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 (
33

It is not uncommon for palygorskite t o be the dominant or only clay in Tampa limestones. Part, probably all, of the palygorskite in the limestones is derived from reworking of the underlying material, as indicated by the presence of clay clasts and clay grains. The dolomite is basically a micrite and intramicrite with areas of larger, well-developed spherical crystals of 5-10 pm in diameter. The intraclasts in the limestone are also rnicrites but have not been dolomitized. The limestone is basically an intramicrite with spar filling the interior of fossils. The distribution suggests the lower section (dolomite) was originally deposited in a low-energy tidal-lagoonal environment. The overlying limestone micrite was apparently deposited in a more openmarine environment where calcite was precipitated and the Mg and Si concentrations were not high enough for palygorskite to form. The sequence in the Lake Talquin core (Fig. 9) demonstrates the environmental and mineralogical history of the Appalachicola Embayment. As transgression of the Oligocene surface started, palygorskite and pebbly dolomite (Tampa) were deposited in a restricted tidal-lagoonal environment. This interval is followed by sediments deposited during the Early Miocene (Torreya). The initial deposition is a montmorillonitic, oyster-rich, sandy coquina indicating brackish to marine conditions. The marine influence continued t o increase till near the middle of the section. Some detrital palygorskite was deposited with these fossiliferous sands and clays. The upper section represents a regressive sequence. The marine shelly sediments are overlain by worm-burrowed micritic limestone presumably deposited in a quiet shallow-marine to open-lagoonal environment. Continued regression is indicated by the upward gradation into a dolomitic sand containing palygorskite, then into a palygorskite clay and dolomite sequence (containing minor phosphate). These latter sediments were presumably deposited in a restricted, closed lagoon. The overlying sand is presumably a mainland beach. The phosphate would appear t o be authigenic as it only occurs in the most restricted environment. The pinch-out of the carbonate facies can be seen in the generalized crosssection (Fig. 10) extending from Lake Talquin north to the Climax and Fall Cave outcrops on the west flank of the Trough (southwest Georgia). In the Caves the fuller’s earth zone is believed t o rest directly on the Tampa. A similar situation exists t o the northeast of Lake Talquin. The contours of the isopach map (Fig. 11)of the clastic facies (sand and clay) generally follow the total Miocene thickness contours in Georgia except that the relatively thick lense of clastic sediments in the Trough is divided, near the middle, by a thin interval. This is compensated for by a relatively thick limestone section (Fig. 6). As will be shown later this is a critical hinge area and is probably related t o the Ocala High. In the Appalachicola Embayment and the eastern portion of the panhandle clastics form a relatively minor part of the total thickness. An isopach map of the sand shows maximum thickness in the north-

34

0-

12-

v)

a W c W

24-

-

r 2

-

I

c a

-

-

W

n

36-

48-

O y s t e r s abundant

LAKE

TALOUIN C O R E W 6890

-

Fig. 9. Transgressive-regressive sequence in core from northeast edge of Appalachicola Embayment, Florida.

eastern half and the southwestern quarter of the Trough, the southeastern portion of the Atlantic Embayment and northern portion of the Appalachicola Embayment. In most wells the sand is described as fine to coarse (Herrick and Vorhis, 1963), but there is some variation. Many wells in the southwestern portion of the Trough are described as fine, and fine to medium sand. Sands in the updip edge of the northeastern portion of the Trough tend t o

35 W 68.90

2

GG S 494

Climaa cove

Fell cove

O-

W

1

f 30 c L W

60J

Fig. 10. Generalized southwest-northeast cross-section showing the wedging-out of the transgressive-regressive sequence shown in Fig. 9. Left edge is same well as shown in Fig. 9. Primary palygorskite is restricted to Tampa dolomite and fuller's earth zone. Intervening marine sediments contain montmorillonite.

be coarser (medium to coarse). In most of Florida, the Lower and Middle Miocene quartz sand is described as fine to coarse and fine to medium (Carr and Alverson, 1959;Espenshade and Spencer, 1963;PirMe et al., 1965;and our own analysis).

/

100' counfour I n t w v a l

mo'

,303N

.2oxy

Fig. 11. Isopach map of the Miocene clastic sediments. Maximum thickness occurs slightly south of the present course of the Altamaha River and at the confluence of the Ocmulgee and Oconee Rivers (formiqg the Altamaha River).

36

Gravel and pebble-size material (quartz, chert, and feldspar) is relatively abundant in the upper portion of the section on the west flank of the Trough and on parts of the Ocala High. Puri and Vernon (1964)report gravel is present in the eastern portion of the Florida panhandle. They consider them deltaic (leaf- and log-bearing), and have given them and their equivalent finer-grained alluvial sands the informal name Fort Preston (Middle Miocene) and Misccosukee (Upper Miocene), though they may be as young as Pliocene. Scattered pebbles and thin beds of pebbles occur throughout the Miocene section in a n o r t h s o u t h belt extending through the Atlantic Embayment and the northeastern portion of the Trough. The quartz and feldspar pebbles are commonly associated with phosphate pebbles. The Atlantic Embayment appears to have been a fairly high-energy estuary during much of the Miocene. Descriptions of sands by Herrick (1961)and ourselves indicate that there are two mineralogically different types which appear to be nearly mutually exclusive. One type is described as phosphatic, the other as arkosic. Fig. 12 is an isopach map of the interval from the top of the Miocene to the first occurrence of phosphate or T.D.

\

Thickness of i n t e r v a l the l o p o f t h e M i o c e n e

to t h e First O c c u r r e n c e o f Phosphate too'

-

countour I n t w v a t /OO'.JO.~H 2OIM

Fig. 12. Maximum thicknesses of terrestrialderived sediments (no phosphate) are related to present river systems. Contours west of no-phosphate line indicate thickness of entire Miocene section.

37

In the Atlantic Embayment all wells have phosphatic sands and the sands usually contain interbedded limestone. To the west in the Trough area the lower sands are usually phosphatic and the upper arkosic. These phosphatic sands commonly contain shells rather than limestone beds. Farther t o the west the sands become completely arkosic. In the southwest the sands are neither arkosic nor phosphatic. The thickest nonphosphate interval occurs in the center of the state (Georgia) where the Miocene rivers were probably located. The eastwardextending lobes suggest the location of major rivers (Fig. 12). The abrupt decrease in thickness in the southwest portion of the Trough coincides with a hinge area. There appear to be at least three sandstone provinces. In the southwestern part of the area sands were derived from Cretaceous and Eocene sands to the north. These contained little feldspar or the feldspar was lost during weathering. To the northeast the shore line was closer t o the igneous and metamorphic rocks of the piedmont, and the sands have a higher feldspar content. The phosphate is related to the Atlantic Ocean and its distribution presumably reflects the effects and distribution of Atlantic waters as opposed to Gulf of Mexico waters. The presence of arkosic sands overlying phosphatic sands presumably reflects the withdrawal of the sea in Late Miocene time and/or regional uplift in the southern Appalachians. The arkosic sands probably were transported by the ancestral Ocmulgee, Oconee, and perhaps Ogeechee Rivers. The abrupt northeastern trend of the Ocmulgee River before it joins with the Oconee to form the Altamaha may be due to its being deflected by the northeastern extension of the Ocala High in post-Miocene time. The X-ray analyses indicate both K- and Na-feldspar are present. K-feldspar is dominant throughout the Miocene section in the western portion of Georgia, and in the Florida panhandle, particularly in the Appalachicola Embayment. Na-feldspar becomes relatively more abundant to the east and in peninsular Florida. On the continental shelf and on the northeastern flank of the Ocala High it is as abundant or more abundant than K-feldspar. This suggests the K-feldspar-rich suite was largely transported by the local piedmont and perhaps Ocala High streams and the Gulf currents and the Na-feldspar-rich suite by Atlantic longshore currents from the north. The clay and equivalent sands (Lower Miocene) in the commercial clay mining area in north Florida and southwest Georgia appear t o contain only K-feldspar as does the Tampa limestone to the west of the Trough area. Nafeldspar may be unstable in the environment in which palygorskite forms; however, it appears that the source area is the controlling factor. K-feldspar is relatively abundant in the clays and sand overlying the southern palygorskite beds; Na-feldspar is present in some samples. The northern commercial clay beds (Middle Miocene) contain both types of feldspar and the overlying montmorillonite clays and sands only K-feld-

N LL W

Level

4 13..

m Fig. 13. Lithologic crosssection extending from St. George Island through the Appalachicola Embayment, down the axis of the Trough to the Savannah River. For location of wells see Fig. 24. Blank areas are predominantly carbonate rocks. Carbonates predominate to the southwest, and sand and clay to the northeast.

39

spar, These clay beds are largely detrital in origin and were deposited in a more normal-marine environment than the clay beds to the south. Detrital (grains and clasts of clay) palygorskite-rich clay in the Savannah area also contains both Na- and K-feldspar. The data is too limited to determine if all clay beds containing detrital palygorskitesepiolite grains contain both types of feldspar, but available data suggest this is the case, and that the reworking was done by Atlantic currents. As the Atlantic Ocean withdrew, sediments (containing K-feldspar) from the western source were deposited on top of the Atlantic sediments. Major phosphate deposits appear t o be restricted t o the flanks of the Ocala High and the northeastern edge of the Atlantic Embayment. These presumably are areas that were exposed t o shallow-water high-energy environments for considerable lengths of time, allowing the phosphate grains to be concentrated. The updip western edge of the Miocene sediments in Georgia is composed primarily of clay rather than sand. It is not clear why this should be, but much of this material is described as having weathered fragments of the underlying limestone and may be a residual clay. The clay has a high kaolinite content. Cross-sections are useful in showing the lithofacies relations in the Miocene. Fig. 13 is a southwest-northeast cross-section down the axis of the Trough. The general picture is a southwest limestone facies interfingering with a clastic facies which appears t o be predominantly sand and then clay and sand to the northeast. The top of the Suwannee limestone appears as an arch through much of the section. As will be shown in the next section this is due to post-Early Miocene movement of the Ocala High. The lower massive limestone, presumably Tampa, is relatively uniform in thickness through the Southwestern portion of the trough and across the Ocala High and thins to the northeast where interbedded sands and limestones and sandy coquina beds are present in the lower part of the section. The interval thins and becomes more clastic northeast of the basin. Maximum thickness occurs in the vicinity of the Altamaha River. The Lower and Middle Miocene sediments also have this lateral facies change. The carbonate-clastic boundary migrated southward from Late Oligocene to Late Miocene time.

This Page Intentionally Left Blank

Chapter 3 REGIONAL CLAY-MINERAL DISTRIBUTION FLORIDA

A study of the Miocene of peninsular Florida by Reynolds (1962) indicated that the carbonate rocks of the Tampa and Hawthorn are largely dolomite. The dolomites and dolomitic clay beds contain a montmorillonitepalygorskitesepiolite clay suite. Sepiolite is ‘almost exclusively’ confined to the dolomites and montmorillonite generally predominates in the less-dolomitic carbonates and pure clay beds. Southward in the South Florida Basin montmorillonite and illite become the only clays present and the carbonates are largely limestones. The clastic Hawthorn sediments in the eastern portion of the peninsula commonly contain a montmorillonite-illite clay suite with minor amounts of palygorskite, though some of the lower silt units contain the montmorillonite-palygorskite-sepiolite suite. Outcrop samples from peninsular Florida collected for the present study tend to confirm these generalities, though in detail there are significant excqtions. The fossiliferous, marine Tampa limestone from near the city of Tampa contains palygorskite with a lesser amount of montmorillonite. Tidal, dolomitic intramicrites contain only palygorskite. To the east, in the Lake Wales area, Miocene dolomites contain essentially pure palygorskite (and abundant cristobalite and phosphate grains). At a quarry near Ocala the Tampa is a pure calcite limestone; some samples contain montmorillonite and sepiolite in approximately equal amounts and the clay fraction of others is composed almost entirely of sepiolite and palygorskite. These latter sampIes are biomicrites and most of the clay exists in rounded transported grains. The overlying clays and sands (considered t o be Lower Miocene) contain sepiolite and opal, in varying ratios, with relatively little palygorskite. Laterally equivalent clay beds a few miles away consist largely of montmorillonite and opal with minor amounts of palygorskite and sepiolite. The latter two clays are present in spherical, sand-size clay grains. These clay beds contain diatom fragments and leached sponge spicules. This major change in the clay-mineral suite over a short distance suggests that the various depositional environments have a relatively small size. Similar relations can be seen farther north and are described in detail. Farther north near Gainesville (Devels Hoper) the basal sediments (Tampa?) are relatively rich in palygorskite and sepiolite; the Lower Miocene contains predominantly montmorillonite and minor amounts of sepiolite and palygorskite in both the dolomites and limestones and clay beds. The overlying Upper Miocene-Pliocene sediments contain only montmorillonite. The clays in scattered samples of pre-Tampa limestone from peninsular Florida are largely montmorillonite with little or no palygorskite or sepiolite. These

42

latter clays are restricted to the Tampa (Upper Oligocene) and younger sediments. A large number of samples were studied from wells, outcrops and mines extending from 40 km east of Tallahassee t o the westernmost portion of the Florida panhandle (Fig. 14). In the eastern portion of the cross-section the kaolinitic Upper Miocene Miccosukee sands overlie the Lower Miocene Torreya and Hawthorn Formations. This is underlain by the Upper Oligocene Tampa Formation which pinches out in the easternmost well or slightly t o the east of it. Palygorskite and sepiolite are apparently absent in the eastern continental sediments and montmorillonite is the dominant clay. To the west, palygorskite and sepiolite become abundant in several intervals. In the eastern section apatite is present throughout the Lower Miocene and much of the Oligocene Tampa and Suwannee. Westward, in the Trough, it is restricted to the upper portion of the Lower Miocene. Montmorillonite is the predominant clay in the Suwannee limestone. Kaolinite is abundant in the central panhandle area. The kaolinite was presumably supplied by a different, relatively small, source area. Pure palygorskite is present in the dolomitic upper portion of the Tampa as far west as the central panhandle (Fig. 14). In the Trough area and slightly t o the east palygorskite and sepiolite are dominant in the upper portion of the Lower Miocene Torreya-Hawthorn. A second less-pure interval occurs near the middle of the Lower Miocene. This could be the top of the Tampa. South and slightly west of Tallahassee (Sopchoppy) the Sopchoppy limestone (Chipola or older), which was deposited in a restricted-marine environment (Huddlestun, 1976), contains palygorskite and trace amounts of montmorillonite. The chain clays and the Torreya Formation pinch out near the Appalachicola River. West of the river the section is largely Middle Miocene Chipola and Shoal River. Montmorillonite is the predominant ciay but illite, and t o a lesser extent kaolinite, become progressively more abundant westward. The lateral change shows the influence of a western, possibly Mississippi River, source. South of the line of the cross-section and near the center of the Appalachicola Embayment (W-7574, St. George Island) the Torreya is absent (Banks and Hunter, 1973). The bottom 6 m of the core consist of fossiliferous limestone which contains only palygorskite and may be Tampa. The clays in the overlying marine ChiPola(?), Middle Miocene and Pliocene, are predominantly montmorillonite but with varying amounts of kaolinite and illite. Palygorskite and sepiolite, apparently detrital, are relatively abundant in the lower part of the section. The stratigraphic correlations indicate that near the end of the Early Miocene there was a significant uplift to the east (Ocala High) and subsidence to the west, with the hinge area being near the Appalachicola River. The commercial palygorskitesepiolite clay beds were formed in the eastern area near the close of the Early Miocene during the final stages of the marine

43 Gooda I

Fig. 14. East-west crosssection extending from western Florida panhandle to 40 km east of Tallahassee on the western flank of the Ocala High, Blank areas indicate no claymineral data was obtained. Horizontal-stripe pattern = montmorillonite, stippled pattern = palygorskite.

regression. If the Torreya (restricted marine) is the time equivalent of the Chipola (open marine) the differential movement would have occurred slightly later (early Middle Miocene). Farther to the east, extending on out t o the Blake Plateau (Weaver, 1968), the vertical distribution of clay minerals is similar to that in the area between Tallahassee and the Appalachicola River. In the offshore wells, in the Florida panhandle, and in the Savannah River area, all of which have faunal data, sediments rich in palygorskite and sepiolite are restricted t o the Lower Miocene and Upper Oligocene (Tampa equivalent). It seems reasonable t o conclude that between the faunal control areas indigenous palygorskite- and sepiolite-rich beds are restricted to this time interval. Overlying and underlying sediments are montmorillonitic but in some places the overlying sediments contain detrital palygorskite and sepiolite.

44 ECHOLS COUNTY, GEORGIA

A series of core holes slightly north of the Georgia-Florida border provides information on the lateral facies change and tectonic movement on the east flank of the Ocala High. An east-west cross-section extending from eastern Lowndes through Echols County is shown in Fig. 15. In Echols County the upper portion of the section consists of 21-31 m of sand which thickens to the east. It has been considered to be post-Miocene, but a portion of it may be Upper Miocene. The top of the Suwannee does not appear to have been definitely penetrated, but several of the wells are probably within 6 m of the Suwannee on the basis of nearby deep wells. Echols-2 (Fig. 15) appears t o have penetrated the most complete Miocene section. The lower 3 m are a pebbly white limestone rich in palygorskite. This is probably near the base of the Lower Miocene. This interval is overlain by two clay beds. The lower is approximately 9 m and the upper 6 m thick, separated by approximately 1.5 m of dolomite. Angular clay and dolomite clasts are common in both beds; however, the two beds are distinct. Sepiolite is commonly the dominant (30-50%) clay in the lower bed, along with varying amounts of palygorskite and montmorillonite. Calcite shell fragments and phosphate grains, are present. The upper bed contains no sepiolite or shell fragments, relatively little phosphate, and is composed predominantly (commonly over 90%) of palygorskite. Opal-cristobalite is abundant and sponge spicules and diatoms are present. Clay-wise these two beds are similar t o those mined to the west on the western flank of the Ocala High. Both clay beds show evidence of considerable reworking and resemble a clay conglomerate with a clay matrix. The dolomite, diatom, and spongespicule content increase toward the east and northeast (seaward). Both beds were apparently deposited in coastal environments that were periodically subjected t o tidal invasion. The lower bed was probably formed landward of the upper bed. The top of the lower bed can be traced east as far as EC-7 and west to EC-1. The upper bed is penetrated from EC-1 to EC-7. It becomes more sandy and montmorillonitic to the west. Since the Oligocene Suwannee outcrops at the Withlacoochee River (boundary between Lowndes and Brooks Counties) is seems unlikely that the clay beds were continuous across the Ocala High. On the east flank of the high (east of EC-3) the upper clay bed is overlain by a highly mixed unit consisting of clay, dolomite, opal-cristobalite clasts and phosphatic sand. The montmorillonite content is relatively high in this facies (Fig. 15). This interval thickens to the east and was presumably largely derived from locally reworked material and material derived from farther up on the Ocala High. To the west this interval appears t o be represented by a poorly sorted sandy interval which contains clay clasts, white dolomite veins, leached white phosphate, montmorillonite and minor weathered palygorskite. It

LO 5

.ygorskite, Fig. 15. East-west crosssection of the eastern flank of the Ocala High. Light stipple pattern indicates sepiolite with palygorskite, darker stipples palygorskite with no or minor sepiolite; horizontal pattern represents montmorillonite; dense-gray represents kaolinite. kaolinite. 0-C equals opal-cristobalite.

CI1

46

appears to be a weathered zone which in part is the equivalent of the submarine to tidal reworked zone t o the east. In the western portion the weathered zone (-3 m thick) is overlain by approximately 6 m of montmorillonitic fine sandy clay and clayey sand. The reworked interval t o the east is overlain by similar sediments. These probably represent a minor marine transgression over the weathered-reworked (tidal) surface. The interval from the top of the lower clay t o the top of the upper weathered interval has a relatively constant thickness. However, in EC-2 to the south of the east-west line of section the interval is thinner and to the north EC-12 the interval is thicker. Thus, there is also a n o r t h s o u t h component t o the regional tilt, as would be expected, with most of the variation apparently coming above the upper clay bed. This same general interval contains abundant opal-cristobalite, some sponge spicules, and a few diatoms. The diatoms consist largely of the bottom-dwelling species Mellosiru suleutu which commonly lives in very shallow water. NaOH solubility studies show that from 25 t o 55% of the opal-cristobalite-rich material is readily soluble. During periods of high pH silica is taken into solution and later precipitated at periods of low pJ3. The presence of opal-cristobalite intraclasts indicates that solution and precipitation occurred before burial. The top and bottom of the interval containing opal-cristobalite (X-ray) appear t o be correlatable time horizons. The base coincides with the base of the upper clay bed and the top coincides with the top of the weathered zone t o the east and is slightly lower in the west. It is missing from the lower weathered horizon in the western part of the section. If the opal-cristobalite forms primarily from silicious ,organisms it should form a reasonable, for the Miocene, time marker. In the east the disappearance of the silicious organisms coincides or is near the top of the marine sand and clay zone and the base of a tan-dolomite facies. In the tan dolomite palygorskite ranges from 5 t o 75%. Dolomite intraclasts are abundant and clay clasts relatively common. The dolomite grains vary from spherical to subrhombic. At least some of the palygorskite appears t o have grown in place (see SEM discussion). The depositional environment is probably tidal to schizohaline. The clay suite suggests the dolomite is Early Miocene in age. To the west the equivalent interval appears to be a 3-m thick weathered (clay cutans and dolomite veins) montmorillonitic sand. The montmorillonite is highly disordered and appears t o have been weathered. In LO-5 this interval is represented by a coarse-grained sand. A shore line appears t o have existed in the vicinity of EC-1 and EC-2. Brooks (1966) described 9 m of fossiliferous beach and shallow-marine sediments outcropping in the Alapaha River at Statenville, approximately 7 km west of EC-3. He considered this t o be Late Miocene. The section rests

on dolomite. The cross-section indicates that this is probably an extension of the tan dolomite. The clay mineralogy (75% palygorskite, 25% montmorillonite) can be used t o trace this interval into EC-3 where the interval contains more clay and has worm burrows. Eastward the equivalent interval is apparently a gray phosphatic sand. A suggested upper boundary for the Upper Miocene is based on a change from Na- and K-feldspar to only K-feldspar and the presence of minor amounts of sepiolite and palygorskite in the lower portion of the sand and its absence in the upper portion. The sand is apparently shallow open marine. The presumed Upper Miocene equivalent t o the west of the outcrop represents a more continental sequence and consists of montmorillonitic clayey sand with irregular patches and veins of dolomite. In LO-5 organic-rich mud (marsh?) and an upper clayey sand with clay cutans (soil) are present. Thus, the Upper Miocene shows an ideal environmental sequence extending from soil-marsh through lagoon-beach t o shallow marine. Phosphate grains (mostly black) are common in the lower clay bed, but scarce in the upper. Phosphate and sepiolite commonly occur together. Phosphate grains and pebbles extend from the base of the Lower Miocene to near the surface but are relatively rare in the two clay beds. Much of the shallow phosphate is leached to various shades of tan, gray and white. The phosphate content, in the upper sandy interval (Upper Miocene-Pliocene), increases from east to west, though the trend may be more northeastsouthwest as the shore line appears t o have changed from a general n o r t h s o u t h direction to a NW-SE direction in Echols County. Reworked phosphate appears to have been concentrated in the littoral environment flanking the eastern edge of the Ocala High. High concentrations (commercial t o near-commercial) are present in the upper sand interval and conglomeratic sandy intervals in the tan dolomite contain up to 50% phosphate grains. The Occidental phosphate mine occurs approximately 38 km southeast of the Statenville outcrop. The palygorskite-rich pebbly phosphate sequence (9 m thick) lies on top of a tan-dolomite sequence (FGS well). According to Brooks (1966) mammalian teeth in the commercial zone are of Pliocene age. Nearby, at White Springs, the Pliocene phosphatic sands rest on top of fossiliferous dolomatized sands of Chipola age (late Early t o early Middle Miocene) (Brooks, 1966). It appears that the Upper Miocene phosphatic interval in Echols County pinches out t o the south where a phosphatic Pliocene sand rests directly on the Lower Miocene. The increased reworking from Late Miocene t o Pliocene time could account for the higher phosphate content in the section in the Occidental Mine. In this general area both the Upper Miocene and Pliocene consist of beach and shallow-marine sands with a relatively high content of phosphate. The phosphate concentrate was derived from reworking and winnowing of the older Miocene, Oligocene, and Eocene sediments outcropping on the fringes of the Ocala High. There appears t o have been little deposition during the

48

Middle Miocene. It was probably during this time that much of the finegrained phosphate was dissolved and reprecipitated to form the larger particles which were reworked and concentrated by the Upper Miocene and Pliocene seas. The occurrence of dolomite underlying the phosphate-rich sands suggests that the carbonates provide the high-pH conditions that caused the precipitation of phosphate from the overlying acidic pore water. Structural cross-sections using various horizons as horizontal data give some idea of the movement of the Ocala High. A cross-section using the top of the opal-cristobalite zone as a horizontal datum indicates there was little relative movement during deposition of the two clay beds and only a relatively mild differential movement (west up; east down) between deposition of the upper clay bed and the start of deposition of the tan dolomite. During deposition of the dolomite and Late Miocene the movement was accented. A southwest-northeast section indicates that the relative downwarping to the northeast during the Early Miocene was probably slightly larger than that t o the east. In addition there appears t o be more of a facies change in this direction. The interval between the top of the upper clay bed (palygorskite) and the base of the tan dolomite consists largely of palygorskite-rich (generally greater than 90%) mixed clastic clay, sand and solomite in the northeastern well (EC-12). REGIONAL CROSS-SECTIONS

An attempt was made to extend the Echols County cross-section eastward on to the continental shelf and west t o the Appalachicola River (Fig. 16). The tan dolomite can be traced t o the inner edge of the shelf. The interval between the base of the tan dolomite and the top of the Oligocene Suwannee and Eocene Ocala carbonates is rich in palygorskite but the two clay beds could not be traced using cutting samples. The lower portion of the Miocene section of JOIDES-1 is similar t o the lower portion of the Echols County Miocene. A lower sepiolite-containing interval is overlain by a palygorskite-rich interval. This is overlain by a 18-m section containing a clay suite with significant amounts of palygorskite, sepiolite and montmorillonite. This may be equivalent t o the reworked interval and/or the tan dolomite. The lower portion of the Miocene also contains diatoms (and radiolarians) as does the section in Echols County. It is not clear whether the palygorskitesepiolite on the continental shelf is detrital or the shelf was covered with very shallow water. In JOIDES-2 (on the outer edge of the shelf) Upper Miocene rests on Lower Miocene (coarse sand) with the Middle Miocene missing. The boundary is marked by an abrupt increase in the montmorillonite content of the clay suite; minor illite is present in most samples. At approximately the same interval there is a similar change in the clay suite in JOIDES-1. This suggests a similar stratigraphic sequence. It is difficult to determine the age of the tan dolomite. There appears t o

Fig. 16. East-west crosssection extending from Jim Woodruff Dam (Appalachicola River) across the Ocala High on to the continental shelf. Clay beds are indicated. The upper bed (Paly.) contains only palygorskite. The lower bed ( S e p . )contains both palygorskite and sepiolite.

