Soils of the Past An introduction to paleopedology
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Soils of the Past An introduction to paleopedology
Dedicated to Ken and Wendy Retallack,for letting me be
Soils of the Past An introduction to paleopedology
Gregory J. Retallack University o j Oregon Eugene USA
SECOND EDITION
b
Blackwell Science
02001 GregoryJ. Retallack Published by BlackwellScience Ltd Editorial Offices: Osney Mead, Oxford OX2 0FL 25 JohnStreet.LondonWClN2BS 2 3 Ainslie Place, Edinburgh M3 6AJ 3 50 Main Street,Malden MA02148-5018,USA S4University Street, Carlton Victoria 3053, Australia 10,rue CasimirDelavigne 75006 Paris, France Other EditorialOffices: Blackwell Wissenschafts-VerlagGmbH Kurfurstendamm 57 10707Berlin, Germany Blackwell Science KK MG Kodenmacho Building 7-1 0 Kodenmacho Nihombashi Chuo-ku.Tokyo 104.Japan Iowa State UniversityPress A Blackwell ScienceCompany 2 12 1S.State Avenue Ames, Iowa 50014-8300,USA The right of the Author to be identifiedas the Authorof this Work has been asserted in accordance with the Copyright,Designs and PatentsAct 1988. Al1rightsreserved.Nopartof this publication may be reproduced, storedinaretrievalsystem. or transmitted. in any formor by any means. electronic, mechanical, photocopying, recording or otherwise. except as permitted by the IJK Copyright,Designs andpatents Act 1988. without the prior permission of thecopyright owner.
First published 1990 by Unwin-Hyman Reprinted 1991by Harper CollinsAcademic Secondedition 2001 by Blackwell Science Ltd Setby Best-setTypesetter 1,td..Hong Kong Printed and Bound in Great Britain at the Alden Press Ltd, Oxford and Northampton
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Marston Book Services Ltd PO Box 269 Abingdon.Oxon OX144YN (0rders:Tel: 01235 465500 Fax: 0123 5 465 5 5 5) USA Blackwell Science. Inc. CommercePlace 350MainStreet Malden. MA02148-5018 (Orders:Tel: 800 759 6102 781 388 8250 Fax:7813888255)
Canada Login Brothers Book Company 324 Saulteaux Crescent Winnipeg, Manitoba R3J 3T2 (Orders: Tek2048372987) Australia Blackwell Science Pty Ltd 54University Street Carlton. Victoria 3053 (Orders:Tel: 3 93470300 Fax: 393475001) A cataloguerecord for this title is available from the British Library ISBN 0-632-053 76-3 Libraryof Congress Cataloging-in-publication Data Retallack. Greg J. (GregJohn), 1951Soils of the past. an introduction to pdleopedology/ Gregory J. Retallack. -2nded. p. cm. Includes bibliographical references and index. ISBN 0-632-053 76-3 1. Paleopedology. I. Title. QE473.R473 2000 5 5 2 's -dc2 1 00-04544s For further information on Blackwell Science,visit our wehsite: www.blackwell-science.com TheBlackwell Sciencelogois a trade mark of Blackwell ScienceLtd, registered at the United Kingdom Trade Marks Registry
Contents
Preface to the second edition, vii Preface to the first edition, ix Acknowledgments,xi
Part 1: Soils and paleosols
Zeolitiiation and celadonitiation of volcanic rocks, 9 7 Coalificationof peat, 97 Kerogenmaturation and cracking, 98 Neomorphism of carbonate, 99 Metamorphism, 100 Common patterns of alteration, 100
Paleopedology,3 Soilson and under the landscape, 7 Soils and paleosols on the landscape, 7 Quaternarypaleosols, 9 Paleosols at major unconformities, 1 0 Paleosols in sedimentary and volcanic sequences, 11
Featuresof fossilsoils, 13 Root traces, 13 Soil horizons, 19 Soil structure, 24 Soil-formingprocesses, 3 7 Indicators of physical weathering, 3 7 Indicators of chemical weathering, 41 Indicators of biological weathering, 50 Common soil-forming processes, 5 7 Soil classification,63 FA0 world map, 63 US soil taxonomy, 64 A word of caution, 76 6 Mapping and naming paleosols, 77 Paleoenvironmental studies, 77 Stratigraphic studies, 8 3 Deeply weathered rocks, 86 7 Alteration of paleosolsafter burial, 87 Burial decompositionof organic matter, 89 Burial gleization of organic matter, 9 0 Burial reddening of iron oxides and hydroxides, 90 Cementation of primary porosity,9 1 Compactionby overburden, 93 Illitization of smectite. 95
Part 2: Factors in soil formation
8 Modelsof soil formation, 105
9 Climate, 108 Classificationof climate, 109 Indicators of precipitation, 112 Indicators of temperature, 11 8 Indicators of seasonality, 120 Indicators of greenhouse atmospheres, 12 5 10 Organisms, 128 Tracesof organisms, 129 Tracesof ecosystems, 145 Fossil preservation in paleosols, 153 11 Topographicrelief as a factor, 160 Indicators of past geomorphologicalsetting, 160 Indicators of past water table, 164 Interpreting paleocatenae, 166 12 Parent materialas afactor, 171 General properties of parent materials, 172 Some common parent materials, 176 A base line for soil formation, 179 13 Time as a factor, 183 Indicators of paleosol development, 18 5 Accumulation of paleosol sequences, 194
V
vi
Contents
Part 3: Fossil record of soils 14 A long-termnaturalexperiment in pedogenesis, 207 15 Soilsof otherworlds,209 Soilsof theMoon, 209 Soils of Venus, 2 1 3 Soilsof Mars, 2 16 Meteorites, 220 Relevancetoearly Earth, 225 16 Earth’s earliestlandscapes, 22 7 Oxygenation of the Earth’s atmosphere, 2 3 1 Differentiation of continental crust, 2 39 Precambrian scenery, 242 17 Early lie on land, 246 Did life originate in soil?248 Evidence for early life in paleosols, 2 56 MotherEarthor heartof darkness?260 18 Large plants and animalsonland, 263 Evidence of multicellularorganismsin paleosols, 2 6 5
How did multicellular land organisms arise?2 72 Biological innovation or environmental regulation?275 19 Afforestationof theland, 280 Early forest soils. 2 82 A diversifying landscape, 2 85 A finer web of life on land, 29 1 Theshapeof evolution, 298 20 Grasses in dry continental interiors, 300 Early grassland soils, 303 How did grasslandsarise?3 12 Evolutionaryprocesses, 3 1 4 2 1 Human impact on landscapes,31 7 Humanorigins, 320 Early human ecology, 324 Atamedlandscape, 328
Glossary, 3 34 References, 3 52 Index, 3 9 5
Preface to the second edition
In the years since first publication of this book, paleopedology has grown from a long childhood into gawky adolescence.Paleopedology’sinfancy was well captured on the opening page of Vladimir Nabokov’s (1955) Lolita, in which it is offered as the epitome of an obscure scientiic interest of Humbert. Now, it is no longer a surprise to find a paleosol or a paleopedologist. Emphasis now is on interpretation of large suites of paleosols, for example, tracking past fluctuations in atmospheric carbon dioxide from the isotopic composition of carbonate nodules in paleosols. Such isotopic studies of paleosols demonstrate that they really were soils of the past. Their message about former environments and ecosystems goes beyond their surface appearance. The study of these remarkable rocks is now in a phase limited mainly by human ingenuity. Isotopy, cathodoluminescence, magnetic susceptibility, X-radiography and microtomography are opening new vistas into the formerly hidden world of paleosols. The first edition of this book was mainly ideas and questions. This edition is devoted more to procedures
and answers. The way of paleopedology is currently being mapped out on several fronts. Global change, coevolution, mass extinctions and comparative planetary geology are some of the currently important topics informed by paleosols. In pursuit of these broader objectives, procedures for recognition and study of paleosols are becoming routine. Much of the first edition outlining such procedures has now been consigned to tables. I have also written another book (Retallack 1997a) as a source book of terminology and procedures for professionals. Here, however, emphasis remains on what paleosols can tell us of the way the world works. The theory and issues of paleopedology continue to grow in the quirky, sometimes upsetting and sometimes inspiring, manner of adolescence. In another 10 years, perhaps the field will have settled into comfortable middle age. For the moment, however, as the Chinese proverb has it, we live in interesting times. GregoryJ. Retallack Eugene, Oregon, 2000
vii
Preface to the first edition
Landscapes viewed from afar have a timeless quality that is soothing to the human spirit. Yet a tranquil wilderness scene is but a snapshot in the stream of surficial change. Wind, water and human activities constantly reshape the landscape by means of catastrophic and usually irreversibleevents. Much of this change destroys past landscapes, but at some times and places, landscapes are buried in the rock record. This work is dedicated to the discovery of past landscapes and their life through the fossil record of soils. A long history of surficial changes extending back almost to the origin of our planet can be deciphered from the study of these buried soilsor paleosols. Somerudiments of this history, and our place in it, are outlined in a final section of this book. But first it is necessary to learn something of the language of soils, of what happens to them when buried in the rock record, and which of the forces of nature can be confidently reconstructed from their remains. Much of this preliminary material is borrowed from soil science, but throughout emphasis is laid on features that provide most reliable evidence of landscapes during the distant geologicalpast. This book has evolved primarily as a text for senior level university courses in paleopedology: the study of fossilsoi1s.Itisnottheusualviewof thissubjectfromthe perspectivesof soil science,Quaternaryresearchor land use planning. It is rather the view of an Earth historian and paleontologist. Compared with the elegant outlines
of a fossil skull or the intricate venation on a fossil leaf, fossil soils may at first appear unprepossessing subjects for scientific investigation. These massive, clayey and weathered zones are fossils in their own way. Their identification within a classification of modern soils presupposes particular past conditions, in the same way as the lifestyle that can be inferred from modern relatives of a fossil species of skull or leaf. Particular features of paleosols also may reflect factors in their formation in the same way as ancient diet can be inferred from the shape of fossil teeth, or former climate from the marginal outline of a fossil leaf. This book is an exploration of the idea that paleosols are trace fossils of ecosystems. Examplesin this book are drawn largely from my own work on fossil soils, some of it not yet published elsewhere. Theoretical concepts have been borrowed more widely from allied areas of science including geomorphology, coal petrography, plant ecology, astronomy and soil science,to name a few. The fossil record of soilsis a new focus for integrating existing knowledge about land surfaces and their biota. Paleopedology remains an infantdiscipline,hungry for theory anddata of the most elementary kinds. This book is one attempt to partially quell the growingpains. GregoryJ. Retallack Eugene, Oregon, 1989
ix
Acknowledgments
This book on paleosols would be slender indeed without extensive borrowing of facts, experiments, ideas and inspiration from allied areas of science. I have been fortunate to be able to draw upon the wise counsel of prominent sedimentologists(J.R.L. Allen, A. Basu, D.R. Lowe, R.M.H. Smith, E.F. McBride andL.J.Suttner),paleontologists (R. Beerbower,A.K. Knoll,J.W. Schopf andR Shipman), geochemists (G.G. Goles, J.M. Hayes, H.D. Holland, W.T. Holser and T.E. Cerling) and soil scientists (P.W. Birkeland, S.W. Buol, L.D. McFadden, P.F. McDowell, L.R. Follmer. D.L. Johnson and A.J. Busacca). Among the emerging cadre of paleopedologists concerned with rocks older than Quaternary it is a pleasure to acknowledge stimulating discussions with D.E. Fastovsky, M.J. Kraus, W.R. Sigleo, VTP. Wright and S.G. Driese. Last and certainly not least, many of my ideas
have been reshaped by students at the University of Oregon (E.A. Bestland, D.P. Dugas, C.R. Feakes, P.R. Miller, J.A. Pratt, S.C. Radosevich, G.S. Smith, G.D. Thackray, E.S. Krull, J.G. Wynn and N.D. Sheldon). They gave real meaningtotheSocraticdictumthat theunexaminedlife is not worth living. Photographs, specimens and other illustrative materials were generously and promptly provided by H.J. Anderson, J.B. Adams, T.M. Bown, J. Gray, R. Greeley, M.J. Kraus, S.C. Morris, NASA Space Science Center and C.J. Percival. Others graciously acquiesced in my adaptation of their published work. For several fine photographs, I thank Sean Poston. The writing would not have proceeded nearly so much to my satisfaction without the happy home life created by Diane, Nicholas and Jeremy
xi
Part 1
Soils and paleosols
~~
2200 million year oldpaleosol (lightcoloredzone) near Waterval Onder, SouthAfrica.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 1 Paleopedology
Paleopedologyisthe study of ancient soils,andis derived from an ancient Greek word ( m b v ,m h v )for ground. It has nothing to do with pedestrians (Latinpes,pedis) or . of the past, pediatricians (Greek R C ~ ~~S ,a t h s )Soils either buried within sedimentary sequences or persisting under changed surface conditions, are the main subject matter of paleopedology. In this book, it is seen as a historical perspective on soil genesis and as a way of reconstructing the geological history of land surfaces on Earth. Soils, like organisms, sediments and surface environments, have changed over the past 4566 myr of recordedEarth history. The concept of fossil soils can be traced back to the Scottish physician James Hutton (1795). His insistence on arguing past causes from those that can be observed today was a prerequisite to makiig a connection between soils of today and those of the distant geological past. Red rocks along angular unconformities exposed along the course of the River Jed and at Siccar Point, southeast of Edinburgh, he regarded as comparable to surface soils and sediments on the modern landscape (Fig. 1.1).’From this it will appear, that the schistus mountains or vertical strata of indurated bodies had been formed, and had been wasted and worn in natural operations of the globe, before horizontal strata were begun to be deposited in these places. . .’ (Hutton 1795, Vol. 1, p. 438). These ideas were reiterated in John Playfair’s (1802) Illustrations of the Huttonian Theory of the Earth, which, because of its conciseness and clarity of expression, was more influential than Hutton’s original two volumes. Playfair also cited a 1799 record of a fossil forest in Lincolnshire, now covered by tidal flat sediments. This is the oldest record of a Quaternaryfossil soil. The oldest record of buried soils within consolidated sedimentary rocks was the ‘dirtbeds’ (Fig. 1.2)and fossil stumps reported in latest Jurassic limestones of the Dorset coast by Webster (1826) and popularized by Buckland (1837) in a volume widely known as the
Bridgewater Treatise. Other fossil forests were discovered in the late 19th century, and their stumps and associated fossil plants described, but little was made of their substrates as fossil soils. Examples include the Eocene Sequoia forests of Yellowstone National Park, USA, and Carboniferous stumps of tree-lycopsids at Clayton (Yorkshire)and in Victoria Park (Glasgow),both UK. A summary of these early discoveries of pre-Quaternary fossil soils was given in introductory chapters of Albert C. Seward’s monumental work Fossil Plants (1898). He appreciated the significanceof fossil soils as evidence for the immensity of geological t i e and as indicators of past worlds. Study of the paleosols themselves had to await the development of soil science. During the late 19th century, buried soils also were recognized w i t h i sur6cial deposits of loess and till. These ‘weathered zones’, ‘forest zones’ and ‘soils’, as they were variously termed, were found in Russia by Feofilatkov (in the 187Os, as recounted by Polynov 1927), in the midcontiiental USA by McGee (1878). andinNewZealandbyHardcastle(1889).Bytheturnof the century such observations had been used for stratigraphic subdivision of glacial deposits (Chamberlain 1895). Despite these discoveries,the origin of paleopedology as a discretefield of inquiry can be traced back to the late 19th century development of soil science (Tandarich 81 Sprecher 1994). Since classical times, soils have been studied from the point of view of plant nutrition. It was not until 1862 that the Saxon scientist Fredrich A. FalIOU 6rst published the term ‘pedologie’for the study of soil science, as opposed to what he termed ‘agrologie’,or practical agricultural science. The foundations of modern soil science were laid by Vasily Dokuchaev with a detailed account of the dark, grassland soils of Russia (1883).This monograph demonstrated that soils could be described, mapped and classified in a scientific fashion. Furthermore, their various features could be related to environmental constraints, of which climate and
3
4
Chapter1
Figure 1.1 Angular unconformity between Lower Silurian (430 Ma) Hawick Rocks and Upper Devonian (360 Ma) Upper Old Red Sandstone along the River Jed in southeastern Scotland (from Hutton 1795).
vegetation were consideredespeciallyimportant. By the early part of the 20th century there was an established scientifictradition of research on soil geography,classification and genesisin Russia, as summarized in the influential general works of K. D. Glinka (192 7). In the course of these early Russian investigations, certain soils were found to be anomalous in that their various features did not fit the general relationship between soil type and their climate and vegetation. It had long been suspected that these were very old soils, perhaps products of past environments. In 1927, Boris Polynov summarized Soviet observations of this kind. His short paper, which introduced the term paleopedology, can be considered the foundation of this branch of science. Polynov included the study of four kinds of
POI llund
materials within paleopedology. ‘Secondary soils’ encompass those formed by two successive weathering regimes, such as grassland soils degraded by the advance of forests after retreat of glacial ice. ‘Two-stage soils’were recognized to have an upper horizon of recent origin, but deeper horizons of more ancient vintage. ‘Fossil soils’ were defined as soil profiles developed on a surface and subsequently buried. Polynov’s final category of ‘ancient weathering products’ included the redeposited remnants of soils such as laterites and china clays. Polynov and his colleagues established a logical framework for the paleoenvironmental study of paleosols of all ages that continues in Russia. Modern soil science in North America can be traced back to Eugene W. Hilgard’s (1892) monograph
Afururtre furmoilon.
Sectioii of the Drrt-led in the Isle of Portland, thewing the rrbterraneata remuitcs of’un a n h i t Forest. D e la Berhr.
Figure 1.2 ‘Dirtbeds’ (paleosols)in a stratigraphic section through the uppermost Jurassic (Tithonian, 150Ma) Purbeck Formation on the Isle of Portland, Dorset, England(fromBuck1and 1837).
Paleopedology
on the relationship of soil and climate. This, and the first Large-scale mapping and classification of North American soils by Milton Whitney (1909),were largely independent of comparable research carried out by Russian soil scientists. Soviet influence first appeared in the work of Curtis Marbut. particularly his monumental soil survey of the USA (1935). Paleopedology also was introduced into North America through a Soviet connection. Constantin Nikiforoff completed doctoral studies at the University of St Petersburg in prerevolutionary Russia, but by 1943 was a scientist with the US Department of Agriculture (USDA) Soil Conservation Service,when he published a short essay outlining the role and scope of paleopedology. A supporting study of paleosols in the same journal of 1943 by Kirk Bryan and Claude Albritton made clear the practical application of such studies. Ideas on the classification and origin of soils have been especially useful for studies of Quaternary stratigraphy and geomorphology. Such studies are now conducted in most parts of the world. coordinated by a Commission on Paleopedology established in 1965 at the Seventh Congress of the International Association for Quaternary Research, in Denver, USA. An early result of the commission’s activities was the publication, in a volume of research papers edited by Dan H. Yaalon (1971). of recommendations for recognizing and classifying paleosols. Mapping units for Quaternary paleosols have been incorporated into official stratigraphic codes (e.g. North American Commission on Stratigraphic Nomenclature 1982). Modern research on Quaternarypaleosols can be found in books and journals on soil science, geography, archeology and Quaternaryresearch (Catt 1990:Holliday 1992,1994: Follmeretal. 1998). In contrast to a steady level of interest in Quaternary paleosols, studies of older paleosols have been slow to gain momentum. In many sequences now known to contain them in abundance, paleosols were not recognized or their features were explained as diagenetic phenomena. Little was made of those few cases where paleosols were explicitly recognized (Barrel1 191 3: Collins 192 5: Allen 194 7; Thorp & Reed 1949).Beginning in the 1960s, interest in pre-Quaternary paleosols has been increasing on several fronts, as they were discovered in many nonmarine sedimentary sequences (Retallack 1997a) and even in deep-sea cores (Ford
5
1987; Holmes 1992). The study of paleosols is especiallycompatiblewith the overall aims of sedimentology to reconstruct ancient environments and geological processes.Paleosols now regularly appear in accounts of sedimentary geology (Esteban & Klappa 1983; Wright 1986a) and of weathering processes (Martini & Chesworth 1992;Ollier&Pain 1996). Paleontological research has always been concerned with reconstructing past biotas, their ecology and the ways in which they are preserved in the rock record. Paleosols can be regarded as both trace fossils of past ecosystems and as preservational environments formanykindsof fossils(Reta1lack1975,1976,1977). Paleontologically oriented accounts of paleosols are appearing now in a variety of books and journals concerned with paleoecology, paleoclimatology, and other paleontological subjects (Retallack 1991a; Parrish 1998;Stanley 1998). A final area of geological sciences now contributing to paleopedology is geochemistry.The chemical study of weathering by Samuel S. Goldich (1938), from which was derived his well-known mineral stability series, dealt with Cretaceous paleosols in Minnesota. Paleosols also were used as indicators of atmospheric conditions and the nature of weathering processes in the very distant geological past (Sharp 1940; Sidorenko 1963). Such studies are now more common in journals and books concerned with geochemistry and Precambrian geology (Holland 1984;Schidlowskietal. 1992). Looking toward the future, fossil soils may be used as stratigraphic markers in the continuing inventory and mapping of the geological resources of this and other planets. Particular features of paleosols may aid in locating especially valued resources. For example, changes in degree of development and mineral content of paleosols reflect former time for formation and degree of waterlogging, and can be used to guide exploration for petroleum, coal and uranium ores. Coal and uranium accumulated in parts of the landscape where groundwater was poorly oxygenated. Uranium is more likely to be found in and near paleosols that are variegated and that formed on the margins of uplands of uraniferous granites (Kimberley1992).Coal is associated with paleosols that are gray and formed in swamps, rather than paleosols that are red and were formerlywell drained (Retallack & Krull 1999). Paleochannei sandstones, located by following lateral trends in paleosol
6
Chapter1
development and waterlogging, can be local petroleum reservoirs (Kraus&Bown 1993). Fossil soils also provide historical validation for theories about how soils form. The geological history of soils can be viewed as a long-term natural experiment in which many fundamental conditions of soil formation, such as vegetation and atmospheric composition, have changed. Information from fossil soils can strengthen ideas about how soils form and how they should be classified. Because such ideas form the basis for much agricultural and engineering activity and their modification
in the face of accelerating global change, it is all the more important that they have a 6rm scientific basis (Retallack1996a). Finally, fossilsoilsare evidencefor reconstructing past terrestrial ecosystems and environments. They can be used to bring particular times and places into sharper focus as evidence independent of associated fossils and sedimentary structures. They are a record of the evolution of ecosystems and of their interaction with environments on land, and provide a perspective on our placeonEarth.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 2 Soils on and under the landscape
Difficultiesin defining ‘fossilsoil’ arise not so much from its fossilnature as from confusion over what is meant by soil. A fossil soil or paleosol, l i e other kinds of fossils, is the remains of an ancient soil. It may have been buried by later deposits or it may be at the surface but no longer actively forming in quite the same way. The word soil on the other hand, like other commonly used words such as love and home, means different things to different people. For farming purposes, soil is fertile, loose, tillable ground. In engineering specifications,soil is anymateria1 that can be excavated without recourse to quarrying or blasting. By both these practical definitions, unaltered sediments, such as dune sands or flood silt, are regarded as soil.Yet some extremely altered soil materials, such as hard laterites, are not. To some soil scientists, soil is the medium in which vascular plants take root (Buol et al. 199 7). This narrow definition includes rock crevices supporting large plants. Yet it excludes rocks colonized by plant-like microbes and rootless plants such as mosses and liverworts. It also excludes hard-setting parts of soils and parts of soilsbelow the level of roots (Hole 19 8 1). Taking an even wider view, what are we to call lunar, Venusian and Martian land surfaces that have been altered in place by surficial processes?The Moon and Venus are lifeless. Mars probably is also. Should calling their altered land surfaces soils necessarily imply that life is present? Or should a n alternative label such as regolith imply that they are barren of life? Similar problems beset naming fossil soils older than Ordovician, before the advent of large land plants. If one takes the view that soils are a medium of plant growth, then both the terms Precambrian soil and Precambrian regolith beg the question of the antiquity and nature of life on land. Like the question of life on Mars, the origin of life on land is an important scientific issue in its own right and should not be confused by semantic considerations. For paleopedology a wider delinition of soil is needed. In this book I take soil to be material forming the surface
of a planet or similar body and altered in place from its parent material by physical, chemical or biological processes. This is close to Nikiforoff’s (1959) concept of soil as the ‘excited skin of the subaerial part of the Earth’s crust’, and its more functional description by Colin et al. (1992)as a ‘geomembrane filter’. Soil can be envisaged as a zone of interaction between the atmosphere and crust of a planetary body. A flux of energy from the Sun, water, snow or living creatures contiinually alters rocks or sediments into what we call soil. A part of the problem with defining soils is their intrinsically intermediate nature. Soils are complex zones of interaction between sediment or solid rock and the ecosystem or atmosphere. Because of the varying levels of interaction down from the land surface, it is usual to study soils in profiles that show the layering of alteration, or soil horizons. Solum is a technical word for that part of the soil profile most altered by soil processes (Fig. 2.1).In many cases it is riddled with roots of plants. The soil solum also may be dark, red, clayey or massive, and so is very different from its underlying parent material. Weathered material between the solum and underlying sediment or bedrock has a mix of soil and inherited features, and is called saprolite, or alterite (Nahon 1991). Saprolite may be soft, oxidized, clayey or otherwise altered l i e a soil, but not to the same extent as the soil solum. Some saprolites show clearly the bedding, schistosity and deformation of parent rock. The distinction between saprolite and solum is a matter of degree of alteration.
Soils and paleosols on the landscape Soils blanket most of the landscape except for areas covered by rivers or lakes or areas freshly uncovered by erosion and human excavation. Conditions of sunshine, moisture and other soil-forming factors vary in different parts of the landscape, and so do the soils forming there. The fundamental unit of soilis acolumn of soilmaterial,
7
8
Chapter2
Figure 2.1 Technical terms for soils and their relationshipto landscapes.
or pedon, of the kind that could be dug out of the wall of a trench. Soils vary in such complex ways that few pedons are exactly alike. Some pedons are sufficiently similar that they are recognized as a discrete kind of soil, different from others in the area. One or more of these similar pedons covering an area of ground is called a polypedon (Soil Survey Staff 1993). These are like tiles in a mosaic of soils over the landscape. The assemblage of polypedons mantling the landscape is called a soilscape. Soils form over many years until covered by sediment or removed by erosion. In river valleys, for example, deposits of sand and silt left behind by an especially powerful flood, or the rubble of a landslide,may cover soilsand drive off or destroy plants and animals (Fig. 2.2). Deep burial is a common way in which paleosols are formed. The covering deposits provide a surface for recoloniza-
relict soil
pedol i t h
ll
II
A
n
sedimentary relict
exhumed soil
1
buried soils
-
Figure 2.2 Technical terms for fossilsoils (paleosols)and their
relationto sediments(fromRetallack 198 3a; reprintedwith permission from the GeologicalSociety of America).
Soils on and under the landscape tion by plants and animals. and soil begins to form anew. Soon the bare surface is dry and cracked and small plants are taking root. With the advent of plants come worms and other burrowing invertebrates as well as herbivorous mammals. In humid regions, these early successional plants and animals are followed by shrubs and ultimately there is reconstituted the kind of woodland and soil that existed before. As a soil is burrowed, penetrated by roots and otherwise altered, the ripple marks and bedding planes of original alluvium are progressively destroyed. Such sedimentary relicts persist in many weakly developed bottomland soils and paleosols. Their persistence is a reflection of the degree to which the soil has been able to form and does not negate the identification of these materials as soils or paleosols. It is not difficult to recognize that paleosols developed during times of peace and quiet between deposition of thick sedimentary layers. Sequences of paleosols can become very difficult to decipher, however, if the intervening sedimentary layers are too thin to separate them effectively (Woida & Thompson 1993: Olson & Nettleton 1998). A common situation that may cause confusion is overprinting of the upper horizon of an older soil by development of a lower horizon of a younger soil with a slightly higher surface. The older near-surface horizon, then, is not genetically related to the younger subsurface horizon. From the point of view of the younger soil,it can be considered a pedorelict of the older soil. Other common pedorelictsare nodules or clasts of pre-existingsoilswithin sedimentson which later soilshave formed.Distinguishing these older nodules or clasts from nodules or clods of the younger soil may be difficultunless they have sharp, ferruginized or truncated boundaries. Some sedimentary layers are very distinctivebecause they are composed entirely of a particular kind of soil material. Such a bed, withclasts of soil minerals and appearing l i e a soil but with Sedimentary organization, may be called a soil sediment (Catt 1998) or pedolith (Gerasimov 1971). The term pedolith was originally coined by Erhart (1965) for redeposited laterites of Tertiary age. These remain a good example because such locally derived soil material forms brightly colored red beds distinct from enclosing alluvium. Pedolith is useful only for such clear cases, where soil-derived sediments are distinct, because most sediment ultimately is derived from soils and so is pedolithic in some sense.
9
Surface soils can also be considered relict paleosols if they are of a kind completely different from those forming under prevailing conditions. Good examples of relict soilsare lateritic and bauxitic paleosolsremaining at the surface in the deserts of Australia from early Tertiary deep weathering (OIIier & Pain 1996). The true soils of deserts are not lateritic, bauxitic, red, clayey and deeply weathered, but thin, sandy and little altered from the color of their parent materials: as can be seen from soils of eastern Oregon, southern California, Nevada, the Middle East and Central Asia (Birkeland 1999).The term relict paleosol should be used only for such clear anomalies, because environmental change can be rapid, and most soils can be considered relicts of different past conditions to some extent.
Quaternarypaleosols By current estimates (Van Cowering 1997), the Quaternary geologicalperiodisthepast 1.8myr.Thousands to millions of years are the time spans it takes to form soils, so those of Quaternary age are important evidence for the way in which soils form. The study of factors in soil formation usually involves a carefully constrained analysis of soils and paleosols of varying age or situation. In the earthquake-prone Transverse Ranges of California,river terraces are uplifted and tilted by folding that is still continuing (Fig. 2.3A). The youngest surfaces are those nearest the streams, where they are still disturbed by annual floods. Older terraces, dated by radiocarbon and other means, are at higher levels. This stepped landscape includes successively older surfaces at higher levels and these in turn bear progressively better-developed soils (Fig. 2.3B). Many such sequences of alluvial terraces and their soils have been studied to document changes in soil formation with t i e (Harden 1990). Another favorite landscape for such studies is areas around the terminus of retreating glaciers (Birkeland 1992). This kind of research provides basic information about the way in which soils form that is vital to the interpretation of paleosols. Studies of Quaternary soils and paleosols reveal clearly what complex things soils are and how many factors enter into their formation (Johnson & Watson-Stegner 1987; Phillips 1993). Not all these complications are relevant to interpretation of older paleosols. Among these are human impacts such as increased incidence of
10
Chapter2
Figure. 2.3 Tectonicallytilted Pleistocene (c. 160-200Ka) terraces disrupted at CulbertsonFault (Rockwellet al. 19 8 5 ) on left skyline and theTertiaryrocksof SantaPaula Ridge in centralTimberCanyon,north of Santa Paula, California,USA (A), and a thick, well-developedsoil (TypicPalexeralf in US taxonomy)on the centraltilted terrace (B): soil horizons are indicated by standard shorthandand tape is 2 mlong.
fires,forest clearance and paving. Some of the difficulty in understanding Quaternary paleosols also has been exacerbated by focusing study on those found in outcrops or shallow trenches rather than in deep boreholes. Paleosols of uplifted terraces or stable continental regions are likely to have been influenced by a greater variety of weathering regimes than those subsiding below the zone of weathering shortly after formation.
Paleosols at major unconforrnities Many unconformities show evidence of paleosols. Not alldo because some have been scoured clean by fluvialor marine erosion before being covered by later sediment. Examination of geological maps for unconformities remains a productive method for locating paleosols, especially in Precambrian rocks, where they are difficult to recognize otherwise. Paleosols at major geological unconformities include thick, well-differentiated layers of rockenriched in ferricoxide(laterite),in alumina (bauxite), in silica (silcrete)or in calcium carbonate (calcrete). The origin of these distinctive materials is a complex issue for at least two reasons. First, they take so long to form that conditions originally encouraging their formation are almost certain to have changed in some way before their burial and preservation (Vasconceloset al. 1992: Chadwick et al. 1995). Second, these are all indurated and weather-resistant materials that can withstandsubsequent erosionalevents (Thiry 1999;Valeton 1999).A brief consideration of some of the leading theories for the formation of one kind of duricrust servesto illustrate some of these complexities. Thick (lorn) laterites are thought not to form within the soil solum, but within deeper and thicker zones of
saprolite below (Oilier & Pain 1996). Especially appropriate sites for the accumulation of ferric oxides to such a concentration are places on the side of plateaux where groundwaters enriched in iron dissolved in swamps within depressions on the plateau are oxidized within a zone of seepage around the plateau margin (McFarlane 1976). Lateritic profiles may reach considerable thickness in such geomorphologicalpositions (Fig. 2.4).Such zones of concentrated ferric oxide are soft and easily excavated within the ground, but once exposed to air they become indurated like a brick.Their excavation and drying for construction stone on the Indian subcontinent is the source of their name from the Latin later (lateritis in genitive) for brick. Indurated laterites may armor hillsides against further erosion or may persist as pebbles of conglomeratic material similar to pisolitic or nodular original laterite (Tardy & Roquin 1992). Once formed, laterites are very persistent. From this brief account of laterites, they can be seen to involve more than just soil formation. Erosional landscape lowering, reorganization of pre-existing soil horizons, changing flow of groundwater and progressive modification of the landscape also occur (Valeton 1999). The overlapping effects of so many processes make interpretation of such old surfaces of weathering difficult and controversial. Interpretation of buried examples is confounded by additional difficulties. Along major erosional unconformities it is difficult to be sure that the entire pre-existing profile has been preserved. Modern duricrusts form the surface of many landscapes because the soil under which they formed has been eroded. Many paleosols at major unconformities are likely to represent saprolite or other deep layers rather than the surface solum (Schau &Henderson 1983). A
Soils on and under the landscape
11
Figure 2.4 Laterite (darklower zone) with pockets of chinaclay (white) andmottled zone (variegated middlezone) of a Miocene paleosol (Plinthic Paleudult of US taxonomy) overlain unconformably by sand capped with a modern soil (Hapludollof US taxonomy) at north end of Long Reef beach near Sydney, Australia (laterites discussedby Faniran 1971).
second difficulty with unconformities is the way in which they juxtapose materials of very different chemical and physical characteristics. Commonly clayey impermeable paleosols are overlain by gravelly or sandy, permeable fluvial deposits. Passage of groundwater through overlying sediments could substantially alter the underlying paleosol with effects becoming less marked downward from the unconformity in a manner difficult to distinguish from former soil formation (Pavich& Obermeier 1985). Examples of formerlywelldrained paleosols that are mineralized with sulfide or uranium minerals characteristic of reducing environments (Mossman & Farrow 1992) are indications of such modillcations.
Figure 2.5 Alluvial-volcaniclastic sediments including a long sequence of superimposed paleosols of Eocene and Oligoceneage in the Pinnacles area of Badlands National Park, South Dakota, USA (from Retallack 1983a; reprinted with permission from the Geological Societyof America).
Despite these problems, paleosols at major unconformities often present soil formation so extreme as to be unmistakable. The accumulated alteration of ages is not easily erased by later events of lesser duration. As evidence of the antiquity and geological history of deep weathering and of duricrusts, they are of interest in themselves.
Paleosols in sedimentary and volcanic sequences Paleosols are abundant in some sedimentary and volcanic successions (Fig. 2.5). In many cases paleosols have masqueraded under nongenetic terms such as red
12
Chapter2
beds, variegated beds, tonstein, ganister and cornstone. Unlike conglomerates and sandstones, which are readily identiable as hardened gravel and sand, these other rock types do not resemble modern kinds of sediment.Their similarityto modern soilswas unrecognized as long as they were compared with sediments and rocks rather than soils. Ganisters, for example, are rocks found in Euramerican Carboniferouscoal measures. The word was coined by Cornish miners for hard, silicified quartz sandstone that is so chemically and physically inert that it is used for lining furnaces. Ganisters commonly are penetrated by carbonaceous root traces and underlie coal seams. It is now recognized that these were upper horizons of moderatelywell-drained soils.Their quartz-rich composition was produced in part by destruction of easily weathered associated minerals and their silicification by diagenetic mobilization of accumulated plant opal (Retallack 197 7 ;Gibling&Rust1992). Another kind of rock that seemed puzzling from a Sedimentaryperspective has been called cornstone: red rock riddled with yellowish nodules of calcium or magnesium carbonate, irregular to rounded in shape and several centimeters in diameter. Marine rocks commonly contain calcareous, sideritic and other kinds of nodules presumed to have formed after burial of the sediment (Boggs 1995), and cornstone also was thoughttobeproducedbyburialdiagenesis.Someof the classical cornstone sequences of the Old Red Sandstone in Britain contain fossil freshwater fish but no marine fossils. From this it could be argued that they formed after burial, but in lake sediments. Careful evaluation of the associated sandstones revealed that cornstones were more commonly associated with rocks thought to have formed in ancient rivers and that they resemble calcareous nodules of modern soils of dry climates (Allen 1986a).Carbonate nodules formed by burial diagenesis are less complexlycracked, less micritic and do not show displacive fabrics. The differences between calcareous nodules of marine rocks and of paleosols are now sufficiently well established that nodular paleosols can be
recognized as evidence of low sea level within shallow marinelimestones (Wright 1 982). Several general problems with the study of paleosols in sedimentary successions are apparent from these two examples. Foremost among these is their alteration upon deep burial. Cornstone was most easily understood because it is indurated and similar in appearance to modern soil nodules. Few modern soils are colored as brightly as red beds nor are many soils hard and flinty like ganisters. The likelihood of some changes after burial should not be taken to mean that all features of these rocks formed during burial. Careful attention to relationships with unquestionably original structures,such as fossil root traces and burrows, allows discrimination between original features of apaleosol and those formed during burial. A second problem is caused by the confused boundaries of some paleosols in sedimentary successions as a result of overlapping of successive paleosols. Erosion and redistribution of soil material also can be confusing. These complications can obscure the expression of soil horizons or other features that would be more obvious indications of paleosols. A third problem has been the application of inappropriate conceptual models to the interpretation of these rocks. Many nonmarine sedimentary rocks include evidence of soil formation as well as sedimentation. Each needs to be considered in interpreting geologicalhistory. Despite these problems, paleosols in sedimentary rocks are promising because of evidence for paleoenvironments that they contain. Evidence from fossil soils can be used not only to validate interpretations based on other lines of inquiry, but also to frame new kinds of interpretations. For example, fossil soils can be used as evidence of former vegetation against which can be assessed the degree of adaptation of limbs and teeth of associatedfossil vertebrates (Retallack 1991a,b).Because of their unique problems and potential, the study of paleosols in ancient sedimentary successionsis developing a research tradition of its own, distinct from that of paleosols at major unconformities and from that of Quaternary paleosols and soils.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 3 Features of fossil soils
Compared with cross-bedded sandstone or coarsely crystalline granite, paleosols at first sight may seem massive and featureless. Despite nondescript first impressions, paleosols do have distinctivefeatures. For the most part these are characteristics also found in modern soils. Yet many paleosols are no longer loose, cracked and at the land surface. Important differences result from compaction and alteration upon burial, which change many of the diagnostic chemical properties of modern soils such as pH. Eh and base saturation. Thus, identification of paleosols can be a problem for both geologists and soil scientists alike (Wright 1992a; Birkeland 1999). There are three main kinds of features by which paleosols may be recognized in the field and from laboratory studies: root traces, soil horizons and soil structure. Using these and other observations it is generally possible to distinguish paleosols from unaltered sedimentary deposits, volcanic flows or zones altered by faulting. Paleosols can be altered by groundwater, by hydrothermal activity or by metamorphism, and in these cases often have a mix of features difficult to disentangle. Alteration of paleosols after burial is the subject of a later chapter. This one is concerned with criteria to determine whether a rock is part of a paleosol or something else entirely. Root traces Fossil roots or root traces are one of the best criteria for recognition of paleosols in sequences of sedimentary rocks. They are evidence that plants once lived there, and that, regardless of the rocks other features, it was once a soil. A gray shale with clear bedding planes may look like an ordinary sedimentary deposit but the existence of a few fossil root traces in growth position means that it was once a soil.The fossil record of roots is now well documented (Pfefferkorn & Fuchs 1991; Bockelie 1994: Retallack 1997a; Elick etal. 1998) and
will be reviewed after considering general aspects of fossilroots. The top of a paleosol can be recognized as the surface from which root traces emanate (Fig. 3.1). Concentrations of other trace fossils such as burrows also can be used because they record periods of reduced or no deposition during which sediment was extensively modified at the surface. There are situations when sedimentation keeps pace with burrowing and vegetative growth, but irregularities in depositional processes are such that perfect balance between depositional disturbance and vegetative colonization is seldom attained. Usually there are zones of more than usual density of root traces and burrows that can be interpreted as horizons close to the top of a paleosol. Under favorable circumstances the original organic matter of a fossil root may be well preserved. Even if only a trace of roots is preserved by an infill of clay or calcite, there are several distinctive features by which they can be recognized. Unlike other trace fossils such as burrows, most root traces taper and branch downward. They also are very irregular in width. Large, nearvertical root traces characteristically have a concertinalike outline because of compaction of surrounding sediments. Outwardflexuresof theconcertinaare located at large lateral roots extending out into the matrix. Despite these characteristics the distinction between root traces and burrows is not always easy. Root mats may spread laterally over hardpans or around nodules. Some kinds of roots, such as the pneumatophores of mangroves, branch upwards and out of the soil (Jenik 1976). Furthermore, a range of soil invertebrates, especially ants, termites and worms, form complex branching burrow systems (Ratcliffe & Fagerstrom 1980) that may have collapsed in places. The distinction between root traces and burrowsis further blurred by soil invertebrates such as cicadas, which burrow around and into roots to feed on them (Retallack 1976). Another potentially confusing case is the preference of roots for the soft fill of fresh 13
14
Chapter3
Figure 3.1 Calcareousrhizoconcretions emanatingfrom the top of a mangal paleosol (Aquept)in the Upper Eocene (35 Ma) Birket elQarunFormationnearMadwar elBighal, Faiyum depression, Egypt (photograph courtesyof ThomasL.Bown).
burrows rather than hard clayey soil matrix (Retallack 1991a). Other structures that could be confused forroot traces include gas escape structures (Neumann-Makhau 1976) and tubular masses of soil fused by lightning strikes (Daly et al. 1993; Gifford 1999). These latter, called fulgurites, are lumpy masses of glass with exotic high-temperature minerals very different from ordinary soil matrix. Gas escape structures, such as those forming the conduit to sand volcanoes in alluvium coveringmethane-generating organic matter, are not so copiously branched or pervasive as root traces. In most cases, wispy tubular structures forming a n irregular, dense network within nonmarine rocks are root traces. One limitation on the use of root traces for recognizing paleosols is that they have not been found in rocks older than Silurian, when the first vascular land plants appeared (Retallack 1992a). There are burrows of invertebrates in paleosolsas old as Ordovician (Retallack& Feakes 1987). For paleosols older than mid-Ordovician root traces and burrows are of no use for identifying ancient soils.
Kinds of roots Fossil root traces are most easily recognized when their original organic matter is preserved. Paleobotanical research has now unearthed fossil examples of most of the major kinds of root now found (Retallack 1997a).
Roots are downward growing plant axes. They have numerous fine branches or rootlets (Fig. 3.2). Both roots and rootlets are anatomically simple, unlike the anatomical diversity seen in aerial parts of plants. Usually a central cylinder (stele) of elongate woody cells (tracheids) is separated by a zone (cortex) of equidimensional fleshy cells (parenchyma) from a tough outer rim of thick-walled cells (epidermis). The central woody cylinder and tough outer rim withstand decay longer than the intervening zone of soft cells. Partly decayed root traces may show a central dark woody streak and a carbonaceous epidermis separated by a zone filled with mud, calcite or other minerals where the cortex has decayed(Reta1lack1976).Withfurtherdecayeventheepidermal and stelar organic matter are replaced by other materials, but a concentric pattern of replacement may remain. Root hairs are individual elongate epidermal cells found in zones near the tips of fine rootlets. Because of their increased surface area compared with that of older parts of root systems, they are especially significant in gathering water and nutrients from the soil. Silurian and Devonian land plants, l i e living mosses and liverworts, lacked true roots, but had functionally comparable rhizoids. Both root hairs and rhizoids are so small and delicate that they are preserved only by exceptional circumstances, such as cellular permineralization (Retallack1997a). Various kinds of roots are distinguished by their
Features of fossil soils
15
Figure 3.2 Fossil root and rootlets from the Miocene MolallaFormationon High Hill near ScottsMills, Oregon,USA. Their original organic matter has been weakly ferruginized (RetalIack specimen R261).
pattern of branching and anatomical structure (Raven 1999).Many plants have a single, thick vertical rootortaproot, asinDevonianEddyu(Beck 1967).Carrots and parsnips are familiar modern plants that have tap roots modified into large underground structures for the storage of carbohydrates. Another kind of root system is seen in living grasses (Gramineae) and quillworts (Isoetes).These have fibrousroots radiating from a thickened stem base known as a corm or rhizomorph, as in Triassic Pleuromeiu (Retallack 1 9 9 7 ~ )If. the roots arise from the stem of a plant rather than its base, they are called adventitious roots. These may arise from rhizomes, which are stems lying in or along the ground. They also may anchor stems scrambling above ground (runners or stolons) as in modern strawberries and CarboniferousCullistophyton (Rothwelll975). Adventitious roots also form prop or stilt roots connecting erect stems and branches to the ground, as in Cretaceous Weichselia (Alvin 1971). In tree ferns, such as Carboniferous Psuronius (Morgan 1959). a very weak stem and leaf bases are completely enclosed by a mass of fine adventitiousroots.Theymay lookliketreetrunks, but these masses of roots and leaf bases are best called false stems. A variety of specialized structures of roots also are known from the fossil record. For example, tubers are underground storage organs branching from roots and rhizomes, as in the common potato and Cretaceous Equisetites(Watson&Batten 1990).Some plants of waterloggedsoilshaverootletsextending verticallyinto theair and these can be fossilized by deposition in swamp and intertidal habitats (Whybrow & McClure 1981).These
peg roots (pneumatophores) may play a role in allowing access to air for root respiration. For similar reasons, some plants of waterlogged habitats have thin-walled openings to theinside of the root (aerophores)or spongy parenchymatous tissue (aerenchyma). The most obvious of these aerating adaptations are large hollow cavities (lumina) found in roots of some swampland plants, such asPermian Vertebruriu (Retallack&Dilcher 1988). Rooting structures such as these not only indicate the existence of paleosols. but are also evidence of particular soil conditions. From the known fossil record of roots, most kinds of rooting structures have been in existence since Carboniferous time (Pfefferkorn & Fuchs 1991). Extinct woody plants whose aerial parts were very different from those alive today showed surprisingly modern kinds of roots. Presumably this is because functional constraints on root evolution have been more important than phylogenetic constraints. Useful paleoecological interpretationscan be made from fossil roots bycomparison with modern studies of root ecology and arrangement. Patterns of root traces While digging for fossil root traces it is useful to consider their arrangement, because this may provide evidence for former drainage, vegetation types and originally indurated parts of a paleosol. Because roots need oxygen to respire they seldom penetrate permanently waterlogged parts of soils. Laterally spreading (or tabular)
16
Chapter 3
Figure 3.3 Tabularrootsystemof alarge, extinct arborescentIycopod (Stigmaria ficoides) of Early Carboniferousage (Namurian.320 Ma) in the Lower Limestone Coal Group, Victoria Park, Glasgow,Scotland. The scale is 1 m for foregroundonIy.
root systems are characteristic of plants growing in swampy ground Uenik 1976) and are common among fossil stumps in sedimentary rocks of lowland environments (Fig. 3 . 3 ) . On the other hand, welldrained paIeosols may be deeply penetrated by root traces. Under wooded grassland, deeply penetrating and stout roots of trees and shrubs are scattered among a diffuse network of fine (<2 mm diameter) grass roots (van Donselaar-ten Bokkel Huinink 1966). In vegetation of drier climates the pattern of roots becomes shallower and more irregular as vegetation becomes more sparse andclumped. Under tallgrass prairie anetworkof
3 loll grass prairie
fine roots may extend as deep as 2 m below the surface (Fig. 3.4). Under short grass prairie grass roots are clumped under individual tussocks and interspersed with tubers and other rooting structures of desert perennials. Documentation of rooting patterns in rnodern soils is a laborious process that involves digging, erecting a supportive net for the excavated roots, and then carefullywashing them out (Weaver 19 19,1920). Such research provides data against which fossil root systems can be compared. Patterns of root traces also are clues for distinguishing original features of paleosols formed after burial.
Figure 3.4 Scale drawings of excavatedroot systems oE (A)short grass prairie near Colorado Springs,Colorado,U S A (B) lowland tall grass prairie near Lincoln, Nebraska, USA; and (C) mountain forest near Pikes Peak. Colorado,USA(afterWeaver 1919, 1920; withpermissionfrom thecarnegie [nstitutionof Washington).
Features of fossil soils Root traces run around hard parts of soils such as pebbles, nodules andcemented horizons (Haasis 1921). Abundant subhorizontal root traces deep within a profile are the best line of evidence for recognizing originally lithified horizons (fragipans and duripans) in paleosols now entirely lithified. Similarly, the avoidance of nodules by roots and burrows closely approaching them is evidence that the nodules were an original part of the soil. Not all root traces avoid nodules because in their early stages of formation nodules may be unindurated chemical segregations (Gile et al. 1966, 1980). With time they become indurated and better defined. As clear primary features of paleosols, root traces are guides to other original features of former soils. Rhizoconcretions Rhizoconcretionsform in soilsbecause of the local environment created by roots. Water is taken in by roots because of the tendency for soil water to dilute their cell sap across semipermeablecell membranes (by osmosis) and because the whole plant maintains a negative pressure (water potential) within thin, water-conducting tubes (xylem)by loss of water from the leaves (transpiration). Nutrients are taken in with water and their uptake is enhanced by a variety of substances exuded by roots and their surrounding mucigel zone (rhizosphere) rich in bacteria and fungi. Many nutrient cations (Ca*+,Mg*+, K+, Na+) are reIeased by dispIacement with hydrogen ions (H+)in mildly acidicsolutionsmaintainedbyorganic acids and by carbonic acid arising from dissolved carbon dioxide of microbial and root respiration. Other nutrients such as iron are dissolved in a reduced state (Fe”) by organic reductants such as caffeic acid or are fixed in particularly favorable molecular sites (chelated) by large organic molecules such as EDTA (ethylene diamine tetra-acetic acid). This is not to say that the rhizosphere is always or uniformly acidic or reducing as was once thought. Most of the time it is near neutral in pH and Eh, allowing for normal aerobic respiration and nutrient uptake by both roots and associated microbes (Richards 1987). Conditions can change over short periods of time. Heavy rainfall may cause temporary waterlogged,reducing and acidic conditions. Long periods of nutrient starvation may induce dramatic increases in the production of exuded reductants or chelates over periods of afew hours (Olsenetal. 1981).Thenet ef-
17
fect of root action is thus to deplete the adjacent soil of nutrients. Zones of depletion are not especially prominent in soils because of the continued elongation of roots. The most actively adsorbing part is the zone of root hairsjustbehindthegrowingpointoftheroot.This root apexmayelongateatratesof morethan 6 cmday-l; about 2 cmday-’ istypical (Russell 1977).Becausethey are so transitory, detecting the effects of fossil rhizospheres is not simple. Nevertheless,the effectsof roots can be striking. Soils rich in calcium carbonate are widespread in desert regions, where rainfall is insufficient to leach it from the soil, and also in humid coastal dunes and beaches, where sand includes numerous grains of broken seashells. With repeated cycles of wetting and drying in such friable, sandy soils, root margins alternately become wet and acidic (dissolvingcarbonate) and then dry and alkaline (precipitating carbonate). Underthese conditions roots become heavily encrusted with concentric layers of very fine-grained, low-magnesian, calcite (Fig. 3 . 5 ) . These calcareous rhizoconcretions may become so thick and unyielding that the root dies and the hole remaining is filled with other materials (Esteban & Klappa 1983 ) . Rhizoconcretions also form by biologically mediated calcification of cells within roots (Alonso-Zarza et al. 1998a). Combinations of intracellular and extracellular calcification create what can be termed rhiolite: rock formed by the activity of plant roots (Wright et al. 1995).Iron mobiliied in the drab ferrous state from minerals within a wet rhizosphere may be oxidized to yellow or red ferric oxides near the roots to form ferruginous rhizoconcretions. Calcium carbonate and iron oxyhydrates are the most common materials encrusting roots in soils, but many other substances also form rhizoconcretions and soil nodules (Brewer 1976).
Drab-haloedroot traces A common feature of root traces in paleosols is a bluish gray or greenish gray halo extending out into the paleosol matrix (Fig. 3.6). Such drab root traces can be formed in several ways. Two of these are uncommon in paleosols and can readily be recognized using textural relationships. First, drab root traces and burrows can form in red beds or horizons when drab material from a higher horizon is washed down into a lower horizon
Chapter3
18
Figure 3.5 Calcareous rhizoconcretions exhumed from Holocenecoastal sand dunes north of WandaBeach. New South Wales, Australia (Retallack specimens P2606A, -C, -D). Scaleinmillimeters.
(Fig. 3.7). Such structures are called krotovinas, tonguing or glossic features. The boundary between drab and red material of a krotovina is a sharp discontinuity between materials of different texture, unlike the diffuse contactinmost drab haloes (Retallack1997a). Second,drab haloes form around root traces in clayey, periodically waterlogged soils. Anaerobic bacterial activity in stagnant water around roots, burrows and cracks in the soil can cause chemical reduction of those surfaces, leaving the interiors of the soil clods oxidized. Such surficial reduction from brown to gray color may
l~llllJ1 I
also result in mineralization with birnessite, pyrite or sphaerosiderite (Retallack 1997d, 1999b). These features of surface-water gley are found in very clayey lowland soils with impermeable subsurface horizons. Such variegated soils have drab-haloed root traces accompanied by drab burrows and soil cracks. Surface-water gleying does not explain soils in which only the root traces have drab haloes. A third possible origin for drab-haloed root traces is as a result of reduction and mobilization of iron within the rhizosphere. The highly reducing conditions implied by
Figure 3.6 Fossilroot traces (2 mmdiameter central streak of pale yellow, 5Y 7/4), drab halo (light gray, 5Y 7/2) and ferruginized matrix(darkred, 2.5YR316)fromthe subsurface (Bt) horizon of a paleosol of late Eocene (38 Ma) age in the basal John Day Formation, near Clarno, Oregon, USA (specimen R4 17;Getahun & Retallack 1991; Retallacketd. 2000).Scaleinmillimeters.
Features of fossil soils C
D
19
E
SOIL
Figure 3.7 Processesandproductsinthe formation of drab-haloedroot traces and superficially similarfeatures:(A) krotovinas: (B) remnants of the rhuosphere;(C) surfacewater gley;(D)reduction of soil by anaerobic decay of buried roots: (E) dehydrationof ferric hydroxidemottles.
SOIL KROTOVINA gruy sand
the halo would be too poor in oxygen for normal respiration of the deeply penetrating roots commonly haloed in this way. Furthermore, water and nutrient uptake is most active near the tips of rootlets, where there are zones of root hairs, whereas drab haloes are found around fossil roots too large to have borne root hairs. A fourth explanation for drab-haloed root traces is that they are reduced areas of anaerobic bacterial decay of organic matter buried within paleosols. This burial gley origin of drab-haloed mottles is likely for paleosols in which the whole surface horizon is drab, from dispersed organic matter there (Fig. 3.8A), and in which there are comparable amounts of total iron in both drab and red areas (Kraus & Aslan 1993; Retallack et al. 2 000). Fifth and finally, the contrast between drab-haloed root mottIes and surface horizons and the reddish remainder of many paleosols may have been enhanced by dehydration of yellow and brown ferric oxyhydrates to brick-red hematite during deep burial. However, it is unliiely that original oxidation occurred at depth, especially within impermeable clayey paleosols. This well-established diagenetic change (Retallack 1991b) may have enhanced the color contrast of drab-haloed root mottles, but is unlikely to have been their cause. If drab-haloed root mottles can be regarded as reflect-
1
ORIGINAL SURFACE-WATER' RHIZOSPHERE GLEY brawn clay
red clay
BURIAL GLEY gray clay
BURIAL REDDENING
::&,,
ing the former rhizosphere or anaerobic decay of buried organic matter, then the drab-haloed root traces represent the last crop of plants in the paleosol. Former roots and their rhizospheres decay rapidly once they die in well-drained soils. Drab-haloed root systems are especially useful in distinguishing between forest, woodland and wooded grassland of the past from paleosols (Retallack 1983a, 1991a; Retallack et al. 2000).
Soil horizons Soil horizons of paleosols vary considerably depending on the conditions under which the soil formed and was subsequently buried. There are, however, some general features of horizons useful for recognizing them even in highly altered paleosols. Horizons of paleosols are distinct from many kinds of geological layering in that the top of the uppermost horizon of a paleosol is usually truncated sharply whereas boundaries between lower horizons and underlying parent material commonly are gradational (Fig. 3.8). Exceptions to the general rule of diffuse contacts below the sharp top of paleosols are common enough to deserve special consideration. Some lowland soils receive thin increments of sediment through which vegetation continues to grow. Their tops are overthickened
20
Chapter3
Figure 3.8 Sharpupper contact and gradational lower contact of soil horizons in two paleosols: (A)Long Reef clay paleosol (probablya Hapludult)of Triassic age (Scytho-Anisian,245 Ma) in the Bald Hill Claystoneat Long Reef, New SouthWales, Australia (scalebar graduatedininches,paleosols discussedby Retallack 199 7d);(B) modern grassland soil (upperleft only) and two comparableChogo clay paleosols (probablyHaplustolls)of Miocene age (14 Ma) in the Fort Ternan Member, mainexcavationof FortTernanNational Monument, Kenya (hammerfor scale,paleosols discussedby Retallack 1991a).
by sedimentation rather than sharply defined by erosional truncation, and could be interpreted as cumulic paleosols. But there are usually one or more layers with roots and burrows denser than elsewhere. The bioturbated zone closest to the top of a paleosol is best taken as its surface and others above that as the tops of additional, very weakly developed, younger soils (Retallack 1997d). Other exceptions to the generally diffuse boundaries of paleosol horizons are sharp contacts occasionally found within profiles. For the most part these are relict beds from sedimentary parent materials not yet obliterated by soil formation. Associated sedimentary features such as ripple marks or load casts allow confident identification of relict bedding (Retallack&Krull1999).There also may be erosional surfaces within profiles where a pre-existing paleosol has been substantially eroded and soil developmenthas proceeded on an additional layer of sediment (Retallack1997d). Such cases may bedifficult to detect in the field if the erosional contact has been obscured by subsequent soil formation. Thin lag depositsof weather-resistant pebbles such as quartzmay be cluesto complexities of this type of soil and paleosol. Stone lines are common in very ancient soilsof stable geologicalsettings (Johnson et al. 1987). Without such field indica-
tions the true complexity of a paleosol may not become apparent until petrographic or chemical studies reveal discontinuities.
Kinds of horizons Although paleosol horizons are rather varied, a mental image of common kinds of soil horizons can be useful in recognizing them in the field. Some kinds of paleosol horizons are so striking that they have attracted specific geological names. Cornstones (Fig. 3.8B), for example, arenodularcalcareous horizons (orBkin the shorthand of soil science) and ganisters are silicified, near-surface, sandy horizons (or E horizons of soil science). Successions of paleosols with gray-green, organic surface (A) horizons and red to purple, clayey subsurface (Bt) horizons (Fig.3.8A) may form strikinglyscenic sequences as gaudy as a barber pole or candy cane (Retallack 1997a). Some Precambrian paleosols have surface (A) horizons of a distinctive lie-green color (Mossman & Farrow 1992). Somekinds of paleosol horizons may prove to be extinct, but for most the horizon nomenclature of soil science is appropriate. The variety of horizons found in modern soils are labeled with a shorthand system of letters and numbers
Features of fossil soils A horizons
profiles above C
5
Figure 3.9 Typicaldepthsfor differentkinds
of soil horizonsas revealedby their frequency distributionin North American soils (compiledfrom Marbut 193 5).These are scales appropriatefor observation of paleosol horizons.
i 2
I5
I$
I >
3
(such as A and Bt). Laboratory studies may force one to change a horizon designation, but it is best to provisionally identify paleosol horizons in the field. Field assessment of horizons may determine how a paleosol is sampled,and also its interpretation and identificationin a modern classification of soils. The field codes for horizons are descriptive rather than genetic in orientation, and the current scheme is simpler than that found in older publications (Table 3.1). Horizons are defined on the basis of the materials composing them. For buried soils, it also is useful to have an idea of typical thicknesses of different kinds of horizons in modern soils. Most surface horizons of soils (A and E) are <50cm thick and subsurface (Bt and Bk) horizons are <2m (Fig. 3.9). Unaltered sediments, duricrusts and saprolites. on the other hand, may be much thicker and reflect the operation of processes other than soil horizon differentiation (Thiry 1999; Valeton 1999). Describing soil horizons
To characterize a paleosol in a way that is amenable to interpretation, horizon thicknesses, grain size, color, reaction with acid and thenature of horizon boundaries must all be recorded in the field. Some observations will prove more important than others in understanding a paleosol, but it is difficult to anticipate which features these will be. Thus, it is useful to have a comprehensive standardized form of horizon description (Retallack 1988a). Especially useful is graphical
10
20
30
, 40
p l o b d and with no A horizon below plow line
E horizons
.
, 5
p
,, 1,5
,
Bt horizons 10
21 Bk
1 5
n=32
field logging in a large, square-ruled notebook as is the custom in making geological sections of sedimentary sequences. Like the interpretative shorthand used to describe soil horizons (Table 3.1). these various symbols and style of representation may at first appear intimidating. With familiarity, such schemes become a rapid way of summarizing and comparing profilecharacteristics. For this reason all graphical profiles in this book show uniform symbols. Information on grain size is needed for classification and interpretation of many soils and paleosols. A profile showing mean grain size, separate from the column with lithological symbols is a useful feature of graphical profiles of paleosols. Grain size can be estimated in the field using a grain-size comparison card or sediment samples. It can later be reassessed by laboratory studies of samples of the horizons, which should be clearly labeled with an indication of their former orientation in the paleosol. The simplest way of marking orientation is to draw or scratch a large (2 cm diameter) circle on the upper surface of the sample or sample container in such a way that it is parallel to the ancient land surface. Such samples can be made into thin sections oriented vertical relative to original bedding (Murphy 1986; Tate & Retallack 1995),and then point-countedunder apetrographic microscope to determine more precisely the amounts of sand, silt and clay (Fig. 3.10) or of diagnostic features such as clay skins (Murphy & Kemp 1987). Common components can be determined to an accuracy of about 2%when 500 points are counted (Murphy 19 8 3; Friedman&Johnson 1996).Alternative methods
22
Chapter3
Table 3.1 Descriptive shorthand of soil and paleosol horizons. Category
New term Description -
Master horizons
0
A E
B K C R Gradations between master horizons
AB
Subordinate
a b
BA EIB
C
e f g h i k m n 0
P q r S
t V
W X
Y Z
Surfaceaccumulation of organic matter (peat, lignite, coal) overlying clayey or sandy part of soil Accumulation of humified organic matter mixed with mineral fraction. Occurs at the surface or below an 0 horizon Underliesan 0 or A horizon and is characterized by lessorganic matter, less sesquioxides(Fe,O, and Al,O,), or less clay than the underlying horizon. This horizon is usually light colored as a result of abundant quartz Underliesa n 0,A or E horizon and shows discernibleenrichment in clay, carbonate, sesquioxides(Fe,O, and Al,O,) or organic matter Subsurface horizon so impregnated with carbonate that it forms a massive layer Subsurface horizon more weathered than bedrock but lacking degree of weathering of A, E. B and K horizons Consolidated and unweathered bedrock
OZd term 0
A A2
B K C R
Horizon with some characteristics of A and of B, but with A characteristics dominant As above, but with B characteristics dominant A horizon predominantly liieB horizon but with tongues of E horizon
A3
Highly decomposed organic matter Buried soil horizon (usually redundant for sequences of paleosols) Concretionsor nodules Organic matter intermediate in decomposition(between a and i) Frozen soil,with ice wedges or other evidence Evidence of strong gleying,such as pyrite Illuvial accumulation of organic matter Slightlydecomposed organic matter Accumulation of carbonates less than for K Evidence of cementation, such as root deflection Accumulation of sodium, halite, columnar peds Residualaccumulation of sesquioxides Plowing or other comparable human disturbance Accumulation of silica Weathered or soft bedrock Illuvial accumulation of sesquioxides Accumulation of clay Pliithite Colored or structural B Fragipan character (cemented with clay and silica and avoided by roots) Accumulation of gypsumorcrystaIcasts Accumulation of other salts or crystal casts
-
Source: AdaptedforpaleosolsfromSoilSurveyStaff (1975,1998).
B1 A&B
b cn -
f g h -
ca m sa P si ox
ir t -
X
cs sa
Featuresof fossil soils
23
100%clay (less than Zum)
Figure 3.10 Soil texturalclassesof Soil
Survey Staff (199 8). Filled circleswith numbersshow clay percentages for critical boundaries:open circles with numbers show diagnosticsand percentages (komRetaUack 199 7, A Colour Guide to Paleosols:with permissionOJohnWiley&SonsLtd).
of passing unconsolidated soil through sieves or through a column of water and then weighing each segregated size fraction are techniques used in soil science (Klute 1994).They are of limited usefulness in evaluating grain size of lithi6ed paleosols because there is no easy way of disaggregating them to their original grain-size distribution. One should be aware also that the silt-clay size criterion ( 2 pm) of soil science is different from that widely usedin sedimentarygeology (Kraus 1997). Color can be estimated in the field using a standard color chart such as the one produced by Munsell Color (1975).This is arranged to reveal changes in the primary colors such as red and yellow (hues),in the degree of lightness of the color (values)and the degree of grayness of the color (chroma).It is best to consider only the hue initially, to find the correct page of the charts. Then the color can be determined within a grid of value and chroma. Well-indurated and metamorphosed paleosols may hold their color well, but Iittle-aItered clayey paleosols of the kind widespread in scenic clayey badlands of Mesozoic and Cenozoic rocks in the western USA change color upon exposure to air and on lab0 ratory storage. Commonly they become paler (higher Munsell value)after a few hours of drying.After several months of laboratory storage, greenish gray parts of paleosol samples may become more yellow (warmer in
Munsell hue) because of oxidation of iron-bearing minerals. Thus, it is best to record color on fresh rock within a few minutes of exposure. The calcium carbonate content of paleosols can be determined in the laboratory, but should also be estimated in the field by applying drops of dilute (10% of standard 1 M solution) hydrochloric acid from an eye-dropper bottle.The degree of reaction with acid can be divided into five readily observablestages (Table 3 . 2 ) proportional to the amount of calcium carbonate present. This property is used for classifying soils, as well as for distinguishingcalcium carbonate from dolomite or chert. Another feature of horizons that should be recorded in the field is the nature of their contacts with adjacent horizons. Two aspects of the contact are of interest: whether one horizon passes into another within a narrow or broad vertical distance and whether the contact is laterally planar or somehow disrupted (Table 3.3). Abrupt or broken boundariesmay lead to suspicionthat the profile contains erosional discontinuities. Diffuse planar contacts, on the other hand, may represent genetically related horizons of a single soil. The transition zones between diffuse horizon contacts may seem so thick as to warrant a separate horizon name. When assessing a profile withespeciallydiffuse horizon boundaries it is best to take an overview of the profile to decide
24
Chapter3
Table 3.2 Field scale of acid reaction to approximate carbonate content. Class
Carbonate (wt %)
Soundof reaction (holdclose)
Reaction with diluteacid
Noncalcareous Weakly calcareous
<0.5 0.5-1
Calcareous
1-5
Strongly calcareous
5-10
Very strongly calcareous
>10
None Faintly increasing to slightly audible Faintly increasing to moderatelyaudible Easily audible, heard away from ear Easily audible
Acid unreactive, may form inert bead Little movement within the acid drop, which could be flotation of dust particles as much as bubbles Numerous bubbles,but not coalescingto form a froth Bubblesforming a white froth, with bubbles up to 3 mm diameter, but drop not doming upward Drop vigorouslyfrothing and doming upward with some bubbles up to 7 mm diameter
Source: Retallack(l997a).
which parts of the profile are most distinct, and thus constitute the main horizons, and which are merely intergrades between them. Horizons also may be characterized by laboratory analyses of hand specimens. Chemical analyses for major and trace elemental composition may be useful. For lithfied paleosols it is best to use whole-rock chemical analyses by standard geological methods such as atomic adsorption, X-ray fluorescence,inductively coupled plasma fusion spectroscopy or neutron activation analysis. Full chemical analyses were once fashionable in the study of modern soils (Marbut 193 5) and this important data base of analyses is occasionally supplemented with additional published analyses (Soil Survey Staff 1975). Determination of materials extractable by Table 3.3 Soil horizon boundaries. Category
Term
Description
Sharpness
Abrupt Clear Gradual Diffuse
<2 cmwide 2-5cm 5-1 5 cm >15 cm
Smooth Wavy
Nearly aplane Undulating, pockets wider thandeep Undulating, pockets deeper than wide Partsof the horizondisconnected
Lateral continuity
Irregular Broken
Source: SoilSurveyStaff (1993).
solvents, such as iron by sodium dithionate, is a much more common approach of soil science (Sparks 1996). This readily soluble fraction of the soil is most susceptible to diagenetic change upon burial of a soil, and this limits the use of these analyses of paleosols. A useful adjunct to chemical analyses is to determine the density of samples by the standard method of weighing sampIes coated in paraffin, to prevent water infiltration, both in and out of water (Klute 1994). Using density measurements, chemical analyses in weight per cent can be converted to values in grams per cubic centimeter, to gain an inventory of absolute chemical differences between horizons (Chadwick et al. 1990; Brimhall et al. 1991). Staining for feldspars (Houghton 19SO), scanning electron microscopic observation (Mumpton & Ormsby 1976; Smart & Tovey 1981;Sudo et al. 1981;Krinsley & Manley 1989),X-raydiffractometerstudiesof clayminerals (Klute 1994) and X-radiographs to reveal partly concealed structures (Bouma 1969) are just a few of a growing number of laboratory techniques for understanding particular aspects of paleosol horizons.
Soil structure Soils may appear to be fragmented, featureless or massive compared with, for example, cross-bedded dune sands or strongly cleaved slates. But l i e these geological structures, soils have characteristic structures of their own. These structuresare developed to different degrees and progressively overwhelm pre-existing structures of parent material such as sedimentary bedding, metamorphic foliation or igneous crystal outlines.
Features of fossil soils With experience,both the characteristic features of soil structure and their contrast with other structures of enclosing rocks allow paleosols to be readily recognized. Structures of paleosols often can be discerned in geological descriptions by such terms as ‘massive’, ‘structureless’, ‘jointy’, ‘slickensided’, ‘veined’, ‘mottled’ or ‘nodular’. Such language does not serve adequately to characterize soil or paleosol structures. On the other hand, technical jargon for these features used by soil scientists can be intimidating. In this account, the technical terms used are based on the system outlined by Brewer (19 76). Common words are here suggested as equivalent terms. In some cases these ordinary terms may prove inadequate, but they serve to introduce the subject.
25
andpaleosols arecalled cutans. They could also be called clod skins because it is the cutans and the voids with which they are often associated that define the fundamental units of soil structure, the individual clods or peds. Peds may be of various sizes ranging from large prisms occupying most of the thickness of the soil to small granules the size of sand grains. Large peds may be made up of smaller peds. Burrows and root traces are two examples of an additional general class of tubular features found in soils: pedotubules. Such a general term may be needed for objective description of these features when their origin is not obvious. Another general class of soil structures are glaebules, or naturally hardened soil lumps. These are masses of material that have a distinctive mineralogical and chemical composition. Calcareous nodules of aridland soils are a good example. Crystals of minerals such as gypsum are not regarded as glaebules, and also are found within soils. Pedotubules, glaebules and crystals have adistinctive and often also simple mineralogical composition unlike the clayey material making up most of the soil. On a microscopic scale this fine-grained soil matrix (plasma) consists of clay and amorphous iron and aluminum oxides, which support small grains of rocks and minerals (skeleton grains) inherited from the parent material of the soil. The process of weathering tends to create plasma from skeleton grains. Minerals such as plagioclase and hornblende are especially prone to alteration.
Structural elements Perhaps the most striking feature of soils is their intricate system of open cracks and hollows (Fig. 3.11). These open spacesmay form an interconnected network (packing voids),small irregularly shaped pockets (vugs) or near-spherical holes (vesicles).In paleosols such open spaces are often crushed out of existenceby compaction under overlying rocks. Fortunately, some indications of former cracks remain because of modification of soil material where water and air could circulate within the soil.These surfaces can be modified in various ways, e.g. by encrustation with washed-down clay or staining with iron oxide. Such irregularly planar features in soils
void
concretion
intersecting bright clay fabric (S-matrix, bimoseplc plosmic
lay skin (orgillan and lluviation cutan)
Figure 3.1 1 Structural units of soilsin hand specimens and petrographic thin section.
Chapter3
26
Quartz, on the other hand, is resistant and is a common skeleton grain in highly weathered soilsand paleosols. Much can be done to characterize soil structure in the field, but a complete description requires microscopic examination. Petrographic thin sections are useful for documenting differences between horizons of soils and paleosols as well as their microfabric. For this purpose, some of us (Nahon 1991; Retallack 1997a) prefer the original terminology of Brewer (1976), as a middle ground between the simpler system of Bullock et al. (1984) and the much more complex terminology of Brewer & Sleeman (1988). Emphasis on voids in the latter system is unworkable for paleosols, and simplicity of the former system does not address some soil features critical for paleosol studies.Yet another terminology for microscopic alteration of grains and crystals (Delvigne 1998) can be used to supplement these comprehensive systems.
Peds Peds are aggregates of soil: the clods of earth between cracks, roots, burrows and other soil openings. Soil peds commonly can be crushed by hand and are not as indurated as rock. They are persistent clods that tumble
I
TYPE
I
I
I
ANGULAR @LOCKY
SUBANGULAR BLOCKY
tabular and elongate with flat elongate with horizontal to land top and vertical domed top and to land surface vertical to surface surface
mquant with ShOrF nterlocking edger
quant with dull iterlocking edges
E.B~.K.C BI initial disruption of swelling and relict bedding. shrinking on accretion of wetting and cementing material drying
Bt racking around oats and burrow* welliny and hrinking on vetting and trying
81 s for angular ilacky, but with nore erosion and leposition of ioterial in cracks
PLATY
PRISMATIC
COLUMNAR
I
DESCRIPTION I
I
IUSUAL HORIZON^ L MA~N CAUSES
loosewhen one digs into the soil. Peds may initially seem so irregular and formed by such a random concatenation of processes as to be unworthy of careful scientific scrutiny.Even a little experience with soilsis sufficientto dispel such doubts. Consider, for example, the very different domed columnar peds of soils around salt pans of desert regions compared with the fine granular peds found under suburban lawns. Each reflects particular kinds of soilsand environments. Compaction and other alteration after burial of paleosols makes peds in them difficult to recognize. The pattern of cracking that determines which chips of rock come loose under the hammer from modern outcrops of lithified paleosols in most cases has more to do with jointing or other features of burial or modern weathering than with the distribution of voids in the former soil. Thus, care must be taken in paleosols to identify peds by their surrounding cutans. Soil peds are classified by their size, angularity and shape (Fig. 3.12). Platy peds, for example, are thin but extensive laterally. They are often formed by the initial disruption of relict bedding in weakly developed soils or where clay or other materials form laminar layers on top of a fine-grained impermeable lower layer of a soil. Prismatic and columnar peds are taller than wide and
I
1
very thin
1 very thick-I0
I
I
I
I
GRANULAR
CRUMB
I
I
Rn as for prismatic, but with greater erasiw by percolating water, and greater
A ictive bioturbation ind coating of soil vith films of clay, esquioxides ond rgonic matter
A as for granular, includinq fecal pellets and relict soil clasts
ery fine < 0 5 c n
ery fineG05crn
ery fine-dmm
very fine
me 0 5 to I cm
ine 0 5 to I cm
ine I to 2 mm
fine I to 2 mm
nedium I to 2 cm
nedium I to 2 cm
iedium 2 to 5 mm medium 2 to 5 mm
oarse 2 to 5 cm
oorse 2 to 5 cm
mmlvery c o a r s e ~ l O c m ~ v ~ r ~ ~ o ~ r s ery e ~ coorsesScn lOcn
ery coarse > 5 cn
veryfine
veryfine
Figure 3.12 Aclassitlcationof soilpeds (sirnpliedfrornSoilSurvey Staff 1975).
Features of fossil soils may extend through a considerable portion of a soil. They form in clayey soilsby swelling and shrinking associated with wetting and drying. Someespeciallyexpandable clays such as sodium smectites form bulging tops to the peds (McCahon & Miller 1997). These bulges distinguish columnar from prismatic peds. Blocky peds (Fig. 3.1 3) are irregular in shape and usually not quite equidimensional. They may have angular interlocking faces. Where there has been erosion or coating of ped margins they appear less sharp and subangular. A distinctive kind of blocky ped is formed in soils with highly smectiticclays,in which diagonalstresses are associated with theswellingof clayasittakesinmoistureaftereach rain. These lentil or wedge peds are separated by slickensides (Krishna & Perumal 1948; Gustavson 1991). Granular and crumb peds are differentiated on the sharpnessof theirinterlocking edges,as are angular and
-
27
subangular blocky peds. Both granular and crumb peds are smaller and more equidimensional than blocky peds. Theyarefoundinupperportionsof soils,whereas blocky peds are more often encountered in subsurface (Bt)horizons.The gradationin structurefromblocky to granular and crumb peds may reflect the relative role of soil organisms in churning the soil and producing organic compounds, such as polysaccharides, that bind ped surfaces. Granular ped structures are common in the surface horizons of grassland soils. So are crumb peds, which may in part be fecal pellets of earthworms and other creatures (Pawluk & Ball9 85). A distinctivekind of crumb ped called spherical microped comprises sandsized spheres of deeply weathered, ferruginized clay abundant in some tropicalsoils (Mermut et al. 1984). Spherical micropeds, and blocky and prismatic peds are superficially similar to clay clast breccias or con-
28
Chapter3
glomerate beds, to tectonically formed fitted breccias or tessellated pavements,and to mineraliedvein networks or boxworks. These other geological structures commonly are associated with high-temperature minerals, marine fossils or other indications that they did not form in a soil.The diagnostic features of peds in paleosols are complexly altered bounding surfaces, peds within peds, traces of former surfaces annealed within peds, and cracks between peds that widen in a way that indicates the peds were loose enough to rotate slightly (Fig. 13.13).
Cutans Cutans are modified surfaces of peds, clasts or crystals. They vary greatly in composition, which is the main basis for their classification. In the terminology proposed by Brewer (1976), various kinds of cutans are given latinate names indicating their composition, followed by the suffix '-an' to indicate that they are cutans. Argillans, for example, are clay skins. Ferrans are ironstained surfaces, and so on (Table 3.4). Similar cutanic features can form also during diagenesisand metamorphism of paleosols. Such later features are usually thicker, more coarsely crystalline and related to other distinctivemetamorphic textures such as kink folding or schistosity In contrast, cutans of the original soil before metamorphism are irregular, thin and fine grained. The chemical composition of cutans thought to be original can be important guides to former chemical conditions in paleosols (Fig. 3.14). Noncalcareous, nonclayey, ferruginous cutans (ferrans), for example, indicate acidic
Table 3.4 Terminologyfor cutans. Term
Composition
Argillan Ferran Mangan Organan Sesquan Silan Skeletan Soluan
Clay Sesquioxidesof iron Oxidesorhydroxidesof iron andmanganese Organic matter Sesquioxidesof iron and aluminum Silica Clastic grains such as quartz or feldspar Solublesaltssuch as gypsumor calcite
Source: Brewer (1976).
and highly oxidizing conditions as would be found in well-drained,sandy soilsof humid climates.This kind of information can be useful for interpreting and classifying paleosols. Subcutanic features show a fixed relation to a ped surface within the soil, but are not restricted to that surface. Two kinds are recognized. Neocutans (deep clod skins) are extraordinarily thick cutans that often lose intensity of development over a zone of centimeters away from the surface. Quasicutans (halo clod skins) also are thick and show a relation to a surface but are not right at that surface. They form a diffuse zone of alteration beneath the surface, which faithfully follows its outline. Neocutans form in similar ways to other kinds of cutans and have the same range of composition. They may reflect intense development of a cutan. Some quasicutans may form as neocutans partially destroyed by later modification of the surface; for example, a deeply ferruginized surface of aped, formed under oxidizing conditions, may then become reduced and gray in color nearest the surface because of surfacewater gleying. The origin of cutanic features may be complex, but generally falls into three categories: those formed by washing down of material into cracks (illuviation cutans), those formed by progressive alteration inward from a surface (diffusion cutans) and those formed by differential shear forces within the soil (stress cutans). These three genetic categories of cutans may be combined in confusing ways, but often are distinct. Illuviation cutans, for example, show sharp contacts with adjacent soil peds and also may be laminated from successive additions of material washed in. These are the kinds of cutans that are most clearly diagnostic of paleosols. Somewhat similar are clay cements filling secondary porosity created by dissolution during deep burial. Clay cements seldom show the grain-size variation between layers, lateral continuity of layering or correlative banding between broken pieces seen in illuviation cutans. Diffusion cutans, on the other hand, do not consist of material different from that of the peds but rather of modified ped material. They have only one sharp boundary on the outermost surface. the inner surface being a gradational change into unaltered soil material. Acommon exampleis a thick ferruginizedsurface of peds. These neoferrans are most intense and opaque at the surface, but less deeply stained toward the center
Features of fossil soils
HEMAT ITE LlMON ITE MNOXIDES SILICA
CHAMOSITE' CoIcil* Phe.phWmI*
29
:
;
--. --_I:---
--_--_ --_-.-.
~
I
PEAT PYRITE
Figure 3.14 Stablemineral associations under different conditionsof aeration (Eh) and acidity (pH)in sediments and soils (from Krumbein&Garrels1952,JourndofGeology 60,Fig.8;reprintedwithpermissionfrom University of ChicagoPress).
- 0.3
I
of theped. Stresscutansareless welldefined. Acommon indication of stress cutans is the striated and smeared surfaces called slickensides. These form in clayey soils where peds are repeatedly heaved past one another by swell-shrink during wetting and drying episodes.Slickensides also form in paleosols purely by the crushing of peds one against another during compaction after burial. Thus, by themselves slickensides are not compelling evidence of shrink and swell behavior of paleosol clays. Unlike the slickensides associated with faults, those formed in soils or during compaction of paleosols are randomly arranged along diffusezones rather than uni-
directional and concentrated in narrow bands (Gray & Nickelsen 1989). On either side of the highly birefringent, striated clay of the slickensided surface, the transition outward to unoriented clay is gradational with decreasinglyabundant wisps of highly birefringent clay (Brewer 1976).
Glaebules Glaebules are naturally segregated lumps of soil material formed from the same wide variety of materials as cutans (Brewer 1976; Nahon 1991: Delvigne 1998).
30
Chapter3
Common examples are the calcareous nodules of desert soils. Glaebules range from highly irregular to almost spherical in shape and also vary in their distinctness and internal structure, but can be distinguished from cutans by their nonplanar shapes and more distinct outlines. Two common kinds of glaebulesare nodules and concretions. These two terms are widely confused as synonyms, but for soil science are distinguished on the basis of their internal structure. A nodule is massive internally whereas a concretion contains concentric layers. This difference can be discerned by breaking them open or making a thin section through them (Fig. 3.15). Internal structure may have significance for the way in which they form. Continuous growth or recrystalliiation forms nodules whereas concretions form by discontinuous, often seasonal, growth. Some nodules have a system of cracks radiating from their center. These septarian nodules appear like miniature turtle shells. They form when nodules change volume,for example,by drying out irreversibly. This may happen when siderite nodules, formed in stagnant water in lowland soils, are exposed to air (Retallack1997d).Bothnodules and concretions are hard and have a sharp and distinct outer boundary. They contrast with another kind of glaebule, called mottles, which are diffuse patches of the same kinds of materials. Mottles may in some cases be early stages in the differentiation of nodules or concretions, but for the most part are irregular patches of discol-
oration. A final kind of glaebule is local aggregations of clay with sharp boundaries. In some cases it may be clear that these are fragments of older shale or claystone in the parent materialof thesoil (lithorelicts),fragments of other soils in the parent material (pedorelicts). fragments of clay-filled cavities in the soil (illuviation argillans),pockets of minerals locally weathered to clay (clay nodules) or feces of soil-ingesting animals (fecal pellets). For the cases where the origin of such features is not so clear that they can be assigned to one of these genetic categories, the terms papule, clay gall or alteromorph are convenient, descriptiveterms (Brewer 1976; Delvigne 19 9 8). Glaebules can be described by their composition, which may be sesquioxidic,manganiferous, calcareous, sideritic, pyritic or siliceous.These compositions reflect particular soil conditions (Fig. 3.14) provided they can be shown to have originated in the soil. Glaebulescan be described as spherical, ellipsoidal, tuberose (irregular l i e a potato) or irregular (variable, but with rounded edges unlike a cutan). Mottles vary in the distinctness of their boundary with surrounding soil (Table3.5). Glaebules are abundant and conspicuous features of many soils and paleosols, but are not diagnostic of them. Marine sedimentary rocks also contain glaebules. In some cases nodules in marine rocks are thought to have formed within chemical irregularities in the sediment close to the sea floor in a way analogous to
Figure 3.1 5 Thin section of a calcareous nodule (A) and ferruginousconcretion(B). ThenoduleisfromhorizonBtof thetypeLal clay paleosol (OxicHaplustalf),and the concretionfrom the type Pila clay paleosol (AquicHaplustoll),both in the Upper Miocene (8 Ma)Dhok PathanFormationnear Kaulial village,northernPakistan (Srnithsonian specimens353779and353821:Retallack 1991a).Scalebarsare l m m .
Features of fossil soils Table 3.5 Description of glaebules and mottles. Aspect
Category
Description
Distinctness
Very sharp Sharp
Transitiontomatrixin
Diffuse
Transitionto matrix over 1-5 l ~ l l ~ l Transitionto matrix in >5 mm
c.1mm
Very diffuse Contrast
Faint Distinct
Prominent
Abundance
Sue
Recognizableonly on close inspection Readily seen,differing by at least two Munsell hues, cbromas or values Obvious, with hue, chromaor value severalMunsel1units apart
Few Common Many
<2% of exposed surface
Fine Medium Coarse
<5 mmdiameter on surface 5-1 5 mmdiameter >15 mmdiameter
2-20% of exposed surface >20% of exposed surface
Source: SoilSurveyStaff (1975).
their formation in soils, but they also may form during burial. Glaebules also form in cooling volcanic tuffs (thunder eggs), on ocean floors (manganese nodules), in shallow lakes (lime balls), and around springs (tufa: Boggs 1995). Thus care must be taken to determine whether glaebules were formed in the parent material of a soil, in the soil itself or later during burial of a paleosol. Sometimes this can be established from their relationship with original features of the soil such as root traces and burrows, and from their general chemical compatibility with other original features of the paleosol. Abundant nodules in a rock may spark suspicion that it was a soil, but care must be exercised in interpreting them. Crystals Crystals occur in some kinds of soilsand paleosols,especially in cracks and cavities such as root channels and burrows (Maglione 1981;Hartley &May 1998). Crys-
31
tals can form tubular aggregates (crystal tubes), nodular masses (crystal chambers), spherical aggregates of radiating crystals (spherulites) or sheet-like aggregates (crystal sheets), or occur as single isolated crystals embedded in matrix (intercalary crystals). In some cases, crystals enclose grains of the soil. Instead of appearing uniform, translucent and sharp-edged like normal crystals, these look like a piece of sandstone or siltstone, but in the shape of a crystal. A common example of these is so-called ‘sand crystals’ of gypsum (MacFayden 1950). In very old saline paleosols, sabkhas or playas in which original crystals have been strongly modified by diagenesis or metamorphism, only the impression or filled cavityof crystalsmayremain (Smith 1990).Theformof altered crystals in some cases may be sufficientlydistinctive to allow identification (Delvigne 1998). Cubes after halite commonly have a concentric muddy fill and slightly in-bowed sides. These are common examples of crystal pseudomorphs (Haude 19 70). Crystals are found mostly in alkaline and salty soils. Like nodules, their composition may reflect a particular soil chemistry (Fig. 3.14).Calcite and gypsum are found in soils of arid climates. One indication that crystals are an original part of a paleosol is displacivefabric: the way in which peds, nodules or other coherent parts of the soil have fallen or rotated into the crystal-filled region (Fig. 3.16C).Crystals also form during burial of a paleosol, in joints and veins, or in cavitiesformed by dissolution of soluble parts of a paleosol. Crystals can be common but are not diagnostic features of paleosols.
Pedotubules Burrows and root traces are prominent tubular features in soils and paleosols, but are not in all cases easy to distinguish from fulgurites, water-escape structures or other tubular features. A convenient nongenetic term for all such tubular features of soils and paleosols is pedotubule (Brewer 19 76). Within this system pedotubules are classified according to their compositionand contrast with surrounding material (Table 3 . 6 ) . Tubular features are common in soils but not diagnostic of them. Many burrows in soils are similar to those found in lake and ocean bottoms (Buatois et al. 1998).Burrows in paleosols may be valuable indicators of paleoenvironment, especially when sufficiently distinctive that their makers can be identified. Burrows of
32
Chapter3
A. MOSEPIC
C. CLINOBIMASEPIC
-
0.1 mm
-
0.1 mm
6. SKELSEPIC
D. OMNlSEPlC
-
0.1 mm
-
0.1 mm
Figure 3.16 Sepic plasmic (bright clay) microfabrics.all under cross-polarizedlight and with bar scales of 0.1mm. (A)Mosepic plasmic (streakybright clay) fabric from clayey surface (A) horizon of a paleosol (Sul6hemist) in the midcretaceous (9 5 Ma) DakotaFormationnear Russell,Kansas, USA(Retallack&Dilcher1981a;RetallackspecimenR101). (B) Skelsepicplasmic(grainlining bright clay)fabric from the subsurface (C) horizon of the Ogi silty clay loam light-colored variant paleosol (Fluvaquentic Eutrochrept) in theOligocene (29 Ma) SharpsFormationinBadlandsNationalPark, SouthDakota, USA (Retallack 1983a: Indiana Universityspecimen 15621).(C)Clinobimasepicplasmic(trellis-lie bright clay) fabricin the subsurface (Cg) horizon of a paleosol (Aquent) in the mid-Cretaceous(9 5 Ma) Dakota Formation near Fairbury, Nebraska, USA (Retallack&DiIcher 198l a: Retallack specimenR110). (D) Omnisepic plasmic (wovenbright clay)fabric in the clayey subsurface (Bt)horizon of a paleosol (Alfisol)in the mid-Cretaceous(95 Ma)DakotaFormation near Russell,Kansas, USA (Retallack&Dilcher 198l a: Retallackspecimen R94).
earthworms, millipedes, beetles, bees, termites, ants and rodents can be distinguished when suitably preserved (Retallack 1984,1990;Hasiotis & Dubiel1995; Hasiotis & Demko 1996a).None of these creatures can tolerate waterlogging and their depth of penetration into a paleosol may indicate the former level of the water table. As clearly original parts of a paleosol. the relationship of burrows to nodules and crystalsmayreveal their time of formation. The degree of deformation and crushing of burrows may be indicators of compaction in paleosols. Much can be learned from careful study of burrows.
Microfabric Some kinds of microscopic structure are characteristic of soils and paleosols (Brewer 1976: Fitzpatrick 1984; Delvigne 1998). Especially distinctive is the appearance of the he-grainedpart of the soil (plasmicfabric)in thin sections viewed under crossed nicols. Its appearance varies somewhat with the magnification used and the orientation and thickness of the section. Most plasmic fabrics are described from observations made at total magnifications of about 100-250 times natural size. Forworkinwhichit isimportant todistinguishpaleosols
Features of fossil soils from sedimentary rocks, it is best to cut the thin sections vertical relative to paleosol horizons and to the bedding planes of enclosing sedimentary rocks. In addition, thin sect,ionsof paleosols must be ground more carefully and Table 3.6 Terms forpedotubules.
Group
Term
Definition
Fill
Granotubule
Filled with clastic grains and little clay Filled with pellet-lie clasts of clay and grains Filled withclay and grains without orientation Fill of clay and grains showing curved layering
Aggrotubule Isotubule Striotubule Matrix
Orthotubule Metatubule Paratubule
Very similar fabric and compositionto soil matrix Differentfrom matrix, derived from another horizon Differentfrommatrii. u d i e any part of profile
Source: Brewer (1976).
33
thinner than usual for most rocks so that enough light can penetrate the clayey soil matrix, which usually is more opaque than mineral grains such as quartz (Murphy 19 86;Tate & Retallack 1995). The increased abundance of cutans obscuring original sedimentary, metamorphic or igneous textures as a soil develops is expressed on a microscopic scale by the development of bright clay fabric (sepic plasmic fabric: Collins & Larney 1983). Bright clay is highly oriented, and has a high birefringence under crossed nicols (Fig. 3.16,Table 3.7). In contrast, weakly oriented clay appears dull and randomly flecked (asepic). Different kinds of bright clay fabrics reflect increasingly extensive areas of oriented clay formed by washing into cracks, by internal stresses in the soil,and by surface weathering of soil peds. The degree of development of bright clay fabric is due in part to time availablefor soil formation and in part to the intensity of soil-forming processes such as stresses imposed by wetting and drying, filling and closing of cracks, andother alterations (Brewer & Sleeman 1969). Woven bright clay fabric (omnisepic fabric) forms most rapidly in swelling clay soils of seasonal climates, but also eventually in very old soils (Holzhey et al. 1974).
Table 3.7 Termsfor plasmic microfabric.
Term
Subdivision
Definition
Asepic
Argillasepic Calciasepic Silasepic Undulic Inundulic Isotic Crystic
Mainly clay and lacking highly birefringent streaks A mixture of clay andclay-shedcarbonate, andlacking highly birefringent streaks Many silt and sand grains, and lacking highly birefringent streaks Not quite isotropic,very dark under crossed nicols Not quite isotropic under crossed nicols,but cloudy,with large irregular isotropicpatches Isotropic (likeopal) or opaque ( l i e hematite) and thus black under crossednicols Individual euhedral crystals can be discerned
Sepic
Insepic Mosepic Vosepic Skelsepic Masepic Trimasepic Bimasepic Lattisepic Cliobimasepic Omnisepic
Smallisolated patches of highly birefrhgent plasma Partly adjoining highly birefringent streaks Highly birefringent plasma associatedwith void walls Highly birefringent plasma around skeleton grains Highly birefringent streaks forming an extensivecriss-crossingnetwork Network of highly birefringent streaks in three preferred directions Network of highly birefringent streaks in two preferred directions Network of highly birefringent streaks in two preferred directions and at a right angle Network of highly birefringent streaks in two preferred directions and at alow angle All highly birefringent, orientedplasma, with a ‘woven’appearance
Source: Brewer(1976).
34
Chapter3
A. ARGILLASEPIC
C. CRYSTIC
-0.1
mm
-01 mm
B. SILASEPIC
-0.1
mm
D. INUNDULIC
-0.1
mm
Figure 3.17 Asepic (dull clay) microfabrics.All under cross-polarized light and with bar scales of 0.1mm: (A)argillasepic (flecked clay)fabric in alacustrine oil shalein theMiddleEocene (45Ma) Green River Formationin Ulrichs Quarry near Kemmerer, Wyoming,USA (RetallackspecimenP6969B); (B)silasepic(grainy) fabric from the near-surface (E) horizon of apaleosol (Aquept) inthePennsylvanian (310Ma)FountainFormationnear ManitouSprings, Colorado,USA(Suttner&Dutta 1986;Retallack specimenP6185);(C) crystic (crystalline) fabric of fibrous andsparrycalciteinthesubsurface (Bk)horizonof apaleosol (Natrustol1)intheupperPliocene(3.2Ma)HadarFormationatalocalityALX333 4kmnorth-west of thejunctionof theKada Hadar and Awash rivers. Afarregion, Ethiopia (Radosevichet al. 1992;Taieb specimen GO): (D)inundulic (cloudy)fabric from the surface (A)horizon of a paleosol (Fibrist)in the mid-Cretaceous (95 Ma) Dakota Formation near Bunker Hill, Kansas, USA (Retallack&Dilcher1981a; Retallackspecimen R108).
Lattice-likebright clay fabric (lattisepicfabric) forms in the subsurface clayey (Bt) horizons of soils where clays are under moderate confining pressure (McCormack& Wilding 1974). Most kinds of bright clay fabric are characteristic of soils and paleosols. Important exceptions are cavity-filling and grain-coating bright clay, which also are found in sediments (Scholle 1979). Coated grains may form by pressure of compaction of hard grains against a clayey matrix or by rolling or encrustation in a variety of sedimentary and soil environments. Cavity-lining bright clay can form during late diagenetic dissolution of grains or fossils as readily as
within original cavities of the soil. A more highly oriented, but otherwise omnisepic plasmic fabric is found in paleosols metamorphosed to greenschist facies and beyond (Retallack& Krinsley 1993). Asepic microfabrics (Fig. 3.1 7; Table 3.7) are not diagnostic of soils, but can reveal much about conditions of their formation. Usually sediments also show some microscopic indications of bedding or cleavage and can be regarded as laminated (stria1fabrics).Laminated fabrics can be unidirectional (mistrial), as is usually the case for bedding, or bidirectional (bistrial) as in metamorphic rocks with both bedding and cleavage, or with
Features of fossil soils
Figure 3.18 Grain fabrics. All underplanepolarized light and with a scale bar of 0.1 mm, from the type Naranji clay paleosol (Psammentic Haplustalflin the upper Miocene (7.5 Ma) Dhok PathanFormation nearKhaur, Pakktan(Retallack 1991a):(A) granular fabric from subsurface (C) horizon (Smithsonianspecimen 353837); (B) intertextic fabric from subsurface (By) horizon(Smithsonianspecimen353831):(C) agglomeroplasmicfabric from subsurface (Bt) horizon (Smithsonianspecimen 353830); (D) porphyroskelicfabric fromsurface (A) horizon (Smithsonian specimen 353828).
C, AGGLOMEROPLASMIC
Table 3.8 Termsfor soil microfabric. ~
Term
Description
Porphyroskelic
Plasmaformsmatrixinwhichskeleton grains are set like phenocrysts in porphyritic rock Plasma forms incomplete or local matrix to skeleton grains Skeletongrains more prominent than plasma, which forms intergranular braces and pockets Skeletongrains all touching with little or no plasma in the interstices
Agglomeroplasmic Intertextic
Granular
-
A GRANULAR
Source: Brewer (1976).
two cleavage planes. Crystals are widespread in sediments, especially evaporites, limestones and dolomites, as well as in alkaline soils of dry climates (Scholle 19 78). Pyrite and siderite of swampland soils are often crys-
0.1 mm
-0.1
mm
B. INTERTEXTIC
D.PORPHYROSKELIC
-0.1
-0.1
35
mm
mm
talline. Opaque microfabric(onekind of isotic fabric)appears black in petrographic thin section. Opaque fabrics are produced by iron oxides and hydroxides,which form in well-drained soils. Isotropic microfabrics (another kind of isotic fabric) appear clear under ordinary light but black under crossed nicols. This kind of fabric is found in materials replaced by isotropic minerals such as opal or zeolites, found in soils formed in volcaniclastic materials. Few soil fabrics are completely isotropic, but some are almost so. Somber fabric (undulic) and cloudy fabric (inundulic) may be the result of the drying of highly flocculatedcolloidal material such as mixtures of clay with organic matter or iron oxides andoxyhydrates. These dull fabrics are common in the clayey portions of swampland soils and their paleosol equivalents, the seat earths of coal seams. They also can be found in clayey matrices of mud flows and other deposits of weakly orientated clay. Another aspect of soil microfabric (Fig. 3.18, Table 3.8) is the proportion of original mineral grains and rock fragments (skeleton grains) to fine-grained
36
Chapter3
material (plasma). Point counting to quantify these proportions can be useful. Porphyroskelic fabric can be the original texture of the parent material or it can be the end result of soil formation in which many mineral grains have been broken down to clay by weathering.
In microscopic as well as in readily visible features, soils and paleosols have features inherited from their parent material, as well as features acquired during soil formation. Many of these soil features are useful for both interpreting and recognizing paleosols in the rock record.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 4 Soil-forming processes
Implicit among the features by which soils are recognized and classified are the processes that formed them. Such processes are more readily observed in modern soils than in paleosols in which they ceased to act long ago. Nevertheless, processes often can be inferred from features of paleosols or from suites of paleosols. The reconstruction of processes gives meaning and connection to features and kinds of paleosols. Soil-forming processes can be reduced to physical, chemical or biologicalcomponents of weathering. That this distinction is an artificial one can be seen from the example of plant roots, which are biological structures but produce exudates with chemical effects and grow in girth to physically force apart the soil. Nevertheless, a large body of literature on soil biology, chemistry and physics is testimony to the way in which these traditional scientific divisions have permeated soil science (Bohn et al. 1985; Eisenbeis & Wichard 1987; Marshall et al. 1996). Important to the interpretation of modern soils is the measurement of physical,chemical and biological features of soils such as linear expansion of clays, pH of soil solutions and the distribution of soil organic matter. Few of these features can be measured directly from paleosols altered by burial. Other sources of information pertinent to these indices of soil formation must be sought, with due regard for common kinds of alteration after burial.
Indicatorsof physical weathering Various physical processes play a role in forming soil from underlying rock or sediment. Granite and schist form deep within the Earth then expand and break into joints when unroofed by erosion and exposed at the surface. Initial loosening allows circulation of water, which turns into ice in winter or precipitates salt in a dry season. Both of these changes result in expansion, which forces cracks wider. Thermal expansion caused by forest fire may be sufficient to crack open large rocks.
Some kinds of clay swell when moist and contract when dry (Marshall et al. 1996). Effects of such physical processes can be observed and in some cases quantified in paleosols. Physical features of paleosols are estimated by using empirical relationships or constants established by work on modern soils and calibrated by features that are measurable in paleosols.
Loosening Expansion of rocks and sediments accompanies the development of joints and soil cracks during weathering. This results in a lower bulk density, measured in grams per cubic centimeter (Klute 1994). Most rocks have a density more than twice that of water. They range from - ~more) high values for peridotite and basalt (3 g ~ m or ) . have to lower values for granite (c. 2 . 5 g ~ m - ~ Soils densities close to l . O g ~ m - ~ They . range from a low of 0.17 g in organic horizons to a high of 1.85 g cm-3 in subsurface layers (Fig. 4.1). Paleosols vary over this entire range for rocks and soils depending upon the extent to which alteration during burial has changed them into rocks. The original soil density of most paleosols cannot be determined because of compaction, cementation and other changes after burial. In somecases, however, compaction can be calculated from the folding of formerly straight and erect features of the soil such as skeletans or quartz veins. A compaction factor can be calculated from the present length of the vein by the shortest straight distance compared with the former length of the vein along its contortions (Table 4.1). Compaction factors also can be estimated from the behavior of comparable materials at the l i l y depth of burial (Caudill etal. 1997;Sheldon&Retallack2000).
Fluid flow The amount of water and gases that a soil can hold is 37
38
Chapter4
Table 4.1 Estimatingoriginal density,moisture equivalents andporosityof paleosols (fromBodman&Mahrnud1932; Birkeland 1999).
where symbols are: C, diagenetic compaction ratio (fractionof original thickness of soil as a result of burial) M , moisture equivalent (%) 0,porosity (x) P, volume (%) T,thickness (cm) v,volume (cm3) W, weight (g) X, proportion (fraction) p, density (g
and subscriptsare: a, sand c, clay f, buried or fossil soil i, silt m.n, for a given mineral s, original soil
related to the open spaces between the grains of the soil (Marshalletal. 1996).Of interestistheamountof water that can move through a soil in a given time (permeability), the amount of water that can be held in a soil (porosity)and the amount of water that can be retained by surface tension in the smallest cavities of the soil when the large pores have drained out (field capacity). None of these characteristics can be measured effectively in buried soils in which the cavities have been crushed or filled during burial. Some idea of former field capacity can be gained from paleosols by studying their grain-size distribution as established by point counting thin sections. Sandy soils are more porous and permeable, but have a lower field capacity than clayey soils. Field capacity can be approximated using a n empirical equation based on the percentageof sand, siltandclay(Table4.1). A more accurate estimate of former porosity can be obtained if the mineralogical composition and former density of the paleosol are known. The proportions of different minerals can be estimated by point counting thin sections.Originaldensity of thepaleosolcan beesti-
Soil-formingprocesses mated from its present density using compaction factors. Porosity is then calculated from the proportions of the various minerals, the density of the particular minerals from mineralogical reference works (Deer et d. 1992)andtheformerdensityof thepaleosol(Table4.1). A general idea of water flow and retention in soilsmay be gained from the distribution of root traces, burrows and soil structures. If root traces and burrows sidle along rather than penetrate a particular layer, that may have been an impermeable and hard barrier. Ferruginous concretions indicate oxygenated, free-flowing water. Sideritenodules. on the other hand, indicate stagnant, poorlyoxygenated water (seeFig.3.14). Clay swelling
The amount of physical expansion as a result of wetting of clay can be measured by changes in length or density Most clays expand to some extent when wet. This behaviorismostmarkedinsmectite (28 k Z%linearextensibility), and less so in illite (20 k 1%) and kaolinite (12 k 2%),as indicated by experiments (Nettleton & Brasher 1983) and numerical simulation (Karaborni at al. 1996). Burial compaction, dewatering and cementation of paleosols alters the swelling properties of clays. Their mineralogical composition also changes because of
regular mukkara
I
39
cations lost during dewatering. A well-known diagenetic alteration of clay is the formation of illite from smectite (Mora et al. 1998). Unraveling the amount and nature of such changes can be difficult and this compromises the use of typical values of expansion for different modern clays to interpret paleosols. Despite these difficulties, structural effects of clay heave can remain visible in paleosols. The most striking of these features is the hummock-and-swale microtopography called gilgai microrelief (Paton 1974). These natural undulations are thought to arise by swelling and upward buckling of the soil along deeply cracked hummocks. This may be exacerbated by dry soil falling into cracks, causing hummocks to heave further following the next rain. Hummocky subsurface differentiation is called mukkara structure (Fig. 4.2), and it has been recognized in very ancient paleosols (Retallack 1986a). Small-scale structures also may reflect former clay heave. Subangular blocky peds in heaving clay soils are often sheared by a conjugate system of slickensidesinto lentil and wedge peds (Krishna & Perumal 1948; Gustavson 1991).Microfabricalsomayshowarelationship to linear extensibility of clays, which is greatest (up to 12%)for woven bright clay (omnisepic plasmic) fabrics but minimal (up to 3.8%) for flecked bright clay (insepic plasmic) fabrics (Holzheyet ul. 1974).
distorted mukkaro
I
zone of dislorted mukkara
yellowish brown clay-
Figure 4.2 Mukkara structure in soilsof the Darling Downs, southeastern Queensland, Australia (fromGeoderrna 11,T.R.Paton 1974, Originandterminologyfor gilgaiinAustralia, p. 223; reprintedwithpermissionfromElsevierScience).
