 4 m 

jwide, about twice as long as they were wide, and about 

|20 cm in maximum depth. There were from 1 to 5 

.or 6 mounds and craters in most depressions (fig. 35.3).

A thin film of brown clay covered lower areas of 

'the depressions. Some of the mounds of sand were 

jalso partly veneered with clay on the inner side of their 

,craters. One breached crater with an outer diameter

3 METERS

FIGURE 35.3. Plan showing typical distribution of sand mounds 

and craters and their relation to fissures and depressions 

in a cultivated field in the Boca Sur area.

of 30 cm was veneered not only in the center but also 

in the gap of the crater wall (fig. 35.2Z>). Mounds 

in the deepest parts of depressions were completely 

filmed by clay.

Sand mounds and craters were found also outside 

the sunken areas. These were not veneered with clay. 

Some of them were roughly alined in groups of three 

or more, parallel to the predominant fissures or actually on the fissures. Other sand mounds were of seemingly random orientation (fig. 35.3).

Fragments of clay from 2 cm in diameter to chunks 

as much as 20 cm across and 7 or 8 cm thick, evidently 

spewed out by water under pressure, were sparsely 

strewn over the field. Chunks of clay were found as 

far away as 1.2 m from the nearest fissure or crater.

According to Armando Gaete, a local resident, cracks 

and holes of the same types were present on the mudflat between the field and the river to the north, during 

the first days after the earthquake. The features had 

been largely obliterated by the tides before the authors 

visited the area. Nevertheless, a 6 by 9 m depression, 

elongated approximately normal to the coast, was still 

observable on June 4 in the tidal flat immediately north 

of the field. Chunks or "bombs" of brown soil or clay 

lay on the surface of the depression, and nearby; the 

largest was 40 cm wide and 20 to 25 cm thick.

In the spit just northwest of the field, wider and 

deeper cracks than those of the field were visible. 

Their average strike was about N. 30° W., or approximately parallej to the bar itself. These fissures were 

confined to the northeastern or inland side of the sandbar, where the dune sand overlies clayey soil containing much vegetal matter. Chunks of clay were strewn 

over the sand near some of the northwest fissures. 

One northwest fault zone 2 to 3 m wide was represented by a depression which had been filled with 

slumped sand to a depth of about 50 cm to 1 m below 

the surface. The northeast rim of the depression was 

45 cm lower than the southwest rim; about 30 cm 

probably represented vertical displacement. In one 

place the form of the fault-zone depression, in plan, 

was that of an ellipse with two circles joined at one 

end (fig. 35.4); the width of the depression was about 

1 cm and its deepest part was 60 cm below the surface. A thin film of brown clay coated the gray sand 

to about 15 cm above the bottom of the depression.

ERUPTIONS AND THEIR CAUSE

The earthquake took place at daybreak, and the 

people who lived on higher ground to the south did not 

immediately see its effects on the low-lying field and 

near the river. At about 8:00 a.m., fountains of water 

were observed to be emerging from fissures and circular
B134 GEOMORPHOLOGY, GLACIAL GEOLOGY, AND GLACIOLOGY

FIGURE 35.4 Plan of fault-zone depression in the sand 

spit northwest of a cultivated field in the Boca Sur 

area.

holes in the field to a maximum height of 50 cm, and 

the activity continued until midday (Jose Badilla Diaz. 

local resident, oral communication).

The following is proposed as a mechanism for the 

eruptions. A sudden increase in artesian pressure 

caused by compressional earthquake waves ruptured the 

relatively impermeable layer of brown clay that had 

had a confining effect (Parker and Stringfield, 1950, p. 

455) on the water-soaked sand underneath. Confined 

water was thus able to discharge through the fissures 

formed; the velocity of flow was sufficient to carry considerable sand and finer material in suspension. Presumably the chunks of clay observed around the vents 

were ejected by water under high pressure during the 

early stages of the fountaining. The resulting fountains, with their tendency to produce rapid depletion

of the head, lasted for 4 hours or more, so that large 

quantities of water were discharged. This suggests that 

the sand was very thick, and capable of storing large 

volumes of water and of releasing it rapidly.

Once the upthrown mass was in the air its behavior" 

was similar to that of the eruptive column of a volcano: 

the light particles (chiefly droplets of water) tended to 

be thrown farther than the heavier particles (chiefly 

grains of sand), which formed ridges and mounds along 

and around the eruptive vents. Clay that was carried 

in suspension by the erupted water later veneered areas 

where the water had ponded.

* REFERENCES

Casertano, L., 1962, Sui fenomeni sismo-vulcanici del sud de 

Chile: Naples, Annali dell' Osservatorio Vesuviano, v. 4, 

ser. 6, p. 1-30.

Ericksen, G. E., Fernandez, C., J., and Silgado, E., 1954, The 

Cuzco, Peru, earthquake of May 21, 1950: Seismol. Soc. 

America, v. 44, p. 97-112.

Galli, O., C., 1961, Geologia del area de San Vicente en relacion 

con el proyecto de construction de un puerto: Institute de 

Investigaciones Geologicas open-file report, Santiago, Chile.

 1962, Morfologia de los depositios cuaternarios costaneros 

de la region de Concepcion: Institute de Investigaciones 

Geologicas open-file report, Santiago, Chile.

Galli O., C., and Sanchez R., J., 1960, La geologia y los efectos 

de los terremotos en mayo de 1960 en Concepcion y alrededores (informe preliminar) : Institute de Investigaciones Geologicas open-file report, Santiago, Chile.

Parker, G. G., and Stringfield, V. T., 1950, Effects of earthquakes, trains, tides, winds, and atmospheric pressure 

changes on water in the geologic formations of southern 

Florida : Econ. Geology, v. 45, p. 441-460.

Richter, C. F., 1958, Elementary seismology: San Francisco, 

W. H. Freeman, 768 p.

Segerstrom, K., I960. Eruption of water, sand, and clay, resulting from the earthquake of May 21, 1960, near Concepcion, 

Chile [abs.] : Geol. Soc. America Bull., v. 71, p. 1972.

Veyl O., C., 1961, Contribution al conocimiento de la geologia 

regional de la Provtncia de Concepcion (informe preliminar) : Minerales, Ano 15, no. 72, p. 21-71, Santiago, Chile.

Watanabe, T. and Karzulovic K., J., 1960, Los movimientos 

sismicos del mes de mayo de 1960 en Chile: Anales de la 

Facultad de Ciencias Fisicas y Matematicas, Universidad 

de Chile, Santigao, Chile, Institute de Geologia, v. 17, 

pub. 14.
Article 36 

PARADISE DEBRIS FLOW AT MOUNT RAINIER, WASHINGTON

By DWIGHT R. CRANDELL, Denver, Colo.

i Abstract. The Paradise debris flow blankets the Paradise 

[Park area on the south flank of Mount Rainier. It is roughly 

'correlative in age with the 4,800-year-old Osceola Mudflow on 

ithe northeast flank of the volcano; both probably originated as 

[debris avalanches from a former summit of Mount Rainier 

[volcano.

Studies of surficial deposits in Mount Eainier National Park have revealed an unconsolidated till-like 

deposit of Eecent age on the south flank of Mount Rainier volcano. This deposit is inferred to be a debris flow

.that originated in an avalanche from the upper slopes

|or summit of the volcano. Because it veneers a rather 

wide area at Paradise Park and Paradise Valley (fig.

136.1), it is referred to as the Paradise debris flow.

I Description. The debris flow consists of angular and

I subangular rock fragments as large as 8 feet in diameter 

in a purplish-gray matrix of sand, silt, and clay, which 

oxidizes to yellowish brown. The rock fragments are

| nearly all andesite from Mount Rainier volcano, but a 

few are of granodiorite that crops out in the upper

i Nisqually River valley. The debris -flow is 1 to 5 feet 

thick in most outcrops, but a thickness of as much as 15 

feet is not uncommon. Owing to its thinness, the debris 

flow rarely forms constructional topography, but simply 

veneers ridges, knobs, and depressions in older material 

(fig. 36.2). At Reflection Lakes, however, the debris 

flow is thick enough to form a hummocky surface, and 

the lakes occupy closed depressions in it.

Distribution and volume. The Paradise debris flow 

forms a blanket that is almost continuous in an area 

bounded on the north and east by Panorama Point and 

Mazama Ridge, and on the west by late Recent moraines 

that extend along the east side of Nisqually Glacier. 

Southward, the debris flow crops out discontinuously 

down the Nisqually River valley. It has been recognized as far downstream as the mouth of Tahoma Creek, 

about 4 miles down valley from Longmire, and it may be 

one of several debris flows that occur in the valley as 

far downstream as National, 13 miles below Longmire. 

A large lobe of the debris flow overlies glacial drift 

of late Wisconsin age in the Reflection Lakes area. The

flow entered the Reflection Lakes depression through 

a saddle at an altitude of between 5,100 and 5,200 feet 

in Mazama Ridge that is about 300 feet above the floor 

of the adjacent Paradise Valley.

The debris flow occurs at a higher altitude at Paradise Park than it does on the opposite side of the Nisqually River valley. Whereas the deposit is at an altitude of as much as 6,600 feet near Panorama Point, it 

is not present on a late Wisconsin cirque floor at an 

altitude of about 5,900 feet on the valley wall directly 

opposite.

Outcrops along the Entrance Road half a mile east 

of Ricksecker Point indicate the maximum depth of 

the debris flow while it was in motion. There, at an 

altitude of about 4,300 feet, remnants of the Paradise 

deposit lie about 800 feet above the present floor of the 

Paradise River valley. The deposit is not found above 

about 4,250 feet at Ricksecker Point itself. Thus, while 

the debris flow was moving, its surface sloped westward 

in the Paradise River valley.

The volume of the debris flow preserved upstream 

from Longmire probably is not more than about 30 

million cubic yards. An estimate of its volume while 

flowing, based on heights it reached on valley sides and 

on the probability that the fluid material was draining 

from one area while the crest was still passing points 

farther downstream, is 400 to 500 million cubic yards.

Remnants of the deposit that are preserved high on 

valley sides are not believed to be erosional remnants 

of a fill that solidified in the valley to that height. 

Rather, they are interpreted as veneers left after the 

crest and most of the flow had moved downvalley. The 

postglacial age of the deposit indicates that its remarkable height on valley walls is not the result of the 

debris flow having moved on the surface of glacial ice 

in the valley.

Age and origin. At Paradise Park the debris flow 

is separated from till of late Wisconsin age and older 

bedrock by colluvial deposits and ash layers that include 

ash layer O (Crandell and others, 1962), which is

ART. 36 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B135-B139. 1963.

B135
B136 GEOMORPHOLOGY, GLACIAL GEOLOGY, AND GLACIOLOGY

121°45'

46°55' 

46-45'

FIGURE 36.1. Reconstructed areal extent of Paradise debris flow on south side of Mount Rainier, and Osceola Mudflow on north side. Glacier margins are those surveyed in 1910-13; many ^rnall glaciers are not shown.
CRANDELL B137

FIGURE 36.2. Paradise debris flow veneering till of late Wisconsin age at Paradise Lodge parking lot. Highest point in view 

is Point Success; Gibraltar Rock is at upper right. Nisqually Glacier heads in summit snowfield adjacent to Gibraltar 

| Rock; Wilson Glacier is to the left.

thought to be about 6,500 years old. The debris flow is 

overlain at most places by ash layer Y, which probably 

is about 3,200 years old. In a road cut at the west 

abutment of the Entrance Road bridge across the Nisqually River, carbon collected from a layer between ash 

layer Y and the debris flow has an age of 4,000±250 

years (W-1116) (Crandell and others, 1962).

The coarse, unsorted texture of the Paradise deposit 

could be explained by several modes of origin. Some 

likely ways in which debris of this nature could be deposited are by a glacier, by local sliding and flowage 

of weathered ash and rock debris, by direct falling of 

material explosively erupted by the volcano, and by 

avalanching from high on the flank or from the summit of the volcano. Avalanching is the most probable 

origin; evidence that seems to rule against other origins 

is presented below.

A wide range in particle size in the debris flow suggests the possibility that it is glacial drift deposited by 

Nisqually Glacier when it was thicker and broader than 

at present. However, the presence of ash layer O 

(Crandell and others, 1962) at the ground surface 

immediately beyond late Recent moraines of Nisqually 

and Paradise Glaciers indicates that the glaciers have 

not extended beyond these moraines in about the last 

6,500 years. In addition, distribution of the debris flow 

suggests that it was not formed by Nisqually or Paradise Glaciers, because in order for these glaciers to have 

covered Paradise Park and Paradise Valley, they also 

would have covered areas where the debris flow is not 

present.

Thfe debris flow also somewhat resembles colluvium 

that could be formed by local sliding and flowage of 

rock debris mixed with volcanic ash. Occurrence of the 

debris flow as a veneer on some ridgetops, as well as on 

valley floors and slopes between, rules against this
B138 GEOMORPHOLOGY, GLACIAL GEOLOGY, AND GLACIOLOGY

origin, as does the restriction of the deposit to the areas 

immediately adjacent to the Nisqually and Paradise 

River valleys.

That the Paradise deposit might have resulted from 

an explosive eruption of Mount Rainier seemingly is 

denied by its distribution. Ejection of rubble from high 

on the south side of the volcano in an explosion sufficiently violent to have covered Paradise Park and 

Paradise Valley with debris would also have left deposits in the area upslope from Panorama Point.

The absence of any topographic evidence of a source 

for the deposit near McClure Rock suggests that the 

debris flow originated high on the volcano in the form 

of a debris avalanche. Such an avalanche would have 

accelerated rapidly as it traveled down Nisqually Glacier. At a point west of McClure Rock, the top of the 

east valley wall becomes low enough to have permitted 

a large avalanche to surmount it. Owing to the steep 

slopes bounding the west sides of Wilson and Nisqually 

Glaciers, and to the arcuate path of the avalanche, the 

debris was not deposited very high on the west wall of 

the Nisqually River valley opposite Paradise Park. 

Absence of the deposit on Mazama Ridge is best explained by lack of either momentum or thickness sufficient for the avalanche to surmount the ridge.

The avalanche could have been either wet or dry when 

it covered the Paradise area; in either condition it probably would have had sufficient momentum to flow over 

topographic obstacles. Even if it started dry, it probably soon became wet in transit and then formed a debris 

flow.

Possible correlation and common source of Paradise 

debris -flow and Osceola Mud-flow. The Paradise debris 

flow is lithologically similar although less clayey than 

the 4,800-year-old Osceola Mudflow in the West Fork 

and White River valleys on the northeast side of Mount 

Rainier (Crandell and Waldron, 1956). The Osceola 

moved 45 miles down the White River valley and spread 

into the Puget Sound lowland in a lobe 20 miles long 

and as much as 10 miles wide. An investigation of the 

mudflow in the lowland suggests an average thickness 

of about 20 feet and a volume of perhaps as much as 

1.3 billion cubic yards (Crandell, 1963).

From its distribution at Mount Rainier, Crandell and 

Waldron (1956) concluded that the Osceola Mudflow 

had an eruptive source in a vent or vents high on the 

volcano, but neither an exact source nor mode of origin 

was proposed. Subsequently, Crandell (1963) suggested that the mudflow originated in one or more phreatic explosions that ejected clay-rich debris formed 

within the volcano by hydrothermal alteration of wallrock along old conduits and fissures. It was thought

that as this material settled on the flanks of Mount 

Rainier, it moved on downslope as a mudflow.

The highest outcrop of the Osceola thus far recognized is at the 9,700-foot summit of Steamboat Prow, 

but it must have originated somewhere above this point, 

for it flowed over Steamboat Prow into Glacier Basin, 

apparently as a debris avalanche.

If the Paradise debris flow and Osceola Mudflow 

originated as debris avalanches high on the volcano, 

a source area of large volume is required, and it seems 

likely that the missing summit of Mount Rainier might 

in some way be involved.

Russell (1898) first pointed out that the present summit of Mount Rainier consists of the higher of two cones 

built within a large crater, the rim of which is preserved at Point Success and between Liberty Cap and 

Russell Cliff. High points on this older crater rim 

indicate that the summit of the volcano was removed 

above an altitude of about 14,000 feet. The old summit 

crater had a diameter of about 1.25 miles, measured 

from Liberty Cap to Point Success. Its formation was 

attributed both by Russell and by Matthes (1914) to 

explosive decapitation of the volcano, but Coombs 

(1936) proposed that the shape of the old summit area 

resulted from "* * * intermittent explosive activity, 

probably breaching first one side of the crater and then 

the other." He also suggested that the old crater subsequently had been modified by glacial erosion.

Fiske and others (1963) conclude that the missing 

upper part of Mount Rainier represented the loss of a 

volume of rock of more than 800 million cubic yards. 

They point out that if the summit of the volcano had 

been removed in a violent explosion, explosion breccias 

should be found around the flanks. As these were not 

recognized, they conclude that the summit was removed 

either by vertical subsidence into underlying magma, by 

outward slumping and flowage of solfatarized and softened rock in the central part of the volcano, or by headward glacial erosion into a weak core.

The estimated volume of the Osceola and Paradise 

deposits requires a source area of at least 1.7 billion 

cubic yards. If it is assumed that the missing summit 

cone had a basal diameter of 2,200 yards and a height 

of 500 yards, it would have had a volume of between 600 

and 700 million cubic yards. If the cone had a summit 

crater, as suggested by Coombs (1936), the estimated 

volume of the summit cone should be reduced accordingly, perhaps by a fourth or a third. In addition to 

removal of the volcano above an altitude of about 14,000 

feet, it seems likely that the east crater wall also was 

destroyed, and a huge cirque-shaped depression was 

formed which opened to the east between Gibraltar
CRANDELL B139

feock and Russell Cliff. The floor of this depression 

|may have been as much as 1,500 feet below the high 

[points oh the rim at Liberty Cap and Point Success. 

^Although an estimate of the size of this postulated deipression is hardly more than a guess, the depression may 

nave contained as much as a billion cubic yards of material. Fiske and others (1963) point out that the young 

icone that forms the present summit of the volcano was 

built on the eastern rim of the old summit crater; thus, 

fchis cone has largely filled the wide low breach inferred 

!to have been present in the eastern rim.

I propose that the former summit of Mount Rainier 

was removed principally by very large avalanches that 

originated in hydrothermally altered and weakened 

rocks. The avalanches were probably caused by phrejatic explosions of the volcano, but could have been 

caused by earthquakes or even slope failures due to 

oversteepening by glacial erosion. An avalanche deiscending the south slope of the volcano was confined by 

! the high sides of the Nisqually River valley until it 

'reached a point west of McClure Rock, where it flowed 

'Over the east valley wall, veneered Paradise Park and 

jParadise Valley, and extended downvalley as a debris 

'flow. The Osceola Mudflow probably originated in a 

comparable avalanche or several nearly contemporaneous avalanches fromlhe summit of the volcano and 

I moved into the valleys of West Fork and White River.

If parts of the avalanches on the northeast side of the 

volcano were initially dry, dust rolling up from them 

may have been deposited on adjacent upland surfaces, 

accounting for the layer regarded as an air-laid facies 

of the Osceola Mudflow (Crandell and Waldron, 1956).

A mudflow or debris flow comparable in age to the

Paradise and Osceola deposits has not been seen in the 

Carbon or Cowlitz River valleys; the .valleys heading 

in Tahoma, Puyallup, and North and South Mowich 

Glaciers have not yet been investigated. The absence 

of such a deposit in the Carbon River valley is consistent 

with the presence of a high, unbroken rim of the old 

summit crater between Liberty Cap and Russell Cliff, 

and its absence in the Cowlitz River valley may have 

resulted from a predominantly northeastward movement of avalanches on the east side of the volcano. The 

valley of Tahoma Creek downstream from South Tahoma Glacier contains a debris flow of large volume that 

postdates ash layer Y, and thus is less than about 3,200 

years old. This debris flow perhaps was formed by an 

avalanche derived from the western part of the former 

summit.

REFERENCES 

Coombs, H. A., 1936, The geology of Mount Rainier National

Park: Washington Univ. Pubs, in Geology, v. 3, no. 2, p.

131-212. 

Crandell, D. R., 1963, Surficial geology and geomorphology of the

Lake Tapps quadrangle, Washington: U.S. Geol. Survey

Prof. Paper 388-A. [In press] 

Crandell, D. R., Mullineaux, D. R., Miller, R. D., and Rubin,

Meyer, 1962, Pyroclastic deposits of Recent age at Mount

Rainier, Washington: Art. 138 in U.S. Geol. Survey Prof.

Paper 450-D, p. D64-D68. 

Crandell, D. R., and Waldron, H. II., 1956, A Recent volcanic

mudflow of exceptional dimensions from Mount Rainier,

Washington: Am. Jour. Sci., v. 254, p. 349-362. 

Fiske, R. S., Hopson, C. A., and Waters, A. C., 1963, Geology

of Mount Rainier National Park, Washington: U.S. Geol.

Survey Prof. Paper 444. [In press]

Matthes, F. E., 1914, Mount Rainier and its glaciers: Washington, U.S. Dept. Interior, 48 p. 

Russell, I. C., 1898, Glaciers of Mount Rainier: U.S. Geol. Survey

Ann. Rept. 18, pt. 2, p. 349-115.

680231 O 6S -10
Article 37

PRELIMINARY REPORT ON GLACIOLOGY AND GLACIAL GEOLOGY 

OF THE THIEL MOUNTAINS, ANTARCTICA

By BJORN G. ANDERSEN, Oslo, Norway

Work supported ~by the National Science Foundation

Abstract. The climate of the Thiel Mountains, 500 km from 

the South Pole and 1,100 km from open water, is continental 

and is characterized by a mean annual temperature of -36 °C 

and an annual snow accumulation equivalent to 17 cm (6.7 

in.) of water. The flow of the continental ice sheet is very 

slow, probably less than 5 m per year. High-level striae and 

erractics indicate that the ice was formerly at least 500 m 

thicker than at present.

The Thiel Mountains of west Antarctica are situated 

between lat 85°00' S. and 85°30' S. and long 85°00' W. 

and long 95°00' W. (fig. 37.1). A program of geologic and topographic mapping of the area by the 

Geological Survey was started during January in the 

austral summer of 1961 and concluded between November 10, 1961, .and January 20, 1962, the following 

summer. The fieldwork was carried out using motortoboggans from a base camp (Camp Washington) supported by aircraft of the U.S. Navy. This report summarizes the results of studies of glaciology and glacial 

geology carried on by the author as part of the progo- E.

FIGURE 37.1. Index map of Antarctica showing the location of 

the Thiel Mountains.

gram. Data were collected on temperature variations 

in air, rock, and ice; snow stratigraphy; rates of glacial 

movement; and geomorphology. These data show that 

the average annual snow accumulation in the past 12 

years is equivalent to about 17 cm of water, the glacial 

plateau ice moves 1 to 5 m per year, and the ice cover 

in the Thiel Mountain area has been at least 500 m 

thicker than at present.

The Thiel Mountains consists of a northwest-trending bedrock and ice escarpment 150 km long and 500 m 

high. Hypersthene-quartz monzonite porphyry is predominant along the scarp, but biotite granite, metasedimentary rocks, and sedimentary rocks are also found 

(Ford and Aaron, 1962). Ice flows over the scarp 

from a high southern plateau about 2,200 m above sea 

level, to a lower northern plateau about 1,700 m above 

sea level (fig. 37.2). Several nunataks rise as high as 

800 m above the lower ice plateau in the southeastern 

part of the mountains, and gently sloping glaciers flow 

from the upper to the lower plateau through wide gaps 

between the nunataks (fig. 37.2). A cirque glacier 

terminating in a small ice-free valley was noted on one 

of the higher nunataks on the ice escarpment.

Both the upper and lower ice plateaus have gently 

undulating surfaces in most places. The margin of 

the lower plateau adjacent to the base of the mountain 

scarp, however, is a foothill belt, 1 km to 3 km wide, 

characterized by a rugged ice topography with abundant broad, closed depressions 10 to 100 m deep, which 

are commonly floored by blue ice.

The Thiel Mountains are about 500 km from the 

South Pole or 1,100 km from open sea and therefore 

have a continental climate. The mean annual temperature at Camp Washington on the lower ice plateau 

is 36°C as determined by measurement of snow 

temperature at depth. The mean air temperature for 

December 1961 was 16°C; the highest temperature 

recorded was 9°C. Temperature data are not available for the austral spring and winter in the area.

ART. 37 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B140-B143. 1963.

B140
B141

:GTTBE 37.2. Aerial view of the Thiel Mountain^, looking southeast. Length of scarp in foreground is approximately 5 milea

Photograph by U.S. Navy.

he annual snow accumulation for 1961 was 45 cm 

.(equivalent to 16 cm of water). During the period 

'November 10, 1961, to January 20, 1962, snow fall was 

recorded on 20 days, with a total accumulation of 10 

to 15 cm. During this 75-day period the wind was 

predominantly from the south. The maximum 

[recorded velocity was between 40 and 50 knots, but for 

30 days the maximum daily wind velocity was less than 

5 knots. The wind accounts principally for the distribution of snow on the plateau surfaces and for the 

[fact that it is generally well packed and de.nse, 

1 Continuous temperature measurements were made on 

a selected bedrock exposure in the vicinity of Camp 

[Washington for a 7-day period at the end of December 

1961, and for a 3-day period in early January 1962. A

typical daily thermograph is shown in figure 37.3. A 

Taylor recording thermograph equipped with two 

identical bright-silvered elements was used. The elements were buried in sandy gravel derived by weathering from the rock, one element (.4) to a depth of less 

than 1 cm, the other element (B} to a "depth of 4 to 5 cm. 

During 6 days of the total 10-day interval, in 24-hour 

daylight, the temperature at A ranged from 12°C to 

28°C during the hours the station was in direct sunlight, 

and dropped to 13°C while the station was in the 

shadow produced by nearby mountains. The corresponding temperature range at B wTas 1°C to 7°C in 

sunlight and -12°C to -13°C in shadow. The sharp 

dips of the thermograph near the maximum curve are 

interpreted to be the result of temporary cloud cover.
B142 GEOMORPHOLOGY, GLACIAL GEOLOGY, AND GLACIOLOGY

FIGURE 37.3 Diurnal temperature variations on bedrock at 

foot of a north-facing cliff about 10 miles west of Camp Washington. A, Temperature at rock surface with element covered 

by less than 1 cm of sandy gravel. B, Temperature with element covered by 4 to 5 cm of sandy gravel.

The air temperature at Camp Washington varied from 

 11°O- to 19°C during the period of measurement. 

Temperatures recorded in joints in the bedrock at several places in the Thiel Mountains also showed a considerable variation on both sides of the freezing point. 

Snow was commonly seen melting rapidly on sun-heated 

rock surfaces, and running water was seen on several 

occasions.

Three pits, 2 m, 4.5 m, and 5 m deep, were excavated 

in the plateau surface at Camp Washington for study 

of snow stratigraphy and temperature measurements. 

Coarse-grained porous snow (firn) and fine-grained 

hard-packed snow are interlayered in the upper parts 

of the pits. With increasing depth in the lower parts 

of the two deeper pits the snow becomes coarser, more 

equigranular, and more layered. The layers of coarsest 

firn presumably were formed during the summer, and 

correspondingly the thickest layers of hard-packed 

fine-grained windblown snow were deposited during the 

winter. Stratigraphic interpretations based on alternations of grain size and porosity are not completely 

straightforward, however, for thin interbedded layers 

of hard-packed snow and firn were observed that may 

have formed during a single summer. The snow 

stratigraphy of the 5-m pit, the corresponding snowdensity curve, and a preliminary interpretation are 

shown in figure 37.4. The pit exposes 12 years of snow

accumulation, and the average annual accumulation during this time is equivalent to about 17 cm of water.

Continuous temperature measurements were made 

with metal-expansion thermometers and a thermohm 

during the excavation of the snow pits and the boring 

of 2 holes 6.5 m and 10 m deep. These measurements 

showed a decrease of temperature with depth to a depth 

of 10 m. At and below that depth the temperature was 

 36 °C, which is presumably the mean annual temperature at that locality.

The rate and direction of movement of the ice sheet 

on the lower plateau near Camp Washington were determined in January 1961 and January 1962 by measurement of a stake network 10 km long surveyed from 

triangulation stations on bedrock. Preliminary calculations indicate a northeasterly flow at a rate of between 1 and 5 m per year. A similar survey of stakes 

placed upon 3 of the largest glaciers flowing across 

the escarpment indicates a maximum flow of about 1 m 

per month on a slope of 5° to 10°.

Glacially sculptured rock surfaces, glacial striae, and 

erratics found on nunataks in the Thiel Mountains at 

elevations of up to 500 m above the present ice sheet

50 

100

150

in 

2 200

UJ

Z

z 250

o

z

3C 

t

° 350 

400 

450 

500

"~v...!..; " T i i i 

" 

_:.-..

"f ':.....

"""" --..

....................v

* - . .

,.. ...........: .

""'" .

i i i * i i

WWW

8fc**ffl 

ffla^SS-s?

mxtm^t 

5SR^53S

raSKSBSBSSW

^^fetfe-^-J:

ksMWfe*

 ~

BOWSER

l»t>l-t><f

} 1960-61

} 1959-60 £j

> 1958-59 *

_i

< 3

\ 1957-58 z

1 *

k956-57 &

\ i 

co

1955-56 j=

< 

1954-55 u

i1953-54 £ 

co

1952-53 §:

\ 1951-52 

1950-51 

1949-50

300 350 400 450 500 

SNOW DENSITY, IN GRAMS PER LITER

FIGURE 37.4 Stratigraphy and variation of density of snow in 

5-m-deep pit at Camp Washington. Large dots represent observed densities. In stratigraphic column, an unbroken line 

represents a distinct break, usually with a thin, hard crust; a 

dashed line represents a less distinct break; and a hatched 

area represents a bed of very porous coarse-grained summer 

snow usually grading downward into hard-packed snow. 1961- 

62 represents summer season only.
ANDERSEN B143

indicate that the ice was once 500 m thicker than at 

present. This observation is significant in estimating 

former ice thicknesses in the interior parts of Antarctica, because the Thiel Mountains are close to the center 

of the continental sheet. Aughenbaugh (1961) reports 

a thickness 250 m greater than today in the Duf ek Massif, and Anderson (1961) reports a thickness that was 

200 m greater in the Sentinel Mountains. Similar observations in most of the coastal areas of Antarctica are 

further evidence that the ice sheet was once thicker and 

more extensive than it has been in recent historic time.

Observations in areas of bedrock indicate that the 

effects of repeated freezing and thawing are the most 

important weathering agents, and that frost wedging is 

a highly significant mode of mechanical weathering 

iji parts of the Thiel Mountains. Frost wedging is especially noticeable where the rock is highly jointed and 

mountain slopes are covered by angular blocks. Many 

rock faces in contact with snow surfaces are unusually 

steep or undercut, probably as a result of intensive frost 

wedging at the edge of the snow.

Several hill slopes covered by a mantle of brown 

silty to sandy gravel containing many angular blocks

are characterized by a patterned ground surface consisting of circles of poorly sorted material surrounding 

distinctly domed central areas. The circles are closely 

spaced and range from 1.5 to 5 m in diameter; the domes 

are 0.3 to 1 m high. Trenches across several of these 

circles showed that the domes are frozen ground covered by a blanket of unfrozen rock debris about 0.3 m 

thick on the highest part of the domes and about 0.5 

m thick in the depressions between the domes. Most 

of the larger blocks are segregated in the interdomal 

depressions, but otherwise sorting is poor. This particular type of patterned ground is not reported frequently in the literature; it is, however, probably common in high latitudes.

REFERENCES

Anderson, Vernon, 1961, Geographic features observed on the 

Marie Byrd Land traverse, 1957-1958: International Geophysical Year Glaciological Rept. 4, chap. 6, p. 143-149.

Aughenbaugh, N. B., 1961, Preliminary report on the geology of 

the Dufek Massif: International Geophysical Year Glaciological Rept. 4, chap. 8, p. 155-193.

Ford, A. B., and Aaron, J. M., 1962, Bedrock geology of the Thiel 

Mountains, Antarctica: Science, v. 137, no. 35S2, p. 751-752.
Article 38

INFLUENCE OF SNOW COVER ON FROST PENETRATION

By DANIEL B. KRINSLEY, Washington, D.C.

