Geochemistry of the Phosphoria Formation 

at Montpelier Canyon, Idaho: 

Environment of Deposition

U.S. GEOLOGICAL SURVEY BULLETIN 2023-B
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Chapter B

Geochemistry of the Phosphoria Formation 

at Montpelier Canyon, Idaho: 

Environment of Deposition

By D.Z. PIPER and M.D. MEDRANO

U.S. GEOLOGICAL SURVEY BULLETIN 2023

THE PHOSPHORIA FORMATION:

ITS CEOCHEMICAL AND BIOLOGICAL ENVIRONMENT OF DEPOSITION
U.S. DEPARTMENT OF THE INTERIOR 

BRUCE BABBITT, Secretary

U.S. GEOLOGICAL SURVEY 

Gordon P. Eaton, Director

Any use of trade, product, or firm names

in this publication is for descriptive purposes only

and does not imply endorsement by the U.S. Government

Text and illustrations edited by George A. Havach

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON : 1994

Free on application to 

U.S. Geological Survey, Map Distribution 

Box 25286, MS 306, Federal Center 

Denver, CO 80225

Library of Congress Cataloging in Publication Data

Piper, David Z.

Geochemistry of the Phosphoria Formation at Montpelier Canyon, Idaho : 

environment of deposition / by D.Z. Piper and M.D. Medrano.

p. cm. (U.S. Geological Survey bulletin ; 2023) (The Phosphoria 

Formation its geochemical and biological environment of deposition ; ch. B)

Includes bibliographical references.

Supt. of Docs, no.: I 19.3:20238

1. Geochemistry Idaho Montpelier Canyon Region. 2. Phosphoria 

Formation. I. Medrano, M.D. II. Title. 111. Series. IV. Series: The Phosphoria 

Formation its geochemical and biological environment of deposition ; ch. B. 

QE75.B9 no. 2023-B 

[QE515]

557.3 s dc20 94-30601 

[551.9796] CIP
CONTENTS

Abstract Bl 

Introduction Bl 

Analytical techniques B3 

Results and discussion B6

Detrital fraction B6

Marine fraction B7

Biogenic fraction B12 

Hydrogenous fraction B18 

Conclusion B23 

References cited B24

FIGURES

1. Map of Northwest-Central United States, showing extents of phosphorite 

deposition and of correlative facies during the middle Permian B2

2. Chart showing lithologic section of the Phosphoria Formation at Montpelier

Canyon B4 

3-11. Plots of:

3. Interelement correlations in samples analyzed by more than one 

technique B5

4. Detrital fraction versus major-element-oxide contents in Permian rocks of 

the Phosphoria Formation B7

5. Detrital fraction versus minor-element contents in the Phosphoria 

Formation at Montpelier Canyon B9

6. Organic-matter versus minor-element contents in the Phosphoria Formation 

at Montpelier Canyon BIO

7. Apatite versus rare-earth-element contents and Ce anomaly in the 

Phosphoria Formation at Montpelier Canyon B12

8. Minor-element versus contents in modern plankton B13

9. Nutrient and metallic-ion concentrations versus depth in the Pacific 

Ocean B14

10. Minor-element versus Cu contents in the marine fraction of the Phosphoria 

Formation at Montpelier Canyon B17

11. Cr versus V contents in the marine fraction of the Phosphoria Formation at 

Montpelier Canyon B22

TABLES

1. End-member major-element-oxide and minor-element contents in sedimentary 

fractions of the Phosphoria Formation B8

2. Correlation coefficients between major sedimentary fractions and bulk minorelement contents and between major marine fractions and minor-element contents 

in the Phosphoria Formation at Montpelier Canyon, Idaho Bll

3. Minor-element contents in organic matter and seawater B15

4. Half-cell reactions, assuming standard-state conditions B20

Contents III

Geochemistry of the Phosphoria Formation at 

Montpelier Canyon, Idaho: 

Environment of Deposition

ByD.Z. Piper anc/M.D. Medrano

Abstract

The Permian Phosphoria Formation of southeastern 

Idaho, western Wyoming, southwestern Montana, and northeastern Utah was deposited in a broad ocean-margin basin, 

which was tectonically quite stable but underwent several 

transgressive-regressive events, owing possibly to glacial 

advances and retreats. Throughout most of its depositional 

history, intense upwelling supported exceptionally high primary productivity in the photic zone of the water column, 

resulting in a high flux of organic debris to the sea floor. In 

contrast, the flux of detrital material was low. Thus, the 

lithology of the rocks is dominated by marine phases carbonate fluorapatite, biogenic silica, calcite, dolomite, and 

organic matter. The rocks also contain high concentrations 

of minor elements Cd, Cr, Cu, Mo, Ni, rare-earth elements 

(REE's), V, and Zn. Partitioning of the minor elements into 

these host phases is largely a function of diagenesis.

The geochemistry of the water column at the time of 

deposition, however, is more likely to be understood by examining the partitioning of minor elements among the following 

source fractions: a detrital fraction, a marine biogenic fraction, 

and a marine hydrogenous fraction. Interelement relations of 

several major-element oxides and minor elements show that 

the detrital fraction has a composition closely approaching that 

of shale standards, requiring that this fraction had a terrigenous 

origin. Interelement relations of Cu, Mo, Zn, and, possibly, Cd 

and Ni within the marine fraction resemble those in modern 

marine plankton. Rates of accumulation for these minor elements furthermore required accumulation rates for organic 

debris on the sea floor similar to that at present on the Peru 

and Namibian Continental Shelves, areas of high primary productivity. Our interpretation of these relations is that these 

minor elements accumulated as biogenic debris, derived from 

the photic zone of the Phosphoria Sea.

The contents of Cr, REE's, and V in the marine fraction 

cannot be attributed solely to a biogenic source. Instead, 

these elements accumulated from seawater by inorganic precipitation and (or) adsorption onto other particulate phases.

Manuscript approved for publication, April 13, 1994.

Their marine chemical properties and distributions in the 

ocean today require their precipitation and (or) adsorption 

onto settling particulate matter, under bottom-water conditions of denitrification. Their accumulation rates can be used 

to estimate the residence time of bottom water on the shelf 

versus the rate of bacterial respiration, which is determined 

by the rate of settling organic matter from the photic zone 

again, a function of primary productivity. The distributions of 

these and the biologically derived minor elements indicate 

that the composition of seawater during the Permian was 

virtually identical to its composition today.

Of the major-element oxides and minor elements, only 

Cr, V, and REE's that is, the minor elements with a hydrogenous source show a difference in their distribution in 

rocks deposited during the proposed glacial and interglacial 

episodes. Their distributions suggest that the bottom water 

was more reducing, albeit still denitrifying, during glacial 

than interglacial times. Because bottom-water respiration is 

governed by the balance between the rates of settling of 

organic matter and advection of bottom water, the distribution of these minor elements suggests greater photic-zone 

productivity or lower bottom-water advection during glacial 

episodes; we favor enhanced productivity.

INTRODUCTION

The Phosphoria Formation, in the Northwestern United 

States (fig. 1), was first examined in detail by Mansfield 

(1927). The formation was later carefully described and 

measured by McKelvey and others (1959). Samples were 

analyzed for their minor-element contents (Gulbrandsen, 

1966), particularly uranium (Gulbrandsen, 1960; Swanson, 

1961). Thousands of samples were collected from natural 

outcrops and tens of freshly dug trenches. Although the 

formation failed to yield the sought-for uranium, it continues 

to be a major producer of phosphate and vanadium. Since 

1911, it has yielded some 140 million t of phosphate rock as 

of 1980 (Cathcart, 1991), with approximately 524 billion t of 

resource remaining (Sheldon, 1989). In addition to reserves 

of phosphate and vanadium, the Phosphoria Formation is also 

a major source rock for oil in Permian and younger rocks of

Geochemistry of the Phosphoria Formation at Montpelier Canyon, Idaho: Environment of Deposition B1
the Northwestern United States (Claypool and others, 1977; 

Maughan, 1984;Thornburg, 1990).

The environment of deposition of the Phosphoria Formation has been interpreted as one of exceptionally high 

ocean upwelling and biologic productivity (McKelvey, 

1967), a hypothesis first proposed by Kazakov (1937) to 

explain the high content of seawater-derived phases 

organic matter, carbonate fluorapatite, biogenic silica, and 

carbonates (Calvert and Price, 1983; Froelich and others, 

1988; Glenn and Arthur, 1988) in this type of deposit. The 

modern continental shelves of Namibia and Peru, where 

primary productivity of organic matter averages as much as 

200 mg/cm2 per year (Chavez and Barber, 1987; Berger and 

others, 1988), are commonly cited as modern analogs of the 

Phosphoria and other seas within which large amounts of 

sedimentary phosphate accumulated. The high rate of primary productivity is driven by upwelling of cold, nutrient 

(principally NOJ and PO4~~)-rich water into the photic 

zone from approximately 80- to 200-m depth (Sverdrup and 

others, 1942; Wyrtki, 1963; Calvert and Price, 1971; Codis112° 108° 104°

42°

40° -

Phosphorite 

deposits

EXPLANATION

Gypsum-anhydrite 

deposits

400 Kilometers 

I

Halite 

deposits

Figure 1. Northwest-Central United States, showing extents 

of phosphorite deposition and of correlative facies during 

the middle Permian (Wordian), along with paleolatitudes 

from Sheldon (1964). More recent work by Scotese and 

McKerrow (1990) gives 10° N. latitude line as 25° N.

poti, 1980). The resulting sediment is strongly enriched in 

organic matter, carbonate fluorapatite, biogenic silica, and 

carbonates (Glenn and Arthur, 1988; Froelich and others, 

1988). These same marine fractions dominate the bulk composition of the Phosphoria Formation (McKelvey and others, 1953, 1959). Conodont and brachiopod assemblages of 

the Phosphoria Formation furthermore indicate a cold-water 

depositional environment (Wardlaw and Collinson, 1986).

The dissimilarities between modern shelf environments and the Phosphoria Sea, however, must also be considered. The Peru Shelf is approximately 10 km wide, 

bordered by the Peru-Chile Trench to the west and the 

Andes Mountains to the east. It is an extremely active tectonic area. This narrow shelf is swept by one of the strongest surface currents in the ocean, the Peru Current. As a 

result, the sea floor is eroded in some areas but shows 

high sediment-accumulation rates in other areas. Although 

the shelf off Namibia is much wider (Hart and Currie, 

1960), it, too, is swept by one of the ocean's great currents, the Benguela Current. The sea floor exhibits a sedimentation regime as complex as that on the Peru Shelf, of 

erosion, sediment reworking, and primary deposition (Calvert and Price, 1983). On both the Peru and Namibia 

Shelves, nodular phosphate seems to represent a major 

proportion of the phosphatic fraction (Baturin and others, 

1972; Burnett and others, 1980). Hardgrounds are also 

abundant (W.E. Dean, oral commun., 1993), and pelletal 

phosphate is much less common (Burnett and others, 

1988).

In contrast, the shelf of the Phosphoria Sea was several hundred kilometers wide and gently sloping. Surrounding landmasses to the east were low lying (Paull and 

Paull, 1986), providing minimal detritus (Cook, 1968) to 

this broad shelf. The outer west boundary of this oceanmargin sea was characterized by a remnant of the Antler 

Uplift and, farther to the west, by an island arc (Stevens, 

1991), which would have shielded the Phosphoria Sea 

from such great ocean current as those that sweep the 

modern phosphogenic provinces of the open shelves of 

Peru and Namibia. The sea-floor bathymetry and oceanographic setting resulted in a minimal, but relatively uniform, average accumulation rate of the detrital fraction of 

0.5 mg/cm2 per year, comparable to its accumulation rate 

in the pelagic environment of the ocean today. The phosphatic fraction, which is virtually totally pelletal, accumulated at a comparable rate. Sediment thickness increased 

rather continuously from east to west and southwest 

(McKelvey and others, 1959). To the east, primarily in 

western Wyoming and northern Utah, coeval carbonates of 

the Park City Formation intertongue with the Phosphoria 

Formation. These rocks are interpreted as having been 

deposited on a shallow shelf area. Somewhat farther to the 

east, evaporites and continental red beds of the Goose Egg 

Formation accumulated (fig. 1). Thus, the Phosphoria Sea 

was a gently sloping, broad foreland (or sag) basin,

B2 The Phosphoria Formation: Its Geochemical and Biological Environment of Deposition
behind an outer island arc, basin, and highland (Antler 

Uplift).

