Miocene of the S.E. United States: a model for chemical sedimentation in a peri-marine environment

DEVELOPMENTS IN SEDIMENTOLOGY 22

MIOCENE OF THE S.E. UNITED STATES: A MODEL FOR CHEMICAL SEDIMENTATION IN A PERI-MARINE ENVIRONMENT

FURTHER TITLES IN THIS SERIES

1. L.M.J.U. V A N S T R A A T E N , Editor DELTAIC AND SHALLOW MARINE DEPOSITS 2. G.C. A M S TUTZ, Editor SEDIMENTOLOGY AND ORE GENESIS

3. A.H. BOUMA and A. BROUWER, Editors TURBIDITES 4. F.G. TICKELL THE TECHNIQUES O F SEDIMENTARY MINERALOGY 5. J.C. INGLE Jr. THE MOVEMENT O F BEACH SAND 6. L. V A N D E R PLAS THE IDENTIFICATION O F DETRITAL FELDSPARS 7 . S. DZULYN S K I and E.K. W A L T O N SEDIMENTARY FEATURES O F FLYSCH A N D GREYWACKES 8. G. L A R S E N and G.V. CHILINGAR, Editors DIAGENESIS IN SEDIMENTS

9. G.V. CHILINGAR, H.J. BISSELL and R. W. FAIRBRIDGE, Editors CARBONATE ROCKS 10. P. McL. D. DUFF, A. H A L L A M and E.K. WA L TO N CYCLIC SEDIMENTATION 11. C.C. R E E V E S Jr. INTRODUCTION T O PALEOLIMNOLOGY 12. R.G.C. B A THURS T CARBONATE SEDIMENTS AND THEIR DIAGENESIS 13. A.A. M A N TE N SILURIAN REEFS OF GOTLAND 14. K.W. GLENNIE DESERT SEDIMENTARY ENVIRONMENTS 15. C.E. W E A V E R and L.D. P O L L A R D THE CHEMISTRY O F CLAY MINERALS 16. H.H. RIEKE III and G.V. CHILINGARIAN COMPACTION O F ARGILLACEOUS SEDIMENTS 17. M.D. PICARD and L.R. HIGH Jr. SEDIMENTARY STRUCTURES O F EPHEMERAL STREAMS 18. G.V. CHILINGARIAN and K.H. WOLF COMPACTION OF COARSE-GRAINED SEDIMENTS 19. W. SCHWARZACHER SEDIMENTATION MODELS AND QUANTITATIVE STRATIGRAPHY

20. M.R. W A L T E R , Editor STROMATOLITES 21. B. VELDE CLAYS AND CLAY MINERALS IN NATURAL AND SYNTHETIC SYSTEMS

DEVELOPMENTS IN SEDIMENTOLOGY 22

MIOCENE OF THE S.E. UNITED STATES A MODEL FOR CHEMICAL SEDIMENTATION IN A PERI-MARINE ENVIRONMENT

CHARLES E. WEAVER and KEVIN C. BECK School o f Geophysical Sciences, Georgia Institute of Technology, Atlanta, Ga. (U.S.A.)

Reprinted from Sedimentary G e o l o g y , Vol. 1 7 Nos. 112

ELSEVIER SCIENTIFIC PUBLISHING COMPANY Amsterdam - Oxford - New York 1917

ELSEVIER SCIENTIFIC PUBLISHING COMPANY 335 Jan van Galenstraat P.O. Box 211, Amsterdam, The Netherlands Distributors for the United States and Canada: ELSEVIER/NORTH-HOLLAND INC. 52, Vanderbilt Avenue New York, N.Y. 10017

ISBN: 0444-41568-8 Copyright 0 1977 by Elsevier Scientific Publishing Company, Amsterdam All rights reserved. No part of this publication may be reproduced, stored in a retrieval system, or transmitted in any form or by any means, electronic, mechanical, photocopying, recording, or otherwise, without the prior written permission of the publisher, Elsevier Scientific Publishing Company, Jan van Galenstraat 335, Amsterdam Printed in The Netherlands

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ACKNOWLEDGEMENTS This research was supported largely by the National Science Foundation, Grant GA-1330. Thanks are extended t o Dr. Sam Patterson, U.S.G.S., Dr. Jack Williamson, Englehardt Minerals and Chemical Co., and the various personnel of the Georgia and Florida Geological Surveys for supplying samples and information. Paul Huddlestun of the Georgia Survey was particularly helpful in supplying paleontologic and stratigraphic data and discussing many of the Miocene problems. We are grateful t o the patient secretaries who plowed through draft after draft, particularly Dianne Clark, Jeannie Greene, and Barbara Haas. We also wish t o acknowledge Robert E. Dooley for doing the drafting.

CONTENTS

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V

CHAPTER 1. INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Theproblem . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Nature of palygorskite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

1 1 1

CHAPTER 2 . FRAMEWORK . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Regional . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Georgia-Florida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Structural and isopach maps . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Lithofacies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

5 5 7 12 16

ACKNOWLEDGEMENTS

CHAPTER 3 . REGIONAL CLAY-MINERAL DISTRIBUTION . . . . . . . . . . . . . Florida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Echols County. Georgia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Regional cross-sections . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Trough . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

..

41 41 44 48 63

CHAPTER 4. MINES . LOWER MIOCENE . . . . . . . . . . . . . . . . . . . . . . . . . . . Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . La Camelia Mine (MC-1) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Clay mineralogy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Texture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Composition of sand grains. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Thin-section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Heavy minerals . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Interpretation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . La Camelia Mine (MC-2) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . La Camelia Mine outcrop . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Adjacent mines . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . MidwayMine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Gunn Farm Mine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Chesebrough Mine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Attapulgus, Georgia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

71 71 71 73 77 82 84 88 88 89 92 97 98 100 100 101 103

CHAPTER 5 . MINES - MIDDLE MIOCENE . . . . . . . . . . . . . . . . . . . . . . . . . . . Cairo Production Company Mine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Adjacent cores . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Cherokee Mine area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Waverly Mine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

105 105 111 111 113 116 117

VIII

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119

CHAPTER 7. ELECTRON MICROGRAPHS . . . . . . . . . . . . . . . . . . . . . . . . . . . MC-1Core . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Other locations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Dolomite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Calcite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Other Lower Miocene clays . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Middle Miocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Soil . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Grains and pebbles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Amorphous silica . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

123 123 138 138 146 147 157 159 160 171

CHAPTER 8 . CHEMISTRY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Silicates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Thermodynamic calculations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Phosphate . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

177 177 188 197

CHAPTER6.TEXTURE

.

CHAPTER 9 OVERVIEW . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 201 Palygorskite in the Ocean . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 201 ‘Marine’ sedimentary rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 201 Deep-sea occurrences . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 203 Gulf of Mexico and western Atlantic . . . . . . . . . . . . . . . . . . . . L . . . . . . . 203 East Atlantic Ocean . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 206 Indian Ocean . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 208 RedSea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 208 Mediterranean Sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 209 Global distribution . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . :209 Distribution . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 209 Paleolatitude . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 211 Role of continental drift . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 214 Palygorskites in space and time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 216 Temporal distribution . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 216 Environment and source material . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 220 APPENDIX . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . X-ray . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Quantitative X-ray analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

225 225 226

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227

REFERENCES

Chapter 1 INTRODUCTION THE PROBLEM

The Miocene sediments of the southeastern United States are unique in having large deposits of palygorskitesepiolite (commercial deposits occur in SW Georgia and north central Florida) and phosphate. The original intent of the study was t o determine the conditions under which the chain-structure clays formed. It soon became apparent that this was as much a stratigraphic and petrologic problem as a geochemical problem, and that t o understand the controlling chemical processes it was necessary to determine the origin of the carbonate and phosphate as well as the silicate minerals. It was necessary to determine what was unique about the paleogeographic environment in order to determine the factors controlling the geochemistry and mineralogy. This has become increasingly important with the discovery of abundant palygorskite and sepiolite in the deep-sea deposits and in a wide variety of Holocene, Tertiary and Mesozoic deposits. As the stratigraphy of this area is extremely complex and fossils are scarce it seems like an ideal problem for a clay-mineral stratigrapher. Approximately 170 wells and outcrops and in the neighborhood of 3,000 samples were described and X-rayed. Chemical analysis, SEM and TEM examination, textural and thin-section studies were made on a large number of samples. With these data we have been able t o unravel much of the stratigraphy and synthesize the geologic history and complex interplay of environmental conditions in a classic shallow-water hinge area separating two major bodies of water - the Atlantic Ocean and the Gulf of Mexico. We were able t o determine the physical and chemical conditions under which the authigenic minerals - palygorskite, sepiolite, smectite, phosphate, carbonates, opal-cristobalite - formed. I

NATURE OF PALYGORSKITE

Most of the pertinent literature is discussed in the body of the text, so only a brief introduction will be presented. Palygorskite has a fibrous texture and a chain structure. The structure proposed by Bradley (1940) is that of a 2 : 1 layer structure with five octahedral positions. (four filled); four Si tetrahedra occur on either side of the octahedral sheet. These structural units alternate in a checkerboard pattern leaving a series of channels between the structural units. These channels contain water molecules. None of the chemical analyses approach that of ideal formula for the halfunit cell: Si8Mg,020(OH)2(OH2)4* 4H20. Only four of the octahedral posi-

2

tions are actually filled. The H' content of the octahedral sheet is so speculative that it is impossible to make a reasonable calculation of the layer charge. The tetrahedral aluminum ranges from 0.01 t o 0.69 per 8 tetrahedral positions which is similar to the range for low-aluminum montmorillonites. Octahedrally coordinated aluminum, and total A1203,is less than that found in the montmorillonites. The magnesium content of the octahedral sheet and MgO is 2-4 times as abundant as in montmorillonite. The iron contents are similar. The proportions of divalent and trivalent ions in the octahedral position are approximately equal (Weaver and Pollard, 1973). Sepiolite is a lath-shaped magnesium-rich clay mineral with a structure similar to palygorskite. Nagy and Bradley (1955) and Brauner and Preisinger (1956) have proposed structures that differ only in detail. The one structural arrangement has nine octahedral sites and the other only eight (as compared to five for palygorskite). Both structures have channels on both sides and top and bottom of each ribbon, which contain water molecules (zeolitic water). Additional water is bound t o the edge of the ribbons and OH occurs in the structure proper. The ideal structural formula for sepiolite based on the Nagy-Bradley model is (Sil2)(Mgg)030(OH)6(OH2)4 - ~ H z Oand , (Si12)(Mg,)030(OH)4(OH2)4 8H20for the BraunerPreisinger model. Most sepiolite calculated structural formulas indicate a minor amount (0.04-1.05) of A13+and/or Fe3+substituting for Si4+(11.96-10.95) in the tetrahedral sheet. The magnesium-rich sepiolites have a relatively consistent composition. A1203,Fe203and in some the MgO samples Mn203, are commonly present in amounts less than 1%; content ranges from 21 t o 25%. Mg fills 90--100% of the occupied octahedral positions. The analyses indicate sufficient cations to fill approximately eight (7.74-8.14) octahedral sites. This would fill all the sites allotted by Brauner and Preisinger and leave one vacant site if the Nagy-Bradley structure is correct. Most of the cations in the octahedral positions are the large variety; Al, a smaller cation, is relatively uncommon. This is in contrast t o palygorskite, where A1 commonly fills half the octahedral positions. As Wiersma (1970) has published a thorough review of the literature only a sampling of available references will be mentioned. Although palygorskite is relatively rare, it forms in a variety of ways. Caillere (1951), Stephen (1954) Christ et al. (1969) and Bonatti and Joensuu (1968) describe hydrothermal origins, the latter under marine conditions. Muir (1951), Millot (1953), Grim (1953) Barshad et al. (1956) Loughan (1959) and Rateev (1963) have indicated that it can form in lagoons, playa lakes or evaporatic basins. Roaeks et al. (1954) and Millot (1964) have described lacustrine origins. Millot (1970) proposed a marine origin for many of the African deposits. Elgabaly (1962) described attapulgite in the desert soils of Egypt, A1 Rawi et al. (1969) in the arid regions of Iraq, and Van den Heuvel (1964) in the desert soils of New Mexico, U.S.A.

3

Thus, the physical environment is not restrictive, though the chemical environment must be. In these aspects palygorskite resembles the smectites. Sepiolite is commonly associated with palygorskite, montmorillonite, dolomite and magnesite. It is reported t o have formed in lacustrine environments of high basicity (Longchambon and Morgues, 1927; Millot, 1949), pluvial lakes (Parry and Reeves, 1968), in highly saline evaporitic environments (Yarzhemskii, 1949), under basic marine conditions (Millot, 1964), by hydrothermal alteration of serpentine (Midgley, 1959), by hydrothermal alteration in a deep-marine environment (Hathaway and Sachs, 1965), and by the solution of calcite and phlogopite (Lacroix, 1941). Sedimentary sepiolite is most commonly formed, along with dolomite, in highly alkaline and schizohaline evaporitic environments. Sepiolite apparently forms under more alkaline conditions than does palygorskite and where Si and Mg concentrations are high and A1 low.

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Chapter 2

FRAMEWORK REGIONAL

The area of study straddles the boundary between the Atlantic and Gulf Coastal provinces (Fig. 1).The sediments reflect the interplay of the tectonic and sediment-transport features of the two areas. The stratigraphy is further complicated by the fact that the tectonic movements, controlling transgressions and regressions, in the two regions were independent of each other. The Gulf Coast Geosyncline was filled with a thick wedge of Cenozoic clastic sediments, whereas the Atlantic Coastal provinces, and south Georgia and Florida received a relatively thin blanket of sediments, much of it being calcareous (Murray, 1961;Gibson, 1970). During the Paleocene, Eocene, and Oligocene sedimentation in the Gulf Coastal Plain was similar to present-day sedimentation in deltas, barrier islands, bays, marshes, and a continental shelf. Maximum thickness is on the order of 4.5 km or larger (Rainwater, 1960). Equivalent-age sediments in the northern and central Atlantic Coastal Plain consist predominantly of thin beds of glauconitic sand (approximately 100 m) and, in the shelf areas, planktonic foraminifera1 oozes with little detrital material (Gibson, 1970). The Atlantic Coastal Plain beds become progressively more calcareous southward, and in southern Georgia and Florida shallow-water limestones predominate. The limestones grade westward (Florida panhandle and southern Alabama) into the clastic facies of the Gulf Coastal provinces. On the continental shelf t o the east of Florida and Georgia limestones were deposited in shallow-marine and coastal-lagoon environments (Joides, 1965). Montmorillonite is by far the dominant clay mineral in all areas. However, in the Gulf area detrital illite, kaolinite, and chlorite are relatively abundant, and in the other areas opal-cristobalite is common. Glauconite, probably formed from the montmorillonite, is abundant in the Atlantic Coastal Plain sediments. Zeolite (largely clinoptilolite and heulandite, with associated opal-cristobalite) is present in the continental-shelf carbonates of both the north and south Atlantic shelf (Weaver, 1968;Gibson, 1970)and in the relatively thin clastic sediments of southern Alabama (Reynolds, 1966). The Gulf Coast continued t o be a major depocenter during the Miocene. In southern Florida about 275 m of limestone was deposited. The Miocene becomes thinner and more sandy northward into southern Georgia. On the continental shelf and Blake Plateau 21-24 m of calcareous silts and silty calcareous oozes were deposited. In the central and northern Atlantic Coastal Plain detrital sands and clays are as much as 30 m thick. Neither regional nor local correlations of the Miocene are well established. The older Miocene is poorly represented in the northern Atlantic Coastal

-..

G

APPROXIMATE EDGE OF CONTINENTAL SHELF

1'

Irn -

-

Irn >"0'./_.,.,,

I

'.

*.,.v

1

,cn -

0

.

,

,

."

'

. ....... . ' . '

Fig. 1. Generalized thickness of Cenozoic deposits in the Atlantic and Gulf Coastal provinces. Triangle outlines areas studied. (After Murray, 1961.)