Abstract. Frost penetration is most effective through the 

thinnest snow cover, and proceeds more rapidly in gravel and 

sand than in silt or clay. Observations in Alaska, Wisconsin, 

and U.S.S.R. suggest that an increase of 1 to 4 inches in the 

thickness of a 6-inch snow cover significantly dampens the fluctuation of the ambient air temperature.

Three thermocouple cables were installed in the vicinity of Buffalo Center, Tanana Valley, Alaska, during 

September 1961. Each cable contained thermocouples 

at the ground surface and at depths of 6, 12, 18, 24, 36, 

and 48 inches. The sites were inspected during the first 

and third weeks of October, and at about 2-week intervals from November 24 to February 17,1962. Soil temperatures at the thermocouples were read directly from 

a calibrated potentiometer connected to the thermocouple cable. In addition to the thermocouple data, air 

temperature and snow depth were recorded at each 

site.

Site A, three-tenths of a mile south of the Little 

Gerstle River and approximately 50 feet above the 

stream channel, is adjacent to milepost 1388 of the 

Alaska Highway and is at an altitude of 1,300 feet. 

This well-drained area has been burned over, but currently is densely vegetated with black spruce, white 

spruce, and willow. The trees range from 15 to 25 feet 

in height, and the understory consists of low bushes and 

a mat of berries and grasses. Beneath the 2-inch-thick 

organic mat, glaciofluvial silt with several scattered thin 

gravel layers overlies coarse outwash gravel at a depth 

of 10 feet. No permafrost was found, and the water 

table is at least 20 feet below the ground surface (on the 

basis of exposed gravel pits in the vicinity).

Site B, 14 miles from site A and at an altitude of 

1,220 feet, is located on a sand dune adjacent to milepost 1402. The dune is composed of medium sand and 

is stabilized by a dense forest of poplar and birch 20 

to 30 feet in height. A 2-inch-thick organic mat composed of berry plants and vegetable fibers covers the 

sand, which is underlain by outwash gravel at a depth 

of 20 feet. No permafrost was found, and the water 

table is at least 25 feet below the ground surface (on 

the basis of exposed gravel pits in the vicinity).

Site C, 19 miles from site B and 33 miles from site A, 

is 300 feet west of the headquarters building at Fort 

Greely and is at an altitude of 1,280 feet. Dense stands 

of white spruce and birch, 20 to 30 feet in height, and 

a 2-inch-thick organic mat cover the site. The soil consists of 3 feet of eolian silt overlying coarse outwash 

gravel. Permafrost was not found, but it has been 

found previously in discontinuous irregular bodies at 

least 25 feet below the ground surface in the outwash 

gravel at Fort Greely (Holmes and Benninghoff, 1957, 

p. 169). The ground-water table has been found to 

range from 184 to 215 feet below the ground surface 

(Holmes and Benninghoff, 1957, p. 173).

The average air temperature during September 1961 1 

was 43.4°F, a departure of -2° from the 1943-60 

mean.2 The snow cover on September 30 was 1 inch 

thick. The average temperature during October 1961 

was 19.5°F, a departure of -5.7° from the 1943-60 

mean. By October 31, the snow cover was 19, 15, and 6 

inches thick at sites A, B, and G, respectively. September was a climatologically normal or average month, 

but October wag colder and had more snowfall and residual snow than the average October of the 17-year 

record. There were two periods of below-zero temperatures during October, each preceded and accompanied by snowfall. These periods occurred during 

the second and last weeks of the month.3

Ice appeared on the Tanana River and Jarvis Creek 

(adjacent to Fort Greely) on October 12 (U.S. Weather 

Bureau, 1962), which is normal for the area. Simultaneously, most of the ponds froze and flow ceased in 

several creeks. The soil at site C was frozen to a depth 

of 3 inches but it was not frozen at sites A or B.

During the third week in November, snow cover at 

sites A, B, and C was 19,15, and 6 inches thick, respectively. These thicknesses increased 2 inches at each

1 All air temperatures were recorded at the Fort Greely recording 

station, 0.6 mile from site A.

2 U.S. Army Signal Corps Meteorological Team, 1961, Climatology 

chart: Fort Greely, Alaska, Nov. [Duplicated report]

3 U.S. Army Signal Corps Meteorological Team, 1961, Monthly climatological summary: Fort Greely, Alaska, Sept., Oct. [Duplicated report]

ART. 38 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B144-B147. 1963.

B144
p PO"

MEAN DAILY AND MEAN MONTHL

IN DEGREES FAHR

I I I I 

ft W ho l-1 

O o O O 

Thermocouples

O O

D>

m

O

I I I I o +

g 8 5 5

DISTANCE FROM GROUND SURFACE, IN INCHES
B146 GEOMORPHOLOGY, GLACIAL GEOLOGY, AND GLACIOLOGY

site during the second half of December, and then 

decreased 3 inches at each site during January and 

February (fig. 38.1). The snow was light, dry, and 

free of packing or crusting at all sites. November was 

6.3° colder than normal (1943-60 mean); December 

was 17.2° colder than normal; and both January and 

February were 2.0° and 6.4°, respectively, warmer than 

normal.4

The air temperature at site C was generally within 

1° of the air temperature at the Fort Greely recording 

station (fig. 38.1). Above 13°F the air temperature 

at site C was always 2° to 3° lower than the air temperature at site A or B (which were generally identical); 

below 13° F it was always 2° to 3° above the air temperature at site A or B. All observations at the thermocouple sites were generally made during the hours 0800 

to 1200.

The mean temperature at the snow-ground interface 

for the period November 24 through February 17 was 

26.1°, 19.6°, and 1.1°F at sites A, B, and C, respectively. The mean air temperature during the period 

November 24 through February 17 was 8.3°F. According to Lachenbruch (1959, p. 29), "In the Arctic the 

mean winter temperature of the snow surface is generally probably somewhat lower than the corresponding 

air temperature because of the high emissivity of snow 

for long wave lengths." If it is assumed that the airsnow interface had a mean temperature (November 24 

through February 17) of approximately 9.0°F, then 

the mean temperature difference between the air-snow 

and snow-ground interfaces was 35.1°, 28.6°, and 10.1° 

at sites A, B, and C, respectively.

Frost penetration proceeded most rapidly at site C, 

which had the thinnest snow cover (fig. 38.1). Curve 

C steepens considerably after, the 32° isotherm passes 

through the eolian silt at a depth of 36 inches and enters' 

coarse outwash gravel. This is in accord with the observations made at Dow Field, Bangor, Maine (U.S. 

Army Corps of Engineers, 1947) and at Portland, 

Maine (Fuller, 1940) that frost penetration was faster 

in gravel and sand than in silt and clay. It seems 

reasonable to assume that curve B would have been 

closer to curve A if the material at site B had been silt 

rather than medium sand. Conversely, i f there had been, 

an identical snow thickness at all sites, curve B would 

have been below curve C at depths less than 36 inches, 

but would have crossed curve C somewhere below 36 

inches.

Observations at site B indicate that although its snow 

cover was 2.3 times as thick as that at site C, the insulating effect of the snow was 2.9 times as great as that at 

site C (table 38.1). The snow cover at site A was 2.9

* Idem, Nov. 1961-Feb. 1962.

TABLE 38.1. Summary of ratios of snow thickness, mean temperature differences between the air-snow and snow-ground interfaces, 

and mean frost penetration

Site

A.. ______________________

B -

C

Ratio of mean 

snow thickness 

(fig. 38.1)

2. 9 

2.3 

1.0

Ratio of mean 

temperature 

difference 

between 

snow-cround 

and air-snow 

interfaces 

(insulation)

3.5 

2.9 

1. 0

Ratio of mean 

frost 

penetration 

(Nov. 24- 

Jan. 7) 

(&?,. 38.1)

1.0 

1.5 

2.2

times as thick as that at site C, but the insulating effect 

of the snow was 3.5 times as great as that at site C. 

Although visual inspection of the snow at the three 

sites did not reveal any density differences, layers, or 

crusts, the existence of these differences cannot be discounted in the absence of more precise measurements. 

Examination of table 38.2 discloses the existence of a 

significant break in the insulating effect of snow at

TABLE 38.2. Differences in temperature between the air-snow and 

snow-ground interfaces at site C

Date

Nov. 24 __ ________

Dec. 11_. __________

Dec. 26-_--_---_-__

Jan. 7___ -.__

Jan. 20____-_- ___-

Jan. 27...- __ ___

Feb. 17 

Mean temperature

of air-snow interface, and total fluctuation within

period (degrees

Fahrenh

-9.71 

18. 3J

-45. 71

-26. 7 

7.3

-26. 7

-. 3

eit)

>28. 0

53.0

Snowdepth

(inches)

6 

6

8

7 

7

7

5

Temperature of

snowground

interface, 

and total

fluctuation

within

period

(degrees

Fahrenheit)

18J 14

-51

0 8 of8

3j

4

Difference

in temperature

between 

air-snow

and snowground

interfaces

(degrees

Fahrenheit)

13.7

.7

40. 7

26. 7 

9.3

29. 7

4. 3

thicknesses between 6 and 7 inches. This break has 

been referred to as the "critical thickness" of snow, 

which may be defined as the amount of snow cover required to significantly dampen the fluctuations of the 

ambient air temperature. The critical thickness of 

snow has been observed by others (table 38.3).

The insulating quality of snow is an expression of 

its thermal conductivity. This value is difficult to 

measure, and it depends on density, structure, texture, 

and interstitial air flow (Bader, 1962, p. 58). In addition, the thermal conductivity of snow is sensitive to the 

thermal properties of the underlying material (Lachenbruch, 1959, p. 24). In view of all these variables and 

the widely scattered observations (table 38.3), it is remarkable that the critical thickness of snow appears to 

occupy such a narrow range of values.
KRINSLEY BUT

TABLE 38.3. Critical thickness of snow observed in this and 

other studies

Source

Krinsley 

(this 

report) . 

Pruitt (1957). 

Kabanov 

(1937). 

Atkinson 

and Bay 

(1940).

Location

Buffalo Center, 

Alaska.

Fairbanks, Alaska _ 

U.S.S.R___. ______

La Crosse, Wis____

Mean 

annual 

temperature 

(degrees 

F)

25. 7

24. 7

46.3

Minimum temperature in year 

of observation

-61 (1961) 

-55 (1956)

-38 (1936)

Critical thickness 

(inches)

6-7

6-8 

8-10

10

The data summarized in tables 38.1 and 38.2 suggest 

that the transfer of heat through dry snow is reduced 

abruptly when a critical thickness is attained. The 

measurement of the exact rate of reduction at critical 

thickness and the explanation for this phenomenon are 

worthy of further study.

REFERENCES

Atkinson, H. B., and Bay, C. E., 1940, Some factors affecting 

frost penetration: Am. Geophys. Union Trans., v. 21, pt. 

3-B, p. 935-951.

Bader, Henri, 1962, The physics and mechanics of snow as a 

material: U. S. Army Corps Engineers Cold Regions Research and Engineering Lab., Cold regions science and 

engineering, pt. 11, sect. B, 79 p.

Fuller, H. U., 1940, Studies of frost penetration: New England 

Water Works Assoc. Jour., v. 54, no. 3, p. 275-281.

Holmes, G. W., and Benninghoff, W. S., 1957, Terrain study of 

the Army Test Area, Fort Greely, Alaska: U. S. Geol. Survey, Military Geology Branch, v. 1, 287 p.

Kabanov, P. G., 1937, Soil temperature and warming of winter 

crops [abs.j : Am Meteorological Soc. Bull., v. 18, p. 65-66.

Lachenbruch, A. H., 1959, Periodic heat flow in a stratified medium with application to permafrost problems: U. S. Geol. 

Survey Bull. 1083-A, 36 p.

Pruitt, W. O., Jr., 1957, Observations on the bioclimate of some 

taiga mammals: Arctic, v. 10, no. 3, p. 131-138.

U. S. Army Corps of Engineers, 1947, Report on frost investigation, 1944-45: Boston, Mass., New England Division, 57 p.

U. S. Weather Bureau, 1961, Climatological data Alaska: v. 47, 

no. 10, Oct. 1961, p. 156-171.
Article 39

ORIGIN AND NATURE OF THE PROBABLE SKELETAL FUZZ ON THE MOON

By CHARLES R. WARREN, Washington, D.C.

Work done in cooperation with the Office, Chief of Engineers, Department of the Army

Abstract. The most probable origin of skeletal fuzz on the 

Moon is that it is a residue from vesicular rocks whose surface is skeletonized by sputtering caused by the solar wind. 

Such fuzz, blackened by radiation damage, would have most 

of the known physical properties of the surface material of 

the Moon.

The surface material of the Moon has the property 

of strongly retrodirecting light that falls on it; that 

is, sunlight incident on the lunar surface is backscattered (reflected back in the direction of the light 

source, the Sun) much more strongly than it is §cattered in other directions. The conclusion that light is 

retrodirected is based on three otherwise anomalous 

facts: the extreme steepness of the photometric function (the curve showing the change in brightness of 

the Moon with the phase) ; the shapes of the photometric-function curves for individual points on the 

Moon, which have maximums at or near full moon 

(Diggelen, 1959) instead of at local noon as theory 

predicts; and the lack of limb darkening on the full 

Moon.

To give the observed retrodirected light, the surface 

material of the Moon probably must have an extremely 

complex structure (Hapke, 1962). Analysis of the 

requirements for this complex structure (Warren, 

1963) shows that the Moon is probably covered, to a 

depth of several millimeters or centimeters, by a highly porous, open-textured, skeletal fuzz that consists at 

least in part of randomly oriented linear units. These 

linear units either anastomose in a three-dimensional 

mesh, or branch, as in snowflakes or reindeer moss; in 

either case they may support nodes as in a TinkerToy structure, or may not.

The skeletal fuzz evidently grows or develops at the 

Moon's surface, for it is present in craters and on their 

rims, which consist of material thrown out from some 

depth, yet it is so porous that it must be too fragile 

to survive a crater-forming explosion, or even to exist 

at depth, where it would have to support the weight

of a thick overburden. The layered structure of the 

lunar surface (Gibson, 1961) also shows that the fuzz 

must be a surface feature.

Three general types of origin may be considered for 

the lunar surface fuzz. (1) It might form primarily 

from new material, perhaps burying a prefuzz surface 

that remains essentially unaltered: new material arrives on the Moon not only in the form of meteoroids, 

including meteorites and micrometeorites, but also as 

individual atoms, which might condense onto the surface from the solar wind. (2) It might form chiefly 

from material already present at the Moon's surface, 

the texture opening out without quantitatively significant gain or loss of matter: the meteoroids that bombard the lunar surface from time to time must greatly 

modify it. (3) The fuzz might form by the removal 

of material, etching out and skeletonizing the surface: 

the Moon's surface continuously loses material as the 

protons of the solar wind sputter it away.

The solar wind is a flux of ionized particles moving 

away from the Sun that has been shown by the recent 

satellite and space-probe data to be continuously present in interplanetary space. Most of the particles are 

protons, though a relatively few helium nuclei and 

even fewer heavier atoms are present. As the protons 

approach the Moon, they are not deflected into belts 

comparable to our Van Alien belts because the Moon 

has essentially no magnetic field (see Markov, 1960, 

p. 365). Neither are they absorbed by air, because 

the lunar atmosphere is so tenuous that the mean free 

path of an atom or molecule in it is much longer than 

the total thickness of the atmosphere. Thus nearly 

all the protons strike directly on the Moon's surface. 

Laboratory experience shows that many of the protons 

must sputter (eject) an atom from the surface, in most 

cases with a velocity above the escape velocity from the 

Moon. Because very few of these atoms are slowed 

by collisions with atoms of the lunar atmosphere, they

ART. 39 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B148-B152. 1963.

B148
WARREN B149

must escape from the Moon and be permanently lost 

to it.

The long-term average rate of this loss of matter by 

sputtering is open to question because of uncertainties 

in the average intensity of the solar wind and the 

energies of its particles, in the efficiency with which 

various kinds of particles at various energies sputter 

various materials, and in the nature of the materials 

actually present on the Moon. However, the best estimate appears to be that over a period of time the Moon 

loses more matter than it gains. Wehner and others 

(1960) estimate the net loss at 40 meters of rock per eon 

(109 years), or 1 millimeter every 25,000 years, and the 

actual rate may be higher than this if the mechanism 

accounts for the fact that the Moon's surface appears 

to have been eroded to a depth of more than a kilometer 

since the oldest recognizable craters were formed 

(Baldwin, 1949; Gold, 1955).

On this basis, the third possible process of forming 

the skeletal fuzz on the Moon, by etching and skeletonizing the surface, seems more likely than the first, by 

adding new material. However, the process of opening out the texture by meteoroid impacts -may be an 

important contributing factor.

The sputtering process might conceivably produce a 

skeletal structure by building metal whiskers from the 

residue of sputtered silicates, but metal whiskers would 

have a much higher electrical conductivity than the 

surface material of the Moon, which appears to have 

electrical characteristics "similar to those of pumice or 

dust" (Mayer, 1961, p. 454). I therefore suggest that 

the most probable origin of the skeletal structure is by 

skeletonizing vesicular rock.

Suppose that a vesicular rock is exposed on the surface of the Moon. Sputtering by the solar wind slowly 

etches the substance of the rock. In a few thousand 

years the thin walls between adjacent pairs of vesicles 

will be sputtered away entirely, leaving as a rodlike 

body the thicker material between three adjacent 

i vesicles, where three such thin walls converged. If the 

| vesicles have the close-packed arrangement,1 the corners 

between four adjacent holes will be left as tetrahedral 

nodes, where four rods converge. The result will be 

a type of Tinker-Toy structure, a skeletal mesh of thin 

rods joining at thicker nodes. As the sputtering con- 

, tinues, the rods and nodes will be reduced further, but 

at the same time the skeletonizing process will penetrate 

! farther. By the time all the rods supporting a given

» The three-dimensional distribution of the vesicles in vesicular rocks 

has not been studied in detail, to my knowledge. The spacing of nucleation centers may control this distribution, rather than the minimum surface-tension energy as in a foam, because the viscosity of a 

silicate melt is much higher than that of the liquid in a foam.

node are gone, perhaps the node will be the size of a 

grain of dust rather than of sand. The dust grain will 

drop into one of the holes in the surface, but presumably 

such particles are not numerous enough to clog up the 

texture, and the skeleton can maintain its structure.

The only other reasonable explanation for the extremely complex structure of the Moon's surface that I 

have seen proposed is Hapke's suggestion (1962) that 

the surface consists of "fairy castles" of dust. However, 

the dust "castles" that Hapke has produced in the laboratory are believed to owe their complex structure to 

electrostatic charges, and would not be likely to remain 

stable for thousands of years under monthly cycles of 

ionizing radiation. Moreover, if the particles were 

only a few microns in diameter, as Hapke's suggestion 

requires, they probably could not have the opacity 

necessary to cast dark shadows; dust is normally lighter 

in color than the rocks from which it is derived, and 

the low electrical conductivity indicates that the lunar 

surface material cannot be metallic.

The hypothesis that the skeletal fuzz on the Moon is 

formed by the sputtering of vesicular rocks requires 

that the lunar surface rocks be vesicular. Three general 

classes of vesicular rocks might be expected on the 

Moon: (1) Vesicular volcanic rocks such as pumice, 

scoria, or pumiceous tuff may be present. (2) Impactite, a vesicular material found around meteorite 

craters on the Earth, is doubtless present. Sharonov 

(1960, p. 359) attributes to Madame Sytinskaya the hypothesis that the entire lunar surface is covered by a 

"meteor slag" of this origin. (3) Vesicular material 

from planetesimals that became aggregated to form the 

Moon in the first place may be present. Comet nuclei 

and many meteoroids in interplanetary space may have 

an extremely porous texture produced by sublimation of 

interstitial ices (including perhaps frozen methane, 

ammonia, and carbon dioxide as well as water). The 

fact that meteorites on the Earth do not show such 

porosity might be because the texture is too insubstantial to survive frictional ablation by passage 

through the air, or too fragile to survive impact on the 

ground. On the Moon, however, some porous texture 

might survive, especially if the interstitial ices were 

present at impact to provide support, because impact 

on the Moon would involve essentially no atmospheric 

friction and might occur at less than one-fourth the 

velocity of impact on the Earth: the escape velocity 

from the Moon, and therefore the minimum velocity of 

infall on the Moon, is 0.213 that on the Earth.

All three of these types of vesicular rock are very 

probably present on the Moon. If the material flooring 

the maria is ignimbrite (welded tuff), as suggested by
B150 ASTROGEOLOGY

O'Keefe and Cameron,2 its surface contains much 

pumice; if it is basalt, as supposed by many earlier 

writers, its surface consists of scoria or vesicular basalt. 

The uplands, on the other hand, must have large 

amounts of impactite and may have much porosity inherited from cometlike planetesimals as well. Isostatic 

evidence suggests that the terra material may be of low 

average density compared with the mare material; 

mapping by the Army Map Service (1962) indicates 

that the Central and Southern Highlands (Hackman 

and Mason, 1961) stand an average of 4 to 6 km higher 

than Mare Imbrium, where the mare material is probably thickest.

If their surfaces were skeletonized by the solar wind, 

any or all of these vesicular rocks would have most of 

the known physical properties of the lunar surface 

material. For example, they would retrodirect the sunlight; they would have the observed infrared thermal 

properties, including two anomalous ones; the thermal 

observations at radio wavelengths would be explained; 

and the electrical properties of the lunar surface, including the inferred layering, might be expected.

In infrared thermal properties, the skeletonized 

vesicular rocks can be expected to have the low thermal 

inertia reported by Sinton (1961, p. 438) because both 

the density and the thermal conductivity are low. The 

density is low because most of each cubic centimeter 

(at least 75 percent, and more likely 90 or 95 percent) 

consists of vacuum. The thermal conductivity is low 

because this vacuum is sufficiently subdivided to reduce 

heat transfer by radiation, and little heat can be transferred by conduction along the thin rods.

Three infrared thermal properties of the lunar surface, two of which are anomalous, can be explained 

as resulting from the thermal isolation of the individual 

nodes in skeletal fuzz. First, at full moon, as Pettit 

observed (1961, p. 415), the infrared radiation temperature drops off more slowly from the subsolar point to 

the limb than is predicted for a smooth, slowly rotating 

sphere. This is explained as being because most nodes 

of the skeletal fuzz near the limb that are visible from 

the Earth are in full sun, and are not much cooled 

by heat conduction to elements that are not in the sun; 

they are therefore nearly at the equilibrium temperature 

for a node in full sun radiating to a hemisphere of 

space. The temperature measured at the subsolar point 

is higher chiefly because at that point the bolometer 

receives radiation not only from nodes at the surface, 

but also from the deepest nodes to which the Sun's rays 

penetrate. A node in the interior of the fuzz receives 

as much solar heat as one near the surface, but its

2 J. A. O'Keefe and W. S. Cameron, paper presented at COSPAR 

meeting, May, 1962, entitled "Evidence from the Moon's surface features 

for the production of lunar granites".

efficiency in disposing of this heat by radiation is 

reduced by the fact that only a small part of the area 

to which it can radiate is at the cold radiation temperature of space. Thus, to reach equilibrium and reradiate 

heat as fast as a node receives it from the Sun, the 

deep node must radiate at a higher temperature. The 

infrared temperature measured is a mean between the 

high temperatures of the elements (rods and nodes) in 

the interior and the lower ones of those at the surface.

Second, the fuzz explains the anomaly noted by 

Pettit (1961, p. 412) that the infrared temperature of 

the subsolar point depends on the angle of view; the 

measured temperature is roughly 50 °K lower when the 

subsolar point is at the limb, at quadrature, than when 

it is at the center of the apparent disk, at full moon. 

This is explained because at the limb the bolometer 

measures radiation only from the cooler surficial fuzz 

elements, and not from the deeper and hotter elements.

Third, as Strong and Sinton noted (Sinton, 1961, p. 

439), during the penumbral stage of a lunar eclipse the 

limb areas lose a much larger proportion of their heat 

than the center of the disk, in spite of the fact that they 

were already cooler before the onset of the eclipse. 

This is explained as being because the elements seen at 

the limb are those near the surface, which cool most 

efficiently by radiation during the penumbral phase.

The observation that the radiation temperature of 

the subsolar point depends on the wavelength at which 

it is measured is generally interpreted (for example, 

Gibson, 1959) to indicate that as the wavelength is increased the radiation measured comes from increasing 

depths beneath the surface. This implies that the surface material is partly transparent to longer waves, 

though it is opaque to optical and infrared waves. A 

skeletal fuzz would be expected to have this property.

The known electrical properties of the lunar surface 

material can also be explained. The low overall electrical conductivity, similar to that of "pumice or dry 

sand" (Mayer, 1961), is to be expected for the fuzz because it is mostly vacuum and because the surface of 

such matter as is present is highly broken up, so that 

path lengths are great and resistance is increased. Gibson has inferred (1961) that a subsurface layer has a 

significantly higher electrical conductivity than the 

0.5-cm layer at the surface. This inference might be 

explained by saying that below 0.5 cm only the horizontal cell walls have been breached, the vertical ones being shielded from the protons of the solar wind by fuzz 

elements above them.3

3 This explanation may be quantitatively inadequate, however, and 

a further explanation, also based on the sputtering of vesicular rock, 

will be offered in a report that is in preparation.
WARREN B151

The remaining known physical properties of the 

lunar surface material are not necessarily inherent in . 

skeletonized vesicular rock. For example, the Moon is 

dark colored, about as black as a domino or a rich garden soilf whereas skeletonized vesicular rock could as 

easily be light colored. None of the properties appear 

to be incompatible with the inferred material, however, 

and some of them are easily explained by slight extensions of the hypothesis. Thus if the material of the 

vesicular rocks has the property of darkening slowly 

on exposure, not only the black color but also the relations of the lunar rays would be explained. And if the 

vesicular rock types present in the maria differ from 

those in the uplands, their various differences in observed properties could result.

The Moon is so much blacker than most terrestrial 

rock materials, including desert soils, that some process of blackening at the surface seems probable. This 

may be some form of radiation damage, such as by ultraviolet or X-radiation or by protons of the solar wind.

Provided this darkening process is slow, the observed 

relations of the lunar rays are explained by the skeletonized-vesicular-rock hypothesis. A lunar ray or ray 

crater is an area where fragments thrown from a craterforming explosion have either buried the darkened surface material or plowed "it up and turned it over, exposing the underlying lighter colored material. If the 

fresh material is vesicular, it will quickly be skeletonized and will eventually be darkened, in its turn, so 

that the ray will disappear into the background, as those 

formerly radiating from Eratosthenes have done (Shoemaker and Hackman, 1963). However, the darkening 

will be differential. At an intermediate stage, when 

the ray is partly darkened, the deeper fuzz elements, 

shaded from the Sun most of the time, will still be relatively bright when those near the surface, which are in 

sunshine all 14 days of the lunar day, have become well 

blackened. At this intermediate stage, only the blackened surface elements can be seen when the Moon's surface is illuminated by sunlight at large phase angles. 

Near full moon, however, one can see the deeper elements that the Sun reaches at no other time, and that 

are therefore little darkened. Moreover, in the fresh 

rock, where skeletonizing is still proceeding actively, we 

will see at full moon many vesicles in the interior of the 

fuzz to which sputtering has only recently broken 

through, and which have as yet scarcely begun to darken. These two factors would account for the observed 

fact, often considered anomalous, that the lunar rays 

are bright on the full Moon but fade out and may disappear entirely at large phase angles, near the terminator (the twilight line marking the edge of the sunlit portion of the Moon).

The differences between the maria and the uplands 

are also reasonable. If different types of vesicular rock 

are present, the difference in color could reflect a difference in susceptibility to the blackening process. The 

difference in blackenability may apply to the rate as 

well as to the ultimate degree; lunar rays on terra material are brighter than those across maria. Other differences, such as the generally greater degree of polarization of light from the maria, could result from 

differences in chemical composition, in texture, or in 

scale size of the fuzz.

In summary, the extremely open, porous texture of 

the skeletal fuzz that probably covers the Moon is more 

likely to be produced by removal of material by etching 

out the surface, than by adding to it. It could be produced by metal whiskers, but these would lack the observed low electrical conductivity. The geologically 

most reasonable explanation of the fuzz appears to be 

that it is vesicular rock skeletonized by the solar wind. 

The rock could be volcanic (for example, tuff or scoria) 

in the maria, and of impactite or planetesimal origin, 

or both, in the uplands. Such skeletonized vesicular 

rocks would have most of the known physical properties of the lunar surface, and would have no properties 

that appear to be incompatible with the results of measurements yet made on the Moon.

REFERENCES

Army Map Service, 1962, Lunar map 1:2,500,000, Mare Nectaris-Mare Imbrium. 

Baldwin, R. W., 1949, The face of the Moon: Chicago, 111., Univ.

Chicago Press, 239 p. 

Diggelen, J. van, 1959, Photometric properties of lunar crater

floors: Recheches astronomiques de 1'Observatoire d'Utrecht, v. 14, no. 2,114 p. 

Gibson, J. E., 1961, Lunar surface characteristics indicated by

the March 1960, eclipse and other observations: Astrophys.

Jour., v. 133, p. 1072-1080. 

Gold, Thomas, 1955, The lunar surface: Monthly Notices Royal

Astron. Soc. [London], v. 115, p. 585-604. 

Hackman, R. J., and Mason, A. C., 1961, Engineer special study

of the surface of the Moon: U. S. Geol. Survey Misc. Geol.

Inv. Map 1-351. 

Hapke, B. W., 1962, Second preliminary report on experiments

relating to the lunar surface: Cornell Univ. Center for

Radiophysics and Space Research, CRSR 127, 41 p. 

Markov, A. V., ed., 1960, Conclusion, p. 365-378 in Luna: Moscow, Gosudarstvennoye Izdatel'stvo Fiziko-Matematicheskoy

Literatury. 

Mayer, C. H., 1961, Radio emission of the Moon and planets,

chap. 12 in Kuiper, G. B., and Middlehurst, B. M., eds.,

Planets and satellites, The solar system, v. 3: Chicago, 111.,

Univ. Chicago Press, p. 442-472.
B152 ASTROGEOLOGY

Pettit, Edison, 1961, Planetary temperature measurements, 

chap. 10 in Kuiper, G. B., and Middlehurst, B. M., eds., 

Planets and satellites, The solar system, v. 3: Chicago, 111., 

Univ. Chicago Press, p. 400-428.

Sharonov, V. V., 1960, The nature of the lunar surface, chap. 9 

In Markov, A. V., ed., Luna: Moscow, Gosudarstvennoye 

Izdatel'stvo Fiziko-Matematicheskoy Literatury. p. 331-364.

Shoemaker, E. M., and Hackman, R. J., 1963, Stratigraphic 

'basis for a lunar time scale, in Kopal, Zden£k, and Mikhailov, Z. K., eds., The Moon 'Symposium No. 14 of the 

International Astronomical Union: London and New York, 

Academic Press, p. 289-300.

Sinton, W. M., 1961, Recent radiometric studies of the planets 

and the Moon, chap. 11 in Kuiper, G. B., and Middlehurst, 

B. M., eds., Planets and satellites, The solar system, v. 3: 

Chicago, 111., Univ. Chicago Press, p. 429-441.

Warren, C. R., 1963, The lunar surface material: Science, v. 140, 

p. 188-190.

Wehner, G. K., Laegreid, N., and Stuart, R. V., 1960, Study of 

sputtering of materials: General Mills, Inc., Rept. 2133 for 

Geophysics Research Directorate, Air Force Cambridge Research Center, Air Research and Development Command, 

U.S. Air Force.
Article 40 

PILLOW STRUCTURES OF SUBMARINE BASALTS EAST OF HAWAII

By JAMES G. MOORE and RONALD K. REED, Hawaiian Volcano Observatory, Hawaii, 

and U.S. Coast and Geodetic Survey, Washington, D.C.

Abstract. Submarine photographs from two camera sites 

east of the island of Hawaii on the submerged part of the east 

rift zone of Kilauea Volcano show fresh lava flows with welldeveloped pillow structure. The sates are at depths of 10,800 

and 14,400 feet and are offshore 32 and 38 miles. Photographs 

from a third site on the summit of Bushnell Seamount, at a 

depth of 5,200 to 6,200 feet, show in general a surface that appears old, but one young, fresh lava flow with pillow structure 

was photographed.