Ip

W

50

be a hiatus preceding its deposition suggesting it may not be part of the Lower Miocene. It could represent the Middle Miocene but the absence of Middle Miocene on the continental shelf and to the west of Echols County makes this suspect. However, the area between Lowndes County and the shelf could have existed as a shallow semi-closed basin during the Middle Miocene. Dolomite would be expected to form in such an environment. It probably is Chipola in age (late Early to early Late Miocene). In Echols County phosphate pebbles occur at the base of the sand overlying the tan dolomite. On the continental shelf (JD-1) phosphate pebbles occur at the very base of the Upper Miocene and about 6 m above the Oligocene-Lower Miocene contact. In the Trough phosphate pebbles are concentrated at the Lower Miocene-Middle Miocene boundary and also at the Miocene-Suwannee boundary. Thus, phosphate pebbles commonly characterize the boundaries of the major Miocene time units. Phosphate and quartz pebbles occur throughout much of the Miocene section in the Atlantic Embayment. To the north of Echols County and on the east flank of the Ocala High an opal-cristobalite ( 0 4 ) zone, with a few diatoms and sponge spicules in some wells, occurs in the Lower Miocene. This is probably correlative with the O-C zone in Echols County. The palygorskite and ,sepiolite distribution is also similar, with sepiolite being restricted t o the lower section below the O-C zone. Some sepiolite also occurs on the top of the 0% zone as in the reworked interval in EC-12. In these northern wells palygorskite extends up to the base of the Middle Miocene (Shoal River) diatomsponge-spicule zone. In Echols County it extends to the top of the tan dolomite. This further suggests the tan dolomite is of Chipola age.’The upper portion of the Lower Miocene (and Chipola) in these northern wells consists of fossiliferous sands and dolomite with phosphate pebbles. This appears t o be a seaward equivalent of the tan dolomite. The Miocene section thickens abruptly a few kilometers northeast of Echols County. A ‘cherty’ zone, in some wells containing a few sponge spicules and diatoms, is probably correlative to the O- C zone in Echols County. The lower part of the section contains intervals that are montmorillonite-rich as well as palygorskitesepiolite-rich. This, along with the presence of shells indicates a more normal-marine environment. The various eastern facies are difficult to trace across the crest of the Ocala High (Withlahoochee River). The uppermost Oligocene (Tampa) occurs west of the arch but appears to be absent or very thin east of the arch (Fig. 16). It thickens and becomes more sandy westward (Chattahoochee facies). Varying amounts of palygorskite are present in these sediments with the more marine (St. Marks facies) containing more montmorillonite. Whether this interval actually thickens in the Trough area cannot be demonstrated from the well data along this line of section. The Tampa outcrops in southwest Georgia, west of the Trough. At Fall

51

Cave (NW Grady County) and Climax Cave (eastern Decatur County) the Tampa consists of rounded limestone pebbles (minor dolomite) in a clay matrix. The lower Tampa and Suwannee contain montmorillonite and perhaps traces of palygorskite. Near the top of the Tampa only palygorskite is present. The limestone grades into an overlying clay bed approximately 2 m thick. The clay bed (Lower Miocene Torreya or Hawthorn) contains whitish clay clasts and layers in a gray-green clay matrix. The whitish clay is essentially pure palygorskite and the gray-green clay contains palygorskite and minor montmorillonite. Both the Tampa limestone and the overlying clay contain abundant palygorskite clay clasts indicating shallow-water reworking. The clay bed is overlain by a montmorillonitic sand (Hawthorn) containing a few palygorskite clay grains. Farther west at the southwestern corner of Georgia (Jim Woodruff Dam) the Chattahoochee formation is more clastic containing sandy limestones (some fossils) and clayey sands with a few dolomite beds. Pure palygorskite occurs in dolomite beds in the upper part of the formation (lower part was not exposed). Microscopic examination of the residue indicates it is present as silt-size clay grains and presumably detrital in origin. In the overlying marine Torreya Formation montmorillonite is the predominant clay, though most samples contain small amounts of palygorskite. Nearly pure palygorskite clay clasts (up to a foot in diameter) occur in a dolomitic limestone bed. In several Torreya dolomite beds montmorillonite is the predominant clay. The Lower Miocene clays reflect lateral environmental changes. Montmorillonite is dominant on the continental-deltaic Ocala High area, sepiolite is dominant at the edge of the Trough. In the Trough and to the west palygorskite is abundant in tidal-lagoonal sediments (including the commercial clay beds). Farther west in the more marine facies montmorillonite again becomes dominant. In the line of section the Trough was not a trough but a shallow embayed area. Both east and west of the Suwannee High, Upper Miocene is present (continental to delta on the west and shallow marine t o the east). Montmorillonite is the dominant clay mineral (kaolinite in some of the sands). Fig. 1 7 contains a sequence of cross-sections illustrating the structural evolution along the Georgia-Florida border. While the Suwannee limestone was deposited in much of the area a high existed in the eastern portion (Charleston County) of the area. Maximum deposition probably occurred in the Trough area though this is not certain. Uplift continued and migrated westward during the latest Oligocene (Tampa). Most of the area east of Brooks County was probably above water level. The sea t o the west was very shallow and much of it tidal. It is not known whether the Trough (near the FloridaGeorgia border) was present. During the Early Miocene the Florida panhandle t o the west of the crosssection was elevated and few if any sediments were deposited. Shallowmarine and tidal-flat palygorskite-rich sediments were deposited in a narrow

52

FLA

IGA

A\ ’

GA

EN0

’

ATLANTIC SUWANNEE

SEA LEVEL

SUWANNEE

.EVEL

SEA

END

LOWER

LEVEL

MIOCENE

SEA L E V E L UPPER M I O C E N E

E N D UPPER

MIOCENE

Fig. 1 7 . Eastwest cross-sections paralleling and near the Georgia-Florida border. Same sections as Fig. 16. Structural high migrates from east to west. Movement appears to be due to tilting. Only minor, if any marine deposition occurred during the Middle Miocene. The tan dolomite may be Middle Miocene. C.B. indicates palygorskite clay beds.

53

basin west of the Suwannee High and on a broad shelf east of the High. Montmorillonitic continental sediments were deposited on the High. During the Middle Miocene the western portion of the Florida panhandle was invaded by the sea. Most of the area to the east was above sea level except for the Georgia portion of the Trough. During the Late Miocene downwarping or tilting was renewed in the eastern area and/or the western area was elevated. This possibly is the result of a mild uplift of the southern Appalachian Mountains. There was a fair amount of structural activity during the Oligocene and Miocene. The accretion of the drifting Florida block to the main continent may not have been completed till near the end of the Middle Miocene. A southwest-northeast cross-section was constructed (Fig. 18) extending down the axis of the Ocala High. Immediately t o the east and west of this high the Miocene section thickens but is still relatively thin. The Miocene clay and sand (very fine) on top of the Suwannee High is iron-stained and contains angular weathered fragments of white limestone, phosphate and chert. The clay suite is montmorillonite and kaolinite. The kaolinite is higher than in the post-Miocene sediments. This material appears to be a leached residue from the Suwannee. The sands and clays off the flanks are light gray and contain fragments of white limestone. This mpterial appears t o have been deposited while the soil was developing on the high areas. Palygorskite and particularly sepiolite are abundant in these sediments. In three core holes in this area (BR-1, LO-6, BE-1) the first occurrence of palygorskite and sepiolite occurs at 49 m above sea level and at approximately the same horizon as the phosphate pebbles and the bottom of the residual soil zone. The upper portion of this consists of clay clasts and vertical fractures with clay skins suggesting a mild weathering and reworking. The phosphate pebbles, residual soils, and end of palygorskitesepiolite deposition (with weathering and reworking) indicate a significant regional hiatus. On the extreme northeast flank of the Atlantic Embayment (EF-3) a soil forms the boundary between the Middle Miocene Shoal River and the Lower Miocene palygorskitesepiolite deposits. The montmorillonitic marine Middle Miocene contains an abundance of diatoms and sponge spicules and can be traced southwest t o the northern edge of the Ocala High and northeast into South Carolina (Coosawatchie Clay; Abbott, 1974). As in the Trough area the lower massive limestone (Tampa and/or Lower Miocene) extends northeast past the edge of the Ocala High but ends south of the present Altamaha River (GGS-50). The boundary of the northern thick edge of this limestone extends northwestsoutheast, becoming progressively farther removed from the river towards the southeast, though a thinner limestone bed extends t o near the edge of the river. To the northeast this lower interval consists of coarse sand t o pebble phosphate and quartz, shell fragments, and thin beds of dolomite. The environment was presumably highenergy shallow-marine t o estuarine. The carbonates and the fringing clastic sediments commonly have a mont-

3N3301W lSOd

IZC s99

M 3N3301W Y 3 M O l S3lll31dS 30

Fig. 18. Southwest-northeast crosssection extends from west flank of Ocala High (near Tallahassee, Florida) down the nose of the High, across the Atlantic Embayment to South Carolina. Lower blank area is Suwannee limestone containing montmorillonite. The upper blank area is post-Miocene sand with a relatively high kaolinite content. Middle Miocene sediments contain diatoms and sponge spicules (D). It is absent south of the northeast edge of the High. Horizontal pattern = rnontrnorillonite; stipple pattern = palygorskitesepiolite.

55

morillonite-rich clay suite. However, palygorskite is predominant in GGS-425 on the north edge of the Ocala High. The carbonate in this well is largely tan dolomite, and it appears likely that the dolomite and palygorskite formed in a lagoon-like environment between the Ocala High area t o the south and a carbonate barrier fringing the Atlantic Embayment. After deposition of these lower sediments the environment was largely deltaic-estuarine. The Lower Miocene opal-cristobalite-diatom-spongespicule zone found t o the west in the Trough area can be traced over much of the Atlantic Embayment. This is believed to be equivalent t o the opalcristobalite zone in Echols County. North of the Altamaha River, sand is much more abundant and coarser than t o the south. Phosphate and quartz pebbles are common throughout the section. This presumably reflects the presence of highenergy marine conditions between the northern extension of the Ocala High and the shallow-marine area in northeast Georgia. Farther west, palygorskite is the predominant clay mineral in the upper portion of the Lower Miocene except immediately north of the Altamaha River. The palygorskite on the southwest and northeast flanks of the Embayment, where dolomitic clay is abundant, is probably authigenic and that in the center detrital. The palygorskitesepiolite-rich clay suite, with varying amounts of montmorillonite, extends into South Carolina where it has a maximum thickness of about 33 m (Heron and Johnson, 1966); it is also associated with dolomite and minor clinoptilolite. On the western portion of a cross-section extending through the Atlantic Embayment and out onto the shelf (see Fig. 19) Lower Miocene and probably Upper Miocene (continental) sediments are present. On the shelf only Lower and Upper Miocene is present. Correlation between these two areas is highly speculative. These wells contain a basal montmorillonitic carbonate and coarse-clastic section overlain by gray dolomitic clay and clayey dolomite and coarse sand t o pebble-size quartz and phosphate. The dolomite interval contains opal-cristobalite as well as a few diatoms and sponge spicules and has a relatively high content of palygorskite and sepiolite. The two units are mineralogically and lithologically similar t o the Lower Miocene section on the north flank of the Embayment. The overlying, more clastic, pebbly interval could be the eastern equivalent of the more open-marine Middle Miocene diatomsponge-spicule zone to the west, with nondeposition occurring on the shelf. Alternately, this interval could be the strand-line equivalent of the Upper Miocene on the continental shelf. A linear deposit of coarse sand and gravel (Trail-Lake Wales Ridge) extends nearly 330 km along the eastern side of the Florida peninsula. This can be traced northward into southeast Georgia. Many speculations have been made concerning the origin of this material but that proposed by Brooks (1966) and Alt (1974) best fits the data in the present study. They suggest the coarse sand was deposited as a marine barrier bar or spit, probably during the Late Miocene and earlier, and that it becomes progressively younger t o the south. The coarse sands and gravels, in the upper part of the

56 GGS 273

PLIOCENE UPPER MIOCENE

LOWER MIOCENE

-

OLIGOCENE

1 3 0M

24KM

Fig. 19. Northwest-outheast cross-section extending from Dodge County, through the Atlantic Embayment and out to the continental shelf. M.M. = Middle Miocene; U.M. = Upper Miocene; D = Diatoms and sponge spicules.

section, along the Georgia Coast (see Fig. 19) are probably a northern extension of these deposits and may range in age from Middle to Late Miocene. The Trail-Lake Wales Ridge sands and gravels are probably a longshore extension of the coarse sands and gravels deposited in the northwest-trending estuarine facies in the center of the Georgia Embayment. Fig. 20 shows a cross-section extending from the Effingham-Screven County boundary through Chatham County to the Atlantic Coast, paralleling the northeast flank of the Embayment. The Miocene section is an offlap sequence. The mineralogic and lithologic correlations are similar to the paleontological correlations of Huddlestun (1973).Updip the lower half of the section consists of montmorillonitic, marine fossiliferous sands and clays. Phosphate, quartz, and clay pebbles are abundant at the base and at several other intervals. Within this interval is a correlative zone containing clinoptilolite and biotite (in addition to montmorilionite). The upper one-half to onethird of the montmorillonitic interval, overlying the zeolite zone, contains minor palygorskite and sepiolite and is apparently a transitional interval. The montmorillonite section is earliest Miocene in age (Huddlestun, 1973) and possibly equivalent to the Torreya. Phosphate and quartz pebbles occur at the base of the Miocene and occur unconformably on the Oligocene. The section appears to represent the earliest Miocene marine transgression. This

57

Fig. 20. Northwestsoutheast cross-section along the Savannah River extends from northern Effingham County t o near the coast. The Miocene shows an off-lap sequence. The uppermost sand unit is post-Miocene. The top of the Oligocene in the easternmost well may be at the top of the coquina or the top of the underlying massive limestone. Light stipple zone contains sepiolite; dark stipple palygorskite. Eastern two wells are not plotted t o scale. GGS-394 is 26 km east of EFF-3 and GGS-717 is 19 km east of GGS394.

is one of the few areas in the southeast where appreciable clinoptilolite was encountered. Farther updip the basal Miocene (Aquitanian) has a restricted fauna (Huddlestun, personal communication, 1975) and contains palygorskite rather than montmorillonite. The oldest Miocene polygorskite facies lies in a northeastsouthwest band along the western flank of the Trough. The montmorillonite is overlain by a palygorskite-rich interval, which in turn is overlain by a sepiolite-rich interval. Huddlestun says the limited fauna in these sediments is indicative of a restricted, fresher than normal marine,

58

environment. Evidence indicates the more southern chain clays were also deposited in a brackish-water (schizohaline) environment. The interval between the montmorillonite facies and palygorskitesepiolite facies contains significant amounts of both clays, and presumably represents the transition from open-marine t o brackish-bay or estuarine conditions. The chain clays in the transition interval are presumably detrital and were derived from the more shoreward brackish-water environment. The interval containing the chain-structure clays is apparently Early Miocene in age (probably late Early Miocene) and represents the regression of the sea near the end of the Early Miocene. The concentration of sepiolite in the upper portion suggests it forms, in preference t o palygorskite, in a less marine, more landward environment - possibly alkaline lakes. The sepiolite content of this zone decreases seaward (southeast and southwest) which also suggests a relation to salinity. The phosphate (BPL) concentration is apparently controlled by the regression. In Eff-3 phosphate is at a maximum at the base of the Lower Miocene and systematically decreases upward as the interval becomes less marine (Fig. 21). The lowermost marine Miocene pinches out between Eff-6 and Eff-3 (Fig. 20) and apparently does not occur southeast of Eff-3. A montmorillonite interval occurs at the base of the Miocene over much of the Trough and Embayment and appears t o be equivalent to the lowermost Miocene in Eff-6 to Eff-10. A high area apparently existed in southeastern Effingham and Chatham Counties during this time and sediments were not deposited. The palygorskitesepiolite zone extends the length of the cross-section, though the sepiolite zone appears to be missing in the southeasternmost well and the sepiolite content is quite low in the adjacent well (GGS-394). This probably represents a more marine influence rather than a new time-rock unit. The sepiolite facies does not extend to the southwest into the Embayment. The upper Lower Miocene palygorskitesepiolite section has its maximum thickness between Eff-3 and GGS-717 (Savannah). This reflects the presence of the Beaufort High t o the east and the Ridgeland Basin (Heron and Johnson, 1966) between the seaward high and Eff-3 which is located on the western edge of the basin. A Lower Miocene section identical to that in GGS-394, but thicker (75 m), occurs due south in McIntosh County (GGS-88). This indicates the ‘high’ area in eastern Chatham County is not extensive. In Eff-3 the Lower Miocene palygorskite-sepiolite section is overlain by 24 m of montmorillonitic sediments and 6 m of kaolinitic sand (Fig. 21). The basal foot contains abundant phosphate sand and pebbles. This is overlain, at 28.5-30 m, by a 1.5-m thick organic-rich clayey silt which has the features of a paleosoil. This soil is overlain by a dark-gray sandy clay containing an abundance of phosphate grains and pebbles. This grades upward into 13 m of gray, slightly calcareous silty clay and approximately 6 m of

59 B PL

CLAY MINERALOGY 2

4

6

8

1

1

1

1

%

Fig. 21. Core hole Effingham No. 3 (see Fig. 20) showing lithology, mineralogy and phosphate content. Bone Phosphate of Lime (BPL)data from Furcron (1967). Z = illite; B = biotite; K = kaolinite; M = montmorillonite; P = palygorskite; and S = sepiolite.

clayey sand with a quartz-pebble zone near the middle. These sediments contain Middle Miocene (Shoal River) foraminifera (Huddlestun, 1973).Biotite is relatively abundant throughout this interval. The heavy-mineral suite indicates the biotite is metamorphic rather than volcanic. Diatoms and sponge spicules occur throughout most of the dolomitic-clay interval (16-24 m). Opal-cristobalite occurs above and below this interval indicating some solution of the silicious organisms during the initial transgressive phase and near the final phase of the following regression. The diatoms are similar to those in the Middle Miocene Calvert Formation of Maryland (Abbott, 1974). The diatoms can be traced to the northeast and southwest but do not

60

occur in wells t o the northwest and southeast. The distribution of diatoms and sponge spicules suggests they were deposited in an elongated trough or gulf open t o the Atlantic Ocean to the north (Fig. 22). Though the eastern boundary could be caused by deeper-marine conditions, the sediments do not suggest this. The more reasonable explanation is that this was an enclosed area, similar in many respects to the Gulf of California, though the eastern boundary was probably a shallow littoral zone where sand and phosphate were reworked but deposition was minor. Calvert (1966) has calculated that diatoms comprise an average of 21% by weight (over 50% in some areas) of the silty clays in the Gulf of California. Some of the Si is supplied by the rivers but most, along with necessary phosphates and nitrates, is supplied by upwellings from the deep ocean t o the euphotic zone where it is extracted. In the Eocene, diatoms are abundant in the Alabama Gulf Coast (Wise and Weaver, 1973) and in Georgia. However, in the Miocene they are apparently restricted t o eastern Georgia, eastern Florida, and the continental shelf and

Thickness of Middle Miocene

Diatoms

Sponge Spicule Zone Conlour IblerYol in Feet 100'. 30.5M

Fig. 22. Map showing distribution and thickness of Middle Miocene diatoms and sponge spicules. Facies extends eastward through South Carolina and connects with the Atlantic Ocean.

61

are absent or rare in the Florida panhandle. The distribution patterns are probably related t o the change from east- to west-flowing currents to northward-flowing currents caused by the closing of the Thetys. A period of nondeposition, weathering, and mild erosion occurred sometime towards the end of the Early Miocene and during the beginning of the Middle Miocene (Chipola time?). This hiatus is extensive and can be traced over much of the Georgia Coastal Plain. Correlation above the Lower Miocene is difficult in the Chatham County area where commercial phosphate deposits occur. The relative distribution of phosphate in the Lower Miocene of the commercial area (slightly east of Savannah) is similar t o that in Eff-3. The absolute values (except for the ore) are also similar (Figs. 23 and 21). The phosphate-ore zone lays on top of the palygorskite-rich Lower Miocene in the same position as the Middle Miocene soil horizon in Eff-3. Bone Phosphate of Lime (BPL) values in the ore zone commonly range from 20 t o 40% as opposed to 3 4 % in the soil in Eff-3.

DEPTH I N METERS

8

I

1

I I

-

N

l

l

I

L I

N

I

I

I

I

I

I

I

0 I

0

05 N <

0

E

$ 2

m

s:

r

m 0 00

E* 0

ae

0 - N

W

T i r

--p

-m -m

-0 W D P N N W W O N !B W 0

$

w o o w 0 k

Fig. 23. Composite section from area containing commercial phosphate, east of Savannah. BPL data from Furlow (1969).