40
Chapter4
Fire heating
Freezing
Charcoal remains in the soil after a fire and is remarkably resistant to decay and chemical degradation (Retallack 1998a).Fossil charcoal has been found in paleosols and sediments as old as earliest Carboniferous (Cope & Chaloner 1985). In paleosols this old, it can be difficult to distinguish it from unburnt wood that has been coalified during burial. There are, however, a number of differences between charcoal and coal fragments (Table 4.2). The abundance of charcoal can be quantified by point-counting thin sections or by washing samples of known volume. A second indicator of fire is the pattern of stone spalling produced by forest fires over stony soils. The heating of stones creates concentric fractures. Curved exfoliated shells of rock are thrown off the littleweathered, rounded core of the boulder. In the absence of fire, on the other hand, rocks are more deeply weathered chemically and biologically. They may maintain an irregular shape covered with lichens and a rind of strongly weathered rock. Such differences between fired and unfired rocks are used to assess fire frequency in bouldery soils of alpine regions (Birkeland 1999).
Frigid periglacial climates create a variety of soil structures. The most distinctive of these are ice and sand wedges. Sand wedges are sharply tapering, vertically bedded, sand-filled structures (Fig. 4.3). Ice wedge casts are similar in outline but filled with slumped and horizontally layered sediment. Both sand and ice wedges have a strong taper that is very different from the deep, often contorted cracks of mukkara structure. Ice and sand wedge casts commonly are found in much coarser-grained material than the clayey soilsusual for mukkara structure. A variety of other periglacial features are known in soils and paleosols.String bogs, rock glaciers and pingos are a few examples among many (Washburn 1980: Karte 1983: Van Vliet-Lanoe 1985). Each forms in a particular modern climatic regime and this can be inferred for fossil examples (Willams 1986: Retallack & Alonso-Zarza 1998;Krulll999).
Table 4.2 Criteriafor distinguishingcharcoal from coaMed wood.
Fossil charcoal
Codified wood
Equant shape Sharplybroken or rounded ends Black and opaque Broken surface fibrous Broken surface showingcells No middle lamella between cell walls as seenwith SEM Resistant to oxidation Found inoxidized and gleyed paleosols Glows on burning
Usually elongatesplinters Irregularor frayed ends Brown to black Broken surface conchoidal Broken surface structureless Middle lamella visible between crushed cells under SEM Easily oxidized Found ingleyed paleosols only Burns with abright flame
SEM. scanningelectron microscope. Sources: Harris (1981);Cope&Chaloner (198 5).
Figure 4.3 Near-verticalsandstonewedge penetrating quartziteof the Cattle GridBreccia,both of LateProterozoic age(680Ma) intheMtGunsonMine.SouthAustra1ia (photographcourtesyof G.E. Williams).
Soil-forming processes
cal weathering. A working familiarity with common values for the weight per cent of oxides of major elements andof waysof manipulatingsuchanalyticaldata can beuseful for interpretingancient chemical weathering from paleosols. Minerals also are present in limited variety in soils compared with exotic crystals listed in mineralogical compendia (Deer et al. 1992). A few silicates (quartz, muscovite, microcline), clays (smectite, kaolinite and illite), carbonates (calcite, dolomite), sulfates (gypsum), oxides (hematite) and hydroxides (goethite) are the commonest of soil minerals. These minerals have characteristic optical properties and fields of chemical stability (see Fig. 3.14). Their textural relationships anddegree of etching observedin thin sections can be guides to the nature and severity of chemical alteration in paleosols. They preserve in an arrested state the progress of chemical reactions because chemical alteration in soilsis slow and seldomreaches thermodynamic equilibrium.
Indicatorsof chemical weathering Soil-forming chemical reactions are for the most part of four main kinds: hydrolysis, oxidation, hydration and dissolution (Table 4.3).The reverse reactions of alkalization, reduction, dehydration and precipitation also occur within soils and within paleosols during burial. The problem of recognizing chemical reactions in paleosols would be daunting were it not for such a limited array of common reactions. The problem also is simplified by the fact that most soil consists of only a few elements and minerals. Eight elements make up 98%by bothweight andvolumeof averagecrustalrocks and soils. These are (with average weight per cent in parentheses, after Mason & Moore (1982)):0 (46.60), Si (27.72), A1 (8.13), Fe (5.00), Mg (2.09), Ca (3.63)’ Na (2.83) and K (2.59). Trace elements may be important for interpreting many aspects of paleosols, but it is these major elements that definethe net effect of chemiTable 4.3 Common kinds of chemical reactions during weathering. I. Hydrolysis 2NaAlSi,O, albite
+
11. Oxidation
Pel+ ferrous ion
+
2Fe2+ ferrous ions
+
Fe3+ ferric ion
+
+
IV. Dissolution CaCO, + CO, calcite carbon dioxide
’/202 oxygen
+
CaSO, anhydrite
+
+
11H,O water
AI,Si,O,(OH,) kaolinite
+
2NaC sodium ions
+
+
4C0, carbon dioxide
+
2HC0,+ bicarbonate ions
e- (partialreaction) electron toother element
4HC0,+ bicarbonate ions
In. Dehydration 2FeOOH 4 Fe,O, goethite hematite CaS0,.2H,O gypsum
+
2C0, carbon dioxide
+
4H,O water
4
Fe,O, hematite
H,O water
+
H,O water
Source: Garrels&Mackenzie(l971).
2H,O water Ca2+ + calcium ion
41
2HC0,bicarbonate ions
6H,O water
4H,SiO, silicic acid
42
Chapter4
Acidification Hydrolysisis the reaction of carbonic acid with a cationrich mineral grain to produce clay and cations. Clay accumulates in the soil and cations are washed out or taken up by plants in solution. As an example, we may consider the hydrolysis of albite (Table 4.3). In essence, hydrolysis involves the displacement of cations by hydronium. It is the main chemical reaction resulting in the destruction of silicate minerals (Chesworth 1992; Bland & Rolls 199 8). A widely used average measure of the prevalence of hydrolysis reactions in modern soils is their acidity as measured by pH. This is technically the negative logarithm of the hydronium ion (Ht) activity. It is usually measured by a chemically treated electrode that develops an electrical potential with respect to a standard electrode when placed in a solution of equal amounts of soil and distilled water (Sparks 1996).This approach is not suitable for paleosols in which the soil has been so compacted and cemented that it is not possible to resurrect its originally reactive surfaces. Studies of Quaternary paleosols (Simonson 1941) have shown that pH changes substantially on burial of a soil or subsidence below the water table. Fortunately, the pH of paleosols can be estimated within broad categories from the stability fields of minerals thought to have been original (seeFig. 3.14)as well as by biological and structural features. The extreme range of soil pH from 2.8 to 10 can be subdividedinto six general categories (Baas-Beckinget al. 1960). Extremely acidic soils (pH 2 . 8 4 . 5 ) are noncalcareous and consist largely of peat, pyrite, quartz or a deeply weathered clay, such as kaolinite. They show little evidence of burrowing or plant decay. Organic acids are produced by plants and microbes, carbonic acid by dissolution of high concentrations of soil carbon dioxide, and sulfuric acid by oxidation of sulfides. Extremely acidic soils contain a dearth of exchangeable cations such as Ca2+,Mg2+,Nat and Kt and are dominated by hydronium (Ht) along with aluminum (A13+)through its reaction with water to produce Al-hydroxy ions (Al(OH)2tand Al(OH)2+). Moderately acidic soils (pH 4.5-6.5) are similarly noncalcareous and base poor, but may contain a wider range of easily hydrolyzed minerals such as feldspar and of base-rich clays such as illite or smectite. Limited bio-
logical activity may be evident from the presence of fecal pellets or burrows. Plant material in moderately acidic soilsshows some signs of decay. Near-neutral soils (pH 6.5-8) are weakly calcareous. Apatite, birnessite and siderite also are stable under these conditions, along with a variety of easily weatherable minerals. These soils are 70-90% saturated with bases (Ca2+,Mg2+,Na+and Kt) depending on the kind of clay present. They have no exchangeable cationic aluminum (A13+)and little exchangeable hydronium (Ht). Fecal pellets, burrows and other evidence of biological activity may be abundant innear-neutral soils. Moderately alkaline soils (pH 8-8.5) are calcareous with free crystals and nodules of calcite or other carbonate minerals. Their clays are cation rich (largely with Ca2+and Mg2+).Pelletoidal structure from the copious activity of soil invertebrates is common. Alkaline soils (pH 8.5-9) are similarly calcareous, nodular and pelletoidal. Their clays tend to be smectitic with exchangeable cations including sodium and potassium (Na+ and K+). These are swelling clays that may create mukkara structure (Fig. 4.2). Gypsum, anhydrite, dolomite and zeolites also are characteristic of alkaline soils. Extremely alkaline soils (pH 9-11) may be sodium saturated and contain halite, gypsum and other evaporite minerals (Chen 1997). These soils have base-rich clays such as smectites.Domedcolumnar ped structures arecharacteristic(McCahon&Miller1997).Roottraces and other evidence of life in these soilsare usually sparse because these soils form in very dry and saline desert basins. These broad categories of pH can be refined by calculating mineral weathering ratios within a profile. The abundance of easily hydrolyzed minerals compared with hydrolysis-resistant minerals increases with pH, among other factors. Such ratios are more meaningful whencompared with theratioof themineralsin theparent material and when there is independent evidence of the time over which the profile developed. A commonly used mineral weathering ratio is quartdfeldspar. Microcline and plagioclase feldspar may be distinguishable from quartz by their twinning, but it may be necessary to stain a thin section with potassium dichromate (Houghton 1980) to count proportions of orthoclase and quartz. Because feldspar is much more readily hydrolyzed than is quartz, this ratio is higher in acidic
Soil-formingprocesses soilsthan in alkaline. Another mineral weathering ratio is the ratio of zircon and tourmaline to pyroxene and amphibole. The latter two minerals are much less resistant to weathering than are zircon and tourmaline. These minerals do not occur in sufficient abundance to be counted in thin sections. They should be concentrated from known volumes of soil for counting. These mineral weathering ratios are based on the observed differential stabilities of minerals in soils (Fig. 4.4). Among mineral grains of sand and silt size, microcline (a potassium feldspar)is more resistant to weathering than is albite (a sodium feldspar). Both feldspars and m&c minerals (olivine, pyroxene, amphibole and biotite) are more readily hydrolyzed than is quartz (Goldich 1938).Quartzisnot so weather resistant when fine grained, however, because of its increased surface area (Jackson et al. 1948). From these general relationships a variety of mineral weathering ratios can be devised to suit particular profiles. For metamorphosed paleosols it may be worthwhile to calculate mineralogical composition from a chemical analysis. Various recipes and computer programs are available for calculating mineralogical modes of igneous rocks from a chemical analysis, but these are based on minerals and chemical reactions inappropriate for soils. For paleosols, once the likely minerals originally present have been established, their proportions can be calculated by ‘reacting’mole fractions of each major elemental oxide on a large balance sheet (Garrels & Mackenzie 1971). Such methods for calculating modes have been computerized (Reynolds 1985; Schmitt 1999).
An alternative approach is to use chemical analyses of paleosols to obtain molecular weathering ratios, which vary greatly with soil type (Fig. 4.5).These are calculated by dividing the weight per cent of each relevant oxide by its molecular weight and then adding or dividing as specified by the particular ratio. Molecularweathering ratios useful for assessing former extent of hydrolysis reactions are aluminalbases, alkaline earthslalumina, silicalalumina and bariumlstrontium (Table 4.4). Values of aluminalbases can reach 100 or more in strongly developed soils (Oxisols and Ultisols), but areless than two in moderately developedfertile soils (Alfisols and Mollisols).The ratio bariumhtrontium is an indicator of leaching, degree of free drainage and time of development. Strontium is the more soluble of these two otherwise chemically similar trace elements. The bariumlstrontium ratio ranges from near 10 in acidic, sandy soils (Spodosols),to near two in most rocks andsoils (Retallack 1997a). Another technique using chemical analyses is to calculate the loss of cationic bases in a soil or paleosol comparedwithits assumedparentmaterial. Acid(orproton) consumption in milliequivalents is assumed equal to each mole of alkalis (Na+ and K+) lost plus each halfmole of alkaline earths (Ca2+and Mg2+)lost (Table4.5). Acid consumption of soilsis usually about a milliequivalent per square meter per year under humid temperate forests (Retallack 1996b). A more rigorous method of analyzing chemical change is to normalize chemical compositionsto a constituent thought to have remained stablein both thesoil and itsparentmaterial(Table4.6). This is a more complexprocedure that relies on a variety
STABILITY SERIES sand- and silt-size mineral grains __ ---
CO
K feld~ar
muscovite ’
least eastly
quartz-
WEATHERING INDEX clay-size mineral grains
- -- -_--.‘ .,..Z.calcite, 1.gypsum. holite apatite
* -
most easily weathered
Figure4.4 Relativestability of mineral grains of silt size or larger (left-hand side)and of clay-sizeminerals (right-hand side)under conditions of weathering in soils (data from Goldich 1938;Jacksonetd. 1948).
43
feldspar
~
- -
,
* ,‘ ~
‘
‘,3.olivine, pyroxene 4. biotite, glauconite 5 albite, pnorthite, volcanic glass 6. quartz, cristabalite ,/Z muscovite, seridte 8. vermiculite 9. smectite 0. kaolinite halloysite, allophdne II. gibbsite, boehmite 12. hematite, goethite 13. anatase, rutile. zircon
44
Chapter4 niolecular weathering ratios field observations b,Q CaOtMEO MA?, 2 K,O AI,O, SiO, Munsell color b \ percent grain size c Q soil horizon 50 2 4 0 01 020
2
cm
'*"*
dc.'"
a
'4)
Bt
a
C
a
LLU
-
m
ULTISOL (Typic Paleudult) Cecil sandv loam. Milne
amar
"1
gravel
n
b CA
sand silt
E ul clav-
gg granular peds 0
calcareous nodules
L
-
rm
ALFISOL (Tvoic Haoludalfl Miami silt loam. WestDo
MOLLISOL (Udic Haplustoll)
Decatu
LIlI
fl
rrrr
- Moody silt loam, Moo co.,SI
a
.Laveen loam, Buckeye,
LLLL
L
Figure 4.5 Profiles,grain-suedata andmolecular weathering ratios forselectedNorth American soils (from Retallack 1997, A ColourGuide toPaleosols;withpermission0 JohnWiley&SonsLtd:datafromMarbut 1935;Davis 1972;Hartman 1977; Shively 1983:Kunze 1989).
Soil-forming processes Table 4.4 Major kinds of chemicalreactions in paleosolsand molecular weathering ratios used to approximatethem.
1. H y h o b s i s
alumina/bases = AI,O,/(CaO +MgO + K,O alumina/silica = AI,O,/SiO, barium/strontium =Ba/Sr
45
+Na,O)
II. Oxidation ferrous/ferric iron =FeO/Fe,O, total iron/alumina= (Fe,O, +FeO)/AI,O, total iron and manganese/alumina= (Fe,O, +FeO+ MnO)/AI,O,
HI. Hydration silica/sesquioxides= SiO,/(Fe,O,
+A1,0,)
n! Saliniaation akalis/alumina= (K,O +Na,O)/AI,O, soda/potash=Na,O/K,O soda/alumina=Na,O/AI2O3 V. Atomic weights Si=28.09 A1=26.98 0=16.00 Fe=55.85
of assumptions about parent materials (Brimhall et al. 1991;Beauvais& Colin 1993 ) . Oxidation Oxidation reactions are those in which an element suffers electron loss when forming a compound. The reverse reaction of reduction is more easily memorized from an apparent oxymoron: ’reduction is electron gain’. Reduction occurs in waterlogged soils and in buried soils, but for well-drained soiIs oxidation is the rule. Oxygeninthe atmosphereis themostpowerfulnaturally occurring electron scavenger and elements combining with oxygen surrender their electrons to its orbitals. Of the major elements making up most Earth materials, manganese and iron can readily spare an electron, and of these two elements iron is much more abundant. At its simplest level the oxidation of iron can be expressed by a partial reaction for iron (Table 4.3). Minerals containing iron in the ferrous valence state (Fe”), such as olivine or siderite,tend to be gray or green in color. Ferric (Fe3+)iron, on the other hand, is found in minerals, such as goethite and hematite, that have yellow, brownandredcolors (Hurst 1977).Whenliberated from silicate minerals by hydrolysis, ferrous cations of iron are readily soluble in water and often are washed
K=39.10 Na=22.99
Ca=40.08 Mg=24.32
Ba=137.3 Sr=87.6
out of the profile. Ferric iron cations are relatively insoluble in water and so remain as oxides and oxyhydrates. The formation of hematite from olivine is a two-step process of hydrolysis followed by oxidation (Table4.3). An overall measure of the extent of oxidation and reduction reactions in modern soils is Eh. This is the electrode potential in volts induced by soil solutions. Low negative Eh is an indication that the soil can readily donate electrons and is reducing. A high positive Eh, on the other hand, is found in oxidizing soils with a high demand for electrons. The principal oxidizing agent in modern soils is free oxygen from the atmosphere. Its exclusion by waterlogging creates reducing conditions (Vepraskas & Sprecher 1997). Because the water content and pore structure of paleosols is so altered by burial, it is not possible to establish their former Eh by direct measurement. Like pH, Eh is altered substantially with burial or subsidence of a paleosol below the water table (Bohnet al. 198 5). The former Eh of paleosols can be reconstructed within broad limits from mineral assemblages thought to have been original (see Fig. 3.14).Because chemical reactions of mineral transformation that define these categories also are pH dependent, it is not possible to give boundary values of Eh categories in millivolts. However,
46
Chapter4
Table 4.5 Acid and oxygen consumptionof soils and paleosols.
Table4.6 Formulaerelatingvolume, thicknessandchemicaI changesin soil and paleosol horizonsduring soil formation
and burial compaction.
B=2p,(O.O1783C+0.02481M+0.01062K
+O.O.O613N)/lOO I=p(O.O1392F)/100 TI= [(D,+D, - 1 )-(DI+D,+ I)I/2 A=ZT,(B,- B,)
o=ZT,(I,-I,)
symbols: A, acid consumptionof profile (equiv.cm-2) B,base content of sample(equiv.cm-’) 1. reduced iron of sample (equiv.~ m - ~ ) C.M.K.N.F. CaO, MgO. K,O, Na,O, FeO (wt X ) (respectively) 0,oxygen consumptionof profile (equiv.cm-I) D, depthto sample (cm) T,thicknessrepresentedby sample(cm) p, bulkdensity ( g ~ m - ~ ) subscripts: i. for sample i i + 1,for sample or surfaceabove i i - 1,for sampleor parent material below i p, for parentmaterial constants: 0.01783. etc.,element weight fractionof oxides 100,percentage adjustment Sources: Holland (1984); Retallack (1996b); Bland & Rolls (1998).
there are three recognizable categories of Eh: reducing, intermediate and oxidizing. These correspond to waterlogged, wet and well-drained soils (Baas-Becking et al. 1960).This may not always have been the casebecause atmospheric oxygen may have been lower in the distant geological past (>2200Ma: Rye&Holland 1998). when well-drained soils would have been reducing. This may be a problem for interpreting the paleodrainage and former atmospheres of such ancient paleosols, but it does not affect their inferred redox status. Reducing soils (strong negative Eh) are usually bluish or greenish gray. The iron in their minerals is in the ferrous state (Fez’), which is naturally drab colored. Additional grayness may be due to organic matter (Retallack 1994a).Thesearethesoils that accumulate peat, which on burial and diagenesis is converted to coal. One bright
variables:
E , strain, or volume change comparedwith initial volume
p. dry bulk density (g ~ m - ~ ) z. transportedmass fraction, gains or losses or concentration
ratio of a chemicalconstituentcomparedwith parent material ( g ~ m - ~ ) C, chemicalconcentrationof a constituent (wt X ) F, compactionratio or fractionof originaIthicknessas a result of burial compactionor of soil formation R, amount of chosen stableconstituent (wt%) S, weight of a chemicalconstituentnormalizedto a stable constituentremainingfrom parent material (g cm-’) T.thickness(cm) W, weight (g) v. volume (cm3) X. weight of achemical constituentin a given volume (g ~ m - ~ ) subscripts: f, for the buried or fossil soil p. for the parent material s, for the originalsoil Sources: Brewer (1976):Chadwicketal.(1990).
mineral associated with reducing soils is pyrite, often in the form of framboids (Fig. 4.6).In natural exposures of paleosols pyrite may be destroyed by modern oxidation. I t can still be detected as dustings of its weatheringproduct: a light yellow powder of jarosite, which has a characteristic smell Like rotten eggs (Kraus 1998). Another distinctive mineral of reduced soils is vivianite
Soil-formingprocesses
sphoerosiderite
Figure 4.6 Some common constituents of coals and waterlogged paleosols.In some cases (botryococcoidalgae, megaspores, cuticles, coaliEedwood)specimens show marked compaction and in other cases (sphaerosiderite,pyrite framboids and sclerotinite) little change as a result of compaction.
pyrite frornboid
botryococcoid algal colony
leof cuticle
(Bloomfield 1981).This is white as nodules or replaced arthropod carapaces when reduced, but turns a distinctive bright blue color when mildly oxidized.Strongly reduced soils also contain well-preserved plant remains because lack of oxygen curtails the activities of aerobic microbial decomposers. The proportion of amorphous compared with structurally preserved plant material, or of fossil roots compared with shoots, is auseful indicator of the degree of former oxidation of a paleosol. The activity of large animals and plants in reducing soils also is limited. Stronglyreduced soils contain few burrows and fecal pellets. Root traces do not penetrate them deeply, but spread in a tabular pattern (see Fig. 3.3). Soil structures also are weakly developed in reducing soils. Original sedimentary structures such as bedding and ripple marks may persist. On a microscopic scale this may be reflected in striated (mistrial plasmic) or in dull cloudy (undulic and inundulic) fabrics (Brewer 1 976). Soils of intermediate redox status (near-neutral Eh) may be either partly or periodically waterlogged. Many swamps experience a season dry enough for forest fires that introduce charcoal into the peat. Charcoal is a distinctive component of some coal (Fig. 4.6, Table 4.2). Seasonally dry soils also may be strongly mottled. During gleyed conditions of a wet season, hydrolysis liber-
47
fungo1 sclerotium
megospore
chorcoal (burnt wood) coolified wood (unburnt)
ates drab ferrous iron (Bennett et al. 1991). During the dry season this may be oxidized and fixed as orange goethite (McCarthy et aI. 1998). Two common kinds of mottling associated with seasonal waterlogging can be recognized (Fig.4.7). Usually the lower part of a paleosol below the water table is more gleyed than the upper part. In very clayey soils impermeable layers may form a barrier to downward percolation of water (a perched water table), which can lead to greater waterlogging of the upper than the lower part of the profile. The difference between groundwater and surface-water gley is most apparent around burrows and root traces, which are more reduced than the soil matrix in surface-water gley, but more oxidized in groundwater gley (PiPujol & Buurman 1997). Different kinds of mottling must be interpreted with care for paleosols because drab-haloed root traces (see Figs 3.6 and 3.7) may be a consequence of burial rather than reflecting original conditions. In addition to these mixedredox indicators, soils of intermediate oxidation contain more mildly reduced minerals and lack the sulfideminerals of extreme reduction. Characteristic of soils of intermediate redox status are minerals such as birnessite and siderite (Moore et al. 1992).Sand-sized spherical aggregates of siderite crystals (sphaerosiderite) are common in paleosols of coal
48
Chapter4
GROUNDWATER GLEY
0
silt. .. .. . .. sand ,
a -
Bray clay
SURFACE-WATER GLEY brown clay
iron, stain
(>:-’
measures and may be products of original soilformation at intermediate Eh or of gleization attendant on burial (Ludvigsen et al. 1998). Siderite is readily oxidized to goethite or hematite. All three minerals may be present in the same paleosol as a result of an original soil regime of fluctuating oxidation, or oxidation of originally gleyed minerals in the modern outcrop. Fluctuating oxidation in the original soil would be likely if the paleosol also showed a mixture of well-decomposed and recognizable organic matter or patterns of root traces that include both deeply penetrating and tabular root systems. Swamps that are periodically dry have very low rates of peat accumulation. Paleosols formed under such conditions have little or no coal (Retallack 1994a). Fecal pellets and burrows may be present and also some crude blocky and platy ped structures. Microfabrics of soils of intermediate redox status may include bright clay (sepic plasmic fabric). Oxidized soils (high positive Eh) are warm colored with iron oxyhydrates such as brown ferrihydrite and yellowish brown goethite, and oxides such as brick-red hematite. The variation in color of these oxidized minerals can be related to light absorption by varying proportions of 0x0- and hydroxyl-groups (Rossman 1996). Hematite is especially common in paleosols because of dehydration reactions accompanying burial (Blodgettet al. 1993). Organic matter is not noticeable in oxidized
siderite, pyrite
Figure4.7 SchematicmodeIs for the formationofgroundwatergley(1eft)and surface-watergley (right).
soils except near the surface and around living roots. Dead organisms in well-drained soils are decomposed and recycled by a community of microbial decomposers. Although organic matter may be scarce, there may be abundant evidence of life in oxidized soils in the form of burrows, fecal pellets and copiously branching, deeply penetrating root systems. Well-drained soils also have well-developed soil structure including peds. illuviation clay skins and sepic plasmic microfabrics. Mineral weathering ratios are not especiallyuseful for quantifying soil oxidation because iron oxide minerals are too fine grained and dispersed to count easily. Oxidation state of a paleosol can, however, be chemically characterized using molecular weathering ratios (Table 4.4), by normalizing the relevant chemical analyses to their abundance in parent material (Table 4.6), or by calculating equivalents (i.e.electrons) used to oxidizereducediron (Table4.5).Ferric toferrousironratiosare an effective proxy measure of oxidation state. In many methods of bulk chemical analysis the two oxidation states of iron are not distinguished. The weight per cent value from such analyses may be cited as ferric (Fe,O,) or ferrous (FeO) iron. An independent method of determining the abundance of one of the oxidation states is needed: for example, potassium dichromate titration (Maxwell 1968).The abundance of the other oxidation state can be estimated by molar difference. When such additional analytical data are not available, the total
Soil-formingprocesses iron or total iron plus manganese values may be useful guides to the degree of oxidation of a paleosol. The oxidized minerals of these elements are insoluble and accumulate in soils, whereas their reduced cations are lost from gleyed soils in groundwater flow. Total iron to alumina ratios calculated for a variety of North American soils are mostly less than 0.4 (Marbut 1935). The ratio can reach 1.2 in ferruginous (spodic) horizons, 1.9 in clayey (argillic) horizons and 1.5 in red and deeply weathered (oxic)horizons. Hydration Hydration or dehydration reactions involvethe addition or loss,respectively, of water that is structurallypartof a mineral. Hydrated soil minerals such as goethite and gypsum are formed by hydration reactions and also can dehydrate tononhydrous minerals such as hematite and anhydrite (Table 4.3). The degree of hydration of a soil can be approximated by identifying and quantifying the abundance of these minerals using Mossbauer spectroscopy for iron oxyhydrates (Dixon & Weed 19 89) and X-ray diffractionfor clays (Klute 1994).Moredirectestimates of hydration can be gained by weighing ovendried soil after heating to temperatures that have been shown to drive off different volatile fractions. These volatile materials include organic matter and carbon dioxide as well as structural water (Sparks 1996). The degree of hydration of paleosols could be determined in the same way were it not for dehydration reactions following burial and compaction. Gypsum, for example,isdehydrated toanhydrite (Boggs 1995). Ferrihydrite and goethite are converted to hematite (Blodgett et al. 1993). Smectiteis altered to illite, then sericite and ultimately muscovite (Mora et al. 1998). The transition from acicular low-density gypsum to orthorhombic high-density anhydrite may be obvious from the relict crystal outlines of the formerly larger crystals of the hydrated mineral. For fine-grained minerals such as hematite and illite, however, there is no such guide. Exactly how hydrated they may have been is difficult to estimate, but hydration is likely to be proportional to the abundance of these minerals. Molecular weathering ratios also may be a crude guide to former hydration of minerals in a paleosol (Table4.4).Theratioof silica to sesquioxideswillbehigh in quartz-rich paleosols with few hydrated minerals.
49
Conversely, the ratio is low in clayey ferruginous soils, which may have been full of iron oxyhydrates and hydrated clays. Salinization The disappearance of minerals into solution and the opposite effect of precipitation of soluble salts, are widespread in soils. Dissolution is a reaction in which a compound disaggregates into its constituent ions in water like a cube of sugar in tea. Hydrolysis reactions also produce ions in solution, but during hydrolysis a new insoluble product, such as clay, is produced. A common dissolution reaction in soils is the dissolution of limestone in a weak solution of carbonic acid derived from atmospheric carbon dioxide and water. In modern soilsa useful indicator of dissolution is the ionic concentration in soil water. A good indicator of salinization is the saturation conductivity of the soil. This is measured using a Wheatstone bridge and electrical conductivity cell on a vacuum-filtered extract of saturated soil paste (Klute 1994). The conductivity of the paste increases with salt content. Neither of these direct measurements is useful for paleosols, but both dissolution and salinization may leave observableeffects.Etching and pitting of mineral grains can be observed in thin sections or scanning electron micrographs of paleosols,but it is difficultto distinguish the effects of dissolution during soil formation from former hydrolysis in the soil or from comparable diagenetic changes after burial. Dissolution is more usual in carbonate or sulfideminerals than in silicates and it may be especiallyobvious in calcareous or evaporitic paleosols. Karst topography is a striking geomorphological example of limestone dissolution. The irregular ferruginized dissolution planes in limestone (stylolites)are another example on the scale of hand specimens (Bathurst 1975). Salinization in paleosols is more amenable to analysis than is dissolution. Salts common in soils of very dry climates include halite, gypsum and mirabilite. Also characteristic of very saline soils are the phyllosilicates palygorskite and sepiolite.X-ray diffractionto determine relative abundance of these relative to other minerals may provide a useful index of salinization. On burial of a soil, most evaporite minerals are subject to alteration by dissolution of salts and by dehydration of zeolites and
50
Chapter4
hydrated phyllosilicates. However, the highly alkaline conditions under which they form in soils also promote dissolution of quartz and other forms of silica, which may cement indurated horizons of a soil. Such aridland chert may be firm enough to preserve undeformed pseudomorphs of the original crystal outlines of soluble salts (Barleyet al. 1979) and brittle enough to break into stratabound brecciatedzones after their dissolution and removal (Bowles & Braddock 1963). Deformation of sedimentary layering by salt (haloturbation) as puffy grounds can also be distinctive (Malicse & Mazzullo 1996). In very ancient rocks such pseudomorphs and breccias constitute the main evidence for former evaporites. Molecular weathering ratios are not especially useful for assessing dissolution. Salinization may be approximated by ratios of alkalis to alumina, soda to potash or soda to alumina (Table4.4). Salinizationcan be indicated bysoda/potashratios greaterthanoneorthisratioincreasing up-profile. Sodium is generally a more reliable indicator than potassium because it is more soluble than potassium and also less susceptible to diagenetic alteration (Retallack 1991b).
Indicatorsof biologicalweathering The influence of organisms in soil formation is so pervasive that it is difficultto draw the line between biological and other weathering processes.This is particularly true for soil microbes, which rapidly exhaust soil oxygen and coat soil peds with slimy polysaccharides (Hiebert & Bennett 1992). It is partly for this reason that soil physics and chemistry have remained so empirical in orientation. Much can be learned by applying thermodynamic chemical models or the theory of fluid mechanics to soils,but predictivephysicochemicalmodels continue to be elusive because of complications imposed by soil biology. From an organism’spoint of view, including our own agricultural perspective, the most important biological processes are those affecting productivity of the soil. Three of these will be considered here. First, humification is a biologicalprocess by which organic matter from once-living creatures is reincorporated into the soil for further use by organisms. Second,nutrient utilization is the degree to which essential elements are used within soil ecosystems. For example,magnesium from the soil is
a vital constituent of chlorophyll, which makes plants green and allowsthem to harness theenergy of sunlight for the photosynthetic manufacture of organic matter. Third. bioturbation is the degree of soil churning by organisms. The activity of roots and burrowing animals has profound consequences for soil structure. Humification Progressive decay at the surface of modern soils can be followed from the time a leaf is shed from a tree to its attack by successive waves of microorganisms on the ground and its final disappearance as amorphous dark organic matter. Chemically speaking, this is a n oxidation in which complex reduced organic compounds are converted to more simple ones such as carbon dioxide. Physically, it is a comminution of large structures into fine-grained amorphous organic matter. In practice, however,it is a series of life activities of successive waves of decomposingorganisms. Soft watery cell contents are consumed first by a variety of soil microarthropods and worms. Then the decomposing remainder is broken down further by fungi and bacteria. Only certain kinds of microbes can break down cellulose cell walls. Very few can destroy lignin, which commonly remains as a woody leaf skeleton after the rest of the leaf material has been destroyed. A good deal of the material of the leaf is metabolized by soil organisms and then by organisms that feedonthem. Someorganicmatterisreleasedto the soil in asoluble (humic. fulvicor other acids)or insoluble form (polysaccharides).The process of humification in modern soils can be characterized by the degree of destruction of identifiable plant fragments as well as by analysis for the amount of humic or fulvicacids or amorphous organic carbon in the soil. Such analyses may involve treatment with an alkali to separate soluble and insoluble fractions of organic matter followed by treatment of the soluble part with acid to separate humic from fulvic acids. Total organic matter can be determined by the loss in weight during combustion in a furnace of a soil sample already treated with acid to remove carbonate. An alternative method for estimating content of organic matter is the Walkley-Black titration procedure, in which ferrous ammonium sulfate neutralizes chromate utilized in the oxidation of organic matter within a solution of potassium dichromate in sulfuric and orthophosphoric acid (Sparks 1996).
Soil-formingprocesses Variations in the isotopiccompositionof organic carbon within soils and paleosols may also reflect humification (Sharpenseel &Pfeiffer1997). Each of these techniques for estimating the degree of humification of modern soils also can be used for paleosols, although not always with meaningful results because of modification of organic matter during burial. Diagenetic alteration affects fine-grained organic matter more than recognizable plant fragments so that the larger pieces are an especially useful guide to the degree of humification experiencedby a paleosol. Three general stages of humification can be recognized in leaf litter and surface horizons of soilsand their fossilized equivalents. This kind of fossil leaf locality is very differentfrom those formed within the fine layers of lacustrine shales or incorporated in river and delta channel deposits. In fossil litters, leaves are mixed with roots of various kinds that pass downward into the paleosol. The plant remains show varying stages of comminution, skeletonization, nibbling by insects, clumping because of the gluing effect of decompositional polysaccharides and creasing from the activity of leaf miners and fungal hyphae. A large range of such features distinguish a fossil moder humus of the kind found under broadleaf oak-hickory forests of moderately neutral to alkaline soils. A mor humus, on the other hand, shows little evidence of humification. This kind of humus is characterized by a high proportion of intact leaves and other plant remains. A typical mor humus is the carpet of pine needles found under conifer forest. The needIes are protected from decay by abundant phenols and resins within the needles and acidic soil conditions hostile to decomposingmicrobes. A mull humus, at the other extreme, contains little in the way of identifiable plant fragments. Most have been broken down into he-grained organic matter intimately admixed with clay of the soil. Mull humus is formed by the activity of productive decomposer ecosystems, which may leave abundant evidence in the form of small ellipsoidalfecal pellets.These often dominate the microfabric of mull humus. The rich, dark surface horizons of grassland soils are a good example of a mull humus. These three kinds of humus found in well-drained soils can be recognized in paleosols (Retallack 1976: Wright 1983) and quantified in them by point-counting thin sections or estimating from fossil plant collections the proportions of amorphous organic matter, structureless plant
51
fiber and pIant fragments with clearly recognizable cellular structure. Depending on which of these predominate, fossil humus can be classified as mull, moder and rnor. A comparable division of the thick peaty surface horizons of wetland soils also can be applied to the interpretation of humification in coal-bearing paleosols. Fibric peats are those formed in such stagnant or acidic bogs that more than a third of the plant fragments retain recognizable cellular structure. Sapric peats, on the other hand, contain less than a third of recognizable plant fragments and appear structureless and fine grained in thin section because of extensive decay in periodically well-drained or alkaline bogs. Hemic peats contain intermediate amounts of partly decayed plant fiber. These three kinds of peat observed in modern peaty soils may be distinguished in coals of paleosols using techniques of coal petrography (Diessel 1992). A terminology different from this approach of soil science has been used to describe coals and also can be used to assess humification in coal-bearing paleosols. A variety of coal measure rock types can be distinguished in the field (Table 4.7).but are better characterized microscopically (Fig. 4.8).Coal is very dark to opaque in thin section, and polished thick sections examined under a microscope by reflected light are more informative. Polished surfaces are also easier to make than thin sections because many coals are brittle and fractured (Laubach et al. 1998). Mineral grains are a minor component of coal. Most coal is made up of plant fragments called macerals, which can be classified by optical and shape characteristics (Table4.8).The vitrinite maceral group is gray and yellowish white with low relief under the reflected light microscope, and light orange to dark red in transmitted light. The inertinite group is light gray to white in reflectedlight and nearly opaque in transmitted light, but not nearly so bright or opaque as pyrite. The liptinite group is generally dark gray to black in reflected light and yellow to yellowish brown in transmitted light. Each kind of liptinite maceral has a distinctive shape that can be a clue to its identication (Fig. 4.6). Point counting of thin sections or polished surfaces for decayed macerals (collinite, micrinite and semifusinite) vs. undecayed macerals (vitrinite, sporinite) can yield quantitative data on the degree of humification of a coal (Diessel 1992;Parrish 1998). Paleosols also can be analyzed for total organic car-
52
Chapter4
Table 4.7 Rock types of coalmeasures. Rock type
Description
Vitrain
Bright coal: vitreous to subvitreous luster, conchoidal fracture, brittle, may contain a few ( ~ 1 0 %thin ). (<5 mm) dullcoal bands Bandedbrightcoal: mainlybrightcoalcontainigcommon( 10-40%),thin(<S mm) dullcoalbands. even fracture Banded coal: thin (<5 mm)bright and dull coal bands in proportions of 40-60% each Bandeddullcoal: mainlydullcoal withcommon (1040%),thin (<5 mm) bright bands, uneven fracture Dullcoal: mattluster anduneven fracture, maycontainsome(
Clarain Duroclarain Clarodurain Durain Fusain Cannel coal spores Torbanite Shalycoal Coaly shale Carbonaceousshale Shale Source: Diessel(1992).
bon. Unfortunately,such analyses cannot automatically be assumed to reflect original organic matter distribution. Studies comparing modern surface soils with equivalent buried Quaternary paleosols have shown that paleosols have only a fifth to a tenth of the organic matter originally present (Retallack 199lb). Some loss of organic matter is due to stripping of vegetation and erosion of the soil surface shortly beforeburial. Moreimportant for soils that do not subside into a chemically re-
ducing zone below the water table is consumption of organic matter by microbes deep within the overlying soil. Because of these losses of organic matter after burial, the shape of the organic matter profile rather than its magnitude is more important for interpreting paleosols (see Fig. 7.3). Steadily declining organic matter with depth is the rule with grassland soils, but a subsurface accumulation may be found in woodland and forest soils (Stevenson 1969).
Figure 4.8 Broken fragments of semi-fusinite (light gray, upper left)flanked by vitrinite (gray)with fragments of fusinite (white, right) in a polishedsection viewed under reflectedlight of the upper Elkhorn Number 3 coal fromLetcher County,Kentucky,USA (photographcourtesy of J. Crellig). Scale bar represents 50pm.
Soil-formingprocesses Table 4.8 Coal macerals and their origin. Group
Maceral
Origin
Vitrinite
Telinite Collinite
Wood Decayed wood
Inertinite
Fusinite Semifusinite Macrinite
Sclerotiite Inertodetritinite
Charcoal Degradedcharcoal Coarse (1040pm) soot or decay Fine (c10pn)soot or decay Fungal sclerotia Fragmentedinertinite
Liptiite
Sporinite Cutinite Suberinite Resinite Alginite Liptodetritiiite
Sporesand pollen Plant cuticles Plant corkycellwalls Plant waxes and resins Algae Fragmentedliptinite
UV fluorescent
Fluorinite
Lenses of plant oil. decay Streaky,groundmass decay Cavity-fillingdecay
Micrinite
Bituminite Exsudatinite
53
metamorphism of the greenschist facies (Summons & Hayes 1992). Many paleosols formed in well-drained parts of the landscape or subsequently highly metamorphosed contain negligible amounts of organic carbon. Fossil root traces and burrows may provide evidence that it was once present despite very low analytical values. In these cases trace elements characteristically complexed with organic matter in soil may provide a useful proxy indicator of the former profile of organic matter. Copper (Cu), chromium (Cr), nickel (Ni) and zinc (Zn),tend to follow organic matter and clay in soils.Their variation within a profile can be normalized to trace elements characteristically stable in soils, such as lead (Pb) and zirconium (a). Phosphate (P,O,) also may show a distribution parallel to that of organic matter in soils (Smeck 1973) and a number of rare earth elements, such as yttrium (Y) and lanthanum (La), commonly follow phosphate. These various trace elements can be assayed by methods such as neutron activation analysis and inductively coupled plasma fusion spectroscopy. A growing data base of trace elements in modern soils is increasing our understanding of their abundance and behavior during weathering (Wedepohll969-78; Aubert & Pinta 1977; Kabata-Pendias&Pendias 1984).
Source: Diessel(l992).
Nutrient consumption Paleosols with low amounts of organic carbon can retain sufficient carbon so that humification can be assessed from depth functions of carbon isotopic composition (Krull&Retallack2000). Within the range of values found within a soil or paleosol profile, the isotopically heaviest carbon is in the subsurface clayey (Bt) horizon, where humification is most marked (Fig. 4.9). The isotopicenrichment (measured as 6l 3C0,g)of the Bt horizon relative to other parts of the profile is significant (2%0) in weakly developed soils (Inceptisols), more marked (So/,,) in strongly developed soils (Alfisols, Ultisols) and very marked in rice paddy soils (7-8%0). Peaty soils (Histosols) show slight depletion ( 1-3YoO) at the surface, and swelling clay soils (Vertisols) have comparable isotopic composition throughout the profile (Sharpenseel & Pfeiffer 1997). Carbon isotopic composition of organic matter is resistant to alteration after burial, until alteration of D/H ratios during
Biological productivity of soils is determined in part by the availability of nutrient elements needed for metabolism. For large plants the most needed nutrients (macronutrients) are hydrogen (H), carbon (C), nitrogen (N), oxygen (0).magnesium (Mg),phosphorus (P). sulfur (S), potassium (K) and calcium (Ca). Lesser amounts of the following elements (micronutrients) also are necessary for large plants: boron (B), chlorine (Cl). vanadium (V), manganese (Mn), iron (Fe), copper (Cu). zinc (Zn) and molybdenum (Mo). Hydrogen, carbon and oxygen are readily available from air and soil as molecular oxygen, carbon dioxide and water. There is also nitrogen as molecular nitrogen in the air, but large plants are unable to fix it in that form. Nitrogen, phosphorus and sulfur are obtained in solution as cations such as nitrate, sulfate and phosphate derived from other formsof theseelements by soilmicrobes.Theother plant nutrients are obtained by roots as ions in solution resultingfrom hydrolyticweathering (Stevenson1986).
54
Chapter 4
field observations
YOorganic
6'3C -25
molecular weathering ratios
carbon
-1:
+
-%OL CsO + MgO Na,O K,O
+
peFcent grain size
Ba/Sr
10
LLLL
50
:lay
Figure 4.9 Variation in carbon isotopic compositionand organic carbon content of a strongly developed pakosol (Sombrihumult) from theLower Triassic (245 Ma) FeatherConglomerate inthe Allan Hills,VictoriaLand,Antarctica,showingmaximally heavy isotopicvalues(613C,,,)inthe clay-enrichedsubsurface(Bt)horizon,thought to reflectmaximal humificationthere (seealso Fig. 10.1 1:fromKrull&Retallack 2000:with permission from the GeologicalSociety of America).
The nutrient requirements of animals include the macronutrients of plants together with sodium (Na) and chlorine (CI). Micronutrients of animals include also fluorine (F), silica (Si). chromium (Cr),nickel (Ni), cobalt (Co), arsenic (As), selenium (Se), tin (Sn) and iodine (I). Most of these are taken in as food but some (notably Na and Cl) may be obtained from salt licks and in some cases by eating soil (Jones& Hanson 198 5). Most of the macronutrients provided by soils form positively charged ions (cations) in solution. A good overall indicator of their abundance in soils is cation exchange capacity (CEC). This has been traditionally determinedby displacing cationswith ammoniumchloride and measuring its abundance by titration. Cation exchange capacity is a measure of freely available
cationicbases,mostlyCa2+,Mg2+,K+andNa+.Twoother cations abundant in soils (A13+and H+) generate acids and these may be assayed by similar displacementtechniques to measure exchange acidity. Base saturation is another measure of soil nutrient availability. It is calculatedas thepercentageof thetotalcations that are basic. Available phosphorus, nitrogen and sulfur also may be measured in soils using similar displacement techniques. Available phosphorus, for example, may be taken as that extracted with a 0.5 M solution of sodium bicarbonate (Sparks 1996). These techniques for soils are not suitable for determining the base status of paleosols, in which nutrients formerly available on the surfaces of weathered grains are the most liiely to have been recombined during bur-
Soil-forming processes
55
100
80 BOUND H AND AI 60
w
W 2 40
I Q
Figure 4.10 Relationshipbetween soil acidity and cation exchangecapacity(fromBirkeland 1999;reprintedwithperrnissionfromOxford University Press).
::20 W
0
3
ial. However, CEC and base saturation are loosely linked with soil pH (Fig. 4.10) and weatherable-mineral content (Nettleton et al. 1998). Extremely acidic soils (pH 2 4 . 5 ) are noncalcareous and have a high exchange acidity related largely to A13+.Both their CEC and base saturation are low. Moderately acidic soils (pH 4.5-5.8) are similar but their exchange acidity is due largely to hydronium (H+).Weakly acidic soils (pH 5.8-6.5) may have rather high CEC and base saturation (70-90%) depending on the kind of clay present. Higher CEC values in the range for each kind of mineral represent finer grain sizes (Table 4.9). Thus noncalcareous paleosols with appreciable amounts of montmorillonite are likely to have been fairly highly base saturated. Near-neutral soils (pH 6.5-8) are weakly calcareous with little remaining exchange acidity.Their CEC is moderate to high depending on the kind of clay present, and their base saturation is close to 100%.This also is true of alkaline soils (pH 8-1 1).Excess calcium and magnesium may be obviousin the form of powder and nodules of carbonate in moderately alkaline to aIkalie soils (pH 8-10). In extremely alkaline soils (pH 9-1 l),sodium dominates the exchange complex and an excess of this element may be detectable in traces of salts or soil structures such as domed columnar peds (McCahon & Miller 199 7). Although nutrients are abundantly availablein extremely alkaline soils,these are soils of dry. evaporitic climates in which water is so scarce as to limit growth. The availability of other nutrients such as nitrogen, phosphorus, sulfur and chlorine seldom can be assessed from mineralogical observations because these nutrients are usually only minor constituents of soils that are actively recycled by microorganisms (Stevenson 1986).
5
4
6
7
8
9
I0
SOIL pH
Table 4.9 Cation exchangecapacity(CEC) of common soil materials. Mineral
CEC(mequiv.per 1008)
Humus Vermiculite MontmorilIonite Saponite Nontronite ILlite Kaolinite Hydrous oxides
150-300 150 80-100 69-86 5 7-64 1240 2-1 5 4
Sources: Grim (1968):Ross (19 89).
Rock salt, gypsum, guano phosphates such as struvite and nitrates such as sodaniter may be abundant in desert soils (Mueller 1 9 68), but these soils are often too dry to be biologically productive. Similarly, sulfur in pyrite and phosphorus in vivianite are not easily available to plants because these minerals are stable under conditions more anoxic than most organisms can tolerate. To some extent, the availability of phosphorus in paleosols formed under more normal conditions (near-neutral pH and Eh) can be approximated by the abundance of apatite and bone fragments. Phosphorus is a critical limiting nutrient in old growth ecosystems, but communities early in ecological succession are more limited by nitrogen (Vitousek et af. 1997a; Chadwick etaf.1999). Chemical methods for approximating CEC and base saturation include those already outlined as proxy indi-
56
Chapter4
cators of former pH, such as the molecular weathering ratio of bases to alumina (Table 4.4). Depletion of individual nutrient elements can be calculated from chemical analyses and bulk density measurements (Brimhall etal. 1991;Bestlandetal. 1996).Acidificationestimates from chemical data also can be aguide to nutrient status of paleosols (Table4.5).Analysesfornitrogen, phosphorus and sulfur, and for various trace elements also are useful. Phosphorus commonly is present in significant amounts (0.02-0.5%,averaging 0.05%:Dixon&Weed 1989) in the mineral fraction of soils. All three may be detectable in the organic fraction of paleosols. Utilization of iron in paleosols can also be approximated by magnetic susceptibility,which is the ability of a material to enhance an applied magnetic field. Microbial enrichment of iron in the surface of desert and Precambrian paleosols (Aridisols,Inceptisols) may produce a strong peak of magnetic susceptibility at the surface of paleosols (Soreghan et d. 1997). In forested paleosols (Alfisols, Ultisols) the utilization of iron results in lowered susceptibilityat the surface and other patterns (Oches&Banerjee 1996;Maher &Thompson 1999). Bioturbation Thedevelopment of soil structureis related to the degree of reworking of the soil by organisms, or bioturbation. From their inception in colonizing barren soil. plant communities show a slow increase in biomass, usually building to a sustainable level (Kimmins 1997). Roots, nematodes and other soil organisms break soil into smaller and smaller peds. These are then coated in clay and slimy polysaccharides so that their natural tendency to shrink and swell is subdued. The net effect of plant succession is to change a rocky or alluvial subsoil into an organic, well-structured soil. In other words, plants promote what farmers cat1 tilth. The epitome of tilth is the rich dark soil formed under grasslands (see Fig. 3.8B). The building of biomass not only promotes tilth, but preserves it from disturbance. Both soil erosion and redeposition are mitigated most effectively by the largest and most complex forested ecosystems. In contrast, desert soils show little tilth or biomass and are readily eroded. The dependence of biomass, productivity and stability on increasing rainfall is a prominent feature of modern ecosystems (Fig. 4.1 1)used widely for
rangeland management (Munn et al. 1978; Sala et al. 1988). Bioturbation thus has temporal and ecological components. The degree of destruction of relict bedding may be a guide to the development of tilth in soils during early succession. A comparative scale for the persistence of bedding developed for marine trace fossils (Fig. 4.12) can be used for the evaluation of paleosols. Another proxy indicator of tilth is the percentage of a horizontal line transect (1m is sufficient in some paleosols) occupied by drab-haloed root traces formed by burial. As already argued, these may represent the last crop of large roots in a paleosol (see Fig. 3.7).This measure has proven especially useful for distinguishing paleosols formed under woodland from those of wooded grassland (Retallack 1991a).Other root traces without drab haloes may be present as well, but these represent roots that died and decayed before burial of the soil. Unfortunately, most animal burrows in paleosols do not have haloes or other evidence of their relative age. Estimates of animal activity from abundance of burrows in paleosols is compromisedby the unknown degree of occupation of burrows and the unknown rate of destruction of abandoned burrows. Estimates from abundance of fecal pellets are compromised by similar considerations. Finally, quantification of soil structures and microfabrics can be achieved by point counting thin sections under the microscope or by measuring intersections along line transects. For example, the number of clay skins encountered along a 20 cm transect in the subsurface (Bt) horizon of a paleosol is a useful measure of cutan development (Retallack 1997e). Within aparticularecosystem the degree of bioturbationreflectsthe passageof time, but differentecosystems vary markedlyin biomassand productivity (Table4.10). and thus potential for bioturbation. The measurement of biomass and primary productivity is a difficult and time-consuming process for most natural ecosystems (Lieth & Whittaker 1975). Grasslands or agricultural crops can be mown or harvested and weighed. Because most of the biomass in such ecosystems is in their roots, these also must be excavated, washed free of soil and weighed. Determining the biomass of forests is much more intimidating and seldom has been attempted directly Usually it is estimated from empirical relationships to easily measured parameters such as diameter of
Soil-formingprocesses
Figure 4.1 1 Increased productivity of herbivores and plants, optimal soil carbon storage and mollic structure, and increased depth to calcic horizon withmean annual precipitation in modern African soils:(A) herbivore net production (g m-2yr'live weighkCoeetal. 1976); (Bfnetprimary production (g m-j y-'dry weight: Murphy 1975;Scholes&Hall1996); (C)soilcarbon (kgm->:Zinke eta]. 1984); (D)thicknessof mollic epipedon(cm);(E) depth to calcic horizon of nodular form in unconsolidated parent materials (cm) (D&Ecompiledfrom Brammer 1955;Harrop 1960;ThorpetaL 1960: Anderson 1963a.b;Thompson 1965; Scholz 1968; Bonarius & Mugai 1977; FA0 1977a; Kanaka & Mugai 1977; Mbuvi & Njeru 1977;Siderius&Muriuki 1997; Sogomo et d.1980;Fenger et d.1986; Macmillanetal. 1990;Retallack 1994b).
57
-100-
0 DEPTH OF CALCIC -100 HORIZON (cm)
tree trunks at breast height (about 1.5m above the ground). This measure can be used to estimate the height of former trees from fossil logs, using empirical regression equations (Whittaker & Woodwell 1968; Pole 1999).Forpaleosols.generalecosystemtypecan be assessed from such features as root traces, soil horizons andsoilstructures(Retallack 1997a). An ingenious approach for estimating secondary productivity, or soil respiration, has been devised using the carbon isotopic composition of carbonate within goethite of paleosols (Yapp& Poths 1994).This is based
-200 MEAN ANNUAL PRECIPITATION (mm)
on the observation that respired soil CO, is isotopically lighter (13C depleted) than CO, in the atmosphere. The curve of adjustment of these two values is near the surface of the paleosol (Fig. 4.13) when soil respiration is high. Conversely, isotopically heavy CO, diffusesfurther down into soils with low respiration rates. Common soil-formingprocesses So far in this chapter soil-forming processes have been
viewed from the perspectives of physicists, chemists and
58
Chapter4
Figure 4.12 Fossilroottraces emanating from the top (carbonaceousA horizon)of a paleosol (Aquept)and cutting acrossitsrelict beddingin theMiddleTriassic(23OMa),Tank Gully CoalMeasuresnearMt Potts. New Zealand(to1eft:fromRetallack1979,with permission from the Royal Societyof New Zealand),which is two on a schematicscale for the destruction of bedding by burrows (to right; fromDroser &Bottjer1986: with permission from the Societyof Economic Paleontologists and Mineralogists).
biologists. Soil scientists have yet additional perspectives, recognizing at least 17 distinct soil-forming processes (Table 4.1 l).These processes are thought to be important for understanding the main kinds of soils
cm 0
now found on Earth (Buol et al. 1997).Most classifications of soils implicitly acknowledge these processes although they are rarely used explicitly in taxonomic nomenclature (Bockheim & Gennadiyev 2000). Only
carbon isotopic composition of mole fraction ethite g13CpD8 Cog in goethite 8 -16 -14 0.004 0.OOf
---FIT--
marine fossils mdolostone m o o l i t i c ironstone
IIIIZI]red color
a
brown color
100
bedding
Figure 4.1 3 Isotopic composition of carbon incarbonate and mole fraction CO, in goethiteof a late Ordovicianpaleosol from Wisconsin, showingsurprisinglyhigh soil respiration by adjustmentof soil to atmosphericvalues (fromYapp&Poths 1994: withpermissionfromNature and Macmillan Journals).
Soil-formingprocesses
59
Table 4.10 Productivity,stature and biomass of vegetation. Plantformation
Stature (m)
Rootingdepth (m)
Biomass (kg m-2,dry wt)
Netproductivity (kgm-’yr-’)
Aquatic stromatolites Sabkha stromatolites Microbial earth Seaweedbed Microbial rockland Polsterland Salt marsh Brakeland Marsh Pondweed Mangal Forest Wooded shrubland Dry woodland Swamp Fen Carr Dune binders Oligotrophicforest Rainforest Shrubland Bog Tundra Taiga Desert scrub Heath Fire-proneshrubland Sea-grassbed Wooded grassland Open grassland
0.5 0.5 0.001 18 0.001 0.001-0.1 1 0.001-2 1 20 10 10-100 10 10 3-100 2 3-100 0.1-2 10&100 10-100 0.5-2 1 0.1-1 1-10 10 0.1-2 0.5-2 0.3 0.5-10 0.5-2
0.001 0.001 0.001 0.1 0.001 0.001-0.12 0.001-0.1 0.001-1 0.14.5 0.1 0.001-1 1-5 0.1-1 0.5-2 0.1-0.5 0.1-0.5 0.1-0.5 0.5 1-5 1-5 0.1-1 0.1-0.5 0.1-1 0.5-2 0.1-1 0.1-1 0.1-1 0.001-0.1 0.1-2 0.1-2
0.02-0.16 0.19-0.25 0.04-0.36 0.26-8.00 0.0007-0.49 0.05-0.95 0.2 1-1.9 7 0.02-0.99 0.11-30.00 0.05-2.00 2.2046.90 10.0~5.00 0.40 4.00-2 5.00 30.00-50.00 0.36-1.82 3.11-9.81 0.10-3.00 0.99-88.20 8.00-70.00 0.09-1.20 2.504.00 0.10-5.85 2.5 5-3 5.00 0.10-1.50 0.20-2.60 1.OO-10.34 0.04-8.10 1.50-6.00 0.06-30.00
0.01-0.11 0.03-0.36 0.002-0.20 2.00-12.88 0.02-0.20 0.02-0.28 0.3cb3.70 0.002-0.05 0 . 1 4 1 5.00 0.13-2.00 0.68-2.56 0.65-2.00 0.03-0.39 1.20-1.70 1.30-2.50 0.66-1.40 0.58-0.65 0.01-0.80 0.83-3.28 0.3 1-3.50 0.016-0.40 0.22-2. so 0.07-0.87 0.40-1.40 0.05-1.20 0.10-0.40 0.60-1.99 0.33-2.29 0.40-1.40 0.06-1.50
Source: Retallack (1992a).
the six most common of these processes will be consideredin moredetail here (Fig. 4.14). Gleization Waterlogged, mucky ground with a bluish gray or greenish gray color is called gley, a Russian agricultural term. Gleying or gleization is a genera1 term used to characterize processes that produce these colors and other distinctivefeatures of waterlogged soils. Stagnant
water is exhausted of oxygen by microorganisms and under these anaerobic conditions other microorganisms reduce oxidized minerals. Microorganisms almost always are involved because chemical reduction of oxide and oxyhydrate minerals is very slow at surface temperatures and pressures (Stevenson 1986). Gleization results in the production of minerals containing iron in areduced state (Fe2+)such as drab-colored clays, siderite and pyrite (Kraus 1997, 1998).Gleization also can be associated with the accumulation of organic
60
Chapter4
Table 4.11 Soil-formingprocesses.
Process
Other name
Description
Lessivage Biogeochemicalcycling Andisolization
Argilluviation
Washing down of clay into subsurface horizons Biological enrichment of soilin base cations or other soil fertilization Formation of fertile,low-density,high-porosity, amorphous weathering products from glassy volcanic ash Formation of peat Redoximorphicalterations: particularly chemical reduction of iron and manganese as a result of waterlogging Darkening of soil as aresult of increase inorganicmatter Increased abundance to dominance of sesquioxidesof iron and aluminum Acidicleaching anddestruction of clay with precipitationof humus andiron oxides and hydroxides Leaching of base cations (Ca2+,MgL+,Na+,K+) Deep cracking and heaving of clayey soils as a result of drying and wetting Crackingand heaving of soils as a result of ground ice Accumulation of soluble salts by evaporation Accumulation of carbonates such as dolomite and low-magnesian calcite Accumulation of soda and other alkalis in clays, often producing domed-columnar peds Subsurface accumulation of soda and other alkalisin clays, but surface acidiEcation,givingbleached andtruncatedcolumnar peds Cementation with silica, such as chalcedony and opal Alteration of soil by human activity
Paludization Gleization Melanization Ferrallitization Podzolization
-
-
Hydromorphism -
Lixiviation
-
Vertization Cryoturbation Salinization Calcification Solonization
Vertisolization
Solodization Silicification Anthrosoliation
-
Alkalization -
-
-
Source: Bockheim&Gennadiyev(2000).
matter (paludization), because lack of oxygen discourages the activity of aerobic decomposer microbes. Also discouraged is the activity of large plants and animals, which form distinctive low-diversity assemblages of swamps and marshes. Gleyed soils show little evidence of bioturbation and may have abundant relict structures of parent material. The root systems of trees within them tend to be planar and very near the surface because roots require oxygen to respire. Soil pH may be rather variable. It is often acidic from the organic acids, but can be made alkaline by abundant m d c minerals or nearby limestone bedrock. The process of gleization retards mineral weathering because water flow and removal of weathering products is limited. Thus, easily weatherable minerals tend to persist in gleyed soils. Gleization reflects local waterlogging rather than wider effects of climate and vegetation evident fromother soilforming processes.
Podzoliation Translocation of aluminum and iron oxides, organic matter, or all of these, to subsurface (Bsor Bh) horizons, overall destruction of clay and leaching of exchangeable cations (CaZ+,Mg2+,K+, Na') are the main features of podzoliiation.The result is acolorful, contrastingprofile with a white near-surface (E) horizon largely of quartz over brown, red or black subsurface (Bs or Bh) horizons cemented by sesquioxidesor organic matter, or both. Podzoliiation occurs under acidic (pH <4) and oxidizing conditions in moderately to well-drained soils of humid climates. The characteristic acidic reaction can arise in several ways. Such soils preferentially form on materials initially sandy and quartz-rich in which there are very low levels of weatherable minerals that would hydrolyzeto clay. They also form under particular kinds of vegetation that can tolerate low levels of nutri-
Soil-formingprocesses
GLEIZATION
PODZOLIZATION
LESSIVAGE
FERRALLITIZATION
CALCIFICATION
61
SALINIZATION
Figure 4.14 Common soil-formingprocesses.
ent bases. Typically this is conifer forest or alpine or coastal heath. Such plants contain phenolic and other acidic substances that may be washed out of the leaves by rain (Lundstrom et al. 2000). Their leaves remain in the litter for long periods of time (in a mor humus) because decomposer microbes are sparse in these low-nutrient soils. A final contributing factor to their acidity is a cool and humid climate in which cations released by weathering are more liiely to be flushed out of the profile by groundwater rather than taken up by organisms or claysorprecipitatedas salts (Fitzpatrick1980). Lessivage Differentiationof a near-surface horizon that is leached and lighter in color (E) and a subsurface zone enriched in clay and darker in color (Bt)is the soil-forming process of lessivage. UnIike podzolization. Iessivage results in an illuvial horizon that is clayey (Bt),and also may be reddish with sesquioxidesor dark with organic matter. Evidence for clay illuviation may be sought in the form of clay
skins.These are especiallystriking in thin sections of the subsurface (Bt) horizons (Holzhey et al. 1974; Murphy 1983). Those formed by lessivage tend to have bright clay (sepic plasmic) fabric and resistant grains floating in clayey matrix (porphyroskelic).A podzoliied subsurface (Bs or Bh) horizon, on the other hand, is usually grain supported. Its grains are coated with a thin opaque layer of sesquioxides or organic matter, often with fine desiccation cracks radiating from the grains (de Coninck et al. 1974). Soils formed by lessivage also are chemically distinct. They are neutral to moderately acidic. There is little exchange acidity and it tends to be dominated by hydronium (H+) rather than aluminum ions (A13+)so that clays within the profile are more stable than in podzolized soils. Hydrolytic reactions, nevertheless, remove exchangeable cations (Ca2+,Mg2+,K+, Na') from the weatherable minerals (lixiviation).These can be washed out of the profile and over time this can seriously reduce fertility.The process of lessivage is favored on parent materials such as basalt or shale, which are rich in minerals that weather to clay. Lessivage also is
62
Chapter4
favored by vegetation such as deciduous broadleaf woodland that is low in acid-generating phenolic compounds and resins (Buol et aI. 1997). The fall of broadleaves allows recyclingof nutrients and abundant soil fauna, which is a feature of soilsformed by lessivage. Ferallitization In well-drained soils of wet tropical regions, weathering is intense and deep. It may result in thick, uniform profiles depleted in exchangeable cations. As a result, the soil as a whole is enriched in clay and sesquioxidesin the form of fine-grained, weather-resistant minerals such as kaolinite, gibbsiteand hematite. The process of ferrallitization is encouraged by a long period of soilformation and under ecosystems of great biomass and stability such as tropical rainforest (Nahon 1991).As in temperate woodland soils formed by lessivage, there may be some indication of a clayey subsurface (Bt) horizon, but it is usually thick (>1m) and has diffuseboundaries. Soil reaction is buffered by abundant clay to near neutral or mildly acidic.It is paradoxical that deeply weathered and infertilesoils should be able to support such lush vegetation as rainforest. This is achieved through a nearly complete recycling of nutrients within a biologically activeleaf litter (Sanford 1987).Therootsof tropical trees in ferrallitized soils commonly are shallow and spreading.Thesoilitself may bedeeplyand copiouslyburrowed by ants and termites, which create sand-sized stable spherical micropeds that are characteristic of these soils (Stoops 1983). Calcification
In well-drained soils of semiarid to subhumid regions only the surface of the soil may be moistened by rain before the water evaporates, is used by plants or animals,
or soaks into organic matter and clay. Thus, rainwater may not flow through the soil as it does in more humid climates.In the moist and biologicallyactivesurfacelayers of the soil, exchangeable cations may be hydrolyzed out of weatherable minerals. In subsurface horizons, however, there may be sufficient water to remove some exchangeable cations (Na', K+. Mg2+)from the profile, but not calcium (Ca2+),which accumulates as a distinctive subsurface (Bk) horizon near the depth of average wetting. Compared with ferallitization, lessivage and podzolization, the process of calcification involves little loss of nutrient cations and overall soil pH is alkaline. Soils of this kind can be fairly fertile. They support open grassland or grassy woodland vegetation with abundant soil fauna. In drier regions they support desert vegetation (Jenny 1941). Salinization In very arid climates rainfall may be sufficient to hydrolyze soluble cations from weatherable minerals, but insufficient to remove them from the profile, where they accumulate as surface crusts and crystals of salt such as halite, gypsum and mirabilite (Hartley & May 1998). Salinizationcan also occur by precipitation of salts from the top of a temporarily high water table. Irrigation of aridland soils may exacerbate this process. Salinized soils are generally well drained for most of the year and extremely alkaline (pH 9-1 1).Their high salinity discourages the growth of plants other than small-leaved, evergreen, desert shrubs or cacti. The activity of soil fauna also is curtailed under such conditions. At its extreme development in barren desert playas and coastal sabkhas, salinization may be regarded in part as a sedimentary process, but even there it involves modifications in place that are the hallmark of other soil-forming processes (Calvoetal. 1999).
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 5 Soil classification
Many features of paleosols can be understood by comparison with modern soils in the same way as fossil bones are best interpreted by comparison with bones of similar modern organisms. For this reason, modern soil classifications are important for understanding paleosols. Soil classificationsdate back to the latter half of the 19th century. From the beginning, these classifications had practical aims in planning for agricultural and other land use. Thaer’s classification of 185 7, for example, was based at the highest level on the kinds of crops that could be grown and on soil texture: wheat soils (clay), barley soils (sandy clay), oat soils (clayey sand) and rye soils (sand). Like other kinds of classification in natural history, soil classification provides basic units for extending and extrapolating research results. Most soil classificationshave some theoretical underpinning in soil-formingprocesses and are designedto bring order to themultitude of combinations of differentsoil properties. Such organizational systems are more easily remembered than the raw data on which they are based. Systems of classification also become a way of thinking and communicating akin to a foreign language. The optimal classification for paleopedological comparisons would be one based entirely on observable features of profiles that are relatively resistant to diagenetic alteration. Suitable classificationsin this respect are the simplified paleosol classifications(Fig. 5.1)of Macket al. (1993),andof Nettletonetal. (1998,2000). Classifications of this kind may prove important to paleopedology in the future, but are not discussed further here for two reasons. First, few paleosols or soils have been classified by these schemes (Mack & James 1994; Parrish 1998). Second, the classification of Mack et al. (1993) is a simplification of the US soil taxonomy (Soil Survey Staff 1975. 1998), which is of greater instructional value and practical utility. Also not discussed further for reasons of scope and utility are classifications of South Africa (van der Eyk et al. 1969: Soil ClassificationWork-
ing Group 1991), Australia (Stace et al. 1968; Isbell 1996) and the factual key of Northcote (19 74). Each of these classificationsmay prove useful for particular paleosol studies, but two classificationsparticularly command attention: the soil map of the Food and Agriculture Organization (FA0 1971, 1974, 1975a,b, 1977a,b, 1978a.b. 1979, 1981, 1988) of the United Nations Economic and Scientific Organization (UNESCO) and the soil taxonomy of the Soil Conservation Service(Soi1SurveyStaff 1975,1998)of theUSDepartment of Agriculture (USDA). The FA0 classification is especially good for tropical soils (Richter & Babbar 1991).TheUSsoiltaxonomy, ontheotherhand,isbased largely on soils of temperate climate in tectonically active, volcanic and glaciated terranes (Birkeland 1999). The FA0 maps are useful for h d i n g modern soilscapes analogous to suites of paleosols.The US soil taxonomy is becoming a global vocabulary for soil science, with wide application in agronomy and environmental sciences. FA0 world map The FA0 classification was begun as an international project in 1961. Explanatory volumes and soil maps began to appear in 197 1and the tenth one (VolumeV) was published in 1981.A revised legend and classification has been issued (FA0 1988),but is of little use until the maps also are revised. The maps are continental in scale (1: 5 000 000), but the accompanying volumes make up for this lack of detail with extensivenotes on climate, vegetation, geomorphological positions, parent materials and land uses of the map units. Descriptions of individual soil profiles with analytical and site data are included as examples (FA0 1971, 1974, 1975a,b, 1977a,b. 1978a,b. 1979, 1981).It is an unparalleled global data base for comparison with paleosols. The FA0 classificationretains some traditional names such as Chernozem and Podzol, and adds new names such as Acrisol and Ferralsol.The classification is based 63
64
1
Chapter5
In S,fu M s n a r a l OXISOL~a",enr,"s
1F iElJ I l l ~ i i a l t O n(It
Inlllluble Minaral.i/Camp0unds
ARGILLISOL
s P O D 0 s 01-B
* slay
o r a a n ~ cm a t t e r "d Iron
*cCYm*I.llOn 01 Soluble Mineral.
+ , CALCISOL-CaC03
GVPSISOL-CaSDa
PDOl
on a system of diagnostic horizons and is hierarchical like the US taxonomy, Instead of many hierarchical levels, the FA0 world map has only two. These are the 'major units', which have specific soil names such as Podzol and Acrisol, and 'soil units', which are specified by adjectives, for example, Humic Podzol and Ferric Acrisol. Not all of the 2 6 major units are subdivided,but some include many subunits. The total number of soil units is 106. Such a hierarchical system implies that soils must first be classified within the 2 6 major units beforebeing assigned to subsidiary soilunits.This reflects a judgement that some features of soils-their diagnostic horizons-are more fundamental for their classification than are others. For example, Ferric Acrisols have a strongly ferruginized profile like Ferralsols, but they have less deeply weathered, base-depleted subsurface horizons than Ferralsols and this feature of Acrisols is thought to be more important to their classification than ferruginization. Hierarchical systems thus involve a series of decisions. They are also easy to remember, at least at higher levels, where there are fewer categories. A potential weakness of such a global system of classification is establishing correspondence between the regional schemes of classification and soil surveying from which much of the information was extracted. It is difficult enough to achieve consistent classification in a national soil survey by scientists speaking the same language and followingthe same philosophy of soil genesis developed at a small group of universities.This problem was addressed at length by numerous working parties on the correlation of different national soil classifications. However, not all scientists were happy with all aspects of the resulting classification for the world
IIC
rich
Figure 5.1 Flowchartfor determining soil orders in the paleosolclassificationof Macketd. (1993;with permissionfromthe Geological Societyof America).
map, and it was not adopted by the Soil Conservation Service of the USDA. The original volumes accompanying the soil maps are too detailed to afford a concise introduction to the FA0 classification, but Fitzpatrick (1980) has prepared a well-illustrated introductory summary. Only soil schematic representationscan be offeredhere (Fig. 5.2), arranged in general order of increasing development, with waterlogged soils at the end. US soil taxonomy The classification of the Soil Conservation Serviceof the USDA (SoilSurveyStaff 1975,1998)was toaconsiderable extent the work of Guy D. Smith. It was guided by him through a series of early drafts, which were widely circulated for criticism and comment. The Seventh Approximation was first published in 1960 and a definitive account appeared in 19 75 under the distinctive title of Soil Taxonomy. It continues to grow and be emended, with new editions of Keys to Soil Taxonomy appearing almost annually (SoilSurveyStaff 1998). This classification was meant to be a total break with the past, which, in the USA, was C.F. Marbut's (1935) equally monumental Soils of the United States. The word 'taxonomy' is a Greek equivalent of the more familiar Latinate 'classification'.but this isnot the only change of terminology,The whole taxonomy is based on a new and rigorously defined nomenclature for soils and soil features. The soil taxonomy also marked an attempt to classify soils on the basis of their own measurable properties rather than by the various factors thought to be impor-
Soil classification tant in their origin. This is especiallyfortunate for paleopedologybecause paleoclimates and other genetic factors in the origin of soils can be inferred from paleosols identified objectively with modern soils. The nature of diagnostic horizons and other observable soil features such as calcareous nodules form the basis of the classification. For example, recent versions of the key (Soil Survey Staff 1998)include calcic horizons at less than 1m depth, as well as climatic criteria, in the definition of Aridisols. The soil taxonomy also has hierarchical structure. Thus soils are classified first at a very general level, then assigned to progressively more limited subdivisions. At the highest level there are only 12 soil orders. This is a considerable reduction of initial choices compared with the classifications of the FA0 world map. This does not mean that the US taxonomy is less precise, because each order is divided into suborders, 63 of them in all. Each suborder is divided into Great Groups, of which there are 32 7 (SoilSurvey Staff 1998).These are divided into hundreds of subgroups, then thousands of families and finally, tens of thousands of locally named soil series. This hierarchy of subdivision is built into the nomenclature for soilsin which a distinctivesuffix refers to the soil order, a prefix to suborder and qualifiers refer to subgroups. For example, a Petrocalcic Paleustalf belongs to the order AEsol (suffix ‘alf’)and the suborder that forms in dry climates (prefix ‘ust’),the Great Group that forms on old land surfaces (prefix ‘pale’) and the subgroup that has acalcareous hardpan at depth (petrocalcic).In thisway each soilnameincludes anindication of its position in the hierarchy.The soil names are meant to be formal names that rank higher than series and so are customarily capitalized. This practice has been followed for allthesoilclassificationsin this book, although this is not done for classificationsused in other countries. The three main features of the soil taxonomy+mphasis on observable features, new nomenclature and hierarchical organization-are also the source of some difficulty. The nomenclature is like a foreign language on first acquaintance. It takes time and practice to become fluent. The emphasis on objective features in the soil taxonomy has led to much of it seeming like a legal document complete with specific stipulations, provisos and caveats. Such legalistic strictures often obscure the main idea behind the particular soil type. There are some places where the aim to abandon genetic concepts
65
was not carried far enough. Climatic conditions are the main criteria for distinguishing suborders of Vertisols. The particular information needed about how many days per year the cracks are open is difficultto obtain for modern soils, and impossible for paleosols. Another problem is inherent in the hierarchical structure of the taxonomy, which calls for selective weighting of different features of soils.The diagnostic horizons and the degree of development at which they become diagnostic are clearly specified,but these same features also may be used at lower levels to define intergrades. This can be confusing until one learns to operate strictly within the hierarchy by choosing among alternatives at each level in the precise order specified.For paleosols it may be ddficult to make decisions on diagnostic criteria at high levels and the identification should stop at those levels. Despite these drawbacks, the US soil taxonomy has an undeniable influence on soil classification throughout the world. Although the FA0 decided on its own classification, it incorporated basic ideas and terms from the US taxonomy for horizons and properties diagnostic for classification. A key (Fig. 5.3) and schematic representations (Fig. 5.4) supplement the following written summary of the US taxonomy. Full-color illustrations anddescriptions of soilsandpaleosols using this classification are available (Aandahll982;Barker et al. 1983; Retallack 1997a). Other introductory accounts can be found in most textbooks of soil science (Buolet al. 1997). Diagnostic horizons and properties Fundamental to the soil taxonomy are 19 diagnostic subsurface horizons and eight diagnostic surface horizons (SoilSurvey Staff 1998).Asurface horizoniscalled a n epipedon (plural epipedons). This does not correspond exactly to the A horizon; B horizon material is included in some cases. The mollic epipedon is a soft, dark, humus-rich, finely structured surface horizon found under grasslands, including suburban lawns. High base saturation characteristic of this horizon may be implied for paleosols if they have dispersed or nodular carbonate, abundant burrows and fine root traces, or common easily weathered mineral grains such as feldspar. Characteristic granular or crumb ped structure may be preserved in paleosols (Retallack 1997b), but organic matter is sel-
66
Chapter5
AY
A
C
C
A C
A C
R
SOLONCHAK
FLUVISOL
REGOSOL
LITHOSOL
CAMBISOL
ANDOSOL
ARENOSOL
RANKER
VERTISOL
SOLONETZ
IFA0 WORLD MAP CLASSIFICATION
Y XEROSOL
YERMOSOL
Part I of 2
Figure 5.2 Schematicrepresentationsof climate,vegetation and soil profile form of the major soil units of the FA0 SoiZMupof the WorZd(basedon data from FA0 19 71,1974.197Sa,b,19 77a,b, 1978a,b, 19 79,198 1).(For key to lithologicalsymbols,see Fig. 5.4.)Sunlcloud and hourglasssymbols indicate typical climatesand t i e s for formation. (Continued)
dom preserved at original levels (Retallack 1991b). For this reason care must be exercised in considering the requirement for a mollic epipedonto have more than 2.5% organic matter in the upper 18cm or to have a dark
Munsell color (value and chroma less than 3.5 moist and value less than 5.5 dry). Also difficultto apply to paleosols because of their likely compaction and erosion are the thickness requirements for a mollic epipedon.
Soil classification
I
GREYZEM
I ACRISOL I F A 0 WORLD
PWEOZEM
CHERNOZEM
NITOSOL
FERRALSOL
MAP CLASSIFICATION
-
KASTANOZEM
RENOZINA
1
GLEYSOL
HISTOSOL
I
Part 2 of 2
67
I
Figure 5.2 (Continued)
These are more than lOcm thick if developed on bedrockandmorethan 18cmif there areother soilhorizons within 7 5 cm of the surface or if the soil is loamy or clayey, and more than 2 5 cm if these other horizons are more than 75 cm down or if the soil is on sandy or
gravelly materials. For paleosols, adjustments to the amounts of organic matter and thickness of horizons need to be made based on estimates of burial diagenetic modification. Umbric epipedons are more organic and less finely
68
Chapter5 Freeze-thaw banding,
sand wedges and other periglacial features
n
KEY
SOIL ORDERS AND PETROGRAPHY OF CHARACTERISTIC CHARACTERISTIC HORIZONS HORIZONS
GELISOL gelic materials
relict structures such as beddinq clear relict
ENTISOL ochric epipedon
ANDISOL volcanic ash surface peal more than 40 c m INCEPTISOL cambic horizon
HISTOSOL
Ek-j
granular peds
MOLLISOL mollic epipedon
more than 10% easily weathered minerals such as feldspar mainly clay with
VERTISOL cambic horizon
ARlDlSOL calcic horizon
ALFISOL argillic horizon
nodules common less than 1 m down abundant clay, which is kaolinitic or gibbsitic
fi
ULTISOL kandic horizon
common clay subsurlace horizon iron or organic cemented sandy
OXlSOL oxic horizon
SPODOSOL spodic horizon
Figure 5.3 AsimpliEedkey to the 12 soil orders of the US soil taxonomy,emphasizing field and petrographicfeatures observablein paleosols (modifiedfrom Retallack 199 3a;with permission from the Geological Society of America).
losll3o
iosiaitlv
losIllow
AWONOWl 110s 'S'n lOSllH3A
D U
10SIXO
1OSOlSIH
iosiiin
1OSlaNV
-
10SlJlV
lOSlld33NI
losoaods
10SllN3
3 V
70
Chapter5
structured than mollic epipedons, which they most resemble. In contrast to the mollic epipedon, an umbric one has base saturation of less than 50%,and so is noncalcareous with low reserves of weatherable minerals. The anthropic epipedon may resemble a mollic epipedon, but has been altered by human use. It generally has a less finely developed structure than a mollic epipedon and may show signs of trampling or fertiliation with bone scraps or shell fragments. A high content of phosphate is diagnostic (McDowell 1988). Evidence of artifacts, campfires or hut foundations also can be used to recognize an anthropic epipedon. In theory, the anthropic epipedon also includes other human works such as roadways and railway grades, but not the plaggen epipedon, which is a soil surface created by manuring and plowing. This is a dark organic horizon with a poor structure, but in other ways similar to an anthropic epipedon. Spade or plow marks may be visible as well as pieces of brick or pottery. Like the cultivated field in which they are formed, plaggen epipedons commonly occupy square or rectangular areas of ground on moderatelylevel sites. Histic and folistic epipedons are peat layers, which in paleosols are converted to coal seams. Most peats and coals were histic epipedons, defined as waterlogged for more than 30 days per year, which encourages minerals such as pyrite and siderite. These minerals are lacking in unwaterlogged. folistic epipedons, which include deep mats of pine needles and thick composts of logs and leavesin montane soils. The thickness of apeat required for a horizon to qualify as a histic or folisticepipedon is at least 20cm. For an histic epipedon to qualify as a Histosol, peat on bedrock must be more than lOcm thick, but on sediments, low-density peat of the kind formed under Sphagnum moss must be more than 60cm thick and other kinds of peat more than 40cm thick. The original thicknesses of coaly surface horizons of paleosols can be reconstructed from field or general information on peat compaction. It should be noted that peat at the base of a thick unburied histic epipedon is already compacted compared with peat at its surface. Compaction to formcoalcan be0.03-0.71 times theoriginal peat thickness (Winston 1986; Nadon 1998). By the lowest ratio a seam of woody coal only 1.2cm thick could qualify as an Histosol. Another diagnostic criterion for histic epipedons and Histosolsis the amount of organic matter and clay.The percentage of organic carbon
(0)relative to clay content (c) is given by the following conditions:o>16if c>60,0>8+(~/7.5)ifc<60.Thus, a histic epipedon or Histosolcan have no more than 84% mineral matter and no less than 8% organic matter. In practice, this means that histic epipedons and Histosols are very dark to black with organic matter rather than gray with clay. Almost all economicallymineable seams qualify as fossilhistic epipedons and Histosols. A melanic epipedon is also dark with organic matter (more than 6% organic C) but does not qualify as histic because of mineral content, usually allophane and other amorphous minerals fromvolcanic ash. Like other parts of volcanic ash soils, melanic epipedons have low bulkdensity and highcation exchange capacity. Opaque organic matter and volcaniccomponents in thin section are good guides to melanic horizons in paleosols. A final kind of surface horizon, which accommodates most others, is the ochric epipedon. This is too thin, too light colored or not organic enough to qualify as one of the other kinds. An ochric epipedon may contain organic matter, but it is less intimately mixed with clay and may contain recognizable leaf litter. It also is more likely to be sandy, blocky or otherwise more crudely structured than, for example, a mollic epipedon. Among subsurface diagnostic horizons, the argillic horizon is one of subsurface clay enrichment. Clay skins along the margins of soil peds and root channels are evidence of clay that was washed down the profile (McCarthy et al. 1998). Clay rinds extending into fractured and partly hydrolyzed grains may be evidence of clay formation in place. Neither of these features can be duplicated by mechanical ctay infiltration after burial (Buurman et al. 1998). Cornpactional effects may obscure these features in petrographic thin sections of paleosols. Compaction also alters the degree of clay enrichment and thickness of the clayey horizon in paleosols. The per cent clay ( k ) of the subsurface horizon compared with that of the surface horizon (c) required for an argillic horizon varies with the degree of clayeyness of the soilas a whole: if Oc+ 3; if 15 < c <40, then k > 1 . 2 ~andif ; 40 < c< 100,then k> c + 8. The requirement that an argillic horizon be at least onetenth as thick as all overlying horizons also must be applied to paleosols after considering burial compaction. Kandic horizons are comparable, but not exactly the same, in per cent clay enrichment of subsurface ( k ) compared with surface (c) horizons: if 0 < c < 20, then
Soil classification
k = c + 4; if 2 0 < c < 40, then k > 1 . 2 ~if; 40 < c < 100, then k > c + 8. The main distinction of a kandic horizon is its extremely low cation exchange capacity (CEC < 16 cmol kg-I). This is mainly related to the presence of minerals such as kaolinite, boehmite and gibbsite (see Table 4.9), with very low amounts of alkalis (Na,O, K,O) and alkaline earths (MgO,CaO). Another kind of clayey subsurfacehorizon is the agric horizon,which forms after cultivation.Clay washes into the large cracks opened up by the plowed layer. These wedge-shaped masses of clay may be layered with increments of washed-in topsoil. The agric horizon has a sharp lower boundary at the base of the plowed layer and may have associated human artifacts. The agric horizon is found only in Holocene soils and paleosols. The natric horizon is a clayey subsurface horizon strongly base saturated with sodium. This can be estimated for paleosols from soda to potash molecular ratios greaterthanone(Retal1ack1997a),andfromcolumnar or prismatic structure with a sharp upper boundary at the topof the horizon (McCahon&Miller 1997). A sombric horizon is like a subsurface version of an umbric epipedon. The dark organic matter of a sombric horizon is not associated with iron and aluminum oxides as in a spodic horizon, nor is it base saturated as in a natric horizon. Sombric horizons are not found under an albic horizon. They are found mainly in moist soils of high plateaux and mountains in tropical to subtropical regions. A spodic horizon is a sandy, usually quartz-rich subsurface horizon cemented by amorphous iron and aluminum oxides, or organic matter, or different layers and combinationsof these.Thisdark,cementedhorizon must be laterally continuous and at least 2.5 cm thick. There are also requirements for the amount of iron and aluminum as extractedby ammonium oxalate, but this cannot reliably be determined for well-lithified paleosols (Retallack1991b).Thespodic horizoncan beidenti6ed in thin section by complete grain coatings of opaque, amorphous sesquioxides and organic matter, which commonly have numerous radial cracks as if they had shrunk around the grains on drying (de Coninck et al. 1974).Ortstein is an especially welldeveloped spodic horizon with more than 50%opaque cement, Spodic horizons form in humid, acidic soils in which clay is destroyed, and thus have very little clay. Red or organic clayey subsurface horizons are better re-
71
garded as argillic,oxic or sombricthan spodic (Retallack 1997d). The placic horizon is a black to dark reddish hardpan cementedby amorphous iron and manganese oxides, or by an iron and organic matter complex.They are mostly thin (2-10mm). but are rather variable in thickness (14Omm) and lateral continuity (Retallack et al. 2000). Some are brittle and break into angular segments. Others appearwavy,with twoormoreof thembifurcating and anastomosingmore or less parallel to the ground surface. In thin section they show a massive opaque cement enclosingthe clastic grains. Also, unlike spodic horizons,theyarefoundinclayeysoilsand atvarious depths in the profile other than below a sandy eluvial (E) horizon. The cambic horizon is a mildly weathered, slightly clayey or oxidized subsurface horizon more altered than underlying material, but lacking the distinctiveproperties and degree of development required for other kinds of subsurface horizons. This mild weathering is best judged by comparison with underlying parent material or saprolite.The cambic horizon may appear more massive and structured, show less sedimentary or other relict features, seem colored more yellow, brown or red as a result of oxidation,or be less calcareousor salty. Oxic horizons are so highly weathered that they have few exchangeable cations remaining (CEC < 16cmol kg-'). This is reflected in the abundance of kaolinite and other 1 : 1 lattice clays, trace amounts ( 4 0 % )of weatherable minerals, such as feldspar, and molecular ratios of bases to alumina close to zero. Although very clayey,thereisusuallylittleevidenceof clay skins.Unlike argillic horizons,which show a subsurfacepeak of clay enrichment, oxic horizons are normally deep (atleast 30 cm) and show near-constant amounts of clay with depth or a very diffuse zone of subsurface clay enrichment. Stable, sand-sized, spherical, opaque micropeds arecharacteristic of oxic horizons (Stoops 1983). An albic horizon is a light-colored and sandy layer from which clay and oxides of iron and aluminum have been leached leaving naturally light-colored minerals such as quartz and feldspar (Dumanski & St Arnaud 19 66). Clay and sesquioxides washed out of the albic horizon may be just below in an argillic or spodic horizon.The albichorizon may be just at the surface and commonly is near the surface,below a thin organic and rooted horizon. The light color of an albic horizon is
72
Chapter5
often conspicuous by contrast with overlying and underlying materials. These are the pastel shades in the upper left-hand corners of Munsell color charts (value generally greater than four and chroma less than three, with some exceptions noted by Soil Survey Staff 1998). Some soils show prominent tongues of sandy albic horizon tapering downward within the underlying clayey subsurface (argillic, kandic or natric) horizons. These are called glossic horizons. A fragipan is a dense subsurface hardpan of clay. Commonly it is mottled and has a prismatic structure. Groundwater perched on top of a fragipan forms a reduced and drab-colored surface on the hardpan and its prismatic peds (Lindbo& Veneman 1993). Some fragipans may be buried soil horizons, but others are probably horizons altered by permafrost in forested humid climates (Retallack&Alonso-Zarza 1998). A petroferric contact is a strongly ferruginized upper surface to bedrock at the base of a soil profile. Its close relationship with bedrock contacts and small amounts of organic matter distinguish it from spodic and placic horizons. Specimens from petroferric contacts are heavy, dark red, and rich in iron (often more than 30wt% Fe,O,). They are extensive in forested, or onceforested, soils of tropical and subtropical regions (Ollier &Pain 1996). Plinthite is a term for a particularly distinctiveform of laterite. Plinthite is an horizon of soil (notthe whole soil) formed in place (not redeposited) and has the unusual property of drying irreversibly on exposure to air. It is a material rich in iron with scattered red mottles of hematite and goethite in a matrix of highly weathered, light-coloredclay (usually kaolinite). Hardened, vesicular, pisolitic or brecciated laterites are not included in this definition, although these and other lateritic soil materials can be derived from plinthite by drying or redeposition. Plinthite is thought to form deep within forested soils in humid, tropical to subtropical climates (Valeton1999). A distinctive kind of subsurface horizon found in marine-influenced waterlogged soilsis the sulfuric horizon. This is either flecked with golden specks from pyrite or is a dull yellow color from jarosite formed by the oxidation of pyrite. Sulfides are fixed bacterially in these soils from sulfate. The sulfuric horizon is common in soils of mangal and salt marsh (Rabenhorst & Haering 1989; Rabenhorst 1990).
Subsurface horizons can also become cemented with calcium carbonate. These are called calcic horizons when the carbonate is in the form of powder or isolated nodules, or petrocalcic horizons when extensively cemented to form a continuous brittle layer within the soil.These soil carbonates are generally micritic and are petrographically more complex than carbonate formed during burial (Drees & Wilding 1987; Mack & James 1992). Calcic and petrocalcic horizons have complex dissolution and cavity-fillingstructures. The remaining clastic grains characteristically have nibbled edges where replaced by carbonate. There may also be displacive fabric where carbonate has filled holes left by rotated or fallen grains. Calcic and petrocalcic horizons are found in aridland soils in which carbonate is not leached effectively by available water (McFadden et al. 1991).
Gypsic and petrogypsic horizons are similar to calcic and petrocalcic horizons, but their cementing material is gypsum (Hartley & May 1998). Salic horizons are cemented with salts that are even more soluble than gypsum, including mirabilite and halite. These form in very arid climates such as the margins of desert and playalakes (Calvoetal. 1999). Entisol (incipient soil) The main feature of this order of soils is a very slight degree of soil formation, either because of a short time available or because of exceedingly unfavorable conditions. Entisols may be penetrated by roots and show some mineral weathering and surface accumulation of organic matter, but the original crystalline, metamorphic, or sedimentary features of their parent materials remain little altered by soil formation. Entisols are thus as variable as their parent materials, which range from alluvium, till and sand dunes to a variety of rocks. Their topographic settings are also variable. Most are found on young geomorphological surfaces such as floodplains and on steep slopes where erosion removes soil material as it is formed.The climatesin which they form also vary. Those forming in humid, warm climates where soil formation is rapid are younger than those formed in dry or cold climates. Early successional vegetation of grasses and other herbs and shrubs is characteristic of Entisols. Some of these soils on steep rocky slopes and along streams support trees. For paleosols the presence of root
Soilclassification traces is diagnostic of Entisols because in other respects they are little altered from their parent material. For paleosolsof deserts and of Ordovician and older geological age, before evolution of roots, a peak in magnetic susceptibility can be a useful indication of Entisols (Soreghan et al. 199 7). Some Entisols are too stony, infertile or poorly drained for cultivation. However, large areas of Entisols in alluvial bottomlands are cultivated for a variety of grain and vegetable crops, and grassed over for pasture (Buolet aI. 1997). Inceptisol (young soil) These soils represent a stage in soil formation beyond that of Entisols but still short of the degree of development found in other soil orders. They may have accumulation of clay in a subsurface horizon, but it is insufficient to qualify as an argillic horizon, which is diagnostic for Alfisols and Ultisols. Similarly, they may have organic matter at the surface but not so thick or abundant as in Histosols. Although varied as precursors of a variety of soils, a typical Inceptisolcan be imagined as having a light-colored surface horizon (ochric epipedon) over a moderately weathered subsurface horizon (cambic horizon). These soils have developed to the extent that some relict features from their parent material may be difficult to detect within the profile. These primary igneous, metamorphic and sedimentary structures normally take some time to be obliterated entirely. In humid to subhumid climates this may be only a few thousand years, and in drier climates tens of thousands of years. Inceptisols form in low-rolling parts of the landscape in and around steep mountain fronts. In sequences of alluvial terraces they form at intermediate positions between Entisols nearest the stream and better developed soils farther away from the stream. The parent material of Inceptisols is as varied as that of Entisols, but does not include vitric volcanic ash of Andisols. The climates of Inceptisols also are varied and their vegetation ranges from forest to tundra. The shrubby woodlands of pole trees that form during recolonization of disturbed ground by forest are especially characteristic, as are open woodland and wooded grassland. Many Inceptisols offer excellent natural grazing, and they can be cultivated to improve pasture and grow a variety of vegetable and grain crops (Buoletal. 199 7).
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Andisol (volcanic ash soil) These soils of volcanic ash are equivalent to Andosolsof the FA0 (1974,1988)classification. Siliceous volcanic ash consists of bubbles or shards of volcanic glass with a high internal surface area. This weathers rapidly to a mix of amorphous weathering products such as imogolite and base-rich clays such as smectite.These soilsthus are highly fertile, rich in organic matter, and have unusually low bulk density (Tan 1984; Shoji et al. 1993). These properties, along with imogolite and smectite, are commonly altered during burial, when distinctive minerals such as celadonite and clinoptilolite form (Retallack at al. 2000). The best guide to Andisols among paleosols is at least 60%recognizable pyroclastic fragments in thin section. One exception is the highly unusual volcanic ashes of carbonatites and nephelinites, which weather more l i e calcareous beach sand than siliceous volcanic ash (Retallack et al. 1995). Andisols form in a variety of geomorphological settings in and around volcanoes.Their overall profile form is most like that of Inceptisols, and l i e these Andisols are converted to more deeply weathered soils, such as Oxisols and Ultisols, with time. Andisols are economically important, fertile soils in tropical regions, widely used for subsistence or marketgardens, terracedforricepaddyor left for bamboo forestry. Histosol (peatysoil) These are organic-rich soils with thick peaty horizons, which form in cool, well-drained localities (folistic epipedon) or low-lying, permanently waterlogged parts of the landscape (histic epipedon). The main process in their formation is accumulation of peat, which means that organic matter is produced by growth of vegetation fasterthanitisdecomposedin the soi1.Thebreakdownof organic matter is inhibited by waterlogging because the most effective microbial decomposers require oxygen and this is rapidly depleted by microbes in stagnant groundwater (Falini 19 65). Sediment or rock underlying the peat may be altered by weathering (Bennett et al. 1991), and in some cases shows weathering from an earlier round of soil formation (Gardner et al. 1988). Associated with peat accumulation may be leaching or formation of gley minerals, such as pyrite or siderite, overprinting any prior soil or sedimentary features.
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Typical rates of accumulation for woody peat are 0.5-1 mm per year (Falini 1965: Moore & Bellamy 1973). Rates are much lower in swamps that are drained for a part of the year, as this allows seasonaldecomposition of peat. Rates are much faster under herbaceous vegetation such as marsh grass, moss and algae. The plant species of Histosols are usually low in diversity and restricted to suchwaterloggedsites (Retallacketal. 1996). Histosols support bog, swamp and marsh. These soils can be drained for cultivation, but are best left alone for specialty timber cutting, rough seasonal grazing or to maintain water quality. Vertiiol (swelling clay soil) These uniform, thick (at least 50cm), clayey soils have deep, wide cracks for the dry part of the year. This cracking may produce a hummock-and-swale topography (gilgaimicrorelief)and its subsurface expressionof adisrupted, festoon-shaped surface horizon (see Fig. 4.2). Pavements, fences and trees may be unbalanced by the strong shrinking and swelling action of smectitic clays in these soils. Other kinds of clay also are found, although they are less common (Paton 1974).Most Vertisols are found on materials of intermediate to basaltic composition. They may form in only a few hundred years on claystone, shales or mark of smectitic composition. It takes thousands of years for them to form on limestone,volcaniclastic sandstone and basalt. Vertisols are found mainly in flat terrane at the foot of gentle slopes.Their climate is dry enough and their vegetation sparse enough that alkaline reaction and good reserves of exchangeable cations can be maintained (Drieseet al. 2000). These soils occur in regions with subhumid to semiarid climates (180-1 5 2 0 m mean annual precipitation) with a pronounced dry season. Their vegetation ranges from grassland to open woodland. Wooded grassland is common. These soils offer excellent natural grazing, and with irrigation can be made to produce rice, cotton and sorghum.
Molliol (grassland soil) These soils have a well-developed, base-rich surface horizon of intimately admixed clay and organic matter. Abundant fine root traces and crumb ped structure
are characteristic (Pawluk & Bal 1985). Subsurface clayey (argillic or Bt), calcareous (calcic or Bk), or gypsiferous (gypsic or By) horizons also may be present, but are not definitive of the order. The characteristic surface horizon is created by fine root systems of grassy vegetation and the burrowing activity of diverse populations of soil invertebrates. Mollisols are found under grassland vegetation in subhumid to semiarid climates. Most are found in low, rolling or flat country. They form under a wide range of temperatures from the equator to the poles and in lowlands to high mountain meadows. They also form on a variety of parent materials but especially on base-rich sediments such as clay, marl and basalt. In dry regions these soils are used mostly for open-range grazing. In wet regions they are cultivated for wheat and maize and also can produce a variety of vegetable crops (Aandahl 1982). Aridisol (desert soil) These are soils of arid to semiarid regions. Rainfall in such regions is insufficient to leach soluble salts, and these soils have shallow (less than 1m deep) calcareous (calcic,petrocalcic or Bk), gypsiferous(gypsic,petrogypsic or By) or salty (salic or Bz) horizons. These cements form large nodules or continuous layers. The surface horizons of Aridisols are light colored, soft and often vesicular (McFadden et al. 1998). Subsurface horizons not cemented with salts, carbonates or sulfates may be similarlyfriable and silty, but many Aridisolshave clayey (argillic or natric) subsurface horizons. Both clay and carbonate in these soils are thought to be derived from weathering and flushing down the profile of extremely fine-grained dust of feldspar and other easily weatherable minerals, rather than the complex processes of weathering found in forested soils (McFadden et al. 1991).Aridisols are found mostly in low-lying areas because steep slopes in arid regions tend to be eroded back to bedrock andEntisols of badlands.Theparent material of Aridisols is varied. Unconsolidated alluvium, loess and till are common parent materials. Vegetation of Aridisols is sparse and includes various prickly shrubs and cacti. Aridisols can be irrigated for cultivation, but at the risk of salinization. They are best left for sparse nativegrazing(Gi1eetal. 1980).
Soil classification G e l i d (permafrost soil) Gelisols are soils with ground ice or other permafrost features within 1m of the surface. Ice is rarely preserved in paleosols (Sugden et al. 1995),but they may show clastic dikes, freeze banding or other deformations createdbygroundice(Krull1999;Retallack1999a,c).Another clue to Gelisols of the past comes from association with tillites and other glacigenedeposits (Daily & Cooper 1976;Williams 1986).Gelisolsvarywidelyin theirtime of formation from only a few years for frost heave of solifluction deposits, to millions of years on stable Antarctic land surfaces (Bockheim 1990, 1997).They form under polar desert, tundra and taiga vegetation (Bockheim 1995). Although mostly little-weathered boulder clays, Gelisols include a surprising array of histic epipedons, desert pavements, and salic and calcic horizons.These soilshave little agricultural use, and can also be difficult building sites because of problems with ice thawing. Spodosol(sandy forest soil) The diagnostic feature of these soilsis a subsurface horizon enriched with iron and aluminum oxides or organic matter (spodic or Bs horizon). In thin section these opaque cements form distinctive radially cracked, concretionary rims to abundant quartz grains (de Coninck et uI. 1974). Commonly the spodic horizon underlies a bleached, sandy near-surface (albic or E) horizon, although this is not essential for Spodosols.This has been a source of confusionbecause the broadly equivalent Podzols and Podzolic soils in the traditional sense (Marbut 193 5) have been recognized from their albicrather than spodichorizons, and albic horizons are found in both Alfisols and Ultisols in addition to Spodosols. Alfisols and Ultisols have abundant clay in their subsurface horizons, unlike Spodosols.Spodosolscan form in only a few hundred years on quartz-rich sands, but also can form by deep weathering of more clayey parent materials (Bloomfield 1981). They form on hilly bedrock or low, rolling quartz-rich sediments. Spodosols are found principally in humidclimates, in which clay and solublesalts are dissolved and washed out of the profile. Although most common in temperate regions, they also are found in the tropics and near the poles. Conifer forest is their
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most characteristic vegetation, but they also support other kinds of evergreen woody vegetation that can tolerate low nutrient levels and high soil acidity, Spodosols are naturally infertile and used mostly for softwood timber production (Bud et al. 1997).
AUisol (fertile forest soil) These base-rich forested soils have a light-colored surface horizon (ochricepipedon or A) over a clayeysubsurface (argillicor Bt) horizon that is rich in exchangeable cations (base saturation greater than 35%). Such base saturation can be assumed for paleosols when they containnodulesof carbonateina horizon (calcicorBk) deep (more than 1m) within the profile. If these are lacking, fossilAlfisols can be distinguished from otherwise superficially similar base-poor soils (Ultisols) by the abundance of base-rich clays, such as smectite, of easily weathered minerals such as feldspar (more than 10%in the 20-200 pn size fraction) or by molecular weathering ratios of alumina/bases of less than two (see Table 4.4). The prefix ‘alf-’ of Alfisols is derived from a traditional subdivisionof soilsinto calcareous (pedocals)and noncalcareous soils (pedalfers;Marbut 193 5). This can be a source of confusion because some Alfisols are calcareous. Alfisols form on base-rich parent materials in climates and under vegetation that allow maintenance of reserves of mineral nutrients. In general, this means sediments and rocks of intermediate to basaltic composition, climate ranging from subhumid to semiarid and vegetation ranging from grassy woodland to open forest. Their topographic setting and temperatures are extremely varied, although they are not found at the poles or on high mountain tops. When cleared they may be cultivated for a variety of fruit, vegetable and grain crops(Buo1etul. 1997). Ultisol (base-poor forest soil) Thesebase-poor forest soils aresimilar to Alfisolsin overall profile form as they include a well-developed clayey subsurface (argillic or Bt) horizon. Unlike Alfisols, Ultisols are more deeply weathered of mineral nutrients. They do not include calcareous material anywhere within the profile, have low reserves of weatherable minerals such as feldspar (less than 10%in the 20-
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200 pmfraction), and have molecular weatheringratios of alumina to bases of more than two (Table4.4). Basepoor clays, such as kaolinite, and highly weathered a h minous minerals, such as gibbsite, are common in these soils. Their low base status is commonlyrelated to a long period of formation (tens to hundreds of thousands of years). Over such a period they can form on a wide variety of parent materials. They form mostly on older parts of the landscape, such as rolling hills of bedrock, high alluvial terraces, and plateau tops (Markewich et d. 1990). Ultisols form most readily in humid, warm climates. There are some examples in polar and desert regions, but these are thought to be relicts of former climates more favorable to deep weathering. Their natural vegetation is coniferous or hardwood forest. Some also support wooded grassland, which can sometimes be traced to human modification of forest. Most of these soils are used for forestry In tropical regions, some produce pineapples and sugar cane. Some also are cultivated for vegetables and grain, but only after extensive fertilization.