In October 1962, a deep-sea dredging program was 

Undertaken to collect volcanic-rock samples from the 

Ocean bottom adjacent to southern Hawaii. The 

sampling, which was part of a cooperative program 

between the U.S. Geological Survey's Hawaiian Volcano Observatory and the U.S. Coast and Geodetic 

Survey, was done aboard the USCGS ship Pioneer. 

puring the course of dredging, submarine photographs 

ivere taken at three sites (fig. 40.1). Two of the sites 

are on the submarine part of the east rift zone of Kilauea Volcano at depths of 10,800 feet and 14,400 feet. 

The third site, at a depth of 5,200 to 6,200 feet, ig on 

the summit of Bushnell Seamount, a submarine volcanic mountain 75 miles southeast of the island of 

Hawaii.

; The authors wish to acknowledge the help of Captain H. G. Conerly and the officers and men of the 

Pioneer, whose ability and interest made the project a 

success. Mr. Byron Hale of the Coast and Geodetic 

Survey processed the film and provided valuable advice.

CAMERA

The camera system, built by Edgerton, GermeshauSen, and Grier, Inc., consists of 2 cameras mounted 3 

: :eet apart, a light source 6 feet to the side of the cameras, and a sonar pinger for bottom detection. These 

: !our units are contained individually in steel cylindrical 

lousings. The housings are attached to a slotted 

:nounting rack, the whole being suspended on 5/32-inch

oceanographic wire handled on the ship by an electrohydraulic winch.

The 35-mm cameras utilize a Hopkins f/11 lens; no 

shutter is necessary since exposures are made only at a 

depth greater than 600 feet, the lower limit of light 

penetration. The film is exposed at 12-second intervals 

each time the lamp flashes, and the film (on a 100-foot 

roll) is advanced after each exposure by a 6-volt filmdrive motor in synchronization with the light source.

The 100 watt-second light source is supplied by a 

6-volt silver-zinc rechargeable battery. Also within the 

light source housing is a 6-volt timing motor which 

drives a cam to regulate the cycling. The cameras and 

flash are triggered by a delay timer which is set to allow 

sufficient time for the unit to be lowered to the bottom 

(generally 2 hours) before exposures begin.

The sonar pinger consists of a driver and 6-volt silver-zinc battery mounted in the steel cylindrical housing, a pulse transformer contained in a resilient oilfilled housing, and a transducer mounted directly on the 

camera-system frame. The pinger system produces 12- 

kc pulses at precise 1-second intervals. The direct signal and the signal reflected from the ocean bottom are 

received and displayed on the depth recorder in the 

ship's chart room. The difference in arrival time of 

these signals indicates the height of the camera above 

the bottom. The camera is positioned approximately 

8 feet above the bottom during the entire operation by 

the winch operator on the deck, who receives telephone 

instructions from the depth-recorder operator.

Under ideal conditions with the camera 8 feet from 

the bottom, an area 4 by 5 feet is photographed. Actually, many photographs were taken farther from the 

bottom and are underexposed. The ship drift tows the 

camera along the bottom at 1 to 2 knots and hence individual exposures are taken at distances of 20 to 40 

feet.

ART. 40 IN U.S. GBOL. SURVEY PROF. PAPER 475-B, PAGES B153-B157. 1963.

B153
B154 MARINE GEOLOGY

158°

BASE FROM U S NAVY HYDROGRAPHIC 

OFFICE PRELIMINARY SHEET BC 04N

FIGURE 40.1. Map showing submarine topography adjacent to the island of Hawaii, and location of three

ocean-bottom camera sites.

EAST RIFT ZONE OF KILAUEA

Two camera sites were occupied on the east rift zone 

of Kilauea Volcano (fig. 40.1, sites 1 and 2) at depths 

of 10,800 feet and 14,400 feet and distances of 32 and 38 

miles from land. The submarine part of the rift zone 

is marked by a prominent ridge extending directly on 

line with the trend of the rift zone on the island. Because the sides of the narrow ridge slope rather steeply 

(10°-12°), lava would almost certainly flow from vents 

along the crest of the ridge down one side or another. 

Hence photographs and samples of lava flows taken on 

the crest of the ridge are representative of materials 

near a volcanic vent and not of lava which has flowed 

many miles, as would be the case if the lower slopes of 

the ridge were examined.

Most of the photographs show original, unmodified 

surfaces of fresh-looking lava flows. The surfaces of 

the flows are clean, being free of silt, sand, or mangat O J 7 O

nese encrustations. Living organisms have apparently 

not had time to establish themselves well because not a 

single organism was seen in 45 photographs. On Bushnell Seamount, on the other hand, many photographs 

show living creatures.

A most significant feature of the lava flows on the east 

rift zone is the dominance of bulbous or cylindrical 

masses of lava which average a foot or two in diameter

(fig. 40.2, p. B156) and strongly resemble the pillows so 

common in ancient submarine lavas (Tyrrell, 1929, p. 

37-39). These subspherical masses are presumably produced when an elastic skin forms on the surface of the 

flow as it is chilled by contact with water. Internal 

pressure causes the skin to expand and produce the pillow, at which point it may rupture and feed a new 

pillow.

Similar bulbous masses are not uncommon on the surfaces of some subaerial flows and have been called blisters, squeeze-ups, or toes (Nichols, 1939, and Wentworth and Macdonald, 1953). However, the welts on 

subaerial flows are generally scattered over the surface 

or occur on the margins of pahoehoe flow units, whereas 

those shown in the photographs are closely packed together, both side by side and on top of one another and 

comprise virtually the entire lava flow. Moreover, the 

areal distribution of the submarine pillow lavas is much 

greater than that of subaerial flows showing bulbous 

squeeze-ups. At camera sites 1 and 2, 76 percent and 

58 percent, respectively, of the photographs show pillow 

structures; the other photographs show rubbly surfaces. In contrast, only a small percentage of subaerial 

lava flows shows bulbous structures similar to pillows.

Fifteen dredge hauls of basalt were recovered from 

the east rift zone of Kilauea and are presently under-
MOORE AND REED B155

going petrographic and chemical study. All of the 

basalt is fresh, with glassy outer rinds and none is 

palagonitized.

BUSHNELL SEAMOUNT

The surface at the top of Bushnell Seamount (fig. 

40.8, p. B157), at a depth of 5,200 to 6,200 feet, is markedly different from that of the east rift zone of Kilauea 

and appears much older.

Sandy surfaces are a common feature in many of the 

photographs, and rocky outcrops commonly show no 

original volcanic flow surfaces but instead are covered 

with characteristic botryoidal manganese encrustations.

! Living bottom dwellers are common and appear in 41 

percent of the 86 photographs taken. In addition the 

tracks of swimming or crawling animals are visible in 

nearly every photograph showing a sandy bottom.

1 Most of the organisms are branching coelenterates, or 

delicate echinoderms affixed to the bottom. Brittlestarlike creatures and siliceous sponges are also common, and two photographs show an eel-like fish about 

ajfoot long.

Among the apparently old and weathered rocks of 

Bushnell Seamount, one young lava flow ~with welldeveloped pillow structure was photographed (fig. 40.3, 

D). In the 86 photographs obtained, of the seamount, 

6 consecutive photographs show this young flow. The 

line of traverse across the flow was approximately 180 

feet long, assuming 30 feet between photographs. 

Hence Bushnell Seamount has had relatively recent 

volcanic activity, and an extensive dredging operation 

could no doubt recover fresh lava. Two dredge hauls 

on the seamount, however, recovered only altered and 

manganese-encrusted basalt.

REFERENCES

Nichols, R. L., 1939, Squeeze-ups: Jour. Geology, v. 47, p.

421-425. 

Tyrrell, G. W., 1929, The principles of petrology, 2d ed.: New

York, B. P. Dutton and Co., 349 p. 

Wentworth, C. K., and Macdonald, G. A., 1953, Structures and

forms of basaltic rocks in Hawaii: U.S. Geol. Survey Bull.

994, 98 p.

680231
wI *

or os

 Lk ^^^^BHHH^H

FIGURE 40.2. Surfaces of submarine lava flows of east rift zone of Kilauea Volcano. Camera site 1, depth 10,800 feet: A, fresh lava-flow top showing pillow structure; B, fresh lava-flow top showing elongated pillows draped on pile of pillows; 0, fresh lava-top showing pillow structures. 

Camera site 2, depth 14,400 feet: D, fresh lava-flow top showing pillows with trace of pahoehoe ropy structure. Area of photographs 

approximately 4 by 5 feet.
FIGURE 40.3. Surface features near summit of Bushnell Seamount. Camera site 3, depth 5,200 to 6,200 feet: A, Echinodermlike animal attached 

to bottom composed of manganese-coated pebbles; area approximately 2 by 2% feet. B, Animal tracks on sandy bottom with ripple marks; 

area approximately 4 by 5 feet. C, Eel-like fish above pebbly bottom; area approximately 4 by 5 feet. D, Fresh lava-flow top showing pile of 

pillows; brittle star in bottom center; area approximately 4 by 5 feet.

td
Article 41

ESTIMATES OF WORLD BAUXITE RESERVES AND POTENTIAL RESOURCES

By SAM H. PATTERSON, Beltsville, Md.

Abstract. A new review of world bauxite reserves indicates 

that total world reserves are 5.7 billion tons and that marginal 

and submarginal resources are 8.7 billion tons.

A new review of world bauxite reserves, listed by 

country in the accompanying table, indicates that total 

world reserves are 5.7 billion tons and that marginal 

and submarginal resources are 8.7 billion tons. The 

world reserves are more than triple those known a 

decade ago (Fischer, 1953). The increase is due to 

vigorous exploration for bauxite, which has resulted 

in the discovery of several important deposits, including the very large Weipa deposits of the Cape York 

peninsula, Australia, and the deposits near Boke, Kindia, and Fria in the Republic of Guinea.

Approximately 35 percent of the world's bauxite 

reserves are in Australia, 20 percent in the Republic 

of Guinea, 10 percent in Jamaica, 5 percent in Hungary, 5 percent in Yugoslavia, 4.5 percent in Ghana, 

and 4.5 percent in Surinam. Of the marginal and 

submarginal resources, 45 percent are in the Republic 

of Guinea, 13 percent in Australia, 11 percent in British 

Guiana, 11 percent in Cameroun, and 11 percent in 

Communist China. Other countries possessing adequate reserves and other resources to maintain significant bauxite production for several decades include 

France, Greece, and India.

The United States contains less than 1 percent of 

the world's bauxite reserves. It is in a somewhat more

favorable position in regard to the size of its submarginal resources of bauxite, but these resources cannot be used profitably now because they consist chiefly 

of low-grade bauxite in Hawaii and Oregon and deeply 

buried, thin deposits in Arkansas and Georgia.

In 1961 domestic consumption of bauxite was 8.6 

million tons and world consumption was 29 million 

tons (Higbie and others, 1962). At this rate of consumption, domestic reserves of bauxite would last less 

than 6 years if the foreign sources now supplying most 

of the United States' requirements were cut off. Known 

world reserves are adequate to last more than 175 years 

at the current rate of depletion, and if prices and technology adjust to permit mining of the marginal and 

submarginal resources, they would supply the world's 

needs for an additional 200 years or more.

The estimates are based on information from more 

than 200 geological and mining publications, trade 

journals, and U.S. State Department Foreign Service 

dispatches. The limitations of this publication do not 

permit the listing of the sources of information; however, a bibliography will be included in a U.S. Geological Survey publication in preparation.

The author is indebted to Miss Mary E. Trought, 

Mr. Richard D. Wilmot, and Mr. Kenneth B. Higbie, 

of the U.S. Bureau of Mines, whose suggestions led to 

obscure reference data for several countries.

ART. 41 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B158-B159. 1963.

B158
PATTERSON B159

Estimated total reserves and other resources of bauxite, in millions 

| of tons l

[L, large; S, small]

Drth America: 

United States:

Southeastern States

Oregon __ ___ ____ _____

Hawaii __ __ _________

Total (rounded) _______________

entral America: 

Costa Rica___ _ ____ ___ _______

Mexico __ ____ _________

Panama ____ _ _________ _

Total (rounded) __________

aribbean Islands: 

Dominican Republic __ _ __ _ ___

Haiti_-_--.______ _ __ _ _________

Jamaica _ _ _ ___________

Total (rounded) __________ __

outh America: 

Brazil _____ __ _______

French Guiana __ ______

Surinam _ __ _____ ___

Total (rounded)- _ __ _____

Jurope:

France _ _

Greece

Hungary _ _

Italy__--.___ _____________________

Poland. __________________________

Rumania _

Spain _ __ _ _

U.S.S.R- ___-__-_-_---.__________-

Total (rounded) ____ __ _

irica:

Cameroun ___ _ _

Congo, - _ _ _

Ghana _ _ ______

Guinea __ _ __ ____

Eeserves 2

49

1

50

QfON

S

60

25

600

690

40

150

250

440

2

70

84

300

24

20

7

3 100

290

900

S

254

1.100

Marginal and 

submarginal 

resources

65

25

85

125

300

50

25

80

40

400

440

200

1,000

70

150 mo

1,520

190

100

Of)

S

320

10

985

L(?)

2.400

Estimated total reserves and other resources of bauxite, in millions 

of tons 1 Continued

Africa Continued

]VTf»T*f»^pf»

Rhodesia _ ________ _______

Total (rounded) _-__ ___

Asia:

Indonesia ____ _______ _

Malay ____ _ ______ ___

Turkey.--.-------- ----------- _ _

Vietnam (North) _ _ __ _

Total (rounded)____--_--------

Oceania: 

Australia _ - ____ ____

Fiji. __________________ __ --_____

Palau_ _______________-_--_---_--_

Total (rounded) for world- _. __

Reserves 2

S

S

1,350

150

. 1

58. 1

25

7

10

5. 6

9.3

.3

270

2, 060. 3

3

2,060

5,760

Marginal and 

submarginal 

resources

25

L

20

2. 4

60

2

3,500

1,000

200

10+

16

40

9.5

28

65

1,370

1,190

.6

. 2

20

. 5

1,210

8,740

1 Most figures are in metric or long tons; however, many estimates used in compilation failed to designate type of tons used.

2 Measured and indicated reserves that in some degree have been inventoried in 

terms of commercial enterprise and could be used under the economic and technologic conditions existing in 1963. 3 Rough estimate based on geologic inferences. Much low-grade bauxite is included that would be classed as marginal and submarginal resources in other 

countries.

REFERENCES

Fischer, E. C., 1953, Sources and reserves, Chap. 3 in Bauxite: 

U.S. Bur. Mines Materials Survey 3, 59 p.

Higbie, K. B., Sullivan, A. C., and Trought, M. E., 1962, 

Bauxite: U.S. Bur. Mines Minerals Yearbook 1961, 

p. 309-331.
Article 42

LEAD REFERENCE SAMPLE FOR ISOTOPIC ABUNDANCE RATIOS

By MARYSE H. DELEVAUX, Washinston, D.C.

Abstract. Relative isotope abundances for lead reference 

sample GS-4, determined by the lead iodide method, have a 

standard deviation of 0.2 percent for the ratios Pb^/Pb20*, 

Pb^/Pb'04, and Pb^/Pb204. This reference sample is available 

for interlaboratory comparison. Comparative data are given 

for reference samples T-1003, NBS 200a, and CIT Shelf Lead.

A number of lead reference samples have been used 

for interlaboratory comparisons of lead isotope abundance ratios. In recent papers (Richards, 1962; Doe, 

1962), data are given for a number of reference samples such as Lament sample 120N, PbSO4, originally 

analyzed by A. O. Nier; Toronto sample T-1003^ and 

National Bureau of Standards sample NBS 200. The 

U.S. Geological Survey has been using a lead reference 

sample, GS-4, which has also been analyzed by a few 

other laboratories (Farquhar and others, 1953) and 

which is available to interested spectrometry laboratories. Analytical data on this sample accumulated 

over a period of 8 years are presented here together 

with some abundance ratios for T-1003, NBS 200a, and 

CIT Shelf Lead. Sample GS-4 is spectrographic lead 

obtained in 1950 from the Johnson Matthey Co., Ltd.

Comparative results for abundance ratios on lead 

reference sample GS-4 are given in table 42.1. The

sample was run periodically with groups of samples, 

and the tabulation is arranged to show .significant 

changes in the mass spectrometer or in technique. The 

mean values shown in the first line of table 42.1 represent 23 separate analyses during the interval MarchOctober 1962. These were made on a 12-inch 60° Niertype mass spectrometer with the measurements as lead 

iodide. The iodide is volatilized in a furnace and ionized by electron bombardment. The Pb208I+1 signal is 

approximately 2 X 10~" amperes measured with a vibrating reed, but without an electron multiplier, and 

with an acceleration potential of 4,500 volts. Changes 

in the instrument made just before this period included 

a new emission regulator and an expanded-scale circuit 

for the recorder which were designed by W. R. Shields 

of the National Bureau of Standards. The data (table 

42.1) clearly indicate that these changes greatly improved the precision. The resolution of the lead iodide 

peaks was essentially doubled and at present, expressed 

as the height of the Pb207 peak divided by the height 

of the well between the Pb207 and Pb208 peaks, is approximately 250 to 300.

Improved electronic circuits and use of the expanded 

scale have reduced the necessary number of scans and

TABLE 42.1. Average abundance values for lead reference sample GS-4

Number of determinations

23-.--_______________

8__ _____.__.__._ ___.

5_ _..____ __ ..

29-_-»---_-_____-____

28_ _-_-______-__-_

12._--__-_______.___.

204

1. 451 ±0. 0030

1. 000 

1. 445 ± .0054

1. 451 ± .0068

1. 463 ± .0144

1. 457 ± .0097

1. 446 ± . 014

206

23. 62 ±0. 017

16. 27 ± . 033 

23. 68 ± . 05i

23. 64 ± . 056

23. 62 ± . 055

23. 60 ± . 070

23. 62 ± . 034

207

22. 52 ±0. 014

15. 52 ± . 035 

22. 44± .054

22. 47 ± .044

22. 50 ± . 044

22. 53 ± . 068

22. 57 ± . 045

208

52. 41 ±0. 023

36. 11 ± . 079 

52. 43 ± .08i

52. 43 ± .080

52. 41 ± . 08o

52. 41 ± .085

52. 36± .059

Remarks

Mar. to Oct. 1962.

Emission regulator and expanded scale, Pbl+i. 

Oct. 1961 to Feb. 1962.

Ion pumps, Pbl+1 . 

June to Aug. 1961.

New focus and electronics, Pbl+1 . 

Nov. 1958 to Sept. 1960. Pbl+1 and Pb+1 . 

Nov. 1956 to Aug. 1958.

12-in. spectrometer, Pbl+1 and Pb+1 . 

1954-56.

6-in. spectrometer, Pb+1 .

NOTE: s (standard deviation) =±

/»-<« 

A/____«_'

ART. 42 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B160-B161. 1963.

B160
DELEVAUX 

TABLE 42.2. Comparative results for reference samples by Pbl+1 and other methods

B161

Sample No.

T~4 003. ---__----

^BS200a_. ______

CITShelf._______

Number of determintions

3

2

3

204

1.457

1. 457 

1.458

1.456 

1. 525

1. 532 

1.429

1.437 

1.440

206

23. 52

23. 42 

23. 47

23.45 

22. 50

22.49 

23.95

23.92 

23. 93

207

22. 61

22. 62

22. 58

22. 61 

22.64

22. 61 

22. 31

22. 28 

22.30

208

52.41

52. 51 

52.49

52. 48 

53. 33

53. 37 

52. 32

52. 37 

52.33

206/204

16. 14

16. 07 

16. 10

16. 11 

14. 75

14. 68 

16.76

16. 64 

16. 62

207/204

15. 52

15. 52 

15.48

15. 53

14. 85

14. 76 

15. 61

15. 50 

15.49

208/204

35.98

36. 03 

35.99

36. 05 

34.97

34. 84 

36. 60

36.43 

36. 34

206/207

1. 0400

1. 0356 

1. 039

1. 037 

. 9938

.9947 

1. 0736

1. 0728 

1. 0732

Remarks l

Pbl+1 ; this paper.

Pb+!; CIW (Doe, 1962). 

PbMe3+1 ; Can 

(Richards, 1962). 

PbMe3+1 ; BC 

(Kollar et al., 1960). 

Pbl+1 ; this paper.

PbMe3+1 ; Can 

(Richards, 1962). 

Pbl+1 ; this paper.

Pb+1 ; CIT (Chow and 

McKinney, 1958). 

Pb+1 ; CIW (Doe, 1962).

: Symbols: BC, University of British Columbia; Can, Canberra, Australia; CIT, California Institute of Technology; CIW. Carnegie, Institution of Washington.

also simplified the calculations for an analysis. With 

the expanded scale a total of 8 mass scans, 4 up and 4 

down, are made, and 4 abundance ratios are read directly from the chart. Results for sample T-1003 measured as Pbl+1 (table 42.2) are in good agreement with 

British Columbia measurements as PbMe3+1 (Kollar, 

and others, 1960). Similarly, results for Caltech Shelf 

Lead as Pbl+1 (table 42.2) are in good agreement with 

determinations as Pb+1 (Chow and McKinney, 1958; 

Doe, 1962).

There are two fundamental differences between the 

lejad iodide technique and the other techniques which 

might be considered advantages. The mass range for 

Pbl+1 is 331-335, considerably greater than in other 

nethods, and, in addition, the ratio of PbI+1/Pb+1 is 

about 3, so that the isotopic composition of the material 

loaded in the mass spectrometer is measured rather than

some breakdown product. Effects due to thermal fractionation and to chemical breakdown are minimized.

REFERENCES

Chow, T. J., and McKinney, C. R., 1958, Mass spectrometric determination of lead in manganese nodules: Anal. Chemistry, v. 30, p. 1499-1503.

Doe, B. R., 1962, Relationship of lead isotopes among granites, 

pegmatites, and sulfide ores near Balmat, New York: Jour. 

Geophys. Research, v. 67, p. 2895-2906.

Farquhar, R. M., Palmer, G. H., and Aitken, K. L., 1953, A comparison of lead isotope analysis techniques: Nature, v. 172, 

p. 860.

Kollar, F., Russell, R. D., and Ulrych, T. J., 1960, Precision intercomparisons of lead isotope ratios, Broken Hill and 

Mount Isa: Nature, v. 187, p. 754-756.

Richards, J. R., 1962, Isotopic composition of Australian leads: 

Jour. Geophys. Research, v. 67, p. 869-884.
Article 43

SYNTHESIS OF LIEBIGITE

By ROBERT MEYROWITZ, DAPHNE R. ROSS, and ALICE D. WEEKS, 

Washington, D.C.

Abstract. A new procedure for the synthesis of liebigite, 

Ca2UO2(CO3)3-10-llH2O, is described and the discovery of what 

is believed to be a new uranyl tricarbonate,

K2Ca3 ( UO2 ) 2 ( CO3 ) e

is noted. The quantitative chemical analyses of two samples 

of synthetic liebigite are given.

The synthesis of liebigite, Ca2UO2 (CO3 ) 3 -10-llH2O, 

has been recorded previously. Hedvall (1925, p. 226) 

reported that Blinkoff 1 synthesized

CaUO2 (CO3 ) 3 -10H2O

by treating a suspension of calcium uranate with carbon dioxide under a pressure of 10 atmospheres. 

Bachelet and others (1952, p. 565, 576) reported that 

calcium uranyl carbonate can be prepared by double 

decomposition and by a "synthetic" method. The double-decomposition procedure consists of mixing solutions of sodium uranyl carbonate and calcium chloride 

or calcium nitrate; no further details are given. In 

the "synthetic" method, carbon dioxide gas is passed 

through a water suspension of uranic oxide hemihydrate and calcium carbonate. After 24 hours a stream 

of air is passed through the mixture, forming additional precipitate. The filtrate of this mixture is concentrated by evaporation in free air, and greenish-yellow 

crystals form. The water content is given as 10 molecules of water, of which 8 molecules are lost at 90° C, 

and all 10 molecules of water are lost at 100°C.

In 1953, M. E. Thompson (oral communication, 1953) 

synthesized liebigite by a procedure similar to that used 

by Axelrod and others (1951) for the synthesis of bayleyite. Successive crops of crystals were obtained by 

allowing the solution containing the proper salts to 

evaporate at room temperature. The initial crops of 

crystals contained potassium, which was subsequently 

shown, as described below, to be in the compound

1 Blinkoff, C., 1906, BeitrSge zur Kenntnls kondensierter Uranylverbindunge: Dissertation, Univ. of Bern, Switzerland.

K2Ca3 (UO2 ) 2 (CO3 ) e 9-10H2O. Crystals were finally 

obtained which were identified as liebigite by their Xray powder-diffraction pattern (Frondel and others, 

1956). The chemical analysis of the synthetic liebigite 

is given in column 1 of the accompanying table.

Various procedures for the synthesis were tried first 

by one of us, using potassium carbonate, uranyl nitrate, 

and calcium nitrate as reactants. Although 4 batches 

of pure liebigite were obtained, the formation of 2 other 

distinct substances was observed in preliminary experiments.

Recently we have synthesized liebigite using sodium 

carbonate instead of potassium carbonate. The procedure is probably similar to the "double decomposition" procedure of Bachelet and others (1952, p. 565) 

for which no details are given. The procedure for the 

synthesis of liebigite using sodium carbonate is as 

follows.

A solution containing 5.02 g UO2 (NO3 ) 2 -6H2O(0.01 

mol UO3 ) in 10 ml of water is added slowly with constant stirring (using a magnetic stirrer) to a solution 

containing 3.18 g anhydrous Na2CO3 (0.03 mol CO2 ) in 

100 ml of water. The pH of the resulting solution 

varies from 8.0 to 8.2. A solution containing 4.72 g 

Ca(NO3 ) 2 -4H2O (0.02 mol CaO) in 10 ml of water is 

added slowly with constant stirring to the uranyl carbonate solution. The pH of this mixture is 6.7 to 6.8. 

Dilute Na2CO3 solution is added dropwise with constant stirring until the pH of the solution is 8.0 to 8.2. 

Within a few hours liebigite will begin to form. The 

beaker is sealed with plastic film so that no evaporation 

takes place, and the solution is allowed to stand for 2 

weeks. The crystals range in size from 1 to 2 mm up to 

5 to 10mm. The crystals are detached from the sides 

and bottom of the beaker and washed by decantation 

with water. Most of the excess water is removed by 

rolling the crystals on absorbent paper. The crystals

ART. 43 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B162-B163. 1963.

B162
MEYROWITZ, ROSS, AND WEEKS B163

are then air-dried and are identified as liebigite by their 

X-ray powder-diffraction pattern.

The procedure described above is superior to the 

K2CO3 procedure in that no precipitate forms when the 

Na2CO3 and UO2 (NO3 ) 2 solutions are mixed; and more 

important, no other crystalline phase forms at pH 8.0, 

as was experienced using K2CO3 . If the concentration 

of Ca(NO3 ) 2 is halved, andersonite, Na^aTJOaCCOs)^ 

6H2O, is formed (Meyrowitz and Boss, 1961).

Liebigite sample SRM-9 (column 2 of table) represents a batch prepared by one of the procedures using 

potassium carbonate.

In many of the preliminary experiments with potassium carbonate, a large amount of well-formed greeni$h-yellow crystals was obtained alone or together 

with liebigite. Its X-ray powder-diffraction pattern 

matched that of the first crop of crystals containing 

potassium which M. E. Thompson prepared as deScribed above. However, it could not be matched with 

any other known X-ray powder-diffraction pattern, including those of the uranyl carbonates. We believe that 

these crystals represent a new uranyl tricarbonate having the formula K2Ca3 (UO2 ) 2 (CO3 ) 6 9-10H2O. Further X-ray study of this material is planned.

Chemical composition of synthetic lieUgite, in weight percent 

[Analyst: Robert Meyrowitz, U.S. Qeol. Survey]

Constituent

CaO--.--, -

UOt 

KiO. 

COi 

HsO . 

Total. -

1

Sample 

AW-176-53 

(1954)

15.4 

40.4 (') 

18.0 

26.8 

100.6

2

Sample 

SRM-9 

(1958)

15.5 

39.1 (2) 

17.7 

27.2 

99.5

3

CasUO2 

(CO3)s- lOHzO 

(theoretical)

15.8 

40.3 

.0 

18.6 

25.3 

100.0

4

Ca2UOn 

(COj)»-llHK>

(theoretical)

15.4 

39.3 

.0 

18.1 

27.2 

100.0

5

Ca2UO2 

(CO3)3-12HaO 

(theoretical)

15.0 

38.3 

.0 

17.7 

29.0 

100.0

i Semiquantitative spectrographic analysis showed that potassium was absent. 

8 Semiquantitative spectrographic analysis of ignited sample showed potassium 

content to be 0.0002.

REFERENCES

Axelrod, J. M. Griinaldi, F. S., Milton, Charles, and Murata, 

K. J., 1951, The uranium minerals from the Hillside mine, 

Yavapai County, Arizona: Am. Mineralogist, v. 36, p. 1-22.

Bachelet, M., Cheylan, E., Dovis, M., and Goulette, J. C., 1952, 

Preparation et proprietes des uranylcarbonates: Soc. Chirn. 

France Bull., 1952, p. 565-569.

Frondel, Clifford, Riska, Daphne, and Frondel, J. W., 1956, Xray powder data for uranium and thorium minerals: U. S. 

Geol. Survey Bull. 1036-G, p. 119.

Hedvall, J. A., 1925, Beitrage zur Kenntnis der komplexen 

Uranylcarbonate: Zeitschr. amorg. v. allgerneine Chemie, v. 

146, p. 225-229.

Meyrowitz, Robert, and Ross, D. R., 1961, The synthesis of 

large crystals of andersonite: Art. 113 in U. S. Geol. Survey Prof. Paper 424-B, p. B266.
Article 44

SEPARATION OF TELLURIUM FROM IRON AND GOLD 

USING TRIBUTYL PHOSPHATE AND ETHER

By C. E. THOMPSON and H. W. LAKIN, Denver, Colo.

Abstract. The separation of traces of tellurium from iron by 

extracting the iron as a thiocyanate complex with a tributyl 

phosphate-ethyl ether mixture is satisfactory for milligram 

amounts of iron. In addition, milligram amounts of gold can 

be separated from traces of tellurium by the method.

A procedure for separating tellurium from iron and 

gold is frequently required in the analysis of geologic 

material for tellurium.

Tributyl phosphate (TBP) has been reported by 

many investigators as a useful solvent in separation 

procedures for many elements in halogen acid and 

thiocyanate systems.

Hikime and others (1959) used the extraction of 

ferric thiocyanate from hydrochloric acid solutions of 

less than O.l^V concentration with TBP to separate 0.35 

to 35 mg of Fe from 250 to 350 Mg of Te (IV). The 

small fraction of the tellurium extracted by the TBP 

was recovered by a O.LV HCL wash of the organic 

phase.

Hikime and others (1959) and Inarida (1958) reported that tellurium (IV) is extracted as the chloride 

complex by TBP at hydrochloric acid concentrations 

above 0.5^V.

Melnich and Freiser (1953) have reported the extraction of ferric thiocyanate with a mixture of 

TBP and carbon tetrachloride from slightly acid 

solutions.

However, no quantitative data have been given for 

the separation of relatively large amounts of iron from 

trace amounts of tellurium. The separation of gold 

and tellurium by this procedure has not been reported. 

The work reported here is an evaluation of this method 

of separation.

TBP forms with dilute hydrochloric acid an emulsion 

which separates slowly. No emulsion is formed, however, if a solution of 150 ml of TBP in 350 ml of ethyl 

ether is used. This TBP-ether solution separates very

.rapidly from the aqueous phase and extracts a larger 

percentage of the ferric thiocyanate from the aqueous 

solution than either of its components or a mixture of 

TBP and carbon tetrachloride.

Five grams of ferric chloride in 30 ml of 0.5^ HC1 

containing an excess of ammonium thiocyanate was 

completely extracted by ttfo 30-ml portions of the 

TBP-ethyl ether solution. However, when a similar 

amount of iron was extracted from a 0.47V HC1 solution containing 10 /ug of Te, the tellurium was also extracted. It was evident from this that either the tellurium was being extracted as a chloride complex or as 

a thiocyanate complex by the TBP-ether solution.

The extraction of the chloride complex of tellurium 

(IV) by the TBP-ethyl ether solvent from various concentrations of hydrochloric acid was tested, using tellurium-129 as a tracer. The activity of 5-ml aliquots 

of the extracts was measured by a scintillation counter 

using a well-type 1.5-inch thallium-activated sodium 

iodide crystal detector. The data in table 44.1 show 

that no tellurium was extracted from the 0.57V acid; all 

of it was extracted from 2N HC1. From these data, 

it is seen that the TPB-ether solutions act the same way 

as has been previously reported for 100-percent TBP 

in the extraction of the tellurium chloride complex.