62

Palygorskite is abundant in the lower portion of the ore and montmorillonite in the upper portion. The basal phosphate was apparently locally derived by reworking the underlying clays and the upper was transported by marine currents. In Eff-3 the BPL values systematically decrease upward in the marine Shoal River (Middle Miocene) reaching a minimum near the top of the diatomaceous marine clay bed. The overlying 6 m of sand and clayey sand shows an increase in BPL. This sand is distinctive in that it contains a minor amount of clinoptilolite. A considerable portion of the phosphate consists of shell fragments rather than rounded grains. It is not clear whether this upper phosphate sand is part of the Middle Miocene or is younger. It rests directly on top of the diatom zone. In the Middle Miocene of North Carolina a three-foot diatomaceous clay bed occurs near the middle of the 12-m commercial phosphate unit (Gibson, 1967). Thus, the 21-m Middle Miocene sequence in Eff-3 is similar t o that in North Carolina except it contains less phosphate and zeolite. In North Carolina the phosphorite bed rests on the Middle Eocene and in Georgia on the Lower Miocene. If the time of origin and concentration of phosphate is related t o periods of nondeposition, as it appears t o be in many places, then the North Carolina material had a much longer time t o form and be concentrated. The transgression by the Middle Miocene sea was apparently restricted to an elongated trough or gulf (Fig. 22). The age of the sediments overlying the Lower Miocene to the southeast of Eff-3 is difficult t o establish. As diatoms and sponge spicules are not found in GGS-394 and zeolite occurs near the base of the montmorillonite zone it appears that the upper portion of the Middle Miocene in Eff-3 correlates with the basal dolomitic sand and clay unit in GGS-394 (Fig. 20). The phosphate-ore zone (clayey sand and sandy clay) to the east may represent the period of time during which the soil interval in Eff-3 was formed but it seems likely it represents the entire Middle and Upper Miocene and was never more than a few meters above or below sea level. A comparison of Figs. 21 and 23 shows the Middle Miocene diatomaceous marine clay in Eff-3 is missing from the well containing the commercial ore. The phosphate distribution in Eff-3 suggests the two phosphate zones on either side of the marine clay converge in the area of the composite well. It further indicates that in the Middle Miocene, as marine sediments were being deposited in the trough area, continued reworking and concentration of phosphate occurred seaward in the Savannah area. In summary the sediments in the cross-section (Fig. 20) reflect a Lower Miocene marine transgression over Oligocene carbonate rocks. The presence of Jacksonian rocks (Eff-6) flanked by Vicksburgian rocks (Huddlestun, 1973) suggests a high existed in this area. The transgression was followed by a regression and the development of a thin continental facies during the end of the Early Miocene and the beginning of the Middle Miocene. The area was again invaded by the sea during

63

the Middle Miocene though the axis of the depositional trough moved to the southeast. The seaward flank of the trough apparently remained near sea level during much of the Middle and Late Miocene and the depositional ‘basin’ migrated farther to the southeast as the western flank was mildly uplifted. TROUGH

Fig. 13 is a lithologic cross-section extending from the Appalachicola Embayment (St. George Island), through the center of the Trough t o the Savannah River. Fig. 24 shows the distribution of the clay minerals. In the southernmost well there is a relatively thick section of Middle Miocene (Huddlestun, 1976) which is absent in the northern wells. Palygorskite is the only clay in the bottom 3-5 m (dolomite) of the core. This interval is presumably Torreya or Tampa. The overlying shelly limestone contains varying amounts of palygorskite, kaolinite, and illite in addition to montmorillonite. From Lake Talquin (W-6890) north the lithologic correlation appears to be relatively good. Limestone and dolomite are predominant in the lower half of the section till near the middle of the Trough. The upper portion is generally white and the lower portion various shades of cream to light gray t o light tan. Shell fragments increase t o the south. The lower portion appears t o be equivalent t o the Tampa and the upper portion to the Lower Miocene Torreya. Palygorskite with varying amounts of sepiolite is the dominant clay in most of this lower interval. However, there is considerable variation suggesting there were major environmental differences in this area. Northeast of Cook County (GGS-25) the basal limestone thins and becomes more gray in color. This grades into a shelly sandy facies containing quartz and phosphate pebbles and gravel, The clay in both facies is predominantly montmorillonite with only minor amounts of palygorskite. These sediments were apparently deposited under shallow-marine conditions as opposed to the tidal-lagoonal environment t o the southwest. They are apparently the time equivalent of the lowermost Lower Miocene sediments identified in northern Effingham County (Fig. 20). GGS-159 (Berrien County) contains no palygorskite in the lower limestone unit. This well is relatively high up on the Ocala High and the limestone is either Suwannee in age or part of the more marine northern facies. Following deposition of the shallow-water Tampa deposits, the area was invaded by the sea during the Early Miocene (Torreya). A wedge of marine sands and fossiliferous limestone sediments was deposited in the southwestern portion of the Trough. Montmorillonite predominates (Fig. 24). Shells become more abundant t o the southwest and in W-6890 much of the section consists of sandy coquina. Further north montmorillonitic clays and dolomite become predominant. The commercial palygorskite clay beds were deposited during the regression near the end of the Early Miocene. Phos-

Q,

4

1-ZOKM

Fig. 24. Southwestqortheast crosssection extending from St. George Island through Embayment and "rough to Savannah River. Section shows distribution of clay minerals. Blank areas in lower part of section indicate no data. In upper part of section kaolinite is the predominant clay in sample from blank area.

65

phate is relatively sparse or absent southwest of central Colquitt County (GGS-175). In this area (GGS-175) the Torreya appears t o be represented by white limestone, though the limestone may be older. Palygorskite is abundant northeast of this area. This appears to be the area of transition between a southwest facies and a northeast Atlantic facies. GGS-159 is on the edge of the Ocala High and apparently contains no Lower Miocene. There may have been a physical barrier in this area separating the two bodies of water. The northeast facies consists largely of montmorillonitic sand and clay containing varying amounts of shell fragments, chert, quartz, feldspar, phosphate, and clay pebbles. The clay content increases toward the southeast. The depositional environment appears t o be nearshore marine to littoral with coarse clastics supplied by stream in the northwest portion of the area. Near the center of the northeast portion of the Trough (GGS-157) the Lower Miocene consists largely of dolomitic clay and sandy clay with thin beds of dolomite and few phosphate pebbles. There sediments were deposited in a lower-energy environment apparently seaward of the littoral zone or in a sheltered area. The sediments deposited during the regressive phase near the end of the Early Miocene are similar to those to the southwest in that palygorskite and sepiolite are the predominant clays. The sediments are more sandy and dolomitic than those to the southwest indicating different environmental conditions existed. The data suggest an extensive fairly shallow, brackish bay rather than a closed lagoon. This clays zone extends well out into the Trough and Embayment (and continental shelf) suggesting shallow-water conditions prevailed over much of the area, though the common occurrence of clay pebbles indicates much of this clay is detrital and was probably derived from the coastal areas. Fig. 25 is an isopach map of the Lower Miocene. The top of the Lower Miocene is based on the base of the Middle Miocene diatom zone, the top of the palygorskite zone, and the reworked pebble horizon. Contours were constructed including and excluding the thick basal limestone fringing the northern edge of the Ocala High. The latter interpretation is favored. The pattern shows thins in the area of the Ocala High and Ridgeland High (Savannah area). A relatively wide northwestsoutheast aligned thick area parallels the Altamaha River. Another narrow thick trough occurs along the west flank of the Ocala High. This trough terminates slightly southwest of the Georgia-Florida border. Fig. 26 shows, very generalized, the regional distribution of the major lithologic types. Wells near the boundaries commonly contain both types of sediments. The low-energy sediments are located on the east and north flank of the Ocala High and in the NE flank of the Atlantic Embayment. The coarse facies occurs between these two areas and extends seaward parallel to the strike of the Altamaha River, suggesting the river had a similar location but

66

Thickness of Lower Mibcene

20KM

Fig. 25. Isopach map of the Lower Miocene. Dashed 200-foot (61-m) contour indicates pattern if the basal limestone is included in the Lower Miocene.

did not extend as far to the east during the Early Miocene. The area was probably a high-energy estuary. There is a reentrant to the southwest which may indicate the location of the northward-flowing Suwannee River. There is a relatively abrupt change in lithology at the confluence of the Ocmulgee and Oconee Rivers. Cross-sections indicate that the basal limestone (36 m thick) ends abruptly in this area. This suggests the limestone may be Tampa equivalent and existed as an escarpment, controlling the shore line. This is also suggested by cross-sections farther to the west and to the east. GGS-236, in western Coffee County and GGS-25, northern Cook County (northwest flank of Ocala High), have a high clay content (palygorskite), but it is not dolomitic, and has a low phosphate content. The clay is white to cream colored rather than the light-gray color of the dolomitic clay. To the west, in the Trough, montmorillonite is predominant. Apparently no major streams entered this area and it was not readily available t o the open ocean. Palygorskite apparently formed in this restricted environment fringing the Ocala High. This general area and its southwestern extension was probably the source of the detrital Middle Miocene clays t o the southwest.

67

LOWER

MIOCENE

Fig. 26. Distribution of major lithologic units in the Lower Miocene. Mixed area to the southwest contains beds of palygorskite but overall lithology is complex. Palygorskite is major clay in dolomitic sediments.

Palygorskite along with sepiolite is the major clay in the Lower Miocene north of central Colquitt County, except in those wells near the Altamaha River (Fig. 24). Most of these clay minerals grew in place on the shallowwater (tidal-lagoonal) flanks of the basin and were carried into the higherenergy environment as pebbles, grains, and individual needles by waves and tidal currents. It is difficult t o be certain that palygorskite was not also formed in a normal marine environment. However, wherever it was possible to make detailed studies there was no indication that palygorskite formed in a normal marine environment. As much of the palygorskite was formed on the extensive tidal flats and tidal lagoons it is not surprising that much of it was eroded and transported seaward. Clay is by far the dominant lithology in the Middle Miocene diatom-

68

sponge-spicule interval and montmorillonite is the dominant clay mineral except in the southwest portion of the Trough (Fig. 24). In the Trough, extending from northern Cooke County t o southern Grady County, palygorskite and sepiolite are relatively abundant in the lower portion of the interval. Palygorskitesepiolite clay grains and pebbles, and phosphate grains are abundant; all of the clay is probably detrital. The distribution of clay facies in the area of transition between the northeast and southwest facies can be explained by taking into account the regional tectonics (Fig. 27). During Tampa deposition the Ocala High was above sea level (GGS-421 and 425). As uplift migrated from east to west, as demonstrated farther south, the western edge of the Ocala High was elevated (GGS-159) and the eastern edge depressed. Palygorskite was deposited in restricted environments on the flanks of the high. At the beginning of the Middle Miocene the area was downtilted to the east and/or north (or perhaps, block-faulted) and a normal marine environment invaded some of the previous high area. The Middle Miocene sediments, as defined by diatoms and sponge spic-

GGS25

1

GGS159

I

GGS421

I

BE-1

GGS 425

I

S L END OF TAW%

S L E N 0 OF -0WER M.OCENE

S.L. END MIDDLE MIOCENE

Fig. 27. Northwest-southeast cross-sections extending across Trough (left) and northeast end of Ocala High (right) showing three states of development. Blank areas indicate no clay-mineral data.

69

ules, ends fairly abruptly somewhere in the southern half of Grady County, suggesting a land barrier existed in this area. The palygorskitesepiolite clay beds t o the south, east, and west of the southwesternmost marine Middle Miocene tongue are quite different from those associated with this tongue and appear to have been deposited under different environmental conditions and at a different time (late Early Miocene). Fig. 28 shows the general distribution of the Lower Miocene chain clays. The two palygorskite beds in the Lake Talquin well (W-6890), apparently of Late Torreya age, can be traced northward into the commercial mining area of Florida (Midway, La Camelia, Chesebrough, Gunn Garm, etc.). Farther north only one clay bed is present. It contains no sepiolite and is presumably the upper bed or its equivalent. This bed can be seen in GGS-494 (southwestern Grady County), Block N mine (southeastern Decatur County), in Climax Cave (east central Decatur County), and Fall Cave (northwest Grady County). The clay beds have a relatively high content of palygorskite and sepiolite (80-100%) throughout much of the area occupied by the Trough and Embayment. These clay minerals are only moderately abundant in a northwestsoutheast band paralleling the Altamaha River. This was apparently an

Fig. 28. Map showing distribution of palygorskite and sepiolite in the Lower Miocene. Montmorillonite is the dominant clay in unlabeled areas. Dotted line indicates location of concentration of detrital palygorskite in Middle Miocene sediments.

70

estuary during the Early Miocene, and the clays are largely detrital. Sediments relatively rich in sepiolite appear to be most abundant along the flanks of the Ocala High and the northwestern flank of the Embayment. Montmorillonite is dominant on the Ocala High and all dong the northwest flank of the Trough and Embayment. Most of these latter montmorillonitic sediments are marine or near-marine and are probably mostly older than the chain-clay deposits. The sediment overlying the diatomsponge-spicule zone is largely sand and gravel And is probably the equivalent of a band of coarse, gravely sand that has been described by Olson (1967) in outcrop as the Ashburn Formation. At Ashburn, Georgia, the outcrop consists of coarse sand and gravel beds interlayered with beds of relatively pure clay (but containing coarse sand grains). Clay clast are present but cross-bedding was not observed. Kaolinite is the only clay present. Most of the kaolinite was probably formed by in-place weathering. Microscopic examination reveals the possible presence of a few highly etched sponge spicules and diatoms and some dolomite rhombs. The depositional environment appears t o have been littoral with appreciable tidal effect.

Chapter 4

MINES - LOWER MIOCENE INTRODUCTION

Thin, relatively pure palygorskite clay beds, 1-5 m thick, are restricted to the area slightly north of the Georgia-Florida border and south to the Lake Talquin area (see Fig. 4 and Map 1). The main commercial interval commonly contains two clay beds separated by a sand, shell, dolomite, or soil bed of variable thickness. These two clay beds (identified in the Lake Talquin well) appear to be relatively continuous throughout the north Florida area of the Trough. Thin, discontinuous beds occur above and below the main clay horizon but they generally contain abundant montmorillonite. LA CAMELIA MINE (MC-1)

A core (9 m) from the La Camelia palygorskite mipe (Engelhard Minerals and Chemical Corp.) in north Florida was studied in detail to determine the vertical variability and the relations of the various parameters. The structural, textural, mineralogical and chemical data indicate there are two major depositional cycles represented within the section. The sediments deposited during the two cycles differ in detail but in general are similar. The environments of deposition grade from shallow marine to lagoonal to tidal flat to soil. Fig. 29 is a generalized lithologic description of the core. There are two pure clay beds (0.5-1.5 m and 6.0-7.3 m) that consist of relatively pure, parallel-laminated clay. The clay beds were deposited during two periods of regression separated by a period oE
72 Meters Marine Cloy Burrowed

Logoonol

Tidal

Marine Sand

Soil

Supra Tidal

Lagoonal

Tidal Marine Sand

Fig. 29. Lithology of MC-1core from La Camelia Mine, Florida. Blank intervals contain relatively pure homogeneous palygorskite clay beds. General environments and direction of shore-line movement is indicated. Core diameter 10 cm.

This interval is overlain by a sandy bed which is a greenish clayey sand with irregular mottles of white sand and worm tubes. Over this is a white sand containing pelecypod shells. The mixing in the lower part of the sand is apparently due t o both burrowing and current action. There is a gradual transition from sand to dolomitic clay ( 3 m). A pure clay bed containing patches of dolomite extends up to 2.4 m. Dolomite then becomes predominant and there is 0.9 m of clayey dolomite. The dolomitic bed is overlain by 1 m of pure clay. The lower one-third has an irregular, massive appearance with some irregular vertical fracture surfaces. This clay bed grades into an interval which consists largely of the same type of clay, but is heavily burrowed and infilled with a coarser clayey sand

73

Fig. 30. M u d cracks in palygorskite clay infilled with sandy montmorillonitic clay (7.6 m).

(Fig. 32). This is the top of the core. It is overlain by a clayey sand similar to the type that occurs in the burrows in the top of the core. Thus on gross lithology, there appear t o be two similar, but distinct depositional units. Both units appear to be topped by a hiatus, and both start with a sand bed of probably marine (littoral) origin. The lower unit is characterized by mud-cracked and locally reworked sediments and the upper unit by dolomitic beds.

Clay mineralogy Palygorskite is the predominant clay in the section. Montmorillonite is second in abundance, followed by sepiolite. Illite and mica are present in

74

Fig. 31. Sample from soil zone showing ped surface with clay skin (cutan) and burrow filled with sandy clay (4.9 m).

minor amounts throughout the section. The estimated relative clay-mineral content, based on X-ray patterns of oriented slides, is shown in Fig. 33. The clay-mineral suite is closely related to the lithology and thus presumably depositional environments. In addition, there are significant differences between the upper and lower depositional units. In the pebbly zone (6-5 m) the pebbles have a high palygorskite content. The clayey-sand matrix is composed largely of montmorillonite (Fig. 33). Small (2-5 mm) rounded tan clay grains are similar in composition to the large blocks of apparent mud-crack origin. Sepiolite is present throughout this interval in the clay pebbles and grains but not in the matrix. The clay

75

Fig. 32. Palygorskite clay containing burrows filled with sandy montmorillonitic clay and lightcolored fragments of palygorskite clay (0.3 m). (Top is to the right.)

suite of the matrix is similar to that of the overlying montmorillonitic sandy soil. Montmorillonite comprises over 90% of the clay suite in the soil zone. Sepiolite has a maximum concentration at the bottom of the soil zone and is not present in the overlying sediments. This is the only interval that contains more sepiolite than palygorskite. A detailed study of the bottom part of the soil zone indicates that the clay minerals are inhomogeneously distributed. The greenish sandy clay is composed almost entirely of montmorillonite, with some biotite. Some small white pebbles have a composition similar t o the underlying large pebbles (palygorskite > montmorillonite > sepiolite). Also present are some small tannish grains (2-5 mm) and a thin tannish coating on the vertical fracture surfaces. In both of these types of samples montmorillonite is the dominant clay and sepiolite is more abundant than palygorskite. Thus the ‘sepioliterich’ suite has a distinctive occurrence. The distribution suggests the clay may be secondary and has formed by postdepositional leaching of the upper

76 Meters

WILYGORSKITE

75;-j

MONTMORILLONITE

--

SEPlOLlTE

9b

5FT50

I= BURROWS M-MATRIX P i QESOLES

Fig. 33. Graph showing distribution of major clay minerals in MC-1core.

part of the soil interval and growth in the bottom. Tannish grains in the upper portion of the soil zone are composed almost entirely of palygorskite, sometimes with dolomite. These grains must have been added from adjacent overlying sediments. Burrows in the 4-4.8 m montmorillonite-rich interval are composed largely as palygorskite, indicating clay has been worked down from as much as 1.5 m above. The burrows contain no dolomite which suggests that either the animals selectively by-passed the dolomite rhombs or that the dolomite was formed at a later stage.

77

In the overlying shelly-sand zone palygorskite increases and becomes relatively abundant within the sand. However, the mottles and pebbles of greenish clayey sand in this interval have a high montmorillonite content similar t o that of the underlying interval. Thus some of the mixing is probably due to current reworking of the lower material into the upper, rather than burrowing which would cause a downward mixing. The large shells in this interval (3.3-3.6 m) have been converted to dolomite and much of the palygorskite is secondary. This interval is discussed further in Chapter 7. Palygorskite is at a maximum (90%) in the dolomite and upper clay bed. There is a slight decrease in palygorskite in the clays of the uppermost burrowed zone. The burrows are filled with a sandy montmorillonite clay derived from the overlying montmorillonite.

Texture Twenty-one samples were sieved at 0.5-4intervals. The samples were disaggregated by pounding with a vertical motion. This tended $0 break the nonclay minerals from the clay with a minimum of damage. The material coarser than 4.38 4 was considered to be the ‘sand’ fraction. Microscopic examination indicates most of the clay ‘broke down’ and passed through the finest sieve. Fig. 34 shows the amount of nonclay material and its composition. The sands at the base of each unit contain 6 0 4 0 % sand, most of which is quartz. The clay beds and dolomite bed contain from 0.3 to 7.2% sand. Sorting values, means, etc. are not too meaningful as most samples contain several minerals, some of which are detrital and some authigenic. The calcite was dissolved from the sieved fractions of several samples and frequency curves constructed for the calcite (plus phosphate) fraction and the quartz fraction (Fig. 35). The sample from the overlying reworked zone (7.6 m) contains 24.6% nonclay material which consists of 59% calcite (plus minor value for the quartz component phosphate) and 41% quartz. The sorting (q) is 0.51 which Folk (1968)classes as moderately well sorted. The value for the combined samples is 1.33 or poorly sorted. The calcite, apparently authigenic, has a coarsely skewed size distribution. The detrital quartz in the lower interval, whether it comprises 67% or less than 1% of the sample, is moderately sorted with a well-defined symmetrical mode at 3.2-3.5 9. As the depositional environments indicate a wide range of energies were operative, the source area for the quartz sand must have had a uniform texture throughout deposition of the lower unit. The series of frequency curves (5.0-3.0 m) in Fig. 36 shows the textural transition from the lower interval t o the upper interval. In the soil zone between 5.0 m and 3.0 m the proportion of the sand-mode characteristics of the lower interval decreases and a newer, coarser, but similarly sorted, quartz-sand suite increases. The amount of clay, phosphate and calcite grains is not sufficient t o account for the new mode and most of the new, coarser

78 Meters

Fig. 34. Amount and composition of the coarser than 4.4 (I material. The bars in the left graph show the total amount of greater than 4.4 (I material. The vertical connecting line indicates the amount of quartz in each sample. The right graph shows the composition, in percent, of each sample. Mean and sorting values are also listed.

mode is made up of quartz sand, presumably from a metamorphic source (as indicated by thin-section and heavy-mineral studies). At 3.8 m, where the marine mottled sand and shelly-sand zone starts, the mode characteristic of the lower interval is completely gone and the new mode at 2.8 4 is dominant; however, this is accompanied by an even coarser new mode at 2.3 4. This coarse material is made up largely of clay grains,

79 100

90

80

70

--

60

59% CALCITE 41% QUARTZ BULK

I-

= r 0

=

g 50 IY

6 40

UUARTZ

30

20

10

0 1

2

' 0

5

Fig. 35. Frequency distribution curves of the matrix sandy clay sample from 7.6 m. Broken-line curve is for total sample. The other two curves show the distribution for quartz and calcite grains.

though there is appreciable quartz in this size range. The transgressing sea reworked some of the lagoonal palygorskite deposits and formed sand-size clay grains. In the clay-rich interval which starts at about 3 m, the two quartz-sand modes characteristic of the soil interval (5.0-4.0 m) reappear and comprise the bulk of the nonclay material. Above this interval the amount of quartz is minor; also, it is much finer, being mostly in the silt range. The quartz in the upper clay bed is distinctly finer grained than that in the lower clay bed (Fig. 37); the coarser mode in the 1.1-m sample is due to calcite. Extremely fine-grained quartz also characterized the underlying

80

0 Fig. 36. Series of frequency curves of samples from the transition interval between the two major depositional units showing the systematic development of coarser modes with decreasing depth and increasing marine influence. Dashed lines show suggestive distribution of two populations.

dolomite. The coarser fraction in this interval is also calcite. The amount of quartz and the grain size increases in the top 30 cm of the upper clay bed where some burrowing is evident. The coarse mode at 3.1 4 is presumably the quartz sand which the burrowing organisms have brought down from the overlying sand bed. The 3.8-4 very fine sand is the quartz sand that was originally deposited with the clay. In the samples which show evidence of reworking (lower interval) and burrowing (upper interval) the quartz in the matrix clay tends to be coarser grained than in the indigenous or reworked material. As might be expected,

81 101

91

81

71

61 I-

r

3 2

r U

5

51

I IE Y

41

31

21

/

11

I

Fig, 37. Frequency curves of the material from the upper and lower clay beds. The model values for the shallower samples must be in the fine siltaize range.

the currents that accomplished the reworking were probably of higher energy than those in the environment in which the mud cracks formed and carried coarser detrital material. The burrows in the more clayey material contain sandy coarser material carried down from the overlying beds. Frequency curves for the quartz, phosphate, and clay grains (latter two based on microscopic examination of the sieved fractions) indicate the mode of the clay grains is from 0.2 t o 0.5 o units coarser than the quartz mode (Fig. 38). This difference is what would be expected on the basis of the difference in specific gravity. In the soil zone (4.5 m) both the quartz and clay grains have a bimodal distribution (Fig. 39). This presumably reflects limited sorting action and

82 100

I

90

-

80

-

70

-

60

-

50

-

40

-

I

I

I

I

I

I

a RI

1 z a I

a W

z IL I-

z

W W

n

\:,

\

\

Fig. 38. Frequency curves for the quartz, clay and phosphate grains from 5.5 m sample.

may be due to deposition by flood waters as is suggested from other lines of evidence. One clay mode is coarser and the other finer than the two central sand modes. It is possible that the finer grains are fecal pellets or were created during the formation of the soil.

Composition of sand grains The sieved fractions of the greater than 4.44 fraction (nonclay) was studied with the microscope and by X-ray diffraction. Quartz is present in all samples and is generally highly polished though much of the quartz in the sand samples from 4.1 to 3.0 m have thin clay coating (discussed in Chapter 7). Varying amount of K-feldspar are present in all but the upper 0.6 m. Well-rounded, white to light-tan phosphate grains are present in amounts ranging from less than 1--15% of the greater than 4.44 fraction in all s a m ples but the top burrowed zone. The phosphate is two to three times more abundant in the lower clay bed than in the upper clay bed. The origin of the phosphate is not clear. The clayey-sand interval at 3 m contains hollow

83

I

2

4

3

5

6

7

8 Fig. 39. Frequency and cumulative curves for quartz and clay grains from the soil interval (4.5 m). The soil is believed to have developed on river flood-plain deposits which would account for the poor sorting.

cylinders that appear to be worm tubes, with the walls of the cylinders made of light-tan apatite. This same interval contains paper-thin flakes of material that also appears to be apatite (larger sheets of phosphate occur at 4 m). Phosphate is relatively abundant in the sandy interval (shallow marine) between the two basic depositional Qnits. Some of this phosphate is, at least indirectly, organic in origin. This suggests that some of the round phosphate grains could be locally deposited fecal pellets. Calcite, in amounts ranging from 2 to 78%, is present in nearly all but the upper 0.6 m of the section. From the bottom up to 2.7 m the calcite grains are generally subequant in shape with some clear rhombs and a minor amount of rice-shaped grains. The dolomitic unit, from 2.7 to 1.5 m, contains an abundance of distinctive rice-shaped calcite grains that are apparently authigenic in origin. The calcite in the upper part of the dolomitic zone and the upper clay bed is largely subequant in shape. Well-rounded clay grains are fairly abundant in the central sandy interval and in the uppermost burrowed interval. The clay grains are probably more

84

abundant than indicated in Fig. 34 as some probably broke up and others are so coarse as not to be included in the sieve analysis. Visual examination of the samples indicates clay grains range from silt size up t o 3 cm or more in diameter. Pyrite, in the form ,of rosettes and elongated ovals, occurs in the upper clay and dolomitic zone (2.3-0.5 m). This suggests that reducing conditions existed at some stages during deposition of this interval.