We have analyzed samples from Montpelier Canyon 

(McKelvey and others, 1953), southeastern Idaho (fig. 2), 

for their major-element-oxide and minor-element composition (Medrano and Piper, 1992). The aim of this study is 

to establish the marine geochemistry of the environment 

of deposition. The great abundance of marine mineral 

phases (for example, biogenic silica, organic matter, and 

apatite) in the Phosphoria Formation suggests that the 

marine minor-element content was only slightly diluted by 

the detrital fraction and, possibly, not erased by postdepositional alteration. Thus, the marine fraction of minor elements might address such aspects of the depositional 

environment as (1) primary productivity in the photic zone 

and (2) redox conditions of the bottom water of the Phosphoria Sea. We show that the Permian Ocean must have 

had a composition virtually identical to that of the modern 

ocean, whose composition is maintained by the addition of 

minor elements through both hydrothermal activity (Von 

Damm and others, 1985a, b) and continental denudation 

(Martin and Maybeck, 1979). Such an ocean was the 

source for the marine constituents of this deposit and, 

quite likely, nearly all similar sedimentary deposits (Holland, 1979; Piper, 1994). It is unnecessary to invoke an 

unusual or exotic seawater chemistry to account for the 

composition of this deposit.

Determination of the seawater chemistry of deposition 

of ancient deposits, on the basis of their current minorelement contents, is complicated by the hydrodynamics of 

O2-depleted waters and the coarseness of most sampling 

schemes. The chemical properties of bottom water can 

change quite drastically over a period of seasons to years 

(Sholkovitz and Gieskes, 1971; Anderson and Devol, 

1973; Codispoti, 1980), as well as over longer time periods, largely in response to changes in seawater advection. 

Probably, no sampling of sedimentary rocks can identify 

such short-term changes as surely occurred in the 

past. Sampling schemes of ancient sedimentary deposits 

unavoidably average ocean conditions at the time of deposition over hundreds to thousands of years. For a sedimentary rock that accumulated at a rate of 2 mg/cm2 per year, 

a sample that is only 1 cm thick represents several hundred years of deposition. The integration imposed by this 

aspect of most investigations limits our ability to sort the 

ancient record. We might see in a single sample the geochemical signal of several different conditions, the sum of 

hundreds of years, rather than the details of season-to-season and year-to-year events. As a result, an individual 

sample might contain a record of contrasting bottom-water 

conditions. For example, it might have both high Mo and 

high rare-earth-element (REE) contents, whereas such a 

cooccurrence would seem to be excluded by the geochemistry of these minor elements in modern marine environments (Piper and Isaacs, in press).

A final complexity is introduced by the changes that 

sedimentary deposits undergo during diagenesis (Shaw and 

others, 1990) and, later, during burial and uplift (Isaacs, 

1980). Indeed, in some rocks, the record of deposition can 

be erased beyond our ability to decipher it from the current rock composition. The abundance, mineralogy, and 

composition of the host phases for minor elements that 

survive these changes or, more accurately, that are 

formed as a result of these changes (for example, apatite, 

glauconite, pyrite, bauxite, palygorskite, opal-CT, and so 

on) largely record these postdepositional environments 

within which alteration occurred. Such minerals can indicate the extent to which alteration proceeded (Vieillard 

and others, 1979). This aspect of the history of a formation is interesting from a scientific, as well as economic, 

point of view because it can either add to or subtract from 

the value of a deposit. Still, its added effect is to create an 

increasingly dimmer view of the depositional environment. 

Nonetheless, the marine minor-element signal of the Phosphoria Formation lends strong support to our interpretation 

of primary productivity in the photic zone and redox conditions of the bottom water at the time of deposition.

Acknowledgments. The senior author has had many 

discussions through the years with the geologists who contributed much to our current understanding of this and 

similar deposits. Most influential have been two former 

colleagues at the U.S. Geological Survey, Dick Sheldon 

and Bob Gulbrandsen. They introduced me to the Phosphoria Formation, pointed me in the direction I first took, 

and made every effort, sometimes without success, to 

urged me back on path when I wandered far. Reviews of 

the manuscript by B. Murchey, J. Herring, G. Breit, and 

(}. Piper greatly improved its organization and clarity.

ANALYTICAL TECHNIQUES

The analytical techniques used in this study, the analytical data, and their precision and accuracy were discussed by Medrano and Piper (1992). Quite briefly, the 

accuracy of these analyses is established by analyzing a 

standard rock along with each group of 25 to 40 samples; 

precision is established by repeatedly analyzing a single 

sample, or standard (Baedecker, 1987). An alternative 

approach to accuracy and precision is to analyze groups 

of samples by more than one technique (fig. 3). Plots of 

the major-element-oxide analyses measured by X-rayfluorescence (XRF) and inductively coupled plasma/ 

atomic-emission spectroscopy (ICP-AES), and REE analyses measured by ICP-AES and ICP-mass spectroscopy 

(ICP-MS), demonstrate this approach. Our reason for 

showing these plots is to demonstrate that analytical error 

introduces only a small error in the stoichiometric calculations which we make throughout this report and on 

which our conclusions are strongly based.

Geochemistry of the Phosphoria Formation at Montpelier Canyon, Idaho: Environment of Deposition B3
EXPLANATION Dinwoody Fin. 

1.2m

Limestone

Sandstone

Chert

o

s

<u r^^

<a 

X 

oo 

w*

SL

Phosphate 1Q 4 m 

rocks

Argillaceous 

limestone

Phosphatic 

chert

Calcareous 

mudstone

Calcareous 

chert

Cherty 

limestone

Argillaceous

phosphate

rocks

Phosphatic 

mudstone

Calcareous 

phosphate 

rocks

Phosphatic 

limestone ;

c

Quartzite

"Interbedded" 

limestone and chert

5 METERS

I

QJ

u

I

43.2m

'U-12

U-7 

-U-4

-R-22

IZII

 P-157

<N

- P-87

 P-148

 P-146

 P-141 

P-140

>/

'P-131 .p l2g 1fM.Q lU4.yrnm

 P-105

 P-104

 P-103

 P-100

P-82

P-71 

P-69

128.7 m

Figure 2. Lithologic section of the Phosphoria Formation at Montpelier Canyon, showing sample locations. Inferred glacialinterglacial episodes are shown immediately to left of column (labeled G-1, IG-1, and so on).

B4 The Phosphoria Formation: Its Geochemical and Biological Environment of Deposition
-y = -0023 + 0.94xR= 1.00 

---- y = -0.19 + l.lxR= 1.00

2 4 6 8 10

A12 O3 CONTENT (ICP-AES), 

IN WEIGHT PERCENT

faZ 

«W 

XU

si 35

o^ 30

25

20

2.5

- - - y = -0.77 + 1 .Ox R= 0.99

25 30 35 40 45 50

CaO CONTENT (ICP-AES), 

IN WEIGHT PERCENT

x x~'

uw

2 0.5

0.5 1 1.5 2

MgO CONTENT (ICP-AES), 

IN WEIGHT PERCENT

z 

wH ^£

I £g 20 

c£o3 15

<~ 10 -y = 1.4 + 0.95x R= 0.99

----- y = 0.27 + 0.97x R= 0.99

10 15 20 25 30 35 40

P2 O5 CONTENT (ICP-AES), 

IN WEIGHT PERCENT

700

c^O

O, J 500 

n_ Hi 400

200

100

0

- - y = 2.8 + 1 .Ox R= 0.99

100 200 300 400 500

La CONTENT (ICP-AES), 

IN PARTS PER MILLION

Z f O °* 100

cu 

3£ 50

- - - y = 3.4 + l.lx R= 0.98

0 50 100 150 200

Ce CONTENT (ICP-AES), 

IN PARTS PER MILLION

350

^;Z 300

ft, J 250

100 150 200 250

Nd CONTENT (ICP-AES), 

IN PARTS PER MILLION

300 350

^Z* 30

c« O

^J

0,J 25 

-S

H« M

Z K

w fc

H«3 15 

ZHO« 

U< 10

>£ 5

//

5 10 15 20 25

Yb CONTENT (ICP-AES), 

IN PARTS PER MILLION

30 35

Figure 3. Interelement plots, showing correlations of elements, determined by analysis of samples by more than one 

technique. Open symbols, wet-chemical data; solid symbols, XRF data.

Geochemistry of the Phosphoria Formation at Montpelier Canyon, Idaho: Environment of Deposition B5
Plots of individual oxides show an excellent correlation, better than 0.99 in all cases. Clearly, the precision of 

these analyses approximates the reported values of less 

than 1 weight percent for XRF (Taggart and others, 1987) 

and seems to be better than that for ICP-AES (Lichte and 

others, 1987). The variation from the 1:1 line of about 5 

percent (figs. 3A-3C) must somehow reflect the inaccuracy of the analyses. The close agreement of 1.9 percent 

for the XRF and wet-chemical P2C>5 data about the 1:1 

relation (fig. 3D) suggests that the ICP-AES analyses contribute most to the inaccuracy (the XRF and ICP-AES 

analyses were made in 1991, and the wet-chemical analyses in about 1952). We should note, however, that both the 

wet-chemical and XRF data can also have an indeterminate inaccuracy. Because ICP gives both major-elementoxide and minor-element results and all analyses show a 

high precision, we elect to use the ICP data throughout the 

following discussion.

The REE analyses by ICP-AES and ICP-MS show a 

correlation and agreement with the 1:1 line that suggest a 

precision for the minor-element analyses of better than the 

approximately 5 percent based on ICP-AES analyses 

alone (Lichte and others, 1987).

RESULTS AND DISCUSSION

We calculate the mineral composition of the Phosphoria 

Formation at Montpelier Canyon from its major-elementoxide composition by a normative scheme (Isaacs, 1980; 

Medrano and Piper, 1992). Appropriate stoichio-metric values and major-element-oxide contents are substituted into 

the following formulas: A12C>3X6.4, SiO2-detrital SiO2, 

(P2O5-detrital P2O5)/0.406, (MgO-detrital MgO)/0.219, 

[CaO-(detrital CaO+apatite CaO+dolomite CaO)]/0.56, and 

organic carbonxl.7, to give the amounts of detritus, biogenic 

silica, carbonate fluorapatite, dolomite, calcite, and organic 

matter, respectively, in each sample (see Medrano and Piper, 

1992, table 4). The accuracy of our calculations depends 

on the assumption of a constant composition for each fraction 

and the validity of the stoichiometric values assigned to 

each fraction. The sum of these fractions in the Phosphoria 

Formation at Montpelier Canyon is 100±2 percent for 35 

of 41 samples and ±4 percent for all 41 samples (Medrano 

and Piper, 1992). The close approach to 100 percent and 

the random scatter about 100 percent support our choices 

of the above stoichiometric values.

Detrital Fraction

The partitioning of minor elements in the detrital host 

phase can be critical to the discussion of the geochemical 

properties of the depositional environment. We attach such 

importance to the detrital fraction because the seawater 

that is, the marine contribution of minor elements is calculated as the difference between the bulk minor-element 

composition of each sample and the minor-element contribution of the detrital fraction. Because the abundance of 

the detrital fraction is based solely on the A12C>3 content, 

the uncertainty in calculating the abundance of detritus in 

each sample is the sum of the uncertainties in the determination of A12C>3 contents, approximately 5 percent (fig. 

3A), and the uncertainty of the stoichiometric determination of detrital A12C>3 contents. The sum of all host fractions is 100 percent, however, supporting our selection of 

the A12O3 content of the detrital fraction. Also, strong correlations exist between the abundance of the detrital fraction and TiO2 and K2O contents (Medrano and Piper, 

1992), and lines of best fit extrapolate to zero (fig. 4). Furthermore, Al2O3 :TiO2 and A12C>3:K2C) ratios are similar to 

those of the World Shale Average (WSA) and the North 

American Shale Composite (NASC), all of which supports 

our interpretation of a uniform composition for detritus. 

Thus, our conclusion is that the uncertainty in our calculation of the detrital fraction is less than 10 percent.

Minor elements, which are hosted predominantly by 

the detrital fraction of many, though certainly not all, sedimentary rocks, include Sc, Ga, Li, and, possibly, Th and 

Co. The strong correlations between the abundance of the 

detrital fraction and Ba, Co, Ga, Sc, and Li contents (table 

1) supports the results for major-element oxides, of a uniform composition between samples. Their convergence 

toward standard shale values (figs. 5A-5D) further supports a terrigenous origin for the detrital fraction. Implicit 

in this interpretation is that the detrital fraction has 

remained nonreactive: It is both source phase and current 

host phase.

It would be a mistake, however, to assume that the 

detrital contribution of those minor elements which do not 

correlate with A12O3 can be determined merely from the 

composition of WSA or NASC. The variations in the composition of the rocks analyzed in many different laboratories to obtain the averages of these two standards 

(Wedepohl, 1969-78; Gromet and others, 1984) demonstrate variations in the composition of sedimentary rocks, 

even of similar lithology, and require that any comparison 

with other rocks is merely a first step in determining the 

detrital contribution of minor elements to the total inventory in any single deposit. In this study, the detrital contribution of this group of minor elements is determined from 

the minimums of the detritus-element plots (figs. 5E-5//). 