7

Plain and appears to be marginal marine to nonmarine (Gibson, 1967). The Late Miocene was a period of major transgression over the northern Atlantic Coastal Plain (Yorktown Formation) but apparently not in the Gulf Coastal Plain. Gibson (1970) believes that the ‘fairly heavy’ influx of detrital material over the entire Atlantic margin in the Late Miocene reflects uplift in the Appalachian area. Montmorillonite is the dominant clay mineral in the Miocene coastal plain and shelf sediments. Chlinoptilolite is abundant in the northern area and palygorskite and sepiolite in the southern. Phosphate is relatively abundant throughout the Atlantic Coastal Plain and shelf. In Florida phosphate was believed t o have been deposited during the early Middle Miocene and reworked and concentrated in younger sediments. Actually, apatite is relatively abundant in rocks at least as old as the Eocene. The North Carolina deposits were deposited at the same general time. In northeastern Georgia phosphate is present throughout the Miocene section but is most concentrated in the Middle to Upper Miocene sediments. Gardner (1926) from a study of the molluscan fauna of the Florida panhandle found that the upper Lower Miocene t o lower Middle Miocene Chipola fauna was a subtropical assemblage. The fauna in the overlying Oak Grove and Shoal River sediments indicates progressively colder water, ‘foreshadowing’ a radical fall in temperature at the close of the alum bluff epoch (Middle Miocene). Tho authigenic chain-structure clays are largely restricted t o the subtropical Lower Miocene and Upper Oligocene, suggesting temperature was a factor in its formation. The sparsity of carbonate rocks in the Middle Miocene sediments of Georgia might also be a reflection of a colder climate. The lower Middle Miocene phosphate sediments (Pungo River Formation) farther north in North Carolina have a cold-temperature water foraminiferal assemblage (Gibson, 1967). Montmorillonite and clinoptilolite are abundant and palygorskite is not present. GEORGIA-FLORIDA

A lithofacies study by Chen (1965) of the Paleocene and Eocene sediments of Florida showed two distinct facies. The sediments of the Florida panhandle are largely clastic sands and clays while those of peninsular Florida are shallow-water carbonates and evaporites. Chen believes little mixing occurred because the two provinces were separated by the Suwannee Channel (Jordan, 1954). Another possible factor is that the regional slope of peninsular Florida was so small that little detritus could be transported there. The southern Appalachians apparently had a relatively low relief and did not contribute a large amount of sediment t o the south and southeast. The Florida Platform was submerged during Paleocene and Eocene times, except in late Middle Eocene and late Late Eocene times when most of the northern and central portions of the platform were emergent and subjected

8

t o nondeposition and subaerial erosion (Chen, 1965). The emergence is demonstrated by the unconformable relation between the Ocala Formation and the overlying and underlying beds. The Upper Eocene of Florida and much of south Georgia and Alabama is represented by the relatively pure shallow-water Ocala limestone (Puri and Vernon, 1964). The limestone interfingers updip with the argillaceous Cooper Marl and Barnwell Formations (arkosic) (Herrick and Vorhis, 1963) and westward into the Yazoo Group of clastic sediments. In south Georgia and Florida the Ocala is overlain by another shallowmarine limestone sequence which is referred to as the Suwannee or Oligocene undifferentiated. The Tampa and equivalent formations have been considered to be of Early Miocene age, but recently Huddlestun (personal communication, 1975) has acquired faunal evidence to indicate it is Oligocene in age. The Tampa units consist of tidal, restricted-environment sediments that may in part be a facies equivalent of the marine Suwannee sediments. The clay minerals tend t o support this interpretation. During Miocene time there were four dominant structural elements within the south Georgia-north Florida area (Fig. 2). The major positive feature is the Ocala Uplift or Ocala Arch. There are differing opinions as to time and extent of structural activity (for brief review see Olson, 1966).' Some believe activity t o have dated from before the Late Eocene and persisted into the Early Miocene. Brooks (1966) presents data that indicates the arch was first uplifted and eroded somewhat prior to early Late Miocene time. In southwest Georgia and the eastern portion of the Florida panhandle Miocene sediments are thicker than t o the east or west. The general alignment of the thickest interval is northeast-southwest (Fig. 2). This area has been called, among other things, the Chattahoochee Embayment, Appalachicola Embayment, and Gulf Trough (for review see Patterson and Herrick, 1971). The southern portion has the shape of a broad embayment and the northeastern portion that of a narrow trough. The trough is presumably a remnant of the Suwannee Channel and will be referred to as the Trough, as it did not connect the Gulf and Atlantic provinces during the Middle Miocene. Murray (1961) believes the basin (Appalachicola) has existed since the Early Paleozoic. The fourth major structural feature is the Atlantic Embayment, encompassing most of eastern Georgia and northeastern Florida (Jacksonville Basin). The nomenclature and stratigraphic sequence used in this study is shown in Table I. The justification for this classification is explained in the text. The Late Oligocene t o Early Miocene sea transgressed over the eroded surface of older limestones, incorporating some of the underlying material. A sandy, silty clay facies was deposited in the western Florida panhandle and t o the northwest, generally becoming more sandy t o the north and northwest. These shallow-marine to brackish-water sediments were named the Chattahoochee Formation (Puri and Vernon, 1964).

9

H

20 K M

CONTOUR / N T E f f V A L

/00'=30.5 M

Fig. 2. Generalized Miocene paleogeographic map showing two positive areas separated by two depocenters. The Suwannee Uplift is an older feature than the Ocala Uplift. Narrow sill separated Appalachicola Embayment from trough area of Atlantic Embayment.

Seaward to the south, east, and northeast a sandy limestone and dolomite were deposited. This lithofacies has been named the St. Marks Formation by Puri and Vernon (1964).It is commonly referred to as the Tampa Formation. The St. Marks was considered to be deposited in a deeper-marine environment than the Chattahoochee. To the south in peninsular Florida the lower portion of the St. Marks (or Tampa) contains only limestone. The uppermost part of the section contains green-clay interbeds. To the north clay beds are present throughout the St.

10

TABLE I Stratigraphic nomenclature used in the present study Series

Formation

/

Miocene Upper

/

Choctawhatchee Upper Miocene clastics

Middle

/ 1

Shoal River Chipola Torreya *

Lower

marine facies on continental shelf

/

/

/

Hawthorn

/

I

/

Oligocene

/

Tampa

Upper Middle Lower

Chattahoochee

Suwannee

Eocene

/

/

St. Marks

Ocala ~~

~~

* Torreya may be lateral equivalent of Chipola. Marks. In southwestern Georgia it is a sandy limestone and contains some interfingering beds of sand and clay in the updip direction (Herrick and Vorhis, 1963). Whether equivalent-age sediments are present in the Atlantic Embayment of Georgia and the continental shelf has not been determined. The Lower Miocene of the inner continental shelf consists of a phosphatic clay containing several layers rich in diatoms and radiolarians (JOIDES, 1965). The outer shelf contains 42 m of coarse-grained quartzose sand and sandy silt. The Blake Plateau contains 33-48 m of foraminiferal calcarenitic sand and silty calcareous ooze. Puri and Vernon (1964)considered the Hawthorn to be Middle Miocene. Espenshade and Spencer (1963) considered the Hawthorn Formation to represent the Lower and Middle Miocene and described it as having an upper quartz sand and phosphorite unit and a lower phosphatic dolomite unit. Based on faunal studies along the Appalachicola River, Banks and Hunter (1973)have suggested that post-Tampa, pre-Chipola sediments are a distinct unit (marine to brackish) and suggested the name Torreya Formation. Huddlestun (1976)dates the Chipola as late Early Miocene to early Middle Miocene (approximately straddling the boundary between the Burdigalian and Langhian). The Torreya contains Atlantic Ocean fauna that has more of an affinity to the Chipola than the Tampa fauna. The Torreya Formation has a relatively high content of carbonate rocks. Banks and Hunter (1973) restrict the Hawthorn to the overlying (pre-Chipola) more clastic facies. Though they could not obtain satisfactory palaeontologic data they placed

11

the commercial fuller’s earth beds of Florida in their restricted Hawthorn Formation but suggest the material represents a new formation of preChipola age or an upper member of the Torreya Formation. The Hawthorn is extremely complex lithologically . Near Gainesville, Florida, the Hawthorn (-20 m thick) contains beds of clay, sand, limestone, and dolomite mixed in all proportion but more calcareous in the lower portion (Torreya Formation) and usually containing phosphate and fossils. In north central Florida (Hamilton County) the section (30m thick) is equally complicated and is described as having thin-bedded palygorskite clay interbedded with sand and phosphorite, burrowed clayey sands, clayey calcilutite, cobbles of clay and broken limestone, alternating beds of dolomite, palygorskite, and phosphate--quartz sand. Some of the dolomite has ripplemarks and mud cracks, and pebbles occur in cross-bedded phosphate--quartz sand (Pun and Vernon, 1964). The sedimentary structures, along with oysters, indicate shallow-water strand-line deposition (as is indicated for all clayey beds rich in primary palygorskite). In Georgia, Herrick and Vorhis (1963)describe the Hawthorn Formation as a phosphatic, very sandy, locally fossiliferous and cherty, micaceous clay interbedded with scattered tongues of fine- to coarse-grained arkosic, phosphatic sand. In the Trough (Georgia) the sediments which lie on top of the Tampa limestone and which contain commercial fuller’s earth beds in the upper part are called Hawthorn (Siever, 1964;Gremillion, 1965). The present study indicates they are Middle Miocene (Shoal River equivalent). The Middle Miocene alum bluff stage sediments, deposited unconformably on the Tampa, are divided into seven lithofacies by Puri and Vernon (1964). They include the Chipola lithofacies in the alum bluff stage, and consider it the downdip facies of the Fort Preston deltaic and prodeltaic sands. However, the Chipola is now consigned t o the upper Lower-lower Upper Miocene (Huddlestun, 1976). Following deposition of the Chipola-age beds a period of erosion was thought to have occurred in northwestern peninsular Florida and extreme southwestern Georgia. Sea level was lowered and continental, deltaic, and prodeltaic sands and clays of the Fort Preston Formation were deposited (Olson, 1966). Similar sediments were deposited in the eastern portion of the Florida panhandle (Pun and Vernon, 1964). The Middle Miocene Shoal River Formation and Oak Grove sand are the marine equivalents of the Fort Preston Formation (Huddlestun, 1976),and occur in the central panhandle between the two deltaic facies. They consist of very fine to medium-grain sand, clay, and argillaceous shell marl. The Upper Miocene Choctawhatchee Stage of Pun and Vernon (1964) contains several lithofacies and biofacies. Most of these biofacies have now been placed in the Middle Miocene or Pliocene. Only the Arca zone of the western panhandle is considered t o be marine Upper Miocene (Akers, 1972). Over the remaining portions of north Florida and Georgia the Upper Mio-

12

cene, where present, is a clastic continental t o littoral facies. The Upper Miocene clastics and Miccosukee Formation, 24-30 m thick, are believed t o be remnants of the Middle Miocene delta. Sands and clays of this facies have been reported in northwest peninsular Florida and southwest and southeast Georgia (Olson, 1966). STRUCTURAL AND ISOPACH MAPS

Detailed isopach and structural maps including both Georgia and Florida are rare. In an effort to obtain some stratigraphic background information maps were synthesized from a number of published sources and from data acquired in the present study. ‘State-line-faults’ are abundant but general trends can be established. Data were not readily available t o construct region structural maps on ‘top’ of the Oligocene (Suwannee). In addition, there is considerable difference in where the top is picked by various people. More and better data are available on the top of the Eocene (Ocala limestone) and a structural map was constructed (Fig. 3) using the data of Moore (1955), Herrick and Vorhis (1963), Chen (1965) and Brooks (1966). Chen’s top is approximately 30 m higher than those picked by others and has been correspondingly lowered. As the Oligocene in the area of interest is seldom much more than 30 m thick the structural map of the top of the Eocene is generally similar to that of the top of the Oligocene. The Atlantic and Appalachicola Embayments, and the Trough are well delineated. High areas outline the Ocala Arch and the southwestern extension of the southern Appalachians and/or the Chattahoochie Anticline. The northern extension of the Ocala Arch serves t o separate the two structural lows. The structural contour map of the ‘top’ of the Oligocene (Suwannee limestone; Herrick and Vorhis, 1963) shows similar features. Structural contour maps of the ‘top’ of the Oligocene for southwest Georgia and adjacent parts of Florida have been constructed by Siever (1964) and Gremillion (1965). Their picks of the ‘top’ of the Oligocene are reasonably compatible, and both pick the top considerably lower (30-60 m) than Herrick (1961). As the structural configuration is pertinent t o the origin of the commercial palygorskite deposits an attempt was made t o construct a more detailed structural map by combining the data of Gremillion and Siever (Fig. 4). Siever (1966) interprets the east flank of the depression as a fault, but this does not materially change the generalized structural pattern. A cross-section parallel t o the axis of the depression is shown in Fig. 5. The maximum structural depression occurs in the southwestern Georgia portion of the trough. This low area is separated from the Appalachicola Embayment by a relatively high area in the form of a sill. On the basis of the available well control it is not possible to determine the northeastsouthwest width of the sill. The commercial primary-palygorskite beds appear t o be related t o the sill. The

13

w 50 KM

TOP OF O C A L A CONTOUR

INTERVAL

/00‘=30.5M

Fig. 3. Structural map of the top of the Eocene Ocala limestone. (Georgia data largely from Herrick and Vorhis, 1963.) Cross-hatched areas indicate Ocala missing.

bed shows the effects of postdepositional tilting t o the southwest. The presence of Lower Miocene fossils in mined clay beds on the Georgia-Florida border (Olson, 1966) suggests the top of the Oligocene and Tampa is relatively shallow in this area. Structural contours of the top of the Tampa limestone in southwest Georgia and northern Florida are similar t o those of the Oligocene (Gremillion, 1965;Siever, 1966). The significance of these maps and isopach maps of the Tampa are somewhat questionable as the review reveals that seldom do any two geologists pick the same top for the Tampa. The same is true for the ‘top’ of the Oligocene. Differences can be 50 m or more. ‘The Miocene of peninsular Florida was deposited upon an eroded surface composed of Eocene and Oligocene limestones. Deep valley and hills, sinks, lake basins and structural features had developed upon this surface and these irregularities were covered by sediments formed during the Early and Middle

14

-

TOP SUWANNEE (OLIGOCENE)

CONTOUR /NT€Rv4L-/oo' .JOJ

M

I5 K M

Fig. 4. Structural map of the top of the Suwannee limestone. Dots show the location of the palygorskite mines. Solid dots indicate primary Lower Miocene clays ( 1 = Midway; 2 = Chesebrough; 3 = La Camelia; 4 = Gunn Farm; 5 = Block N) and cross-dots indicate reworked Middle Miocene clays ( 6 = Cairo; 7 = Cherokee; 8 = Waverly). Dashed contour lines show alternative interpretations.

Miocene' (Vernon, 1951). Much of the Ocala Uplift remained above sea level during Miocene time. Structural contours of the Oligocene (top of Suwannee) in Leon County, Florida (Hendry and Sproul, 1966) and south central Georgia (Herrick and Vorhis, 1963) suggest that the Ocala Uplift extended into southern Georgia during a portion of Miocene time. The Atlantic Embayment area of Georgia also has a very thin and erratic thickness of Oligocene (Herrick and Vorhis, 1963), suggesting that it may have been exposed to subareal erosion. The Suwannee limestone is commonly dolomitized and silicified (particularly in the upper portion) suggesting there has been some subaerial exposure. The thickest section of Oligocene occurs in the Trough suggesting that deposition may have been continuous from Oligocene to Miocene in the area. A thicknessdistribution map of the Miocene plus the Tampa and Tampaequivalent sediments (Fig. 2) was constructed using the welldescription data of Moore (1955), Godell and Yon (1960), Herrick (1961)' Yon (1966),

15

Sea L

-30

z

-60

L

u

C

- 90

L

-120

5

4J

0

-150

NE-SW C r o s s S e c t i o n Down C e n t e r o f Trough

NW-SE Cross Sect ions

I5 KM

Fig. 5. Cross-section parallel to the axis of the Trough shown in Fig. 4 . Location of palygorskite clay beds is shown. A , B and C are cross-sections at right-angles to the axis of the Trough showing the configuration of the top of the Suwannee Formation.

Hendry and Sproul (1966), Marsh (1966), our own data, and the map of peninsular Florida by Vernon (1951). As both the top and bottom of the Miocene are picked differently by different people the details of the map are subject to some error but the general pattern is reasonable. The intervals include the Tampa and equivalent sediments. Three major depocenters are evident: the Atlantic Embayment, the Appalachicola Embayment, and an embayment in the western Florida panhandle. The first two depocenters are connected by a long, 'narrow relatively thick band. The Ocala High extends well into Georgia and appears t o partially divide the Atlantic Embayment into two parts. The northwestern portion is elongated and in line with the Trough. The Miocene fauna in the Florida panhandle and Appalachicola Embayment is similar t o that in the Gulf of Mexico, though Banks and Hunter (1973) found only Atlantic fauna in the Lower Miocene sediments of the upper portion of the embayment. Atlantic Embayment fauna is similar t o the Atlantic Ocean fauna. The significance of the Trough fauna is more obscure, but the presence of diatoms in the Trough suggests an affinity for the Atlantic. The Ocala High generally trends northwestsoutheast. However, the isopach map indicates it has a general north-south trend in northern Florida and a northeastsouthwest trend in Georgia. Thus, the High has an arcuate shape. This may be the result of interaction between tectonic activity in the peninsular arch of Florida and the Appalachian Mountains. Brooks (1966) presented convincing evidence that during the Miocene, following the uplift

16

of the Ocala High, the Suwannee River flowed northeast into the Atlantic Embaymen t. The Miocene tectonic pattern suggests that most of the detrital sediment in the southwestern portion of the Trough and the northeastern portion of the Appalachicola Embayment (areas where commercial clay beds occur) was derived from the Ocala High. All other depositional areas, with the exception of the northeastern flank of the Ocala High, received their detritus from rivers draining the piedmont or from marine currents. Detritus coming from the Ocala High or from the Atlantic Ocean was probably not transported across the Trough in Georgia. In Florida the Appalachicola River and Embayment acted as a barrier t o prevent eastern sediment from reaching the western panhandle. The presence of different mineral suites tends to confirm this. Further, Tanner (1966) believes that in the panhandle area the Miocene longshore currents flowed to the east. LITHOFACIES

Using the well descriptions of Herrick (1961) and our own description of 120 wells an attempt was made to obtain an idea of the distribution of various lithofacies in the Miocene of Georgia. Map I shows the location of the wells, mines, and outcrops studied (inset plate). Limestone and dolomite are largely restricted to the lower part of the Miocene and most are presumably the equivalent of the Tampa, though in the southern and northeastern portion of Georgia some are present in the Lower Miocene. An isopach map showing the distribution of carbonate rocks is shown in Fig. 6. A thick, long, narrow bank of unfossiliferous, dense dolomite occurs in the southwestern one-half of the Trough. This is apparently a bank deposit fringing the Ocala High. The carbonate interval to the west contains limestone intraclasts and calcareous sand, and has a relatively large content of interbedded sand. This would appear t o be a seaward facies. In the northeastern portion of the Trough shelly sands and coquina are abundant (Fig. 7). The Atlantic Embayment is partially separated from the Trough by two thin carbonate sections. The southern one (Ocala High) was a high area and the northern one appears t o have been a shoal area (fossiliferous limestone). The carbonate in the embayment area is largely tan t o brown dolomite that was deposited in shallow lagoons or tidal flats. The configuration of the isopach contours would tend t o suggest the area was partially closed to the Atlantic and open t o the Gulf. Limestones comprise the bulk of the sediments in the Appalachicola Embayment and the southwestern portion of the Trough with the proportion increasing (up t o approximately 90%) towards the Gulf of Mexico. Much of this carbonate sequence is younger than that to the north. The carbonate rocks are described as having a wide variety of colors: light brown, honey, white, cream, buff, and various shaded of gray. X-ray analyses

pp. 11-28

J-2

0

WELL LOCATION MAP

29

T o t a l Llrnestone Thickness ( M i o c e n e and T a m p a ) Counrour l n r s r v o ~i n Feel /00’.305H

m

Fig. 6. Isopach map of the Miocene and Tampa carbonate rocks.