TABLE 44.1. Relation of extraction of tellurium (IV) by TBPethyl ether solution to hydrochloric acid concentration 

[5 y Te129 in 25 ml of aqueous solution extracted with 25 ml of organic phase]

Normality of HC1

0.5------------------

1.0_-_. --------------

1.5---_~_ --------

2.0------------------

Activity of 5 ml of 

organic extract 

(counts per 5 min)

1,790 ±35 

3, 463 ±41 

5, 290 ±53 

6, 793 ±54

Counts due 

to Tei2», 

background 

subtracted 1

22 

1,695 

3,522 

5,025

Te (IV) 

extracted 

(percent)

0. 0 

33. 6 

69.9 

100.0

i Activity of 5 ml of original aqueous solution with background subtracted = 

5,040±53 counts per 5 min (background =1,768±34 counts per 5 min).

ART. 44 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B164-B165. 1963.

B164
THOMPSON AND LAKIN B165

The influence of ammonium thiocyanate on the extraction of the tellurium (IV) is demonstrated in table

2. Increasing amounts of ammonium thiocyanate 

were added to 40-ml portions of 0.4/V HC1 containing 10 

/ng of Te129. Each of these solutions was extracted with 

£5 ml of TBP-ethyl ether solution. The tellurium 

content of the organic phase was determined by measuring its activity, the TBP-ethyl ether phase was then 

washed with three 20-ml portions of water, and the activity was again measured. It is evident from these 

data that tellurium is extracted from hydrochloric acid 

solutions containing high concentrations of ammonium 

thiocyanate as the thiocyanate complex and is not readily recovered by a water wash of the organic phase. 

Large amounts of iron, therefore, cannot be separated

TABLE 44.2. Influence of ammonium thiocyanate cencentration 

on the extraction of teuurium (IV) from 0-4-N HC1 solution

NH4SCN (grams)

0.5______.___ -_____-__--_

1

1.5. __._..__.. ___________

2.5__________-__________-

5____._ __________________

Activity of 5 ml of the TBPethyl ether counts per 2 min, 

background subtracted >

Extract from 

0.4ZVHC1

71 

120 

223 

479 

3,001

After 

washing

35 

46 

92 

345 

3,242

Te«9 

remaining in 

organic phase 

(percent)

<2 

<2

2.5 

10 

90

i Activity of 5 ml of a solution of 10 7 of Te»w in 25 ml=3,469 counts per 2 min with 

background subtracted (background « 700 counts per 2 min) .

from traces of tellurium because an excess of thiocyanate is necessary for the extraction of iron to be 

quantitative. The procedure is useful, however, for the 

separation of small amounts of iron from tellurium. 

For example, 150 mg of Fe has been quantitatively 

separated from 5 jug of Te by this procedure.

Gold forms a thiocyanate complex which is extracted 

from dilute hydrochloric acid solutions containing relatively low thiocyanate concentrations. Ten milligrams of gold has been successfully separated from 10 

jug of Te in this manner. Thus, this separation procedure can be used to advantage in the preparation of 

gold telluride ores for analysis when the iron content is 

not too high. Also, tellurium has been separated from 

gold used as a collector in the precipitation of tellurium 

from solutions of sulfide minerals and from the iron 

coprecipitated with the tellurium.

REFERENCES

Hikime, Seiichiro, Hoshida, Hitoshi, and Uzumasa, Yasumitsu, 

1959, Separation of Te (IV) and Fe (III) by extraction 

methods using tributyl phosphate: Hokkaido Univ., Sapporo Bunseki Kagaku, v. 8, p. 531-535.

Inarida, Mariko, 1958, Separation of Te from L by extraction 

with 100 percent TBP: Toyko Univ. Nippon Kagaku Zasshi, v. 79, p. 721-723.

Melnich, L. and Freiser, H., 1953, Extraction of metal thiocyanate complexes with butyl phosphate. Iron thiocyanate: 

Anal. Chemistry, v. 25, no. 6, 856-859.
Article 45

DETERMINATION OF TRACES OF BORON IN HALITE 

AND ANHYDRITIC HAUTE ROCKS

By F. S. GRIMALDI and F. O. SIMON, Washinston, D.C.

Abstract. A direct method for the determination of 0.1 

ppm or more of boron in halite and anhydritic halite rocks is 

presented. The sample is dissolved by fuming with sulfuric 

acid in an open quartz beaker at a sufficiently low temperature 

to allow reflux action. No boron is lost by volatilization. The 

determination is completed spectrophotometrically with dianthrimide. Possible losses of boron through volatilization, surface adsorption, and polymerization are assessed. When a 

borate solution is evaporated to dryness in etched quartzware, 

boron is irreversibly adsorbed on the surface, and low recoveries result.

A number of difficulties are connected with the determination of traces of boron. It is common knowledge 

that boron can be lost by volatilization during the 

evaporation of aqueous or acidic solutions of boric 

acid. Wakamatsu (1958) suggests that boron is not 

volatilized from fuming sulfuric acid solutions, but 

that it is converted to a form (possibly a pyroborate) 

that does not react with carminic acid, a generally reliable reagent for the spectrophotometric determination 

of boron. In a series of carefully executed experiments, Feldman (1961) reviewed the situation and concluded that volatilization is slight (<3 percent) from 

solutions of H3BO3 in H2O, HNO3 , H2SO4 , and HC1O4 

at about 75 °C unless evaporation is carried to dryness or 

near dryness; losses from HCI solutions can occur early 

in the evaporation; and that losses from H2SO4 can 

occur at above 228 °C but are negligible with sufficient 

refluxing during the evaporation.

Our laboratory has had to determine very small 

amounts of boron in almost pure halite rock or mixtures 

of halite and anhydrite containing up to 10 percent 

anhydrite. A simple procedure suggested itself conversion of NaCl to Na2SO4 and dissolution of CaSO4 

by heating with H2SO4 , followed by addition of dianthrimide for color development provided the dianthrimide method tolerates appreciable amounts of sodium salts and that no losses of boron occur during the 

preparation of the solution for analysis.

During the course of the investigation and in the 

preparation of standard working curves, a hitherto unobserved difficulty arose. When standard aqueous solutions of sodium tetraborate were evaporated in quartz 

and color was developed with dianthrimide, variable 

absorbances were obtained for solutions containing the 

same concentration of'boron. For some, only about 20 

percent of the expected absorbance was found. This 

difficulty was traced to surface effects and occurred only 

when etched quartzware was used in the experiments. 

This article presents evidence for this conclusion and 

discusses other experiments that form the framework 

for the method here described.

REAGENTS AND APPARATUS

l,l'-dianthrimide stock solution A: Dissolve 0.4 g in 100 ml of 

concentrated H2SO4. Mix, transfer solution to a plastic bottle, 

stopper, and store in refrigerator.

l,l'-dianthrimide working solution B: Dilute 10 ml of stock 

solution to 100 ml with concentrated H2 SC>4. Store as above.

Standard boron solution, 1 ml equals 50 fig of B2O3 : Dissolve 

0.07226 g of anhydrous Na^Or in water and dilute to 1 1 in 

a volumetric flask. Store in a plastic bottle.

Quartz beakers (General Electric Co., Catalog Q7A) : Beaker 

No. QBB 25T, 25-cc capacity, 57-mm height; beaker No. QBR 

50T, 50-cc capacity, 72-mm height. Both sizes were used. 

Beakers referred to as etched were those that had been in 

use in the laboratory for a number of years and had been 

etched through some past contact with solutions containing 

fluoride.

Spectrophotometer, Beckman Model DU: 1-cm corex cells were 

used in the study.

GENERAL PROCEDURE

1. Transfer 1 g of finely pulverized sample to a quartz beaker. 

Add 0.2 ml of water.

2. Add carefully and in small portions a total of 5 ml of 

ItSOt. Addition of acid should be slow enough to prevent too 

vigorous an evolution of HCI gas.

3. When the reaction subsides, grasp the beaker with Blair 

tongs and heat carefully over a burner until no more HCI fumes 

are given off, again being careful that too vigorous an evolution of HCI does not occur during the heating.

ART. 45 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B166-B168. 1963.

B166
GRIMALDI AND SIMON B167

j 4. Bring to fumes of H2 SO* and fume gently from 1 to 3 min. 

pleating should be at a low enough temperature so that most 

of the H2 SO* condenses about one-half and no more than twothirds of the way up from the bottom of the beaker. Cool.

5. Add 5 ml of H2 SOi, 0.2 ml of water, and 10 ml of dilute 

(lianthrimide solution. Mix well.

6. At this time prepare a blank, using 0.2 ml of water, and 

!! standards containing 10 fig of B2O3 . Add 10 ml of H2SO* and 

10 ml of dilute dianthrimide solution, and mix.

7. Place blank, standards, and sample in a controlled temperature oven and heat for 2 hours at between 90° C to 95° C. Cool.

8. Determine the absorbance of each at 625 m/t in 1-cm cells 

against water as reference.

9. Determine the concentration of boron from the straight line 

relating the average absorbance of the standards and blank 

with boron concentration.

RESULTS AND DISCUSSION

Effects of etched and unetched quartz beakers. 

The experiments made and the results obtained are 

listed in the table. In each experiment, quadruplicate 

tests were made using 0.2-ml portions of a standard 

boron solution containing 10 /*g of boron trioxide. The 

experiments were designed to determine losses by volatilization, by surface effects, and by inactivation through 

polymerization. The following conclusions can be 

drawn.

1. With smooth unetched quartz, good results are obtained independent of the treatment given the 

; sample.

2. With etched quartz low results are obtained if either

| H3BO3 or Na2B4O7 solutions are evaporated to dryness. With H3BO3 , boron is lost by volatilization;

I with Na2B4O7 losses are due to irreversible surface 

effects. The various treatments tried (experiments 

7, 8, 9,10, 11, and 13) in an attempt to prevent or 

recover boron loss by adsorption on etched beakers 

failed.

3. As long as H3BO3 or Na2B4O7 solutions are not 

evaporated beyond a point where some liquid remains (about 0.2 ml), no boron is lost either by 

volatilization or through surface effects (experiments 2, 4, 6, 15, 16, 17, 18, 19, 20).

! 4. No boron is lost by volatilization on fuming up to 5 

minutes (maximum tested) with H2SO4 in the 

presence or absence of NaCl, as long as the temperature of heating is low enough to allow reflux 

action (experiments 12,14,17,18,19, and 20). 

5. No evidence was found for inactivation or polymeriI zation of boron. However, if polymerization occurs, the polymer is not converted to active species 

by heating and hydrolyzing with water (experiments 11 and 13).

Experiments with etched and smooth quartsware

Experiment

Quartz

Treatment

before

evaporation

Evaporation

Treatment after

evaporation

Recovery of

boron

H3BO3

1 .

2 do .... Add 5 ml

water.

Add 10 ml H2SO4 and

10 ml dianthrimide.. 

Heat 2 hrs at 93°C. 

 -do 

3 _ ...

4 ...

5 ...

6. 

7. .

8 .

9.. 

10 

11.. ...

12 

13 

14 

15 

16 

17 

18 

19 

20 

Unetched -

...do ...

... do ....

...do _ 

 do ... .

... do... ....

 do 

... do 

Unetched.

Unetched.

Unetched.

Unetched . 

Etched. -.

Unetched .

None ..

Add 5 ml

water.

Add 5 ml

water. 

None _ 

... do 

...doAdd 50

mg 

NaaCOa 

None _ .

... do ... 

... do 

... do 

... do... ...

 do 

... do 

... do ... 

... do -

... do ...

DrynessDryness..

... do ...

... do 

... do 

... do.... -

... do 

... do... ...

 do -

oration. 

 do... 

... do 

 do 

... do 

 do -

As in experiment 1 ....

..... do ..... ... 

 .do ... .... .... ...

Add 0.2 ml water, 

and warm. Then 

as in experiment 1. 

Add5mlH2SO4 and

fume 1 to 5 min. 

Add 5 ml HzSO4 

and 10 ml dianthrimide. Heat 2 hr 

at 93°C. 

AddlgNaCl. Then

as in experiment 8. 

As In experiment 1 ..

AddlgNaCl and 

5 ml H2SO4. 

Fume 1 to 5 min. 

Cool. Add 0.75 ml 

H2O. Digest % hr 

on steam bath. 

Add 5 ml H2SO4 

and 10 ml dianthrimide and heat at 

93°C for 2 hr. 

As in experiment. 11. 

As in experiment 11,

except no NaCl 

added. 

 .do... ... - ... 

As in experiment 1 ....

.....do...... ........

Add 5 ml H2SO4 ,

fume 1 to 5 min. 

Add 5 ml H2SO4 

and 10 ml dianthrimide. Heat for 2 

hr at 93° C. 

 do ... ........ 

As in experiment 9 

__ do __ ._ 

Do.

roughly dependent on 

degree of 

etching.

20-50 percent. 

Do.

Do.

20-60 percent.

15-60 percent.

Complete. 

15-60 percent.

Complete. 

Do.

Do. 

Do.

Do. 

Do.

Do.

Effect of sodium and calcium salts. Application of 

the procedure to a set of blanks and to one of 10 /tg of 

B2O3 , with members of each set containing from 0.25 to 

1.25 g of NaCl in increments of 0.25 g, showed that the 

procedure tolerates at least 1.25 g of NaCl. The average absorbance of the blanks against water was 0.164 

with a standard deviation of 0.006; the average absorbance of the samples was 0.404 with a standard deviation of 0.007. These values are not significantly different from those obtained on similar tests of blanks 

and standards without added NaCl, and either with or 

without preliminary fuming with H2SO4. The pro-
B168 ANALYTICAL TECHNIQUES

cedure was also found to tolerate at least 0.1 g CaSO4 

without interference. Other elements were not tested 

both because the dianthrimide method is almost specific 

for boron and because such elements that might interfere are not present in evaporites in sufficient concentrations to cause disturbance.

Boron content of halite and anhydritic-halite rocks. 

A sample of almost pure halite rock from Tatum salt 

dome, Lamar County, Miss., was found to contain 0.16

ppm of boron; 3 samples of anhydritic halite from the 

same locality contained 0.21, 0.29, and 0.34 ppm of 

boron.

REFERENCES

Feldman, 0., 1961, Evaporation of boron from acid solutions 

and residues: Anal. Chemistry, v. 33, p. 1916-1920.

Wakamatsu, S., 1958, Loss of boron by volatilization in the determination of boron: Japan Analyst, v. 7, p. 309-313.
Article 46 

DETERMINING DENSITY AND POROSITY OF TUFF CONTAINING ZEOLITES

By D. D. DICKEY and E. F. MONK, Denver, Colo.

Work done in cooperation with the U.S. Atomic Energy Commission

Abstract. Standard methods of density and porosity determination do not give reliable results for certain zeolitized 

Southern Nevada tuffs. To obtain the complete saturation 

needed for accurate determinations of density and porosity, a 

powdered sample must soak 3 or 4 days in distilled water and 

cores 1 inch long and 1 inch in diameter must soak more than 

20 days.

During the course of determining density and porosity of some volcanic rocks, it was noted that a group 

Of natural-state samples gave values of more than 100- 

percent saturation when run by standard methods. In 

Order to find what modifications were required to obtain true saturation, three specimens of zeolitized tuffs 

were selected for experimentation. These tuffs, which 

are part of the Survey Butte Member of the Piapi Canyon Formation of early Pliocene or younger age, were 

collected from underground workings at Rainier Mesa 

in southern Nye County, Nev. X-ray analysis by A. O. 

Shepard showed the major constituent of the tuff to be 

blinoptilolite, a zeolite. Other constituents are cristobolite, quartz, feldspar, and minor amounts of montfaiorillonite.

Standard procedures for determining the dry bulk 

density can be used for these rocks. The method congists of cutting a specimen of regular shaje from the 

rock and calculating its volume from the dimensions 

'(the mercury-displacement method can also be used to 

obtain volume). The oven-dried weight divided by this 

volume yields the dry bulk density.

In an effort to obtain the grain density and porosity, 

the samples were treated by the water-saturation method that has been generally successful for rocks with interconnected pores. This method consists of drying a 

rock sample at 105°C until a constant weight is attained. 

The sample is then put in a dessicator, pumped to a 

vacuum of 1 millimeter of mercury and covered with 

de-aerated distilled water and left under a partial

vacuum for 24 hours. The saturated sample is weighed 

in air, and weighed while suspended in water. From 

these measurements the porosity, grain density, and dry 

bulk density can be calculated (Meinzer, 1923, p. 12). 

The porosity thus determined for the tuff samples appeared to be too low. The samples were again placed 

under water and the saturated and suspended weights 

determined 5 days later, and again at intervals up to 31 

days. It was found that about 21 days was required 

for core samples 1 inch long and 1 inch in diameter to 

reach equilibrium (fig. 46.1).

Commonly the grain density is determined by the 

powder method described by the American Petroleum 

Institute (1960). Essentially this method consists of 

measuring the weight and volume of a powdered sam15 20 

TIME, IN DAYS

FIGURE 46.1. Grain density determinations of three core samples of tuff by the water-saturation method (temperature 

corrected).

ART. 46 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B169-B170. 1963.

B169
B170 ANALYTICAL TECHNIQUES

pie of rock suspended in a liquid. The porosity in percent is then equal to G-D/G X100, where G is the grain 

density and D is the dry bulk density. Trouble was 

immediately encountered using this method. It is usually necessary only to add the powdered rock sample to 

the liquid, wet the grains of powder by swirling the 

bottle and then place it under a vacuum for a few minutes to remove any air introduced with the sample. 

Several tests of the zeolitized tuff showed that air bubbles continued to escape from the powdered rock even 

after the sample had been kept for a few hours in a 

vacuum. The air bubbles did not entirely disappear 

until the sample had been soaked for 2 days under a 

partial vacuum. A grain-density determination was 

made at this time. The samples were kept under water

and additional grain-density determinations were made 

periodically for 14 days. These results indicated that 

3 to 4 days of soaking in water was necessary to insure 

complete penetration of the zeolites by water. Experiments using this method but with fluids other than water yielded data that were peculiar to these particular 

fluids and that differed appreciably from the porosity 

and grain densities obtained for water-saturated samples.

REFERENCES

American Petroleum Institute, 1960, Recommended practice for 

core-analysis procedure: Am. Petroleum Institute RP40, 

sec. 3.59.

Meinzer, O. E., 1923, The occurrence of ground water in the 

United States: U.S. Geol. Survey Water-Supply Paper 489.
Article 47

LAND SUBSIDENCE IN THE ARVIN-MARICOPA AREA, 

SAN JOAQUIN VALLEY, CALIFORNIA

By BEN E. LOFGREN, Sacramento, Calif.

Work done in cooperation with the 

California Department of Water Resources

Abstract. Land subsidence is affecting 425 square miles of 

farmland in the Arvin-Maricopa area, at rates to 0.5 foot per 

year. The subsidence is related directly to declining groundwater levels and results from the compaction of unconsolidated 

alluvial and lacustrine deposits to depths of at least 1,000 feet.

Land subsidence is continuing in three principal 

areas of the San Joaquin Valley, Calif., as ground-water 

levels continue to decline. About 3,000 square miles of 

irrigable land is affected by the subsidence, which locally exceeds 1 foot per year. As shown in figure 47.1 

the Arvin-Maricopa subsidence area is at the extreme 

southern end of the San Joaquin Valley and is the 

smallest (425 square miles) of the 3 subsiding areas.

As in the Los Banos-Kettleman City and the TulareWasco subsidence areas (Poland and Davis, 1956), 

most of the subsidence occurring in the Arvin-Maricopa area is related directly to the decline of groundwater levels. Accelerated pumping of ground water 

for irrigation, especially since 1940, has caused water 

levels throughout the area to fall at rates ranging from 

5 to 15 feet per year. With the reduction of artesian 

head in the saturated unconsolidated alluvial and lacustrine deposits, more of the stress of the overburden 

load is transferred to the grain-to-grain contacts within 

those deposits. This transference of stress causes compaction of the deposits and the resultant subsidence of 

the land surface. Although near-surface subsidence related to the wetting of low-density deposits is known 

to occur in the southern part of the area and subsidence 

due to tectonic readjustment is suspected along the 

northern flank of Wheeler Ridge, only the subsidence 

attributable to compaction caused by water-level decline is considered in this article.

Subsidence was first recognized in the Arvin-Maricopa area in 1953 following the destructive 1952 ArvinTehachapi earthquake. Bench marks that had re124° 120° 116°

42'

38'

34'

*'. \

T--N O

V^V4- \

lettleman

100 200 MILES

FIGTJBE 47.1. Areas of land subsidence in the San Joaquin 

Valley, Calif.

mained at relatively constant elevations during earlier 

surveys of 1926, 1935-39, 1942, and 1947 gave evidence 

of a broad area of subsidence in the valley when resurveyed in 1953 by the U.S. Coast and Geodetic Survey. 

Because the Tehachapi Mountain area south of the 

White Wolf fault was upwarped during the 1952 period 

of tectonic activity, Whitten (1955) ascribed the downwarp in the valley alluvium north of the fault to this 

same readjustment. However, an analysis of subsequent bench-mark releveling in 1957, 1959, and 1962 

in relation to the earlier leveling indicates that subsidART. 47 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B171-B175. 1963.

680231 O 63 12 B171
B172 ENGINEERING HYDROLOGY

ence in the Arvin-Maricopa area was not due to the 

1952 earthquake but, instead, was related to the waterlevel decline that began about 1940 and has continued 

at an accelerating rate to the present time.

Subsidence of the land surface at a point 8 miles 

west of Mettler is shown in figure 47.2 by a change in 

the elevation of bench mark LS-2 between 1942 and 

1947 and by subsequent changes in the elevation of 

bench mark J-824, which replaced LS-2 in 1947. (For 

the locations, see fig. 47.6.) The subsidence at this location began during the period 1942-47 and has been 

continuous since 1947. Also shown in figure 47.2 is a 

hydrograph for well 11N/21W-14D2, which is about 1 

mile southeast of the bench-mark location. The water 

level in the well declined about 270 feet between 1949 

and 1962, and probably had been declining for several 

years before 1949. Since 1957 the subsidence at bench 

mark J-824 has been at a rate of about 1 foot for each 

120 feet of water-level decline in well 11N/21W-14D2. 

This ratio is considerably less than the ratios (1:10 to 

1:30) determined for the other two subsiding areas in 

the San Joaquin Valley (Poland and Davis, 1956).

£ 1.5

Bench mark J-824

ark LS-2 - '

Well 11N/21W-14D2

300

450

1940 1950 1960

FIGURE 47.2. Land-surface subsidence and water-level 

decline, 8 miles west of Mettler.

The relation of subsidence to water-level decline 2 

miles west of Mettler is shown in figure 47.3. Subsidence 

at bench mark A-303 was first recorded in 1942 and has 

continued at an accelerating rate to the present time in 

direct response to the rapid decline of ground-water 

levels. Water levels at this location fell from less than 

20 feet below the land surface in 1929 (California Division of Water Resources, 1931) to about 130 feet below 

the land surface in the summer of 1946, a rate of about 

7 feet per year. During the period 1946-62, the water 

level in well 11N/20W-9A1 fell from about 130 feet to

Q 2.5

3.5

4.0

Bench mark A-303

Well 11N/20W-9A1

1940 1950 1960

UJ ui 

u.

Z

200 ££

ui

$ 

250 g

t ui

300 Q 

350 

400 

450

FIGURE 47.3. Land-surface subsidence and water-level 

decline, 2 miles west of Mettler.

about 415 feet, an average rate of 18 feet per year. 

During the period 1947-53, the rate of subsidence at 

bench mark A-303 was 0.16 foot per year and during 

the period 1959-62 was 0.30 foot per year. Thus, the 

ratio of subsidence to water-level decline increased between these two periods from 1:112 to 1: 60, respectively.

The series of profiles in figure 47.4 show the amount 

of land subsidence that occurred during the period 

1926-62 along U.S. Highway 99 from Grapevine to 

Bakersfield. The profiles are based on periodic releveling of bench marks by the U.S. Coast and Geodetic Survey. For bench marks south of the White Wolf fault 

which were elevated during the 1952 earthquake, the 

1947-53 differences in elevation were adjusted by the 

author to eliminate the effect of this tectonic uplift. 

Thus, all the subsidence shown in the profiles and subsequent maps is due to the decline of ground-water 

levels. As shown, most of the subsidence along the 

line of profiles, locally amounting to more than 6 feet, 

has occurred since 1942 when heavy pumping began in 

the area.

Subsidence profiles along the Maricopa road from 

Mettler to Maricopa are shown in figure 47.5 for the 

period 1935-39 to 1962. Since 1942, the maximum 

subsidence along this line has been 5 feet. However, 

since 1957, little subsidence has occurred along the 

western third of the line.
LOFGREN B173

GRAPEVINE

| 1926 base

BAKERSHELD

FIGURE 47.4. Profiles of land subsidence along U.S. Highway 

99, Grapevine to Bakersfield, Calif., 1926-62.

FIGUEE 47.5. Profiles of land subsidence along the Maricopa 

road, Mettler to Maricopa, Calif., 1935-39 to 1962.

Even though the U.S. Coast and Geodetic Survey ran 

precise levels in the Arvin-Maricopa area as early as 

}926, it was not until 1957, with the financial cooperaijion of the California Department of Water Eesources, 

tjhat a network of bench marks was established throughput the area. Eeleveling of this network in 1959 gave, 

lor the first time, sufficient area! control to map the extent of subsidence. Based on differences in the elevation of bench marks during that 2-year period, the lines 

of equal subsidence shown in figure 47.6 (p. B174) indicate that nearly 200 square miles (within the 0.2-foot 

line) was subsiding, at a maximum rate of about 0.5 

1 oot per year.

The possibility of tectonic movement in the ArvinMaricopa area during the period 1957-59 is suggested 

tyy the close spacing of the lines of equal subsidence 

4orth of Wheeler Eidge and by the west-northwest

alinement of the deepest part of the subsidence trough. 

However, as the lines of equal subsidence pass smoothly 

over the trace of the White Wolf fault (where the last 

measurable movement occurred in 1952), there is no direct evidence of tectonic activity during the 1957-59 

period.

Land subsidence during the 3-year period 1959-62, 

shown in figure 47.7 (p. B175) is based on tentatively 

adjusted 1962 leveling data of the U.S. Coast and Geodetic Survey. During this period, subsidence continued 

at a maximum rate of 0.5 foot per year and the area of 

subsidence increased to 425 square miles (within the 0.2- 

foot line). The northward expansion of the subsidence 

area suggests that the effect of concentrated pumping 

north and west of Mettler is moving northward beneath 

the Kern and Buena Vista Lake Beds (dry).

One of the compaction recorders of a type described 

by Lofgren (1961) is located 8 miles west of Mettler 

and 2 miles north of well 11N/21W-14D2. Measurements of compaction at this site since June 1960 indicate that about three-fourths of the subsidence can be 

accounted for by compaction in the unconsolidated 

water-bearing deposits above a depth of 810 feet; therefore, one-fourth of the compaction occurs at depths 

greater than 810 feet. Inasmuch as most of the ground 

water pumped in this area is from wells 600 to 1,000 

feet deep, the compaction probably extends downward 

to a depth of at least 1,000 feet.

REFERENCES

California Division of Water Resources, 1981, San Joaquin 

River Basin: California Dept. Public Works, 656 p.

Lofgren, B. E., 1961, Measurement of compaction of aquifer 

systems in areas of land subsidence: Art 24 in U.S. Geol. 

Survey Prof. Paper 424-B, p. B49-B52.

Poland, J. F., and Davis, G. H., 1956, Subsidence of the land 

surface in the Tulare-Wasco (Delano) and Los BanosKettleman City area, San Joaquin Valley, California: Am. 

Geophys. Union Trans., v. 37, no. 3, p. 287-296.

Whitten, C. A., 1955, Measurements of earth movements in 

California, pt. 1 of Oakeshott, G. B., ed., Earthquakes in 

Kern County, California, during 1952: California Dept. Nat. 

Resources, Div. Mines Bull. 171, p. 75-80.
B174 ENGINEERING HYDROLOGY

0246 8 10 MILES

EXPLANATION

Line of equal subsidence, in feet Trace of White Wolf fault Line of bench-mark control by the

U. S. Coast and Geodetic Survey

FIGUEE 47.6. Land subsidence in the Arvin-Maricopa area of the San Joaquin Valley, Calif., 1957-69. Shaded areas

are deformed rocks bordering the valley.
LOFGREN B175

R. 23 W

0 2 4 6 8 10 MILES

FIGURE 47.7 Land subsidence in the Arvin-Maricopa area of the San Joaquin Valley, Calif., 1959-62. Heavy contours are lines of equal subsidence, in feet. Shaded areas are deformed rocks bordering the valley.
Article 48

WATERPOWER INVESTIGATIONS OF LAKES IN ALASKA

By ARTHUR JOHNSON, Washington, D.C.

Abstract. Many lakes in Alaska situated at high altitudes 

and drained by streams of steep gradient offer natural storage 

needed for hydroelectric-power development. Topographic 

mapping of the surrounding slopes and the underwater relief 

provides essential information for project planning.

The streams of Alaska are characterized by pronounced seasonal variation in flow, and storage is necessary to equalize the flow for hydroelectric-power development. Lakes, in general, provide favorable opportunities for the development of storage and, when at 

high elevations and drained by streams with steep 

gradients, provide favorable powersites.

The Geological Survey, in the investigation of prospective waterpower sites in Alaska, has mapped 27

lakes. This mapping included soundings on most- of 

the lakes for control of underwater contours, which are 

used to determine and compare the storage capacity of 

a lake below normal elevation with that obtainable by 

building a dam at the lake outlet and raising the level 

above normal elevation. Additional soundings were 

made to determine approximate maximum depths.

The soundings were made from boats with wire line 

and weight, or by electronic depth sounder. One lake, 

Eklutna, was surveyed in the winter of 1947 by making 

soundings through carefully located holes in the ice. 

The lakes for which maps had been published as of 

December 31, 1962, are listed in the accompanying 

table.

Dimensions of lakes in Alaska mapped for waterpower investigations

Lake

Eklutna.. _ _. _

Chakachamna _ - _

Ptarmigan. __ ___ __

Cooper. .

Crescent- . _ __

Lost. _ _ __

Nellie Juan __ __

Bradley __ _ __ _

Seldovia

Silver _ _ _

Chilkoot___--___ _______

Crater

Long _ ___ _____

Scenery __ __ ____

Swan___ _ _ __

Ruth___________________

Virginia. _ _ _

Anita _ _

Kunk__________-______---

Olive- _ ________________

Baranof-__

Takatz

Kasnyku____ _ _ _

Location

20 miles north of Seward _ _ _ _

25 miles north of Seward__

_-__do_ __ _ ._____._ __ -___

. do

_ _do_ ___--_---____._-______ . __ _

-__-_do_________ ________________________________

Island.

Island. 

20 miles north of Port Alexander, Baranof IslandLength 

(miles)

7.0

15. 0

4. 5

5. 7

6. 0

1. 7

4.8

3.5

1.0

3. 0

4. 4

2. 3

5.0

1. 4

5. 6

2. 8

1.7

. 6

2. 0

1. 0

1. 3

. 8

3.0

1. 8

1.8

1. 8

3. 2

Maximum 

width 

(miles)

1 0

2 0

5

1 0

5

4

6

1 3

2

7

1 0

5

7

4

6

4

6

2

7

3

5

4

5

4

5

. 4

.7

Area 

(acres)

3,256

16, 780

636

2,050

1,370

370

1, 130

1, 570

127

978

1, 720

503

1,320

286

1,250

544

570

55

636

142

238

158

727

410

403

301

969

Altitude 

(feet)

868

1, 142

755

1, 168

1,454

1,920

1, 189

1,090

427

306

29

1,022

814

1,865

544

957

1,514

1,353

104

1,370

275

438

145

211

905

595

374

Maximum 

observed 

depth (feet)

206

380

204

400 +

291

130

231

250

0)

278

282

414

470

319

459

235

300 +

91

179

210

150

152

295

258

474

373

877

' No soundings made.