In every case, this procedure gives a value that is similar 

to the minor-element content of WSA (table 1).

Although the error in our estimate of the minorelement contribution from detritus is large for some 

minor elements, it introduces little uncertainty to the 

minor-element contents of the marine fraction. The detrital contribution of Cr, Cu, Ni, Mo, REE's, V, and Zn is 

small relative to the contributions from the marine fracB6 The Phosphoria Formation: Its Geochemical and Biological Environment of Deposition
tions. For example, the abundance of the detrital fraction 

averages approximately 35 percent. The uncertainty in

0.5

OA

0.3

~ 0.1

EXPLANATION

-- - Interglacial

--«-- Glacial

10 20 30 40 50 60 70

10 30 40 50 60 70

 Z

Z 

W 

H 

Z 

O 

U

4

3.5

3

2.5

2

1.5

1

0.5

0

10 20 30 40 50 60 

DETRITAL FRACTION, 

IN WEIGHT PERCENT

70

Figure 4. Detrital fraction (Al2 O3 x6.4) versus major-elementoxide contents in Permian rocks of the Phosphoria Formation. Dashed curves, least-squares lines of best fit; solid 

curves, World Shale Average (WSA; Wedepohl, 1969-78) 

and North American Shale Composite (NASC; Cromet and 

others, 1984). Clacial-interglacial episodes determined from 

lithology (Sheldon, 1984) and supported by 818O values of 

PO;S~ (Piper and Kolodny, 1988). See figure 2 for identification of glacial and interglacial events.

estimating the marine fraction of La from its bulk content 

minus its detrital content is, on average, approximately 

±{(0.05x32 ppmx0.35)+[0.02xbulk content (in parts per 

million)]}, where detritus is considered to contain 32 

ppm La (table 1). The uncertainty clearly is dominated by 

the 2-percent uncertainty in the ICP-AES measurement 

of bulk La content (fig. 3£).

There is nothing surprising about the result that the 

detrital fraction in the Phosphoria Formation had a terrigenous origin of uniform composition. The lithology and 

stratigraphy suggest that the Phosphoria Formation accumulated on a stable ocean margin (McKelvey and others, 

1959; Swanson, 1973). However, the conclusion that the 

detrital fraction had the same approximate composition 

(figs. 2, 4, 5) during previously proposed glacial and interglacial times (Sheldon, 1984; Piper and Kolodny, 1988) is 

somewhat unexpected. Nonetheless, these results allow us 

to dismiss any significant error in the minor-element contents of the marine fraction, introduced by possible major 

variations in the composition of the detrital fraction of the 

Phosphoria Formation. Also, these results diminish the 

possibilities of a major local hydrothermal input to the 

Phosphoria Sea or of severe alteration during weathering.

Marine Fraction

The marine fractions of Cu, Cr, and, less so, several 

other minor elements correlate with the organic-matter 

content (fig. 6; table 2). Partitioning of minor elements 

into organic matter was reported by Leventhal (1989), 

Odermatt and Curiale (1991), and Piper and Isaacs (in 

press) for the Monterey Formation; by Swanson (1961) 

for several Paleozoic organic-matter-enriched shales; and 

by Nissenbaum and Swaine (1976) for modern marine 

sediment. Lewan (1984), Moldowan and others (1986), 

and Odermatt and Curiale (1991) further examined the 

fractionation of metals between the various phases of 

organic matter. For example, Lewan (1984) interpreted the 

V/(Ni+V) ratio in crude oils to reflect the Eh, pH, and 

sulfide activity of the depositional environment.

The REE's are partitioned into apatite (fig. 7), similar 

to their partitioning in other phosphate deposits (Wright 

and others, 1987; Piper, 1991). Additional host phases 

include glauconite and metal sulfides (Glenn and Arthur, 

1988; Piper and others, 1988), predominantly iron sulfides. 

They form various minerals, but pyrite is the stable end 

product (Goldhaber and Kaplan, 1974). Pyrite is present in 

many sections of the Phosphoria Formation (Sheldon, 

1963), but only in trace amounts in Montpelier Canyon 

rocks (McKelvey and others, 1953). We have also detected 

trace amounts of sphalerite by microprobe.

Partitioning of minor elements into these host phases 

and, particularly, into organic matter must occur during 

diagenesis and, possibly, during later alteration of the

Geochemistry of the Phosphoria Formation at Montpelier Canyon, Idaho: Environment of Deposition B7
Table 1. End-member major-element-oxide and minor-element contents in sedimentary fractions of the Phosphoria Formation

[Major-element-oxide contents in weight percent; minor-element contents in parts per million. 

Values for the detrital fraction of the Phosphoria Formation represent minimums, or single trends, 

obtained from plots of element or element-oxide versus detritus contents (see figs. 4, 5); values 

for the detrital fraction in parentheses are for the World Shale Average (WSA) or North American 

Shale Composite (NASC). Dashes, no data available. Other minor-element contents for the detrital fraction of the Phosphoria Formation and WSA, in parentheses, are as follows: As, 10 (10); 

Ba, 360 (640); Ga, 18 (21); Li, 45 (60); Sc, 16 (13); Sr, 210 (125); Y, 35 (26). Data sources: 

WSA, Wedepohl (1969-78); NASC, Gromet and others (1984); biogenic SiO;, Elderfield and 

others (1981); calcite, Palmer (1985); organic matter, Sverdrup and others (1942), Elderfield and 

others (1981), and Brumsack (1986)].

Detrital fraction

Phosphoria Formation 

(WSA or NASC)

.. . ,. . , . Authigemc Marine biogenic fractions , ° ° fraction

Biogenic , . 

Si02 Calcite 0r§aniC Apatite 

matter r

Major-element oxides

P 2f\U5 

TiO2 - 

K20 - 

CaO -----

Fe203 -~- 

A12O3 

SiO2

.16(0.16) 

.77 (0.48) 

3.98 (3.2) 

3.9 (3.09) - 56.0 

6.75 (4.81) 

15.6 (15.5)

58.5 (58.4) 100

1 74 40 7 1 . / H tv/. /

55.5 

.050

Minor elements

T a

r-e ____

pr_.__. _..

MfJ _______

<\m

En ______

/~* A

TH_

Dy______

Ho -

Fr _______

Tm 

Yb -

T 11

Pn

Cd - 

Co - 

Cr

Mo ------

Mj _ _______

<\f>

IT- - - 

v

Zn -

33 (31.2) 

66 (66.7)

28 (27.4) 

5.59 

1.18

.85

2.85 (3.06) 

.456 

36 (35.0) 

.2 (0.2) 

11 (19) 

83 (83) 

2.6 

44 

1.0 

3.0 

110(110) 

110(110)

5.80 

8.21

6.71 

1.56 

.40 

1.78

1.69 

.86

0.128 

.105

.106 

.019 

.004 

.022

.024

.015 

.700

0.14 

.23

.30 

.01

11.0 

12.0 

1.0 

2.0 

2.0 

7.4 

3.0

3.0 

110.0

-

~

originally deposited marine phases. Minor elements commonly are enriched by a factor of more than 10 in the

residual organic fraction over their concentrations in modern planktonic debris. For example, the slopes of the

B8 The Phosphoria Formation: Its Geochemical and Biological Environment of Deposition
14

O 12

Oft.

eu 4

Z1-4

2

0

>,-o' o o o

250

O 200

§s

II

Ucri 100

W * 150

Ha; 

zw

Z

50

0 10 20 30 40 50 60 70 10 20 30 40 50 60 70

14

1"

»J 10 i*la 8

; W

) W2 ^

>* 4 

Cu

£ 2

 o,.' ^ -»o» 

o x ' - ooo

P' 00

0 10 20 30 40 50 60 70

Z

o

J

a;

a,

t»

<'

a.

Z* "'

IUU

80

60

40

20

n

F

o

'

-

o

'

o

0 .

0 . 

 8 .

 . .

 o

O O 0* f <P 9 %

o 8 , . . o WSA , | 

0 10 20 30 40 50 60 70

350

| 300

J 

P J 250

is

<fA 200

Ucc 150

nf f"1

100

50

^- --6»

O t ft' .0°

"'92-

-O ,'~0

_l i l i_____i_____i_

0 10 20 30 40 50 60 70

20

15

HO; 

zw 10

ft. 5

Z

o- - "o o

O ,0'''

0 10 20 30 40 50 

DETRITAL FRACTION, 

IN WEIGHT PERCENT

3000

Z

2 2500 -

. ^ M

z; S 2000 -

z

1500 -

S g* 

N ^ 1000 -

a. 

Z "" 500

60 70 10 20 30 40 50 60 

DETRITAL FRACTION, 

IN WEIGHT PERCENT

70

Figure 5. Detrital fraction versus minor-element contents in the Phosphoria Formation at Montpelier Canyon. Same 

symbols as in figure 4. See table 2 for correlation coefficients.

Geochemistry of the Phosphoria Formation at Montpelier Canyon, Idaho: Environment of Deposition B9
80

Ig 70

£w eo R u

W« r

-«W 50

cc O

W3r 20 OK 

2 10

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60

40

20 . .

.-- «

8 10 12 8 10 12

1 -TV

Z 120

O

H J 100

H a 80

8*60

T3 fid

w<; 40

PL,

2 20

0

R .

0

o

o _

0 

 0

o _ 0 

o o .

0 

<b ,°A : - , . ,° . , °

10 12

400

Z 350 

O

g i 3oa

U3^ 250

SU 100

Z

^ 50

0

*' *'

10 12

TVUU

Q 3500

H J 3000

H « 2500

® ft" 2000

fc^

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z 100°

500

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o ^ ,'.---' .

.*, ,-£ '-'' *o

o o -"'-'' 

- o °S^'?~''''

.'"^° . ° 1 , 1 ,

1200

2;

2 1000

Has 

zw

800

600

400

Cu

200

10 12 10 12

200

Z 

O 

.3 150

wS

^50

D

2 4 6 8 10

ORGANIC MATTER CONTENT,

IN WEIGHT PERCENT

12

3000

Z

22500

j

H J 

g § 2000

1 500

^ 500

-TD 

2 4 6 8 10

ORGANIC-MATTER CONTENT,

IN WEIGHT PERCENT

12

Figure 6. Organic-matter versus minor-element contents in the Phosphoria Formation at Montpelier Canyon. Same 

symbols as in figure 4, except for figure 6C, where solid line separates glacial from interglacial group of samples. 

See table 2 for correlation coefficients.

BIO The Phosphoria Formation: Its Geochemical and Biological Environment of Deposition
Table 2. Correlation coefficients (A) between major sedimentary fractions and bulk minor-element contents and (6) between ma

contents in the Phosphoria Formation at Montpelier Canyon, Idaho

[Values >0.393 are significant at the 1-percent level]

A

Dolomite 

Calcile -

Apatite- -

Organic matterDetritus -----

Biogemc SiO2-

A« 

As ------

Ba - 

Cd 

Ce 

Cr - -

Cu 

Ga 

La 

L, - 

Mo 

Nd - - 

Ni - -. 

Pb ... ..... 