indicate the brown- and honey-colored varieties are protodolomites; others may be limestones or dolomites. The variously colored limestones have relatively distinct distribution patterns. The tannish-colored dolomites, which may comprise as much as 30 of the section, are largely restricted to the eastern portion of the Georgia coastal Plain (Fig. 7) with the maximum thickness being in southeastern Georgia. Through much of this area the tan dolomite comprises the major portion of the carbonate rocks. In southwest Georgia the colors are more variable (white to brown, gray to light brown). The carbonates to the west of the tan-dolomite facies (in the Trough) are primarily white dolomites and fossiliferous limestones. Some white limestones also occur in the area where tan dolomite is predominant. Wells in the northeastern half of the Trough contain intervals of sandy coquina up to 60 m thick (Fig. 7). Gray limestones occur in the thin northeast portion of the Georgia Coastal Plain on the northern edge of the Atlantic Embayment. Relatively thick sandy coquina beds are apparently restricted t o the northeastern portion of the Trough and the Appalachicola Embayment. They are not reported in the approximate 260-km interval separating these two areas but fossils are abundant in the northeastern one-third of this interval and are rare or absent in the southwestern two-thirds. Two ‘shell banks’ occur in

30

\

\

\

Miocene C o r b o n a t e s Tan White Shelly S e d i m e n t and Coquina . ... .... > 5 0 % C a r b o n a t e

---

Conrour fnrervof In Feel

-

100'. 30.5M 20KM

Fig. 7. Map showing thickness of various types of carbonate rocks in the Miocene. Tan carbonate is largely coarse dolomite deposited in shallow brackish water.

the Atlantic Embayment. Fossils in general are unreported in southern Georgia and at least the northern part of peninsular Florida and in much of the eastern part of the coastal plain. Maximum thickness of fossiliferous limestones lies on trend'with and includes the coquina interval. X-ray analyses of the carbonate rocks do not indicate any clear-cut pattern for dolomite and calcite. In general the tan and brown carbonates are dolomites. Thus dolomite is predominant in the Georgia Embayment area. Calcite becomes more abundant in the northeastern Trough and Florida panhandle area where the limestones commonly have a white t o gray color. Dolomites in this area can also have a gray to white color. Limestones, largely bioclastic, become abundant in the Appalachicola Embayment. In peninsular Florida both limestones and dolomites are present. Most of the dolomite is a protodolomite and has 2-5 mole % excess CaC03. The dolomite will be discussed in more detail in another section. Dolomite occurs as anhedral, subrounded spheres and well-developed

31

rhombs (Fig. 8). Both types commonly have hollow (dark) centers. The former type is identical to those found by Muller and Irion (1969) in a recent salt lake in Turkey. Rounded dolomite crystals are commonly associated with evaporites, but Folk and Land (1975) suggest they are more likely to form in schizohaline conditions where the salinity fluctuates from hypersaline to brackish-fresh, i.e. coastal evaporitic lagoons occasionally flooded by fresh water, or a meteoric fresh-water table moving through a hypersaline environment after burial. The fresh water presumably dissolves many of the evaporite minerals. The spherical dolomite is most abundant in the tan-dolomite facies and

Fig. 8. Upper picture shows euhedral dolomite with dark centers. Tan Miocene dolomite from Echols County. Lower picture shows spherical dolomite from Tampa Formation west of Trough. Light material in background is palygorskite. White bars equal 0.05 mm.

32

in the shallow-water dolomites in extreme southwest Georgia. Large euhedral dolomite rhombs, commonly hollow, appear to comprise most of the thick whitedolomite sequence in the southwestern portion of the Trough. The white dolomite is apparently secondary. Alteration possibly occurred at the end of the Early Miocene when the general area was elevated. As we will show with additional data, the dolomite-palygorskite facies apparently represents a restricted or brackish environment and calcitemon tmorillonite and dolomite-montmorilloni te a normal-m arine environment. On the west flank of the Trough in southwestern Georgia some of the Tampa consists of sparites, micrites and intrabiomicrites, and contains mostly montmorillonite. Also present are fine-grained. dolomite beds in which the dolomite is subspherical and commonly has hollow (dark) centers. The dolomites commonly contain pure palygorskite but some contain montmorillonite. The palygorskite is present as sand-size grains and stringers. Pebbles of palygorskite can be seen in the outcrop and some units are burrowed. Some of the dolomite beds contain abundant clasts with a palygorskite clay matrix. Tampa outcrop samples from the southeastern flank of the Trough are indicative of both high- and low-energy environments (biosparite, intrapelbiosparite, intramicrite, and micrite). The dominant clay mineral is palygorskite, and no evidence of clay grains was seen in thin-sections. Further to the east the carbonate facies consist largely of tan dolomite, and may occur in both the lower and middle portion of the Miocene section. The dolomite crystals vary from spherical to rhombic. Most samples contain both spherical and subrhombic crystals. The crystals have hollow centers. Dolomite clasts are present in some samples, indicating that the dolomite was probably formed before appreciable burial. Varying amount of quartz, feldspar, and phosphate are present, usually in minor ( S p > M. 'P>>M. 3 P >> M, Qtz. 4P>>M.

181

12

10

8

0

rn

E $

6

4

2

Fig. 103. Relation of MgO to

A1203

in Miocene clays. Soil samples are relatively low in

Mg .

from 7 to 9 and the palygorskites from 1 to 1.5 (samples with a relatively high sepiolite content have ratios less than 1).Samples containing both clays lie at intermediate positions. This suggests a mixture of two end members with relatively fixed compositions. The soil samples plot below the line and appear to be deficient in Mg or Al. This is apparently due t o their high Fe content. When A1203 is plotted versus MgO + Fe203, the soil samples plot relatively close to the linear trend based on the two end points. Fe203and A1203have positive correlation. The montmorillonites contain approximately 1%more Fe203 than the palygorskite samples. The samples from thin soil zones, regardless of their mineralogy, contain 2% more Fez03 than the other samples. This tends to confirm the identification of the soil zone and indicates a secondary mobilization of Fe and its precipitation lower in the section, possibly as Fe-rich montmorillonite. The organic material present was presumably involved in the transport. The relatively constant increase in Fe, independent of the clay-mineral suite, is possible evidence that the bulk of the clay minerals in the soils is inherited. A plot of K 2 0 vs A1203shows a fair amount of scatter but in general the K20 content increases as the A1203 (montmorillonite) increases. More detrital illite (-200 m.y., unpublished data) was deposited in the montmorillonitic marine facies than in the tidal-lagoonal facies. Correcting for the illite content would cause a slight decrease in the amount of tetrahedral A1 for both clay minerals, but little change in the octahedral layer. The presence of detrital illite in the palygorskite-rich beds indicates it is more stable than montmorillonite under the conditions in which palygorskite forms. This is also true of the environment in which phosphate forms (Weaver and Wampler, 1972).

182 Meters

Analyses were made for Li, Sr and Ba (Table 111) t o determine if additional environmental information could be obtained. Tardy et al. (1972) showed that there was a direct but scattered relation between Li and Mg in MC-1 core in order that the chemical data could best be related to the mineralogic data. In general AlzOJand MgO are inversely related (Fig. 104).The AI2O3/Mg0 ratio is relatively high (5) in the basal sands and decreases to a value of approximately 1.5 in the lower clay bed. The pure palygorskite has a ratio 1.0. The ratio reflects variations in the relative amounts of montmorillonite the clay minerals, the evaporitic-facies minerals (sepiolite, stevensite and

183

*

hectorite) having a high (400-6000 ppm) Li content and a relatively high Li/Mg ratio. Actually, there is a vague inverse relation for samples containing less than 300 ppm Li. None of the Miocene samples had more than 90 ppm Li, which tends t o confirm that the environment in which the clays formed was not hypersaline. Miocene palygorskites show a suggestion of an inverse relation between Li and Mg with some outlier values. The marine montmorillonites contain an average of 39 ppm Li and palygorskite 24 ppm Li. Tardy et al. (1972) found a similar difference (69 vs 51 ppm Li, though if one anomalously high value is excluded the montmorillonite average changes from 69 to 46). The clays with the lowest Li values (Tardy et al., 1972) are kaolinites and montmorillonites formed by fresh-water weathering (average 22 ppm Li). The low values for the Miocene palygorskites suggest that they were formed in water of less than normal salinity. The samples from the.Miocene soil have approximately twice as much Li as the mineralogically equivalent nonsoil samples. This is further indication that arid, evaporative conditions existed during formation of the soil. The sepiolite-rich (-50%) sample, F5-7, contains only 67 ppm Li compared with an average of 621 ppm Li reported by Tardy et al. Thus, it is probable that the sepiolite was not formed under hypefsaline conditions. The Sr and CaO show a good linear relation - Sr (ppm) = 11.7 + 30.1 CaO!%); r = 0.97 - with the exception of two samples, FF-1 and FH-37. (These samples are weathered, and also show anomalously high Ba.) Both Sr and Ca are primarily in phosphate grains. The palygorskite samples have a lower content of Sr and Ca than the montmorillonite though there is considerable overlap. The average Ba value is 123 ppm (excluding the anomalous samples FF-1 and FH-37). Aside from the anomalous samples there is a direct relation between Ba and Al2O3 (some scatter) and a somewhat poorer relation between Ba and KzO. The montmorillonites contain 2-3 times more Ba than the palygorskites. The average shale contains 580 ppm Ba (Turekian and Wedepohl, 1961) or 800-900 ppm for the nonquartz components. Most of this Ba is probably in the illite which comprises the bulk of the shale. The clay component of the average shale contains 5.5% KzO (Shaw and Weaver, 1965). The Ba vs K 2 0 graph of the Miocene data (less than two micron fraction) shows 1.0%KzO is equivalent to approximately 175 ppm Ba. If all the Ba is in the illite it would contain 962 ppm Ba. The similarity of this value t o that of the clays in the average shale suggests much of the Ba is probably in the illite. In summary, the palygorskites contain less Li, Sr and Ba than the marine montmorillonite clays. This is interpreted t o mean that the palygorskites were formed in waters of less than normal salinity and have a lower apatite and illite content than the marine montmorillonite clays. Chemical analyses were made of the clay fraction of a series of clay samples from a core (MC-1) from the La Camelia Mine (Fig. 104) and from a

184

TABLE V Chemical composition of clay from a Cherokee Company Mine Core A

*

Feet

Si02

A1203

Fe2 0 3

MgO

CaO

Na2O

K2O

18 22-25 25-30 30-31 31-35 44-48 59-62 65-69

69.50 67.00 63.50 66.50 70.50 68.20 71.90 69.90

14.13 14.62 14.10 12.90 12.15 11.50 10.70 8.95

2.16 4.79 5.55 6.68 4.80 3.44 3.77 5.08

0.20 1.94 4.36 5.88 4.61 5.27 3.40 4.29

0.15 0.39 0.40 0.34 0.42 2.01 1.56 1.28

0.12 0.22 0.23 0.26 0.26 0.26 0.57 0.22

0.52 1.16 1.04 0.98 0.92 0.88 1.35 1.02

* Chemical analysis by G . Banchero core from the Cherokee Mine (Table V) in the Ochlocknee area. The analyses were made of material scraped from the slides used for X-ray analyses of the plus feldspar (Alz03)and palygorskite, sepiolite and dolomite (MgO). In the mud-crack intervals (5.2-5.5 m) the palygorskite-rich pebbles have a low ratio and the matrix montmorillonite has a high ratio. The ratio is highest in the montmorillonitic soil zone, being a maximum at 4.6 m near the center of the zone. The ratio systematically decreases upward, reaching a minimum (0.5%)in the dolomitic zone. The ratio is lower, and the absolute amount of Mg is higher, in the upper section than the lower section. Even though the Mg-rich clay sepiolite (Al2O3/Mg0= 0.1) is relatively abundant in the lower clay bed, the upper clay bed has a larger percentage of MgO. Apparently Ca was more available during deposition of the upper sediment and Mg entered dolomite. The maximum amount of Fez03 (6%) occurs in the lower portion of the soil zone where the organic material is concentrated and is relatively high through the soil and overlying sand zone. This could be interpreted as evidence for a soil leaching profile. CaO values are relatively constant at l-2% except for the doIomitic zone where 8--10% is present. The K,O distribution is similar t o that of A1203. The upper sequence has KzO values close to 0.5%. The values for the lower sequence average 0.75%. Higher values, 1.0 t o 1.795, occur in the soil zone where the mica content is relatively high. The chemistry closely reflects the mineralogy and tends to confirm the environmental interpretation based on other data. In order to obtain some idea of the vertical variation in the composition of the interstitial water and to see if it was in equilibrium with the mineral suite, 10 g of dry clay were washed with distilled water and the water analyzed. The distribution of cations is fairly uniform throughout the section. Ca is the dominant cation, comprising 48-62% of the cation suite (all data on ppm basis). There is a slight but systematic decrease in Ca with depth. Mg values range from 21-27% and also decrease slightly with depth. Na (5--14%) and K (7--16%) are present in approximately equal amounts and both increase slightly with depth, the Na more than the K.

185

The high Ca content and the decrease with depth suggest that the calcite (largely in shells) in the overlying section controls the chemistry of the downward percolating rain water. The Ca/Mg ratio of 2.5 suggests that some of the calcite filling fractures and coating bedding planes may be secondary. The process could have been going on since Miocene time. Chemical analyses of bulk samples from the Cherokee Mine (Fig. 54) also show a close relation to mineralogy. The A120JMgO ratio is high in the upper montmorillonite zone and low in the palygorskitesepiolite zone. A1203/Fe203and Fe20/Mg0 show a similar distribution. Ca is most abundant in the lower two-thirds of the core, as is apatite. The A1203/Fe203and Fe203/Mg0values indicate these clays have a higher Fe content than the clays t o the south. Calculations using the A1203/Mg0 values for pure montmorillonite (7.0) and palygorskite (1.O) indicate that the Cherokee samples should contain between 60 and 80% palygorskite. These values are high on the basis of the X-ray data. Making allowance for the sepiolite content and using the ratio of A1203/(Mg0+ Fe203)the calculated values for palygorskitesepiolite can be lowered to 4040%. This range appears to be more realistic. X-ray analyses of the clay fraction probably cause an over estimate of the montmorillonite content. Analyses of one of the clay pebbles from this area (FG-16, Table 111) show that, in addition t o the high phosphate content, it also has a relatively high Fe content and contains some organic material. In these latter two characteristics, and in the relative abundance of sepiolite, the pebbles resemble the soil clays. Thus, as suggested by other lines of evidence, the pebbles may have been derived from arid soil or lacustrine deposits flanking the sea. Chemical analyses of samples from the Cairo Production Company Mine and the Waverly Petroleum Mine (Gremillion, 1965) are similar t o those for the Cherokee Mine. Calculations indicate maximum palygorskite content is approximately 50%. The trace-element data, carbonate morphology, and other lines of evidence suggest that, though the palygorskite clay formed in a supratidalshallow-lagoon environment, conditions were probably not predominantly hypersaline but brackish or schizohaline (hypersaline diluted by the periodic influx of fresh surface or ground waters). Studies by Siffert (1962) have shown that sepiolite can form in dilute solutions of Si and Mg. As yet, palygorskite has not been synthesized in the laboratory. The presence of appreciable A1 in the octahedral sheet of palygorskite apparently is the factor that makes it difficult t o synthesize. However, palygorskite is more abundant than sepiolite in nature. Si and Mg are relatively mobile, compared t o Al, and pose no problem. The source of A1 is the major problem. Even under the basic pH conditions in which palygorskite forms, Al solubility is low. To form palygorskite an Al-containing mineral precursor is apparently required. In the present study, in those instances where the clays are obviously formed from solution (in the soils and secondary vein deposits) sepiolite is

186

commonly the predominant clay. Sepiolite is relatively common in soils formed under arid conditions where Si and Mg are mobilized relative to Al. Montmorillonite is present in both the marine and continental sediments and it is inconceivable that it was not present in the supratidal-lagoonal environment. The amount of detrital quartz in the palygorskite clay beds is extremely low but it is present, along with detrital mica and heavy minerals, indicating detrital material was supplied t o the site in which the palygorskite was formed. Minor montmorillonite is present in nearly all the palygorskite clay deposits and this could represent the entire detrital clay phase. However, montmorillonite is the only apparent source of A1 and much more must have been present in the original mud. The oxide composition of a typical montmorillonite and palygorskite (with minor montmorillonite) was recalculated (Table VI) to show the ratio of cations in a 160-oxygen-ion standard cell (Rarth, 1948). The difference between the two clay minerals and the difference between the two when the A1 in the montmorillonite is held constant (with the assumption that the montmorillonite is altered to palygorskite) gives the same results. Si, H, Mg, Na and K must be added t o montmorillonite t o form palygorskite, while Al, Fe, and Ca remain essentially constant. The first three ions are the most significant. These ions are readily available in the mixed fresh and marine waters present in the environment of formation. Montmorillonite can be formed in both fresh and normal-marine waters but apparently is not formed in brackish water. Montmorillonite commonly forms from Si-A-rich solid material (largely volcanic) with Mg being partially supplied from solution. Palygorskite contains appreciably more Mg than montmorillonite. Palygorskite is commonly formed from an intermediate mineral montmorillonite rather than by direct precipitation. TABLE VI Ions in 160-oxygenrock cell of palygorskite and montmorillonite

1

2

A(1-2)

3

A (3-2)

51.60 12.32 2.47 11.15 1.45 2.48 1.42 56.29

47.44 Si 19.12 A1 3.40 Fe 3.44 Mg 2.72 Ca 1.50 Na 1.05 K 48.71 H

+ 4.16 -6.80 -0.93 + 7.71 -1.27 + 0.98 + 0.37 + 7.58

80.08 19.12 3.83 17.30 2.25 3.85 2.20 87.36

+ 32.64

139.19 ~

127.38

~

1. Ions in 160-oxygen rock cell of palygorskite. 2. Ions in 160-oxygen rock cell of montmorillonite. 3. Recalculated ( 1 ) holding A1 constant to 19.12.