Oxisol (tropical deeply weathered soil) These are deeply weathered soils with texturally uniform protiles, no more than trace amounts of easily weathered minerals, and dominated by kaoliitic clays or other base-poor oxides such as gibbsite or boehmite. They have only traces of easily weathered mineral grains such as feldspar, total bases in chemical analysis make up only a few weight per cent, and molecular weathering ratios of alumina/bases are 10ormore (oxic horizon). Deeply weathered mottled horizons (plinthite) also are found in these soils. A stable microstructure of sand-sued spherical micropeds of iron-stained clay is characteristic (Stoops 1983).Theadvanced weathering of these soils is due in part to their great age, often amounting to tens of millions of years. They are found mostly on stable continental locations on gentle slopesof plateaux, terraces and plains. Their development is especially favored by humid tropical climates, where weathering is most intense. They are known also in arid and cool climates, where they are usually found to be relict soils or to have formed on highly weathered sediments (Ollier & Pain 1996).Their natural vegetation is rainforest. Large areas of these soils on Precambrian rock in tropical regions are now covered with wooded
grassland, but were initiated as forested soils in the distant geological past (at least Miocene: Valeton 1999). Oxisols can be used for rough grazing and for forestry. Some produce tree crops such as cocoa, coffee, sugar cane and tropical fruits. A word of caution
These outlines and schematic representations of the FA0 and US soil taxonomy may be a useful initial guide, especiallyfor use with paleosols. However, such outlines do not do justice to these complex and evolving classifications. There is no substitute for carefully reviewing original accounts and comparing a paleosol in detail with one of therepresentative described soil profiles.The FA0 classification is particularly useful for finding, through its unique system of shorthand and indexing, comparablemodern soilscapesto suites of paleosols.For example, the Triassic paleosol of Fig. 4.9 (see also Fig. 10.11)is deeply weathered chemically,drab colored and enriched with clay in the subsurface IikeHumicAcrisols (FA0 code Ah), comparable with map unit Ah3-2bc of the Cowlitz-Chehalis lowlands of Washington state, USA (FA0 1975a) and map unit Ahl-3a south of Auckland, New Zealand (FA0 1978b). Even with excellent field and laboratory data for a paleosol it may not be possible to identiiy it within a modern soil classification. Some paleosols can be identified only to suborder level in the US classification, whereas others can be keyed to subgroup level. For example, the only AKsols with a continuous calcareous horizon at depth are Petrocalcic Paleustalfs. Fossil Vertisols, on the other hand, can seldom be identied beyond order level because of criteria involving duration of open cracks not preserved in paleosols. In addition, some soil types may be extinct, particularly those of the Precambrian formed under oxygen-poor atmospheres (Rye& Holland 1998). At first appearance, these classificationsand their terminology may seem intimidating. My recommendation is toplungeinandmemorizethe essential features of the 12 soil orders of the US soil taxonomy. With use they become familiar, and the similarities with other classifications are more striking than the differences. Facility with the language of soil classification is useful in many walks of life related to agriculture, land planning and politics.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 6 Mapping and naming paleosols
Paleopedology is a field science. Its objects of study are too bulky to be brought back into the laboratory in their entirety, and must be characterized and sampled outdoors. Field work commonly involves identifying different kinds of paleosols and establishing their relationships with each other and the enclosing sediments or rocks. These dual activities of mapping and naming are a necessary first step in the study of paleosols. Field observations determine how a paleosol is sampled and later analyzed. More sophisticated indoor studies refine or test hypotheses developed in the field, but seldom supplant them. Thus, it pays to have a logical plan for mapping and naming paleosols. and this plan will vary according to the aims of the study. Three common aims are considered in this chapter: paleoenvironmental interpretation of paleosols, use of paleosols as stratigraphic marker horizons and portrayal of paleosols on geologicalmaps. In the face of a burgeoning scientific nomenclature for different kinds of fossils,rock layers and soils, it may fairly be asked whether there is a need to map and name paleosols. The formal mapping and naming of fossils, rocks and soils often becomes an end in itself. Their nomenclature is regulated by such long-winded and legalistic documents as the International Code of BotanicalNomenclature(Vossetal.1983)andlnternational Code of Zoological Nomenclature (International Commission on Zoological Nomenclature 1995), the International Stratigraphic Guide (International Subcommission on Stratigraphic Classification 1994), numerous local stratigraphic guides (North American Commission on Stratigraphic Nomenclature 1982) and the USDA Soil Survey Manual (Soil Survey Staff 1993)..Thenamed entities are redefined, emended and debated. If this is the result of naming and mapping, should not paleosols continue to be studied on an informal basis? It is also artificial to bestow biological binomial Latin names on such things as fossil tracks and trails (Bromley 1996). and on the acid-resistant walls of plant spores and
pollen (Traverse 1988).Yet experience in ichnology and palynology has shown that common names, acronyms or descriptive alphanumeric terms have failed to yield replicable results. Names that are concise, constructed along a pre-established set of rules, and based on real and well-illustrated examples are needed. In this way, a large data base that is internally consistent can be built and r e h e d over the years by different investigators. Mapping and naming of paleosols has not yet been legalizedto theextent of fossils,rocks and soils, but there are signs that this is happening. The North American Stratigraphic Code (North American Commission on Stratigraphic Nomenclature 1982) includes provision for formal naming of paleosols as stratigraphic marker horizons. Perhaps there will follow a national registry for such names, as is maintained by the US Geological Survey for stratigraphic names (searchable online at http://ngmdb.usgs.gov/Geolex).Other approaches to paleosol mapping are accumulating in scientific publications. These other names must be regarded at present as informal. The distinction between formal and informal names can be indicated by capitalization. The Sangamon Geosol, for example, is a formal name, whereas the type Gleska clay paleosol is not. Until there is an officially endorsed code for naming paleosols there is scope for a variety of approaches to mapping and naming them.
Paleoenvironmentalstudies My own interest in paleosols has been concerned with interpreting ancient environments of soil formation. Each type of paleosolrepresents adistinct paleoenvironment. There are no special difficulties in naming a paleosol in this context if it involves a single profile at a geological unconformity. If work is extended to characterize variation in paleosols along an unconformity or in numerous paleosols interbedded with fluvial deposits, then many distinct names may be needed in a small
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area. In a study of paleosols in Eocene-Oligocene (39-29 Ma)volcaniclasticalluviumof thePaintedHills, central Oregon, 435 successive paleosols of 1 7 kinds were recognized in 430m of stratigraphic section (Retallacketal. 2000).In theSiwalikGroupof Pakistan, India and Nepal, there is on average a paleosol for each meter of its 6 km thickness (Retallack 1991a; Quade et ul. 1993, 1995; Tanaka 1994. 1997). For paleoenvironmental studies of such sequences a separate name is not needed for each paleosol, only for each kind of paleosol insofar as it represents a discrete subenvironment of past landscapes. Such studies are similar in concept to modern soil mapping except that the soils of many superimposed landscapes must be named and mapped rather than those of just one soilscape. Although modern soil mapping units were not designed with paleosols in mind, they are serviceable for paleopedological studies. The soil units defined by the US Department of Agriculture (Soil Survey Staff 1993) are used for naming and mapping variations in soil across the landscape (Fig. 6.1). In addition to numerous published soil surveys, data on soil series are now searchable online (at http://www.statlab.iastate.edu/soils/osd). Soilscapesare considered to consist of a number of irregularly shaped polypedons that interlock like a large mosaic (see Fig. 2.1). A pedon is the smallest volume of soil that can be recognized and identified. It may be a substantial area of soil (10m2 or so) if soil horizons are laterallyvariablein thickness. Usuallyapedon is the profile visible in a small pit (some 1m2 or less in area). The task of the soil mapper is to compile information from a number of soil pits so that areas of substantially similar soil may be delimited from areas with different soils. The basic unit of mapping is a soil series,which includes one or more polypedons scattered over the landscape. A soil series is based on field and laboratory characterization of a particular pedon and named after a nearby geographical locality,e.g. the Bellpine Series (Fig.6.1). An individual pedon can be named for the texture of its surface horizon, e.g. the Bellpine silty clay loam. With local experience,this name conjuresup an image of a thick, red, clayey soil (a Xeric Haplohumult) that forms on moderately sloping,well-drained hills of marine volcaniclastic sandstones under mixed conifer forest dominated by Douglas fir (Pseudotsugumenziesii). This soil mapping system also allows for natural variation and relationships between soils. Minor fea-
tures of the soil such as small ferruginous nodules may be used to define soil variants, such as the Bellpine silty clay loamnodularvariant.Slight variations in thickness of horizons may be used to define phases, such as the Bellpine silty clay loam eroded phase. Soil series also can be grouped into larger units or associations. The Bellpine Series, for example, is part of an association of soil series formed on moderately sloping sedimentary bedrock. These are different from alluvial soils of basaltic gravel in the nearby valley bottoms. For paleosols the problem of mapping and naming is conceptually similar. Paleosols may be characterized as they are encountered in a stratigraphic sequence by examining cliff outcrops, excavating badlands or logging drillcore. Each recognizable paleosol profile can be assigned to a limited number of generalized kinds. The term ‘series’of soil science is confusing when applied in geological settings because of the long-established use of Series as chronostratigraphic units (North American Commission for Stratigraphic Nomenclature 1982). The alternative term pedotype has been coined for a distinctive kind of paleosol (Retallack 1994a), and may prove useful for soils as well. Pedotypes are ideal for reconnaissance mapping and logging of paleosol sequences (Fig.6.2).Foreachpedotype, aprofileshouldbe sampled and described in detail as a reference standard. With laboratory studies to determine the texture at the ancient surface, it may be possibleto assign a name to an individual paleosol,e.g. the type Gleska clay paleosol. As in soil mapping, it is best if the type profile of a paleosol pedotype is distinct and typical rather than a composite profile or an intergrade between different pedotypes. This in the end is less critical than its careful and unambiguous characterization, so that later investigators can find and assess it for themselves. This kind of mapping and naming has been applied to sequences of paleosols as measured in the field or from drill cores. The lateral distribution of different kinds of soils across an ancient landscape can be reconstructed from a vertical section by using a common kind of geological inference often called Walther’s Facies Law (Boggs 1995). This states simply that different kinds of sediment deposited side by side in nature will be preserved one on top of another in a sedimentary sequence. Stream channels, for example, migrate laterally across their floodplains as they erode the outside of their meanders. Their levees and point bars inside the meander follow this lateral migration so that point-bar deposits
Mapping and naming paleosols
RIVERWASH Fluvent
CAMAS Fluvent ic Hoploxeroll
NEWBERG
Fluventic Hoploxeroll
CLOOUATO
Cumulic Ultic Hoploxerol I
Figure 6.1 Soil series and vegetation along the McKenzie andwdlamette rivers near Eugene in the southernWillamette Valley of Oregon,USA. KeytolithologicalsymbolsisgiveninFig.6.2 (datafrompatching198 7).
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Chapter6 ClOStlC
grain size
develooment colcoreousness hue
..
KEY
. lbreccio red claystone E *A
groy claystone Elbreccia F' ....1 . sandstone Aa
0
siltstone
claystone
I=I2 & g @ cross bedding u scour ond fill IVI ripple morks contact Eerosional ferrugintzed Esurfoce red color brown color drab haloed root tmcos root troces
Elcalcoreous nodules bonasand teeth coprolites snoils burrows hackberry Elendocarps
w
rTf 0
Figure 6.2 Pedotypesrecognized in a measured stratigraphic section through Eocene and Oligocene (35-29 Ma) alluvial rocks in the Pinnacles area of Badlands National Park, South Dakota, USA (seealso Fig. 2.5). Positionof paleosols is marked by boxes whose widthcorresponds todegreeof development (Table13.1),caIcareousness by reaction withdiluteacid(Tab1e 3.2) and hue froma Munsellcolor chart (data from Retallack 1983a: Terry 1998).
Mapping and naming paleosols become overlain by levee and then floodplain deposits in a characteristic fining-upward sequence. The relationship of paleosols in a sequence to point-bar, levee and
81
floodplain sediments can be a guide to their former distance from streams and distribution across alluvial landscapes (Fig.6.3).Suchfacies interpretations should
Figure 6.3 A reconstruction of pedotypes and vegetation during early Oligocene time in the Pinnacles area of BadlandsNational Park, SouthDakota. USA. LithologicalsymbolsareasinFig. 6.2 (fromRetallack 1983a: reprintedwith permissionfromthe Geological Societyof America).
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be restricted to a single genetic package of sediments and not extend across a major unconformity. Such unconformities can be discerned by the presence of paleosols that are especially thick or well developed, or so closely superimposed that they overlap in a confusing manner (Retallack1998a). These techniques for mapping and naming soils can be applied also to the study of lateral variation of paleosols across a single buried land surface. This straightforward way of reconstructing soilscapesis complicated by difficulties in determining relative ages of different parts of the landscape. Let us consider, for example, the modern rock-bound alluvial landscape of the Willamette River in central western Oregon (Fig. 6.1). The bedrock unconformity may have been covered by a soilscapewhen 6rst it was excavated,but progressive infillingwith alluvium hasresulted incontinuedmodification of soils high in the landscape long after the lowest soils were covered by river deposits.Thus,the buried soils on the unconformity between modern alluvium and bedrock represent conditions of the distant past whereas higher soils represent a combination of past and present soil-forming conditions. On the present land surface also, the soils are of very different age. Those closest to the stream are the youngest and in some cases are forming on flood deposits only a few years old. Soils of higher terraces are somewhat older and soils on hilly bedrock are the most ancient. Because landscapes are renovated in this patchwork fashion, lateral mapping of paleosols is not necessarily a good guide to soilscapes of the past. There are rare cases where time planes can be identified using paleomagnetic reversals (Behrensmeyer & Tauxe 1982) or volcanic ash beds (Burggraff et al. 1981). These allow reconstruction of an ancient soilscapein detail. Another problem in naming paleosols is finding a name that is distinctive and appropriate, and that is not already used for a stratigraphic unit, modern soil series, geomorphologicalsurface, distinctivefossil faunule or other natural feature. In remote regions there may be few place names of any kind that can be used (Retallack & Krull1999). Even in well-populated regions the diversityof paleosolswithin small outcrops may quickly exhaust availablelocal names. In my own studies, I have used descriptive terms from native languages such as Lakota Sioux (Native American: Retallack 1983a) and Dholuo (Kenyan: Retallack 1991a). Even with a
distinctive name, a paleosol should be clearly labeled a paleosol, to prevent confusion with modern soil series, e.g. the Bellpinesilt loam vs. the Gleska clay paleosol. The term ‘soil facies’ has been used (by Birkeland 1999) in a similar sense to pedotype, as outlined here. Soil facies are lateral subdivisions of soil stratigraphic units, just as sedimentary facies can be lateral subdivisions of sedimentary formations. Pedotypes, in contrast, can be side by side, superimposed or unconnected with one another. Unlike sedimentary facies, both soil facies and pedotypes represent conditions during nondeposition rather than during sediment accumulation. Another comparable term is ‘pedofacies’(Bown & Kraus 1987;Kraus 1987),definedas ‘laterallycontiguous bodies of sedimentary rock that differ in their contained laterally contiguous paleosols as a result of their distance (during formation) from areas of relatively high sediment accumulation’ (Fig. 6.4).This differs from Birkelands concept of soil facies and my concept of pedotype, in that it includes sedimentary rocks of specific lateral relationships and paleoenvironments. This definition has also meant that pedofacies cannot be recognized in weakly developed and cumulic paleosols (Kraus 1992, 1997. 1998; Kraus & Aslan 1999). ‘Soil facies’ and ‘pedofacies’both have connotations in sedimentary geology that are avoided by the term pedotype.
s
E
D
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~
~PEDOFACIES ~ ~ ‘ yellow
m i n t e r b e d d e d m brown sand
red
PEDOTYPES OVw
llrJon BOttOmlY
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GEOSOLS
Rivers’de LXIIIUUUUUULIUUII Farmville Upton
r.n.rlrnll Hilldale
Figure 6.4 Therelationshipbetween sedimentary facies, pedofacies,pedotypes and geosols (fromRetallack 1998b;with permission from Quaternary International).
Mapping and naming paleosols
83
guish from one another. Paleosols also have been used as stratigraphic markers in pre-Quaternary rocks (Ortlam 1971: McCarthy et al. 1998). Geosol is the formal term now recommended for such soil stratigraphic units by the North American Commission on Stratigraphic Nomenclature (1982).A geosol is not a soil or paleosol, but rather a whole soilscape that can be recognized as a laterally extensive stratigraphic horizon. GeosoIs are named from localitiesor areas, e.g. Sangamon Geosol from the county including Springfield, Illinois (Follmer 1978). From the example of the southern Willamette Valley already considered (Fig. 6.1), it can be seen that the relationship of a soilscape to landscape development is
Stratigraphicstudies An additional use for paleosols is as stratigraphic marker horizons: distinctivehorizons that can be traced to establish the relationship in time and space between different sedimentary units. Paleosolshave proven especially useful for establishing relative age within Quaternary sediments that are too old for radiocarbon dating and too young or poorly fossiliferous for establishing their age paleontologically(Fig. 6.5). Quaternary paleosols may be distinctive, red, clayey (Follmeret al. 1979) or stronglycalcareousrocks (Busacca 1989).Theyform natural divisions between deposits of alluvium, loess and till that may look so similar as to be difficultto distin-
I
I
ILLINOIS
SIERRA NEVADA
I
WISCONSINAN
Figure 6.5 A correlation of Quaternarysoil stratigraphicunits (geosols)and sedimentary formationsbetween centralIllinoisand the WesternsierraNevada, California,USA (data fromFollmeretaZ.1979;Birkeland1999).
~ weak
strong
~
~
e
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~
loess ~
~
~ till ~
O
l
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seldom as simple as the ‘layercake’ arrangement of laterally extensive layers ideal for stratigraphic studies. Typical complications can be codified to define different kinds of geosols (Fig. 6.6). Geosols may be found either as soils at the surface or buried in stratigraphic sequences. Those at the surface are called relict geosols when lateral mapping or other techniques demonstrate that they formed under soil-formingconditions different from those of the present. Buried geosols, on the other hand, have clearly ceased soil formation. A welldeveloped, buried geosol of a stream terrace may splay into several, weakly developed buried soils interbedded within near-stream deposits. Where this local multiplication of more or less similar paleosolscan be tracedinto a well-developed and distinctive geosol, it is more productive for stratigraphic purposes to recognize it as a divided geosol, rather than name each one separately. On the other hand, regional mapping may show that one or more distinctiveand stratigraphically useful geosols are amalgamated in especiallystable parts of the landscape. These are recognized as a compounded geosol. These distinctions work in practice because for stratigraphic purposes emphasis is laid upon those paleosols that are distinctive and laterally extensive. Other paleosols can be accommodated within the general framework of these well-characterized paleosols without formal names. Lateral variants of geosols could be named pedotypes in the same way that lateral variants of sedimentary formations are calledsedimentary facies.These two different approaches to the same materials emphasize different features of paleosols. Geosols are recognized from features thought to reflect time. Within sequences deposited by successive advance and retreat of glaciers this includes the degree of developmentof a paleosol insofar as it reflects the time between glacial advances.
It also includes the degree of calcareousness, clayeyness or red hue as this is related to regional climatic fluctuations. In contrast, pedotypes are recognized by a wider variety of features that reflect soil-forming processes specific to a particular site. These include patterns of root tracesinsofar as they relate tolocalvariationsinformer vegetation, or drab hues and peat accumulation as indications of locally waterlogged parts of the landscape. The use of paleosols in mapping and correlating Quaternary sediments is a well-established scientific tradition (Chamberlain 1895),now applied throughout the USA (Fig. 6.5), Europe, Africa and China (Paepe & van Overloop 1990). The term ‘geosol’ was proposed by Morrison (1967) to replace a previous usage of ‘soil’. Soil has such varied meaning to soil scientists, engineers, farmers and geologists that it was deemed inappropriate for this special use. A geosol is not a soil in the senseof apedon, aconstruction material, an agricultural resource or a zone of weathering: rather, it is an assemblage of paleosols representing an ancient landscape. The need for a new term was independently recognized by Brewer et al. (19 70). who proposed the term ‘pedoderm’for much the same reasons. Pedodermis officially recognized in Australia (Stratigraphic Nomenclature Committee 1973). but geosol is now more widely used. The main theoretical difficulty with the use of geosols for stratigraphic mapping is the assumption that they are of the same age everywhere. Morrison (1978) has argued that geosols represent geologically brief soilforming intervals, when climate was especially conducive to soil formation. These soil-forming intervals were envisaged to punctuate longer periods of time during which climate was too cold or dry for soil formation. The opposite view that soil formation proceeds at a con-
Figure 6.6 Concepts and terminology for the recognitionof geosols (fromMorrison19 78: reprintedwith permission from Geoabstracts Inc.).
Mapping and naming paleosols
85
Figure 6.7 Weatheringprofiles,their stratigraphic relationshipand distribution in south-westernQueensland,Australia (basedon dataof Senior&Mabbutt19 79).
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stant rate unless interrupted by sedimentation or erosion is equally unjust. The true position probably lies somewhere between: the rate of soil formation varies considerably with climatic conditions, degree of prior development and other factors (Birkeland 1999).With the continued applicationof thermoluminescence, uranium series and accelerator radiocarbon dating, geosols are commonly found to have very different ages from those expected from correlation (Feng et al. 1994; Hart &Peterson 1997; Richardson et al. 1997). Geosols are not markers of time planes as ideal as ash beds. They are imperfect, yet useful stratigraphic markers. Many regions have a shortage of names unused for other natural features that might be appropriate for geosols. This problem is not as serious as for naming more numerous paleosol pedotypes because weakly developed or laterally impersistent paleosols or the individual profiles of a divided geosol are seldom so stratigraphically significantthat they need to be named. Geosols are named only when they are mappable over a wide area. The Sangamon Geosol, for example,has been traced through much of the midcontinental USA from Ohio toTexas. Some soil scientistswho find the proliferation of geosol names confusing prefer to adapt preexisting stratigraphic names. The paleosols on top of the Tahoe Till (Fig. 6.5), for example, could be called the post-Tahoe Geosol and those underneath it the preTahoe Geosol. A similar practice has been adopted informally for naming Precambrian paleosols at major geologicalunconformities, for example,the pre-Pongola paleosol(Grandstaff etal. 1986).
Deeplyweathered rocks Somezones of deep weathering are so thick and distinctive that they cover larger areas of ground than geological formations (Fig. 6.7). They are not the same as geological formations because their lower boundary is diffuse and because several different thick weathered zones may overlap substantially. Nor are they the same as geosols,which consist of soil solum. Deep weathering profiles, in contrast, are mainly saprolite and thick duricrusts: laterites, bauxites, calcretes and silcretes (Milnes
1992; Thiry 1999; Valeton 1999). The soil solum seldom is preserved in these deep weathering profiles, which can reach hundreds of meters into their parent materials (Ollier&Pain 1996). A way of portraying these deeply weathered zones on geological maps has been suggested by Senior & Mabbutt (1979), who proposed using a locality name together with the simple and succinct term ‘profile’.It may also be desirableto use a term descriptive of the profile, e.g. the Curalle silcrete profile of south-western Queensland, Australia (Fig. 6.7). Terms other than the locality name should not be capitalized because at present these names are informal. Nevertheless, they should be defined according to standards appropriate to other stratigraphic units. Several features of a profile should be characterized for adequate definition: the locality of a reference section, a description of the lithologiesandthicknessof theparent andalteredrocks within the reference profile, an estimate of the thickness and other variations in the profile elsewhere,and indications of its age, distribution and relationships with enclosing rocks. Advantages of this system lie in its adherence to general principlesfor establishing other stratigraphic units. This approach clarifies the only record of much geological time in some continental regions. With such difficult and controversial rock materials it is important to have agreed names for examples independent of theories of their origin. Such a reconciliation of opinion may be overoptimistic because deep weathering profiles often combine in exaggerated form many of the problems of paleosols and pedotypes. Profiles may vary considerably along the same land surface depending on parent material, age and other factors. It may be difficultto attribute variations to any particular paleoenvironmental factor such as climate because this may have changed many times over the long period of time during which the profiles formed. Deep weathering profiles may have continued to form in some places long after they were buried or partlyerodedin anothecTheexistence of such complexities makes it necessary to have generally agreed nongenetic procedures for mapping and naming.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 7 Alteration of paleosols after burial
A part of the problem with unraveling alteration of paleosols after burial is clarifying terminology for these changes. Diagenesis is the term used in studies of sedimentary petrology to describe alteration after deposition. Diagenesis thus includes soil formation,which on Earth occurs at 1bar of pressure and temperaturesfrom 84°C to -88°C (Kimmins 1997). Additional changes afterburial pass into what would be called metamorphic alteration at temperatures of 200Y or more (Fig. 7.1). Because of the natural gradient of increased temperature downward in the Earth’s crust, there is a pressure limit to diageneticalteration of less than 7 kbar of lithostatic pressure, which is found at depths of about 30km within the crust. Metamorphic alteration proceeds at more extreme temperatures and pressures where new mineral assemblages form at the expense of the old ones androckstructureisreconstitutedbythegrowthof new crystals. Destruction of primary features of soils and imposition of metamorphic foliation, schistosity and crystalline texture l i t interpretation of paleosols. However, many diagnostic features of paleosols such as root traces, soil horizons and soil structure survivesuch changes. These structures have been found in paleosols metamorphosed to zeolite, prehnite-pumpellyite and lower greenschist facies (Retallack 1992b,c). With greater metamorphic alteration, the only remaining indication of a paleosol may be a highly aluminousbulk composition and minerals such as kyanite, sillimanite, garnet or corundum (Barrientos & Selverstone 1987; Dash et al. 1987). With only this compositional and mineralogicalinformation it may be difficult to distinguish between paleosols and zones of hydrothermal alteration. Although diagenesis is defined as alteration of sediments after deposition, it sometimes is taken to mean only alteration after burial. Mere documentation of dissolution of grains and their coating with clay may provide evidence of diagenesis, but does not in itself demonstrate whether it occurred in the soil or after its
burial (Mausbachetal. 1982).Much informationabout soil formation of the past is masquerading under the guise of diagenetic studies of nonmarine rocks. The distinction between alteration occurring during soil formation and that attendant on burial can be made by establishing the relationship of mineral grains and other features with those parts of the soil thought to have been original,such as root traces. Field relationshipsprovide the most s t r i n g and unequivocal illustrationsof the relative time for formation of soil and burial diagenetic features. Sand-filled burrows approaching,then sidling around siderite nodules have been found (Fig. 7.2), unequivocally indicating that the nodule or some precursor of it was there when the burrows were excavated. Nodules of this kind also are formed in marine and lacustrine rocks (Boggs 1995), and commonly are assumed to be features of burial diagenesis. The following additional features of pedogenic nodules may prove useful in distinguishing them from burial-diagenetic nodules (Retallack1983a; Downing & Park 1998). Fossil root traces may sidle around them, but if penetrating the nodule are better preserved and more fully inflated than in the surrounding matrix. Other fossils such as skulls,turtle shells and hollow bones alsomaybe less compactedwithinnodules than outside them. This is evidence that the nodules formed before there was any appreciable compaction and that they subsequently resisted crushing better than the surrounding matrix. Original pedogenic nodules also may contain more easily weatherable minerals such as pyroxene and olivine than are found in the surroundmg matrix. This indicates that nodules formed during soil developmentprotected these minerals from further weathering. Nodules of paleosols may form horizons at a fixed depth below an ancient surface. Finally, nodules of the same kind within paleosols and as pebbles in associated fluvial paleochannels provide evidence that the nodules were formed in the soil and eroded out of the banks of streams. 87
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..
im
200
300
LOO
500
600
700
800
km 10 15
M 25
30 35
100
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600
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700 OC
Cross-cutting relationships in petrographic thin sections also provide evidence for the sequence of mineral development in paleosols. Cross-cutting relationships may be difficultto see within clay-sizedminerals,but are clear between larger crystals. Cross-cutting relationshipsin very fine-grained materials can be made obvious by back-scatter electron microscopy (Krinsley& Manley 1989; Retallack & Krinsley 1993). Diagenetically formed crystals may be large enough to cut across large structures such as burrows. If, on the other hand, the burrow wall truncates or has rotated crystals, then the burrow can be assumed to be of about the same age as or
Figure 7.1 Pressureandtemperature conditionsof diagenesisand various metamorphic mineralassamblages, includinghigh-temperature aluminosilicates (kyanite,sillimanite)commonin metamorphosedpaleosols(fromWinkler 1976:with permission from Springer-Verlag).
younger than the crystal and the crystal an original part of the soil (Retallack 1985.1993b). Long sequences of events can be reconstructed by stringing together such observations. In the search for order, however, it should not be overlooked that a particular mineral may be of different age in different parts of the paleosol. Care and patience are needed to assemble a representative paragenetic sequence. Other kinds of evidence also may be usefuI in unraveling the sequence of diagenetic modifications.For example, the strontium isotopic composition of clay coatings to grains of Miocene to Oligocene sandstones from the
Figure 7.2 Burrows (metaganotubules)of cicada-like insectssidlingaroundsiderite nodulesin a very weakly developed paleosol (Aquentof Warriewoodpedotypeof Retallack 1997d)fromtheEarlyTriassic (245 Ma).GosfordFormationnearAvoca, New SouthWales,Australia.Thecoinis28 mmindiameter (fromMcDonnelll974; with permission from the Geological Societyof Australia).
Alteration of paleosols after burial Great Plains of North America provides clues to their age (Stanley&Faure1979).Themain sourceof thelight isotope (*%r) is the radioactive decay of rubidium (86Rb).The ratio 87Sr/s6Srcan be used for radiometric dating, but clay coats arenot suitablyfreeof contamination from groundwater for accurate dating. Nevertheless, strontium isotopic ratios demonstrated younger geological age of samples collected higher in the sequence. This is consistent with their origin during or shortly after deposition rather than by equilibration with groundwater at a much younger date. A variety of geochemical approaches to dating paleosol features are now being exploited (Vasconceloset al. 1992; Rasbury etal. 1998). Another method for establishing the age of alteration seen in paleosols is by considering how consistent a feature is with known kinds of alteration after burial. Various features of paleosols can be compared to assess whether their conditions of formation, such as Eh and pH (Fig. 3.14). are internally consistent. Well-known kinds of alteration after burial also can be considered. The rest of this chapter constitutes an exploration of these.
Burialdecomposition of organic matter Studies of Quaternarypaleosols and equivalent surface soils in central North America have shown that soon afterburial, paleosols lose up to an order of magnitude of organic carbon as determined by the Walkley-Black technique, but paleosols may preserve the general trend
I
2
89
of their depth function for organicmatter (Fig. 7.3).This loss of organic matter was found only in paleosols similar to well-drained surface soils, not in paleosols that were peat-rich and waterlogged (Stevenson 1969).The lost organic matter may have been metabolized by aerobic microbial decomposers that were part of the later ecosystem forming soil on the sedimentary increment that buried the paleosol. Additional loss of organic matter may occur during deep burial and the generation of oil and gas (Surdam & Crossey 198 7). Many paleosols contain much less organic carbon than one would expect considering their fine structure and abundant root traces and burrows that indicate former biological productivity For example, Miocene paleosols from Kenya and Pakistan (Retallack 199la), have very low analytical values of Wakley-Black organic matter (nomore than0.65wt %). Generally comparable surface soilsof grasslands on the Serengeti Plain of Tanzania have organic carbon contents of up to 9 wt % (de Wit 1978;Jage 1982).Even tropicalforestedsoilsof the Songhor area of Kenya have organic carbon values of up to 5.64wt % (Thorp et al. 1960) and in those of the Indo-Gangetic Plains of India and Pakistan organic carbonisup to 2.68wt% (Murthy et d . 1982). Loss of organic matter in paleosols may change their color so that it is less turbid in appearance (higher Munsell chroma) than for the original soil. Loss of organic matter also makes the paleosols more prone to badlands weathering in the modern outcrop, compared with original soil clays stabilized by roots and organic coatings around soil clods.
weight percent organic corbon
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modern soil
05 f
P
=
c
modern and buried soils
I
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c b u r i e d soil
Figure 7.3 Comparisonof organic carboncontent determined by Wakley-Black titrationin some paleosolsand comparable surface soilsof the mid-westernUSA. (A)Yarmouthian(600Ka) paleosol and surface AlbaqualE @)Yamouthian(600Ka) and surfaceHapludalf;(C) Farmdalian ( 3 0 Ka) paleosol and surfaceHapludalf;(D) Wisconsinan(14Ka)paleosoland surfaceUdiiluvent (modifiedfromStevenson1969;with permission from Williams & Willrins Co.).
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Burialgleizationof organic matter Bluegray or green-gray color of horizons, of mottles or of haloes around root traces provides a striking contrast to red parts of many paleosols. In some profiles, the light-colored surface horizon and root traces extending down into a reddish subsurface horizon are superficially like the arrangement in some forested soils (AEsols, Spodosols and Ultisols), in which a sandy near-surface horizon (eluvialor albic horizon) has been bleached and passes down irregularly into a subsurface horizon (illuvial or argillichorizon) that is clayey and brown (see Fig. 3.8A). Root channels and burrows form natural conduits for leaching and for the penetration of the lightcolored sandy material deep within the profile, in the form of distinctivestructures called glossic features (Soil Survey Staff 1975) or krotovinas (Buol et al. 1997). A few paleosols of this kind are known (Retallack1976). Bluish and greenish colors are found in clayey and peaty waterlogged soils (Vepraskas & Sprecher 1997), and are a conspicuous feature of groundwater gley. These drab colors also are found in clayey soils on which stagnant water ponds for a good part of the year, as features of surface-water gley (Bouma et al.1990). In both cases, anaerobic microbes in poorly oxygenated water have reduced brown and red iron oxides and hydroxides from the ferric state to a drab-colored ferrous state. Because ferrous iron is much more solublethan ferric iron, such gleyed layers or spots in soilscommonly are strongly depleted in iron compared with other parts of the soil. Gleyed soils commonly also show other evidence of iron remobilization, such as nodules or layers of vivianite, pyrite, siderite and amorphous iron-manganese oxides and hydroxides(Mooreetal.1992;Shotyk1992).Many gleyed soils have peat-rich surface horizons, and wellpreserved tabular root systems and leaf litters. Such drab paleosols rich in fossil plants are widely known, especially in coal measures (Hughes et al. 1992; Retallack 1994a). Drab horizons and mottles also can form by burial gleization (see Fig. 3.7).This is the chemical reduction of iron hydroxides and oxides by anaerobic bacteria consuming organic matter buried with the soil below or near the water table. Reduction haloes around buried organic matter can form in only a few hundred years (Allen 1986b;MathewsonetaZ. 1992).Burialgleization is an especially appealing explanation for drab-colored
surface horizons and drab-haloed, deeply penetrating, root traces in clayey paleosols with red oxidized subsurface horizons, and lacking textural differentiation of horizons or evidence of former waterlogging. There are manysuchpaleosols (Retallack 1983a, 1991a, 1997a; Blodgett 1988).Geochemicalstudiesrevealthat insome cases total iron is depleted in drab compared with red parts of the profile, but other examples have equivalent amounts of chemically reduced iron remaining in drab areas(Retal1ack 1976,1983a;Kraus&Aslan 1993),as might be expected during alteration in a slowly permeable system soon after burial. It is likely that a variety of anaerobic microbes common in waterlogged soils play a role in early burial gleization. The inorganic chemical reduction of iron oxides is thermodynamically difficult under burial conditions milder than amphibolite grade metamorphism (Thompson 1972), as shown by redand purple-colored and red-green mottled paleosols altered to the greenschist facies of regional metamorphism (Retallack 1985; Feakes & Retallack 1988).The original color of parts of paleosols chemically reduced by burial gley was probably brown, orange or yellow, darkened with organic matter, as is usual for the surface horizons of well-drained soils. Some paleosols lack any evidence of burial gley. Among Miocene paleosols of Kenya, those formed in alluvial lowlands around the base of carbonatitenephelinite cinder cones and stratovolcanoes (forexample, at Songhor) had abundant evidence of burial gley, but those formed on nearby volcanic footslopes (for example, at Koru and Fort Ternan) did not show burial gley of either root traces or horizons (Retallack 199la). Such observations support the idea that subsidence below water table is necessary for active microbial reduction of buried organic matter. In contrast, paleosols buried and remaining above water table lost organic matter by aerobic decomposition.
Burialreddeningof ironoxides and hydroxides Many paleosols are a 6re-engine red color (Munsell hue 5YR or redder), but most soils are brownish red. The degree of redness of soils is due mainly to the nature and grain size of iron oxide and hydroxide minerals (Hurst 1977; Morris et al. 1985; Blodgett 1988),formed by oxidation during weathering of iron-bearing minerals in
Alteration of paleosols after burial the parent material. These form brown hydroxides such as goethite at first, and then dehydrate and recrystallize to coarser grain size with further weathering to the red iron oxide, hematite. Red soils are found mainly in tropical regions and on land surfaces hundreds of thousands to millions of years old (Blodgettet al. 1993). There are many brick-red paleosols that appear neither to have formed under a tropical climate nor to have beenso strongly developed(Retallack199 7d).Forthese, a more likely explanation of their red color is burial dehydration and recrystallization of iron hydroxides to hematite. This process during burial is petrographically indistinguishable from the formation and maturation of iron hydroxide minerals in soils. Iron-bearing minerals are mantled with weathering rinds and wisps of opaque red hematite crystals that extend out into the matrix. In a few Quaternary paleosols and comparable soils of central North America for which this change has been assessed, the change in color during shallow burial for only a few tens to hundreds of thousands of years amounted to 2-3 Munsell hue units, from lOYR to 7.5YR or from 7.5YR to 5YR. No change in hue was noticed in drab-colored (SY) paleosols compared with similarwaterlogged soils (Simonson 1941;Ruhe 1969). Comparable time scales of burial reddening are evident from paleomagnetic studies of red paleosols of Miocene age from Pakistan (Retallack 1991a).A magnetic component from pigmentary hematite is magnetized in the direction thought to have prevailed during soil formation within several paleosols in a sequence deposited over the time of field reversal temporally constrained to less than a few thousand years (Tauxe & Badgley 1984, 1988).Rubillcationof iron hydroxidesproceedsinafew tens of thousands of years regardless of whether soils are buried or remain exposed. An alternative and unlikely hypothesis is that red paleosols and paleochannels were oxidized by groundwater during burial from an originally drab color (Walker 196 7). Oxidizing groundwaters are uncommon, and largely found within actively recharged, sandy aquifers. Most groundwater is chemically reducing because of oxygen scavenging by microbes and by abundant mafic minerals, such as the schist and hornblende found in the Pakistani paleosols discussed above. These Miocene paleosols also are too clayey to have permitted extensive groundwater flushing. Many of these clayey paleosols are a brilliant red, but interbedded sandstones, includ-
91
ing parent materials to red paleosols,have remained tan and bluegray in color. The opposite coloring of red sandstones within gray claystones would be expected if groundwater oxidation had played a role: and I know of no such occurrences. For these paleosols, as well as for other sequences of interbedded red clayey paleosols and green-gray sandstones (Retallack 1976,1997a;Bown & Kraus 1987). burial reddening by dehydration and recrystallization of iron hydroxides is likely, whereas oxidation of the paleosols during burial is not.
Cementationof primary porosity Transformation of soft friable soil into an indurated paleosol within sedimentary rocks is in part due to the precipitation of cements of calcite, gypsum, hematite and silica. Cementation is also a widespread process in soils at the surface, giving rise to petrocalcic, petrogypsicand petroferric horizons and duripans (of Soil Survey Staff 1998),respectively,for the various cements mentioned above. It may not always be easy to determine whether a particular cement formed during soil formation or burial. A common kind of soil cement is micritic, lowmagnesium calcite, which can be seen to replace pre-existing grains of the soil,so that the grains are deeply embayed (caries texture), and appear to float within a micritic matrix (Fig. 7.4A: Esteban & Klappa 1983). Such zones of replacive micritization in sequences of shallow marine limestones have been interpreted as paleosols and evidence of subaerial exposure (Sear11989: Goldstein 1991). Unlike these replacive fabrics, displacive calcite fills cavities opened by the expansion or rotation of large clods of soil or the cracking out of clods or grains. Displacive fabrics are characteristic of soils, because rotations and expansions of volume are rare under confining pressures of deep burial environments (Braithwaite 1989). Another microfabriccharacteristic of soils and paleosols is a loose aggregate of randomly arranged calcite needles. Void-filling needle-fibercalcite is thought to be produced in soils by fungi (Wright 1986b; Loisy et al. 1999). In contrast, calcitecements formedduring burial or in groundwater tend to be sparry or bladed, cavity-lining crystals (Fig. 7.4B).They also tend to beconcentrated in coarse-grained parts of sandy to gravelly paleosols, or within open cavities such as cracks and root channels.
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Chapter7
Figure 7.4 Pedogenic (A) and groundwater (B) calcitefrornChogoclay paleosols (A) and their nephelinite grit parent material (B) in the middleMioceneFortTernanMember. KerichoPhonolites,at FortTernan,Kenya (scalebars represent 1 mm) (fromRetallack 1991a;withperrnissionfromOxford UniversityPress).
In general, they form solid cemented beds or large nodules with smooth margins, which are much less irregular or complexly mineralized than pedogenic nodules (Retallack 1991b). This account of pedogenic vs. groundwater and burial calcite cements presents useful caricatures, but real situations may be complex. For example, late Miocene paleosolsin Pakistan (Retallack199l a ) include subsurface horizons and diffuse nodules with micritic carbonate replacing clay skins and grains, and filling etch pits in grains. In addition to these pedogenic cements, there also are multiple generations of sparry, pore-filling cements, especially in paleochannel sandstones interbedded with clayey paleosol-rich sequences. Also obvious in paleochannel sandstone are detrital grains of limestone and re-sorted calcareous nodules. Such a complex mix of detrital, pedogenic and groundwater carbonates is common in surface soilsof seasonally dry, riverine plains, such as the Indo-Gangetic alluvium of northern Pakistan and India (Courty & FBdoroff 1985; Courty 1990). Fortunately, there are other techniques that can be brought to bear on this problem. The carbon isotopic compositionof pedogenic and early burial carbonate cement is generally lighter than for calcite cement formed during deep burial or limestone (Searl 1989; Goldstein 1991).Calcareous rhizoconcretions commonly have a sparry fill of the root hole with isotopically lighter carbon than the micritic sheath, presumably as a result of
decay of organic matter in the shallowlyburied paleosol as its porosity was occluded by cement (Driese & Mora 1993).Such burial cements also show relatively low luminescence under cathode rays (Marshall 1988; Searl 1989).This is because, unlike rain and surface waters, groundwaters tend to be chemically reducing and offer more iron than manganese to carbonate cements, and a high ironhanganese ratio within calcite suppresses luminescence. Other carbonate cements include siderite, dolomite and dawsonite. Siderite occurs as nodules and small spherulesinunderclaystopeatysoils(Mooreetal. 1992) andin comparablepaleosols (Browne &Kingston 1983; Leckieetal. 1989: Witzke&Ludvigsen1994).Dolomitic crystals, nodules andlayers may be foundin highly alkaline, dry, salty soils (Botha & Hughes 1992). Siderite, dolomite and dawsonite also are common minerals of burial diagenesis, especially when filling secondary porosity created by acidic fluids associated with kerogen maturation and cracking during deep burial (Surdam & Crossey 1987; Baker et al. 1995). Silica is a common cement in desert soils and playa lakes of high pH conditions that markedly enhance silicasolubility for the short periods when water is available in such environments (Summerfield1983: Chadwick et al. 1989).Such an origin is likely for small silica nodules or for silica pseudomorphs of gypsum roses or other desert salts (Retallack 1983a). Siliceous sinters also form in the immediate vicinity of volcanic hot springs
Alteration of paleosols after burial (Jonesetal. 1998).Voluminouscementsof silicain paleosols more probably form during deep burial, in part as a consequence of dewatering and illitization of smectitic clays (Weaver 1989) and intergranular pressure solution of sandstones (Houseknecht 1990).Some strongly silicified surface horizons of paleosols could have derived much silica from plant opal, which is an easily mobilized form of silica compared with silica in quartz or clay minerals (Retallack 1976; Tarnocai et al. 1991). This remobilization of plant opal must also have been a phenomenon of burial, because such silicified surface horizons, or ganisters as they are widely called (Percival 1986; Gibling & Rust 1992). preserve fossil root traces that were unimpeded by cementation. Hematite, goethite and other iron oxides and hydroxides are common cements in soils, and are particularly abundant in the reddish subsurface horizons of acidic sandy soils (the spodichorizon of Spodosols: Soil Survey Staff 1998). These soil horizons have a characteristic micromorphology of concentric layers of iron-stain around grains and radiating desiccation cracks locally disrupting the stain (de Coninck et al. 1974).The whole horizon may be deflected around tree stumps (a basket podzol)or redeposited within deep burrows or root holes (a krotovina); features also known from paleosols (Retallack 1976. 1985). Arguments have been advanced (byWalker 196 7) for extensive cementation of red beds by oxidation of iron-bearing minerals during burial, presumably by oxidized groundwater. Most of these arguments establish that oxidation was in place without addressing timing, which could be either during soil formation or during burial. For red clayey paleosols of my experience,burial reddening of ferric hydroxides is likely, but not burial oxidation, because of associated paleochannel sandstones remaining unoxidized, among otherreasons (Retallack 1991b). Gypsum is a common cement in desert soils (Porta & Herrero 1990; Chen 1997) and evaporite sediments (Warren 1989).Many so-called ‘deposits’of playalakes and coastal sabkhas form by growth of gypsum and other salts in place within pre-existing sediments.This is the hallmark of soil formation (as in Salids and Gypsids of Soil Survey Staff 1998), rather than the layer-bylayer accumulation of a precipitate that would be better regarded as a chemical sediment. Dissolution of salts on burial appears to have been more widespread than their precipitation. Many evaporite beds, salty paleosols, and
93
salt crystals in paleosols have been dissolved into brecciated layers (Bowles & Braddock 1963) or replaced by other minerals during burial (Barleyet aI. 1979).
Compaction by overburden Burial of paleosols results in compaction as the void spaces, fossils and pore water are crushed by the weight of overburden. The compaction of originally loose soil clods against one another creates a complex pattern of slickensides superficially similar to the more regular slickensides of swelling-clay soils of seasonally dry climates (Gray&Nickelsen1989).Changesin thicknessas a result of compaction can also compromise interpretation of former rainfall from depth within a paleosol to the horizon of calcareous nodules (Retallack 1994b), and reconstruction of former physicochemical weathering (Retallack 1986a). The best indicators for degree of compaction suffered by paleosols are clastic dikes: cracks in the soil that were more or less vertical and subsequently were filled with contrasting material. With compaction, clastic dikes are deformed ptygmatically, so that the amount of compaction can be calculated by the ratio of the vertical height of a clastic dike to the distance around undulations of the dike. Unfortunately, such features are found in few soils and paleosols,for example, the swelling-clay soils or Vertisols and their buried equivalents (Retallack 1986a;Caudillet d.1997). Other features useful for estimating compaction include fossil skulls, turtle shells, snails, logs, burrows and the spacing of uncompacted laminae within nodules compared with laminae outside the nodules. Fossil logs remain as wide as their former diameter, but are commonly crushed during burial to a thin band (Walton 1936; Briggs & Williams 198 1).Unfortunately, these various kinds of fossils can collapse before burial, because of decay in and on a soil. There are similar problems for calculating the compaction of paleosols from their bulk density,usually estimated by the clod method (Klute 1994). compared with the bulk density of comparable surface soils. Bulk density varies considerably within modern surface soils. Another approach is to point count the number of planar, concavo-convex and suturedcontactsbetween grains in sandstones, because sutured contacts become more abundant at greater depths (Fig. 7.5). This kind of analysis is restricted to a
94
Chapter7
70~
7 5
tangential conmc1s
60-
8c 50-
long conlacls
40-
0
3020-
- -;-.:.
a \*-=-=- a\, .’
concouoconuer cantocts
IO-
sutured conlacts
500
1000
’\
1500 DEPTH (ml
1
’
‘ <’;.-,
1
2000
3ooo
2500
Figure 7.5 Relationshipbetweendepthof burial and types of grain contact in Mesozoic sandstones of Wyoming,USA (modifiedfrom Taylor 1950,AAPGO 1950:withpermission from the American Association of Petroleum Geologists whose permission is required for further use).
0
Table 7.1 Burial compaction equation for paleosols.
5
C=-S/{[F/e(D’g]- 1) where: Cis compaction (as a fraction) Dis depth of burial (-km) Fis initial porosity (per cent pores) Kis acurve-fitting constant S is solidity (per cent solid grains or reciprocal of porosity)
10
15 20 __
--
+Retallack (1994b)
*
Caudill eta/. (1 997) --+- Sheldon & Retallack (2000)
0.5 i 1.5 compaction (as a percentage of original thickness)
0
Figure 7.6 Compactionof Vertisols as a fraction of original thicknesswith increasing depth of burial (km),using the Compaction formula recommended here (filledtriangles, Table 7.1) comparedwith less satisfactoryformulae of Retallack(1994b)and Caudill et nl. (199 7) (Sheldon& Retallack2001: with permission from the Geological Society of America).
with the following material-dependent values: S
F
K
Sediment
Soil
0.69 0.62 0.5 1
0.31 0.38 0.49
0.12 0.17 0.27
Hardclay Moistclay Sandyclay
0.37 0.30 0.06
0.63 0.70 0.94
0.52 0.71 2.09
Sand Vitricash Peat
Vertisol, Oxisol Alfisol, Ultisol Entisol,Inceptisol, Aridisol,Mollisol Spodosol,Gelisol Andisol Histosol
Sources: Nadon & Issler (199 7); Sheldon & Retallack (2001).
particular mineral, and quartz is best, because different kinds of grains deform differently under compaction (Retallack 1994c;Stone & Seiver 1996). For most paleosols, compaction is best estimated by comparison with standard compaction curves (Fig. 7.6) now available from the statistical study of large numbers of boreholes in sandstones, claystones and coals
(Zimmerman 1990; Sheldon & Retallack 2001). These curves vary with differences in physical properties of soil materials (Table 7.1). Quartz sandstone is less readily compacted than lithic sandstone, calcareous or pyritic shale less compacted than carbonaceous shale, and
Alteration of paleosols after burial peat formed from wood (coal rich in vitrinite) less compacted thanpeatformedfrommossesor leaves (coalrich in exinite).In peaty soils,there can be considerable compaction at the base of the organic horizon, even before burial (Elliot 1985; Nadon 1998). Similarly, dry footworn soils may have been compacted before burial to a greater extent than marine sediments and lowlandsoils, which formed most of the rock sequences used to construct compaction curves. One way of using these curves is to measure theporosity of the sample, or preferably the solidity or per cent solidgrains (Baldwin&Butler 1985),and use this figure as an index of compaction to establish a depth of burial. Considering the magnitude of likely errors and the way compaction curves steepen with depth, this approach is of limited usefulness. More consistent results can be gained by estimating the depth of burial by geological methods, and then estimating likely compaction from a curve or equation using an appropriate value for solidity, porosity and a curve-fitting constant (Table 7.1).
Tertiary
lllitization of smectite Many clayey paleosols are predominantly illitic in composition, and in many soils illite or interlayered illite-smectite is the main clay mineral, especially young soils of desert regions with strong wet-dry seasonality (Robinson & Wright 198 7). Considering the probably tow demand for potassium of microbes and early land plants, unlike that of forests, illite-rich soils could have been much more widespread in the distant geological past than they are today (Fig. 7.7). Although these explanations have some appeal, many illitic paleosols lack evidence of sodium or calcium enrichment within salty or carbonate-rich horizons characteristic of desert soils (Retallack199 7d; Retallack&Alonso-Zarza1998).Nor do they show the cracking and undulation of the surface (mukkara structure of Paton 1974) found in soils of wet-dry seasonal climates (Retallack & Mindszenty 1994). Many paleosols of Mesozoic and Paleozoic age with copious large root traces of the kind formed
10
20
percentage of clay fraction 30 40 50 60 70 80 SMECTITE
Cretaceous
Jurassic Triassic Permian
-
Carboniferous Devonian Silurian Ordovician
Cambrian
-
Figure 7.7 Relativeabundance of clay mineralsin shales throught i e based on 40000analysesfromNorthAmerica(from Weaver 196 7;reprintedwith permissionfrom Springer-Verlag).
95
ILLITE
90
96
Chapter7
under forests are predominantly illitic (Retallack 1985, 1997d,e). An additional explanation for the illitic composition of many paleosols is the alteration of smectite to illite during deep burial. The overall chemical effects of this alteration aredisplacement byK+ofCa2+,Mg2+andNa+, which are lost with water of hydration, alumina and silica (Land et al. 1997). This has been regarded as a solid-state replacement of interlayer cations with preservationin illite of most of the 2 :1layer structureof precursor smectite. However, X-ray diffraction and transmission electron microscopy studies show that it proceeds by destruction of small crystals of smectite by interstitial fluids, followed by the neoformation and growth of increasingly large crystals of illite, a process called Ostwald ripening (Eberl et al. 1990; Lanson & Champion 1991).The source of added potassium may have been minerals such as microcline, as shown by loss of microcline and concomitant illitization in deep boreholes (Hower et al. 1976). Microcline and other potassium-rich minerals may have been hydrolyzed by solutions rich in carbonic acid generated during kerogen maturation and cracking (Hover et al. 1996), but a case also has been made for theillitization of paleosolsby groundwater containing potassium leached from surface soils (Nesbitt&Young 1989).Pervasive localpotassium enrichment also occurs around hydrothermally altered veins (Worden & Rushton 1992) and certain kindsof igneousintrusion (LeBas 1977). Unlike local hydrothermal and igneous potassium enrichment, diagenetic illitization is limited in its thoroughness. The ionic exchange between microclinerich sandstones and smectitic claystones may be restricted to distances of only a few tens of centimeters: gradients of illitization have been observed away from a likely source bed (Altaner et al. 1984). Shaly paleosols can remain unillitized under pure quartz sandstone (Kimberley & Grandstaff 1986). Furthermore, some classical studies of illitization have been shown to reflect changing provenance of sediments rather than burial alteration (Bloch et al. 1998). Isotopic studies of illite and calcite in Pennsylvanian clayey paleosols indicate surprisinglylimitedillitization(Moraetal. 1998).Inany case, burial illitization seldom goes to completion. Remnant smectiteiscommon(Howeretal.1976).Illitization may proceed during a specific window of burial conditions and then slow or stop, as indicated by radiometric dating of shales in boreholes (Morton 1985). The
process may be retarded by dewatering during burial and scarcity of readily altered potassium-bearing minerals. Illitiiation of shales becomesreadily discernible when shales are buried at depths between 1.2 and 2.3 km and reach burial temperatures of 55-100°C (Weaver 1989;Sarwar&Friedman 1995).Theoretically,however, the transformation of smectite to illite could occur at much lower temperatures over very long periods of geological time: for example, Ordovician shales that are illitiized but only shallowly buried (Bethke & Altaner 1986). One way to demonstrate illitization is through bulk chemical analysis of paleosols. Illitization should be suspected if there is a surficial increase in amounts of potash, without a concomitant surficialincrease in soda or local enrichments in lime found in desert soils (Retallack 1986a). This anomalous combination of chemical depth functions can be displayed graphically in a triangular plot of carbonate-free recalculatedmolar proportions of lime and soda, vs. alumina, vs. potash (Fig. 7.8). Unillitiizedprofiles onsuchaplotshouldfallon a line heading toward the alumina corner with increased weathering or proximity to the surface of the
gibbsite
// 1
hornblende clinopyroxene
No20 t nan-carbonate CaO
K20
Figure 7.8 Chemical analyticaldatafor a surface soil on the ToorongoGranodiorite,Mt Baw Baw, Victoria,Australia (Elled circles) and for a pre-Huronian(2450Ma)paleosol on the Vale Marie Granite near LakeTimiskaming,Quebec,Canada (6lled squares)plotted with the compositionof selected minerals (opentriangles)on an A-CN-Kdiagram (from Rainbird et a!. 199l,lournn!o~Geolog~98,Fig. 12A;withpermissionfrom ChicagoUniversityPress).
Alteration of paleosols after burial profile, whereas illitied parts of the profile will deviate toward the potash corner (Nesbitt 1992). Another approach is to use chemical analyses to calculate the potash-alumina stoichiometry of the clay fraction, which can in some cases approach the composition of muscovite rather than the usual composition of illites foundin soils(Feakes&Retallack1988).Yetanotherapproach is to calculate the normative mineralogical composition of a paleosol from chemical analyses, using minerals judged likely to have been in the paleosol from petrographic and other observations (Garrels & MacKenzie 1971). This can be carried out using a spreadsheet or specially designed computer programs (Reynolds 198 5). For illitized paleosols it may be necessary to use illite compositions unrealistically close to those of muscovite to solve the equations (Retallack 19 86a). Yet another approach is to evaluate illite crystallinity from the sharpness of the l O A peak on X-ray diffractometer traces. Illites formed during burial have sharp narrow peaks, whereas those formed in soils have broad low peaks. Anumber of indicesof peak sharpness can be used (Frey 1987): for example, the Weaver index is the peak height above background at l O A divided by the peak height at 10.5 A and the Kubler index is the width in degrees 20 of the l O A peak at half the level of its height. The onset of illitization is marked by an increased proportion (more than 4 : 1)of illite interlayers within smectites, but only a slight improvement of illite crystallinity.Illitization is very marked, and little smectite remains, at the transition from diagenesis to metamorphism, which corresponds to about 5-7 km of burial or about 200°C in burial temperature, and to a Weaver index in excess of 2.3 or a Kubler index less than 0.42OA29.
Zeolitization and celadonitizationof volcanic rocks Many paleosols in volcaniclastic successions are now rich in zeolites such as clinoptilolite and in the distiinctive lime-green, ferrous, illite-like clay, celadonite. Zeolites are common in hot springs, cooling volcanic rocks and alkaline lakes (Bargar 1994).but not in soils,where they are readily destroyed by weathering (Jacob 81Allen 199 3). Zeolitization and celadonitiiation of volcaniclasticrocks has been regarded as a process of burialdiagensis akin to illitization,in which pore fluids introduce
97
alkali cations at depths of 400-1200m and temperatures of 2 7-5 5°C (Hay 1963; Lander & Hay 1993). A problem with such models is the mass transfer required to convert volcanic rocks to zeolites, as delicate volcanic bubbles and shards indicate little mass transfer or compaction (Moncure et al. 1981). An alternative idea is that zeolitization and celadonitization represent crystallization by Ostwald ripening during deep burial of amorphous weathering phases such as imogolite common in volcanic soils such as Andisols (Retallacket al. 2000). By this view, some soils (Andisols) are predisposed to zeolitization and celadonitiiation during burial, whereas others (AHsols, Vertisols, Oxisols) are not.
Coalificationof peat Very few peats accumulate from rafted plant debris (Wagner & Pfefferkorn 1997). Almost all thick coal seams are formed by alteration during burial of peats that accumulated under swamp, marsh, or other kinds of vegetation that can tolerate waterlogged ground. For soil science, peats formed in place are called histic epipedons and are the most characteristic feature of a kind of soil called a Histosol (of Soil Survey Staff 1998). In a sense then, most coal seams are paleosols, although their original profileform is altered considerablyby compaction and chemical alteration of the peat during deep burial. The chemical alteration of peat during burial is primarily an enrichment in carbon with the loss of water, nitrogen and other volatile materials. As a result, a soft, dull, brown peat is transformed into a brittle, microfractured, shiny, black coal (Diessel 1992). A commonly used chemical index of coaliication is total volatile matter (% VM). A common physical index is the reflectance of a particular component of coal, the macera1 vitrinite. which is derived from wood fragments (%VRO).These indices, together with the geological age of the coal, can be used to infer the maximum paleotemperature experienced by the coal (Fig. 7.9). Estimated paleotemperatures can be useful for reconstructing geothermal gradients related to burial depth, or to hydrothermal or igneous activity. Coaliication also affectsorganic materials other than peats. The browning and then blackening of normally white to gray conodonts with increased depth of burial is useful for estimating depth of burial of marine sediments (Epstein et al. 1977). Two techniques used
98
Chapter7
0.001
2
- Scale
0.m 0.01 002
0.05
0.1
02
0.5
1
in this manner in nonmarine rocks are blackening of fossil pollen and spores (Batten 1996; Marshall & Yule 1999) and the increased reflectance of isolated pieces of organic matter with increased burial (Diessel 1992). These techniques require local calibration for specific sedimentary basins, but the information they provide on burial conditions can be invaluable in sorting out the effects of other alterations after burial.
Kerogenmaturationand cracking A widespread chemical reaction during soil formation is hydrolysis of aluminosilicate minerals to clay and cations in solution. Basic cations are displacedby hydronium in weak solutions of carbonic acid, created principally by respiration of carbon dioxide by soil organisms. Similardestruction of grains also can occur during deep burial by acidic solutions generated during decarboxylation of buried organic matter (kerogen), its maturation to oil, andultimate cracking to natural gas (Surdam & Crossey 198 7). Secondary and primary porosity can appear similar under the microscope, with deeply etched grains or hollow clay skins remaining from destruction of the grain within. Characteristic of secondary porosity are oversizedvugs, with pronounced elongation and alignment, and cutting sharply across other grain boundaries (Schmidt&McDonald19 79).Thesekindsof cavities are not common in secondary porosity,but are very distinct
2
5
10
Figure 7.9 Karweil's (19 75) nomogram for estimatingpaleotemperatureduring burial from vitrinite reflectance (% Rmaxj or volatile matter (% VM on a dry ash-freebasis)and geological age of coals, showing recalibration of Bostick(1971:fromDiessel1992:with permission from Springer-Verlagj.
from primary pores in soils, which do not tend to cut sharply across grain boundaries and are not so well aligned. Soil pores also tend to be altered around the margins with clay skins or oxidation rinds. In most paleosols, primary pores are crushed during burial compaction, so that abundant inflated vugs may indicate secondary porosity. Nevertheless, some root holes and vugs in soilsfortiied by cements during soil formation or early after burial can withstand compaction. Secondary pores also may be distinguished from primary pores by cementation with minerals stable under conditions of Eh and pH that do not make sense for the host paleosol: for example, siderite cement in a red, hematite-rich paleosol. Cathodoluminescence and isotopic studies of cements may be useful for distinguishing between the primary pores of a n ancient soil andporosity imposed during deep burial (Marshall 1988; Sear1 1989). For some paleosols, the influence of decarboxylation and kerogen maturation and cracking can be ruled out by evidence of shallow burial, such as low vitrinite reflectance or poor illite crystallinity. Abundant oil, wet gas and organic acid are generated from organic source rocks at burial temperatures beyond those needed to create vitrinite reflectance greater than 0.5%, at which pointcoalsreach bituminousgrade(Tab1e 7.2). Kerogen cracking to dry natural gas and carbon dioxide is greatest when vitrinite reflectance exceeds 2%. as in higher rank bituminous coal and anthracite (Tissot & Welte 1984).This 'oil window' is typically at burial tempera-
Alteration of paleosols after burial
99
Table 7.2 Correlationof thermal maturation indices for paleosols and sediments (fromBatten1996). U t e crystallinity Burial temperature ("C)
30-65
80
120-170
170-180
Coalrank Peat Brown coal Lignite SubBituminous HighVolatile Bituminous
MediumVolatile Bituminous LowVolatile Bituminous Semi-anthracite Anthracite
f200
Vitrinite reflectanceR,
Palynomorphcolor
0.2
Pale yellow
0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0 1.2
Yellow Orange Pale brown Light brown
Weaver
Kubler
Hydrocarbongeneration
Dry gas
1.0
1 .0
Wet gas Medium to light oil
Dark brown 2.0
0.60
Condensates
1.5 Very dark brown
2.0 3.0 4.0
Drygas+H,S,CO, Bitumen pluggingporosity Black (opaque)
Graphite
2.3
0.42
Poor porosity
Note: crystallinity is based on height of X-ray diffraction l0A peak (H), with Weaver index as H,,,,,,/H,,,,,,Kubler index as width of Hat half height in "28.
tures of 50-1 50"C, or depths of 1-5 km,depending on local geofogical age, tectonic history and geothermal gradient. Assessing whether a particular paleosol has been affected by fluids generated during kerogen maturation and cracking can be as complex as finding oil, and relies on the same kinds of information.
Neomorphismof carbonate Many paleosols are strongly calcareous, and calcite is prone to neomorphism to coarse grain size during burial (Bathurst 1975). Neomorphic calcite may be distinctive in thin section, with its equiangular crystal junctions, cross-cutting relationships or other features (Table 7.3). Neomorphism can be especially obvious when it affects fossil shells, such as those of snails, as it converts their original fibrous or prismatic microstructure to a
sparry mosaic. Similarly,it also has been assumed that calcareous nodules inPaleozoic pdeosols with scattered replaced and embayed grains were recrystallized to sparry calcite from the more common micritic texture of soil nodules (Retallack 1985). Isotopic and cathodoluminescence studies are now being applied to the study of such nodules (Sear1 1989). Neomorphism also explains why pore-fillingneedle-fibercalcite common in aridland soils and Quaternary paleosols (Courty 1990; Guo & Fedoroff 1990). has been found in so few pre-Quaternary paleosols (Wright 1986b: Alonso-Zarza et al. 1998a). Evaluation of burial conditions using such indices as vitrinite reflectance or illite crystallinity also may be useful. Recrystallizationis generally not extensiveunder conditions of burial more shallow or cool than those required for extensive illitization, or coali6cation beyond bituminous grade.
100
Chapter 7
Table 7.3 Distinction between cementationand neomorphism. Cementation
Neomorphism
Detrital grains touchingin places Sharpcontactsbetween grains andcrystals No crystals cuttingacrossboundaries of detrital grains
Detrital grains with abnormallyloosepacking Wavy or 'nibbled and irregular contactbetween grains and crystals Some crystalscuttingacross boundaries of detrital grains or other original features Crystalswith inclusionsor relict structures Crystalsnot showingregular relationship to grainor void boundaries
Crystals lacking inclusionsor relict structures Crystals forminglayersconcentricto surfaceof originalcavity Contacts of crystal faces at odd and different angles
Contactsbetween crystal facestendmg to form equal angles
Note: adaptedfrom Folk (1965).
Metamorphism The various kinds of alteration already outlined pass into metamorphism beyond the point of extensive illitization, kerogen cracking to natural gas and coalification to anthracite grade. This transition is passed at temperatures somewhere in excess of 200°C or depths of 7 km, whichever comes first (Fig. 7.1). A precise boundary is difficult to specify In practice, it is best to label a rock metamorphic if it has newly formed structures such as schistosity,or minerals such as clinozoisite. Many metamorphosed paleosols have now been reported. Especially common minerals in metamorphosed paleosols are corundum, andalusite and sillimanite, which are to be expected in metamorphic rocks derived from such aluminous protoliths (Grandstaff et al. 1986; Dash et a2. 1987; Golani 1989; Luc et al. 1989; Krois et al. 1990). Illitized clays in paleosols may be altered during metamorphism to sericite, chlorite and muscovite (Retallack& Krinsley 199 3 ) . Even with extensive metamorphism it is surprising how well the chemical signature of soil formation survives. Chemical depth functions similar to those of soils in oxides other than potash have been widely reported in paleosols metamorphosed to greenschist facies (Grandstaff etal. 1986;Melcher 1991;Mossman&Farrow 1992)and in some cases to amphibolite facies (Barrientos & Selverstone 1987). Hematite and red color is preserved in paleosols metamorphosed to greenschist facies (Retallack 1985), and is not converted to gray magnetite until amphibolite facies (Thompson 1972). Textural and chemical alteration as a result of meta-
morphismis in most cases easily distinguished from that caused by soil formation. However, it can be difficult to distinguish highly metamorphosed paleosols from hydrothermally alteredrocks (Hollandet al. 1989).
Common patternsof alteration The great array of potential alterations during burial that could compromise interpretation of paleosols can be intimidating, but as research proceeds, common patterns of alteration are emerging. Burial decomposition, burial gleization, burial reddening and cementation can significantly alter the appearance of paleosols very shortly after burial. This is not to deny the importance of alteration during deep burial, such as illitization. coalication, kerogen maturation and recrystallization. Many physical features of paleosols such as horizonation and root traces persist through these deep burial alterations well into greenschist facies of metamorphism. It is not until the amphibolite metamorphic facies that physical structure is completely obscured by recrystallization. The effects of burial on soils can be summarized by considering a common pattern of alteration (Fig. 7.10) of weakly to strongly developed soils under forest or woodland vegetation in well-drained parts of lowland landscapes and subhumid to humid, nonfrigid climates. The original soil may be clayey,brown and friable,with a surface (A horizon) dark gray with decaying organic matter over a brown subsurface (Bt horizon) enriched in clay and penetrated by stout root traces. The effects of burial decomposition, burial gley and burial reddening may convert this to a strange caricature of a soil: a clay-
Alteration of paleosols after burial
A
Holocene
B
Eocene
c
101
Devonian
Figure 7.10 Hypotheticalalteration of a woodland soil by burial gleization of organic matter,burial reddening of iron hydroxide and oxide minerals, compactionand illitiation of smectite.Theprofilesillustrated(all from USA) are (A)a surface soil in Iowa, (B) early Eocene (52 Ma) paleosols in the type area of the Sand Creek Facies of the Willwood Formation of Wyoming, and (C) Late Devonian (365 Ma) PeasEddyclaypaleosolofNewYork (fromRetallack1991a:withpermissionfrom AnnualReviewsInc.).
stone bed with very low amounts of organic carbon, a former surface horizon of green-gray claystone that penetrates as haloes around root traces into a brick-red clayey subsurface horizon. A subtly colored gray to brown soil has thus been converted to a green-red mottled bed as gaudy as a candy cane. There are many such paleosols (Retallack 1983a, 1991a, 1997a,d).Withadditional burial and metamorphismto lower greenschist facies, such a profile may darken so that its surface horizon is dark grayish green and its subsurface purple, its clays may coarsen to silt size with the development of illite and chlorite, and its irregularly orientated soil structures become overprinted by schistosity. Paleosols of this kindalso areknown (Retallack 1985). Alteration of paleosols during burial is of interest in itself, because soils are a distinctive class of materials that behave somewhat differently from sediments dur-
ing burial. Untangling the alteration of paleosols after burial fromthealterationcaused by soilformationisalso an important prerequisite for interpreting paleoenvironments from paleosols. The environmental significance of the amount of organic carbon, of oxidation state and of clay mineral composition of soils is well established. Such studies of surface soils must be used with care to interpret paleoenvironments from paleosols, because these features of soils are prone to alteration after burial. Fortunately,there remains an impressive residue of soil features in paleosols that can be used for environmental reconstruction. Physical features such as nodules, root traces and burrows are particularly robust in the face of alteration after burial. These deserve scrutinyto determinethe extent to which paleosols as natural essays on landscapes of the past have been smudged and torn by alteration after burial.
Part 2
Factors in soil formation
Holocene soils and paleosols (foreground) and Oligocenepaleosols (background ridge), BadlandsNational Park, SouthDakota,USA.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 8 Models of soil formation
Different definitions of soils reflect different approaches to them. Soils present a different challenge to farmers and to engineers. Soil scientists have a different view again, or rather several differentviews of soils (Johnson & Watson-Stegner 1987). These varied views of soils are, in effect, different theoretical bases for their study. So far, in describing various features and kinds of soils, they have been considered as if they were natural history specimens. Like the trilobites, cowries or narwhal tusks in a gentleman’sspecimen cabinet from Victorian England, the variety of soils found in nature can be described and classified. Fundamental aspects of soil formation can emergefrom such comparativestudies.The quest for a ‘natural classification’,for example, aims to reveal distinct processes and products of soil formation. Mapping soils as objects of natural history also is useful for guidingeffectivehuman use of the landscape.Incontrast, soils also can be viewed as open systems,as energy transformersand as environmentalproducts (Fig. 8.1). Each of these different views of soils, as detailed below, elicits different kinds of researchquestions,experiments and observations. Soils can be envisaged as open systems to the extent that they represent a boundary between earth and air through which materials move and are changed. Four basic kinds of fluxes can be imagined: additions, subtractions, transfers and transformations (Simonson 1978; Anderson 1988). Additions include mineral grains brought in with airborne dust. Organic matter also is added to the soil in the form of leaf litter. Subtractions include surface erosion of mineral and organic matter. Mineral matter is lost in solution in groundwater, and organic matter is lost through cropping by organisms, which then migrate elsewhere. Transfers include movements of material within a soil profile. An example is the downward leaching of clay to produce a clayey subsurface (Bt) horizon in Alfisols and Ultisols. Relocation of soil material in burrows is another example. Transformations involve changes in composition
and form of soil materials. Organic matter, for example, is recycled through a variety of forms ranging from decayed plant material to parts of living soil organisms. The formation of clays from easily weathered minerals such as plagioclase is another example.Flux models aim to quantifymovementsof material within and around a soil, such as the rate of calcium carbonate accumulationindesertsoils (Machette 1985). Soils also can be viewed as energy transformers:that istosay,asabodyof materialchangedby thecontinuing efforts of natural processes (Runge 1973).The pattern of energy flow through soils can be as complex as the flowof material.The primary source of energy in soils is heat from the Sun, which far surpasses (by about 7500 times)heat flow from deep within the crust generated by decay of radioactive elements. Energy also can be gained, in a sense, from deposition of sediment on the soil, or from the addition of groundwater or rain, because these new mineralsand water are capableof altering preexistingmaterials, or doing work within the soil. Less apparent, but no less important, is energy gained by exothermic mineral alterations, transformation of organic matter, biological growth, friction, wetting and thawing of soil. Solar energy is lost by radiation, reflection and evaporation. Other energy sinks include erosion,drainage,drying or freezing. Energy also can be transformed within a profile, by conduction, convection, condensation, evaporation, percolation and chemical reaction. In addition to energy flow within a soil, there is also impedance of energy flow, such as a rate-limiting factor for a chemicalreaction or restricted groundwater movement because of low permeabilityof the soil. All of these processes and the amount of energy available to drive them are what create a soil profile. For example,a mathematical model for the development of calcic horizons in desert soils can be based on fluxes of materials, together with their rate of chemical reaction under different physical conditions (McFadden et al. 1991). 105
106
Chapter8
/?
of work done
TS,
of climate
of subtractions MATERIALS FLUX MODELS
= f(A,Te,To,S)‘
done to o parent moterial ENERGY PROCESS MODELS
+
.~(w,o)~
323
of topogrophic relief ENVIRONMENTAL FACTOR MODELS S = f(Cl,O,R,P,T)
Figure 8.1 Three mathematicalmodelsof soil formation.
Soilsalsocan be viewed as an environmentalproduct, molded over time from whatever materialwas available, by climate,organismsand geomorphological processes. It would be ideal if there was a one-to-one correspondence between features of soils, such as clay content, and a particular environmentalfactor,such as rainfall. Unfortunately, clay content, l i e many other features of soils, also is dependent on the temperature of the soil, the nature of minerals available to be weathered and the time available for weathering. Soil clay content is thus less l i e an automated rain gauge than a poorly calibrated ‘synthetograph’,which averages out anumber of overlapping modifications (Phillips 1993 ; Huggett 1998). The aim of environmental factor models is to tease apart these separate influences on the soil. The multitude of specificinfluenceson soil formation can be reduced to five main factors: climate, organisms, topographic relief, parent material and t i e (Jenny 1941; Buol et al. 1997). These classical five factors are most easily remembered by the acronym ‘CLORPT’.This is a useful set of categories for mentally considering all aspects of the formation of a soil or paleosol in the field.
More importantly, CLORPT provides a theoretical framework for devising natural experiments to investigate processes of soil formation. Because soil formation is a multivariateprocess, study of any one of the factors in isolation is necessarily limited just to those cases where all the other factors are constant, or at least nearly so. To study climate,for example,what is needed is a number of soils formed in different climates, but as similar as possible in their ecosystem, topographic setting, parent material and period over which the soil formed. A well-known example is the mathematical formula relating depth to calcic horizon (D, in -cm) and mean annual precipitation (P,in mm) in grasslandsoils developed on postglacial loess of the midcontinental USA (Jenny 1941; Retallack 1994b, 2000; Royer 1999):P=418.5 - 2.3 3 5D.Thegroupof soils used to establish this relationship is called a climosequence,and mathematical relationships relating soil features to climate are called climofunctions.Biofunctions,topofunctions, lithofunctions and chronofunctionsare useful for quantifying the effects of organisms, topography, parent material and t i e available for soil formation. The
Models of soil formation approach assumes that these factors are independent, but it may be difficult to find situations where vegetation, for example, is uniform over regions of different climate. It is questionablealso whether time can be considered an active factor in the same way as rainfall or root growth. Nevertheless, each of the factors does have a recognizable effect, and with such a strategy of dividing to conquer, it is possible to make some sense out of the complex multivariate process that is soil formation (Yaalon 1975;Buolet d.199 7). Each of these approachesto the studyof modern soils provides information valuable for the interpretation of paleosols, even though it is not possible to study paleosols in quite the same way because the fluxes of materials, flow of energy and environmental factors
107
controlling them ceased to act long ago. From materials flux models comes information about rates of development of soil features. From energy process models comes an understanding of the relationships between variables in soil formation. From factor function models comes information about the role of selected environmental influencesin soil formation.More and better information about the past will be gleaned from paleosols as such studies of Quaternary soils gain in scope and precision.Althougheach approachisof value, thisbook emphasizesthe environmentalfactor approach because there is a tremendous amount of information about soil formation already available in this form and because the five factors offer a systematic framework for paleoenvironmental interpretation of paleosols.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 9 Climate
During the founding period of soil science, climate was considered one of the most important factors in soil formation. There are widespread acidic sandy Spodosols in temperate regions, but in the tropics red clayey Oxisols are common. Such observations encouraged the concept of zonal soils (Glinka 1927).These are soil types restricted to a particular climate. Other soils, for example Entisols, were azonal because of limiting local circumstances, such as a short time of development, so that climatic effects were not sufficiently marked. Some generalizations of the zonal concept of soils are still regarded as valid. However, a more detailed understanding of soil geography has revealed further exceptions. The tendency to incorporate climatic information within soil classifications has been countered by efforts to base soil classification on observablefeatures of soils. In the US taxonomy, units as broad as the order Aridisol were defined by climatic criteria (Soil Survey Staff 1975).This was unfortunate, because paleoclimaticinterpretations cannot be made from paleosols identified using paleoclimatic data. Fortunately, Aridisols have been redefined (Soil Survey Staff 1998) as soils with a calcic horizon at a depth of less than 1m. Paleosols classified using such criteria may be compared with modern examples to gain some idea of paleocliiate. This method of paleoclimatic inference is rather l i i using a particular species of fossil leaf to infer a paleoclimate similarto that of a related living species.Another way of interpreting paleoclimate is to examine the size, shape and marginal outline of fossil leavesas physiological indicators of paleoclimate,independent of the leaf species involved (Wolfe et al. 1998). It is this kind of interpretation of paleosols and features of paleosols, independent of their identiication, that will be attempted in this chapter. The terms climate and weather sometimes seem interchangeable in common usage but have distinct scientific meanings. Weather is the record of rainfall, temperature and humidity as reported daily in news108
papers and on television. Climate, on the other hand, is the average of data compiled from weather reports, usually summarizingrecords from at least 30 years of observations. Climatic data are based on a particular weather station, but these stations are generally chosen to reflect conditions of the surrounding region (Muller 1982). Local frost hollows and exposed high ridges are to be avoided for regional weather stations because they have their own microclimate. These local climatic deviations, such as that within a hollow tree or behind a rock, are significantly different from regional climate and can be especiallyimportant for small animals and plants. Soil climate is a specialkind of microclimate. It refers to the conditions of moisture, temperature and other climatic indices within soil pores. In well-drained soils, the soil climate is a somewhat muted version of regional climate, with extremes of variation damped out. In waterlogged soils, however, soil climate is unrelated to regional climate. The temperature and oxygenation of waterlogged soils depend more on local rates and pathways of groundwater flow than on atmospheric conditions. Waterlogged soils can be found in desert oases as wellasinwetforests(Lottes&Ziegler1994).Suchdifferences between soil climate and regional climate can be a source of confusion for interpreting paleosols. For example, the transition from Carboniferous coal measures, with their drab paleosols,to Permian red beds (290Ma), with their variegated red and orange paleosols, reflects improved drainage of paleosols as well as regional climatic drying (Retallack 1995a; West et al. 1997).Estimates of other kinds of soil climate are now finding their way into soil classifications,models for soil formation and studies of soil biology In few cases, however, have observations of soil conditions continued for periods of 30 years or more, so that in many instances the available data are really an account of soil weather. Because of this and because the conditions of human life and agricultural production are more easily related to regional climate, the climate above ground is most
Climate used in studies of the relationship between soil features and climate. It is this growing data base of climosequences of soils and climofunctions of soil features that forms the primary subject matter of this chapter.
Classificationof climate Some climatically sensitive features of soils can be related clearly to particular climatic variables, but even the best of these lack precision. Soils are not as sensitive for recording climatic conditions as meteorological instruments, but climate can be interpreted from them within fairly broad categories.Suitablybroad categories are provided by a number of classifications of climate, but only a few of these are reviewed here.
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One of the most influential large-scale classifications of climate was devised in 1918 and modified over the next two decades by the German meteorologist Vladimk Koppen (Trewartha 1982). He recognized five main groups of climates, corresponding to the main kinds of terrestrial vegetation. Each kind of climate is designated by letters, as a kind of meteorological shorthand. Uppercase letters are used for the main climatic groups and lower-case letters for subsidiary climatic features (Table 9.1). One large climatic group includes dry climates. Humid climates are divided into climatic groups based on temperature. On a global basis the main categories of climate formdistinct latitudinal belts (Fig. 9.I). Subdivisions of the main climatic groups are based on a variety of features, mainly temperature in the case of dry and
Table9.1 Koppen's classificationof climates (after Trewartha 1982).
A Af Aw
Am B BW
BS
C
cw cs Cf
Tropical wet: hot rainforest;coldestmonth >18"C (64'F) Tropical ever-wet: driest month >60 mm, nonseasonal Tropical wet-and-dry: distinct dry season in low sun period or winter: at least 1month with <60mmprecipitation Monsoon climate: short dry season, but total rainfall keeps ground moist year-round to support forest Dry; excess of evaporation over precipitation;permanent streams do not originate in B climates Arid or desert climates (fromGerman Wiiste for desert) BWh mean temperature over 18OC(64°F) (Germanheiss) BWk meantemperatureunder 1 8 T (64'F) (German kalt) Semiaridor steppe climate (from Steppe or dry grassland) BSh mean temperature over 18°C (64°F) (Germanheiss) BSk mean temperature under 18°C (64°F) (Germankalt) Warm temperate to subtropical:temperate forest:warmest month mean >1O"C (SOT) and coldest month mean <18"C (64°F) but>-3"C (26"F)fdefinesequatorwardlimitof frozengroundand snowfor>l month) Winter dry:at least 10t i e s as much rain fallsin the wettest month of summer as in the driest month of winter: alternative is 70%or more of mean annual rainfall in summer 6 months Summer dry; at least three times as muchrain falls in the wettest month of winter as in the driest month of summer, with <30 mm (1.2 inches): alternative is 70%or more of mean annual rainfall in winter 6 months No distinct dry season: differencebetween rainiest and driest month less than for C w and Cs,and the driest summer month >30mm(1.2 inches)
Df
Cooltemperate:coldsnow-forest: coldestmonthmean<-3"C(26"F)andwarmest monthmean>10"C(5O0F).Thislatter temperature coincideswith the poleward limit of forest Dry winter: at least 10times as muchrain in the wettest month of summer as in the driest month of winter Dry summer: at least three times as much rain in the wettest month of winter as in the driest month of summer, which has <30mm(1.2 inches); alternativeis 70%ormoreof themeanrainfallinthewinter 6months No distinct dry season; differencebetween rainiest and driest month less than for Dw and Ds
E ET EF
Frigid,mean temperature of the warmestmonthO"C (32'F) Perpetual frost, meantemperatureallmonths
D Dw DS
Climate frigid climates, and seasonality for the other categories (Table 9.1). Mediterranean climates (Cs), for example, have verylowsummerprecipitation.Lack of moisturein the growing season means that these climates are more difficult for plants and less encouraging for soil formation than other climates with comparable mean annual temperature and precipitation. There have been many attempts to improve, modify and replace Koppen’sclassification of climates. For soil science,it is perhaps not the best classificationbecause it subdivides extensively on the basis of temperature and seasonality, which are less significant for soils than is precipitation. From this perspective the classification of Holdridge (1947) is more useful because it relates particular kinds of vegetation to the amount of mean annual precipitation, potential evapotranspiration ratio and mean annual biotemperature in a straightforward manner. Mean annual biotemperature is a climatic index based on temperature records, adjusted for the observation that large vascular plants become physiologically dormant at temperatures of less than 0°C and more than 30°C. Mean annual biotemperature can be calculated from hourly or daily temperature records by substituting 0°C for all temperatures Iower than that, and 30°C for all higher ones. Holdridge’sclassification was devised and works best for tropical and subtropical vegetation (Holdridge et al. 1971). The typical vegetation for each climatic type is a useful concrete image of that climate, but it can be misleading. Vegetation of an unusual soil, of waterlogged sites or of a distinct microclimate may reflect regional climate less faithfully. Another disadvantage of this classification for paleopedological studies is that few of these plant formations have a long geological history. Even if Ordovician (440Ma) paleosols formed in a climate of the wet forest life zone, they could not have supported trees because these had not yet evolved. Some of the problems with climatic classifications stem from their attempt to place limits on a complexly moving atmospheric system interacting with obstacles in the form of continents and mountain ranges (Parrish 1998).Large-scale paleocliatic patterns can be reconstructed for a variety of continental configurations of the geological past by using atmospheric circulation models. Such reconstructions can be useful for understanding the wider paleogeographic context of paleosols. The simplest circulation pattern that can be
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imagined is one in which the Earth is a flat, featureless sphere, either all water or all land. In this case, there would be a zone of westerly wind flow in middle latitudes, consisting of a succession of cyclones (lowpressure centers, with clockwise circulation in the southern hemisphere) and anticyclones (high-pressure centers, with clockwise circulation in the northern hemisphere). Wind directions would converge to run easterly near the equator, and they would also be easterly near the poles. This general circulation pattern is somewhat modified by intervening continental masses. Meridionally orientated continents with marginal mountain ranges astride the westerly circulation gather clouds and precipitation on the windward side, leaving large areas of dry climate, the so-called ‘rain shadow’, downwind. These are found east of the Rocky Mountains of North America, the Andes of South America, and even the Southern Alps of the much smaller land massof NewZealand(Fig. 9.1). Other climatic effects are created by large continents orientated latitudinally. Large land masses cool more in winter and warm more in summer because they are remote from heat circulated by ocean currents. The large landmass of Eurasia. for example, develops a highpressure cell during the winter because its air is generally colder than that of the Indian Ocean, and this deflects the intertropical convergence zone southward (Fig. 9.1).In summer, however, its high-temperature (or low-pressure)cell attracts the intertropical convergence zone northward over India and southern China, bringing a season of onshore winds and torrential rains known as the monsoon. Intertropical regions of the Americas and northern Australia also have monsoonal circulation as, presumably, did other large latitudinally orientated land masses in the geological past. Summerdry subtropical climates, as in the Mediterranean region, the pacificNorth-west of theUSA, southernsouth Africa and south-eastern South Australia, are a special case of a similar phenomenon. These are produced in the downwind (usually eastern) end of summer highpressure (low-temperature) cells developing over a cool ocean. Each of these classifications and general circulation models is based on climates that have been observed in historical times. Over geological times, there is evidence for considerable variation in such fundamental constraints as the amount of solar energy received on Earth
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(Parrish 1998). For example, a wobble in the angle of the Earth’s rotational axis with regard to its plane of orbit around the Sun produces the precession of the equinoxes on a cycle of 2 3 kyr, and variation in obliquity of the axis of rotation on a cycle of 42 kyr. There are changes in the eccentricity of the Earth’s orbit also on time scales of lOOkyr (Hays et al. 1976). Over the past 1.8myr paleoclimate has fluctuated dramatically on comparable cyclicities, with large ice sheet expansion and frigid climate alternating with warm climate (Nilsson 1983). There also were earlier torrid ages punctuated by ice ages during the Carboniferous-Permian (390 Ma), Ordovician-Silurian (443 Ma), Late Precambrian (900-600 Ma) and Early Precambrian (about 24 50 Ma). These long-term paleoclimatic cycles could also be cued by astronomical forcings,but forcingson all time scales are so small that amplification by the climate system here on Earth is needed to explain the severity of past climatic change. The growth of polar ice caps promotes glacial conditions by reflecting sunlight back out into space. On the other hand, solar heat also can be retained on Earth by highly absorptive gases, such as carbon dioxide (CO,) and methane (CH,), which trap heat in the same way as the glass of a greenhouse. Both interglacials of the last 1.8myr and torrid ages of the last 1000myr appear to have been times of elevated atmospheric CO, (Jouzel et al. 1993; Berner 1997) and a spread of forests into polar regions (Harden et al. 1992; Retallack 1999b). Thus paleoclimatic change was not simply a matter of continental orientations or astronomical forcings,but of a series of historical changes so complex that additional information from paleosols is needed.
Indicatorsof precipitation Many soil-forming chemical reactions occur in dilute solutions, within pore spaces intermittently supplied with water by rain. Even when soil pores are full, these reactions differ from those in open water, such as rivers and lakes. The wetting and drying of soil changes the volume of small water films available for reactions, removing or concentrating reacted products. It is no wonder then that freely drained soils show more profound chemical weathering than waterlogged soils, or that freely drained soils of humid climates are more altered than those of dry climates.
Numerous chemical and mineralogical features of soilscan be related to mean annual rainfall. but only the most reliable soil features for interpreting paleosols are discussed here. Some features of modern soils, such as organic matter and clay content, show strong correlation with rainfall (Jenny 1941). Unfortunately, organic matter seldom is preserved sufficiently abundantly in well-drained paleosols (see Fig. 7.3). Magnetic susceptibility has been used to infer rainfall of Pleistocene paleosols (Maher & Thompson 1999), but more ancient (Pennsylvanian-Permian, 290 Ma) paleosols show such slight susceptibility enhancement that the Pleistocene transfer functions cannot be applied (Soreghan et al. 1997). Soils of more humid climates tend to be more clayey and more red-colored and have fewer easily weatheredminerals, but this also is true of soilsin hotter climates and soils of greater age (Birkeland 1999). Soils of great age also show conflicting indicators of paleocliiate, reflecting Quaternary fluctuations in precipitation (Chadwick et al. 1995). Conflicting indicators, competing factors and preservational problems compromise many paleoclimatic interpretations from paleosols.
Histosols, peats and coals The histic epipedon or peats of Histosols, which become the coals of sedimentary rocks, are mainly found in wet climates with an excess of precipitation over evaporation (Fig. 9.2), although peats are very broadly distributed (Lottes & Ziegler 1994). Thin, high-ash coal beds from desert oases, and lacustrine or coastal marshes are sometimesfound in sequences with caliche and gypsum where mean annual precipitation is 2 50-1 300 1~1111. High to moderate, or variable, ash-content coals, also with variable sulfur content, from planar (rheotrophic) mires, salt marshormangalarefoundlargelywithinthe bounds of 1300-2500 mm mean annual precipitation. Low-ash, low-sulfur coals of raised (ombrotrophic) mires are found in very rainy (>2500 mm mean annual precipitation) and everwet climates (Gyllenhaall99 1; Parrish 1998). Karst In rainy climates, limestone is dissolved and sculpted in outcrop (karst) and at the base of the soil or between
Climate
.-
. I -
Figure 9.2 Modern occurrenceof nonfrigid peats (histicepipedons)according to mean annualprecipitationand latitude (data from Gyllenhaal1991:Parrish 1998).
geological formations (cryptokarst). Buried karst landscapes are filled with sediment that shows some signs of alluvial or colluvial abrasion, whereas cryptokarst is filled only with soil, residues of dissolution or collapse breccias. The two are not always easy to distinguish because material filling true karst topography may include soil and the rubble of rockfall that at intervals were altered by soil formation as karst depressions
n = 196
1000
1000
2000
3000
4000
2000
3000
4000
mean annual precipitation (mm)
WerefilledfFig.9.3)(James&Choquette 1988).Thereis also the possibility of karstification of dolostone driven by waters enriched by gypsum dissolution, rather than the usual action of carbonic acid on carbonate (Bischoff et al. 1994). Nevertheless, fossil examples are now known of bothpaleokarst (Leary 1981; Coxon & Coxon 1997)andcryptokarst (Wright 1981). An unusually grotesque form of spongy and jagged
Thin
t Figure 9.3 Climaticsignificance of karst and calcrete landforms (fromJames& Choquette 1994:with permissionfrom the Geological Association of Canada).
113
ARID
SEMI-ARID
WET
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Chapter9
karst, called black phytokarst, is created by endolithic cyanobacteria, algae and lichens (Viles 1987). It appears to be restricted to humid tropical climates now (Folk et al. 1973), and probably was also in the geologicalpast (Folk&McBride1976). Pedocals and pedalfers Calcium carbonate is a common component of soils in dryclimates, butisnotformedinsoilsof humidclimates. Soils have been divided (byMarbut 193 5) into two main kinds: those with free carbonates (pedocals) and those lacking free carbonates (pedalfers).Apart from mountain soils, most soils west of the prairie-forest transition zone near Lincoln, Nebraska, in the midwestern USA, are pedocals, and most soils east of that l i e are pedalfers. In the midcontinent, the dividing l i e is at about 500 mm mean annual rainfall in the cool climate (5-6°C mean annual temperature) near Red Lake Falls in north-western Minnesota, and about 600mm in warmer climate (22OC mean annual temperature) near San Antonio in southernTexas (Birkeland 1999).Rainfall is more effective for weathering when it falls mainly during the mild winter in the Mediterranean climate of Israel, compared with the summer-wet climate of monsoonal India (Royer 1999; Retallack 2000). Although calcareous paleosols are a reliable indicator of dry climates, and noncalcareous paleosols of wet climates, it is not possible to specify the dividing line accurately without information on other aspects of paleoclimate.
Depth to calcic horizon Soils with free calcium carbonate usually have a distinct calcareous layer or calcic (Bk) horizon (Fig. 9.4). The depth of this horizon below the surface reflects the depth of wetting of the soilby availablewater.This depth to the calcic horizon should not be confused with the depth of carbonate leaching, which inpaleokarst andpedalfers is more clearlyrelated to time for development than paleoclimate (Retallack 1997b). In dry regions, the calcic horizon is closer to the surface than in wetter regions (Fig. 9S).Thisrelationship holdsonlyfor soilsof moderate development (nodules of carbonate rather than wisps or layers), in unconsolidated parent materials (loess, alluvium). It also works for soils and paleosols of seasonal warm climates provided the depth is measured to abundant (rather than solitary) nodules (Retallack 1991a) and within the low points of gilgai microrelief (Driese et al. 2000).It has not been shown to be valid for petrocalcic horizons, calcretized bedrock, dolocretes, carbonate collars, or hillslope soils prone to erosion (McFadden et al. 1998; Retallack 2000). Nevertheless, the relationship between depth to calcic horizon and mean annual precipitation has wide applicabilityto soils of tropical to frigid, monsoonal to mildly seasonal, anddesert to forest climates. Three difficultiesarise in applying this relationship to the interpretation of former rainfall from paleosols. First, a paleosol may have been eroded before its burial by later deposits. This can be assessed from root traces
Figure 9.4 Calcareousnodules (white) formingcalcic horizons (Bk) in Upper Oligocene paleosols of the Poleslide Member of the Brule FormationinBadlandsNational Park, SouthDakota,USA.Theerodedpipeof the mushroom-shapedconcrete benchmark in the center foregroundis 4 cm in diameter.
Climate
115
0-
-50-100-
-1 50-
-
-200-
J
~
0
=
r, Q
Figure 9.5 Depthof thecalcic horizon insoils formed under differentregimes of mean annualrainfall from a compilation of 3 17 soils &omallcontinents (fromRetallack 1994b;with permission from the Soil Science
al -250-
U
-300-
P = 139 6 -6 388D -0.01303d P = precipltation (mm) D = depth ( cm) r = 0.79 (I = f141 mm
-3507 1
Societyof America).
and soil structure in the surface of the paleosol. Signscant erosion may be suspected if a paleosol is overlain by cross-bedded sandstone of a paleochannel rather than by silts or clay that settled from ponded floodwaters. A second problem is compaction of paleosols after deep burial. This is more significant for clayey than for sandy paleosols, and can be calculated using standard equations from geological estimates of burial depth (Caudillet al. 1997; Sheldon&Retallack 2001). A third problem is higher past levels of CO, in the atmosphere, which is predicted by modeling to create deeper calcic horizons (Marion et al. 1985; McFadden et al. 1991). There are methods of assessing atmospheric CO, from the carbon isotopic composition of pedogenic carbonate (Cerling 1991). This is a serious problem only in the extreme greenhouse periods of the JurassicCretaceous, Ordovician-Silurian and perhaps also early Precambrian (Schwartmann & Volk 1991; Berner 1997;Ekartetal. 1999). Application of the calcic-precipitationrelationship to Eocene-Oligocene paleosols of central Oregon (Retallack et al. 2000) did not require recalibration for atmospheric CO,, nor for erosion in this aggrading alluvial sequence, but did require adjustments for compaction afterburial (Fig. 9.6).Applicationsof thisrelationshipto Pliocene ( 5 Ma) and Mississippian (330 Ma) paleosols madesimilar assumptions (Mack&James 1992;Caudill
100
200
300
400
500
I
600
700
mean annual rainfall (mm)
800
= I
-=I 900 1000
et al. 1996), but atmospheric CO, may be important for Cretaceous (9 8 Ma) applications (Mack 1992). Chemical composition Soilsof wet climates have fewer alkali and alkalineearth elements normalized to silica and alumina than soils of dry climates. This is because both alkali (Na+,K+) and alkaline earth cations (Ca2+,Mg2+)are progressively stripped from soils by hydrolytic weathering in wet climates. Surprisingly robust relationships can be found between the molecular ratio of bases/alumina and mean annual rainfall in soil subsurface (Bt) horizons (Fig. 9.7).This relationship was compiled from the soils of the conterminous USA analyzed by Marbut (1935) and includes mainly moderately to strongly developed soils on felsic and fluvioglacial parent materials under temperate arid to humid climates, and vegetation ranging from sagebrush desert scrub to humid deciduous and coniferous forest. Outliers to the relationship include very weakly developed soils of coastal dunes and streamsides, and very strongly developed midwestern soils on paleokarst or high terraces of the Carolina piedmont. Nevertheless, the relationship is surprisingly strong, and has proven stronger than any other derived from these data, including sophisticated algorithms to factor in parent material.
116
Chapter9 EOCENE
Age
I DUCH. 1
UINTAN
Ma
CHADR. 35
40
45
I OLIGOCENE IOR.\WHI.]ARIK. 30
-
terminal Eocene climatic drying
-A
~
mean annual 1000precipitation -
(mm)
500100
-B
"
" - I
I
'
"
"
"
"
"
/Ifores>
regional vegetation 50 ("'7 paleosols)
A
100
4c ,p-early '
:d
a
'
'
.
'
"
successional
"
"
"
"
'
rise of weeds1
local vegetation 50 (% paleosols) 03-
P E
a 2L E
.E 5
strong
aQ)
a 5 moderate E" $2 weak > o f very weak :s I
fossiliferous levels study areas
This relationship (Fig. 9.7) is unsuitable for paleosols of very dry climates (with evaporites) or those of very wet climates (with common kaolinite), but is applicable to many paleosols when one considers that modal annual precipitation on land is 746 mm (Baumgartner & Reichel 1975). It is also unsuitable for paleosols of very weak or very strong development and for those altered during burial of chemical composition by illitization (Mora et al. 1998). In a recent application to Eocene-Oligocene paleosols of Oregon, USA, for which illitization is not evident from X-ray
terminal Eocene destabilization
-
Figure 9.6 Climaticdrying over geological t i e (Ma)is indicated by depth of calcic horizon (open circles)and chemical composition of paleosols (filled circles) in a sequence of Upper Eocene to Oligocene paleosolsin the Painted Hills. central Oregon, USA(A).Alsoshownarechangesin vegetation through time from root traces and profile form of the paleosols (B and C) and concomitant changes in ecosystemduration fromdevelopment of thepaleosols (D) (from Retallacketal. 2000;with permission from the GeologicalSocietyof America).
diffractometry, paleoprecipitation indicated by chemical composition of subsurface (Bt)horizons proved to be very similar to that indicated by depth to the calcic (Bk) horizon (Fig. 9.6).
Clay minerals The kinds of claysformedin soils also represent products of hydrolysisof weatherable minerals and can be related to the amount of rainfall available to soils (Folkoff & Meetenmeyer 1987). Grain size and mineral composi-
Climate 4
117
I
p = -7598 + 1300 r = 0.7 o = 174 mm 80 -I
*.e
-
2 c
60-
20
-Z on J
gibbsite
diospore
boehmite
L
'
400
800 1200 mean annual precipitation (mm)
Figure 9.7 Relationshipbetween mean annual precipitation and baselalumina molecular ratio of subsurface (Bt) horizon in 127NorthAmericansoils (fromReady&Retallack1995; with permission from the Geological Society of America).
tion of parent materials, temperature, seasonality of rainfall and time for formation of a soil are competing factors determining clay composition. For paleosols there are additional problems determining which clays were products of soil formation, as opposed to inherited from a parent material or altered after burial. The bright clay fabric, fine grain size and poor crystallinity of soil clays compared with alluvial clays may be substantially alteredduringdeep burial (Retallack1991b). With these cautionary issues in mind, clays of wetter climates generally have a 1: 1 rather than 2 : 1 layer structure, are less rich in cations and lower in the general weathering sequenceof clay-sizedminerals(seeFig. 4.4).In basaltic soils of tropical, continuously wet parts of Hawaii, USA, for example, smectite is the dominant soil mineral below about lOOOmm annual precipitation, kaolinite between 1000 and 2000mm, and iron oxide and alumina above 2000mm (Fig. 9.8). Similar results have been reported for tropical, alternating wet and dry climates (Sherman 1952). Progressively more base-poor clays in wetter climates also have been observed in Californian soils under a warm temperate (mean annual temperature of 10-15.6"C) Mediterranean climate. Some differences observed can be related to coarser-grained parent materials of felsic composition, such as granite, compared with finergrained rocks of mafic composition, such as diabase (Barshad 1966). In both, however, smectite is a major component of soils that receive less than 500mm mean annual rainfall, and kaolinite and halloysite dominate
kaolinite
smectile
3000
1000 2000 3000 mean onnual precipitation precipitation (mm)
Figure 9.8 Clay mineralcomposition of soilsformedunder differentregimes of mean annual rainfall in the equable, tropicalclimate of Hawaii (adaptedfromSherman 1952: with permission from the American Institute of Mining Engineers).
soils of wetter climates. Iron and aluminum oxides predominate in climates with more than 2000mm mean annual rainfall. Similarchangesinamountandkindof clay inancient sequences of paleosols can be regarded as evidence of paleocliatic change (Retallack 1986b; Fiorillo & McCarty 1996). Sodium smectites of dry climates produce adistinctive soilstructure of domedcolumnar peds extending throughout much of the solum (natric horizon: McCahon &Miller1997). At the other extreme, the clay mineral kaolinite and bauxite minerals boehmite and gibbsite are indicators of very humid ever-wet climates(PriceetaZ. 1997;Parrish 1998). The fibrous clay minerals palygorskite and sepiolite are found in soils of very arid regions (Botha & Hughes 1992;Calvo et aI. 1999). InMorocco. forexample, palygorskite makes up the entire clay fraction of soils that receive less than 300mm mean annual rainfall. These mineraIs are not stable during burial in paleosols, although they have been recognized in some very ancient examples (Watts 1976). Palygorskite may be hydrated to sepiolite, and both minerals are leached by rainwater and groundwater. Evaporite minerals Evaporite minerals are mainly salts, readily soluble in
118
Chapter9
water, that accumulate in soils of very dry climate. The most common of these minerals in soilsis gypsum, usually found in the crystal habit of needles or laths (selenite). Other common evaporite minerals are cubic halite and fine needles of sylvite or mirabilite. Evaporiteminerals form where there is a source of salts, such as ashyvolcanic materials or in areas affected by sea spray. The most important constraint on their formation, however, is a climate in which evapotranspiration exceeds precipitation (Chen 1997). Like calcic horizons in soils,the gypsic horizon (By) is shallower the drier the climate. In Israel it is within 2 m of the surface in soils receiving 290mm mean annual precipitation, within 20cm in soilsreceiving 100mm, andwithin2 cmin thosereceiving less than 50mm (Dan & Yaalon 1982). Although fossil gypcretes are known (Hartley & May 1998),salts such as gypsum are readily dissolved by groundwater. All that remains of many ancient evaporite layers are pseudomorphs of crystals (Barleyet al. 1979)or zones of breccia where country rock has collapsed into the dissolvedsalt layer (Bowles&Braddock1963). Desert pavement In climates too dry (lessthan 300 mm mean annual precipitation) for abundant plant cover, soils are armored by a natural pavement of stones, formed by a combination of lifting by dust infiltrating into cracks and of eolian deflation of materials finer than gravel. Such stone lines can be striking in eolian successions (Wright et al. 1991), but more is needed than merely oversized clasts as indicators of desert pavements. Also found are windsculpted stones (ventifacts),with a characteristic vuggy or keeled shape. Other associatedfeatures of desert pavements are carbonate collars (rims of calcite around stones at ground level), rock varnish (amorphous iron-manganese crusts on top of stones) and vesicular structure (roundedvugs: McFaddenet al. 1998).
Indicatorsof temperature Although water is the medium for most soil-forming processes, the rate of these reactions also is controlled by temperature. According to Van t’Hoff’s temperature rule of chemistry,for every 10°Crise in temperature the rateof achemicalreaction increases byafactorof two to three. As surface temperatures over which water re-
mains liquid range from 0 to 84°C in soils, there is the potential for a differenceof two orders of magnitude in reaction rates between frigid and tropical soils. This tallies well with the observation that most tropical soils are much more deeply weathered than polar soils (Birkeland 1999). Despite these dramatic differences, it is in many cases difficult to disentangle the complementary effects of rainfall and time for formation from those of temperature in promoting weathering. This particularly applies to the amount of clay, redness of hue, plinthite formation,depthof weathering,depletionof weatherableminerals in soils and dissolution of limestone. Each of these features of soils is more pronounced in soils of warm climates, suchaslateriticandbauxiticsoilsandtowerkarst landscapes (Jennings 1985; Bardossy & Aleva 1990; Parrish 1998).Tropical regions, however, were less disrupted by glaciation periods in the past 1.8myr, and also were continuously unfrozen, unlike many soils of cool temperate climates (Paton et d.1995).The competing effects of rainfall and time for formation compromise the paleoclimatic interpretation of these features of soils, which can form in cool climates of temperate rainforest (Tayloret al. 1992 ;Bird&Chivas 1993). Sphericaf micropeds and termite nests A distinctive feature of some tropical soils, such as Oxisols (of the US taxonomy) and Ferralsols (of the FA0 classification),is the abundance of spherical micropeds in petrographic thin section (Stoops 1983). These are small (10-100pm) spherical aggregates of claystone. Usually the clay is kaolinitic and strongly stained red and opaque with hematite, or pink with gibbsite. There may also be small grains of quartz within the peds, but few other minerals are found. The micropeds are both physically tough and chemically inert. Although made of clay, they bestow a porous texture to the soil, so that it seems sandy to the touch (Nahon 1991). With compaction in a paleosol,the micropeds may not be obvious in hand specimen, but they can be revealed in petrographic thinsections (Retallack 1991a). Oxic spherical micropeds are thought to originate primarily as fecalpellets of ants and termites (Mermut et al. 1984). Complex nesting structures are even more diagnostic of termites (Fig. 9.9). Termite nests are now known as far back as the Triassic (230Ma: Hasiotis &
Climate
119
surrounding galleries &calcareous rind ,central chamber
Figure 9.9 A fossil termite nest Termitichnus qatrani Bown 1982 from the Upper Eocene, JebelQatraniFormation,FayumDepression, Egypt (fromBown1982;withpermission from Elsevier,Amsterdam).