Thin-section Thin-section studies indicate the basal clayey sand has approximately 2% K-feldspar. The sand is composed largely of plutonic quartz and a minor amount of embayed quartz. Minor phosphate is also present. The sand grains

Fig. 40. a. Channel argillan in clay at base of soil (MC-1, 5.2 m). White bar equals 0.05 mm. b. Channel argillan and grain argillans from near base of soil (MC-1, 5.0 m). White bar equals 0.1 mm. c. Craze planes in sandy clay from near base of soil (MC-1, 4.9 m). White bar equals 0.1 mm. d. Omnisepic plasmic fabric near middle of soil (MC-1, 4.4 m). White bar equals 0.01 mm.

85

Fig. 41. a. Thick skew plane argillan lining void near middle of soil ( M C - 1 . 4 . 4 m). Dark vein is a void. White bar equals 0.05 mm. b. Mosepic plasmic fabric near top of soil (MC-1, 4 . 1 m). White bar equals 0.1 mm. c. Thick grain argillans (montmorillonite) coating quartz grains at top of soil ( M C - 1 , 4 . 0 m). White bar equals 0.05 mm.

are floating in a clay matrix, and there is little grain-to-grain contact. The depositional environment was probably shallow marine or tidal. The clay in the lower clay bed is very well oriented parallel to the bedding and was deposited in a low-energy environment. Near the top of the bed thin (2-4 grains thick) horizontal laminae of quartz grains are relatively common. These laminae also contain some clay and phosphate grains. The clay grains appear to be montmorillonite. The matrix of the laminae consists of unoriented palygorskite clay similar to that in the clay bed. The detritus in these laminae was presumably carried into the low-energy environment (lagoonal) by wind or relatively gentle periodic water currents. The matrix material in the pebbly (palygorskite) zone overlying the clay bed consists of clayey (montmorillonite) sand and sandy clay. K-feldspar, phosphate, and clay grains and relatively common. Large, embayed, quartz grains are common. The embayments probably indicate dissolution rather than a volcanic origin (Cleary and Conolly, 1972). A few composite, metamorphic quartz grains are present in this interval and in the overlying sands.

86

Fig. 42. Clayey (palygorskite) dolomite. Some have hollow (dark) centers and some carbonate (apparently dolomite) nuclei (MC-1, 2.1 m), White bar equals 30 pm.

These are not present in the basal sand which suggests there was a partial change in the source area. Thin-sections in the soil interval from 5.0 t o about 4.0 m show a systematic change. The lower interval contains an abundance of palygorskite clay grains (50% of rock) in a matrix of montmorillonitic sand. Upward the amount and size of palygorskite grains decreases, the amount and size of quartz increases, and the amount of montmorillonite appears to increase. Cutans (argillans) of montmorillonite are common, lining grains, channels, voids and joint planes. They appear t o have been formed by illuviation, though TEM pictures suggest some are authigenic. There is a suggestion that some of the palygorskite has been altered to montmorillonite, but in most

87

Fig. 43. Vague ped structure near top of upper clay bed (MC-1, 0.6 m). White bar equals 0.5 mm.

cases the montmorillonite fills thin veins and fractures in the palygorskite grains, SEM pictures of quartz grains show that some grains contain abundant solution pits. The plasma (reorganized colloidal material) fabric varies from mosepic t o skelsepic to omnisepic (Brewer, 1964). A selection of these features is shown in Figs. 40 and 41. They appear t o be identical to those found in modern soils (Brewer, 1964). The dolomitic, sandy clay (2.9m) contains scattered rhombs of dolomite. The rhombs are euhedral and many have dark centers that are apparently voids. A few aggregates of fine dolomite are present but most occur as isolated rhombs. Thin laminae contain quartz, phosphate, and calcite grains

88

that are rice-shaped. Similar calcite grains (-10%) occur scattered throughout the clay. The clay is relatively well oriented. In the overlying clayey-dolomite interval (1.5-2.3 m) the dolomite has a fair amount of grain-to-grain contact but the rhombic shape is maintained. In addition t o dark centers some of the larger rhombs have carbonate cores (Fig. 42), probably dolomite according t o X-ray data. The rhombs range in size from less than 0.005 mm t o 0.025 mm with most being 0.015 mm. The clay in the overlying clay bed is highly oriented in a horizontal direction except for the upper part (0.5-0.8 m) which has a patchy pattern (peds) with the clay in the patches being well oriented (Fig. 43). A few thin veins (channel argillans) of montmorillonite are also present. This interval appears to have been reworked, probably by burrowing organisms, and a minor amount of clay has been carried by water from the overlying montmorillonitic clay. The burrows at the top of the clay (Fig. 32) contain sandy montmorillonitic clay and pebbles and grains of palygorskite. Much of the burrow material and some of the palygorskite pebbles have a dark-brown stain which appears to be a thin film of montmorillonite.

Heavy minerals In order t o determine if volcanic material was a major source material, the heavy minerals were collected from the lower (7 m) and upper (1m) clay beds and the soil zone. The lower clay bed contains significant amounts of zircon, tourmaline, rutile, apatite, staurolite, kyanite, and sillimanite. The heavy-mineral suite suggests a mixed metamorphic-igneous source with, probably, relatively little volcanic material. The heavy-mineral suite from the soil zone is similar except that the metamorphic suite is relatively more abundant. Thin-section examination also indicates metamorphic quartz is more abundant in this section. This new source material presumably accounts for the occurrence of the bimodal sand in this interval (Fig. 36). Most of the heavy minerals from the upper clay bed are opaques. A few grains of tourmaline are present. There is little evidence that volcanic material was a major source of the sediments in this section.

Interpretation It is evident from the study of this one core that these Miocene sediments were deposited in a shallow-water environment near the strand line. In general the montmorillonitic sandy intervals appear t o be of shallowmarine origin. The horizontal-bedded, clay-rich palygorskite beds must have been deposited in a quiet lagoon. The dolomitic beds were deposited in a

89

similar environment, though some is replacement dolomite and probably formed epigenetically . The pebble and mud-crack beds represent lagoonal deposits that were later reworked by currents coming from either the seaward or landward direction. The vertically oriented montmorillonitic, organic, sandy clay bed is a soil formed on fluviatile sediments. Burrowing is evident throughout, but is particularly important in the sediments closing the end of each depositional cycle. The depositional cycles started off with a sand or sandy shell bed which acted as barriers. Shallow-water lagoons developed behind these relatively thin barriers. In the older cycle the lagoon was sometimes evaporated t o near dryness and mud cracks developed. During the early stage the barrier was breached and clayey marine sands were mixed with the mud clasts. Near the end of the cycle fresh-water currents probably did the reworking. This situation seemed t o be characteristic of the lower depositional unit. This lower unit is topped by a soil zone developed on fluviatile sediments, suggesting the overall unit is regressive. This regression is followed by an abrupt transgression (shelly marine sand) which might reflect only a minor lateral shift of environments. The upper depositional unit shows little reworking ,and a relatively thick lagoonal sequence (dolomite and palygorskite clay bed), which suggest a more permanent barrier plus a decrease in available Si and Al. This interval appears to end in a shallowing regressive environment (burrows and reworking), followed by a relatively abrupt marine transgression (Fig. 29). Thus, both depositional units appear t o represent a seaward migration of a shallow-water fluviatile-lagoon-barrier sequence. Whenever the migration was interrupted by a sustained marine transgression a new cycle began. The lithology indicates that the lower unit was probably more effected by physical energy (clay clast) and the upper by chemical energy (dolomite). The detailed mineral and chemical differences between the two units likely reflect only minor environmental and source differences; however, from a practical standpoint these differences can be used to identify the different clay beds. LA CAMELIA MINE (MC-2)

A second core (8.8 m) was obtained approximately 4 km northwest of the first core. The sediments in this core were deposited in a slightly different but related sequence of environments. The lower 5.8 m, apparently equivalent to the interval below the soil zone in the first core, is composed largely of montmorillonite sand (Fig. 44). The upper and lower portions have minor amounts of palygorskite and sepiolite. Sepiolite-rich pebbles are common in the lower meter. This latter interval also contains worm burrows and 5--10% phosphate. The sand has a major mode, 3.1 @, similar to the sand in the bottom of the first core; however, this sand has a minor coarse mode (2.0-

90 CLAY

MINERALOGY %

sa

I1 5

10.5 6 5

0 3 86

37 3.9

78

35

69 5 77 2

2.83 2.77

52 0 77 3

3.2 2.66

43 I

3.48

pY 1% 76

2.78

67 0

27

72.0

2 57

Fig. 44. MC-2 core, 4 km northwest of MC-1, La Camelia Mine. Blank area represents massive, relatively pure clay.

2.6 qj) suggesting slightly higher energy conditions in the northern portion of the area. The lower 1 m of sand is overlain by 0.6 m of relatively coarse sand (coarse Mo 1.5-2.0 qj predominant). Some cross-beds are present. The overlying sand is finer grained with the major Mo ranging from 3.0 t o 3.5,which is similar t o the sand in the lower clay and clay-clast interval in the southern core. This sand is fairly massive, but contains a few clay patches and clasts. Phosphate is scarce but the upper portion contains 5--10% of sand-size palygorskite grains. Thus the interval below the soil horizon in the northern well represents a higher-energy environment than that in the south. The sands may have been deposited as a beach or sandflat as compared t o the more tidal flat-lagoonal conditions to the south. The basal sand is overlain by approximately 2.1 m of complex sediments, presumably equivalent t o the soil and shell interval (3.3-5.2 m) in the southern well, The lower 0.6 m consists of irregularly mixed and burrowed sand and white clayey dolomite. Some calcite cement is present. The lower 15 cm consist of horizontally bedded white rhombic dolomite (the only bedded dolomite in the core). This may just be a lens. The white material overlying this occurs as irregular patches, clasts, and thin coatings. It contains hollow rhombs of dolomite and abundant palygorskite and sepiolite with only minor montmorillonite. SEM pictures indicate the material in the

91

Fig. 45. Incipient clay peds near top (1.8 m ) of MC-2core. White bar equals 0.1 mm.

white coatings is probably authigenic. The matrix clay is predominantly palygorskite, but appreciable sepiolite and montmorillonite are present. This interval is overlain by about 1m of horizontally laminated clay with thin laminae of coarse granular calcite in the lower portion and sand in the upper portion. Montmorillonite decreases upward in the section, but never attains the low values found to the south. The grain size is quite fine (Mo= 3.5 @), and the upper portion contains a few diatoms and sponge spicules. This interval is overlain by a 0.5-m thick clayey shell (calcite) bed. Sepiolite occurs in the lower portion of the shell bed but does not occur any higher in the section. Thus, sepiolite is restricted to the lower depositional unit as it is in the first core.

92

The cored section does not have a well-developed soil zone similar t o that in the southern core, but the mixed lithology and the development of veins and coatings of dolomite and thin laminae of granular calcite suggest some subaerial exposure (3.3-5.5 m). As we will show this zone is quite extensive in north Florida and has a highly variable character. The shell bed is overlain by 0.8 m of well-laminated, light-gray clay (commercial zone) which in turn is overlain by 2.4 m of fairly well laminated gray, partially iron-stained (goethite) clay with thin sand laminae and a few small patches of sand. Palygorskite is most abundant in the lower 0.8 m of the clay bed. Montmorillonite and kaolinite systematically increase in the overlying 1 m and then remain constant for the final 1.5 m. There are 3 m of sandy overburden. The top of the clay bed appears to contain incipient clay peds outlined by argillans (Fig. 45), suggesting mild weathering. The problem is whether this relatively low palygorskite content is due to weathering or is depositional. The sand content of the clay ranges from 6 to 12%. This, plus the sand laminae, indicates a slightly higher energy environment for this clay than its southern equivalent. The nonclay component is mostly fine quartz and a few heavy minerals. The iron-stained areas have the same clay suite as the unstained gray areas. SEM pictures d o not suggest the palygorskite fibers are altered. The decrease in palygorskite in the northern core probably is a result of deposition in a different, probably higherenergy environment. Diagenetic modification of the detrital clays did not progress as far as in the MC-1core. LA CAMELIA MINE OUTCROP

Over a period of several years several sections of the mine were sampled. The first section sampled was near the center of the mine and is similar to the first core (Fig. 46). The two clay units are separate by a 2-m interval (12-14 m). The upper 1 m of this interval is a fossiliferous, montmorillonitic sand. (In other areas a 1-m oyster bed (coquina) occurs at this interval.) This is underlain by 0.6 m of gray montmorillonitic clay and 0.3 m of brown pebbly (0.5-2 cm) clay. The upper part of this brown zone contains dark-brown montmorillonite clay pebbles or peds and a sandy matrix of similar composition. The matrix contains small rounded white grains rich in palygorskite and sepiolite. The clay in the lower portion of the tan interval has the same composition as the overlying rounded white pebbles and presumably they were derived from the underlying material. Some pebbles contain more sepiolite than palygorskite. Vertical worm burrows are present in this interval and associated with them are white, elongated (3 mm) fecal pellets composed only of quartz and palygorskite. Either the worms are very selective in their concentrating or they have transported clay from the overlying beds. This interval is equiva-

93 CLAY

MINERALOGY

L A CAMELIA

MINE

Fig. 46. Outcrop section in La Camelia Mine showing composition of clay fraction. 121 4 m interval is equivalent to the soil zone separating the two clay beds.

lent t o the clay soil in the MC-1 core. It is underlain by approximately 1m of pebbly or crumbly gray clay with tan montmorillonite films and then 1m of well-laminated gray clay. This is the lower commercial clay bed. The pebbly interval is extensive throughout this mine and several other mines. In general, the pebble size increases downward. In some instances the material looks like mud-crack clasts that have been only slightly moved. In other areas, and lower in the interval, it looks more like soil macropeds (compound particles, clusters of primary particles, which are separated from adjoining aggregates by surfaces of weakness). Further, the pebbles are commonly coated by a thin clay film (montmorillonite). These films are present in soils where they are called plasma separations. Thin-section examinations confirm the ped texture of the pebbly or crumbly interval. It is difficult t o establish if any of the montmorillonite is replacing the palygorskite, but SEM pictures suggest it is not. The shell-soil’horizon is overlain by approximately 2 m of gray laminated palygorskite clay (minor montmorillonite, no sepiolite) which contains lenses of white dolomite and thin layers of mud-cracked dolomite intimately mixed with the clay. The mud-crack structure indicates the dolomite is penecontemporaneous rather than syngenetic. Layers and veins (filled mud cracks) of grainy rice calcite are also present in this interval. These layers and veins are rich in montmorillonite and low in palygorskite. Montmorillonite was presumably the original detrital clay supplied to the mud-cracked depositional site by invading waters and the Ca/Mg ratio must have been

94

high enough to allow calcite t o form and to preserve the montmorillonite. The dolomite zone is overlain by 1.5 m of well-laminated gray clay (commercial bed) which grades upward into a 0.5 m of crumbly, nonlaminated clay and then to 0.3 m of pebbly clay with a calcareous sandy matrix. In the uppermost zone some pebbles are dolomitic and contain pure palygorskite. The calcareous sandy matrix has a high montmorillonite content, as does the granular calcite filling the mud cracks. This upper sequence is similar to that in the first core and also suggests a hiatus and some form of reworking at the top of both the lower and upper clay beds. The pebbly layer comprising the top of the upper clay bed is overlain by approximately 7.5 m of clay, sand, and clayey sand with several thin beds of fossiliferous limestones. Fossils in the equivalent interval in the Gunn Farm mine are assigned a Chipola age (Olson, 1966), but may be Upper Torreya (Lower Miocene). The clay in this interval is almost entirely montmorillonite. Some palygorskite sand-size grains are present. Phosphate grains are present through most of the interval. The sands are similar texturally to the lower sands associated with the clay beds. They are well sorted with a well-defined mode ranging from 3.0 to 3.4 4. In some parts of the mine, this upper montmorillonitic interval rests on a palygorskite-rich dolomite bed equivalent to the upper clay bed. The two are separated by a 5-cm layer of fine sand which contains a relative abundance of coarse, black phosphate grains. The sand layer rests unconformably on the dolomite and appears to have been derived from it by leaching. The relatively large size of the phosphate grains suggests mild marine reworking. This unconformity is equivalent to the clay-on-clay unconformity seen in other parts of the mine. In what is apparently the western portion of the mine the brown-soil zone (Fe and organic material) can be traced for at least a quarter of a mile. Over most of this interval it is composed largely of montmorillonite with minor amounts of palygorskite and sepiolite with the sepiolite being more abundant, relative to palygorskite, than in the underlying clay. The sand content becomes relatively high and the material has a mottled appearance. The underlying clay is similar to that previously described; palygorskiteand sepiolite-rich pebbly clay grading down into laminated clay. Westward dolomite lenses become increasingly abundant both above and below the soil zone. Sepiolite is present in and below the soil interval and not above. At one point in the westernmost portion of the mine the brown interval gradually thickens from 0.3 to 1.5 m, and the brown color fades out downward. Vertical burrows, 1 cm in diameter and 0.3 m or more long, are abundant. They are filled with quartz sand and clay grains and contain only montmorillonite. The brown interval is predominantly montmorillonite with a slight increase in palygorskite and sepiolite in the lower section and in the top of the underlying gray clay.

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Starting approximately 1 m below the top of the brown zone there is a concentration of white, rounded, pebble-like particles and irregular flat particles (5-15 mm in diameter) which are composed predominantly of palygorskite and contain more sepiolite than the bedded clay. Minor montmorillonite is present but this may be contamination. Below the tan interval this whitish, pearly-appearing palygorskite is present as a thin coating on the gray clay. SEM studies show the white palygorskite-sepiolite material is a secondary clay skin. Van den Heuvel (1964)found sepiolite and palygorskite in the form of sand- and silt-size aggregates and as linings on channel walls in a calcareous soil from New Mexico. The data is fairly convincing that these clays formed in place in the soil. The white particles in the Miocene soil always contain more sepiolite than any of the other clays and occur in the lower portion of the soil. This s u g gests a secondary origin. It is difficult to visualize how these large white particles could form in a soil, but then chert nodules and concretions of various compositions form easily enough in muds. The presence of dolomite or calcite particles may maintain the high pH conditions necessary for the precipitation of sepiolite. The brown zone has a higher sand content than the ,palygorskiteclay beds. East of the thick (1.5 m) brown area the material equivalent to the lower 1.2 m is a palygorskite-rich pebbly or ped clay with a relatively low sand content. Some brown clay extends down into this interval. This suggests that the variation in thickness of the brown zone is, in part, controlled by original lithologic and environmental differences, and is related to a montmorillonitic sandy facies. It is unlikely the montmorillonite formed from the alteration of palygorskite. That the montmorillonite is largely detrital, rather than altered palygorskite, is suggested by the fact that in some areas of the mine the montmorillonitic brown sand-clay is overlain by approximately one meter of montmorillonitic light-gray sandy clay with marine calcareous fosssils. In the westernmost portion the entire upper interval is apparently represented by a montmorillonitic sandy-clay and clayey-sand facies. Unfortunately this interval was not sampled in detail. To the southeast, in the Midway Mine, the brown-soil zone (relatively rich in Fe and organic material) is only 15 cm thick, and is composed of pebbly clay containing palygorskite, much sepiolite, and only trace amounts of montmorillonite. The brown stain is independent of clay-mineral type and if it represents a soil horizon it indicates that at least in some areas weathering was not severe enough to alter the palygorskite to montmorillonite. In any event there was a widespread change of conditions at the end of deposition of the first clay bed. The upper clay bed in the La Camelia Mine is much more erratic than the lower one. In addition to being a well-laminated clay and highly dolomitic, it also contains layers (up to 0.6 m thick) of grainy calcite which contain pure palygorskite. The thin-section of this material (Fig. 47) shows the paly-

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Fig. 4 7 . Spar calcite separated by thin ribbons of optically oriented palygorskite. La Camelia Mine. White bar equals 0.1 mm.

gorskite is all in optical continuity and consists of long lenses and ribbons periodically separated by layers of spar-calcite grains (1-2 mm). Most grain boundaries are smooth but a few are serrated, suggesting replacement or interference. Some grains contain small fragments of palygorskite. Residues of the calcite grains are composed almost entirely of palygorskite, in contrast to the montmorillonite in the calcite grains associated with mudcracked clay. Though there is some suggestion of replacement it seems more likely that there was a rhythmic deposition of calcite and palygorskite. First calcite would precipitate, thus increasing the Mg/Ca ratio and allowing palygorskite to form. The other possibility is that the moist palygorskite dried and developed fractures parallel to the bedding and calcite formed in these fractures. Similar rocks occur below the lower clay bed in some areas. One sample which was collected from a clayey, very fine sand in the montmorillonite ‘overburden’ interval contained a mixed-layer montmorillonite-kaolinite A good example of lateral variations in mineralogy was observed in a pit (cut NS) in the southwesternmost part of the mine. At the northern end of the west wall of the mine there is a gray, sandy (30% greater than 48 pm)

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section, 6-9 m thick. The lower portion contains mostly montmorillonite with sepiolite (minor) being more abundant than palygorskite and illite. This grades upward into a gray, sandy clay with palygorskite more abundant than montmorillonite (20-3076). On the opposite wall of the mine (15 m away) the lower sandy, gray clay (or clayey sand) contains only montmorillonite and minor illite. This is overlain by a sandy shell bed with large worm burrows. This shell interval is the lateral equivalent of the palygorskite-rich, sandy clay on the west wall and has the same clay-mineral suite. Thus, within an interval of approximately 15 m the sandy palygorskite clay grades into a marine or brackish-water shell-rich sand which presumably acted as a barrier. Laterally (along the west wall) the palygorskite content of the upper clay increases and the montmorillonite decreases to less than 10%. The sand content decreases to a few percent. This relatively pure clay continues for approximately 100 m and changes from a very light-gray, fairly well-laminated, clay into a medium-gray blocky clay with small patches of sand. Over an interval of about 3 m this grades into a brown clayey montmorillonitic sand with small patches of calcite and with only minor palygorskite and kaolinite. This in turn grades gradually into a white sandy fine-grained calcareous material containing slightly more kaolinite. This sand may be effected by recent weathering and the calcite may have come from dissolved shells. For about 30 m the section is missing, but was probably largely sand (barrier or channel?); on the other side the palygorskite-rich upper clay bed is present and continues for at least 50 m. Thus in this area of the La Camelia Mine.the lower clay bed is quite sandy and largely montmorillonitic, the soil zone is not present, the upper clay bed grades laterally into sandy clays and sands, and the montmorillonite content is closely related to the sand content as it is in other areas. These factors strongly suggest that the clay distribution is environmentally controlled and not seriously modified by weathering. The sediments appear t o have been deposited in a coastal deltaic influenced environment. ADJACENT MINES

About 1.5 km north of the northern portion of the La Camelia Mine is a small mine with a thin, slightly sandy, palygorskite-rich clay bed. Worm burrows are abundant. A thin layer of small shells is present. Granular calcite is abundant. The calcite grains contain a montmorillonite and illite. Approximately 3 km t o the southeast of the La Camelia Mine is a mine with 1.2-1.5 m of palygorskite-rich clay with no sepiolite. The bottom 0.3 m is well laminated and contains worm trails and vertical tubes. This grades upward into pebbly and poorly laminated clay with a sandy-clay matrix. Granular-calcite grains occur in both vertical cracks and horizontal layers and are associated with a relatively high montmorillonite clay suite. The montmorillonite and quartz contents increase slightly upward. The clays and quartz contents in the worm burrows in the lower portion are similar to

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those in the overlying interval. The upper 0.3 m or so of crumbly clay has a high montmorillonite content. There is a fairly abrupt contact between the clay and the overlying montmorillonite sand. Relatively well-rounded clay pebbles, from 0.5 t o 15 cm in diameter, are abundant in the sand. They usually contain a minor amount of palygorskite in addition t o montmorillonite. The pebbles were presumably derived from the top of the underlying clay bed. MIDWAY MINE

In the Midway Mine, approximately 20 km southeast of La Camelia, the lower clay bed is well developed (Fig. 48). The base of the lower clay (base of the mine) contains thin (-15 cm) beds and lenses of soft dolomite and hard, granular calcite. Thin-section studies indicate the calcite is identical to that found in the upper portion of the La Camelia Mine (alternating layers of spar calcite and oriented palygorskite). The generally smooth boundaries of the calcite grains and the orientations of the clay suggest that there has been no replacement, only infilling. The dolomite is rhombic with most having dark centers. Grains and patches of palygorskite are abundant. The clay bed contains more sepiolite and less montmorillonite than in the La Camelia Mine. Most of the clay is massive t o laminated but the upper 0.3 m consists of angular to round clay pebbles (peds) in a tannish, silty, clay matrix. In the lower portion they range from 0.5 to 8 cm in diameter. The size decreases to 0.5-1.2 crn at the top. The upper 1 5 cm has a dark-brown to black organic stain similar to that La Camelia Mine. The clay suite itself is similar to that in the

9-

6-

3-

a

0-

Fig. 48. Cross+ection of Midway Mine, Florida. Only lower clay bed is present.