215.99

-

+ 0.43 + 13.86 - 0.47 + 2.35 + 1.15 + 38.65

187

When volcanic or similar material is exposed to a basic solution containing Mg, montmorillonite is the initial stable phase. With continued exposure to Mg and Si, palygorskite may be formed. Apparently when appreciable A1 is present it is difficult for much Mg t o be incorporated in the octahedral layer and a two-step (i.e., intermediate mineral) process is necessary. A low cation concentration and relatively high concentrations of Si and Mg are apparently required for the formation of palygorskite from montmorillonite. Indirect evidence suggests near-tropical conditions are necessary for the conversion. Greene-Kelly (1955) has demonstrated that when Mg-saturated montmorillonites are heated at -300” for a few hours they lose their ability to expand. He speculated that the Mg migrated into the octahedral layer. Later infrared studies (Tettenhorst, 1962) suggested the Mg migrated into the hexagonal holes in the tetrahedral sheet. In any event Mg is unique among the more abundant cations in nature in its ability t o enter the montmorillonite layer. Environments in which Mgz+ has a high activity and the temperature is high would tend t o favor such a reaction in the natural system. The energy requirements for physically inverting Si tetrahedron, as required for a montmorillonite -+ palygorskite conversion, in the solid state would seem t o be large. It should also be noted that Edelman and Favejee (1940) proposed a structure for montmorillonite in which every other Si04 tetrahedron is inverted. It has not been proven that this interpretation is incorrect. It would be expected that some such structure would allow for an easier transformation of a sheet-structure t o a chain-structure clay. Siffert (1962) and Wollast et al. (1968) were able to precipitate sepiolite from solutions (25”C, 1 atm total pressure) saturated with amorphous silica. Christ et al. (1973) carried out detailed sepiolite synthesis experiments at temperatures from 51 t o 90”C.Siffert used solutions with Si/Mg molar ratios of 1.43 and 0.70. Sepiolite precipitated at pH = 8.5. With increasing pH the amount of Mg in the precipitate increased and Si decreased. Wollast et al. added sodium metasilicate to sea water and varied the pH by adding NaOH. Sepiolite precipitated at pH 8.2-8.3. At pH values above 9 Siffert’s experiments produced smectite and talc. Preisinger (1963) was able t o precipitate sepiolite readily at temperatures below 80°C from both marine and ‘fresh’ waters (in the absence of Al) over a wide range of Si/Mg ratios (-0.05 to >1.4) in the pH range 8-9. At higher temperatures and pH smectite (stevensite or saponite) or talc formed. In the Miocene of the SE U.S. detrital Al-montmorillonite is so abundant that there is seldom enough ‘excess’ Mg and/or Si to allow appreciable sepiolite t o form, except in some of the carbonate rocks and soils. Mg and Si presumably react with the montmorillonite t o produce palygorskite. When the detrital A1 phase is ‘deactivated’ sepiolite can then precipitate. One of the problems is to determine what conditions cause the solid montmorillonite phase to become so unstable that the solid-solution reaction can occur. Sepiolite is commonly considered t o be related to dolomite and palygor-

188

skite to the nondolomitic sediments. In the Miocene, sepiolite is relatively abundant in the dolomite rocks (Tampa) which contain relatively little clay. In these instances, there is simply not much detrital Al-clay available. Also the environment is more suggestive of hypersaline conditions. Both factors favor the formation of sepiolite. However, where detrital clay minerals are abundant there is little or no sepiolite associated with the dolomite. Where montmorillonite is abundant and the solution contains appreciable amounts of Ca as well as Mg, the equilibrium association is apparently palygorskite and dolomite. When an Al-silicate phase is present, Mg combines with this material in preference t o precipitating as sepiolite (montmorillonitesepiolite association is relatively rare but does exist). THERMODYNAMIC CALCULATIONS

The compositions o f calcite, dolomite, and sepiolite can be written in terms of Ca0-Mg0-Si0z-C02-H20, and their stability fields can be expressed usefully in terms of C02(,,, Ca!zq,, Mg?zq,, H4SiO&aq),and pH. This dissociation of calcite is generally expressed by : C ~ C O ~ (= ,CaZ+ ~ ~+~CO3~ ~ ~ )

where brackets indicate activity. The activity of carbonate ion can be related to P(CO2) by:

Thus we express the dissociation of calcite by: log[Ca'+] + 2 pH = log Kcal - logP(CO2) - log(K(CO2)K(H2CO~)K(HCO;)) A similar expression obtains for the dissociation of dolomite. Values of Kcal and Kdol are taken from Langmuir (1971) for 15", 25", and 35"C, extrapolated as necessary. The constants in the C02 system, at these temperatures, are those tabulated in Stumm and Morgan (1970, pp. 148, 149),and are compatible with Langmuir's data. The dissociation boundaries of calcite and dolomite are expressed in Fig. 105 in terms of (log[Mg2'] + 2 pH) and (log[Caz+] + 2 pH) for temperatures of 15", 25", and 35" C at log P(COz) = -3.5 and for 25" C at log P(C02) equals -2.5 and -1.5. Christ et al. (1973) have studied the dissolution of sepiolite in aqueous solution at temperatures of 51", 70", and 90°C. They express this dissolution in terms of:

'c

189 1

171

A o

...........................

13

-

......................... CALCITE

12 SOLUTION II

10

9

10

II

12

13

14

15

16

log [Co"] t 2pH

Fig. 105. Aqueous dissociation relations of calcite, dolomite, and sepiolite in terms of log [Mg%]+ 2 pH and log [Ca"] '+ 2 pH at log [H4SiOz] = -3, -4, and -5, and log P(CO2) = -1.5, -2.5, and -3.5. Dashed lines (- - - - - -), 35OC; continuous lines (), 25OC; dotted lines (. . . . . .), 15OC.

This can be expressed also as: log[MgZ+]+ 2 pH = (log Ksep - 3 log[H,SiO$] - 4 log &)/2 where K , is the dissociation constant for water. Values of K , are taken from Barnes et al. (in Clark, 1966 p. 404), while values for Ksep are the extrapolated values of Christ et al. (1973). The sepiolite dissolution boundary, under different conditions of [H4SiOE] and at 15", 25", and 35"C, is presented in Fig. 105. Note that the Ksep chosen was that of well-crystallized material. A more poorly crystalline sepiolite, such as that synthesized by Wollast et al. (1968),requires a higher value of (log[Mg2+]+ 2 pH) for precipitation, although there is some doubt that the values reported by Wollast et al. represent equilibrium (Christ et al.,

1973). The interrelations between calcite, dolomite, sepiolite, and an aqueous solution at 25°C are shown in terms of (log[MgZ'] + 2 pH) and (log[Ca2'] + 2 pH) at several values of P(C02) and [H4SiOl] in Fig. 106. In this diagram lines represent equilibrium between a mineral and associated aqueous solution, while line intersections represent equilibrium between a mineral pair and the associated aqueous solution. We could also consider the dissociation boundary of magnesite (possibly nesquehonite) which would plot on a horizontal line in Fig. 106. Thermodynamic data are in some doubt, but the boundary is close to: log[Mg2+]+ 2 pH = 10.5 - log P(CO2) i.e., log[MgZ+]+ 2 pH = 12, 13, or 14 at logP(C02) = -1.5, -2.5, or -3.5, respectively.

190 16

,

1 4 \ : ; V 4 ' 0 ,

I

12

n

N

+

SOL 10

8

10

.

CAL 12

16

14

-3.0

10

16[

8

10

12

14

16

log [ H 4 S 1 0 ~ ] -~2 . 6

14

109 [~.a"]t 2 p~

Fig. 106. Stability relations of calcite, dolomite, and sepiolite at 25OC at log [H4SiOl] -2.6, -3, -3.5, and 4,and l o g P ( C 0 2 ) = -1.5 (dotted lines), -2.5 (dashed lines), and -3.5 (continuous lines). =

The appearance of sepiolite, not magnesite, places' a lower limit on [H4SiOz] in the environment - for any realistic P ( C 0 2 ) .In addition, the lack of a sepiolite-calcite association may place an upper limit on [H4SiO:], again dependent on P ( C 0 2 ) . The relationships between palygorskite and its montmorillonite precursor and the equilibrium aqueous solutions are complicated by the more complex chemistry of the mineral and the paucity of thermodynamic data. Field and chemical data, presented earlier, strongly suggest that the palygorskite formed directly from the montmorillonite, without significant solution and reprecipitation, and that both A1 and Fe are conserved in the montmorillonite-palygorskite reaction. We have based our calculations on a palygorskite (=PAL) composition ~ . + 4 8 / n ( M g 1 . 8 3 F e 0 . 3 6 A l 1 . 6 6 )( A ~ o . z S ~ ~ . S ) ~ Z O ( O H ) Z and a montmorillonite (=MONT) composition ~.+,/,(Mgo.,Feo.,Alz.,) (Alo.2Si7.8)020(OH)4.(Here X"' refers to exchange cations; Na', Mg", and/or Ca" are expected.) These structural formulae are based on those of Table IV,in particular of samples F5-13, FF-1, and FE-21, but are slightly modified to maintain the same tetrahedral site chemistry and the same Al/Fe ratio, thus allowing Fe t o be conserved when a reaction is written t o conserve Al. Such a reaction is: 5 PAL + 15.6 H' + 24.4 water = 3 MONT + 7.65 Mgz++ 0 . 3 / n p ' H4 SiOz

+ 15.6

log K p A k M o N T = 7.65 log[Mg2'] + 0.3/n log[X"'] + 15.6 pH + 15.6 log[ H4Si02]

191

To evaluate K p A h M O N T we require values of the standard Gibbs free energies of formation (AFF) of the species involved (log K = -AFi/1.364 a t 25"C, the temperature considered in the calculation). The AF; of water and appropriate dissolved species are tabulated by Tardy and Garrels (1974). Those for the Na, Mg, and Ca exchange forms of palygorskite and montmorillonite were calculated by their procedure of summing empirical factors of the structural formulae. They have shown that their method gives results within normal experimental error for montmorillonites, and our values (-2493.75, -2488.19, and -2496.34 kcal/mole for the Na, Mg, and Ca forms, respectively) are believed to be reasonable. The calculated values of A q (kcal/ mole) of Na, Mg, and Ca exchange forms of palygorskite are -2352.73, -2348.91, and -2354.51, respectively. In order to check the validity of the Tardy and Garrels method for palygorskites we calculated AFfo for a paly(Singer and gorskite of reported composition Mg,.3,Alo.46Feo.24Si4.~101~.~ Norrish, 1974). The calculated value is -1128.59 kcallmole, and the value determined by dissolution studies is -1131.25 kcal/mole (-1129.64 kcal/ mole if a correction of -3.5 kcal/mole is made for A T of aluminum-containing species, as suggested by Tardy and Garrels). The discrepancy is trivial, particularly when one notes that their reported composition has an impossibly high Si/O ratio, is not charge-balanced, and does nut consider exchange cations. Using these AFfo values we calculated K P A G M ~ N T for the reaction with X = Mg: log K p A G M O N T = 7.8 (log[Mg2+]+ 2 pH + 2 log[H4Si02]) = 44.85 i.e., log[Mg2+] + 2 pH + 2 log[H4Si02] = 5.75 This states that in, for example, sea water (log[Na'] = -0.48, log[MgZ+]= -1.87, log[ Ca2+] = -2.62) the palygorskite-montmorillonite equilibrium ( X = Mg) occurs at a value of (pH + log[H4SiOg]) = 3.81. If other exchange cations or combinations of exchange cations are used (pH + log[H4Si02]) = 3.81 f 0.01. Thus, provided that waters have (exchange) cation activity ratios similar t o that of sea water, the Mg system provides a reasonable guide to palygorskite formation, and this will be used henceforth. The dissolution of palygorskite and montmorillonite in alkaline waters, where Al(0H); is the dominant dissolved Al species, may be described by the reactions: PAL + 17.18 water + 2.28 H' = 0.18 Fe203(c)+ 2.07 Mg' + 7.8 H4Si0i + 1.86 Al(0H); and : MONT + 20.5 water = 0.3 Fe203(c)+ 0.85 Mg' +. 7.8 H4Si02 + 3.1 Al(0H); + 1.4 H'

192

PALYGORSKITE

-6

-5

-4

log H

[ 4

-3

sio,

-2

I'

Fig. 107. Stability fields of palygorskite, montmorillonite, and aqueous solution at 25"C, log [Al(OH)i] = -5.5.

Here we follow the suggestion of Tardy and Garrels (1974) in regarding the Fe as forming a poorly crystalline oxide (Fez03(+ AFF = -170.0 kcal/ mole). The values of log K for the reactions as written are:

The relations among palygorskite, montmorillonite, and an aqueous solution are shown in Fig. 107. Sepiolite and palygorskite deposits appear t o correlate with times of higher than normal environmental temperatures. Thermodynamic data on palygorskite at temperatures other than 25°C does not exist, but Christ et al. (1973) do provide such data for sepiolite. The sepiolite-aqueous solution and sepiolite-dolomite boundaries, plotted in terms of pH, log[Mg2'], etc. (Fig. 105), show an expansion of the sepiolite field at increasing temperatures. However, in a natural aqueous system, other interrelated reactions (e.g., dissociation of water) cause the solubility of sepiolite to increase with increased temperature (Christ et al., 1973). Any temperature effects favoring formation of sepiolite at higher temperatures must be indirect. In the range 0-60"C direct temperature effects on [Mg2'], [H4SiOg], and/or pH in sea water, due to changes in activity coefficients and complexing, are minor, and large changes in pH and ZCOz have trivial effects on [Mg2+] and [H,SiOi] (Lafon, 1969). Moreover, equilibrium with calcite places severe limits on sea water pH (drop of -1 pH unit from 8 t o 60°C). The activity coefficient of dissolved silica is essentially unchanged by moderate changes in temperature or salinity. (The activity of Mgz+ drops about 6% from simple concentration increase with a 3X increase in salinity). Palygorskites appear to have been formed by two mechanisms, direct pre-

193

cipitation from solution and by solid state transformation of a preexisting mineral (montmorillonite). Singer and Norrish (1974) document a probable example of the first mechanism, where palygorskite forms crusts on the surface of soil peds in Australia. We find branching fibers of palygorskite growing into voids, obviously precipitated from solution. However, field evidence and comparison of chemical analyses of purified samples suggest that the bulk of our palygorskite has formed from montmorillonite. The detrital montmorillonite and the palygorskite both have low tetrahedral A1 and essentially the same octahedral A1 and Fe. They differ in that the palygorskite shows a (mechanistically reasonable) increase in Mg in the (once dioctahedral) octahedral sites. Such a change in octahedral occupancy causes lattice strains which lead to the sheet- t o chain-silicate inversion (Weaver and Pollard, 1973). Transformation from other phyllosilicates (e.g., those with high tetrahedral Al) would appear t o involve chemical modifications which are energetically unfavored at low temperatures. Thermodynamic calculations (25" C) have been made for three reactions palygorskite-aqueous of direct concern; montmorillonite-palygorskite, solution, and sepiolite-aqueous solution. The stability-field boundaries for these reactions are defined by: log[Mg2+]+ 2 pH + 2 log[H,SiO!]

= 5.75

0.69 log[MgZ'] + 0.76 pH + 2.6 log[H4Si02] + 0.62 log[Al(OH);] = -10.70 and : log[Mg*+]+ 2 pH + 1.5 log[H4SiO:]

=

7.95

respectively. In all cases the chain silicates are favored by an increase in one or more of [Mg2+],pH, and [H4Si02]. Palygorskite also requires an appropriate input of A1 (and Fe), either inherited directly from the precursor clay or taken from solution, and whether palygorskite or sepiolite is formed will depend largely on factors (notably the detrital input of aluminosilicates, aluminum hydroxides, etc.) which determine the availability of dissolved Al. Otherwise, at levels of dissolved A1 commonly found in rivers and soils (few tens to hundreds of ppm Al), the palygorskite-aqueous solution and sepiolite-aqueous solution boundaries differ little on a pH-log[Mg'+]log[H,SiO~] plot. In the case of the solid state transformation montmorillonite-palygorskite the octahedral A1/Fe3+ratios in the two minerals are so close that there is very little direct dependence on activities of dissolved Al and Fe species. (Naturally, these activities cannot fall to such low values, as in a dynamic leaching environment, that palygorskite dissolution occurs). Singer and Norrish (1974) have determined a free energy of formation of their palygorskite, and they find that the occurrence or nonoccurrence of (authigenic, in situ) palygorskite in their soils can be predicted quite accu-

194

rately on the basis of its theoretical dissolution equation and chemical analyses of .soil water extracts. (In fact, their prediction is excellent if one assumes that the palygorskite cutans were originally montmorillonitic and considers the appropriate montmorillonite-palygorskite boundary.) Thus the thermodynamic data appear t o be a reasonable guide for predicting field conditions of palygorskite formation. Although there is no firm evidence that palygorskite forms in the ocean, and evidence presented elsewhere in this paper (fauna, associated dolomite, etc.) suggests that a brackish environment is involved, the study area was peri-marine, and it is convenient to consider ocean water as a starting point for discussion. According to our calculations the chain clays are stable under conditions near those of sea water (pH = 8.1, log[Mg2+] = -1.87, log [H4SiOs] = -4.7). A t the listed values of [Mg2+]and pH sepiolite requires log[H,SiO$] = 4.25 (around 3.0 ppm Si02 in sea water, assuming r(H4Si00,) = 1.13) for stability with respect to aqueous solution. (Effects of temperature and crystallinity changes are discussed elsewhere.) Palygorskite should form from montmorillonite at log[H,SiOB] 2 -4.29 (around 2.7 ppm Si02). At this value of [H4SiO:] palygorskite should be stable with respect to aqueous solution for log[Al(OH);] 2 4 3 (-0.1 ppm Al). Thus, from the point of view of thermodynamic calculations, only slight modifications of normal sea-water conditions are required to form sepiolite and palygorskite. However, if this were true these minerals should be more common. The calculations indicate the chain silicates are formed by an increase in [MgZ+], pH, and [H4Si0:]. Field observations indicate they are also favored by less than normal salinity and by high temperature. In our samples palygorskite has formed from montmorillonite under conditions of high [MgZ+]and pH. An alternative product might have been corrensite, and so it is of interest to explore the relation between montmorillonite, palygorskite, and corrensite. Corrensite is described as a regular 1 : 1alternation of chlorite and a 2 : 1 clay. The 2 : 1 portion is generally considered to be vermiculitic. Bradley and Weaver (1956) assign the formula M'MgsA13Si6020(OH)loto ideal corrensite. They describe a corrensite (Bradley and Weaver, 1956) with calculated structural formula indicating high interlayer charge, and analyses of other corrensites (Weaver and Pollard, 1973, table L) do suggest that the (tetrahedral) charge is high. However, the 2 : 1 component of the Bradley and Weaver corrensite expands to 17 A with ethylene glycol. All chemical analyses, and calculated structural analyses, appear t o be based on impure samples. We have considered three possible corrensites: (1)'ideal' corrensite, as indicated above, with exchange Mg; (2) Bradley and Weaver's (1956) corrensite, with formula approximated as Mg'o:&Mg7. &I. 2Si5.s&. sOZO(OH)IO; and (3) a hypothetical corrensite, formed by replacing half of the exchange Mg in our montmorillonite by a (Mg,A1)2(0H)6brucitic layer of appropriate composition to balance the charge. For simplicity, the Fe3+ in the formula was replaced by Al, leading to a composition M&~:7sMg3.1sA13.8sSi7.&lO20

195

16

-

I

I

-6

-5

-4 log H S O o L4.41

-3

1

I

-2

Fig. 108. Stability relations among simplified (Fe-free) palygorskite and montmorillonite and various corrensites at 25OC. Continuous lines (-), "ideal" corrensite; dashed lines ( - - - - - -), Bradley and Weaver corrensite; dotted lines (. . . .), hypothetical corrensite. See text for details.