Dubiel 1995; Hasiotis & Demko 1996b), and termite wings as old as the Cretaceous (120Ma, Labandeira & Sepkoski 1993). No ground-dwelling termites are currently known in climates cooler than the 8°C mean annual isotherm (Retallack199la). Oxygen isotopiccomposition As temperature rises the activity of the common light isotope of oxygen (l60)increases relative to the heavy isotope ("O), and the ratio of these isotopes in soil minerals or the skeletal biominerals of soil creatures reflects their temperature of formation(Hays & Grossman 1991).These isotopic measurements are made as parts per mil (%o) negative (isotopically lighter or with more l60)or positive (isotopically heavier or with more l80) relative to a standard, a fossil belemnite from the Peedee Formation of South Carolina (PDB) or Standard Mean Ocean Water (SMOW).The principal difficulty with estimating paleotemperaturefrom oxygen isotopic valuesis that the isotopic composition of the water from which themineralformedmust also beknown.Theoxygenisotopic compositionof water is sensitive to salinity, evaporation and other factors in addition to temperature. Original water compositions can be inferred from the oxygen isotopic composition of aragonite in bivalve or snail shells,or less directlyfrom the apatite of vertebrate bone or teeth. However, different biominerals are subject to different vital effects that alter the oxygen isotopic composition. Finally,oxygen isotopic compositionof soil minerals can be reset by elevated temperatures and fluid migration during burial and metamorphism.Such burial alteration has confoundedseveral recent attemptsto infer paleotemperaturefrom oxygen isotopic composi-
entrance gallery
tion of pedogenic carbonate (Quadeat al. 1993,1995: Mora at al. 1998). Even if absolute temperature determinations remain model dependent, oxygen isotopic data frompaleosolscan be useful for revealingrelative changes in temperature through time, such as the brief greenhouse warming at the Paleocene-Eocene boundary (55 Ma) in the Powder River and Bighorn Basinsof Wyoming, USA (Kochetal. 1995).Theoxygen isotopic compositionof kaolinitefrom Australiandeeply weathered paleosols ranging in age from Permian to Neogene also shows a trend of warming compatible with evidence of northward continental drift of Australia (Fig. 9.10). Permafrost structures Soilsof frigid climates (mean annual temperature
120
Chapter9
Post mid-Tertiary residual clays
Pre mid-Tertiary residual clays
&b..adb
. ..
Pre late-Mesozoic residual clays
Permian residual clays DD
DD
d D
I
I
I
I
I
I
I
I
I
+7
+9
+11
+13
+I5
+17
t19
+21
+23
Somepermafroststructures in paleosols couldbe confused with mukkara structure of Vertisols (see Fig. 4.2) or with earthquake-induced liquefaction features. Frost-heave cracks are generally wider and less deformed, are not associated with lentil peds and distributed slickensides, and are seldom in soils as clayey as Vertisols (Paton 1974). In addition, permafrost deformation is filled from within or above the soil, unliie earthquake liquefaction features, such as craters, conical sand blows and sill-like extensions of dikes, which originate from paleosols and sediments well below the soil (Obermeier 1996).
Indicatorsof seasonality The seasonal availability of water and temperature may have an appreciable effect on the development of soil features. The winter rains in deserts of Israel are more effective at leaching carbonate into a subsurface calcic horizon than are the rapidly evaporated summer rains of India (Yaalon 1983).Seasonalityof rainfall, of heat, of dust influx and of other agencies of soil formation also are important to the clayeyness, redness and base saturation of soils. Each of these features,however, is controlled by other factors that make it difficult to tease out the contribution of seasonality. The nature and degree of seasonality is better revealed by other features of soils, only the most diagnostic of which are reviewed here.
Figure 9.10 Oxygen isotopiccomposition of kaolinitic deeply weathered paleosols from Australia reflecting increasingly warm paleocliates through geologicaltime (from Bird&Chivas 1993;with permission from the GeologicalSocietyof Australia).
Mukkara and gilgai Gilgai microrelief and mukkara subsurfacestructure of swelling clay soils (see Fig. 4.2) are indications of a ciimate with pronounced dry and wet seasons. They are not found in extremely arid or humid climates.Mean annual rainfall of gilgai soils in Australia ranges from 180 to 1520mm. Some moisture is needed to promote clay formationand provide a contrasting wet season.In drier climates these soils form on parent materials already clayeyandsmectitic(Paton 1974;Patonetal. 1995).On the other hand, clays become more deeply weathered and lose their swelling properties in humid climates. This is in part the result of hydrolytic loss of cations, but alsoisrelatedto the conditioningandstabiliiation of soil with organic matter and roots of forest vegetation in wetter climates. These paleoclimatic generalizations may not have held during the Precambrian, before the advent of substantial land vegetation or before the development of the present atmosphere(Fig. 9.13). Concretionsand argillans Concentricbands of concretions and strong banding of clay skins (argillans)may reflect seasonal differences in the chemical condition of the soil (Nahon 1991: Delvigne 1998). Well-drained soils of subtropical, monsoonal climates have both calcareous and ferruginous concretions (Fig. 3.15B: Sehgal & Stoops 1972). The calcareouslayers may have accumulatedunder alkaline
<+4
<+1 <-2
<+3
<4
<-1
<-1 to<-3 <+1 <-2
<1200
Thin snow cover
>400-800
> 50 to 500 Rapid cooling early winter with thin snow Rapid cooling Rapid cooling, continental 1000to>7000
2600 to>7000
Freeze index (“Cday yr-l)
Continental Continental
Continental
Gentle cooling
1500->7000
1000->7000 2700-7000 2300-27000 7O(tSOOO 1000-5000 800->7000
1000->7000
70S-3800
<-35 <-8
+10t0+20
Other climatic indication
<+3
<-12 to<-20
Sand-wedge polygon Seasonal frost crack polygon Tundra hummock (high latitude) Earth hummock (thufur) Seasonal frost mounds Micro-hummocks Mudpits,nonsorted circles Palsas Closed-systempingos Open-system pingos String bogs Rock glaciers Features from thaw of ice-rich permafrost Sortedcirclesand stripes (4m) Sorted circlesand stripes (>1m) Miniature sorted forms Microforms of gelosolifluction
<-25t0<40
Mean temperature Mean annual warmest month (“C) precipitation (mm)
2300 to >7000
<-4 to <-8
Ice-wedgepolygon
Mean temperature oldest month (“C)
<-lo
Air mean annual temperature (“C)
Periglacialfeature
Table9.2 Climaticinferencesfromperiglacialfeatures(Williams 1986).
200-1000
300-2200 <100-1500 250-1 700 lOO(t3000 300-1000
200-1000
50S3000
100-1500
1000-2000
100-1000
Thaw index (OCday yr-l)
122
Chapter9
6”C,,,
26
-24
Stratigraphic section of the Weller Coal Measures formation, Allan Hills
PDB W o o )
-22
Reconstruction of mid-Permian landscapes in the Victoria basin
20
m
80
Discontinuous permafrost of taiga climate
.-cC
E
60
b
r
40
t
B
20
E
Continuous permafrost of periglacial tundra climate
Figure9.11 Restorationof landscapesduringmid-Permian(260-2 70Ma)deglaciation with depositionof Weller Coal Measures in the Man Hills,Victoria Land, Antarctica.Deglaciation also resultedin a less stronglyfluctuating carbon isotope composition of organicmatter (S13Corz)(fromKrull1999;withpermissionfromPalaios).
conditions of dry periods but were corroded as ferruginous layers formed during the wet season. Monsoonal tropical soils and paleosols also show unusually diffuse calcic horizons with calcareous nodules scattered throughout theentire profile (Retallack 1991a). Patternsof root traces In seasonallydry wooded grasslands,grasses and trees
have a copious surficial network of roots that is active during wet parts of the year. During the dry season, however, leaves of grasses wither back to their root stock, but some trees obtain moisture from deep within the soil by means of especially stout, deeply penetrating roots called ‘sinkers’ (Van Donselaar-ten Bokkel Huinink 1966). This pattern of a dense near-surface networkof finerootswith a few deeply penetrating stout roots isrecognizableinpaleosols (Retallack199la).
Climate
123
Figure9.12 Deepclastic dike in acoalof the mid-PermianWeller Coal Measures of the AIlan Hills, Antarctica,interpreted as polygonalpatterned ground. E.S. b u l l and hammer for scale.
A comparable bimodal distribution of roots also is seen in soils of seasonally dry swamps (see Fig. 11.5). These are distinct from permanently waterlogged swamps even to the casual observer wading through the wetlands of Louisiana. One only sinks to the knees in the water of a seasonally dry swamp because the peat is thin over more firm clay or sand. In permanently wet swamps, on the other hand, it is difficult to walk in the soupy peat. Unlike the tabular root systems typical for permanently waterlogged swamps, seasonally dry ones may have some deeply penetrating roots for the dry season, as well as tabular roots for the wet season. Soils of seasonally dry swamps also have little or no surface peat because this decays during the dry season after it has accumulated during the wet season. Both kinds of swamp soils have the drab color and gley minerals typical of poorly drained soils. A serious complication for the interpretation of seasonality in swamp soils is the tendency of such lowland regions to be slowly subsiding sedimentary basins, so that a seasonally dry swamp subsides below water table to become a permanently waterlogged soil. This is common in Louisiana today (Coleman 1988), and also was common in the past because the seat earths under many coal seams contain both deeply penetrating and tabular fossil roots (Gardner et al. 1988; Retallack 1999b). In such situations the earlier seasonally dry paleosol can often still be discerned despite overprinting by the later permanently waterlogged one because soilformation that would alter
the older paleosol is curtailed below permanent water table.
Charcoal Fossil charcoal, or fusinite as it is termed by coal petrographers (see Table 4.2), is a common component of paleosols and can be evidence of fire in woody vegetation. Occasional fires rage in a wide range of climates, exacerbated by human carelessness. Occasional fires, however, do not produce charcoal with such regularity that it accumulates in soils faster than it decays. In nature, fire is most common in soils and vegetation of seasonally dry climates. The scrubby chaparral vegetation of California hillsides and the comparable maquis of south-western Europe and mallee scrub of southeastern Australia are renewed every few decades by burning during the long, alpost rainless summers they endure (di Castri et al. 1981).Even swamps, such as the Florida Everglades, burn extensively after short electrical storms during their long dry season. Many seasonally dry grasslands also burn during a dry season. Not all fires are recorded in paleosols, but if charcoal is common, then fire and dry seasons probably were frequent (Retallack 1991a).
Carbon isotopic composition Soil organic matter has a carbon isotope composition
124
Chapter9
roil an Hekpoort Basalt
I
//
\ I
yellowish sericitic schist weathered basalt greenish gmy, crystalline, quartz-berthierine rock
WEST
n
Waterval Onder clay
R
-5
/
cross-bedded quartzite’ black sericitic schist EAST
1 10 5 I
-
-
-
20 m
Figure 9.13 ReconstructedearlyProterozoic (2200Ma)paleosolinthe 1owestDwaalHeuvelFormationnear WatervalOnder, Natal,SouthAfrica.withanannotatedfieldsketch(below)showinginterpretationofitsmukkarastructure(fromRetal1ack 1986a: with permission from Elsevier,Amsterdom).
Climate
(6l lCo,g)that reflects the annual average composition of vegetativebiomass, but soil carbonate in the same soil or paleosol has a carbon isotopic composition (613C,,rb) that reflects dry-season soil dissolved CO, from root respiration and microbial decay of organic matter. The difference between these two values (813Corg- 813Ccarb) within individual soils and paleosols reflects the magnitude of seasonal difference in precipitation, particularly in strongly seasonalmonsoonal tropical climates (Wang &Follmer 1998).Thedifference(813Corg- 813Ccarb)also is related to (i) diffusion of atmospheric CO, depending on its atmospheric partial pressure and (ii) equilibrium fractionation of CO, for the temperature of formation. Seasonality changes can be observed as relative shifts in isotopic values from successions of paleosols, or modeled by making assumptions about diffusion and fractionation. Indicatorsof greenhouseatmospheres Methane and CO, are currently trace gases in our atmosphere (0.00017 and 0.035 vol. %, respectively). Both gases also are human pollutants, produced in unnaturally large amounts by agricultural and fossil fuel industries. As large molecules they are of concern as greenhouse gases. Like the panes of a greenhouse, an atmosphere enriched in CO, and CH, is transparent to solar radiation, but reflects back to the surface solar energy that would otherwise escape to space. Both CO,
Figure9.14 RelationshipbetweenP3Cof soil carbonate and inferredpast atmospheric CO, partial pressure (multiplesof present atmosphericlevel),showing values for differentcarbonate phases in a single Upper Devonian (360Ma)paleosol from Pennsylvania (fromDriese&Mora1993; with permission from the International Association of Sedimentologists).
125
and CH, increase in abundance within bubbles in polar ice cores formed during interglacial stages, and both gases decrease during glacial stages of the past 1.8myr (Jouzel et al. 1993).These gases also have been analyzed from bubbles within amber in paleosols (Landis et al. 1996), but such records are limited by the rarity of amber. Both gases also leave distinctive isotopic traces within paleosols, particularly in their carbon isotopic composition. Soil profiles show a strong gradient in both the abundance and isotopic composition of CO, and this depth function can be exploited in paleosols as an isotopic CO, paleobarometer (Ceding 199 1, 1999). Carbon dioxide near the surface of the soil is present in only trace amounts and has a heavy isotopic composition (positive or enriched in I3C). Deep within the soil, CO, from root and animal respiration is isotopically lighter (negative and depleted in 13C) and CO, is more abundant than in theatmosphere(byasmuchas 110times;Keller&Wood 1993: Amundson et al. 1998).The actual value in the soil depends on a mixture dominated by isotopically light CO, from soil biota when atmospheric partial pressure of CO, is low, but dominated by isotopically heavy CO, diffusing from the atmosphere when atmospheric partial pressure of CO, is high. The isotopiccomposition of soil CO, is preserved in calcite and other soil minerals within paleosols and this can be used to estimate former atmospheric abundance of CO, by assuming typical isotopic values of atmospheric and soil gases (Fig.9.14).
0-
1 I
SPAR I
ZOCONCRETIONS 1
I
3,000
6,000
I
ppmV past atmospheric partial pressure of carbon dioxide
126
Chapter9
0" 8000 0
e
-g
f
6000
4000
n
'C 0
2 2000
m
500
400 300 200 millions of years ago
100
Figure9.15 Atmospheric CO, partial pressure (multiples of present atmosphericlevel: PAL and pA or micro-atmospheres = ppmV) of the past SOOmyr,as estimated by a sedimentmassbalance model (curve and shading, after Berner 1997) and confirmedby the isotopic composition of paleosolcarbonate (opencircleswithstandarddeviation) (afterRetallack 1993b: Yapp&Poths1994: Ekartetal. 1999).
field observations
Use of this model is compromised by a number of competing variables. It assumes, for example, a modern range of biological productivity and soil respiration. A different family of curves must be used for desert paleosols (Fig. 9.14), which can be identified from gypsum crystals or shallow calcic horizons (Retallack 1994b). The model also assumes a well-aerated, nonwaterlogged soil. Yet another family of curves must be used for gleyed paleosols (Fig. 9.14), which can be identified from high chroma mottles and minerals such as siderite (JCraus & Aslan 1993; PiPujol& Buurman 1997). Another assumption concerns photosynthetic pathways of plants (seep. 304) in thepaleosol, usually assumed to be C,, although isotopically heavier C, and CAM pathways are common in Neogene plants (<7Ma: Bocherens et al. 1994; Koch 1998), and may also explain anomalous Permian and Carboniferous values (c. 300Ma in Fig. 9.1 5). Finally, the model assumes simple physical diffusion of gases withminimal biologicalinfluence,as is best
-
% organic carbon
613C
-35
-25
0 10 20 3C
weathering ratios . A12(3,
CaO + MgO Na,O + K,O
Ba/Sr
+
1
3
5
u
1
0
4
8
percent grain size
50
Figure 9.1 6 Unusually light carbon isotopic composition of organic matter and selected pedologicalvariables in a LowestTriassic paleosol(250 Ma, above) compared with an uppermost Permian paleosol (below)at Graphite Peak, centralTransantarctic Mountains, Antarctica (fromDull&Retallack 2000: with permission from the Geological Society of America).
Climate represented by micritic pedogenic nodules. The plantrespiration-dominated gases of rhizoconcretions and the sparry cements formed during burial by the last phases of microbial decay can give very different isotopic values than nodules in the same paleosols (Fig. 9.14; Retallack 1 9 9 2 ~ )Nevertheless, . early results of this CO, paleobarometer match the general rise and fall of previously modeled atmospheric CO, over the past 500myr(Fig.9.15). Atmospheric methane can also leave an isotopictrace in paleosols in the form of unusually depleted or light isotopic values. Organic matter from plants is typically
127
light by -20 to -30%0, but only methanogens can fractionate between "C and 13C to produce methane with values of less than -36o/w to as low as -8OD/oo. Such isotopically light methanogenic CH, can then be fixed in soil organic matter by methanotrophic bacteria (Hayes 1994). This is the best explanation for extremely depleted isotopic compositions of organic matter in earliest Triassic paleosols of Antarctica (Fig. 9.16).Methane was probably a n important contributor to the postapocalyptic greenhouse following that greatest of all extiictions in the history of life, at the end of the Permian (Retallack& Krulll999; Krull & Retallack 2000).
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 10 Organisms
Climate and vegetation are so closely linked in the modern world that these two controls on the degree and kind of soil formation may be difficultto tease apart. The role of vegetation in soil formation can be understood from the example of artificial growth chambers, such as the climatron of the MissouriBotanical Garden in St Louis, USA. This enormous growth chamber was built with heating and plumbing (a microclimate), soil material (parent material) and landscaping (topography). The planting of seeds and seedlings,however, has markedly affected the quality of its soils, independent of these other factors. Its plantations have become more lush, bulky and resistant to perturbation. This experiment in recreating natural habitats has counterparts in nature in the form of plant succession on surfaces disturbed by fire, flood, landslides, volcanic eruptions or human activities. There are close parallels between the successionalcolonization of asurface by organisms andsoildevelopment, as was seen during recolonization of the volcanically devastated Indonesian island of Krakatau after 1883 (Fig. 10.1).There was a buildup of biomass, accumulation of reserves of organic matter, maintenance of stability, and mitigation of diurnal and seasonal fluctuations in water availability and in temperature (Thornton 1997). On geological time scales, the evolution of more bulky and complex plant communities during Paleozoic time (Retallack 1997e) provides another example of vegetation as an independent variable in soil formation. These successional and evolutionary changes in vegetation can be contrasted with plant zonation and migration. Zones of different kinds of vegetation develop along environmental gradients, such as the successively more windblown habitats near beaches and hilltops on the remnant peaks of Krakatau (Fig. 10.1). In these cases the effects of vegetation on soil formation are mixed with those of topography. Zonation of plants over larger regions controlled by regional climatic differences, or the migration of plant communities as climate 128
changes, also are situations in which it is difficult to disentangle the effects of vegetation as an independent variable. In assessing the effects of vegetation alone, the critical question is this: what would the soil have been like under different plants? Large plants are only the most obvious part of the organisms that form soils. Many other creatures also play a role in their formation. Fungi, for example, in intimate association with the roots of many vascular plants, make available nutrients such as nitrogen and phosphorus in forms that can be utilized by their host plants. Fungi also play a role in decay of leaf litter, returning organic matter to the soil. Earthworms aid in mixing soil minerals with organic matter, thus improving its fertility and stability for other organisms. These general effects of organisms are critical to understanding soil development as a whole. Here, however, attention is focused on particular effects of organisms from which their presence in paleosols can be inferred. Some kinds of burrows and fecal pellets are distinctive of particular organisms. Not only particular organisms but also different kinds of ecosystems can be interpreted from features of paleosols. Different plant communities create distinctive patterns of root traces, soil structures and overall profile form. These features reflecting different kinds of ecosystemsprovide a useful overall assessment of the influence of organisms on a particular paleosol. Quantification of such general effects of organisms is difficult because their degree of expression is related as much to the nature of organisms as to the time available for soil formation. Time for formation can be constrained by radiometric dating and other methods in the study of Quaternary soils, and such studies provide a basis for interpreting the time for formation of more ancient paleosols. A hidden assumption in such interpretations is that rates of soil formation in the past were comparable with those found today under comparable kinds of vegetation. Emphasis is laid here on the nature rather than
Organisms m
1883
1886
1906
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SUCCESSION 1
Figure 10.1 Plant successionand zonationfollowing the 1883 volcanicdevastationof the Indonesian island of Krakatau (data fromThornton 1997).
degree of soil formation that can be attributed to general ecosystem types. Within such wide interpretive limits, fossil soils can be viewed as grounds for interpreting the evolution of terrestrial ecosystems, independent of the fossil record of plants and animals (Retallack 1997a,e). Much is already known about land organisms from the study of spores and pollen (Traverse 1988). fossil plant remains (Taylor & Taylor 1993) and bones and teeth (Benton 1997). These various kinds of fossils are found in deposits of lakes and streams but also are preserved in fossil soils. It is clear, considering likely ecosystems supported by particular paleosols, that only a small fraction of the potential kinds of fossils of organisms that once lived in a paleosol are actually preserved in it (Retallack 1998a). Fossil soils are preservational environments as well as former living environments, and consideration of their fossils is revealing incompleteness of the fossil record. Both paleosols and fossils may provide supportive evidence for a particular kind of biota. If these two lines of evidence are not in accord, however, this is not necessarily an error. It could instead reflect preservational biases in the kinds of information recorded in fossils and paleosols. Alternatively, it could be that particular kinds of soils in the past supported a biota that seems anomalous by modern standards (Retallack 1997d).Even when the fossils found in paleosols are well understood, morecan be learned of their preservation, ecology and evolution from the paleosols on which they lived.
Traces of organisms A footprint in shale or a burrow in sandstone is a fossilized record of past life, but one very different from a fossil of the usual sort, such as a bone or leaf. Footprints and burrows do not represent part of an organism, but are a record of an organism’s activities. Whereas a bone or a leaf may be distinctive enough to be identified as a particular species, it is seldom possible to identify the makers of a footprint or burrow with such precision. On the other hand, trace fossils are found in the place where the animal lived, rather than transported and selectively preserved in a burial environment, like many fossil bones and leaves. This important advantage of trace fossils as evidence for past animal behavior has long been recognized, especially in the interpretation of marine trace fossils (Bromley 1996). It is only recently, however, that the rich record of trace fossils in paleosols has begun to berevealed (Buatoiset al. 1998). Trace fossils have been studied in the same way as other fossils, even to the extent of giving them L a t i binomial names. This is more a matter of convenience than of taxonomic consistency. A system for naming the great array of trace fossils now known was needed to communicate between interested scientists. Informal naming systems of letters or numbers tended to be ignored without the rules of priority and other archival procedures demanded by the International Code of Zoological Nomenclature (International Commission on Zoological Nomenclature 1995 ) .It is, of course, rec-
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ognized that trace fossils are not actual fossils of organisms, and that in most cases their makers may never be discovered. Some admission of this is implicit in the proposal that generic names of trace fossils uniformly be distinguished by the suffix ‘-ichnus’,from the Greek word for track (Bromley 1996). Classifications of trace fossils have been attempted, based on preservational types or behavioral groups, rather than evolutionary relationships inferred by classification of other kinds of fossils. Many trace fossils are preserved best in, and on, thin beds of sandstone interbedded with shale. Sandstone is a durable casting medium. Whenweatheredalongthecontactwithshale, it reveals trace fossils more clearly than those entombed only by sandstone or by shale. Preservational types of trace fossils can be classified according to whether they are on the top, bottom, within or beyond the casting sandy bed (Hantzschel1975). Such distinct beds with trace fossils along zones of lithological contrast are known in paleosols (Retallack 1976) but generally are rare and so of limited value in characterizing terrestrial trace fossils. Another way of classifyingtrace fossilsis by the kind of activity they represent, such as feeding burrows, dwelling burrows and crawling marks. Such systems also are of limited value for trace fossils in paleosols, in part because so much more is known about the behavior of terrestrial animals. The complex burrow systems of bees, for example, are used for egg laying, feeding larvae and for dwelling. More appropriate for trace fossils in paleosols are common English terms used in studies of behavioral entomology, such as burrow, hive, cell, gallery and chimney. The naming of trace fossils in paleosols will continue to rely on procedures established in the study of marine trace fossils, using the same ichnogenera in many cases. Detailed interpretation is better based on the enormous amounts of information gathered by soil biologists, as summarizedin the ensuing paragraphs. Microbes This common abbreviation for ‘microbiological organism’ is a suitably vague term for a variety of organisms that seldom leave distinctive traces in soils but nevertheless have profound effects on soil formation. Microbes include a host of unicellular creatures,such as bacteria, algae and protoctistans (Margulis & Schwartz 1982).
Their effects on soils are best considered under general roles and nutritional groups (Table 10.1)rather than along taxonomic lines. Basic requirements of microbes include a source of energy, which can be either from the Sun (phototrophs) or chemicalcompounds (chemotrophs),and a source of carbon, which can be from carbon dioxide (autotrophs) or other organic compounds (heterotrophs). These metabolic reactionscan be drivenby the reducing power of electronsfromorganicmatter (organotrophs)or from minerals in the environment (lithotrophs). Because reduction and oxidation reactions are central to many metabolic processes, organisms also can be classified according to their relations with oxygen. Those that must use oxygen as the terminal electron acceptor for energy conversion (aerobes)can be distinguished from those that cannot use oxygen in this way (anaerobes), and those that can use either aerobicor anaerobicmetabolic pathways (amphiaerobic).From this brief outline it can be appreciated that microbes influence the pH and Eh of soils as a consequence of their livelihood (Baas-Becking et al. 1960).This would not be especially noteworthy if microbes were merely responding to predetermined physicochemical conditions, but this is not entirely the case. Chemical and physical equilibrium in soils is constantlythwarted by the metabolic activitiesof organisms, especially of microbes. Depletion of oxygen in stagnant water within soils, for example,occurs more rapidly than would be expected by simple diffusion under a given temperature and pressure because of the oxygen demand of aerobicmicrobes. Once oxygen is depleted, these microbes die or migrate elsewhere and a new wave of anaerobic microbes determine further changesin the chemistryof the soilwater. Weathering is demonstrably accelerated and guided by microbes (Hiebert& Bennett 1992). Another general effect of microbes is an increase in soil stability.Their binding effect is not as obviousas that of the roots of large plants but is nevertheless measurable. Experimental washing of barren soil slopes with water consistently produces greater erosion than hosing down comparable slopes armored with undisturbed microbial scums (Booth 1941). Individual soil clods (peds) also are more stable in soils inoculated with microbesthan thoseleft alone (Metting1987).Microbes resist soil erosion in several ways. Some microbes promote the hydrolytic comminution and cavernous
Table 10.1 Kinds of microbes, their metabolic requirements and roleinsoils (fromPelczaretal.1986). ~_______
General role
Kind of organism
Example genus
Energy source
Carbon cyclers
Algae
Chlamydomonas
Sunlight (PI
Aerobic
Primary producer in normal-wet soils
Cyanobacteria
Nostoc
Sunlight (P)
Aerobic
Primary producer in normal-wet soils
Purple nonsulfur bacteria
Rhodospirillum
Sunlight (PI sometimes organic compounds ( C)
Methanogenic bacteria
Methanobacterium
H, (C)
Anaerobic CO2 (A) sometimes formate (H)
Aerobic sporeforming bacteria
Bacillus
Organic Organic compounds compounds (0) (C)
Aerobic Organic compounds
Fermenting bacteria
Clostridium
Protoctistans
Amoeba
Nitrogen Nitrogencyclers Exing bacteria Root nodule bacterioids
Denitrifying bacteria
Sulfur cyclers
Organic compounds (C) Organic compounds ( C)
Electron donor
Organic compounds (0)
sometimes H2S(L)
Carbon source
Oxygen relations
Comments
Organic Amphiaerobic Primary producer in compounds swamps (H) sometimes co, ( 4
(W
Organic compounds
Anaerobic Organic compounds
(0)
(H)
Organic compounds
Aerobic Organic compounds
( 0)
(H)
Creatorsof ‘swampgas’ (CH4) Decomposer of organic compounds in normal-wet soils Decomposerin swamp soils Predator in normal-wet soils
Azotobacter
Organic N, ( L ) compounds (C)
Aerobic Organic compounds (H)
Creates ammonium (NH4+) in normal-wet soils
Rhizobium
Organic compounds (C)
Aerobic Organic compounds
Supplies ammonium (NH4+)to host plant
Pseudomonas
Sulfur-reducing Desulfovibrio bacteria
(H)
NO, (C) sometimes organic compounds ( C)
N2 ( L ) sometimes organic compounds
Organic compounds (C)
so42-( L )
H,SorS(L) sometimes organic compounds (0)
Sulfur bacteria
Chromatiurn
Sunlight (PI
sulfur metabolizing bacteria
Thiobacillus
S,FeS(C)
(0)
co, (
Anaerobic 4 sometimes organic compounds (0)
Releases nitrogen (N,) to atmosphere from swampy soils
Anaerobic Organic compounds
Creates ‘rotten egg gas’(H2S) and pyrite in swamp soils
(H)
s,so42-
Pe203(L)
Amphiaerobic Remobilizes co, ( 4 sulfur in poorly sometimes oxygenated organic soils compounds (H)
CO, ( 4
Aerobic
Remobilizes sulfur in normal-wet soils
Trophic groups: A, autotrophic; C, chemotrophic; H,heterotrophic; L, lithotrophic; 0,organotrophic; P, phototrophic.
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weathering of sand and silt grains of the soil to colloidal materials, such as clay, calcium carbonate and iron hydroxides (Fisk et al. 1998), and these act as cements. Mucopolysaccharidesare produced by many microbes to aid motility, and these substances also bind soil particles together (Foster 1981). It is these materials that maintain much of the distinctivefine structure of soils, or ‘tilth’. Nodular and concretionary structures in soils also may be produced by microbes. Laminated surfcia1limestone crusts are formed by coccoid cyanobacteria in the same way as the marine and intertidal structures called ministromatolites (Krumbein & Giele 1979; Reams 1990). Photosynthetic microbes in these crusts promote the precipitation of carbonate by using carbon dioxide dissolved in water, thus neutralizing its acidity. Heterotrophic bacteria also can promote local calcite cementation by release of ammonia or bicarbonate (Castanier et al. 1999).Bacteria thus become encrusted in lime to form carbonate microrods (0.5-2pm; Loisy et al. 1999). Subsurface filamentous structures in caliche nodules also are formed around microbes (Klappa 1979a). Another product of microbial metabolism is the distinctive sand-sized aggregates of tiny balls of pyrite known as framboids (seeFig. 4.6).These are produced in waterlogged soilsby reduction of ferric iron, sulfate and sulfur, coupled with the breakdown of organic matter, by microbes such asDesulfovibrio (Altschuleretal. 1983; Garcia-Guineaetal. 1997). A final distinctivechemical segregation widely attributed to microbes is rock varnish. These thin (usually 20pm, ranging from 2 to 500pm) dark coatings of
amorphous iron and manganese oxides (such as birnessite) are widely known as desert varnish. A more general term is needed because comparable features are found in soils of humid climates, as well as in rivers and lakes. In polished sections, the thin varnish coats commonly are lumpy (Fig. 10.2B), in part reflecting the shape and position of nearby microbes (Krinsley 1998).They also may have fine internal ministromatolitic layering (Fig. 10.2A). Varnish characteristically is developed on the tops but not the bottoms of pebbles and soil peds: a distribution also seen in very ancient examples (Retallack 1997a). Opinions differ on the way in which the iron and manganese is introduced and the exact nature of the microbes involved (Perry & A d a m 1978; Krumbein & Jens 1981; Margulis et al. 1983). There is, however, general agreement that these distinctive structures are biogenic in origin. The activity of microbes also can be inferred from the condition of fossils of larger organisms in paleosols. Large plants and animals are so dependent on microbes that one can reasonably infer that they must have been present wherever there are larger fossils. Trees, for example, depend upon a variety of bacteria within the mucigel surrounding their root hairs to obtain essential micronutrient elements. Many termites and ruminant animals have microbial communities within their guts that allowthe breakdown of chemically inert portions of plant food, such as cellulose. From the presence of large fossil root traces, termite galleries and the cake-shaped coprolites (fossil feces) of ruminants, one could infer that these various kinds of microbes were present (Retallack1990).Thepreservational style of fossils also provides clues to the activity of microbes. Fossil leaves
Figure 10.2 Micrographs of apetrographic thin section(A) and from a scanningelectron microscope (B) of rock varnish from near Phoenix,Arizona, USA (photographs courtesy of JohnB.Adams). Scale bar represents 20pm.
Organisms found in paleosols commonly show a range in decomposition comparable with that found in the leaf litter of modern soils (Retallack 1976). Some leaves are coriaceous, dry and curled, whereas others are wilted and threadbare. The veins of lignin in a leaf commonly are the last to resist microbial decay, and these may be fossilized as a kind of leaf skeleton. The lack of fossil flesh attached to bones, except under unusual circumstances of high salinity,freezing or lack of oxygenation, also is a testimony to the effectiveness of microbial decomposition in soils. Direct evidence of microbes in paleosols is their fossils permineralized in chert (Horodyski & Knauth 1994) or calcite (Trewin & Knoll 1999), or trapped in amber (Poinar et al. 1993). Fossilizable hard parts of soil-dwellingthecamoebans are small (20pm diameter) siliceous or organic, sometimes calcareous, tests, often bedecked with sand grains and other debris (Medioli & Scott 1988). Thecamoebans also live in lakes and the sea, so it is difficult to discern whether their tests, l i e those of diatoms and foraminiferans, might not have blown in from elsewhere, or have been part of the aquatic parent material of a paleosol. There is also the problem that microbial fossils may be contaminants introduced during surface exposure or deep burial (Andersonetal. 1998;Fisketal. 1998).
Fungi Many kinds of fungi can be regarded as microbes, but they are singled out here because of their distinctive
Figure 10.3 Needle-fibercalciteinatubular void remaining from a root trace, forming a septum separatingsepta1filling calcitefrom an area of random needle-fiber calcite,from the lower Carboniferous(262 Ma), HeathersladeGeosol, on the Gully Ooliteat Three CliffsBay,South Wales, UK (from Wright 1986b: with permissionfrom the International Association for Sedmentology).Scale bar represents0.1 mm.
1 33
effect on soils. Fungi have long tubular cells (hyphae) that are larger (2-10pm diameter, <3 cm long) than bacteria (0.5-2 pm diameter, d 0 p m long) but smaller than unicellular algae ( 1 0 4 0pm) and protoctists (30-50pm). Hyphae form a loose network (mycelium) that thoroughly penetrates soil and leaf litter, in addition to forming the solid fungal bodies recognized as mushrooms, toadstools and brackets. Hyphae, mushrooms and spores of fungi are preserved especially well in permineralized peats and coal balls (Taylor& Osborn 1996; Taylor et al. 199 9). Traces of hyphae may be preserved in calcareous paleosols. Many mushroom-like fungi (Basidiomycetes) have hyphae encrusted withneedle-fiber calcite. Irregular grain coatings of needle-fiber calcite preserved in modern and fossil soils (Fig. 10.3) also may be evidence of fungal hyphae (Monger et al. 1991; Alonso-Zarza et al. 1998a). Other kinds of endolithic fungi generate acids that etch an impression of hyphae into the surfacesof bothcarbonateandsilicategrains(Ross&Fisher 1986). Fossil root traces of vascular plants may have associated evidence of fungi. Needle-fiber cement associated with root traces in calcareous nodules may have been precipitated by fungi during fungal decay of the root, or fungal activity associated with the rhizosphere (Wright 1986b; Alonso-Zarza et d. 1998a). Many fungi, particularly those of the groups including yeast (Ascomycetes)and mushrooms (Basidiomycetes),enter into a mutualistic relationship with plant roots known as a mycorrhiza. These can form an external sheath
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(ectomycorrhiza)to roots with some hyphae penetrating the outer (or epidermal) cells of the host plant. A more intimate relationship is found where some of the living cells of the host are invaded by hyphae (endomycorrhiza). Fungi benefit from such mutualistic relationships by gaining carbohydrates and other foodstuffs from the host plant. The plant benefits because fungi are capable of mobilizing important but difficult-to-obtain nutrients, such as phosphorus, nitrogen and potassium (Simard et al. 1997; van der Heldjen et al. 1998). Fossil endomycorrhizae have been found in silicified plants and may have played a role in nurturing the most ancient landplants (Simonet al. 1993;Taylor&Osborn 1996).Mycorrhizal roots are very different in morphology from those with which fungi and other microbes are less intimately associated. They are stouter and more abruptly ending, branch at shorter intervals and at higher angles than ordinaryroots (Richards 1987), and can berecognizedinfossilroots (Pitt etal. 1961). Some resting stages of fungi also are distinctive.Small woody balls with vermiform hollows common in peats and coals are similar to sclerotia (see Fig. 4.6), the resting stages of fungi (Moore et al. 1996).Enigmatic small balls and tubes of radiating calcite crystals found in modern and ancient caliche (ichnogenus Microcodium) have been regarded as casts of fruiting or resting stages of soil fungi or lichens (Klappa 1978), but may also be calcified root cells (Wright et al. 1995; Alonso-Zarza et al. 1998a). Effectsof fungi also can be recognized in fossils associated with paleosols. Impressions of leaves in fossil leaf litters commonly are thin and flaccid, as if internally decayed. Fine networks of creases on fossil leaves may represent hyphae, and masses of leaves closely adpressed together in fossil leaf litters are similar to those bound by fungi on modern soils (Retallack1976). Some forms of fungal decay of wood, such as dry rot or pocket rot, alsoaredistinctiveinpermineralizedwood(Taylor& Osborn 1996).
green algae, such as Trebouxia or Pseudotrebouxia, or a cyanobacterium, such as Nostoc. A much greater diversity of fungi and actinobacteria have entered the lichen partnership, and lichens are properly classified from this phycobiont component. Unlike the algal or cyanobacterial part of the lichen, these are morphologicallydistinct fromfree-living fungior actinobacteria (Hale 1983). Lichens grow on bare rock or tree trunks and have been regarded as pioneers in the early successional colonization of new land surfaces. They are, however, very slow growing. Studies of long-term lichen growth on boulders of till abandoned by retreating glaciers have shown that lichen colonies grow in diameter 0.10.03 mmyr-l (Calkii & Ellis 1984).Thus, lichens differ from early successional plants, which grow and reproduce quickly. Lichens are better regarded as organisms able to withstand conditions hostile to most other forms of life(Grime 1979). Traces of lichens may be preserved in calcareous soils and paleosols (Klappa 1979b). Individual thalli of lichens are patchy in distribution, and the acids generated by lichens form solution pits in limestone bedrock or around soil nodules exposed at the surface. This miniature karst topography may show outlines of the thallus. Lichens also may be encrusted to form a banded structure called a lichen stromatolite. Their laminae are somewhat irregular and less laterally extensive compared with those of stromatolites formed by films of cyanobacteria. The banding of lichen stromatolites may be interrupted by small channels created by rhizines (rootlikebundles of hyphae). In sandstone, endolithic lichens produce distinctive leached zones under a surfxial silicified crust that is prone to flaking (Friedmann & Weed 1987). Endolithic lichens also contribute to the spongy weathering of liestonecalledblackphytokarst(Folketal. 1973;Viles 1987). Other distinctive patterns in sandstones may also be fossil lichens, but silica-permineralization is needed to see definitive cellular details (Retallack 1994c;Taylor et al. 1997).
Lichens These hardy organisms are plant-like in their sessile, somewhat leafy bodies, but they are not true plants. They are a symbiotic association of fungi or actinobacteria, and unicellar photosynthetic microbes. In most lichens the photosynthetic organism (phycobiont) is a
Liverworts and mosses These nonvascular plants are both bryophytes. but are distinctfromeachother (Fig.10.4).Mosses have slender axes with numerous small, helically arranged leaf-like structures. Liverworts, on the other hand, are flat
Organisms
1 35
Figure 10.4 Afossilliverwort (A)Marchantites tennanti andmoss (B) MuscitesguelescinifromtheUpperTriassic(225Ma),Molten0 Formationnear BirdsRiverSiding,SouthAfrica.Scalebarsrepresent 1cm (AisBP/2/4658aandB isBP/2/4722,photographs courtesyof H. M. Anderson).
thallose plants, resembling aquatic algae more than other kinds of land plants. Both mosses and liverworts have arudimentary system of cellswith thickened walls (hydroids) for water transport, unicellular hairs (rhizoids)that function as roots, and thin patches of cells (in liverworts) or stomates (mosses and anthocerote liverworts) that allow gas exchange. They lack the waterconducting tracheids and true roots of vascular land plants (Kenrick & Crane 1997). Mosses and liverworts live by aerobic photosynthesis like most plants. Although some species tolerate arid sand dunes and deserts, mosses and liverworts are most abundant and diverse on moist forest floors, and in bog soils of high latitudes and altitudes (Richardson 1981).Sphagnum moss, for example, forms thick peats, which have low bulk density and high water retentivity compared with woody peat. Sphagnum can formHistosolsevenover rock outcrops or on moderately sloping ground. The weathering of minerals under waterlogged soils is somewhat limited,but the strongly acidic pH under Sphagnum bogs can promote podzoliation in moderately drained sites. More obvious evidence of mosses and liverworts is their preservation in carbonaceous shales and permineralizedpeats of Histosols (Anderson&Anderson 1985; Smoot &Taylor1986;Edwards et al. 1995a).
Vascular plants Vascular plants are named for their conducting tissues of woody, elongate cells whose walls bear distinctive
helical or annual thickenings (tracheids), and which form the veins of leaves and the wood of trees. Vascular plants are not just woody plants, however. They include most plants of common human experience on Earth today, ranging from tiny filmy ferns to giant redwoods. Like mosses and liverworts, vascular plants photosynthetically reduce carbon dioxide to organic matter, using chlorophyll as a catalyst and releasing oxygen to the air. These metabolic processes of leaves and aerial shoots depend on the energy of sunlight, and are not found in subterranean root systems. Roots gather water and nutrients, and their respiration releases carbon dioxideinto soil solutions to create carbonic acid. These and other chemical effectsof roots dominate the weathering of most modern soils. Different kinds of vegetation vary in their weathering effects, as well as in their characteristic patterns of root traces and of soil horizons. Such observations have been gathered together in a separate and following section concerning the interpretation of ecosystemsfrom paleosols. Only those features of soils and paleosols pertinent to interpreting the nature and distribution of individual plants are addressed here. A dramatic illustration of the role of an individual plant in soil formation is the peculiar overthickened eluvial horizon called a ‘basketpodzol’,found in some Spodosols (Retallack 198 1).The term ‘podzol’is meant here in its original Russian sense of ‘under ash’, for the white, quartz-rich eluvial horizon that is thickened under individual trees. Studies of the chemical variation in soils
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under individual trees (Zmke 1962) and experimental work establishing the rapidity and mechanisms of podzolization (Liindstrom et al. 2000) are indications that eluvial horizons of basket podzols are leached of clay, iron and other materials through the action of a variety of chemicals produced by plants, particularly phenolic compounds washed out of leaves by rain. Some cryptokarst features, such as rounded solution hollows,may be a comparable phenomenon. These basin-lie depressions are corroded into limestone bedrock beneath an individual plant because of the acid generated by respiration of roots and associated microflora (Jennings 1985). Like fossilized stumps, basket podzols and cryptokarst solution hollows can be evidence of the spacing of trees in ancient forests (Retallack 1976; Decourten & Bovee 1986). Another structure that can be an indicator of forest density is the disturbance of soils associated with windthrows. Most trees have a radiating array of sudcial roots as well as some deeply penetrating ones. When uprooted by wind, a hemispherical mass of soil is supported by roots vertically in the air, and then forms a mound as the roots decay (Schaetzl&Follmer1990,Norman et al. 1995).The pit and mound in the soil surface created in this way is called a ‘cradle knoll’. Irregular surface horizons of many paleosols could be a reflection of such windthrows. Similar features could be produced by the catastrophic clearing of vegetation by a mudnow or flood deposits, but the holes remaining in the soil after such violent uprooting are more symmetrical than usual for cradle knolls and also are prone to additional scouring before the energy that uprooted the trees is spent. Fossil stumps are in some cases permineralied within the soil in which they grew (Ammons et al. 1987). Indifferentlypreserved fossil tree trunks also can be guides to spacing and structure of ancient forests (Calder et d. 1996; Pole 1999). Subcylindrical calcareous nodules near the surface of paleosols may represent stumps degraded by termites and other decomposers (Kraus 1988; Retallack 1991a;Retallacket al. 2000). Calcrete deep within soils can be formed by tree roots (Rossinsky etal. 1992). Flatworms and other small soft-bodied creatures A variety of multicellular, small (<2 mm), soft-bodied
creatures contribute to the stability of soil peds and may play an important ecological role in soil communities but leave few distinctive traces in soils. In this group are flatworms (Platyhelminthes), rotifers (Rotifera). gastrotrichs (Gastrotricha) and ribbon worms (Nemertina); all have a very limited fossil record (Gray 1988). Nematodes These worm-like creatures are extremely abundant and widespread in soils. Their lifestyles include predators, microbial feeders, omnivores, herbivores and plant parasites. The majority of nematodes are less than a few millimeters long. They serve to circulate organic matter within the soil. Their movement on thin films of water plasters over the surfaces of peds and so stabilizes soil structure. Even large nematodes tend to move by wriggling through cracks in the soil rather than by creating distinctive burrows (Poinar 1983). Nematodes leave sinuous surface trackways, but insects also leave very similar trace fossils (Metz 1987). Galls formed on leaves, twigs and roots of vascular plants are known as fossils, but these also are created by insects (Boucot 1990; Labandeira 1998).Nematode galls on plants are distinguished by their ‘giant cells’ uones 1981). Damage to fossil arthropod cuticle also has been blamed on nematodes, and fossil nematodes themselves are found in amber(Gray 1988;Poinar1992).
Mollusks Brackish adapted mollusks, including oysters (Boucot 1990) andmussels (Retallack&Dilcher 1981a,b),have been found in mangrove paleosols formed within the intertidal zone. On land the principal soil mollusks are snails and slugs. Snail shells in paleosols can be useful paleoenvironmental indicators (Solem & Yochelson 1979; Pickford 1995). Alsofossilizableis the distinctive slime-bordered, flat, undulating trail of snails and slugs, reported in marine sediments (Aulichnites, Scolicia and allied ichnogenera in Hantzschel 1975). The scrapedout feeding trails made by snails feeding on endolithic lichens in limestone bedrock (Shachak et al. 1987) also could be potentially recognized in paleokarst. Slugs and snails eat other animals, fungi, algae and leaves. Fossil leaves with nibble marks partly repaired with callus by
Organisms
13 7
the plant (Phagophytichnus of Hantzschell9 75;Labandeira 1998)could equally be produced by herbivorous insects and their larvae, as by snails and slugs. Annelids Segmented worms (Annelida) are common in soils but have a sparse fossil record (Morris et ul. 1982;Gray 1988),and are best represented by clitellate cocoons (Manum et ul. 1991). Best known are earthworms (Lumbricidae)and potworms (Enchytraeidae). Leeches (Hirundinea) are found within leaf litter, but most leeches and leech fossils are aquatic (Gray 1988). The brown-to-pink lumbricid worms of calcareous, grassland and garden soils commonly are 2-5 mm in diameter, but some, such as Megascolides australis of Australia, attain diameters of 2 cm and lengths of 1.4m (Lee 1985).Earthworms feed on decomposingleaf litter and other organic detritus in and on the soil. They may gather surficial mounds (‘middens’)of leaves that they particularly favor. They also ingest soil as they excavate their burrows, and this is intimately mixed with organic matter in their ellipsoidal fecal pellets (Rusek 1985). Small (0.5-2mm) flask-shapedto circular bodies of calcitealsoareexcretedinthefecesof earthworms (Darwin 1881).The granular and crumb structure of grassland soilsis thought to be largely a product of earthworm activity (Pawluk & Bal 1985).Fecal casts at the surface of the soil may form tall (to 15 cm high, 4cm diameter) erect pipes or slurry-like masses. Fresh earthworm burrows are near-circular in outline, with scattered eltipsoidalfecalpelletsatthe bottom(Bar1ey 1959;Westetal. 1991).Once abandoned, the burrows are flattened and deformed into a digitate structure (Fig. 10.5).The walls may be lined with clay, organic matter or iron hydroxides. Both collapsed and intlated burrows have been found in paleosols (Retallack 1976;Meehan 1994), and several ichnogenera (Oligichnos, Eduphichniurn of Bown & Kraus 1983) accommodate earthworm burrows. Enchytraeids are generally white in color and smaller (0.1-5cm long) than earthworms. They are not, as is commonly supposed, young earthworms. Whereas few species of lumbricids can tolerate a soil pH <4, enchytraeids are often abundant in acidic woodland litter and peat (O’Connor1967).Like lumbricids,potwormsmove by peristaltic action and ingest bacteria, fungi, decom-
Figure 10.5 Sketchof petrographic thmsectionof afossil earthworm burrow, includingfecal pellets. The specimenis from the surface (A) horizon of theTurrimetta clay slightly erodedvariant paleosol,aPibrist, in theLowerTriassic (245 Ma),basal Newport Formationsouth of Bilgola, New SouthWales, UK (fromRetallack 1976;reprintedwith permission from the GeologicalSocietyof Australia).
posing plant fragments and mineral soil. They do not leave well-defined burrows but use existing soil cracks. Their cocoons are covered with sand grains and organic debris cemented by mucus and are potentially fossiliiable. Potworms contribute to the enormous numbers of very small (120-200pm)fecal pellets found in some soils,but their feces aredifficult to distinguish from those of mites and springtails (Rusek 1985). Velvet worms These creatures are now a minor component of soil faunas and have been placed in their own phylum (Onychophora),but their long fossil record (Ramskold& Chen 1998) and molecular biological relationships (Ballard et al. 1992)are evidence that they are basal arthropods. Their activities leave little trace in soils. They are predators, capturing and then externally digesting insects and other creatures by means of slime strands. They have soft flexible bodies, susceptible to desiccation, and so are found mainly in leaf litter or mosses of woodland soils in humid regions. Some speciesretreat to burrows during the heat of the day or a dry season, but theseareshallow cracksin thesoilrather than elaborate constructions (Endrody-Younga& Beck 1983).
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Water bears These small rotund creatures with paired appendages have been placed in the phylum Tardigrada, but like velvet worms, are probably basal arthropods (Kinchin 1994). Some tardigrade fossils were large, but most liv. ing water bears are microscopic (50-1200 p)Tardigrades pierce plant cells, rotifers and nematodes with their stylets and suck out their body fluid as food. They have remarkable powers of resistance to extremes of desiccation,of temperature (withstanding a range from 1 51"C to -2 70°C) and of X-rays (surviving 24 h exposure to 5 70 000 roentgens). They can become dormant for as long as a hundred years, in a little modified barrelshapedform (Kichin 1994).These tuns are potentially fossilizable in organic soils. Crustaceans Barnacles, lobsters and shrimps are abundant and diverse in oceans and lakes but also common in mangal and swamp soils. The considerable diversity of crustaceans found in mangal soils is reflected in the great variety of burrows and other traces found in them. Barnacles may beinevidencefromsegmentedcircularscars they leave on roots, stems and rocks (Boucot 1990). Ghost shrimps and crabs excavate complex branching systems of burrows (ichnogenus Thalassinoides), in some cases distinguishedby finely striatedwalls (Spongeliornorpha) or walls lined with small balls of clay (Ophiomorpha; Frey et al. 1978). Crabs and shrimps also excavate inclined tubes with ovoid basal living chambers (ichnogenera Psilonichnus, Pholeus and Macanopsis), and tubular helical burrows (Gyrolithes; Hantzschel1975).Thefecalpellets of crabs andshrimps also may be distinctive in their regular arrangement of internal cavities(Favreina and other coprolite ichnogenera; Boucot 1990). Crustaceans found in freshwater, waterlogged and seasonally flooded soils include cladocerans (Cladocera), copepods (Copepoda), ostracods (Ostracoda), sand-hoppers (Amphipoda) and crabs and crayfish (both Decapoda: Gray 1988; Babcock et al. 1998). Extinct lipostracan branchiopods (Lepidocaris:Fig. 10.6d) have been found in permineralized peaty paleosols (Histosols). In the drainage of the Mississippi River of North America as far north as the Canadian border,
5cm
,
Figure 10.6 Reconstruction of semiaquatic and terrestrial arthropodsfromtheLowerDevonian(Seigenian.410Ma) Rhynie Chert of Scotland.The plants includean extinct nonvascular stomatophyte (A, Aglaophytonmajor)and a vascular zosterophyll(B, Asteroxylon mackei),and the animals includeapachygnathidmite(a,Protacaruscrani), trigonotarbidspider (b,Pdaeocharinus sp.),liphistidspider (c, Pdaeoctenizacrassipes),branchiopodcrustacean(d. Lepidocaris rhyniensis),collembolan(e,Rhyniellapraeciirsor)and a possible thysanuran (f, Rhyniognatha hirsti) (fromKuhne& Schliiter 198 5; reprinted with permission from EntomologicaGenemles).
tunnels of crayfish (such as Procambarus gracilis and Orconectes imrnunis) in soils may extend as deep as 5 m into the permanent water table (Thorp 1949). The entrance to these burrows is marked by a short chimney formed of balls of clay Similar structures have been reported from paleosols (ichnogenera Carnborygrna, Scaphichniurn;Hasiotiset al. 1993). Woodlice (Oniscoidea: Isopoda) also are marine crustaceans, but many of them live on dry land. They are prone to desiccation once out of water. Many species roll into a ball or cluster together to block their pseudotracheae duringdry conditions. Becauseof this sensitiv-
Organisms ity to dry conditions, woodlice are most common in organic horizons of grassland soils or the leaf litter of deciduous woodlands. They are omnivorous, feeding on dead wood, fungi, leaves and carcasses of insects and mammals (Warburg 1993). Some woodlice form complex burrows (Shachak 1980). Fossil trails have been attributed to terrestrial isopods (Brady 1947).The exoskeleton of woodlice is hardened by biogenic accumulations of calcium carbonate like that of other crustaceans, and can be persistent as bleached skeletons at the surface, in paleosols (Morris 19 79) and in lake beds (Gray 1988).
Millipedes and centipedes Millipedes (Diplopoda)have two pairs of legs for every segment and are herbivores and detritivores (Hopkin & Read 1992).The pill millipedes (Glomerida)are ecologically l i e woodlice,which they resemble. Flat-backmillipedes (Platydesmida, Polydesmida and Chordeumida) also are restricted to moist leaf litter and organic soils. Extinct millipedes such as Archidesmus (Fig. 10.7) may have been similar. One group of flat-backed millipedes (Polydesmida)produce distinctive small (2.5mm diameter) disk-shaped feces of clay and sand, with a deep
Cambmpodusgractlts (Middle Cambrian)
Waukesha myriapod (Early Silurian)
Stonehaven millipede (Late Silurian)
Figure 10.7 Reconstructionsof early Paleozoic myriapods,from top Curnbropodus grucilis (Cambrian)a scutigerimorph centipede (DelleCave&Simonetta1991). unnamedWaukeshafossil (Silurian)a lithobiomorphcentipede (Mikulicet ul. 1985). unnamedStonehavenfossil(Si1urian) an earlypolyzoniidmillipede(Almond 1985), Kurnpecurisforfurensis (Devonian)an extinct aquatickampecarid(A1mond1985), Archidesmus rnucnicoli (Devonian)an early platydesmidmillipede(Almond1985), 'Kurnpecuris' tuberedata (Devonian)an early platydesmidmillipede (Almond19 8 5), unnamed Battery Point fossil (Devonian)an early spirobolidmillipede (Shearet ul. 199 6).
Kampecafis fodamnsrs (Early Devonian)
''
139
Archldesmusmacntcolr (Early Devonian)
"Kampecaris"tubemulafa (Early Devonian)
Battery Point millipede(Early Devonian)
-
5 mm
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impression down the middle from their anal flaps (Kiihnelt 1976).These are sometimes lined with clay or organic matter from the enclosing peritrophic membrane (Rusek 1985).Noneof these flat-backmillipedes areefficient burrowers. If they burrow at all, they tend to use existing spaces in the soil. In contrast, many round-back millipedes (Spirobolida, Stemmiulida, Polyzoniida, Siphonophorida, Iulida, Spirostreptida) are active burrowers, and such a habit can be inferred for comparable fossils (Stonehaven and Battery Point millipedes of Fig. 10.7). Their burrows are round in cross-section, l i e their bodies, and lined with a thin rim of pressed clay (Paulusse & Jeanson 1977). The fecal pellets of some round-back millipedes (Iulida) are cylindrical and have concave and convex ends where cut off by anal flaps (Kiihnelt1976). Millipede trackways are also distinctive (Johnson et al. 1994).Millipede skeletonsare sometimes found in paleosols (Leakey 1952; Jell 1982), presumably because of unusually strong biomineraliation. The exoskeleton of some millipedes has up to 5 5 wt % calcium carbonate and calcium phosphate, and only 34wt% chitin and lOwt% protein (Neville 1975). Other millipede fossils are found in nearshore and lake beds (Hannibal1997) and amber (Poinar 1992). Centipedes (Chilopoda)have only one pair of legs per segment and are fast-movingpredators. They live in leaf litterandbarkof trees, aswellasinthesoil(Lewis 1981). They use existing soil cracks and burrows rather than excavating distinctive burrows of their own. An allied group are 'glasshouse centipedes' (Symphyla). These small white myriapods feed on living roots, decaying plant matter and microorganisms.The pauropods (Pauropoda) are another centipede-like group of animals with branched antennae and only 9-10 pairs of legs. They eat decaying plant matter and are found in loose, loamy grassland soils (Kiihnelt 1976). Neither 'glasshouse centipedes' nor pauropods are known to leave distinctive traces in soils. The fossil record of centipedes is poor, and surprisingly, includes fossils in marine rocks (Cambropodus and Waukesha myriapod of Fig. 10.7). Horseshoe crabs Chelicerates, including horseshoe crabs (Xiphosura), are marine. Horseshoe crabs congregate on beaches to spawn, and small extinct horseshoe crabs may have
ventured high into the tree canopy (Fisher 1979). Like aglaspids, eurypterids and other extinct arthropods, horseshoe crabs may have been part of some salt marsh and mangal communities (Stmmer 1977; Gray 1988).Fossil trackwaysattributedtothesecreaturesare known from marine and lacustrine deposits (Boucot 1990). Scorpions Most scorpions (Scorpiones) live in cracks in the soil, but some large tropical species excavate sloping burrows up to 80 cm long (Kiihnelt 1976). Fossil tracks similar to those of scorpions have been found (Brady 1947). Many fossil scorpions have external gills, and were probably aquatic (Kjellesvig-Waering1986;Shear etal. 1996). Spiders Spiders (Arachnida) are active predators and most of them live above the soil, but several species excavate burrows. Australian trapdoor spider (Anidiops villosus) burrows have round lids (Main 1985). Other trapdoor spider burrows have stout inner doors, socks and chambers. Their silk linings are unlikely to be preserved in most paleosols, although fossil cocoons and spider webs areknownfromamber (Boucot 1990). Mites Most mites (Acarini) are microscopic (Fig. 10.6a), but these are the most abundant and diverse group of soil animals. Different kinds of mites parasitize and eat live prey, feed on plants, fungi, bacteria, algae and lichens, and decomposedead organisms. They leave few obvious traces in the soil other than small (140-200pm). smooth, ellipsoidalfecalpellets(Rusek 1985), and these can be difficult to distinguish from pellets of other small soil creatures (Labandeira et al. 1997; Labandeira 1998).The fossil record of mites is sparse (Gray 1988; Boucot 1990). Springtails and other wingless insects A variety of silverfish-lie insects are found in soils (Thysanura, Diplura, Protura). Different species eat live
Organisms prey, fungi, algae and decaying vegetation. They live in crevices in the soil, leaving small (30-90~m),irregularly formed, fecal pellets (Rusek 1985).The springtails (Collembola)are the most abundant anddiverse of these wingless insects. Their common name refers to a large spine in their tails, which can be flexed to effect a quick escape. They are small (<2 mm) and notable for their ability to tolerate climatic extremes ranging into arid and polar soils, where few other soil organisms are found. Collembola also are the most ancient fossil insects known (Fig. 10.6e).
Butterflies and moths Lepidopteran galls and mines in leaves and bark fragments have been found as fossils (Crane&Jarzembowski 1980),althoughotherinsects also formgalls andmines (Labandeira 1998). Fossil twig-covered cases of the larvae of case-worm moths (Lepidoptera, Psychidae) have been foundin amber (Boucot 1990).
Lacewings Most larvae of this group of predatory flies (Neuroptera) are aquatic, but one group (Myrmeleontidae)has larvae that construct funnel-like pits in the soil. These ‘ant lions’ lie in wait, buried at the bottom of the funnels, for their prey to stumble into the loose walls (Kiihnelt 1976). Such structures are potentially preservable in paleosols.
Caddisflies Most of these insects (Trichoptera) have aquatic larvae that construct larval and pupal cases from pebbles, shells or whatever is at hand. Fossil cases of this kind are common in lake deposits (Williams 1988; Boucot 1990). A few modern species of caddis fly larvae are completely terrestrial and live within leaf litter. Some of the fossil cases made completely of leaves (ichnogenus Folindusia) could have been of this kind.
Crickets Most crickets, grasshoppers and related insects (Orthoptera) are herbivores that leave little trace in the soil.Their feeding, however,can be obvious among fossil
141
plant remains (Beck & Labandeira 1998).Two groups, the mole crickets (Gryllotalpidae) and pygmy mole crickets (Tridactylidae),burrow in moist soil near lakes and streams (Ratcliffe & Fagerstrom 1980). Burrows of pygmy mole crickets have simple tear-shaped living chambers either at the end of vertical passages (likethe ichnogenus Macanopsis;Hiintzschell9 75) or terminating horizontal galleries, similar to those made by bugs, cicadas, beetles and wasps.The burrows of mole crickets have spacious (1cm diameter) galleries that branch and join to create a n irregular network just below and parallel to the ground (similar to the ichnogenus Protopalaeodictyon;Hantzschell9 75). Bugs Among the bugs (order Hemiptera), there are a few kinds that burrow: shore bugs (Saldidae), toad bugs (Gelastocoridae) and burrower bugs (Cydnidae). Most of these burrows are simple inclined shafts with a large living chamber at the bottom (ichnogenus Macanopsis; Hantzschell9 75). The burrower bugs do not dig access shafts to the surface. Their burrows are just under the surface of the soil, under stones or logs, in sand or in mold near the roots of tussock grass. Often the roof of the burrow is cracked through to the surface and partly collapsed after their passing (Ratcliffe & Fagerstrom 1980). Cicadas (suborder Homoptera, superfamily Cicadoidea) are locally abundant soil burrowers. Mature cicada nymphs excavate long, straight burrows directly up toward the surface from their subterranean root-feeding areas at depths of 25-50cm into the soil (Ratcliffe& Fagerstrom 1980). The larvae and nymphs spend many years (upto 1 7 in some species) at this level deep within the soil creating burrows with conspicuous backfills (ichnogenus Taenidiurn; O’Geen & Busacca 2001). During the spring of some years, they emerge in such great numbers that the soil is riddled with simple vertical burrows. Similar behavior may be of great antiquity among insects (ichnogenus Skolithos; Retallack1976). Scale insects (Coccoidea)are bugs, and their characteristic waxy patches may be preserved on fossil leaves (Boucot 1990). Aphids (Aphidoidea) also are bugs, which form galls on leaves and twigs that also may be fossilized (Scott etal. 1992).
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Beetles Beetles (order Coleoptera) are the most diverse group of insects associated with soil, but only a few kinds of beetles leave distinctive burrows in the soil (Ratcliffe & Fagerstrom 1980). For example, adult tiger beetles (Cicindelidae)are predators. They excavate simple shallow burrows in order to retreat from nightly or winter cold or from rainy or hot weather. Their larvae also are predatory and excavatesimple burrows that may be vertical, and have right-angled bends to subhorizontal galleries. They range in depth from 2 to 125cm. A distinctive feature of these burrows is their meniscate backfills of clay and silt (ichnogenus Taenidium: Hantszchel 1975). Another distinctive feature of tiger beetle burrows is ellipsoidal pupation chambers scattered along the burrows. Rove beetles (Staphylinidae)feed on algae, decaying vegetation and other insects. Algal-feeding species are especially conspicuous in sandy soils of beaches and streamsides.Theirburrows may be complex,with multiple entrances and chambers situated above and doubled back over their entrance burrows. Ground beetles (Carabidae) are a large group of predatory burrowing beetles that make simple vertical and horizontal burrows. They also make a distinctive kind of burrow, with radiating and branching subterranean galleries (ichnogenus Megagrapton: Hantzschel 1975). Both larvae and adults of variegated mud-loving beetles (Heteroceridae) excavate horizontal galleries that meander and branch just below the surface of soils. These herbivorous beetles are especiallycommon along the shores of streams and lakes. Scarabs (Scarabaeidae) construct simple vertical or inclined burrows (ichnogenus Macanopsis) in which to spend anight or to over-winter,or simplevermiform burrows in dung pats (Chin & Gill 1996).Burrows made to provision larvae are more elaborate and varied (ichnogenera Coprinisphaera, Devincenzichnus, Madinaichnus, Martinezichnus, Monesichnus, Teisserei, Rebuffoichnus: Johnston etul. 1996;Genize & Laza 1998;Duringeret al. 2000: Genizeet al. 2000).Freshleaves,decayingleaf litter, dung and humus are buried by various species, each with specific dietary preferences. These food reserves support the development of larvae, which hatch from eggs buried alongside the food and also pupate under-
ground.The young adult beetles emerge at the surface to feed. The behavior of scarabs can be considered in order of increasingly elaborate parental care. Little care is invested in future generations by small dung beetles such as Aphodius. Their life cycle is short and egg production copious enough that a few eggs are liely to survive to adulthood in a large dung pat. Other species, such as Dichotomius carolinus (Fig. 10.8A), pack an egg, food supply of dung and partition of soil successively, end to end, in simple near-vertical burrows. More elaborate still are burrows made by dung beetles such as Geotrupes stercorarius (Fig. 10.8B). This speciesexcavates a long vertical burrow beneath a large pat of dung. Working back toward the surface, successive lateral chambers are constructed. Each one is supplied with an egg and a large wad of dung before being sealed with earth. Trace fossil burrows of this kind have been referred to the ichnogenus Pallichnus (Retallack 1984). Other species, such as Phanaeus palliatus, not only bury dung and eggs in subterranean chambers, but also sculpt a thick clayey rind around both (Fig. 10.80),to deter predators and parasites. Trace fossils of this kind have been referred to the ichnogenus Coprinisphaeru. Other beetles cut spheres of dung from a large dung pat, which are then rolled away for shallow burial. This behavior of Scarabaeus sacer was seen by ancient Egyptians as a metaphor for the rising and setting of the Sun, and a sign of good luck that the Universewas still in operating order. Beetle activity can be inferred from other fossils associated with paleosols. Many kinds of beetles (Scolytidae, Buprestidaeand Cerambycidae)bore into wood. Similar borings and fecal pellets have been found in permineralized wood (ichnogenera Dekosichnus, Palaeobuprestus, Palaeoipidus, Palaeoscolytus, Xylonichnus; Genize & Hazeldine 1995: Labandeira 1998). The engraver beetles (Scolytidae)make a long central gallery occupied by the female beetle. Radiating galleries are excavated by numerous larvae as they feed outward from the central gallery. This pattern of boring is especiallyobviouswhen confined to bark, because this may break open to reveal the entire pattern (Boucot 1990). Carrion beetles (Dermestidae)are effective in defleshinglarge carcasses. Colonies of these insects are maintained by museums and biologicalsupplyhouses for cleaning skeletons.Dermestid beetle activity in paleosols can be recognized from bored fossil bones (Kitching 1980) and from burrowsfilledwith bone chips (Saik 2000).
Organisms
143
A. Oichotomius carolinus D. Phonoeus polliotus O&OCrn
dung
B.
pocked earth
soil
Geotrupes stercororius
Figure 10.8 Brood burrows of dung beetles (Coleoptera,Scarabaeidae)below the dung they use to provision their larvae: (A) Dichotomiuscurolinus, fromcuernavaca,Mexico;(B) Geotrupesstercorurius fromnear Saalfeld.East Germany:(C) Coprisarmutus from Salazar,Mexico: (D) Phunueuspdliutusfrom Ocoyoacac,Mexico (A, C, D adapted from Halffter & Matthews 1966: B from von Lengerken 1954).
Termites Large organized societies of termites (Isoptera) have a profound effect on the decomposition of wood, humus, lichens and fungi in soils of tropical to warm temperate climates. Earthen termite mounds up to 1 0 m high dot the landscape of many tropical grasslands. The internal structure of these termitaria can be complex, including a centrally located system of chambers for the queen and larvae, large fungus gardens, and extensive passageways for ventilation. These kinds of structures are striking when preserved in paleosols (see Fig. 9.9; ichnogenera Archaeoentornonichnus, Terrnitichnus: Bown 1982;Hasiotis&Dubiel1995). Apart from organized structures of termitaria, soils and paleosols also may be riddled by a network of galleries leading deep into the soil to water, or ramifying in all directions to sources of organic matter. Many of the small spherical micropeds that dominate many tropical soils and paleosols are pellets of termites. Moundbuildingtermitesfashion small (12 5-750 pm)spherical to ovoid pellets from earth and saliva, and expel them as feces (Mermut et al. 1984). In contrast, fecal pellets of dry-wood termites are short rods, with a hexagonal cross-section (Rohr et al. 1986; Rosefelds & de Baar
1991). Termites are unique in having extremely alkaline midguts (pH 11-12.5, Brune &Kahl1996),which has the effect of strongly desilicating their sesquioxiderich fecal pellets. Termite nests are large structuresthat alter the physical properties of soils.Their presence may be discernible long after they were invaded and torn apart by aardvarks or reworked by burrows of other soil animals. The aerating effect of termite galleriescan result in a n appreciably more alkaline soil pH within a termitarium, compared with the soil beyond. In some tropical countries, caliche from the center of termite mounds is an important local source of agricultural lime fertilizer (Thorp 1949). Ants This group of social insects (Hymenoptera, Formicidae) also produce complex nests. Harvester ants (such as Pogonornyrmex rnaricopa) create large mounds, up to 76 cm high, that consist of particles of sand and granule size, coarser grained than the surrounding soil (Cole 1968). Extensive networks of burrows attributed to ants (ichnogenus Attaichnus; Hasiotis & Demko 1996a) and colonies of fossil ants associated with fossil leaves
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(Wilson &Taylor 1964) have been found in paleosols. Ants also have a modest fossil record in shale and amber (Agosti et al. 1997). Like those of termites, the nests of ants aid in mixing of organic matter, aeration, circulationof WaterandraisingthepHof soil(Ho1e 1981). wasps Many of these predatory social insects (Hymenoptera, Superfamilies Pompiloidea, Vespoidea and Sphecoidea) excavate burrows for their larvae. They bury their eggs with an anesthetized insect or spider, on which the hatchling larvae feed until pupation. The larval cells of wasp burrows are generally tear-shaped. They terminate a simple inclined entrance burrow or are arranged laterally along the entrance burrow (Ratcliffe& Fagerstrom 1980;Bown 1982).More complex andcopiously branching nests are made by more highly social species (Evans&Eberhard 19 70). Wasp nests constructed above ground also have been preserved in paleosols. Leaf galls known from fossil leaves (Scott et al. 3992; Labandeira 1998) are created by wasps, flies and aphids. Clay nests (ichnogenus Chubutolithes) like those of mud-dauber wasps (Auplopus militaris) also have been found fossilized (Freeman et al. 1991). Complex multicellular carton nests of the kind constructed by highly social wasps also have been found (Wenzell990). Bees Bees (Hymenoptera, Apoidea) can be regarded as specialized hairy wasps that feed their larvae with pollen and nectar collected from flowers or with other glandular secretions of plants. Burrows of solitary bees are similar in overall plan to those of wasps, but can be distinguishedby two distinctiveand fossilizable features: an elaborate and highly polished wall to the larval cells and clayey covers with concentric or helical grooving (Fig. 10.9).These also are features of multicellular nests of bees (ichnogenera Celliforma, Ellipsoideichnus, Palmiruichnus, Kose~lichn~s: Genise & Born 1996). Simple shallow burrows are made by solitary bees, and more complex structures by social species, but there are exceptions to this generalization (Elliott & Nations 1998;Genise& Hazeldine 1998). Other evidence of bees associated with paleosols in-
Figure 10.9 Fossil bee brood cells (CellijortnuspirlJer) showing spiralclosures,from theMiddle Eocene (48 Ma),Bridger Formation,near Mountainview,Wyoming, USA (Retallack photographof typespecimens of Brown 1934, Smithsonian Institutionno.340878).Scale at bottomisinmillimeters.
cludes fossil leaves with neatly cut circular holes of the kind made by leaf cutter bees (Boucot 1990), and shallow borings in wood, like those of xylocopid bees (Hasiotis et al. 1996). Fossilized wax honeycomb of the kind made by social honey bees also has been found in cave earth (Stauffer 1979). Fish A variety of fish fry are found within mangal soils. They persist in small shallow pools and burrows as the tide goes out. Some fish, such as the Australian mudskipper (Periophthalmusgracilis,among others), climb out of the water andonto the exposedmudtlat andmangroveroots (Nursalll981). Fish may also be important in swamp and marsh soils (Gray 1988). In the black water (varzea) swamps of Brazil, fish eat copious quantities of large seeds and aid dispersal of these plants (Gottsberger 1978). Lungfish burrow into the bottoms of seasonally dry lakes. Their lariat-shaped burrows are distinctiveand known in very ancient lakeside paleosols (Boucot 1990). Amphibians Spadefoottoads (Scaphiopushammondi)bury themselves deeply in desert soils for extended periods (about 9
Organisms months) while conditions are dry (Ruibal et al. 1969). They excavate simple vertical shafts 2-5 cm in diameter with an inclined living chamber at the bottom (ichnogenus Mucunopsis; Hantzschel 1975). Extinct amphibian footprints have been found as well as skeletons fossilized in aestivation burrows (Gray 1988; Boucot 1990). Limbless amphibians such as the caecilians (Gymnophiona)burrow through soil continuously, like the earthworms and insects on which they feed.
Reptiles
Snakesandtortoisesretreatfromextremesofweather in shallow, sloping burrows. During cool weather, many individuals and species of snakes huddle together in the same burrow. This may be an explanation for the rare occurrence of interwoven, articulated skeletons of snakes in paleosols (Breithaupt & Duvall 1986). Other snakes (Ophidea), lizards (Lacertilia) and snake-like lizards (Amphisbaenia) excavate long, narrow burrows (Kuhnelt 1976). Extinct reptiles also appear to have burrowed and otherwise modified soils. Articulated skeletons of the therapsid Diictodon have been found within helical burrows in paleosols (Smith 1987). Dinosaur footprints are widely known in lake-margin sediments (Lockley 199 1). These enormous creatures probably had a profound effect on mixing and compaction of soils. Shallow nests of dinosaur eggs and hatchlings have been found within calcareous paleosols of dry upland regions (Retallack 1997f; Carpenter 1999).
Birds Many kinds of birds build nests in and on the ground, but these have only rarely been encountered as fossils (Boucot 1990; Emslie et ul. 1998). Especially distinctive nests in soil are the salt-encrusted pedestal nests of flamingos (Phoenicopterus ruber), large (3.5 m diameter by 2 m high) incubation mounds of Australian mallee fowl (Leipou ocellutu), and pebble nests of penguins (Pygoscelis udeliae). Other traces of birds include fossil footprints, which are especiallywell known in lake-margin sediments. Local concentrations of bones of rodents and other small mammals in paleosolsmay represent regurgitated pellets of owls and other raptorial birds (Mellet 1974).
14 5
Mammals Many kinds of mammals, such as moles, live their whole lives within the soil and seldom come to the surface. For others, burrows provide shelter and a safe place to raise young. Mammalian burrowsrange fromsimple inclined dens to complex systems of burrows found in prairiedog towns (Meadows & Meadows 1990; Burns 1996). Among the most impressivemammalian burrows found in paleosols are the deep (up to 2.75 m) helical entrance shafts and elongate living chambers (ichnogenus Duimonelix) thought to be the work of extinct beaver-like rodents (Fig. 10.10). Simple shallow dens of carnivorous mammals also have been found in paleosols (Hunt et ul. 1983:Retallack 199 1a). Mammals also affect soils in other ways. Their fossil feces (coprolites) are commonly found in paleosols (Retallack 1998a). Herbivore coprolites are seldom found in well-drained paleosols because they are so readily decomposed. Carnivore coprolites are quickly bleached of organic matter and become cracked and powdery upon exposure, but their phosphatic composition ensures their preservation in well-drained alkaline soils in numbers much greater than herbivore coprolites. Fossil footprints have been found in weakly developed paleosols (Boucot 1990). More subtle effects of large mammals have been noted in studies of buffalo wallows: areas locally trampled by North American Bison (Polley & Collins 1984). Compared with soil outside the wallow, trampled soil inside is more compact, clayey, alkaline, less moist and poorer in phosphorus. As a result, wallows are overgrown with different kinds of herbs and grasses that can tolerate such different soil, as well as heavy grazing. Similar effects on soilhave been documented for East African elephants and hippos. Elephants destroy trees, converting bushlandintograssland (Vesey-Fitzgerald 1973). Effects of vegetation clearance and the creation of hard infertile surfaces are seen in a most extreme form in human influence on soils (Keller 1999). Traces of ecosystems Just as a fossil burrow or coprolite can be regarded as a trace fossil of a n organism, so can a fossil soil be regarded as a trace fossil of a n ecosystem. Few of the individual species of an ecosystem or their interactions leave a
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Figure 10.10 Trace fossilsof rodents, beetles, ruminants and plants all drawn to the same scale, from the Miocene (2 1Ma) upper Harrison Formation, at Eagle Crags, near Harrison, Nebraska, USA. The large corkscrew burrows Daimonelix circumaxillaris (A-D) are thought to have been excavatedby the fossil beaver Pdaeocastorjbssor(A) occasionallyfossilizedwithin them. Smaller helical burrows (E) were probably excavatedby the endoptychine gopher Gregorymys. Narrow burrows associated with the larger ones (C-E) may have been made by dung beetles, because flattenedcoprolites of ruminants (F)have associatedburrows of the same size and rounded dung-lie boli (G and H) also are found. The surface of all the burrows is encased in fossil root traces (L) sometimes silicifiedto the extent that cellular structure is preserved (M) (fromRetaUack1990; with permission fromElsevier,Amsterdam).
discernible trace in paleosols,but their more general effects may be preserved. For example,the kind of vegetation may be evident from the pattern of root traces, sequence of soil horizons and other features of paleosols (see Figs 5.2 and 5.4). The general features of vegetation, such as stature and spacing, determine what botanists call a ‘plant formation’ (not to be confused with rock formations of geology). A plant formation is distinct from acommunity or association,in that it is not defined by the particular species present, as are some ecological categories of vegetation. Forest, woodland and grassland are examples of plant formations. The
wooded grasslands of North America, South America, Africa and Australia all have small-leaved and often many-stemmed trees scattered in more extensive grass, although the trees and grasses in each region are different species.It is plant formations, rather than plant communities, that are reflected best in soils. It may not be possible, without supporting paleobotanical evidence, to determine from a paleosol whether the trees were oaks. acacias or eucalypts. On the other hand, it is possible from paleosols to distinguish woodland and grassland. Plant formations can be reconstructed from fossil plant remains only under exceptional circumstances.
Organisms Many fossil plants were transported far from where they lived and mixed in lakes and rivers, where they were preserved (Gastaldoet al. 1995). Information on the relationship between plant formations and soil features is gathered primarily from observations of modern soilsand the vegetation they support. For example, the idea that Mollisols are the natural soils of grasslands and Alfisols the soils of woodlands has been established by detailed studies of the Illinois forest-prairieecotone (Jenny 1941)andbyglobalsoilmapping (FA0 1971, 1974. 1975a,b, 1977a,b, 1978a,b, 1979, 1981; Buol et al. 1997). Such observations are not without exceptions. Much tropical grassland grows on red, noncalcareous, base-depleted and clayey soils (Oxisols of the Soil Survey Staff 1975, or Ferralsols of FA0 1971, 1974, 1975a,b, 1977a,b, 1978a,b, 1979, 1981).These stable landscapes on Precambrian bedrock received their soil profiles from rainforest of at least Miocene age (20Ma; RetaIIack 199la). Where grasslands grow on Quaternary alluvium or volcanic ash in Africa (de Wit 1978;Jager 1982),their soils are gray, yellow or brown, calcareous, nutrient-rich and silty (Inceptisolsor Mollisols).A comparable case is open grasslands growing on Alfisols in theTexas High Plains. These soils were at least 250000 years in the makiig, and there is evidencefrom fossil pollen, fossil beetles and isotopic composition of soil organic matter (see Fig. 20.5) for pinyon (Pinus edulis) woodland on these soils as recently as early Holocene (HoIIiday 1990, 1995). Yet another exception is Douglas fir (Pseudotsuga menziesii) forests growing in soils with mollic epipedons in the hills around my home in Eugene, Oregon, USA. The widespread Mollisols of the WillametteValley of Oregon (seeFig.6.1) are thought to be in part aproduct of Native American use of fire. Since European settlement, cultivated fields continue to be burnt to encourage spring growth of grasses, but fire restrictions in other parts of the landscape have allowed reforestation (Johanessen et a]. 1971). If these soils became paleosols, their profile form would be evidence of a long time of formation under grassland rather than historical invasion by conifers. Similarly,the red soils of Africa and Texas are evidenceof past woodland and forest rather than recent grassland. These situations could be viewed as confusing the true relationship between vegetation and soils, and they have taken some scientific ingenuity to
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unravel. In each case, however, soilspreserve a record of thevegetation that was most significant for its formation in the long run. Early successional vegetation In ecological succession, the first to colonize bare ground are usually herbaceous forms such as lichens, mosses, grasses and ragweed (Kimmins 1997). These plants are widely regarded as weeds in gardens because of their ability to establish themselves on bare earth and their lack of showy flowers. Early successional soils and paleosols (Entisols and Inceptisols) retain most of the original features of their parent material, such as the bedding and ripple marks of flood alluvium, despite its incipient disruption by root traces and burrows. Such weakly developed soils are common in and around paleochannel sandstones of ancient alluvial successions (Bown & Kraus 1987;RetaIIack & Krull 1999). Very weakly developed paleosols with calcareous rhizoconcretions in ancient deposits of beaches (Ettensohn at al. 1988), desert dunes (Loope 1988; Loope et al. 1998) and loessites (Soreghan et al. 1997) may have supported dune-binding early successional vegetation. Forest and woodland These closed-canopy formations may be distinguished by stature: woodland under 10m high and forest taller than 1 0 m (White 1983).Large root traces are obvious indicators of forests and woodlands. The large biomass of trees also has substantial effects on differentiation of soil horizons into a sandy, bleached near-surface (eluvial or E) horizon and a clayey, sesquioxidic or humic subsurface (illuvialBt,BhorBs) horizon asinSpodosols, Alfisols and Ultisols (see Fig. 3.8A). The lateral continuity of these horizons is an indication of the degree of canopy closure, because widely spaced trees promote the development of cradle knolls and basket podzols (Schaetzl & Follmer 1990). The thickness of the horizons is related to biomass because of the depth and density of root penetration and the flux of weathering chemicals produced. Forested soils generally are about a meter thick, whereas those of woodlands are thinner (see Fig. 3.9). Such well-differentiated paleosols are
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common in floodplain facies of ancient fluvial deposits (Retallack1983a, 1991a, 1997a).
at high paleolatitudes also are known (Fig. 10.11; Retallack 1997d).
Rainforest
Oligotrophic forest
These tall forests of humid tropical regions have two or more canopy levels, a great diversity of plant speciesand a profusion of vines and epiphytes (Leigh 1998). Their soils (Oxisols and Ultisols) are thick, clayey, deeply weathered of bases, and in some cases lateritic and bauxitic. Many have little soil structure apart from abundant sand-sized spherical micropeds. In many cases the large roots of buttressed tree trunks do not penetrate the infertile substrate, but form a surface root mat (Sanford 1987). Generally similar paleosols are widespread at major geological unconformities that may have supported rainforest or a number of other kinds of vegetation during their long period of formation (Ollier & Pain 1996). More convincing examples of ancient rainforest paleosols have been found in ancient fluvial deposits (see Fig. 19.5; Retallack & German-Heins 1994;Bestlandetal. 1996). Humid forestsof southern beech (NothoJagus)in Chile and New Zealand and of hemlock (Tsuga)in the northwestern USA are sometimes termed temperate rainforest. These form soils with well-differentiated clayey subsurface (Bt) horizons that are noncalcareous and often kaolinitic (Ultisols).Comparable paleosols formed
Quartz-rich soils low in nutrients (Quartzipsamments, Dystrochrepts,Spodosols)are now widespreadin humid temperate and alpine regions under needle-leaf forests. In tropical regions such sandy, low-nutrient soils commonly have a very deep horizon of iron and humus enrichment (Bs) and support dwarf angiosperm forest (Nahon 1991).Some sandy paleosols are identical with modern Spodosols (Pomerol1964), but Carboniferous to Triassic paleosols with quartzose surface horizons (ganister or E) commonly also have sideritic gleyed subsurface horizons (Bg) indicating less strongly acidic pH than under modern oligotrophic forests (Retallack 1997d). Heath The original vegetation for which the term heath was coined is dominated by heather (Calluna vulgaris) on thin, sandy, oligotrophic soils (Spodosols) in northwestern Europe. The term also has been used for similar low (<1m tall), dense, woody vegetation of thin, sandy, waterlogged soils (Aquods) of Australia and South Africa (Specht 1979). Paleosols of this second type are
Figure 10.1 1 Type Johnclaypaleosolfrom theLowerTriassic(245 Ma),Feather Conglomerate,in the Allan Hills, Antarctica. This is interpreted as a temperate rainforest soil,Ultisol (USDA)or Acrisol (FAO),on the basis of its large root traces,subsurfaceclay enrichmentand deepweathering (chemical data on this paleosol are presented inFig. 4.9: seeRetallack&Krulll999).E.S. Krullfor scale.
Organisms known from fossil leaf litters to have supported heath of extinct seedferns (Retallack 1977,1997d).
Dry woodland In many tropical and southern continents, dry woodlands (sometimes called savanna woodlands) are widespread on ancient (Cretaceous and early Tertiary) land surfaces (Cole 1986). Their soils include red, kaolinitic, lateritic or bauxitic horizons from ancient humid weathering together with superimposed calcic (Bk), gypsic (By) or silicic (Bq) rhizoconcretions and other features from their current round of soil formation in dry continental interiors (Nahon 1991).The effects of dry woodland are clearer in alluvial deposits, where their soils commonly have well-differentiated subsurface clayey (Bt) horizons, and deeper horizons of calcareous nodules (Bk). Calcareous and siliceous rhizoconcretions are important evidence of dry woodland from paleosols (Retallack 1991a). In some cases, twigs, pollen and permineralized stumps are preserved in paleosols of dry woodland (Francis 1986), their microbial decay limited by aridity. Wooded grassland This is a plant formation of solitary trees scattered in grass. The term savanna is broadly synonymous, but includes also what I would call grassy woodland (Cole 1986).Wooded grassland soils and paleosols have a mix of features found in soils of woodland, in which tree canopies cover more than 40% of the ground, and of open grassland, in which tree canopies cover less than 10%of the ground (White 1983). Surface horizons of wooded grassland soilsare organic and have a fine structure (granular or crumb). Their subsurface is moderately calcareous, with a shallow (30-60 cm) calcic horizon. These features of subhumid to semiarid grassland soils are mixed with others, such as weakly developed, leached, near-surface (eluvial or E) and clayey subsurface (illuvialor Bt) horizons typical under woodlands. Similar calcareous paleosols (Mollisols and Inceptisols)commonly preserve bones of mammals also showing a mix of woodland and grassland affinities (Retallack1988b, 1991a). Root traces are important guides to ancient wooded grassland. Most roots of grasses are less than 2 mm in
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diameter, although their rhizomes. running along or just below the surface, may belarger. Abundant fineroot traces, together with a few substantial ones, are typical of wooded grassland soils. In paleosols this may be especially clear because root traces have drab haloes formed by burial gleization (seeFig. 3.7), which is thought to reflect the last crop of roots before burial. Charcoal also is typical of wooded grassland, because much of it is prevented from reverting to woodland by frequent fires (Cole 1986). Wooded shrubland This is a plant formation with the general appearance of wooded grassland, but with few if any grasses. Their place is taken by woody shrubs or other herbs. Examples are the sage (Arternisia tridentata)and juniper (Juniperus occidentdis) vegetation of central Oregon (Franklin & Dyrness 1973) and saltbush (AtripZex vesicuna) and myall (Acacia sowdenii) vegetation of central Australia (Beadle 1981).Soils of wooded shrubland have shallow hardpans and calcareous nodules, and surficial salt crusts and stony pavements (Aridisols).Root traces in these soils and paleosols are more sparse and clumped than in wooded grassland, and there is no hint of the organic, crumb-structured surface layer (mollic epipedon) of grassland soils. Before the evolution of grasses, wooded shrubland may have been widespread insemiaridandaridregions(Reta1lack 1985,1997e). Fire-prone shrubland Another kind of dense shrubby vegetation is found in fire-proneregions, with a summer-dry (Mediterranean) climate. Chaparral is the name given to this plant formation in southern California. Similar kinds of vegetation in southern Europe are called maquis and garrigue, in Chile matorral and in Australia mallee (di Castri et d. 1981). These grow in shallow stony soils (Entisols and Inceptisols). Some have thin clayey (Bt) or calcareous (Bk) subsurface horizons. The common charcoal in thesesoilsispreservedinpaleosols(Harris 1957,1981). Open grassland Treeless grassland vegetation is known by several regional synonyms, including Russian steppe, North
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American prairie and Argentine pampas (Archibold 1995).The term grassland is best used generally to include such formations, as well as wooded grassland, parkland (grassy glades in woodland) and groveland (woodlandpatches in grassland).The soils of grasslands are united by their distinctive well-structured organic surface horizons (mollic epipedons: see Fig. 3.8B). This horizon is formed by the multitude of fine roots and abundant soil fauna associated with grasslands. This distinctive surface horizon may be laterally continuous and deep (a meter or more) in grasslands of subhumid climates, such as the tall-grass prairie of the western Great Plains of North America. The surface horizon is shallower and laterally variable in thickness under tussock grasses of semiarid climate, such as the short-grass prairie of the western Great Plains and intermontane rangelands of the western USA (Ruhe 1984). Abundant, fine (<2mm diameter) root traces and small (<4mm)crumb peds are especially diagnostic of grassland paleosols, but their dark color and organic matter are seldom preserved in paleosols (Retallack 1997b). Also useful is supporting evidence from grass phytoliths (Piperno 1988), grass cuticles (Morley & Richards 1993), fossil grasses (Dugas & Retallack 1993) and mammal bones and teeth (Jacobset al. 1999). In addition, an isotopicallyheavy isotopiccomposition(S13C)of both paleosol organic matter and pedogenic carbonate can be taken as evidence for the C, photosynthetic pathway (see p. 304) found largely in tropical grasses since the late Miocene (7Ma),or as evidence for CAM plants such as desert saltbush (Atriplex) or aquatic quillworts (Isoetes: Cerling et al. 1997a; Koch 1998). In a Pleistocene loess sequence near Kahlotus, Washington, isotopically heavy values represent saltbush (Atriplex) deserts, also indicated by Aridisols with shallow calcic horizons and abundant cicada burrows (Fig. 10.12). Grassland ecosystemsat this high latitude (47"N) show light isotopic values of C, grasses, also indicated by Mollisols with abundant fine root traces, crumb peds, earthworm castings and grass phytoliths (Busacca 1989,1998). Shrubland Many desert regions support a steppe-likevegetation of low-growing bushes that includes few grasses. Exam-
ples are the bluebush (Maireana sedifolia) shrublands of inland Australia (Beadle 1981) and the creosote bush (Larrea cuneijiliu) shrublands of Argentina (Evenari et al. 1986). Individual small shrubs dot the landscape, often in intriguingly regular patterns (Klausmeier 1999),but there is much bare ground exposed. Root traces in these soils are sparse and strongly clumped under individual bushes. There may be evidence of wind scouring between the plants and of dunelike accumulations of wind-blown silt around them. These soilsalso tend to be stony, sometimeswith adesert pavement of interlocking rocks. They also may be encrusted with salt and cemented with calcareous nodules. This vegetation is best recognized in paleosols by the size and distribution of calcareous rhizoconcretions (Loope 1988), but isotopic and trace fossil evidencemayalsobeuseful (Fig. 10.12). Brakeland Herbaceous vegetation of well-drained soils is now mostly grassland, but there are areas of herbaceous vegetation with rhizomes, runners and roots not so densely turf-forming as in grasslands. These have been called brakelands, after fern brakes of bracken (Pteridiurnaquilinum), although this is commonly an early successional phase of recovery fromfieldfies (Retallack 1992a).Examples of permanent brakelands are Tillandsia cushions of the fog loma desert of Namibia (Rauh 1985) and Dryas integrijilia cushions of Devon Island, Arctic Canada (Bliss 1977). The soils of these examples are evaporitic-calcareous and permafrost-deformed, respectively, but brakelands of the geological past identified on the basis of their rhizome and root traces were widespread in less severe environments (Retallack 1992a). Polsterland This is a formation of well-drained soils consisting of patches (pokers) of multicellular plants, such as lichens, mosses and liverworts, lacking roots or rhizomes. The best-known examples are from till plains near Mt Rigby, Antarctica (Campbell & Claridge 1987) and on Elephant Island, near the Antarctic Peninsula (O'Brien et al. 1979), where their soils are thin and
Organisms KAHLOTUS, WASHINGTON calcareousness development hue
earthworm cicada pellets Durrows c carbon isotopes oxygen isotopes g! E marine grain size 50%
ilt [:::lclaystone sand,silt
clasts
clay B v o l c a n i c ash
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mcrumbpeds
s
-6
-4
, -1
-5 , - 3
i - root traces
cicada burrows
brown colour
E
P 8
isotopic stage
1 -12 Ka
2
3 -59 Ka
]081calcareous nodules ~
calcareous mottles
Figure 10.12 Asequenceof Quaternarypaleosols(0-lOOKa) in PalouseLoessnearKahlotus,Washington,USA(section 36T13N R33E,FranklinCo.),showing(leftto right) field section, thermoluminescencedates,paleosoldevelopment,paleosol calcareousness,paleosol Munsell hue, sand-silt-clayproportions,6' 3C of pedogenic carbonate,6180of opal phytoliths. abundanceof earthwormpellets,and abundance of cicadaburrows (fromBusacca1998).Thepaleosolsare interpreted as Aridisols of desert shrubland and MolIisols of C, open grassland.
weakly developed. Although currently rare, polsterlands may have been much more common and widespread in the distant geological past (Retallack 1992a).
Desert scrub Deserts have a sparse distribution of woody and succulent plants (Evenari et al. 1986). Some of the plants of these formationsare peculiar in appearance,such as the African candelabra cactus (Euphorbia candelabra) and the North American saguaro cactus (Trichocereusgiganteus) and boojum tree (Idria colurnnaris).There is a good deal of bare earth exposed. Their soils are similar to those of desert shrubland but for the large root traces
and very shallow calcic horizons found under desert scrub.
Microbial earth Salt pans, sand dunes androck surfacesof desert, alpine and boreal regionsappearvirtually barren of vegetation but support fairly complex communities of microbes (Campbell 19 79).In some cases the microbes are abundant enough to discolor the ground green or purple. Apart from a few lichens, other large organisms are excluded from these communities by extreme dryness, saltiness,windinessor cold.The soils of microbialearths show little profile differentiation, except for surface features such as rock varnish (Krinsley 1998),ministro-
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matolites (Fig. 10.2)andvesicular structure (McFadden et aI. 1998). Microbial earths may not have been confined to difficult habitats before the advent of vascular plantsonland(Retal1ack 1992a).
Microbial rockland In the Dry Valleys of Antarctica, the top few millimeters of sandstone and dolerite house communities of algae and lichens (Friedmann & Weed 1987) that create a weathering rind like a miniature Spodosol. The small vase-shaped solution hollows of algae in limestone can be so abundant as to blacken the white rock (Folk et al. 1973; Viles 1987), and likely fossil examples are known (Folk&McBride 1976).
tend to be acidic and organic and may show permafrost deformation (Retallack 1999a). Bog British high-latitude mossy vegetation of Sphagnum is a good example of bog. The term bog is used generally for vegetation of rootless plants such as mosses and liverworts, in acidic, freshwater, waterlogged ground (Gore 1983).Acharacteristicof these soilsis thesurficialpeat layer, which is an important source of domestic fuel in the northern British Ides and in Scandinavia. Moss peat forms a dull coal, lacking vitrinite and fusinite of the kind found in coals formed under swamps. Fossil mosses are sometimes well preserved in ancient bogs (Smoot & Taylor 1986).
Taiga and krummholz Taiga is a stunted woodland found in polar regions. Krummholz is a comparable vegetation of gnarled and windswept asymmetric trees (‘flagtrees’)found near the snowlinein alpineregions (Archibold 1995).Bothkinds of plant formation grow in soils with distinct nearsurface leached (E) horizons and subsurface clayey, organic or ferruginous (Bt, Bh or Bs) horizons, like other woodlands. Most are acidic, sandy and organic, because little moisture is lost to evaporation, and rates of decomposition of leaf litter are low in cold climates. Also limited is the activity of soil fauna and evidence of these organisms in the form of burrows and fecal pellets. Most distinctive of these soils are traces of permafrost or frost heave, such as patterned ground or ice wedges (Retallack1999c).
Tundra and alpine fellfield Shrubby and grassy vegetation of high mountains and polar regions can be distinguished from that of dry to desert regions by the abundance of mosses, lichens and a variety of dwarfed vascular plants, including some with extremely reduced and clumped shoots (‘cushion plants’). Some grasslands are found also in tundra and alpinefellfields(Bliss 1977),but soilsunder suchvegetation are thin and weakly developed, lacking the thick surface (mollic epipedon) and calcareous subsurface (Bk)horizons of grassland soils in warmer regions. Like soils of krummholz and taiga, these soils of cool climate
Marsh The terms bog and marsh are sometimes used interchangeably, but marsh is here taken as vegetation of rhizomatous or rooted, herbaceous plants of acidic freshwater, waterlogged ground (Gore 1983). If waterlogging persists for most of the year, these soils accumulate a surficial layer (0horizon) of peat, in which plant material may remain recognizable as cuticles (Krassilov 1981), pollen (Batten 1973) or permineralizations (Riceet al. 1995). Fen This waterlogged vegetation is similar to marsh but for its neutral to alkaline groundwater (Gore 1983). The natural acidityof decaying plant material ismitigated in fens by a variety of factors: groundwater draining from limestone, carbonate-king organisms such as charophytes and snails, rainfall in the subhumid range and a pronounceddryseason. Fenpeats and coals can bemore shaley than those of marsh, and contain caliche nodules, charophyte oogonia, ostracod valves, and snail shellsnotfoundinmarsh (Wright&Platt1995; Metcalf 1996).
Salt marsh Herbaceous intertidal vegetation of rooted and rhizomatous plants is found mostly at high latitudes,
Organisms beyond the latitudinal distribution (currently from 30"N to 45"s) of mangal (Chapman 1977). Salt marsh is restricted to muddy bays and estuaries that are protected from wave action. Small amounts of peat may accumulate in salt marsh soils, but most organic matter is utilized by a diverse marine fauna including oysters, mussels, polychaete worms and crabs, which leave a variety of distinctive trace fossils and burrows in salt marsh soils (Freyat aZ. 19 78). Other distinctive features of salt marsh soils are nodules of pyrite and marcasite formed by reduction of marine sulfate by anaerobic bacteria(Altschu1eretaZ. 1983).Saltcrustsarefoundinsalt marsh soils above mean tidal level. Swamp Forestsand woodlandsof waterlogged ground are called swamps. A well-known example is the bald cypress (Taxodium distichum)swamps of theFloridaEverglades (Ewe1 & Odum 1984). Considerable thicknesses of peat accumulate under permanently waterlogged swamps. These peats are transformed during burial into coals of high calorific value, full of woody fragments (vitrinite and fusinite).Themineralportionof the soilunderthepeatis dark with organic matter, drab with reduced minerals and may contain well-preserved large root traces (Phillips& DiMichele 199 7). Seams of woody peat and vitrinite-rich coal can form from allochthonous accumulations of plant matter on shores and river banks, but these are uncommon and thin compared with coals formed under swamps (Wagner&Pfefferkorn 199 7).Peat doesnot accumulate under seasonally dry swamps, because of dry-season decay, but their soils and paleosols may include fossil trunks, well-preserved carbonaceous root traces and siderite nodules like the underclays of coals (Gastaldo 1986). Carr
Swamps are acidic in reaction like marsh, but carr is woody wetland vegetation growing in alkaline waters (Gore 1983). The neutralization of soil acidity is achieved in ways already outlined for fen vegetation. The woody peats of carp can be distinguished by the presence of calcareous shells or nodules of carbonate minerals, such as siderite or calcite. Coal balls are cal-
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careous nodules common in some Late Paleozoic coals, and may have formed in this way (Scott et al. 1996). Noncoaly gray paleosols with calcareous nodules and large carbonaceous root traces may represent seasonally dry carr (Retallack&Dilcher 1988). Mangal The confusion of tangled roots and gnarled stems of trees growing in the intertidal zone are the basis for the name mangal. This term is now widely used to describe the plant formation rather than the term mangrove, which refers to a particular kind of plant (Chapman 1977). Many mangroves have distinctive and potentially preservable features, such as prop roots and air roots that aid stability and respiration in waterlogged substrates. Other plants of mangalvegetation are not so distinctive but have physiological tolerance for the anaerobic and saline condition of mangal soils. Mangal soils are similar in many ways to salt marsh soils. Both are drab colored, and contain pyrite or marcasite nodules (Retallack & Dilcher 198 la,b; Phillips et al. 1994). They are riddled with burrows and other traces of marine organisms, such as oysters and mussels (Boucot 1990;Galli 1991).Root tracesalsoarepreserved wellin mangal soils, and they are larger and more varied in morphology than those of salt marsh soils (DiMichele etal. 1987).
Fossil preservationin paleosols Perhaps the most obviousevidence of organisms inpaleosols is fossils, which are abundant in some paleosols, because under suitable conditions they are concentrated on ancient land surfaces. Paleosols also provide evidence for past life where no fossils are preserved. Bones,for example, are not preserved in acidicpaleosols. nor leaves in oxidized profiles, yet ecosystems of considerable biomass, such as tropical rainforest, grow insuch noncalcareous red soils. Such thick red soils may represent breaks in sediment accumulation of many thousands to millions of years. Other kinds of soilsreflect the time scales of plant succession, of the order of tens to hundreds of years. Information from paleosols on preservational biases against fossilization and on the duration of ancient ecosystems may be useful to check or enlarge inferences drawn from the fossil record of life
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on land. Paleosols ground hypotheses of former ecosystems both as a preservational filter and as a basis for natural associations of fossils (Retallack 1983a, 1994a).Evidence from paleosols is especially useful because it is independent of the fossils themselves. Paleosols provide evidence of laterally interacting ecosystems on ancient landscapes, comparable with the perspective of landscape ecology (Forman & Godron 1986). Preservation as a function of soil chemistry Ecosystems provide a pool of potential fossils, but for various reasons, few of these ultimately &id their way into the soil. Most are eaten or decayed, then trampled or weathered at the surface until unrecognizable. Fragments of bone or leaves also can be mixed with other remains by the action of wind and water. The moving and mixing of fossils from different communities is, however, more of a problem for fossils within deposits of streams and lakes than for those within paleosols. Once an assemblage of fossils is buried or trampled into a soil, it is subject to a new round of destructive processes.Trampling and physical compaction by swelling of wet clay result in cracking and distortion of fossils. Further microbial decay and chemical alteration provide a final filter for what ultimately is preserved in paleosols, as outlined in the following paragraphs. The preservation of plant leaves and other fossil organic matter in paleosols can be related to the degree of oxidation of soils. This is not so much because the fossils are directly oxidized but because the most effective microbial decomposers in soils have an aerobic metabolism. For this reason, organic matter does not accumulate in well-drained soils. In soils where the availability of oxygen is limited by stagnant groundwater, thick deposits of peat form. These waterlogged soils also are the ones most liely to preserve fossil leaves and pollen. Paleobotanical research has unearthed a great deal of information about wetland vegetation but rather little concerning plants of dry soils.The activity of decomposer microbes is limited to a lesser extent by soil acidity or pH. Spores and pollen may be preserved in well-drained soils if they are strongly acidic (pH <4) or strongly alkaline (pH >9). Charcoal is appreciably more resistant to decay and it has longer residence time
in soils than other kinds of organic matter (Retallack 1998a). Plants also produce mineralized hard parts that can be fossilized in paleosols. Much of the coarse texture of grass and its abilityto form hay when the leaves are dead is due to encrustation with silica in the form of opal. Small opal bodies (phytoliths) of characteristic shapes for particular species are also found within the cells. These may accumulate in soils as the plants die or are burnt (Dixon&Weed 1989;Gibliing&Rust1992).Plant opal is more soluble than quartz at a pH of <9; both dissolving at more elevated pH. Woody pits (endocarps) of hackberry fruits (Celtis occidentalis)contain 2 5-64% by dry weight of calcium carbonate and 2-7% silica (Yanovsky et al. 1932).These ‘stones’of fruits accumulate in alkaline soil. Rainwater may have a pH as low as 5.7 before it is considered polluted or acidic rain (BaasBecking et al. 1960), capable of dissolving calcareous phytoliths on the surface of soils. Also dissolved in acidic soils are shells of land snails, calcite-impregnated carapaces of millipedes and woodlice,andegg shellsof dinosaurs and birds (Carpenter 1999). Dissolution of snail shells proceeds rapidly once the organic outer coating (periostracum) decays aerobically.Millipede exuviae are even more ephemeral because few species are heavily calcified and many millipedes consume their molted exoskeletons (Neville 1975). Birds, snails and millipedes also tend to be more abundant and diverse in calcareous soils (Hopkin & Read 1992;Gravelandetal. 1994). Bones and teeth are composed of a variety of calcium phosphate minerals (Dixon & Weed 1989). which also are dissolved in acidic soils. They can persist under more acidic conditions (down to about pH 6) than the calcium carbonate of snail shells. In seasonally dry subtropical Zimbabwe, human bones in 700-year-old graves were well preserved under alkaline (pH 6.2-7.9) termite mounds but had been completely destroyed under adjacent acidic (pH 4.1-5.4) soils (Watson 1967). The so-called ‘bog people’ are corpses interred in acidic bogs of north-western Europe for thousands of years (mostly from about 100 BC to AD 500). They show varying degrees of bone decalcification and, in one case, complete loss of bone within well-preserved skin and other soft tissues (Glob 1969). Teeth persist in acidic conditions for a longer time than bone because they are denser and less porous and so present a smaller
Organisms internal surface area to dissolving solutions. Bone cracking and flaking is partly a process of degreasing and desiccation (Behrensmeyer 1 978), but also reflects dissolution in the natural acidity (pH 5.7) of unpoltuted rain. Smaller bones have a higher surface-to-volume ratio than large bones and so are more readily dissolved. This may be the reason why bones of small and young animals are so rare in association with those of adults in many paleosols (Gordon & Buikstra 198 1; Carpenter 1999).On theother hand, boneisoftenbestpreservedin calcareous nodules (Downing & Park 1988; Retallack 1983a). Coprolites,or fossil feces, have compositions as varied as the diets of the creatures that produced them. Coprolites of organic matter, such as plant fiber or hair, are preserved in anaerobic, waterlogged environments, but they also can be preserved by freezing, permafrost, excessive acidity in peat bogs or extreme desiccation in deserts. All these situations serve to inhibit the activity of aerobic microbial decomposers (Heizer & Napton 1969).Coprolitesof birds and other carnivores contain appreciable amounts of bone and other broken-up materials of phosphatic composition (Chin et al. 1998). preserved under the same kinds of alkaline conditions as bones. In alkaline-oxidizingpaleosols, where organic coprolitesof herbivoresare destroyed, almost all the preserved coprolites may be those of carnivores, despite the much greater original abundance of herbivores (Retallack198 3a). From these considerations, each kind of fossil can be consideredto be chemicallystable under certain general conditions of pH and Eh (Fig. 10.13). This model predicts only the most usual case. Depending on their pH and Eh, paleosols tend to be dominated by the kinds of fossils favored for preservation. This is not to say that an occasional fossil skeleton overwhelmed by a flood may not be found within a noncalcareous paleosol. Such fossils will be rare and show other signs of being exceptional, such as full articulation of the bones. This model explains the common field observation that bones, insects, and leaves are seldom found preserved in the same deposit or paleosol. Preservation of all three requires the geologicallyunusual combination of both alkaline and reducing conditions, as found in a few unusual deposits of lakes (Grande 1980; Schaal & Ziegler 1992) and paleosols (Reisz et al. 1982; Francis 1986; Retallack 19921).