B176

ART. 48 IN U.S. GBOL,. SURVEY PROF. PAPER 475-B, PAGES B176-B178. 1963.
JOHNSON B177

CRATER LAKE 

Jteservoir level with dam, max elev 1130Jt

1000

a BOO L_

z 

- 600

g

£ 400

u

UJ

200

Natural lake level, elev 1022 ft \DAM

.~-JJ!!!*i!S£(2&J^^

x Speel 

Power house"^-, Arm MEAN

3 2 1 

DISTANCE, IN MILES

FIGURE 48.1. Profile of Crater Creek and section along approximate centerline of Crater Lake.

Most of the lakes presumably are of glacial origin, 

either owing to morainal damming or to overdeepening 

by glacial scouring; however, specific detailed geologic 

studies of the lake areas have not been made.

Several of the lakes discharge over bedrock lips, 

which is suggestive of an origin by glacial scouring. 

Brew and others (Art. 28) show that many of the linear 

topographic features of southern Baranof Island, including deep lake-filled valleys, are controlled by joints 

;and faults in the metamorphic and igneous rocks. Glacial erosion was most effective along these planar zones 

of weakness. The bottoms of some of the lakes are well 

below present sea level, as shown by the data in the accompanying table. For example, at Deer Lake on 

Baranof Island, the deepest observed point was 503 feet 

below sea level.

The application of a map showing contours both 

above and below lake level in the evaluation of potential 

waterpower is illustrated in the following example.

Crater Lake, near Speel Arm and about 40 miles 

southeast of Juneau, is 2 miles long, has a maximum 

jwidth of half a mile, and has a surface elevation of 

1,022 feet above sea level. The outflow from the lake 

is over a bedrock lip into Crater Creek, which discharges into tidewater at Speel Arm only 1 mile from 

the lake (fig. 48.1). Underwater contouring was completed to the 800-foot contour (222-foot depth) and the 

[700-foot contour was completed from about the center 

of the lake to the outlet. The maximum depth observed, 

along a line of soundings taken at approximately quariter-mile intervals along the centerline, was 414 feet, 

inear the midpoint of the lake. A depth of more than 

;200 feet was reached in a distance of less than 300 feet 

from the lake outlet. The underwater slopes at the 

head of the lake were not quite as steep as those along 

the sides; however, a depth of 222 feet was reached 

jwithin 700 feet of the head of the lake. Figure 48.2

800 750 700 650

AREA, IN ACRES 

600 550 500 450 400 350 300

1200

1120

O

JZ 1040

960

880

800

Reservoir capacity 71,000X 

acre-ft, elev 1130 ft

Capacity^

Lake surface, elev 1022 ft

lean reservoir level with 

drawdown, elev 944ft

-Reservoir capacity 71,000 

acre-ft, elev 853 ft

60 40 20 0 20 40 60 80 100

Below lake level I Above lake level

CAPACITY, IN THOUSANDS OF ACRE-FEET

FIGURE 48.2. Relation between area and capacity, Crater Lake.

shows the area and capacity of Crater Lake both above 

and below its normal surface elevation.

An analysis of the streamflow data indicates that 

complete utilization over a period of years is not practicable, but that a storage capacity of 71,000 acre-feet 

would provide a regulated flow equivalent to 90 percent of the mean flow (Johnson, 1960). This capacity 

could be obtained by building a dam at the lake outlet 

to raise the lake surface 108 feet or by tapping the lake 

by a tunnel and drawing it down 169 feet below its 

normal elevation. The elevation of the reservoir at 

median capacity would be 1,081 feet with a dam or 944 

feet with drawdown, a difference of 137 feet. The 

powerhouse would be at sea level, and the foregoing 

figures represent the gross mean head available for the 

two methods of development. A development by building a dam to raise the lake level would therefore have 

a head, and consequent power capacity, 15 percent 

greater than a development based on drawdown. This
B178 ENGINEERING HYDROLOGY

information enables project planners to determine 

whether or not the 15-percent greater capacity would 

justify the greater cost of building a dam as compared 

with the cost of a development based on drawdown. 

Development by a dam would provide a fuller use of 

a potential resource. Actual development must, however, be based on the most favorable economic alternative.

At Crater Lake, where the lake surface is at an elevation of over 1,000 feet, the difference in power between the 2 methods of development was shown to be 

15 percent. However, for lakes at lower elevations, and 

with consequently lesser potential heads for power, the 

percentage difference between the two methods of development would be greater. This difference would increase as lake elevations decreased, and a point would

be reached where the difference would be so great as 

to preclude consideration of development by drawdown. 

The Eklutna Lake power project, 30 miles northeast 

of Anchorage, with a rated capacity of 30,000 kilowatts, 

is an example of a development by drawing a lake below 

its natural level. The normal elevation of this lake is 

868 feet. A tunnel 4^/2 miles long, with an intake elevation at 800 feet, 68 feet below normal lake level, conveys the water to a powerhouse along the AnchoragePalmer highway. The turbines are at an elevation of 

25 feet, giving a gross head of 843 feet. The tailrace 

from the powerhouse discharges into Knik Arm.

REFERENCE

Johnson, F. A., 1961, Waterpower resources near Petersburg 

and Juneau, southeastern Alaska; U.S. Geol. Survey WaterSupply Paper 1529, p. 66.
Article 49

GROUND WATER IN THE NAHUD OUTLIER OF THE NUBIAN SERIES, 

KORDOFAN PROVINCE, SUDAN

By HARRY G. RODIS1 and WILSON ISKANDER,2 Khartoum, Sudan 

Work done in cooperation with the Geological Survey Department of Sudan

Abstract. Wells tapping aquifers in the Nahud outlier of the 

Nubian yield water of good quality for domestic, agricultural, 

and livestock needs. Evaluation of hydrologie data indicates 

that withdrawals at the present rate could be continued almost 

indefinitely without significant depletion of the water supply. 

However, additional development should be preceded by detailed 

investigation.

Rocks of the Nubian Series occupy a large isolated 

basin in the west-central part of Sudan (fig. 49.1). 

These rocks, referred to as the Nahud outlier of the 

Nubian Series, underlie a developing agricultural and 

livestock-raising area. Because the climate is semiarid, 

most of the water to meet local needs is obtained from 

wells. The small quantity of surface water collected in 

hafirs (excavated ponds) and hollow tabeldi trees during the rainy season is not sufficient to sustain local needs 

during the remainder of the year.

An investigation was begun in 1961 to determine the 

ground-water potential of the Nahud outlier. This 

article, which presents the preliminary results of the 

investigation, is based on an evaluation of geologic and 

ground-water data collected by earlier workers and on 

new data collected by the authors during a recent field 

reconnaissance. The study is being made by the Geological Survey Department of Sudan and is under the 

direct supervision of Mahmoud Ahmed Abdulla, Director of the Geological Survey Department.

The Nahud outlier underlies an area of about 4,000 

square miles. It extends eastward from the town of 

Nahud for about 100 miles and is about 65 miles wide 

from north to south. The surface, which has a topographic relief of more than 250 feet, is formed largely 

by gently to moderately undulating fixed sand dunes 

("qoz" sand) that are covered in part by tall rank

1 On deputation from the U.S. Geological Survey to the Geological 

Survey Department of Sudan under the auspices of the U.S. Agency 

for International Development.

8 Geological Survey Department of Sudan.

10 0 10 20 30 KILOMETERS

EXPLANATION

Granite and gneiss of 

Precambrian age

Contact

Margin of the principal 

zone of saturation

.Ca

Water yard containing one or

more producing wells

Ca indicates water sample collected

for chemical analysis

o 

Dry well

Abandoned well

FIGURE 49.1. Geologic sketch map of the Nahud outlier of the 

Nubian Series showing the location of wells drilled since 1920 

and -the areal extent of the principal zone of saturation.

grasses, shrubs, and sparsely spaced trees. There is no 

surface runoff into or out of the outlier; all surface

ART. 49 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B179-B181. 1963.

B179
B180 GROUND WATER

runoff within the outlier drains into numerous clayey 

swales. Except for possible subsurface outflow, the 

Nahud outlier is a complete hydrologic unit.

The climate is typical of the low-latitude steppe. 

Precipitation, "which averages 15 inches annually, falls 

mostly in the summer (May to October) during storms 

of high intensity and short duration. The mean daily 

air temperature ranges from 71°F to 113°F; the average annual air temperature is 80°F. The mean daily 

relative humidity ranges from a low of 21 percent during the winter to 75 percent in the summer.

Reports on previous ground-water investigations in 

west-central Sudan are regional and are listed at the 

end of this article. These reports were especially useful 

in constructing the geologic sketch map of the Nahud 

outlier (fig. 49.1). Prior to this investigation no attempt had been made to evaluate the ground-water potential of this area.

The Nahud outlier, an erosional remnant of the formerly more extensive Nubian Series of Mesozoic (Cretaceous ?) age, is both underlain and surrounded by granite and gneiss of Precambrian age; there are no known 

outcrops of the basement rocks within the margin of the 

outlier (fig. 49.1). Although they locally yield small 

quantities of water from cracks and fissures, the Precambrian rocks generally are not considered to be a 

source of water supply.

The rocks of the Nahud outlier are composed mainly 

of loosely consolidated, thin- to thick-bedded sandstone, 

siltstone, and mudstone; commonly the sandstone is 

clayey, silty, or gravelly. As these strata were deposited in a continental or marginal marine environment, 

changes in facies and bedding are common. Dipping 

very gently toward the center of the basin, the rocks 

of the Nahud outlier thicken from a featheredge at the 

margin of the basin to about 500 feet in its central part. 

Where saturated, the more permeable rocks yield water freely to wells.

In most places the Nahud outlier is veneered by late 

or post-Mesozoic laterite and by Quaternary clay and 

fixed sand dunes. These surficial deposits, which have 

an average thickness of about 50 feet, are of no importance as a source of water supply.

The principal zone of saturation within the Nahud 

outlier underlies an area of about 2,000 square miles 

(fig. 49.1) and has a maximum thickness of about 100 

feet. The layers of more permeable sandstone within 

this zone are the principal aquifers, and the water they 

contain is under slight artesian pressure. Because the 

more permeable sandstones grade both laterally and vertically into less permeable rock, individual aquifers generally are of limited areal and stratigraphic extent. 

Studies of well logs and water levels indicate that the

principal zone of saturation has a specific yield of 5 percent^and a usable storage capacity of about 64,000 acrefeet per foot of saturated material.

The altitude of the piezometric surface of the ground 

water in the basin is about 1,675 feet above sea level, or 

between 100 and 250 feet above the altitude of the 

piezometric surface in other sedimentary basins in the 

region. Because the Precambrian rocks in the area are 

known to contain cracks and fissures down to depths of 

at least 500 feet below land surface, it is possible that a 

small amount of water flows out of the outlier into the 

surrounding basement rocks.

Thirty-four successful wells have been drilled since 

1920 at 15 locations, known as "water yards," within the 

area underlain by the principal zone of saturation (fig. 

49.1). Each water yard has 2 or more closely spaced 

wells that are generally less than 300 feet apart. Each 

well is equipped with a diegel-powered piston pump, 

and in most water'yards the pumps are operated at 

nearly full capacity. The wells range in depth from 

about 300 to 500 feet, are 6 to 12 inches in diameter, 

and are finished with slotted-pipe screens. Reportedly, 

the wells are pumped an average of 10 hours per day 

throughout the year and individual well yields average 

1,000 gallons per hour. Eight dry wells were drilled 

outside the area underlain by the principal zone of saturation (fig. 49.1).

The depth to water, as measured in wells about once 

a year since 1932, ranges from about 200 to 400 feet 

below land surface. Although rough or approximate, 

the measurements suggest no significant net decline or 

rise in water levels during this period.

Withdrawals from the principal zone of saturation 

are increasing steadily; in 1961 the total draft is estimated to have been 125 million gallons, or 383 acre-feet. 

Even if no recharge were to occur, withdrawals at this 

rate would lower the general water level by only 1 foot 

in 167 years. Obviously, an increase in the withdrawal 

rate would lower the water level the same distance in a 

proportionately shorter time. The estimate of waterlevel decline in response to withdrawals is conservative 

because the zone of saturation undoubtedly receives 

some recharge from precipitation.

Sufficient hydrologic data for a quantitative evaluation of recharge are not available. As the rains generally are of short duration, it is believed that much of 

the precipitation evaporates soon after it falls and that 

most of the water entering the soil is transpired by 

vegetation. It also is believed that little, if any, of the 

water that accumulates in the swales infiltrates the tight 

clayey soil underlying them. Only where large quantities of surface runoff are stored in hafirs for long 

periods is there much opportunity for infiltration to
RODIS AND ISKANDER B181

tjhe zone of saturation. However, because the Nahud 

Outlier has such a large surface area, a very low average 

annual rate of recharge would result in a significant 

increment to storage. For example, if no allowance is 

made for natural discharge from the zone of saturation, 

a recharge rate of as little as 0.004 inch per year 

throughout the area underlain by the principal zone of 

saturation would more than sustain the present rate of 

withdrawal. Discharge of ground water by evapotranspiration is nil because of the great depth to the 

zone of saturation. Therefore, subsurface outflow is 

the only natural discharge that must be taken into 

consideration in evaluating the ground-water potential 

of the Nahud outlier.

The results of chemical analyses of water from eight 

wells (fig. 49.1) are summarized in the table below. 

The values shown are believed to be generally representative of the water in the principal zone of saturation. The low dissolved-solids content of the water inc.icates that there is little soluble material in rocks of 

the Nahud outlier and that some subsurface outflow 

into the Precambrian rocks does occur. The temperature of the water sampled for analysis ranged from 

91°F to 93°F.

Not included in the table are the results of the chemical analyses of water from wells in Nahud. There the 

dissolved-solids content of the water ranged from 1,804 

to 1,944 parts per million and the amounts of calcium, 

sulfate, and chloride were considerably higher than the 

riaximums given in the table for these constituents. 

Geologic data, together with the water-quality data, inSummary of chemical analyses of water from eight wells tapping 

the zone of saturation in the Nahud outlier of the Nubian Series

[All analyses made by the Ministry of Health, Republic of Sudan, 

per million, except as indicated]

Results in parts

Silica (Si02)_. ......-___--_-_-.____

Sodium (Na)_ _ _ _

Sulfate (S04 )-----'- ----------------

Chloride (Cl). .___-______--_-_.___.

Fluoride (F)_ ______________________

Nitrate (N03) ---------------------

Hardness as CaCOs: 

- Total __ __-__-___--_-_--___

Specific conductance --micromhos per 

cm at 20°C--

Range

5-60

25 - 85

10 - 22

4-41

0-10

0-30

.3- 1.0

.6- 3. 1

112 -260

60 -264

0-20

125 -400

Average

23

60

16

18

9

11

.5

1. 1

185

182

10

293

dicate that the wells in Nahud tap a zone of saturation 

that is not hydraulically connected with the principal 

zone of saturation in the outlier.

REFERENCES

Iskander, W., 1963, Geology and hydrology of sheet 54-L:

Sudan Geol. Survey, Dept. [In press] 

Kleinsorge, H., and Zscheked, J. G., 1958, Geologic-hydrologlc

research in the arid and semi-arid zone of the western

Sudan: Sudan Geol. Survey Dept. open-file report. 

Kleinsorge, H., Zscheked, J. G., and Iskander, W., 1963, Geology

of sheets 54-P and 54-O: Sudan Geol. Survey Dept. [In

press] 

Sanford, K. S., 1935, Geological observations on the north-west

frontiers of the Anglo-Egyptian Sudan and the adjoining

part of the southern Libyan Desert: Geol. Soc. London

Quart. Jour., v. 91, p. 323-381.
Article 50

CONTAMINATION OF GROUND WATER BY SEA-WATER INTRUSION

ALONG PUGET SOUND, WASHINGTON, AN AREA HAVING ABUNDANT PRECIPITATION

By GRANT E. KIMMEL, Tacoma, Wash.

Work done in cooperation with the

State of Washington Department of Conservation, Division of Water Resources

Abstract. Some wells at Hyada Park, on an upland coastal 

area near Tacoma, Wash., produce water that, with continued 

pumping, becomes too saline for use. Although precipitation 

is abundant, recharge to the body of fresh ground water beneath 

the area apparently is limited by the low permeability of the 

glacial till that mantles the upland.

As much as 2,000 feet of unconsolidated sedimentary 

rocks of late Tertiary and Quaternary age underlies 

the upland areas that border Puget Sound. These deposits locally contain highly permeable aquifiers, and 

large quantities of water are pumped from the many 

wells tapping them. Because precipitation is abundant 32 to 64 inches per year it has been assumed 

that the ground-water reservoir, in general, is amply 

recharged. For that reason, the possibility that the 

supply of fresh ground water in the Puget Sound 

region is vulnerable to sea-water intrusion has been considered remote.

The first instance of marked intrusion of sea water 

was in the summer of 1959 when 1 of 6 public-supply 

wells at Hyada Park, just north of Tacoma, Wash., 

began to yield detectably salty water. Since then, each 

time this well has been pumped, the water has become 

too saline for use. Three of the other public-supply 

wells now yield moderately salty water, and one yields 

water that is slightly saline. As yet there has been no 

detectable increase in the chloride content of water from 

the sixth well. In the hope of finding an additional 

supply of uncontaminated water, a well was drilled into 

deeper water-bearing beds; however, the well was 

abandoned when the water was found to be highly 

saline.

The purpose of this article is to summarize the available geologic and hydrologic information that pertains

to this first known occurrence of salt-water intrusion in 

the Puget Sound region. An understanding of the 

factors that govern saline intrusion in the Hyada Park 

area would serve as a guide for the future development 

of ground-water supplies in that and similar areas along 

the coast of Puget Sound.

Hyada Park is on a westerly projection of an upland 

block that extends about 25 miles south from Seattle 

and is about 9 miles wide at the latitude of Hyada Park 

(fig. 50.1). The block is bordered on the west and north 

by Puget Sound and is separated from other uplands to 

the east and south by broad trenches occupied by the 

Green and Stuck Rivers, respectively. The two trenches 

are believed to have been formed during the ice age by 

the erosive action of glacial ice and (or) glacial melt 

water rather than to have been cut more recently by 

the streams now flowing in them.

At Hyada Park, glacial till lies at or near the surface, 

as is true of the entire upland block of which Hyada 

Park is a westerly extension. Beneath the till is a deposit of sand and gravel, largely glacial drift, which is 

as much as 500 feet thick under the upland and contains 

thin discontinuous layers of silt and clay.

Precipitation, about 36 inches per year in this part of 

the Puget Sound region, is the sole source of recharge 

to the ground-water reservoir beneath the upland block. 

Most of the precipitation occurs in the period October 

through April. As the evaporation and transpiration 

rates are lowest during that period, most of the precipitation on the upland sinks into the glacial till or 

runs off, although some is stored temporarily in swampy 

depressions and lakes on the till surface and infiltrates 

or is disposed of by evapotranspiration. Because glacial 

till is characterized by a generally low coefficient of

ART. 50 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGE'S B182-B185. 1963.

B182
KIMMEL B183

50.1. Map of the Puget Sound region showing the location of the Hyada Park area.

bermeability, the quantity of water infiltrating through 

it to the zone of saturation in the underlying highly 

permeable deposits is smaller than at many other places 

in the Tacoma area.

The six public-supply wells at Hyada Park are shown 

in figure 50.2 as LI, L2, PI, P2, P3, and Nl. The depth 

of well L2 is unknown, but. the others are known to exto depths ranging from 60 to 170 feet below mean 

level. Well N2, the deep test well that was abandoned, reached a depth of 329 feet below mean sea 

evel. The drillers' logs of wells P2, P3, and Nl- indi- 

£ate that these wells tap deposits consisting chiefly of 

$and and gravel, but apparently the individual strata 

iire not continuous from well to well. However, at least 

k>me of the zones tapped by these wells are known to

R.3E.

CONTOUR INTERVAL 50 FEET 

DATUM IS MEAN SEA LEVEL

FIGTJEE 50.2. Topographic map of the Hyada Park area showing 

the location of public-supply wells.

be hydraulically interconnected because the water level 

in well PI is lowered about a foot when well P3 is 

pumped at a rate of 200 gpm (gallons per minute) and 

the water level in well N2, at the time drilling had 

reached a depth of 262 feet below mean sea level (157 

feet below the bottom of the perforated interval in 

Nl), was affected by the pumping of well Nl. In well 

Nl, water-level fluctuations due to ocean tides have an 

amplitude of at least 1 foot; it is reported that the 

water level in well LI also fluctuates with the tide.

The yield of wells LI and L2 was not measured but is 

estimated to be about 200 gpm each. The measured 

yield of the other four public-supply wells is as follows: 

PI, 40 gpm; P2, 80 gpm; P3, 215 gpm; and Nl, 575 

gpm. The wide range in yield probably reflects differences not only in the transmissibility of the aquifer but 

also in well construction and pump capacity. A pumping test of well P3, the casing of which is perforated in 

the interval 30 to 60 feet below mean sea level, indicated 

a coefficient of transmissibility of 10,000 gallons per day 

per foot for the deposits tapped. The specific capacity 

of this well is 8 gpm per foot of drawdown.

The static water level in both wells LI and P3 was 

about 6 feet above mean sea level in both 1960 and 1961, 

and the static water level in well PI was reported to be
B184 GROUND WATER

10 feet below mean sea level when drilled. Throughout 

the drilling of well N2, the water level in it remained at 

about sea level, but it is not certain that the water level 

in the well was representative of the static water level 

in the aquifer in the immediate vicinity of the well. 

The sparse information on static water levels does not 

warrant generalizations on the direction or steepness of 

the hydraulic gradient. During pumping, the water 

level in each well is drawn down below mean sea level.

The chloride content of most of the ground water in 

the Puget Sound region is between 3 and 10 ppm (parts 

per million). However, as may be seen in the table 

below, the chloride content of the water from 4 of the 

public-supply wells (LI, L2, and P2, and Nl) has exceeded 200 ppm on at least one occasion and in well Nl 

has been as high as 1,050 ppm. During a 5-hour period 

of pumping well Nl, the chloride content of the water 

increased from 54 to 980 ppm. Samples taken from 

wells LI, L2, and P2 seem to indicate a progressive increase in chloride content during the summer, which is 

the period of heaviest pumping, but the evidence is not 

conclusive because the wells probably were not pumped 

continuously at a constant rate.

From the evidence, it cannot be determined whether 

the general position of the interface between salt water 

and fresh water has changed significantly. Instead, the 

normal position of the interface may be near enough to 

the bottom of the wells that, when the wells are pumped, 

the position and shape of the salt-water-fresh-water 

interface may be distorted only temporarily as the result of pumping (fig. 50.3).

The vulnerability of the ground water in the vicinity 

of Hyada Park to contamination by sea-water intrusion 

undoubtedly is related to the low rate of recharge to the 

ground-water reservoir. In turn, the low rate of recharge in spite of abundant precipitation probably is a 

function of the low permeability of the glacial-till mantle through which the recharge must infiltrate. Unless 

the rate of recharge can be increased artificially, an

Chloride content of water from wells at Hyada Park, Wash., and 

approximate elevation of water source

Well No.

LI

Pl.__----------------

P2 ----------------

Nl

Date collected

June 4, 1960

July 5, 1960 

Aug. 11, 1960 

Sept. 11, 1960 

July 30, 1959

Dec. 3, 1959 

July 5, 1960 

Aug. 11, 1960 

Sept. 11, 1960 

July 30, 1959

June 4, 1960 

July 5, 1960 

Aug. 11, 1960 

Sept. 11, 1960 

July 4, 1961 

July 30, 1959

June 4, 1960 

July 5, 1960 

Aug. 11, 1960 

Sept. 11, 1960 

Julv 4, 1961 

May 13, 1961

July 4, 1961 

July 30, 1959

Dec. 3, 1959 

June 2, 1960 

Oct. 31, 1960

Chloride 

content 

(ppm)

139

231 

241 

302 

93

32 

145 

240 

113 

18

17 

18 

19 

19 

19 

399

8.3 

220 

770 

57 

22 

1 9.5

9.7 2 54-980

535 

1,050 

880

Position of 

bottom of 

well or of 

perforated 

interval 

(feet below 

sea level)

170

Unknown

60

85

30-60

70-105

1 After 8 hrs. pumping at 215 gpm.

2 Continual rise in chloride content during 5 hrs. pumping at 200 gpm.

unlikely possibility in the Hyada Park area, the problem becomes one of determining the sustainable freshwater yield of the ground-water reservoir and of limiting withdrawals to that amount. Quite possibly, the 

sustainable fresh-water yield on an annual basis actually 

exceeds the present rate of withdrawal, but to produce 

fresh water without saline contamination at the most 

efficient rate will require careful planning of the number, depth, spacing, and pumping schedule of the wells. 

Also, additional facilities for storage may be needed, so 

that some of the water needed for times of peak demand 

can be pumped during periods of low use.
KIMMEL B185

Land surface

Water

Fresh

surface

water

Salt water

WELL NOT BEING PUMPED

Land surface

Fresh water

Salt water

PUMPING AT LOW RATE OR EARLY 

STAGE OF PUMPING AT HIGH RATE

Land surface

Fresh

*/'

/

' water

Salt water

PUMPING AT HIGH RATE

FIGUEE 50.3. Sketch illustrating possible distortion of the interface as the result of pumping.
Article 51

INFLUENCE OF LAND-SURFACE CONDITIONS ON GROUND-WATER TEMPERATURES 

IN SOUTHWESTERN SUFFOLK COUNTY, LONG ISLAND, NEW YORK

By E. J. PLUHOWSKI and I. H. KANTROWITZ, Mineola, N.Y.

Work done in cooperation with the

Suffolk County Board of Supervisors and Suffolk County Water Authority

Abstract. Variation in absorbed solar radiation among differing surface environments in southwestern Long Island, N. T., 

results in lower ground-water temperatures under wooded areas 

tban under cleared areas. Shade and a layer of organic material on tbe ground account for tbe lesser amount of solar 

radiation absorbed in tbe wooded area.

Data collected during 1959-61 from 13 observation 

wells in the Champlin Creek basin of southwestern Suffolk County, N. Y., indicated that in the summer, 

ground-water temperatures at shallow depths under 

residential, deforested areas may be 3° to 5°F higher 

than those observed under wooded areas (Pluhowski 

and Kantrowitz, 1961). Twelve of the observation 

wells were screened about 5 feet below the water table 

at depths ranging from 8 to 17 feet below land surface; the remaining well was screened about 55 feet 

below the water table at a depth of 60 feet below land 

surface. Because most of the wells were in an area 

where the temperature of the ground water possibly 

was affected by cesspool effluent, it was decided to make 

a similar study in two areas where such recharge was 

less likely to be a factor. The results of that additional 

study are reported in this article.

The sites selected were the Sawyer Avenue and 

August Road pumping stations of the Suffolk County 

Water Authority. These pumping stations are about 7 

miles west of Champlin Creek and, at both, the water 

table is about 5 feet below the land surface and the general direction of ground-water movement is southward. 

Five l^-inch drive points were installed at Sawyer 

Avenue and three at August Road. Both sites are hydrologically similar to the Champlin Creek basin, each 

being underlain by glacial outwash deposits having a 

thickness of about 80 feet and a coefficient of permeability of about 1,000 to 1,500 gallons per day per 

square foot.

The Sawyer Avenue site is about 900 feet down the 

hydraulic gradient from the nearest houses. Between it 

and the houses is an athletic field which is devoid of 

trees and shrubbery except for a small stand of trees 

about 30 feet upgradient from the group of 5 observation wells used at this site. The screens in the wells were 

12, 15, 44, 67, and 72 feet below the land surface.

The August Road site is in a heavily wooded area that 

is virtually in its native state. Until a few years ago, 

the woods extended about 2,500 feet upgradient from 

the site, but now there are houses as near as 400 feet. 

The well screens in the observation wells were 15, 38, 

and 77 feet below the land surface.

Temperature measurements were mada biweekly during 1961-62. Each well was pumped briefly before the 

water sample was collected in a thermos bottle, and the 

temperature of the sample was determined immediately 

with an expanded-scale thermometer.

As shown by the graphs in figure 51.L4, the maximum temperature of the ground water to a depth of 

about 75 feet beneath the cleared Sawyer Avenue site 

is 2° to 4°F higher than the maximum temperature of 

the ground water at equivalent depths beneath the 

wooded August Road site. This relation is similar to 

that previously noted in the Champlin Creek basin, and 

suggests that in cleared areas the absence of shade and 

lack of an insulating layer of organic material, such 

as is common on the ground in wooded areas, increases 

the opportunity of the soil to absorb solar energy. 

Thus, the temperature of the soil in cleared areas is 

higher than that in wooded areas, particularly during 

the summer months when the emission of solar radiaART. 51 IN U.S. GBOL. SURVEY PROF. PAPER 475-B, PAGEiS B186-B188. 1963.

B186
PLUHOWSKI AND KANTROWITZ B187

tion is at peak intensity. As a result, a greater amount 

of heat is exchanged during the summer between the 

fcone of aeration and the zone of saturation in cleared 

areas, and the temperature of the shallow ground water other factors being equal becomes correspondingly higher than in wooded areas. Furthermore, the 

higher soil temperature in cleared areas results in either 

a greater warming or a lesser cooling of any water that 

infiltrates deep enough to become recharge to the zone 

of saturation. This factor, obviously, also tends to increase the differences in maximum ground-water temperatures between wooded and cleared areas.

The differences in minimum ground-water tempera- 

^ures at equivalent depths in the cleared and wooded 

areas (fig. 51.15) is much less than that for maximum 

temperatures; in fact, the differences are not significant. 

JDuring the winter, the low angle of the sun, short 

^ength of day, higher percentage of cloudiness, and ocbasional snow cover tend to minimize the differences in 

j:he amount of solar energy absorbed in the wooded and 

pleared areas. Consequently, during the cooler months, 

the rate of heat exchange between the zones of aeration 

and saturation under both types of land surface is vir2 20 

<r

D

0 30

z

5 40

jrl 50

Js6?

80

90

100

10

10 

< 20

J Q 

LJ 40

l!LUJ 60 ^ ffl 70 

80 

90 

100

Woode

10

54 56 58 60 62 64 

MAXIMUM TEMPERATURE, 

IN DEGREES F 

A

44 46 48 50 52

MINIMUM TEMPERATURE,

IN DEGREES F

B

50

2 4 6 8 10 12 14 16 18

ANNUAL RANGE IN TEMPERATURE,

IN DEGREES F

C

40 

50 

60 

70 

30 

90 100 [

Cleared

51 52 53 54 55 

MEAN ANNUAL TEMPERATURE 

IN DEGREES F 

D

FIGURE 51.1. Relation of ground-water temperature to depth 

below land surface of wells at the August Road (triangles) 

and Sawyer Avenue (circles) pumping stations, Babylon,

j N.Y., 1961-62.

 680231 0 63 13

tually the same and the differences in minimum groundwater temperatures are correspondingly slight.

Theoretically, recharge from cesspool effluent should 

tend to raise ground-water temperatures, particularly 

in winter. However, no such effect could be detected 

at the August Road site, even though cesspool effluent 

from the housing developments upgradient from the 

site probably moves about 1 foot per day and, consequently, has had sufficient time to reach the observation wells. The slight differences in the minimum temperatures at equivalent depths beneath the two sites, and 

the fact that the maximum temperatures at the August 

Road site are lower than those at equivalent depths below the Sawyer Avenue site, indicate that any heat 

added to the ground water by cesspool effluent upgradient from the August Road site is dissipated before the 

effluent reaches the observation wells.

As shown in figure 51.1£7, the annual range in the 

temperature of the ground water decreases with depth. 

At the cleared Sawyer Avenue site the annual range 

decreases from about 17°F at a depth of 12 feet to about 

3.5°F at a depth of 72 feet, and at the wooded August 

Road site from about 9°F at a depth of 15 feet to about 

2.5°F at a depth of 77 feet. It is estimated from these 

data that the annual range of ground-water temperature is about 6°F greater in the cleared area than in the 

wooded area at equivalent shallow depths and about 

1°F greater at a depth of 70 feet. The difference in 

ground-water temperature fluctuations probably continues to decrease with greater depth and at some unknown depth becomes undetectable. Below the depth 

at which the. exchange of heat between the zone of aeration and the zone of saturation is an effective cause of 

temperature fluctuations, the dominant factor controlling ground-water temperature is the outward flow 

of heat from the interior of the earth. Because the 

rate of outward flow of heat in any given locality is 

relatively uniform, the temperature of the ground water 

at any specific depth greater than about 100 feet remains 

nearly constant throughout the year.