Sc 

Sr - --- 

V ----- -----

Y - 

Yb 

Zn - ----- -

B

Dolomite- -

Calcite -

Organic matterBiogenic SiO2-

Marine Ag 

Marine As 

Marine Ba -

Marine Cd 

Marine Ce 

Marine Co 

Marine Cr 

Manne Cu 

Marine Ga-- 

Marine La -

Marine LiMarine Mo -----

Marine Nd - -

Marine Ni 

Marine Pb 

Marine Sc -

Marine Sr -

Marine V -

Marine Y -

Marine Yb 

Marine Zn --

Dolomite

1

17

- 36

 25

- 18

- 10

-24 

-28 

-29 

-.02 

-.41 

- 19

- 31 

-35 

-22 

- 33 

-.17 

-.16

-.31 

- 18 

-24 

-.38 

-32

-.12 

-32 

- 33 

- 21

Dolomite

1

.17

-25

-.10

-24

-.27

-29

-02

-.37

.07

-.31

-.33

02

-.32

.06

-.15

-.29

-.17

-.12

-.28

-30

-.12

- 31

-.31

-.20

Calcite

1

- 34

- 18

- 39

- 10

- 36 

- 38 

-.47 

19 

-60 

-25

- 33 

-.31 

-34 

-.43 

- 35 

- 16

-41 

- 26 

-.59 

-.60 

- 18

- 12 

-.40 

-42 

-20

Calcite

1

- 18

-.10

- 36

-.34

-31

.19

-43

.41

-.33

-26

.28

-.40

.16

-15

-.36

-24

-.33

-.34

- 15

-.10

-.39

-.38

-.19

Apatite

1

18

- 51

-41

25 

12 

- 32 

13 

08 

-48

21 

26 

-41 

.56 

-30 

04

51 

- 13 

05 

17 

68

24 

63 

58 

-05

Apatite

18

-41

.25

.21

.18

13

53

39

22

34

34

.60

.36

05

55

-.09

.60

.72

70

.26

.65

.62

-.04

Organic

1

19

- 41

65

.58 

.13 

16 

.17 

.02

.87 

.89 

28 

.43 

36 

.49

32 

.66 

.26 

60 

67

20 

29 

.28 

53

Organic 

matter

1

-.41

.65

.58

-06

.16

.06

-29

87

.87

.10

.42

.12

48

.29

68

.11

.51

.64

.19

.28

.25

.53

Dctntus

1

03

.26 

.49 

87 

- 10 

.62 

.80

.33 

20 

91 

-05 

67 

.25

-03 

.56 

.64 

.61

-22

-04 

- 18 

-.10

47

Biogenic 

Si02

1

-.33

-40

.14

-.29

-.08

06

-.40

-.30

-.18

-26

-.01

-.23

-.21

-.30

-.25

-.41

-.56

-.14

-.21

-.17

-.30

B,ogemc 

S,02

1

- 33 

-37 

10 

-29 

- 04 

14

-40 

-29 

05 

- 25 

03 

-23

21 

- 28 

- 18 

- 31 

-57

- 14 

-21 

- 18 

-29

Manne 

Ag

1

72

-09

48

17

-31

.77

80

-.10

41

.02

.51

.31

.61

45

.48

.44

.40

.30

28

.67

Ag

1 

71 

16 

48 

30 

12

77 

83 

25 

43 

34 

51

34 

61 

59 

60 

46

41 

31 

32 

66

Marine 

As

1

-20

27

.27

-35

.63

.63

-.24

.39

-.27

.43

32

67

27

.42

43

15

28

.25

.67

As

1 

31 

24 

52 

34

64 

67 

43 

39 

29 

44

34 

71 

59 

68 

39

15 

25 

26 

70

Manne 

Ba

1

- 18

34

- 16

.01

-.11

- 11

.19

.06

-.14

.16

- 19

-.11

08

-.05

-01

.23

.25

-.20

Ba

1 

-18 

67 

62

28 

11 

77 

06 

64 

14

.06 

36 

52 

.58 

-21

-03 

- 02 

05 

29

Manne 

Cd

1

-01

-.03

.30

29

.04

.11

.09

.15

07

.08

.41

.20

13

64

.05

.04

.31

Cd

1 

-.08 

-.23

.29 

.27 

 09 

10 

-.05 

.15

06 

.07 

23 

.08

13

.64 

.04 

04 

31

Manne 

Ce

1

.06

.20

.18

.06

.84

.02

-.06

87

-.05

.39

.60

.40

-.08

.87

.85

-.01

Ce

1 

48

38 

26 

58 

63 

40 

12

67 

33 

64 

76 

16

- 10 

57 

61 

31

Manne

1

-42

-.13

.44

.05

.44

-.25

09

-.41

.10

.17

.24

-.04

.15

.13

-.41

Co

1

10 

07 

70 

- 13 

44 

16

- 07 

49 

43 

42 

-20

- 18 

21 

- 16 

36

Manne 

Cr

1

.85

.10

.56

16

38

44

57

.19

.58

58

.14

.40

37

.51

Cr

1 

88 

45 

58 

56 

39

47 

58 

43 

76 

60

.15 

41 

.41 

52

Manne 

Cu

1

.17

.53

26

.48

.41

.61

.29

.62

.65

.22

.42

.39

.55

Cu

1 

28 

.52 

41 

51

40 

.66

.38 

64 

.62

22 

39 

.39 

.61

Manne 

Ga

1

.28

64

.08

.26

-.15

08

.39

.40

-.06

.31

.29

-.14

Ga

1 

08 

74 

31

09 

56 

57 

67 

-08

-08 

- 05 

02

47

Manne 

La

1

.26

.10

.97

.14

42

.85

.65

-.02

.97

.94

.12

U

1 

15 

11

97 

17 

30 

65 

64

-03 

97 

95

15

Li

1

-.08

.21

-.28

.07

.36

.27

.09

.26

.25

-.23

Li

1 

22

11 

40 

38 

56 

-.06

00 

01 

06 

35

Marine 

Mo

1

.06

.76

15

.24

.20

.37

.06

.05

.79

Mo

1

.08 

.76 

33 

.39 

.20

38 

06 

.08 

.79

Manne 

Nd

1

.08

.39

.80

.60

-.05

.97

.95

.07

Nd

1 

.13 

29 

.63 

.58

-.07 

.96 

.95 

11

Manne

Ni

1

.02

.20

.27

.23

.05

.04

.87

Ni

1 

48 

60 

27

24 

05 

.08 

88

Manne 

Pb

1

.42

.31

.56

.45

.44

.12

M
minor-element/organic-matter curves (fig. 6) are unrelated 

to the composition of modern marine planktonic organic 

matter, a relation similar to that observed for other depos400

Z 350

O

-3 300

200

100

50

35

Z 30

O

H J 25 

20

20 40 60 80 100

* H

>«< 10 

O.

5

O* « *-° 2 

20 40 60 80 100

-0.2

J -0.4

^

O -0.6 

Z

^

« -0.8

-1.2

----- y = -0.118 + -0.4191og(x) R= 0.873

----- y = -0.283 + -0.3081og(x) R= 0.807

20 40 60

APATITE CONTENT, 

IN WEIGHT PERCENT

80 100

Figure 7. Apatite versus rare-earth-element (REE) contents 

and Ce anomaly in the Phosphoria Formation at Montpelier 

Canyon. Ce anomaly is defined as log {(3xCe/83)/[(2xLa/ 

41)+(Nd/38)]}, where each REE is normalized on an elementby-element basis to World Shale Average (WSA; Piper, 

1974). Same symbols as in figure 4.

its (Calvert and Price, 1983; Piper, 1991; Piper and Isaacs, 

in press).

The minor-element/organic-matter relations probably 

do not change drastically during settling through the water 

column. Fisher and Wente (1993, p. 671) concluded that 

Am, Ag, and Sn are "* * * retained sufficiently long, even 

by decomposing cells, to suggest that phytoplankton sinking as aggregates at rates of 100 m day" 1 would effectively 

transport these metals hundreds of meters out of oceanic 

surface waters." We might conclude from this finding that 

the current high minor-element content of sedimentary 

organic matter reflects uptake of minor elements by the 

residual organic phase within the sediment. This interpretation parallels that advanced to explain the relation 

between C and S in organic matter of modern sediment 

(Mossmann and others, 1991). Also, apatite, which precipitates not from seawater but within the upper few centimeters of the sediment (Kolodny, 1981; Burnett and others, 

1988), must acquire its minor-element signal and, particularly, its REE signal from sediment pore water. Thus, the 

partitioning of minor elements among the different marine 

host phases of sediment (residual organic matter, apatite, 

and so on) reflects pore-water conditions, that is, diagenetic 

conditions. The original source phases (planktonic organic 

matter, opal-A, CaCO3, and hydrogenous minor-element 

precipitates and adsorbed phases) must be evaluated to 

decipher the seawater environment of deposition.

Biogenic Fraction

The marine source phases are (1) planktonic debris 

and (2) hydrogenous precipitates and adsorbed phases. The 

biogenic contribution of minor elements is determined, 

first, by assuming a stoichiometry for plankton; then, the 

residual minor-element content is assigned to the hydrogenous source. Minor-element contents in plankton vary 

considerably (Martin and Knauer, 1973; Eisler, 1981; Collier and Edmond, 1984), owing to inconsistencies introduced during sample collection and analysis, but also to 

such natural factors as variations among species, geography, growth rates, and seawater pH (Eisler, 1981). The 

care with which the samples were collected and analyzed 

(Martin and Knauer, 1973; Collier and Edmond, 1984; 

Sherrell, 1989) requires that the reported variations in the 

composition of plankton (fig. 8) are real, rather than artifacts of sampling and analysis. The results indicate that we 

can never expect to obtain a single value for any minorelement content in individual samples of plankton but 

must settle for an average value, or best estimate. We 

might add that significant variations also are reported for 

the NOJ and PC>4~ contents in plankton (Sverdrup and 

others, 1942; Collier and Edmond, 1984); these ions constitute the major nutrients (Redfield and others, 1963), 

whose functions in plankton growth are well established. 

Given these variations, we chose to use the estimates of

B12 The Phosphoria Formation: Its Geochemical and Biological Environment of Deposition
Brumsack (1986) rather than introduce a new set of values 

(table 1).

The magnitude of minor-element variations in plankton (fig. 8) can easily discourage further consideration of 

the data and their application to geochemical investigations of sedimentary rocks. However, two important 

aspects of the marine environment require that the contribution of organic debris to the accumulation of minor elements be examined. First, rates of primary productivity 

throughout the ocean margin are high (Berger and others,

1000

Z;

zw

100

10

0.1

EXPLANATION

o Mixed plankton 

» Phytoplankton 

D Zooplankton

 Suspended matter o 

(0-100m depth) o o

Plankton

0.1 10 100 1000

1000

z

o

H 100

10

a- 1 

z

0.1

0.1 10 100 1000

10*

100

10 Plankton

0.1 1 10 100

Cu CONTENT, 

IN PARTS PER MILLION

1000

1988), resulting in a large flux of organic matter to the sea 

floor (Calvert, 1990; Pedersen and Calvert, 1990; Calvert 

and others, 1991). Second, a significant amount of the 

organic matter produced in the photic zone accumulates on 

the sea floor of ocean margins, owing to shallow water 

depths (Reimers and Suess, 1983). Thus, this fraction of 

sediment on ocean margins should represent a major sink 

for many minor elements in the oceans today (Collier and 

Edmond, 1984) and in ancient oceans as well.

The distributions of Cd, Cu, Cr, Ni, V, and Zn in the

^_ _

water column (fig. 9A), relative to those of PO4 , NO^, 

and Si(OH)4 (fig. 95), reflect the importance, if not predominance, of plankton (that is, organic detritus) as a carrier of many metals from the photic zone to the sea floor. 

More importantly, the distributions of these elements provide a critical evaluation of the minor-element stoichiometry of plankton. NO^ and PO4~ are the limiting nutrients 

to phytoplankton productivity in the photic zone of the 

ocean today (Sverdrup and others, 1942; Broecker and 

Peng, 1982; Codispoti, 1989), and Si(OH)4 may be a limiting nutrient in some ocean-margin areas (Dugdale and 

Goering, 1970; Nelson and others, 1981). NO^ and PO^" 

(and Si(OH)4) are extracted from seawater in the photic 

zone during photosynthesis and returned to the ocean at 

depth by bacterial respiration, and only a small portion of 

the initial particulate form survives to the deep-ocean sea 

floor. As a result of their cycling through the water column, the NC»3 and PO4~ concentrations in surface water 

are relatively low and can even approach zero; their concentrations increase sharply below the photic zone to the 

thermocline and then remain relatively constant below 

approximately 500-m depth.

Many minor elements show a similar depth profile. The 

variation in Cd content with depth parallels that of PO4 

concentration (Boyle and others, 1976), more so than those 

of the other minor elements. Zn and Ni depth profiles resemble that of Si(OH)4 (Sclater and others, 1976;Bruland, 1983). 

For Zn, the Zn-Si(OH)4 relation in seawater (Broecker and 

Peng, 1982) corresponds to a Zn content in opal of approximately 60 ppm. The measured Zn content of opal averages 

only 5 to 10 ppm (Martin and Knauer, 1973), much less than 

its bulk content in plankton of 110 ppm (table 1). Thus, the 

high Zn content in plankton (Collier and Edmond, 1984) and 

its much lower measured content in the siliceous fraction of

Figure 8. Minor-element versus Cu contents in modern 

plankton. Phytoplankton and zooplankton values are for 

Monterey Bay, Calif., and the North Pacific Ocean (Martin 

and Knauer, 1973); mixed plankton values are for the North 

Pacific Ocean (Collier and Edmond, 1984); and suspendedmatter values are for two samples from photic zone of the 

California Current (Sherrell, 1989). Curve labeled "plankton" 

represents a best estimate of element:Cu ratio of plankton 

(Brumsack, 1986); A-seawater values are ratios of 8-seawater 

values listed in table 3.

Geochemistry of the Phosphoria Formation at Montpelier Canyon, Idaho: Environment of Deposition B13
plankton (Martin and Knauer, 1973) indicate that Zn and, 

probably, Ni are associated predominantly with the soft parts 

of organisms, similar to Cd.