155
Preservation as a function of time Thedestruction and accumulationof fossils inpaleosols takes time. In the calcareous and saline soilsof Amboseli National Park, Kenya, it has been suggestedthat 10000 years would be needed for the accumulation of bones of large mammals to a density of 100 bones per 1000m2 (Behrensmeyer 1982a). In acidic soils of Zimbabwe, in which human skeletons in 700-year-old graves had been destroyed, the skeletons of farm animals only 20 years old showed evidence of corrosion and flaking (Watson 1967).This is not to imply that the accumulation or destruction of USA, fossils in paleosols is constant. Soil development and ecological succession go hand in hand. The differential preservation of different kinds of fossils can be related to the time over which the paleosol formed, insofar as both soil chemistry and community composition change during plant succession (Retallack 1998a). For example, in the Lower Eocene (56 Ma) Willwood Formation of Wyoming, USA, fossil leaves are found in nearstream shales and shaley paleosols (Entisols)and in coal-bearing paleosols (Histosols) and thus reflect only early successional and swamp vegetation (Wing 1984).Fossil snails and mammal bones are found mainly in a variety of well-drained paleosols formed under stable floodplain forests over 102-105years (Bown & Beard 1990). Only fossil hackberry (Celtisphenacodorurn)pits in the paleosols and rare and poorly preserved juglandaceous (walnut family) pollen grains in nearstream shale remain as paleobotanical evidence of these stable floodplain forests. The forests in which the mammals lived are thus poorly represented by fossil plants. It would be a mistake to assume that the diet of the fossil mammals is represented by the well-preserved leaf assemblages or that this kind of vegetation was an important selective pressure in their evolution. The fossil plant assemblages represent an ephemeral and local kind of vegetation. Regional vegetation covering most of the region for most of the time is better represented by the paleosols. Completeness of the fossil record It has long been appreciated that the fossil record provides a very biased impression of the past. But without evidenceindependent of fossils,it is difficultto determine how biased. Paleosols provide independent
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Chapter 10
red clay brown clay
(=I relict bedding Iron cancretions
@ snail shells X opal phytoliths 6 calcareous phytoliths b leaves and fructifications
evidence of the kind needed. Regardless of whether they contain fossils or not, they are themselves evidence of a former ecosystem. Because most current land communities produce leaves, charcoal, phytoliths, land snails, bones, teeth and coprolites, and so did many ancient communities, then all of these categoriesof fossil should be preserved in every paleosol of a sequence if it had a ‘complete’fossil record, Real sequences of paleosols fall far short of this expectation for reasons such as differential preservation as a result of paleosol chemistry and timeof formation (Retallack 1998a).Thiscrudemethod
Figure 10.13 Theoretical&-pHstability fields of commonkinds of terrestrialfossils preserved in paleosolsand representative kinds of soils.Soilcolumnsare about 2 m thick, and height for forest vegetation is about lOm(fromRetaUack1998in TheAdequacyof the FossilRecord, editedby S.K. Donovan& C.R.C. Paul:reprinted with permission,0 JohnWiley & Sons).
of quantifying completeness of the fossil record allows assessment of the preservational potential of different kinds of paleosols. The most complete record of fossil plants is found in paleosols of swamps (Histosols and gleyedsoilsof other orders).The most completerecord of fossil mammals is in well-drained alkaline paleosols of dry climates (Alfisols, Mollisols and Aridisols). These findings provide an explanation of the common observation that fossil plants are best preserved in coal measures and fossil mammals are most common in calcareous red beds.
Organisms Paleosols also may be useful for assessing how complete a representation of a former community is an assemblage of fossils. Fossil leaf litters preserved where they fell on the soil provide the most unbiased record of former vegetation. These contain leaves showing great variation in their degree of decay, as well as organs of differing robustness, such as flowers, fruits and logs, that would be segregated if the remains had been transported any great distance. Such fossil leaf litters are especiallycommon in mangal, swamp and early successional lowland paleosols (Retallack 1977; Retallack & Dilcher 1981a; Gastaldo et ul. 1995). Attritional accumulations of fossil bones within paleosols on which the animals lived present a truerpicture of past animalcommunities than those special cases where herds of animals are overwhelmed by volcanic ash or other catastrophes. Such catastrophic assemblages may be beautifully preserved and attract admiration in museum displays, but often reflect circumstances, such as seasonal migrations or aggregation around a salt lick, that are as special as those surrounding their preservation. Attritional assemblages within paleosols, on the other hand, provide a time-averaged view of what was most abundant over the period of formation of the soil (Behrensmeyer1982a). This is not to say that the fossil record of communities in soils is unbiased. The variable preservation of leaves in leaf litters is evidence that some leaves are being destroyed preferentially (Ferguson 1985). In my own observations of the leaf litter of a living scribbly gum and red gum (Eucalyptus-Angughora) woodland near Sydney, Australia, the high lignin and phenol content of the two dominant tree species favored their preferential preservation over a host of other perennial trees and shrubs with smaller and more fleshy leaves. Such differential preservation may be the reason why fossil leaf litters and some other fossil plant assemblages in nearby Triassic (245 Ma) rocks are dominated by large, robust leaves of Dicroidiurn zuberi, whereas associated shale deposits contain a greater diversity of leaves (Retallack 1976,1977,1997d). Studies of the abundance of modern bone on soils of Amboseli National Park, Kenya, have shown that the animals died and their bones were preserved where they lived, but there is a strong bias against preservation of small bones (Behrensmeyer et ul. 1979). This can be attributed to the higher surface-to-volume ratio of
157
smaller bones, which are thus more prone to acidic dissolution in rain and soil solutions than are larger bones. Indeed, the slope of the line relating abundance to the sizeof animalsis0.68, veryclosetothevaluethatwould be expected from surface-to-volume scaling (0.67). Such size biases can be seen also in early Oligocene (32 Ma) fossil assemblages from Badlands National Park, South Dakota, USA, where bones are generally regarded as beautifully preserved (Fig. 10.14). This bias against small animals can be contrasted with the situation in living populations of mammals, in which smaller animals are generally more abundant than larger ones (Damuth 1982). Using this yardstick, the fossil assemblage of the more calcareous silty savanna paleosol (Conata pedotype of Fig. 10.14) is more representative of past communities than of the less calcareous clayey gallery woodland paleosol (Gleska pedotype). The size bias in eachassemblage hasbeenaddressed (byDamuth 1982) using the correction factor from studies at Amboseli, with the result that Conata and Gleska assemblageslook more l i e modern assemblages. Interestingly, chevrotains (Leptomeryx evunsi) and rabbits (Pulueolugus huydeni) are still well above the line of equal biomass in the Conata assemblage, as are oreodons (Merycoidodon culbertsoni) and three-toed horses (Mesohippus buirdi) in the Gleska assemblage, confirming the apparent dominance of these taxa in the raw collections. In this case, then, taphonomic corrections did not overturn a prior impression of ecological dominance of these species. These difficulties in interpreting paleocommunities from fossils in paleosols pale in comparison with the problems of resurrecting fossil assemblages from deposits of rivers and lakes. Studies comparing leaves in modern lakes with the vegetation growing around the lake have shown that leaves of understorey shrubs are under-represented in lakes compared with leaves of canopy trees (Ferguson 1985). Decay, abrasion and mixing of leavesfrom different communities means that ecological reconstruction of the surrounding vegetation is difficult from the material in the lake alone. The movement of fruits and other plant parts in large deltas is even more complex (Sheihing & Pfefferkorn 1984). Although the fossil record of lakes and rivers is biased toward lowland vegetation l i e that of fossil leaf litters, some far-travelled fragments of upland vegetation do find their way into lakes and rivers. The rain of pollen
Chapter 10
158
A
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into the center of large lakes and into the ocean is more representative of regional vegetation than that found closer to shore or in a paleosol (Chaloner & Muir 1968). Without such records, nothing at all might be known of the taxonomic composition of vegetation in welldrained inland soils.Lakes also may preserve a more precise record of temporal variation in vegetation, such as the successional eutrophication of small lake basins
Figure 10.14 Restorationof interfluve savanna on Conata pedotypepaleosols (Andic Ustochrepts,A) and of streamside gallery woodland on Gleska pedotype (Petrocalcic Paleustalfs,B) during Oligocene (32 Ma) deposition of theScenicMember of theBrule FormationinBadlandsNational Park, South Dakota,USA, with plots of the abundance (numberof specimens)for various speciesof fossil mammals in collectionsfrom each kind of paleosol (C andD. respectively),and after correction for size-relatedbias caused by acidicdissolutionof bone (EandF, respectively)(fromRetallack 198 8b: reprinted with permission kom the Societyof Economic Paleontologistsand Mineralogists).
(Smiley81Rember 1981).Paleosols, on the other hand, represent a view of vegetation averaged over the time of development of the soil. There are, then, advantages to studying paleosols and fossil plants together. Land animals are preserved in lakes but are not as common in them as fossil plants and fish (Grande 1980; Schaal&Ziegler 1992). Most fossil land vertebrates are found in paleosols or in deposits of streams. Bones found
Organisms in streams are mixed from a variety of natural communities in the watershed and re-sorted from geological deposits exposed in stream banks (Behrensmeyer 1982a). Although there is some mixing of mammals and time-averaging of remains in paleosols, in neither case is it so extreme as in assemblagesof stream deposits. Fossil assemblagesfrom paleochannels are nevertheless of value in providing an impression of regional vertebrate diversity, particularly of small mammals. In sequences in which paleosols are unfossiliferous because they are too acidic, streams may provide the only evidence of past vertebrate life. In the Miocene (8.3Ma) upper Dhok Pathan Formation of northern Pakistan, bone is more easily collected from paleochannels and streamside sediments than from paleosols. There is a variety of paleosols formed under swamp woodland, tropical deciduous forest, early successional woodland and seasonally wet wooded grassland (see Figs 11.7 and
159
11.8). Such habitat diversity is reflected in the high diversity of fossil mammals, but one cannot be sure from the fossils alone how they were distributed through the mosaic of likely habitats. Fossils in paleosols show, for example, that the chowsingha-like antelope (Elachistoceras)was more common in gallery forest,but nilgai-like antelopes (Selenoportax and Miotragoceras) were more common in floodplain monsoon forest (Retallack 1991a). Paleochannels and paleosols offer complementary but distinct perspectives for reconstructing past communities. Paleosols and their trace fossils provide a source of evidence for ancient land ecosystems that is independent of fossils. Conflicting indications of habitats from fossils and from features of fossils are not necessarily mistakes of interpretation, but may instead point the way to fundamental questions about the role and evolution of organisms on land.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 11 Topographic relief as a factor
Variation in the nature of soils with topography can readily be appreciated by comparing thin, rocky soils of mountain tops with thick, fertile soils of grassy lowlands. But even in relatively featureless lowlands, the nature of soil varies profoundly depending on whether it is well or poorly drained. This aspect of soil formation is not completely independent of other factors because vegetation, microclimate and age of land surfaces vary in different parts of the landscape. In small areas, however, other factors may be limited to such an extent that variation in soils across the landscape constitutes a true toposequence, and features of the soilscan yield reliable topofunctions. The undulating till ridges abandoned from the last glacial advance (CaryTill, 1 4Ka) in northcentral Iowa, USA, provide examples of topofunctions that are all themore striking because of the subtle topography involved a relief of only 5m from ridge top to boggy bottom. Soils from the top of the ridge to the bog show marked decreases in gravel and mean grain size and increases in thicknesses and amounts of clay, organic matter and carbonate (Fig. 11.1). As an example of a topofunction, the increase in thickness (y) of the soil with distance from the summit down the slope (x) can be expressed by a polynomial equation: y=1.41-0.91x+0.49x2-0.034x3(Ruhe 1969). Other landscapes are rather more intimidating and inspiring, as when one is gazing over row upon row of alpine ridges and peaks. Such bold landscapes also can be resolved into particular elements characterized by distinctive slope-related processes. On steep alpine slopes, for example, vegetation is sparse. The soils are eroded by snow melt and churned by frost heave. Rocks fall from cliffs above. Such processes result in thin, shallowly rooted, little-weathered and rocky soils, adjusted to the local environment of mountain slopes. The scale and scope of such processes in many such landscapes do not lend themselves to strict analysis as topofunctions because of the great variation in climate, vegetation, parent materials and age of land surfaces up and down 160
mountainsides. A less rigorous term for lateral variation in soils across such landscapes is ‘catena’.Derived from theLatin for ‘chain’,its links aremetaphors for adjacent soil types. The usual unit for a soil of a particular kind found in a particular subenvironment of the landscape is the soil series (see Fig. 6.1) or pedotype (seeFigs6.2-6.4).
Indicatorsof past geomorphological setting Within modern landscapes, the following variation in features of soils is found on descending from high to low elevations. Soil profiles are thicker, more organic, more moist, more strongly colored, less differentiated into horizons, more alkaline, more likely to contain soluble salts, lesslikelyto have well-developed hardpans of truncated horizons and more likely to be formed on alluvial or colluvial parent materials (Buol et al. 1997). These are all major differencesin soil character. They are, however, relative differences. They may be observable in paleocatenae but call for exceptional lateral exposure to be discernible. Differences in age of different parts of paleocatena make their interpretation difficult. Only features that independently indicate a specific geomorphological setting are considered here. Soil creep Soil creep can be envisaged as a kind of landslide but with a slow, steady movement of the surface soil rather than a rapid movement of a large mass of soil and rock. Like the fence posts that gradually tilt downhill, rocks and other erect objects in the soil curve toward a prone position near the surface of the soil. This constitutes the best evidence of soil creep in paleosols (Retallack & Mindszenty 1994). Such indications of creep imply a former slope, as well as the former direction downhill. The steepness of the slope may be difficult to determine
Topographicrelief as a factor
till
hill loom
bogsilt
muck
- _ - depth to carbonate
-5 2
n
8 3
1 61
-
gravel (greater than 2 mm)
^ I E
50%
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Figure 11.1 Downslope variation in percentagegravel,clay. mean grain size. organic carbon and carbonatein a soil catena on gently undulating Wisconsinan ( 14 Ka) ridges of CaryTill around a bog near Jewel], north central Iowa. USA (adaptedfrom Walker 1966:with permissionfrom Iowa StateUniversity Press).
carbonate
- 1 E - 2
40
60
80
eometric mean grain size (pm)
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becausecreep can be initiated on slopesas gentle as 2.5", especiallyduring the summer melting of frozen ground (Washburn 1980). The leaning and bowing of tree trunks on slopes, often attributed to soil creep, is a more complex phenomenon involving snow damage of young trees, and optimization of light and support (Alexander 1997). Unconformities Sedimentary sequences may preserve the lower parts of rugged landscapes of the past. A rise in sea level or the dammingof avalleybylandslidecanleadtoburialof the landscape in sediment. Fossilized landscapes of this kind are widely known and include sea stacks, coastal clifs (Dott 1974),boxcanyons (Leary 1981),radialdrainage (Stewart et al. 1986),glacialvalleys (Andreiset al. 1986)
.
40
.
.
60
distance ( r n )
80
20
40
60
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80
and fjords(Herbert 19 72). The most spectacular examples of these are preserved at the contact between rocks different in age by millions of years, such as sandstone overlying granite (Figs 11.2 and 11.3). There also are disconformities within sedimentary sequences. Paleochannels of ancient streams are the most obvious of these because filled with contrasting materials, usually sandstone and conglomerates, compared with the clayey floodplain deposits. Nevertheless, stream-channel depositsaccumulate within floodplains in such a way that the erosional disconformity at their base is not a true reflection of paleotopography at any one time. Certain features of sediments, such as the height of low-angle cross-sets of alternating shale and sandstone formed in point bars and levees can be used as a guide to local topographic relief around stream channels, as can the tracing of individual paleosols over
162
Chapter 11
A.
SLIOCH
-
8.
F
I KH
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I KH
~LEVISIAN
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Figure 11.3 Interpretedpaleogeographyof north-west Scotlandat 8 1OMa, during deposition of large alluvialfans of theTorridoniansequencefrom crystalline highlandsto the west, now riftedoff intoGreenland (fromWilliams 1969, lournal oj’ Geology 77,Fig. 17;reprintedwith permission from the University of ChicagoPress).
F = FAULT
Figure 11.2 Geologicalsectionsdrawn with no vertical exaggeration of paleotopographic relief onkwisiangneiss (>1450Ma) and the unconformablyoverlyingTorridonian(810 Ma) alluvialdeposits in Slioch and Quinag Mountains,north-westScotland (from Williams 1969,lournalofGeology77,Fig. 2; reprinted with permission from the University of Chicagopress).
levees and point bars into the floodplain (Clemente & Perez-Arlucea 1993;Willis &Behrensmeyer1994). Unconformities that represent erosionalgullies or terraces of floodplains are more subtle than sandy or conglomeratic paleochannels, but can be recognized from associated paleosols (Bestland 1997). Each paleosol represents an unconformity, the duration of which is reflected in its degree of development.Even longer periods of nondepositionmay be marked by abrupt changes in the nature and spacing of paleosols. In Badlands National Park, South Dakota, USA, each of the rock unitsrecognizedonthe basis of distinctivecolor,nodules and other features of the paleosols in them has weakly developedwell-spaced paleosols at the base, and strongly developed, closely superimposed, sometimes overlapping and eroded paleosols at the top. These differences represent a kind of equilibriumbetween subsidence and sedimentation attained near the end of deposition of each unit until a period of gully erosion and fluvial terraces ushered in a time of more rapid sediment accumulation (Retallack 1986b). Such ‘ageing upwards’ sequences of paleosols are recognized also in other sequences (Fig. 11.4).The abrupt boundaries between such genetic units represent erosionallandscapes.
Erosional planes Some paleosols have erosional planes within them as well as on top. These tend to form in paleosols developed in higher parts of the landscape and so prone to erosional truncation followed by addition of new material. Such features are difficult to interpret, even in soils (Johnson & Watson-Stegner 1987). The abrupt and
Topographicrelief as a factor
163
Figure 1 1.4 Paleogullyinto strongly developedsequenceof paleosols(dark colored),filled with alluviumincluding weakly developed paleosols (lightcolored) in theUpperTriassic(230Ma),Chinle FormationinPetrified Forest National Park, Arizona, USA. The hill inthe foregroundis 11 mhigh(fromKraus&Middleton1987: with permission from the GeologicalSociety of America).
laminated tops of many hardpans, such as petrocalcic horizons, also have been thought to be the result of erosion back to that level (Klappa 1979b),although corrosion by perched groundwater is also a viable explanation (Machette 1985). Evidenceof erosion also is providedby lines of stones along an erosional plane in a clayey or silty soil. It takes a powerful stream or flood to transport pebbles, compared with silt and clay. They would be anomalous in a muddy or silty stream and would be accompanied by other pebbles in a gravel stream. Stone lines are evidence of erosional breaks in well-drained soils in areas of moderate relief (Ruhe 1 959). This does not necessarily mean a hilly landscape because appreciable relief may develop around large streams during the dry season. Another way in which stone lines can form is by the burrowing activity of gophers because stones in the soil fall through burrows to a zone beneath the nests that is less actively bioturbated (Johnson et aI. 1987).Suchstonelines, however, areirregularinshape and often festooned under individual nests. Erosional stone lines, on the other hand, form more planar surfaces. Gilgaimicrorelief Gilgai is not only a form of topographic relief on a small scale, but distinctivekinds of gilgai form on level and on sloping land. There are two main kinds of gilgai (Paton 1974): those in which the depressions are equant in plan (nuramgilgai)and thosein which they are elongate
(linear gilgai). Nuram gilgai is common on level areas (see Fig. 4.2). This kind of relief can itself be complex, with shelves and inner depressions, in places up to 3 m deeper than the mounds. Linear gilgai has not been found to have such great relief but forms perpendicular to contours on gentle slopes of between 15" and 3 O (see Fig.9.13).
Salt crusts Solublesalts may form in soilsof dry climatesin avariety of geomorphologicalpositions, but the thickest surficial crusts are generally found in lower parts of the landscape. Soluble salts accumulate where ponded waters evaporate. These salts would continue downhill and out to sea were they not so close to the surface in the shallow water tables of desert playas, supra-tidal flats and overirrigatedlowlands (Jenny 1941).Bedsof saltminerals in sedimentary sequences commonly have been interpreted as subaqueous precipitates of saline lakes and seas, but careful mapping of associated sedimentary facies has shown that some have formed in supratidal mudflats (sabkhas), l i e those around the Persian Gulf (Evanset al. 1 973), or dried-out desert playas, like Lake Magadiin southernKenya (Eugster 1969).Mudcracked dolomitic mark riddled with rosettes of trona (NaHCO3.NaC0,.2H,O) in the Eocene (48Ma) Green River Formation of Wyoming, USA, once considered lacustrine. are now thought to have formed at the surface of a dried-out lake (Eugster & Hardie 1975) and can be
164
Chapter 11
considered a kind of desert paleosol (Solonchak of FA0 or Salid of USDA). Examples of marine-influenced evaporitic supratidal paleosols also are known from the rock record (West 1975).
Cumulic horizons Although soil formation and sediment accumulation are in some ways antagonistic processes, small amounts of windblown dust and floodborne alluvium do accumulate in soils. The rate of influx of these materials in well-drained soils is commonly so slow that they are incorporated into the fabric of the soil by roots and burrows and do not form distinct surficial layers (Muhs 1983: Mason &Jacobs1998). In lowland soils, on the other hand, sedimentary layers may form at the top of the profile as vegetation grows progressively upward. These are called cumulic horizons (Soil Survey Staff 1975). There is a fine line between soils with thick cumulic horizons and ordinary sedimentary deposits of lakes and streams. The distinction is largely a matter of root and burrow density, which reflects the relative contribution of pedogenic, as opposed to sedimentary, processes. Cumulic horizons in paleosols are evidence that they formed in low-lying parts of the landscape.
Indicators of past water table The geomorphological setting of paleosols may not always be clear, but their relationship to the water table often is obvious and has a clear bearing on position in landscapesof thepast.The watertableis that levelbelow which the earth is permanently saturated. It separates a surface region of active leaching and more or less oxidizing chemical alteration from a stagnant region of little chemical change below. The water table may seem level, when exposed as a pool of water at the bottom of a well. On a broader scale, however, it is not tabular. The water table is deeper on hillslopes than in valley bottoms, and follows elevation to a certain extent. It may be perched on shaly impermeable subsurface layers or can be deeper within friable sandy materials. It is closer to the surface in humid forested regions than in barren deserts, and varies in depth with the availability of moisture through the year, and even daily. There are three possible situations for a soil with respect to the water table (Fig. 11.5).It can be entirely above the water table,
well drained crocks
seasonally wet root traces
waterlogged peat
humus siderite or
Figure 11.5 Characteristic patterns of root traces, cracking, peds, nodules, color, coal distribution and relict bedding in formerlyweI1-drained,seasonally wet and waterlogged paleosols.
in part or wholly within the zone of water-table fluctuation, or entirely below the water table. Soils of the first kind are normal. The second kind is floodprone or seasonallyswampy soil.The third kindis apermanentlywaterlogged soil of swamp or marsh. A variety of features allow discriminations between these general kinds of paleosols. Some of these features of paleosols are prominent enough to be detected by satellite-based remote sensing (Kraus 1992). Soil horizons The nature of horizons formed above the water table, within the zone of its fluctuation and below it are distinct. Below the water table, stagnant groundwater allows persistence of chemically reduced minerals, such as siderite and pyrite. These iron-bearing minerals, like the clays that contain them, are drab and gray with iron in the ferrous state, and with admixed organic matter (Kraus 1997; Retallack 1999b). Because water in this part of the soil is slow moving and carries little material away, weathering of minerals below the water table is slight (C horizon). Within the zone of water-table fluctuation, the alternation of oxidizing and reducing conditions produces strongly developed nodules or mottles of gleyed (Bg)
Topographic relief as a factor horizons (Kraus 81Aslan 1993; McCarthy et al. 1998). Streaks of red and orange (high chroma mottles) are strikingin soils of this kind (seeFig. 4.7). Most paleosols have some of this character, however, because of color changes upon burial, as a result of burial gley and reddeningof ferrichydroxideminerals(seeFig.7.10).Some caremustbetakentodisentanglefeaturesof theoriginal soil from those imparted by conditionsof burial (PiPujol &Buurman 1997). Above the zone of water-table fluctuation are formed most other kinds of soil horizons, such as clayey subsurface (Bt)horizons (Cremeens &Mokma 1986).Characteristically these are yellow, brown or red with iron oxide and hydroxide minerals, such as hematite and goethite. These warmcolorsalsomay be alteredduring burial (see Fig. 7.10),but washed-in (illuvial)clay skins and deeply penetrating root traces remain as evidence that paleosols were welldrained (Retallack 1991a.e). Calcic horizons differ in form, depending on whether they formed above or below the water table. Groundwater calcrete, precipitated below the water table, can form a simple cement that fills the open spaces between larger grains (see Fig. 7.4B; Mann & Horwitz 1979). Caliche, on the other hand, forms above the water table in aridland soils and has a more complex appearance (Wieder&Yaalon 1982;Retallack 1991a).It is usually micritic and can be seen to replace and surround clastic grains. Even such weather-resistant grains as quartz may show irregular dissolved margins (caries texture) and abnormally loose packing when viewed in petrographic thin section (seeFig. 7.4A). In hand specimen, calicheis generally nodular to irregular in shape. Multiple generations of cement and local colloform or sparry fillings of cavities may indicate that the nodule could move slightly within the soil (displacivetexture: see Fig. 3.17C). Root traces The leaves of green plants require carbon dioxide to produce organic matter and oxygen, but roots are aerobic respirers. Even with special aeration structures, such as aerenchyma,they do not grow far below the water table. This is most obvious in waterlogged soils, in which roots spread out laterally rather than vertically (seeFig. 3.3). The water table indicated by root traces in a paleosol is not alwaysinthesameplaceasindicatedby soilhorizons
165
becauseof two commoncomplications.Where the soil is waterlogged for most of the year, except for a short dry season,root traces may extend deeply into a gleyed horizon that is dry for a part of the year. In addition, it could be that fossil roots were preserved from a time of better drainageby the general rise of the water table with continued subsidence and sediment cover over the soil (Coleman 1988).These two situations can be resolved from the relative proportions of horizontal vs. vertical roots and their quality of preservation.
Burrows Afewcreatures, suchascrayhhandcrabs, liveinwaterfilled burrows (Thorp 1949).but most soil animals are air-breathers(Buatoiset ul. 1998).Earthworms,insects and mammals do not burrow below the water table. Where their burrows are dense,their lower termination can be an obvious indicator of the usual level of water table (Retallack1976).
Soil structure Cutans are most characteristic of those parts of soils above the water table (Brewer 1976). Especially common in well-drained parts of soils are clay skins (illuviation argillans),washed down into cracks from higher in the profile. Cutans formed above the water table also tend to be yellow, brown and red in color with sesquioxides, like other well-drained parts of soils. In contrast, waterlogged soils and parts of soils have weakly developed cutans that tend to be dark with organic matter or manganese stain. Cutans are themain evidenceinpaleosolsof peds (soil clods) because open spaces are crushed out of paleosols during burial. Peds are largely a phenomenon of welldrainedupper portionsof soils, where thesoilcanshrink and swell and where there is transport of materials. Waterlogged soils and parts of soils, in contrast, have weakly expressed soil structure.
Soil nodules By their chemical composition,soil nodules and concretions reflect the availability of oxygen at the time they formed (see Fig. 3.14). Pyrite, marcasite and siderite nodules are characteristic of permanently waterlogged
166
Chapter 11
soils (Altschuleretal. 1983;Mooreetal. 1992). Goethite and hematite concretions and nodules are found in welldrained soils. Concretionary banding can, in some cases, reflect periodic change in chemical conditions (seeFig. 3.15).Iron deliveredto concretions in areduced state can be oxidized in concentric bands in ferric concretions. Complex nodules with reduced centers, oxidized rinds and septarian cracking also may reflect fluctuations of oxygenation inpaleosols (Nahon 199 1). In each of these cases, care must be taken to ensure that the nodules are an original part of the soil rather than a phenomenon of deep burial or recent exposure. Microfabric Soil structures are just as diagnostic of the influence of water table when examined in petrographic thin sections as with the naked eye. The general microfabric of soils is very different under waterlogged and welldrained conditions. Bright clay fabric (sepic plasmic fabric: see Fig. 3.16) is a feature of well-drained parts of soils (Brewer & Sleeman 1969). These highly birefringent segregations of orientated clay form in much the same way as clay skins: by washing down cracks or by the shearing associated with clay swelling. Waterlogged soils, on the other hand, have microfabrics similar to the parent material of a soil (asepic)or massive, nearly isotropic, fabrics (undulic and inundulic: see Fig. 3.17). These massive fabricsmay be the consequence of flocculation of unoriented clay and organic matter (Brewer 1976).
Interpretingpaleocatenae A soil catena can be characterized easily enough as the soils along the present land surface. Understanding how that particular suite of soils formed is not so easy. Nor is it easy to determine which of a series of interdigitated paleosols in a sequence constitute an actual ancient landscape. Soils are very much a part of the landscape,and their history isjust ascomplex andintriguing. Lateral variation in paleosols Karst topographydevelopedon limestoneor dolostoneis one of the most irregular of landscapes. Paleokarsts. particularly those of the humid tropics, may show con-
siderable topographic relief,steep cliffs and caves (James & Choquette 1988; Coxon & Coxon 1997). The limestone surface is usually little altered beneath a sharp weathering front. However, the filling material may range from manganous to bauxitic in composition, depending on the degree of drainage of the karst depressions (Bardossy 1982). An added complexity in the evolution of karst topography is the collapse of limestone caves.The fillingmaterial in this case is a complex mix of brecciated fragments of roof rock, talus cones of breccia, silty deposits of streams and shaly deposits of lakes. These can be dissolved and eroded during cave formation for another generation of filling material. Such complex sequences of cave fill have been reconstructed in detail for fossiliferouscavedeposits of Swartkrans and other comparable early (1-2 Ma) human-ancestor fossil sitesinsouth Africa (Brain 198 1;Butzer 1983). Paleosols also vary laterally at major geological unconformities. A variety of paleosols have been documented in granite and amphibolite below the Torridonian sedimentary sequence (8 1 0 Ma) in the north-western corner of Scotland (Retallack & Mindszenty 1994).Regional topographic relief here has been measured to be at least 1OOOm. Hills of Lewisian gneiss have been exhumed from under their Torridonian cover by more recent erosion (Fig. 11.2). On a regional scale, paleosols are thicker (3 m) on the undulating pediment to the south-west, whereas paleosols are thin (only 10cm) on hilltops and valley walls. These differences, to some extent, reflect differences in age of the surfaces, including a period of glacial valley cutting (Williams & Schmidt 199 7). These differences could be substantial because the gneiss is dated at 2 600-2 700 Ma and was metamorphosed at about 1300-1 600 Ma, whereas the cover sequence accumulated at 8 1 0 Ma. Lateral variations in paleosob of alluvial sequences have been widely observed but are difficult to document and interpret (Kraus 1987, 1996, 1997: Wright 1992b). Some of the especially obvious paleosols have been mapped laterally in Miocene (8 Ma) alluvium of the Dhok Pathan Formation of the Siwalik Group in northern Pakistan (Fig. 11.6).My own examination of these sequences has revealed many more paleosols than this(Fig. 11.7).but thisdoesnot alterthedemonstration that paleosols vary in age laterally.This is revealed by the truncation of a geosol by a reversal of the Earth’s mag-
Topographic relief as a factor
blue-gray sandstone
gray paleosols
0buff
167
2 km
sandstone =red
paleosols
m
IOm
Figure 11.6 The lateraldistribution of especially well-developedpaleosols,and channeldeposits of an ancestral Indus River (bluegray system)and of drainage of the Himalayanfoothills (buff system),and its relationship to a paleomagneticreversalin the Upper Miocene (8.3Ma),DhokPathanFormation,near Khaur, northernPakistan. Abbreviations at the top represent individual measured sectionsused to make this diagram(fromBehrensmeyer&Tame1982: with permission from theInternationa1 Organizationfor Sedimentology).
netic field (Tauxe& Badgley 1988).The varied paleosols along the time plane revealed by this paleomagnetic reversalare a better guide to former catenae of soils than the lateral tracing of paleosols in the sequence. An ash bed deposited in a short time would be another useful indicator of an ancient land surface. Such a situation is known in Plio-Pleistocene (1-2 Ma) alluvial deposits east of Lake Turkana in north-western Kenya (Burggraff et al. 1981).Without such exceptionalindicators of time planes, paleocatenae cannot be reconstructedprecisely. Vertical variation in paleosols Another way of characterizingpaleosols is by measuring a vertical section of the kind used in stratigraphic logging. The past geomorphological setting and relationship to water table of each paleosol can generally be established from the various criteria already outlined. With information from associated sediments, such as sedimentary structures of paleochannel or lake deposits, these can be assembled into a general model of ancient landscapes.It may at 6rst seem paradoxicalthat the lateral relationships of paleosols can be recon-
structed from their vertical relationships, but this is an old technique of sedimentologicalinterpretation, usually called Walther’s Facies Law (Boggs 1995). Simply stated,the idea is that depositsformed side by side in nature are usually preserved on top of each other because of lateral shifts in the environments in which they form. A falling sea level, for example,results in a superposition of river, delta and beach depositsover thoseof theocean. A lateral shift of a meandering stream channel away from its point bar results in the superposition of soils in the silts and clays of its levee that are thicker and better developed the further they are from the stream and the higher they are above the paleochannel deposits. The key to this common kind of geological inference is the genetic relationship between different parts of the landscape.This technique cannot be used across major unconformities that separate unrelated depositional systems. Major unconformities are obvious when they separate granitic basement from overlying sediments, but profound breaks within sequences of paleosols also can be detected by the degree of development,spacing anderosion of paleosols (Retallack 1998a). The Miocene (8.3Ma) upper DhokPathan Formation near Khaur in northern Pakistan is a genetically related
grain size
development calcareousness
hue
KEY ***a
claystone breccia sandstone siltstone claystone
I\wl E E n g
a
ripple marks
red color
BHuRAe
brown color
m
drab haloed root traces
SARANG
PlLA
ferruginous concretions
p J calcareous nodules __ 0ferruginized surface F;1root traces snails burrows
[dJ ;
and
fossils not
mlinof trench section
BHuRAF
SARANG BHURA SONITA
Figure 11.7 Adetailedstratigraphic sectionshowingdifferentkindsof paleosolsin theupper Miocene (8.3Ma),middleDhok PathanFormation, inKaulial Kas, near Khaur, Pakistan (data from Retallack 1991a).
Figure 11.8 Reconstructedpaleoenvironment of northernPakistan during the Late Miocene (8.3Ma) deposition of the Dhok Pathan Formation of the Siwalik Group (fromRetallack 199 la: with permission from OxfordUniversity Press).
170
Chapter 11
sequenceof paleosols (Fig. 11.7)that can serve as a basis for reconstructing an ancient landscape independently of the lateral relationships already discussed (Fig. 11.6). A carefullyexcavated and recorded stratigraphic section of these deposits has revealed six kinds of paleosols.Two of these (Pandu and Khakistari pedotypes in Fig. 11.7) are drab colored and associated with deposits of the ancestral Indus River (blue-gray sandstones), and the other four pedotypes are associated with deposits of streams draining the Himalayan foothills (buff sandstones). The drab paleosols are interpreted as waterlogged lowland soils formed under tropical forest. The better drained, perhaps seasonally dry, foothill streams were flanked by very weakly developed paleosols with sparse or small root traces, like that of early successional vegetation. These (Sarang pedotype) have abundant relict bedding and reflect the most disturbed nearstream areas. Another kind of paleosol (La1pedotype), forming floodplains of the foothills drainage system, is thick and red. with broad clayey B horizons and numerous large root traces. These may have supported tropical
monsoon forest, similar to those found in modern Uttar Pradesh. Weakly developedreddish brown (Bhura pedotype) and red (Sonita pedotype) paleosols are intermediate in development and represent streamside woodland at intermediate stages of plant ecological succession. Two other pedotypes (Kala and Pila) have fine root traces and granular structure (near mollic),as well as indications of seasonal waterlogging in the form of a subsurface horizon stained with manganese (a placic horizon of Soil Survey Staff (1975) for Kala), and local ferric concretions (for Pila). Both probably supported lowland wooded grasslands in floodplain swales. This is not a precise reconstruction of a particular landscape but rather a general concept of the relationship between observed sedimentary environments and paleosols (Fig. 11.8). A part of the charm of landscapes is that they are as individual as people, with a combination of inherited and acquired characteristics. Soils are an important part of landscapes, and paleosols provide a partial narrative of landscape evolution.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 12 Parent material as a factor
The rock or sediment on which soilsdevelopis its parent material. It is a starting point for the process of soil formation. In the early stages of their formation, soils are not much different from their parent materials. With time, fewer and fewer features of the parent material persist, and ultimately the soil takes on a n identity of its own. The nature of parent material is a base line that must be known in order to assess the amount of soil formation that has taken place. Perhaps the clearest case where parent materials are known is soil developed on igneous rock, such as granite. In unweathered granite, there are no clay or carbonate minerals. Even small traces of these around the margins of crystals in petrographic thin sections are a clue that the rock is weathered. Small amounts of clay, carbonate and iron oxyhydrates may be found in a transitional horizon (C or saprolite) between fresh bedrock and the soil proper. This may look like parent material in its crystalline texture and other features but is considerably altered compared with true parent material. Material above this transitional horizon (in horizons A, E and Bt) may show more characteristic soil textures and structures. Such soils with a deep transitional saprolite to unweathered bedrock often are assumed to have formed on fresh bedrock. This may be a fairly safe assumption if especially weatherresistant parts of the soil, such as quartz veins, persist to the top of the profile (Fig. 12.1) as a relict structure. There could also be chemical relicts of granitic structure, such asdomains of ferric beidellite (asmectite clay) after hornblende, mixed with aluminous beidellite after plagioclase (Nahon 1991). But what about the addition of airborne dust, landslidedebris or floodborne silt on top of bedrock? These additional layers of very different materials can be easily obscured by soil formation. Even more difficult to detect would be additions of salts or other materials from groundwater solutions. All these doubts that could be harbored about the
parent material of soils on bedrock apply with equal force to soils developed on alluvial parent materials. In addition, however, the differentiation of soil horizons in alluvium could be as much a reflection of the alluvial deposition of sandy over clayey layers as the result of the washing down of clay into a subsurface (Bt) horizon during soil formation (Fig. 12.2). If the clayey subsurface horizon was primarily sedimentary, then the soil would be regarded as young and weakly developed. If. on the other hand, the clay is entirely pedogenic, then the soil would be interpreted as old and strongly developed. In most cases parent material is an independent variable in soil formation. Igneous, metamorphic and volcanic rocks are formed in places and by processes removed from the Earth’s surface. Sediments are often parent material for soils and are derived from soils, but the ways in which sediments are sorted and distributed are so variable that these, too, are independent of soils. Geological maps of many regions reveal a variety of bedrock and sediment that is reflected in variation of soils formed on them, or lithosequences of soils. Certainly there are situations where one rock type is more weather resistant than others and so forms steeper slopes, like those of metamorphic hornfels around large granitic intrusions (Ollier 1965). These weather-resistant ridges of rock also may encourage different microclimatic conditions and vegetation and have younger cliffs and scarps than nearby deeply weathered rocks. Thus, despite the abundance of lithosequences, in only a few of them are all of the other soil-forming factors well enough controlled to yield useful lithofunctions. The sheer chemical and physical variation of soil parent materials also is a problem for studies of their role in soil formation. The roles of some general properties of parent materials in soil formation are becoming well understood, and these will be considered beforecommon kinds of parent materials are described. 171
172
Chapter 1 2
,
PARENT MATERIAL OF GRANITE ? granite
quartz vein
0
sond
y
sill m r e d clay
If parent material were a natural resource and soil a commercial product, economy would dictate the manufacture of soil from parent materials most like it in bulk properties. Some parent materials rapidly and with little energy are converted to soil, whereas others are resistant to weathering. Parent materials
'lay
r
OR OF SEDIMENTARY COVER ?
General propertiesof parent materials
SOIL HORIZONS 7
T
can thus be viewed in a functional way by their properties that promote soil formation. This view has the virtue of glossing over the great variation of parent materials and focusing only on key properties. As in other considerations of factors of soil formation, those parent materials observable in paleosols are stressed over those only easily demonstrated in modern soils.
OR SEDIMENTARY BEDS 7 ripple
Figure 12.1 Soils developeddirectlyon bare granite can appear somewhat similar to soils developedon a sedimentary cover on granite, although these latter may show quartzveins truncated well below the surface and a sharper base to the clayey subsurface (Bt) horizon.
brown 'lay
Figure 12.2 Soils developedon fairly uniform alluvium can appear similar to those developedon a less homogeneous sedimentary sequence, although these latter may have more sharply definedlithological boundarlesandmoreobvlousrellctbedding.
Parentmaterialasafactor Uniformity Few parent materials are completely uniform in composition or structure. Almost always there is some irregularity, some foliation, veining, jointing or layering that in some cases aids and in other cases hinders soil formation. Some sedimentary layering promotes soil formation, as in a silty cover on bedrock (Fig. 12.1) or a sandy over a clayey layer of alluvium (Fig. 12.2).Inbothcases,friable surficial material has been created by nonpedogenic means. Other situations of surficial cementation or of finely interbedded sequences of clay and sand could be imagined that were not so conducivefor soil formation. In some cases, nonuniform parent materials may be difficult to detect in soils and paleosols. Anomalous minerals may be a clue. A few grains of primary minerals that are not found in the parent material can provide an indication of later additions. Quartz, for example, does not occur in basaltic phonolite, nor does olivine in granite. Grain-size irregularities also may be an indication. Formation of clay skins and oxidized grain coatings may obscure in the field what are obvious sedimentary layers of different grain size under the petrographic microscope. Perhaps the most obvious of grain-size irregularities are stone lines: very thin zones of scattered pebbles much larger than other grains in the paleosol. Stone lines often can be seen to lie on erosional planes (Ruhe 1959). These and any other sharp contacts in a soi1 should be grounds for suspicion that the soil developed in several stages on different layers of parent material (Johnson & Watson-Stegner 198 7). Unusual horizons that do not seem to fit normal sequences of horizons also can be grounds for suspicion: two distinctly colored clayey subsurface (Bt)horizons, for example,or a buried zone of leaf litter and stick debris of a former surface (A) horizon. Another source of nonuniformity that aids soil formation isjoints: deepcracks in large bodies of igneous or other hard rocks formed by contraction on cooling and expansion following the unloading of exposure at the surface.These deep cracks allow the deep penetration of air and water so that joints become more deeply weathered than the rock they enclose. Progressive weathering of blocks of rock defined by joints leads to the formation of Liesegang banding and also of unweathered corestones with concentric rims of weathered rocks (Au-
173
gustithis & Ottemann 1966). Liesegang banding may form in a similar way or by reaction fronts in slowly diffusing pore waters (Ortoleva et al. 1986). This onionskin weathering is common in soils and paleosols (Williams 1968;Boucotetall974). The effect of foliation, folding and veining can be more complex. Quartz-rich parts of gneissic foliation, quartzite beds folded into schist and quartz veins through granite all resist weathering more than their matrix. These commonly stand out as low ribs on the ground or within a soil or paleosol (Retallack & Mindszenty 1994). These situations are not difficult to identify because the nonuniformity commonly is oriented at a high angle to soil horizons. Induration The formation of soil proceeds more rapidly on materials already friable and loose, such as sediment or volcanic ash, because spaces between the grains allow free movement of weathering fluids. Roots and burrows readily penetrate soft materials, but are modeled around corestones, nodules or other originally hard parts of the soil (Haasis 1921). In paleosols in which the whole profile has been buried and altered to solid rock, features such as root traces and burrows are evidence of formerly unindurated parts of a profile. This also may be apparent from abrupt erosional discontinuities on the bedrock and a composition that clearly was originally hard, such as basalt. Despite such alteration during burial, a scale of friability devised for soils can also be useful for study of Tertiary and Quaternary paleosols (Table 12.1). The way in which indurated bedrock is converted to
Table 12.1 Scalefor strengthof soils andpaleosols (Ollier1965). Very strongly indurated Strongly indurated Moderatelyindurated Weakly indurated Friable
Hammer tends to bounce off rock Easilybrokenwith a hammer Broken by kick with boots,but not by hand Brokenby hand,nodisintegration in water Runs throughlingers, disintegratesin water
174
Chapter 12
soil varies considerably.Foliated or jointed rocks may be loosened by preferential weathering along these planes of weakness. Crystalline rocks, on the other hand, are loosened by destruction of those minerals, such as feldspar, most susceptible to weathering. This local unevenness of soil formation can be seen in both soils andpaleosols(Grandstaff etal. 1986; Nahon 1991). Effects of induration on soil formation are especially well illustrated by soils and paleosols developed on limestone and calcareous sand. Soils formed on indurated limestone and marble develop by chemical dissolution through the action of acidic rain and groundwater. Soil material accumulates from insoluble residues of the limestone or sediment washed or blown in from nearby. This noncalcareous soil material is usually sharply divided from limestone by an abrupt weathering front uennings 198 5). Numerous fossilized examplesof these ‘Terra Rossa’ (Rhodustalf, Rhodudalf) soils are known (Wright & Wilson 1987). In contrast, soils formed on calcareous sands, with grains of oolites, foraminifera or shell fragments, do not show such an erosional surface and may be more deeply and thoroughly weathered. There may be nodular and other complexly cemented zones, but not an abrupt and distinct difference between limestone and soil (Wright 1982).Similardifferencescan be seen between Miocene paleosols formed on hard carbonatite volcanic flows, as opposed to friable carbonatite tuff, in south-western Kenya (Retallack 199 1a).
A . 10.000 years ,----//,-grus m
Grain size The size of crystals of igneous rocks or of grains of sedimentary rocks controls the rate of weathering in several ways. The larger the grains or crystals, the larger are the spacesbetween them opened by weathering. This allows greater and deeper flow of weathering fluids.The differential weathering of coarse- and fine-grained parent materials is especially well seen in soils and paleosols that formed on till or conglomerate containing pebbles of different lithologies. Individual pebbles weather first on the outside. This clayey and oxidized zone gradually extends inward as a weathering rind. In very old (>lo0 Ka) tills of theeastern SierraNevada of California,USA, granitic boulders were weathered to a loose sediment of angular coarse grains called grus. Coarse-grained pebbles and cobbles of porphyritic andesite are weathered to the core, but those of fine-grained basalts have weathering rinds only 2 mm thick (Fig. 12.3). Similarly.paleosols on fine-grained amphibolite dikesare thinner than those developed on the coarse-grained granitic rocks along major unconformities (Retallack & Mindszenty 1994).In sedimentary sequences also, grain size plays a role in the flow of water and air through the soil. Welldrained gravelly or sandy soils and paleosols are more deeply weathered than well-drained clayey ones (Retallack 1991a; Birkeland 1999). Very fine-grained patches of amorphous aluminosilicates derived from weathering or microbial activity can protect the large
8. 140,000years m
C. more than 200,OOOy
weathering rind A-andesitic 8-basoltic m G-granitic
--
soil
Figure 12.3 Differentialdevelopmentof weathering rindsover time in bouldersof different internal grainsize in Pleistocenetills near LakeTahoe,California,USA (adaptedfromBirkeland1999:with permissionfrom OxfordUniversityPress).
Parent material as a factor grains they encrust from further weathering (Nugent et al. 1998). Crystallinity Resistance of minerals to weathering is to some extent conferred by the rigidity, and proportion of defects in, their crystal structure (Eggleton 1986). One of the most stable of minerals is quartz, which is a threedimensional network (tectosilicate)of silica tetrahedra (SiO,). Phyllosilicates with an open and expanded (2 :1) structure, such as biotite, muscovite and smectite, are more easily weathered than those with a nonexpanded (1: 1)structure, such as kaolinite. Minerals in which silica tetrahedra are in chains, such as hypersthene or augite, or isolated,such as olivine, are even more readily weathered. Volcanic glass lacks crystalline regularity, and like window glass, can be regarded as a chilled fluid. This process of chilling, along with bubbling of frothing volatiles, is responsible for the sharp-edged fragments of the rims of bubbles called volcanic shards. Soils formed of such glassy volcanic airfall materials as ash and pumice have distinctive properties. In the humid temperate climate of the North Island of New Zealand, the half-life of andesitic glass in soils has been estimated at only 7000 years (Neal1 1977). The initial alteration products are amorphous materials, such as allophane, imogolite,opaline silica and humus complexeswith iron andalumina. In time, clay minerals form. Even when extensively weathered to clay, ashy soils retain a low bulk density and a porous structure because of cavitiesleft by weathered volcanic shards. Soils developed on glassy volcanic materials may be more clayey, humic or deeply weathered than soils formed on more highly crystalline materials under otherwise comparable conditions of weathering. Mineral composition Minerals are in effect chemical compounds and this also plays a role in their vulnerability to weathering. Chemical compounds vary in energy of bonds between their constituent atoms (Keller 1954) and in the size and charge of participating ions (Mason & Moore 1982).The strongest bonds are those between Si4+and 0 (3110k~alrnol-~) and A13+and 0 (1793 kcalmol-'), but the bonds between common basic cations (Ca2+,
175
Mg2+,Fe3+,K+ and Na+) and 0 are weak (839, 912, 919, 322 and 299 kcalmol-l, respectively). For these reasons the most stable of minerals are quartz (SiO,) and gibbsite (A1203).Within the framework of silica and alumina in other aluminosilicates, large ions (Ca2+, K+ and Na+) are less easily dislodged by hydronium (H') during hydrolysis than are small ones (Mg2+,Fe2+). Once in solution, however, the ratio of charge to ionic radius (or ionic potential) is such that most of these ions remain in solution (ionic potential of K+ = 0.7 5, Na+= 1.O, Ca2+=2 .O, Fez+=2.7. Mg2+=3.0), whereasother ions (Fe3+=4.7andAI3+=5.9) arereadily precipitated as hydroxides. These differences in ionic potential account for the ready weathering and incongruent dissolution (hydrolysis)of olivine, hypersthene and calcic plagioglase, as well as the accumulation of ferric and aluminous hydroxides in well-drained soils. They also are reflected in mineral stability series (see Fig. 4.4), which are based on empirical observations of the persistence of minerals in soils and paleosols (Goldich 1938). Granitic rocks contain much feldspar and quartz, withfewmalicminerals, suchasbiotiteand hornblende. Clayey soils developed on granite may contain only quartz and microcline, with feldspar and m d c minerals all altered to clay. This is as much the case for modern deeply weathered soils on granitic rocks as for paleosols on such parent materials (Fig. 12.4; Grandstaff et al. 1986). Basalt, on the other hand, consists IargeIy of sodic and calcic plagioclases, with common m&c minerals such as olivine, and few weather-resistant minerals such as quartz. Soils on basalt weather to clay without many grainsremaining, and this also is the case for paleosols on basaltic rocks (Mossman & Farrow 1992). Sandstones show similar differences in weathering according to their mineral composition. Sandstones rich in volcanic rock fragments weather deeply to clay, whereas those rich in grains of quartz or chert weather tosandysoils(Retal1ack 1997d). Chemical composition A simplified chemical model of weathering can be constructed on the assumption that gains and losses of material in soils are largely a consequence of two common classes of soil-forming chemical reactions. First, weatherable ions are released from silicate minerals by
176
Chapter 12
Figure 12.4 Petrographic thinsectionof the surface horizon showing largely quartz and microcline in sericitic matrii (A),compared with the fresh and more diversemineral compositionof gneissic parent material (B) in an Upper Proterozoic (8 10Ma) paleosol underTorridonianalluvialdeposits near Sheigra,Scotland (AisRetallackspecimen R3 18 andB is R330;Retallack &Mindszenty 1994).Scale barsrepresent 1 mm.
acidic hydrolysis in carbonic acid formed from soil carbon dioxide. Second, iron is fixed in soils if oxidized by soil oxygen but washed out of the profile if it remains in the reduced state in which it occurs in most silicate minerals (Holland 1984: Rye 81Holland 1998). Acid consumption of soils and paleosols can be approximated by the relative amounts of major oxides of alkali and alkaline earth elements (CaO, MgO. Na,O, K,O) remaining in a soil or paleosol compared with its parent material. Similarly, oxygen consumption of soils and paleosols can be approximated by the relative amounts of reduced iron (FeO) remaining compared with the parent material. Both quantities can be calculated in moles, but the formulation I prefer is equivalents of hydrogenor electrons (seeTable4.5). Acid (or proton) consumption varies considerably in soils or paleosols, as does potential consumption of parent materials. Rocks rich in carbonates or weatherable bases require more acid for weathering than those low in these constituents. The most acidic environments are those of humid climates under biologically productive ecosystems such as conifer forest (Brook et al. 1983). where even limestone and basalt can be rendered noncalcareous and base poor, given time. In less humid and more sparsely vegetated regions acid production is lower, but soils of quartz sand remain acidic into climates as dry as subhumid. In an analogous fashion, oxygen consumption of soils or paleosols also varies, as does potential consumption of parent materials. Reduced ions of iron (Fez+)are solu-
ble and generally lost from soils, whereas oxidized iron (Pe3+)is~ed(asFe,0,)andimpartsredcolortosoilsand paleosols. Greater amounts of oxygen are required to redden a soil developed on iron-rich parent material, such as sideritic shale or basalt, than to redden one on iron-poor material, such as sandstone or granite. There is more than enough oxygen in our current atmosphere to redden a wide array of parent materials in all but the most stagnant swampy soils, and a parent material effect on oxidation is likely to have been significant only under early Precambrian reducing atmospheres (Holland 1984),if then(0hmoto 1996).
Some common parent materials Potential parent materials for soils include a wide range of rocks and sediments. Many of these are very limitedin distribution and do not figure prominently in the formation of soils. Here attention is focused on widespread kinds of parent materials for soils and the kinds of soil features that are peculiar to them.
Till Heaps of rocky rubble left by glaciers are called till, or tillite when compacted or cemented to form a rock. Till is incompletely milled rock ranging in size from rock flour, produced by grinding of rocks entrained in glacial ice, to huge boulders transported on top of the ice. Although physically fragmented, these rocks are little altered by chemical weathering. In composition, till is extremely
Parent material as a factor variable because it consists of various rocks found in mountainous collecting basins of glaciers. Pleistocene tills of eastern New York state were derived from the quartz-rich sandstones and granites of the Adirondack and Catskill Mountains, but those in the western part of the state are from the limestones and dolostones around Niagara Falls. Sandy acidic soils (Spodosols) under conifer forest have developed on the noncalcareous tills of the eastern mountains, but clayey neutral-toalkaline soils (Alfisols)under broadleaf forest in calcareoustillsinthewesternhills(Cline 1953).
Loess The sparsely vegetated terminus of glaciers is commonly dusty with fine-grained sediment blown in from floodplains and glacial lakes. This windblown silt. or loess, consists of finely ground rock fragments that are little weathered. Loess is very uniform in grain size becauseof transport by wind. From the point of view of plant and animalnutrition, loess can be an ideal parent material. It is rich in unweathered minerals already finely comminuted to expose their nutrient cations. The natural fertility of loess soils is also assisted by their ability to retain water (high field capacity). The soil-forming processes of decalcification and lessivage prevail, resulting in the formation of dark clayey soils (Mollisols and AEsoIs) in regions of moderate rainfall. The high fertility and diverse biota of loess soilsgive them great agricultural potential, as demonstrated in the North American midcontinent and the central western Russian steppe (Fehrenbacher et al. 198 6).
Alluvium The shales, sands and gravels of rivers are derived in large part from soils of their drainage basins. Unlike till and loess, this material is not merely comminuted but also has been altered by weathering. Some easily weathered minerals, such as olivine, are readily destroyed in soils and do not commonly contribute to alluvium. In many regions, however, this loss of potential fertility for alluvial soils is not severe. In arid regions and in hilly areas of temperate climates, soils of drainage basins are not so thick and well developedthat alluvium is strongly deleted in nutrient-rich minerals. As a result, the flood-
17 7
waters of the Nile River of Egypt and of the Mississippi River and its tributaries in North America leave a legacy of nutrient-rich silt for streamside agriculture after each flood (Gerrard 1987). Near streams, there are mainly young soils (Entisols and Inceptisols). Further from streams where disturbance of flooding is less frequent, other kinds of soils (Aridisols, Mollisols, Vertisols. Alfisols and Ultisols) form according to regional climatic conditions (Walker & Butler 1983; Busacca & Singer 1989). By contrast, wet, tropical regions of low relief have soils that are thick and deeply weathered, with mainly weather-resistant minerals, such as quartz, hematite and kaolinite (Sanchez & Buol 1974; Schwartz 1988). Young soils (Entisols and Inceptisols) flank tropical rivers also, but soils away from streams include clean quartz-rich sands of abandoned stream channels (Spodosols) supporting stunted broadleaf forest and red clayey kaolinitic soils (Oxisols)supporting tropical rainforest.This extreme situation is not found everywhere in tropical regions. Little-weathered alluvium from highland regions is carried by some tropical lowland streams. Volcanoes also can renew the landscape and contribute little-weathered minerals. Other areas are too dry for such extreme weathering. Nevertheless, it is useful to consider the great variety of alluvial parent material as a continuum from material almost as rich in weatherable minerals as loess and till to materials as deeply weathered and inert as some strongly developed tropical soils (Oxisols;Nahon 1991).
Marine sediments Noncalcareous shales and sandstones deposited in and around the ocean can be almost as variable in composition as alluvium. The thickest accumulations of marine sediments in deltas, continental shelves and submarine fans are deliveredby rivers. Once in the sea, this material is not weathered appreciably further. Removed as they are from surficial weathering, marine sediments tend to reflect tectonic setting during their accumulation rather than their paleoclimate. This is clearly seen in marine sandstones, whose mineralogy is readily determined by petrographic studies. Sandstones formed around volcanic mountain chains have principally volcanic rock fragments and ash. Those from fold mountain ranges, on the other hand, consist to a large extent of sedimen-
178
Chapter 1 2
tary and metamorphic rock fragments (Dickinson & Suczek 1979;Boggs 1992). Soils formed on these kinds of sandstones include abundant weatherable minerals that maintain soil fertility and promote the formation of clayey soils (Vertisols and AHsols, depending on climate and vegetation). On the other hand, sandstones formed around continental block mountains or from stablecontinental areas are rich in quartz and kaolinite. Soils formed on these materials tend to be acidic and sandy (Spodosols)or clayey (Ultisols)in regions of reasonable rainfall. Marine shales are more base rich and uniform than alluvial shales for a variety of reasons. Inert clays, such as kaolinite and chlorite, accumulate close to the seashore. Base-rich clays, such as smectites and illites, on the other hand, are carried further out into deep ocean basins (Chamley 1989). In addition, many marine shales have become more potassium rich during deep burial because of dissolution of associated potassium feldspar and muscovite (Weaver 1989). Smectitic marine shales favour the development of heavytextured clayey soils that are prone to waterlogging by perched water in lowland situations (Aqualfs, Aquolls and Histosols)or to cracking and self-plowingin seasonal climates (Vertisols). Many marine sediments are more uniform and laterally extensive than alluvial deposits, which are a complex interdigitation of paleosols with lake and channel deposits. Deltaic and submarine fan deposits approach the complex lateral variation of alluvium. In general, marine shales and sandstone are found in discrete beds that are thick and laterally extensive. Even after uplift, deformation and exposure, tracts of a single kind of marine sediment may cover large areas of land. The monotonous black shales of the Upper Cretaceous (70-83 Ma) Pierre andBearpaw Shales of North America, for example, are parent material to large areas of MolIisols and Aridisols inNorth and SouthDakota andMontana, USA (Aandahll9 82).Soil types developed on such extensive parent materials also may be widespread,and individual profiles show little evidence of internal variation in parent material. Schist Most schists are derived from marine shales because these accumulate to greater thicknesses than nonma-
cine shales. Easily weathered micas, particularly ironand magnesium-rich chlorites, are the principal mineral of schists. The main weather-resistant material of schists is quartz, which is finer grained and more easily weathered than quartz in sandstones. There is a gradation in composition from schist to metasiltstone to quartzite comparable with the variation from shale to siltstone to sandstone from which they were derived. The schist end of this range, however, has the potential to form red fertile clayey soils rich in nutrients (Alfisols) given sufficient rainfall. The cleavage planes of schist and intercalated quartzites are often at high angles to the land surface rather than parallel to it like the bedding of some marine shales. This fabric allows deep leaching, unhindered by shallow impervious layers (Marron & Popenoe 1986). Limestone Many limestones are made of skeletons of marine animals, such as foraminifera, coccolithophores. pteropods. mollusks, echinoderms and corals. Such fossils are obvious in petrographic thin sections of limestone (Scholle 1978), and can be used to distinguish parent material from soil carbonate. Dolostone also may contain fossils, but is less soluble than limestone and so less easily karstified or micritized. Limestone and dolostone can by definition contain almost 50% of other minerals. Soils formed by the dissolution of limestones consist of these residual minerals separated from bedrock by a sharp and irregular unconformity (as in Rendolls and Orthents). Insoluble residues from limestone may include weather-resistant material such as kaolinite. gibbsite and hematite, but these are buffered from extreme acidity by nearby limestone (Scholten & Andriesse 1986). Some clayey impurities in limestone are smectitic, and such marls weather to produce swelling clay soils (Vertisols),such as theHouston Black Clay of the Gulf Coast of Texa,USA (Drieseet al. 2000). Granitic rocks Here the term granitic rocks is used in a general sense to include a variety of coarse-grained igneous and metamorphic rocks, principally consisting of quartz, potassium feldspar, muscovite and hornblende. The tectonic setting of granites in stable continental areas, or cra-
Parent material as a factor tons, predisposes them toward the formation of thick and deeply weathered soils, and so does their grain size, which is usually coarse. The rapid weathering of feldspar, hornblende and biotite in these rocks leaves a loose aggregate of angular grains called grus (Isherwood & Street 1976; Dixon & Young 1981; Bouchard et al. 1995). Soils of very ancient granitic landscapes, such as are widespread in Brazil, West Africa and central Australia, may be much more clayey and red (Ultisolsand Oxisols;Ollier & Pain 1996). Basaltic rocks These volcanic rocks have fine grain size as a result of chilling of lava flows on eruption. They are rich in iron and magnesium, and in minerals such as pyroxene and olivine.They also have moderatelyhigh amounts of calcium,potassium and sodium,mainly in feldspar.Most of their silica is in minerals other than quartz. Most basaltic rocks have little free quartz (tholeites and andesites), and in many there is none (phonolites).Compared with soils developed on granite, those on basalt tend to be more clayey from decay of feldspar, more fertile from release of nutrient cations and more shallow becauseof lower porosity of both rock and soil. Because of their high iron content, freely drained soils developed on basalt are mostly reddish in color from iron hydroxides (AEsols)although basaltic soils under grasslands (Vertisols)or marshes (Histosols) may be gray or black from admixed organic matter. Many Precambrian paleosols on basaltic rocks are a distinctive lime-green color (GreenClays of Retallack 1986c)presumablybecause of theveryweaklyoxidizingatmosphereat thattime(Rye& Holland 1998). Ultramaficrocks These are all characterizedby unusually high concentrations of minerals rich in magnesium and iron (mafic minerals),such as pyroxene and olivine. Some of these rocks, such as peridotite and pyroxenite, are coarse grained and form large intrusions. Others, such as serpentinite, are fine grained and micaceous and crop out as narrow bands along major fault traces. The exchange complex of soils on ultramafic rock is dominated by Mg2+,with low levels of the plant macronutrients K+ and CaZ+.There also are high levels of trace elements,
179
suchasCrandNi, toxictoplants(Brooks1987).Soilformation may be limited on ultramafic rocks compared with adjacent, more densely vegetated rocks. In northern California and south-western Oregon, USA, for example, the open shrubby vegetation of serpentinite soils contrasts strongly with the dark conifer forests on adjacent diorites and metamorphosed volcanic rocks, and serpentinitesoils are thinner and more rocky than those nearby (Alexanderet al. 1990). Volcanic ash The chemical composition of volcanic ash ranges widelytoincludethatof bothbasalticandgraniticrocks. Some volcanic ash contains crystalsand rock fragments of equallyvaried composition,but also common is noncrystallineglass(Fisher&Schmincke1984).Scoriaand pumice are two kinds of fine-grained volcanic products riddled with bubbles formed by gases escaping the rock during eruption, in a similar way to frothing of an uncapped beer bottle. In thin sections, shards of volcanic glass are isotropic to polarizing light and bounded by concave curved surfaces (Fig. 12.5). The high surface area of these loosely packed, easily weathered noncrystalline shards bestows such characteristic properties to soils of volcanic ash that they are commonly classified in distinct categories (Andisols of US taxonomy and Andosolsof FA0 l971,1974,1975a,b,1977a,b,1978a,b, 1989, 1981). These soils have a low bulk density (<0.9 g cm-3),high water and organic matter retention, weakly developed soil structures and a high content of amorphous alteration products, such as allophane, imogolite, opaline silica and complexes of Al, Fe and humus (Tan 1984).Abundant shards and other alteration minerals such as zeolites can be used to identify paleosols as Andisols (Retallacketal. 2000). A base line for soil formation
The role of parent material is most apparent from studies of soils formed under similar conditions on different parent materials or lithosequences, which provide a basis for understanding the role of parent material as a factor. The generalizedrelationships derived from such studies,or lithofunctions,can be used to infer the effects of parent material on paleosols. However, parent material also can be considered a basis for understanding
180
Chapter 12
Figure 12.5 Unreplaced volcanic shards (black)observed in petrographicthin section under crossed nicols,of a caliche nodule (lowerBk horizon) of type Samna silty clay loam (UstollicEutrandept),in the upper Oligocene (29 Ma) SharpsFormation. Badlands National Park, SouthDakota.USA (Indianauniversityspecimen15682). Scale barrepresents0.1rnm.
soil formation in a more literal sense. Because parent materialis thestartingpointfor asoil, itsdegreeof development is measured by the amount of change compared with parent material. This way of defining parent material is logically necessary for parent material to be considered an independent factor in soil formation (Jenny 194 1).There is a difficulty, however,because the precise parent material no longer exists. Instead, its nature must be estimated from nearby materials. Such estimates typically are based on four critical assumptions that should be recognized as such and carefullyassessed in evaluating particular soils and paleosols. Assumption 1:parent material is fresh A first simplifying assumption is that the material identified as a proxy for the parent material is in fact chemically and physically representative. Saprolite, for example, cannot be regarded as an accurate reflection of the parent material of a forested soil on granite, but saprolite and soil could be regarded as parent material to a cultivated soil that formed after clear-felling and erosion of the forested soil. In the first case, saprolite and fresh granite are distinguishable by their texture in petrographic thin section. In the second case, as with soils developed on sediments and on other soil materials in general, one may have to be content with quantifying only soil development beyond initial soil formation (Feakes & Retallack 1988). Both cases imply specific assumptions about soil genesis.
Assumption 2: parent material was uniform
A second common assumption is that parent material was uniform in composition within the soil profile. Only if this were true can the properties of the material below the profile or elsewhere be accepted to represent the parent material of the entire profile. As already discussed, there are grounds for suspecting that few rocks or sediments actually are uniform enough to be an accurate reflection of the true parent material of a profile. We may consider, for example, how difficult it is to detect a thin layer of windblown dust on top of a granite withinathickclayeysoil (Fig. 12.1). Assumption 3:one constituent is stable A third simplifying assumption is that at least one constituent of the parent material has remained unaltered by weathering. This is a necessary standard or benchmark for calculating volume, weight and thickness of parent material that has weathered to create a soil or paleosol, which can be quantified as soil strain (Figs 12.6 and 12.7). This is also a helpful assumption for calculating gains and losses of minerals or chemical elements, which can be quantified as mass transfer (Figs 12.6 and 12.7). The main difficulty with this assumption is that no single constituent is immune to weathering over the wide range of conditions encountered in nature (Gardner 1980).Quartz is a fairly stable mineral, especially in soils with pH < 9. Zirconium and
Parent material as a factor
D
0.2
0.4
D UPSE 4' LOSS
0
C O U PSE &
0.6
ADDITION
i
i
MASS TRANSPORT FUNCTION Ei.u =
pp = pH. = Ci.p
=
ci.u =
7j.r =
strain of weathered product according to immobile element density of parent material density of weathered product concentration in wt % of the immobile element in the parent material concentration in wt % of the immobile element in weathered product mass transport function of any element in weathered product
Figure 12.6 Strain and mass transfer diagram modified for use with chemical data on paleosols.This diagram explains the field and equations used to construct the workedexample of Fig. 12.7.The stippled band indicates strain induced by burial compaction to 75430% of original thickness (after Bestland et al. 1996;withpermission from theGeologicalSocietyof America).
tourmaline are stable over a wide range of Eh and pH, but may be present in amounts too small to be an effective standard. Among chemical components, alumina (AI,O,) is a widely used standard constituent because it is fairly immobile between pH 4.5 and 8 and also is rather abundant, largely in clay For acidic soils, silica (SiO,) can be used, but even this is lost in large amounts at pH >9. For alkaline soils, titania (TiO,) is useful, although this is susceptible to loss in extremely alkaline or acidic conditions and if present in volcanic glass rather than a weather-resistant mineral such as rutile or ilrnenite. Trace elements usually stable in soils over a widerangeof conditionsincludePbandZr. but these are not always sufficiently abundant to provide a useful standard.
181
Assumption 4: volume change is proportional to thickness and density A fourth simplifyingassumptionis that loss of volume of soils and also their compaction during burial are related simply to density or thickness change. This may seem contrary to common sense because volume and density are measures of three dimensionswhereas thickness is a measure of only one dimension. However, observations on a variety of materials, including fossil plants of known shape (Walton 1936;Briggs & Williams 1981), suggest that under conditions of static vertical load, the horizontal or cross-sectional area of sediments, soils and fossils is maintained by pressure at the side. In the absence of folding, thrusting or other lateral compression, the pressureexertedby sediment on either side is at least equal to that exerted by overburden. Similar arguments can be made from volume and density loss during soil formation because weakening of a uniform soil by weathering decreases downward from the surface in such a way that one piece of soil has asimilar strength to another beside it. Some studies make the even bolder assumption that volume has not changed (isovolumetricmethodof Bland&Rolls1998),althoughthisisplausible for very few parent materials, such as quartz-rich granite. Problems arise if soils are not uniform because of deep cracks, corestones, hard nodules, burrows or other lateral variations in density. These difficulties are mitigated if heterogeneity is minor or if samples of the soil studied are large enough to include this variation. A fist-sized specimen may be representative of the microscopic crack system or very small nodules, but not of larger features. The assumption that thickness change is representative of volume and density change is useful because direct evidence for thickness change is easier to find in soils and paleosols. The degree of compaction, or fraction of the original thickness as a result of soil formation, can be estimated in the case of soils formed on granite with near-vertical quartz veins by the degree of buckling of the weather-resistantveins within the soil. Compaction of a fossil soil as a result of burial may be indicated by other features, such as the convolution of desiccation cracks or mukkara structure that were vertical and straight in the original soil (Retallack 1986a; Caudill etul. 1997).
182
L
S
yl
Chapter 12
IlATION
011AlION
$/T
AODlTlOh
0
0.0
CO(LAPS1
I
AODlTlOh
0.4
.
-0.6
CO(LAPS1 &
-0;0 -4.6 -0.4 -0.2 O
0.8
i l ; . TAI
i:! .0.4
D l l A SE
-0.8 -0.4
DILATION
7Si
C
DILATION
OIIATION
0.0
DILATION
0.0
-0.2 -0.4
-0.6 COLLAPS
TCa Figure 12.7 Strainand mass transport in several deeply weathered paleosols developedon Eocene (44-34 Ma) rhyoliteand colluvium of the Clarno FormationinBrown Grotto,Painted Hills. Oregon,USA.These pedotypes all lost volume and mass during weathering and include Nukut (Hapludox),TiIiwal (Kandiudox)and Apax (Dystrochrepton pedolith) (fromBestland et aI. 1996: with permissionfrom the Geological Society of America).
Quantifyingsoil and paleosol development These four simplifying assumptions allow detailed analysis of changes during soil formation and burial of soils using a variety of mathematical relationships (seeTables 4.5 and 4.6). A commonly used formulation (Brimhalletal. 1988,1991)calculatesvolumechanges in soils or paleosols (or pedogenic strain) as a separate variable from losses in mass of individual chemical constituents (or mass transport). This formulation is well suited for paleosols because the strain related
to burial by overburden can be visualized easily (Fig. 12.6). Application of this technique to strongly developed paleosols on fresh igneous parent materials has given striking trends of soil volume reduction and losses of alkali and alkaline earth elements typical for tropical humid climates (Fig. 12.7; Retallack & Mindszenty 1994). Such weathering trends are less marked in alluvial and eolian paleosols (Driese et al. 2000), which in some cases show net dilation with fluvial-eolian additions to the paleosols (Retallack 199%).
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 13 Time as a factor
As in physics, chemistry and biology, it is common also in soil scienceto study changes in soil featureswith time as the independent variable,or x-axis, in a graphicalrepresentation. These kinds of relationships are called chronofunctions. Some processes of soil formation, such as podzolization, are rapid enough to be studied by laboratory experiments (Liindstrom et al. 2000). Most soil-forming processes are too slow for such an approach, and their chronofunctions are inferred from soils of different age. Such chronosequences should be comparablein climate,organisms, topography and parent material.The best known kinds of chronosequences are flights of terraces abandoned on valley walls by downcutting streams (Harden 1982: Markewich et al. 1990)and concentricmorainesdumped successivelyby retreating glaciers (Hall & Michaud 1988; Birkeland 1992).Chronosequencestudies are of use in estimating rates of active tectonic deformation and recurrence intervals of earthquakes (Fumalet al. 1993).The security of large permanent engineering works, such as dams and nuclear power plants, depends on such studies of soils (Shlemon 1985). Soils form on time scales ranging from ecological (days,years)to geological (millions of years).The initiation of few soils has been recorded: for example, tombstones showing the date that they were erected. Studies of tombstone weathering demonstratedifferentrates of weatheringfor different conditions (Rahn 1971). However, those tombstones so weathered that they could be regarded as soils no longer have any recognizable evidence of their age. Ancient ruins can be covered in weakly developed soils whose age may be known by other artifacts. For example, 41 cm of mollic epipedon was createdby earthworms on a Roman villa at Abinger, England, and dated by coins to AD 133-367 (Darwin 1881).In history, archeologyand geology, there may be such uncertainties about the numericalage of materials that it is convenient to give periods of time distinctive names. The Julian dynasty of ancient Rome, the Pale-
olithic culture of stone tools and the Pliocene epoch of geological time are examples. The order of such named periods can often be determined from the order of superposition of artifacts or fossils in sedimentary layers: the older ones beneath the younger ones. Numerous relative time scales based on a variety of cultural and fossil materialshave been devised for global and local use (Fig. 13.1).These relative time scalesseldom are precise. Furthermore, the indicatorsof a time interval may not have changed at the same t i e in all areas. Paleolithic cultures still exist in remote parts of the world despite the technological revolutions that allow you to read this book. Fortunately,several methods for determining ages in years have been devised as a supplementand calibration of these relative time scales.Theradioactivedecayof the carbon isotope 14Cin wood, charcoal and sea shells can be used to date soils as old as 30Ka by conventional means of assay in a mass spectrometer,with allowance for production rate (Stuiver & Reimer 1993) and stratigraphic considerations (Biasi & Weldon 1994). and as old as 200 Ka by using a particle accelerator to assay the very small amounts of the radioactive isotope remaining (Rucklidge 1984). The accumulation of the cosmogenic isotope "%e also has promise for dating clayey soils as old as 200Ka (Pavich et al. 1986). For time scales of millions of years accurate methods of age determination include single-crystal laser-fusion 40Ar/39Ardating of feldspar (Renne et al. 1995) and WPbdating of zircon (Bowringet al. 1998).Thesemethods have been used mainly on volcanic tuffs and lavas, but U/Pb dating of paleosol carbonate has also yielded good results (Rasbury et al. 1998). Calibration of the geologicalt i e scalegains in accuracy each year as new dates become available (Gradstein& Ogg 1996).Radiocarbon dating of late Quaternary rocks and sedimentsis now so commonplace that the Holocene-Pleistocene boundary has been arbitrarily set at 1 0Ka. The time it takes to form various soil features varies 183
184
Chapter 13
PERIOD
EPOCH
NORTH AMERICAN PND MAMMALSTAG1
_____
OLD WORLD CULTURAL STAGE
Holocene
NEOLITHIC
1,*
MESOLITHIC
Ka
Mcqdolenion
_ _ _ _ _
Sol!treonAuriqnocian cornplax
Quoternary
E
c
5
I
Roncholobrson
H
Mourterion
z
Sangomor
ii
_~
Acheulian lrvinglanian
C -132 Ka
lllinoion (glociol)
formouthion Konson _(giociol) _ _ _ -Acorja!
.g
/ Oldowon
Pliocene
Bloncon
est oustrolopithecines
Hemphillion Tertiary
Miocene
Clorendanion Kenyapithacus
Borstovion Hemingfordion Arihreeon Oliqoeene Eocene Paleocene Late Creloceous Eorly Jurossic Triossic Permian - Gxboniferou DevonionTr, 3rdovcion __ Cambri -Sinion
oldest multicellular animals oldert multicellulor plants
.Ripheon +iuronion .Rondion -Swozian -Isuon +lodean
oldest morine microbes oge of Earlh,Moan and Solor System
oldest redimentory rocks
- 4 Ga - 4.5 Ga
Time as a factor
18 5
very strong
-
1
-
d
2
very weak
lnceplisol n16o1
16z A
yeors of soil formation
B
lb3
Ib4
165
ib6
of soil forrnatlon
Figure 13.2 Schematic representationof the t i e s needed to attain various properties of soils (A)and orders of soilsrecognizedby the SoilConservation Serviceof theUS Departmentof Agriculture (B) (modifiedfromBrkeland1999: withpermissionkom Oxford University Press).
considerably.Organic surface horizons form in only tens to hundreds of years, but it takes millions of years to form a horizon as deeply weathered as an oxic horizon (Fig. 13.2).Deeply weathered red soils (Oxisols)on Precambrian rocks now exposed in arid regions of Africa, Brazil and Australia bear evidence of middle Tertiary (>20Ma) humid weathering (Nahon 1991; Ollier & Pain 1996). Some of these bauxitic soils could be approaching the limit of possible weathering. It is theoretically plausible, but difficult to prove, that soil properties reach a steady state, envisaged as a kind of equilibrium between climatic and other forces shaping the soil vs. erosion and other forces destroying it (Phillips 1993; Huggett 1998).The notion of steady state in soils has a direct parallel in the idea of climax vegetation (see Fig. 10.1).Just as ecologistsare questioning the idea of equilibrium in undisturbed vegetation, so too is the concept of steady state of soils coming under close scrutiny. It could be that climatic variation over the past several hundred thousand years has been such that rates of soil formation have been negligible during ice ages compared with soil-forming intervals. It could also be that the odds of disturbance by climatic change, erosion or burial are such that very few soils overreach a usual degree of development. It could also be that apparent steady state of soils reflects difficultiesin estimating the age of soils. With the greater array of dating techniques now available, studies of change in soil through t i e (Harden 1 9 9 0 Birkeland 1992) indicate that while some soil properties reach a kind of steady state, many do not. Several soil features do not develop appreciably
until late in soilformation when someliiiting condition or threshold has been exceeded (Muhs 1984). Few soil features vary linearly with t i e , as would be ideal for unique estimates of the t i e required for the formation of paleosols.
Indicatorsof paleosoldevelopment Research on soils in chronosequences provides an essential data base for the interpretation of paleosols. An ideal indicator of time for the formation of a paleosol would be present in all soils, be easy to measure and show a linear relationship with time, without indications of steady state. There does not appear to be any such soil feature. It is possible, however, to formulate a qualitative scale of paleosol development (Fig. 13.3, Table 13.1). This scale works along three separate tracks: the degree of calcareous horizon development for aridland paleosols, the degree of clayey subsurface horizon development for paleosols of humid climates and the degree of peat accumulation for waterloggedpaleosols. The scale corresponds to order-of-magnitude changes that can be recognized in the field. More detailed scales (Follmer 1998) and comparison with appropriate soil chronosequences (Birkeland 1992) can yield more accurate estimates of time over which paleosols formed, as outlined in the followingparagraphs. Calcic (Bk) horizon development
In dry climates calcium carbonate may be liberated by
Figure 13.1 (Facingpage)A geologicaltime scale,plotted on a logarithmic scale to emphasize detailed understanding of mammalianevolution,humanculturalstagesandglacialchronology(datafromGrotzingeretal. 1995;Gradstein&Ogg1996; Van Couvering 1997;Bowing et d.1998;Prothero 1998).
18 6
Chapter 13 10’ years very weak
lo3 years weak
lo4 years
moderate
10’ years
strong
lo6 years very strong
Clayey 1 subsurface 2
Entisol
Bt
-+
argillic horizon 2
Peaty surface
2 4
4
Entisol histic epipedon + Histosol -, 0 = 0-4 cm 4-40 cm 40 cm 4-40
sand
clay
peat
roots
peds and cutans
,40
mcarbonate
the limited weathering of minerals or accumulate in soils from incoming windblown dust, but available moisture from rainfall is insufficient to completely remove it from the soil (Fig. 13.3;Tables 13.1 and 13.2). Within soilsdeveloped on silty and sandy parent materials, the carbonate is noticed first (in very weakly developed soils) as a fine white powder lining soil peds (clods) and as slender tubular structures, such as coatings of roots and fungal hyphae. With time, discrete hard nodules form within the calcareous horizon (weakly developed). In petrographic thin sections, the growth
Soil andpaleosoldevelopment stages forcalcic (top),argillic (middle) and histic (bottom)regimes(fromRetallack 1998 in The Adequacy of the Fossil Record, edited by S.K.Donovan&C.R.C. Paul;withpermission, 0 JohnWiley&Sons). Figure 13.
of micritic carbonate can be seen to replace matrix and skeleton grains, which are ‘nibbled’ and irregular around the edges (caries texture) and abnormally loosely packed (Wieder & Yaalon 1982). Nodules become larger and eventually coalesceto form a solid horizon toward the top of the calcareous zone (moderately developed).This solid layer can form a barrier to further percolation of water (Liu etaI. 1994),which is thus confinedtoflowalong thetopof thecementedzone toforma sharply truncated and laminated top (strongly developed). Older soils again have an even thicker crust, in
T i e as a factor
187
Table 13.1 Stagesof paleosoldevelopment (Retallack 1988a).
Stage
Feature
Very weakly developed
Little evidence of soil development apart from root traces: sedimentary, metamorphic or igneous texturesremain fromparent material
Weakly developed
With surface rootedzone (A horizon) as well as incipient subsurface clayey, oxidizedor calcareous, or surface organic horizon, but not developedto qualify as USDA argiuic. spodic or calcic horizon or histic epipedon
Moderately developed
With obvious subsurface clayey, oxidizedor calcareous horizon, or organic surface qualifying as USDA argillic,spodic, calcic. or histic; calcic horizon to stage I1 (nodules)
Stronglydeveloped
With especiallythick, red, clayey or calcareous subsurface horizons,or thick peats andcoals. or unusually strong soil structure, or calcic horizon to stages II-V (solidlayer)
Very strongly developed Unusually thick subsurface horizons, or surface coals or peats, or calcic horizons of stage VI(pisolitic, brecciated);found mainly at major geologicalunconformities
places brecciated and pisolitic, where erosion and large roots have broken the crust (very strongly developed: Alonso-Zarza et al. 1998b). Broadly similar stages of development have been documented in soils on gravelly alluvium. In the younger examples of these soils (weaklydeveloped),carbonate formsencrustationsonthebottoms of the clasts. The amount of carbonate for each morphological stage and the t i e it took to accumulate have been establishedfordesert soilsaroundLas Cruces,New Mexico,
USA (Fig. 13.4).The average rate of pedogenic carbonate accumulation in this cool desert region (mean annual precipitation of 204 mm and mean annual temRates of carperature of 15.5OC) is 0.26g~m-~kyr-l. bonate accumulation may be less than half this (about 0.09 gcm-2 kyr-I) in regions of wetter and cooler climate (near Boulder, Colorado, USA, with mean annual precipitation of 3 7 6 4 7 2 m m and mean annual temperatureof 9.7-11.O"C)andabouttwicetheLasCruces value (0.5 1g cm-2kyr-l) in warmer and marginally
Table 13.2 Stages of carbonate accumuIationinsoiIs (Machette 1985).
Stage
Soil developedingravel
Soil in sand, silt or clay
I
Thin discontinuous carbonate coatings on, and under, clasts
Dispersedpowdery and filamentous carbonate
I1
Continuous coating around and between clasts: more discontinuous carbonate beyond main horizon
Few to common carbonate nodules and veinlets with powdery and filamentous carbonate between nodules
III
Carbonate a continuous layer envelopingclasts, less pervasive carbonate beyond main horizon
Carbonate acontinuous layer of coalescingnodules. some nodules and powdery carbonate beyond main horizon
IV
Upper part of solid carbonate layer with a weakly developedplaty or lamellar structure capping less pervasively calcareous parts of the profile
V
Platy or lamellar cap to the carbonate layer strongly expressed,in places brecciated and with pisoliths of carbonate
VI
Brecciationand recementation as well as pisoliths common in association with the lamellar upper layer
188
Chapter 1 3
I
L A S CRUCES, NEW MEXICO (mean annual temperature=15 5'C, mean annual precipitation=204 mm)
2,200-4.600 vrs Oraan I
8,000-15.000 yrs Isaackd Ranch
25,000-75,000 yrs
2 5 0 , 0 0 0 - 4 0 0 , 0 0 0 vrs Jornada
h
? r i m varian
mkmm;rtci percent K E Y
cloy
sand
percent
duy:..
carbonate
other sample $ystalline carbonate paleosol of Jarnada 11 surface
Figure 13.4 Distributionof carbonate and clay in surface soils and paleosols (indicatedby arrows) of different age,formedon alluviumnearLasCruces,New Mexico,USA (compiledfromdataofGileetal.1980).
wetter climate (Roswell,New Mexico, with mean annualprecipitation of 320-3 5 5 mm and mean annual temperature of 15.3-17.2"C: Birkeland 1999). Thus, natural variation in rates is less than the several orders of magnitude in time spans. For interpretation of paleosols, there also are problems in examples older than Oligocene (>33 Ma)of differentatmospheric and biologically induced concentrations of carbon dioxide (seeFig. 9.1 5). Nevertheless,mathematical models are available to estimate carbonate accumulation in soils (McFadden et al. 1991. 1998). Some of these could be modified to make more accurate estimates of the time for formation of calcareous horizons in very ancient paleosols. Clayey subsurface (Bt) horizon development Like the accumulation of calcareous horizons, enrichment of clay in subsurface horizons also takes time. Clay is formed by weathering of parent material and of windblown dust and is washed down into subsurface horizons. Several morphologicalstages of clayey subsurface (Bt or argillic) horizons can be recognized comparable with those of calcareous horizons. These are most readily appreciated in soils formed on mixed sandy, silty and clayey alluvium (Fig. 13.3; Table 13.1). At the outset, there may be little evidence of a clayey subsurface horizon, only bedded alluvium penetrated by root traces
(very weakly developed). With time, the soil becomes cracked into recognizable peds (clods), and some of these are defined by clay skins washed down the profile (weakly developed). The clay skins are at first widely spaced, and much relict bedding remains. In the next stage (moderately developed),clay skins (argillans) and peds (clods) are more pronounced, and the total clay enrichment has exceeded that required by the definition of an argillic horizon (Soil Survey Staff 1975). Some traces of relict bedding may remain within peds at this stage, but it may only be apparent from subtle grain-size variations obscured by clay enrichment in hand specimens. With further development (strongly developed), the total amount of clay and of clay skins increases further so that the clayeysubsurface horizon is now distinct and thick, and contains no trace of relict bedding. With additional development (very strongly developed) the clayey horizon gains in thickness to well in excess of 1 m. It is common in soilsof this stage to have a thin to nonexistent surface (A) horizon, lost to erosion.These are soils of very old landscapes (pre-Pleistoceneor many millions of years old) and paleosols of major geological unconformities (Ollier 1991; Ollier&Pain 1996). These stages of development can be observed also in petrographic thin sections, which can be point-counted to quantiiy clay accumulation by volume (Murphy 1983; Murphy 81Kemp 1987) and then converted to
Timeasafactor
189
S I N JOAQUIN VALLEY, CALIFORNIA (mean annual temperature: 16'C,mean annual preciprtation= 410 mm)
ZOO y r s post-Modestolll
3,000 yrs post-Modesto II
10,000 yrs upper Modesto
40,000 yrs lower Modesto
130,000 yrs 250,000 yrs 330,000 yrs upper Riverbank middle Riverbank lower Riverbank
660,000 yrs Turlock Lake
2 KEY percent
50
PM14
Osarnple
R30 M46
0 MI2
-3
R33
0
5111
R32
L
TI1
Figure 13.5 Distributionof clay in surface soils of different age on alluviumof the Merced River,San JoaquinValley, California. USA(cornpiledfromdataof Harden 1982).
weight with information from paleosol density (see Table 4.6). As clay forms by hydrolysisfrom pre-existing skeleton grains, the grains lose their sharp crystal or erosional outlines and clear color and become embayed, cloudy and murky in appearance. The murkiness stems from fine-grained clays, which, when randomly orientated, tend to block the passageof light through the thin section. Clay that is washed down cracks, however,is deposited and squeezed into preferred orientations, which appear bright and iridescent under cross-polarizedlight. The scattered planes of bright clay (sepicplasmic fabric) are characteristic of clayey subsurface (Bt) horizons. The streaks of bright clay are sparse and disconnected (insepic) at first (in very weakly developed soils). With time (moderately developed), the clayey horizon has more and more bright clay (mosepic and masepic) until there forms a complex woven fabric of bright clay (omnisepic) in very old soils (strongly developed see Fig. 3.16). The development of clayey subsurface (Bt) horizons has been calibratedin somedetail along alluvial terraces in the San Joaquin Valley of central California (Fig. 13.5) and the coastal plain of the south-eastern USA (Fig. 13.6). The rate for formation of clayey subsurface horizons in the dry cool San Joaquin valley (mean annual precipitation 410 mm, mean annual temperature 16°C) is slow compared with that in the southeastern
USA (mean annual precipitation 1120-12201~1, mean annual temperature 15-20°C). In addition to climate, the texture of the parent material also plays a role in the rate of development of clayey subsurface horizons. Clayey horizons comparable with those formed on silty and clayey alluvium dated at lOOKa in Ohio (Lessig 1961) formed in only 40kyr on till in areas of comparable rainfall and temperature in Pennsylvania (Levine&Ciolkosc1983).Oneof themost rapidly formed clayey subsurface horizons documented is the accumulation of 1.7% additional clay in soil on porous loess thrown up only a hundred years ago byrailroad construction in a region of cool humid (mean annual temperature 9.61"C and mean annual precipitation of 8 4 7 1 ~ 1 )prairie near Cedar Rapids, Iowa (Hallberget al. 1 978).Such studies on the controls and time involved in forming clayey subsurface horizons are gaining in quantity and quality (Birkeland 1992), and should be amenable to modeling similar to that devised for carbonate horizons.
Peat accumulation The formation of peaty surface horizons (histic epipedons)ismoreaprocessof accumulation thanof differentiation comparable with calcareous and clayey subsurface horizons (Fig. 13.3).The rate of accumula-
190
Chapter 1 3 A. Solum thickness
1
x = -343446~+ 491 5 r = 0.89
A
I
5
h
I
I
400,000 800,000 years B. Argillic horizon thickness
v
c
0
.-
5
x = -196090~+ 4844 r = 0.90
400,000 C. Total profile clay
800,000 years
x = -46784~+ 7068 r = 0.98
400,000
800,000 years
Figure 13.6 Chronofunctionsfor solumthickness(A),argillic (Bt) horizon thickness (B) and totalclay (C) on alluvialterraces in the CoastalPlain and Piedmont physiographicprovincesof the south-easternUSA (datafromMarkewich et al. 1990).
tion of peat dependsupon a balance between the supply of organic debrisfromswampvegetation and itsdestruction by aerobic decay. In well-drained soils,peat does not accumulate because leaf litter is never submerged in stagnant water, and it is destroyed by aerobic decay at rates greater than its production (Fig. 13.7).The rate of accumulation of peat in this case is zero. In subsiding swamps, however, peat accumulates around the waterlogged roots of the vegetation that provides the organic debris.Therate of accumulationof peatinswampsis not
limitless, because peat-forming vegetation cannot grow undergreatdepthsof stagnantwater(Fig. 13.7).Mosses and marsh grasses can cope with rapid subsidence and generate peat, For swamp trees, which must maintain some aeration of their roots, the rate of peat accumulationisscaledclose to their growthrate. Considering these constraints on rates of peat accumulation, it is useful to consider qualitative stages of peat accumulation based on experience with modern soils. Peaty surface horizons (histic epipedon) in the classification of the US Soil Conservation Service (Soil Survey Staff 1975) must be at least 60cm thick for mossy peats and 40 cm for other kinds of peat to qualify as a Histosol. These figures reflect practical experience with typical peat thicknesses encountered in nature, in the same way as the definitions of calcic and argillic horizons. Unlike the other horizons with characteristic morphology, stages of peat accumulation may be based on original thickness. An order-of-magnitude scale for woody peats has proven useful: very weakly developed (less than 4cm), weakly developed (4-40cm), moderately developed (40cm to 4m), strongly developed ( M o m ) and very strongly developed (more than 40 m). The virtue of simplicity in such a scheme can be seen by the necessity for further corrections for compaction in applying this scale to paleosols (see Fig. 7.6). For example, bituminous coals may be compacted 0.05-0.1 times former thickness (Elliott 1985; Nadon 1998). A moderately developed paleosol in this case would have a surficialcoal 2-20 cm thick. Someestimates of peat accumulation provide a quantitative perspective on the problem (Falini1 96 5).Taking an unlikely maximum production of fresh vegetable matter in a hypothetical tropical forest of about 50 kgm2yr-l, together with observations that living plant matter is about 80%water and that pure peats contain 50 k g ~ r of - ~dry fuel, yields arate of peat accretion of about 20cmyr-l. This is an exceedinglyunlikely maximum because it is based on assumptions of unrealistically high rates of subsidence and negligible losses to decay and herbivory.This rate can be whittled down to a more realistic one if there is evidence for the amount of mineral matter in the coal, of decay in the coal from the proportion of degraded materials (macerals collinite and semifusinite),and for the rate of growth of plants in the form of fossil stumps whose age can be estimated
Time as a factor live oak forest (Ouercus virginiona I
191
bald cypress swamp (Taxodium distichurn)
maiden cane flotant (Panicurn hemitomon)
backswamp
levee
0
white sand
Figure 13.7 Soils in differentparts of the Mississippi Delta plain, near Baton Rouge, Louisiana,USA (datafromColeman 1988).
from growth rings. It is not uncommon to find at least half of acoal sample to consist of partly decayedorganic matter or mineral matter, thus suggesting that a more likely maximum rate of peat accumulation may be closer to lOcmyr-'. This may be reasonable for peats formed under moss and marsh grass, which can grow upwards at this rate. Peat formed under trees, on the other hand, probably accumulates at rates of no more than 1cmyr-l. and where there are large stumps thousands of years old, at rates of about 0.1mmyr-'. These theoretical expectations are close to observed rates of accumulation in nature (Moore & Bellamy 1973). Peat of Sphagnum moss in a bog north-west of Lake Rigsjon in southern Sweden has been radiocarbon dated at 1 8 points in a thickness of 4 m above buried birch stumps. It accumulated at rates of 0.3-1.6 m m y r l (Nilsson 1964). Peats formed under bald cypress (Taxodium distichurn) forest in Okefenokee Swamp of south-western Georgia have accumulated to a thickness of 5.9m in 7000 years (Cohen 1985), giving a
VILCU
lake
brown clay
gray shale
root traces
siderite nodules
water table
long-term accumulation rate of 0.84 mmyr-l. From suchexamples, rates of peat accumulation of 0.5-1 mm yr-l are common for woody peats, which dominate the fossil record (Retallack et al. 1996). At these rates, the maximum number of years for a very weakly developed peat would be 40-80 years, but 400-800 years for a weakly developed peat, 4000-8000 years for a moderately developed peat, 40-80 kyr for strongly developed peat and even more for very strongly developed peat. These estimates are of the same order as those for formation of calcareous and clayey subsurface horizons. They may have to be tailored for the interpretation of paleosols after consideration of compaction, botanical nature, amount of degraded plant matter and mineral content of coals.
Mineral weathering Soil development proceeds as new soil minerals are created in place of older minerals. It can thus be quantified
192
Chapter 1 3
)oints cores
I
I
weathered
2 to 3
somewhat weathered 1 to -1
-2to -3
highly weathered -4 to -5
149 to 50
49 to -49
-50 to -149
-150 to -250
C
I
D
B
A
by studying the disappearance of those minerals that once made up the parent material. The gradual loss of mineral grains can be seen under the petrographic microscope by the progressive etching of weatherable minerals (Fig. 13.8). As weatheringproceeds, the grains lose their original sharp edges from the interlocking crystal faces of igneous parent rock or the smoothing of abrasion during sedimentary transport. Many features of mineral weathering can be recognized and quantified (Table 13.3). The status of minerals with respect to these stages and the depth in the soil to which
Figure 13.8 Variationinetchingof grainsof hypersthene from soilsdevelopedon late Pleistocene (10-25 Ka) till (A, B) and middle Pleistocene(135-145Ka) till(C,D)inthe eastern SierraNevada, California,USA, illustrating relative weathering classes (imagesfromBirkeland 1999;reprinted with permission from Oxford UniversityPress: classesfromTejan-Kellaetal.1991,inwhose schemeofTable 13.3 points/scoresof illustratedgrainsare2/117forA,0/46forB. -3/-85 for C, -5/-16 1for D).
minerals of these stages extend are two ways of quantifying mineral weathering in soils of different age. In alluvium of the San Joaquin Valley of California. hypersthene is slightly weathered in l O K a soils, weathered in 130Ka soils and highly weathered in 600Ka soils. Pyroxene in these same soils is slightly weathered at lOKaand highlyweatheredat 130Ka.Acompilation of such data from soil chronosequences of wetter and of warmer climate reveals that in both cases alteration is more rapid than in the San Joaquin Valley (Birkeland 1999).
Table 13.3 Weathering of mineralgrains (Tejan-Kellaetal. 1991). Measure
Feature
Points
Area
Degree
Freshness
High microtopographic relief Clean cleavage faces Arc-shaped/parallel steps Sharp edges,angular grains Conchoidalfracture, V-pits
+1 +1 +1 +1 +1
10 10 10 10 10
f5 f5 f5 f5 f5
Solution Scaling, surface roughness Orientedor randometchpits Subdued or rounded edges Hairline cracks
-1 -1 -1 -1 -1
10 10
f5 f5 fS *5 f5
Weathering
~~
10 10 10
Score
+so +so +so +so +50
-50 -50 -50 -50 -50
~
Note: weathering classes (Fig. 13.8)canbequantiiedaseither surnof pointsorsumof scoresfromarea timesdegree(out of lOand 5, respectively).For chronosequences, average points or scores from at least 30 grains should be used.
Timeasafactor Another way to quantify such observations is to compare the abundance of easily weathered minerals with those that are weather resistant. Commonly used mineral-weathering ratios include quartz/feldspar, which can be estimated by point-counting petrographic thin sections, preferably after staining for feldspar (Houghton 1980).Anotherratio, of zirconandtourmaline to amphibole and pyroxene, can be calculated from heavy mineral separations (Bland & Rolls 1998). Care must be taken to examine the separates or thin sections for evidence of grain etching, rather than burial alteration or depositional sorting of the minerais. Estimation of both these mineral-weathering ratios for soilson tills of various age in central Iowa, USA, show progressively deeper and more thorough weathering with time (Ruhe 1969). Additional similar studies of other chronosequences in different climatic, topographic and other environmental conditions are needed as a data base for the interpretation of mineral weathering in paleosols. Weathering rinds Pieces of rock forming the pebbies and bouIders of gravel and till represent locally indurated and less permeable parts of some soils, and they weather slowly from the outsideinward.The marginalzones of clay and
Figure 13.9 Increasedthickness of weathering rinds of andesite and basalt clasts withageof their containingtills andin regions of different mean annual precipitation(fromColman 1986;reprinted with permission from Academic Press).
193
red sesquioxideson pebbles are called weathering rinds. Thickness of rinds can be measured from a large sample of boulders broken open in the field. It is best to sample all the boulders a fixed distance below the surface so as to avoid the sur6cial effects of fire-spalling, sand-blasting or wear from animal tracks. Differential rates of development of weathering rinds with rock type are strikingly seen in lithologically heterogeneous tills, such as those around Lake Tahoe in the eastern Sierra Nevada of California,USA (seeFig. 12.3). Weathering rinds are slower to form on finegrained basalt than on coarse-grained andesite. Coarsegrained granitic rocks do not develop such a distinctive rind but are deeply weathered so that grains are loosened to a friable grus. Over the past 200 kyr in this cool desert region (mean annual precipitation 785 mm, mean annualtemperature 5.8"CatTahoeCity), granitic and andesitic boulders have been deeply weathered, but cores of fresh rock remain in basaltic pebbles. These results are at the slow end of the spectrum of rind development (Truckee in Fig. 13.9),which can proceed at rates three times as fast in wetter climates. Combinations of features In attempts to assess overall soil development, a number of indices have been devised to reduce combinations of
50
100 150 TIME i?O' yr)
200
250
194
Chapter 13
Table 13.4 Buntley& Westi(1965) andHurst (1977)colorscores. Munsellhue Hurst H' Buntley and Westii score
10R 10 7.0
2.5YR 7.5 6.0
5YR 15 5.0
7.5YR 17.5 4.0
lOYR 20 3.0
1.25Y -
2.5Y -
2.5
2.0
5Y -
1.o
T= P.L . c?' where Tis totalironcontent (Fe,O,,wt%) P is Hurst hue L is Munsell value Cis Munsell chroma
soil features to a single value. By using enough timedependent features, it is hoped to swamp competing effects of other soil-forming factors in creating the features on which the index is based. The soil development index of Harden (1982) provides an example of the technique applied to a wellunderstood chronosequence of soils in alluvial parent materials of the Merced River in the San Joaquin Valley of California, USA. The index is derived from field estimates of eight soil properties: redness (or rubification). total texture (or clayeyness), clay flms, structure, dry consistence, moist consistence, darkness (or melanization) and chemical reaction (pH). This last feature is measured in both soil and parent material using a standard field meter, but others are quantified in various ways. Total texture, for example, is quantiied by assigning 10 points for each line of the textural triangle (see Fig. 3.10) that separates the horizon fromitsparentmaterial, 1Opointsforincreasesinstickiness and 10 for increases in plasticity. These values are then added together and normalized by dividing them by the maximum attainable value. Similar normalized values for all eight properties are averaged to calculate the development index for each horizon. These horizon indices are then averaged to calculate overall soil development index. For some features, such as profile structure and melanization, steady state was achieved within lOkyr on the Merced River chronosequence. For other features, steady state does not seem to have been attained after half a million years. The Buntley and Westin index (1965)is yet another composite that reduces several color properties to a single value. This index has proven useful for soils at high
latitudes (Bockheim 1990), and can be correlated with total iron content (Table13.4). Despite what appear to be a number of arbitrary assumptions in the formulation of soil development indices, they include ingenious ways to quantify field observations.They are unlikely to be applied to paleosols in their present form, however, because many of the properties on which they are based are not preserved in paleosols. Diagenetic dehydration of ferric hydroxide minerals (see Fig. 7.10), destruction of soil organic matter (see Fig. 7.3), and cementation irreversibly alter the degree of rubification, melanization, stickiness and plasticity of a buried soil compared with its precursor soil.
Accumulation of paleosol sequences Soils are clues to the way in which landscapes and sediments form though time. Paleosolsprovidesimilarinformation about the formation of ancient landscapes and accumulation of nonmarine sedimentary sequences, as illustrated by the following examples based on Eocene-Oligocene (38-29 Ma) paleosols of Badlands National Park, South Dakota, USA (see Figs 6.2 and 6.3).
Hillslope development There are several different views on the development of hillslopes over time (Fig. 13.10). The idea of slope decline was popularized by William Morris Davis at the turn of the century. By this view, hillslopes lose their steepnessbecause the tops of hills are eroded more energetically than their lower slopes,which are protected by
Time as a factor
.
I.