Although the mean annual temperature of the shallow ground water at both sites decreases with depth (fig. 

51.1Z>), the mean annual temperature of the ground 

water below some unknown level increases with depth in 

accordance with the natural geothermal gradient. The 

mean annual temperature of the shallow ground water 

at the Sawyer Avenue site is about 1.5° higher than at 

the August Road site. At the latter, it is more nearly 

equivalent to the mean annual air temperature (51°F) 

of the area.

The insulating effect of the zone of aeration is evident from the fact that the annual range in air temperature on Long Island may be as much as 100°F, whereas
B188 GROUND WATER

ground water at the shallowest depth has an annual 

temperature range of only 1Y°F. Moreover, it has been 

demonstrated that the ground-water temperature is 

unaffected by the alternate passage of warm and cold 

air masses (Pluhowski and Kantrowitz, 1961, p. 101- 

102). By acting as a buffer between the ground-water 

reservoir and the atmosphere, the zone of aeration also

causes the extremes in ground-water temperature to lag 

about 2 months behind the extremes of air temperature.

REFERENCE

Pluhowski, E. J., and Kantrowitz, I. H., 1961, Hydrology of the 

Babylon-Islip area, Suffolk County, Long Island, N. Y.: 

U.S. Geol. Survey open-file report, p. 99-104.
Article 52

SEASONAL CHANGES IN THE CHEMICAL QUALITY OF SHALLOW GROUND WATER 

IN NORTHWESTERN ALASKA

By ALVIN J. FEULNER and ROBERT G. SCHUPP, Anchorage and Palmer, Alaska 

Work done in cooperation with the U.S. Air Force, AlasJcan Air Command

Abstract. Marked increases in mineralization of ground 

water during the colder months are accounted for by simple 

concentration by freezing, and by reduction in dilute recharge 

Curing winter.

Data on the.quality of shallow ground water at a 

site in northwestern Alaska indicate that marked systematic changes occur throughout the year. These 

changes are believed due to dilution of mineralized 

ground water by melt water and precipitation in the 

summer and by concentration of mineral constituents 

in the water through freezing during the .winter. The 

,data consist of analyses of a series of water samples 

collected during the period June 1961 to July 1962 at 

a remote site near the Bering Sea in northwestern 

Alaska, just south of the Arctic Circle. The water 

supply was developed during the late summer of 1961 

by the U.S. Air Force; the analyses were made to rate 

the water for use as boiler feed and to determine what 

treatment, if any, would be required to make it suitable 

for that use.

Although permafrost is known to extend to a depth 

of 1,000 feet or more in this part of Alaska, water at 

the site described here remains fluid throughout the 

year in a permeable zone most likely fractured rock 

along a high-angle fault which trends northward into 

the nearby mountains (fig. 52.1). The occurrence of 

springs in summer marked by ice mounds in winter suggested the possibility of developing a year-round water 

supply from this permeable zone. The temperature of 

the spring water was reported to be 35°-36°F in August 

1959. The explanation of the year-round fluidity of 

the water has not been determined. It is unlikely that 

the water rises from a great depth or has been in contact 

with incompletely cooled igneous rocks; if such were the

167° 58'

FIGURE. 52.1. Sketch map showing fault, spring, and watersupply installations; datum mean sea level.

case, the water probably would be more highly mineralized than it is. More likely, the water is derived from 

precipitation and snowmelt in the nearby mountains, 

and because of its relatively high velocity through the 

permeable zone it does not freeze completely.

ART. 52 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B189-B191. 1963.

680231 O 63- -14 B189
B190 GROUND WATER

A water supply was developed by excavating a trench 

into the permeable zone, placing perforated pipe along 

the bottom of the trench, and backfilling the excavation. 

Water entering the perforated pipe flows by gravity 

through a lateral pipe nearly 400 feet long to a sump 

(fig. 52.1) from which it is pumped for use. To prevent freezing of the water enroute to the sump, steampipes were laid along the full length of the lateral pipe, 

which was laid in the bottom of a trench excavated by 

blasting frozen bedrock.

During the coldest months the inflow of water to the 

sump declines, because recharge to the permeable zone 

virtually stops and because of freezing inward from the 

boundaries of the permeable zone. The discharge to the 

sump has ranged from several hundred gallons per minute for short periods to a low of 1,000 gallons per day in 

late May 1962. The area lacks a well-defined drainage 

system; the nearest stream, which flows only during the 

warmer part of the year, is about 600 feet away and 

downslope from the permeable zone; thus recharge is 

chiefly from infiltration of snowmelt and rain.

Water samples were collected by Air Force personnel 

for chemical analysis by the Geological Survey in June, 

September, and December 1961 and in January, February, April, May, June, and July 1962. Of these, the first 

and second were collected before installation of the 

water-supply system was completed. All the analytical 

results are given in the accompanying table, and selected 

results are shown in figure 52.2.

As may be seen by examination of the table and figure 

52.2, the chemical quality of the water changed systematically during the period of sampling. For example, the sulfate content was 5 ppm (parts per million) 

or less in June and September 1961, increased to more 

than 40 ppm by late April and early May 1962, and 

dropped to 6 ppm by early July 1962. The noncarbonate hardness showed a similarly wide range as low as 

6 ppm in September 1961 and 5 ppm in July 1962, but 

about 10 times as great in April and May 1962. The

JUNE JULY AUG.SEPT.OCT. NOV. DEC. JAN- FEB. MAR. APR. MAY JUNE JULY

1961 1962

FIGURE 52.2. Changes in the chemical quality of the water during the period June 1961 to July 1962.

differences between the summer and winter concentrations of magnesium, silica, and dissolved solids, in the 

summer and winter carbonate hardness, and in the summer and winter specific conductance follow a similar 

trend.

The mineral constituents most likely to show a wide 

range in concentration probably are determined by the 

kinds of water-soluble minerals with which the water 

has come into contact. At this particular site, where 

the summer and winter concentrations of sulfate differ 

so widely, the ground water has percolated through 

partly metamorphosed thin-bedded limestone intruded

Mineral constituents, in parts per million, and other characteristics of ground water at a site in northwestern Alaska

Date of

collection

1961

June 20

Sept. 14

Dec. 11

1968

Jan. 15

Feb. 12

Apr. 5

Apr. 20

May 10

June 12

July 3

Silica

(SiCh)

4.6

3.7

2.9

3.6

4.9

7.8

5.4

5.4

5.5

4.7

Iron

(Fe)

0.05

.00

.02

.02

.02

.02

.02

.02

.02

.07

Manganese

(Mn)

0.00

.00

.02

.00

.00

.01

.00

.00

.00

.00

Calcium

(Ca)

12

9.2

15

17

23

33

31

35

19

14

Magnesium

(Mg)

1.4

1.2

1.8

1.3

3.0

4.9

3.9

3.9

1.5

1.2

Sodium

(Na)

6.9

6.1

6.7

7.2

8.1

10

11

11

7.5

6.8

Potassium

(K)

0.3

.4

.1

.1

.5

.6

.6

.7

.6

.4

Bicarbonate

HC03

27

26

42

39

42

53

53

56

55

43

Sulfate

(SOO

5.0

4.0

10

15

26

39

42

46

9.0

6.0

Chloride

(01)

16

11

10

11

16

25

21

21

11

9.0

Fluoride

(F)

0.6

.6

.8

1.0

1.3

.7

1.0

.8

.4

.5

Nitrate

(NOs)

0.4

.2

.6

.3

.9

2.9

1.7

1.9

1.9

.8

Dissolved

solids

(residue

on evaporation

at 180°C)

60

49

69

76

105

150

131

154

83

64

Hardness as

CaCOs

Carbonate

36

28

45

48

70

103

94

103

54

40

Noncarbonate

14

6

10

16

36

59

51

58

9

5

Specific

conductance (micromhos

at 25°C)

106

92

122

138

180

258

253

270

146

107

PH

7.2

7.1

7.1

7.2

7.8

7.4

7.3

7.7

7.4

7.3

Color

0

5

0

0

0

5

0

0

5

0
FEULNER AND SCHUPP B191

by granite. Pyrite, pyrrhotite, fluorite, and sphalerite 

are common in the mineralized zones along the contact 

of the limestone with the granite. Sulfate compounds 

formed during weathering of these and other sulfide 

minerals are dissolved readily in the ground water. 

The relative enrichment of the ground water in calcium 

and sulfate in the later winter months probably is due 

to the fact that recharge from precipitation ceases and 

the discharge from the permeable zone consists almost 

entirely of ground water that has long been in contact 

with mineralized rocks. The recharge from precipitation, which has a relatively high sodium and chloride 

content owing to proximity to the Bering Sea, is cut off 

during the winter as soon as the ground freezes. The 

increases noted in sodium and chloride probably are 

due to simple concentration by freezing whereas the 

larger increases in calcium and sulfate are due in part 

to concentration by freezing but also to the fact that 

the winter discharge has been in contact with the rock 

materials longer. Sparse data from another Air Force 

station in Alaska, where the bedrock consists wholly 

of limestone, indicate that there the greatest seasonal 

differences in the chemical quality of the ground water 

are in the concentrations of calcium and bicarbonate 

and in the carbonate hardness, and are caused by a 

combination of the processes similar to those described 

above. R. M. Waller (U.S. Geological Survey, oral 

communication, 1962) reported observing similar winTtime increases in mineral content in ground water 

owing beneath a frozen stream in another part of

laska.

As shown by figure 52.2, the mineral content of the 

water was lowest during the summer, increased gradually during the fall, increased relatively rapidly during the winter and early spring, and then declined very 

rapidly in the late spring to about the same level as 

that of the previous summer. Concurrent with the 

more rapid increase in mineralization was a marked 

decline in the yield of the water-supply system. The 

decline in yield presumably was caused partly by a 

freezing inward from the sides and downward from the

top of the aquifer, thus reducing the cross sectional area 

through which the water could flow. This freezing 

accentuates the normal seasonal decline in yield, which 

is due to lack of recharge from precipitation in the 

winter and to depletion of the water in storage in the 

aquifer. The increase in mineral content was caused 

partly by selective concentration of mineral matter in 

the unfrozen water. The freezing accentuates a normal 

seasonal increase in mineral content, which is due to a 

reduction in the diluting effect of recharge from precipitation. Thus, in an aquifer 20 feet thick, freezing 

extending 15 feet into it may result' in a doubling or 

tripling of the dissolved-solids content in the water that 

remains unfrozen. The rapid decline in the mineral 

content of the water in the late spring is related to the 

thawing of the ice within the aquifer, plus infiltration 

of surface melt water which recharges the aquifer. 

The double peaks on several of the curves in figure 

52.2 probably indicate that at least one slight thaw 

preceded the main thaw in the spring of 1962; that 

the thawing proceeded by stages is confirmed by airtemperature data. As the water level in the sump 

is reported to have risen markedly on May 28, 1962, it 

may be reasonable to assume that the maximum mineral 

concentration of the ground water was reached shortly 

before that date.

Although documentation is sparse, probably other 

shallow ground-water supplies in subarctic and artic 

areas are subject to similar marked seasonal changes in 

chemical quality. Several factors such as the depth 

of freezing before the snow cover becomes well established, the total moisture content of the snow cover, 

and the length of time the ground is snow covered 

may affect the amount of increase in the concentration 

of dissolved minerals somewhat. However, it seems 

unlikely that the extremes would differ more than 10 

percent from the values reported here. The observed 

changes are believed sufficiently representative to provide a basis for predicting the type of treatment that 

boiler-fed water may require at different times during 

the year.
Article 53

CHANGES IN GROUND-WATER MOVEMENT AND BANK STORAGE 

CAUSED BY FLOOD WAVES IN SURFACE STREAMS

By H. H. COOPER, JR., and M. I. RORABAUGH, Tallahassee, Fla., and Tacoma, Wash,

Abstract. Analytic solutions for the changes in groundwater head, ground-water flow, and bank storage that result 

from the passing of a flood wave in a surface stream are determined from equations defining asymme ric stage oscillations. 

Computations for a variety of conditions are presented.

A flood wave passing down a surface stream causes 

changes in ground-water movement and bank storage 

in an aquifer that is hydraulically connected with the 

stream. If the stage hydrograph of the flood wave 

can be described approximately by a mathematical 

function, analytic solutions for the changes in groundwater movement and bank storage can be derived.

Many flood-wave stage hydrographs are approximated by one of the family of curves defined by

 cos at), when 0<t<r (1)

where h0 is the maximum rise in stage, t is the time 

since the beginning of the flood wave, r is the duration 

of the flood wave, w STT/T, and 5 and N are constants. 

The constant 8 determines the degree of asymmetry. 

Curves for 5=w and 5=0 are shown in figure 53.1.

FIGUKE 53.1 Stage hydrographs described by equation 1 when 5=0 and

The constant N serves to make the curves of the 

family peak at the height h0 and is defined by

N=-r 1

The solutions for the changes in ground-water head, 

ground-water discharge into the stream, and bank storage are based on the following assumptions: (1) The 

stream channel has vertical banks and extends to the 

bottom of the aquifer, (2) in unconfined aquifers the 

slope and fluctuations of the water table are small in 

comparison to the saturated thickness (that is, the 

transmissibility of the aquifer remains virtually constant), (3) the stream does not overflow its banks, 

(4) the changes in ground-water head are due solely 

to the passing of the flood wave, (5) the water table 

or piezometric surface is initially horizontal and at the 

same height as the stream stage. The last assumption 

is made without loss of generality because, under the 

preceding assumptions, the solutions can be superposed 

on the initial condition.

Aquifer bounded by valley wall. Consider first an 

aquifer of finite width / bounded on one side by the 

stream and on the other by a vertical valley wall, as 

shown in figure 53.2. The differential equation governing nonsteady one-dimensional flow of confined 

ground water (Jacob, 1950, p. 333) is

_ (2)

where h is the ground-water head at time t since the 

beginning of the stage oscillation and at distance x from 

the stream bank, T is the coefficient of transmissibility 

of the aquifer, and 8 is the coefficient of storage of the 

aquifer. The conditions on h as given in equation 1 

and as shown in figure 53.2 are

h(x,G)=0, when 0<x<l, (2a)

ART. 53 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B192-B195. 1963.

B192
2 t*

OQ

CD* CO

"

- o to !-*> P 55

p co co |

o ^ P h"

8 g £ ^

&3 gr. +

TO B' i "* U

IIp ci,

O <!

l-h ^<

*J§«.

IV

oo

*

 1 <-

C3

xJ"

tol '"'

+"

^ J

^ C5 CO

o e1 CO P

Si Si

8 SP- &

JS J2

p.

e-t- §

ES*rrt *rrt

<^-* <^*

fefe

*fYir *TifTr 

«rnr JVT

1 1

C5 O

CO w 

to _

+ + 

CO CO 5' 5'

tstr

>r«v

N*

1.

crcj

>-" P

£' g 

P-> H P-

|cf^

CD cr p

TOCD

thp

tO

 Cb

d& 

crcj rB

g: o"

CDd

o.

(g

0

CQ

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>Q

e. 
B194

equation 8 yields

ANALYTICAL HYDROLOGY

and

4.

sin (o)t-\-<p arctan

and

X

Semi-infinite aquifer. If the aquifer is not bounded 

by a valley wall, the geometry of the aquifer is that of 

the semi-infinite half plane #>0. For this condition, 

solutions are obtained only for 5=0 that is, for the 

case where the stage hydrograph is sinusoidal. The 

conditions on h are given by equations 2, 2a, 2c, and

lim h(x,f)=0, when t > 0.

Z->co

The solutions satisfying these conditions are

(11)

* COS .-»

and

+- r e-»f sin far J-~\ -^-2 du \ (12)

7T Jo L V ffj W2 + C02 J y

r=^ / erfcfl-erfc 2 \

- r°

7T JO

From equations 5, 12, and 13 it can be shown that

4n-l

Qt<r='

7rn~ 1 2 II 1 (2m+l)

(13) V !

(14)

m=l

2

xJ4n-l"|

a J

II (2ro+l)

TO = 1

' (15)

By integrating equations 14 and 15 and substituting 

in equation 8, it is found that

h0

-

4n+l

n

and

4n+l

2 _/'0,.,/_

H (2m+1)

TO=1

(17)

Computations. Computed curves representing the 

ground-water flow and the bank storage caused by the 

sinusoidal stage oscillation described by equation 1 when 

8=0 are shown in figure 53.3. In finite aquifers the 

bank storage declines very rapidly when /? is large and 

slowly when /? is small. For /?^5 almost all the bank 

storage will have returned to the stream after one 

flood-wave period (£=T), and for /?^0.5 almost all 

of it will have returned after two flood-wave periods. 

(A value of /3=0.4 would correspond, for example, 

to T=10,000 ft2 per day, #=0.10, r=10 days, and 

Z=1,000 ft.) On the other hand, in semi-finite 

aquifers the bank storage declines very slowly. The 

remaining bank storage after 10 flood-wave periods is 

13.9 percent of the maximum and that after 100 periods 

is 4.3 percent.

REFERENCE

Jacob, C. E., 1950, Flow of ground water, Chap. 5 in Rouse, 

Hunter, Engineering hydraulics: New York, John Wiley & 

Sons, Inc.
COOPER AND RORABAUGH B195

Ground-water flow into stream

Semi-infinite aquifer and j8=0.1

3.0 3.5 4.0 4.5

FIGURE 53.3. Ground-water flow and bank storage resulting from stage oscillation defined by equation 1 when

5-0.
Article 54

PREPARATION OF TYPE CURVES FOR CALCULATING 7/5 

OF A WEDGE-SHAPED AQUIFER

By I. S. PAPADOPULOS, Denver, Colo.

Abstract. Selected type curves have been prepared for computing T/8 (coefficient of transmissibility/coefficient of storage) 

for a wedge-shaped aquifer from observed water levels. The 

procedure for computing the type curves is explained, and use of 

the curves for obtaining T/8 is outlined.

A wedge-shaped aquifer (in plan view) in Portage 

County, Wis., contains unconfined water. It is bounded 

on two sides by rivers, the Plover and the Little Plover, 

which meet at an angle of about 72°, as shown in figure 

54.1. The water level in the aquifer rises suddenly each 

spring owing to recharge from melting ice and snow 

and generally declines throughout the remainder of each 

year as the aquifer discharges water to the rivers. If 

the water stage in the rivers and the head along an arc 

at an assumed radius of influence are considered to be

FIGURE 54.1. Schema of a wedge-shaped aquifer in Portage 

County, Wis.

constant, a value for T/8' (coefficient of transmissibility/coefficient of storage) can be determined by analyzing the rate of decay of the recharge slug. The method 

of analysis is graphical and is based on type curves prepared from the analytical solution by Jaeger (1942, p. 

532, eq. 16) of a heat-flow problem in which the boundary conditions are identical to those of flow from the 

wedge-shaped aquifer. The applicability of Jaeger's 

solution was suggested by R. W. Stallman, of the Geological Survey.

Jaeger's solution, as applied to the ground-water 

flow system, may be expressed as follows:

where

(1)

(rak, m\

\ a ) r

m=l

tJt(a*, mt)dt, (2)

and in which

a=radius of influence, or radial distance from 

the confluence of the bounding streams to 

the boundary of constant head, in feet;

Tt b=7A8Sa2' '

Jv (x) ihe vth ordered Bessel function of the first 

kind;

k=-

r=radial distance from the confluence of the 

bounding streams to any point in the 

aquifer, in feet;

S coefficient of storage;

ART. 54 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B196-B198. 1963.

B196
PAPADOPULOS B197

s=water-level drawdown at time t at any point

in the aquifer, due to decay of recharge

slug, in feet; 

s0 =instantaneous water-level rise caused by

recharge slug, in feet; 

T= coefficient of transmissibility, in gallons per

day per foot; 

£=time since instantaneous change in water

level, in days; 

afcTO =the mth positive root of Jfc (z)=0;

6=polar angle to any point in the aquifer, in

radians; and 

00 := total angle of the wedge-shaped aquifer, in

radians.

As ground-water systems similar to the one shown 

in figure 54.1 are fairly common, an extensive evaluation of the function F(l>, r/a, 00 , O/Oo) would have practical application in the field of hydrology. If done 

manually, the multitude of computations needed for 

such an evaluation would consume prohibitive time; 

therefore, an extensive evaluation is being programmed 

for the electronic digital computing equipment of the 

Geological Survey. Until the results of that evaluation 

become available, the function can be evaluated numerically for use in analyzing data collected at specific 

locations. For the benefit of those who would like to 

make immediate use of equation 1, the procedure followed in computing values for the required type curve 

of S/SQ versus l> for each of the three observation wells 

shown in figure 54.1 is decribed in the following 

paragraphs.

Preliminary computations for the ground-water flow 

system depicted in figure 54.1 indicated that k equals 

2.5 (2n+1). For each value of n, the summation on m 

was computed and multiplied by the sine term. The 

sum of all the products, one for each n, gave the value 

of the function.

As the summation on m is the most complicated 

part of the computations, the procedure followed is 

described step by step below.

1. The roots akim of Jfc (z)=0, for 0<z<50, were 

obtained by preparing a graph of Jk(x) versus x, values 

for which were either computed from the following 

equation given by McLachlan (1955, p. 193, eq. 54)

(3) 

or taken from tables of Jk (x) given by Watson (1962,

p. 740). For x>50, the following equation given by 

McLachlan (1955, p. 30, eq. 4) was used:

(M -l) 4(M -1) (7/J-31) 

<xk, m ~P gp 3(8/3)3

32(M -1) (83M2 -982M +3779) 

15(8/3)6

64(M -l)(6949M3 -153855M2+1585743M -.6277237)

105(8/3)7

where

(4)

 |)x, and

2. Values of Jk(raktTn/a) were obtained from either 

equation 3 or the specially prepared graph of Jk(x) 

versus x.

3. Values of Jk-i(<xktTn) were obtained from either 

equation 3 or a specially prepared graph of J^-^x) 

versus x.

4. The integral in equation 2 was put to a new form 

by substituting z for (ak>mt), which yields -the following:

r tjk(ak, mt)dt=4- r*'

JO ak,m Jo

For afc>OT <50, the integral in this new form was made 

to depend on Fresnel's integrals, tables of which are 

given by Watson (1962, p. 744), and on Bessel functions of the (jo+^)th order and first kind, by the 

following two formulas given by Jaeger (1942, p. 532, 

eq. 19, 20) :

/»z /»z

zJk (z)dz=k 

Jo Jo

(6)

For afc>OT>50, the integral was evaluated by using 

equation 6 and the following equation, which was presented by McLachlan (1955, p. 91, eq. 7) :

Jo r JT

sm __ __

1 1

cos x- Tr-

(8)

After all the individual terms in the summation on m 

had been evaluated, the value of

k-i (<xk, m)], f1 tjk(ak, mt)dt (9) 

Jo
B198 ANALYTICAL HYDROLOGY

was computed and multiplied by the exponential factor. 

To insure accuracy to 2 decimal places in the final tabulation, the summation was continued until convergence 

to 4 decimal places was obtained. After the numerical 

values of expression 9 had been computed, the evaluation of the function for different values of 6 was a very 

simple task because 6 is contained in only the exponential 

factor. By inspection, the upper limit of 6 was determined to be K; for this value, the exponential term becomes zero to four decimal places, even for the smallest 

value of ak>m. As 6 was decreased, the number of terms 

in the summations on m and on n increased. The 

value of equation 1 became very small at 6=10~3, for 

which n had to be given values up to. 4 for convergence 

to 4 decimal places. For an assumed value of T/S= 

5X105, a curve of s/s0 versus 6 within the range 10~3<C 

6<Cl is sufficient for an analysis of a 1-year decay of 

water stage. Values of 6=1X10-*, 2X10'*, and 

5X10"*, where i=l, 2, and 3, gave enough points for 

smooth curves. Because most of the terms computed 

in the evaluation of equation 2 were independent of the 

location of the well, they could be used directly for 

computing the type curves of the other two wells.

Figure 54.2 shows the type curves for the three 

observation wells shown in figure 54.2. The dashed 

part of the curves (in the range 10~4<6<10~3) for 

wells Pt 374 and Pt 357 may not be accurate because 

it was extrapolated from the trend of the curves in 

the range 10-3<6<10~2. The range 10-4<6<10~3 

corresponds to the early stage in the period of decay. 

Because the observed data for that stage may not be 

representative, it is not advisable to rely on results 

obtained from the part of the curve that corresponds 

to that stage.

To find T/S, the observed s/s0 values for each well

are plotted versus log t. Each of the resulting curves 

is superposed on the corresponding type curve, holding 

s/s0=l in coincidence. A match point is selected, the 

corresponding b and t values are noted, and T/S is 

calculated from

T/S=7A8ba2/t. (10)

10

FIGTJBE 54.2. Type curves of water-level decay for three selected observation wells in a wedge-shaped aquifer.

REFERENCES

Burington, R. S., 1958, Handbook of mathematical tables and 

formulas : Sandusky, Ohio, Handbook Publishers, Inc., 296 p.

Jaeger, J. C., 1942, Heat conduction in a wedge, or an infinite 

cylinder whose cross-section is a circle or a sector of a 

circle: Philos. Mag. and Jour. Sci., v. 33, no. 222, p. 

527-536.

McLachlan, N. W., 1955, Bessel functions for engineers : London, 

Oxford Univ. Press, 239 p.

Watson, G. N., 1962, A treatise on the theory of Bessel functions, 

2d ed.: Cambridge Univ. Press., 804 p.
Article 55

EFFECT OF THE INJECTION SCHEME ON THE SPREAD OF TRACERS 

IN GROUND-WATER RESERVOIRS

By AKIO OGATA, Honolulu, Hawaii

Abstract. Commonly the dispersion of a tracer or contaminant injected into an isotropic granular medium, in which the 

regional flow is unidirectional, is attributed primarily to microscopic velocity variations and ionic diffusion. This is true provided the injection scheme neither disturbs nor alters the preexisting flow regime. The effect on the spread of tracers, caused 

by injection disturbance, is discussed using two examples. The 

first example concerns injection from an elliptic source-and the 

second concerns injection from a point source. These examples 

demonstrate that the spread of tracers caused by these methods 

of injection may be much larger than the spread due to diffusion.

The spread of tracers or contaminants introduced into a ground-water system is attributed to dispersion 

and convection. Dispersion phenomena in isotropic 

media exist primarily because of microscopic variations 

in velocity within each capillary tube and what is commonly called molecular diffusion. Some laboratory 

investigations of dispersion, both transverse and parallel to the direction of flow, have been examined. In 

explaining the observed phenomena, no mention is 

made of the spread of tagged particles that is relatable 

solely to the potential distribution in the existing flow 

field. When the analysis concerns one-dimensional dispersion in the direction of flow, consideration need not 

be given to the natural spread caused by flow from an 

injection source, or an existing discontinuity in the flow 

field. However, the effect of the existing injection 

scheme becomes important in analyzing dispersion 

transverse to the direction of flow.

In the following discussion it is assumed that the 

porous medium is isotropic, that conditions are such 

that two-dimensional analysis is valid, and that the flow 

of ground-water is unidirectional.

The differential equation describing mass transport 

from a tracer source into the surrounding groundwater region is analogous to the heat conduction or the 

diffusion equation with a moving source. One of the 

boundary conditions commonly postulated is that the 

head at which the tracer is injected is negligible relative

to the ground-water head near the injection point. -The 

following discussion attempts to show that the hydrodynamic conditions governing the flow system near 

the injection source must be considered before any 

realistic appraisal can be made of the magnitude of 

spread due to the diffusive system.

Laboratory analyses (Jacques and Vermeulen, 1957) 

and statistical analyses (de Josselin de Jong, 1958; 

Saffman, 1959) show that transverse diffusion is proportional to the magnitude of the average fluid velocity. 

This is not surprising since the statistical model assumes 

existence of a microscopic velocity transverse to the direction of the head gradient. The prime difficulty is 

in analyzing the laboratory data from which the angle 

of spread is determined. In a short laboratory model, 

alteration of the flow system due to the injection scheme 

may not have reached an equilibrium state before 

measurement is made. Under such conditions laboratory data may tend to show a continuous spreading at 

a given angle, which approximates that of a diffusive 

process but which may be wholly attributed to hydrodynamic conditions.

Two examples will serve to illustrate the spreading 

process. The first example involves a tracer or contaminant source occupying an elliptical region of high 

permeability within a region of distinctly lower permeability in which the initial areal flow can be described by parallel flow lines. This exemplifies conditions in a laboratory model, where the cross-sectional 

area of an injection well may be a relatively large part 

of the surrounding region. The second example consists of a point source from which a tracer is injected 

into a region of parallel flow. This typifies conditions 

in the field where the wells are essentially points in 

relation to the overall region.

In the first example it is postulated that parallel 

areal flow exists at some distance from the elliptic discontinuity. Mathematically, the flow of a tracer from

ART. 55 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B199-B202. 1963.

B199
B200 ANALYTICAL HYDROLOGY

an elliptic source of high (for example, infinite) permeability is equivalent to the flow into an elliptic 

region of infinite permeability except for the reversal 

in direction. Furthermore, the analysis of flow into 

such an elliptic region is the same as the analysis of 

flow around an elliptic cylinder provided the potential 

theory holds. Bear in mind, however, that the potential functions and the stream functions are interchanged in the two systems. Inasmuch as most books 

on classic hydrodynamics illustrate calculations of flow 

around an elliptic cylinder, this method will be used.

An analysis of irrotational flow around an elliptic 

cylinder is presented by Milne-Thompson (1950, p. 

161-166). Only the more pertinent expressions are 

reproduced here to give continuity to this discussion.

The generalized expression for the complex potential 

of flow past an ellipse, given in elliptic coordinates, is

w=u(a-\-b) cosh [£ £0 ia]. (1)

The relation between the elliptic coordinates and the 

cartesian coordinates as well as the pertinent notation 

are shown in figure 55.1. By separating equation 1 

into its real and complex parts, the potential function, 

(j>, and the stream function, \f/, are determined to be

and (f>=u(a-\-b) cosh (£ £0) cos (57 0;), 

$=u(a-\-b} sinh (£ £0) sin (17 a). (2)

To determine the spread of the tagged particles, as 

shown in figure 55.2, the value of the limiting velocity 

potential, 0, which just touches and skirts the ellipse, 

must be found. It should be noted that in the groundwater or the conjugate system, the limiting velocity 

potential is the limiting streamline which touches the

Stagnation 

point

FIGURE 55.1. The geometry of flow past an ellipse.

elliptic well portrayed by the elliptic solid cylinder. 

In the following analysis the functions <£ and ^ in 

equation 2 will be interchanged so that no confusion 

as to the system considered will exist.

The geometry (fig. 55.1) of the problem indicates 

that the value of the limiting stream function can be 

most readily obtained at the stagnation point. The 

stagnation point corresponds to the point where the 

complex velocity potential is zero, hence

dw_dw d% _u(gjrb) sinh (£ £0 io) 

dz d dz c sinh £

=0.

From the above expression the stagnation point is 

determined to be at £=£0 and ri=a or ri=ot-\-Tr. Substituting these values in equation 2 gives \f/=±u(a-\-b) 

for the values of the limiting streamlines.

At a large distance from the ellipse where £»c, the 

complex potential given in equation 1 may be approximated by

«?=«'

In other words, the relation takes a form similar to that 

for the complex potential of a stream past a circular 

cylinder.

As z becomes very large, that is, as 0 ><» , the second 

term in equation 3 approaches zero. Hence, separating 

the real part from the imaginary part gives the equation for the velocity potential some distance from the 

ellipse; that is,

\l/=u[x cos a+2/ sin a]. (4) 

Substituting the value ±u(a-\-b) for ^, the equation of

Zone of uncontaminated 

water

Asymptote

Zone occupied by tagged 

or contaminated water

/6 - angle of attack 

. t/-areal flow

FIGURE 55.2. The effect of injection from an elliptic source.
OGATA B201

the asymptotes is

a? cos a+y sin a=±(a+b). (*)

In a similar manner, the equation of the zero streamline can be determined to be

X cos a + y sin a = 0. (6)

By analytical geometry and use of equations 5 and 6, the 

ultimate width of the streamtube between ^=0 and \]/=u 

(a + b) can be shown to be (&+ b) , or the full width between the limits \]/=±u(a+b) is determined as 2 (a+ b ). 