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Fig, 49. Ped structure and argillans (white) in upper portion of lower palygorskite clay bed, Midway Mine. White bar equals 0.1 pm.

underlying clay except sepiolite is more abundant. Thin films of clay and worm burrows extend down below the brown zone. It is difficult to see the pebble character in thin-sections of the brown zone but it is very evident in the underlying material where the relatively pure clay clast and peds are separated by silty clay and argillans (Fig. 49). The brown zone is thinner than in La Camelia and the clay is largely palygorskite rather than montmorillonite. Sepiolite is relatively abundant in both areas. Round phosphate grains are relatively abundant in the lower clay bed. The material overlying the brown-soil zone is quite variable. Sand and clay are mixed in wide variety of ratios and textures; clay pebbles and patches of

100

calcite are present (Fig. 48).The lower 1.5 m of the interval overlying the soil zone contain roughly equal parts palygorskite and montmorillonite. Above this montmorillonite is predominant. Phosphate grains are common throughout the interval (-6 m). Thus, the sepiolite-containing lower clay bed in its structural and textural features is similar to that in the La Camelia Mine. The upper clay bed, which has an erratic distribution even in the La Camelia Mine, is missing and is replaced by a more clastic, sandier facies. GUNN FARM MINE

In the small Gunn Farm Mine (Fig. 4) about 3 km north of the La Camelia Mine the lower clay bed (3 m) contains appreciable sepiolite, in addition to minor montmorillonite. In some samples, sepiolite is more abundant than palygorskite. The upper 0.6 m of the clay bed is made up of clay clasts and contains a brown stain similar to that seen in other mines near the top of the lower clay bed. This is overlain by approximately 1.5 m of grayish, very sandy clay to clayey sand with a clay suite similar to that below. This is the highest occurrence of sepiolite and probably is equivalent to the top of the soil zone seen in other areas. This is overlain by 0.6 m of montmorillonitic sand with clay clasts and burrows. The sand is overlain by 1.5-1.8 m of palygorskite-rich clay which contains no sepiolite. Montmorillonite is more abundant than in the lower clay bed. A few oyster shells occur in the upper portion. The clay contains an abundance of relatively coarse-grained calcite. This unit is probably equivalent to the upper clay bed at La Camelia. The next interval consists of approximately 4 m of interbedded coquina (three layers) and montmorillonitic fossiliferous sand. The fauna (pelecypods, gastropods, foraminifera, ostracods, etc.) in this interval has been described as Chipola age, but may be older (Huddlestun, personal communication, 1975). It represents nearshore, brackish, shallow-water conditions (Olson, 1966). This faunal interval appears to be the same as that in La Camelia (on top the upper palygorskite clay bed). CHESEBROUGH MINE

The westernmost commercial mine occurs approximately 1.5 km south of Quincy (Fig. 4). The palygorskite clay bed (only minor amounts of montmorillonite) is approximately 5 m thick. Phosphate grains are present. Thin dolomite lenses and beds are present and most of the section contains megafossils. A thin (0.5 m) irregular fossiliferous sand bed occurs near the middle of the clay bed. This section contains more shells than sections in the other mines, suggesting it was deposited closer to a normal marine environment. It is not clear whether the whole section is equivalent to the upper clay bed or

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whether the lower portion, below the sand interval, is a more marine equivalent of the lower clay bed. The palygorskite clay is overlain by approximately 6 m of fossiliferous (pelecypods, ostracods, and manatee) clayey sands, sandy clay, and thin limestone beds. Montmorillonite is dominant with minor amounts of palygorskite. The vertical sequence suggests a transition from a tidal-lagoonal environment upward to an open-water brackish or marine environment. SUMMARY

In the southeastern portion of the mining area the lower clay bed is commonly underlain by dolomite and limestone beds which contain abundant palygorskite. Northward, near the Georgia-Florida border, montmorillonitic sand is present. The lower commercial clay bed, characterized by the presence of sepiolite, appears to be relatively extensive and continuous but grades laterally from clean palygorskitesepiolite clay to sandy montmorillonitic clay and clayey sand. Dolomite lenses are present in the first of these facies. The palygorskite-rich facies usually contains peds and/or pebbles, particularly in the upper portion, indicating some weathering and reworking after deposition. The sandy facies lies to the west of the clay facies and appears to have been a major north-south barrier sand (parallel t o the Trough and Ocala High) behind which lay a relatively large lagoon in which the clay was deposited. To the southwest (seaward) dolomite was deposited on top of the lower clay unit, indicating lagoonal conditions still prevailed. Over most of the area a brown organic soil zone was established and a pedal structure developed. To the east (landward) the soil developed on the palygorskite clay bed and there was little mineralogical change except for the development of some secondary sepiolite. Westward the top of the lagoonal clay is overlain by a montmorillonitic sandy clay with soil features. The increased sand content of this material indicates that higher-energy conditions were introduced to the area. The sepiolite distribution suggests there was some subsequent weathering of the montmorillonitic unit. Farther west the brown zone thickens and becomes more complex. The interval is represented in some areas by 1.5 m of montmorillonitic sandy clay with abundant worm burrows. This increased thickness appears to be at the expense of the underlying palygorskite clay bed. Nowhere does sepiolite extend above the brown zone. In some locations the brown color is missing and the entire lower palygorskite clay bed may be replaced by a montmorillonitic sandy clay or clayey sand. On top of the barrier sand separating the lagoon from open-marine conditions the soil interval appears to be represented by a gray, crumbly, calcareous, palygorskitemontmorillonite clay with thin secondary films of dolomite, palygorskite, and sepiolite. In other areas the brown interval grades upward into a gray sandy mont-

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morillonitic clay soil overlain by fine sand, usually containing shells. This may be overlain by a shell or sandy shell bed 0.6-4.5 m thick. The sand and shell sequence presumably represents the end of subareal or near-subareal weathering, and marks the beginning of a shallow-marine and brackish-water transgression. During deposition of the lower clay bed a large sand barrier island and lagoon apparently was present. The clay in the lagoon was fairly extensive but locally interrupted by the deposition of tidal delta sands and wash-over fans. As the lagoon filled subareal alteration occurred. This varied in character throughout the lagoon and barrier area. In the eastern portion the soil developed on the clay bed proper. To the west fluvial montmorillonitic sandy clays were deposited on top of the lagoonal deposits. This may indicate a slight lowering of sea level and/or the diversion of a nearby river into the area. The humic acids in the soil apparently were deposited with the sediments. Many of the present-day rivers in the southeast have such a high content of organic material that they are colored dark brown. The isopach map (Fig. 2) indicates an elongated thick interval in this area running perpendicular t o the shore line. This could represent a Miocene river deposit. The river would have kept the lagoonal waters relatively fresh. The suggested beach and river alignment is in general agreement with the cross-bedding inclination directions obtained from outcropping Miocene sands (Tanner, 1955). Some sands have a general northeastsouthwest pattern suggesting deposition by littoral currents. Others have a west and southwest orientation and are presumably stream deposits. The sea then reinvaded the low-relief area, a sandflat was developed, and localized shallow enclosures were formed, commonly by the development of sandy shell barriers. The resulting clay beds formed in these enclosures are of more local extent than that of the underlying lagoonal clay. The upper palygorskite clay bed (usually laminated) is well developed at Lake Talquin, near Quincy, in portions of the La Camelia Mine, and slightly east of La Camelia. It appears to extend into Georgia where it is present in the area of the Block N Mine and at Climax and Fall Caves. The lower portion usually has a high content of lenses and thin beds of dolomite. Eastward (landward) the dolomite is not present and this interval is represented by a continental mixed sand and clay facies which contains appreciable amounts of montmorillonite. To the west and northwest of the palygorskite bed the clay becomes sandier and the montmorillonite and kaolinite content increases. It eventually grades into montmorillonitic sand, presumably marine, and shell deposits. There appears to be a brief hiatus after the deposition of this upper clay bed, indicated by the presence of a burrowed thin crumbly and pebbly zone and, in some areas, a thin residual phosphatic sand layer. This was followed by an eastward movement of the open-marine conditions and the deposition of montmorillonitic sand and clays and brackish-water fauna (Chipola or

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Late Torreya age). This facies is restricted to the northwestern portion of the area (Quincy, La Camelia, and Gunn Farm areas). East and south of this area fossils are absent and the highly mixed sediments resemble continental deposits. Sepiolite is present in the soil and the lower clay bed but not in the upper clay bed. This may indicate fresher-water conditions existed during deposition of the lower sediments. This is suggested by the fact that regionally sepiolite occurs landward of palygorskite deposits. However, this could also be indicative of periodic hypersaline conditions near to shore. A'ITAPULGUS, GEORGIA

Approximately 11 km northeast of the La Camelia Mine, (Fig. 4) thbre are several small mines which contain a light-tan blocky palygorskite clay differing in general appearance from the laminated to pebbly clays to the south. Stratigraphic correlation (Fig. 5) indicates they are equivalent to the lower Miocene clay beds of Florida. At the bottom of the clay bed is a light-tan granular limestone, which in some areas is composed of interlocking sparry-calcite grains (0.5-2 cm) containing silt- and sand-size grains of quartz, K-feldspar and palygorskite. Other portions contain thin laths of parallel-oriented palygorskite units scattered among the calcite grains (Fig. 47). In the more granular material the oriented palygorskite is abundant and the calcite grains (rice-shaped) are scattered through it. The presence of worm trails suggests the original material was a carbonate mud (micrite). The quartz grains have a Mo of 2.8 $ and the lower-density clay grains 3.3 $, indicating little, if any, current sorting. This also suggests the clay grains are fecal pellets. The approximately 3 m thick clay bed was sampled at two different locations. In one section the clay bed is composed predominantly of palygorskite (-95%) with variable, but minor, amounts of montmorillonite. The uppermost portion of the clay bed has slightly more montmorillonite. The lower meter of clay contains thin sand laminae and channel-shaped patches (-25 cm). The well-sorted channel sand (Mo= 3.3) is partially cemented with calcite. The clay contains only 1.0% coarser than 400 mesh very fine sand (Mo = 3.9 $), similar to that in the Florida mines, with approximately one percent clay and phosphate grains. A few sponge spicules and diatoms are present. Some flat rounded clay clasts (-15 mm) occur associated with the sand, and one spoon-shaped patch with a concentric dehydration pattern was found The tannish clay overlying the sand interval contains irregular lenses of white clay with gradational boundaries. The white clay is nearly pure palygorskite. The amount of nonclay material (1.1%) is similar in both clays. The nonclay material in the white clay is similar to that in the tan except it contains a few sponge spicules and diatoms. The detrital material appears to have been uniform. Therefore the local variations in clay mineralogy indicate

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there may have been variations in water chemistry. The white clay may have floated in from the edge of the lagoon. Approximately 0.6 m above the top of the sand is a 2-10 cm thick bed of rice calcite with large (1-5 cm), horizontal clay chips. About 0.3 m above the calcite bed is a 0.3 m thick bed of soft white dolomite. The dolomite consists of thin-walled, partially formed rhombs. This is overlain by 1 m of clean, blocky clay. In this area there is considerable lateral variation over relatively short distances. In other portions of the mining area the entire section above the basal calcite bed is a clean blocky tannish clay with no sand laminae, calcite, or dolomite, and a relatively low palygorskite content. The sand content of the clay decreases from 34% at the base to 0.8% at the top of the section, whereas the palygorskite decreases from 90% to 40% from bottom to top. This trend (high sand-high palygorskite) is the opposite of that found in the Florida area. The sand is bimodal (3.5 4 and 4.0 4). The coarser mode value is similar to that found for the La Camelia samples. The sand represented by the finer mode is more abundant. The Mz values for the clay samples from both sections are larger (3.5 4 to 3.8 4) than for the La Camelia clays. The fact that both sand populations were present and the finer one is relatively abundant in the Block N clays suggests these clays were formed in a lower-energy environment than some of the La Camelia and other Florida clays. Mineralogically, texturally, and structurally these Georgia clays resemble the upper clay bed in the La Camelia area and have little in common with the lower clay bed. The weathered montmorillonitic clay immediately on top of the palygorskite clay (0.8% sand) contains 30.1% sand, and the sand is coarser than that in the palygorskite clay bed, suggesting the difference is a depositional rather than a weathering feature. The lower half of the clay bed in this area contains more sand than the upper half and contains phosphate grains which are not present in the upper half. This vertical change in depositional conditions is presumably equivalent to that seen in the other section where the lower portion contains sand channels and clay clasts and the upper portion dolomite. The vertical changes are not identical in the two sections, but both indicate a change to a lowerenergy environment during deposition of the upper portion of the clay bed. The thin clay beds in Climax and Fall Caves, to the north, contain the only pure palygorskite found in this study. The clay is relatively clean and does not contain phosphate grains, sponge spicules, or diatoms. It must have formed in a shallow, protected area near the edge of the lagoon.

Chapter 5 MINES - MIDDLE MIOCENE CAIRO PRODUCTION COMPANY MINE

The only other commercial mines in the area are 50 km to the northeast in the Ochlocknee-Meigs area (Fig. 4). Data from well cuttings indicate palygorskite is present in the intervening area but not as the dominant component. The regional correlation studies indicate these sediments are marine Middle Miocene. The Cairo Production Company (CPC) Mine, in northwest Thomas County, contains approximately 6 m of commercial blocky t o platy clay (Fig. 50). Approximately 3 m from the bottom there are four t o six thin beds of fine sand, ranging from less than 1 t o as much as 15 cm thick. The sand contains an abundance of well-rounded, light-tan phosphate grains and clay grains ranging from sand size up t o 10 cm in diameter. Disseminated clay and phosphate grains are also present in the clay. Thin (0.1 mm) laminae of very fine sand occur throughout the clay bed: They probably were deposited by wind action. The clay grains and pebbles have a distinct composition. Montmorillonite comprises 20% or less (zero for some), and both palygorskite and sepiolite are present (either may be predominant). Sepiolite tends t o be more abundant in the pebbles and palygorskite in the sand- and silt-size grains. Both-size particles contain more apatite than is found in the matrix clay. The basal clay is composed of approximately 50%palygorskite and sepiolite (palygorskite dominant) and 50% montmorillonite. Montmorillonite and kaolinite increase towards the top. The mined clay bed is overlain by 12 m of interbedded clay and sand with sand becoming more abundant upward in the section. The lower 7.5 m of clay are composed of 90% montmorillonite, 10% kaolinite, and a minor amount of illite. The kaolinite increases upward, becoming dominant 3 m from the surface. About 4 m from the surface there is a 1.2-msand bed that contains an abundance of biotite and muscovite. The mica is scattered through the sand and also forms relatively pure thin laminae. Microscopic and X-ray examinations show that both micas are relatively fresh, indicating that it is unlikely that the underlying clay beds have been appreciably weathered. The feldspar distribution has been discussed. Both Na- and K-feldspar are present in the palygorskite-sepiolite clay bed and only K-feldspar in the overlying material. In addition t o quartz and feldspar, the palygorskite-sepiolite clay beds contain phosphate and clay grains, sponge spicules, silicoflagellates, ebridians, and diatoms. The phosphate and clay grains are most abundant in the

106 CLAY MINERALOGY 20

40

60

80

x

I

% Sd -

l!!L

3.3 3.3 3.I

2.9 38.0 53.8 332

'

2.5

3.6 3.3

65.4

3.0

0.6

3.0

6

u)

L Y

W I-

w I

9

M

z I l-

Mica and K Feldspar

a W

G

12

7

7

Feldspar

M

15

I

O8

.

5

7

Phosphola Groins

3.5

11.8

3.1

52.7

2.7

3.4

ie

CAIRO PROD. CO.

Fig. 50. Lithology and clay mineral distribution in section from Cairo Production Company Mine, NW Thomas County, Georgia. Top horizontal line near 15 m is composition of clay pebble, lower bar is for sandy matrix. From left to right, montmorillonite, palygorskite, sepiolite.

lower portion of the bed and sponge spicules and diatoms in the upper portion. In thin-section, most of the diatoms are broken and appear to comprise 30-50% of the clay (Fig. 51). The sponge spicules are commonly filled with phosphate and many are pitted and etched. Both the needle-shaped spicules and the bean-shaped microscleres are commonly present (Fig. 52). Sponge spicules and microscleres are nearly as abundant in the overlying 7.5 m montmorillonite clay-rich interval as in the palygorskite bed, but only a few diatoms were found. Phosphate is absent and only trace amounts of clay grains are present in two samples. A few sponge spicules and diatoms occur in the sand near the top of the section.

107

Fig. 51. Photomicrograph of thin-section of palygorskite clay from Cairo Production Company Mine showing several species of diatoms. White bar equals 0.1 mm.

Dr. William H. Abbott of the South Carolina Geological Survey studied samples from this mine and kindly provided the diatom species assemblage information in Table 11. The regional significance of the diatoms has been discussed earlier. The age is Middle Miocene, confirming the Middle Miocene age based on foraminifera1 studies (Huddleston, 1973) near the Savannah River. The diatom assemblage indicates the palygorskite clay unit was deposited in a marine environment. However, Diploneis crabo, which favors a high salinity, systematically increases (3-15% of the assemblage) from the base t o the top of the palygorskite clay section. This suggests that circulation became progressively more restricted in this closed end of the narrow Middle Miocene Gulf. The more restricted assemblage in the overlying montmorillonitic sediments suggests a brackish-water environment. The diatoms in the basal portion of the uppermost sand indicate a nearshore marine environ-

108

Fig. 52. A picture of the nonclay fraction of palygorskite clay from near base of bed, Cairo Production Company Mine. Shows sponge spicules and microsclere (bean-shaped objects). Light-gray grains are clay grains. White bar equals 0.05 mm.

ment. The assemblage is similar t o that found in modern sediments, suggesting the uppermost sand sequence may be Pliocene in age. The amount of material coarser than 4.38 @ ranges from 0.5 t o 7.5% in the clay beds in both the lower and upper unit; the sands contain from 45 t o 65%. The sand sample (14.9 m) is unimodal with a M o of 2.75 @ and is very well sorted (uI - 0.35). The modal values of the material in the overlying sediments are generally coarser than those below. The nonclay in the lowermost clay sample (0.6% sand) is very poorly sorted with a coarse mode at 1.8 @ and one below 4.0 @. The size distribution and modal values of the sediments above the paly-

109 TABLE I1 Known ranges and ecology of diatom species found in the clays of the Cairo Production Company Mine, NW Thomas County, Georgia as determined by William H. Abbott Diatom species

Age

Environment

Actinocyclus ehrenbergii A. ehrenbergii var. tenella Act in op ty chus sp lend ens

Miocene to Holocene Miocene to Holocene Mid-Miocene to Holocene

A. undulatus

Late Cretaceous to Holocene Miocene to Holocene Miocene to Holocene

neritic planktonic neritic planktonic cosmopolitan-temperate-littoral cosmopolitan-temperate-neritic planktonic neritic-temperate neritic-li ttoral-benthictemperate marine-subtropical

Aulacodiscus argus Biddulphia aurita

Late Cretaceous to Holocene

B. tuomeyi Coscinodiscus apiculatus C. arcus * C. elegans * C. excentricus

Late Cretaceous to Holocene Mid-Late Miocene Late Cretaceous to Holocene Eocene to Holocene

C. marginatus C. monicae * C. obscurus C. ocu lus-iridis C. perforatus C. rothii Cy matogonia amblyocerus * Diploneis crabo Diploneis sp. Dimerogramma sp. Epithemia sp. Fragilaria sp. Fragilaria sp. Grammatophora sp. Hyalodiscus sp. Melosira complexa * M. sulcata Melosira sp. Nitzschia sp. 87 Nitzschia sp. 88 Nitzschia sp. 90 Plerosigma affine var. marylandica * Pseudauliscus spinosus Pyxidicula crutiata Pyxilla sp.

Mid-Miocene Mid-Miocene Eocene to Holocene

plank-

Eocene to Holocene

cosmopolitan-oceanic-pelagic-temperate planktonic-temperate

Mid-Miocene Miocene to Holocene

marine (favors a high salinity)

Miocene Cretaceous to Holocene

*

cosmopolitan-pelagic tonic oceanic-temperate

Mid-Miocene Mid-Miocene

benthic (neritic plankton)neritir-littoral-tychopelagic

8

(continued on p. I 1 0)

110

TABLE I1 (continued) ~

Diatom species

Age

Environment

Rhaphoneis amphiceros

Early Miocene to Holocene Miocene Early to Late Miocene Early to Late Miocene Early to Mid-Miocene Early to Mid-Miocene Early to Mid-Miocene Miocene to Holocene

marine

R . angusta * R. elegans * R . gemmifera * R . immunis * R . obesula * R . parilis R . surirella Rhizosolenia cf. barboi R . styliformis S. turris Stictodiscus kittonianus Thalassionema nitzschoides Thalassio thrix longissima Triceratitum interpunctatum *

*

oceanic-temperate neri tic-planktonic-temperatesubtropical Mid-Miocene Miocene to Holocene

neritic

Miocene to Holocene

oceanic-pelagic

Lower Mid-Miocene

* Extinct. gorskitesepiolite clay bed are similar t o those in the palygorskite section in Florida. The values for the palygorskite-sepiolite bed ( M o 4.0 > 2.7 @) are significantly lower. This is due in part, but not entirely, t o the abundance of sponge spicules and diatoms in the finer sizes. The sands in the thin sand layers in the palygorskitesepiolite bed are very well sorted; that in the clay proper is less well sorted (bimodal) with sorting decreasing upward. This upward trend probably reflects the upward increase in sponge spicules and diatoms. Von Engelhardt (1940) found that it was possible t o distinguish between water- and wind-laid deposits by determining the grain-size (mode in mm) ratio of mineral grains of differing density such as quartz and garnet (light and heavy). The ratio is larger for water (more effective separation) than for air, but always larger than one (1.9 vs 1.5). The clay grain/quartz ratio for the grains in the clay and sand beds ranges from slightly less than 1.0 t o 1.5 which is similar t o the values found in the La Camelia Mine. The distribution of the phosphate grains is similar t o the clay grains. Some of the sand layers have large, 0.5-1 cm and larger, rounded, generally flat, chip-shaped clay grains that were not included in the analysis. The sand-size grains are more nearly spherical. Thus, there are really two modes for the clay, a pebble mode and a very fine sand mode, with the quartz mode ranging from very fine t o fine sand. Closer examina-

111

tion of the samples indicates the round clay pebbles, along with angular chips, are concentrated in distinct horizons in the sand near or at the top. These horizons have a high clay grain/quartz Mo ratio as would be expected for water-deposited grains. The flat clay pebbles were probably rafted or rolled. Due t o their shape, they would not be expected to necessarily form a continuous grain-size range. As the chips break into smaller pieces they would become more equant. The size would be controlled by the minimum diameter of the original flat pebble. The sandy zones, which are more numerous farther north, probably were deposited by periodic storms. The storms transported both sand- and pebblesize material from the shoreward region to deeper water. There they were deposited with little sorting. The major source of the sand was probably well-sorted Lower Miocene sand similar to that in the Lower Miocene sediments t o the south. SUMMARY

There are two broadly different depositional marine units, each of which can be divided into two subunits. The presence of phosphate and clay grains and Na- and K-feldspar in the lower unit suggests the source of the detritus was from the east or southeast (Ocala High). The absence of phosphate, clay, and Na-feldspar grains, and the presence of abundant K-feldspar, mica, kaolinite, and a zircon-rutile-tourmaline suite in the upper unit suggest the source was from the west or northwest (southern Appalachian Mountains). The diatoms indicate the lower unit was deposited in a marine environment and the upper unit (montmorillonite) probably in a bay environment. The differences between the two units are not so much due t o a change of environments as a change in the source of sediments from one side of the trough t o the other. During deposition of the lower unit, sediment may have come from both sides of the basin. Eventually, as regression commenced, the western detritus predominated. The palygorskitesepiolite bed has a lower and upper unit. The lower unit, with the highest palygorskitesepiolite content, is relatively rich in clay and phosphate grains and the upper unit in sponge spicules and diatoms (particularly the hypersaline species). This may represent less connection with the open-marine environment for the upper portion and the influx of less detrital material from an eastern source. The overlying unit can be divided on the basis of gross lithology, into a clay-rich lower portion (bay environment) and a sand-rich upper portion (nearshore marine). The upper stand-line sand was apparently deposited during the final stage of regression during the Late Miocene or Pliocene. ADJACENT CORES

Scattered cores from a series of shallow wells 5 km southeast of the Cairo Production Company. Mine show that the upper portion of the section is

112

similar to that in the mine. The top interval consists of 3-6 m of sand containing montmorillonite partially weathered to kaolinite. This is underlain by 1-9 m of montmorillonite clay (-10% kaolinite and minor illite) with variable amounts of sand. Minor amounts (10--20%) of palygorskite were found at 9-12 m beneath the surface. The palygorskite clay section is as much as 4.5 m thick but is absent in some wells. This is underlain by at least 6 m of gray, clayey, phosphate sand which contains minor amounts of palygorskite. In several holes the section was almost entirely sand. The diatom and sponge-spicule distribution is similar t o that in the CPC mine but they appear t o be less abundant. CPM

3

4

2

y-,: . ;.