SLOPE DECLINE
SLOPE REPLACEMENT
PARALLEL RETREAT
so11development colluviurn
- .- - former slopes
( W. M. Davis)
(W. Penck)
(L.C. King)
soil depth badrock
Figure 13.10 Contrasting views on the development of hill slopes(based0ndataofYoung 1972).
a colluvial mantle. Soils on a declining slope should be of similar thickness and degree of development everywhere, but a little thicker and better developedin the cumulative colluvial mantle just above the base of the slope. An alternative view of parallel retreat promulgated by Lester C. King emphasizes the constant nature of geomorphological processes on different segments of slopes. By this view, soils within the slope are of approximately the same degree of development,but exceptionally well-deveIoped and thick soils persist on the divides. A final view of slope replacement, advocated by Walter Penck, is somewhat conciliatory. By this scheme, the steepest portions of hillsides and thinnest soilsare in mid-slope.Slope replacement and parallel retreat can be responses to tectonically rapid uplift or stream downcutting, whereas a more even soil mantle is created when rates of uplift and climate change are more evenly matched to rates of soil formation and plant colonization (Kooi& Beaumont 1996). Studies of soils on modern slopes show that each of
195
these models of slope development is feasible under different conditions. One well-known study on postglacial moraines under sod grassland around Jewel1bog in tectonically stable north-central Iowa, USA (see Fig. 11.1) supports a model of parallel retreat with modest slope decline (Ruhe 1969). Post-glacial moraines clad in desert grassland on the arid fault scarp flanks of the eastern Sierra Nevada of California contain soils better developed in downslope positions than in intertluves, compatible with slope decline (Hallet & Putkonen 1994). In the actively uplifted and forested Oregon and California Coast Ranges, however, soil formation varies inversely with hillslope curvature, as in the upper parts of slope replacement (Heimsath et al. 1997; Roering etal. 1999). Ancient hilly landscapes also can be reconstructed from their paleosols. A part of a hilly landscape is buried at the base of the sequence of middle Tertiary alluvial deposits in Badlands National Park, South Dakota, USA (Fig. 13.11).A thick well-developed paleosol (Yellow Mounds pedotype) with an oxidized subsurface (C) horizon reaching as deeply as 2 7 m is present in many places at the disconformity between Late Cretaceous marine rocks and the overlying alluvium of late Eocene to Oligocene age (38-29Ma). By Late Eocene time (37Ma), when this low rolling area was covered in humid tropical forest, it was a n ancient alluvial landscape, as indicated by deposition of the Chamberlain Pass Formation and formation of strongly developed Wetaandhteriorpaleosols (Terry 1998).Thissituation changed during latest Eocene time (36Ma), when a major stream draining the emergent Black Hills incised deeply (50m in section of Fig. 13.11)into Cretaceous marine rocks, well below the zone of earlier Eocene alluvium and weathering. The red clayey Interior paleosol and its thick yellow saprolite were truncated by steep dopes around this deeply incised valley. Some destabilization of the older soil surface on the high drainage divide may be indicated by bedding in deeply weathered materials above the undisturbed clayey subsurface (Bt) horizon of the Interior paleosol (in measured section of Fig. 6.2). The persistence of a very strongly developed paleosol on upland divides and terraces through a cycle of deep valley cutting, contrasted with weak development of paleosols within the paleovalley, supports the idea that this partial ancient landscape formed by parallel retreat. The nature of the paleosols on the inter-
196
Chapter 1 3 0
0
20 4 0 k m
U Oligocene alluvium
mmmD bte Eocene paleosols
a
Precornbrian rocks
-
r+, synsedimentary fouit and downthrown side
NORTHEAST
SOUTHWEST Cheyenne
Corrol
Indian
Scenic
20 miles
0
0
5
10
Sage Creek
Pinnacles ore0
elevotim in 100s f t , m X I 0 0 verticol exa(lgerotion
20 kilometers
fluves indicates that abrupt climatic drying to a more open forest than before, rather than tectonic uplift, was the cause of this paleovalley incision (Retallack 1986b, 199 2d). Alluvial architecture Just as buildings are composed of discretestructural elements, such as brick walls, stone columns and glass windows, river deposits can be regarded as being built from rock types of characteristic appearance.The paleochannel facies is the most striking and best understood architectural element, with its cross-bedded gravel and sand. Other sedimentary facies of point bars, floodways, levees and oxbow lakes can be contrasted with pedogenic facies of floodplain soils (Clemente & Perez-Ar-
Figure 13.1 1 Outcropsof deeply weathered Upper Eocene paleosols (Interiorand Yellow Mounds paleosol zone including ChamberlainPass Formation),of Eocentdlligocene ( 3 8-29 Ma) alluviumof the White River Group (Chadronand Brule Formation)and distributionof synsedimentary faults (above),and reconstructed cross-sectionof a deep paleovalley (below)in south-westernSouth Dakota and north-westernNebraska, USA (fromRetallack 1988b:reprintedwith permission from the Society of Economic Paleontologists and Mineralogists).
lucea 1993;Willis&Behrensmeyer 1994; Kraus 1997; Kraus & Aslan 1999). Many different architectures of paleochannels and paleosols can be imagined (Fig. 13.12). In meandering streams under a regime of high accommodation space, channels are narrow because they are restricted by lush growth of vegetation and clayey banks. A very different architecture is produced by meandering streams in a regime of slow subsidence or sea-level fall leaving little sediment accommodation space. Under such conditions clay and silt are lost downstream, and the sequence is largely sand and conglomerate. A third architecture is created by steep stream gradients and channels that are broad, shallow and braided by islands of sediment. Such braided streams create sequences dominated by sand and gravel under a wide range of sediment
Timeasafactor
197
HIGH ACCOMMODATION LOW ACCOMMODATION
braided streamsunder different regimes of sediment accommodation.
mgravel u s a n d
accommodationregimes. The most distinctive architecture of the four possibilities is the asymmetric paleochannels of meandering streams separated by thick floodplain deposits. The other three situations have proven difficult to distinguish using evidencefrom paleochannelsalone (Bridge 19 85).However,evidencefrom paleosols can help insofar as their degree of development and spacing in sedimentary sequences reflect overall accumulation rates and longevity of floodplain surfaces. Each of the rock units of Badlands National Park, South Dakota, USA, shows a different kind of alluvial
Figure 13.13 Sheetsandstonesof Oligocene (32 Ma)paleochannelsinthe ScenicMember of the Brule Formation,Pinnaclesarea, BadlandsNationalPark,SouthDakota,USA.
=clay
soils and paleosols
architecture,and this is in part the reason they were distinguished as formations. Let us consider, for example, the sandy Scenic Member of the Brule Formation of Oligocene age (32Ma: Fig. 13.13).Are thesedepositsof braided streams or of meandering streams that coalesced because of low rate of sedimentaccommodation? The degree of development of associated fossil soils provides a clue. The lower part of the Scenic Member contains numerous weakly and very weakly developed paleosols whereas the upper part contains numerous strongly developed paleosols, like the upper part of the underlying Chadron Formation.These general impres-
198
Chapter 13
sionscan be considered in more detail by assigning times of formation to each paleosol by comparison with modern soils whose age is known. Minimal estimates can be used in deference to concepts of soil-forming intervals and steady state. Alternatively, maximal estimates can be used, to minimize potential gaps in the record (Retallack 1998a). Once ages for each paleosol are assumed they can be added together to obtain the duration of a particular sequence. The thickness of the sequence divided by this estimate of time gives rate of sediment accumulation. Estimated rates based on both minimal and maximal soil duration turn out to be an order of magnitude greater than those based on radiometric and paleomagnetic estimates of the duration of these rock units. A part of this discrepancy may be blamed on the conservative nature of the estimates used, but a greater part results from the incompleteness of these sequences. It is notable that rates of sediment accumulation estimated from paleosols show the same relative change as those calculated by radiometric and paleomagnetic means. The value of paleosolestiiates is that they can be applied to sequences not amenable to direct paleomagnetic or radiometric dating, and compared with other paleoenvironmental indicators (Fig. 13.14). The lower portion of the Scenic Member, for example, accumulated at much greater rates than the middle or upper portion. Thus, these sandy paleochannels formed in loosely sinuous and partly braided streams, unlike the asymmetric paleochannels of meandering streams found in the upper part of the Scenic Member and in the underlying Chadron Formation.
Completeness of the rock record Sedimentary successions are undoubtedly incomplete records of past events and times, but just how incomplete are they? How long did it take each sedimentary layer to accumulate?And how long was it between deposition of the layers? These are important questions for estimating the resolution of sedimentary successions for rates of evolutionary or geomorphological change. They are difficultquestions to answer for most sedimentary rocks, but not intractable for successions of paleosols. Flooding events that cover soils and initiate soil formation are known to be rapid (days or weeks) compared with the time it takes to form a moderately developed soil (many thousands of years). Thus, t i e s of deposition can be regarded as insignificant, and times between events of deposition can be estimated from the degree of development of paleosols. The record of time provided by a paleosol could be obscured by being incorporated in a better-developed paleosol. Also, the age of a paleosol could consistently be underestimated if the features used to estimate its age were reaching a kind of steady state or represented brief soil-forming intervals. Both these difficulties are inherent is using chronosequences to interpret times over which paleosols formed (Retallack 1998a). They are minor sources of error compared with those introduced by large-scale erosional cutting followed by periods of filling. The erosional disconformity between the base of one of these depositional units could represent enormous amounts of time but appear little different from the disconformityat the top of each paleosol. These cut-
28 n
1
I
local rock m sediment accumulotion Epoch magneticanomalies
0---
5010 I.(
12 13 140
L
0
10 20 30
'C mean annual lrmprature from reptiles 0 and leaves
no mommolion genera living 0 ond erlinqu6hed &
Figure 13.14 Factorscontrolling fluvialdepositionin ameasuredsectionof the Pinnaclesarea (seeFig. 6.2).Badlands National Park, SouthDakota,USA(fromRetallack,Gregory,0 1992 inEocene-Oligocene CZimaticandBioticEvolution,editedbyD.R.Prothero & W.A.Berggren;reprintedwith permission fromPrincetonUniversity Press).
Time as a factor tiig and filling cycles may be revealed by abrupt changes insuccessionsof paleosols (Retallack 1986b, 1998a)or by lateral mapping (Kraus& Middleton 1987;Bestland 1997). Such disconformities particularly confound efforts to distribute geological time proportionally using relativepaleosoldevelopment (Bown &Kraus1993a.b). Withdue regard for these important sourcesof error, paleosols may be useful records of the temporal resolution and completenessof nonmarine sequences. Resolution is a question of scale. Fine resolution for successionsof paleosolsishundredsof years, andcoarse resolution is millions of years. Completeness, on the other hand, is a measure of the reliability of a sequence at a given resolution. For a complete sequence at a resolution of 1000 years, there should be on average a bed, paleosol or other record forevery intervalof 1000years. In general, a succession of paleosols can be regarded as complete at a resolution that is equivalentto the average time of formation of each paleosol. Fractional or per cent completenesscan be estimated for shorter resolutions, but a general impression of completenessalso can be gained by casual inspection. Successions complete at fine resolution will have numerous weakly developed paleosols with abundant relict bedding and little pedogenic clay, structure or color differentiation. Successions complete at coarse resolution would have few, strongly developed paleosols, which were clayey, well structured and strongly differentiated into red or dark colors. A more complex way of estimatingcompletenessof a succession of paleosols or of sediments alone is by comparing rates of sediment accumulation of a sequence with rates usual for that environment and time span, as estimated from a large compilation of rates (Sadler 1981; Sadler & Strauss 1990). Rates used in such compilations include those from direct observation of sediment accumulating from floodwaters, from measurement of survey points, from radiocarbon dating and from other longer-term forms of geophysical dating. The rates decline for estimates made over longer time spans because of correspondingly longer gaps in therecord. Medialratesdefineanexpectationof therate for a complete sectionat a particular resolution. These various estimates of completenessand resolution may be compared by considering their application to the succession of paleosols in Badlands National Park, South Dakota, USA, and coeval paleosols in the
199
PaintedHillsareaof centralOregon,USA(Fig. 13.15).A simpleway to estimateresolutionfrompaleosolsis to average the estimatednumber of years for each paleosol in the sequence.Fromthesevalues, the probability that the succession is completefor intervalsof 1000years can be estimatedby the fraction of 1000yearsover the time for which it would be complete. For the Chadron Formation above the Interior paleosol in Badlands National Park, for example,there is a 1:5 chance that a given 1000year interval will be represented by sediment capped by a paleosol, but the chances are only l :20 using rates of accumulation determined from radiometric dating of these units (Retallack 1998a). Despite their disparity, eachmethod preserved the same ranking of formations: Poleslide Member of the Brule Formation more complete than Scenic Member, and these more complete than the Chadron Formation above the Interior paleosol. The general appearance of these rock units also bears out this ranking. Paleosols of the ChadronFormation developed on the same ashy parent material as the other units, but they are clayey, with green surface and pink subsurfacehorizons. The Poleslide Member, on the other hand, is silty and more uniformly yellow and brown in color, reflecting its less-developed paleosols. Similar calculations have been made of the coeval succession of paleosols in the Painted Hills of central Oregon (Fig. 13.15). Here also the relative completeness of stratigraphic units increases up section from the uppermost Big Basin Member (correlativewith Chadron Formation), to middle Big Basin Member (correlative with Scenic Member) and upper Big Basin Member (correlative with Poleslide Member),and in each successive unit the paleosols become less and less developed (Bestland et al. 1997). Also possible is comparison of the two successions:more complete for the uppermost Eocene (localChadronian)succession in South Dakota, but more completefor the lower Oligocene (Orellanand Whitneyan)successionin Oregon. Sequencestratigraphyof paleosols Before the advent of accurate informationon the timing of events of deposition, sedimentation was widely considered cyclical. Familiar examplesrange frompoint-bar cycles (Allen 1965)to glacio-eustaticcyclothems(Cecil 1990; Tandon & Gibling 1994) and the tectonosedimentary ‘pulseof the earth’ (Grabau 1940).In geomor-
200
Chapter 13 calcareousness development CT m
ta
aconglomerate, breccia KEY
=]sandstone
3 claystone, IRIKAREEAN
shale
tuff
a
w andesite
corestones calcareous nodules planar bedding
\
red root traces bones and teeth snails calcareous phytoliths coprolites
J=j NHIT-
logs
[v1leaves
.IEYAN\
BADLANDS, SOUTH DAKOTA +-
PAINTED HILLS, OREGON
Figure 13.1 5 A comparison of the incompletepaleosol successions in Badlands National Park, South Dakota, USA, with the more completepaleosolsuccession of the same geological agein the PaintedHillsUnit, John Day FossilBedsNational Monument, Oregon,USA (GomRetallack 1998 in The Adequacyof theFossilRecord,editedbyS.K. Donovan&C.R.C.Paul: withpermission, 0 John Wdey &Sons).
Timeasafactor phology, thereis theerosionalcycleof Davis (1899)with its youthful, mature and senile stages of valley and river form. More recent quantitative studies of changes in landscapes and rivers (Schumm 1977) indicate that erosion and sedimentation are not as regularly cyclical as is sometimes assumed, andmay be more productively considered episodic. In recent years episodicity of sedimentation has become incorporated within a new framework entitled sequence stratigraphy (Emery & Myers 1996;Miall 199 7).The fundamental tenet of sequence stratigraphy is that sedimentary facies form repetitive sequences genetically related to environmental shifts such as sea-level change. Paleosols can be important clues to sequence stratigraphy and its causes at a variety of scales (Wright & Marriott 1993; Miller &West l993,1998;McCarthy&Plint1998). The finest level of episodicity preserved is the way in which one paleosol is found on top of another in long sedimentary successions. This appears to be a record of periodic destruction and sedimentation over former ecosystemsand their soils.The degree of developmentof paleosols in such sequences provides crude estimates of the recurrence interval of such events. For flood discharges, and perhaps other geomorphologicalevents as well, events of long recurrence intervals also are events of greater magnitude. Both recurrence and magnitude are important elements for understanding how sedimentary sequences are put together. Regular recurrence intervals are evidence of a landscape in a kind of dynamic equilibrium between forces that preserve it, such as slow subsidence and dense vegetation, and forces that destroy it, such as high subsidence rates and sparse vegetative cover. Erratic recurrence intervals are evidence of a system out of balance and seeking a new equilibrium. Both recurrence intervals and magnitude revealed by paleosols are clues to the nature of the event and whether it was a normal part of a sedimentary system like the avulsion of a stream channel or something else entirely,such as a major climatic change. Not much can be done to analyze the causes of superposition of individual paleosols in Badlands National Park because the completeness and resolution of this sequence is inadequate. Such studies are better undertaken in successions of paleosols that can be dated with greater precision, such as Quaternary paleosols of Kansas (Fengetal. 1994).Washington(Busacca 1989), CzechRepublic (Kukla&Cilek1996).andChina(Wang
201
et al. 1998). In these cases, glacially related climatic change on time scales of tens to hundreds of thousands of years has led to widespread alternation through time of soils and ecosystems. It is conceivable that similar processes acted in creating the sequence of paleosols in Badlands National Park, but the cause of each event of paleosol superposition remains beyond resolution. A broader pattern of episodicity in Badlands National Park can be related to cutting and filling cycles (Fig. 13.16).Because each cutting cycle ushered in different soil-forming conditions, each sequence of paleosols has a different appearance. Many of these features already have been used to establish rock formations in the area. When stage of development of each paleosol is plotted against stratigraphic level (Fig. 13,15),the times of erosion stand out as dividing sequences of closely superimposed, better-developed paleosols from those of distantly spaced, poorly developed paleosols. The lower parts of each filling cycle were times of rapid sediment accumulation and an erratic pattern of a recurrence of paleosol types. These appear to be t i e s of disequilibrium when the system had yet to reach a balance between controlling forces. The upper parts of each sequence, however, have numerous paleosols showing comparable degrees of development. These are too regular to be due to chance and appear to reflect a kind of dynamic equilibrium. During these relatively stable times, thevegetation and other environmental conditions are filtering possible destructive events for those of a particular magnitude and imposing a crude kind of stability on the fluvial system. Such progressive changes in paleosols provide examples of geomorphological thresholds: of a system driven to a breaking point every few million years. The resolution of the sequence is fortunately adequate to address the causes of these environmental crises. The factors controlling fluvial systems include the same ones important to soil development, such as t i e , initial topographic relief, base level, local geology, climate and vegetation, as well as upland drainage net, hillside morphology, downstream deliveries, channel behavior and pattern of deposition (Schumm 1981). Many of these features can be quantified in ways already outlined in other parts of this book. Only a summary of the most important factors is outlined here (Fig. 13.14). By plotting the sedimentary sequence against real geologicaltime, as estimated paleomagnetically,the cutting phases can be seen to represent as much time as filling
202
Chapter 13 CUT
FILLING
FULL
CUTTING ROCK UNITS Sharps Formation
Poleslide Member
CITJ Scenic Member Chadron Formation Chamberlain Pass Formation 13
VEGETATION open rangeland wooded rangeland woodland
@ Z Jforest
Figure 13.1 6 A reconstructionof changes in vegetation and landscape during several Late Eocene and Oligocenecutting and fllingepisodesthat created therockunitsnowexposedinBadlandsNationalPark. SouthDakota,USA (fromRetallack 1986b; reprintedwith permission fromthe Society of Economic Paleontologists and Mineralogists).
phasesof whichthereisarecord.Thecuttingphasescan be recognized from breaks in sediment accumulation and in the downward limit of channel incision. Some also were times of extinction and range truncation of local fossil mammals (Prothero & Whittlesey 1998). In one case, a cutting cycle coincides with an increase in far-traveled alluvium (quartz in C horizons of these dominantly volcanic-ashsoils),perhaps implicatingtectonic uplift in the quartz-rich source terrain as its cause. In another case, an especially thick volcanic ash is present, perhaps an indication of volcanic destabilization. Each cutting phase, however, corresponds to a period of drier climate and sparser vegetation (Fig. 13.14). Although contributing factors were instrumental in some erosional phases, and each was different in some way, climate can be identied as the forcing mechanism in each case. There is independent stratigraphic, sedimen-
tological, isotopic and paleobotanical evidence from Oregon for climatic fluctuations at similar times (Bestland et al. 1997). Paleosols provide insight into the robustness of landscapes in the face of external environmental change and of the global integration of environmental systems represented by these other sources of information about climate. Interpretation of time for formation of paleosols opens the way to a variety of studies concerning historical aspects of landscapes and life. This and other comparable examples can be generalized to a sequence stratigraphic model for paleosols (Fig. 13.1 7). Incomplete successions comparable with the Badlands of South Dakota are common and show a sawtooth pattern of environmental change. Paleosols within each sequence tend to show increasing developmental age, climatic drying and geomorph-
Time as a factor
203
calcareousness development hue
calcareousness hue development
UPLAND c-----C DEPOCENTER
I
I SAWTOOTH PATERN INCOMPLETE SEQUENCES
**
Ir aging upwards
drying upwards draining upwards
High Stand Systems Tract
f
/
Regressive Systems Tract
Systems Tract
/ Low Stand Systems Tract
1
IDEALIZED SINGLE SEQUENCE SINUOUS PAlTERN LITHOLOGICAL COMPLETE SEQUENCES channel strongly developed sandstone paleosol Ir aging younging =coal or lake B w e a k l y developed dry +, wet climate shale paleosol draining +, ponding claystone root traces
0
F.f;l
* *
-
Figure 13.17 A model for sequencestratigraphyof paleosols (fromRetallack 1998 in The Adequacy of the Fossil Record, edited by S.K.Donovan&C.R.C.Paul: withpermission.0JohnWiley&Sons).
ologicaldrainage upwards within the sequence,until interrupted by a disconformity at the base of the next sequence with weakly developed paleosols of wetter climates and poor drainage. Only with exceptionally high rates of sediment accumulation are sinuous patterns of these environmental trends preserved, for example in some parts of the John Day Formation of Oregon, USA (Fig. 13.15) and in the Siwalik Group of Pakistan (seeFig. 11.7).A common explanation of such sequence stratigraphic changes in continental margin deposits is global sea-level change (Emery & Myers
1996; Miall 1997). Although Oligocene sequence boundaries in both South Dakota and Oregon have been shown to coincide in time with sea-level changes (Bestland et al. 1997), these successions are a long way from the sea. These paleosol successions were probably coordinated with sea-level change by means of global paleoclimatic changes (Retallack 1998a). Although each local succession has its own paleoenvironmental history, the ultimate causes of such paleocliatic episodicity are potentially amenable to study using globally distributed paleosol data.
Part 3
Fossil record of soils
Eocene red paleosols above paleochannelsandstonesand conglomeratesof the Clarno Nut Beds,Hancock Field Station,CentralOregon,USA.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 14 A long-term natural experiment in pedogenesis
In assessing the effects of a particular factor in soil formation, field situations are sought in which most of the variables are comparable except the factor chosen for study. In many cases this is difficult. The effects of climate, for example, may not be easy to tease apart from those of vegetation. Studies of the role of particular factors in formation of paleosols should select examples where other factors of paleoenvironment and diagenetic alteration were comparable. This may seem a daunting enterprise, but it is in principle no different from interpretative paleoecology or sedimentology. There are many questions of general scientific interest concerning conditions of soil formation that can be addressed only using paleosols. In the geologicallydistant past, for example, climatic conditions were not always linked with the same kind of vegetation as that of today. The fossil record of soils can be considered a long-term natural experiment in which many of the factors of soil formation operated in fundamentally different combinations. The fossil record of soils is long. The oldest paleosols yet studied indetail are 3 500 myr old (Buick et al. 1995: Rye &Holland 1998). We still can only speculate about soil formation on the early Earth back to its origin at some 4566Ma, as indicated by radiometric dating of the Moon and meteorites (Birck & Allegre 1988). Such speculation becomes better informed with each passing year as very ancient paleosols are better understood, as computer modeling of soil formation becomes more sophisticated and as the soils of other planets and moons come under scientific scrutiny. The processes of soil formation on the Moon and Venus are very different from those on Earth. Surface processes on these other heavenly bodies provide new models for soil formation against which liely soil formation on the early Earth can be reassessed. Major events in the history of soils on Earth are already apparent from the rock and fossil records: the appearance of lie on land, the oxygenation of a
primeval reducing atmosphere, the advent of large plants and animals on land, the evolution of forests, the spread of grasslands and human modification of the land. Their surficial effects can be anticipated only to a limited extent from what is known of analogous Quaternary soils because they remain in essence historical events. Evidence from paleosols is especiallyuseful in establishing when they occurred. The evolution of grasslands in dry continental interiors is poorly known from the fossil record because swampy and wet conditions are neededfor the preservationof plants (Retallack 1998a). Fossil teeth adapted for grazing tough grass presumably postdate the appearance of grasslands, andsuchecosystems leave little trace in sediments. Grassland paleosols, however, can be distinctive(Retallack 1997b). Paleosols also can be expected to provide insights into how surface environments and ecosystems evolved. Organisms and their environments are bound in a complex web of interactions that has evolved through time. The interactions and selectivepressures at the earliest evolution of what are now major ecosystem types may have been very different from the factors that have sustained those ecosystems. A comparable phenomenon of biological evolution is pre-adaptation in the evolution of major new biological structures. Feathers, for example, may have evolved primarily for insulation or display, and only later been pressed into service for flying (Chatterjee 1997: Ostrom 1997). Similarly, the origin of a major ecosystemsuch as open grassland may have had little to do with their historical ability to cope with frequent fires set by humans (Retallack 1997b). Questions of the origination of different kinds of soils and ecosystems can be addressed by the study of paleosols. As a list of major events in Earth history, those cited above are rather few. This is not to deny the importance for soilsof other events such as the Cretaceous-Tertiary extinction of dinosaurs (Retallack 1994a) or the greatest of all mass extinctions at the Permian-Triassic boundary (Retallack 1999b). There were profound 207
1 208
Chapter 14
r
b
L
,
I r
'
I
I
I
I
1
green clay bauxite karst calcrete
silcrete ' z GELISOL VERTISOL OXISOL AND~SOL HISTOSOL
Figure 14.1 Geologicalrange of soil features
4 and soil orders of the Soil Conservation Serviceof theUS Departmentof Agriculture. $~~~~ Entisols andInceptisolsare presumed to have SPODOSOL
I
_-
A.-. .__ I
PRECAMBRIAN 3 2 1 thousands of millions of years before present
4 MOLLISOL
W W I
changes associated with these events that did much to fashion subsequent ecosystems, but they were not reflected in fundamental or persistent changes in the nature of soilson Earth. Like organisms and communities, soils have evolved over the geological history of the Earth. For the most part, this evolution has consisted of the addition of new kinds of soils as new ecosystems and environments appearedonEarth (Fig. 14,1).Therehasbeenanoverall diversification of soils with time because ancient kinds of soils have persisted in other regions of a n increasingly heterogeneous world. Extinct kinds of soils are theoretically possible,but the only extinct soils known are Green Clay paleosols of early Precambrian time. These are green, clayey, alumina-rich and base-poor paleosols developed on iron-rich parent materials such as basalt. Their composition may reflect lower oxygenation of the atmosphere because basalts now weather to red, ironrichsoils (Rye&Holland1998).Therarityof extinct soil types could reflect the vastly greater amount of information available on modern soils compared with those of the past. Considering difficulties in unraveling the alteration of paleosols after burial, it has been easier to
beenprecursorsof other Precambrian paleosols,eventhough difficult to recognize among Precambrianpaleosols because they lack root traces and burrows.
identify a paleosol with a modern soil type rather than make more controversial claims that it represents a unique kind of extinct soil. The progress of paleontology has been similar. Leafless, branching stems (Psilophyton princeps) in early Devonian rocks were originally compared with modern whisk ferns (Psilotum nudum). Large teeth (Iguanodon rnantelli) from early Cretaceous rocks were compared with those of modern lizards (Iguana iguana). It took much research to establish that these fossils represented extinct kinds of organisms: trimerophytes (Taylor & Taylor 1993) anddinosaurs(Bent0n1997).Atfirst,the nature of fossil plants and animals was understood by means of modern analogy With more copious collections of these fossils and more sophisticated methods of studying them, they were better assessedon their own terms. A history of plants and animals on land could then be reconstructed from the beginning. Comparable understanding of the fossil record of soils has not yet been secured. It is likely, however, that paleopedology will become as important to the development of theoretical soil science, ecology and historical geology as those studies have been to paleopedology.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 15 Soils of other worlds
The fertile imaginations of writers of science fiction have seldom devised life forms of other planets that were totally unlike those of Earth. Most extraterrestrial creatures of fiction are recognizably like insects or humans, complete with legs, eyes, ears and mouths. Imagined landscapes of other worlds also owe much to human experience with deserts, jungles and cities here on Earth. Life beyond Earth, although eagerly anticipated, remains to be demonstrated. Our present understanding of landscapes of the Moon, Mars and Venus, on the other hand, confirms the old adage that fact is stranger than fiction. Soils on the Moon, for example, are produced principally by micrometeoroid bombardment and those of Venus form by what would be regarded as metamorphic alteration on Earth. These strange landscapesformed by sand blasting and glazing have no clear analogs on Earth. These other systems of soil formation broaden our intellectual horizons on surficial processes. Asecondreasonforthestudyof soilsof otherworldsis that they may give clues to conditions during the first 700 myr or so of Earth history unrecorded in sedimentary rocks. The oldest well-preserved sedimentary rocks are some 3 500 myr old in the Barberton Mountain land of South Africa (Lowe & Byerly 1999) and the Pilbara region of Western Australia (Buick et al. 1995). Rocks of 3800Ma near Gahaab (Nitk) in south-west Greenland were once sediments, but are now strongly metamorphosed (Rosing et al. 1996). The oldest known igneous rocks on Earth, at 4000Ma, are the Acasta Gneisses of Canada (Bowring & Housh 1995). Even older single crystals of zircon, dated at 4276Ma, were eroded from igneous rocks of that age into younger (3350Ma) sandstones in which they are found near Narryer, Western Australia (Compston & Pidgeon 1986).On the Moon and Mars there are land surfaces as old as 4000Ma and rocks perhaps as old as 4500Ma (Greeley 1994). Inclusions rich in calcium and aluminum (CAI) within meteorites are thought to be the first solid objects formed in the Solar System, and have
an age of 4566Ma (Birck & Allkgre 1988).Thus ideas concerning soil formation on the primordial Earth depend on information about extraterrestrial soils. One problem for the study of altered planetary surfaces is what to call them. They are called soils in this book. This common English term also is used in many scientific accounts of lunar, Venusian and Martian surfaces. Other reports show preference for the term regolith, which includes a greater variety of surficial materials such as soil, sediment and saprolite. Soil can be distinguished from saprolite by the intensity of its alteration and from sediment by its formation in place without appreciable lateral transport. Gradation is another vague term for all the sedimentary. pedogenic and other surficial processes that create regolith from bedrock. This vagueness of definition has commended both terms for geological reconnaissance of planetary surfaces based on images from orbiting satellites (Greeley 1994).The terms regolith and gradation deliberately gloss over roles of soil formation, sedimentation and biological processes, but the teasing apart of these various processes is the main purpose of this account of extraterrestrial soils.
Soils of the Moon On the Moon, soil-forming conditions are very different from those on Earth. Most striking is the Moon's thin atmosphere and rarity of free water at the surface (Table 15.1). Micrometeoroidsand cosmic radiation bombard a surface unaltered by chemical weathering (Fig. 15.1). Seasonal and daily variation in temperature is extreme. At the Apollo 1 7 site, the soil reached temperatures above boiling (111"C) by day and well below freezing (-17l0C)bynight(Taylor 1982). Life has not yet been discovered on the Moon despite exploration by automated vehicles and astronauts. Life as we know it on Earth would be very unlikely in most areas of the Moon now explored because water and car209
210
Chapter15
bon are extremely scarce, although remote sensing has revealed water ice at both lunar poles (Feldman et al. 1999). The surface of the Moon has a topographic relief of 16km.Mare basins produced by the impact of large Table 15.1 Comparison of planetary bodies (Greeley1994).
Diameter (km) Mass (Earth=1) Density ( g ~ m - ~ ) Surface temperature (“C) Main atmospheric gases Atmospheric pressure (bars) Siderealperiod of rotation (years) Equatorial surface gravity (cm SKI)
Moon
Mars
Venus
Earth
3476 0.0123 3.34 c.80
6794 0.1074 3.95 ~.-50
12100 0.8150 5.27 c.470
12 756 1 5.52 c. 2 0
None
CO,
CO,
0
0.008
93
1
0.075
1.88
0.616
1
162
371
890
978
meteoroids and filled with flood basalts are on average 3 4 k m below old cratered highlands, and the South Pole-Aitken Crater is an astounding 2 500 km in diameter (Nozetteetal.1994). Comparedwith those on Earth, a limited array of rock types has been discovered on the Moon. Many are basaltic and ultramafic in composition. At the other compositional extreme are anorthosites and anorthositicbrecciasof the highlands.These are richer in alumina(Al20,,9-13wt%)andlime(CaO,lO-l2wt%)than basalts, but have silica (SO,, 3 8 4 6 wt %) within the basaltic range. Lunar basalts have some free quartz, but compared with basalts on Earth they are rich in iron (FeO, 16-2Owt%) and titania (TiO,, 0.3-13wt%) and poor in soda (Na,O, O.l-O.Swt%) and potash (K,O, 0.02-0.4 wt%, but in someup to 1.11wt%). Radiometric dating of highly brecciated rocks from the lunar highlands indicates that most of the anorthositic crust was differentiated by 44004500Ma. A part of this early differentiation may have been in magma oceans created during origin of the Moon by impact with Earth of a Mars-size object at a time of heavy meteoroid bombardmentin an early solar
Figure 15.1 SurfaceprocessesontheMoon(adaptedfromEglintonetd.1972: withpermissionfrom W.H. Freeman&Co.).
Soils of other worlds system littered with rocky debris (Lee et al. 1997). The rate of bombardment has declined since then and the relative ages of different surfaces can be reconstructed from crater densities visible in images of the lunar surface. Mare-filling lavas range in age from 4000 to 2000 Ma. Craters as old as 2000Ma still have fresh outlines and clear rays of ejecta. The persistence of such features is an indication that rates of erosion, soil formation and tectonic change on the Moon are all much more sluggish thanon Earth.
Soil composition The unconsolidated material at the surface of the Moon is typically fine grained and very poorly sorted. Its porosity is high (41-70%)and density low (0.9-1.1 g c m 3 ) compared with the density of its constituent particles (2.9-3.2g~m-~).There are three main kinds of soil components: rock fragments, mineral grains and glass particles. Rock fragments are the most variable in size, ranging from silt-sized grains to large boulders (Fig. 15.2).
Figure 15.2 Lunar soil near the Apollo 15 site on the margin of Mare Imbrium,overlooking Hadley Rille: acollapsedlavatube that containsconspicuoustalus blocks. The Apollo deep core was drilledon the plain in the left distance,3.8 kmfkom the astronaut(NASA photographAS1 5-85-01 145 1 courtesyof EJ. Doyle and National Space ScienceData Center).
2 11
Fragments of highland anorthosite and mare basalt have been mixed by a number of large impacts. Mineral grains constitute most of the silt size fraction (0.064.03 mm). This is because of the fine grain sizes of crystals in lunar basalts. Plagioclase and pyroxene are present in most soils in substantial %). Olivine constitutes only a amounts (a few to 40~01. per cent or so and is sporadic in occurrence. Much less common are grains of ilmenite, spinel, metallic particles (kamacite and taenite), phosphide (schreibersite) and sulfide(troilite). Glass particles are of two main kinds: homogeneous glasses and inhomogeneous glass-bonded aggregates called agglutinates. Both are variable in size, ranging down to fine silt size (<2 pm) and up to 1mm in the case of agglutinates or 2 cm in the case of some homogeneous glass particles. Homogeneous glass particles vary in shape from spheres to dumbbells and teardrops (Fig. 15.3).These homogeneous glasses are thought to be volcanic ash or rock melted by the heat of meteoroid impacts. Agglutinates are highly irregular masses of glass and
212
Chapter 15
Figure 1 5.3 A teardrop-shapedhomogeneousglassparticle from the lunar soil near the Apollo 1 5 site and HadleyRille (NASAphotographS72-52 308 courtesy of F.J. Doyle and NationalSpaceScienceDataCenter).Scalebar represents 50pm.
crystals held together by bridges of glassy cement between mineral grains and rockfragments. Some agglutinates have a ring or bowl shape, like a miniature crater. These are thought to have formed where impact melt has spread outward to cement surrounding grains of soil (Fig. 15.4). Other less distinctively shaped agglutinates may have been cemented by scattereddrops of impact melt or may be parts of agglutinates broken by later impacts. The chemical composition of lunar soils is very close to that of nearby parent material. Soil formation on the Moon is a physical rather than chemical process. Nevertheless, lunar soil grains have thin (20-50 p)amorphous rinds formed by local cooling of vapor from impacts, and these are enriched in Si and S, and depleted in Mg, Ca, A1 and Ti compared with the rest of the grain (KellerLGMcKay 1993).
Soil development The unconsolidated material at the surface of the Moon has been estimated from the geometry of craters to reach 16 m in thickness and from seismic observations attheApollostations toreach 12.2m.Not allthismaterial can be considered soil in the sense espoused here. Processes acting to form the soil include micrometeoroid bombardment, mixing as a result of solar-induced charge separation, churning associated with temperature changes and solar wind sputtering. Even the most deeply penetrating of these, micrometeoroid bombardment, affects only the uppermost few millimeters.
Figure 15.4 Ring-shapedagglutinatefrom the lunar soil at the Apollo 12 site inMare Cognitum(NASAphotograph S71245 75 courtesy of D. McKay,F.J. Doyle and National Space Science Data Center).Scale bar represents 5 0 pm.
Larger impacts broadcast a blanket of ejecta that interrupts soil formation. In the Apollo 15 deep drill core, large impacts are recorded by erosional planes and by large rock fragments, including numerous rocks from outside the mare basin in which the core was drilled (Fig. 15.5). Beds in the core can be considered a long sequence of impact-depositional events, some of which were modified by soil formation, in a similar fashion to a sequence of alluvial paleosols in a fluvial sedimentary sequence on Earth. Layers on the impact-cratered Moon are not so laterally continuous as on Earth, and cannot be correlated between cores a few kilometers apart (Basuetal. 1987). Developmental stages of lunar soilsare based on their degree of reworking by micrometeoroids. The starting material is a coarse-grained and poorly sorted blanket of impact ejecta. Under micrometeoroid bombardment this is broken down to a more even and smaller grain size, and the proportion of agglutinates increases (Fig. 15.6). Micrometeoroid bombardment also adds meteoritic metal and reduces iron (Fez') in silicatesto metallic iron. This effect can be quantified rapidly and nondestructively by measuring the intensity of ferromagnetic resonance (I,) and concentration of nonmetallic iron (FeO) to calculate a surface exposure index. In the Apollo 1 5 deep drill core both ferromagnetic index and agglutinate abundance can be used to recognize especially welldevelopedpaleosols(Fig. 15.5).
Soils of other worlds graphic mean qram Slte % oggiuflnates
125 6 2 5 3 I 3 p m
cm
relalwe ferromagnetic resonance
(90-150pm)
I,/FeO
213
non-marelmare
I C rack fragment ratio
v
0
0
v
v
I
v 0
v v
0
a a
T
breccia
planar surface
sit1 ond sand
erosional surface
v
interpreted paleosol
0 interpreted
larpe impact debris
Figure 15.5 An interpretationof the mostprominent buried soils in the Apollo 15 deep drill core,based on soil maturityindices (mean grain size,per cent agglutinatesand ferromagneticresonance)and on evidencefor large impacts (non-marehare rock fragmentratio:data from Heiken et al. 19 73,1976).
It takes a long time for a lunar soil to darken with metal and agglutinates. The Apollo 1 5 deep drill core penetrated at least seven recognizable paleosols in the upper 2.4m of some 5 m of unconsolidated material overlying marebasalt, datedat 3300Ma(Taylor 1982). There also is a general enrichment in agglutinates and ferromagnetic index toward the top of the core that may reflect increased redistribution by large impacts of micrometeoroid-modified soil material through time. Nevertheless, each paleosol could have been hundreds of millions of years in the making. Comparably slow rates of soilformation on the Moon have been confirmed also by calculations based on the rate of influx of micrometeoroids of a size range capable of forming agglutinates (McKay& Basu 1983).
Soils of Venus Venus is a permanently clouded planet. Its surface re-
mained a mystery until landings of the Venera probes, and until radar mapping from orbiters and Earth-based observatories.The climate of Venus is so hostile that few landers have been able to operate for more than a few hours (Table 15.1). The upper cloud deck, 70-56km of above the surface, is an aerosol (300 particle~m-~) small (0.1-2.8 pm) crystalsanddropletsof sulfuric acid. The bottom cloud layer, 5 0 4 8 km above the surface, is similarin composition,but isdenser (400particlesm-’). Below that the atmosphere is clear enough that images of the Venusian surface could be made under natural lighting. The cloud layers trap and circulate heat so effectively in a kind of greenhouse effect that there is little difference in heat between the equator and the poles or between day and night. At the level of the lower clouds the temperature is close to 72°C and the pressure about 1bar. Temperatureand pressurein the highlands (elevations of 10-12 km) are oppressive (374°C and 41 bars) but discernibly lower than in the lowlands (465°C and
2 14
Chapter 1 5
49
I
39
OI
15.3% AVERAGE
29 -
36.57.
BLACK GLASS
ORANGE AND
14 A
I
9
5pm
I
.~
AGGLUTINATE CONTENT 46.9%
I
J
I
I
3+ M, 49 125pm 6 2 . 5 ~ MEAN GRAIN SIZE
Figure 15.6 Stagesin the developmentof lunar soil, as shown by changes in abundance of agglutinates,mean grainsize anditsstandarddeviation(fromMcKayetul.1974;with permissionfrom PergamonPress).
96 bars: Warner 1983).Windsof this thickatmosphere are effective in transporting dust and sand, as can be seen from numerous wind streaks in radar images of the Venusian surface (Greeley et al. 1994). The chemical composition of the Venusian atmosphere is mainly carbon dioxide (97 vol. %) with small amounts of nitro%), and gen (1-3vol. %) and water vapor (0.1-0.4~01. traces of helium (250ppmV), neon (6-250ppmV). argon (20-200 ppmV), sulfur dioxide (240ppmV) and oxygen (60 ppmV) (Barsukovet al. 1982). At the high temperatures on the surface of Venus, neither liquid water nor organic compounds are stable. No trace of life or biological processes has been detected. Venus has the kind of atmosphere that would be predicted on the basis of equilibrium thermodynamics, very unlike the peculiar, biologically influenced atmosphere of Earth (Lovelock 1979;Lenton 1998). Radar mapping of Venus has shown that it has less topographic relief than Earth despite its comparable size. Some 60% of the planet is within 500m of the modal planetary radius. The highest point is Maxwell Montes, which rises 10.8km above the modal radius. The lowest point is Diana Chasma, some 2.9 km below this datum. Images transmitted by Venera landers 9 and
10 depict bouldery slopes around a large volcanic upland, but Veneras 1 3 and 14 landed in broad plains (Fig. 15.7). The soil at the Venera 1 3 site is comparable in chemical composition with alkaline melanocratic basalt and that at theVenera 1 4site with tholeiitic oceanic basalt of Earth (Basilevsky et al. 1985). Their compositions are reasonable considering the location of the alkaline rocksnearalargeriftvalleyand of thetholeiiticrockson a plain near a large shield volcano: tectonic locations where such basaltic rocks would be found on Earth. Volcanic activity is widespread on Venus (Crumpler et al. 1993). There are also arcuate asymmetric trenches like those of subduction zones on Earth (Sandwell & Schubert 1992). In other areas of Venus there are elevated regions of about the size of Australia and Antarctica (Ishtar and Aphrodite Terra, respectively), perhaps formed over mantle plumes (Phillips & Hansen 1998). These include banded regions superficially similar to orogenicbeltsonEarth (Zuber&Parmentier 1995),and extensiveridged plains that may reflect auniquely Venusianformof thermaldeformation (Solomonetal. 1999). Someform of crustal resurfacing also is indicated by the low density of craters on Venus. Only stony meteoroids larger than about 80 m in diameter and iron meteoroids larger than 30 m survive passage through its thick atmosphere to impact with the surface. The smallest surviving meteoroids would make craters about 3 0 0 4 0 0 m in diameter for stony meteoroids and 150-200m for irons. Despite this limitation, age estimates based on the well-calibrated impact frequencies on the Moon have indicated that Venusian surfaces are noolderthanabout 500Ma(Price&Suppe1972;Kaula 1995).
Soil composition Information on the chemical composition of Venusian soils was relayed back to Earth by Venera landers, which were equippedwith miniaturizedx-ray fluorescence analytical laboratories. Calculationsbasedon the observed impact and deformation of the landers calibrated with experimental dropping of comparable landers on Earth (Basilevsky et al. 1985) indicate that the surface material at both sites is low in density (1.2-2.5g~m-~)and highly porous (SO-60%). This is a surprise because the ground in transmitted images (Fig. 15.7) looks like
Soils of other worlds
2 15
Figure 15.7 Images of the surfaceof Venus relayedby Soviet Venera 9, Venera 10,Venera 1 3 and Venera 14 (photographscourtesy of R.J. Allenby andNASANationa1Space ScienceData Center).
either pahoehoe lava or indurated ripple-marked siltstone. Also unusual is the very low electrical resistivity of thematerial( 73-89 ohms)andthedarkercolorof the loose dust compared with the lithified-looking fragments. Dry rock or dust should have a high resistivity and dust should be lighter colored than rocks from which it was derived. These meager data and observations are permissiveof a number of interpretations, each of which is based on such a complex series of assumptions that they should not be taken too seriously.Although the chemical composition of the soils at the Venera 13 and 1 4 sites is basaltic, their density, resistivity and general appearance is more like that of siltstone. Cementation similar to that of a potter’s glaze would be expected on Venus, considering the high temperature and pressure. A cement of light-colored salts around and bridging silicate grains could explain the lighter color of the slabs compared with the dust, their low electrical resistance and low bulk density, The fretted, glazed and pitted appearance of these indurated-looking rocks could be due to sublimation or melting. Theoreticalmodels for the mineralogical composition
of Venusian soils also have been proposed based on the available partial analyses and other information about the surface environment. Especially critical to such calculated reconstructions, yet poorly constrained, is the abundance of oxygen and water at the surface. In the model of Barsukov et al. (1982) both H,O ( 5 x 10-3v01. % ) a n d O , ( 3 . 3 ~ 1 0 - ~ ~ v%)areassumedtobe ol. scarcecomparedwithSO,(1.85~ 10-2vol.%)andC02 (96 vol. %). Under such conditions sulfate cement would be more abundant than cements of hydrated silicates or silica. The mineral assemblages favored on basaltic soils like those at the landing sites would be primarily albite (28 vol. %) andanorthitefeldspar(24vol. %), withlesser amounts of clinoenstatite (13 vol. %), pyrite (1Ovol.%), microcline (1Ovol.%), anhydrite (8vol. %) and quartz %). Different mineral assemblages would be fa(7~01. vored at higher elevations, where it is thought there are slightly greater amounts of gases produced photochemically in the atmosphere, such as SO,, CO and 0,. Here basaltic rocks would be altered to albite (27vol. %), anhydrite (18 vol. %) and quartz (17 vol. %), with lesser clinoenstatite (13 vol. %), microcliie feldspar (1Ovol.%), hercynite (10vol. %) and pyrite ( 5 vol. %).
2 16
Chapter 1 5
Soil development
Assuming the partly theoretical model for soil composition outlined above, Venusian soil formation consists of solid-phase mineral transformation together with cementation of loose parts of the soil with calcium sulfate.These processes of soil formation in lowland soilson basaltic terrains are interrupted by thin layers of windblown dust from the highlands. The degree of soil development can be assessed by the extent of glazing and cementation of the originally dusty surface. In both the highland andlowlandchemicalmodelsfor basaltic soils, weathering produced quartz and feldspar at the expense of more matic minerals so that soils become more siliceous and aluminous than their parent materials. Weathering on Venus also may produce sulfates and sulfides, which may be buried and recycled in sediments (Warner 1983). The tendency toward more felsic sur6cial composition and crustal recycling of the volatilesulfur are general features of geochemicalcycles onEarth. Amajorgeochemicaldifference betweenEarth and Venus is the instability of carbonates and organic matter at the surface. These cannot accumulate and be recycled into sediments of the crust, so that carbon has accumulated in the atmosphere as carbon dioxide almost to the exclusion of other gases. By both these chemical and physical criteria the Venera 1 3 site has a less well-developed soil than that of the Venera 14 site. Such a relative age of these landscapes also would be expected from their geomorphological position. The Venera 1 4 site is farther out into a large basin (Nauka Planitia) than the Venera 13 site, which is in rolling foothills of an extensive highland (Phoebe Regio). It is unclear how much time is involved in converting a dusty surface to one that is as glazed and corroded as that visible at the Venera 14 site. Considering the high temperature and pressure on Venus, a thin layer of dust could be altered more rapidly than the hundreds of millions of years likely for formation of comparably thin soils on the Moon. If Venusian soil formation can be compared with pottery glazing or experimental simulation of rock metamorphism, the times involved in altering such thin dust layers may be only hours. Even within this theoretical framework for Venusian surface conditions, other kinds of weathering can be envisaged on bedrock rather than in depositional land-
scapes and on granitic rather than basaltic parent materials. Soil formation on Venus is clearly unusual compared with that on Earth. Only time and more information from space exploration will demonstrate exactly how unusual. Soils of Mars The Martian atmosphere has regional and seasonal weather patterns, dust storms and ice caps like those on Earth, but conditions are far fromEarth-like.The atmosphere of Mars is two orders of magnitude thinner than that of Earth (Table 15.1).This contrast may be related to Mars' smaller size and gravity, which allows gases to escape into space. What gases have remained are largely carbon dioxide (95 vol. %) with minor amounts of nitrogen (3 vol. %), argon (2 vol. Yo) and only traces of oxygen (0.13 vol. %), carbon monoxide (0.07~01.YO)and water vapor (0.03 vol. %).Although water vapor is scarce, it is close to saturationin the thin atmosphere, and large fluvial channels attest to its flow in the distant geological past. Most of the water remaining on the planet is frozen in permafrost and in polar ice caps. The low surface temperatures of Mars are due to its thin atmosphere and large distance from the Sun. The Viking landers recorded northern summer temperatures at a low northern latitude (ChrysePlanitia) varying from -90°C at dawn to a maximum of -30°C at noon. Temperatures to the north in midlatitudes (Utopia Planitia) were 5-10°C cooler. In Chryse Planitia there was a slow seasonal cooling of about 22°C as the northern winter approached(Carr 1981). These weather patterns and supposed 'canals' of Mars have stimulated the idea that the planet was alive with a highly technological civilization.Now that there are thousands of detailed photographs of the Martian surface, this possibility seems remote. The canals were probably optical illusions of early telescopes. Highresolution images of the channels and chasms are more like fluvial and tectonic features than engineering works (Baker 1982).Organic compounds were not detected by gas chromatography or mass spectrometry in the Viking landers. Nor was any life seen in images of the Martian surface. The life-like experimental results on the Viking landers were thus most surprising. In the pyrolytic release experiment, small amounts of radioactively labeled carbon monoxide and carbon dioxide were
Soils of other worlds incorporated into organic matter that was synthesized in contact with illuminated Martian soil within the lander. In the gas exchange experimentoxygen was rapidly released from Martian soil samples incubated in a humid environment.In the labeled release experimenta labeled nutrient solution of simple organic compounds was broken down to carbon dioxide when in contact with Martian soil. The progress of the first two experiments was little affected by high temperatures (ranging to 140°C and 9O"C, respectively) that would have curtailed a similar metabolic response of microbes. In contrast, the breakdown of organic matter was reduced (by 70%)at moderate temperature (50°C)and ceased at high temperature (160°C). These life-like results have been duplicated in the laboratory under simulated Martian conditions using iron-rich montmorillonite asasubstituteforMartiansoil(Banin 1986). Although there may not be any current life on Mars, there have been reports of possible microfossils within a Martian meteorite (ALH84001) found in Antarctica, where it arrived 13kyr ago, after long exposure to cosmic rays in space following impact blasting off Mars at 16Ma (McKay et al. 1996).The rock has been dated at 120Ma, but was shock metamorphosed at 4500 about 4000 Ma, and secondary carbonate containing the possible microfossils was created by aqueous alteration at 3900 ? 40Ma (Borg et al. 1999). Small chains of cell-like structures, organic geochemicals (polycyclic aromatic hydrocarbons) and associated lozenges of magnetite (250 pm across)have all been comparedwith these features of magnetotactic bacteria and nanobacteria on Earth (Folk & Lynch 1997). However, each of these lines of evidence has been disputed individually, to with the cell-like structures too small (0.02-0.1 p) be viable or perhaps artifacts of electronmicroscopy, the organic compounds isotopically and racemically like terrestrial contaminants, the magnetite with nonbiogenic screw-dislocations,and the carbonate zoned l i e that formed at high temperature (Bada et al. 1996: McSween 1997;Barratetal. 1998;Julletal. 1998).Neither the carbonates, nor other minerals, of meteorite ALH84001 have the petrographic appearance of soils or paleosols, but are instead selvages and fracture zones like those of a metamorphically and hydrothermally altered igneous rock (Scott et al. 1997). Furthermore, the isotopic composition of sulfides and sulfates in SNC meteorites does not indicate biological fractionation
*
2 17
(Farquhar et al. 2000). Although the demonstration of past life on Mars remains unconvincing,the prospect of such a discovery will ensure continued investigation. Mars is a topographically rugged planet, ranging from elevations of 2 7 km above the planetary radius at the top of the large volcano Olympus Mons to 4km below its radius in the floor of the Hellas Basin (Greeley 1994). There also are deep canyons with cliffed walls. Even in the relatively flat regions chosen for the Viking and Sojourner landers, there is local relief around boulders, dunes andchannels (Fig. 15.8). The chemicalcompositionof Martian soil androcksis more or less basaltic (Rieder et 41. 1997). Steep-sided composite volcanoes and cinder cones, and fretted and smooth plains of pyroclastic materials are rare on Mars. Most of the volcanoeshave a topographicallylow profile and are mantled with very long lava flowscharacteristic of basaltic lavas, in some cases perhaps as ultramafic as komatiites.Some kinds of basaltic and ultrama6c meteorites (SNC for original finds at Shergotty, Nakhla and Chassigny)represent rock samples from Mars that were ejected from the surface of Mars by large impacts,stored in solar orbits and then captured by the gravity of Earth (Vickery&Melosh1987).All of these but Antarcticmeteorite AM86001 are anomalously young (1300Ma) compared with other meteorites.They have an unusual oxygen isotope composition, negligible magnetic paleointensity and evidencefor long (0.5-10 myr) exposure to interplanetary cosmic rays that rule out most sources otherthanMars(Wood&Ashwall981). The surface of Mars includes several heavily cratered plains dating to 35004000Ma. Thus the planet has been assumed to lack plate tectonicorotherresurfacing, although there may have been tectonic activity early in its history, as indicated by magnetic lineations like those of sea-floor spreadingin Earth (Connerneyet al. 1999). The abundance of craters on runoff and outflow channels, shorelines and glacial moraines is an indication that all these features created by streams, oceans and glaciers date back to 2 500-3 500 Ma (Bakeretal. 1991). There are also many younger parts of the planet around volcanoes, craters, scarps, eolian dune fields and polar layered terrains (Malin et al. 1998). Soil composition Viking landers carried X-ray fluorescence spectrome-
218
Chapter 15
Figure 15.8 US Viking images of the Martian surfaceat (A) Chryse Planitia (NASAViking 1 event number 11A079)and(B) Utopia
Planitia (Viking2 event number 2 1A024) (both courtesy of RE. Arvidson andNational SpaceScienceDataCenter).
ters and Sojourner Rover had an alpha proton X-ray spectrometer, which found Martian soil to be remarkably uniform in chemical composition at these distant sites. Magnets of the landers extracted loose mineral grains with the magnetic properties of maghemite. Both carbon dioxide and water were detected by the spectrometers after heating soil samples to 500°C and these bound volatiles could amount to a few per cent of the sample. The Viking labeled release experiment demonstrated that unheated fine-grained material contained an oxidant strong enough to decompose organic matter to carbon dioxide. To these experimental data can be added information from images of the areas around the sites (Sharp &
Malin 1984; Smith et al. 1997). All three sites show boulders surrounded by what appears to be wind-blown material of fine silt to sand size (10-100 pm diameter). Orange pebbles that were analyzed had a low density (1.1k 0 . 1 5 g ~ m - andahigh ~) sulfitecontent (6.5-9.5 wt%) like the 6ne material. Although these looked like rocks, they probably are fragments of a cemented soil layer. The blue-gray and green-gray, vesicular and coarse-grained boulders have the chemical composition of basaltic and ultramafic lavas, and may include conglomerates with clasts of that composition (Rieder et al. 1997). Freshly exposed deep layers of the fine-grained soil are a similar blue-gray color to the rocks from which they were presumably derived (Fig. 15.9). The overall
Soils of other worlds
2 19
Figure 15.9 Areconstructedsoilprofileatthe Vkmg 1 landmg site on Mars (basedon data of Sharp&Malin1984)
reddish color of the surface is due to a thin (1mm) oxidized surface layer of dust. Oxidized layers also can be seen within freshly excavated dust between the boulders. The exposed surfaces of some of the rocks also show a dusty oxidation rind, but others show a gray surface. Beneath the landers where the loose dust and rock were most disturbed by touch-down, there is an orange layer. This is cracked into fragments several centimeters across in a way that suggests it is only a centimeter or so thick. This layer is a salic horizon (of Soil Survey Staff 1975) considering its high CaO (5.0-5.3 wt %) and SO, (6.5-9.5 wt %) content. Additional evidence comes from low-temperature alteration minerals in SNC meteorites (McSween 1994). The nakhlite meteorite Lafayette, for example, contains smectite, calcite, gypsum and ferrihydrite. Thus Martian soils consist largely of clay and evaporite minerals. Using these assumptions and the chemical analyses at the Viking 1or Chryse site (Bell 1996), the soil composition there probably has nontronite as the main clay mineral (47v0l.Y~)with lesser amounts of montmorillonite (17 vol. YO)and saponite (15 vol. %). In this model the evaporitic minerals were kieserite (13 vol. %) and calcite (7vol. %), and there were traces of heavy minerals such as rutile and maghemite. The salic horizon does not differ greatly in composition from the loose dust and is presumed to be more thoroughly cemented with kieserite and calcite.
Soil development Theparent material forthe assumedsmectitesandevaporites on Mars appears to have been underlying basaltic or ultramafic igneous rocks or breccias. If one of the least mafic of the SNC meteorites (Shergotty) is taken as a parent material of the Martian soil. the only differences clearly beyond the experimental errors of the
Viking and Sojourner analyses are an enrichment of S and loss of Ca in the soil. An assumption that known Martian soils formed from materials like the more mafic SNC meteorite Chassigny gives an impression of more profound weathering: increases in Si, A1 and Ca, and losses in Fe and Mg. Even so. Martian soils are shallow and the degree of weathering is slight compared with Antarctic desert soils on Earth (Campbell & Claridge 1987). Leaching of salts and formation of clays by hydrolysis in water is another implication of the hydrated smectitic and evaporitic composition of the Martian soil. The formation of these materials on the present Martian surface is unlikely to be rapid or widespread because of extremely low temperatures. The runoff and outflow channels of Mars provide evidence of some surface water before 2500Ma. It could be that most of the salts and clays of the Martian soil formed at that time and have been blown around or held in a deep freeze since then. Other less effective and local sources of water can be imagined. Volcanic eruptions or large meteorite impacts through ground ice could generate large amounts of water for relatively short periods of weeks or months. More effective long-term alteration could be achieved by volcanically heated groundwater insulated from the Martian surface by frozen soil. It has also been found that hydrolytic weathering reactions can occur in an unfrozen monolayer of water molecules around mineral grains at temperatures well below zero in Antarctica (Ugolini 1986). Although the effects of volcanoes. impacts and of monolayer water films would be extremely slow compared with weathering under surface water, these processes cannot be discounted because of the thousands of millions of years available for the alteration of some Martian landscapes. A paradoxical feature considering that the Martian atmosphere has very little oxygen is the thin sur6cial fer-
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Chapter 1 5
ruginized layer of the soil responsiblefor the red color of the planet. Deeper and more pervasive oxidation would be expected if Mars had a more oxygenated atmosphere at some time in the geologicalpast. A likely mechanism is direct photo-oxidation under ultraviolet light. This process has been shown experimentally (Braterman et al. 1983)to beeffectiveinacidictoneutralaqueoussolutions in which the oxidation of iron is coupled with the reduction of water to hydrogen gas. On Mars this process can be imagined to have occurred slowly over thousands of millions of years within water films or other local sources of water, or the oxidation could be a remnant of that distant geologicaltime when there was free water on Mars. Photo-oxidation affects only a thin surface layer because the penetration of ultraviolet light is curtailed by surficialaccumulation of opaque oxidized minerals. Using these ideas concerning soil formation on Mars, the degree of development of soils at the three landing sites can be compared. At the Viking 2 (Utopia) and Sojourner (Ares Vallis) sites the boulders are angular and lying on the surface rather than embedded in the soil. They are blue-gray in color and show few red or greenish gray weathering rinds. The salt crust at Ares Vallis is smooth and thin (probably 5mm near rock Scooby Doo; Rover Team 1997).The salt crust at Utopia is smooth to lumpy and a centimeter or so thick (Mutch et al. 1977). These are less developed soils than that at the Viking 1(Chryse)landing site, where the salt crust was exposedandcracked, but not penetrated. Here, halfburied boulders have extensive greenish gray and red alterationrinds, andmost boulders are rounded and pitted (Binderetal. 1977). Interestingly,suchdifferencesindegree of weathering do not correspond to the presumed geological age of the three sites. The well-developed Chryse soil is on an extensive plain near where outflow channels end and the poorly developed Ares Vallis soil is within the outflow channels. The age of both sites from crater counts is about 3200Ma (1200-3800Ma; Hartmann et al. 1981). The Utopia site, with intermediate soil development, is in a region dated at 1800Ma (2300-600Ma), on the debris apron of the large (104km) impact crater Mie only 17Okm away. Fluvial outwash and erosion before 2500Ma in Ares Vallis may have destroyed soil development there. Nevertheless, soil formation is clearly slow and ineffectiveon Mars.
Soil formation also is slow in cold deserts of the Antarctic Dry Valleys on Earth, where annual temperature is a little less than -1O"C, summer maximum is 30°C and mean annual precipitation is 45mm. Here comparably thick (4 cm) salic hardpans andnontronitic clays are found only in soils more than 2 myr old. Oxidation rinds (primarily rock varnish) are not found on boulders exposed for less than 800kyr (Campbell & Claridge 1987). Similar hydrolytic formation of clays, salinization and oxidation may have been possible on Mars before 2500Ma when there was surface water available to create channels.The lander sites are on land surfaces that approach that age and now endure much colder and drier climatic conditions than in Antarctica. Thus,itislikelythatthesearerelictcolddesert(Salid)paleosolsformedunder conditions l i e those in the Antarctic Dry Valleys and then deep frozen for many thousands of millions of years with only minor alteration in the form of surficial oxidation and of added eolian dust and ejecta from impacts.
Meteorites In the long history of human fascination with meteorites, they have been regarded in many ways. Some meteorites were worked for iron and others were worshipped as divine gifts. Their fiery passage through the atmosphere as a meteor or shooting star was widely regarded as an important omen. It is now known that meteorites are fragments of rock orbiting the Sun that fall within the influence of Earth's gravity.They are still esteemed as evidence for the nature and origin of our Solar System. Three main kinds of meteorites are recognized (Table 15.2). The iron meteorites consist mainly of metal. Stony irons contain a mixture of iron metal and silicate minerals. Stony meteorites consist of silicate minerals with only small amounts of metal. Stony meteorites can be divided into those that have a chemical composition approaching that of the Sun (chondrites) and those that have a chemical composition more like basaltic igneous rocks (achondrites). Most of the chondrites contain distinctive spherical or ellipsoidal mineral grains called chondrules (Fig. 15.10) from which they are named. Achondrites have an igneous or brecciated texture lacking chondrules. Different kinds of chondrites are most effectively discriminated by their chemical composition
Soils of other worlds
22 1
Table 15.2 Different kindsof meteorites (fromWasson 1984). Class I. Stones Carbonaceouschondrites (CV)
Carbonaceouschondrites (CO)
Carbonaceouschondrites (CM)
Carbonaceouschondrites (CI)
Ordinary chondrites (LL.L, H)
Enstatite chondrites (K,EH)
Achondrites(EUC, HOW, DIO, URE. AUB. SNC) 11. Stony irons
Selectedcharacteristics
Find
AI/Si(atom%): 1 2 metalFe/Si: 0.6-19 C: 8-50 mgg-I H,O: 20-220mgg-' 1 mm chondrules AllSi (atom %): 9 metal FelSi: 2.3-1 5 c: 2-10mgg-' H,O: 3-30mgg-' 0.2 mmchondrule AI/Si(atom%): 10 metalFe/Si: 0 . 1 4 . 5 C 8-2 6 mg g-' H,O: 20-160mgg-' 0.2 mm chondrules AI/Si (atom%):9 metalFe/Si: <0.1 C: 30-50mgg-' H,O: 180-220mgg-' lacking chondrules Al/Si(atom%):7 metalFe/Si: 2.7-52 C: < 10mg g-' H20:<30mgg-' 0.4mm chondrules AIlSi (atom %): 5 metalFe/Si: 47-72 C: < 10mgg-' H20:<30mgg-' 0.5 mmchondrules Basaltic and ultramafic igneous rocks
Mesosiderites (MES) Pallasites (PAL)
Brecciaswith iron and minerals (mainlypyroxene) Breccias with iron and minerals (mainlyolivine)
In. Irons Irons
Mainly metallic
Totals
and secondarily by petrographic features. Especially diagnostic chemical properties are the abundance of volatileelements most common in the Sun (mainly C, H, N and S) over refractory elements most common in
3
Fall
%
0.03
%
8
1.1
1
0.1
6
0.8
0
0
14
2.0
0
0
5
0.7
438
45.0
558
78.9
8
0.8
13
1.8
15
1.5
64
9.0
14 33
1.4 3.4
6 2
0.8 0.3
459
47.3
31
4.4
971
100
707
100
Earth's crust (mainly Al, Ca and Ti). Also used is the degree of oxidation (inverselyproportional to the ratio of metallic iron to silica, among other measures). By a widely accepted system of shorthand, many kinds
222
Chapter 15
Figure 15.10 Petrographic thinsectionof AUende carbonaceous chondrite under crossed nicols. showing barredchondrules and irregular areas of poorly characterized phases (Arizona State University Center for Meteorite Studiesspecimen 8 18).Scale bar represents 1mm.
of meteorites are named from the initials of a locality where a particularly well-known example was found. For example, CM meteorites are named from the carbonaceous Murchison meteorite found near the town of that name in Victoria, Australia. Other shorthand names of meteorites are contractions of names, such as EUC for eucrites, or acronyms, such as SNC for shergottites-nakhlites
Carbonaceouschondrites Unlike rocky or metallic meteorites, these rare chondrites are so soft and friable that they can be crushed between the fingers. Inaddition tochondrules and mineral grains such as diamonds that formed at high tempera-
tures (1300-1 SOOOC), from impact, lightning, shock waves and electromagnetic forces (Ott 1993; Weidenschilling et al. 1998). carbonaceous chondrites have a matrix of minerals that formed at very low temperatures (O-SSOC) in the presence of water. This includes a variety of iron-rich smectitic and serpentine-like clays (Tomeoka & Busek 1988), calcic micas (Keller & Busek 1991)and organic molecules including amino acids (Endress et al. 1996). Other evidence of hydrous alteration is framboids and plaquettes of magnetite (Kerridge et al. 1979), and veins of epsomite, gypsum, calcite, breunnerite and siderite (Richardson 1978). Some carbonaceous chondrites also show high-temperature (2 50-300°C) aqueous minerals such as talc (Brearley 1997, 1999; Hutcheon et al. 1998). indicating hydrothermal or metamorphic alteration. The low-temperature phases were not created by reaction with air and water after falling to Earth. Many of these meteorites have a thin protective fusion crust formed at high temperature during passage through the atmosphere. In some cases they have been recovered very soon after they fell. Moreover, the matrix material has been dated radiometrically using 207Pb120hPb, at no more than 60myr younger than associated inclusions rich in Ca and A1 (CAI),dated at 4566Ma (Birck &Allegre 1988).The 129Xe/'271 radiometric ages of magnetite in carbonaceous chondrites aretheoldestknown,butclose(within 1myr) tothoseof similarly dated minerals in other chondritic meteorites (Lewis & Anders 1975). The decay of 53Mnto 53Crin
Soilsof otherworlds carbonate is evidence that it crystallized no more than 50 myr after formation of the parent body (Endresset al. 1996). The source of carbonaceouschondrites is likely to be small (about 500km diameter or smaller)asteroids beyond the orbit of Mars. The spectral reflectance of carbonaceous chondrites is similar to that of some (C type) asteroids, such as 130 EIectra (Cruikshank & Brown 1987)and253Mathilde(Veverkaetd.1997).Theseare periodically perturbed into unstable (about 10myr duration), elliptical,Earth-crossing orbits from their more permanent orbits in the asteroid belt between Mars and Jupiterand within the Oort cloud at the outer reaches of the Solar System. Two divergent views on the origin of carbonaceous chondrites can be caricatured as frozen solar nebula (Wasson 1984) and earliest soil hypotheses (Bunch & Chang 1980).By the frozen solar nebula hypothesis the variouskinds of chondritesformed by coolingand differentiation of a solar nebula of dispersed gas and dust. By this view the nebula became chemically and physically differentiatedaccording to distance from the Sun. The most highly differentiatedmeteorites, stones and irons, formed closest to the Sun on magmatically differentiated asteroids. Among the carbonaceous chondrites, the volatile-poor refractory ones (CV) formed closest to the Sun (near the orbit of Jupiter)and the more volatilerichones (CI)atthefarthestreachesof theSolarSystem. An appealing featureof this model is that carbonaceous chondrites represent the least-altered products of nebular condensation and the most volatile-rich ones (CI) can be used as a baseline for assessing chemical evolutionof planetary bodies. By the earliest soil hypothesis, however, some carbonaceous chondrites include nebular and interstellar componentssuch as chondrulesand diamonds,but owe their peculiar bulk compositionto surficialalteration of planetismalsduring their earliest differentiationand degassing. Such soils may have been widespread in the earliest phases of formation of the Solar System, but have survived only on those planetary bodies too small to have undergone significantinternal magmatic differentiation and too far from the Sun for surficial modification. An appealing feature of this model is that carbonaceous chondrites represent fragments of the oldest land surfaces in the Solar System at a stage even earlier than possible ancient relict soils of Mars.
223
The most convincing evidence for the origin of carbonaceous chondrites as soils comes from their petrographical and mineralogical features. If they represented frozen solar nebular material, then the grains and chondrules would be little altered and have sharp boundaries. In contrast, carbonaceous chondrites contain pseudomorphs of chondrules, clasts and aggregates replaced by calcite, septechlorite clays and iron oxides, which imply extensive chemical alteration in aqueous solution (Bunch & Chang 1980).There are also cross-cutting veins, cavity-lining concentric zones of colloidal material (colloform structure),interlocking crystal growths of septechlorite clays, clayey alteration along cleavage planes of mineral grains, and mesh-like bright clay microfabrics (Iattisepicplasmic fabric).These are all more like alteration of materials by water in a soil than individual reactions between dispersed gas. liquid and dust. If we assume for the moment that carbonaceous chondrites are pieces of primeval soils, their conditions of formation can be reconstructed (Chang & Bunch 1986) under the familiar headings of climate. organisms, topographic relief, parent material and time. Temperatures were probably within the range of liquid water and below the upper limit for stability of gypsum (0-58OC). Water was freely available at times, but not always. Shrinking and swelling as a result of wetting and drying may have produced new microfabric (lattisepic plasmic fabric) and the several generations of veins filled with salts in some carbonaceous chondrites. From comparison of carbonaceous chondritesshowing different degrees of alteration, there appears to have been depletion of Ca2+,Na+,SO,2- and C1- from the primary minerals and chondrules to form clay by hydrolysis and salts by salinization. The hydrolytic solutions were presumably acidic, but strong evaporation and salinization most of the time would have given the soils an overall alkaline pH like that of desert soils on Earth. Weathering solutions also produced hydrous sulfide in colloidal materials, presumably by alteration of pyrite also present. Magnetite plaques may be aqueous alteration products of iron hydroxide. Conditions could not have been very oxidizing because a variety of organic molecules are present. From these considerations and the overall similarity of carbonaceous chondritic clays and salts to those of Martian soils, an atmosphere predominantly of carbon dioxide is liely.
224
Chapter 15
A case continues to be made for biological origin of the organic matter in carbonaceous chondrites (Folk & Lynch 1998),but especiallygood evidence against a biological origin are the even mix or only slight imbalance of left- and right-handed versions of asymmetric molecules, which in organisms are more consistently of one kind for any particular compound, and also isotopic compositions far from those of terrestrial organic matter (Cronin & Pizzarello 1997; Engel & Macko 1997). Many experimental studies (Rao et al. 1980) have shown that wet, clayey, salty,friable planetary surfaces could have produced a variety of organic compounds without the aid of living creatures. Topographic relief of the parent body is unlikely to have been very great because salts were not separated from the clays by groundwater flow to any great extent. Instead, they were deposited in cracks nearby so that the overall composition of carbonaceous chondrites remains l i e the nonvolatileportion of the Sun. Refractory carbonaceous chondrites (CV such as Allende) can be regarded as possible parent materials for volatile-rich chondrites (CM such as Murchison). The main differences between the two kinds of meteorites are the great gains in volatiles (H,O, CO,, C and SO,) in the supposed soil (Murchison) compared with parent material (Allende).These may have accumulated in the soil as they were degassedeither directlyduring accretion or later during impact melting. By this line of argument, the best developed of these soils would be represented by the most volatile-rich (CI) of the carbonaceous chondrites. The planetary parent materials altered to form them could have been as refractory as ordinary chondrites. It is also possible to envisage these different kinds of meteorite materials as soil horizons, with the least chemically altered ordinary chondritic materials at the base of a profile and the most volcatile-enriched materials at the surface in contact with the atmosphere and hydrosphere. Yet another course of alteration that was isochemicalcan be attributed to metamorphism on meteorite parent bodies. Some limits on the time for soil formation during this earliest period of soil formation can be gained from radiometric dating of carbonate veins, which postdate chondrules and inclusions (CAI)of carbonaceous chondrites by no more than 50myr (Endress et al. 1996). Slow development is liely if their atmosphere was thin.
Mesosiderites and howardites These two kinds of meteorites are stony irons and achondritic stones, respectively. Both are breccias with theirclastsmainly of basaltic andultramaficrocks. Both are thought to have formed on the surface of a planet or asteroid as soil breccias formed by the impact of meteoroids, including iron meteoroids in the case of mesosiderites (Jain&Lipschutz1973;Bunch 1975). Basaltic and ultramafic rock fragments of mesosiderites and howardites are the same as those found in other kinds of achondritic stony meteorites such as eucrites and diogenites. The mineralogical and chemical composition of these igneous rocks is compatible with an origin by partial melting of a source region of chondritic composition followed by varying degrees of fractional crystallization (Stolper 1977). In contrast to basalts of the Moon and Earth, eucrites and associated rocks appear to have formed within an asteroid or planet previously undifferentiated magmatically. Radiometric dating of all these different kinds of igneous meteorites has yielded ages of about 4530Ma: close to the origin of Solar System igneous activity. This early origin together with low pressures needed for the formation of eucritic magmas and the similar spectral reflectance of these meteorites to that of observed asteroids (especially 4 Vesta) has led to the idea that eucritic parent bodies were of asteroidal dimensions (about 500 kmdiameter; Binzel &Xu 1993). The fragmental texture of howardites and mesosiderites is similar to that of lunar breccias (Bunch 1975). They also resemble lunar soils in containing solar rare gases, crystals riddled by tracks of solar flare particies, impact-generated gIasses, and micrometeoroid craters. These features provide evidence of soil formation largely by meteoroid impact in a very thin atmosphere. Impact not only pulverized the surface but also added material to the soil.The trace element composition of howardites cannot be matched by different mixtures of basaltic (eucritic) and ultramafic (diogenitic) fragments like those obvious in thin sections, but can be explained by addition of about 2-3 wt % of material similar to CM carbonaceous chondrites (Chou et al. 1976). In mesosiderites there is abundant metal ( 17-80 wt %; Mason & Jarosewich 1973), presumably added by the impact of iron meteorites. The volume of these iron bodies and mixing of basaltic and ultramafic
Soilsof other worlds rock fragments are indications of more violent impact and a more rugged, cratered topography than outlined for the otherwise similar impact-generated soils of the Moon. Soil formation on eucritic asteroids may have been as violent, as rapid and as ancient as that envisaged for the period (before 4000Ma) of intense bombardment evident in the lunar highlands (Hartmann et al. 1981).
Relevanceto early Earth A new scientific discipline such as planetary geology offers many excitements: the joy of unanticipated discovery, the tantalizing search for patterns among data still in many ways inadequate and the evaluation of a variety of plausible hypotheses. The likely naivety of present views of planetary evolution and extraterrestrial soil formation (Fig. 15.11) can be appreciated more fully when expressed in the simple and unqualiied form of a creation myth as in the followingparagraphs. The visible universe began its present expansion from a primeval fireball at some 12 000-1 5 000 Ma. Most of
the matterreleased by that initialBig Bang condensed to hydrogen upon initial cooling. Eddies of hydrogen contracted gravitationally to form the first stars, lit by nuclear reactions producing helium. In the dense interiors of very large stars complex nucleosynthetic reactions were possible, producing heavy elements such as carbon, nitrogen, sulfur, silicon, aluminum, calcium and potassium, and small amounts of really heavy elements, iron and nickel. These large stars were unstable and prone to explode as supernovae, suffusing these heavy elements through space. One such cloud of rock, dust and gas began contracting gravitationally by about 4600Ma to form our Sun and Solar System. Most of the matter fell toward the center of the nebula, but some larger pieces and an associated sheet of dust and debris continued in orbit around the growing but still cold Sun. Both the future Sun and numerous orbiting planetismals swept up nearby dust and debris. Grinding, impact and shock waves associated with planetary accretion created the distinctive spherical chondrules still found in many meteorites. The originally chondritic Sun, planets and planNO ATMOSPHERE
porslble evolutionary pathways
carbonaceous chondrites
Venus
THICK ATMOSPHERE
Earth
0
sIIIcote sand siiicote crystals
Figure 15.11 Aschematiccomparisonof soil formation onEarth,Moon,Mars, Venus and some hypothetical meteoriteparent bodies.
lo,~blrock
bosoltic rocks
225
metoiiic iron
a [OoOo01
0 0 impact mlt spheres
chondrules
sulfate
oxldotion rind
ogglutinotes
root troces
226
Chapter 1 5
etismals continued to evolve depending on their mass and the influence of nearby bodies.The smallest masses (c50km diameter) changed little because of their small reserves of radiogenic heat and volatiles. Those that escaped the subsequent differentiation and ignition of the Sun are occasionally perturbed or fragmented from their remote orbits to land on Earth as chondritic meteorites. On larger planetismals (50-500 km diameter), gravitational compaction of planetary interiors, radiogenic heating and impacting meteoroids released volatiles (H,O, CO, and other gases containing S and N) to the surface. Here chondrule-rich pre-existing materials were hydrolytically altered to clay and organic matter. Salts, such as gypsum and calcite also formed upon evaporation of these volatiles into the thin atmosphere of these small planetismals. In many instances the mass of these planetismals was insufficientto check the loss of volatilesto space and this earliest phase of soilformation ceased when their internal reserves were exhausted. These most ancient of soils, dating back 4566Ma, are now preserved in deep freeze as carbonaceous chondrites from asteroids beyond the orbit of Mars. Larger planetismals (500-1000 km)also developed a clayey initial soil from the early release of volatiles, but continued heating and differentiation of their interiors led to the formation of volcanoes. The 6rst products of melting within these chondritic planetismals were unusual suites of volatile-rich basaltic and ultramafic rocks: now represented by eucrite and diogenite meteorites. An early phase of clayey and salty soil development on these small planetismals was obscured by volcanic eruptions. Once internal volatiles were exhausted, volcanism also ceased and the surfaces of these atmosphereless volcanic planetismals were remolded by a period of intense meteoroid impacts before 4000 Ma. These early impact-generated soils are now represented by howardite and mesosiderite meteorites. On even larger planetary bodies (1000-5000 km)a greater variety of more differentiated volcanic rocks came to dominate the surface, but their mass was still inadequate to retain an atmosphere. After a n early period of intense, large-body bombardment, soil formation became dominated by micrometeoroid impact as it still is on the Moon.
Larger planetary bodies still (5000-10 000 km) proceeded through the chondritic, carbonaceous and eucritic phases of surfaceevolution toanextended phaseof basaltic and ultrama6c volcanism. There also was sufficient mass to retain a thin atmosphere. This protected the ancient carbonaceous soilssomewhat from destruction by micrometeoric bombardment. There was not, however, a sufficiently thick atmosphere to sustain weathering at the rates found during this early period of degassing and the relict soils of that early time were modified slightly by subsequent photo-oxidation that destroyed their organic matter. Such oxidized, clayey and salty soils are now known on the ancient land surfaces of Mars. On yet larger planetary bodies such as Earth and Venus ( 10000-1 5 000 km)thick atmospheres of volatiles accumulated. A combination of volcanism and surface weathering that still continues created such discrepanciesin rock composition and density that gravitationally driven crustal recycling was initiated. On Earth this takes the form of plate tectonics. Crustal recycling on Venus remains poorly understood. The ancient record of the earliest carbonaceous and clayey soils on both planets was destroyed by subsequent events. On Earth, biological and weathering systems resulted in burial of carbon dioxide as organic matter and limestone, and the burial of sulfate in salt deposits. In this way both these common and chemically aggressive volatileswere kept out of the atmosphere, which came to consist mostly of the inert gas nitrogen. Sulfate salts may be precipitated and recycled on Venus, but there is no evidence on Venus of a comparable mechanism for recycling carbon dioxide. Over millions of years this gas has built to high pressure. Like a greenhouse, it has retained incoming solar radiation to produce high surface temperatures that make it increasingly difficultfor limestone to precipitate or for development of organisms to bring it under control. The preceding selection of ideas on the evolution of soil formation in our Solar System is only one of a number of plausible combinations of currently argued hypotheses.The purpose of this account is less to establish a particular history than to reveal the potential of such studies. The rapidly accumulating data on meteorites and planetary surfaces offerfertilenew ground for a paleopedological point of view.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 16 Earth’s earliest landscapes
It has long been suspected that during the earliest phases of Earth’s geological history its surface environment was very differentfrom the one we enjoy now. Surprisingly, Precambrian paleosols now known are similar in many ways to deeply weathered soils at the land surface today. Nothing in the Precambrian rock record has yet been found comparable with the soils rich in glass and agglutinates on the Moon or the glazed and metamorphosed soils thought to be forming on Venus. Known Precambrian paleosols are neither as oxidized and sulfate-rich as the soils of Mars, nor as carbonaceous and sulfate-rich as carbonaceous chondrites. By this standard, Precambrian paleosols are not as different from modern soils as they could be. Some continuity in soil-forming processes over the past several thousand million years also is evident from information about Earth’s surface environments gleaned from other lines of geological research (Fig. 16.1).Evidence of liquid water is provided by the oldest metamorphosed sedimentary rocks on Earth, dated at 3800Ma, in south-western Greenland near GBthaab (Nfik) (Rosing et al. 1996). These include metamorphosed remnants of conglomerates, quartz sandstones, limestones or dolostones, claystones and the distinctive cherty and iron-rich rocks called banded iron formation. From these early sediments it can be assumed that there were rains, running water and large lakes or oceans. Even at this early time there was extensive weathering by hydrolysis.This kind of weathering can be inferred by comparing the proportion of bases (CaO+MgO + K,O + Na,O) normalized to acommon igneous value of alumina ( 15.6%)for sedimentsof various ages compared with the igneous and other rocks of their presumed source terrains (Fig. 16.1A). Surprisingly, very ancient Precambrian sedimentary rocks show near-modern degrees of weathering even though (a) vascutar land plants did not exist to contribute acidic weathering solutions and (b)these sediments have not been eroded and
redeposited (or ‘recycled’)to the extent of recent sediments (Holland 1984).The main weathering acid then as now probably was carbonic acid. Other evidence for abundant carbon dioxide in the Precambrian atmosphere is provided by carbon isotopic ratios: 13C compared with the more common l*C. In carbonate rocks this ratio has remained surprisingly constant over recorded geological time (Fig. 16.1B). Carbon dioxide may have been present during the Precambrian at higher levels than its present concentration (3.4 x 10-4A or atmospheres), but could not have been as abundant in the atmosphere as it is in some tropical rainforest soils (3.7 x A, or 110times present atmospheric level or PAL: Brook et al. 1983) because this would have led to greater weathering of shales than has been observed in Precambrian sedimentary rocks and surface heating approaching that found on Venus (Kasting 1987).Even at maximal likely concentrations carbon dioxide was a minor component of the atmosphere (Rye et al. 1995). Nitrogen gas (N,) is the main atmospheric component now and the main form of nitrogen released by volcanoes (Holland 1984). Some evidence for its presence in past atmospheres is provided by the near-modern isotopic ratio of I5N to I4N in marine organic matter as old as2500Ma(Fig. 16.1D). Surface temperatures were within the limits of liquid water (0-100OC) to allow formation of sedimentary rocks. Temperature was at times close to zero, as indicated by tillite, varved shales, dropstones and other evidence of glaciations at least as far back as 2 500 Ma (Fig. 16.1E:Hoffmanetal. 1998). Anotherlimittopaleotemperature is provided by pseudomorphs of gypsum as old as 3500Ma(Barleyetal. 1979)becausethismineralinverts to anhydrite at temperatures higher than 58°C. It co-precipitateswith halite at temperatures below 18”C, which is a maximum likely temperature for joint occurrence of these minerals (Fig. 16.1F) as old as 2000 Ma (Schopf 1983).Temperatures below freezing and above boiling also are inimical to life, yet there is evidence of
22 7
228
Chapter 16 years x lo9 4 IHADEAN
3 ARCHEAN
I
I
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base loss during 1 sediment formation mole/kg ASouth Africa
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mean global temperature
-
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range of evaporite minerals
-t
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I PHAN. 1
(0-58OC)L gypsum gypsum t halite (0-18°C)I I
J
columnar stromatolite 2 forms
G
a bundonce
-
of micrabiol fossils L 1
n
J
I
4
3
2
I
present
Figure 16.1 Geologicalindicesof hydrolytic weathering (A),of organic fixation of atmosphericcarbonandnitrogen(B,C.D. G, H) and of moderate temperatures (E. F, G. H) well backinto recorded geologicalhistory. (A) Loss of acid titratable bases (CaO+MgO+K,O +Na,O; in mol kg-’) of average shales and sandstones compared with their likely source regions(datafromHolland 1984);(B andC) ratioof carbonisotopes(’3Cand’2C)relative to Peedee belemnite standard in limestones and dolostones(B) andin fossilkerogen(C) showing spread of data, as well as selected less metamorphosed materials (fromDes Marais etd. 1992:DesMarais 1997);(D)ratiosof nitrogen isotopes( I 5Nand I4N)relative to modern air in fossilkerogen (data from Schopf 1983): (E) qualitative assessment of torrid and glacialphases of Earth’s paleocliate (after Frakes 1979); (F)stratigraphic range of gypsum and of gypsum associatedwith halite, and the likely surface temperatures thattheseimply (after Schopf 1983); (G) number of described forms of columnar stromatolites(afterAwramik1971);(H) abundance of microbial fossil occurrences (after Schopf 1983).
Earth’searliest landscapes organically bound stromatolites (Fig. 16.1G) and of microbial fossils (Fig. 16.1H)as old as 3500Ma (Schopf 1993: Lowe 1994). In summary, weathering by acidic aqueous solutions, organic fixation of atmospheric carbon and nitrogen and moderate temperatures (0-S8OC) extend well back into recorded geologicalhistory. Other aspects of Earth’s surface environments have changed considerably over the sweep of recorded geological history, as revealed by peculiar kinds of rocks abundant in the distant geologicalpast and now seldom formed(Fig.16.2).Someof thesechangescanbeunderstood in terms of a primeval reducing atmosphere becoming oxidized.Other features are compatiblewith the developmentof large granitic continents from an ultramafic primitive crust of the Earth. The two processes may be interrelated. Increased production of oxygen by photosynthetic organisms would have been sustained by nutrients liberated by the weathering of increasingly large continental masses. Dilution of initially high oceanic salinity by continental runoff would also have allowed increased photosynthetic productivity in the sea (Knauth 1998). The consumption of oxygen by reduced iron produced at mid-ocean ridges would have become less important as continental weathering became more extensive (Veizer et al. 1982). Evidence for oxidation of the atmosphere is provided by the distribution of iron, uranium and sulfur in sedimentary rocks. In the present oxygen-rich atmosphere, iron is oxidized to yellow, brown and red minerals. Under reducing conditions, iron-bearing drab-colored minerals are formed. No red beds are known older than about 2000Ma (Fig. 16.2E). A distinctive kind of laminated iron-rich sediment called banded iron formations was abundant from 1700 to 2 500Ma. but also is known as far back as 3800Ma and as recently as 1OOOMa (Fig. 16.2E). One interpretation of these rocks is that they reflect a stage in Earth history under reducing atmospheric conditions when iron was removed from igneous rocks on land or in mid-oceanic ridges in a ferrous state. Once in the iron-rich ocean it was precipitated by locallyoxidized surface waters, by photo-oxidation or by biologically mediated fixation in banded iron formations (Holland1984).Otherinterpretationsinvokemagmatic and hydrothermal activity (Barley et al. 1997). Since 2000 Ma, however,much iron has been oxidized,and by 1000Ma most iron remained in soils and other continental red beds. Similarprogressive oxidation steps may
229
be reflected in changes in the ratio of oxidized iron (Fe,O,) to reduced iron (FeO)in sediments over geological time (Fig. 16.2D), a selective culling of organic carbon isotopic data for samples least affected by metamorphism (Fig. 16.1C) and increased diversity of microfossilsandstromatolites (Fig. 16.1G andH). Uranium is another element with very different solubilities in different oxidation states. Oxidized uranium (LJb+) is found in yellow minerals such as carnotite, which are readily soluble, whereas reduced uranium (U4+) is found in insoluble metallic-looking minerals such as uraninite (UO,). These reduced uranium minerals are readily oxidized in the present atmosphere and are washed into groundwater. When they reach a locally reducing region in underground sandy beds, they may precipitate as a local uranium ore of the roll type (Fig. 16.2E) known as old as latest Carboniferous (295 Ma). Of similar origin are the rich ores of uranium in veins or breccias associated with major unconformities and their paleosols as old as 2000Ma (Kimberley 1992). Some of the most productive uranium ores date back to 1900-3100Ma. These are fluvial conglomerates and sandstones containing abundant grains of unoxidized uraninite that are rounded and concentrated in layers like the placer deposits of modern streams. The persistence of uraninite in near-surface environments may indicate a chemically reducing atmosphere before 1900Ma. Other indications of changes in the atmosphere come from studies of sulfur. The oxidized form of suIfur (S6+) in sulfates such as gypsum is more soluble than the reduced form (S2+)in sulfides such as pyrite. Rounded grains of pyrite also are found within uranium-bearing fluvial placers of 2450-2900Ma and are rare subsequently However, casts of gypsum crystals have been found in sedimentary rocks as old as 3 500Ma. Isotopes of sulfur provide additional clues. Microbes producing sulfidein stagnant parts of the ocean and swamps select for isotopically lighter sulfur (with negative 634Svalues: rich in 32S).However, the evaporation of seawater to produce sulfates shows a very small enrichment of the heavy isotope (positive 634Svalues: rich in 34S). The divergence in isotopic composition of sulfide and sulfate can be traced back to at least 2500Ma, with marked increases in divergence at 2100 and 1OOOMa (Fig. 16.2C). This has been interpreted as evidence for oxygenation of the atmosphere to levels where terrestrial
lo9
years x
ARCHEAN
1
2
20
I
PROTEROZOIC
.
00volume percent cantinental cover 40
U
3
4
I HADEAN I
present
I PHAN. 1
shale
volcanic and volconiclastic
L
1
3 K20/Na20 weight percent
2
I
B
sulfotes and sulfides
20-
5 su I dri=
30-
c -40. sediments
W
,
D - - bonded no r i
formation
relative abundance of sediments and ares
nan-marine red beds uranium uranium in veins below unconformities uranium in redax fronts of Sedimentary sulfide ores
E
ratio of quartz monzonite to quartz diorite
relative abundance of igneous rocks
F
'
1
4
3
2
i
present
Figure 16.2 Indicesof change in crustal and atmospheric composition through geological time. (A) Changingproportions of Sedimentary and other rocks covering continental regions (after Ronov 1964): (B) ratio of weight per cent potash (K,O)to soda (Na,O) in sedimentary and inigneousrocks (afterkgeletal. 1974): (C) sulfur isotope ratios (34Sand 32S)relative to troilite (FeS,) from the Canyon Diablo meteorite, in sulfide and sulfate minerals in sedimentary rocks (range and common values after Can6eld & Teske 1996;Canfield 1998); (D)theratioof weightpercent oxidizediron (Fe,O,) to reduced iron (PeO)in sedimentary rocks and ores(afterRonov&Migdisov1971); (E) abundance of various sedimentary rocks and ores(after Meyer 1985;Barley&Groves 1992):(F)abundanceof differentkindsof igneous rocks (afterEngeletal. 1974).
Earth’s earliest landscapes weathering provided unprecedented amounts of sulfate to the ocean (Canfield 1998)and stimulated large-scale formation of sedimentary sulfide ores (Fig. 16.2E). Alternatively,the isotopic and ore data could reflect lower oceanic temperatures, allowing increased rates of biologicalsulfatereduction(Ohmoto&Felder 1987).Either way, Proterozoic increases in effective marine microbial dissimulatory sulfate reduction would have aided oxidation of oceans and air. Another fundamental change recorded in Precambrian rocks was the development of large granitic continents from an originally ultramafic and basaltic crust. The engine of crustal development operating now is plate tectonics. Recycling of rocks begins below the deep trenches of the world’s oceans, where cold basaltic oceanic crust plunges under its own weight beneath the continental margins. The melting and mixing of this basaltic material with more granitic material of the lower crust produces igneous rocks that are generally richer in silica, alumina, lime and potash, and poorer in iron, magnesia and soda than oceanic basalt (Rudnick 1995). When exposed to weathering, erosion and redistribution by streams and wind, this material is again chemically altered, generally to enrich silica and alumina at the expense of alkalis and alkaline earths. This great rock cycle is not complete because the low-density, silica-richcontinental crust continues to ride well above the high-density, iron- and magnesium-rich oceanic crust and mantle. The rock record on land is biased toward these lighter crustal rocks so it is a surprise to find that continental rocks olderthan about 3000Ma are for the most part mafic volcanic rocks, intrusions and sediments derived from them (Fig. 16.2A and F). These are the Archean greenstone belts: narrow, highly deformed sequences of sedimentary and volcanic rocks wrapped around domes of less mafic kinds of crystalline rocks. The world in those days can be imagined as an Indonesian archipelago of global proportions. Granitic rocks high in potash and silica are found in very old greenstone and granulite terrains, but these become much more common at roughly 2000Ma (Fig. 16.2F). This corresponds in time to a marked jump in the potash to soda ratios of igneous and sedimentary rocks (Fig. 16.2B). Sedimentary sequences deposited on thicker and more stable continents of these granitic rocks have remained in many cases undeformed. Quartz sandstone, limestone and dolostone of these stable conti-
2 31
nents are much more common than earlier in the rock record (Fig. 16.2A). The appearance of these platform sequences in any particular region may be abrupt because major unconformities and structures commonly separate Proterozoic sedimentary basins from Archean granulites and greenstone belts. With more widespread radiometric dating of these different kinds of rocks it is now known that the advent of platform sequences on the present continents occurred at different times in different regions over a time span from 3000 to 1000Ma. Island arcs today can be ranked into evolutionary stages from small andesitic arcs l i e the Marianas to larger and more complex ones like Indonesia. Volcanic arcs l i e Japan approach small continents in size and complexity. These ultimately will be swept up into the more massive continental mass of Asia as have other arcs before them. Although the formation of continents is an unfinished process, the evolution of the earliest ones heralded new kinds of environments. With these new kinds of landscapes came new kinds of soils whose fossil record may provide additional evidence for the nature and pace of early crustal evolution.
Oxygenation of the Earth’s atmosphere The idea that the atmosphere of the early Earth was reducing and became oxygenated through photosynthetic and other activities of microbes is widely accepted (Schopf & Klein 1992). But is it correct? Could red beds older than 2000 Ma have been chemically reduced by metamorphic or hydrothermal alteration? Could there have been a very arid climate or exceptionallyhigh rates of sedimentation to preserve detrital grains of uraninite in river deposits?These and other objections have been raised by those who believe that there was appreciable atmospheric oxygen back to the beginning of the Precambrian rockrecord (Ohmoto 1996). As a record of the interaction between the atmosphere and surface sediments and rocks, paleosols are coming to play a more prominent role in the examination of this long-standing geological problem (Rye & Holland 1998).This is not to say that interpretation of such ancient paleosols is without problems. Several difficulties arise in considering the Iiiely oxidation state of Precambrian paleosols. For example, paleosols could have been chemically reduced by metamorphism or clay diagenesis. Their oxidation state may not represent
232
Chapter 16
atmospheric conditions because they were too clayey or impermeable for air circulation, too rich in organic matter or other reducing materials, or too waterlogged (Ohmoto 1996, 1997). Fortunately, there are Precambrian paleosols for which these concerns can be set aside.
Pre-Huronian (2450Ma) paleosols of Ontario, Canada Among the most thoroughly studied paleosols are profiles developed on Archean granite and greenstone at the unconformity below 2450Ma fluvial sediments of the Huronian Supergroup, exposed mainly in mine workings north of Lake Huron, Ontario, Canada (Mossman&Farrow1992;Schmitt 1999).Oneof these profiles developed on greenstone has been called the
Denison paIeosoI after a nearby mine (Fig. 16.3). It consists of a green sericitic rock, lighter in color and finer grained than its parent greenstone. Its clay content (now sericite)was greatest near the surface, moderately high to a considerable depth and negligible within the parent greenstone. The clayey surface horizon can be interpreted as an At horizon. The less clayey subsurface horizon is a Cr horizon showing relict crystalline fabric, l i e saprolites of modern deeply weathered soils (see Fig.2.1). Formation of clay (now sericite) from actinolite, feldspar and, to a lesser extent, from chlorite, appears to have been the major soil-forming process.There is an increase in the amount of quartz at the expense of the same minerals. Alumina and silica vary in abundance with clay and quartz content, as they would in a modern soil. The leaching of magnesia from the profile is much
Figure 16.3 Diagrammaticsketchesof Precambrian paleosolson mafic and felsic
@
ag{z>
PYrrhotlte
mbpk$acsky
gpJtg'om- asandstone
~ a m p h i b o f , t e ~ ; 7 g r e e n s t ocorestones ne~
parent materials at major unconformities before (B) and after (A)the early Proterozoic (2000Ma)rise of atmosphericoxygen. (A) Sheigra and Stacapedotypesbeneath the 810 MaTorridonianSupergroupnear Sheigra, Scotland (B) Pronto andDenisonpedotypes beneath the 2450 Ma Huronian Supergroup near Elliot Lake, Ontario, Canada (datafrom Mossman&Farrow1992;RetaUack& Mindszenty 1994).
Earth's earliest landscapes aswould beexpectedina humidclimate.Thesamecou1d be said of lime, although there was very little of this in the parent material of the soil. There was even less soda in the parent material and it varies little within the profile. Potash appears to have accumulated. Total iron is strongly depleted from the profile. There is a little more oxidized iron (Pe3+)than reduced iron (Fez') near the surface of the profile, but below that iron is mainly reduced. Another paleosol along the same pre-Huronian unconformity, but developed on pink alkali granite, has been called the Pronto pedotype (Figs 16.3 and 16.4). This has a clayey, light green surface (At) and a deep, light-colored subsurface (Cr) horizon. The main soilforming process appears to have been the formation of clay (now sericite) from chlorite and feldspar. Chemical variation in alumina, silica and magnesia is similar to that observed in the Denison pedotype. The amount of potash is almost constant with slight subsurface accumulation and slight surficial depletion. There are low values of lime throughout theProntopedotype and soda shows some surficial depletion like that seen in soils of humidclimates. A notable feature of theProntopaleoso1 is the suficial enrichment of total iron, largely in the
Figure 16.4 Early Proterozoic (2450Ma), Pronto pedotypepaleosol exposed in glacial pavement north-eastof Quirk Lake, near Elliot Lake, Ontario,Canada.The paleosol dips vertically and strikes toward the back left of the photograph. Overlyingconglomerate is to the right,clayey surface of the paleosol is erodedas a gutter, and corestonesof saprolite are to the left. Hammer isfor scale.
23 3
oxidized state in the upper part of the profile, although not to the extent that the paleosol is visibly red or ferruginized. These two paleosols on the same ancient land surface are both drab colored, and include pyrite and uranium minerals. They are surprisingly different in their degree of oxidation and amount of total iron. Just as modern soils are known to be products of many interacting factors, the interpretation of Precambrian paleosols must also take into account complications such as metamorphism, clay diagenesis, original permeability, organic matter content, paleodrainage and parent material. Metamorphic environments are mostly reducing, so it is possible that the lack of red paleosols older than about 2000Ma is because many rocks of this great age have been strongly metamorphosed or hydrothermally altered (Ohmoto 1996). However, brick red paleosols have been preserved in the Upper Silurian Bloomsburg Formation of Pennsylvania, USA, despite its metamorphism to lower greenschist facies (Retallack 19 8 5 ) .This is a higher grade of metamorphism than sufferedby the pre-Huronian paleosols (Mossman & Farrow 1992). Furthermore, the preservation of pyrite and uranium minerals is evidence against pervasive metamorphic or
234
Chapter 16
hydrothermal alteration of originally oxidized rocks (Schmitt 1999). Changes in the composition of clays during burial may further obscure the original nature of paleosols. Particularly problematic is possible diagenetic enrichment of potash from the dissolution of muscovite and potash feldspar in deep groundwater before compactional dewatering, as described by Hower et aJ. (1976), although this classic study of this process has been challenged (Mora et al. 1998; Bloch et al. 1998). Other explanations for potash enrichment include less intense leaching of potash before the advent of vascular land plants (Weaver 1989) and evaporative enrichment in a dry climate (Wright & Robinson 1988). This last explanation is unlikely for preHuronian paleosols considering their abundant clay, depth of weathering, lack of evaporite minerals or crystal casts and depletion of soda in both profiles. The inconsistent behavior of potash, enriched in the Denison paleosol and depleted in the Pronto paleosol, undermines the likelihood of pervasive potash metasomatism and its possible alteration of the oxidation state of both paleosols. Furthermore, there is no clear relationship between potash and the much more abundant iron within the profile (Denison)that would have been most affected. Potassium metasomatism has been found to have affected Late Ordovician paleosols, but it did not discernibly altertheirredcolororoxidizediron(Feakes& Retallack 1988).In this case modeling studies indicate that potash metasomatism is unlikely to have greatly affected the oxidation state of these paleosols (Schmitt 1999). Drab-colored soils with iron-bearing minerals containing reduced or ferrous iron are formed today in anaerobic waterlogged parts of the landscape (Ohmoto 1996, 1997). This explanation for drab Precambrian paleosols has general appeal because waterlogged soils are more abundant than oxidized ones in many sedimentary environments and so would more probably be preserved in the rock record. Poor drainage seems unlikely for the pre-Huronian paleosols, however. Their depth and degree of developmentis well in excessof that found in modern waterlogged soils.In the Denison pedotype, clay skins deep in the profile (6 m down in the compacted paleosol) are evidence of very low water table. In both Pronto andDenison pedotypes,corestone weather-
ing indicates good drainage. It is possible that the clay and soil structure of both Denison and Pronto paleosols formed under freely drained conditions, but were then reduced under waterlogged conditions shortly before burial by overlying fluvial deposits. If this were the case, however, it should have affected these paleosols in a more similar way. Another possibilityis textural controlof the oxidation state of thesepaleosols. Modern clayey soils are less permeable to air and water than sandy ones (Bland & Rolls 199 8). In the case of the pre-Huronian paleosols, however$the fine-grained parent material (greenstone) is weathered to a greater depth than the coarse-grained parent material (granite): the reverse of the usual relationship in modern soils. The Denison pedotype may have been more tectonically sheared than Pronto paleosols, but parent material grain size was not of overriding importance. Nor was the texture of the resulting soil because the surfaces of Denison and Pronto paleosols do not differ greatly in the amount of clay. Even in the present oxidizing atmosphere, complexes of clay and stable organic matter in some well-drained Vertisols can impart a drab color. Neither of these paleosols is as dark, organic or slickensided as such modern soils. Low amounts of organic carbon with the isotopically light composition characteristic of life have been detected in a Denison paleosol (Mossman & Farrow 1992).There probably was microbial life in these paleosols, but even forested modern soils above such thick saprolites are not chemically reduced by biological activity (seeFig. 2.4). A final consideration is the amount of iron present in the parent material. Under present oxygenic conditions most iron released by weathering is oxidized.Thus, the reddest soils form on iron-rich parent materials such as basalt and greenstone. Curiously, the situation is reversed with the pre-Huronian and other Precambrian paleosols. Those developed on iron-rich rocks, such as the Denison paleosol, are less oxidized and have lost iron compared with those developed on iron-poor rocks, such as the Pronto paleosol. Such a relationship is easiest to understand under conditions of a very weakly oxidizing atmosphere. Generally speaking, soil formation can be considered as a competition between carbonic acidic hydrolysis and oxidation by soil oxygen. Under atmospheres very low in oxygen, hydrolysis would release
Earth’searliest landscapes iron from iron-rich parent materials at a rate greater than it could be oxidized, so that it would be lost from the profile. On the other hand, even low amounts of oxygen may be sufficientto oxidize small amounts of iron slowly released by hydrolysis of iron-poor parent materials, with the result that iron is fixed within the profile (Holland 1984). Thus, none of the known circumstances for producing chemically reduced modern soils can be applied easily to these Precambrian paleosols. Differential oxidation of different parent materials under low concentrations of oxygen seems most likely, and has been confirmedby studies of a variety of other paleosols older than 2200Ma (Rye & Holland 1998). Archean and Early Proterozoic drab paleosols are sufficiently distinctive that I have proposed (Retallack 1986c)calling them Green Clays. They are deeply weathered yet fall within the weakly developed soil order Inceptisol because they lackdiagnostic horizons of other soilorders (seeFigs 5.3 and 5.4). Green Clay paleosols appear to be a record of conditions much less oxygenated than the present atmosphere. Pre-Torridonian (8 10 Ma) paleosols of north-west Scotland Another group of paleosols is developed above Archean biotite gneiss, amphibolite and microcline pegma-
Figure 16.5 Sheigrapaleosol(bleached and reddenedzoneabout1 m thick to right)under Torridonian(810 Ma) alluvialfan deposits, and Staca paleosol at the same unconformity but on arnphibolite(left-handside) near the hamlet of Sheigra, north-westScotIand.The white tape extending downfrom the unconformity,left of center,is 2 mlong (from Retallack 1986c; reprintedwith permission fromBlackwell ScientificPublications).
235
tite at the unconformable contact with Proterozoic (8 10Ma) alluvial fan deposits of theTorridonian Supergroup (Figs 16.3, 16.5 and 16.6) in north-western Scotland (Williams 1968: Retallack & Mindszenty 1994).These former soils mantled low hills and a piedmont flanking a fault mountain range of considerable topographic relief (see Figs 11.2 and 11.3) now rifted and drifted from Scotland to Greenland (Williams 1969).The Staca pedotype is developed on amphibolite and the Sheigra pedotype on biotite gneiss along this same geological unconformity (Fig. 16.5). The surface (At) horizons of both paleosols are stained purple-red. Underlying this is a zone of mottled and light-colored rock including large corestones of little-altered parent material. Both profiles are more clayey toward the surface, where there are crude, platy peds, but there is a relict crystalline texture throughout. Quartz and microcline persist in the weathered part of the profile, but biotite, plagioclase and hornblende of the parent rocks have been extensively altered to clay, Clay formation is reflected in the decreased silica and increased alumina toward the surface. Lime and soda are depleted toward the surface, but potash increased toward the surface in both paleosols. Total iron is enriched in the surface horizons compared with parent material. Ferric iron (Fe,O,) is well in excess of ferrous iron (FeO)throughout both paleosols,incontrast to the reduced state of iron in both parent materials.