Thus the maximum spread between the limiting streamlines of the ground-water system is constant and independent of the flow direction. An interesting aspect is 

that any tracer fluid introduced from a finite well at 

zero velocity will ultimately take up a stream width of 

2(a+2>), as indicated in figure 55.2, a great distance 

from the region where the tagged water particles 

originate.

In the special case of a circular cylinder (that is, 

a b) the width of the stream at some distance from the 

source will be 2Z> where D is the diameter of the cylinder. Thus if a circular region of infinite permeability is 

present in the ground-water regime the ultimate spread 

of the tracer will be of the order 2Z>. The spread will 

occur without any dispersive effect and will be attributable solely to the existence of the discontinuity.

Another example as pointed out by H. R. Henry 

(written communication, Aug. 16, 1961) is to let &-»0, 

which then approaches the flow out of a line segment 

of infinite permeability. In that event, regardless of the 

angle of attack 6 (fig. 55.2), the tracer will assume a 

maximum band width equal to the length of the line 

segment.

In idealizing any large flow system in an actual field 

environment, any injection well can be considered a 

point source. The spread of a contaminant from an 

elliptic source is practically the same as that from a 

point source.

In practice, even in small laboratory experiments, it 

is virtually impossible to maintain dynamic conditions 

in which the head of the source is exactly the same as 

the head of the surrounding fluid at the point of injection. If the injection velocity is large in relation to the 

velocity of the ground-water system, the zone of spreading in an infinite aquifer may be extremely large. This 

spreading would be present without the existence of any 

mechanism which tends to disperse the injection fluid. 

In fact, the spread due to the strength of (amount of 

flow from) the point source may be extremely large compared to dispersion. This mechanism is best illustrated 

by analyzing the point-source injection system in a uniform flow field.

Physically, the point source may be regarded as an 

elliptic well reduced to zero diameter. In this analysis 

the head in the well is increased so that the head in the 

source will be greater than the head in the surrounding 

flow field. The complex potential of this two-dimensional system is given by Milne-Thompson (1950) as

w= U2 mlnz, (7)

where m is the strength of the source (ft3 per sec per ft 

of length of the source).

The stream function, determined from equation 7, is

= uy m tan" 1 y/x. (8)

The stagnation point can again be obtained by setting 

dw/dz=Q, and the value of the dividing streamline 

(^= m?r) can be readily determined. The equation 

of the outside or limiting streamline is given by the 

expression

TOTT m. _, y m , M y= tan 1 -= (TT 0). 

u u x u

(9)

Letting a?-» <» (or 0-»0) in equation 9, the value of the 

asymptote aproached by the limiting streamline is 

y=mir/u. The geometry of the flow field, showing the 

limiting streamline separating the region occupied by 

the tracer or contaminant from the region of uncontaminated water, is shown in generalized dimensionless 

coordinates in figure 55.3. The equation of the 

asymptote indicates that for an increase in the ratio 

between the strength, m, of the point injection source 

and the velocity, u, of the surrounding fluid there will 

be a corresponding increase in the maximum spread of 

the injected fluid. In other words, a tenfold increase 

in the ratio in/u produces a tenfold increase in the 

maximum spread.

Zone of uncontaminated water 

___ / Asy m ptote

7T

E

S >- c «>

Limiting streamline

Zone occupied by tagged 

or contaminated water

Source of strength m (ft 3 per sec per ft)

10 15 20

xu 

m

25 30

FIGURE 55.3. The effect of injection from a point source of 

strength m.
B202 ANALYTICAL HYDROLOGY

In conclusion, the injection scheme by which tracers 

or contaminants are introduced into an aquifer must 

be considered in order to interpret the results obtained. 

The neglect of the velocity component in the direction 

transverse to the regional flow will give reproducible 

results near the source; however, the description of the 

mode of mass transport may be erroneous. This is 

especially true in short laboratory models where the 

spread caused by the injection source has not reached 

equilibrium before measurements are taken.

The concentration distribution cannot be obtained 

from an analysis of the flow system. However, the 

preceding analysis and discussion points out that the 

presently calculated mass transfer transverse to the direction of flow may be due principally to the injection 

scheme used. Because an equilibrium condition has

not been established near the source, the mass transfer 

described by the diffusion equation may give a good 

approximation if the lateral velocity component is included in the dispersion coefficient.

REFERENCES

Jacques, G. L., and Vermeulen, Theodore, 1957, Longitudinal

dispersion in solvent-extraction columns Peclet numbers

for ordered and random packings: Calif. Univ. Radiation

Lab. Kept 8029,152 p. 

Josselin de Jong, G. de 1958. Longitudinal and transverse

diffusion in granular deposits: Am. Geophys. Union Trans.,

v. 39, no. 1, p. 67-74. 

Milne-Thompson, L. M., 1950, Theoretical hydrodynamics: New

York, MaeMillan Co., 600 p. 

Saffman, P. G., 1959, A theory of dispersion in porous media:

Jour. Fluid Mech., v. 6, pt. 3, p. 321-349.
Article 56

EFFECT OF BASE-LEVEL CHANGES ON BEDDING DEVELOPMENT 

IN A LABORATORY FLUME

By ALAN V. JOPLING, Cambridge, Mass.

WorJc done in cooperation with Harvard University

Abstract. In a series of flume experiments, the water discjmrge and sediment input were kept constant while the level 

of the downstream weir was raised at a moderately rapid rate. 

Ripple and dune bed forms were preserved in the aggrading 

deposit, and the bedding developed was either irregular, undtulose, lenticular, trough shaped, or festoonlike.

The polygenetic development of bedding under disequilibrium conditions of sediment transport in a laboratory flume is described qualitatively in this article. 

Specifically, consideration is given to the effect of a 

rising base level on the-development of bedding in a 

rjipple or dune regime. Water discharge and sediment 

i^iput* are kept constant while base level is raised 

either continuously or incrementally by an adjustable 

\jreir at the downstream end of the flume. The generalized observations recorded here apply to sediment 

mixes composed mainly of sand sizes, and they relate 

particularly to the effect of a moderately rapid rise 

of base level.

\ As an initial postulate, consider that conditions of 

equilibrium transport prevail before the base level is 

ifaised. For these conditions there is no net aggradation 

Or degradation. However, a departure from equilibrium caused by a rise of base level results in a zone 

of gradually varied flow. This zone may extend to the 

entrance section of the flume. The transport system 

then adjusts by aggrading the bed toward a new profile 

Of equilibrium which is graded to the raised weir level.

For small departures from equilibrium caused by a 

Slowly rising base level, the mechanics of bedding forijnation generally conforms to the description by Ein- 

^tein and Chien (1953). The movement of ripples or 

(jlunes during the process of aggradation results in a

1 Sediment introduced from a dry-feed hopper; the apparatus has 

been described by the author (1960) in an experimental study on the 

Mechanics of bedding, Harvard Univ. Ph. D. thesis, 358 p.

deposit characterized by a sequence of regular, undulatory, or irregular laminae with sporadic "pockets" of 

ripple or dune cross-lamination. Here the mechanics 

of deposition is related to the operation of form-drag 

processes (Jopling, 1963).

In contrast to this mode of deposition, a large departure from equilibrium triggered by a rapid and relatively large rise of base level initiates a new cycle of 

deposition and culminates in a tabular deposit of aggradational ("deltaic") crossbedding. This type of 

crossbedding is characterized by topset, foreset, and 

bottomset components, and it has been described in 

several laboratory studies on delta building, including 

those of McKee (1957) and Jopling (1962).

For the intermediate case considered in this article, 

the departure from equilibrium is caused by a moderately rapid rise of base level of either a steady or incremental nature. Under such conditions there is a progressive increase in water depth concomitant with a 

decrease in stream velocity. The transport system is in 

a state of disequilibrium, and aggradation is at a maximum in the upstream reach of the channel where the 

backwater effect is least in evidence.

The configuration of the bed is continually remolded 

as successive ripples or dunes override and partly assimilate their predecessors in the depositional sequence 

(fig. 56.1). Some ripples or dunes may even be buried 

intact in this aggrading sequence. For periodic incremental rises of base level the longitudinal profile of the 

bed may divide into several segments or reaches, thereby imparting a steplike appearance to the profile. Each 

segment represents the establishment of a rudimentary 

transport slope more or less commensurate with the 

local flow conditions in the downstream direction.

In summary, therefore, the superposition of ripples 

or dunes results in a complex bedding pattern charART. 56 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B203-B204. 1963.

B203
B204 EXPERIMENTAL HYDROLOGY

Rising water level

Zone of varied and 

unsteady flow Flow direction

FIGURE 56.1. Complex bedding formed by the superposition of 

ripple structures in response to a rising base level.

acterized by the partial preservation of the original 

bed forms. Bedding deposited under these conditions 

of disequilibrium may be undulose, irregular, troughshaped, festoonlike, hummocky, or disconformable, 

and it is characterized by numerous intercalations of 

ripple or dune cross-lamination. From a genetic point 

of view, this type of bedding has both form-drag and

aggradational affinities. Indeed, polygenetic bedding 

deposited in this manner can be regarded as a transitional phase in a genetic series linking bedding of formdrag origin with bedding of topset-foreset-bottomset 

origin.

REFERENCES

Einstein, H. A., and Chien, N, 1953, Transport of sediment mixtures with large ranges of grain sizes: California Univ. 

Inst. Eng. Research, MRD Series, No. 2,40 p.

Jopling, A. V., 1962, Mechanics of small-scale delta formation 

a laboratory study: Natl. Sci. Found, and Office of Naval 

Research Natl. Coastal and Shallow Water Research Conf., 

1st, Baltimore, Md., Tallahassee, Fla., Los Angeles, Calif., 

1961, Proc., p. 291-295.

 1963, Genetic classification of cross-bedding [abs.]: Geol. 

Soc. America Spec. Papers, no. 73, p. 183.

McKee, E. D., 1957, Flume experiments on the production of 

stratification and cross-stratification: Jour. Sed. Petrology, 

v. 27, p. 129-134.
Article 57

FACTORS INFLUENCING THE SURVIVAL OF ESCHERICHIA COLI 

IN DETERGENT SOLUTIONS

By C. H. WAYMAN, J. B. ROBERTSON, and H. G. PAGE, Denver, Colo.

Work done in cooperation with the Federal Housing Administration

Abstract. Coliform bacteria seem to survive longest at low 

temperatures and neutral and alkaline pH in detergent solutions. Phosphate, carboxymethylcellulose, and other constituents of household detergents contribute to bacterial survival. 

Survival of bacteria in low concentrations of ABS is not greater 

than in distilled water, and high concentrations of ABS are 

toxic to bacteria.

Experimental studies on the saturated flow of sewage 

water through soil columns (Page and others, 1963) and 

field determinations of detergents (alkylbenzenesulfonate) and bacteria in ground water (Nichols and 

Koepp, 1961) suggest that soils do not totally remove 

the bacteria of infiltrating sewage.

Because sewage and septic-tank waste water contain 

both coliform bacteria and detergents, a study was made 

to determine the influence of temperature, pH, and certain components in commercial detergents on the survival of EscJiericJda coli, one of the common bacteria 

normally present in human intestines and indicative of 

fecal contamination in water.

The literature lists but few previous studies of bacterial survival. Carlucci and Pramer (1960) indicate 

that Escherichia coll have optimum survival rates in 

acid water (pH 5), in saline water, and in water containing inorganic nutrients, (NH4 ) 2SO4 and (NH4 ) 2 

HPO4 , and organic matter (glucose, peptone, sewage 

volatile solids, and cysteine). Dyadichev (1959) points 

out that the chemical quality of the water, temperature, 

pH, organic nutrients, hydrologic conditions of a 

water basin, and other factors influence survival of 

EscJierichia coli, but he provided no experimental documentation.

Changes in biochemical or physicochemical properties 

of solutions produce consequent changes in properties 

at the bacterial-cell-membrane-solution interface. 

These changes can promote growth or cause death to

bacterial colonies. In this study, changes in bacterial 

reactivity were measured indirectly by determining the 

length of time a specific number of bacteria survived 

at varying temperatures, pH, and concentrations of 

detergent component.

Bacterial solutions were prepared by adding pure 

cultures of bacteria (E schericMa coli) , sodium hydroxide or hydrochloric acid (for adjustment of pH), and 

detergent components to sterilized, distilled and deionized water. Bacteria were added to different solutions at temperatures to cover the range usually found 

in ground water, and at selected pH to include the acid 

through basic range as shown in the table below.

Concentrations, temperatures, and pH of solutions investigated

Solution

Distilled water _

10 ppm ABS i_ - _ -

100 ppm ABS____

1,000 ppm ABS.._

5 ppm ABS plus

10 ppm PO4 -_-

10 ppm ABS plus

10 ppm PO4 ----_-

10 ppm ABS plus

10 ppm PO4 plus

1 ppm CMC *.._.-

12.5 ppm commercial detergent.. _

10°C

X

X

X

X

X

pH4

20°C

X

X

X

X

X

35°0

X

X

X

X

X

10°C

X

X

X

X

X

X

X

X

pH7

20°C

X

X

X

X

X

X

X

X

35°C

X

X

X

X

X

X

X

X

10°C

X

X

X

X

X

pHIO

20°C

X

X

X

X

X

35° C

X

X

X

X

X

1 ABS=alkylbenzenesulfonate.

2 CMC=carboxymethylcellulose.

The prepared solutions were placed in a constanttemperature water bath (±0.1°C) and were sampled 

periodically by a sterile technique. One milliliter of 

solution sampled in duplicate was removed from each 

of the preparations and mixed with McConkey's agar.

ART. 57 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B205-B208. 1963.

680231 O - 63 - 15 B205
B206 EXPERIMENTAL HYDROLOGY

The agar plates were incubated at 37°C±1°C for 72 

hours, and the number of bacteria remaining in the 

sampled solution was determined by the plate-count 

method (American Public Health Assoc. and others, 

1960). Two experimental difficulties that may have influenced these studies are (1) the inability to place 

identical starting amounts of bacteria in each test solution and (2) the possible variation in the age and 

virility of the different stocks of bacteria used. Either 

of these variables might influence the bacterial survival 

rate, but what that influence is, if any, is unknown to 

the authors.

Figure 57.1 shows the survival time of bacteria in 

solutions of varying pH and ABS concentration at 

10°C. Bacteria apparently can survive for many days 

in distilled water at low temperature without any 

available nutrient. In solutions containing 10 ppm of 

ABS, bacteria survive longer at a pH of 7 than at either 

pH 4 or 10. In alkaline media (pH 10) bacteria initially 

seem to experience a certain degree of shock and many 

die off rapidly; however, the surviving bacteria apparently recover and multiply temporarily. Neutral 

solutions (pH 7) containing as much as 100 ppm of 

ABS are not toxic to bacteria, whereas solutions containing 1,000 ppm are toxic. The rapid death rate at 

the higher concentration of ABS may be attributed to 

the high degree of surface activity at the interface between the bacterial membrane and the solution.

5000

0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 

TIME, IN DAYS

FIGUEE 57.1 Survival time of bacteria in solutions of varying 

pH and concentration of ABS, at 10°C.

Changes in permeability, caused by the high degree of 

surface activity at the interface, probably alter normal 

metabolic processes of the bacteria.

Figure 57.2 shows survival time of bacteria in solutions of varying pH and ABS concentration at 20°C. 

In solutions containing 10 ppm of ABS, maximum survival was at pH 7 and pH 10. At pH 7, solutions containing either 100 or 1,000 ppm of ABS were highly 

toxic to bacteria; however, the toxicity is less pronounced than that in solutions containing 10 ppm of 

ABS at a pH of 4.

Figure 57.3 shows the survival time of bacteria in 

solutions of varying pH and ABS concentration at 

35°C. Bacteria do not persist at 35°C regardless of 

pH or ABS concentration.

Figure 57.4 shows survival time of bacteria in solutions containing both ABS and phosphate at 10°C. At 

this temperature, phosphate acts as an inorganic nutrient for bacteria through a pH range of 4 to 10, although 

a pH of 4 seems most detrimental to survival. Similar 

experiments at 20°C and 35°C indicated a maximum 

survival time of 4 days for bacteria at a pH of 7.

Figure 57.5 shows the survival time of bacteria in 

solutions containing 10 ppm of ABS, 10 ppm of phosphate, and 1 ppm of CMC (carboxymethylcellulose). 

Maximum survival times are in solutions of pH 7 and 

pH 10 at both 10°C and 20°C. The small amount of 

CMC seems to be a very effective nutrient. Low pH 

and high temperature again seem most detrimental to

1000

500

100

j 50

0 1 2 3 45 6 7 8 9 10 11 12 13 14 15 

TIME, IN DAYS

FIGURE 57.2. Survival time of bacteria in solutions of .varying 

pH and concentration of ABS, at 20°C.
WAYMAN, ROBERTSON, AND PAGE B207

5000

1000

500

100

10

TIME, IN DAYS

FIGURE 57.3. Survival time of bacteria in solutions of varying 

i pH and concentration of ABS, at 35 °C. Curve 1, 1,000 ppm 

[ ABS, pH=7; 2, 10 ppm ABS, pH=10; 3, 10 ppm ABS, pH=4;

4, 10 ppm ABS, pH=7; 5, 100 ppm ABS, pH=7; 6, distilled

water.

urvival. The bacterial revival indicated by the curve 

si 10°C and pH 10 probably would have receded within

, few days (similar to the effect shown in figs. 57.1 and

7.2) if the test had been continued. 

Figure 57.6 shows the survival time of bacteria in

olutions containing 12.5 ppm of a commercial household detergent (approximately 40 percent ABS plus 

ijmknown builders). In general, the optimum conditions for bacterial survival were in solutions at a pH of

20,000

3 4 7 8 9 10 11 12 13 14 15 16 17 

TIME, IN DAYS

FIGURE 57.4. Survival time of bacteria in solutions of varying 

pH and concentration of ABS and 10 ppm of PO*, at 10°C. 

Solid line, 5 ppm of ABS plus 10 ppm of PO*; dashed line, 

10 ppm of ABS plus 10 ppm of PO*.

7, regardless of temperature, or in alkaline media (pH 

10) at 20°C. At low pH and at 35°C and pH 10, the 

solutions were highly toxic to bacteria.

Bacteria not removed from infiltrating water by soils 

constitute a potential pollutant in ground water. This 

study has indicated that conditions contributing to survival of coliform bacteria in detergent solutions are:

1. Low temperature of infiltrating solution (less than 

20°C).

2. Neutral pH (about 7) and in some solutions alkaline 

pH (about 10).

3. Phosphate, carboxymethylcellulose, and other constituents of commercial detergents.

In general, bacteria seem to disappear logarithmically with respect to time. Low concentrations of 

ABS (10 ppm or less) do not prolong growth any more 

than distilled water. High concentrations of ABS (100 

to 1,000 ppm) are toxic to coliform bacteria; the toxicity, which kills the bacteria, probably is due to changes 

in permeability of the bacterial membrane.
B208 EXPERIMENTAL HYDROLOGY

8,9 10 11 12 13 14

FIGURE 57.5. Survival time of bacteria in solutions containing 

10 ppm of ABS, 10 ppm of POi, and 1 ppm of carboxymethylcellulose, at varying pH and temperature. Solid line, pH 10; 

short dash, pH 7; long dash, pH 4.

Studies are now in progress to determine the survival 

rates of coliform and other types of sewage bacteria 

in solutions containing either a branched-chain ABS or 

a straight-chain ABS.

REFERENCES

American Public Health Association and others, 1960, Standard 

methods for the examination of water and wastewater: 

llth ed., New York, Am. Public Health Assoc., Inc., p. 626.

5000

56789 

TIME, IN DAYS

10 11 12 13 14

FIGURE 57.6. Survival time of bacteria in solutions containing 12.5 ppm of household detergent, at varying pH and temperature. Solid line, pH 10; short dash, pH 7; long dash, 

pH4.

Carlucci, A. F., and Pramer, D., 1960, An evaluation of factors 

affecting the survival of Escherichia coli in sea water: 

Applied Microbiology, v. 8, p. 243-254.

Dyadichev, N. R., 1959, Survival and proliferation of typhoid, 

paratyphoid, and dysenteric bacteria in water: Gigiene i 

Sanitariya, v. 24, p. 11-15.

Nichols, M. S., and Koepp, E., 1961, Synthetic detergents as a 

criterion of Wisconsin ground-water pollution: Am. Water 

Works Assoc. Jour., v. 53, p. 303-306.

Page, H. G., Wayman, C. H., and Robertson, J. B., 1963, Behavior of detergents (ABS), bacteria, and dissolved solids 

in water-saturated soils : Art. 237 in U.S. Geol. Survey Prof, 

paper 450-E, p. E179-E181.
Article 58

EFFECT OF DETERGENTS ON THE VISCOSITY OF WATER 

CONTAINING BACTERIA AND CLAY IN SUSPENSION

By C. H. WAYMAN, H. G. PAGE, and J. B. ROBERTSON, Denver, .Colo.

Work done in cooperation with the Federal Housing Administration

Abstract. Sewage and household waste water containing detergents but not suspended solids probably travel at the same 

rate as natural water for a given temperature. Addition of 

detergent to water containing clay generally results in a reduction of viscosity, although at high pH the viscosity generally 

increases.

Experiments with detergent solutions have suggested 

that detergent components are very important in flocculating and dispersing clays and therefore have a direct 

relation to the viscosity of clay-water suspensions. 

Viscosity in turn is a controlling factor in the velocity 

of the flow of fluids. Thus, knowledge of the effect of 

tl le addition of detergents to clay-water suspensions can 

ba important in the fields of open-channel flow and sedimentation, flow in pipes, underground flow through 

fractures and solution openings, and in drilling-mud 

technology.

In this study the viscosity of solutions containing detergents (ankmic, cationic, or nonionic), detergent 

builders (such as polyphosphates, silicates, carbonates, 

silfates, and carboxymethylcellulose), bacteria (Esc^ierichia coli}^ and suspended solids (kaolinite, illite, 

or montmorillonite) was determined with the use of an 

Qstwald type viscometer at various concentrations, temperatures, and pH values.

The effect of temperature on viscosity is illustrated 

by viscosity measurements of solutions (without suspended solids) containing as much as 500 ppm (parts 

per million) of anionic, cationic, or nonionic detergents. 

IVfeasurements indicate that the solutions move only 

about two-thirds as fast at 5°C as at 20°C. Neither 

changes in the concentration of detergent and in pH 

(3 to 10) nor as much as 1X108 bacteria (EscJierichia. 

coli] per ml (milliliter) affect the viscosity of these 

solutions. The viscosity of solutions containing as

much as 500 ppm of anionic, cationic, or nonionic detergents at 20°0 and a pH of 6.8 is about equal to the 

viscosity of pure water for the same conditions of temperature and pH. These results indicate that dissolved 

solids and bacteria have little influence on the viscosity 

of solutions for the conditions specified above.

To determine the effect of suspended and dissolved 

solids on viscosity, three series of clay suspensions were 

prepared using only the <2-micron size fraction of three 

API (American Petroleum Institute) clays: montmorillonite (No. 26), illite (No. 35), and kaolinite (No. 2). 

The concentration of clay for the illite and the kaolinite 

series ranged from 0.0 to 4.0 g (grams) per 100 ml of 

deionized water (pH 7) and for the montmorillonite 

series from 0.0 to 3.4 g per 100 ml of deionized water. 

The viscosity of each suspension was measured at 20° C. 

Each series was repeated, but either 10 ppm of the 

anionic detergent, alkylbenzenesulfonate (ABS), or 10 

ppm of commercial detergent (Tide) was added to each 

clay suspension. This commercial detergent consists 

of about 40 percent ABS and 60 percent unknown detergent builders.

Figure 58.1 shows the relation between viscosity and 

concentration of clay and detergent components for the 

3 API clays at pH 7 and 20°C. Neither as much as 

4.0 g of kaolinite per 100 ml of deionized water nor 10 

ppm of ABS or other components in commercialdetergent preparations significantly affects the viscosity 

of the kaolinite suspensions. The viscosity of illite and 

montmorillonite suspensions begins to increase at concentrations of 1.3 and 0.1 g of clay per 100 ml of deionized water, respectively; components in commercial 

detergents tend to cause a slight decrease in the viscosity 

of these suspensions. Similar families of curves have 

been obtained for montmorillonite, attapulgite, and

ART. 58 IN U.S. GBOL. SURVEY PROF. PAPER 475-B, PAGES B209-B212. 1963.

B209
B210 EXPERIMENTAL HYDROLOGY

EXPLANATION

Pure-clay suspension

10 ppm of ABS added 

10 ppm of Tide added

0.5 1.0 1.5 2.0 2.5 3.0 

CLAY CONCENTRATION, IN GRAMS PER 100 MILLILITERS

3.5

FIGURE 58.1. Relation of viscosity to clay concentration and .commercial-detergent components, at 20°C, and pH 7.

kaolinite suspensions in solutions of Calgon, sodium 

chloride, and hydrochloric acid (Searle and Grimshaw, 

1959).

Changes in pH and temperature have marked effects 

upon the viscosity of clay suspensions. These effects 

were studied on pure clay suspensions consisting of 1.5 g 

of clay per 100 ml of deionized water. The pure clay 

suspensions were compared to clay suspensions containing 100 ppm of ABS or 200 ppm of Tide. The pH of 

the suspensions was adjusted by adding hydrochloric 

acid or sodium hydroxide; the observed changes in viscosity with changes in pH and temperature are shown 

in figures 58.2 to 58.4.

Figure 58.2 shows the effect of pH changes and of 

addition of detergents on the viscosity of kaolinite suspensions. At 10° C, a maximum viscosity for a pure 

kaolinite suspension occurs at a pH of about 4. Between pH 4 and pH 7, ABS and other components in 

detergent formulations cause a marked decrease in the

viscosity. This decrease is less intense at 20°C than at 

10°C. In acid suspensions of kaolinite, ABS and other 

ions seems to neutralize positive charges on the clay surface and thus to deflocculate (to decrease the viscosity 

of) the suspensions.

Figure 58.3 shows the effect of similar treatment of 

illite suspensions. At 10°C, maximum viscosity of the 

illite suspension is at pH 4 and pH 10; a minimum 

viscosity is at pH 7. Addition of detergent ions tends 

to flocculate (to increase the viscosity'of) this clay suspension through a wide range of pH. At 20°C, viscosity 

for similar tests is less than at 10°C, and the viscosity 

varies less as the pH changes than at 10°C.

Figure 58.4 shows the effect of similar treatment of 

montmorillonite suspensions. At 20°C, a minimum 

viscosity occurs at pH 7; detergent components 

markedly reduce the viscosity between pH 4 and pH 7. 

At 10°C, the reverse occurs; that is, the detergent-components increase the viscosity or flocculate the suspen-
WAYMAN, PAGE, AND ROBEETSON B211

FIGUBE 58.2. Relation of the viscosity of kaolinite suspensions 

to pH and commercial-detergent components, at 10°C and 

20°C.

FIGUBE 58.3. Relation of viscosity of illite suspensions to pH 

1 and commercial-detergent components, at 10°C and 20°C.

Sion. Earlier studies on purified montmorillonite 

[(Street, 1958) indicated the general trend observed at 

10°C, but not at 20°C.

4.40-

4.00 -

3.60 -

g 3.20 rt 2.80 -

2.40 -

2.00 -

Montmorillonite + 200 ppm Tide 10°c 

(read right scale)

FIGUEE 58.4. Relation of viscosity of montmorillonite suspensions to pH and commercial-detergent components, at 

10°C and 20°C.

The results of the viscosity tests on montmorillonite 

suspensions are difficult to understand. However, previous investigators have indicated that the viscosity of 

clay suspensions can be explained in terms of electrokinetic phenomena with use of the Gouy-Freundlich 

diffuse double-layer concept and van der Waals attractive forces at the clay particle-solution interface. 

The concentration of the clay suspension (Coughanour 

and Norton, 1949), surface area of the clay (Johnson 

and Lawrence, 1942), and electrolytes in solution (Hauser and Hirshon, 1939; Speil, 1940; Johnson and Norton, 1941; and Kingery, 1951) influence electrical 

properties at the clay particle-solution interface and 

produce changes in the viscosity of clay suspensions. 

The role played by surface tension alone, neglecting 

ionic effects, is of dubious importance in the viscosity 

of clay suspensions. Schwartz (1952) and Kingery and 

Francl (1954) suggest that wetting agents (detergents 

and surfactants) decrease the viscosity of clay suspensions. However, a detailed study of wetting agents in 

clay suspensions (Ormsby, 1960) indicated that surface tension seems to have only a minor effect upon the 

viscosity of clay suspensions. Several variables, some 

of which could not be considered in this study, influence 

the viscosity of clay suspensions. Apparently, a greater 

number of these variables affected the viscosity of montmorillonite suspensions to a greater extent than they 

affected the viscosity of suspensions of either kaolinite 

or illite.
B212 EXPERIMENTAL HYDROLOGY

Commercial-detergent components in kaolinite or illite clay suspensions slightly decrease and slightly increase the respective viscosities of these two clay 

suspensions as compared to the viscosity of water at 

the same temperature. At 10° C and pH 7, commercialdetergent components in montmorillonite suspensions 

can increase the viscosity to as much as 25 times that of 

water. This study has shown that the viscosity of clay 

suspensions depends on the composition of the clay, pHy 

temperature, and electrolytes (builders) in detergent 

formulations.

Because sewage and household waste water containing detergents, but no suspended solids, have a viscosity 

similar to that of pure water, these solutions probably 

travel at the same rate as water, for a given temperature. However, waste water containing clay in suspension, especially montmorillonite, can be as much as 

25 times less mobile than similar water containing no 

suspended solids, for the conditions mentioned in this 

article. Thus, the flow of clay-water suspensions can 

be greatly altered by the presence of detergent components which may have significant influence in many 

areas of fluid-flow studies.

REFERENCES

Coughanour, L. W., and Norton, F. H., 1949, Influence of 

particle shape on properties of suspensions, pt. 9 of Fundamental study of clay: Am. Ceramic Soc. Jour., v. 32, p. 

129-132.

Hauser, E. A., and Hirshon, S., 1939, Behavior of colloidal suspensions with electrolytes: Phys. Chemistry Jour., v. 43, 

p. 1015-1036.

Johnson, A. L., and Lawrence, W. G., 1942, Surface area and its 

effect on exchange capacity of kaolinite, pt. 4 of Fundamental study of clay: Am. Ceramic Soc. Jour., v. 25, p. 

344-346.

Johnson, A. L., and Norton, F. H., 1941, Mechanism of deflocculation of the clay-water system, pt. 2 of Fundamental 

study of clay: Am. Ceramic Soc. Jour., v. 24, p. 189-203.

Kingery, W. D., 1951, Note on effect of fluoride ions on a clay 

suspension, pt. 11 of Fundamental study of clay: Am. 

Ceramic Soc. Jour., v. 34, p. 242-244.

Kingery, W. D., and Francl, J., 1954, Drying behavior and 

plastic properties, pt. 13 of Fundamental study of clay: Am. 

Ceramic Soc. Jour., v. 37, p. 596-602.

Ormsby, W. C., 1960, The role of surface tension in determining 

certain clay-water properties: Ceramic Bull., v. 39, p. 

408-^12.

Schwartz, B., 1952, Note on effect of surface tension of water 

on plasticity of clay, pt. 12 of Fundamental study of clay: 

Am. Ceramic Soc. Jour., v. 35, p. 41-43.

Searle, A. B., and Grimshaw, R. W., 1959, The chemistry and 

physics of clays and other ceramic materials: New York, 

Interscience Publishers, Inc., p. 458 162.

Speil, S., 1940, Effect of adsorbed electrolytes on properties of 

monodlsperse clay-water systems: Am. Ceramic Soc. Jour., 

v. 23, p. 33^38.

Street, N., 1958, Viscosity of clay suspensions : World Oil, v. 147, 

p. 151-156.
Article 59

ADSORPTION OF THE SURFACTANT ABS35 ON MONTMORILLONITE

By C. H. WAYMAN, J. B. ROBERTSON, and H. G. PAGE, Denver, Colo.

Work done in cooperation with the Federal Housing Administration

Abstract. A radiochemical-tracer technique was used to study 

the adsorption of ABS (alkylbenzenesulfonate) on montmorillonite. This adsorption is influenced by length of the alkyl 

chain, by pH, by phosphate in solution, and by amount and type 

of ionic salts in solution.