The distribution of these minor elements in seawater, 

relative to that of PC>4~, can be used to evaluate their 

measured contents in plankton (table 3). The NOJ: PC>4~ 

ratio,

6-seawater[ NO^ ] =

NO3 (deep ocean) - NO 3 (photic zone)

3- (1)

PO4 (deep ocean) - PO^ (photic zone)

based on the chemical profiles of these two ions in seawater (fig. 9), equals the average N:P atomic ratio of 

plankton of 15:1. This agreement indicates that metabolic 

and advective processes dominate their distributions in the 

deep ocean. If this conclusion is true for the minor elements as well, 5-seawater values for minor elements also 

should equal the metal:P ratios in plankton. Although the

two values for Cd closely approach each other, the 8- 

seawater values for Cr, Ni, Se, and V are higher, and for 

Zn lower, than the mean element:? ratios in plankton. The 

8-seawater value for Cu also resembles the Cu:P ratio in 

plankton, although it is 2.5 times higher than the ratio in 

plankton when the Cu concentration in bottom water (fig. 

9A) is used in the calculation, rather than the Cu concentration at 2,000-m depth (table 3).

Boyle and others (1977) explained the distribution of 

Cu in the deep ocean by a combination of biologic and 

nonbiologic processes; the nonbiologic processes involve 

scavenging of Cu throughout the water column by settling 

particles and recycling of Cu into bottom water from oxic 

pelagic sediment. These and other nonmetabolic processes 

might also explain the high 8-seawater values for Cr, Ni, 

Se, and V, but their marine chemistry has not been examined as thoroughly as that of Cu. Nonetheless, the similarities of the seawater depth profiles of many metals to those 

of NO3~, PO^~, and Si(OH)4, and the fact that their 8- 

seawater values are within a factor of 2 to 3 of the average 

compositional ratios in plankton (fig. 8; table 3), suggest

g

1000

2000

3000

4000

5000

*-«

EXPLANATION

-o P043 '

+ N03- 

0 Si(OHL

+ o 

+ o

0.1 1 10 100 1000

NUTRIENT CONCENTRATION, 

IN MICROMOLES PER KILOGRAM

\ a te * c B \ I

\ 1

» 1

 O 

» O

 O 

» 0

 o

» o

 o

» o

- » o

EXPLANATION

 o

o Mo

» V

o Cu

D Cd

+ Zn 1 %

* K>

a +'3

D O +

DO +

DO +

DO +

DO H

DO +

DO +

D O +

D 0 +

D 0 +

D O +

D 0 +

1

0.01 0.1 1 10

Mo and V CONCENTRATION, 

IN MICROMOLES PER KILOGRAM 

Cd, Cu, and Zn CONCENTRATION, 

IN NANOMOLES PER KILOGRAM

Figure 9. Nutrient (A) and metallic-ion (B) concentrations versus depth in the Pacific Ocean (from Bruland, 1983, and Collier, 

1984, 1985).

B14 The Phosphoria Formation: Its Geochemical and Biological Environment of Deposition
Table 3. Minor-element contents in organic matter and seawater

[Seawater content is the approximate elemental concentration for the Pacific Ocean at 2,000-m 

depth. Data sources for minor-element contents in organic matter are listed in table 1, and in seawater in table 4]

Element Organic matter Element:? ratio 6-seawater Seawater 

content (ppm) (organic matter) (weight ratio) content (ppb)

/~1 J

Co 

r1 ,, _ __ _

Mo - -----

Mi

TT

V 

7n

REE's (La) -

2.0

11

2

1.5

3.0

3.0

110

.6xlO"3

.26x1 0~3

3

40x1 0~3

14.00xlO~3

1 3x1 0~3

3

1 1x1 0~3

3

1 5x1 0~3

1 1x10~3

6.1xlO"3

0.10

OO19

.22

1 0

10.60

.59

.13

a no

i on

52

that the elemental composition of plankton approximates 

the average composition of marine organic detritus settling 

into the deep ocean. Ni is a major exception; it has a 6- 

seawater value 5 times larger than the Ni:P ratio in plankton (table 3). Owing to the coarseness of our sampling, as 

discussed above, only the average 6-seawater value is germane to our study, and this should be the value given by 

seawater profiles.

Nonbiologic processes, which likely influence the 

distributions of many metals in the deep ocean, might 

contribute much less to metal distributions in the shallower continental shelf and slope environments. This 

reduced influence is suggested by the Cu distribution. Its 

6-seawater value, based only on the uppermost 2,000 m 

of seawater (l.lxlO~3 ), is very close to the Cu:P ratio in 

plankton (1.4x10 ). If we consider its concentration in 

bottom water, then its 6-seawater value is 3.5xlO~3 .

The distributions of Mo, U, and REE's in seawater 

are little influenced by the biologic cycle. Mo and U concentrations are constant with depth (Broecker and Peng, 

1982; Collier, 1985; Emerson and Huested, 1991). Mo is 

mildly bioreactive (table 3), but its uptake by plankton in 

the photic zone and the remineralization of planktonic 

debris at depth by bacteria are simply insufficient to measurably affect the vertical distribution of Mo in the water 

column (fig. 9A). The content of Mo in plankton is only 

/6 that of Cd, but the Mo concentration in seawater is 

approximately 100 times greater (table 3). Thus, oxidation 

of organic matter at depth increases the Mo concentration 

above its surface-water value by less than 2 percent, somewhat less than the precision of the analyses (Collier, 

1985). This result is analogous to our inability to detect

the change in Ca2+ concentration (Broecker and Peng, 

1982), forced by phytoplankton precipitation of CaCC>3 in 

the photic zone for example, by coccolithophorids and 

dissolution at depth.

U and REE's apparently are even less bioreactive. U, 

similar to Mo, is conservative in seawater. Although REE 

concentrations in seawater increase with depth and are relatively high in handpicked diatom tests (Elderfield and 

others, 1981), their seawater profiles are explainable 

largely in terms of nonbiologic processes (de Baar and 

others, 1985). Such an explanation suggests that the REE 

contents are very low in silica, possibly similar to their 

contents in biogenic calcite (Palmer, 1985), despite the 

elevated values of single analyses.

We present the minor-element contents of plankton 

below as single, well-defined values that we consider to be 

the best averages, as suggested by Brumsack (1986). 

However, they certainly are not well defined and unlikely 

to be single values. Nonetheless, the successful application 

of this approach to ancient rocks (Piper, 1991; Piper and 

Isaacs, in press) lends support to the values recommended 

by Brumsack (1986).

Before considering the minor elements, we calculate 

the accumulation rate of PC>4~ in the Phosphoria Formation, relative to the current extraction of PO4 from 

the photic zone of the ocean by photosynthesis. The duration 

of deposition for the Meade Peak Phosphatic Shale Member 

of the Phosphoria Formation (fig. 2), based on the conodont 

and brachiopod zones of Wardlaw and Collinson (1986) 

and the time scale of Harland and others (1990), was 

approximately 7.2 m.y. The amount of P2®5 m tne Meade 

Peak Member, at Montpelier Canyon, is 7.65 kg-m P2O5

Geochemistry of the Phosphoria Formation at Montpelier Canyon, Idaho: Environment of Deposition B15
per kilogram of sediment (McKelvey and others, 1953), 

giving a mean accumulation rate for P2O5 of 0.23 mg/ 

cm per year. Assuming a composition for settling organic 

matter which equaled that of modern plankton, the rate 

of accumulation of organic matter necessary to deliver 

P2O5 to the Phosphoria Sea floor would have been 14 

mg/cm2 per year.

We can compare this accumulation rate with that for 

organic matter on the Peru Shelf, an area of high primary 

productivity with a sediment composition (Burnett and 

Froelich, 1988) similar to that of the Meade Peak Phosphatic 

Shale Member (McKelvey and others, 1953). Primary productivity of organic matter on this shelf is approximately 

200 mg/cm per year (Chavez and Barber, 1987), as high 

as that of any region in the oceans (Berger and others, 

1988). Of this amount, approximately 15 to 45 percent settles out of the photic zone (Dugdale and Goering, 1970; 

Von Bockel, 1981; Baines and others, 1994). A widely varying amount, as much as 8 percent or approximately 16 mg/ 

cm per year, escapes oxidation in the water column altogether and accumulates on the sea floor (Reimers and Suess, 

1983). This amount corresponds closely to the required 

value for the Phosphoria Formation of 14 mg/cm2 per year 

and to the amount predicted for a water depth of 500 m 

(Suess, 1980).

It may not be entirely surprising that primary productivity in these two environments, separated in time by about 

260 m.y., is similar. Could the rate of primary productivity 

of 200 mg/cm2 per year represent an approximate upper limit 

for the photic zone, both past and present? The depth of the 

photic zone on the Peru Shelf today is only approximately 

30 m (Von Bockel, 1981; L.A. Codispoti, oral commun., 

1992), limited by the high content of particulate organic 

matter in the water. Higher productivity would yield more 

organic matter within any volume of water in the photic zone 

but also limit even more the depth to the base of the photic 

zone. As this depth decreases relative to that of the mixed 

layer, phytoplankton growth per area of sea surface must 

decrease, owing largely to mixing below the photic zone. At 

some photic-zone versus mixed-layer depth, mixing will 

carry phytoplankton into the aphotic zone at a rate that must 

limit total productivity and the flux of organic matter to the 

sea floor. The boundary conditions are a photic zone of 0-m 

depth and a photic zone of a depth equal to the mixing depth, 

or to the depth of the sea floor. Both the Phosphoria Sea and 

the Peru Shelf might have achieved a self-limiting maximum 

productivity of organic matter, approaching 200 mg/cm2 per 

year. We note that the average productivity of the Phosphoria 

Sea remained near this inferred maximum, if indeed it is a 

maximum, for about 7.2 m.y.

The ratio of the depth of the photic zone to that of 

the mixed layer must surely influence other aspects of the 

flux of organic matter to the sea floor. For example, the 

loss of organic matter as the fecal debris of grazing zooplankton, relative to the loss of phytoplankton from the

aphotic zone through mortality, might decrease with a 

decrease in the above ratio. We can only speculate how 

the loss of phytoplankton by mortality versus that by 

grazing might influence the overall flux and composition 

of accumulating organic matter. Our assumption is that 

the two different loss terms will not change the composition of organic matter.

The minor-element contribution from settling organic 

matter (plankton) can be evaluated in the same way as for 

P2O5 . Consider Cu: If the average marine Cu content in the 

Phosphoria Formation is 95 ppm (fig. 6) and the stoichiometry of accumulating organic matter was the same as that of 

modern plankton (table 3), then the accumulation rate of 

organic matter required to deliver Cu to the sea floor was 14 

mg/cm2 per year. Calculations based on the marine Mo and 

Zn contents (fig. 6) that is, in excess of the terrigenous 

contribution (WSA, fig. 5) and their contents in plankton 

(table 1) give rates of 16 and 17 mg/cm2 per year, respectively. Cd requires a slightly lower rate.

The accumulation rate of organic matter required to 

deliver Ni to the sea floor is approximately 40 mg/cm2 per 

year, or 3 times higher than that based on the contents of 

P2O5 and other trace elements. Although a similar result was 

obtained for the Monterey Formation, a phosphatic formation 

of Miocene age (Piper and Isaacs, in press), no simple explanation seems adequate to account for such a high rate. The 

8-seawater[Ni] value suggests that Ni content of organic 

matter is 5 times higher than our stoichiometric model allows 

(table 3). It seems unlikely that the mean Ni content of 

plankton could be 35 to 40 ppm without such high contents 

being observed (Martin and Knauer, 1973, Collier and 

Edmond, 1984). Although only 25 ppm is required to bring 

the Ni content of plankton into agreement with the relation 

between Ni and the above minor elements and P2O5, this 

value still seems high to us. Its disagreement with the accumulation rates of P2O5, Cu, Mo, and Zn is unfortunate 

because Ni should provide the best estimate of primary 

productivity (see discussion below).

An alternative approach to examining the contribution 

of minor elements from organic matter is provided by ratios 

of the minor elements (fig. 10). The major reason for this 

approach is the similarity in the marine geochemistry of 

several of the minor elements. They precipitate as sulfides 

(Jacobs and others, 1985, 1987; Landing and Lewis, 1991) 

and should be retained in the sediment as metal sulfides 

during and after early diagenesis. Phosphate, in contrast, 

is retained as carbonate fluorapatite (Burnett and others, 

1988; Froelich and others, 1988). This difference could result 

in fractionation between PO|~ and the minor elements during diagenesis and weathering, and explain the absence of 

correlation between apatite and the bioreactive minor elements (table 2). The minor elements, in contrast, should 

be less fractionated from each other, possibly allowing for 

a comparison between the interelement ratios in rocks, plankton, and seawater through simple x-y plots.