. .

(OH),,. Values of AFF (kcal/mole), calculated by the method of Tardy and Garrels (1974) for these three corrensites, are, respectively, -3345.05, -3370.66, and -3148.07. The montmorillonite and palygorskite considered in the reactions are simplified (Fe3+ replaced by Al) varieties with calculated AFF = -2549.96 and -2385.76 kcal/mole, respectively. As an example, considering the 'ideal' corrensite (=COR):

2.22 COR + 10,08 H4SiO$ + 25.32 H'

= 3 PAL

+ 12.66 Mgz++ 40.92 water,

and :

3.7 COR + 1.2 H4Si02 + 57.8 H'

=

3 MONT +'28.9Mgz++ 43.8 water.

These reactions, at equilibrium, are described, respectively, by: log[MgZ+]+ 2 pH - 0.80log[H4SiO:] = 16.19 and : log[MgZ+]+ 2 pH - 0.04 log[H4SiO:! = '13.41 The paly gorskite-corrensite-montmorillonite relations are illustrated in Fig. 108. The same relations are also illustrated using the alternative corrensites. In many instances there is field evidence that suggests corrensite is formed from illite via an intermediate stripped illite. A suggested process is one which leaves the tetrahedral sites of an original illite unchanged, adds sufficient Mg to (di-) octahedral sites for approximate balance of the tetrahedral charge, and leads t o a regular 1 : 1alternation of brucitic and low-occupancy

196

interlayers. (Although corrensite is generally described as a regular 1 : 1 alternation of chloritic and vermiculitic layers, ethylene-glycol treatment generates a 314 basal spacing, indicating that the 2 : 1componect is probably a smectite.) We have chosen compositions of an Fe-free illite and its corrensite successor in a Mg-rich environment as:

KO.E(MgO.35All. 69 )Si3.43AO.5 7 0 1O(OH12 and Mg3(Mg1.5A13.38 )si6.8 6 4 1.14020(OH)10 The standard Gibbs free energies (kcal/mole), calculated by the Tardy and Garrels (1974) method, are -1318.73 and -3216.83, respectively. For the illite-corrensite reaction we write: 2 illite + 3.8 Mgz++ 6 water = corrensite + 1.6 K' + 6 H' The illite-orrensite

stability field boundary is then described by :

pH + 0.63 log[Mg2+]- 0.27 log[K'] = 8.17 The value of this expression in sea water is about 7.5, and only small increases in pH and/or Mg would be required to favor corrensite over illite. There appears t o be a mutual antipathy (at low temperatures) between clinoptilolite and palygorskite, with palygorskite being the fresher-water mineral and clinoptilolite being the more saline equivalent. To investigate the possible controls on this we can write a palygorskite + clinoptilolite reaction. A simplified clinoptilolite formula, ignoring structural water and varying amounts of substitution of K, Mg, and Ca for Na, is NazAl2Si,O2, + (Deer et al., 1963). There is some uncertainty as t o the value of x . Values ranging from 7 t o 10.5 are suggested. We have arbitrarily chosen x = 8.5. The reaction is then:

4 PAL + 6 Na' + 1 6 H' + water = 3 CLIN + 6.5 H4Si04+ 11Mg2+ log I( = 11log[MgZ+]+ 16 pH - 6 log[Na'] + 6.5 log[H4Si04] We thus see that palygorskite is favored over clinoptilolite at high pH, high [H4Si04], and a high ratio of [MgZ+]"/[Na'] '. (This general conclusion is not altered by selection of other values of x within the range 7 < x < 10.5; only the degree of dependence of [ H4Si04] changes.) A numerical value for log K requires knowledge of the standard free energy of formation of clinoptilolite. Using Tardy and Garrels' (1974) technique we calculate (for x = 8.5) a value of -2284.3 kcal/mole, assuming the Na is nonexchangeable, or -2296.9 kcal/mole, assuming the Na is exchangeable. For an ocean water (pH = 8.2, pNa = 0.48, pMg = 1.87) the calculated value of log[H4SiO:] for clinoptilolite-palygorskite is -2.11 or -7.91, depending on whether clinop-

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tilolite Na is exchangeable or nonexchangeable. The thermodynamic data is, admittedly, extremely imprecise, but it is compatible with a marine clinoptilolite facies and a high pH, high H4Si04paralic palygorskite facies. PHOSPHATE

In the Miocene of the SE U.S. the distribution of the indigenous phosphate is closely, but not exclusively, related to the distribution of the diatoms, largely Lower and Middle Miocene. Both occur in the Atlantic more so than the Gulf provinces. The phosphate has a general negative relation to calcite and dolomite, The deep-ocean currents are apparently the major source of the Si and P. Diatoms have a relatively high P content, 0.7%P in the ash (Vinogradov, 1953). In a study of phosphate deposits from the floor of the Peru shelf, Burnett (1974) reported that in interstitial waters of shallow cores, the distribution curve for dissolved silica was similar t o the PO4 curve. The PO4 content was well over two magnitudes above the apatite saturation value for ocean water. Diatoms are abundant in these sediments and he observed authigenic apatite on diatom frustules. He concluded that diatoms were the source of P which precipitated as apatite in the pore waters. A similar situation, though a shallow-water version, apparently existed in the Miocene. Gulbrandsen (1969), in a thorough discussion of the formation of marine apatite, showed that the precipitation of apatite rather than calcite is favored by increased pH, temperature, and PO, concentration. The evaporation G f sea water causes the coprecipitation of apatite and CaC03 which are both assumed to be 'in equilibrium with sea water. The addition of phosphate, causing an increase in the HPO;-/HCO; ratio, would allow apatite to be precipitated alone. An increase in temperature causes a decrease in the solubility of apatite (calculated). As C 0 2 solubility is inversely related to temperature and t o pH, temperature and pH are directly related. Both decrease the solubility of apatite. Gulbrandsen (1969) concluded that the optimum conditions favoring the formation of marine apatite are oxygenated water that is warmer, of higher pH, and of higher salinity than normal sea water, and an extraneous supply of phosphate (from organisms). These conditions are best obtained in shallow waters derived from nearshore upwelling cold ocean waters. A low rate of supply of detritus and a warm arid climate are also favorable factors. These conditions, with the possible exception of high salinity, existed in the southeast Atlantic Coastal Plain area during much of the Early and Middle Miocene. High salinity may be a questionable requirement for the formation of apatite, particularly if it forms by replacement. In the Miocene of the SE US. phosphate grains and pebbles occur in shallow-marine, brackish, and continental environments. In most instances it is difficult to determine which is primary and which has been transported. The presence of clay-phosphate pebbles with relatively high concentrations of

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sepiolite and palygorskite suggests that some of the apatite was deposited in waters of less than normal salinity. Experiments of Martens and Harris (1970) showed that the Mg in sea water inhibited the precipitation of apatite. They concluded that a Ca/Mg ratio higher than 4.5-5.2 was necessary for apatite to precipitate from solution. Thus, it is possible the sepiolite and apatite are coprecipitated. The formation of sepiolite, with Si obtained from the diatoms, would increase the Ca/Mg ratio, allowing apatite to form. The phosphorous would also be supplied by the diatoms. However, a more plausible process is that of replacement of preexisting minerals, probably clay minerals in this instance. Pevear (1966) has suggested that the Atlantic coastal phosphorites were deposited in estuaries and that much of the phosphorus was supplied by rivers. From a study of the phosphorites from the South African continental margin, Parker (1975) concluded they were formed by the replacement of lime muds in a shallow-water lagoonal--estuarine environment. Later they were reworked and transported t o the continental shelf. The Miocene phosphate pebbles most commonly occur in extremely shallow-water sediments and associated with unconformities, suggesting they were probably formed in coastal environments. The concentration of P on the Atlantic Coast and not on the Gulf Coast suggests that upwelling currents rather than rivers were the dominant source. Phosphatization of limestone and lime mud is an accepted method of concentrating phosphate (Parker, 1975), but in the Miocene of southeastern United States clays appear to have been the more common host rock. In the present Atlantic coastal area the only sizable concentrations of clay occurs in the estuaries, lagoons and marshes, with the shelf area being nearly devoid of clay. Pomeroy et al. (1965) have shown that estuarine muds act as reservoir (buffer) for phosphate and as the phosphate content of the overlying water increases much of it is adsorbed in the muds. As the phosphate in the water is decreased it is replenished from the mud. If this adsorption mechanism occurred in shallow water where the mud was periodically dried the phosphate could become deactivated by combining with or replacing the clay. Thus, the accumulation of phosphate could continue by periodic wetting and drying. Further, phosphate could be adsorbed by the clays directly from solution and a high concentration of organic matter would not be necessary. Various clay minerals will adsorb phosphate at lower pH values and lower concentration than that necessary for the precipitation of apatite. The similarity in size and configuration of the phosphate and silicate tetrahedra makes replacement and epitaxial growth rather easy. Weaver and Wampler (1972) suggested that such an adsorption mechanism, accompanied by clay solution, could account for the formation of some phosphates. In their samples they found that regardless of the nature of the matrix clay, illite was the only clay mineral preserved (or formed) in the phosphate grains. Other studies indicate that illite adsorbs less phosphate and is presumably

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TABLE VII Chemical analyses of phosphate and residue

S i02 A1Z03 Fe203

CaO MgO

K2 0 Naz 0 p2 0 5

Soluble

Soluble MgO

(2)

Residue * nonorganic (3)

(4)

(5)

41.8 18.0 6.0

67.90 20.50 1.24

28.88 16.62 8.17

48.18 27.72 13.63

27.5 1.3 5.5

0.44 6.19 3.69

40.05 0.0 6.27

10.46

99.99

99.99

Bulk

Nonapatite

(1)

4.50 1.94 0.64 50.50 2.96 0.14 0.59 33.0

-

-

94.27

100.1

100.0

* Corrected for 3.48%CaO and 1.35%P2O5. more stable in the presence of phosphate than the other clay minerals. The present study shows that partially phosphatized clay pebbles contain both the chain clays and montmorillonite. When phosphatizition is complete these clays are presumably destroyed. The concentration of clay minerals in the clay-phosphate pebbles is sepiolite > palygorskite > montmorillonite. This may reflect the high pH required for apatite formation and the effect of pH on the stability of these clays but their Mg content may also be a factor. Martens and Harriss’ (1970)data shows that Mg inhibits the precipitation of apatite. Perkins (1947)found that Mg decreased the amount of phosphate fixed by kaolinite. Thus, in a mud containing clays with varying Mg contents it might be expected that phosphate would replace the low-Mg clays first. (In fact, most sand size phosphate grains may have been clay-rich fecal pellets in which the clay minerals had been converted, in large part, to amorphous silicates, facilitating the adsorption of phosphate.) In order t o obtain some idea of the nature of the residue and the amount of dissolution that may have occurred, chemical analyses were made of the bulk cream phosphate and the HC1 residue (Table VII). The insoluble residue comprised 4.97% of the total sample. Of this, 70.56% consists of the oxides in column 3. CaO and P 2 0 3 comprise 3.59% of the residue. The remaining 25.85% is largely organic material. X-ray analysis of the residue indicates that K-feldspar, illite, and quartz are present. The chemical data confirms the X-ray data. The oxides of column 3 total 10.77% in the bulk analysis. Thus, only 32.5% of these oxides are present in the insoluble residue (silicate fraction). The remaining material (7.27% of bulk sample) in the soluble fraction, is presumably present in the apatite. Column 4 shows the composition of the soluble material and column 5 the composition assuming all the Mg is present as dolomite, though some of this is probably in the apatite. It does

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seem unlikely that so much Si, A1 and Fe would be obtained from seawater. All of the K is present in the silicate minerals which tends to confirm the stability of these minerals under conditions of high concentrations of PO4. The abundances Si > A1 > Fe is similar to that found in the clay minerals. It is suggested that this material was largely obtained by the replacement and dissolution of clay minerals by the phosphate and incorporated in the apatite structure. X-ray of the residues of cream and gray phosphate grains in NE Florida indicates only illite is present. However, there is considerable difference in the morphology of the clays. The material in the residue of the cream phosphate consists of bundles of fibers or laths, whereas that in the gray phosphate is flakes. The Miocene black phosphate pebbles and grains consistently have a higher Fe20, content than the cram pebbles. The latter could have formed by replacement of sepiolite-palygorskite (with a relatively low Fe content) and the former by replacement of montmorillonite (high Fe). This would also imply that the dark and light phosphate formed in different environments. The black type (montmorillonite) would form under marine or estuarine conditions and the white type (sepiolite-palygorskite) under restricted brackish and schizohaline conditions.

Chapter 9

OVERVIEW PALYGORSKITE IN THE OCEAN

‘Marine’ sedimentary rocks In the southeastern United States primary palygorskite grew only in environments (bay, lagoon, lake, or soil) where the salinity was less than normal sea water. Palygorskite in marine sediments is either detrital or secondary (by post-depositional circulation of brackish waters). We see no overwhelming reasons to believe this restriction is not universal. On the basis of studies of French deposits Millot (1970) was of the same opinion. However, after studies of the North African deposits he thought that palygorskite could also form under marine conditions. We will not make any attempt to evaluate the ‘marine’ palygorskite deposits in detail but will mention a few. In the eastern Sudan Basin the palygorskite occurs in Middle Eocene sediments which ,are underlain by montmorillonitic marine limestones (reefs and glauconite) deposited during transgression and overlain by Upper Eocene and Oligocene continental lacustrine sediments, relatively rich in kaolinite (Millot, 1970). Thus, it is possible that the palygorskite was deposited during a transition period (beginning of regression) when the water was brackish much of the time. To the east in Senegal Occidental the Lower Eocene pure palygorskite clays were deposited on a karst topography formed on an anticline of Paleocene limestone. They are overlain by palygorskite-containing marine calcareous clays and then limestone (Wirth, 1968). In this instance the palygorskite was deposited at the beginning of or preceding transgression. The sequence suggests that the water was brackish. The descriptions of the various lower Tertiary deposits of west Africa (Millot, 1970) indicate that for nearly every age that contains pure palygorskite clay (and no marine fossils) there is an equivalent age of fossiliferous marine sediments containing montmorillonite. Millot (1970) summarizes that in the western Africa basins phosphates and glauconites are intimately mixed. They occur in beds alternating with palygorskite, but do not occur in the palygorskite beds. As the former two minerals are almost certainly marine the alternation strongly suggests the palygorskite is not formed under marine conditions. Millot observed that there is commonly a seaward progression from kaolinite to montmorillonite to palygorskite to sepiolite. (The latter two may be reversed.) However, the assumption was made that the latter two minerals formed in the center of the basin. Actually, it appears that this sequence applies to the edge of major basins, and farther seaward, in the

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Fig. 109. Idealized transgressive-regressive sequence showing distribution of clay minerals in a marine to brackish to continental sequence under conditions where palygorskite forms. Palygorskite forms boundary between marine and continental environments. Continental montmorillonite zone often not present.

open-marine environment, montmorillonite will be encountered. Millot’s sequence would be restricted t o relatively small closed-lacustrine or brackish basins, not connected t o the open ocean. Fig. 109 is an idealized sketch illustrating the distribution of kaolinite, palygorskite, and montmorillonite in a regressive-transgressive sequence. Montmorillonite can occur both in the open-marine and in the mainland beach-marsh-continental environments. Kaolinite is a major constituent of some of the African palygorskite deposits. This suggests that some of the deposits are detrital with the palygorskite being derived from the edge of the basin. Paquet and Millot (1973) found that palygorskite is inherited and well preserved where the rainfall is less than 100 cm per year. If it can be preserved through the soil-forming process it can surely be preserved during transportation - either by water or air. The larger deposits commonly occur in partially and intermittently closed, shallow depressions flanking the open sea. The conditions are ideal for large amounts of palygorskite to be carried seaward. Further, slight changes in sea level will produce an intimate interlayering of marine and brackish deposits. It is too commonly assumed that because a thin bed of limestone or clay is marine the adjacent clay bed is also marine. In the past it has often been assumed that the presence of dolomite indicated a marine environment. As has been discussed most limpid dolomite probably formed under brackishwater conditions. The same arguments can be applied to the opal-cristobalite or chert frequently associated with palygorskite deposits.