73MAboveSea Level

.

. . ..

. :. .' . .

.

'.,

.

'

.1

..

. .. . . .. . ... .. .. .,. ..

\

In W K

,

W +

. ,. .

. ....

MONTMORILLON I T E

2

T

a W

..

m 0.1 km

. . .. .

. .. . ..

, ..

Fig. 53. Crossisection from Cairo Production Company Mine (CPM) east-ward, cores 4, 3, and 2. Large indicates 4 0 4 0 % palygorskite and small -20%.

113

The thickness of the clay unit decreases from 1 2 to 1m over a distance of 1.5 km. The clay body is lens-shaped with the axis of maximum thickness aligned NE-SW, generally thinning t o the south and west and possibly continuing to thicken to the northeast. The thickness of the palygorskite portion of the clay bed increases to the northeast. This may be a local pattern, but it is very near the NE-SW aligned shore line and probably represents the eastern termination of the Middle Miocene marine clay bed. In three core holes directly 0.8-1.6 km east of the CPC mine the palygorskite clay bed, 9 m thick, thins t o 4.5 m (Fig. 53). The total clay interval (upper montmorillonite and underlying palygorskite clay beds) thins from 1 5 to approximately 9 m and becomes sandier to the east. The vertical diatom and sponge spicule distribution is similar to that in the CPC mine. Both marine clay beds thin towards the eastern shore. CHEROKEE MINE AREA

Three kilometers north and slightly west of Ochlocknee, five cores were obtained from the property of the Cherokee Company. The wells lay in a general south to north direction. The general section and mineralogy is similar t o that in the Cairo Production Company mine (Fig. 50) but appears to be sandier. Data for one core is shown in Fig. 54. The cross-section (Fig. 55) shows that there are four general lithologic zones; a lower sand, a palygorskite clay bed, a montmorillonite clay bed, and an upper sand. The top of the basal sand has a relief of approximately 3-4.5 m and resembles a beach or offshore bar. However, the clay content is relatively high. Some of the clay is in the form of silt-sized grains. The sand becomes finer downward, which is characteristic of a beach. The abundant clay suggests a low-energy environment but the clay probably filtered down from the overlying mud. A dark-tan-colored sand zone, 0.6-0.9 m thick, occurs near the top of the basal sand in 3 cores (B, A , and D ) . Sepiolite occurs above this zone but not below. Vertical clay cutans occur 1.2-1.8 m below the tan sand in cores B and A (others did not go deep enough). A ped structure can be seen in core B . SEM studies show the development of authigenic palygorskite in the cutan zone. It appears fairly certain that the upper portion of this basal sand was weathered and a soil profile developed. This presumably marks the boundry between the Lower and Middle Miocene. Diatoms and sponge spicules are present in the clay but not in the basal sand. The upper part of the sand contains clay grains and pebbles and phosphate grains. The sand is similar t o the sand laminae and thin layers in the clay bed. The nature of the top of the basal sand suggests reworking and transgression by the sea over a weathered beach. The palygorskite clay bed has four sand-clay pebble zones. The three lower zones consist of 1.2-1.5 m thick intervals that contain thin sand

114

6

9

co

n

w t

w

I

5

12

r

t-

a

w n 15

18

Fig. 54. Core A (see Fig. 5 5 ) from Cherokee Company Mine, approximately 2.5 km northwest of Cairo Production Company Mine. Horizontal lines are used to indicate composition of clay pebbles. Line at 16 m indicates pure palygorskite. Lowest line is for clay cutan. Sequence from left to right is montmorillonite, palygorskite, sepiolite, illite.

laminae and beds up to 0.3 m thick. Clay grains and pebbles are common and blocky clay clast (mud-crack chips) are present in the thicker intervals. The upper zone consists of abundant flat clay clasts and a silty clay matrix containing clay grains. The clay grains and pebbles are similar to those reported previously. Fine clay grains and phosphate grains (mostly clay) occur throughout the palygorskitesepiolite clay and sand interval, making up from 5 to 70% of the coarser than 30 pm fraction. The distribution suggests that much of the palygorskite and sepiolite in the clay is present as fine detrital grains, though some may be in the dispersed state. Both the clay grains and pebbles are strongly cemented and difficult to grind. The clay does not disperse well and only poor X-ray patterns are obtained. Apatite, present in both the grains and pebbles, presumably acts as a cementing agent.

115 B

a

C

,

,. .

D

.

. . . . _ , . _. , ' . '. -.

.

0

15

CHEROKEE

CO. MINE

Fig. 55. N o r t h s o u t h line of section (4 cores) through Cherokee Company Mine, NW Thomas County, Georgia.

Sponge spicules, microscleres, and several varieties of diatoms are abundant, though diatoms appear to be less abundant than in the CPC mine. They are not present in the basal sand and the lower sandy zone, which suggests this latter unit may be part of the basal sand. They are relatively abundant throughout the palygorskitesepiolite interval. Only sponge spicules and microscleres are present in the overlying montmorillonite clay bed. They become less abundant as the surface sand layer is approached. Many of the clay grains contain sieve-shaped diatoms and sponge spicules, suggesting they may be fecal pellets. The upper unit consists of a lower massive marine montmorillonitic sandy and silty clay (less than 10% kaolinite) and an upper clayey sand (Fig. 55). Kaolinite is more abundant in the sand. The feldspar distribution appears t o be similar t o that in the Cairo Production Company Mine (Na- and K-feldspar in the palygorskite interval and only K-feldspar in the overlying beds). The elevation on top of the palygorskite bed appears to be relatively constant (61 m f 1.5 m in this general area), which suggests that bottom relief is real. The clay bodies are generally lens-shaped. They were apparently deposited as the Middle Miocene sea gently transgressed over a series of Lower Miocene beaches, barrier islands, cheniers, or similar features, with

116

the clay being concentrated in the sheltered or lee areas. Such an environment would also be ideal for the accumulation of diatoms and sponge spicules. Two samples of palygorskite were broken in 0.5-2.5-cm size angular fragments and rolled, with water, in a roller mill. After one hour (approximately 3 km transport) the laminated sample was well rounded. The massive sample required several hours t o round. Thus, the well-rounded clay pebbles and grains need not have been transported very far. In a wave or surf environment they could be rounded within a few meters of their point of origin. As the sea transgressed over the Lower Miocene clay deposits, fragments were rounded and incorporated in the marine montmorillonitic muds. Dispersed fibers and fiber bundles were also presumably incorporated in the marine sediments. The sandy zones in the clay bed were presumably formed when current velocities were relatively high. WAVERLY MINE

The northernmost mine in the area is 10 km northwest of Ochlocknee (Waverly Petroleum Products Company). The palygorskitesepiolite clay bed increased in thickness, and the sepiolite content increased with depth from approximately 50 to 70%. The clay is more blocky near the top and more platy near the bottom. At 7.5 m from the base of the clay is a 1.5 m sand bed containing clay lenses and round t o angular clay pebbles. The sand itself, at least in part, is composed largely of relatively coarse-grained clay grains with lesser amounts of quartz. Diatoms and sponge spicules are abundant, both in the clay grains and pebbles. Biotite is present. Clay grains, phosphate grains, sponge spicules, and diatoms are also present in the clay beds. Foraminifera occur near the base. Trace amounts of dolomite are present in some samples. Aside from being thicker this clay is similar t o the clay bed near Ochlocknee. Gremillion (1965) describes two wells, 23 and 26 km t o the ENE of the Waverly Mine, that encountered the top of the palygorskitesepiolite zone at about the same elevation. The palygorskite zone is approximately the same thickness as in the Waverly Mine. In the western wells the section has 7.5 m of clay and 4.5 m of sand; in the eastern well it has 1.5 m of clay, 9 m of sand. Palygorskitesepiolite pebbles are present and more abundant in the western well. Tan chert, presumably opal-cristobalite, is present in both sections, being much more abundant to the east. In the same vicinity (approximately 8 km t o the north) well GGS175 has a 58 m diatomspongespicule section that is predominately sand, much of it coarse to very coarse. Clay pebbles are present at the base of the interval, but no palygorskite zone is present. This appears to mark the northeastern boundary of the palygorskite clay bed, though minor palygorskite and sepiolite are present at the base of the diatomsponge-spicule zone in GGS25, 32 km t o the northeast and farther up on the Ocala High.

117

DISCUSSION

Fig. 56 is a cross-section approximately parallel to the axis of the Middle Miocene trough. The top of the diatomsponge-spicule zone is used as a horizontal datum. The cross-section shows the presence of a southern barrier near the Georgia--Florida border and the location of the authigenic Lower Miocene palygorskite clay beds to the south relative to the Middle Miocene detrital ones to the north. The cross-section indicates that the palygorskite zone in the OchlockneeMeiggs area occurs at the base of a sill-like extension of the Middle Miocene section. Though such a configuration might be construed to indicate an authigenic origin the data given in preceding pages indicate that deposition is detrital with the palygorskite and sepiolite being derived from the flanks of the sill. In the Middle Miocene farther north, where only montmorillonite is present, there is a thin interval that contains chert and very coarse to pebbly quartz. This interval appears to be at the same approximate hodzon as the top of the palygorskite-sepiolite bed in the sill area. Size analysis of the CPC mine samples indicates that the coarsest size quartz occurs at the base of the montmorillonite clay bed overlying the palygorskite clay bed. This, along with the change in the faunal suite, presumably r

oa’ E

GGS

GGS

GGS

194

205

SS

5s

1564 75

25

GGS 236

I

I

I I

II

I

1

GGS

GGS I57

Q Quartz

Fig. 56. Southwest-ortheast cross-section, down center of Trough, extending from the Georgia-Florida border to the center of the Atlantic Embayment. Lower unit is largely limestone. Lower Miocene primary-palygorskite clay beds occur in upper left portion of section. Middle Miocene detrital beds occur slightly to the northeast on a sill-like area.

118

reflects a general change in environmental conditions (within the Middle Miocene or at the start of the Late Miocene) in which higher-energy conditions prevailed. The change in feldspars indicates that an uplift to the west probably caused the eastward regression of the sea. There is a belt of sepiolite-rich clay (Fig. 28) that fringes the west flank of the northern extension of the Ocala High (east flank of the Trough). Ten kilometers t o the southeast of Ochlocknee, along the east bank of the Ochlocknee River is an outcrop containing 1.5 m of pebbly sepiolite with minor amounts of montmorillonite and palygorskite. This outcrop and its northeastern extension probably represents a shore area of the Lower Miocene Ochlocknee sea and was the source of some of the detrital clay grains and pebbles supplied t o the Middle Miocene sea. Lower Miocene palygorskite-rich clays occurred along the southern and western flanks of the Middle Miocene Trough, extending at least as far north as northern Grady County, and may also have been a source of detrital palygorskite. The northern extent of the detrital palygorskitesepiolite clay in the Middle Miocene is apparently controlled largely by the Ocala High. This area remained high or was uplifted and was the source of much of the detrital clay. Farther north towards the Atlantic Embayment the Lower Miocene palygorskitesepiolite clay zone is capped by a thin montmorillonite shallow-marine sequence (sand, clay and phosphate pebbles) and was not available as a source for the Middle Miocene sediments.

Chapter 6 TEXTURE The grain-size data of the Lower Miocene samples were plotted in a number of ways t o see if there was any relation to environment. Fig. 57 shows a plot of percent coarser than 48 pm versus M,.The marine montmorillonitic sediments and the lagoonal (or low-energy environment) palygorskite-rich sediments plot in two mutually exclusive areas. Palygorskite clays contain less than 10% coarse material (sand) with a mean size finer than 3.0 4. The montmorillonitic sediments generally have more than 30% sand and the sand size increases linearly as the amount of sand increases. A number of samples lies outside these two areas. In most cases the reason is evident and confirms the legitimacy of the boundaries. The montmorillonitic soil samples plot outside the montmorillonite area due t o the presence of ‘extra’ clay that was added by the soil-forming process. Theoretically the amount of secondary clay could be calculated (5--10% minimum). The sample labelled 1 is from an interval transitional between the marine and restricted environment. In addition, it has a high phosphate content. When only the quartz fraction is considered, the sample plots close to the palygorskite area. Sample 2 is actually a marine sand and originally contained montmorillonite that was later partially altered to palygorskite. Sampies 3 and 4 are from an interval transitional between the restricted and marine LOWER MIOCENE 100

I

I

I

I

I

I

I

I

O1.2

. . . I

I

20

I

I 28

I..

. A . 3.6

.

Fig. 57. Plot of M, (in 4 units) versus % sand (>48 pm). Large dots = marine montmorillonitic sediments. Small dots = lagoonal palygorskite clays. Circled dots = soil. For meaning of numbers see text.

120 Middle M i o c e n e

0 24

28

32

36

38

Mz

Fig. 58. Plot of M, (in 9 units) versus % ’ sand (<48 pm) for Middle Miocene sediments. Large dots = montmorillonitic sediments. Small dots = palygorskite-rich sediments. Circled dots = sandy beds in palygorskite clay.

environment. The samples contain mud cracks filled with marine montmorillonitic sand. The marine influence dominates the textured data. Sample 5 is from the very base of a palygorskite clay bed and has a relatively high (-20%) montmorillonite content. The textural data suggest this sediment was deposited under marine conditions during the final withdrawal of the sea. The montmorillonite was later altered t o palygorskite as the environment became restricted. The overlying clays have the texture typical of sediments deposited in a restricted environment. A few samples plot above the montmorillonite (marine) area. They have little in common except they may all have been deposited under conditions where winnowing was more effective than normal. Fig. 58 is a similar plot of textural data from the Middle Miocene, mostly the Cairo Production Company Mine. There are a number of things of interest. Most of the montmorillonitic sediments have the same M, vs % sand relations as the montmorillonitic Lower Miocene sediments. However, three samples fall in the palygorskite field. These samples contain less than 4% quartz and indicate that physically restricted environments existed in the Middle Miocene but the chemistry and/or temperature did not favor the formation of palygorskite. Also, even though the palygorskite clays are detrital and were deposited in a marine environment, the textural data indicate the environments were of extremely low energy. They were adjacent to highenergy environments, as indicated by the presence of sandy layers that have the textural characteristics of open-marine conditions. The overall grain size is smaller than for the Lower Miocene. This indicates either generally lowerenergy conditions or a difference in source material. In part, the finer size is due to the presence of diatoms and sponge spicules which are concentrated in the finer sizes.

121

A plot of percent finer than 48 pm and the phi value of the major mode shows that sediments having less than 5%sand have a mode larger than 3.4 Cp. Those with more than 80% sand have modes less than 3.0 Cp. In between 5 and 80% sand there is a systematic decrease of 0.5 Cp in the M o value. The plot confirms the general relation of energy (-% sand) and the size of the sand grains. The effect is relatively minor except at energy-extremes beach vs restricted lagoon. Eighty percent of the samples have their major M o between 3 and 4 @I (very fine sand). This is because the samples are all from clay beds and associated sediments that were deposited in low-energy environments. Very coarse sand and gravel are relatively abundant in other Miocene sediments.

This Page Intentionally Left Blank

Chapter 7

ELECTRON MICROGRAPHS MC-1 CORE

Transmission electron microscope (TEM) pictures were made of both replicated and dispersed samples. Scanning electron microscope (SEM) pictures were made of rock fragments. SEM pictures of the mud-cracked clay chips which occur in a zone (7.7 m) on top of the basal sand in the MC-1core (Fig. 29) indicate the chips contain scattered, 4-10 pm, euhedral dolomite crystals in a matrix of long fibers of palygorskite and sepiolite (Fig. 59). Some short fibers coat quartz grains. The presence of long fibers in the thin mud chips and short fibers in the

Fig. 59. SEM picture of clay chip from mud-crack interval near base MC-1core (7.7 m). Large euhedral dolomite rhombs occur in matrix of long fibers of palygorskite and sepiolite. White bars = 1 pm.

124

Fig. 60. Bundles of palygorskite fibers in lower clay bed (MC-1). Montmorillonite occurs as thin 0.5-1.0 pm patches. White bar = 0.1 pm. (TEM)

thick, overlying clay bed (Fig. 60) suggest that the more evaporatic environment favors the development of long fibers. Other samples will be described which tend to confirm this relation. The palygorskite in both the upper and lower clay beds mostly consists of 1 pm fibers. Montmorillonite occurs as thin patches 0.5-1.0 pm in diameter (Fig. 60). The orientation pattern is more likely due t o authigenic growth than sedimentation. Various lines of evidence indicate the palygorskite formed from the montmorillonite but there is nothing in these micrographs t o suggest either clay mineral has formed by alteration of the other. The two types of clay are intimately mixed and it is not obvious which, if any, might be authigenic. The restricted size of the montmorillonite areas, the random distribution, and lack of any continuous sheets suggest the montmorillonite may be floccules that settled among the palygorskite fibers. However, the chemical data suggest it is an authigenic Mg-smectite. The random orientation of the palygorskite aggregates is more apparent than real. Some pictures suggest that in a given plane the palygorskite fibers are nearly parallel, but the next layer of fibers may be oriented at a right angle t o the underlying layer.

125

SEM pictures of the smooth vertical faces (cutans) from the lowermost part of the soil zone show a smooth surface consisting of long (greater than 10 pm) parallel fibers (Fig. 61). These fibers may be sepiolite which is abundant in this interval. A sample about 0.3 m higher, above the sepiolite zone, has thin continuous sheets of smectite (saponite or nontronite?) with scattered unoriented palygorskite fibers on the vertical faces. High-magnification pictures (Fig. 62)

Fig. 61. Vertical cutans from near base of soil interval (MC-1) showing subparallel alignment of long fibers which are probably sepiolite. White bar = 1 p m in both cases. (SEM)

126

Fig. 62. Cutan consisting largely of thin laths of smectite. White bar = 0.1 pm. (TEM)

show that the smectite consists of thin laths, 0.05 pm wide and 0.25 pm long. Many of the laths are oriented parallel to one another and are draped over the coarser palygorskite fibers. The montmorillonite flakes in the clay beds are subsequent in shaps. The smooth, glossy vertical faces are apparently clay skins (cutans) which have been deposited from water percolating downward. It is difficult to determine if this is due t o a simple translocation of clay particles or authigenic growth. The occurrence of sepiolite, the parallel orientation of the smectite laths, and the chemistry suggest that there has been some growth of clay minerals. The marine sandy clay and clayey sand overlying the soil contains thin pelecypod shells which have an interesting mineralogy. A small ribbed shell near the base of the marine sand has been altered to protodolomite. However, no rhombic shape has developed and only trace amounts of short palygorskite fibers are present. The shell consists of overlapping plates (Fig. 63a), presumably similar in shape to those in the original calcite or aragonite shell.

127

Fig. 63. a. Dolomitized shell consisting of overlapping plates of dolomite. b. Partial rhomb constructed of thin plates. c,d. Replaced shell composed of partial, hollow rhombs and palygorskite. e,f. Low-magnification picture showing parallel orientation of partial rhombs in replaced shell. g. Palygorskite layer associated with dolomitized shell. h. Anhedral rhombs and plates of dolomite in center of shell. White bar = 1pm in all cases.

Fig. 64. Replaced shell from base of lagoonal dolomitic sandy clay interval. MC-1,3.3 m. a. Long palygorskite fibers on surface of shell, Both top and bottom surfaces have similar coatings. b,c, and d are views of a vertical fractured face. b. Band of long fibers at base of

shell. Dolomite rhombs occur towards the interior of the shell. c. High-magnificationpicture showing flamboyant growth of fibers near edge of shell. d. Dolomite rhombs and plates near center of shell edge. White bar = 1 pm in all cases. (SEM)

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1

Energydispersive X-ray (EDX) analyses indicate only Ca and S are present in the surface layer. This is a common feature of the Miocene protodolomites. The outer layer is deficient in Mg and, in some, S is relatively abundant. Sulphur is more abundant in this shell than in the overlying shells containing rhombic dolomite. It may be that in the early stages of dolomite formation sulphur or sulphate is incorporated and inhibits the development of the rhombic shape. Fig. 63b shows a clay-coated rhomb from the sand matrix. Six centimeters above this shell is a large thin shell composed of hollow, poorly formed rhombic dolomite, commonly containing only three sides (Fig. 63c, d, e, f, and h). Some particles appear to be simple plates. The shell walls are 0.24.5 pm thick and the rhombs range from 2 t o 1 0 pm, with most being between 2 and 5 pm. Also present are aggregates of 0.1-pm particles which may be precipitated dolomite. In some areas the partial rhombs have an overlapping parallel orientation, probably inherited from the original calcite plates. The dolomite layer is overlain, and perhaps underlain, by a thin burlaplike layer composed of long (greater than 10 pm) palygorskite fibers. The palygorskite is authigenic ,and commonly occurs in exotic patterns (Fig. 63g). A small shell, approximately 30 cm higher in the section and from the dolomitic sandy-clay interval, contains well-developed rhombs and an abundance of palygorskite. The top surface of the shell consists of a mat of long fibers oriented parallel t o the shell ribs (Fig. 64a). Underneath the surface layer are welldeveloped dolomite rhombs and long fibers in various flamboyant arrangements (Fig. 64b). Dolomite anhedral rhombs and plates increase in the interior of the shell (Fig. 64c, d), and the palygorskite occurs as both short and long fibers. Long fibers are abundant at the bottom edge and coat the basal surface. The dolomite in the clay immediately under the shell is mostly in the form of euhedral crystals, but some partial rhombs are present (Fig. 65). The upper shell contains welldeveloped rhombs near the surface, poorly developed ones in the center, and abundant palygorskite. The shell 30 cm lower contains hollow, skeletal rhombs and less palygorskite. The shell 6 cm lower contains no rhombs (but is dolomitized) and contains little or no palygorskite, though incomplete rhombs and palygorskite occur in the sand in which the shell occurs. Thus a gradation exists within a single shell and within the sand interval. It seems probable that the upper shell, which was deposited in a tidal or lagoonal environment, was dolomitized before or shortly after burial. Dolomitization of the lower shells, which occur in a shallow-marine tidal sand, was presumably restricted by the limited availability of Mg. During the early stages of replacement, Mg apparently diffuses into the shell’s calcite plates, preserving the plate morphology. Once the plates achieve a protodolomite composition growth starts at the edge of the plates,

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Fig. 65. Incomplete dolomite rhombs and long fibers in dolomite sandy clay in which the shell shown in Fig. 6 3 occurs. White bar = 1 pm.

forming first two- and three-sided rhombs and eventually six sides, leaving the middle hollow. Growth appears t o be layer by layer. The layers and lenses of dolomite in the lagoonal clay overlying the sand contain well-developed rhombs. The dolomite rhombs disseminated in the clay are relatively fine grained (2-7 pm) and a few have hollow centers. Upward, where the dolomite is predominant over the clay, the rhombs are larger (10-40 pm) and slightly less euhedral; a few (-5-10%) have hollow centers (Fig. 66). The large dolomite rhombs in the dolomite bed have a surface that consists of overlapping, thin, flame-shaped layers, starting from rhomb edges (Fig. 68a, b). The appearance suggests slow sheet-by-sheet growth under near-surface conditions. This would appear t o be the mechanism by which crystals grow once the rhombic shape is achieved. The surface layers are deficient in Mg, but not as much as in the plate-

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Fig. 66. Intimate mixture of euhedral dolomite rhombs and patches of short fibers in lagoonal sediments, MC-1,2 m. Note growth features on surface of rhombs. White bar = 10 pm in length. a,b. Two slightly different types of overlapping, thin, flame-shaped layers on the surface of dolomite rhombs. This pattern suggests layer-by-layer accretion rather than replacement or direct precipitation. White bar = 1pm in both cases. (SEM)

shaped dolomite of the lowermost shell. EDX analyses showed a Ca/Mg ratio of 4 (peak height) for normal dolomite and 13 for the surface of euhedral Miocene dolomite. After treatment with HC1 for 1.5 minutes the ratio of the surface was lowered t o 7. It should be mentioned that this Ca-rich rind, rather than indicating accretion of Ca and later diffusion of Mg inward, may only indicate that the final water in contact with the crystal was Ca-rich. The clay-dolomite configuration suggests that the dolomite formed earlier than the palygorskite. If the Mg/Ca ratio is larger than unity, the ratio in solution will increase as dolomite grows (assuming direct replacement of calcite is not a major factor). If the ratio becomes high enough and sufficient dissolved Si is present the formation of palygorskite would be favored.

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Fig. 67. Lense of soft dolomite mud from lower clay bed, La Camelia Mine. a. Thin dolomite plates arranged in the form of petals in a rose bud. b. Plates, incomplete rhombs and rib-like structures of dolomite. c. Low-magnification picture showing porous, fragile

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texture. d. Anhedral dolomite grains with coating of fine palygorskite fibers (frosted appearance). White bar = 1 pm in all cases. (SEM)

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Fig. 68. Lense of dolomite from upper clay bed, La Camelia Mine. a. Euhedral rhombs and scattered palygorskite fibers. b. Euhedral dolomite grains with a coating of palygorskite fibers. c. Long fibers extending across voids between layered dolomite crystals.