B16 The Phosphoria Formation: Its Geochemical and Biological Environment of Deposition
140

50 100 150 200 50 100 150 200

400

Z 350

O

250

100

50

0

0 50 100 150

Cu* CONTENT, 

IN PARTS PER MILLION

200

-2

" 6

-10

H

\

\

50 100 150

Cu* CONTENT, 

IN PARTS PER MILLION

200

Figure 10. Minor-element versus Cu contents in marine fraction of the Phosphoria Formation at Montpelier Canyon 

during glacial and interglacial episodes. See figure 8 for definition of curves, except for figure 10H, in which curve 

was fitted visually to data.

Geochemistry of the Phosphoria Formation at Montpelier Canyon, Idaho: Environment of Deposition B17
A second reason to consider minor-element ratios is 

that sediment-accumulation rates are seldom or poorly 

known. It might be argued that they are too poorly known 

even for the Phosphoria Formation to permit an evaluation 

of rates of primary productivity, as we did above. The 

agreement between the different calculations, however, 

suggests otherwise. The x-y plots should permit an even 

more precise estimate of elemental contributions from the 

biogenic fraction and identify any possible hydrogenous 

fractions. This procedure is aided when the minor elements exhibit strong interelement correlations. Simple x-y 

plots of minor elements will clearly show those elements 

that are enriched above a biogenic contribution.

This procedure points up the desirability of a single 

master element that defines the marine organic source, as 

A12O3 defines the detrital source. In this discussion, we 

use Cu as the master element for the marine biogenic 

(planktonic) fraction. Several aspects of its chemistry recommend it.

1. The content of Cu in plankton (table 3) exceeds that 

of most other minor elements (Martin and Knauer, 

1973; Broecker and Peng, 1982; Bruland, 1983; Collier and Edmond, 1984; Brumsack, 1986), and so its 

biogenic signal might be strongly imprinted in rocks 

that initially had a high content of organic matter.

2. Cu accumulates in oxic sediment in strong association 

with Mn (Piper, 1988), but it precipitates virtually 

completely from seawater under seawater sulfatereducing conditions as a sulfide (Jacobs and others, 

1985). Its absolute concentration in seawater should 

make the biogenic input, however, greater than or, at 

worst, approximately equal to the hydrogenous input 

of many sedimentary deposits. If we assume redox 

and advection conditions similar to those in the Cariaco Trench (Jacobs and others, 1987) and a Cu concentration in seawater advecting into the basin equal 

to that in seawater at 2,000-m depth (table 3), then the 

accumulation rate for Cu that would precipitate as a 

sulfide is approximately 0.1 |ig/cm2 per year, slightly 

less than its accumulation rate in the biogenic fraction 

of 0.22 u,g/cm2 per year.

3. Cu should be retained in organic-matter-enriched sediment. The current organic-matter content of the Phosphoria Formation (Medrano and Piper, 1992) is high 

in terms of sedimentary rocks, even though it represents only approximately 1 percent of the rain rate of 

14 mg/cm2 per year. It is considerably less than the 

organic-matter content of surface and near-surface 

sediment accumulating currently on the Peru Shelf, 

but it likely exceeded Peru Shelf values at the time of 

deposition. Such a high accumulation rate for organic 

matter allows sulfate reduction in the pore water to 

extend to or near the surface (Froelich and others, 

1988). One result of such pore water chemistry is 

strongly negative 813C values, measured in Peru Shelf

pore water (Glenn and Arthur, 1988) and in the CO§ 

of carbonate fluorapatite in these samples (fig. 10f/) 

and in other samples from the Phosphoria Formation 

(Piper and Kolodny, 1988). Cu is quickly recycled 

into the overlying water under oxic pore-water conditions (Boyle and others, 1977; Fischer and others, 

1986), but it should be retained under sulfate-reducing 

conditions. This interpretation is based on the low solubility of Cu2S or CuS, as reflected by its precipitation under natural seawater sulfate-reducing conditions 

(Jacobs and others, 1985).

4. The marine fraction of Cu correlates with those of several other minor elements in the Phosphoria Formation 

(fig. 10; table 2), allowing a precise comparison of 

minor-element ratios in the Phosphoria Formation with 

those in different modern marine fractions through the 

use of simple x-y plots. The absence of such correlations does not preclude the use of this approach, inasmuch as the strength of present correlations should 

depend more on the composition of host, authigenic 

phases than on that of source phases. For elements that 

fail to correlate with Cu or some other master element, 

many samples may have to be analyzed to document an 

average minor-element ratio. As we demonstrated 

above, the marine fractions of Cu and ?2O5 give similar accumulation rates for organic matter, yet the two 

do not correlate (table 3).

Plots of the contents of minor elements versus Cu 

(fig. 10) reinforce the calculations of their accumulation 

rates. We should have expected this result because both 

schemes are based on the assumption that the composition 

of deposited organic matter was the same as that of modern plankton. Nonetheless, our conclusion from this exercise is that primary productivity in the photic zone 

provided the dominant marine fractions of Cu, Cd, Mo, 

and Zn to the floor of the Phosphoria Sea. Identification of 

the Ni source will require further work on both modern 

and ancient deposits.

Hydrogenous Fraction

The low REE, Cr, and V contents in organic matter 

(table 1) and their relatively high contents in Montpelier 

Canyon rocks require that these elements had a hydrogenous marine source, in addition to their detrital and biogenic sources. Before accounting for these high values, we 

examine the geochemical properties of seawater under 

which these and all minor elements in the modern ocean 

accumulate through inorganic processes. Their behaviors 

have been thoroughly examined under redox conditions of 

oxygen respiration to sulfate reduction. Their precipitation-dissolution under varying redox conditions is driven 

by the bacterial oxidation of organic matter, for which the 

three major electron acceptors are O2, NOJ, and 804". 

The major seawater redox reactions that drive the precipiB18 The Phosphoria Formation: Its Geochemicai and Biological Environment of Deposition
tation-dissolution reactions of individual minor compounds 

can be written as follows:

(2)

(3)

(4)

Oxygen respiration:

10CH2O + 10O2 + 10CaC03 =

20HCOJ + 10Ca2+

Denitrification:

10CH2O + 8 NOJ + 2CaCO3 =

4N2 + 12HCO3 + 4H2O + 2Ca2+

Sulfate reduction:

10CH2O + 5 SO|~ + 5CaCO3 =

5HS~ + 15HCO3 + 5Ca2+.

The formula for organic matter is better represented by

elements)0 QX.

Its minor-element contents are listed in table 1. The hierarchy of these three reactions is determined by the reaction 

yielding the greatest free energy (Froelich and others, 

1979). Thus, reaction 2 proceeds until O2 decreases to a 

concentration such that denitrification yields equal free energy, and so on.

Within several ocean-margin basins, advection is 

weak enough and the flux of settling organic matter great 

enough to allow depletion of successive electron acceptors 

with depth in the water column (Richards, 1965; Emerson 

and others, 1979). These basins include the Gulfo Dulce 

(Costa Rica), Baltic Sea, Black Sea, Cariaco Trench (Venezuela Shelf), and several fiords at high latitudes. Concentration profiles of O2, NOJ, and H2S show that O2 is 

abundant in the surface mixed layer, owing to mixing with 

atmospheric O2 and to photosynthesis, but decreases in the 

water column below the mixed layer, in response to oxygen respiration, predominantly by bacteria. O2 concentration is low at intermediate depth (Emerson and others, 

1979), the depth interval over which NOJ concentration 

decreases in response to denitrification. O2 and NO^ are 

virtually absent in bottom water, that is, in the presence of 

H2S under conditions of sulfate reduction (eq. 4). Sulfate 

reduction continues into the sediment until SO^" is 

depleted in the pore water, at which depth the residual 

labile organic matter is broken down through methanogenesis and fermentation reactions (equations not shown). In 

the Santa Barbara Basin (California Continental Borderland) and Darwin Bay (Galapagos Islands), bacterial respiration in the bottom water does not proceed beyond 

denitrification (Richards and Broenkow, 1971; Sholkovitz 

andGieskes, 1971).

By contrast, the deep ocean exhibits oxygen respiration throughout. Denitrification occurs at intermediate 

depth, in the oxygen-minimum zone (OMZ), but only 

along the Peru and Mexico Shelves in the eastern Pacific

Ocean (Codispoti, 1980) and at a few other localities of 

limited areal extent in the Atlantic Ocean (Calvert and 

Price, 1971) and the Indian Ocean. Although the OMZ is 

present at intermediate depth throughout most of the 

oceans, only in these relatively small areas is the balance 

between advection in the OMZ and organic-matter productivity in the photic zone such that bacterial oxidation of 

settling organic matter drives down the O2 content within 

the OMZ of the water column to a value that promotes 

and maintains denitrification.

On a single cruise to the Peru Shelf in 1976, low H2S 

concentrations were measured in the OMZ (Dugdale and 

others, 1977), indicating that sulfate-reducing conditions 

occur in the OMZ of the open ocean, at least briefly. The 

existence of such conditions is easily understood because 

the maximum O2, NO^, and SO^" concentrations in seawater are approximately 0.36, 0.04, and 28 mM, respectively. From equations 2 through 4, NO^ has only about 

15 percent the oxidizing potential of O2 and only 0.1 percent that of SOj". The NO3 should be relatively quickly 

utilized once O2 is depleted, pushing bacterial respiration 

into the field of sulfate reduction. It is somewhat surprising that H2S has not been detected in the water column 

more often than on this single occasion.

Minor elements that act as electron acceptors include 

Cu2+, CrO2~, Fe(OH)3 , MnO2, MoO^, SeOjf, 

UO2(CO3)^ (or, possibly, UO2(CO3 )3V"), and VOj~. 

Their reduction can be calculated from Nernst equations 

and ordered according to Eh (table 4), where we merely 

omit the half-cell reaction representing the oxidation of 

organic matter:

CH2O + H2O = CO2 + 4H+ + 4ef (5)

If this half-cell is included, as was done in equations 2 

through 4, Fe(III) reduction can be written in the following form:

10CH2O + 40Fe(OH)3 + 60HCO3 + 70Ca2+ = 

70CaCO3 + 40Fe2+ 100H2O. (6)

Its reduction occurs before sulfate reduction (table 4) and 

toward the end of denitrification (Berner, 1980; Bender 

and others, 1989); CrO^"" also is reduced under conditions 

of denitrification (table 4; Murray and others, 1983). For 

completeness, table 4 includes the reduction of other redox 

reactions (for example, the reduction of Re and I), although they are not discussed further here.

The minor-element contents of inorganic paniculate 

phases in seawater reflect these redox conditions. Under 

conditions of oxygen respiration, inorganic precipitates consist mostly of Fe and Mn oxyhydroxides (Landing and Bruland, 1987). Other minor elements are undersaturated in oxic 

seawater, but high concentrations of several minor elements 

in deep-ocean ferromanganese nodules indicate that a small

Geochemistry of the Phosphoria Formation at Montpelier Canyon, Idaho: Environment of Deposition B19
Table 4. Half-cell reactions, assuming standard-state conditions

[Activity coefficients calculated by using the Debye-Hiickel equation (Wagman and others, 1982); a, effective ionic diameter used in the Debye-HUcke! 

equation (from Nordstrom and Munoz, 1985, or estimated from their table 7-4). Seawater concentrations of dissolved species (in moles per kilogram) 

are those at approximately 2,000-m depth. The pH was taken to be 7, although it is closer to 8 under conditions of oxygen respiration and 7.5 under 

denitrification. Data sources for thermodynamic constants and species concentrations in seawaten 1, Bruland (1983); 2, Latimer (1953); 3, Wagman and 

others (1982); 4, Landing and Bruland (1987); 5, Broecker and Peng (1982); 6, Murray and others (1983); 7, Elderfield (1970); 8, de Baar and others 

(1985); 9, Langmuir (1978); 10, Jacobs and others (1985); 11, Collier (1984); 12, Boyle and others (1976); 13, Collier (1985); 14, Emerson and Huested 

(1991); 15, Anbar and others (1992); 16, Koide and others (1986)]

Eh

1.099

.805

.704

.702

.698

.545

.320

.296

.153

.013

-.018

-.040

-.055

-.055

-.137

-.162

-.170

-.175

-.188

-.209

-.210

-.250

-.250

3Co(2J) + 4H20(,) -

2H2O([j -> O_ + <

Mn(2J} + 2H20(1) ->

Half-cell reaction

>C°3°4(s) + 2e" + 8H(aq)

le + 411(3^)

Mn02(s) + 2e-+ 4H(aq)