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Another source of confusion is that in a regressive sedimentary sequence, brackish to fresh-water environments may overlie porous marine sands and limestones, and downward seepage can produce secondary palygorskite (see description of MC-1 core). Sepiolite is less sensitive. It can form under hypersaline conditions (Millot, 1970), highly alkaline conditions, and marine and brackish-water conditions (Hathaway and Sachs, 1965).

Deep-sea occurrences Gulf of Mexico and western Atlantic. Palygorskite is relatively abundant in deep-sea cores from the southern Gulf of Mexico, the Bahamas area, and off the northwest and northeast coasts of Africa. It is apparently detrital. Deep-sea Drilling Project X-ray analyses of Jurassic samples from site 100 Deep-sea Drilling Project Cruise XI, located between the Hatteras Abyssal Plain and the Bahama Platform indicate palygorskite is present (trace to approximately 40%) along with illite-biotite and mixed-layer illitesmectite in the 50 m of Upper Jurassic. This is underlain by 40 m of Middle Jurassic which rests on basalt. N o palygorskite was detected in this material which consists of smectite and either glauconite or poorly crystallized biotite. The samples containing palygorskite do not have a well-developed smectite. Instead they have a mixed-layer illitesmectite. This suggests there was a change in source area. The depositional environment for the entire section was considered to be bathyal. However, Hollister, Ewing, et al. (1972) state: ‘The approximately 50 meters of late Jurassic reddish limestone and calcareous mudstone contain numerous flow structures and clasts which indicate deposition in an active environment.’ It seems reasonable t o conclude that the palygorskite and associated clays are detrital and that they indicate a change in source during the Late Jurassic. The source was probably shallowwater Low Jurassic sediments. Significant amounts of palygorskite (30%) have been found in the Upper Cretaceous from site 97 and site 95 to the northwest of Cuba (Cook and Zemmels, 1973). The senior author has found minor amounts of this clay in the Cretaceous carbonates of southern Florida. Palygorskite has also been reported in the Eocene and Paleocene sediments in wells north of the Yucatan Peninsula (sites 89,94, and 95). Minor amounts are present from the Pliocene-Pleistocene from this area. Only trace amounts were reported in wells in the Caribbean (Rex and Murray, 1970). Sepiolite was reported from the Middle Eocene of one well (site 29), but was not found in three other wells in the immediate vicinity. This suggests a localized hydrothermal origin. A study of the descriptions of the sediments cored in the southern Gulf of Mexico (DSDP Cruise X) provides an insight to the origin of ‘marine’ palygorskite. The Lower Cretaceous at site 95, on the lower flank of the Campeche Bank (2,096 m below sea level), is described as consisting of

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dolomites and limestones containing solution breccia and algal mats. The environment is interpreted as backreef to supratidal. Iron-rich soil zones are also present. Sediments of Late Cretaceous and younger age are bathyal, indicating that major sinking occurred during the Cretaceous. Most of these sediments are considered to be slump deposits from the adjacent shelf. Coarse turbidite layers occur seaward in the deep basin. Thus, it is easy to envisage that the palygorskite in these marine sediments is really slumptransported shallow brackish-water deposits from the Lower Cretaceous of the Campeche Bank. Unfortunately no X-ray analyses were made of the Lower Cretaceous samples, but it is likely that palygorskite is abundant in these rocks and on the Campeche Bank in general. At site 97, in the strait between Florida and Cuba, the Upper Cretaceous contains an abundance of shallow-water clasts and pebbles (pebbly mudstones) suspended in a matrix of mixed deep-water ooze/clay and ‘shallowwater’ debris. Worzel, Bryant, et al. (1973) state that this type of sediment is common to deep-water sediments of many areas. At sites 4 and 5 (DSDP 1;Ewing et al., 1968) in the Abyssal Plain 60 and 100 km northeast of the Bahama Platform pebbly mudstones and turbidites are abundant throughout the Cretaceous section and in sediments as young as Oligocene. At site 98 bioclastic turbidites, as well as shallow-water limestone and perireef deposits, are present in the Upper Cretaceous (Paulus, 1972). This section contains the maximum amount of palygorskite and its clay is probably entirely detrital. During the Early Cretaceous a large barrier reef extended around nearly the entire circumference of the Gulf of Mexico (Paulus, 1972). The southern portion extended along the outer edge of the Campeche Escarpment. The northeastern portion was on the northern coast of Cuba and extended north along the Bahamas and the Blake Escarpment. A major backreef evaporite province existed throughout much of the area. The backreef and associated coastal sediments should contain an abundance of palygorskite. Weaver and Stevenson (1971) reporting on analyses made in 1955 of clays in the Cretaceous carbonates of south Florida noted the presence of a 9 : 1 mixed-layer illite-montmorillonite in the Lower Cretaceous. It now seems more likely that this material was palygorskite. It should be noted that the Cretaceous and Jurassic bathyal sediments contain little or no kaolinite though the climate was tropicalsubtropical. Kaolinite is relatively abundant in continental and shelf sediments of this age. Either kaolinite was not transported to the open ocean or it has had time to be destroyed in the sea water. Water content of the clays range from 20 to 40%(Hollister, Ewing, et al., 1972). Lower Miocene sediments were not found in the hemipelagic continentalrise sediments along the western flank of the Atlantic and in the Straits of Florida, though they are present on the continental shelf off Georgia. This could suggest that, at the closing of Tethys, marine bottom currents were

205

strong enough to sweep the area clean. Middle and Upper Miocene sediments are present. Lancelot et al. (1972) point out that the Middle Miocene to Pleistocene sediments in this area (from Bahamas to New York) have a similar clay-mineral suite. The relative abundance of chlorite, mica, amphiboles, and pollen with northen affinities suggest that most of the clays were transported from the north. Smectite is relatively more abundant in the Oligocene through Upper Cretaceous section. The clay suite in these latter sediments is similar to that of the Miocene on the continental shelf except for the presence of detrital palygorskitesepiolite (derived from coastal Georgia) in the Lower Miocene. The illite and kaolinite content of the Miocene clays increased from Lower to Upper Miocene and seaward from the coast t o the Blake Plateau, showing the increasing influence of the northern current. The influx of clays brought in by southward-flowing (bottom) currents in the Middle Miocene suggests that the closing at Gibraltar allowed cooler northern waters to penetrate far to the south. The clay data tends to confirm the paleontological data (Berggren and Hollister, 1974). Tropical surface currents may have continued to flow t o the north (Gibbson, 1967). Though Middle Miocene sediments are abundant on the continental rise they appear to be missing from the continental shelf off the Georgia-Florida coast (JOIDES, 1965), and have a restricted distribution on the mainland. Phosphate pebbles were formed throughout the area at this time. Many believe the relatively abrupt increase in clastic sediments in the Middle Miocene indicates an uplift of the Appalachians. The distribution of Middle Miocene sediments in the southeastern United States suggests the whole general area, particularly the Ocala High, was bowed up and marine deposition was largely restricted to a narrow trough. This gentle uplift apparently coincides with the collision of Africa and Europe at Gibraltar. Another interpretation is that the Middle Miocene was a time of extensive glaciation and sea level was lowered 70-100 m (Tanner, 1965). Zemmels et al. (197 2) suggest that the relatively abrupt increase in illite and chlorite and decrease in smectite in the post-Oligocene sediments (Lower Miocene missing) may indicate increased cooling. More likely it indicates the influx of sediments transported by cold northern waters. Thus, in this area the change in clays is due to a change in current pattern which may in turn be related t o glaciation. These data also indicates a similar increase in illite and chlorite in passing down from the Kimmeridgian t o the Oxfordian. Does this indicate glaciation during the Oxfordian? It is plausible that tectonic activity, changing current patterns, and glaciation all occurred near the beginning of the Middle Miocene. One would expect the controlling factor to be the tectonic activity which would effect the pattern of the ocean currents which would in turn control the development of glaciers.

East Atlantic Ocean. Palygorskite, and to a lesser extent sepiolite, is relatively abundant (70-90% in many samples) in six wells off the northwest

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coast of North Africa (DSDP Cruise I1 and XIV) (Rex, 1970; Berger and von Rad, 1972). It is most abundant in the Eocene, Paleocene, and Upper Cretaceous sediments. Palygorskite is abundant in the Paleocene and Eocene sediments in the many basins fringing the coast of northwest Africa (Millot, 1970). It is present, but less common, in the Upper Cretaceous. Berger and von Rad (1972) and others concluded the palygorskitesepiolite sediments were formed authigenically in deep water. The palygorskite-sepiolite facies is characteristically a relatively clean clay. It passes abruptly into younger deep-water chalk ooze and is underlain by basalt. Near the base it commonly contains alternating thin beds of coarse, retransported, shelf dolomite and sapropelic clay. Light-gray to white dolomitic palygorskite clay layers alternate with dark organic clay layers. Near the coast thin layers of clean sand occur in the clay. Red and gray banding is common. The clay is nearly devoid of fossils but contains fish teeth and dwarf forams. A mottled structure and palygorskite clay clasts are abundant. (Many of the palygorskite clay beds could be largely clasts.) Volcanic material and chert are abundant throughout the section. Virtually every feature observed in these supposed deep-sea clays, except for the volcanic material and the organic-rich beds, can be observed in the coastal brackish-water palygorskite deposits of the southeastern United States. Further, hiatuses are abundant, and are believed to be caused by strong currents (Berger and von Rad, 1972). The basalts encountered at sites 138 and 1 4 1 have been partially serpentinized. As Hathaway and Sachs (1965) and Bonatti and Joensuu (1968) have shown, palygorskite and sepiolite, along with serpentine, appear to be common subsea hydrothermal alteration features of basalt. One could hypothesize that large amounts of hydrothermal palygorskite and sepiolite are created along the mid-Atlantic ridge, which was nearshore during the Cretaceous and Early Tertiary, and were transported into the deep-water basins. This is a possibility, but it is more likely that most of the palygorskitesepiolite and dolomite had a coastal, brackish-water origin. The problem is whether these sediments were transported from coastal areas or were formed in place in shallow-water environments. Basins containing palygorskite-sepiolite and dolomite, but lacking organic clays and volcanic ash, flank the northwest coast of continental Africa (Millot, 1970) extending from Morocco (and as far north as France) to Angola. Assuming palygorskite and limpid dolomite do not form in normal sea water it is reasonable to assume these minerals were transported from coastal areas into deep-marine environments. (DSDP Cruise XIV authors are willing to transport dolomite and sand from coastal areas but not palygorskite.) Alternatively, the continental rise and abyssal plain could have been areas of brackish, shallow-water sedimentation during the Cretaceous and Paleogene. A study of interstitial waters from several wells in this area (Waterman et al., 1972) showed a systematic increase in NaCl with depth. They concluded that evaporite deposits (probably Lower Cretaceous) existed at depth. If this

207

is true, shallow-water conditions probably existed in this area immediately prior to the deposition of the Upper Cretaceous palygorskitesepiolite. Peterson et al. (1970) point out that the presence of possibly marine Jurassic outcrops of chert in the Cape Verde Islands suggests this is the western flank of a large shallow basin (extending onto continental Africa) that formed during the Jurassic as the Atlantic started rifting. By Early Tertiary time the basin, or series of basins, was downwarped and buried. Palygorskite and dolomite could have formed on the brackish-water flanks of the large basin and on islands in the basin, and could have been transported towards the center. It is not impossible, on the basis of the available data, that portions of the nearby closed basin could have had oceM water sufficiently diluted with fresh water to allow palygorskite t o form in place. The abundance of basalt on the flanks and base of the basin would provide an abnormally high influx of Mg and Si. Quartz, feldspar, mica-illite, chlorite, kaolinite, and some montmorillonite are considered to be terrigenous components. However, palygorskitesepiolite, which was formed in coastal lagoons closer to the edge of the basin, is not considered to be present in the transported suite. This is extremely unlikely, though the amount transported is open to question. The palygorskitesepiolite clays commonly contain 3-4 other clay minerals, including kaolinite. Few definite authigenic deposits contain such a complex clay suite. Further, zeolites are a common component of these clays, but in the southeastern United States zeolite and palygorskitesepiolite are mutually exclusive. Zeolite does not appear to be present in other low-temperature deposits of these clays, although it has been reported in deep-sea hydrothermal deposits (Hathaway and Sachs, 1965; Bonatti and Joensuu, 1968). In most regions where authigenic palygorskite-sepiolite is present the sequence from the edge towards the center of the basin is kaolinite --* montmorillonite + palygorskite -+ sepiolite. If the basin is evaporatic it is chlorite-montmorillonite + chlorite (illite is commonly present in all stages). In any given sample only two or at the most three of these clays are usually present. Many of the clays sampled off the coast of northwest Africa contain all of these clay minerals - certainly a nonequilibrium assembledge. Whether large dolomite rhombs can form under normal-marine conditions is open to question (Folk and Land, 1975). The dolomite in the Messinian evaporite deposits of the Mediterranean is believed to have formed in shallow-water sediments during periodic influxes of meteoric waters (Nesteroff, 1973a). This further suggests that many deep-water dolomites are detrital. The possibility exists that the banded dolomite and clay deposits of the northwest coast of Africa may indicate the periodic influx of meteoric waters into a relatively shallow, partially closed basin, rather than indicate a periodic influx of coastal dolomite.

Indian Ocean. Palygorskitesepiolite has also been found off the east coast

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of Africa (DSDP XXIII, XXIV, XXV). Palygorskite is present in the Upper Cretaceous in wells situated on the Mozambique Ridge (site 249) and the Somali Basin (site 241), located 350 and 550 km off the present coast of Africa. At site 249 the Upper Cretaceous is separated from the montmorillonitic Lower Cretaceous (which rest on basalt) by a major hiatus. It is also separated from the overlying Miocene by a hiatus. The Upper Cretaceous clayey chalk is considered to be an offshore facies of the Grudja Formation which outcrops on the coast. A t site 241 the equivalent-age deposits are turbidites (Girdley et al., 1974). Thus in both areas there is an excellent chance the palygorskite has been transported from the continent. At site 241 similar amounts of palygorskite (10-30%) occur in the overlying Tertiary and Quaternary sediments, suggesting they are detrital. Palygorskite is relatively abundant (30-50%) in the Paleocene and Eocene clayey sediments in the vicinity of the Island of Madagascar, and is present in minor amounts in most of the younger deposits. The age distribution is similar to that off the west of Africa. The available data does not give any obvious suggestion of origin. Abundant volcanic material and terrigenous material is present. Until some evidence of deep-sea growth is presented it is more logical t o conclude that the palygorskite is detrital, as is the quartz, illite, kaolinite and some of the montmorillonite. In fact there is little specific evidence to indicate that much of the montmorillonite is formed in the deep-sea environment. Palygorskite is present, but with erratic distribution, in the cores of northwestern Indian Ocean Lower Tertiary sediments, obtained during Cruises XXIII and XXIV (Matti et al., 1974a, b). The most consistent trend can be observed in the Miocene. The palygorskite content increases systematically from approximately 10% in the vicinity of Madagascar t o 4 0 4 0 % near the southern coast of Arabia. This is also true in the Pliocene and Pleistocene. This suggests a detrital source from Arabia where palygorskite is abundant (Muller, 1961; Wiersma, 1970). Further, the maximum amount (up to 65%) occurs in brecciated zones. The zones are believed t o be due t o slumping. Younger sediments contain turbidite beds (Whitmarsh et al., 1974). Goldberg and Griffin’s (1970) study of ocean-bottom sediments in the same area showed a similar increase in palygorskite (and coarse dolomite) towards the coast of Arabia. They believed the distribution indicated an eolian margin. Red Sea. In the Red Sea palygorskite is relatively abundant in the Pleistocene and Pliocene sediments overlying the Upper Miocene evaporites (DSDP Cruise XXIII). The sediments are described as gray micarb-rich detrital clay nanno-ooze and chalk and a gray micarb-rich nanno detrital silty claystone (Whitmarsh et al., 1974). These sediments also contain large dolomite rhombs, and Stoffers and Ross (1974) concluded that both the dolomite and the palygorskite are detrital. The authors would agree with this interpretation. It is of interest t o note the dolomites associated with the upper part of the evaporite sequence are fine grained and d o not contain palygorskite.

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Mediterranean Sea. Zemmels and Cook (1973) report finding palygorskite in many of the cores of Neogene and Quaternary sediments from the Mediterranean (DSDP Cruise XIII).Nesteroff (1973b) reports he found none but instead found interstratified clay minerals which had probably been misidentified as palygorskite. In any event the sections which are reported to contain significant amounts of palygorskite are described variously as: Lower Cretaceous dolomitic rock fragments, turbidites, pebble clasts, restricted and brackish-water fauna, etc. (Ryan, Hsu, e t al., 1973). Thus, if palygorskite is present, most or all of it is likely detrital. GLOBAL DISTRIBUTION

Distribution The global distribution of the major palygorskite deposits suggests it is temperature-dependent aS well as salinity-dependent. Figs. 110 through 113 show the distribution of palygorskite for various periods of geologic time. Most of the locations are referred to by Millot (1970) and Wiersma (1970). More recent references include Bohor (1975), Martin (1975), and Sartbaev (1975). Deposits range in age from Triassic t o Quaternary, Most of the postMiocene deposits are small, lacustrine-type deposits, and much of the clay is detrital. Palygorskite and sepiolite occur in Carboniferous sediments in Russia (Rateev, 1964; Zaritsky and Orlov, 1973) but the age of these clays is not clear. On a Permian paleogeographic map the Triassic-Jurassic deposits occur in a relatively narrow east-west band extending westward from the western nose of the Tethys (Fig. 110). The palygorskite beds in South Africa occur with Triassic volcanics (Heystek and Schmidt, 1954) but are of questionable age. It is of interest t o note that the Permo-Triassic red bed-evaporite deposits of Morocco and France characteristically contain chlorite, corrensite,

..-

....... .....'