1 Incompletely degradable surfactants injected into 

the ground in household waste water have in many 

areas polluted ground-water supplies. One of the nattural means by which these surfactants are removed 

from waste water is by their adsorption on soil minerals. A previous paper (Wayman and others, 1962) described the adsorption of alkylbenzenesulfonate tagged 

with S35 (ABS35 ), an essential incompletely degradable 

surfactant in detergents, on kaolinite. A similar study 

j in which montmorillonite was substituted for kaolinite 

1 is described in this article.

A radiochemical-tracer technique identical to that 

used in the kaolinite study (Wayman and others, 1962) 

was used in the montmorillonite study. The effects of 

such variables as alkyl chain length (Ci 2 and C15 ), pH, 

phosphate ion, and type and amount of dissolved ionic 

salts (NaCl, OaCl2 , or A1C13 ) on ABS adsorption by 

m'ontmorillonite were studied. ABS can be represented 

in solution as

where K represents the alkyl chain.

The preparation of the montmorillonite and the 

experimental procedures used were identical to the 

procedures used in the study of kaolinite. The montmorillonite used was the standard clay mineral (No. 

26) of the American Petroleum Institute. The composition of the mineral can be represented by

(Si3. 91A10. 09 0. i 6Mgo.33)O10 (OH)2 .

Nan on ( -=-

Microscopic examination indicates that it contains 

about 7.0 percent quartz and feldspar (orthoclase and 

albite) and 2.0 percent glauconite. Its cation-exchange 

capacity is 99.0 milliequivalents per 100 grams, and 

the pH of a suspension of 200 milligrams of the clay 

in 35 ml (milliliters) of distilled water is about 7.5.

Figure 59.1 shows that both pH and length of the alkyl chain significantly influence the adsorption of ABS 

from solutions containing 4 ppm (parts per million) of 

the detergent. The adsorption of dodecyl (C12 ) ABS 

is insignificant at a pH of 4,7, or 10. Pentadecyl (C15 ) 

ABS is adsorbed readily at pH 4 and 7 and to a lesser 

extent at pH 10; however, maximum adsorption at pH 

7 seems to occur in about 7 days and desorption occurs 

therafter. After 14 days, adsorption at pH 7 is only 

slightly greater than adsorption at pH 10.

The dependence of adsorption on molecular weight, 

as found in this study, has been observed by others. 

McAtee (1959) indicated that sodium can be replaced 

in montmorillonite more easily by large quaternary

200 

(/) UJ

S t 150

§3O -I

ggioo

ABS IV ADSORBED, IN 

GRAM OFMONTM PER , 

8 Sgo

.

^^_ ^^^_-_-_ o- ,

: f / ^^ ^

: / /^ ^~~ ^

I / ' a _____ - "

X 4ppm of pentadecyl ABS & PH 7

! ' 4ppm of dodecyl ABS Q pM ~ 1U -

D 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 1

.

30 |-

z

UJ

O

20 UJ

' °K1

° CO 

0 

ABSDSORBED, IN

TIME, IN DAYS

FIGURE 59.1. Influence of pH, length of alkyl chain, and time

on adsorption of ABS by montmorillonite.

ART. 59 IN U.S. GEOL. SURVEY PROF. PAPER 475-B, PAGES B213-B216. 1963.

B213
B214 EXPERIMENTAL HYDROLOGY

compounds than by substances of lower molecular 

weight such as primary amines. Also, Greenland and 

others (1962) showed that peptide adsorption on montmorillonite increases as the molecular weight of the 

peptide increases. Similarly, Smith and Bader (1961) 

demonstrated that the association of amino acids, dibasic acids, and carbohydrates with montmorillonite 

depends largely upon their molecular weight. Thus, 

these data indicate that the free energy of adsorption, 

including physical adsorption, increases as the molecular weight of the adsorbate increases. The enhanced 

adsorption of anionic ABS on montmorillonite at low 

pH can probably be attributed to the development of 

an increased number of positive sites on the clay surface. Other investigators indicate that anionic adsorption on soil minerals is dependent on pH (Schofield, 

1949; Schofield and Samson, 1953; Street, 1956; and 

Holtzman, 1962). Significantly, both dodecyl ABS 

and pentadecyl ABS seem to be adsorbed negatively 

during the initial equilibration period. This negative 

adsorption can be attributed primarily to the high 

water-sorption capacity of montmorillonite, as indicated by White and Pichler (1959), and to the disparities of ionic distribution in the double layer by 

using the Donnan equilibrium or the Gouy theory.

Complex phosphates such as tetrasodium pyrophosphate (Na4P2O7) or pentasodium tripolyphosphate 

(Na5P3O10) are essential components in detergents. 

In waste water these phosphates are probably biologically degraded (Cohen, 1959) or hydrolyzed (Morgan 

and Engelbrecht, 1960) to orthophosphates. The 

specific anionic form of orthophosphate in water 

depends upon the pH of the water. The ionizations 

of H3PO4 have been reviewed by Pitzer (1937) and 

listed as

H3PO4 =H+1 +H2POr 1

HPQ4-2

(1)

#2 =6.2X10- 8 (2)

#3 =10- 12 (3)

At pH 4, H2PO4~1 predominates, but the solution 

contains a minor amount of H3PO4 ; at pH 7, mixtures 

of HPO4~2 and HaPOr1 are in solution. At pH 10, 

HPO4~2 predominates, but the solution contains a 

minor amount of PO4~3 . Because these anionic species 

may compete with ABS for adsorption by montmorillonite, solutions containing orthophosphate were used 

in the study of ABS adsorption by montmorillonite.

Figure 59.2 shows that pH influences the amount of 

ABS that montmorillonite adsorbs from a solution containing 5 ppm of dodecyl ABS and 10 ppm of orthophosphate. In neutral (pH 7) and in alkaline (pH

80

70

60

50

40

30

20

10

0

-10

-20

-30

0 2 8 10 12 14 16 18 20 22 24 26 28 30 32 

TIME, IN DAYS

-40

FIGURE 59.2. Influence of pH and time on ABS adsorption by 

montmorillonite. Test solutions contained 5 ppm of dodecyl 

ABS and 10 ppm of orthophosphate.

10) solutions, small amounts of ABS are negatively 

adsorbed for as long as about 2 weeks, at which time 

positive adsorption begins. In acid solutions (pH 4) 

the amount of ABS adsorbed initially increases with 

time; then, after about 9 days, desorption occurs for 

about 1 week, after which apparent equilibrium is 

achieved. Because the amount of ABS adsorbed from 

solutions containing both ABS and phosphate ions is 

greater than that adsorbed from solutions containing 

only ABS (figs. 59.1, 59.2), phosphate fixation probably 

occurs on the clay either before or simultaneously with 

ABS adsorption. Some investigators (Coleman, 1944; 

Chatterjee and Datta, 1951) suggest that this fixation 

results from active iron and aluminum oxides in montmorillonite. Others (Murphy, 1939; Scarseth, 1935; 

Stout, 1939) believe that phosphate ions exchange with 

hydroxyl ions at the edge of the crystal lattice. Ravikovitch (1934) proposed that phosphate becomes fixed 

by an H2PO4-Ca-clay linkage. Low and Black (1948) 

attribute phosphate fixation on kaolinite to aluminum 

released by decomposition of the clay. A detailed study 

on montmorillonite (Ellis and Truog, 1955) indicates 

that phosphate becomes fixed by the formation of calcium phosphate complexes. Aluminum and phosphate 

in dilute solution at low pH may form complex cations 

such as A1HPO/S A1H2PO/2 , and A1(H2P(X) 2+1 (Salmon and Wall, 1958). Because ABS uptake on montmorillonite increases with phosphate in solution at pH 

4 (fig. 59.2), phosphate probably does not compete with 

ABS for adsorption sites. In this study the heteroionic 

nature of the clay surface could have contributed 

aluminum to the solution for the formation of an 

aluminum phosphate ion; these large ions might have 

combined with ABS to form an aluminum-phosphateABS complex or precipitate, either of which could
WAYMAN, ROBERTSON, AND PAGE B215

lave been adsorbed by the clay. In addition to this 

mechanism, any of the above mechanisms might be used 

;o explain phosphate fixation on montmorillonite be- 

:ore or simultaneously with ABS adsorption.

Figure 59.3 shows the influence of salt concentration, 

valence, and pH on the adsorption of ABS by montmorillonite. Initially ABS is adsorbed on montmorilLonite to a greater extent when salts are in solution. 

Solutions containing highly charged ionic salts (for example, CaCl2 and A1C13 ) at low pH seem to offer optimum conditions for adsorption of ABS. The plotted 

points represent equilibrium adsorption values ranging 

in time from 6 to 10 days. At increasing salt concentrations in the range from 0 to about 1,000 ppm, ABS adsorption increases ; at increasing salt concentrations of 

more than 1,000 ppm (pH 4), ABS adsorption decreases. No single mechanism can be used to account for 

increased ABS adsorption in acid solutions at high salt 

concentrations (to 1,000 ppm). The general trend of 

enhanced adsorption as the salt concentration increases 

might be explained by a lowering of the critical micelle 

concentration. Micellar adsorption should exceed ionic 

adsorption because of increased surface activity of the 

.micelles. Figure 59.3 indicates that greater amounts of 

monovalent salts are required to lower the critical micelle concentration than of divalent or trivalent salts; 

this observation may be attributed to an effect of the 

Schulze-Hardy rule. The combined effects of aluminum hydrolysis and complex formation might explain 

the increased adsorption of ABS in acid solution containing ABS and aluminum ions. Below pH 5, Al+3 

predominates in solution (Miller, 1925). In acid solution one possible hydrolysis scheme (Brosset and others, 

1954) is

6Al+3 + 15H2O=Al6 (OH)+315 + 15H+1 . (4)

Other possible complexes are A1(OH) +2 , A12 (OH) 2+4 or 

A13 (OHV3 . The A13 (OH) 6+3 polymeric complex was 

observed by Ruff and Tyree (1958) . Where clay or soil 

is in solution with AH, hydrolysis becomes more intense 

(Ragland and Coleman, 1960). If the large complex 

can be formed in solution by equation 4 and adsorbed 

on montmorillonite, then the number of positive sites 

can be effectively increased. Hence, ABS could neutralize these positive sites by adsorption in formation of 

a complex such as A16 (OH) i5+3-ABSx. Similar aluminum-hydroxy-ABS complexing has been suggested for 

acid ABS-alum systems ( Sedlander and Gates x ) . The 

lesser adsorption of ABS from acid solutions contain- 

: ing 2,000 ppm of salt than from acid solutions containtaing 1,000 ppm of salt probably is due to the precipi1 Sedlander, N. R., and Gates, C. D., 1962, Effect of alkyl benzene 

sulfonate surfactant on alum floe : Paper presented at Am. Chem. Soc. 

Mtg., Atlantic City, N.J., Sept. 9-14. 1963, p. 1-9.

MOLAR CONCENTRATION OF AICI 3 

0.002 0.004 0.006 0.008 0.010 0.012 0.014

-100

0 250 500 750 1000 1250 1500 1750 2000' 

CONCENTRATION OF SALT, IN PARTS PER MILLION

FIGURE 59.3. Influence of salt concentration, valence, and pH 

on adsorption of ABS by montmorillonite. Test solutions 

contained 5 ppm of dodecyl ABS.

tation of huge micellar aggregates without adsorption 

or inability of these aggregates to diffuse across the 

sausage membrane (Wayman and others, 1962). Unfortunately, data are not available for similar analogs 

of Ca+2 and Na+1 .

This study indicates that the adsorption of ABS on 

montmorillonite depends on the length of the alkyl 

chain and pH. Montmorillonite clay adsorbs more pentadecyl ABS per gram than dodecyl ABS and more 

ABS from acid or neutral solutions than from alkaline 

solution. Initially, adsorption of ABS from a solution containing both ABS and phosphate ions is greater 

than that from a solution containing only ABS; however, some desorption of ABS occurs at a later time. 

Acid solutions (pH 4) containing divalent or trivalent 

salts in concentrations of 1,000 ppm produce optimum 

conditions for ABS adsorption; for salt concentrations 

exceeding this value, adsorption of ABS decreases.

REFERENCES

Brosset, C., Biedermann, G., and Sillen, L., 1954, Studies on the 

hydrolysis of metal ions, pt. 11. The aluminum ion, Al+3 : 

Acta Chemica Scandinavica, v. 8, p. 1917-1926.

Chatter.iee, B., and Datta, S., 1951, Phosphate fixation by clay 

minerals Montmorillonite and kaolinite: Soil Sci. Jour., 

v. 2, p. 224-233.

Cohen, J. M., 1959, Syndets in water supplies: Soap and Chem. 

Specialties, v. 35, p. 9, 53.

Coleman, R., 1944, Phosphorus fixation by the coarse and fine 

clay fractions of kaolinitic and montmorillonitic clays: Soil 

Sci., v. 58, p. 71-77.

Ellis, B., and Truog, E., 1955, Phosphate fixation by montmorillonite: Soil Sci. Soc. America Proc., v. 19, p. 451^54.

Greenland, D. J., Laby, B. H., and Quirk, J. P., 1962, Adsorption 

. of glycine and its di- , tri- , and tetra-peptides by montmorillonite : Faraday Soc Trans., v. 58, p. 829-841.
B216 EXPERIMENTAL HYDROLOGY

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.

Low, P. F., and Black, C. A., 1948, Phosphate induced decomposition of kaolinite: Soil Sci. Soc. America Proc., v. 12, 

p. 180-184.

McAtee, J. L., 1959, Inorganic-organic cation exchange on montmorillonite: Am. Mineralogist, v. 44, p. 1230-1236.

Miller, L. B., 1925, A study of the effects of anions upon the 

properties of alum floe: U.S. Public Health Repts., v. 40, 

pt. 1, p. 351.

Morgan, J. J., and Engelbrecht, R. S., 1960, Survey of phosphates and ABS concentration in Illinois streams: Am. 

Water AVorks Assoc. Jour., v. 52, p. 471-482.

Murphy, H. F., 1939, The role of kaolinite in phosphorus fixation : Hilgardia, v. 12, p. 343-382.

Pitzer, K. S., 1937, The heats of ionization of water, ammonium 

hydroxide, carbonic, phosphoric, and sulphuric acids 

The variation of ionization constants with temperature and 

entropy change with ionization: Am. Chem. Soc. Jour., v. 

59, p. 2365-2371.

Ragland, J. L., and Coleman, N. T., 1960, The hydrolysis of 

aluminum salts in clay and soil systems: Soil Sci. Soc. 

America Proc., v. 24, p. 457-460.

Ravikovitch, S., 1934, Anion exchange; pt. 2, Liberation of the 

phosphoric acid ion adsorbed by soils: Soil Sci., v. 38, p. 

279-290.

Ruff, J. K., and Tyree, S. Y., 1958, Light-scattering studies on 

aqueous aluminum nitrate solutions: Am. Chem. Soc. Jour., 

v. 80, p. 1523-1526.

Salmon, J. E., and Wall, J. G. L., 1958, Aluminum phosphates: 

Chem. Soc. Jour., v. 22, p. 1128-1134.

Scarseth, G. D, 1935, The mechanism of phosphate retention by 

natural alumino silicate colloids: Am. Soc. Agronomy 

Jour., v. 27, p. 596-616.

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.

Smith, J. B., and Bader, R. G., 1961, Preliminary investigations 

of the association of organic material and carbon dioxide 

with sedimentary particles: Texas Agr. and Mech. Coll. 

Proj. 142, Reference 61-8T, Tech. Rept. 3, p. 1-221.

Stout, P. R., 1939, Alteration in the crystal structure of clay 

minerals as a result of phosphate fixation: Soil Sci. Soc. 

America Proc., v. 4, p. 177-182.

Street, N, 1956, The rheology of kaolinite suspensions: Australian Chem. Jour., v. 9, p. 467.

Wayman, C. H., Robertson, J. B., and Page, H. G., 1962, Adsorption of the surfactant ABS35 on kaolinite: Art. 238 in U.S. 

Geol. Survey Prof. Paper 450-E, p. E181-E183.

White, W. A., and Pichler, E., 1959, Water-sorption characteristics of clay minerals: Illinois Geol. Survey Circ. 266, p. 

1-20.
SUBJECT INDEX

[For major topic headings such as "Economic geology," "Engineering geology," "Stratigraphy," see under State names or refer to table of contents]

Page

AB8. See Surfactants.

Al Rose Formation, California, definition.._ B79

Alaska, geochronology, Talkeetna Mountains- 56

geomorphology, Baranof Island-_____ 110

glaciation, Shumagin Islands..______ 106

glaciology, Tanana Valley.________ 144

ground water, northwestern. _______ 183

petrology, Shumagin Islands.______ 106

stratigraphy, Shumagin Islands_ ____ 106

structural geology, Baranof Island....__ 110

surface water, waterpower investigations. 176

Antarctica, glacial geology and glaciology,

| Thiel Mountains._________ 140

Aquifers, calculating T/S....,________ 196

effect of earthquakes..__________ 131

effect of flood waves. ___________ 192

} injection of tracers_____________ 199

salt-water intrusion_____________ 182

Arizona, petrology, Colorado Plateau..___ 48

Arsenic, in basalt._______________ 20

B

Bacteria, survival in detergent solutions___ 205 

badger Flat Limestone, California, definition. 80 

parite, bedded deposits .___._____. 38 

Basalt, arsenic content. ____________ 20 

submarine, pillow structures_______ 153 

(Base level, effect of changes on bedding

development..__________ 203 

(Bauxite, world reserves..___________ 158 

Bedding development, effect of base-level

changes .. _____..... 203 

Bermeja Complex, Puerto Rico, stratigraphy

and structural geology._____ 114 

Beryllium, in mineralized tuft________ 16 

in silicic volcanic glass___________ 18 

relation to fluorine content.._______ 16,18 

Billman Creek Formation, Montana, deflni- 

; tion..________________ 91 

[Berate minerals, theoretical relations____ 24 

Boron, determination in halite and anhydritic 

' halite rocks-___________ 166 

! losses through adsorption on etched

quartzware____________ 167 

Brachiopods, Caradocian, Maine______ 117 

Bryozoans, Caradocian, Maine________ 117 

tBushnell Seamount, submarine basalt flow_ 153

California, geochemistry, Death Valley.___ 28 

ground water, Arvin-Maricopa area. ... 171 

stratigraphy, Inyo Mountains_____.. 74 

pambrian, California, stratigraphy_____.. 74 

fcaradocian, Maine, paleontology_______ 117 

parbon isotope ratios, in hydrothermally 

1 altered carbonate rocks. ___ _. i 

Carboniferous. See Mississippian. 

pedar Mountain Formation, Utah, correlaI tion.________________ 97 

phile, earthquakes, Conception area_____ 131 

hinle Formation, Petrified Forest Member,

petrology...... ..___.__ 43

lay, flocculation and dispersion by detergents. . . ....._____ 209

Page

Clay minerals, in saprolite.. ____ B32

Cokedale Formation, Montana, definition. 89

Colorado, paleontology, Middle Park..___ 120

petrology, Central City quadrangle. __ 35

Colorado Plateau.. .......... 48

stratigraphy, southeastern________. 99 

Colorado Plateau. See specific State name. 

Corals, Caradocian, Maine.._________ 117 

Cordierite, associated mineral assemblages in

metamorphic rocks.________ 35

Cretaceous, Colorado, stratigraphy____ 99

Montana, stratigraphy......_ 86

New Jersey, stratigraphy__.. 1,02 

Puerto Rico, stratigraphy._____.. 114

Debris flows, association with volcanism 135 

Desert varnish, minor-element content. .. 28 

Detergents, effect on viscosity of water.. 209

influence on bacteria survival.. 205

See also Surfactants. 

Dolomite, hydrothermal origin.._______ 1

Earth resistivity measurements, electrical-... 

Earthquakes, eruptions of water and sand....

71

131

Fiords, structural control______ _____ 110

Flood waves, effect on ground-water movement and bank storage.--.---.. 192

Fluorine, in mineralized tuff. -..______ 16 

in silicic volcanic glass. . 18 

relation to beryllium content-..___... 16,18

Frost penetration, effect of snow cover____ 144

Gastropods, Caradocian, Maine..______ 117 

Georgia, clay mineralogy, Georgia Nuclear

Laboratory.____________ 32 

Glaciers, rate of movement, stratigraphy, and

temperature.___________ 140 

Gold, separation of tellurium from__--.-._ 164 

Graptolites, Caradocian, Maine-_______ 117 

Gravity surveys, Nevada, Gold Meadows

stock_________________ 64 

Ground water, effect of flood water. _____ 192 

salt-water instrusion.________.__ 182 

seasonal changes 189 

temperature variations. ________._ 186 

See also under State names.

H

Halite, determination of boron......____. 166

Hawaii, submarine geology, east rift zone of

Kilauea...................___. 153

Hoppers Formation, Montana, definition__ 91 

Hydrothermal chloride solutions, effect on

carbonate rocks ___...____ 1

Idaho, stratigraphy, eastern..______... 93 

Idaho Springs Formation, Colorado, petrology 35 

Infrared photography, detection of buried

structures...____________ 67

Page

Ion-exchange capacity, saprolite. B32 

Iron, separation of tellurium from.....____ 164 

Isotopes, carbon and oxygen, ratio in hydrothermally altered carbonate rocks. 1

Juana Diaz Formation, Puerto Rico, stratigraphy and structural geology__ 114

Karst topography, relation to glaciation 129 

Kelvin Formation, Utah, emendation___ 95 

Kentucky, glacial geology, Ohio River valley. 125 

Kilauea Volcano, submerged east rift zone,

basalt flow...___.... 153

Land subsidence, California, Arvin-Maricopa

area.__._ 171 

Lapilli, accretionary- 42 

Lead-alpha ages, plutonic rocks, Alaska_... 56 

Lead Gulch Formation, California, definition. 74 

Lead isotopes, reference sample._ 160 

Lead-uranium ages, factors affecting 60 

Liebigite, synthesis 162

Limestone, effect of hydrothermal chloride

solutions on 1 

Livingston Group, Montana, redefinition 86

M

Maine, paleontology, Ashland 117 

Mazourka Group, California, definition___ 78 

Metamorphic rocks, cordierite-bearing mineral assemblages. 35 

Miner Creek Formation, Montana, definition. 90 

Minor elements, in desert varnish .. 28 

Miocene, Colorado, paleontology 120

See also Tertiary.

Mississippian, Idaho, paleontology and stratigraphy 93

Montana, paleontology and stratigraphy.. 93

Wyoming, paleontology and stratigraphy. 93 

Montana, mineralogy, Ravalli County 10

stratigraphy, Crazy Mountains basin__ 86 

southwestern.. 93 

Montmorillonite, adsorption of surfactants on. 213 

Moon, surface material 148 

Morrison Formation, Utah, emendation 95 

Mudstone, siliceous, composition and origin.- 45

N

Nevada, economic geology, Shoshone Range.. 38 

geochemistry, Carlin and Pine Valley

quadrangles--- 45 

northeastern.__ . 28 

geophysics, Nevada Test Site --- 64 

mineralogy, Shoshone Range.. 38 

petrology, Carlin and Pine Valley quadrangles.-. 45

Colorado Plateau--_ 48 

Nevada Test Site 52

New Jersey, stratigraphy, Island Beach State

Park...__-__. . 102

New Mexico, geochronology, Ambrosia Lake. 60

geophysics, east-ceutral 67

petrology, Colorado Plateau 48

B217
B218 SUBJECT INDEX

Page

New York, ground water, Long Island, Suffolk County... B186

Nopah Formation, California, correlations.... 78

North Carolina, petrology, Carolina slate belt,

Stanly County... --------_--- 42

Nubian Sandstone, Sudan, ground water 179

Oligocene, Puerto Rico, stratigraphy..___. 114

See also Tertiary. 

Ordovician, California, stratigraphy___ 74

Maine, paleontology._______.__ 117 

Oxygen isotope ratios, in hydrothermally

altered carbonate rocks.___ 1

Pacific Ocean, submarine basalt flows, east of

Hawaii -...- ... ... - 153 

Paleozoic, New Jersey, stratigraphy..____ 102

See also system or period names. 

Park City Formation, paleoecology.-____ 123 

Patterned ground, Antarctica _ _ _ _ ______ 143

Permian, Idaho, Montana, Utah, Wyoming,

paleoecology..__________ 123 

Phosphoria Formation, paleoecology______ 123 

Photogeology, use of infrared photography.... 67

Piapi Canyon Formation, Nevada, petrography.......____-._.-_.__. 52

Pierre Shale, Apache Creek Sandstone Member, Colorado, redefinition-.___ 99 

Pillow structures, submarine basalts, at

extreme depths___________ 153 

Pleistocene, wave-cut platforms... _ ____ 108

See also Quaternary.

Pogonip Group, California, correlations. _ 78 

Polygenetic bedding, development-.. ____ 203 

Ponce Limestone, Puerto Rico, stratigraphy

and structural features______ 114 

Potassium-argon ages, plutonic rocks, Alaska. 56 

Precambrian, New Jersey, stratigraphy __.. 102 

Puerto Rico, structural geology, GuanicaGuayanilla Bay area.-- _____ 114

Quality of water, seasonal changes _____ 189

Page

Quaternary, Kentucky, glacial geology. B125 

New Jersey, stratigraphy__...... 102

See also Pleistocene.

Radium, migration of...___.-..__ 60 

Rare earths, thortveitite. 10 

Resistivity nomographs, in interpretation of

earth resistivity.. 71

Salt-water intrusion, ground water.. 182 

Saprolite, ion-exchange capacity.... 32

mineral content._.. 32 

Scandium minerals, thortveitite.. 10 

Sedimentation, effect of base-level changes 203 

Shedhorn Sandstone, paleoecology.. 123 

Silurian, California, stratigraphy... 74 

Skeletal fuzz, definition and origin._ .- 148 

Snow, critical thickness, definition.. 146

effect on frost penetration.. ..-. -. 144 

Solar wind, effect on Moon.. - 148 

Streams, effect of flood waves on ground-water

movement and bank storage 192 

Structural features, depiction by infrared

aerial photography-. 67

relation to topography___... 110 

Sudan, ground water, Kordofan Province. - 179 

Sunday Canyon Formation, California, definition 83

Surfactants, adsorption on montmorillonite 213

See also Detergents.

Tamarack Canyon Dolomite, California, definition... - 77 

Tellurium, separation from iron and gold... 164 

Tertiary, New Jersey, stratigraphy_ . - 102

See also Miocene and Oligocene. 

Thortveitite, first occurrence in Western

Hemisphere.___ .. 10 

Tin deposits, associated beryllium .... 12 

Topography, relation to structure, 110 

Tracers, spread in ground-water reservoirs 199

	Page

Trilobites, Caradocian, Maine_________ B117

Tuff, accretionary lapillL.____ ..__ 42

beryllium and fluorine content - _____ 16

determination of density and porosity. 169

petrology______________..__ 48

welded, petrography___.__... 52

U

Uranium deposits, age determinations___.. 60

Uranium minerals, liebigite, synthesis..... 162

Utah, geochemistry, Drum Mountains....-.. 1

geochemistry, Spor Mountain.. .. 16

petrology, Colorado Plateau______ 48

stratigraphy, Salt Lake City area 95

Vaughn Gulch Limestone, California, definition-....--_-_______._-. 81

Vegetation, effect on ground-water temperature. 186

Vertebrates, Colorado, Miocene. ______.. 120 

Vesicular rock, on Moon______. _ 148 

Virginia, economic geology, Irish Creek area.. 12 

Volcanic ash, accretionary lapillL. .. ... 42 

Volcanic glass, beryllium and fluorine content. 18 

Volcanic rocks, determination of density and

porosity___.__._ .__... 169 

Volcanism, debris flows.. .. 135

W

Washington, geomorphology, Mount Rainier. 135

ground water, Puget Sound area... 182

volcanology, Mount Rainier. 135

Waterpower investigations, Alaska 176

Wave-cut platforms, Pleistocene 108

Weathering, polar environment 140

Western States, geochemistry 18

See also under State names. 

Wyoming, geomorphology, Gros Ventre Mountains-. 129

glacial geology, Gros Ventre Mountains.. 129 

stratigraphy, western 93

Zeolites, eflect on measurements of density and

porosity of volcanic rocks 169
AUTHOR INDEX

Andersen, B. G.

Page 

B140

B

Bartel, A. J..._.-----_._---._._.__._ 20

Brew, D. A-_._..._._____........ 110

Brown, C. E__________________ 12

Cjadigan, R. A_________________ 48

Qasertano, Lorenzo______________ 131

Cbbban, W. A......_._.__._.__.. 99

Cboper, H. H., Jr____._____._.__ 192

C^andell, D. R............................... 135

CMttenden, M. D., Jr..........._..._.... 95

Ejavidson, D. F......___._.__........ 28

Delevaux, M. H......._.__._.......... 160

Dickey, D. D......___....._............ 169

D|utro, J. T., Jr.........._._.___...... 93

Fpnnelly, E. J....-...__....._....__.. 20

Feulner, A. J___________________ 189

FJscher, W. A.....-. ..................... 67

Ftledman, Irving...______________ 1

Gible, D. J^...... ._...................... 35

Gdli O., Carlos..._._._____.......... 131

G-anger, H. C.....__._._._.____.. 60

G rantz, Arthur..__.................._.. 56, 106

GHffitts, W. R-. .- .............-..... 16, 18

Gttmaldi, F. B............................... 166

Gfrossman, I. G._________________ 114

Hanshaw, B. B............................_ 24

Havens, R. G.......__._____.......... 10

Healey, D. L.______....___.......... 64

Hjiffman, Claude, Jr..____________ 20

H|ant, C. B................................... 28

Islander, Wilson_______________ 179 

Iz|Btt, G. A....-... .-... ..-........... 120

Page

Johnson, Arthur________________ B176 

Jopling, A. V................................. 203

Kantrowitz, I. H.._................_..... 186

Keefer, W. R.........._...._.___..._ 129

Ketner, K. B........_........_._.__ 38, 45

Kiilsgaard, T. H______...._.....__..... 12

Kimmel, G. E...._______________ 182

Krinsley, D. B.......__....._....._._ 144

L

Lakin, H. W__-.____........._....._ 28, 164

Lesure, F. G............_....._____.... 12

Lewis, G. E.__........___._.......... 120

Lofgren, B. E............._._............ 171

Loney, R. A. _...._.-.....-.._...... 110

Levering, T. 8.........__................. 1

M

McCarthy, J. H.............................. 1

Meyrowitz, Robert_._____________ 162

Miller, C. H.... ....- ....... ..... 64

Monk, E. F.................................. 169

Moore, J. G.................................. 153

Mrose, M. E___________________ 12

Muffler, L. J. P.....-............ ........ 110

N 

Neuman, R. B............................... 117

O'Connor, J. T........__........._...... 52

Oda, Uteana..._________________ 28

Ogata, Akio...___.....__.........._.. 199

Page, H. G........................... 205, 209, 213

Papadopulos, I. 8. _______________ 196

Parker, R. L.___.____________.._......... 10

Patterson, 8. H............................... 158

Pluhowski, E. J ........................... 186

Pomeroy, J S.............._............... 110

Powers, H. A...........___.....___.._ 18

Page

Rader, L. F., Jr_.___._____.__ B16, 20 

Ray, L. L.___..........___.._____ 125

Reed, R. K . -. 153

Roberts, A. E.........._.___..._.__ 86

Robertson, J. B.__...........___ 205, 209, 213

Rodis, H. G..... .......................... 179

Rorabaugh, M. I_...._.__.... ...._ 192

Ross, D. C............................. ... 74

Ross, D. E..___.__..._..... .. .. 162

3, W. J..........-..-. ......... 93

Schupp, E. G..____._......_ ...... 189

Scott, G. R _ . . 99

Scott, J. H.. .- .. ...... .. . 71

Seaber, P. R...................-. .. 102

Segerstrom, Kenneth. . 131

Sheffey, N. B........................ ... 56

Simon, F. O..___._____.-__--_ 166

Sims, P. K.. ................... ' 35

Smith, J. F., Jr. .... ..._ - 45

Stem, T. W. . - - 56

Stewart, J. W................. .. 32

Sundelius, H. W................. ..... 42

Thomas, Herman...___.__ _ 56 

Thompson, C. E.___________. . 164

Vecchioli, John......................-. 102

W

Warren, C. R..................... ... 148

Wayman, C. H...._ . .. . 205, 209, 213

Weeks, A. D.................... . ... . 162

Y 

Yochelson, E. L.............................. 123

B219

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