!(aq) + 3H2°(1) ~* IO 3 (aq) + 6e + 6H(aq)

N2(g) + 6H2O(!) -> :

Cr(OH)3(s) + H20(I)

Ce(aq) + 2H2O(I) -»

P-» 2+ , -JTT f\ .

he(aq) + 5M2U(1) ->

Cu+ -> Cu2+ + e~

>MO~ J. 1fta~j- I'JH + ifNU 3 (aq) + lUe + l^M(aqj

-> CrO 4~(aq) + 3e"+ 5H(aq)

Ce02(s) +e-+4H(aq)

Fe(OH)3(s) +e-+3H(aq)

U02(s) + 2HC03(aq) -» U02(C03)i-(aq) + 2e-+ 2H(aq)

2Fe(2q-) + 3H20(1) -)

V204(s) + 4H20(1) -

HS(aq) + 4H20(1) -j

Cu2S(s) + 4H200) -

CdS(s) + 4H20(I) ->

ZnS(s) + 4H20(1) ^

Mo02(s) + 2H20(1) -

MoS2(s) + 12H20(,)

NiS(s) + 4H20(1) ^

FeS(s) + 4H20(I) ^

U02(s) + 3C03-(aq)

Re°2(s) + 2H2O(1) -

HS(aq) + 4H20(I) ->

> Fe203(s) + 2e-+ 6H(aq)

^2H2V04-(aq) + 2e-+ 4H(aq)

> 804 (aq) + 8e + 9H(aq)

^ SO 4 (aq) + 2Cu(aq) + 8e + 8H(aq)

SO 4 (aq) + ^(aq) + ^e + ^H^

SO 4 (aq) + Zn(aq^) + 8e + 8H(a^)

-^MoOj-(aq) +2e-+4H(aq)

-> MoO J-(aq) + 18e-+ 2SOf(aq) + 24H(aq)

SO 4 (aq) + ^(aq) + 8e + 8H(aqj

^O 4 (aq) + rC(aq) + 8e + 8H(aq)

-^U02(C03)f(aq) + 2e-

» ReO4 (aq) + 3e~ + 4H(aq)

, §o^~ + 8e~ + 9H +

a

6

6

3

4

3

4

9

6

4.0 

4.5

6

5.6

3.5 

4

4 

2.5

4 

6

4 

6

4.5

4.5 

5

4 

6

4 

6

4.5 

5.5

4.5

3.5 

4

Seawater concentration in moles 

per kilogram or partial pressure

Co2+ = 2.04x 10'11

O2 = 0.2 atm

Mn2+ = 2.5 x 10"10

r = 5.oxio-9,

IO 3 =4.5x 10~7

NO 3 =3.9xlO~5, 

N2(g) = 0.8 atm

CrO 2~= 4.04 x 10~9

Ce3+ = l.Ox 10~n

Fe2+ = 1.20 x 10~9

ZCu = 3.59 x 10~9, 

Cu+ = 1.80 x 10"9

HCOg =2.47x 10~3, 

UO2(CO3)|" = 1.26x 10"8

See above

H2VO4 = 4.0 x 10~8

HS~ = 7.75 x 10~25,

SO 2-= 2.8 x 10"2, 

Cu+ = 3.59 x 10"9

SO|"=2.8x 10~2, 

Cd2+ = 6.94 x 10~10

SO|"=2.8x 10~2 

Zn2+ = 5.97 x 10"9

MoO|~= l.lx 10"7

SO|"=2.8x 10"2, 

MoO^"= l.lx 10~7

SOj~=2.8x 10"2, 

Ni2+ = 8.01 x 10"9

SO|"=2.8x 10~2, 

Fe2+ = 1.20x 10~9

CO 3 = 1 x ID"4, 

UO2(CO3)^~ = 1.26 x 10~8

ReO4 =4.5x 10"n

HS~ = 2.8 x 10~2, 

SO J~= 4.3x10^

Reference

1,2

3

3,4

1,3

3,5

2,6,7

3,8

3,4

1,3

3,5, 

9

2, 11

3,5

1,3, 

10

1,3, 

5, 12

1,3,5

3, 13

3, 13, 14

1,3, 

5, 10

1,3,4

3,5, 

9

3, 15, 16

3,5

fraction coprecipitates with the Fe and Mn oxide phases 

(Goldberg, 1961) or is scavenged by these oxides (Goldberg, 

1954; Balistrieri and others, 1981; Li, 1981; Clegg and 

Sarmiento, 1989). Pelagic sediment, however, remains a 

minor sink for most metals, except Mn (Lyle and others,

1984). Their high metal contents reflect very low bulk-sediment-accumulation rates rather than high metal-accumulation rates.

Within basins that exhibit sulfate reduction in bottom 

water, Fe2+, Cu+, Cd2+, and Zn2+ precipitate as sulfides,

B20 The Phosphoria Formation: Its Geochemical and Biological Environment of Deposition
f\

004 is reduced and precipitates as Cr(OH)3 (Brewer 

and Spencer, 1974; Emerson and others, 1979; Jacobs and 

others, 1985, 1987), and VC$~ may be reduced to V2O3 

or, more likely, some form of V(OH)3 (Wanty and Goldhaber, 1992). Mo and U also are removed from the bottom 

water (Bruland, 1983; Anderson and others, 1989; Emerson and Huested, 1991), possibly according to the reactions listed in table 4. Although this environment is 

limited in areal extent, it represents an important sink for 

several elements in the ocean today (Kolodny and Kaplan, 

1970; Bertine and Turekian, 1973; Cutter, 1982; Jacobs 

and others, 1987; Haraldsson and Westerlund, 1988; Emerson and Huested, 1991; Klinkhammer and Palmer, 1991). 

The much higher concentration of Mo in open-ocean seawater relative to the other minor elements (tables 1, 3) and 

its low content in detrital phases and organic matter (table 

1) make it possibly the most important minor element for 

identifying rocks that accumulated in a sulfate-reducing 

environment. Its distribution in modern sediment of the 

Cariaco Trench supports this contention (Jacobs and others, 1987).

Of the elements that precipitate under the intermediate 

redox conditions of denitrification, U, V, and Cr should have 

the highest concentrations in sediment, although there are 

possible problems with V and U. The absence of a clear 

signal of U reduction and precipitation in the water column 

under denitrifying conditions (Anderson and others, 1989) 

makes its distribution in sedimentary deposits possibly less 

diagnostic of bottom-water conditions than that of other 

minor elements. Its possible occurrence as UO2(CO3)3~ 

(Goldberg, 1963; Anderson and others, 1989), instead of 

UO2(CO3)2 , an equally thermodynamically stable species 

(Langmuir, 1978), may account for this; UO2(CO3 )3~ 

should be reduced well into the zone of sulfate reduction 

(Eh=-0.21), rather than in the zone of denitrification. The 

precipitation of V occurs near the NO3 / SO^~ -reduction 

boundary, which we define rather arbitrarily by the precipitation of the least soluble minor-element sulfide (Cu2S) 

under consideration. A slight error in either the V or Cu 

half-cell reactions could push the precipitation of V into 

the zone of sulfate reduction, although the close association 

of V with Cr in some phosphate deposits (Piper, 1991; Piper 

and Isaacs, in press) indicates initial precipitation in the less 

reducing environment.

REE's have a unique distribution in O2-depleted seawater environments. They maintain their 3+ valence state 

throughout the marine environment except for Ce, which is 

oxidized to the 4+ valence state under oxic conditions (table 

4). Within the zone of denitrification in the Cariaco Trench 

and the Black Sea, immediately above the zone of sulfate 

reduction (Hashimoto and others, 1983; Codispoti and others, 

1991), the concentrations of the 3+-valence-state light REE's 

(La, Pr, Nd, and Sm) show minimums (de Baar and others, 

1988, German and others, 1991; Schijf and others, 1991). In 

both basins, their concentrations increase sharply at greater

depth, toward the NO3 / SO^ -reduction interface and into 

the sulfate-reducing water. Within the zone of denitrification 

in the OMZ in the eastern Pacific Ocean, the concentrations 

of the 3+-valence-state REE's also show minimums. They 

are scavenged by particulate phases, most likely oxyhydroxides of Mn (German and others, 1993) or Fe (Sholkovitz, 

1993), under mildly denitrifying conditions and released to 

solution under sulfate-reducing conditions. The carrier phase 

in the zone of denitrification in the OMZ of the eastern Pacific 

Ocean is unlikely an Mn phase. The concentration of dissolved Mn shows a maximum, whereas those of the REE's 

show minimums (Klinkhammer and Bender, 1980; de Baar 

and others, 1985; German and others, 1991). Thus, the mechanism whereby the 3+-valence-state REE's are transferred 

to the sea floor within the hydrogenous fraction remains 

problematic, although the redox conditions under which their 

removal to the sediment occurs are well established.

Ce is somewhat more complex than the other REE's. 

Its occurrence as insoluble Ce(OH)4 under oxic conditions 

(table 4) accounts for the negative Ce anomaly (see fig. 7) 

of seawater. Its reduction to the 3+ valence state under 

denitrifying conditions has been proposed to account for 

the increase in Ce concentration and the more positive Ce 

anomaly in the OMZ of the eastern Pacific Ocean and the 

Cariaco Trench (de Baar and others, 1988).

In the Black Sea, however, the trend of the Ce anomaly is different (German and others, 1991; Schijf and others, 1991): the dissolved-Ce concentration in the zone of 

denitrification decreases to a minimum, but so does the 

anomaly. This trend parallels those of the 3+-valence-state 

REE's, which mirror the maximum concentration of particulate Mn. All the REE's are scavenged in the upper part 

of the zone of denitrification, Ce more so than the other 

REE's. This distribution should be expected if Mn is the 

dominant carrier, because Ce4+ reduction occurs at a lower 

Eh than MnO2 reduction. The difference in the distribution 

of the Ce anomaly between the Black Sea and the Cariaco 

Trench places an unfortunate constraint, at present, on the 

uniqueness of the Ce anomaly in unraveling the geochemistry of ancient sediment (Wright and others, 1987).

The low concentrations of Ce and other minor elements 

in seawater severely limit their contributions to the actual 

oxidation of organic matter but contribute significantly to 

the sensitivity with which their contents in the marine fraction 

of sediment might record redox conditions in the water column. Let us consider Mo. Its concentration in seawater is 

approximately 10~7 M, or !/400 that of NO3 . Clearly, it will 

contribute insignificantly to the oxidation of organic matter; 

and of the minor elements considered here, it is the most 

abundant in seawater. Assuming that the bottom-water of the 

Phosphoria Sea was sulfate reducing, and neglecting for the 

moment the conditions necessary to bring about the depletion 

of O2 and NO3 , then for the Phosphoria Formation, which 

had an average bulk-sediment-accumulation rate of about 

1.7 mg/cm2 per year, 35 percent of which was a terrigenous

Geochemistry of the Phosphoria Formation at Montpelier Canyon, Idaho: Environment of Deposition B21
fraction (Medrano and Piper, 1992), the detrital Mo-accumulation rate would be 0.0016 Jig/cm2 per year. If primary 

productivity was similar to that on the Peru Shelf, as we 

demonstrated above, the Mo-accumulation rate within 

organic matter (that is, within planktonic debris) was 0.04 

jig/cm2 per year. Allowing that the hydrodynamics of the 

Black Sea represents that of the Phosphoria Shelf, the hydrogenous Mo-accumulation rate would have been 2 |ig/cm2 per 

year, or about 850 to 2,000 ppm. This value is based on a 

bottom-water residence time ranging from 350 to 1,000 yr 

(Ostlund, 1974; Murray and others, 1991) and an Mo content 

in seawater of 10 ppb, to which we must add the amount of 

Mo added to the bottom water by the oxidation of organic 

matter, or approximately 0.1 fig/cm2 per year. (The adjustment is small for Mo, but it is significant for Zn and the 

other minor elements.) Such an accumulation rate requires 

the Mo content of the Phosphoria Formation to be 40 to 100 

times higher than that observed, if Mo accumulation occurred 

under sulfate-reducing conditions. We could change several 

factors to reduce this value. The thickness of the bottom 

water could be reduced, but doing so would change other 

factors, for example, the residence time of bottom water, 

which, in turn, would increase the flux of Mo to the sediment. 

Indeed, observations of C^-depleted environments of the 

open ocean (Sholkovitz and Gieskes, 1971; Codispoti, 1980) 

suggest that the residence time of bottom water is likely to 

have been a few years, rather than tens or even hundreds of 

years. Even though this calculation is extremely simplistic, 

it suggests to us that the bottom water of the Phosphoria Sea 

could not possibly have been sulfate reducing for any significant period of time. The five samples with high Mo 

contents (fig. 10C) suggest that bottom-water Eh might have 

slipped into the field of sulfate reduction, but only very 

briefly.