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d. Thin sheets of dolomite. White bars in a and b are 10 pm long; in c and d white bar = 1 pm. (SEM)

138 OTHER LOCATIONS

Dolomite Two lenses of dolomite were examined from the western (seaward) portion of the La Camelia Mine. One sample (FG-12) is 120 cm below the soil zone and the other (FG-15) 140 cm above. Both are ordered protodolomite with three mole percent excess CaC03, but the morphology of the crystals is radically different. FG-12 contains isolated plates and subequant, anhedral, hollow dolomite grains (4-10 pm), some made up of thin plates arranged in the fashion of petals in a rose bud (Fig. 67a). Rib-like structures are also observed (Fig. 67b). The overall impression is that much of the dolomite is present as thin plates (less than 0.1 pm thick) arranged in a delicate fashion t o produce a porous aggregate (Fig. 67c). Palygorskite is present as long interlaced fibers between the dolomite particles and coating the particles (Fig. 67d). The dolomite rock is soft and plastic when collected in the wet state and dries to produce a very light rock..In thin-section the dolomite appears as densely packed subequant particles. All rhomboid grains (-10%) have hollow centers which suggest that hollow rhombs represent an intermediate stage between the formation of skeletal grains and the formation of solid euhedral dolomite crystals. In contrast, FG-15 consists of well-developed 10-15 pm (some as small as 1 pm) rhombs (Fig. 68a) with only a few having hollow centers (limpid dolomite of Folk and Land, 1975). Many of the rhombs have the layered, flamelike surface similar to that described previously. Palygorskite coats many of the crystals (Fig. 68b) and occurs as a rope network extending across crevices 10 pm in width (Fig. 68c). Occasionally two- and three-sided rhombs are present, showing a distinct layered (-0.1 pm thick layers) effect (Fig. 68d). Dolomitization is more advanced in this lens than in the lower one. The clay associated with this lens is nearly pure palygorskite, whereas FG-12 contains appreciable montmorillonite. This suggests Mg was less available in the FG-12 site than at FG-15. The dolomite and palygorskite in these samples appear t o have formed penecontemporaneously in a tidal-lagoonal environment. The layered partial and skeletal rhombs apparently represent an intermediate stage in the construction of limpid dolomite crystals. Due to marine regression and development of a soil on top of the clay bed the Mg supply was cut off and the dolomite in the lower lense was never converted t o limpid dolomite. The lower dolomite lense (FG-12) is extremely pure but occurs in a sandy clay containing abundant clay grains and clasts up to 3 cm in diameter. The upper lens (FG-15) is also relatively free of quartz but contains thin horizontal lenses and stringers of palygorskite. The center of the lens appears to be massive but the c’racked edges (dehydration features) have a definite layered appearance. The textural pattern suggests that the dolomite has not replaced

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the palygorskite clay. This is also suggested by the SEM pictures which indicate that much of the palygorskite formed later than the dolomite. Dehydration fractures apparently developed in the dolomite and were later filled with palygorskite. Cook (1973) found subspherical dolomite concretions forming 10-20 cm below supratidal mudflats (silicate mud) in Broad Sound, Queensland. The dolomite is Ca-rich and occurs as subround anhedral crystals. No rhombs were present. The interstitial waters have a high Mg/Ca ratio. The original concretions were presumably composed of calcite and were progressively dolomitized. There is n o obvious explanation why the original concretions formed. It was suggested that mangrove roots or some other organic material rotted to change the pH and facilitate precipitation of calcite. The dolomite lenses in the palygorskite clays may well have had a similar origin with the extent of crystal construction being related t o the length of time the concretion had access to Mg-rich waters. No mechanism for the formation of the original concretion is evident. Small sandy shell beds occura few hundred yards from the dolomite lenses. The dolomite lenses could have been the site of former shell beds, but the lack of quartz and the presence of clay laminae make this seem unlikely. The porous, frothy-textured dolomite is so light that relatively large fragments could have abeen rafted or rolled from adjacent tidal flats into the lagoonal muds. Dolomitization could have continued after carbonate fragments were deposited in the mud. (In the field notes FG-12 was described as a ‘roundish dolomite pebble’.) Approximately 30 m farther west of the area of dolomite lenses the soil zone is overlain by gray-green sandy montmorillonite clay containing irregular white and light-gray calcareous intraclasts from 0.5 to 7 cm in diameter. This appears t o be a reworked interval equivalent t o that described in MC-1 core (3.5-4.0 m) and was formed during a shallow-marine transgression. The white clast consists of fine-grained (3-10 pm) bladed clacite with a few skeletal dolomite rhombs. This is by far the finest-grained calcite found associated with the clay beds and the only calcite that occurs as intraclasts. Similarly-appearing material is commonly dolomite. SEM pictures of the calcite show subrounded discoid particles as small as 1 pm. Long palygorskite fibers are scattered among the calcite crystals (Fig. 69a). At the boundary of the calcite clast and the clay, long palygorskite fibers extend from the clay area into the calcite area (Fig. 69b). The fibers appear to become progressively shorter (0.5-1 pm), narrower and more densely packed with increasing distance from the calcite (Fig. 69c). At a distance of 800 pm, and perhaps less, fibers are not present and only montmorillonite flakes are seen. The SEM pictures strongly suggest that the montmorillonite has altered t o palygorskite with the amount of alteration being most advanced near the calcite clast. It would appear that the fine calcite was originally Mg-calcite or possibly dolomite from which the Mg was released to react with the montmorillonite to form palygorskite (montmorillonite + Mg-calcite or dolomite + palygorskite + calcite). The resulting

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Fig. 69. Apparent example of montmorillonite and Mg-calcite altering to palygorskite and calcite. a. Discoid particles of calcite with long palygorskite fibers. b. Long fibers extending across void between montmorillonite area (right) and calcite area (left). c. Short fibers a few hundred microns from edge of calcite. Farther from calcite only montmorillonite flakes are observed. White bar = 1 pm in all cases. (SEM)

calcite has plate-like crystals which Folk (1974) believes are indicative of calcite formed from low-Mg water. This could have occurred shortly after deposition or any time thereafter. This location appears to be seaward of the dolomite-palygorskite-rich facies. In the MC-2 core, dolomite occurs as thin white coatings and tQin clasts in the clay and sand zone that is equivalent to the soil zone. Most of the crystals are anhedral and hollow. Some have one or more sides missing, and a small percentage is euhedral. Palygorskite and sepiolite occur coating the dolomite and can be seen in the interior walls of hollow particles. The features observed are quite similar to those in the FG-12 dolomite lens. In one area a burlap-like layer of long fibers overlies the dolomite particles (Fig. 70). This suggests that palygorskite was the final mineral formed. The films of dolomite and palygorskite were presumably formed by percolating waters, apparently on vertical pedal surfaces. Some palygorskite appears t o have formed from montmorillonite (Fig. 71). It is pertinent t o note that the shell bed (largely oysters) overlying the soil zone, and most of the shell beds in other areas, contains only calcite. These calcite shells are associated with montmorillonite clays and generally are not overlain by dolomitic sediments. Apparently, where conditions

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Fig. 70. Small vein of anhedral and incomplete (lower right) dolomite grains overlain by a burlap-like mat of long fiber clay. Vein occurs in sandy soil zone, MC-2core, 5 m (Fig. 44). White bar = 10 pm. (SEM)

were close to normal marine and where current energy was high, the Mg concentration was inadequate for dolomitization or the conversion of montmorillonite to palygorskite. SEM studies of the tan dolomite to the east of the Trough area in Echols County indicate that the spherical to subrhombic dolomite grains (5-10 pm) are composed largely of stacks of thin plates similar to those seen in FG-12 (Fig. 67). Spherical dolomite is common throughout much of the Lower Miocene and Upper Oligocene. Pyrite is present as aggregates of 0.30.5 pm particles and may account for the tan color. Sulphur is also present in the dolomite. The relative abundance of sulphur suggests that the original depositional environment may have been, in part, hypersaline and that some gypsum may have deposited. The gypsum could have been decomposed and pyrite and dolomite formed. Short fibers of palygorskite appear to have grown in place and some occur extending from the edges of montmorillonite

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Fig. 71. Clay from soil zone near vein shown in Fig. 70. The palygorskite occurs as thin sheets of parallel fibers suggesting it formed from montrnorillonite. White bar = 1 pm. (SEMI

flakes. It is difficult to decide if this represents palygorskite forming from montmorillonite or merely indicates that the palygorskite fibers cannot be resolved in the flake area. The former interpretation is favored. Electron probe pictures of unbroken hollow dolomite rhombs (100 pm wide) from higher in the section show good rhomb shapes for Ca. Neither Mg or Fe can be detected in the rhombs (the electron beam penetrates 3 pm). When the rhombs are polished to remove the surface layer, Mg can be detected in the center of the rhomb (Fig. 72). This outer rim of Ca-rich, M g deficient material is approximately 3-4 pm thick on the larger rhombs. The Ca and Mg patterns of these rhombs, with large (30 pm) hollow cores, suggest that there are three zones: an outer Ca-rich zone, a central Mg-rich zone (containing Ca), and an inner Ca-rich zone. X-ray studies indicate no calcite is present; the ordered dolomite contains 5.2 mole 7% excess CaC03. Leaching experiments similar t o those described by Peterson et al. (1963)

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Fig. 72. Electron probe X-ray photographs of dolomite grains. Upper two pictures demonstrate deficiency of Mg in the surface layers. (Field width 100 pm.) Lower two pictures taken of grains that were polished to remove the surface layer. Mg is now evident. (Field width 200 pm.) The dolomite appears t o have both an outer and inner zone deficient in Mg.

caused no shift of the 211 X-ray peak, though EDX studies of the leached surface indicate that the Mg/Ca ratio of the surface increased. Initial growth must start by the precipitation of small carbonate particles (authigenic), probably of varying composition, but subsequent growth (diagenesis) proceeds in a planar or sheet form. In some, if not all, instances the accreting material is Ca-rich and probably deposited relatively rapidly. This is followed by the slow migration of Mg towards the interior of the particle.

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In many cases this accretion and migration can proceed from both sides of the hollow skeletal rhomb. Calcium migrating outward as it is replaced by Mg helps create an environment favorable for the addition of more Ca from solution. When the solution is no longer replenished from external sources, migration stops and an outer Ca shell is preserved. Many have suggested that dolomite, palygorskite and sepiolite are formed primarily in hypersaline brines. No evidence was found in the present study to indicate hypersaline conditions existed during the formation of these three minerals in the area of study, though there is abundant evidence of shallow water and periodic subareal exposure. Badiozamani (1973) has calculated that mixing Yucatan meteoric ground waters with up to 30% sea water causes undersaturation with respect t o calcite, whereas dolomite saturation increases continuously. Dolomitization can therefore occur in the range between 5-30% sea water. At 5%sea water the Mg/Ca ratio is slightly less than 1.0. The Yucatan waters are probably similar to the Miocene ground waters draining peninsular Florida. Folk and Land (197.5) concluded that in a hypersaline environment the Mg/Ca iatio must be as high as 5 : 1 t o 10 : 1 for dolomite to form. Because of rapid crystallization and interference from other ions the dolomite is not well ordered and is micritic. When hypersaline water is diluted with fresh water (schizohaline conditions of Folk and Siedlecka, 1974) dolomite can form at Mg/Ca ratios as low as 1 : 1. Because of the low ion content the rate of crystallization is slow and large, ordered, clean (limpid) crystals develop. The latter conditions are also those that favor the development of the wellcrystallized palygorskite with which the Miocene dolomite is associated. Some of the clay consists of thin layers and sheets, suggesting there was periodic influx of water and sediment t o the lagoons. The limited faunal data suggest the lagoons contained brackish water that was periodically diluted further with fresh water. It is possible that the brackish water temporarily became hypersaline after it was introduced t o the lagoons. The dolomite was constructed layer by layer and face by face (constructed dolomite). When the source of Mg was cut off or the Mg was incorporated in competing silicate minerals, construction stopped short of completion (large limpid crystals). When the Mg supply was not interrupted complete rhombs formed and were enlarged by the growth of thin sheets of dolomite starting from the edges of the crystals. Weaver (1975) published a brief note in Geology describing the layer-bylayer growth or construction of dolomite. In the same issue Deelman (1975) described layer-by-layer growth occurring on a crystal of experimentally produced dolomite-like minerals. Growth was produced by periodically interrupted desiccation. The mechanism was similar t o that occurring in the shallow Miocene lagoons where periodic desiccation must have occurred. Nesteroff (1973) shows SEM pictures of ordered dolomite that consists of 2-5 pm rhombs, partial rhombs with plates, separate plates, and rhombs with central voids. This dolomite is similar in many respects to the Miocene

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constructed dolomite. This is believed t o be penecontemporaneous dolomite and occurs in the Upper Miocene (Messinian) evaporate deposits of the Mediterranean. The restricted faunal and isotopic data indicate the dolomite was deposited in shallow water between periods of desiccation. The water was a mixture of marine and rain water. It would appear that the growth mechanism (construction) observed in the present study may be a common process. SEM pictures of a well-laminated clay from the upper palygorskite clay bed in the La Camelia Mine indicate that the individual layers are approximately 3-5 pm thick and contain short palygorskite fibers and scattered dolomite rhombs (4-14 pm). The clay is nearly pure palygorskite. The layered structure suggests periodic influxes of material (clay and water) to the lagoon. The solid transported into the lagoon was presumably montmorillonite. It could have been derived from either the seaward or landward side. The lack of appreciable S suggests that the continental input was dominant. Locally, the dolomite is concentrated in relatively thin irregular patches. These patches are thicker than the clay layers and have irregular surfaces, suggesting that this dolomite formed epigenetically in the bottom muds rather than by direct precipitation.

Calcite As mentioned previously, rice calcite grains are locally abundant. In the layered palygorskite, the calcite occurs in discrete horizontal zones (approximately 1 cm thick). Calcite is abundant, comprising 50% or more of the layer. A few scattered calcite grains occur in the clay immediately underlying and overlying the calcite-rich layers, indicating that chemical conditions which favored the formation of the calcite began and ended gradually rather than abruptly. Beds of coarse calcite spar commonly occur near the top and at the base of the palygorskite clay beds. The distribution of the calcite crystals indicates they formed contemporaneously or shortly after formation of the palygorskite mud. In some intervals the parallel-laminated dolomitic palygorskite clay consits of angular, 0.5-7 cm chips that have formed by dehydration. The vertical mud cracks are filled with the rice calcite, montmorillonitic clay and quartz sand. An elliptical, well-preserved dehydration mud chip was found in a vertical mine wall. SEM pictures of the broken edge of this clay chip showed the presence of abundant irregular lens-shaped voids. The common occurrence of elongated calcite grains irregularly interlayered with layers of palygorskite clay strongly suggests that calcite grew in and enlarged the horizontally elongated pores formed by dehydration, as well as growing in the vertical cracks. Both types of voids are a product of dehydration. The incoming waters deposited montmorillonite and sand in the vertical

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fractures. Calcite precipitated in both the vertical and horizontal voids. In the former the calcite incorporated some detrital montmorillonite but in the latter palygorskite. Folk (1974) has suggested that where the Mg/Ca ratio is more than 2 : 1, the Mg selectively poisons the calcite, inhibiting sideward growth and producing elongated (c-direction) grains. Where Mg is low, more equant or rhombic calcite spar forms. The rice calcite is definitely elongated but not excessively so, and much is equant. Folk reports that some fresh-water calcite grows as basal hexagonal plates stacked so as t o resemble a book of mica. SEM pictures of a thin calcite bed near the top of a palygorskite bed shows this type of morphology. The elongated calcite and equant to book calcite resemble that described by Folk as forming phreatic meteoric cements at intermediate depths below the water table. This suggests that the Miocene spar calcite was deposited from relatively fresh waters. When crude palygorskite is added to water in equilibrium with MgO it will adsorb Mg up to a maximum concentration of 4% MgO. Thus, palygorskite can not only lower the Mg/Ca ratio when it forms from solution, but after its formation and dehydration it can probably lower the Mg/Ca ratio of new water with which it comes in contact. Thus, when fresh fluids are supplied t o the dessicated lagoon some of the Mg is adsorbed bg the pre-existing palygorskite, decreasing the Mg/Ca ratio and allowing calcite t o precipitate. Though the coarse spar is abundant, small stringers of micrite can be observed in the clay voids. The calcite may have precipitated as fine, 1-2 pm crystals, as the Mg content was lowered these recrystallized into spar calcite. Following deposition of the calcite the Mg/Ca ratio is sufficiently lowered so that dolomite and later palygorskite can form. In adjacent areas, probably closer to tidal inlets, the water was more saline and oysters and other brackish fauna flourished and little chemical change occurred in the carbonate minerals (calcite) and clay minerals (montmorillonite). Other Lower Miocene clays Clays outside the mining area show considerable variation. Fragile-appearing fibers are present in a Miocene soil sample from an Echols County core. Many of these fibers are 0.2 pm long and some only 0.1 pm. There is some suggestion they may have formed from the fine montmorillonite flakes with which they are associated, but this is largely speculation. X-ray patterns have weak broad peaks for palygorskite (10-11 A ) and montmorillonite (16-29 A). Palygorskite boiled three hours in 1N HC1 has a similar appearance. X-ray patterns of the treated palygorskite show a broad shoulder extending to 15 A. In another shallow core in northern Florida, X-ray patterns (Fig. 73) indicated that palygorskite breaks down to form a 1 : 1type mineral (7.5-8.2 A and 3.50-3.70 A). Similar spacings were obtained from laboratory precipitates which contained Mg-Al double hydroxides with a relatively high Mg

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content. TEM pictures show normal fibers and thin flakes. The flakes are presumably the 8-8 material. One meter higher (increased weathering) a chloritic material was formed. In the chloritic sample the flakes are coated with small (-0.01 pm) spheres which are presumably amorphous Fe-A1 oxides and hydroxides (Fig. 74). The section is predominantly carbonate and pH values should be relatively high. During the initial phase the palygorskite fibers must lose some of their silica and form a two sheet clay. As

8.0i

- 2-2 e

Fig. 73. X-ray pattern of 3 clayey sand samples from upper 2 m of core. Patterns (bottom to top) illustrate palygorskite altering to a chloritic material. All specimens treated with ethylene glycol.

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Fig. 74. Palygorskite altered to a chloritic material. White bar = 0.1 pm. (TEM)

additional Fe and/or A1 becomes available from the further breakdown of the palygorskite (and other minerals) they combine with the unstable double hydroxide layer t o form a chlorite. A weathered clay immediately on top of an unweathered palygorskite bed appears to show an example of a palygorskite altering to montmorillonite (Fig. 75). Very thin, delicate fibers occur together with thin flakes. The X-ray pattern of an untreated sample has a broad 001 peak with apex ranging from 10.0 t o 12.0 A . The glycolated 001 peak has its apex at 1 7 A but extends, with a decreasing slope, to 8 A . This may be some form of mixedlayer palygorskite-montmorillonite. In any event the palygorskite definitely appears to have altered towards a montmorillonitic material. Thus, in one instance palygorskite apparently weathers t o chlorite and in another to montmorillonite. A number of outcrop samples were examined from west of the Trough in southwest Georgia. At Climax Cave the clay consists of approximately one meter of laminated white (pure palygorskite) and green (palygorskite plus montmorillonite) layers. The white palygorskite consists largely of long irregularly intertwined fibers (Fig. 76a). The green clay is quite variable, consisting of various mixtures of short palygorskite fibers and montmorillonite

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Fig. 75. Possible example of fiberous palygorskite weathering to montmorillonite flakes. White bar = 0.1 pm. (TEM)

flakes (Fig. 76b). Some areas are devoid of fibers. Many of the short fibers occur as thin, parallel stacked sheets rather than bundles. The clay in the overlying clayey sand is largely very thin lath-shaped smectite. Replicas indicate palygorskite is present as long (more than 20 pm) discrete fibers and bundles of fibers (Fig. 76c) which apparently grew from solution. Pictures of dispersed clay show that there is a continuous gradation in shape between small, thin palygorskite fibers and smectite laths and it is not always obvious which is which (Fig. 76d). The general appearance would suggest the smectite probably formed from the palygorskite (during weathering). Both minerals may be primary with the thin laths being saponite. However, the smectite associated with other authigenic palygorskites does not have a lath shape.

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Fig. 76. a. Long palygorskite fibers in white clay from Climax Cave west of Trough. White bar = 1pm. (TEM) b. Green clay interlayered with white clay shown in Fig. 76a; shows sharp boundary between palygorskite (short) to left and montmorillonite to right. Palygorskite may be thin layer on top montmorillonite. White bar = 1pm. (TEM) c. Concentration of long fibers in clayey sand overlying clay bed shown in Figs. 76a and 76b. Long fibers lie on top montmorillonite and apparently grew from solution rather than from alteration of the montmorillonite. White bar = 1 pm. (TEM) d. Picture of dispersed clay from same sample as Fig. 76c. Fibers range from long and thick to short and thin. These latter fibers resemble lath montmorillonite and may be a weathering product of palygorskite. White bar = 0.1 pm. (TEM)

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The sepiolite in the pebbly-clay outcrop southeast of Ochlocknee, on the east flank of the Trough, consists of short, relatively broad fibers. $EM pictures (Fig. 77) of sepiolite-rich clayey sand from the Goode No. 1 well, farther south on the east flank, show that long authigenic fibers are abundant. They appear identical t o palygorskite fibers in other samples. The long-fiber variety of palygorskite is relatively abundant in most sediments except the relatively thick and pure commercial clay beds. TEM pictures of replicas of the Lower Miocene brittle clays from the Block N Mine northeast of Attapulgus, Georgia, show that the clay consists predominately of short palygorskite fibers containing 0.5-1 .O pm diameter patches of montmorillonite. In some areas palygorskite fibers appear to blend into the montmorillonite suggesting one formed from the other. The bundle texture is most common but in some areas the fibers have a parallel orientation and resemble sheets. The horizontal cleavage surface of this brittle clay is extremely smooth

Fig. 77. Sepiolite in clayey sand from east flank of Trough. White bar = 1 pm. (SEM)

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but contains small areas that have a wrinkled appearance. SEM pictures indicate that the smooth portion contains short fibers and the wrinkled areas long fibers (Fig, 78). Chemically they are nearly identical except the short material appears to have slightly more K and less Mg. A sample was collected from the Block N Mine which showed the features of an elliptically-shaped dehydrated puddle (Fig. 79a) which for some reason

Fig. 78. Upper picture shows long fibers that comprise wrinkled clay surface. Lower picture shows short fibers that occur in the area of the smooth clay surface. Both areas on same bedding plain a few centimeters apart, Block N Mine. White bar = 1 pm in both cases. (SEM)

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was preserved. Small (10 cm), subround intraclasts occur in the same horizon. SEM pictures indicate that the main clay consists of short palygorskite fibers (Fig. 79b). However, the dehydrated clasts contain tightly packed long fibers (Fig. 79c). The situation is similar t o that described in the preceding

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Fig. 79. a. Well-preserved desiccation feature found in wall of Block N Mine. Smooth area at left is regular bedded clay. b. Short fibers in the matrix clay shown in Figure 79a. Material is believed to have formed by replacement of montmorillonite. White bar = 1pm. (SEM) C. Long, subparallel fibers in clay clast formed by desiccation (Fig. 79a). Clay is believed to have grown from residual solution. White bar = 1 pm. (SEM) d. Edge view of clay chip formed by desiccation (Fig. 79a). Desiccation produced small subparallel voids as well as vertical fractures. In other samples these voids are filled with calcite. White bar = 10 pm. (SEM)

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paragraphs. The legitimacy of the dehydration origin is confirmed by the presence of small voids parallel t o the surface of the clay chip (Fig. 79d). Voids have not been observed in the massive clays. The distribution suggests the short fibers grow in a soupy amorphous mass, formed by the partial solution and replacement of montmorillonite. The residual fluid collected in slight depressions in the mud bottom and long fibers grew from the residual solutions. This is in keeping with other evidence that suggests montmorillonite alters directly to short palygorskite fibers and the long fibers grow from a more fluid phase. Later calcite grew in the vertical and horizontal dehydration voids. The small amount of carbonate minerals in these northern clay beds and the lack of fossils suggest the waters were different from those farther south. They presumably had a low Ca content and were possibly lacustrine. Some of the fresh water probably had a northern source (piedmont metamorphic rocks) and had a higher Mg/Ca ratio than the fresh waters to the south.

Fig. 80. Middle Miocene clay from Cairo Production Company Mine. Fibers are short and montmorillonite is relatively abundant. Diatoms are abundant. White bar = 1 pm. (TEM)

157

These latter waters most likely came from the area of the Ocala High (limestones) and had a relatively low Mg/Ca ratio. Middle Miocene In the Cairo Production Company Mine and the Cherokee Mine the Middle Miocene clay consists of an intimate mixture of short palygorskite fibers and montmorillonite flakes (Fig. 80). Some areas are predominantly montmorillonite and others palygorskite and sepiolite. Some diatoms are unaltered, others have been partly dissolved though there is no evidence of replacement. Casts of sieve-type diatoms are common and show palygorskite filling the holes in the diatom mold (Fig. 81). The appearance is more that of a diatom settling into a soft, fluid mud rather than clay growth. A few pictures (Fig. 82) indicate that there has been some solution of the diatoms and sponge spicules.