2 36
Chapter 16 molecular weathering ratios
field observations type Sheigra clay
b2QCaO+MeQ K,O AI,O,
point count data percent percent size composition
"lor
I
m
+Na,O+K,O
1
50
0
1-
artz
io
.
.
1 2 3
?327
dark gray (N4)
2
i 5
1
I
dark reddish gray ( 5 R W weathering rind
R328
greenish gray (SGY6/1)
3 t- .
n str
4 y324
5
6
$1
, , ,
.(
If+l biotite gneiss @ amphibolite pegmatite
-
dark gray
'
chloritized I biotite salinization calcificz hornblende &dspar none none conglomerate sandstone claystone
-T-
B
imbrication
- bedding
+.
ripple mark
'
.
1
base lo$ leaching gleization clayeyr moderate moderate moderate none
red color
corestone
brown color
corestone remnant
fermginized
carbonate
Figure 16.6 Columnar section (measured in field),petrographiccomposition (point counting by G.S. Smith) andmolecular weatheringratios (fromanalyses of Williams 1968)of theSheigrapaleosolof Fig. 16.5(fromRetallack&Mindszenty1994;with permission from the Society for Sedimentary Geology).
Earth's earliest landscapes Clay formation and ferruginization appear to have been the principal processes involved in the formation of both paleosols. Dolomite nodules deep within the Staca paleosol are evidence that rainfall was within the subhumid range (550-850mm mean annual precipitation). Although these paleosols may represent a long time of formation at a major unconformity, they are not profoundly weathered or differentiated into horizons and may be regarded as Inceptisols (Retallack & Mindszenty 1994). Other features of these paleosols provide indications of an oxygenated atmosphere. These paleosols are not metamorphosed, and have retained smectite and poorly crystalline illite. as in paleosols buried no more than 2km (Retallack & Mindszenty 1994). Nor were they waterlogged: the bending of weather-resistant pegmatiteveins at the surface of the paleosol has been interpreted as evidence of Precambrian soil creep on a moderate slope (Williams 1968). Pervasive oxidation and great depth of weathering (now 6 m in places) are evidence that they were well drained. Furthermore, the paleomagnetic orientation of the hematite pigment is in a direction expected for Precambrian, rather than subsequent geological time (Williams & Schmidt 1997). The parent materials of both paleosols are coarse grained, and coarse crystalline textures persist through both profiles. Isotopically light organic carbon in the Sheigrapaleosolisevidenceof microbiallife,but this has not caused chemical reduction of the profile (Retallack &Mindszenty1994).Finally,theamphibolite hada high iron content and thus a high demand for oxygen, closer to that of theDenisonpaleoso1than the Prontopaleosol. Yet both Staca and Sheigra paleosols are visibly ferruginked, reflecting a significant advance in the oxygen content of the atmosphere at 810Ma compared with 2450Ma. Calculating atmospheric oxidation from paleosols There are so many Precambrian paleosols pertinent to Precambrian atmospheric composition that the problem is to decide which of these are the most reliable paleobarometers (Ohmoto 1996, 1997; Rye & Holland 1998). Profiles similar to Green Clays of the preHuronian unconformity are known back to at least 3500Ma (Buick et al. 1995). The youngest Green Clay and oldest appreciably oxidized paleosols are of
23 7
interest in signaling turning points in the oxidation of the atmosphere. As far as I am aware, the youngest well-known Green Clay profile is the 2200Ma Hokkalampi paleosol developed on granite and metasediments in the Koli-Kaltimo area of eastern Finland (Marmo 1992). The oldest well-studied paleosols that are visibly ferruginized are the 2070-1920 Ma Drakenstein and Wolhaarkop paleosols found along the unconformity between the Olifantshoek and Transvaal Groups in Griqualand West, South Africa (Rye & Holland 1998). Ferruginized paleosols are more common in younger Precambrian rocks. Although much has been made of paleosols associated with the 2200Ma Hekpoort Basalt of Transvaal, South Africa (Holland & Beukes 1990; Wiggering & Beukes 1990), the profiles described so far at Waterval Onder and Daspoort Tunnel have chemical discontinuities and metamorphic alteration that renders them unreliable paleobarometers (Maynard 1992;Retallack & Krinsley 1993). Even without them, however, there remains evidence of amarkedrise in atmospheric oxygenation at about 2 100Ma. Estimates of the amount of oxygenation can be based on chemical data from paleosols provided one is prepared to make a number of simplifying assumptions (Holland 1984). One needs to assume that there was limited diagenetic or metamorphic alteration, negligible waterlogging, circulationof airto thesoil andlittleeffect of microbes: all of which are open to debate (Ohmoto 1996, 1997). Also assumed are a limited range of weathering reactions common in modern soils so that the amount of oxygen needed to fully oxidize all oxidizable materials and the amount of carbon dioxideneeded to hydrolyzesusceptibleminerals can be calculated from molar proportions of elemental constituents of the parent material (see Table 4.5). In this way the ratio of oxygen demand to carbon dioxide demand can be estimated. Because an atmosphere with a ratio of oxygen to carbon dioxide in excess of the demand of the parent rock would result in iron fixation and one with a lower ratio in iron loss, the ratios for former atmospheres can be bracketed by comparing the distribution of iron in paleosols formed in chemically differentparent materials. For the Pronto pedotype (Fig. 16.4), for example, the ratio of oxygen demand to carbon dioxidedemand of its parent material is about 0.02 5 mol kg-'. Similar calculations based on a number of paleosols (Fig. 16.7)
238
Chapter 16 u)
h
2
0.41 glaciiions
I
AAA
'
AMAAA
4
r
a,
c
2 0.3.
z
CI
Kl
8 0.2-
v
C
m
-D C
2 m
C
a,
2
0
.. - -
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'
Hokkalampi ' - - - '' Mt Roe Pronto h)aWolhaarkoP 8 8Flin Flon 0- - - - - - ville ~~~i~ Drakenstein Sturgeon Falls
3
2 1 billions of years ago
McGillivray Brook -
Figure 16.7 Reconstructedhistory of atmosphericoxygen abundance showing constraintsfrom selected Precambrianpaleosols (opencircles,fromRetallack&Mindszenty1994;Rye&Holland1998) andfromselectedPhanerozoicpedogenicgoethites(Elled circles,fromYapp 1996).The Phanerozoic curve (continuousline)is from the computer model of Berner et aJ. (2000).and the late Precambriancurve (dashedline)isinferred from carbonisotopicdata(Karhu&Holland1996; Hoffmanet al. 1998).Thetriangles representtimes of glacial sedimentation.
have shown that the value of this ratio at the transition between iron loss and iron gain in paleosols older than 2100Ma was about 0.002molkg-' (Rye & Holland 1998).Using physical constants at an assumed moderate temperature (1SOC), this critical value for the ratio can be converted to aratio of partial pressures of oxygen to carbon dioxide of about 1.3 k 0.5 in the atmosphere before 2 100Ma, as opposed to about 600 in the present atmosphere. The partial pressure of oxygen can be determined if one is prepared to make additional assumptions about the partial pressure of carbon dioxide. Considering that acidic weathering of Precambrian sediments is comparable with that seen in modern sediments (Fig. 16.1A), but without the benefit of vascular plants, carbon dioxide must have been more abundant than it is today A), but probably not more than the highest (3.4 x value found in some tropical forested soils (3.7 x A or 11OPAL: Brooketal. 1983).Thereareotherreasons also for assuming a higher concentration of atmospheric carbon dioxide in the distant past, such as the greenhouse effect needed to counteract the low luminosity of the young Sun (Caldeira & Kasting 1992). Such high
carbon dioxide levels are not needed to maintain a greenhouse in view of probably high levels of methane, which are indicated by very light carbon isotopic composition of organic matter at about 2800Ma and before then (Fig. 16.1: Pavlov et al. 1999). Indeed, Rye et al. (1995) have argued that carbon dioxide could not have been more abundant than 1O-1.7 A (60PAL) because at greater concentrations than this iron carbonate (siderite)would beexpectedrather thantheiron silicates (iron-rich chlorite) found in paleosols older than 2 100Ma. This is a risky assumption given the propensity of carbonates for metamorphic decarbonation, and absence of evidence is not always evidence of absence. Nevertheless, with this additional constraint, atmospheric 0, levels before 2100Ma would have been less than 0.002A (0.01PAL), and after 2100Ma 0, would have risen to only0.3 A (0.14 PAL: Rye &Holland 1998). Such a large and abrupt shift is not unusual compared with modeled shifts in the Phanerozoic (Berner et al. 2000) and is compatible with a large observed carbon isotopic shift in organic matter at about 2100Ma (Des Maraiset al. 1992: Karhu &Holland 1996).Documenting such fluctuations from paleosols, and particularly
Earth’s earliest landscapes the late Precambrian history of atmospheric oxidation, remains a challenge for the future. Such calculations have helped to reconcile those who thought that Precambrian atmospheres were anoxic and those who considered them as well oxygenated as today’s. Oxygen appears to have been present in amounts much less than at present, but was far from negligible. It was present in amounts comparable with that of carbon dioxide in the atmosphere today. Both gases were probably minor components of the atmosphere. Presumably nitrogen was a prominent component then as it is now (Brandes et al. 1998), and there were substantial amounts of methane, ammonia and water vapor as well as carbon dioxide in this primeval greenhouse.
Differentiation of continental crust The view that the development of continental crust begins with subduction and partial melting of oceanic crust under oceanic island arcs has emerged only in the past three decades, and its application to crustal differentiation during Archean times is even more recent (de Wit et d.1992: Calvert et al. 1995: Rudnick 1995). By this model, highly deformed greenstone belts were oceanicbasins and the granitic and gneissicdomes were cores of volcanic arcs. The developmentof continents is thus seen as a process of collision and amalgamation of island arc systems.An alternative view of crustal evolution emphasizes the unique nature of many Archean rocks. By this view, a long period of vertical magmatic differentiation or impact melting produced a mobile early crust. Only after this crust thickenedandstabilized did subduction of oceanic crust and other phenomena of plate tectonics begin (Goodwin 1981: Bailey 1999). Arguments concerning the earliest phases of tectonic activity have hinged largely on interpretations of how the various chemically distinct kinds of igneous rocks may have formed and to a lesser extent on evidence from ancient sedimentary environments. Paleosols also may provide pertinent clues. The transition from a world like a global Indonesian archipelago or Icelandic rifted plateau to one with a few places like Death Valley of California, USA, or the Andean altiplano would have had profound consequences for soil formation. According to the view that a thin,
239
primitive crust was downwarped to form greenstone belts and intruded in actively uplifted regions by large plutons, there should be few, if any, strongly developed paleosols in Archean terranes. By the plate tectonic view, the earliest island arcs, once formed,would remain as buoyant crustal masses after volcanic activity ceased. Thus, some well-developed soils might be expected even on Archean rn&c rocks. By both models, the amalgamation of continents during late Archean and early Proterozoic time should have resulted in some areas falling within the rain shadow of mountain ranges or rift valley scarps, and calcareous desert soils (Aridisols) would have formed on granitic plutons and alluvium of stable continental platforms. Some of these soils may have appeared like modern duricrusts such as bauxites, laterites, silcretes and other deeply weathered materials formed on stable continental land surfaces. Extensive mountain ranges and continental regions would have attracted glaciers and promoted the development of periglacial features in soils. Thus, calcareous, very deeply weathered and periglacial paleosols should appear in the rock record at about the same time as the earliest continents. With these general ideas in mind, we will consider the geologicalrecordof these kinds of paleosols and products of deep weathering. Bauxites These alumina-rich rocks are thought to form by longterm weathering of silicate minerals at moderate pH in humid, non-seasonal climates and on stable land surfaces. They also form as a weathering residuum or deposit on depressions in limestone. Numerous examples of these karst bauxites are known ranging back in age to Precambrian (Bardossy 1982).The most ancient examples are highly aluminous rocks associated with 2300Ma karst landscapes on the Malmani Dolomite of South Africa (Button &Tyler 198 1). Bauxites may be even older than this, judging from records of highly aluminous schists containing kyanite, corundum and pyrophyllite. Bauxites have 45 wt % or more alumina, well above the average for igneous rocksof 15,9wt%orforshalesof 14.6wt%(GarrelsL? McKenzie 1971). Alumina enrichment can be caused by hydrothermal alteration as well as by weathering, so that care must be taken in interpreting aluminous rocks
240
Chapter 16
in highly deformed and very ancient terranes. Corundum ores in highly metamorphosed rocks of the Aldan Shieldof Russia as old as 3 500 Ma have been interpreted as former bauxites (Serdyuchenko 1968). Similar explanations also have been advanced for corundum ores of about 2 700 Ma in Zimbabwe and Western Australia, although both now have been interpreted as hydrothermally altered rocks around volcanic hot springs (Schreyer et al. 198 1;Martyn & Johnson 1986). Laterites and Oxisols These strongly ferruginized soil materials also are found on very stable land surfaces. Oxisols are highly weathered, red, clayey soils whereas pedogenic laterites (or plinthites of Soil Survey Staff 1975) are horizons of soils that harden irreversibly on contact with air. Evidenceof hard, ripped-up clasts in associated sediments must be sought todistinguishvery ancient Oxisols fromlaterites. It is difficult to make such distinctions in highly metamorphosed amphibolite to granulite facies, Oxisol-lie rocks, suchas the2600-2900Makhondalitesof peninsularhdia (Dashetal. 1987).Theoldestclearexamples of Oxisols are dated to at least 1660Ma, below the Thelon Sandstone in the Thelon Basin, North-west Territories, Canada (Chiarenzelliet al. 1983)and some 1500Ma beneath the Athabaska Group in the Athabaska Basin of northern Saskatchewan, Canada (Holland 1984).Kaoliniteisanindicatorof thoroughweathering and has persisted in these profiles despite subsequent diagenetic alteration. These paleosols have a surficial hematite-richzone (upto 24 m thick) overlying a transitional horizon (also 24 m thick) and then a chlorite-rich zone (upto 30 m thick) over fresh bedrock. These weathering zones are similar in thickness to the lateritic, mottled and pallid zones found beneath modern deep weathering profiles (see Fig. 6.7). No laterite has been found, although it could have been eroded. These ancient profiles may be the oldest known Oxisols, but cannot yet be accepted as the oldest laterites. Pisolitic laterites dated at 1920-2200Ma have been found at the extensive unconformity between the Olifantshoek and Transvaal Groupsin Griqualand West, South Africa (Gutzmer & Beukes 1998; Rye & Holland 1998). Although laterites are best known from the Cenozoic (Paton et al. 1995; Ollier & Pain 1996). they have a long fossil record (Bardossy& Aleva 1990).
Calcretes A variety of calcareous materials fall within the general term calcrete. Nodular and micritic alteration of exposed marine limestone and dolostone have a long fossil record (Grotzinger 1986; Horodyski & Knauth 1994) perhaps as old as 3400Ma in the Strelley Pool Chert of the Warrawoona Supergroup in Western Australia (Lowe 1983). Groundwater calcretes have only been traced as far back as the Upper Ordovician (445Ma) Juniata Formation of Pennsylvania (Feakes& Retallack 1988) and groundwater dolocretes back to the Upper Triassic (225Ma) Gres de Chaunoy of France (Spot1& Wright 1992). Of greater paleoclimatic significance is the nodular calcrete formed in weakly calcareous parent materials found in dry regions of large land masses. Such nodular calcrete has a long and copious fossil record (Allen 1986a; Kalliokoski 1986), extending back to 1900Ma in alluvial sequences of the Mara Formation in Bathurst Inlet, North-west Territories, Canada (Campbell & Cecile 1981) and to 2600Ma in ultramafic bedrock at Schagen, Transvaal, South Africa (Martini 1994; Watanabe et al. 1998). Silcretes There is no shortage of very ancient cherts associated with paleosols. The critical question in the present context is how much of these rocks formed as continental silcretes, by cementation with solutions from the deep weathering of silicate minerals in soils (Ollier & Pain 1996).Cherty rocks also form where silica is mobilized under highly alkaline conditions of evaporitic and desert environments (Summerlield 1983; Chadwick et al. 1989),andwherehotwaterfromvolcanic hotsprings cools as it flows away from the vent (Rice et al. 1995). These latter explanations both seem more likely than deep weathering for cherty paleosols in the 3400Ma Strelley Pool Chert of the Warrawoona Supergroup near Marble Bar, Western Australia (Lowe 1983 ) and in the 3300-3500Ma Upper Onvenvacht and Figtree Groups near Barberton in South Africa (Lowe & Byerly 1999). The oldest identifiable continental silcrete is much younger: at the basal unconformity beneath the 1760Ma Pitz Formation in the Thelon Basin of the
Earth’s earliest landscapes North-west Territories of Canada (Ross & Chiarenzelli 1984). This silcrete truncates a deeply weathered paleosol formed during a long period of humid weathering before deposition of the Pitz Formation. Many of the silcretes at and above the unconformity include quartz overgrowths and vug fillings of length-slow chalcedony like those found in aridland silcrete. They are, in addition, associated with evaporitic pseudomorphs and eolian dune sandstones. This ancient sandy desert can be traced over some 600km of the Canadian Shield. Karst This irregular dissolutional and corrosional topography characteristic of limestones and dolostones is a very ancient landform (James & Choquette 1988), now recognized as far back as 2700Ma in the Steeprock Group west of Thunder Bay, Ontario, Canada (Schau & Henderson 1983). The soil material filling this paleokarst includes local concentrations of hematite, manganese and chert, like those known from another very ancient karst on the 2300Ma Malmani Dolomite of Transvaal, South Africa (Button & Tyler 1981). Periglacial soils
A variety of distinctive soil features are found in soils of glaciated continents. Fossil ice wedges filled with massive or horizontally bedded sand are found in periglacial regions of severe cold but moderate humidity (seeTable 9.2). Such structures have been found among glacial sediments as oldas 2450Main theRamsay LakeFormation near Espanola. Ontario, Canada (Young & Long 1976). Sand wedges have vertical layering formed by periodic opening and filling, and indicate colder and more arid conditions. Sand wedges have been found in association with glacial deposits of Late Proterozoic age (1000-600Ma) at a number of localities in South Australia (seeFig. 4.3), Scotland and Norway (Williams 1986). This evidence of periglacial paleosols confirms that of associated tillites and varved shales. Furthermore, these early glaciations extended well into low latitudes, as indicated by paleomagnetic studies of paleosols and associated glacigene sediments (Hoffman etal. 1998;Schmidt&Williams 1999).Somecontinents
241
were glaciated as early as 2450Ma. By lOOOMa, there were even colder and more arid glacial climates, perhaps because continents were larger and some places were more isolated from the ocean. Calibrating continental emergence frompaleosols
In summary, deep weathering under humid climate on stable land surfaces extends back to at least 3 500Ma. Silcretes formed on margins of hypersaline lagoons and volcanic hot springs, and calcretes formed by subaerial modification of marine carbonates also may be as old as 3500Ma. These locally dry, coastal environments may have been downwind of large volcanic edfices. Karst can be traced as far back as 2700Ma, and should be older because it forms on limestone anddolostone under the same conditions as inferred for very ancient Green Clay paleosols. The amalgamation of continents of some size and elevation is needed to create pedogenic caliche as old as 2600Ma and desert silcrete as old as 1900Ma. Such continental masses could also have supported glaciers and humid periglacial soils as old as 2450 Ma. Arid continental glacial climates are indicated by ice wedges in periglacial paleosols of 1000-500 Ma. The cooling and drying of continental interiors may have been related to development of rain shadows behind marginal mountain ranges. It also could have been related to changing atmospheric composition from warmer early Precambrian greenhouse conditions to a more oxygenated late Precambrian atmosphere. In view of these varied lines of evidence from paleosols, a likely model of crustal evolution is one in which buoyant nuclei of protocontinental crust were areas of stability back beyond the beginning of the sedimentary record at 3500Ma. This is more compatible with the view that continental crust was generated by subduction and other processes associated with plate tectonics than the view that it was generated by rifting or vertical tectonic activity associated with a thin, primitive crust. The deep weathering of paleosols also supports isotopic evidence for substantial amounts of early continental crust (Bowring&Housh1995). Terrestrial weathering may have played a role in the early chemical differentiation of the Earth’s surface. Archean paleosols are surprisingly thick and deeply weathered, and represent a massive titration of rocks
242
Chapter 16
with an acidic primeval greenhouse (Buick et al. 1995; Rye & Holland 1998). Their weathering created soil and sediment with a composition unlike that of oceanic crust (SIMA) and more like that of continental crust (SIAL): low density, hydrated, siliceous, aluminous (Maynard 1992),and light rare earth enriched (including favoring neodymium over samarium: MacFarlane et al. 1994).The greater density and more mafic chemical composition of fresh basaltic volcanic rocks may have initiated density-driven tectonic accommodation, and ultimately gravitationally driven subduction of old oceanic crust. This deliberately simplified model is worthy of exploration, although it is clear that recycling through magmatic melting and mixing is needed to explain current crustal and mantle compositional inhomogeneity (Bowring & Housh 1995; Rudnick 1995).
Precambrianscenery It is natural to cling to familiar images in trying to reconstruct Precambrian landscapes. Commonly they are depicted in popular books on Earth history as sulfurous and smoky volcanic landscapes. In active volcanic regions of the North Island of New Zealand, of the Big Island of Hawaii or of the interior plateau of Iceland, it is easy to imagine oneself transported back in time to the earliest Earth. But how typical were these environments at that time? Highly aluminous, clayey paleosols as old as 3000Ma are an indication that in some places nothing nearly so dramatic as a volcanic eruption happened in many hundreds of thousands, perhaps millions, of years. Another source of inspiration for Precambrian landscapes of theimagination are desert regions.The sandur plains of Iceland, the alluvial fans of Arizona, USA, and the channel country of south-western Queensland, Australia, are seen as possible models of unvegetated Precambrian fluvial landscapes. Desert dunes and playa lakes of Death Valley in California, USA, and of the Sahara Desert in Africa may be used to complete a mental image of Precambrian wastelands. Undoubtedly there were some braided streams and alluvial fans. There is evidence of evaporitic coastal lagoons back at least to 3500Ma. Yet there were equally ancient clayey landscapes formed under a regime of deep, humid weather-
ing. Extensive desert dunes, calcareous soils and silcretes didnot appear untilsome 1800Ma. In recent years images of the surface of the Moon, Mars and Venus have provided further sources of inspiration for reconstructing Precambrian landscapes. Yet none of these bear especially close comparison with Precambrian paleosols. What little is now known about Precambrian paleosols constitutes a more promising line of evidence for reconstructing soils and landscapes of the past than these other analogs (Fig. 16.8; see also Fig. 9.13). By 3000 Ma areas of active volcanoes, high mountains, braided stream channels, tidal flats and playas would have seemed similar to equivalent modern landscapes. Also similar to today were salty scum soilson limestone, dolostone and evaporites around dry coastal lagoons downwind of large volcanic edifices.In contrast, rolling terrains indicated by Green Clay paleosols would have seemedunusual by modern standards. One can imagine rounded hillocks of deeply weathered gneiss or granite traversed by low crumbling walls of more weatherresistant quartz veins. Massive spines and walls of volcanic intrusions may have towered over undulating clayey terrain of deeply weathered ash and lava. Without the benefit of multicellular vegetation some of these landscapes would have appeared barren and desolate and yet also lacked the harsh rocky outlines and dunes of modern desert landscapes. These curious extinct landscapes and soils may have occupied large areas of tectonicallystableregions.Their deep weathering can be attributed to a humid, maritime climate and high temperature assured by the greenhouse effect of elevated levels of atmospheric water vapor, methane and carbon dioxide. Leaching of silica, alkalis, alkaline earths and iron from these soils would have encouraged the development of dolomites,bandedironformations andcherts in marine and near-marine sedimentary environments. Photosynthetic microbes had appeared by this time, but oxygenation of the atmosphere remained at low levels because the available land area and its nutrient supply wassmallcomparedwiththeoceans(Fig. 16.9). By roughly 2 600 Ma small continents had deveIoped by amalgamation of early island arcs. By this time oxygen was still a minor component in the atmosphere, but it was present in amounts sufficient to oxidize iron released slowly by acidic hydrolysis from iron-poor
Earth’s earliest landscapes
243
Figure 16.8 Areconstructionof Stacaandsheigrapaleosolsin north-westScotlandduringthelatePrecambrian(810Ma; from Retdack &Mindszenty1994;with permissionfrom thesociety for SedimentaryGeology).
granitic parent materials. These weakly oxidized soils were similar in profile form to Inceptisols. Iron-rich basaltic rocks, on the other hand, released too much iron for it all to be oxidized.These iron-depleted Green Clay soils were the last representatives of a kind of welldrained soil that became extinct. Carbon dioxide partial pressure in the atmosphere was lowered as more and more carbon was sequestered in organisms and buried
organic matter. Easing of the greenhouse effect allowed accumulation of ice and snow on large mountains and in polar regions. Associated with these earliest glaciations were the oldest periglacial soils and structures. Vertisols and calcareous Aridisols also appeared in seasonally dry continental interiors. These new kinds of soilswere additions that diversified the landscape. By about 2000Ma large continents had formed by
244
Chapter 16
gray or green clay
red or brown clay
corestones
silcrete
a
mantle
Ivvv] basaltic rocks
loelcalcareous nodules limestone or dolostone
1++'+1 granitic rocks
gypsum crystals greenstone belt other metamorphlc
Figure 16.9 A speculativescenario for the evolution of Precambrian soils,atmosphere and continents. The soil types illustrated are Green Clays (A), salty clay soilsor Salids (B). swellingclay soils or Vertisols(C), karst and drab cave earth or Orthents (D), oxidized incipient soils or Ochrepts (E), red and deeply weathered soils or Oxisols (F). desert soilswith silcretes or Durids (G), and desert soils with calcareous horizon or CaIcids (H).
Earth's earliest landscapes amalgamation of smaller ones. Carbon dioxide had become noticeably less abundant and oxygen was present in appreciablygreater amounts. Most well-drained soils by this time had the warm brown and yellow hues of iron hydroxides. Oxisols and laterites date back to about this time. The interiors of continents isolated from the ocean by marginal mountain ranges developed into extensive deserts, including widespread playas and dune fields.
245
Additional desert soils appeared, including soils with siliceous hardpans. The world was becoming a rather more familiar place. This cartoon-like account of early soil development on Earth is undoubtedly incomplete and may also prove incorrect. But its inadequacies can be addressed by studies of additional paleosols, which now can be seen as direct results of those elusive Precambrian landscapes.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 17 Early life on land
Soils and lie on Earth are so intimately interconnected that it is well to ask: ’Whatis life?’Threegeneral properties especially distinguish living things from inanimate open systems such as flames and eddies. Living things are complexly organized, have self-reinforcing sequences of chemical reactions and can reproduce. At a molecular level, even such simple organisms as bacteria can be considered perpetual motion machines of incredible complexity (Fig. 17.1). As in the crudely constructed machines that extend our own livelihood, several classes of components are needed: instructions, duplicators, assembly controllers and work stations. The assembly controllers, for example, are enzymes. These are long strands of protein folded in such a way thatthey interact withother moleculesto promote aparticular chemical reaction within the general molecular chaos of heat agitation. Not only must individual large molecules fit into the general system, but also parts of these molecules must fit within the larger ones. Sugars and amino acidsare asymmetricalmolecules that can be produced abiotically in left- and right-handed versions that polarize light in a clockwise (D) or counter-clockwise (I.) direction. Remarkably, only 1) enantiomers of sugars and Lenantiomersof amino acids are found in organisms. The origin of such complex and finely tooled machines by chance is such an unlikely event as to be near impossible, but it is the very complexity of organisms thatisvitalforkeepingthemgoing(Monod1971). Organisms also are constantly in motion. Even standing still, they process material and energy in ways that keep these processes going. Such metabolic activities include the breaking down of organic matter to gain energy. This is well demonstrated by the fermentation of sugars by yeast to produce alcohol or the respiration of sugars by animals (Table 17.1). Both processes are called heterotrophic because they rely on organic matter for energy. Another kind of metabolic activity is termed autotrophic because organic matter is formed from more simple inorganic materials using the energy 246
of sunlight and special catalysts. Autotrophic metabolic processes include the anaerobic photosynthesis of sulfur bacteria and the aerobic photosynthesis of green plants (Table 17.1). The exact nature of the metabolic pathways is less important in defining life than the way in which they provide both energy and machinery for the enterprise to continue. Organisms also create self-starting, small machines withcoded instructions of themore complexmachinery and processes needed for their survival. The instruction carriers in life as we know it on Earth are nucleic acids such as deoxyribose nucleic acid (DNA)and ribonucleic acid (RNA). The small reproductive machines may be simply a divided cell or specially designed cells, such as spores of plants and sperm and eggs of animals. Research into each of these aspects of the origin and early history of life has made spectacular progress over the past few decades. Experimental investigations have shown that a variety of carbon-based compounds form spontaneously upon electric discharge or ultraviolet radiation into reducing atmospheres containing such gases as H,, NH,, CH,, H,O, N,, CO and CO, (Miller & Orgel 1974). Suchorganicmatteriswidespreadininterstellar dust and in carbonaceous chondrites as old as 4566 Ma (Anders 1989). Unicellular, organic-walled microfossils in cherts of lake and ocean margins have now been found as old as 3 500Ma in the Warrawoona Group of north-western Western Australia (Schopf 1993). Advances also have been made in understanding the nature of metabolic processes and their fossil record. Photosynthetic production of organic matter is a reducing reaction that can proceed in either anaerobic or aerobic modes (Table 17.1). Both kinds of photosynthesis show a preferencefor the lighter of the availableisotopes of carbon (12C rather than 13C) in carbon dioxide. This fractionation of carbon isotopes is virtually unchanged in fossil marine organic matter back to at least 3 800 Ma (Mojzsisetal.1996;Eileretal.1997). Ametabolicactiv-
Early life on land
24 7
Figure 17.1 A simple cell of the prokaryotic enterobacteriumEscherichia coli,considered as amolecularperpetual motion machine withexamples of several classes of spare parts (datafrom Kornberg 1980;Steitzetd.1982:Nanninga 198 5).
ity known as dissimulatory sulfate reduction is a kind of anaerobic respiration linking oxidation of organic matter to the reduction of sulfate to sulfide (see Table 10.1)with a marked preference for the lighter isotope (32S rather than 34S).This began by at least 2500Ma, judging from the isotopic record of sulfates and sulfides (Canfield&Teske1996). Attempts to produce the molecules of organic reproduction, DNA and RNA, by undirected inorganic synthesis have not yet been successful (Orgel 1997), although constituents of these molecules (sugar phosphates and nitrogenous bases) have been found in the residues of experiments energizing mixtures of reduc-
ing atmospheric gases (Miller & Orgel 1974). The synthesis and maintenance of DNA has been found to be such a tedious and difficult procedure that the smaller molecule RNA has been viewed as a more likely original molecule of life. As a molecule that transcribes the code of DNA and also synthesizes proteins using this information, RNA is more self-sufficient than DNA. Parts of RNAmolecules (introns)cut out of the parent molecule can act as an enzyme: they rapidly destroy the molecule by promoting hydrolysis (Zaug & Cech 1986). Some fossil unicells appear to be dividing,but such records are hardly necessary to make the point that life reproduced from 3500Ma to the present (Schopf 1993).Compared
248
Chapter 17
Table 17.1 Simplified metabolicprocessesof organisms. 1. Fermentation
C,H,,O, sugar
+
2C,H,OH alcohol
11. Respiration C,H,,O, + 0,
sugar
+
oxygen
+
CO, + energy carbon dioxide
6C0, + 6H,O carbon water dioxide
+
IU. Anaerobic photosynthesis light 6C0, + 12H,S + C,H,,O, carbon hydrogen bacteriophyll sugar dioxide sulfide IV. Aerobic photosynthesis light 6C0, + 12H,O + C,H,,O, carbon water chlorophyll sugar dioxide
energy
+ H,O + 12s
water sulfur
+
6H,O water
+
60, oxygen
with the vastness of Precambrian time, the time available for the evolution of life is thus remarkably short: between 4566Ma (formation of the planet and carbonaceous chondrites) and 380Ck3 500Ma (isotopic and paleontological evidence of life). Even less time is available if an early period of devastating meteorite bombardment until about 4000 Ma succeeded in almost sterilizing the planet (Sleep & Zahnle 1998).
Did life originate in soil? Despite these experimental and geological advances in understanding the origin and early evolution of life on Earth, there remain large gaps. The steps between tarry abioticallysynthesizedorganicmatterand a systemthat would sustain synthesis of RNA are difficult to reconstruct by any simple scheme. In the fossil record of life there is a similar gap. Nothing is known about life between 3 500Ma microfossils and 4566 Ma carbonaceous chondrites with their amino acids and other simple organic compounds that can be regarded as a geologicalanalog to laboratory experiments on prebiot-
ic synthesis. Both the experimental and geological studies have significant biases as well as gaps. The fossil record of life has largely come from studies of stromatolites. These are laminated, domal structures formed by sedment-binding microbes in shallow water of seashores and lake margins.This bias toward aquatic life extends also to many experimental studies that have produced organic compounds in solution (Miller& Orgel 1974).This tradition of research is compatible with scenarios (Oparin 1924;Haldane1929)inwhichlifearose in a primeval soup of ancient oceans. Alternatively, one can take a general view of the problem and try to recognize perpetual motion machines other than the carbon-based, organic-compoundrecycling andDNA-replicatingcreatures with which we are most familiar. For example, clay or iron hydroxide minerals could stabilize soils and so keep them in nearsurface environments where they would be supplied with further raw materials by weathering. Soils are a crude example of a perpetual motion machine that could plausibly have nurtured life itself (Anderson & Banin 1974; Bohn et al. 1985). Another perpetual motion machine of a sort is the formation of sulfides andreduced organic matter in hydrothermal vents suppliedby volcanic fluids (Corlissetal. 1981;Bocket al. 1996;Russell & Hall 199 7). An additional hypothesis,popularized by the Swedish chemist Arrhenius (1909), is that life evolved elsewhere in the universe, perhaps as close as Mars, and that Earth was colonized by propagules that could withstand longdistance transport in space (Parsons 1996). A related idea is that Earth was colonized deliberatelywith simple organisms by extraterrestrial civilizations (Crick 1981). Suchviewsremove theproblemof lie’sorigins to another part of the universe, but do not eliminate it. The environments where that lie arose are likely to have been Earth-like in many respects because life now is very well suited to Earth. Thus it remains a useful exercise to investigate how life could have originated from natural causes here on Earth. Soup,spa or soil? In building the bodies of organisms as we know them now, carbon compounds are needed in considerable concentration. This is hard to imagine in the open ocean or other aquatic environments, where enormous
Early life on land amounts of carbon would have been needed to create and feed the first organisms. If there were local concentrations of organic matter in dried tidal pools, on sea foam or as scums on icebergs,these would have to be dehydrated for some time to states sufficiently robust to withstand hydrolysis and dissolution on rewetting (Woese 1980).Furthermore, earlyoceansonEarthmay have been 2-3 times as salty as the ocean today without the ameliorating effect of continental freshwater runoff and salt uptake by organisms (Lovelock 1979; Knauth 1998).There are similar problems with an overabundance of water around volcanic vents, especially submarine ones. To this may be added the denaturing of complex organic moleculesby acidic hot water (Miller & Bada 1988). Soils, on the other hand, are wet by chemically mild rain and groundwater, and dry out after rain and flooding. They are stable surfaces that accumulate organic matter produced by even the slowest and least effective of prebiotic organic syntheses. Organismsneed a variety of nutrient elements. Many of these elements, particularly H. C, N, S, 0 and C1, are widely dispersed in water and atmospheric gases. Probably from the beginning these would have been equally freely available in all of the environments considered. Micronutrient elements may not be so widely dispersed, but are needed in such small quantities that they seldom are limiting. The remaining macronutrients, on the other hand, show distinct partitioning between environments. Of these phosphorus is, and probably was, the limiting nutrient for life because it is needed in much greater quantities than is available in most rocks, soils, sediments and waters (Fig. 17.2). Phosphorus is much more abundant in basaltic (1400mg kg-l) than in granitic rocks, soilsor shales (700-800 mg kg-': Bowen 1979). There is very little in fresh water (0.02 mg kg-') orseawater(0.06mg kg-l). Lifeismoreabundanton the margins of large oceans and lakes than in open waters away from continental sources of phosphorus, which tend to be biological deserts (Lieth & Whittaker 1975). Potassium is strongly retained in clays of soils and shales. Sodium is readily dissolved during weathering and has accumulated in the ocean. The behavior of magnesium and calcium is intermediate. They are lost from many soils, but in dry climates they accumulate in soils.Themacronutrient element composition of organisms is variable. Edible vegetables are especially rich in them compared with other land plants. So are mineral-
249
ized tissues such as bone compared with muscle. In general, however, the pattern of relative abundance of nonvolatile macronutrient elements in land plants is most l i e that of soil, but the pattern in marine algae and fish and land animals is more like that of marine clay. These patterns of relative abundance found in organisms are unlike those for granitic rocks, basalt, fresh water or seawater. Furthermore, the composition of life is compatible with the view that it evolved from commonly available materials here on Earth, or a planet very similar to it, rather than in some different part of the universe. In addition to raw materials for life, stringent conditions are needed to sustain metabolic activity. These conditions may have been even more exacting for the earliest of life forms, before the advent of robust and varied cellular life in which metabolic processes are protected from the outside world by membranes. Temperatures within the range of liquid water, moderate pH and Eh, and protection from cosmic rays and other radiation would beneeded, as well as acomplex structure for separation of reactants and products, and a source of energy. Oceanic environments have appropriate temperatures for themostpart, but wereprobablywarmeron theearly Earth (Schwartzmann & Volk 1991).The Eh of oceans can vary considerably from anoxic bottoms to oxidized surface waters. In reduced atmospheres early in Earth history,the ocean may have been more uniformly reducing and capable of preserving organic matter (Holland 1984).ThepHof oceansis toostablefor theformationof all the compounds needed for life. Fats and some amino acids require acidic conditions whereas other amino acids necessary for life form under alkaline conditions (Miller & Orgel 1974). Furthermore, complex polypeptides such as coenzymes require desiccation for assembly through hydrogen bonding (Keefe et al. 1995). It could be argued that these various compounds were formed and then combined in locally isolated parts of theocean, such assmallpools onicebergsorin tidalflats. But low-temperature or evaporitic concentrations of salts would make useful combinations of these complex molecules difficult. Energy sources useful for synthesizing organic matter and driving metabolic processes include lightning, shock waves, meteorite impact and sunlight. Most of these sources of energy, with the exception of sunlight in visible and infrared wavelengths, would have been more harmful than helpful, especially
2 50
Chapter 1 7
volcanic poses
. ..... . . ... ...... :..,:.... ,....:.....*.:*:. ..*: ,.'.....:.: *
P co K Mg No
KEY
0edible vegetables fresh woter
bone ond shell
P No
4
m o t h e r parts enlarged scale rnorine fish
ronge of values
No
Figure 17.2 Average amountsof mineralmacronutrients(R Ca, K. Mg. Na) inorganism, soils,rocks and natural gases (datafrom Bowen1979).
in shallow water. Micrometeorite bombardment may have been excluded by an early dense atmosphere, but large meteorites may have been very destructive of terrestrial and shallow marine environments (Sleep & Zahnle 1998). Ultraviolet light may have been harmfully intense before oxygenation of the atmosphere and development of the ozone screen. The protection of diaphanous materials including carbonates may have screened some radiation (Sagan &Pollack1 974). Constraints on metabolic processes in volcanic hot springs are very different. The temperature of water gushing from deep oceanic vents may reach 380"C, but there is a rapid decline in temperature over only a few centimeters away from the vent (Spiess et aJ. 1980). Some vent microbes live at temperatures as high as 113°C (Jsrgensen et al. 1992; Pace 1997). Volcanic gases and fluids from vents are also strongly acidic and chemically reducing (von Damm et al. 1985). The re-
ducing power of hydrothermal vents, the strong Eh gradient to lessreducing surroundings and a strongly acidic pH could have helped make some compounds useful for life, such as heat shock proteins (Nisbet & Fowler 1996; Amend & Shock 1998), but not others (Miller & Orgel 1974). Most of these organic compounds would be washed out into the open ocean, but volcanic vents both on land and sea do have networks of compartments that could have separated and slowed reactions. Examples include the porousvuggy and brecciatedaccumulations of metal sulfides around deep-sea vents, siliceous sinter deposits found on land and associated pre-existing vesicular lavas or porous volcanic ash (Humphris et al. 1995; Rice et d.1995; Russell & Hall 1997). Volcanic vents also receive energy in the form of subterranean heat in deep-sea vents and this is isolatedfrom the harmful effects of hard radiation and meteorite impact at the surface. Nevertheless, conditions most favorable for life
Early life on land are some distance from the vent itself, where vents grade into normal marine and soil environments. Soil temperature, pH and Eh are generally more moderate than near volcanic vents. The temperature and pH of soils now vary considerably, but under a carbondioxide-rich early atmosphere soils are likely to have been more uniformly hot and acidic (Schwartzman & Volk 1991). Soil heating as a result of early meteoritic bombardment could also have selected for heat tolerance during the early evolution of life (Sleep & Zahnle 1998).The Eh of soils is usually equal to, or lower than, thatof theatmosphere, whichontheearly Earthislikely to have been reducing (Rye & Holland 1998).Soils also contain an extensive array of compartments in the form of soil pores. Beneath the soil solum there may be additional compartments in volcanic ash, vesicular lava or deeply weathered granite or uncemented sand. There is a greater variety of pH, Eh and porosity in soils within very small spaces than in either volcanic hot springs or the sea.The less extreme and locallyheterogeneous conditions in soilsare more suited to persistence of complex organic molecules. Reactions in dilute solutions of soils are controlled by the intricate geometry of water films and surface properties of mineral grains rather than by bulkchemical and physical properties of the medium. In hot springs chemical reactions are driven to a particular outcome by flow through a narrow zone of rapidly changing temperature and other conditions. In the sea reactions tend to a chemical equilibrium. In soils, however, a variety of complex and reversible reactions are feasible. Wetting and drying of soils promotes dissolution and precipitation of salts, hydrolysisandneoformation of silicate minerals, and swelling and shrinking of clay. Soils are more like oceanic environments than hot springs in deriving most of their energy from sunlight, andwiththiscomesapotentialcostintheformofharmful ultraviolet radiation. Just as in the shallow ocean, however, there are levels where harmful radiation is screened to manageable levels by refraction through diaphanous materials or by reflection within soilpores:yet sufficient heat and light energy remains to drive photosynthesis and weathering reactions (Sagan & Pollack 19 74). Ultraviolet and other destructive radiation would also be attenuated by an early photochemical smog of methane, carbon dioxide and water vapor (Sagan & Chyba 199 7). Life in soils also would be prone to destruction by large meteorite impact until such time
2 51
as it extended to deep soil cracks or other protected sites from where it could recolonize devastated landscapes. Methods of reproduction can be considered in terms of themodernsystembasedonnucleic acids.Theformation and persistence of a simple nucleic acid such as heat-tolerant, hydrolysis-resistant, guanine-rich RNA is difficult to envisage in the ocean. Moreover, if such a fragile self-replicating molecule did arise against all odds, then the problem would be regulating it. In an extended aqueous medium it would either devastate other evolving molecular systems that included suitable nucleotide bases or diffuse away until dissolved.This problem is less extreme in a volcanic vent because life could have arisen in the narrow zone between abioticallyproduced organic moleculeswithin the vent and the homogeneous, nutrient-poor ocean or atmosphere beyond. But this zone is very narrow, with vigorous flow of hot, acidic water that decomposesamino acids in a matter of hours (Miller & Bada 1988). Although hot springs microbes have been considered phylogenetically primitive (Bock et al. 1996; Pace 1997). the predicted RNA composition of their likely common ancestor is incompatible with hyperthermophily (Galtieret al. 1999).Replication is least troublesome in soils,which have numerous small spaces for the assembly of different reaction products. These dry out periodically so that reactions in solution are disrupted in a dehydrated disequilibrium. If the supply of raw materials by rainwater kept pace with the time available for episodes of synthesis in evaporating water films, then reproduction of molecules could continue indefinitely. Many local replicating molecular systems could develop in a single soil adjacent to different minerals from the parent material. In addition to the replication of individuals, the natural selection of different kinds of molecular systems would have been necessary for them to have evolved in any sense like life. The earliest living molecules would have had some property that promoted their survival (phenotype) as well as the information to create that property (genotype). A particular consistency or sliminess of organic matter, for example, could create complex surface foams from wave action in the sea, or bind replicating molecular systems to an especially favored location in the stream of a volcanic vent, or stabilize a soil against destruction by erosion. The greater stability of soils over longer time spans would have been more effective in amplifying the natural selection of initially
2 52
Chapter 1 7
slow and inefficient synthesis of organic matter and early biologicalsystems. Did life originate in an oceanic soup, volcanic spa or soil sludge? The historically acclaimed concept of a primeval soup has stimulated a great deal of excellent experimental work in organic chemistry and other pertinent areas of the physical sciences (Miller & Orgel 1974). Similarly,the idea of life originating in submarine volcanicvents has stimulated much useful research on the microbiology of extreme environments and on microbial phylogeny (Bock et al. 1996; Nisbet & Fowler 1996).However, there has beenlittleexperimentalwork on the origin of life in soils, despite its theoretical claims as a flow reactor that could have nurtured the origin of life on Earth. In the ensuing paragraphs I outline some modeling andexperimental studies of thekindneeded to explore this additional option for how lie might have evolved.
Role of days Because of their complexphysical and chemical properties and fine grain size, that make detailed understanding difficult, clay minerals have figured prominently in several theories about the origin of life (Cairns-Smith& Hartman 1986). Specific roles assigned to clays include concentrators of reactants, catalysts for reactions, templates for assembly of complex molecules, compartment boundaries for reactions, and devicesfor the storage and transfer of energy and information. The ability of clays to absorb organic matter is well known. The expansion of smectiteclays by absorption of ethylene glycol between the aluminosilicate sheets is a standard determinative tool for identifying smectites. I use smectitic clays to pick up spills of cutting oils in the laboratory, and have often found my hands chapped and dry after diggingin smectitic paleosols. Clays also have been shown to aid in the synthesis of parts of organic molecules (monomers) and in their assembly into complexly folded chain-like organic molecules (polymers).For example, cytosine, uracil and cyanuric acid are a few of the monomers that have been formed from carbon dioxide and ammonium hydroxide in the presence of kaolinite clay. Long polypeptideshave been formed from aminoacyl adenylate and oligonucleotides from phospharamidates in the presence of smectite (Rao et al. 1980). Montmorillonite promotes
assembly of activatedmonomersinto RNAoligomers up to SOunitslong(Ertem&Ferris1996).Theexactroleof clays in promoting these reactions is not known with certainty. It could be that they participate in reversible intermediate reactions as a kind of catalyst, that their pattern of charge distribution serves to adsorb and orientate the reacting molecules as a template, or that they have amoregeneralregulatoryeffectas achemostat, for example,by buffering acidity. Clays form a variety of microscopic structures that could be regarded as compartments for chemical reactions. A ‘house of cards’ structure is widespread in clays because of the generally negative charge on the surface of the platelets, but positive charges along the edgesthat develop under neutral to acidicconditions. Long narrow tubes are formed by imogolite, concentrically arranged vesicles by halloysite and flexible sheets by rectorite (Sudo et al. 1981).These kinds of structures are not so well designed as the glassware of a modern molecular biology laboratory or the ribosomes of living organisms. Nevertheless, differences in pH, Eh and chemical contents of compartments could feasibly allow more complex sequences of reactions than in open water. Somekinds of energy can be stored and transferred by clays. Considerable physical forces are unleashed when expanding clays are wet. These forces can shear and compress ped faces or throw the clayey surface of a soil into folds (see Fig. 4.2). Clay expansion also can expel soil solutions to sandier parts of the soil and temporarily seal the subsurface of the soil from the atmosphere. Light energy can be stored in clays when electrons are excited by sunlight into defect positions in the atomic structure of clays. The release of this energy can be triggered by drying out of the clay or by wetting with chemicalssuchas hydrazine(Coyneetal.1984).Whetherthis thermoluminescent energy is powerful or directed enough to do useful biosynthetic work remains to be determined. The ability of clays to store and transfer information has been claimed (Weiss 198 l),but so far not proven to the satisfaction even of those scientists most receptive to the idea (Cairns-Smith & Hartman 1986). The idea remains a theoretical possibilitythat clays could be a kind of genetic crystal in which a specific order of crystal features is replicated by crystal growth (Cairns-Smith 1982). Several kinds of crystal genes can be imagined (Fig. 17.3). A very simple kind consists of mixed layer
Early life on land
reproduction
TWO DIMENSIONAL
253
gram ONE DIMENSIONAL
Figure 17.3 Diagrammatic representationsof Cairns-Smiths(1982) hypothetical concept of clay mineralsas crystal genes,in which physical properties are coded by layer sequences(left)or patterns of ionic substitution(right).
clays in which layers of different mineral composition are arranged in a specific sequence. This kind of crystal gene would grow by neoformation of clay along the edges and reproduce by breakage of complete stacks of the sequence that could nucleate further copies. Another general kind of genetic crystal could carry information as a particular pattern on the sheet-like crystal faces: for example,by substitution of iron for aluminum within the atomic lattice. Many other kinds of patternforming crystal defects also could carry information. Growth of this kind of crystal gene could proceed by addition of new layers that replicated the pattern. Reproduction of these crystals would be achieved by breaking along the cleavage planes to form nuclei for other stacks of the original pattern. The value of such sequences or patterns can be imaginedinsofar as they might affect the bulk properties of clay, such as its swelling capacity in water. A genetic crystal of clay is one that could replicate physical properties useful for its own persistence. Such vital mud has so far eluded discovery, but the concept points the way to fundamental research on the ultrastructure of clays and on its effect on their bulk properties. Role of iron minerals A variety of iron-bearing clays (nontronites), carbonates (siderite) and hydroxides (goethite) are formed by the weathering of iron and iron-bearing silicate minerals. These minerals also are of interest for studies of the origin of life because they form compartmentalized structures and facilitate oxidation-reduction reactions that can result in the production of organic compounds. The crystal forms of fine-grained rust minerals rival those of clays in complexity and variety. They include small hexagonal planes of hematite, sheaf-like bundles
of siderite needles, hollow balls of ferrihydrite (Dixon& Weed 1989) and extensively channeled crystals of akaganeite (Holm 1985). Like comparable structures of clays, these could have separated chemical reactions important to the early evolution of life. Organic compounds also have been formed by photochemical oxidation of iron carbonate (siderite).In these experiments, ultraviolet light photo-oxidized the iron of siderite to ferric hydroxide and reduced its carbonate to formate, formaldehydeand other organic compounds (Joe et al. 1986). The photo-oxidation of iron-bearing clays also is a potential electron source for the reduction of carbon dioxide to organic compounds in solution (Braterman et al. 1983). Oxidation of reduced iron (Fe2’) also has been shown to produce ammonia (NH,) from nitrite (NO,-), a reaction promoted by salt (NaCl),but inhibited by bicarbonate (HCO,-) and light (Summers & Chang 1993). These kinds of redox reactions can be considered an abiotic form of light-aided organic synthesis (‘photosynthesis’) comparable with ferrous iron oxidation by anoxygenic phototrophic bacteria (Widdel et al. 1993). Before the Earth’s atmosphere was appreciably oxidizing and before continental crust became granitic, there would have been more minerals with reduced iron near the surface of the Earth than at present. The photochemical production of organic matter by such surface processes could have created soil materials similar to carbonaceous chondrites. The survival of organic matter and early organisms in the face of continuing destructive ultraviolet radiation would have been facilitated by the coprecipitation of opaque iron hydroxides (Olson & Pierson 1986).These mineral factoriesof abiotic organic matter deserve the kind of dedicated testing that has been lavished on the study of organic synthesis in aqueous solutions.
254
Chapter 17
A scenario of a selfish soil Much of the line experimental and geologicalwork performed on the origin of life in the sea was inspired by the hypothetical scenario proposed by Oparin (1924) and Haldane (1929). Studiesof the origin of life are so full of theoretical possibilities that such internally consistent scientific fables can be a valuable conceptual focus. Similar scenarios for the origin of life in volcanic vents havebeendeveloped(Cor1issetal. 1981: Nisbet&Fowler 1996). A detailed scheme for origin of life in the soil can be envisaged very much like that proposed theoretically by Cairns-Smith (1982). This alternative scenario is detailed here to promote its investigation and as a good story inits ownright. Stage one in Cairns-Smith's scenario is a time of complex clays, which could have been most abundant in stable, intermittently wet, surface environments of moderate temperature, such as soils. Here the growth of crystals would be slow, confused by adjacent crystals and periodicallyinterrupted or reversed. This fickle crystallization is very different from the externally driven crystallization of plagioclase and olivine in a rapidly cooling basaltic lava or the precipitation of quartz crystals from hot vein-filling fluids in a cooling granitic magma. Clays may have been formed initially on Earth by hydrothermal alteration and by weathering during
sunlight
\\\
nI
the earliest degassing and differentiation of the planet. More complex clays would have become widespread as surface environments assumed more of the character of a chemostat. Atmospheric carbon dioxide could have warmed the Earth by a greenhouse effect despite the weak radiation from a young and not yet fully ignited Sun (Caldeira & Kasting 1992).Raising of surface temperature would allow liquid water. This would have been acidic with dissolved carbon dioxide and so weathered the landscape to produce residual clays and soluble carbonates. Carbonates accumulating in oceanic sediments would have buried carbondioxide that might otherwise accumulate in the atmosphere to develop a runaway greenhouse like that on Venus. Such a fragile early chemostat could develop only on planets of a certain sue and distance from the Sun, as the failed chemostats of Venus, Mars and the Moon demonstrate. Nevertheless,such geochemical balances were vital for allowing complex reversible soil reactions. Stagetwo is a time when certain kinds of clays become more common because they have bulk properties useful for their own persistence. Let us consider, for example, Cairns-Smith's (19 7 1) story of the misadventures of four kinds of clay, nicknamed Sloppy, Lumpy, Sticky and Tough because of their wetted consistencies (Fig. 17.4). Imagine that these clays are formed by weathering of mineral grains by rainwater percolating
r","
LUMPY
u
SLOPPY
u
&@@Wl.
windblown dust /I sheetwash
0 TOUGH
Figure 17.4 Differential survival and evolutionof four kinds of clay nicknamed Sloppy,Lumpy,Sticky andTough,each of different consistencieswhen wet, among the mineral grains of a soil (modifiedfrom CairnsSmith 1971;with permission from the Universityof Toronto Press).
Early life on land through a soil. Depending on the parent material and chemical composition of local solutions they form different proportions of smectite interlayers. The Sloppy clay expandsreadily when wet and is washed deeply into the profile during heavy rains. Tough clay, on the other hand, is virtually inert. It forms immobilerinds that surround the minerals from which it weathered and thus prevent further weathering. Because of continuing losses of Tough and Sloppy, the clays that come to dominate the surface soil will be those with intermediate characteristics, such as Lumpy and Sticky. These clays expand during rain so that they plug the soil pores and maintain their position in the soil. When conditions are drier they crack away from the parent mineral grains so that additional clay can be made. Lumpy clay protects against erosion following mild rain, but a less easily dislodged clay such as Sticky may be needed to defend the soil against thunderstorms and sheetwash. In this way there is natural selection of particular kinds of clays for particular environmental conditions. Those clays in the zone of materials and energy transfer at the surface of the soil could potentially persist by preventing soil erosion. In contrast, clays washed out to sea or deep into the soil were not subject to such stringent natural selection. During a third stage of more deeply weathered stable soils, particular clay minerals would have been produced in abundance and dispersed as small crystals. In our mineral melodrama of clay formation in soil (Fig. 17.4),this would promote the formation of Lumpy clay over Sticky because of Lumpy’s superior ability to form propagules that could withstand transport. Differences in physical appropriateness between Sticky and Lumpy clay might be slight, but the slower dispersal of Sticky would lead to soils regionally dominated in the long run by Lumpy. Nevertheless, a variety of clays would be found in differentplaces and levels of soil. Stage four heralds the synthesis of organic matter. The photosynthetic production of organic matter by electrons gained from iron-rich clays excitedby ultraviolet lightmaybeacrudeprocess(BratermanetaZ.1983). However, the organic matter produced could have a sliminess or other physical property that further stabilied the soil against erosion. Even with slow and crude methods of productionin soils, organic matter could accumulate to considerable abundance in time. This positive feedback would have allowed natural selection for the most effective soil stabilizing agent, because less ef-
255
fective organic matter would be washed out to sea or buried away from sources of energy and materials for additional organic synthesis. The production of opaque oxyhydratesby the photo-oxidativephase of such primitive ‘photosynthetic’ reactions could filter destructive ultraviolet radiation from organic surface horizons. Soil material at this hypothetical stage would have appeared similar to carbonaceous chondrites. Stage five is one in which structural complexity of clay minerals allows multistep organic synthesis. Such processes can be imagined within tubes or chains of vesicles in clay or iron minerals. Large multicomponent organic molecules such as polymers of enantiomerically pure sugars might be formed in this way. These might further bind the soil against erosion, thus ensuring their own preferential survival. The soil at this point can be imagined as a crude, abiotic version of a ribosome. By stage six organic polymers are not only binding together mineral grains, but also forming structures of their own. Such structures could include round bodies (proteinoids of Fox & Dose 19 72),as well as sheets and complex networks. Some of the circular organic structures in carbonaceous chondrites (Folk & Lynch 1998) could represent such locallyflocculatedmasses of abiotically produced organic matter. Some of these organic massesmayhavecatalyzedorganicsynthesis:as theearliest ribosome-like particles (Barbieri 1985) or photosyntheticreactioncenters (Pierson&Olson1989). Stage seven sees the appearance of genetic material formed from organic matter. Long molecules such as DNA or short folded molecules such as RNA may have had value as soil-binding agents themselves. Even if less effective as binding agents, their accurate replication independent of adjacent mineral grains would ensure more consistent reproduction of particular structures and physical characteristics. By stage eight organic genes have proven more versatile, more miniaturized and more rapidly assembled than the archaic clay machinery surrounding them. These organic genes already may have begun to parasitize some steps in the archaic clay-catalyzed synthesis of large organic molecules such as proteins. With this newly refined manufacturing ability these complex organic molecules could begin a takeover of this earlier function of soil clays and iron minerals. The outmoded crystal machinery of the soil becomes increasingly out-
256
Chapter 17
paced and outclassed by the speed and efficiency with which organic molecules can commandeer the primary source of energy from the Sun and the primary materials of carbon dioxide,water and nutrient elements. Stage nine is one of increasing complexity of the organic machinery of life as genetic molecules begin coding and producing enzymes and lipid membranes that further facilitatetheir own survival. Someof these independently functioning organic systemscontinued in the ancient mode of photosynthetic production of organic matter. Others could have begun consuming organic matter in the soil as the fist heterotrophs. The bodies of filamentous and amoeboid forms would have further strengthened the ability of the soil to withstand erosion. At this point also early organisms enveloped in protective membranes may have begun to expIoit other environments less favorableto complexbiosynthesis such as theoceanandvolcanicvents.This stagewas achievedby at least 3500Ma, judging from the fossil record of microfossils (Schopf 1993) and of thick, clayey, organicpoorpaleosols (Rye&Holland1998). This detailed story of ever more complex systems would seem highly unlikely were it not that at each step theresultis improvedsoilstabilitythat servestoperpetuate the system in thezone of energy and materials transfer. The subsequent evolution of nucleated cells by the genetic takeover of free-living cyanobacteria (future chloroplasts) and proteobacteria (future mitochondria) by respiring archaebacteria (Margulis 1981) or by methanogens (Martin&Muller 1998).allowed the evolution of larger and more complexorganisms that were more effective at binding the soil. This natural selection was continued with the evolution of multicellular land plants and then of increasingly complex forest communities. By such a view, three-dimensional machines as complex as living things did not evolve by chance, but out of necessity imposed by the arrow of time.
Evidence for early life in paleosols Theoretical studies of the origin and early evolution of life are useful indications of what might have happened. But ultimately there are too many theoretical possibilities, not all of which could actually have happened. Historical records are needed. Paleosols have much to offer as records of past life on land. Unfortunately,the record is severely compromised by difficultiesof preservation.
Organic microfossils, for example, are not preserved in oxidized paleosols. However, they may leave microbial trace fossils in the form of lamellar or ministromatolitic mineral deposits. They also could leave evidence of Po$saccharides or other organic binding in the form of recognizable soil structures. Finally,there may be evidence of life in the distribution of organic carbon and of elements such as phosphorus that are important for life in the soil.
Microfossils Microfossilshave now been found in many Precambrian paleosols, but problems remain in demonstrating their authenticity and role in Precambrian soil formation. Microfossils in paleosols were long ago expected because of the similarity of Precambrian marine and freshwater microfossilsto modern soil microbes (Campbell 1979). For example, Eoastrion is a stellate microfossil with an opaque central body (Fig. 17.5A-D) abundant in parts of the 2000Mashallow marineGunflint Chertof southwestern Ontario, Canada (Barghoorn &Tyler 1965). It is identical to the problematic modern Metallogenium, common in soils and rock varnish, and variously regarded as a giant manganese-king bacterium, a colony of bacteria or their degraded traces (Staley et al. 1982; Margulis et al. 1983). Other Precambrian microfossils such as Eosynechococcus also are similar to modern microbes that colonize bare rocks on land (Golubic & Campbell 1979). Kakabekia (Fig. 17%-H) is another microfossil from the Gunflint Chert similar to modern soilmicrobes (Siege11977). The most ancient of plausible soilmegafossilsare supposed fossil fungi and lichen-like plants in 2900 Ma fluvial rocks of the Carbon Leader in the Witwatersrand Group, South Africa (Hallbauer & van Warmelo 1974; Hallbaueretal. 1977;agefromdeWitetal. 1992).Their carbon is biogenic considering its isotopically light composition (-22 to -27%0 613C,,,), presence of pristane and phytane, and pentoselhexose ratio of unity (Prashnowsky & Schidlowski 1967). Perhaps some of the fine tubular structures may be artifacts of vigorous chemical maceration to prepare the specimens (Cloud 1976),but stout tubular structures also can be seen in untreated thin sections. Another interpretation of these structures is that they are hydrocarbon spindlescreated by migration and deformation during hydrothermal al-
Early life on land
teration (Barnicoat et al. 1997), but this view is difficult to reconcile with their intimate admixture with lowtemperature minerals such as kaolinite and with the biochemicalcomplexity of the organic matter. Angular fragments of chert with a microlaminar appearance and a very light carbon isotopic composition (-40%06~~C,,,)havebeen foundinthe2765 MaMtRoe paleosol of north-western Western Australia (Rye 1998).The laminae also include spheroidal forms, but their preservation is too poor for positiveidentication as cells. The sharp angular margins of the chert fragments are suspicious, because they could be clasts redeposited from aquatic microbialites (Rye&Holland 2000). These problems do not apply to beautifully preserved permineralized nostocalean cyanobacteria found in cherts on paleokarst of the 1300Ma Mescal Limestone of Arizona (Horodyski & Knauth 1994). These microfossils are associatedwith a paleosol, but were these soil biota or opportunistic aquatic microbial colonists of puddles on the paleokarst? Microfossil nostocalean cyanobacteria have been found in cherts filling cracks into basement rocks underlying the 2 100-1800Ma Pokegama Quartzite in
2 57
Minnesota, USA (Cloud 1976). This could have been a crevice fauna within a soil, but it is also possible that these microbes lived in the intertidal or subtidal zone of the sea in which the overlying Pokegama Quartzite was deposited. Well-preserved phosphatiied filaments in Middle Cambrian ( 5 1 5Ma) phosphorites of western Queensland may represent cyanobacterial colonization of subaerially exposed limestone (Southgate 1986). Large phosphatized striated and branching tubular structures represent more complex life forms, such as algae or lichens (Fleming&Rigby1972;Retallack 1 9 9 4 ~ ) .
Possible microbial trace fossils have been found in the 2200 Ma Waterval Onder clay paleosol of South Africa (Retallack 1997a).These are features l i e modern rock varnish. Small (1-0.1mm in width) grains in the surface (A) horizon of the paleosol have botryoidal encrustations of opaque iron and manganese (manganoferrans) on their tops but not on their bottoms. Examination using high-resolution back-scatter
258
Chapter 1 7
electron microscopy has shown that any possible internal details have been obliterated by metamorphic recrystallization (Retallack & Krinsley 1993). There are numerous theories on how modern rock varnish forms, but a biological origin now is widely accepted. Debate continues on what kinds of organisms are responsible: bacteria, algae and fungi are associated with modern varnish(Krumbein&Jens 1981;Staleyetal. 1982). Comparable microstructure that may be a microbial trace fossil has been found in the surface (A) horizon of the Jerico Dam paleosol beneath the 3000 Ma Pongola Supergroup in South Africa. These are microscopic caps of fine-grained andalusite (probably once kaolinite) with internal laminae of amorphous iron and manganese oxides on top of large quartz grains. These small colloidal caps (former manganoferriargillans) are distinct from the sericitic materials of the rest of the paleosol and could not be weathering rinds of the underlying quartz grains (Grandstaff et al. 1986).They are similar in some ways to ministromatolitic forms of rock varnish (see Fig. 10.2A)and could represent microbial colonies under which dust was trapped, weathered to clay andinterstratified with opaque minerals produced by oxidative photosynthesis or chemautotrophy.
Soil structure A striking feature of the Waterval Onder clay paleosol is the very different soil structure of its surface (A and AC) horizons compared with its subsurface (C) horizons (Retallack 1986a).The surface (A) horizon has aplaty structure with scattered crumb peds. The near-surface (AC)horizon has coarse, angular blocky pedal structure defined by clay and opaque oxides (illuviationmanganoferriargillans) washed down the cracks (see Fig. 3.13). These structures are similar to those made in soils, in showing wispy edges, healed cracks, indications of slight rotation in a loose medium and alteration haloes different from those seen in simple, abiotically produced breccias, boxworks, or systems of mudcracks. The distinctive structures of modern soils are created by coatings of polysaccharides and other materials (Foster 1981), which smooth over the edges of peds, bind together groups of peds andlocalize the actionof weathering solutions. Organisms and their slimy organic products serve to stabilize the entire soil as well as individual peds (Booth
1941). Rock varnish (Krinsley 1998) and carbonate crusts (Krumbein & Giele 1979) armor modern desert soils against erosion. Could this be the reason why many Precambrian paleosols are so thick, clayey and deeply weathered? The 3460 Ma pre-Warrawoona paleosol of Western Australia has evidence of weathering to form clay minerals at depths of 50 m (Buick et al. 1995).The Jerico Dam paleosol below the 3000Ma Pongola Group in South Africa has about 50 vol. % clay to a depth of at least 6 m (Grandstaff et al. 1986).The 2450 Ma Denison and Pronto paleosols of Ontario, Canada, also are impressively welldeveloped(seeFig.16.3).Theserepresent landscapes that were stable for hundreds of thousands, if not millions, of years. This is not what one would expect on an abiotic landscape. Under abiological conditions each mineral grain as it was loosened from bedrock by chemical and physical weathering should quickly have been removed by surface erosion if it were appreciably above water table (Schumm 1977). As a consequence, Precambrian paleosols on bedrock should be sandy and thin rather than clayey and thick as some of them were. Under abiotic well-drained conditions also, soils formed on clay deposits should have been thin and shallowly cracked like those forming in desert badlands of the western USA (Schumm 1956). Yet Precambrian paleosols on shales such as the Waterval Onder paleosol (see Fig. 9.13) are not of this kind. Could organic matter have been a glue that held Precambrian landscapes together? Organic carbon Evidence for life in Precambrian paleosols also can be gained by analysis for organic carbon and its isotopic composition.Down-profileattenuation in abundance of organic carbon is characteristic of soils (Stevenson 1986) and so is the preference for the common light carbon isotope (12C rather than 13C) characteristic of photosynthetic organisms (Driese & Mora 1993). Both indices of ancient life are displayed by a paleosol underlying the 2560 Ma Black Reef Quartzite near Schagen, Transvaal. South Africa (Fig. 17.6). The carbon isotopic valuesof thepaleosol(-16 to-I4%06~~C,,,) arenot as light as those of the overlying marine Black Reef Quartzite (-35 to -30% 613C,,,), but such values are found in microbes of hypersaline environments. Such an alkaline and saline soil is compatible with the abun-
Early life on land CaO (wi%)
SIO, (Wt%)
1
10
100
.01.1 1
10 100
ThICr (wi)
104
lo4
10-2
HIC (atomic) 0.04 0.08
0.00
2 59
-2
0 2
5 0 'One
50
100
Mineral abundance (%)
0 0 0 1 0.2 0.3 0.4 Org. C (W%)
-36 -28 -20 -12
-
sl%m (%3
Figure 1 7.6 Mineralogicalandchemicalcharacteristicsof a 2560Mapaleosolfrom Schagen,South Africa. Isotopicallylight carbonis strong evidence for life in the paleosol.The parent material is carbonate-freeserpentiinite and the paleosol is overlain by Black Reef Formation.The solid bar inthe graph for carbon isotopic composition of organic matteris for shales of the Black Reef Formation at other localities.and the nearby circle is for a stromatolite from the Malmani Subgroupat Schagen (fromWatanabe et al. 2000; with permission from Nature and MacmilIanJournals).
dance of dolomite and calcite in the paleosol (Watanabe etal. 1998,2000). Lightening upwards trends in both carbon and oxygen isotopic composition have been found in a paleosol developed on the 1200Ma Mescal Limestone near Young, Arizona, USA (Beeunas & Knauth 1985: Horodyski & Knauth 1994). These isotopic trends are similar to thoseseen below modern soils (Chromusterts) on limestone on the south coast of Barbados (Ahmad & Jones 1969) and in Mesozoic and Paleozoic paleosols ranging back in age to Ordovician (Allan & Matthews 1982:Goldstein 1991;Tobin & Walker 1994).Endolithic microbiota of these limestone paleosols probably included cyanobacteria, algae or lichens l i e those active in blackphytokarst today(Fo1ketaZ. 1973:Viles 1987). Yet another microbiota is indicated by the exceedingly light isotopic composition (-40%0 613Cp0B)of microlaminatedchertchipsinthe 2 765 MaMtRoepaleosol of Western Australia (Rye 1998). Carbon isotopically this light is known to be produced only by methanogens, which become appropriately heavier in
their carbon isotopic composition (Hayes 1994), thus isotopically light methane from elsewhere was probably fixed in the chert chips of the paleosol by methanotrophs. Unfortunately, the chert chips have angular margins, andmaybe clasts from anearbymicrobialmat, rather than part of thepaleosol (Rye&Holland2000). Normal values for modern soil organic matter (-2 5 to -27%0 613C,,,) have also been reported from Precambrian paleosols (Mossman &Farrow 1992: Retallack & Mindszenty 1994). In the case of the 810Ma Sheigra paleosol (see Figs 16.3, 16.5 and 16.6) there is also a downward attenuation in organic carbon content from the top of the profile. Such isotopic values and depth functions are similar to those in modern well-drained soils. A final consideration is the organic-lean composition of all known Precambrian paleosols, which are comparable with many Phanerozoic paleosols in this respect. Even apart from the idea that carbonaceous chondrites are primeval paleosols (Bunch & Chang 1980), the existence of organic carbon in paleosols at Iow levels
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requires the presence not only of producers, but also of consumers, such as actinobacteria or fungi. Without them, soil microbiota in the weakly oxidized atmosphere of the Archean would have created carbonaceous soils. Although these may be present as yet unrecognized among black shales of the Precambrian, the existenceof so many organic-lean paleosols is remarkable. The isotopic composition and distribution of carbon in Precambrian paleosols is indicating that life on land was not only ancient but varied, even as early as Archean time. Hypersaline, endolithic, methanotrophic and decompositional communities all may be represented. Trace elements In addition to information on organic matter in Precambrian paleosols, much also is known about trace elements commonly complexedby organic matter, such as Ba, Cr, Cu, Ni and Zn (Farrow & Mossman 1988).In the 2200Ma Waterval Onder clay paleosol, for example, these elements are all enriched toward the surface. Most of these elements also are enriched in clayey parts of modern soils (Aubert & Pinta 1977),but this is not an explanation for the upper part of the Waterval Onder paleosol because it has a nearly uniform clay content throughout. These enrichments are especially striking by comparison with the immobility or depletion of other traceelements suchasZrandRb.whichalso arestablein modern soils. Another element showing surficial enrichment in the Waterval Onder paleosol is phosphorus. This Scarce macronutrient is depleted at the surface of modern forestedsoils and paleosols because of the requirements of large plants (Smeck 1973). The pattern of surficial enrichment of phosphorus in the Waterval Onder clay paleosol is more like that found in modern grassland soils or microbial earths, in which much of the biomass is within the ground. Surficial enrichment of phosphorus has been reported in many Precambrian paleosols ranging in age from 3000 to 1647Ma (Grandstaff et al. 1986; Miller et al. 1992). Biological activity is also a likely explanation for the local redistribution of iron and manganese in mottled floodplain paleosols (Dystrochrepts) of the 1800Ma Lochness Formation near Mt Isa, Australia (Drieseet al. 1995). Although their complex system of chemically
reduced cracks has been described as planar, some of them appear tubular as if modeled around cyanobacterial or other microbial threads. These paleosols show remarkable similarity to Phanerozoic paleosols and modern soils,in which iron oxides are locally reduced by biological activity under intermittently waterlogged conditions (PiPujol & Buurman 1997; Retallack 1997a).
Antiquity of life in soil Evidence for l i e in Precambrian paleosols is growing apace. Five avenues of research have been outlined here: microfossils,trace fossils, soil structure, organic carbon and trace elements. Great strides have been made along each avenue over the past decade.The idea of life on land as far back as 3 500 Ma is no longer outrageous speculation, but an idea amenable to further testing from the fossil record of soils. Indeed, there is isotopic and geochemical evidence for a variety of microbial ecosystems in Precambrian paleosols.
Mother Earth or heart of darkness? Of the many images returned from exploration of the Moon andplanetsperhaps themost evocativewasa blue, cloud-drapedEarthrisingover the barren, rocky horizon of the Moon (Fig. 17.7). Even from afar there is something unique about the Earth compared with other planets and moons. It is more common for planetary soils to be full of salts and for their atmospheres to be mainly carbon dioxide. If theEarth can be imagined without life and at chemical equilibrium with the gases released by volcanism, then it should have an atmosphere primarily of carbon dioxide (99 vol. %) and minor oxygen (lvol.Yo), and an ocean pasty with salts of sodium chloride ( 35 vol. %) and sodium nitrate ( 1.7vol. %). The present atmospheric composition of mainly nitrogen (78vol. Yo) and oxygen (2 1vol. %) with minor argon (1vol. Yo) and carbon dioxide (0.03 vol. %), and its clean oceans of water (96vol.Y0),a little salt (3.5 vol. %) and only traces of sodium nitrate are products of a series of biogeochemical cycles in which life plays an important regulatoryrole. Not onlyare theEarth’ssurfaceenvironments maintained at this cosmically peculiar composition, but also the persistence and evolution of lie over the past 3 500 myr can be taken as evidence for similar
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EarthrisefrorntheMoon(NA .SA )tograph courtesy of R.J. Allenby and PhC Nalional Space ScienceData Center). Fig1ure 17.7
regulation in the past. The idea that life is in control of the Earth and that Earths surface systems are a kind of cybernetic extension of life is the central concept of the Gaia hypothesis (Lovelock 1979: Lenton 1998). Gaiawasmother Earthin Greekmythology. An opposing view can be named for Erebus, the primeval darkness of Greek myth. By the Ereban view, life arose by the most remarkable of accidents only on those planets that had a narrow range of physicochemical conditions to allow it. Life has persisted in a generally hostile environment only where specific activities proved advantageous for scraping a living from available resources. By the Ereban view, the pervasive influence of life is illusory because life still depends ultimately on gases, liquids androcks erupted as a byproduct of the internal differentiation of the Earth, and on the quality and quantity of radiation from the Sun. This thin rind of the biosphere with its tenuous grip on Earth resources could still be destroyed utterly by impact of the Earth with a large extraterrestrial bolide or by full-scale exchange of nuclear weapons. From this perspective,popular views of the origin and early evolution of life (Schopf & Klein 1992: Knoll & Carroll 1999)are distinctly Ereban in character. By this view life evolved by the longest of chances in the primordial soup of the world ocean or in a submarine hot spring, and was continuously buffeted by larger environmental forces, such as scalding temperatures and
devastating meteorite impact. The productivity of the earliest heterotrophic forms of life was limited by availability of abioticallyproduced organic nutrients such as sugars for fermenters. Once autotrophic photosynthetic organisms had evolved, their productivity was limited by ultraviolet radiation in shallow water, by lack of mineral nutrients such as phosphorus in deep marine environments and by predation from earlier evolved autotrophs. Oxygen released by these struggling early photosynthesizers was initially scavenged by reduced iron-bearing minerals and decaying organic carbon compounds. It was not until these sinks for oxygen were buried and photosynthetic productivity increased in more extensive shallow seas of continentalmargins that oxygenation of the atmosphere became noticeable. This early oxygenation mobilized sulfur from soils as sulfate, but it did not accumulate in seawater because it was buried as sulfide reduced by microbes in the deep ocean and assulfates inevaporitesof shallowrestrictedarmsof the sea. Carbon also was buried in organic carbon and carbonates. Nitrogen accumulated in the atmosphere as a weakly reactive pressure-building gas. Its reduced forms (such as ammonia, NH,) were oxidized to nitrate (NO,-) and this was converted to nitrogen gas (N,) by denitrifying microbes that use nitrate to oxidize organic matter. It was only after some time that ozone accumulated in the stratosphere (10-50 km up) and filtered out harmful shortwave radiation so that microbes couldcol-
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onize formerly barren land surfaces. Modern oxidizing photosynthetic ecosystems thus came to conquer formerly hostile habitats while poisoning with oxygen an earlier reducing fermentative ecosystem that lingers on in swamps, stagnant ocean bottoms and the guts of animals. An alternative perspective has been offered here (loosely based on Cairns-Smith 1982) that is rather more Gaian. By this view life arose in a stable, prebiotic sur6cial system that promoted its own persistence. Acidic weathering in an early carbon-dioxide-rich atmosphere promoted the formation of clayey soils. These became more effective at holding the landscape together as organic matter was produced by the abiotic photo-oxidation of iron-bearing clays coupled with reduction of carbon dioxide. In the complex, reactive, alternately wet and dry cracks and cavities of stable clayey soils, organic matter accumulated to produce materials somewhat l i e carbonaceous chondrites. Increasingly complex organic compounds would be preserved as long as they promoted the persistence of the soil against erosion. Organic matter, microbial scums and tropical rainforest can be seen as a continuum of increasingly effective methods of stabilizing the landscape. Oxidation of the atmosphere from small stable areas of productive soils was minimal until the
development of continents. By 3500Ma. life had begun to colonize sea shores. Soils oxidized by organic photosynthesis had long been cleansed of organic carbon that accumulated in marine limestones and black shales, of nitrogen that accumulated in the atmosphere, and of sulfur that accumulated as marine evaporites.Biological reduction of sulfate to sulfides in the ocean began by 2500Ma, and probably earlier. Methanotrophs in soils may be as old as 2800Ma. Photosynthetic ecosystems have always been the source of primary productivity. Methanogenic, fermentative and respiratory ecosystemsare, and always were, dependent on organic matter produced by photosynthesis and have extended the influence of life to habitats where photosynthesis is not possible. Present geological and experimental evidence does not allow a clear choice between these or other scenarios for the early evolution of life. A better understanding of Precambrian paleosols is a promising and barely tested additional approach to the problem. It will not, however, be the whole answer. The Gaia hypothesis has marshalled evidence that life has some homeostatic influence on modern surface processes. Whether this also was true for the distant geological past will be established only by considering all possible facets of Precambrian environments.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 18 Large plants and animals on land
Although land may have been green with photosynthetic microbes well back in geological time, the appearance of multicellular creatures was a much more recent event: by current estimates Middle Cambrian (510Ma). The age and nature of the earliest multicellular soil creatures remain uncertain because their fossil record is so sketchy.The fossil record of early land organisms is largely known from remains preserved in nearshore aquatic environments where amphibious and aquatic fossils could have been preserved mixed with those transported from nearby land. In contrast, fossil soils offer evidenceof conditions and life activities where the creatures actuallylived. The fossil record of marine multicellular organisms extends much further back than that for continental creatures.Metazoan burrows andfecalpelletsinmarine rocks have been reported as old as 2500Ma. Few of these records can be trusted as reliable because most have proven on closer examination to be clasts, gasescape structures or other abiotically produced features (Cloud & Lajoie 1980).Carbonaceouscompressionfossils as old as 2100Ma have the regularity of form and branching that one would expect of authentic multicellular algae (Han & Runnegar 1992: Hofmann 1992). A diverse assemblage of latest Precambrian (600544 Ma) large, soft-bodied creatures has been found in many parts of the world, but such creatures are best known from theEdiacara Hills of SouthAustralia.These fossils were marine multicellular organismswhose morphology and style of preservation are so unique that their taxonomic affinities remain unclear (Retallack 1 9 9 4 ~ )Burrows . of animals appear earlier than (perhaps 1000Ma),and alongside,these curious Ediacaran fossils (Seilacher et al. 1998), but the first tiny hardshelled marine animal fossilsfirst appear at 5 50 Ma,just a few million years before the Precarnbrian-Cambrian boundary. In Cambrian marine rocks, there is fossil evidence for most phyla of invertebrates and many kinds of calcareous algae (Knoll & Carroll 1999).
In contrast to this ancient record of multicellular marine organisms, the well-documented record of megafossil land plants does not begin until much later (Edwards& Selden 1993). Middle Ordovician (Caradoc, 450Ma) fragmentary stems with leaves, stomates and nontracheophytic conducting tubes (Akdalaphyton) from shallow marine rocks of Kazakhstan (Snigirevskayaet al. 1992) may represent the earliest known fossil moss. Of comparable age are enigmatic, thick (0.5-1 cm), isotopically light (613C,, -23 to -2 5%0),carbonaceousholdfasts or stemsin growth position within the uppermost Gull River Limestone near Ingham Mills, New York (Argast 1992). In Lower Silurian (Llandoverian, 430 Ma) rocks in northern Maine, USA, are slender (1-2 mm diameter), branching, carbonaceous tubes in position of growth, called Eohostirnella (Schopf et al. 1966).Also of this age are the earliest vascular plant, Pinnatiramosus, from China (Cai et al. 1996). In middle Silurian (Wenlockian, 425Ma) rocks of Ireland are found slender branches of the vascular plant Cooksonia, with globular terminal sporesacs full of trilete spores (Fig. 18.1A;Edwards et al. 1993).It has been surprisingto find large vascular plants such as the leafy lycopod Baragwanathiu (Fig. 18.1C)in rocks of comparable age in central Victoria, Australia (Late Silurian: Ludlovian, 420Ma). Small (4 mm long) helically twisted lenticular sporangia of latest Silurian age (Pfidolian,418 Ma) in Wales referred to Torticaulis may be a bryophyte. Also of this age, from Podolia, are calcified egg cells referred to Trochiliscus, which appears to have been an aquatic alga similar to the living stoneworts (Chara andNitella).Latest SiluriantoEarly Devonian (PFidolian to Gedinnian) rocks in the British Isles, Europe and North America have yielded remains of sizable (up to 7 cm long), thalli bearing numerous ovoid sporangia (Fig. 18.1B). These remains, called Parka, are superficially similar to the modern aquatic charophyte alga Colochaete, considered the closest relative to modern 263
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Figure 18.1 Some early land plant fossils:(A)the early vascular plant Cooksoniasp.of C. pertoni of Late Silurianage (Pfidolian, 41 8 Ma) inthe Bertie Waterstonefromnear Cedarville,NewYork,USA (photographcourtesyof H.P.Banks, plant described by Banks 1969);(B) enigmaticthalloidplant,Parka decipiensof Early Devonian age (Gedinnian,405 Ma) in the Dundee Formation at TurinHill.Angus, Scotland (RetaUackspecimenP8384A):(C) a drepanophycanlycopodBaragwunathialongifoliaof Early Devonianage (Emsian,395 Ma) in the Wilson Creek Shale,near Alexandra, Victoria, Australia (Retallackspecimen P473 1A). Scalebars represent 1 cm.
bryophytes and vascular land plants among living algae (Kenrick & Crane 1997). Latest Silurian nematophytes, Prototaxites and Nematothallus. may represent fungi or lichens (Retallack 1 9 9 4 ~ )Trunks . of Prototaxitesreachedthesizeof largetrees (1mdiameter, 2 m length). Lichens are now well known in the Lower Devonian (Siegenian, 410 Ma) Rhynie Chert (Taylor et al. 1997). By this time, there was a greater diversity of fossil charophyte algae, nematophytes, thalloid bryophytes and vascular land plants (Kenrick & Crane 1997), and large multicellular plants were well established on land. Another source of information on early land plants is microscopic organic fragments such as conducting tubes, cuticles and spores that can be liberated from silicate rocks by dissolving them in hydrofluoric acid (Gray 1985,1993;Wellman 1996). Fossil liverworts may be represented by thick cuticles and spores permanently united in a tetrahedral configuration (Tetrahedraletes: Fig. 18.2A), now known as old as early Middle Ordovician (Llanvirnian, 465 Ma) in both Libya and Arabia (Edwards et al. 1995a; Taylor 1995; Strother et aJ. 1996), and perhaps Middle Cambrian (510Ma) in Arizona, USA (Strother et al. 1998). Mosses or vascular land plants (stomatophyte clade of Kenrick & Crane 1997) may be represented by the oldest known trilete
spores (Besselia) from the Late Ordovician (Ashgillian, 445 Ma) of Greenland (N~hr-Hansen& Koppelhus 1988). Vascular land plants (tracheophytes) are better represented by diverse trilete spores (e.g. Ambitisporites: Fig. 18.2B) and tubular microfossils with lignin-like biochemicalcomposition in Silurian rocks (middleLlandoverian. 435Ma and later: Niklas 1982; Gray 1993). Other microfossilssimilar to conidia, ascospores and hyphal fragments of terrestrial ascomycete fungi have been found in these Silurian assemblages (SherwoodPike &Gray 1985; Taylor et aJ. 1999). Both fungal decomposers and vascular land plants were diverse by the end of the Silurian period. The fossil record of possible early land animals is more sketchythan that of megafossilplants (Shearetal. 1996; Shear 1998). Curiously, some of the oldest fossil representatives of animal groups that now are completely nonmarine have been found in marine rocks: Early Cambrian (530 Ma) tardigrades (Kmchin 1994) and velvet worms (Ramskold & Chen 1998), Middle Cambrian (5 10Ma) myriapods (see Fig. 10.7), and Middle Ordovician (450Ma) earthworms (Morris et al. 1982). Other potential land dwellers are assemblages of fossil arthropods found in shales and dolostones that lack a normal marine fauna, such as Ordovician-Silurian ( 4 7 0 4 17Ma) eurypterid faunas of Britain and north-
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265
Figure 18.2 Sporesof early landplants:(A)permanent tetrad of Tetrahedraletesgrayue, probably a Iiverwort spore,from the Upper Ordovician(Ashgillian,445 Ma) ElkhornFormation,in Ohio Brush Creek, Adams County,Ohio, USA (courtesyof J. Gray, specimenGI28 5);(B) trilete spore of Arnbitisporitesdilutus, probablyfrom a Cooksoniu-like plant,from the UpperSilurian (Ludlovian,420 Ma),Burgsvik sandstone,nearBurgsvik,island of Gotland.Sweden (courtesyof J. Gray,specimen G923). Scale bars represent 20pm.
eastern USA, Cambrian (51OMa)eurypteridKodymirus from the Czech Republic, and Ordovician (470Ma) horseshoe-crab-like Chasmataspis from Kentucky, USA (Gray 1988). The latest Silurian (418Ma) eurypterid Baltoeurypterus had gill tracts l i e those of amphibious isopods (Selden 1985). Late Pennsylvanian (305Ma) horseshoe crabs such as Euproops may have climbed high into trees whose leaves are mimicked by their spines (Fisher 1979). On the other hand, scorpionsnow exclusively terrestrial are known as fossils in marine and lacustrine deposits as old as Early Silurian (430Ma), but there is no evidence that they had book lungs allowing excursions onto land until Early Devonian (395Ma; Kjellesvig-Waering 1986; Shear et al. 1996).Similarly, some Siluro-Devonian (41 8 4 0 0Ma) fossil millipedes (Karnpecaris:see Fig. 10.7)have peculiarities of limb design and tail shape that are more l i e those of aquatic arthropods (Almond 1985). Undisputed evidence of the earliest terrestrial animals areMiddleOrdovician (Caradocian,450 Ma) myriapod trackways from the Borrowdale Volcanic group of the British Lake district (Johnson et al. 1994). Diverse trace fossils of the fluvial-eolian Tumblagooda Sandstoneof Western Australia (Trewin& McNamara 199 5) are now known to be Late Ordovician in age (Iasky et al. 1998).Body fossils of terrestrial millipedes are found in the Upper Silurian (PFidolian, 418 Ma) Old Red Sandstone of Scotland (Almond 1985; see Fig. 10.7). A
varied fauna of collembolans, spiders and mites is known from the Lower Devonian (Seigenian,410Ma) Rhynie Chert (see Fig. 10.6), both in Scotland. Bristletails now are known as old as Early Devonian (Emsian, 395Ma) and scolopendromorph centipedes as old as Middle Devonian (Givetian,3 75 Ma). Fossil winged insects are not known until mid-Carboniferous time (Namurian, 325Ma), when they appear in the fossil recordin somediversity(Labandeira 1998).Perhapsthe most disconcertingaspect of this sketchy fossil record of terrestrial animals is that major discoveries are still being made. Understanding of early land animals is based on such exceptional6nds and preservationthat it cannot yet be regarded as a representativesample.
Compared with the imperfect fossil record of continental plants and animals, paleosols are preserved in a wider range of environmentalconditions and are much more abundant sources of evidence for the antiquity of multicellular organisms on land. Fossils in paleosols would be the best line of evidence although these require stringent conditionsfor preservation. Trace fossils such as burrows in paleosols also are useful evidence. Many of the earliest land plants and animals were small and would have left subtle traces. But some potentially
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continental creatures, such aseurypterids and Prototax-
ites were large and would have left obvioustraces in pale-
osols if they ever had lived on land. Although few early Paleozoic paleosols have been examined from this point of view, there already is evidence for large plants and animals on land earlier than would have been expected from their megafossil record and more in accord with microfossil evidence. Liverwort-millipede polsterlands Evidence for nonvascular land plants of Late Ordovician age (Caradoc,45 5 Ma) has been found in red paleosolsof the Dunn Point Formation in the Arisaig area, NovaScotia, Canada(BoucotetaL1974;FeakesetaZ.1989).The paleosols consist of about 1.3m of red, calcareous claystone on top of weathered corestones and columnarjointed flows of andesite (Fig. 18.3).Near the surface of one of the paleosols are irregular swales,about 1m wide and 20cm deep, filled with red shale redeposited from erosion of the paleosol. Small white reduction spots within the mounds between these erosional swales are similartodrab mottles formingduring diagenetic altera-
Uscale in rndters =red
claystone
red shale
o El
gray claystone
calcareous nodules
Figure 18.3 Thickredpaleosolwithsuperficialerosionscours and reduction spots,developed on basaltic andesiteflow inthe Upper Ordovician(Caradocian,45 5 Ma)Dunn Point FormationnearArisaig, NovaScotia.Canada (fromBoucot et al. 19 74; reprintedwith permission of the GeologicalSociety of America).
tion of organic matter buried in oxidized paleosols (see Fig. 3.6).Boucotetal. (1974)suggestedthatthemounds were stabilized against erosion by clumps of nonvascular land plants, which lacked rooting structures substantial enough to leave obvious traces. These red, weakly calcareous paleosols and associated mudflows formed on a hilly, tropical, volcanic landscape in a SeasonalIy dry, subhumid climate. Such vegetation of rootless land plants indry soilshas few modern counterparts. It can be visualized as similar to the moss and lichen cover of desert or alpine boulders. This kind of vegetation has been called a polsterland, after a moss colonyorpolster (Retallack 1992a). Animal trace fossils in paleosols have been identified in the Upper Ordovician (Ashgillian, 445 Ma), Juniata Formation near Potters Mills. Pennsylvania, USA (Retallack & Feakes 1987). These burrows are referable to the ichnogenus Scoyenia(Freyet al. 1984).They are so abundant and well preserved that they had hitherto been regarded as evidence for estuarine or marine incursions into these nonmarine rocks. However, this part of Pennsylvania at this time was at least 260km east of the midcontinental seaway and the sedimentary sequence containing the burrows has many characteristics of a purely fluvial sequence. No marine fossils of any kind have been found in association with the burrows. More convincing evidence comes from the close association of the burrows with soil features (Fig, 18.4). Their increased density toward the upper part of the profile parallels the increased development of soil structures (platy peds), microfabric (skelsepic to skelmosepic), the abundance of clay and the degradation of mica and feldspar. The density of burrowing also corresponds to the degree of weathering, including ferruginization, desilication and clay formation, as reconstructed by comparing normalized major oxidesin a moderately developed paleosol with an underlying poorly developed paleosol. Small calcite and dolomitic nodules with the light carbon isotopic signature of soil respiration (Fig. 18.5) ensheath about half the burrows in the moderately developed paleosol, but are not present around less abundant burrows in the weakly developed paleosol. In some cases the caliche cuts across the burrows and in other cases the caliche is cut by the burrows. The nodules are indications that the burrows were a permanent feature of the soil rather than inherited from a pre-existing lacustrine parent
Large plants and animals on [and
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Figure 18.4 Clayey red paleosols (dark recessive-weatheringzones) and sandstone paleochannels in the Upper Ordovician (Ashgillian.445 Ma)JuniataFormation in road cut near Potters Mills, Pennsylvania, USA. The top of the type Potters Mills clay paleosolis at the base of the upper black band onthescalebar(fromRetallack1985:with permission from the Royal Societyof London).Scalein foregroundis calibratedin feet.
material or added after soil formation had been terminated by flooding. The late Ordovician Scoyenia burrows are roughly tubular and about a centimeter in diameter. Their outer margins are marked by a thin zone of ferruginized clay,
I:
molecular weathering ratios weight ~ e g 0 3 percent __ FeO
00small amounts
which is strongly grooved and smeared (slickensided)by compaction. Ovoid masses within some of the burrows resemble fecal pellets (Fig. 18.6B).Other burrows have bilaterally symmetrical backfill structures of alternating silty and clayey bands that are W-shaped in both
stable isotopic analyses calcite
1
’I -5
dolomite
calcite
-
dolomite
zzj:
-10 -5
B
,
i
isotopic composition of Bald Eagle Formation limestone nearby (not in this section) red color
z!tEy
burrows
Figure 18.5 Columnar section (measured in field),molecular weathering ratios and phosphorus abundance (frominductively coupled plasma emission spectrometry), and oxygen and carbon isotopic composition of carbonates and organic matter (8’ 3C,,) of the type Potters Mills clay paleosol (OxicUstropept. above)and the type Faust Flat silty clay paleosol (Fluvent, below) from the Upper OrdovicianJuniataFormation, near Potters Mills, Pennsylvania, USA (data fromRetallack 1993b).
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Figure 18.6 Fossil burrows of the ichnogenus Scouenia from Late Ordovicianpaleosolsnear Potters Mills. Pennsylvania, USA. (A)In hand specimen,cracked open in bas-relief, showing surface striations and a hint of bilateral symmetry: (B) petrographic thin section of a burrow under crossed nicols showing ellipsoidalmasses (definedby black clay skins),interpreted as fecal pellets: (C) petrographic thin section under crossednicols along axis of burrow showing W-shaped backtill structures: (D) petrographic thin section under crossed nicols transverse to long axis of burrow showing bilateral symmetry of backfill structure (from Retallack& Feakes 1987: reprinted with permission from the American Associationfor the Advancement of Science).Scale bars represent 1mm.
longitudinal (Fig. 18.6C) and transverse sections (Fig. 18.6D) of the burrows. The burrows form a complex network of galleries, with horizontal galleries more common near the surface, but blind-ended vertical shafts deep in the profiles. None of the burrows were observed to branch. The burrows range in diameter from 2 to 2 1mm with numerous narrow modes that are parasitic on broader peaks. Such parasitic modes are common in size distributions of fossil animals that grow in marked size increments, such as arthropods (Retallack 1993b). Considering the associatedcaliche,parasitic modes of the size distribution and W-shaped backfill structures, these burrows were excavated by bilaterally symmetrical organisms that grew in well-defined growth increments and were able to withstand dry soil conditions. Thus, they are unlikely to have been made by earthworms or velvet worms. Of other possible early soil
invertebrates, the desiccationresistance and burrowing abilities of eurypterids, aglaspids, horseshoe crabs and scorpions were probably inadequate for the job. Spiders are not found nearly this far back in the geological record and no living representatives of these creatures make burrows likethese.The fossil burrows are most like those of polyzoniid millipedes(Hopkin&Reed 1992),although such millipedesare known as body fossils only as old as Late Silurian (PEidolian,418Ma; Almond 1985. SeealsoFig. 10.7.) These and other early land animals were probably abundant in these Ordovician paleosols, considering evidence for high soil respiration from the near-surface accommodation of atmospheric to soil carbonate carbon isotopic composition (Fig. 18.5; see also Fig. 4.13). Presumably this high secondary productivity was supported by a substantial primary productivity of liverwort-like plants lacking roots that would Leave obvious
Large plants and animals on land
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traces. Thus, these soils could have supported a polsterland similar to that envisagedfor the Late Ordovician paleosolof NovaScotia (Fig. 18.3). Paleosols of the Juniata Formation in Pennsylvania, USA, formed on alluvial outwash of theTaconic Mountains to the east (Fig. 18.7) at a time of very low global sea level because of glaciation in the present region of the Sahara Desert of Africa (Brenchley et al. 1994). The climate in Pennsylvania at this time appears to have been tropical, considering the diversity of North American marine faunas and their paleolatitude. Rainfall was in the semiarid to arid range and perhaps seasonal if the depth of caliche in the paleosols and its ferruginous bandscanbetakenasaguide(Retal1ack1993b).
vegetation was probably similar to modern desert vegetation of resurrection plant (Selaginella lepidophylla) or alpine and open woodland carpets of clubmoss (Lycopodium deuterodensum). This kind of plant formation has been called brakeland, following the common English expression ‘fernbrake’ (Retallack 1992a). Animal burrows remain common in Silurian and Devonian paleosols (Fig. 18.8).The diversity of land animals by this time is indicated by surprisingly abundant and diverse predatory animals such as the spider-like trigonotarbids (seeFig. 10.6). Coprolitesfullof vascular plant spores are evidence of plant-animal interactions in these early terrestrial communities (Edwards et al. 1995b).
Tracheophyte-trigonotarbidbrakelands
Salt marsh and marsh
The difference between early vascular land plants such as rhyniophytes and nonvascular land plants such as nematophytes is not always easy to determine because they can beexternally similar andlack the large branching root traces of modern vascular plants (Edwards & Edwards 1986).Making thisdistinction isevenmoredifficult from the indistinct traces left by plants in welldrained paleosols. For example, paleosols in the Upper Silurian (Ludlovian, 420 Ma) Bloomsburg Formation near Palmerton, Pennsylvania, USA (Retallack 1985; Driese et al. 1992)are generally similar to those already described from Late Ordovician rocks, but are thicker and better developed (Fig. 18.8).They have distinct bioturbation of their surface horizons: small (1-5mm across), irregular, subhorizontal to vertical, branching tubular features filled with clay (ortho-isotubules) or sand (metagranotubules) different from the surrounding matrix (Fig. 18.9). They are more irregular than tubular burrows and have bothY and H branching patterns like those of the rhizomes of early vascular land plants(Kenrick&Crane 1997). Bioturbation like that of rhizomes and roots also has been documented from latest Silurian-Devonian (418-395 Ma) paleosols in Canada (Boucot et al. 1974; Elicketal. 1998)andBritain(Allen1986a).Bythistime vascular land plants were abundant and widespread. Of the kinds of megafossil plants found in associated deposits, xeromorphic rhyniophytes more probably vegetated these dry soils than thallose liverworts or nematophytes. This dry land, herbaceous, rhizomatous
Early land plants found in wetland paleosols include Eohostimella heathana from the Lower Silurian (Llandoverian 430 Ma) Frenchville Formation near Stockholm, Maine, USA (Schopf et al. 1966). These fossil plants are small (1-2 mm diameter), erect, dichotomously branching, carbonaceous tubes in a dark, finely laminated shale that elsewhere contains nearshore marine fossils. A case has been made that these carbonaceous tubes were organically lined worm burrows (Strother & Lenk 1983). However, their coalified outer layer appears cellular and spinose in places, and its organic chemical composition is more like that of other Silurian and Devonian vascular land plants than that of fossil marine algae (Niklas 1982).The interior of these fossils is not sufficientlywell preserved to reveal whether they were vascular plants, nematophytes or some other group. The laminated, carbonaceous paleosol was little developed and waterlogged (Aquent) and is in a sequence that has yielded marine fossils. This may have been an ancient example of salt marsh vegetation. Another example is the Cooksonia-trigonotarbid community at thelatest Silurian (Piidolian, 418 Ma) locality of LudfordLane,Shropshire, England (Jerametal. 1990; Edwards& Selden 1993 ) . Vascular plants formed marsh soilsby Early Devonian time (Siegenian, 410Ma) as represented by the Rhynie Chert in Scotland (Rice et al. 1995). This silicified peat contains exquisitely preserved remains of various vascular and nonvascular plants, as well as fossil arthropods (see Fig. 10.6). Alternating layers dominated by
271
Large plants and animalson land A. POTTERS MILLS, PENNSYLVANIA (LATE ORDOVICIAN)
B. PALMERTON, PENNSYLVANIA (LATE SILURIAN)
calcareousness hue develom It
calcareousness develoDment hue
!SS
hue
Ir I L >
0 2
m
-1
85
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80
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a c i a y s t o n e breccia m s i l r s l sandstone
aclaystone
~~~~" @ scour and fill
planar bedding
-burrows
0
erosional surface I I I rhizome traces
l??ii root traces drab-haloedroot traces
IZZ3 calcareous nodule 1
red color
Figure 18.8 Columnar sections (measured in field)showing development, calcareousness and hue of paleosols in the (A) Upper Ordovician(445 Ma)JuniataFormation, at Potters Mills, Pennsylvania, USA (B) Upper Silurian (420 Ma) BloomsburgFormation, at Palmerton, Pennsylvania: and (C) MiddleDevonian(375 Ma) OneontaFormation near East Windham, NewYork.
Figure 18.9 Rhizome traces in an Upper Silurian (Ludlovian 420 Ma) paleosol (Faust Flat pedotype) in the BIoomsburg Formation at Palmerton, Pennsylvania, USA, showingY branching(dichotomous. A) andH branching (pseudomonopodial,B) typicalof early vascular land plant rhizomes. Scales graduated in millimeters (GomRetallack 1992d;with permission from theSocietyfor Sedimentary Geology).
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Horneophyton lignieri and by Aglaophyton major may represent successive episodes of flooded marsh followed by better-drained marsh. An ancient bog ecosystem and Histosolis represented by the 1.4m thick Barzass Coal of Middle Devonian (Eifelian, 385Ma) age near Barzass. Siberia. This is a cuticle coal of the enigmatic early liverwort-likenonvascularplant Orestoviadevonica(Krassilov1981). Atiietable Despite the paucity of pertinent studies, a broad timetable is apparent from paleosols for the advent of various kinds of large plants and animals on land. Microbial earths were probably widespread for much of Precambrian time. Paleosols of livenvort-millipede polsterlands appear by late Ordovician time (Caradocian, 450Ma). Salt marsh vegetation of nonvascular land plants is known as old as Early Silurian (Llandoverian, 430 Ma). Tracheophyt&rigonotarbid brakelands appear in well-drained soils by Late Silurian (Ludlovian, 420Ma). Root-like structures are more widespread in paleosols of latest Silurian (Pfidolian, 41 8 Ma) and Devonian age. Primitive vascular and nonvascular plants of freshwater marsh and bog are known from silicified peat of Early Devonian age (Seigenian,410 Ma). The fossil record of soils thus supports evidence from fossils for Middle Ordovician appearance of land plants. However, the fossil record of burrows in Ordovician paleosols and myriapod tracks predates Silurian fossils of land animals. On the other hand, records of fossil vascular plants are known as both spores and megafossilsearlierthan they can beclearly documentedfrompaleosols. With such incomplete fossil records of soils and early land biota, there are bound to be more surprises.