Fig. 110. Permian paleography showing location of Triassic and Jurassic palygorskite deposits. Triassic deposits occur between 15's and 10'N paleolatitude; Jurassic deposits between 10°N and 25'N paleolatitude. Map after Johnson (1973).

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swelling chlorite, and mixed-layer chlorite-montmorillonite (Millot, 1970). This is also true of the Triassic evaporites of Germany (Lippmann, 1956), Spain (Vivaldi and MacEwan, 1957), England (Honeyborne, 1951), and the United States. Though some sepiolite occurs in evaporite deposits it appears that the more common clay mineral is chlorite, in some stage of formation. Thus chlorite is the stable Mg clay in hypersaline water and palygorskite in brackish water. The starting material may also be a factor. Palygorskite is commonly formed from montmorillonite and the various chloritic clays are in part formed from degraded illite (Lucas and Ataman, 1968). Under hypersaline conditions (high pH) Mg hydroxide apparently precipitates in the interlayer position forming chloritic minerals. Where there is less tendency to form hydroxides (lower pH) the Mg ion migrates into the octahedral sheet, increasing layer strain and forcing tetrahedra t o invert forming palygorskite. When a precursor clay is not present Mg and Si may coprecipitate to form sepiolite. Cretaceous palygorskite deposits occur along the flanks of the warm Tethys ocean and its western extension in the Caribbean. Palygorskite is also present in central west Africa (Fig. 111).These latter deposits are presumably related to the opening of the South Atlantic. During the early Creta-

EARLY

CRETACEOUS

I

I

EOCENE

Fig. 11 1. Early Cretaceous paleography showing location of Cretaceous palygorskite deposits. All deposits are apparently Late Cretaceous in age. E a s t w e s t line shows boundary between tropicalsubtropical and temperate fauna during the Maestrichtian (after Davids, 1966). Deposits occur between 15's and 40'N paleolatitude. Map composited from Dietz and Holden (1970)and Berggren and Hollister (1974). Fig. 112. Eocene paleography showing location of Paleocene and Eocene palygorskite deposits. Two deposits in western Africa and one in Spain are Paleocene. Other deposits are apparently restricted t o the Lower and Middle Eocene. X indicates soil deposits probably inherited from Eocene sediments. Middle East deposits are more numerous than ' s and 35ON paleolatitude; Eocene deposindicated. Paleocene deposits occur between 5 ' 5 and 45'N paleolatitude. Map composited from Dietz and Holden (1970) its between s and Berggren and Hollister (1974).

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ceous extensive salt deposits were formed in the Angola-Brazil Basin offshore from the continent. In the late Cretaceous pelagic clays and chalks were deposited (Bolli, Ryan, et al., 1975). This suggests that conditions were more humid during the late Cretaceous when palygorskite was formed on the flanks of the basin. Palygorskite deposits are probably present on the western flank of the basin in South America. The most extensive palygorskite deposits occur in Paleocene and Eocene sediments (Fig. 112). Most of the deposits occur in North Africa and Europe. The distribution is largely controlled by the warm Tethys currents which extended into Europe and also swung south along the northwest coast of Africa (Berggren and Hollister, 1974). The Yucatan deposits are believed to be of Early Eocene age (Bohor, personal communication, 1975). Dorf (1960) concluded that during the Tertiary the maximum temperature in the northern hemisphere occurred during the Paleocene and Early and Middle Eocene. This is the time of maximum development of palygorskite. Berggren and Hollister's (1974) review of paleotemperature data indicates that in the Late Cretaceous tropicalsubtropical surface marine temperatures existed as far north as northern Europe (50"N) but in North America only extended to the south tip of Florida (25"N). It appears that all the Cretaceous palygorskite-sepiolite deposits were formed in the tropicalsubtropical zone which was warped t o the south in the western hemisphere. Tropical conditions extended as far north as southern England during the Early Tertiary but cooling began in the Late Eocene and continued to the PRESENT

Fig. 11 3. Present paleography showing location of Oligocene and Miocene palygorskite deposits. Miocene deposits are slightly more abundant than Oligocene deposits and in Europe occur farther south (Spain vs France). Oligocene deposits occur between 20"N and 50"N paleolatitude. East-west line represents Miocene 21°C isocryme based on faunal data (Cheetham, 1967). Map after Dietz and Holden (1970).

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Pleistocene. From a study of mussels Strauch (1968) calculated the following yearly average temperatures for central Europe: Middle Eocene 27°C; Late Oligocene - 23.5"C;Miocene - 22°C; Pliocene - 17.5"C. Oligocene and Miocene palygorskite deposits appear to be restricted to the Middle East, Europe, and Georgia-Florida (Fig. 113). In Europe, the Oligocene deposits lie farther north (France and Germany) than those of the Miocene (Spain). This distribution is presumably related to the regional cooling trend that started in the Late Eocene.

Pale0 latitude Figs. 114 and 115 show the variation through time of the paleolatitude of the northern and southern boundaries of palygorskite and of kaolinite plus evaporites. Cenezoic and Mesozoic distribution of kaolinite is based on data of Vlodarskaya (1962), Konta (1968) and Murray and Patterson (1975). Also shown is the relative position of central Europe. Recent to Pliocene deposits appear to be restricted to soils and lakes. The northern boundary should be fairly well established but the southern boundary is questionable since the southern hemisphere has been less thoroughly explored than the northern. However, it seems unlikely that undiscovered major deposits exist. Palygorskite occurs in South Africa (Heystek and Schmidt, 1954) and southern Australia (Rogers et al., 1954). Both deposits

Fig. 114. Paleolatitude of northernmost and southernmost major deposits of palygorskite as a function of time. Thin line is paleolatitude Qf central Europe through time. Also shown is location of Holocene and Pleistocene (and perhaps Pliocene) soil and lacustrine deposits. Paleolatitude data from Green (1961) and Berggren and Hollister (1974).

213 90

I

70 60 80

20-

30-

-

40 50

0

I00

200

300

400

SO0

Fig. 115. Paleolatitude distribution of kaolinite-Iaterite (thick) and evaporite (thin) deposits in northern hemisphere as a function of time. Evaporite data from Green (1961). Kaolinite-laterite data from Konta (1968) and others.

are intimately associated with basalt flows and were apparently formed from basalt by alteration by fresh waters. They cannot be considered the equivalent of the marine-related deposits. The age of both is uncertain. The northern limit of palygorskite is relatively fixed with regard to the geography of the continents, and has systematically moved away from the equator as the continents drifted north. This suggests that extremely high temperatures are probably not a controlling factor. The broader 40-50" wide belt in the Upper Cretaceous and Paleogene indicates that favorable conditions were much more widespread than in more recent and older times. The northern kaolinitelaterite boundary is roughly parallel to that of palygorskite, but lies slightly to the north, more so in the older deposits. This distribution suggests that kaolinite-laterite and palygorskite share some of the same requirements for their formation. The distribution of evaporite deposits (Green, 1961) is similar to that for kaolinite, except that during the Paleozoic the northern boundary apparently occurs slightly north of the northern kaolinite boundary. Conditions favoring the formation of palygorskite were widespread during the Paleogene and Late Cretaceous and systematically became limited to a narrow latitude range in the Early Mesozoic. Presumably the required conditions were not present in the Paleozoic, whereas environmental conditions, climate, etc., favorable ,for the formation of evaporites and, to a lesser extent, kaolinite existed throughout the Paleozoic.

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Both the kaolinite and evaporite belts are wider in Permian and older sediments, which would tend t o indicate relatively high temperatures and a relative abundance of hypersaline waters. Shallow shelf and enclosed-basin environments, which are the normal sites where palygorskite forms, would tend t o be evaporitic and favor the formation of corrensite rather than paly gorskite. A comparison of Figs. 114 and 115 shows there is an inverse relation between the width of the paleolatitude of palygorskite and evaporites. The paleolatitude range of the evaporites is most restricted when the range of palygorskite is the most extensive. This is further evidence that, apart from occurrences in soils, palygorskite is formed under humid conditions.

Role of continental drift From Cretaceous until late Early Miocene (Burdigalian, ca. 1 9 my ago) the continents of Eurasia and Africa and North and South America were separated, and the Tethys warm current flowed westward between the northe m and southern continents (Berggren and Hollister, 1974). Tropical to subtropical conditions existed on the shores of the Tethys. This is confirmed by Millot (1970) and others who found abundant kaolinite, bauxite, and lateritic deposits in the continental sediments adjacenk t o the Tethys Sea. The warm Tethys currents transported tropical fauna to the Caribbean-Gulf Coast region. However, as North and South America were farther apart than today and the gap in Central America occurred close to South America the warm current had a relatively moderate effect on the southern part of North America. During the early Miocene (18 my ago) the Tethys was separated into an eastern and western portion by the junction of Africa and Eurasia forming the eastern Tethys (Indian Ocean) and the western Tethys (Mediterranean), ‘Flow of warm, salty Mediterranean water into the eastern Atlantic began,’ (Berggren and Hollister, 1974). Axelrod and Bailey (1969) found that the peak of the Miocene warming trend in several areas coincided with this closing of the Tethys. They suggest that the western flank of the Atlantic would be most affected. The palygorskite deposits in the southeastern United States range in age from Late Oligocene through Early Miocene (Aquitanian and Burdigallian). Thus, they were formed just prior t o and slightly after the closure of the eastern Tethys. If palygorskite is temperature-dependent it would appear that the effective closing of the Tethys and the accompanying modification of the Atlantic currents began as early as Late Oligocene (approximately 25 my ago). Later, at the beginning of the Middle Miocene (14-15 my ago) Europe and North Africa collided at Gibraltar, forming the Gibraltar Sill, and causing an almost complete cessation in faunal interchange between the western Tethys and Caribbean regions. ‘Surface circulation of the Atlantic was significantly modified. The Gulf Stream became a self-containing system. A part

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of the current was deflected northward t o form the North American Drift. The incursion of warm waters into the northeastern Atlantic probably enhanced circulation in the North Atlantic with the extrusion of a greater volume of Artic waters into western Atlantic . . .' (Berggren and Hollister, 1974). This suggested circulation pattern indicates the temperatures of the ocean waters fringing the southeastern United States were at a maximum during the Early Miocene and became cooler during the Middle Miocene and later. It does not seem mere coincidence that formation of palygorskite in the southeastern United States stopped when the circulation between the tropical western Tethys and the Atlantic stopped. Most of the palygorskite in the Southeastern United States was deposited during the time interval between the beginning of separation of the eastern and western Tethys (by the junction of Africa and Eurasia) and the separation of the western Tethys and the Atlantic (by the junction of Europe and North Africa). Tropicalsubtropical pelagic fauna occurs throughout the Miocene of southeastern United States (Huddlestun, 1975). However, Gardner (1926) concluded from a study of benthonic fauna that the Middle Miocene was considerably cooler than the Early Miocene. Abbott (1975) found an abundance of diatoms (Denticulu)in the Middle Miocene of Sduth Carolina which suggested the waters were cooler than at present. The Middle Miocene clays contain abundant opal phytoliths characteristic of prairie grasses which grow in a low-rainfall environment (Abbott, 1975). From a study of the tropical bryozoan Metrurubdotus Cheetham (1967) concluded that during the Oligocene and Early Miocene the 70°F (21°C) isocryme in the southeastern United States occurred at about 31"N (South Georgia). All Lower Miocene palygorskite deposits occur south of 33"N. In Europe the boundary was at about 40-48"N. In the Late Miocene (no samples were examined from the Middle Miocene) the boundary shifted southward t o 26" in America (south tip of Florida) and perhaps to 28" in Africa. It seems to be a reasonable assumption that in the southeastern United States ocean temperatures became cooler and the climate drier in the Middle Miocene, causing the end of the formation of palygorskite and phosphate. The effect may have been direct or indirect. PALYGORSKITES IN SPACE AND TIME

Temporal distribution A possible clue t o the conditions of formation of palygorskite and sepiolite is the antipathetical relation with time of the abundances of these minerals and other important Mg-bearing sedimentary minerals, notably dolomite and corrensite. These minerals all form under near-surface conditions, and obtain most of their Mg from solution. They, together with possible basaltic precursors and Mg-calcites, are a factor in the Mg budget of the

216

oceans. However, montmorillonite, illite, and “normal” chlorite are also important, even though most of their Mg is obtained directly from the volcanic material from which the parent clay formed or was acquired after burial. The concentration of Mg in these clays is relatively small but due to their relative abundance the total amount is large. Fig. 116 was constructed t o show, in a highly generalized fashion, the abundance with time of authigenic Mg minerals (non-recycled) and kaolinite. Palygorskitesepiolite first occurs in any abundance in Triassic sediments. The other high-Mg clay, mixed-layer chlorite-montmorillonite (= “corrensite”) is abundant throughout the Paleozoic and Early Mesozoic. As a generalization, the amount of dolomite decreases in younger sediments while the amount of palygorskite increases. Dolomite and corrensite are both relatively abundant in the Paleozoic. Micritic dolomite forms in hypersaline waters while limpid dolomite forms in brackish waters where [MgB] is relatively low. As much coarse-grained dolomite is formed by post-burial recrystallization of micritic dolomite it is possible that primary limpid dolomite has the same time-stratigraphic distribution as palygorskite. Corrensite and (nonlimpid?) dolomite are both relatively abundant in the Paleozoic. This was a consequence of deposition in shallow epirogenic seas, with hypersaline waters (Shaw, 1964) and (possibly) relatively low concentrations of silica (see below). As lithospheric rifting commenced, and deep oceans developed, more and more Ca was incorporated in CaC03 in the open-marine environment unfavorable for the development of dolomite. 1

I

Cen0z.l

30

-

I

Mewzoic

Paleozoic

Total Carbonate

--- ----- ----Dolomite

c

0

200

300

400

TIME (M.Y.)

,

500

600

Fig. 116. Estimate of relative abundance of some selected minerals through a portion of geologic time. Source of data is discussed in text.

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Corrensite-like clays are relatively abundant in the Paleozoic and Triassic, decreasing in abundance in Jurassic and Cretaceous time as palygorskite starts to increase. Both corrensite and dolomite are relatively rare in Cenozoic sediments. Corrensite may be a red herring. We have caluclated the stability-field boundaries between our palygorskite, our montmorillonite, and a series of different corrensites. Regardless of the choice of corrensite composition, it is favored over montmorillonite by higher [MgB] and pH. The [H4Si0:] effect is minor. For the corrensite-palygorskite reaction the importance of [MgH] and pH is variable, depending on the choice of corrensite composition. However, in all cases high [H4SiO:] favors palygorskite. This might be interpreted as supporting the statement that the data shows an increase, in the Cenozoic and Mesozoic, of SO,-rich Mg-bearing minerals (palygorskite, sepiolite) at the expense of Si02-poor (or poorer) Mg-bearing minerals (dolomite and corrensite). However, it is difficult from the point of view of reaction mechanism to go from palygorskite (or sepiolite), characterized by low tetrahedral Al, to corrensite, characterized by high tetrahedral Al. It is more likely that the precursor of corrensite is illite. From the thermodynamic calculations presented earlier we see that illite is favored over corrensite in sea water. Under evaporitic conditions (suggested in many instances for corrensite formation), which were common in the early development of ocean basins during the Paleozoic and Early Mesozoic we might expect highpH conditions to be common. In addition, with a moderate increase of ionic strength over that of present sea water the activity coefficient ratio y (Mg’*) /y(K)increases markedly, favoring corrensite. Though the amount of corrensite and dolomite is small they make up a significant portion of the authigenic marine-related Mg minerals formed during this time. The major geologic factor controlling the distribution of these minerals (and kaolinite) appears t o be climate - generally arid during the Paleozoic and Early Mesozoic and becoming progressively more humid in the Late Mesozoic and Cenozoic. Temperature is probably a secondary factor. Schwarzbach (1961) compiled evidence t o show that, for the last 500 million years the climate in North America and Europe was mostly warmer and much drier than today’s. He also concluded that, beginning in the Late Mesozoic, rain fall was more akin t o the present day than t o that of the relatively dry Permian and Triassic periods. However, the Carboniferous is considered t o have been damp. Humidity conditions were erratic in the Tertiary, but it would appear that in general humid conditions prevailed, favoring the development of brackish-water environments. The maximum development of kaolinite and laterites occurred during the Cretaceous and Early Cenozoic (Konta, 1968; Millot, 1970; Murray and Patterson, 1975), confirming that humid conditions were widespread in North Africa, Europe and North America. Palygorskite and kaolinite-laterite have a similar distribution (discussed in

218

the section on Paleolatitude), both forming more abundantly during warm, humid times. The coincidence in their formation does not necessarily indicate a direct, common origin. Lacustrine palygorskite would be favored by the increased stream load of Si and Mg released by the accelerated weathering. In the case of coastal palygorskites, an increase in the size of the North Atlantic would increase the amount of rain fall in the coastal area, thus increasing the amount of brackish-water environments. The opening of the North Atlantic would allow Si-rich water t o move into the northern waters from the South Atlantic (Heath, 1974), adding t o Si from oceanic rifts and, in the case of coastal waters, from continental weathering. DSDP cores indicate that, at various times from the Late Paleozoic to Early Cenozoic, vast amounts of salt were withdrawn from the Atlantic into deepocean brine pools and stagnant basin salt deposits. Such episodes would lead t o brackish waters over much of the surface of the oceans (Fischer, 1964). Periods of abundance of palygorskites and of salt removal do not coincide, but we may speculate, in the absence of firm data, that return to a more humid climate and a lag in the return t o normal ocean salinity may allow the development of coastal brackish conditions (favoring palygorskite) t o follow the evaporative episodes. The palygorskite (and sepiolite) versus dolomite antipathy is of prime concern t o us. Pre-Mesozoic palygorskites may have existed and been destroyed by burial diagenesis. Mumpton and Roy (1958) found that under hydrothermal conditions palygorskite will alter t o montmorillonite at 200"C and probably as low as 100°C. However, montmorillonite is also scarce in Paleozoic sediments. In the natural systkm the sequence of burial diagenesis should be as follows: palygorskite -,montmorillonite + MgAdolornite + SiAChert illite-montmorillonite