Fig. 81. Diatom fragment and impressions in clay. Insert is a higher-magnification picture of knobby impressions. White bar in insert = 10 pm; other bar = 1p .

158

Fig. 82. Lower figure shows unaltered and partially dissolved sponge microsclere from Cairo Production Company Mine. White bar = 10 pm (SEM). Upper figure shows a diatom with various degrees of dissolution. Small spheres may be precipitated opal. White bar = 1 pm. (TEM)

In the lower part of the overlying montmorillonite clay, isolated kaolinite flakes are poorly defined. Some of the montmorillonite is lath shaped. Kaolinite and iron-oxide particles occur on rough surfaces; smooth-surface areas contained only montmorillonite. The sequence becomes more continental upward suggesting the kaolinite is detrital. However, the SEM pictures suggest some of the fine kaolinite has altered from montmorillonite.

159

Soil SEM and TEM pictures of the small soil peds from the soil layer in the Midway Mine indicate that most of the clay consists of short fibers. This material is coated with a thin layer of long-fiber material (Fig. 83). The long fibers extend across 7-pm voids and obviously have grown in place. These are presumably clay skins similar to those in the MC-1 core. It is difficult to determine which fibers are sepiolite. EDX analyses suggest the long fibers have less Ca than the short fibers and that they are formed later. As sepiolite is abnormally high in this bed, it appears likely the long fibers are sepiolite. In the western portion of the La Camelia Mine the organic zone is composed largely of montmorillonite and the palygorskite and sepiolite occur as

Fig. 83. Two examples of authigenic long fibers in soil zone, Midway Mine. Fibers, largely sepiolite, occur as cutans. White bars = 1 pm in both cases. (SEM)

160

vertical clay skins on peds and as irregularly shaped white nodules (0.5-1 cm). The skin consists of approximately parallel short fibers. The material beneath the skin of fibers consists of densely packed flakes, presumably montmorillonite. Pictures of a small, white palygorskite nodule (Fig. 84) from near the base of the soil suggest that montmorillonite has altered to palygorskite. Thus, in various areas, the clay skins are composed of long fibers of sepiolite, short fibers of palygorskite, and lath-shaped smectite. This is presumably a function of available Mg and Si and pH.

Grains and pebbles SEM pictures of quartz grains afforded some useful information. Some of the quartz grains in the lower clay bed in the MC-1 core show a fine etched

Fig. 84. Picture of 0.5 cm white palygorskite nodule from near base of soil. Fibers appear to form from montmorillonite flakes which comprise the bulk of the soil. White bar = 1 pm in both cases.

161

Fig. 85. Upper, surface of quartz grain from lower palygorskite clay bed, MC-1 core. Markings suggest etching. Quartz was apparently one of the sources of Si for the formation of palygorskite. White bar = l pm. (SEM) Middle, clay-coated quartz grain from upper clay bed, MC-1 core. (SEM) Lower, palygorskite fibers replacing quartz in upper clay bed, MC-1 core. White bar = 1pm. (SEM)

162

pattern indicating some solution (Fig. 85). Quartz grains in the upper clay bed and dolomitic clay contain either a thin coating of short palygorskite fibers (Fig. 85) or an etched pattern with depressions of the same size and shape as the palygorskite fibers (Fig. 85). Thus, particularly in the upper clay bed, some palygorskite replaced quartz indicating detrital quartz was a source of some of the silica needed for the formation of palygorskite. The pictures provide further proof that the palygorskite is authigenic. The solution features also indicate that the waters had a relatively high pH. In the sand from 3.8 m, slightly below the dolomitic shells, quartz, feldspar, and Ti-oxide grains are all coated with palygorskite (Fig. 86). The samples were dispersed in a blender and then wet sieved, so the clay is bonded fairly strongly to the grain surfaces. The palygorskite appears to be a clay

Fig. 86. Clay coatings (cutans) on feldspar (upper left), Ti02 (upper right) and quartz (bottom). White bars on upper two pictures equal 10 pm, on lower picture 1 pm. (SEM)

163

coating (grain cutan) rather than a replacement and apparently was formed by seepage from above. Some of the quartz grains from the upper part of the soil interval are highly etched (Fig. 87) and contain no clay (embayed quartz is abundant in thin section). Quartz from near the base of the soil is not etched and has little or no clay on the surface. This sequence is what would be expected in a soil profile and suggests alkaline conditions existed during weathering. The quartz grains in the tan clay in the Block N Mine are not coated, but some appear to be etched, Quartz grains from the white palygorskite-rich clay layer commonly contain coatings of short palygorskite fibers. The feldspar grains do not. This, along with other features, indicates that different environmental conditions prevailed within the lagoonal environment. Quartz grains form the Middle Miocene clays in the Cairo Production Company Mine do not have any clay coating and show little evidence of

-

~~

Fig. 87. Etched quartz from upper part of soil, MC-1. White bar on main picture equals 10 pm, on insert = 1 pm. (SEM)

164

solution. Many of the diatom and sponge spicules show evidence of solution (Fig. 82). The feldspars in this section and in all other samples examined show evidence of differential solution (Fig. 88), with Ca remaining and Na presumably being removed. Leaching may have occurred in the source area. TEM pictures of an apatite-containing, sepiolite-rich pebble from the Middle Miocene show it is composed almost entirely of bundles of long fibers (Fig. 89) and a feVt small patches of montmorillonite. Long fibers were not observed in the matrix clay. Diatoms and sponge spicules are much less abundant than in the matrix clay. The surface of the pebble appears to have a thin coating which makes the fibers difficult t o see. Ca distribution pictures (EDX) indicate that the apatite is fine grained and evenly distributed. Small, 0.1 pm and less, eggshaped particles are abundant. These are apparently the apatite grains. The sand-size clay grains have smooth surfaces that show a dense packing of short fibers and fine granular apatite (Fig. 90). Similar size phosphate

Fig. 88. Weathered plagioclase feldspars from clay beds. EDX shows Ca, Si, Al. Leaching presumably occurred in the source area. White bar = 10 pm. (SEM)

165

Fig. 89. Clay pebble from Cairo Production Company Mine. Thin patch of montmorillonite occurs at bottom of picture. Boundaries do not suggest any gradation between the two morphologies. White bar = 1 pm. (TEM)

grains are composed of fine, 0.1 pm, apatite particles. SEM pictures of the fracture surface of a gray phosphate pebble from the top of the Middle Miocene (GGS-425) show it contains diatoms and sponge spicules which have been replaced by apatite (Fig. 91). The matrix phosphate consists of 0.1-0.3 pm particles which range in shape from subcubic to rod-shaped to platy. The particles tend to be larger than those seen in the clay grains (Fig. 90). The replaced diatoms have a variety of textures which may be due to differences in the morphology of the original frustules. One type (Fig. 92) consists of subspherical particles 0.5-1.0 pm in diameter. Another type (Fig. 93) contains 6-pm spheres composed of small (1-2 pm) parallel-stacked tabular crystals. A third type (Fig. 94) contains -0.5 pm and 2-pm subspherical grains. Still another (Fig. 94) contains 2-pm long prismatic crystal along with subspherical aggregates.

Fig. 90. Typical sand-size clay grain from Middle Miocene clay bed. Fibers are short. Fine, -0.1 pm apatite particles can be seen. White bar on grain equals 10 pm.Other white bar equals 0.1 pm.(SEM)

Fig. 9i. Gray phosphate pebble from near the top of the Middle Miocene. Note diatoms and sponge spicules. The former are replaced by apatite. Insert shows fine apatite crystals. White bar = 1Elm in both cases.

167

Fig. 92. Diatom in phosphate pebble (Fig. 91) replaced by 0.5-1.0 pm subspherical apatite particles. White bar on main picture equals 10 pm, on insert 1 pm. (SEM)

EDX spectra indicate some replaced diatom frustules have more Si than the matrix, but P and Ca are more abundant than Si. Si is most abundant in the small spheres (Fig. 92) and minor or absent in the prismatic crystals (Fig. 94).Replacement is presumably more complete in the more crystalline varieties. X-ray patterns indicate sepiolite and palygorskite are the dominant clays, though none was detected in the SEM pictures. The clay suite (sepiolite > palygorskite > montmorillonite) is similar to that commonly found in the clay pebbles. The original matrix material of t4e phosphate pebbles was presumably these clays which have since been replaced by apatite. Pebbles in the Miocene range from nearly pure clay (largely sepiolite and palygorskite) to nearly pure apatite. The two components occur mixed in all proportions. At least some of the phosphate pebbles form by apatite replacement of clay pebbles. The mechanism is discussed in the chemistry chapter.

168

Fig. 93. Diatom, in phosphate pebble (Fig. 91), replaced by apatite plates which are stacked so as t o form 6 pm spheres. White bar on main picture equals 10 fib, on insert 1 pm. (SEM)

TEM pictures of replicas of a jet-black pebble showed that the interior of the grain is composed of tightly packed rod-shaped apatite grains, approximately 0.1 pm in length (Fig, 95). Other black pebbles contain mixtures of rods and flakes. EDX spectra indicate Si, Mg, and A1 are present in the same relative concentrations as in palygorskite. This suggests replacement of palygorskite-rich pebbles. A replica of the surface of a black pebble shows two types of patterns. One displays the fine rods characteristic of the interior of the crystal and is presumably an uncoated area. The other is a non-granular surface with a thin, shingled, wave-shaped pattern that has the appearance of having been applied with a paint brush (Fig. 96). This apparently is an Fe-organic coating. The pattern is somewhat similar to that seen on some dolomite rhombs. How the coating was applied is not clear but it appears to be an

169

Fig. 94. Upper, diatom in phosphate pebble (Fig. 9 1 ) replaced by approximately 0.5 pm and 2.0 pm subspherical apatite grains. White bar equals 10 pm. (SEM)Lower, diatom in phosphate pebble (Fig. 91) replaced by apatite. Some prismatic apatite crystals are present. White bar = 1 pm. (SEM)

external coating rather than something that diffused from the interior of the grain. The pattern suggests the coating could have been added drop by drop from a percolating solution. Earlier studies of the insoluble residue of phosphates from this and other areas (Weaver and Wampler, 1972) showed that the clay minerals consist almost entirely of illite and glauconite. Weathered, amorphous palygorskiteshaped fibers are commonly observed. These are similar in appearance to the weathered fibers observed in outcrop samples (Fig. 74) and further indicate that apatite has replaced clay pebbles and perhaps clay grains.

170

Fig. 95. Jet-black phosphate pebble consists of 0.1 pm rod-shaped apatite grains. White bar = 1 pm. (TEM)

Fig. 96. Polished black surface of phosphate pebble. The material is apparently an Feorganic complex. White bar = 1 pm. (TEM)

171

Fig. 97. Insert at upper right is a low-magnification picture of white vein in clayeysand soil at base of Middle Miocene clay bed, Cherokee Company Mine. Soil probably represents boundary between Middle and Lower Miocene. High-magnification picture shows fibers of palygorskite extending from edge of vein. Round particles are Si, probably opal. White bar = 1p m in main picture and 100 g m in insert. (SEM)

Amorphous silica Near the bottom of core A in the Cherokee Company Mine is a sandy clay soil zone with tan montmorillonite-rich clay skins on the peds. Thin white palygorskite-rich veins occur in this skin (Fig. 97). These may be burrows but appear more like root casts. The vein area contains long, parallel palygorskite fibers which commonly extend across or into voids (Fig. 97). Intimately mixed with the palygorskite are opal spheres, mostly 1 pm in diameter. These spheres are made up of smaller particles, around 0.05 pm in diameter. Individual fibers are relatively broad at the base and become thin and bend at the end extending into the void (Fig. 97). Those extending across the voids thin towards the middle. Fibers appear t o grow from both

172

Fig. 98. Quartz grain on edge of vein (Fig. 9 7 ) encrusted with Si spheres and coated with palygorskite fibers, White bar on main picture equals 10 pm, on insert 1 pm. (SEM)

sides of the void or vein. When two fibers meet they join together and growth proceeds from both edges towards the middle. Some fibers contain small pods of silica and have the appearance of a pussy willow branch. Near the edge of the vein, quartz grains are encrusted with these silica spheres which commonly form a solid crust (Fig. 98). Just outside the vein there are parallel bundles and sheets of large fibers. Interspersed silica spheres coat the surface. Farther from the edge, short (1 pm) fibers occur. They are made up of 0.05-pm spheres arranged in rows (Fig. 99). These are presumably the same basic silica spheres which form the larger spheres in the vein. Apparently these spheres align themselves and form the framework for the formation of palygorskite. Magnesium may have acted ds a flocculating agent. This is another example of palygorskite forming in the Miocene soils. The palygorskite apparently formed from Si-rich solutions in void areas and by a

173

Fig. 99. Short palygorskite fibers adjacent to vein (Fig. 97), apparently formed by coalescing of 0.05 pm Si spheres. White bar in main picture equals 0.1 pm, in insert 1 pm. (SEMI

coalescing of flocculated silica spheres in the more restricted areas. Appreciable Ca and S are present in the vein and nonvein area. These may be present as gypsum (nondetected) or incorporated in the silica particles. As discussed in previous chapters opal-cristobalite is relatively abundant in the Lower Miocene. The material is fairly well crystallized, giving a broad but distinct peak at 4.12 A and a secondary peak between 4.25 and 4.30 A (tridymite). It is commonly associated with dolomite and palygorskite, suggesting a shallow-water, relatively high-pH environment. Diatoms and sponge spicules are present in varying amounts and are presumably the source of much of the Si. . SEM pictures indicate that much of the opal-cristobalite is massive, but bladed spherules (Fig. loo), similar t o those described by Wise et al. (1972) from deep sea cores, are common. The environment of deposition is no

174

deeper than shallow marine, indicating that the bladed spherules are not restricted t o deep-marine sediments. They more likely represent an environment where deposition was extremely slow. Fig. 101 shows clusters of spherules that appear to be similar to the opal-cristobalite-bladed spherules seen in Fig. 100. Both samples are from the same Lower Miocene strata in Echols County. A higher-magnification picture (Fig. 101), and EDX data, indicate the spherules are made of short, stubby palygorskite fibers. Some of the fibers appear t o be composed of an alignment of small spheres similar to those seen in Fig. 99. This and the previous example indicate that where Si is abundant palygorskite can apparently form by the coalescing of small Si (or Si-Mg) spheres. The same series of pictures indicates palygorskite also grows from solution. Earlier pictures showed it formed from montmorillonite.

Fig. 100. Bladed spherules of opal-cristobalite in Miocene clay from Echols County. si was obtained from diatoms and sponge spicules. White bar in main picture equals 10 pm, on insert 1 pm. (SEM)

175

Round nodules of opal are present near the base of the Middle Miocene. One from the Cairo Production Company Mine was described by Pollard and Weaver (1973). It contains well-rounded spheres of opaline silica (Fig. 102) similar to those in precious opal. The spheres occur in the cavities of diatom fragments. They occur as individual spheres and also coalesce t o form rods and sheets. Most of the opal-cristobalite is massive. Diatoms, quartz, and clay minerals (montmorillonite >> palygorskite) decrease in abundance from the edge towards the center of the nodule. The size of the spheres increases in the same direction. The montmorillonite occurs as 0.4-0.5 pm patches of laths. The thin laths are commonly oriented at 60" t o each other. Most of the montmorillonite in the clay proper consists of subequent flakes, though laths are present. Palygorskite occurs as isolated short fibers.

Fig. 101. Cluster of spherules in palygorskite clay from Echols County, Ga. Compare with Fig. 100 froin same area. White bar = 10 pm. High-magnification of spheres indicate they are composed of short palygorskite fibers, some appear to be constructed of small spheres as in Fig. 99; white bar = 0.1 pm.

176

Fig. 102. Spheres of opal in nodule from base Middle Miocene clay bed. Main picture shows sheets of spheres. Insert shows loose spheres inside cavity in a diatom. White bar = 1 pm in both cases. (TEM)

Partical chemical analyses were made of five samples extending from the white edge t o the middle gray portion of the nodul. A1203 (6.0-2.5%), Fez03 (3.00-2.20%), and MgO (2.12--1.20%) all systematically decreased towards the center, with most of the decrease occurring near the edge. Assuming all the AI2O3 is present in montmorillonite the amount of clay decreases from approximately 50% to 25%. The Al/Fe values are low, suggesting that Fe hydroxides may have aided in the precipitation of Si02. The systematically changing chemical ratios, particularly A$03/Mg0, and the hexagonal arrangement of the laths suggest the montmorillonite is authigenic rather than detrital. Thus, even though the Si concentration was high in this Middle Miocene environment, montmorillonite rather than palygorskite formed.

Chapter 8 CHEMISTRY SILICATES

It is extremely difficult to obtain monomineralic clay separates from these Miocene clays. However, the less than two micron fraction was separated (using distilled water) from 22 samples, and chemical analyses were made. Table I11 contains the raw chemical data and the mineral composition of these samples. The structural formulae for 5 relatively pure montmorillonites are shown in Table IV. Four of the samples are from marine clays, and range in age from Early to Late Miocene. Sample FF18 is from the soil horizon separating the two clay beds in the La Camelia Mine. One of the four samples is from the upper montmorillonite clay in the same mine, one from central Florida, one from the Florida panhandle, and one from Echols County, Georgia. The montmorillonite has a very uniform composition over this relatively large area (approximately 16,000 sq. miles). As much of the Ca is present in apatite and some of the K is in illite, an assumed value of 0.33 was used for the exchange cations. Both tetrahedral and octahedral A1 are relatively low but, with the possible exception of FF-18, within the range for montmorillonites (Weaver and Pollard, 1973). Octahedra Fe and Mg are present in approximately equal amounts (0.27 vs 0.26) in the marine samples. The soil clay contains considerably larger amounts of octahedral Mg and Fe. Some of the Mg is present in sepiolite. Other analyses indicate that the soil interval commonly contains a high Fe content. Some of the Fe may be in the authigenic lath-shaped montmorillonite in this interval but iron oxides are also present. Table IV lists structural formulae for some palygorskitprich samples. In addition to montmorillonite and sepiolite, illite is commonly present. The illite can be detected by the presence of a 5-A peak and the fact that the peak width at half-height of the 1 0 . 5 4 peak increases as the K20content increases. Some of the Ca is present in apatite (X-ray). The only high-purity sample is FE-21 from an outcrop on the west flank of the Trough (FE-17 is nearly as pure). The sample contains only a trace amount of Ca. The structural formula is typical for palygorskites with approximately half the filled octahedral positions containing A1 (Weaver and Pollard, 1973). The chemistry of the other samples reflects the presence of the other clay minerals. Where sepiolite is present, the Mg content is relatively high and Si low. Montmorillonite causes a decrease in tetrahedral Si and increase in tetrahedral Al. Montmorillonite, similar in composition t o Miocene material, should cause a decrease in octahedral Mg. The palygorskite samples with minor smectite actually have more Mg than the pure palygorskite. If it is

TABLE I11 Chemical composition of clay fraction ~

Sample

Wt.%

llgk

SiOl

A1203

Fez03

MgO

CaO

NazO

K20

Li

Sr

Ba

EC-2 EC-12 F5-7 F5-13 FE-17 FE-21 FE-29 FE-62 FE-80 (soil) FF-1 FF-18 (soil) FF-19 (soil) FF-54 FG-9A (soil) FG-10 FG-14 FG-16 Peb.

50.9 53.9 51.8 50.2 56.3 60.5 42.7 53.4 56.5 52.0 52t 53.8 57.3 58.9 60.5 58.0 52.0

2.41 4.33 12.68 2.14 8.87 9.56 6.52 6.68 11.57 2.53 3.67 7.55 10.97 2.99 8.10 3.78 9.87 11.35 * 6.90 3.70 2.06 1.50 10.45 3.64 10.21

0.97 1.02 0.56 1.52 1.79 1.34 1.94 0.80 0.85 1.12 1.36 1.53 1.19 1.47 1.02 1.05

1.22 2.82 0.54 1.05 1.32 0.81 0.96 0.71 0.54 0.90 0.94 0.91 0.46 1.03 0.72 1.32 0.64

45 81 67 42 34 15 85 30 57 26 73 55 21 59 49 48 47

110 180 98 175 35 9 360 130 38 270 46 26 68 65 19 58 235

110 145 40 75 95 75 220 60 45 470 300 105 75 170 95 170 130

58.8 58.7 54.9 55.0 56.8

4.20 3.98 3.65 5.71 3.89 3.75 4.01 4.98 5.68 4.96 7.40 6.16 3.83 5.56 4.65 4.35 4.33 4.98 * 4.30 5.71 4.54 2.73 4.23 5.22 3.86

5.08 5.36 2.95 4.70 0.80 0.04 10.50 4.87 2.08 1.39 0.88 0.46 1.87 1.10 0.32 1.71 7.58

FG-16 matrix FH-5 FH-12 FH-37 MCI-5 MCl-18 MC1-23

17.42 14.12 9.51 16.83 12.41 12.15 13.57 13.74 9.85 17.80 15.32 12.35 10.80 17.3 11.65 15.94 10.50 12.03 * 12.92 16.17 18.04 23.38 7 0.85 14.59 10.79

0.54 1.00 1.36 1.36 1.51 1.12 0.68

0.96 0.97 0.79 0.86 0.50 0.77 0.67

43 29 42 15 23 46 28

44 69 57 200 52 69 62

150 150 180 530 55 150 90

57.7

* Corr. for 15% apatite

0.83 1.00 1.02 0.31 1.05 2.01 1.81

1.00

TABLE I11 (continued) Mineral composition clay fraction Montmorillonite EC-2 EC-12 F5-7 F5-13 FE-17 FE-21 FE-29 FE-62 FE-80 FF-1 FF-18 FF-19 FF-54 FG-9A FG-10 FG-14 FG-16Peb. FG-16 FH-5 FH-12 FH-37 MC1-5 MCI-18 MC1-23

98 62 10 96 2

Palygorskite

Sepiolite

Kaolinite

Illite 2

38 51

39

2 95 99

70 30 34 62 3 68 29 99 84 8 8 27 50 19 5 95 95 2 2 14 65 16 76 22 15 65 20 47 24 5 98 95 Mixed-layer kaolinite-montmorillonite 6 93 96 2 2 7 90 3

2 3 1

4 1

4 1

5 2 1' 3

3 1 2 1

assumed that the palygorskite in the clay that consists of approximately 90% palygorskite and 10%smectite (MCI-5) has the same composition as the pure palygorskite (FE-17) it can be calculated that the smectite has the composition: SiOz = 65.7%, Fez03 = 7.8%, MgO = 26.5% and A1203 = 0.0%. (Actually, the Alz03 calculates as a small negative value.) If all the A1203is assumed to be in the palygorskite, the percent of palygorskite calculates to be 87.4% rather than the 90% value based on X-ray data. If the 87.4% is used, the composition of the remaining material is: SiOz = 68.3%, Fe203= 7.4%, and MgO = 24.3%. The structural formula is Xo.33(Mg2.17Feo.33)(Si4.08)010(OH)z. While this formula probably does not represent a specific mineral and is the result of considerable data manipulation the calculations suggest that the smectite present in these palygorskite-rich clay beds is not the detrital montmorillonite but is an Al-free residuum resembling stevensite, that remains from the conversion of Al-smectite to Al-palygorskite. Stevensite is commonly formed in a lagoon-lacustrine environment associated with dolomite and palygorskite (Tardy et al., 1972). Various graphs can be plotted to demonstrate the relation between the two types of clay minerals. A plot of MgO vs A1203(Fig. 103) shows a linear relation with the pure montmorillonites having an AlZO3/Mg0ratio ranging

TABLE IV 0

Structural formula Montmorillonite EC-2-32

Palygorskite

Mixed

F5-13

FF-1

FH-12

FF-18

FE-21

FE-17

FE-80'

FF-54'

MC1-5

MCI-23

1.43 0.33 0.25 2.01

1.46 0.28 0.28 2.02

1.51 0.25 0.22 1.98

1.24 0.42 0.41 2.07

1.66 0.36 1.83 3.85

1.58 0.40 1.79 3.77

0.79 0.55 2.23 3.57

1.20 0.38 2.15 3.73

1.32 0.43 2.08 3.83

1.33 0.38 2.01 3.72

0.11 3.89

0.11 3.89

0.03 3.97

0.11 3.89

0.19 7.81

0.40 7.60

0.71 7.29

0.47 7.53

0.39 7.61

0.36 7.64

0.01 0.45

0.24 0.40

0.57 0.20

0.53 0.39

0.31 0.39

0.52 0.17

Octahedral

Al Fe3+ Mg

x

1.49 0.24 0.28 2.01

Tetrahedral A1 Si

0.09 3.91

Exchange cations Ca Na

' P > 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-

197

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

199

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

200

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

202

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.

203

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

204

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

206

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.

209

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).

211

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.

214

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

215

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.

217

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,

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[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-

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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-

221

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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