Assuming that primary productivity in the Phosphoria 

Sea equaled that on the Peru Shelf today, as seems plausible from our earlier calculation, we can expect the bottom 

water to have been denitrifying. The high contents of Cr, 

V (fig. 6), and REE's (fig. 7), in excess of a detrital and 

marine biogenic source, strongly support the interpretation 

of a bottom-water chemistry similar to that on the Peru 

Shelf. The concentration of marine Cr can be used to calculate the residence time of the bottom water, the reverse 

of the calculation for Mo-accumulation rate. The average 

Cr content in the Phosphoria Formation is 1,200 ppm, giving an accumulation rate of 2.9 u,g/cm2 per year. If the 

denitrifying layer was 500 m thick (its approximate thickness on the Peru Shelf) and the initial Cr content of seawater was 0.21 ppb (table 3), the residence time of bottom 

water was 3.6 yr. A similar calculation based on La content gives a shorter residence time of 0.62 yr. The agreement, which is certainly within reason, can be improved 

by a greater efficiency in the removal of La than of Cr, as 

clearly suggested by their distributions in NO^-reducing 

seawater (Murray and others, 1983; German and others,

1993). However, a residence time of much less than 1 to 3 

yr seems unlikely to us.

This calculation of the accumulation of marine fractions 

on the sea floor demonstrates the balance between primary 

productivity in the photic zone of the ocean and advection 

of the bottom water. The relation is further represented by 

the following calculation. Primary productivity of organic 

matter in the photic zone produces 200 mg/cm2 per year, 

approximately 25 percent of which must be oxidized in the 

bottom water. Assuming an initial NO^ concentration in the 

bottom water of 40 |iM, the maximum residence time for 

the bottom water that precludes sulfate reduction is 2 to 3 

yr. The greatest uncertainty in this estimate stems from the 

absence of information on the initial O2 content of seawater, 

because an unknown amount of the O2 present in the water 

when it left the surface ocean (approx 8 mL/L) was surely 

reduced before it reached the sea floor at the Montpelier 

Canyon site. We have made other rather gross assumptions, 

some stated and some unstated, but the overall agreement of 

these calculations suggests that they give a reasonable first 

approximation of the bottom-water residence time versus 

surface-water productivity.

The distribution of V clearly cannot be treated in the 

same way as we treated those of Cr and La. The content 

of V is much lower in the Phosphoria Formation than that 

of Cr (fig. 6), but, in seawater, its concentration is much 

higher than that of Cr (table 3). There is far too little V 

relative to Cr (fig. 11) for the bottom water to have been 

strongly denitrifying. The reduction of V at a lower Eh 

(table 4) offers one possible explanation. Throughout the 

deposition of much of the Meade Peak Phosphatic Shale 

Member, the OMZ was denitrifying enough to precipitate 

Cr; only for relatively brief periods of time was it reducing to the point that V also precipitated. The high Mo contents in five samples (fig. 10C) suggest, albeit weakly, that

4000 

§3500 

" - 3000 

2500

g S 2000

* H 

£« 1500

2 1 °°°

i i

500

0

EXPLANATION

 Interglacial 

o Glacial

500 1000 1500 2000

V* CONTENT, 

IN PARTS PER MILLION

2500 3000

Figure 11. Cr versus V content in marine fraction of the 

Phosphoria Formation at Montpelier Canyon during glacial 

and interglacial episodes. Seawater curve shows extrapolated 

values at approximately 2,000-m depth (table 3).

B22 The Phosphoria Formation: Its Geochemical and Biological Environment of Deposition
the OMZ might have exhibited sulfate reduction for even 

briefer periods of time. Examination of other sections of 

the formation, such as the Vanadiferous Zone (a discrete 

section within the Meade Peak Phosphatic Shale Member 

with unusually high V content; Love, 1961; McKelvey 

and others, 1986), might show higher Mo contents and, 

thus, evidence of sulfate reduction. In such sediment, the 

V:Cr ratio should increase toward its seawater value of 

8.6.

The V-Cr relation (fig. 11) shows that, on average, the 

bottom water was more reducing during glacial than during interglacial episodes; the sediment deposited during 

glacial episodes has higher V:Cr ratios. The lower La content in these same rocks (fig. 7A) supports this interpretation. These trends raise the question whether primary 

productivity was higher in the photic zone (greater flux of 

settling organic matter) or advection lower in the bottom 

water (longer residence time) during the glacial episodes. 

Glaciation would have lowered sea level, possibly by 75 

m (Piper and Kolodny, 1988). Primary productivity might 

have been higher during these episodes, although we have 

no supporting evidence to offer for this interpretation. A 

probable change in the rate of input of terrigenous debris 

further complicates any interpretation of the distribution of 

any single minor element. Although we are unable to identify the exact change in the environment that led to a 

change in the V:Cr ratio and La content between these two 

populations, the separation of the two populations lends 

support to earlier interpretations of the lithology (Sheldon, 

1984) and stable-isotope data (Piper and Kolodny, 1988) 

that the Phosphoria Formation was deposited during a 

time of glacial events.

CONCLUSION

The high precision of the major-element-oxide and 

minor-element analyses of this study allows us to identify, 

for the Phosphoria Formation, the current host fractions 

for minor elements (detrital material, organic matter, apatite, biogenic silica, dolomite, calcite, and trace-metal sulfides) and their source phases at the time of deposition 

(detrital debris, marine biogenic matter, and marine 

hydrogenous phases). The interelement relations between 

A12O3 and K2O, TiO2, Co, Ga, Li, Sc, and Ba approach 

those of the standard WSA and NASC. These relations 

indicate that the detrital fraction is the sole host for these 

elements and that it had a terrigenous source. The strength 

of these relations allows us to determine the detrital contribution of the other minor elements (Cd, Cr, Cu, Mo, Ni, 

U, Zn, and REE's) from the A12O3 content of each individual sample.

The difference between the bulk content and the 

detrital contribution of minor elements to each sample represents the marine contribution. The interelement relations

between Cu, Mo, and Zn in this marine fraction of the 

rocks approximate those in modern plankton, suggesting 

that these minor elements and, possibly, Cd and Ni had a 

predominantly biogenic source. During Phosphoria time, 

the rate of accumulation of organic matter on the sea floor 

at Montpelier Canyon necessary to provide these elements 

to the sediment was approximately equal to the rate of 

accumulation of organic matter on the present-day Peru 

Continental Shelf, an area of exceptionally intense coastal 

upwelling and high primary productivity in the photic 

zone. Thus, primary productivity must have been similarly 

high in the Phosphoria Sea.

The enrichments of Cr, V, and REE's required accumulation directly from seawater, a hydrogenous input. On 

the basis of their chemical properties and distributions in 

O2-depleted environments of the ocean today, their accumulation from seawater would have been enhanced by 

conditions of denitrification in the bottom water. The rate 

of accumulation of Cr and La in this fraction gives a residence time for bottom water in the basin of approximately 

1 to 3 yr. The relation between these two elements and 

REE's furthermore suggests that the bottom water was 

more strongly denitrifying during earlier-proposed glacial 

episodes than during interglacial episodes of the Permian 

Period.

Sulfate reduction was not established in the bottom 

water for any significant period of time. This interpretation 

is based on the preservation of a strong, seawater-derived 

REE enrichment throughout the Phosphoria Formation and 

the absence of an enrichment of Cu, Cd, Zn, and, particularly, Mo above a biogenic input. Otherwise, this group 

would have accumulated as inorganic sulfides, in addition 

to accumulating in an organic fraction (a slight enrichment 

of Mo was detected, but in only five samples), and REE's 

would not have accumulated from seawater at all but 

would have had solely a terrigenous source. A hydrogenous input of Mo, which could have been 100 times more 

abundant than Cd, would have been easily detected. Sulfate reduction was restricted to the sediment pore water; it 

likely closely approached the benthic boundary layer but 

did not extend into the water column.

No single environment in the ocean today represents 

a perfect analog of the Phosphoria Formation. Chemical 

and biologic processes were similar to those on the Peru 

and Namibia Shelves, both of which exhibit intense 

upwelling, resulting in high primary productivity in the 

photic zone apparently across much of the shelf, and denitrification in an OMZ at intermediate depth (Calvert and 

Price, 1971; Codispoti, 1980). In the section above entitled 

"Introduction," however, we noted possibly strong differences between the Phosphoria Sea and these shelf environments with regard to sea-floor bathymetry, tectonism, and 

proximity to major ocean currents. We still offer no example of a shelf that receives detritus at a rate less than 1 

mg/cm2 per year. The overall bathymetry and isolation

Geochemistry of the Phosphoria Formation at Montpelier Canyon, Idaho: Environment of Deposition B23
from major currents, however, might be represented by the 

East China Sea, although the differences, again, are noteworthy. The Phosphoria Shelf was situated on the west 

coast of the United States, not an east coast of Asia. It was 

at much lower latitude (Sheldon, 1964), within the belt of 

the easterly trade winds. Low rainfall limited the influx of 

terrigenous debris, and the continuous intense winds drove 

upwelling and maintained high primary productivity in the 

photic zone. The low elevation of the landmass to the east 

and the vastness of the shelf further reduced the flux of 

otherwise-diluting terrigenous debris, allowing the accumulation of marine-dominated deposits throughout the 

basin evaporites along the coast and biogenic sediment 

within the basin. The Antler Uplift or a more westerly 

island arc, analogous to the Ryukyu Islands in the East 

China Sea, shielded the shelf from strong coastal currents, 

such as the Peru, Benguela, and Kuroshio Currents, thus 

providing a relatively undisturbed depositional regime for 

the dominantly marine debris that did settle out of the 

water column. The basin floor, itself, was rather featureless and, as Yochelson (1968) concluded from faunal evidence, likely had a maximum depth of approximately 200 

m. A trough oceanward of the shelf, analogous to the Okinawa Trough in the East China Sea, limited the flux of 

detritus from the west into the shallower phosphogenic 

area of the basin. Weak tectonic activity and strong easterlies further limited the input of volcanic debris. Because 

few formations compositionally equivalent to the Phosphoria Formation are present in the geologic record, these 

conditions of sea-floor morphology, seawater chemistry, 

climate, provenance, and tectonism likely have seldom 

coexisted in the past.

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B28 The Phosphoria Formation: Its Geochemical and Biological Environment of Deposition

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Oil and Gas Investigations Charts show stratigraphic information for certain oil and gas fields and other areas having petroleum 

potential.

Miscellaneous Field Studies Maps are multicolor or black-andwhite maps on topographic or planimetric bases on quadrangle or 

irregular areas at various scales. Pre-1971 maps show bedrock 

geology in relation to specific mining or mineral-deposit problems; 

post-1971 maps are primarily black-and-white maps on various 

subjects, such as environmental studies or wilderness mineral investigations.

Hydrologic Investigations Atlases are multicolor or black-andwhite maps on topographic or planimetric bases presenting a wide 

range of geohydrologic data of both regular and irregular areas; 

principal scale is 1:24,000, and regional studies are at 1:250,000 

scale or smaller.

Catalogs

Permanent catalogs, as well as some others, giving comprehensive listings of U.S. Geological Survey publications are available 

under the conditions indicated below from the U.S. Geological 

Survey, Books and Open-File Reports Sales, Federal Center, Box 

25286, Denver, CO 80225. (See latest Price and Availability List.)

"Publications of the Geological Survey, 1879-1961" may be 

purchased by mail and over the counter in paperback book form and 

as a set of microfiche.

"Publications of the Geological Survey, 1962-1970" may be 

purchased by mail and over the counter in paperback book form and 

as a set of microfiche.

"Publications of the Geological Survey, 1971-1981" may be 

purcha sed by mail and over the counter in paperback book form (two 

volumes, publications listing and index) and as a set of microfiche.

Supplements for 1982,1983, 1984,1985,1986, and for subsequent years since the last permanent catalog may be purchased by 

mail and over the counter in paperback book form.

State catalogs, "List of U.S. Geological Survey Geologic and 

Water-Supply Reports and Maps For (State)," may be purchased by 

mail and over the counter in paperback booklet form only.

"Price and Availability List of U.S. Geological Survey Publications," issued annually, is available free of charge in paperback 

booklet form only.

Selected copies of a monthly catalog "New Publications of the 

U.S. Geological Survey" are available free of charge by mail or may 

be obtained over the counter in paperback booklet form only. Those 

wishing a free subscription to the monthly catalog "New Publications of the U.S. Geological Survey" should write to the U.S. 

Geological Survey, 582 National Center, Reston, VA 22092.

Note. Prices of Government publications listed in older catalogs, announcements, and publications may be incorrect. Therefore, 

the prices charged may differ from the prices in catalogs, announcements, and publications.