K

K

>

ioo--20aoc

illite

>200°c

It is possible that most pre-Triassic sediments have been exposed to temperatures higher than 100°C and all traces of palygorskite destroyed. There is an equally good chance that relatively little was ever formed in the Precambrian and Paleozoic and that the evolution of the earth's crust is not entirely cyclic. We will assume the latter explanation, and for simplicity will discuss the case of sepiolite versus dolomite. Sepiolite and dolomite may compete for available Mg according to a reaction of the type: SEP + 3co2 + 4Ca2++15 water = 4DOL + 6H4SiO! + 8H' If these minerals are formed in marine-related waters we may consider their abundance-age pattern in terms of possible variations in the P(C02), pH,

219

[H4Si04], [Mg”] and/or [Ca”] of sea water during the Phanerozoic. (We have already seen that an increase in temperature expands the sepiolite field at the expense of dolomite in a system where the equation above is valid, and where other factors are equal.) In a series of papers Ronov (e.g., 1964, 1968) has examined the continental sedimentary record, and his interpretation of this, in terms of crustal evolution and tectonism--volcanism, suggests a changing ocean (atmosphere) chemistry, notably a general decrease in P ( C 0 2 ) (marked change at the Paleozoic-Mesozoic boundary) and oceanic alkaline earths (particularly Ca). Such changes would favor a redistribution of Mg from carbonate minerals (e.g., dolomite) t o silicates (e.g., sepiolite, palygorskite, and various phyllosilicates). However, Garrels and Mackenzie (1971) were able to reinterpret Ronov’s data in terms of differential weathering rates of the major sediment types and post-depositional alteration. Their analysis requires n o marked changes in the major element composition of the ocean during the Phanerozoic. Others have attempted to assess the limits on sea water variations imposed by observed marine (and marine evaporitic) chemical sediments (notably calcite and gypsum). Lafon’s (1969) calculations show that, with present-day major-element composition, temperature,induced (0-60”C) excursions in pH and inorganic carbon species in calcite-precipitating sea water are minor. Holland’s (1972) analysis of marine evaporites suggests that Ca, HC03, and SO4 have not varied by as much a factor of two during the Phanerozoic. Aluminosilicate reactions (“reconstitution” of weathered debris) in the ocean (or indirectly connected t o the ocean through exhalations from deeply buried sediments undergoing advanced diagenesis or metamorphism) also influence ocean-water composition. Here the factors involved are much more complicated than in the carbonatesulphate system (e.g., Siever, 1968), but again our gross observations on the sedimentary record do not require a changing (major element) ocean composition (Garrels and Mackenzie, 1971). We may expect small compositional excursions as the oceans respond t o changing continental weathering inputs (tectonism, volcaniccontrolled C02, etc.), episodes of submarine volcanism (basalt-sea water interaction), evolutionary changes (locus and nature of biogenic carbonate deposition), etc., and there is some evidence that these have, in fact, occurred (Broecker, 1974). We have shown that the required departure of sea water from its present composition to one favoring sepiolite-palygorskite formation is small. Small evolutionary or periodic fluctuations may in fact have increased the possibility of Mg-silicate versus Mg-carbonate formation in the younger Phanerozoic. However, direct independent evidence for the appropriate changes does not exist.

Environment and source material There is no firm a priori reason t o suggest that the changing nature of authigenic Mg minerals is due t o changes in the major-element chemistry of the oceans. However, there may be changes through time in certain peri-

220

marine environments. These may involve changes in minor-element chemical cycles in the ocean or climate-induced changes (see above) in salinity and continental input to brackish coastal environments. Millot (1970), Wiersma (1970), and others indicate that the Mg and Si necessary for the formation of lacustrine palygorskites around the Mediterranean is supplied by the rivers from a tropically weathered hinterland where kaolinite is formed. The relation may merely be that both minerals are preferentially formed under humid, tropical conditions. The following quotation from Millot (1970) bears repeating. “The Tertiary series in Africa show, with a considerable amplitude, phenomena of sedimentary neoformation. We cannot consider them as curiosities of puny character. The cumulative thickness of tabular silica can reach some tens of meters, phosphatic beds constitute mineral deposits in which one can move about in galeries, beds with 100% palygorskite have a thickness of 500 meters in Senegal. It is a question of another style of genesis for silica and silicates, synthesis in environments of a marked alkaline chemical character. This epoch favorable for chemical sedimentation was again an extraordinarily calm epoch during which the immobile continents were dissolved of the major part of their substance. But towards the close of the Eocene, agitation set in again, climates changed, there was erosion, and the siderolithic facies invaded everywhere. At the same epoch, Europe was going through an analogous history. We can grasp only the appearance of phenomena, and once again the causes reside in the dynamics of the Earth’s crust. It is understandable that Alpine orogenesis provoked the reworking of the weathering mantles in Europe, but we must agree that its effect is just as clear and occurred at the same time in Africa. The rhythms of sedimentation can define the age of these oscillations.” In Europe the close association of lacustrine palygorskite deposits with continental kaolinite-laterite leaves little doubt, as stated by Millot (1974), that terrestrial weathering was the source of the Si and Mg. In the southeastern United States kaolinite is the major clay in the Miocene continental sediments, but the volume is relatively small. Phosphates are generally absent in the lacustrine deposits of Europe and elsewhere, but they are commonly associated with the marginal-marine palygorskites. For example, both phosphate and palygorskite are closely related in North Africa and the Middle East (south edge of Tethys). Both are abundant in the Upper Cretaceous and Eocene, and are scarce in the Lower Cretaceous. The mechanisms and sources for the introduction of SiOz t o marginalmarine palygorskite deposits appear to be different from those for the lacustrine deposits. A variety of factors may lead t o increases (perhaps localized and temporal) in oceanic [H4SiO:], and hence favor the formation of sepiolite and palygorskitk over dolomite. The first appearance of palygorskite (and a marked decline in the abundance of dolomite) occurs in the Ewly Mesozoic. This was a time of initiation of sea-floor spreading, and attendant introduction of silica (and perhaps Mg) into the oceans. Many of the occur-

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rences of palygorskite in the early maximum (see Fig. 116) are associated with land areas (southeastern U.S.A., Spain, North Africa, East Africa) near actively spreading ridges. A second pulse of palygorskite formation began during the Late Cretaceous. At this time oceanic ridges were far at sea. Silica was abundant in the Mesozoic and Cenozoic seas, as indicated by the abundance of silicious organisms, zeolites, opal-cristobalite, and chert. Rad and Rosch (1972) state that much $ofthe oceanic SiOz is derived from volcanogenic materials. However, in a review of the SiOz cycle in the present ocean, Burton and Liss (1973) concluded that volcanics are a minor source, and most of the SiOz is supplied by Antarctic weathering and rivers. Regardless of source, the dominant control of oceanic SiOz concentration is the formation and later dissolution at depth of silicious tests of diatoms, radiolarians, etc. This biological cycle does not lead t o high SiOz concentrations in open-ocean surface waters, but upwellings may introduce large amounts of SiOz and nutrient elements, such as P, t o coastal waters (Calvert, 1966). Silica and P may be retained in shallow-water, semi-enclosed marginal environments as diatom tests and by absorption onto clays, respectively. The coincidence of the pulse of palygorskite formation in marginal-marine areas and an evolutionary spurt of diatoms in the Upper Cretaceous, plus the association with phosphates, suggests this source for these deposits. Our field evidence suggests that the commercial palygorskite beds were formed in a low-lying coastal-lagoonal environment. Sedimentary features (e.g., mud cracks, sepiolite-rich clay pebbles) suggest a shallow-water environment of occasional high energy, subject t o periodic desiccation. A soil horizon divides the two commercial palygorskite beds, and itself contains palygorskite. During periodic marine invasions montmorillonitic clay and minor silt were introduced into these basins. An early event in these marine invasions was the deposition of rice-grain calcite in the desiccation features, raising the Mg/Ca ratio in solution. Apart from this, calcite is absent, and the carbonate associated with the palygorskite (and its precursor montmorillonite) is dolomite. There is considerable evidence (e.g., Folk and Siedlecka, 1974; Folk and Land, 1975) that authigenic dolomite formation is favored by waters of low salinity, and that in such waters dolomite can form at molar Mg/Ca ratios as low as approximately 1 : 1. (See also Hanshaw et al., 1971; von der Borch et al., 1975). Badiozamani (1973) has calculated the activities of appropriate dissolved species in mixtures of sea water and Yucatan ground water (similar in composition to north Florida karst water), and has shown that dolomite is at or above saturation and calcite is undersaturated in a wide range of brackish compositions. Faunal evidence is sparse, but that which we have suggests that, at least for some of the time, the waters associated with palygorskite formation were brackish. There is no direct thermodynamic reason why the formation of palygorskite from montmorillonite should be favored by low salinities. The activity coefficient of MgB shows a marked increase with decreasing salinity from sea-water composition, but this is more than balanced by the dilution effect. Other

222

environmental factors must be involved. Either suitably high pH can be attained or high activities of Mg* and/or H4SiO! can be reached in the lagoonal environment. If we accept the interpretation that palygorskite formed in a brackish environment the activity of MgB cannot have been high. Even karst waters draining N Florida dolomitized limestones do not contain more than about 10meq/lMg (Hanshaw et al., 1971) and present surface karst waters in N Florida are considerably less concentrated. However, for the equilibria palygorskite-aqueous solution (log [ Al(OH),] = -6), sepiolite-aqueous solution, and montmorillonite-aqueous solution only a few parts per million Mg are required at pH -9, log [H4SiO:] -4. We will see that such pH and silica values are reasonable. Although [Mg*] need not be high, a large influx of Mg is required to account for the observed mineralogy. One source is metastable Mg-calcite, and a SEM picture shows palygorskite fibers growing out of the surface of montmorillonite-bearing calcite patches, the fibers becoming stubbier as they pass into the enclosing montmorillonite. However, calcite deposition was minor in the commercial beds. Biotite (with associated zeolite) is relatively abundant (5-2076) in some of the montmorillonitic sediments underlying, overlying and laterally equivalent (marine) t o the Lower Miocene palygorskites, but it. is not present in the palygorskite sediments. However, unreasonably large amounts of biotite would have had t o be destroyed for adequate in situ production of Mg. It would take approximately a 1 : 1 mixture of montmorillonite and biotite t o provide the Mg necessary for the formation of palygorskite. It is more likely that Mg was released from the biotite t o rivers during intense weathering (very warm, humid conditions) in the Piedmont. Periodic flushing with sea water or constant influx of continental waters (Piedmont rivers and/or karst waters from the Ocala High) is the more likely mechanism for introducing Mg. Modest increases in pH (to around 9) and [H,SiO;] (to around are quite reasonable in the environments postulated for the commercial palygorskite beds. In fact, a serious problem is the relative infrequency of palygorskite and sepiolite formation. Shallow-water dolomitic environments, such as the ephemeral lakes around the Coorong (von der Borch, 1965),attain pHs as high as 10 during periods of intense photosynthetic activity, and we might expect similar high pHs in shallow lagoons, etc. in the subtropical to tropical Lower Miocene. Various possibilities exist for attaining high silica concentrations under such conditions. Diatoms, which are abundant throughout much of the Miocene sediments, provide an important means of concentrating readily mobilizable silica. The amorphous silica of their tests has a high equilibrium solubility (log [H4SiOs] -2.6 at 25°C) which increases with temperature and with pH at about pH = 9. W e might postulate that in the Middle Miocene environments, where abundant diatoms were deposited, the intense diatom activity maintained [H4Si0:] at low levels and a cool climate prevented attainment of conditions suitable for later remobili-

-

-

223

zation of this silica. Only detrital palygorskite and sepiolite is found in these sediments. During the Early Miocene, on the other hand, a suitable warm climate might have lead t o extensive diatom dissolution, and consequent rise in [H,SiO!], in shallow environments. Only limited numbers of (brackish water) diatoms are now found in the commercial palygorskite beds, but they are present in palygorskite clays from other areas. Where diatoms are not present detrital quartz and aluminosilicates may be invoked t o provide a silica source. Peterson and von der Borch (1965) describe extensive quartz corrosion in the high-pH Coorong environment. Here the silica is reprecipitated as an opaline gel. They point out the common occurrence of corroded quartz and of cherts in dolomitic sediments (Peterson, 1962). Corroded quartz is present in the palygorskite beds. In the Lower Miocene the silica tends to be incorporated in chain clays rather than precipitated as opal. An important factor in this regard may be the role of organic-matter decay in controlling pH. In a lacustrine environment, such as that surrounding the Coorong, where there is abundant incorporation of decaying orgaic matter into the sediment, low pHs lead to silica precipitation (Peterson and von der Borch, 1965). We note some traces of organic matter in the Lower Miocene clays, particularly in the fossil-soil horizon between the commercial beds, but this is a very intractable colloid, resistant even t o the HF-HN03-HC104 treatment used for mineral dissolution prior t o chemical analysis. Frequent subaerial exposure and desiccation promoted oxidation of organic matter at the surface and prevented extensive organic-matter buildup within the sediment. (Jones et al., 1969, note that in Great Basin lakes there is depletion of C C 0 2 in playa flat interstitial waters relative t o that in the interstitial waters of the more permanent portion of the lakes). Thus Z C 0 2 cannot attain high levels, pH remains high, and silica enters silicates rather than cherts). Inorganic carbon species, generated by organic decay or by other means, may be important in determining whether available Mg enters a carbonate (dolomite) or a silicate (sepiolite, palygorskite). Dolomitization is associated, in many instances, with ground waters (Hanshaw et al., 1971; von der Borch et al., 1975). The high CICO, in such waters favors the formation of carbonate minerals. An important element in some of these theories of groundwater dolomitization is the infiltration or brine reflux addition of sea water t o provide a Mg source. Badiozamani (1973) has calculated the chemical speciation in various mixtures of sea water and a typical karst ground water. He shows that in ground waters with small admixtures of sea water (with the low ionic strengths favoring dolomite nucleation) the increase in [Mg”] may lead t o dolomite saturation but the lesser increase in [Ca”], in conjunction with a decrease in dissolved inorganic carbon, maintains calcite below saturation. This mechanism of dolomitization depends on the system retaining the high CCOz of the ground water. If the mixing occurs in a system open to the atmosphere, as in the case of periodic marine influx into a basin generally dominated by continental surface waters, the P(C02) would remain near

224

the (lower) level of equilibrium with the atmosphere while the [MgZ+]increase .effect would be unchanged. High temperature would also favor CO, loss.

APPENDIX X-RAY The sharpness of the X-ray peaks indicates that most of the palygorskite is fairly well crystallized, though the patterns of some samples show considerable broadening. There is no obvious relation of crystallinity to morphology. The presence of quartz in most samples makes it difficult to determine if there is a 121 (4.27 A ) reflection for the palygorskite. However, it does appear to be absent, indicating that the palygorskite is orthorhombic (Nathan et al., 1970). The 110 spacing of these palygorskites varies from 10.5 to 11.0 8. When palygorskite is treated with 'various cations (Al, Mg, Ca, Na, K, NH4) the spacing ranges from 10.6 A to 10.8 A. Thus, a portion of this variation could be due to variations in the exchange cations, but most of it probably reflects variations in the Al/Mg ratio. Smectite is almost always present in palygorskite clays, though some carbonate residues are relatively pure. Some form of fine-grained mica is present in all, or nearly all, palygorskite samples. K-Ar analyses of the less than two-micron fraction of two palygorskite samples gave apparent ages of 125 m.y. and 260 m.y. confirming that much of the K is present in detrital mica.

0

20

40

80

00

% PALYGORSKITE

Fig. A l . X-ray standard curve based on ratio of area 10.5-w palygorskite peak to 15-A montmorillonite peak.

226 T h e 110 spacing of sepiolite ranges from 12.0 8 to 1 2 . 5 8. The patterns indicate that the laths are poorly crystallized (Brindley, 1959). The 130 reflection (6.4-6.7 8 )is considerably stronger than reported for other sepiolites and may indicate t h e presence of appreciable palygorskite either interlayered or mixed with t h e sepiolite. QUANTITATIVE X-RAY ANALYSIS It is extremely difficult to mix clays of t w o such differing morphologies as palygorskite and montmorillonite and obtain quantitative X-ray data. Mossman e t al. (1967) described a quantitative method, using zinc hydroxide as a n internal standard, for determining t h e composition of clay-mineral suites. An attempt was made t o use this internal standard method. For the pure components, montmorillonite (15 8) has 1.4-1.8 times t h e scattering power of palygorskite (10.5 8).When the t w o components are mixed (along with the standard) t h e apparent difference in scattering power increases to from 3 to 4. This indicates t h a t the thin flakes are coating the long fibers. Much of the lack of linearity is due to t h e Zn(OH)2 flakes. F o r example, as t h e palygorskite content decreases from 100%t o 70% t h e peak area is reduced by one-half. The peak area of t h e Zn(OH)* increases as the montmorillonite increases. Thus, t h e Zn(OH)z is apparently finer than t h e montrnorillonite and settles last. In this case the internal standard does n o t appear t o be of any help. In order t o obtain a usable routine working curve the ratio of t h e peak areas of palygorskite and montmorillonite were plotted vs % palygorskite (Fig. A l ) . The curve is nearly linear between 30 and 9 0 % palygorskite. T h e difference in syattering intensity in this range is approximately 3 . T h e Wyoming bentonite used was Mg saturated and had a n 001 spacing of 1 5 8. The scattering intensity would be greater when the material is glycolated (17 8);however, as the Miocene montmorillonite has a broader peak, commonly extending t o higher 2 0 values, it was thought that a value of 3 was a realistic conversion factor.

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