How did multicellular land organisms arise? Establishing exactly when various groups of multicellular organisms came to live on land has proven difficult. Investigating how it happened is even more forbidding. The diversity and adaptive features of early continental fossil assemblages provide some clues (Beerbower 1985), which now can be assessed against the fossil record of soils. Several separate problems can be identified. What were the evolutionary origins of the earliest
land plants and animals? To what extent were their life styles constrained by environmental factors such as nutrient availability, atmospheric oxygenation, flooding,volcanism or changes in sea level? Origin of Iand plants The closest living relatives of land plants are charophytes (Kenrick & Crane 1997). Most of these living green algae are aquatic. From this and the geologically ancient fossil record of other kinds of algae in marine rocks, it is widely assumed that plants invaded the land from lakesor rivers.They are seen as invading the land as if it were some kind of military campaign for territory. Stepsin the invasion can be envisaged as a technological escalation of military hardware: the acquisition of (a) aerially dispersed spores, coated in rigid, desiccationresistant, sporopollenin walls, (b) a thick desiccationresistant cuticle, (c) tracheids for support and water transport, and (d) intercellular air spaces and stomates for gas exchange. The theoretical difficulties for large aquatic algae evolving such structures have prompted an alternative view that land plants evolved from unicellular or very small multicellular soil algae (Stebbins & Hill 198O).Bythisviewtheparaphernaliathat isneeded for plants to survive on land could have evolved entirely within soils rather than within environments intermediate between water and land such as tidal flats and streamsides. The idea that land plants evolved in soil is strengthened by evidence of complex microbial communities in soils well back into Precambrian time (Horodyski & Knauth 1994; Rye 1998). Additional evidence is the currently earlier record of plant life in well-drained soils (Late Ordovician, 450Ma) than in salt marsh (Early Silurian, 430) or freshwater marsh (Early Devonian, 410 Ma). The limited records available can hardly be regarded as representative, but they do point the way to a source of new information on this question. Indeed, these new discoveries also support some aspects of the traditional view of invasion of the land from the sea or lakes. Difficulties of conquest would have been mitigated by burrows in soils and by pre-existing microbial ecosystems. Microbes would have stabilized the landscape and initiated cycles of nutrient utilization in soils. Burrows could have provided moist and sheltered local habitats. In modern deserts, the burrows of
Large plants and animals on land rodents may be small semi-autonomous communities including algae, fungi, and herbivorous and dung beetles, all protected within the cool moist burrow from a harsh external environment (Halffter & Matthews 1966).The early evolution of vascular land plants in intertidal habitats also is supported to some extent by their possible occurrence in marine rocks geologically older (Llandoverian,430Ma) than evidence for them in paleosols (Ludlovian, 420Ma). The habitats where early land plants evolved will be better understood with further studies of paleosols and theirroot traces. Origin of land animals The early evolution of land animals is a separate problem.There has been littledoubt that arthropods invaded the land frommarine habitats, becauseof theirlongevolutionary history in marine rocks as old as basal Cambrian. In many ways marine arthropods were preadapted for lie on land. Woodlice and crabs are primarily marine arthropod groups that have independently exploited land habitats (Vermeij 1987).Their exoskeleton is effective for support and protection both on land and in the sea. On land it also reduces desiccation. Their jointed limbs allow locomotion, burrowing and predatory feeding both on land and in the sea. Such mobility also allowscomplexmating rituals. The main evolutionary innovations of land arthropods were methods of respiring in air without drying out, such as book lungs and tracheary systems, and methods of liquifying solid food in a preoral chamber (Stmmer 1 97 7). It is possible, on the other hand, that some of the enigmatic Cambrian arthropods found in unusual marginal marine facies, and perhaps even Middle Cambrian fossil velvet worms in deep marine rocks, were washed in or invaded from lakes or land. Indeed, a number of living insects, such as skaters and caddisflies, have apparently invaded marine from fresh waters (Vermeij1987). One could also imagine that the earliest land animals were tiny, interstitial creatures like living tardigrades (Kinchin 1994). Current evidencefrom paleosolsis not adequate to address this issue,but does point to a significantnew line of enquiry. In the Upper Ordovician Juniata Formation near Potters Mills and the Upper Silurian Bloomsburg Formation near Palmerton, both in Pennsylvania, USA, burrowed paleosols are found in inland fluvial facies near the upper part of these formations, as well as in
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near-marine and intertidal paleosols in the lower part of these formations. Drab, shaley deposits of inland lakes have been identified in both sequences, but these so far have proven devoid of trace fossils or other indications of lie (Retallack1 98 5).Avariety of trace fossils of very diferent kinds have been reported from overlying and underlying marine rocks (Friele & Baldwin 1988; Friele et al. 1988:Driese&Foreman1992).Thereis thusnoevidence of comparable trace-makers that might have invaded these paleosols from apparently oligotrophic lakes, or from biotically diverse tropical seas. Whether thiswilldemonstrateinvasionof thelandfromtheseaor lakes remains to be seen from more systematic studies of suchearly Paleozoicichnofacies (Buatoiset al. 1998). Ecology and environments of early Paleozoic life on land The earliest land plants and animals are usually imagined as conservative in their use of water and nutrients, slow to grow and reproduce, and in other ways making do in a generally hostile environment (Edwards& Selden 1993). This view of early fossil land plants is based in part on their apparently low diversity, slow rates of evolution, simple structure and comparison with modern plants with broadenvironmental tolerances, such asliverworts. Slow-moving detritivorous-herbivorous animals, such as millipedes, are expected to be the earliest kind of land animals according to this stress-tolerant view. However, alternative life styles are known in modern plants (Grime 1979). In frequently disturbed parts of the landscape, for example, it may be advantageous to produce copious small and widely dispersed propagules rather than persist in one place or grow to a large size. Such a life style found in modern plants regarded as weeds also could be argued for early land plants given their small, smooth-walled, permanent tetrads dispersed by wind and water, and capable of forming several small reproductive plantlets (gametophytes) that could fertilize each other if no others were nearby The small size of many early continental animal fossils also supports the idea that early land animals were small and rapidly breeding, like many modern creatures regarded as pests. A third life style has been called competitive because it emphasizes the growth of a large body that commandeers resources at the expense of other organisms.This view also could be supported for early land
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plants if large nematophytes (such as Prototaxites) or large eurypterids, scorpions and trigonotarbids played a role in the earliest continental ecosystems.These three general kinds of life styles are responses to nutrientpoor, to frequently disturbed and to nutrient-rich environments, respectively. Evidence of these selection pressures can come from paleosols. One difficulty that could have been faced by early creatures on land was a n atmosphere rich in carbon dioxide with little oxygen and lacking an ozone shield that filtered out strong ultraviolet radiation. Shortwavelength (220-290 nm) ultraviolet radiation could disrupt nucleic acids and proteins if there were no ozone screen, as would be likely at an oxygen level in the atmosphere of less than 0.01 times the present level (Kasting 1987). About 0.02 times the present level of oxygen is needed for the production of cutin in cuticles that protect against desiccation,of phenylpropanoids that act as a screen against ultraviolet radiation, and of lignin as a basic material for structures of water transport and support in plants (Chapman 1985).The atmospheric track of oxygenation since Precambrian time is not constrained by studies of as many paleosols as one would like (see Fig. 16.7),but oxygen probably exceeded 0.02 times present level by Late Precambrian time, if not earlier (Rye & Holland 1998). These preliminary impressions do not favor the idea that strong ultraviolet radiation or low oxygen levels limited the growth of Ordovician land animals or the evolution of chemicals critical to key adaptations in plant life on land. Obtaining water and nutrients are closely related problems for plants. Some nutrients (especiallyN, P and S) can be used only when supplied by microbes as ions in solution (Stevenson1986).Animalsobtainthembyeating plants or other animals. Fossil fungal hyphae and spores as old as Early Silurian (Sherwood-Pike & Gray 1985) show that fungal microbial nutrient procurement systems are of considerableantiquity. Evidence for life in Precambrian paleosols (Horodyski & Knauth 1994; Rye 1998) that have nevertheless remained organic lean is suggestive of even more ancient microbial decompositional systems, including, for example, actinobacteria. Other nutrients (Mg,K, Ca,Na andC1)are obtained by mineral weathering. The water and nutrient resources of Late Ordovician paleosols in the Juniata Formation of Pennsylvania can be reassessed from this perspective.
The depth of leachingof carbonate in these paleosols indicates subhumid conditions, probably with a severely dry season, as in present-day regions of subtropical wooded grassland (Retallack 1993b). The carbonate also is evidence of abundant calcium that could be used in the exoskeletons of arthropods. Detailed chemical studies of one of the profiles has shown a strong surficial enrichment of potassium. This and petrographic observations on early Paleozoic sandstones (Basu 1981) could be taken as an indication that this element was underutilized, as is typical under modern fungi and liverworts (Shacklette 1965). However, the stoichiometrically inordinate abundance of potassium is a clue that its enrichment was in part diagenetic (Feakes&Retallack1988).Someotherelements(Ba,Cr, Sc, V, Zn and Ti) increase in abundance toward the surface, presumably because of their well-known affinity for clay, which was more abundant there. However, others (Li, Nb, Ni, Sr, Y and P) that commonly follow clay and organic matter within soils are surprisingly depleted at the surface. This may reflect loss within vegetation that was not intimately mixed with the soil and left little humus. The clay content,degree of weathering and evidence for nutrient recycling in this paleosol are greater than in many modern desert soils, but less than in soils of wooded grassland. Secondary productivity and soil respiration comparable with such modern ecosystems also is indicated by the abundance of burrows and distribution of carbon isotopes in paleosol carbonate (Fig. 18.5, see also Fig. 4.13). With plentiful potential soil nutrition, moderate water availabilityand lack of such adverse soil conditions as salts or duripans, conditions for growth were not as grim as they could have been. Additional possible limitations for early plants and animals were environmental perturbations. Floods, dry seasons, volcanic eruptions and marine transgressions all can be imagined as destabilizingthe establishment of early continental ecosystems. Such environmental disturbance leaves arecord in the degree of development of paleosols in long sequences. In the Upper OrdovicianJuniata Formation paleosols are sparse in the sequence and for the most part are weakly developed,each representing only a few hundred years of soil formation at most. Caliche-bearing paleosols are less common and represent parts of the landscape that were stable for several thousand years. The presence of ferruginized
Large plants and animals on land caliche concretions indicates a dry season. The dominance of sandstone in thick tabular beds with only local evidenceof asymmetric channel-lie forms is an indication that the JuniataFormation was deposited in braided to loosely sinuous bedload-streams, which may have had a somewhat flashy flood regime (Cotter 1978). Although such a disturbance is more frequent than in ecosystems of humid floodplains, it is not as severe as that found in desert regions. Volcanism also is unlikely to have been a critical limitation considering the evidence for nonvascular plants and thick red calcareous paleosols on andesitic flows of the Upper Ordovician Dunn Point Formation of Nova Scotia (Fig. 18.3).Dramatic sea-level fluctuations also are unlikely to have restrained continental ecosystems. Latest Ordovician glaciation was accompanied by changes of sea level (Brenchley et al. 1994), yet paleosols and microfossilsare evidence for similar continental ecosystemsboth before and after glaciation. In summary, there is evidence from paleosols that Ordovician livenvort-millipede polsterlands were neither as weedy nor as stress tolerant as they may seem by comparison withmodern vegetation.Indeed, in the context of their times, they may have been competitivecommunities compared with coeval microbial earths and rocklands. The mat-penetrating axes of mosses and Cooksonia, and woody logs of Prototaxites came later, and formed even more massive competitive communities, although still dwarfed by modern forests. Similarly, the animals that formed burrows in Late Ordovician paleosols were detritivorous-herbivorous millipedes. Large arachnid carnivores came later in the Silurian. Both animals and plants of soil ecosystems become increasingly specializedthrough time in their competitive adaptations such as body size, in a form of evolutionary escalation (Vermeij 1987).Evidencefrompaleosolsthus supports the notion that terrestrial food chains were built from the ground up: from producers to consumers and from microbes to monsters.
Biological innovation or environmentalregulation? There are two distinct views of the early evolution of continental ecosystems. By the traditional view, oceans were teeming with multicellular life, but conquest of the landwas held at bay by hostile conditions there (Vermeij
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1987). When ultraviolet radiation, flood frequency and moisture variation of soils were mitigated, then pre-adaptedorganisms invaded the land fromlacustrine and marginal marine habitats. Once unleashed from environmental constraints, evolutionary processes filled this new adaptive zone. This view, that I have characterized as Ereban, is similar in some ways to that idea that life itself evolved by chance and diversified only under the constraints of atmospheric and continental growth. A contrasting view based on the Gaia hypothesis (Lovelock 1979;Lenton 1998)isthatliearoseaspartof a self-sustaining system that continued to incorporate evolutionary innovations. By this Gaian view, evolutionary progress was determined not so much by environmental changes as by biological innovations that ultimately created new opportunities. These biological innovations may have been slow to appear because they were technically difficult. Even if they did appear, the biotically modified environment that would select for them might not have been at hand. Land habitats suitably prepared by microbial communities for the colonization by multicellular organisms may have been present in some places and suitably adapted multicellular plants in other places, long before the two happened together to allow exploitation of the land. By the Gaian view, plants and animals arose according to a schedule determined by the pace of the co-evolution rather than by the tempo of environmental change. Although a Gaian or Ereban view of life is most critical in interpreting its earliest evolution, the two views are more effectively assessed from Phanerozoic evidence because that is where the rock and fossil record is least speculative. Perhaps the most conclusive test is the extent to which life, and particularly humanity, controls various surficial cycles of volatile elements today, such as the carbon cycle (Mooneyet af. 1987; Vitousek et al. 1997b). By a n extreme Gaian view biotic control is near total and has been for some time. By an extreme Ereban view it is mfnimal and whatever slight degree of control that can be demonstrated is of geologically recent vintage. There is a range of intermediate positions that seem more reasonable (Fig. 18.10),and more like existingrecords(seeFigs 16.2 and 16,7).Acasecanbemade that life exerted considerable control on atmospheric composition, temperature and other aspects of the environment well back into the geological past (Lovelock
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Chapter 18 4
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near total
DEGREE OF CONTROL
Figure 18.10 A theoreticalconsiderationof extremeGaian and Ereban views of the history of 1ieonEarthover thepast4500
minimal
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3 THOUSANDS OF MILLIONS
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OF YEARS BEFORE PRESENT
1979; Lenton 1998). It is doubtful if this was a total hegemony, however, because there is an observable Phanerozoic consolidation of control by land plants and animals(Fig. 18.11). Blanketing bogs The increased biomass of continental ecosystems is most obvious from waterlogged paleosols where a portion of that productivity was preserved as a new kind of soil:the peaty soilsorHistosols (Fig. 18.11D).The oldest known Histosol is the Lower Devonian (Seigenian, 410 Ma) Rhynie Chert of Scotland (Riceet al. 1995). Only a handful of such permineralized peats of different geological ages are known (Knoll 1985). More typically, ancient Histosols are not silicified. Coaly layers of large nonvascular plants are found in Silurian (Llandoverian toLud1ovian)rocksoftheeasternUSA(Pratt etal. 1978; Strother & Traverse 1979). These do not have the organic matter content or thickness required of Histosols. The most ancient and spectacular nonsilicified Histosol is the 1.4-m-thick Barzass Coal of Middle Devonian age (Eifelian, 38 5 Ma) near Barzass village, Siberia (Krassilov 198 1).Woody coals are not found until much later in the Devonian period. The earliest eiarnple may be a thin (10cm after compaction) coal overlying claystone withlarge root traces intheupperDevonian (Fammenian, 360Ma) Hampshire Formation near Elkiis, West Virginia, USA (Gillespie et al. 1981). More abundant and thicker woody coals are found in rocks of middle Carboniferous and younger geological age (Retallacket al. 1996).
myr.
A tangle of plant bodies in waterlogged terrain would have done much to stabilize them against floods,storms and other perturbations. It also restricted the diffusion of atmospheric gases into stagnant soil water below, thus promoting preservation of organic matter there. The burial of reduced carbon in organic matter results in the release of oxygen to the atmosphere that otherwise might have been used to fuel its decay (Stallard 1998). More oxygen is released in coastal swamps where marine sulfate is reduced to pyrite and other sulfides by microbes. This also prevents the accumulation of toxic sulfur gases in the atmosphere. Denitrificationof nitrate (NO,’-) to molecular nitrogen (N,) is also a microbially mediated phenomenon of reducing swampyenvironments (Stevenson 1986).Theengineof atmospheric oxygenation and nitrogenation that had been set in motion by Precambrian microbial communities thus received a considerableboost from the advent of stable and productive swamp ecosystems. Taming streams Several elements of fluvial deposits can be taken as evidence that life was taking control of some parts of the landscape through early Paleozoic time. There is an increased abundance of biogenic structures compared with purely sedimentary ones, an increase in the amount of clay compared with sand, and an increased complexity of bedding both laterally and vertically (Fig. 18.8). Platy peds in Late Ordovician paleosols are a coarse kind of structure compared with those in geologically
Large plants and animals on land
$? 30-8.SOIL
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BIOTURBATION
a,
U
C. ATMOSPHERIC CO, sediment mass
E
Figure 18.11 EarlyPaleozoicchangesin(A) soil differentiationas indicated by clay content (volumepercentage) and alumina/bases (molecular ratio) of most weathered horizon of calcareous red paleosols; (B) soilbioturbation as indicated by proportion of line transect in paleosols occupied by roots or burrows (x)and by measured rooting depth (cm);(C) atmospheric carbon dioxide levels (PAL) calculated from a sedimentary mass-balance model (curve) and estimated from carbon isotopiccompositionof pedogenic carbonate (blackboxes):(D) maximum coal seam thickness and average thickness of at least 10 consecutiveseams (m);(E) diameter of fossil plant stemsandroots (m); (F) diversityof fossil plants, including both woody and herbaceousclades (number of species);(G) diversityof soil animals (number of families) (fromRetallack 1997e: with permission from the American Association for the Advancement of Science).
'3 10 E
E. PLANT SIZE trunks U
?1
F. PLANT SPECIES first land vascular 100 plants plants
$200 first
c'
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first
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youngerpaleosols.Late Silurian and Devonian paleosols in contrast are more massive and hackly in appearance with abundant and diverse burrows, root and rhizome traces, calcareous and ferruginous nodules and a variety of ped structures (Fig. 18.11B). The Psammosteus Limestone of the Lower Devonian (Gedinnian, 41 5 Ma) Old Red Sandstone in south-western Britain (Allen 198 6a) is a sequence of paleosols each so strongly bioturbated and thick that there is little sign of sedimentary structure. A greater biomass of plants and animals also promoted clay formation in soils (Fig. 18.11A). By stabiliiing soils against erosion, primary minerals are unlikely to escape hydrolytic weathering. The amount of clay in fluvial deposits is related to many factors: the amount of shale or topographic relief in source regions, the rate of sediment accumulation, the amount of rainfall, and the stabilizingeffect of vegetation. If it were possible to control these variables, then differences in clayeyness of fluvial sequences of different geological age might be apparent. This can be attempted in a crude way by comparing Late Ordovician, Late Silurian and Middle Devonian sequences of paleosols all of which formed in alluvial outwash of major mountain ranges of quartzofeldspathic composition in subhumid, seasonally dry paleoclimates (Fig. 18.8). The increased clayeyness through time evident from this comparison is compatible with the increased amount of total clay in the sedimentary rock record through geological time (seeFig. 16.2A). Equally likely is a change in the nature of clays throughgeological time. Highlyweatheredclayssuchas kaolinite should be more abundant after the advent of land plants, which would have depleted nutrient cations more effectively than before (Knoll & James 1987). Increased abundance of smectite at the expense of illite might also be expected after the advent of vascular land plants, because they require more potassium than nonvascular plants. Both kaolinite and smectite are more common after Devonian time and rare before that (see Fig. 7.7). Furthermore, the degree of weathering indicated by alumina enrichment and base depletion increased through the Early Paleozoic in red calcareous paleosols of comparable paleoenvironments (Fig. 18.11A). This trend also is seen from the lack of weathering of potassium feldspars in Paleozoic sandstones (Basu 1981). Potassium enrich-
ment by illitiiation and sericitization of clays during burial (Weaver 1989) has done much to obscure this trend. Increased lateral complexity of bedding in alluvial sequences also accompanied the advent of land plants and animals. Early Paleozoic channels were laterally extensive, broadly symmetrical in cross-section and scoured shallowly into underlying deposits (Cotter 1978). similar to streamchannels indesert regions at present. Such braided streams are choked with more sand and gravel than they can effectively transport and this is deposited as islands in the channel. By mid-Paleozoic time many channels became laterally restricted and asymmetric, with a cut bank on one side scouring deeply into underlying deposits and a conformable sequence of interbedded sandstone and shale on the other side (Fig. 18.8C). As in channels of meandering streams in wellvegetated, humid regions, the lateral spread of these streams is checked by stable, clayey, vegetated banks so that they carry mainly a suspended load of fine-grained materials. A number of loosely sinuous, intermediate stream patterns are known (Schumm 1981). These stream patterns also are dependent on a variety of factors other than vegetation. Meandering stream deposits are known even in Precambrian rocks of 2900Ma, where they appear to have incised into prior-deposited flat-lyingshales(Button8tTyler1979;agefromde Wit et al. 1992).Inmy exampleof alluvialsequenceschosen to keep various other environmental variables constant (Fig. 18.8) there is a transition from Late Ordovician braided and loosely sinuous streams to Late Silurian loosely sinuous streams, then to Middle Devonian meandering streams. Each of these aspects of alluvial architecture reflects increased soil stabilization by plants. Floods are contained not only by the way in which plants hold stream banks together but also by the way in which plants absorb and use rainwater within the catchment area of streams (Schumm 1 977). In well-vegetated regions, streams flow permanently rather than intermittently. Stabilization of stream banks against erosion and increased water availability allow deeper weathering of soils to obtain essential nutrients such as phosphorus and potassium. This positive feedback also is reflected in increased depth and degree of bioturbation, clayeyness and degree of clay weathering through time (Fig. 18.11A and B).
Large plants and animals on land Filtering air Increased primary production and biomass of plant materials on land would also have had consequences for atmospheric gases, such as 0, liberated by photosynthesis and carbon dioxide from soil respiration. Atmospheric carbon dioxide is consumed not only by plants, but also by weathering, because it forms carbonic acid in soil solution that liberates nutrient cations by hydrolysis from minerals such as feldspar, and then is lost to the ocean in solution as bicarbonate (HC0,-'). The increased base depletion through time seen in Paleozoic paleosols (Fig. 18.11A) thus fueled increased plant biomass (Fig. 18.11E).A general drawdown of an Early Paleozoic carbon dioxide greenhouse is also apparent from the carbon isotopic composition of carbonate nodules in paleosols (Fig. 18.1l C ) , which becomes more l i e the isotopic composition of organic matter in the same paleosols through time (isotopicallylighter, or more negative) as this source of carbon came to dominate the isotopically distinct (heavier or less negative) carbon dioxide of the atmosphere (Mora et al. 1991, 199 6;Mora & Driese 1999). In late Ordovician paleosols there is a strong gradient of accommodation between the atmospheric and soil isotopiccomposition within paleosol profiles,indicating high soilrespiration (see Fig. 4.13;Yapp&Poths 1994). With such abundant animal life and yet such meager plant life that no root traces remain (Fig. 18.7),it is perhaps not surprising that the Late Ordovician was a time of maximal greenhouse (see Fig. 9.15).
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By Devonian time, the tables were turned, as root traces and drab haloes after buried organic matter in paleosols become much more abundant than burrows and other traces of soil animals (Fig. 18.11B). The dominance of plants and primary production is aIso evident from the relatively light (more negative) carbon isotopic composition (613C,,,) of carbonate in Devonian paleosols. By Middle Devonian time there also were large trees and thin coals.The drawdown of atmospheric carbon dioxide has been attributed to the midPaleozoic spread of forests and swamps (Algeo et al. 1995; Algeo & Scheckler 1998). However, fossil trees and peats appear belatedly after more than half of the carbon dioxide greenhouse has been mitigated (see Fig. 9.15). Volcanic eruptions and tectonic activity do not appear to have caused this carbon dioxide drawdown because both increased through geological time and so contributed more, not less carbon dioxide (Bluth&Kump 199 1).Thus much of the carbon dioxide drawdown may have been created by weathering and growth of early land plants during Silurian time (Figs 18.8 and 18.9). Although neither woody nor bulky, these plants, by means of chemical defenses, including those inherent in lignin production, were able to dominate animal activity in soils (Retallack 1997e). It is likely that much of this stabilizingsystem was set in motion within soils by microbes during Precambrian and Cambrian time. With the advent of large land plants it gained momentum and direction that was increasingly difficultto reverse.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 19 Afforestation of the land
The evolution of forests was in some ways a continuation of evolutionary processes set in motion with the advent of vascular land plants. Their appearance heralded marked advances in weathering and stability of soils and landscapes. With trees appeared the various kinds of soils with leached, near-surface (E) and clayey and ferruginized subsurface (Bt and Bs) horizons: Alfisols, Ultisols and Spodosols.This increased diversity of soils paralleled a diversification of life into the varied habitatsfoundwithinforests(seeFig.18.11). The fossil record of trees extends back at least to Middle Devonian time (Givetian, 375 Ma), and fossil trunks referred to Callixyfon (Driese et al. 1997; MeyerBerthaud et al. 1999). With characteristic paleobotanical caution against the possibility of wrongfully assigning parts of different plants to the same species, the fossil leaves (Archaeopteris)and spores (Gerninospora) now known to havebeenattached tothe trunks wereinitially given different botanical names (Fig. 19.1). This combination was surprising botanically. Archaeopteris belongs to an extinct group of plants, the progymnosperms, that had wood anatomy similar to that found in modern conifers, but reproduced by means of spores as do modern ferns.The spores of progymnosperms, like those of ferns and most primitive land plants, probably produced small, delicate reproductive plantlets (gametophytes) on which reproductive gametes met in thin films of water. Gametophytes require a moist place or season for effective reproduction. For this reason ferns are most abundant in habitats that are at least periodically moist. Trees with other kinds of trunk construction and reproductive capabilities also appeared during the Devonian period (Gensel& Andrews 1984). Extinct lycopsids included large trees (Eosperrnatopteriserianum) asearlyasMiddleDevonian (Givetian.3 75 Ma), andtree horsetails (Pseudobornea ursina) date back to the Late Devonian (Frasnian, 3 70 Ma; Algeo & SchecMer 199 8). Well-known tree horsetails and lycopsids of Carbonifer280
ous age were more akin to overgrown herbs than modern trees. They had a weak cylinder of wood with much soft-walledtissue (parenchyma) scattered among the wood cells (xylem), and grew into a rigidly determinate structure, without the haphazard appearance of manymodernplants(Eggert 1961,1962). Otherweakstemmed early trees such as the Middle Devonian (Givetian)cladoxyl fern Pseudosporochnus nodosus had isolated strips of radially arranged woody tissue in the trunk rather than a solid cylinder of wood (Leclerq & Banks 1962). Tree ferns such as AustrocZepsis australis as ancient as mid-Carboniferous ( 315 Ma) had false trunks, with small stems bolstered by abundant leaf bases and adventitious roots (Sahni 1932; Morgan 1959). All these early trees were spore-bearing plants that reproduced by means of water-dependent gametophytes. The earliest seed plants, in which the gametophyte was protected from desiccation by enclosure in seed coats, date back to latest Devonian time (Fammenian, 360Ma: Rothwell et aZ. 1989). It is likely that the earliest fossil seeds were borne on shrubs, but seed ferns such as Pitusprirnaeva were large trees by earliest Carboniferous (Tournaisian, 350 Ma: Retallack & Dilcher 1988). There are a number of plausible explanations for the evolution of trees (Fowler & Schindwein 1991). First, they may have grown tall to shade out neighboring plants, because light is a plant's chief source of energy. Second, tall stature could have discouraged competing plants by showering them with mild poisons. Flavonoids are used in this way by many modern trees and the ability to synthesize these substances is widespread among vascular plants, bryophytes and even some aquatic charophyte algae (Chapman 198 5).Third, trees could have evolved to deter herbivory, and this would have been especially effective before the evolution of winged insects. The ability to digest woody tissue has evolved in very few groups of animals, such as termites. Fourth, the evolution of trees allowed the broader
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scatter of spores from greater heights. These are all Gaian views for the evolution of trees. Alternative more Ereban views relate the evolution of trees to environmentaldifficulties.Fifth, it could be that shortagesof nutrients in the soils of unvegetated watersheds stunted the growth of the earliest land plants. Sixth, it could be that the frequency of disturbance by floods, windstorms or landslides was such that few plants could grow for long enough to attain the stature of trees. These are other questions, in addition to the simpler one of when forest ecosystems appeared, that can be tested against the fossil record of soils.
Early forest soils Fossil stumps in growth position are an obvious, but all too rare, indication of ancient forests because they require rapid burial in a waterlogged habitat (Driese et al. 1997). More abundant evidence of early forests comes from surficial disruptions of paleosols as a result of former windthrows, from the presence of charcoal in paleosols, from deeply penetrating large root traces and from the differentiation of leached, near-surface (E) and enriched, clayey or ferruginized subsurface (Bt andBs) horizons. Alfisols The oldest of these base-rich, clayey forest soils may be among Middle Devonian (Givetian, 3 72 Ma) paleosols of the Aztec Siltstone in Victoria Land, Antarctica (Retallack 1997e). Although many Precambrian paleosols are thick and clayey (see Fig. 16.6),the Rosemary paleosolfrom the Aztec Siltstoneis the most ancient that I have seen with a subsurface enrichment of clay that qualifiesas an argillic horizon (Fig. 19.2).Furthermore, thin sections of this clayey horizon show skins (argillans)of highly birefringent clay, washed down into the profile (Fig. 19.3D). as evidence against origin of all this clay as a preexisting sedimentary deposit. Another criterion required for identification as an Alfisol is evidence of fertility,such as high cation exchange capacity. This measure is compromised by alteration after burial, but can be inferred for the Rosemary paleosol from the presence of free carbonate as fine nodules andrhizoconcretions (Fig. 19.3C) and from geochemical data, such
as low molecular ratios of Ba/Sr and alumina/bases (Fig. 19.2). Whether this is indeed the most ancient Alfisol can be tested by applying these criteria to more paleosols. There is much evidence for trees in the Rosemary paleosol from sudicial tree-throw structures and large drab-haloed root traces (Fig. 19.3A). Furthermore, fossil wood and spores from the Aztec Siltstone are evidence that these trees were Archaeopteris-Callixylon (Retallack et al. 1997). Comparable soils today support monsoon forest of sal (Shorea rubusta) in northern India (Retallack 1991a). The paleoenvironment of these ancient forests can be envisagedas alluvial bottomlands of quartzofeldspathic composition in a highly seasonal, subtropical, subhumid climate. Strong paleoclimatic seasonality can be inferred from extensive surficial cracking of the paleosols (Fig. 19.3A) and prominent concentric banding of concretions (Fig. 19.3E). Depth of the calcic horizon of the Rosemary paleosol, corrected for likely burial compaction effects,indicates a mean annual precipitation of 779 f 141mm (using techniques of Figs 7.6 and 9.5). The degree of weathering and ferruginization is compatible with a warm paleoclimate, as is also indicated by the paleomagnetically determined paleolatitude of 40"f 14"(Retallack1997e). Fossil Alfisols are now known from all the succeeding geological periods up to the present (Retallack 1991a, 1994b, 1995a, 1997f; Retallack et al. 2000). AEsols and forests have had a long and intimate relationship.
Ultiils The most ancient of these base-poor clayey forested soils so far documented is the Lykens Valley paleosol in the Lower Pennsylvanian (320Ma) Schuykill Member of the Pottsville Formation of north-eastern Pennsylvania, USA (Gill&Yemane1996).This is a thick (50cm), gray, kaolinitic paleosol with molecular ratios of alumina/bases of up to 3.2, indicating considerable nutrient cation weathering. Furthermore, there is a subsurface (Bt) horizon of clay enrichment and clay skins that qualilies as argillic, as well as a near-surface sandy eluvial (E) horizon. Such differentiation was probably formed over a long period (many tens of thousands of years) under a humid climate. Drab color and
Afforestationofthe land
283
molecular weathering ratios point count data percent percent grain size composition
Field observations
N2Q3 SiO,
2 5
I
wegk red (10R4/3)
A
dusky red (10R3/3)
Bt
weak red (10R4/2)
Bk
weak red (1 OR413)
C
&a
CaO+MgO
N%O+&O 0 5101
Ea Sr
5'p'
dark greenish gray (5G4/1)
0
rock red color
sandstone
relict bedding
claystone
~~$$,~ross
opaque
salinization calcification clayeyness base loss leaching
gleization
IBe] calcareous rhizoconcretions
Figure 19.2 Columnarsection (measuredin field),petrographic composition(frompoint counting)and molecular weathering ratios (frominductivelycoupledplasmaemissionspectrometry)of theRosemary clay paleosol,a Haplustalf in the MiddleDevonian (Givetian,3 75 Ma) Aztec Siltstone onMt Crean,southernVictoriaLand,Antarctica(fromRetal1ack 19971% reprintedwith permissionfrom the AmericanAssociation for the Advancement of Science).
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Figure 19.3 Fieldphotographs (A,B) and photomicrographsunder crossed nicols (C-E) of theMiddleDevonian(Givetian,375 Ma) RosemaryclaypaleosolatMt Crean,southern VictoriaLand, Antarctica:(A)drab-haloed root traces, large root disruptions (at arrow) and sand-filled cracks in the surface of the profile:(B) completeprofile withinterpreted soil horizons: (C) rhizoconcretionin subsurface:(D) clay skins (climobimasepic fabric) in subsurface:(E) ferruginous concretionin subsurface. Hammer is for scale in A and B, and scale bars for C-E all represent lmm(fromRetallack1997e:with permissionfrom the American Association for the Advancement of Science).
associated coals are evidence of a swampy alluvial plain environment. The paleosol also contains deeply penetrating carbonaceous root traces indicative of good drainage and forest vegetation. The root traces are woody in appearance, like those of gymnosperms, and very different from the stout, succulent roots of tree lycopsids and horsetails common in Pennsylvanian paleosols (DiMichele et d. 1996a). Although fossil plants have not yet been reported with the Lykens Valley paleosol, its vegetation probably included seed ferns, which werecommon at that time (Retallack&Dilcher1988). Ultisols are found, although not common, in every succeeding geological period (see Figs 4.9 and 10.11: Retallack 1983a. 1997d;Retallack&Krulll999).Even in modern landscapes, Ultisols are usually on high terraces hundreds of thousands to millions of years old
(Markewich et al. 1990), and such geomorphological positions may explain their rarity in sedimentary deposits of fluvial lowlands. Spodosols The most ancient paleosol with a strongly ferruginized subsurface (Bs) horizon is in the Lower Carboniferous (Visean, 335Ma) Chipping Sodbury Member of the Clifton Down Limestone in Southfields Quarry, near Bristol, England (Vanstone 1991). This paleosol has a thick (50cm) surface (A) horizon of light gray quartzrich sandstone over a subsurface (Bs)horizon cemented with crystalline hematite, which is especially dense in the upper 5 cm. There are thin dark stringers of organic matter andiron-manganese near the surface of the pro-
Afforestation of the land file, and these, together with the lowland depositional setting of the sandy alluvial parent material, can be taken asevidenceof periodic waterlogging, as in Aquods of Soil Survey Staff (1997). The depth and degree of weathering of the paleosol are compatible with a humid paleoclimate. Its forest vegetation, probably gymnospermous, can be inferred from stout woody root traces filled with debrisfrom above. A few other well-documented fossil Spodosols are known (Pomerol 1964). but it is surprising how rare they are as paleosols, considering how common these soils are today, especially at high latitudes. There are thousands of quartz-rich sandy paleosols with leached surface horizons in Carboniferous coal measures of Euramerica (Percival 1986; Gibling & Rust 1992) and in Permian-Triassic alluvial deposits of Australia and Antarctica (Retallack 1997d, 1999b: Retallack& Krull 1999).Yet none of these have been reported with subsurface horizons that would qualify as spodic. There are a few Carboniferous and Triassic paleosols with limited subsurface iron accumulation that might have become spodic with more time for development (Retallack 1992e, 1997d). Many of these sandy Carboniferous-Triassic paleosols have subsurface nodules of siderite or chlorite, indicating a more alkaline soil chemistry than found in Spodosols. Acid-generating compounds responsible for podzolization, such as insect-repellant phenolic compounds, are probably ancientinpaleosols (Retallack1985),but may havegained in potency and concentration through geological time in response to escalating insect assault on vegetation (Retallack 1997d). Origin of forest ecosystems
Although great strides have been made during the past decade in documenting early forested paleosols, it remains difficult to test effectively hypotheses about the early evolution of forests. At face value the record of early forested paleosolssuggeststhat Devonian forests at first colonized fertile bottomland Alfisols. Only later in the Carboniferousdid forests adapt to low-fertility clays as in Ultisols. or quartz-rich sands of Spodosols. The idea that forests evolved only where little disturbed by flooding is unlikely considering the nearcoincident appearance of forested paleosols in alluvial sequences and of fossil logs that could in part have come
285
from upland soils (see Fig. 18.11E). Nor is it likely that nutrient availabilityrestricted tree growth, because the earliest known paleosols under forest were fairly fertile, even calcareous. Another hypothesis not supported by existing evidence is that trees scattered their spores more widely than other plants. Within alluvial floodplains, fluvial transport of spores and gametophytes would have rendered tall stature and wind dispersal unnecessary, The idea that trees evolved to avoid herbivory is supported by early indications of podzolization (Retallack 1985), because phenolic substances produced by the trees also would have suppressed herbivores. Plant competition also could have played a role because early forested soilshad surface horizons of even thickness: an indication of a closed canopy. If it remains true that Ultisols andSpodosols appeared geologically later than calcareous Alfisols, then it could be because increased complexityof forestecosystemsallowed recycling of nutrients from decaying vegetation in soils in which nutrients were scarce. The fossil record of Spodosols also can be expected to reveal steps in the escalation of chemical warfare waged by plants against herbivoresculminating in copious production of phenolic compounds in modern plant formations such as heath and conifer forest (Retallack 1997d). None of these questions can be answered satisfactorilyat present. They are introduced as a perspective for the further examination of mid-Paleozoicfossil soils, plants and animals. A diversifyinglandscape
With the addition of the main orders of forested soils during mid-Paleozoic time, most types of soils now known on Earth were already present. Of the 12 orders of the US soil taxonomy (Soil Survey Staff 1998), only the grassland soils (Mollisols)hadnot yet appeared. All the older kinds of soils persisted along with early forests and their soils as outlined in the following paragraphs. Entiiols These very weakly developedsoils are so broadly defined that they include early stages of developmentof all kinds of soils.Theyprobab1yformedevenduring Precambrian times, but are difficultto detect in Precambrian rocks be-
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cause the degree of weathering is slight and there were no burrowing animals or rooted plants to leave obvious traces. This was no longer the case on subtropical alluvial flats by Late Ordovician time (Ashgillian),when the Faust Flat silty clay paleosol (a Fluvent) formed (see Figs 18.4 and 18.5). Comparable Entisols in alluvial sequences with fossil burrows, root traces and conspicuous relict bedding are common in sedimentary rocks of every younger geological period (Retallack 1983a, 1985,1991a). Fossil Entisols also record the invasion of woody plants into other habitats such as seashores, lake margins, deserts, periglacial regions and rocky uplands. Saltmarsh vegetation paleosols may date back to the Silurian (Schopf et al. 1966; Edwards & Selden 1993), but woody plants comparable with mangroves had invaded near-marine habitats by Devonian time. There is evidence for this in weakly developedpaleosols containing marine fossils and large tree stumps as old as Middle Devonian (Givetian, 375Ma) age in New York, USA (Johnson 1972; Dreise et al. 1997). Other mangalEntisols have been described from the Early Carboniferous (Tournaisian, 3 50 Ma) Calciferous Sandstone near Foulden. Scotland (Retallack& Dilcher 1988)and from the mid-Cretaceous (Cenomanian) Dakota Formation near Russell, Kansas and Fairbury, Nebraska, USA (Retallack & Dilcher 1981a,b), among many others (DiMicheleetnL 1987).Salinecoastalsoilswereanother habitat colonized and stabilizedby large woody plants. Calcareous palustrine paleosols. largely Entisols, have been found in nonmarine limestones of most Phanerozoic periods, with the oldest known examples from the Middle Devonian (Givetian, 380Ma) Oneonta Formationof upstateNewYork,USA (Dunagan&Driese 1999). These distinctive calcareous paleosols include soil crusts, breccias, desiccation cracks, circumgranular spar-filled cracks, and calcareous rhizoconcretions, as well as aquatic fossils such as ostracods, charophytes and pond snails. The Florida Everglades is a comparable modern mosaic of seasonal lakes, wet prairie, and hammocksof hardwoodsandpine (Galli 1991). Sandy and silty Entisols with root traces have been found in eolian deposits of Pennsylvanian and Permian age in Utah, Colorado and Wyoming (Loope 1988; Soreghan et al. 1997).The size distributions of calcareous root traces in these paleosols are indications of early desert shrublands and scrub.
Weakly developed, rocky upland Entisols (Orthents) had certainly begun to be colonized by latest Triassic time (206Ma).Fissures of karst topography of that age on Early Carboniferous limestones near Bridgend, South Wales, are filled with cave earth and charcoalified remains of a fireprone shrubland that was in a sense a forerunner of modern chaparral (Harris 195 7).
Gelisols The oldest periglacial paleosols (Orthels and Turbels) date back to the early Proterozoic, some 2450Ma (Young&Long1976;Williams 1986),butneitherthese nor Late Ordovician (445 Ma) periglacial paleosols (Daily & Cooper 1976) have been found to contain root traces.Themost ancient woody root traces in periglacial paleosols are in the Late Carboniferous (310-3 1 2 Ma) Seaham Formation, near Lochinvar,New South Wales, Australia (Retallack 1999a). Here Gelisols with earth mounds (thufur) also contain fossils of the seed fern Botrychiopsis. The Seaham Formation includes other indications of glacial conditions, such as tillites and varved shales, and has a paleomagnetically determined paleolatitude of 70"s. These paleosolsand low-diversity fossil plant assemblages may have formed vegetation similar to that of modern polar tundra. If woody plants could adapt to such frigid conditions in polar regions, it is likely that high montane forest (krummholz) is of similar antiquity Gelisols of the Late Permian (258Ma) Gerringong volcanics of New South Wales,Australia, include woody root traces and permafrost disruptions, and can be regarded as the earliest evidence of taiga woodland (Retallack 1 9 9 9 ~ )Permafrost . features of these paleosols include ice cracking and ice lenses. Permafrost mixing of relict bedding has given these paleosols the appearance of greater development than their meager degree of chemical weathering, which was presumably limited by their frigid paleoclimate. Other peaty Gelisols of frigid swamps have been found in Permian rocks of Antarctica (see Figs 9.1 1and 9.12; Krulll999).
Inceptisols Many Precambrian paleosols lack diagnostic horizons of modern soils, and despite impressive development, are best regarded as Inceptisols (see Figs 16.3-16.6).
Afforestation of the land Inceptisols are common throughout the fossil record of soils in sedimentary rocks (Retallack 1983a, 1985, 1991a. 1994a. 199Sa, 1999b). An Inceptisol of Late Devonian (Frasnian. 365Ma) age near Hancock, New York, USA, is of interest in revealing the antiquity of the soil-forming process of podzolization (Retallack 1985). The large root traces and a the cumufossilleaf litterof Archaeopterishal~ia~awithin lative surface horizon provide evidencethat the paleosol was forested. Other evidence for forest cover is the differentiation of a laterally continuous, iron-poor, graygreen surface (A) horizon over a purple subsurface (Bs) horizon a little richer in iron. Translocation of iron to subsurface horizons can be mechanical because of channels offered by decaying roots and open burrows. Someof the fossil root traces and burrows in thepaleosol are filled withiron-stainedclay washed down in this way, but the larger ones are filled with drab-colored sand rather than clay. Physical mixing played a minor role in the formation of this subsurface (Bs) horizon because relict bedding has persisted throughout the profile. A morelikelymechanismis the translocationof iron by the chemical action of phenolic and other substances leached from leaves by rain or from the decaying leaf litter.
Andisols These soils formed on ash of volcanic shards and amorphous weathering products have been separated from theorderInceptisols(Soi1SurveyStaff1998)becauseof their distinctivetow bulk density and high fertility.Such soils are probably of great antiquity, probably including a Proterozoic (1500Ma) paleosol developed in the Johnson Shut-ins Ash Flow near Lesterville, Missouri, USA (Blades & Bickford 1976: Retallack 1997a). It is likely that Andisols were colonized by vascular plants and trees during the Paleozoic, but examples have not been reported. Andisols arebetter knownfrom Oligocene (30Ma) volcaniclastic alluvium in Badlands National Park, South Dakota (Retallack 1983a) and the Painted Hills of central Oregon, USA (Retallacket al. 2000). Ostwald ripening during burial of amorphous phases such as imogolite in the original soils may account for the common zeolites. such as clinoptilolite, in these paleosols. Despite this burial alteration, shard textures remaintoindicate that these were Andisols (seeFig. 12.5).
28 7
Histosols Fossil Histosols can be traced at least back to Early Devonian (Siegenian, 410 Ma) time, but the oldest woody coals are of latest Devonian age (Fammenian, 360Ma; Gillespie et al. 1981). With the advent of trees in wetlands, much thicker peats accumulated (Fig. 19.4). Most of the world’s economically minable coal is of Carboniferous and Permian age (Diessel 1992). Paleosob associated with coals have been given a variety of names: underclay, seat earth, fireclay, tonstein and ganister. Some of these represent the mineral portions of Histosols, but many are more properly identified as gleyed representatives of other soil orders (Gardner et al. 1988;Hughes et al. 1992). The present fossil record of Histosols indicates a later (Fammenian, 360Ma) afforestation of peaty substrates than of waterlogged mineral soils or well-drained soils (both Givetian, 375Ma: Retallack 1997e; Driese et al. 1997). If this holds up to further scrutiny it would add weight to the inference from the fossil record of Spodosols that trees were slow to adapt to acidic, nutrientpoor habitats. A number of special features allow trees to persist in swamps. Hollow roots, for example, allow gas exchange in stagnant groundwater, and are found in fossil lycopods such as Lepidosigillaria whitei and Eosperrnatopteris erianum of the Middle Devonian (Givetian, 375Ma; Algeo & Scheckler 1998). Such adaptations to waterlogged conditions distinguish swamp vegetation from that of surrounding, betterdrained soils. Fossil plants of Carboniferous coal measures are long-ranging, morphologically conservative species compared with rapidly evolving fossil plants of shales and sandstones that represent vegetation of beach ridges, levees and other better-drained land (DiMichele et al. 1987). Furthermore, swamp floras have been remarkably stable on geological time scales despite paleoclimatic fluctuations on scales of 10-100 kyr (DiMichele et al. 1996b). Thus, swamps appear to have been colonized late by trees and to have remained evolutionary as well as literal backwaters. Not all early Histosols were as nutrient poor as most modern swamps. A distinctive kind of fossil Histosol found principally in Late Carboniferous coal measures of the midcontinental USA and in mid-Carboniferous coal measures of Europe and the Ukraine contains numerous calcareous or dolomitic nodules. These so-
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GEOLOGICAL AGE
I
CRETACEOUS
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PALEOLATITUDE 30" 60'
10
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Figure 19.4 A geologicalhistoryof Histosolpaleolatitude (left),coal (0horizon)thickness(maximum,and maximum averageof at least 10 consecutiveseams,center columns)and botanicalcomposition(right column)(fromRetallack et al. 1996: with permission from the GeologicalSociety of America).
called 'coal balls' are original features of the soils, preserving their enclosed plants from compaction caused by peat settling and the pressure of rock overburden. The woody vegetation of these alkaline wetlands was thus a carr, rather than swamp (as defined by Gore 1983). Calcareous nodules are extremely rare in modern peats, whichnormally are too acidic for calciumcarbonate to be a stable mineral. A modern calcareous peat from Eight Mile Swamp in south-eastern South Australia is an exception (Stephens 1943; Retallack 1986d). In this soil, l i e is produced by charophyte algae and aquatic snails in a coastal swamp buffered by springs fed with runoff from regionally extensive limestone bedrock in a summer-dry (Mediterranean)subhumid climate. Such factors also may have played a role in the formation of Carboniferous coal balls. They are found in paleoclimates intermediate between humid
and semiarid. Coal balls also have been found in areas where there is earlier Carboniferous limestone and in coastal although not necessarily marine-influenced swamps (Scott et d.1996). Lycopsids and horsetails of Carboniferous swamps may have been less acidifying than gymnosperms that subsequently invaded swamp habitats, and this may have restricted the formation of coal balls in rocks younger than Carboniferous. Comparable escalating effectsof plant leachates on weathering through geological time also can be examined from the fossil record of Spodosols and siderite nodules in paleosols (Retallack 1997d). Vertisols Swelling clay soils are known to be at least as old as 2200 Ma from a n example near Waterval Onder, South Africa (see Fig. 9.13). Their striking slickensides and
~
Afforestation of the land festooned horizons have led to their recognition throughout the geological record of soils (Galloway 1978; Goldbery 1982; Caudill et al. 1996; Driese et al. 2000). Silurian-Devonian examples formed under woody plants (Marriott & Wright 1983; Allen 1986a) show complex fracture patterns resembling lentil peds, which contrast with the simpler blocky structure of the Precambrian WatervaI Onder clay Phanerozoic Vertisols appear to have had a more unstable, smectitic clay than the illitic clays presumed to have dominated the Precambrian profile (Retallack 1986a; Retallack 8.1 Krinsley 1993). These differences in visible effects of physical properties of clay may be related to a number of factors such as amount and seasonality of rainfall, physical conditioning of the soil by roots, and the greater demand for potash by vascular land plants compared with Precambrian microbial earths.
Aridisols Alkaline, calcareous desert soils have a continuous record from the late Archean (2 560 Ma) to the present (see Fig. 17.6; Steel 1974; Campbell & Cecile 1981; Wright 1982; Naylor et al. 1989: Martini 1994). The oldest large root traces in such paleosols now recorded are in the Middle Devonian (Givetian,3 75 Ma) Aztec Siltstone on Mt Crean, Antarctica (Retallack 1997e; Retallack et al. 1997). Late Jurassic Aridisols of Dorset, England, include large permineralied stumps of conifers and cycadeoids (see Fig. 1.2: Francis 1986). Some very ancient Aridisols are distinctive because their carbonate is dolomite rather than the lowmagnesian calcite most abundant in modern Aridisols (Kalliokoski 1975; Campbell & Cecile 1981; Retallack 1985).Therearealsoancient examplesof groundwater dolocretes associated with rhizoconcretions, comparable with some groundwater calcretes (Spotl & Wright 1992). Some of this dolomite in paleosols is probably of diagenetic origin (Ketallack 1992e). In other cases a diagenetic origin of dolomite is not evident (Ketallack 1985; Spotl & Wright 1992). Dolomite is formedin few modern soils,especiallythose that are arid andsaline (Buietal. 1990;BothatGHughes 1992)or on magnesium-rich parent material (Capo et al. 2000). There may have been long-term changes in the relative abundance of dolomite and calcite in paleosols compa-
2 89
rable with long-term changes in oceanic carbonates (James& Choquette 1994; Stanley & Hardie 1999).
Oxisols and duricrusts Very strongly developed soils and indurated products of deep weathering extend well back into Precambrian time, perhaps 3 500Ma (Serdyuchenko 1968; Martyn &Johnson 1986), but would also have been promoted by the stabilizing effect of land plants and of forests. Duricrusts occur below major geological unconformities of all ages, but the most impressive are associated with Devonian and younger unconformities. Phanerozoic lateritic crusts are commonly underlain by intensely leached masses of china clay (see Fig. 2.4). Such lateriticprofilesare widespread on stable,low-relief continents such as Africa and Australia (Nahon 1991; Ollier&Pain 1996).The largest reserves of bauxites are in early Tertiary erosional landscapes in Guyana, Surinam, Guinea, Ivory Coast, Ghana, northern Queensland and the Northern Territory. Some bauxites also are formed in closed depressions of karst topography (Bardossy 1982). Phanerozoic examples of this kind of bauxite also are much thicker and more deeply weathered than the few Precambrian examples known. Similarly,silcrete of the 1850Ma paleosol beneath the Pitz Formation of the Canadian North-west Territories (Ross & Chiarenzelli 1984) forms only thin veins, whereas thick ridges of silcrete are associated with latest Cretaceous and early Tertiary unconformities in central Australia (see Fig. 6.7). Deep continental weathering has become increasingly more profound since Precambrian time. Many Oxisols today support tropical rainforest, which is indicated by large root traces in an Oxisol as ancient as Pennsylvanian (305 Ma) in the Cheltenham Formation, near Drake, Missouri, USA (Retallack & German-Heins 1994). Identification of this paleosol as an Oxisol is confirmed by its high oxidation state (low FeO/Fe,O, molecular ratios), low base status (high aluminalbases ratio) and abundant kaolinite (Fig. 19.5). These features can be taken as evidence of good drainage(at1eastabove 150cmdepth)andaveryhumid climate (more than 2000 mm mean annual precipitation). Such an ancient rainforest was botanically very different from the modern Amazonian selva, as there
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field observations
pi point count dnln 2 pircent percent d grnin size composition a
nioleculnr weathering rntios
AllQ,
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. . .
r
salinization calcificntion clnyeyness .base loss aching none none modernte mnrked nioderak
"d,zp l-Jfossil lags
00
E3
root traces
red color
0
@ rJ&,g
gray color
slickensides
Eleizhion i t depth
iron nodules
Figure 19.5 Columnar section (measured infield),petrographic composition(frompoint-counting thin section) and molecular weathering ratios (fromchemical analysis) of the Farnberg clay paleosol, an Acrudox. from the middle Pennsylvanian (305Ma) Cheltenham Formation, near Drake, Missouri,USA (fromRetallack & Germtin-Heins 1994: with permission from the American Association for the Advancement of Science).
is evidence from associated deposits of vegetation dominated by large-leaved seed ferns. The adaptation of forests to such low-nutrient substrates, presumably by means of nutrient recycling in litter (Sanford 1987) and interception of atmospheric dust (Chadwick et al. 1999). may have taken some time, judging from the later geological appearance of forested Oxisols than forested Alfisols (Retallack1997a). Also amenable to testing in future studies of paleosols are the observations that Atfisols and Ultisols are unknown without evidence of forest, and so may have been created by forest ecosystems, but Oxisols appear to predate rainforest.
A near-modernworld Considering the three new orders of soils that appeared during the Devonian (Alfisols, Ultisols and Spodosols), together with notable changes in other kinds of soils and weathering products, the advent of forestscan be seen as a time of accelerated diversification of soilson Earth. By Permian time vascular plants had established themselves in alpine to mangal, polar to tropical and rainforest to desert environments. These late Paleozoic woody plants were mainly of kinds now extinct. Some features of Paleozoic soils, such as common calcareous nodules in Histosolsand the rarity of Spodosolsat high paleolati-
Afforestation of the land tudes, are unusual by modern standards. Nevertheless, the world and its soils were becoming much more l i e those with which we are now familiar.
Afinerwebof lifeon land The evolution of a terrestrialenvironment structured by trees profoundly affected numerous other creatures on land. The physical complexity of forests has been matched by increasingly complex interrelationships between forest organisms (Vermeij 1987). Althoughforest ecosystems are complex, they have proven robust on geological time scales, with impressive recovery from times of mass extinction (Retallack 1994a: Retallack & Krulll999). Forests were sources of evolutionary novelty, harboring the origin of many groups of plants and animals important to modern ecosystems such as gymnosperms, angiosperms, insects, amphibians, dinosaurs, mammals and primates. Not all of these groups have left clear traces in forested paleosols, but their evolutionary history can be assessed against the paleoenvironmental information offeredby paleosols.
Soil invertebrates A general diversification of invertebrate trace fossils within paleosols through geologicaltime is now obvious from studies of their trace fossils (Buatois et al. 1998). Furthermore, trace fossils have also increased in their depth and density of soil exploitation (tiering) through geologicaltime(compareRetallack 1985andHasiotis& Dubiel1994). The large number of soil ichnogenera now known can be grouped into several ichnofacies, comparable with those proposed by Seilacher (1967), who noted only one nonmarine Scoyenia ichnofacies, but a variety of depth-relatedmarine ichnofacies.His Scoyenia ichnofaciescan be extended back in geologicaltime to include burrows attributed to millipedes in Late Ordovician (445Ma) Aridisols (see Figs 18.5-18.7). Another ichnofacies dominated by small ellipsoidal fecal pellets of invertebrates comparable with the ichnogenera Coprulus and Tibikoia (Hantzschel 1975) is well documented from Early Carboniferous (330 Ma) mollic Inceptisols (Wright 198 3,198 7) and Late Carboniferous (310Ma)
29 1
Histosols (Labandeira et al. 1997). The Skolithos ichnofaciesproposed by Seilacher ( 1967) for worm-burrowed shallow marine sandstones is also common in sandy river deposits,especiallyin EarlyTriassicEntisolsof eastern Australia and Antarctica, where Skolithos is attributedtocicada-liieinsects (Retallack 1997 d Retallack& Krull 1999). The Terrnitichnus ichnofacies (of Smith et al. 1993) includes a wide variety of termite nests and pellets, and is common in Mesozoic and Early Cenozoic Aridisols and Alfisols (seeFig. 9.9: Bown 1982: Hasiotis & Dubiel199 5).The Coprinisphaeraichnofacies includes avarietyof nestingstructuresof bees (seeFig. 10.9) and dung beetles (see Figs 10.8 and 10.10) in Cenozoic Mollisols and mollic Inceptisols (Retallack1990: Genize & Laza 1998: Genize et al. 2000). These fossil nests are evidence of the habitats and evolutionary history of socialinsects. This overview reveals an intriguing correspondence in time of organic-sequestering ichnofacies (Coprulus-Tibikoia, Coprinisphaera) with global climatic ice ages, and of organic-oxidizing ichnofacies (Scoyenia. SkoZithos, Terrnitichnus) with gIobal climatic greenhouses (seeFig. 9.1 5). Such relationships between global change and soil ichnofacies have been proposed in general terms (Algeo et al. 1995: Retallack 1997e),but remain to be explored more thoroughly with trace fossils in paleosols. Early land vertebrates The earliest fossil amphibians such as Elginerpeton are LateDevonian (Frasnian, 365 Ma: Ahlberg et al. 1996: Marshall et d.1999). A traditional view of these early amphibians is that their limbs hauled them overland to permanent water during dry seasons. However, modern and ancient lungfish faced with the approaching dry season burrow and aestivate rather than migrate to a likely death (Boucot 1990). An alternative view that is that fish evolved stout fin supports and ultimately the limbs of early amphibians for skulling in shallow permanent ponds or for linding food in cool moist forests (Ahlberg & Milner 1994). A humid forested paleoclimate is implied by these hypotheses, but a seasonallydry climate by the dry-season migration hypothesis. Devonian paleosols of moist forests and of seasonally dry woodlands have been described (Retallack 1985: Driese
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& Mora 1993),but such observations have not yet been
applied to the testing of these hypotheses in early amphibian-bearing rock sequences. Some of the earliest reptiles, such as the captorhinomorphs Hylonomus and Archerpeton and the pelycosaur Protoclepsydrops, are found in paleosols in the JogginsFormation of Pennsylvanian age in the sea cliffs near Joggins,Nova Scotia (Falcon-Lang 1999).They are preserved with remains of amphibians, millipedes, spiders, whip spiders, eurypterids and land snails in sandstone casts of tree stumps. The hollows left by rotting of the stumps were traps for a variety of animals. The paleosols themselves have not been studied in detail, but appear to have been weakly developed and waterlogged. The associated fossil plants are characteristic of Carboniferousswampy lowlands. Wet forests rich in insects and other invertebrates would have been an important resource for carnivorous early amphibians and reptiles. Dinosaurs Because of their great size, dinosaurs should have had an effect on soil formation. It has been suggested,for example,that tracts of woodland may have been opened to herbaceous plants by migrating herds of dinosaurs, and that their piles of dung would have required an efficient fauna of decomposers (Bakker 1985). Paleosols of the Late Jurassic (149 Ma) Morrison Formation of Utah, Colorado and Wyoming, USA, lend some support to such views (Hasiotis& Demko 199 6b; Retallack 199 70, The gigantic sauropods Camarosaurus and Diplodocus have been found in paleochannel sandstones, but also in red clayey paleosols with shallow calcareous (Bk) horizons, comparable with soils of semiarid woodland (Fig. 19.6). How did such low-productivity soils and plant communities support such large biomasses of herbivorous dinosaurs and termites? Did the giant sauropod dinosaurs have aphysiologyless demanding of vegetation than modern elephants? Was the Mesozoic greenhouse paleoclimate related to such high consumer compared with producer productivity? Such questions remain to be addressed by more comprehensive studies of paleosols associatedwith dinosaurs. Not all paleosols associated with dinosaurs are so organic lean, and not all dinosaurs so large as in the Morrison Formation. Dinosaurs also are found associated with paleosols of humid swamps and forests, of cool
high-latitude woodlands and of dry deserts (Retallack 1994a, 1997f; Retallack&Dilcher 1981a,b).As dominant creatures on land for some 15Omyr, dinosaurs diversified intomost land habitats (Sereno 1999). Angiosperms If angiosperms existed at all during Triassic time they were rare and geographically restricted plants. Especially notable is a small plant of Late Triassic age (Carnian, 225Ma), which is known from superficially palm-lie leaves (Sanmiguelialewisii), and reproductive organs (Cornet 1986)that suggest to me that it was an early gnetalean, and thus in a sister group to early angiosperms. These well-preserved fossils were found with a leaf litter of ferns rooted within a Fluvent in the swale of a meandering stream. Thus it was an early successional, streamside plant. The most ancient definite angiosperm fossil is from Early Cretaceous (Barremian, 125Ma) lake beds of north-westChina(Sunetal.1998;Swisheretal. 1999). Fossil angiosperm flowers, fruits, pollen and leaves are found in force as a continuous fossil record from then up to the present. The earliest angiosperm remains in North America show a predilection for disturbed streamside and coastal environments (Retallack & Dilcher 1986). In the Early Cretaceous (Barremian, 123Ma) Potomac Group near Washington, DC, USA, angiosperm fossils are rare in very weakly developed shaly and sandy Entisols of estuary margins, in pink clayey Inceptisols of levees and in Histosols of swamps. No fossil plants have been found in red, clayey Alfisols of well-drained floodplains nor in widespread, deeply weathered Ultisols or Oxisols developed on Paleozoic basement rocks throughout this region. Judging from the regional pollen rain into lake and river deposits, which is dominated by conifer pollen, it is unlikely that angiosperms were present in these inland soils. By mid-Cretaceous time (Cenomanian, 94 Ma) angiosperm fossils were more common in many parts of the world: such as the Dakota Formation of central Kansas and Nebraska, USA (Fig. 19.7). Beautifully preserved flowers, pollen, fruits and leaves of angiosperms dominate leaf litters of moderately developed, drab, clayey, estuary margin Entisols and coastal swamp Histosols. Also dominated by fossil angiosperms are weakly developed,sandy paleosols (Psamments) of river levees.
Afforestation of the land
paleosol
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Iv( i;;-&dJ./ng- - -
Figure 19.6 Areconstructionof paleosolsin theLate Jurassic(Tithonian,149 Ma)MorrisonFormationatDinosaurNationaI Monument,Utah-Colorado,USA (komRetallack1997c;with permissionkomDmofestInc.).
Curiously,fossil conifershave not been found within leaf litters of these fossil soils although their pollen, shoots and cones are conspicuous in deposits of associated lakes and seas. Probably conifers dominated the vegetation of red, clayey Ultisols on better-drained parts of the floodplain where fossil plants were not preserved. By mid-Cretaceous time then, angiosperms already dominated mangal, swamp and levee vegetation, but had made limited gains in well-drained forests.
This fossil record of soils associated with early angiosperms can be used to evaluate hypotheses concerning their evolutionary origins (Wing & Boucher 1998). A widely held view of the origin of angiosperms is that they were trees of dry tropical uplands with large insectpollinated Magnolia-like flowers and fleshy, animaldispersed fruits. Their rise to dominance according to this concept is explained by Cretaceous migrations of these plants into lowland sedimentary environments
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Figure 19.7 Areconstructionof soilsandvegetationduringmid-Cretaceous(Cenomanian,94Ma) time incentralgansas, USA (datafromRetallack&Dilcher198 la,b).
Afforestation of theland along with interdependent animal pollinators and dispersers. Evidence from paleosols and early angiosperm fossils does not support this view, because the earliest angiosperms were lowland, weedy plants. Such plants today are small and produce numerous small (152 5 pdiameter) pollen grains and (1mm or less) seeds so as to be able to colonize widely scattered sites. Fossil leaves, fruits and pollen of the earliest Cretaceous angiosperms now known in various parts of the world are of this kind. They also are low in diversity and similar in regions as remote from one another as North America, Australia and central Asiatic Russia. It is as if a small group of weedy plants dispersed widely in coastal regions at this time (Retallack & Dilcher 198lb; Truswell 1987). Why did they migrate during Cretaceous t i e if allied Late Triassic plants were also streamside weeds? Wide dispersal and speciation of angiosperms may have been encouraged by fluctuation of sea level in shallow continental seaways characteristic of mid-Cretaceous time and by the browsing pressure of herbivorous dinosaurs (Weishampel& Norman 1989). For weedy plants there would have been great advantage in an abbreviated lie cycle in which pollination and fertilization followed one another in rapid succession, as in modern angiosperms. In contrast, all modern gymnosperms and apparently also the most reproductively sophisticated of ancient seed ferns (Retallack & Dilcher 1988) were pollinated before the gametophyte was fully differentiated. Thus, fertilizationis delayed.Inpinesfertilization occursmany months after pollination. Near-synchronous pollination and fertilization followed by coordinated development of seeds and enclosing structures would have enabled weedy angiosperms to set seed faster than gymnospermous plants that were equally adapted in other ways to early successional habitats. Angiosperms remain the premier early successional plants on land, as can be seen from the rapid appearance of grasses and plantains in vacant lots. Since mid-Cretaceous time angiosperms have diversified to include magnolias, orchids and apples (Crane et al. 1995).In these plants, co-evolutionwith animalpollinators and dispersers has proceeded to extraordinary lengths (Labandeira 1998). This subsequent evolution of angiosperms also can be addressed through the study of paleosols. The most important pollinators of angiosperms are bees, fossilized in amber as old as Late
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Cretaceous (Campanian, 80Ma; Grimaldi et al. 1989) and perhaps also indicated by nests in paleosols and permineralized wood as old as Late Triassic (Carnian, 225 Ma; Hasiotis et al. 1996). By Eocene t i e , fossil flowers with the asymmetric shape characteristic of bee-pollinated blossoms are common (Crepet & Taylor 1985). Fossil bee nests are also more common and diverse in Cenozoic than in Mesozoic paleosols (Retallack 1990; Genize & Bown 1996). The fossil record of bees, their flowersand nests is evidence of a dramatic evolutionary radiation of both bees and blossoms during Cenozoic time. Crisis and recovery Forest ecosystems have persisted since Middle Devonian time (3 75 Ma), but were severely tested during times of mass extinction at the Permian-Triassic (250 Ma) and Cretaceous-Tertiary (65Ma) boundaries (Hallam & Wignall 1997). The end-Permian extinctions were by far the more devastating of the two events, both in the sea and on land (Retallack 1995b), where there was a marked discontinuity in the botanical composition of swamp floras (Fig. 19.4).Sequences of paleosols across this boundary in South Africa (Smith 1995). Australia (Retallack 1999b) and Antarctica (Retallack & Krull 1999)revealmanydetailsofthistimeof crisis (Figs19.8 and 19.9). In Antarctica and Australia very few fossil plants (only two species) and no comparable pedotypes are found on both sides of the boundary. In South Africa only six species of reptiles and amphibians are known to have survived. It was a clean and abrupt break in these terrestrial ecosystems. There is some evidence of environmental acidification at the boundary in Australian paleosol sequences, but not in those from Antarctica or South Africa. However, there is a profound and rapid decline in the carbon isotopic value (S13C)of pedogenic organic matter or carbonate in each of these sequences, as in marine boundary sections, where it is taken as evidence for an abrupt drop in biological productivity (Morante 1996; MacLeod et al. 2000). Early Triassic paleosols support this with indications of low fertility (oligotrophy) and little soil humus (Fig. 19.9). The Permian-Triassic boundary is also marked by a global proliferation of fungal hyphae and spores, as if forests were rotting on a massive scale (Visscher et al. 1996). There was no peat, and thus no coal, formed in humid
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SEQUENCE AND FOSSILS AT GRAPHITE PEAK
Figure 19.8 Field sketch of the Permian-Triassic paleosol sequence and selected fossils at Graphite Peak, central Transantarctic Mountains,Antarctica (fromRetallack&Krulll999: with permissionfrom theGeologicalSocietyof Australia).
lowland basins for perhaps the first 8 myr of the Triassic following extinction of the peat-forming floras (Fig. 19.4). In Antarctic and Australian sequences formed at high paleolatitudes (65-77's) Permian peaty paleosols are comparable with high-latitude soils of modern Siberia, but Triassic paleosols include such deeply weathered, warm-climate soils as Ultisols (see
Figs 4.9 and 10.11).This post-apocalyptic greenhouse probably was a time of elevated atmospheric methane (CH,) and of its oxidation product (CO,), which would explain the exceptionally light (negative) carbon isotopic composition of organic carbon in many earliest Triassicpaleosols (seeFig. 9.16; Ghosh et al. 1998;Krull &Retatlack2000).This warm thick atmosphere at high
Afforestationof theland
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Figure 19.9 Ecosystem changes inferredfrom paleosols across the Permian-Triassic boundary at GraphitePeak, Antarctica:left to right,proportionalrepresentation of vegetationinterpreted from percentage of pedotypesin that stratigraphicinterval,ecosystem stability inferredfrom paleosol degree of development, soil drainage inferredfrom depth of penetrationof fossil root traces (cm),soil fertilityinferredfrom basedaluminamolecularratios of paleosolsper year of interpretedsoil formation,soil humus inferred from chemicalanalysisof organiccarbon (fromRetallack & Krull1999; with permission of the GeologicalSocietyof Australia).
paleolatitudeswas slow to disperse, because normalcarbon isotopic values and coals reappear in Antarctica and Australia only by the Middle Triassic (Retallack & Alonso-Zama 1998). The perturbation that created such a massive methane release and consequent global havoc may have been impact of a large asteroid or comet, as there is evidence of shocked quartz and iridium anomalies at the boundary (Retallacket al. 1998), but a case can also be made for the role of large-volume volcaniceruptionsof the SiberianTraps,which began to erupt at the Permian-Triassic boundary (Bowring et al. 1998). Another mass extinction, this t i e of dinosaurs, is known at the Cretaceous-Tertiary boundary (65Ma), where there is copious evidence for asteroid impact in the form of strong iridium anomalies, abundant shocked quartz. grains, voluminous impact ejecta, fragments of a carbonaceous chondritic meteorite, and a large crater in Yucatan, Mexico (Alvarez et al. 1995; Kyte 1998). Paleosols across the Cretaceous-Tertiary boundary in Montana, USA, indicate much less severe ecosystem disruption than at the Permian-Triassic boundary (Retallack 1994a). There was an abrupt
change from well-drained to swampy soils, but this was not a profound or unusual change for this paleosol sequence. A post-apocalyptic greenhouse at the Cretaceous-Tertiary boundary indicated by carbonisotopic studies and gases occludedin amber lasted only some 50kyr or so (Landis et al. 1996; Arens & Jahren 1998), and was a time of widespreadand abundant fern spores early in the ecological successionof a devastated landscape. Although dinosaurs were extinguished entirely, this was a crisis less profound than that at the Permian-Triassic boundary.Extinctionof many species of evergreen angiosperms left vegetation dominated by conifers and deciduous angiosperms. Small detrivorous mammals, reptiles and amphibians survived, but not carnivores and herbivores. Many fish survived, but not freshwater clams and snails. These patterns of differential extinction are compatible with fallout of acid rain, mainly nitric acid from atmospheric oxidation and sulfuric acid from vaporization of evaporites by the impactor.Acid loads calculatedfrom the chemical composition of boundary beds and associated paleosols agree well with predicted acid generation by asteroid impact (Retallack1996b).
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. intertidal stromatolites . sabkha stromatolites salt marsh _____ mangal I sea arass freshwater stromatolites pond weed
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Figure 19.10 Geological ranges of plant formations, based on paleosols and associated fossils.
The few studies published on changes in paleosols across the Permian-Triassic and Cretaceous-Tertiary boundaries show differences and similarities between these major mass extinctions. Additional studies with wider geographical spread are needed, as well as studies of paleosols across other extinction events during the LateOrdovician (439Ma),LateDevonian (364Ma) and LateTriassic(206 Ma).
The shape of evolution The mid-Paleozoic advent of forests and the variety of soils that they formed was in some ways similar to major adaptive radiations of organisms, such as the diversification of vascular land plants. Soils, plant formations and organisms increased in diversity without extensive replacement of earlier kinds of soils or plants (Fig. 19.10). Ancient microbial ecosystems and soils persisted in other parts of the landscape along with forested ecosystems and their soils. Forest ecosystems
can be considered an adaptive breakthrough, which presented new opportunities for both organisms and soil-forming processes. The shape of evolution of both soils and organisms in such times of innovation is one of diversification. Such spurts of innovation punctuate a longer-term trend of diversification of life (Courtillot & Gaudemer 1996). The increased diversity and biomass of forests more firmly established changes initiated under early land plants. Greater biomass and root penetration more effectively held the landscape against erosion (see Fig. 18.11).Afforestation of watersheds further slowed percolation of rainwater into soils and its flow overland. Forests promoted meandering rather than braided streams and clayey rather than sandy soils (Schumm 1977; Cotter 1978). Not only was more oxygen photosynthetically produced by forests, but there also was a greater potential for burying carbon in wooded swamps (Graham et al. 1995). More thorough weathering also consumed atmospheric carbon dioxide, exporting more
Afforestationof the land bicarbonate and nutrients to the ocean, where it stimulatedmarineproductivity(Bambach1993)andbiomineraliation (Stanley & Hardie 1999). The physical stability of Late Paleozoic woodland vegetationresulted in burial of carbon on an unprecedented scale in the form of coal seams that now fuel industrial economies (Fig. 19.4). Carbon burial, atmospheric oxygenation and paleoclimatic cooling culminated in ice caps on the Gondwana supercontinent (Retallack 1999a,c). A novel new mechanism for burying carbon that appeared with trees was production of charcoal by wildfire.Fossilcharcoalisknownintherockrecordatleastas far back as Early Carboniferous(3 50Ma. Tournaisian) in the Berea Sandstone near Amherst, Ohio (Cope & Chaloner 1985). Charcoalis abundant in Late Paleozoic coals and red beds, and allowed oxygenation of the atmosphere by burying carbon in well-drained as well as swampy environments. Charcoal forms when oxygen
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available is inadequate for complete burning, usually because it cannot diffuse into burning wood sufficiently fast. For wood to burn at all requires at least 13%0, in the atmosphere(0.6PAL (present atmospheric level)).If oxygen were present at such low levels, then wildfires would be rare and require unusually dry wood. On the other hand, if oxygen became much more abundant in the past (as modeling suggests;see Fig. 16.7) then even wet forest would burn extensively and often,thus diluting atmospheric0, with carbon dioxide. Charcoal thus presents a biologically created mechanism for thwarting runaway oxygenation. Other natural regulators of the atmospheric balance of 0, and carbon dioxide include the balance between producers such as trees and consumers such as dinosaurs and the relative areas of carbon-storing swamps vs. carbon-oxidizingdesert scrubs. With trees came more complex, as well as more massive, environmentalengineering(Lenton 1998).
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Chapter 20 Grasses in dry continental interiors
Before the advent of grasses, dry continental interiors would have been vegetated by other kinds of plants. What these plants were like is difficult to say because well-drained soils of desert regions preserve plant material only under exceptional circumstances (Retallack 1998a), such as the urine-impregnated middens of packrats (Neotoma).Oases in deserts where plant fossils may be preserved support plants different from those found in the desert beyond. Thus, the fossil record of plants in dry regions is both meager and biased. Someindications of ancient aridland ecosystems are provided by fossil root traces (Loope 1988), phytoliths (Meehan 1994). bones (Loope et al. 1998) and burrows (Smith 1987) in calcareous red and brown paleosols. Large plants and animals play a conspicuous role in desert ecosystems, but productivity is limited and there is much bare earth exposed. Aridisolsforming in such environmentsdonot appear greatly different fromcalcareous paleosols as old as 2560Ma (Martini 1994). For many millions of years after the middlePaleozoic advent of trees, forests may have graded out through dry woodland of trees that became smaller in stature and more widely spaced toward scrubby desert vegetation. The middle Tertiary advent of grasslands in subhumid to semiarid regions signaled a new kind of ecosystem on Earth and new kinds of soils: Mollisols. Grasslands did not displace woodlands from humid regions nor desert vegetation from very dry regions, but were interpolated between the two older kinds of plant formation. Grasslands are fast-cycling ecosystems of high productivity on short time scales. In woodland soils, nutrients are leached by rainwater and bound up in standing wood, leaving little for large animals. In desert soils, nutrients are redistributed into duricrusts or remain unexploited in little-weathered minerals, because of inadequate soil water. Grasses, however, are fast-growingplants that areentiiely edibleboth for large animals above ground and for soil invertebrates within. Grassland soils have dark, well-structured surface hori300
zons (mollic epipedons), rich in nutrients in a readily available form (see Fig. 4.1 1).As a result they support a great diversity of mammals and other large vertebrates (Bell 1982; McNaughton 1985a). Rather than being communities intermediate between woodland and desert scrub in stature, density and productivity, grasslands were a new and distinctive addition to the array of Earth’s surface environments. Grasses are unique among plants in the many adaptations that enable them to withstand grazing (Fig. 20.1). Some grasses grow most rapidly when grazed (McNaughton 1983),as I suspected as a teenager given the chore of mowing our suburban lawn. In addition to unfolding from a zone of dividing cells at the tip (apicalmeristem) of the stem, grass stems elongate from zones of dividing cells above the nodes (intercalary meristems).The uppermost parts can be eaten without totally interrupting growth. The apical meristem of many grass shoots also is protected from damage because it is hidden down within the sheathing bases of theleaves. Inmostplants. by contrast, thegrowingpoint is at the tip of the shoot, and growth of the shoot is terminated if the tip is eaten. Grasses also have a dense growth close to the ground. It is a modular kind of architecture spread by repetition of the same simple units of runners or rhizomes and clumps of fibrous roots and linear leaves. Much of the plant’s rhizomes and growing leaf bases are buried within the ground where they are protected from trampling and grazing. In contrast, woody plants have most of their edible parts on display above the ground. Grass leaves are armored with a crust and small gritty bodies (phytoliths) of plant opal (Fig. 20.2). These are abrasive to teeth, carcinogenic and promote formation of calculi in the urinary tract of mammals (McNaughton 1985b). Horsetails and some dicots also are encrusted with silica,but few are so heavily armored as grasses. The siliceous skeletons strengthen bamboo and enable dead winter grass to remain standing long
Grassesin dry continental interiors
Figure. 20.1 Adaptivefeaturesof grasses,as shownby western wheatgrassAgropyron smithii.
30 1
western wheatgrass (Agropyron smithii)
after associated herbaceous plants have withered and decayed. Grasses also have insecticidal and fungicidal chemical defenses, such as cyclic hydroxamic acids (Freyet al. 19 9 7). If, despite all these defenses, grasses succumb to grazing and trampling, they have reproductive features that enable them to colonize scattered patches of bare ground. The plant body of grasses is small compared with other plants and their flowers are inconspicuous, small, and green or brown colored.They produce multitudes of small, smooth pollen grains and seeds that are dispersed by wind. This is very different from forest trees
such as Magnolia with their large, colorful flowers pollinated by beetles and their fleshy fruits dispersed by animals. It also is different from the showy bat-pollinated flowers of North American cacti. In contrast to these competitive and tolerant plants, grasses are weedy in their emphasis on reproduction independent of specific pollinating or seed-dispersinganimals (Grime 19 79). These features of grasses can be found also in their fossil record. The oldest fossils of grasses are simple monoporate pollen grains, which can be securely traced as far back as the Paleocene (65 Ma) and perhaps Cretaceous (90Ma: Jacobs et al. 1999). It is not until the
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Figure 20.2 Dumbbell-shapedopal phytolithsin the cuticle of a middle Miocene (14Ma)fossil grass, Cleistochloukubuyis (Dugas&Retallack1993) from theFort Ternan Member,Kericho Phonolite,near Fort Ternan.Kenya (KenyanNational Museum specimenFT 13120).Scale bar represents 20m.
middle Eocene (48 Ma) that fossils from carbonaceous shales show rhizomes, sheathing leaf bases, basal node meristems and the other distinctivefeatures of modern grasses (Crepet & Feldman 1991). The evolution of grasses in dry habitats is recorded by phytoliths in paleosols as old as middle Eocene (48 Ma) in the Sarmiento Group of Argentina (Spalleti& Mazzoni 1978). Highly siliceous fossil grass leaves and fruits are found in the Eocene-Oligocene White River Group of the central Great Plains of North America (Galbreath 1974: Meehan 1994). but are much more diverse and abundant in Miocene and Pliocene rocks of this region (Thomasson 1979). The diversification of these fossils has been claimed to represent the spread of grassland vegetation.It could equally be interpreted as the acquisition of silica bodies as a defense against grazing (Thomasson 198 5) within several evolutionary lineages of grasses that were already widespread.Neogene global spread of grasslands is recorded by both fossil cuticles andpollenof grasses (Fig. 20.3: Leopold&Denton 1987;Moreley & Richards 1993: Martin 1994). The evolution of grasslands also has left an imprint in fossil and modern mammals (Bakker 1983; Janis & Wilhelm 1993).Adaptations to grassland among mammals are especially striking because they are similar in
evolutionary lineages as independent as horses and antelopes. Grassland mammals have teeth with complexly infolded hard enamel and soft dentine that form low ridges on the flat, upper surface for grinding hard siliceous grasses. Their teeth also are high crowned (hypsodont) and in some cases continuously growing so that they continue to function despite wear. The magnificent grinding molars of modern grazing horses can be compared with the low-crowned molars of their ancestors thought to have browsed on leaves other than grasses (Fig. 20.3B). The ancestral tooth type for all mammals is unlike either of these: jagged with sharp cusps. Mesozoic mammals had teeth of this kind, as do many modern insectivores. The large size and elongate limbs of grassland herbivores such as horse and antelope also are distinctive. Their chief protection against predation is fleeing over open ground rather than climbing or hiding within woodlands. Their lower limbs (ulna and tibia) are elongated relative to the upper limbs (humerus and femur) which are heavily muscularized (Fig. 20.3C). The bones of the feet (metatarsals) and hands (metacarpals) are elongate so that they have become major limb bones. The bone (astragalus) connecting the hands and feet to their respective limb bones is ridged and pulley-like, to
Grasses in dry continental interiors
A
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EQUlD HYPSODONTY
c UNGULATE
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D EQUlD SIZE
Jegetation g10
m
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mQ)
2 20 0 ??
30
grass pollen 10 20 3040 50
percent
1
2
3
hypsodonty index
f
+ 0.50.6 0.70.8
metatarsal/femur ratio
200 400 estimated live weight (kg)
Figure 20.3 Paleontological proxies of grassland evolution:(A) abundance of grass pollen and charred grass cuticlein Nigeria (from Morley &Richards199 3); (B), degree of crownheight (hypsodonty)in North American horse molars (fromMacFadden 1992);(C)NorthAmericanungulatecursorialitybasedonmetatarsal/femurratios(fromBakker 1983):(D)North American horse size based on body weight estimated from measurements of molar volume (fromMacFadden 1992).
restrict lateral motion. This contrasts with the smaller toe bones and rounded astragalus of raccoons and humans designed for flexibility rather than speed. The hands and feet of grassland animals are held erect on their toes (unguligrade) rather than at a n angle to the ground as in dogs (digitigrade)or flat to the ground as in bears and humans (plantigrade).The fingers and toes of grassland mammals are reduced in number compared with woodland creatures. Modern horses have only a singletoe capped by a horny hoof. Three-toed horses and other forms relating them to the archaic mammalian design of five fingers and five toes are well represented in the fossil record (MacFadden1992).Mammals of grassland are also larger than most forest mammals, and grasslands include very few mammals in the size range 1-lOkg (Alroy 1998). Grassland mammals appear totally redesigned compared with forest and woodland prototypes. In North America these grassland adaptations are found in fossil mammals as old as Oligocene (33 Ma), when the distinctive 1-10 kg body size gap first appears in mammal assemblages (Alroy 1998). By the Miocene ( 1 7Ma), severalmammalian lineages evolved markedly
hypsodont teeth, cursorial limb proportions and large size(Fig.20.3B;McFaddenetaI. 1991).Hypsodontyand cursoriality of mammalian faunas in other parts of the world also provide a minimum age for grasslands of only late Miocene in Australia (Archer et al. 1994), middle Miocene in central Africa and Eurasia (Retallack 1991a) and early Oligocene in Argentina (Marshall & Cifelli 1990). Neither the fossil record of grasses nor that of mammals is an entirely satisfactory guide to when and how grassland ecosystems evolved. The earliest known fossil grasses appear to have been minor parts of woodland and swamp vegetation, and the earliest hypsodont teeth and cursorial limbs were adaptations to grasslands that existed earlier. These records do, however, constrain the origin of grasses and grasslands to some time during the Tertiary period and this can be assessedagainst independent evidence of paleosols.
Early grasslandsoils The fossil record of soils is in general agreement with evidence from fossil mammals and grasses that grass-
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land biomes emerged no earlier than Eocene.No definite grassland paleosols older than this have been reported. This is not to say that dry soils were not colonized by herbaceous vegetation until the advent of grasslands. Some red calcareous paleosols of Ordovician and Silurian age (Figs 18.4-18.7) supported early herbaceous polsterlands and brakelands in well-drained sites but lacked the intimate admixture of clay and organic matter, the granular structure and pervasive pellets and burrows of soils formed under grassland ecosystems. Pelletoidal fabrics in Early Carboniferous (Visean, 330Ma) paleosols (Wright 1983)are more l i e those of grassland soils, but lack the thickness of Mollisols and have drab colors and associated marine fossils indicating ancient salt marsh vegetation. Consideringevidence from the fossilrecord as well as from modern soils,grasslands evolved as a unique kind of ecosystem during the Tertiary period. Grassland paleosols can be recognized by a variety of features, but especially from abundant, fine (<2 mm diameter) root traces and small (2-5 mm) crumb peds. The fine structure of sod grasslands can be contrasted with the coarse structure of bunch grasslands and woodland soils, and remains obvious in paleosols (Fig. 20.4). Dark color and high organic carbon content also are diagnostic of modern grasslandsoils, but are seldom preserved in paleosols at original levels (Retallack 199 1b).There should also be evidence of high soil fertility in the form of smectite clay, feldspar and calcite. A
final useful indicator of some grasslands is the carbon isotope composition of soil organic matter, pedogenic calcite or mammalian tooth enamel (Fig. 20.5). Most plants use a C, (Calvin-Benson) photosynthetic pathway, which shows a marked preference for the light isotopeof carbon (l*Cratherthan "Corradiogenic 14C),so that carbon isotopicvalues of the plant material are low (613Cp,, -32 to -22%0), whereas tropical and warm temperate lowland grasses use a C, (Hatch-Slack) photosynthetic pathway, which is less isotopically discriminatory, so that carbon isotopic values of the plant are less negative (6l3CPDB -16 to -10%0).This striking difference in isotopic composition is then transferred to bonesandteethof mammalseating theplantsandtopedogenic carbonate, with some adjustments in absolute values for vital effects and for the mixing of atmospheric and respired soil carbon dioxide. There are problems with interpreting grasslands from isotopic values because grasses of cool climates and high elevations, and almost all known grasses geologically older than 7 Ma, used the C, pathway (Ceding et al. 1997a). Furthermore, an additional photosynthetic pathway known as CAM (crassulacean acid metabolism), found in a few desert and aquaticplants, coverstherange found in both C, and C, plants (Koch 1998).Such plants may explain the rare occurrence of isotopic signatures like those of C,plants before thelateMiocene(Bocherensetal. 1994; Kingston et al. 1994; Tanaka 1997; Mora et al. 1998). Because of such problems, isotopic studies need to be
Figure 20.4 Photographsof fossil peds in paleosols at the same scale (mm at bottom) showing nonmollic (A), near-mollic (B) and mollic (C) structure(fromRetallack1997b: withpermissionfromtheSocietyforSedimentaryGeology).
Grasses in dry continental interiors
8 m
B. Pakistan paleosol carbonate
A. Lubbock Lake organic matter 5'
n ! "
E 6-
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0
/i
m
7-
m
m
,"10-
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5-
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305
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UI
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E 15-
1
-15 -10 carbon isotopic composition (613C)
-20
20 J
-15
I
carbonate
1
1 o tooth enamel I
-10 -5 0 carbon isotopic composition (613C)
Figure 20.5 Carbon isotopic composition of organic matter (A) and carbonate (B) in paleosols of Holoceneage from theTexas High Plains (A) and of Neogene age from the Siwalik Group of Pakistan (B). showing a 10% shift to heavier isotopic values that can be interpreted as evidenceof the adventof C, grasslands replacing C, woodlandsof pinyon (A) and sal (B). The Pakistanishift is also reflectedincarbonisotopicshifts in associatedmammaliantoothenamel(datafromHolliday1992;Cerlingetal.199 3; although the former worker offers a different interpretation).
constrained by supporting studies. For example, fossil charcoal and Alfisols, followed by fossil Mollisols and grazing mammals, support isotopic indications of woodland replaced by grassland in the Miocene of Pakistan and Holocene of Texas (Fig. 20.5). Nevertheless, interpretations of past vegetation from isotopic data have proven controversial (Kingston et al. 1994; Holliday 1995). Dry Oligocenerangelands of South Dakota A transition from dry woodland to early rangeland has been documented from paleosols in volcaniclastic Late Eocene to Oligocene (Chadronian to Arikareean) alluvium of the White River and Arikaree Groups (see Fig. 6.2)inBadlands National Park, South Dakota, USA (Retallack 1983a,b).This is an area of spectacular exposures (see Fig. 2.5) about l00km east of the Black Hills and near the geographical center of the Great Plains of North America. Late Eocene dry woodlands are represented by Alfisol paleosols (Paleustalfs of Gleska pedotype), which have laterally continuous, drab, surface (A) horizons, over pink, clayey, subsurface (Bt) horizons, with abundant large drab-haloed root traces. This red-green banding
from paleosols dominates claystones of the Chadron Formation, and bestows their distinctive appearance. Permineralized fossil fruits and wood of walnut (JugZuns),locust (Robiniu)and keaki (ZeZkovu) have been found at comparable stratigraphic levels in nearby Nebraska (Wheeler&Landon 1992). During early Oligocene (Orellan, 33 Ma) fluvial deposition of the Scenic Member of the Brule Formation there was a greater variety of paleosols (see Fig. 6.2).The strongly developed pink and green Gleska Paleustalfs are still found, but have prominent calcareous stringers and are closely associated with sandstone paleochannels. Red Entisols (Fluvents of the Zisa pedotype) show relict bedding and only small root traces of early successional vegetation, and also are interbedded with stream paleochannels. Away from paleochannels are Conata paleosols: Inceptisols (Andic Ustochrepts) with limited differentiation of a surface (A) and subsurface (Bt) horizon, abundant small root traces and only scattered large, drab-haloed root traces of trees, and coarsegranularpedstructure(near-mollicof Fig. 20.4). These may have supported wooded rangeland on floodplains away from streams and their associatedearly successionalandgallerywoodlandvegetation(seeFig.6.3). A similar vegetation mosaic persisted in this region dur-
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Chapter 20
ing LateOligocene(Whitneyan)depositionof additional volcaniclastic alluvium (PoleslideMember of the Brule Formation).Diverse faunas of large mammals thrivedin SouthDakota at this time and their bones were wellpreserved in these calcareous paleosols (see Fig. 10.14). Phytoliths of grasses also are common in rocks of this age in nearby Nebraska and Colorado (Galbreath 1974; Meehan 1994). In later Oligocene times (Arikareean, 29 Ma) ashy alluvium (Sharps Formation) there is only one Inceptisol (Fluvaquentic Eutrochrept of Ogi pedotype) containing large, drab-haloed root traces as an indication of trees. This paleosol is enclosed within sandstones of a deeply incised stream as evidence of persistent streamside trees. Other paleosols at this stratigraphic level have thinner, calcareous (A-Bk) profiles and abundant small root traces. Those of the Pinnacles pedotype were probably Aridisols (Calcids), light colored with a shallow horizon (Bk) of calcareous nodules, platy with coarse granular surface structure and clumped root traces (Fig. 20.6). These near-mollic paleosols are like those that now support bunchgrass and sagebrush in dry intermontane rangelands. Andisols (UstollicVitrands) of the Samna pedotype higher in the sequence, away from paleochannels, have dark brown surface (A) horizons, a laterally contiguous network of fine root traces and thin, calcareous (Bk) horizons, Both Pinnacles and Samna paleosols may have supported open bunchgrass rangelands.
Figure 20.6 The type Pinnaclessilty clay loam paleosol (Calcid)in the Upper Oligocene (28 Ma)SharpsFormation of BadlandsNationalpark. SouthDakota,USA, showing the irregular thickness of the finely structured surface (A) horizon.This soil is interpreted as an earlyopen rangeland soil. Hammer handle is 28 cm long.
Middle Miocene short sod grasslands of Kenya Another transition from paleosols of forest and woodland to those of grasslands can be found in Miocene rocks of the Nyanza Rift, a south-west trending branch graben from the main Gregory Rift of Kenya (Retallack 1991a). During Oligocene and earliest Miocene time (20-25 Ma),forestsWerewidespreadinEastAfricaasan equatorial extension east across the continent from the jungles of Zaire. Their deeply weathered, clayey, red lateritic paleosols can be mapped throughout Kenya and Uganda (McFarlane 1976). Early Miocene (20-1 8 Ma) fossil localities of Songhor, Koru and Rusinga Island in south-west Kenya are within sequences of ashes and lava flows from local carbonatite-nephelinite volcanoes. Paleosols in these sequences have preserved a variety of fossil apes, snails and fruits typical of
tropical forest (Pickford 1986, 1995). Normally these phosphatic andcalcareousfossil hardpartswould bedestroyed in forest soils of humid climate (Retallack 1998a).The Miocene paleosols containing these fossils have remained calcareous despite their thick, clayey red profiles resembling tropical forest soils (Fig. 20.7). Comparisons of paleosols of different degree of development has shown that soil formation depleted carbonate content, but few became noncalcareous because the carbonatite parent material was very calcareous at the outset. These early Miocene red, clayey, blocky structured, carbonatite paleosols probably supported dry tropical forest, like those now found around the peripheryof theGuineo-Congolian basin (White 1983).
Grasses indry continental interiors
A. SONGHOR, KENYA
calcareousness develoDment hue
claystone
0sandstone conglomerate
~~~~~
@ relict bedding slump bedding
307
B. FORT TERNAN, KENYA
calcareousness development hue
brown crumb red color
$~,b&$~~d
zzpd
cross IU.1trough bedding I-1ripple marks
fossil stump
scour and fill
%calcareous nodules
Figure20.7 Columnarsectionsofpaleosolsequences(measuredinfield)fromtheearlyMiocene(20Ma)of SonghorNational Monument andmiddleMiocene(14Ma) of FortTernanNationalMonument, Kenya. Position of paleosolsmarked by boxes whose width correspondsto degree of development(seeTable 13.1).Calcareousnesswas estimated by reactionwith dilute acid (see Table 3.2) and hue with a Munsell color chart.
New kinds of soil formed in this region after the uplift of the western margin of the East African Rift and extrusion of plateau phonolites by middle Miocene time (14 Ma). These are well known in the sequence of paleosols in carbonatite-nephelinite tuffs sandwiched between thick phonolite flows near the village of Fort Ternan, south-west Kenya (Figs 20.7-20.9). In the main paleontologicalexcavation at Fort Ternan, two superimposed paleosols of the Chogo pedotype (Figs 20.8 and 20.9) have thin (10-12 cm), dark, crumb-textured, surface (A) horizons and a shallow (18-2 6 cm) horizon (Bk) of calcareous nodules and stringers. The Chogo paleosols may be identified as Mollisols (Haplustolls) of dry climate, as indicated by their shallow calcic horizons (see Fig. 9.5). Their vegetation was predominantly grassy considering the abundance of fine root traces, fossil grass pollen and calcareous, grasslike stem casts. There also were numerous trees as indicated by stump casts up to 16 cm in diameter, by fossil fragments of thorn-bearing twigs and by pitted fruit stones. These were mainly found in the upper paleosol in the western
part of the excavations, where there may have been an ecotone into a patch of grassy woodland around a small ephemeral watercourse that was excavated there (Retallack 1991a). The eastern part of the upper paleosol and the entire lower paleosol do not show such conspicuous evidence of trees and may have supported wooded grassland. A thicker but also crumb-textured and finely rooted paleosol of the Onuria pedotype, another Mollisol (Calciustoll),has been found capping a mudflow in the headwall of the large quarry at Fort Ternan National Monument. Its vegetation of at least five species of beautifully preserved grasses was discoveredin life position, buried by the overlying nephelinite grits (Dugas & Retallack 1993). This paleosol also has a shallow (19 cm) calcic (Bk) horizon. The depths to calcic horizons in Chogo and Onuria paleosols, corrected for likely compaction, are indications of mean annual precipitationof only282 f 141mm(seeFig. 9.5). Overall the middle Miocene environment around what is now Fort Ternan was probably a mosaic of short
308
Chapter 20
Figure 20.8 Middle Miocene (14Ma) paleosols of the main fossiliferouslayer in the excavationat Fort TernanNational Monument, Kenya.The top of the upper paleosol,the Chogo clay erodedphase paleosol,aMolliso1(Haplustoll),is at the hammer head,and only the top of the Chogo clay ferruginizednodule variant paleosol is exposedat the base of the excavation.The hammer handleis 28 cm long.
sod grassland (Onuria), wooded grassland and grassy woodland (Chogo,Dhero) (Fig. 20.10).This reconstruction of an early grassland mosaic is also supported by associated fossil grasses (Retallack 19926 Dugas & Retallack 1993), fossil grass pollen (Bonnefille 1994), snails (Pickford 1995), antelope (Scott et al. 1999)and carbonisotopic studies (Cerlingst al. 1997b).Although this early grassland would have appeared generally similar to the modern Serengeti Plains of Tanzania, it differed in having grasses with C, rather than C, photosynthetic pathways, and so may have been more productive than the Serengeti (Koch 1998; Jacobs et al. 1999). Late Miocene tall sod grasslands of Oregon Yet another transition from woodland to grassland is recorded in paleosols underlying the Rattlesnake Tuff (Late Miocene, 7.2 Ma) near Dayville, Central Oregon, USA(Retal1ack 1997b).Thelowestpaleosolswithin this 77m sequence are thick (125 cm), clayey, brown, blocky structured profiles with large root traces, comparable with soils of woodland.Paleosolsof theTatas pedotype, without such aprominent clay bulge, and with the abundant fine root traces and crumb peds of grassland soils, appear higher in the sequence here (Fig. 20.11), but still below the Rattlesnake Tuff. These crumbtextured, finely rooted paleosols have diffuse carbonate nodules at depth (103 cm), which when corrected for compaction can be used to infer a former mean annual
precipitation of some 682 f 1 4 1mm. Thisindication of good rainfall and other grassland features of the paleosols mark them as the oldest known paleosols of tall sod grasslands, like those found now in subhumid parts of the Great Plains of North America. Also found in Tatas paleosols were remains of the horse Pliohippus spectans, which has the very highcrowned teeth, slender monodactyl limbs and large body size of modern grazers (Fremdet al. 1997). Pollen of grasses also becomes regionally abundant in the Pacific North-west during late Miocene time (Leopold &Denton 1987).
A timetable A general schedule of grassland appearance and spread has emerged from reconnaissance studies of Neogene paleosolson threecontinents (Fig. 20.12). Desert grasslands and rangelands can be inferred from paleosols with near-mollic structure and many fine root traces as ancient as early Oligocene (33 Ma) in both Oregon and the Great Plains of the USA. The earliest sod grassland paleosols with fine crumb peds and a dense mat of fine rootlets appear during the early middle Miocene (20-14Ma) in North America and East Africa. In each case, these have shallow calcic horizons indicating dry climates and thus short grasslands like those of the western side of the North American Great Plains today. By late Miocene time (7Ma) in North America, Africa and Pakistan the earliest tall grassland ecosystems can
A
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MOLECULAR WEATHERING RATIOS
1
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brown color
relict bedding ripple marks
K Y
5
Figure 20.9 Columnar sections (measuredinfield)and petrographicand chemicalcomposition(estimatedby point counting and inductivelycoupledplasma emission spectrometry)of Chogo andOnuriapaleosols(middleMiocene,14Ma)at FortTernanNationalMonument, Kenya (fromRetallacketd. 1990;withpermissionfromthe AmericanAssociationfor the Advancementof Science).
dark gray (IOYR4/1)
grayish h
brown (IOYR5/3)
bmwn-dork brown (75YR4/41
very dark gray (IOYR3/1)
grayish brown (IOYR5/2)
grayish brown (IOYR4/2)
grayish brown (K)YR5/2)
dive yellow (2.5Y6/6)
light olive brown (2.5Y5/43
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Figure20.10 Areconstructionof soilsandvegetationduringthemiddleMiocene(l4Ma)intheareaaroundwhatisnow Fort Ternan.Kenya(fromRetal1ack199 la: withpermissionfromOxfordUniversityPress).
be recognized from paleosols with crumb structure, fine root traces anddeep calcic horizons. Compared with paleosols of the earlier Miocene,these paleosols of the late Miocene indicate expansion of the grassland-woodland ecotone from the 400 mm isohyet of mean annualrainfalltowetterregionsof the 750mmisohyet (Fig. 20.13).
Such an expansion of climatic range probably also means a dramatic expansion of grassland’s geographical range, regardless of the great local variability in paleoclimate observed in these paleosol sequences (Retallack199 7b). Such a schedule of appearance and expansion re-
Grasses in dry continental interiors
3 11
Figure 20.1 1 The Tatas pedotype paleosol is finelyrooted and crumb textured beneath a white volcanic ash in late Miocene (7Ma) sediments associated with the Rattlesnake Tuff near Dayville,Oregon, USA. This is interpreted as a soil of tall grassland. Hammer is for scale.
conciles differing opinions, arising from different kinds of study, on the antiquity of grasslands. The early Oligocene (33 Ma) advent of desert grasslandsin North America corresponds to mass extinction of forestadapted large mammals (including titanotheres such as Menodus). reptiles, amphibians and snails, and the appearance of abundant grass phytoliths (Retallack 1992a;Meehan 1994).TheearlyMioceneadventof sod grasslands in North America corresponds to dramatic
A
B
OREGON
increases in fossil horse sue, hypsodonty and cursoriality(Fig.20.3; MacFaddenetal. 1991).ThelateMiocene advent of tall grasslandsin North America, East Africa and Pakistan corresponds to isotopic evidence for the appearanceof C, vegetation at each of these places (Fig. 20.5; Cerlingetal. 1997a).Thisalsowasatimeof horse extinction in North America,with only the large,highly hypsodont, monodactyl horses, interpreted as grazers. surviving (MacFaddenet al. 1999).
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200 400 600 800 200 400 600 800 mean annual precipitation (mm) mean annual precipitation (mm] mean annual precipitation (mm) mean annual precipitation (mm) paleoclimatic range of sod paleoclimatic range of paleoclimatic range of estimated precipitation , paleosols inferred from paleosols inferred from and standard deviation grasslands inferred from degree of weathering for mollic paleosols depth to calcic horizon depth to calcic horizon ~
a
Figure 20.12 The distribution of mollic paleosols intime and on a gradient of mean annual precipitation inferred from depth to calcic horizons in the same paleosols of (A)Oregon, (B) South Dakota-Nebraska, (C) East Africa and (D) Pakistan, showing desert grasslands as far back as Oligocene( 3 3 Ma), short sod grasslands back to early middle Miocene (20-14 Ma) and tall sod grasslands back to late Miocene ( 7 Ma).
312 iom
Chapter20 desert
desert
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PLIOCENE
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MIOCENE
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OLIGOCENE
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How did grasslandsarise? Grasslands have several distinct advantages over other kinds of vegetation.These can be used to assess the fossil record of grasslands. Grasses flourish more than trees in dry or otherwise unfavorable, highly seasonal and unpredictable environments (Scholes & Archer 1997). This includes the cool temperate, semiarid Russian steppe, North American prairie and Patagonian pampas, as well as seasonally dry subtropical and tropical wooded grasslands, salt marshes, coastal dunes, streamsides and alpine meadows. Grasses survive in these difficult environments because of their small stature, unspecialied pollination and dispersal mechanisms, and protection of their tis-
800
mollic brown soil
Figure 20.1 3 A scenario for climatic and geographical expansion of grasslands and their soils in the Great Plains of North America(fromRetal1ack 1997b:with permission of the Societyfor Sedimentary Geology).
sues in rhizomes or other underground structures. Few plants are better suited than grasses to impersistent habitats such as seasonally snowy plains, floodprone streamsides and human construction sites. Evidence from paleosols shows that young floodplains and volcanic ashes rich in nutrients that are easily weathered became widespread during the Oligoceneappearance of early rangelands in North America (Retallack 1983a; Bestland 2000). Animal activity also may play a role in maintaining modern grasslands. The grazing and trampling activity of large herds of antelope or horses keeps East African game parks almost as well groomed as a city park (McNaughton 1984). Elephants knock down and tear
Grasses in dry continental interiors apart trees in the dry season to consume their bark and leaves (Owen-Smith 198 7). Plagues of grasshoppers and other insect pests may defoliate large areas, leaving them open for colonization by grasses. Another factor encouraging development of grasslands is competition among plants. Although trees may have an advantage in shading out and chemically poisoning lower-growing plants, grasses can persist by shading out small seedlings of trees. Fewkindsof plantscreatesuchadense andcontinuous ground cover as grasses. Such biologicalfactors may have played arolein the early to middle Miocene advent of sod grasslands, which correspond to marked advances in mammalian grazer size, hypsodonty and cursoriality (Fig. 20.3B-D), andin phytolith abundance anddiversity (Thomasson 1985:Retallack 1992f). Two final factors promoting grasslands are fire and low atmospheric levels of carbon dioxide. Grasses recover quickly from fire because their fire-sensitivemeristems and rhizomes may be underground and because they have little woody tissue that will burn long and hot. Grass leaves that have died back during winter or drought are especially flammable. Ground temperatures 3 cm below grass fires are only 3°C higher than usual for as little as a few seconds as fire rapidly consumes available fuel and moves on (Gillon 1983). In contrast, grass fires are lethal for small shrubs and trees. Many modern grasslands are thought to be maintained by periodic fire (Johannessen et al. 1971). Grass fires are encouraged by atmospheric oxygenation and low levels of atmospheric carbon dioxide. Low carbon dioxide ( 4 0 0 or micro-atmospheres) promotes the spread of plants using the C, photosynthetic pathway, which are mainly lowland tropical grasses (Ceding et al. 1998).The C, pathway is more conservative of carbon dioxide because of an additional photosynthetic step to exhaust unused carbon dioxide in the bundle-sheath cells characteristic of such plants. An atmospheric carbon dioxide drawdown from higher earlier Tertiary levels tolevelsmoreliiethoseof theQuaternary(400pm) has been detectedduring the late Miocene from changes in stomata1 density of fossil leaves (van der Burgh et al. 1993), but is not apparent from deep oceanic isotopic records (Pagani et al. 1999).An upswing in fire frequency is also evident from charcoal in late Miocene paleosols and sediments (Fig. 20.3A). The combination of declining atmospheric carbon dioxide and increased fire frequency may explain the globallysynchronous ad-
3 13
vance of C, grasslands during climatic cooling at about 7 Ma. If we adopt an Ereban view, grasslands were victims of global climatic changes (Janis & Damuth 1990), produced by such forcing factors as mountain uplift (Raymo&Ruddiman 1992)orchangesinoceaniccirculation pattern (Broecker 1997; Ramstein et al. 1997). Mountain uplift, particularly Himalayan, is envisaged as deflecting tropical wind patterns into cooler highlands with glaciers that radiated solar energy back into space, and that ground up rock as a potential sink for carbon dioxide by weathering. A fundamental problem with this last idea is that glacial loess and till on the Tibetan Plateau is in regions too cold for effectiveweathering. Finally, the uplit itself may have been generated by isostatic unloading because of glacial scouring as a result of, rather than producing, climatic cooling (Molnar &England 1990). Oceanic explanations for Cenozoic climatic change draw primarily on the global heat pump driven by sinking of saline cold water in the North Atlantic, that then runs deeply south through to the Southern Ocean and upwells off the coast of Chile, gathering heat and salt through the shallows of the tropical Pacific and Indonesia for surficial return on ocean currents through the Indian Ocean to the North Atlantic. This global thermohaline conveyor does much to distribute heat evenly around the planet, and cooling could be induced by local restrictions to flow such as the closure of the Mediterranean and earlier equatorial oceans with the northward continental driit of Africa and India (Broecker 1997: Ramstein et ul. 1997). Unfortunately, Mediterranean enclosure was already far advanced during the Paleocene, and the main trendof continental drift since then has been a northward drift of most continents, opening, rather than closing, the Southern Ocean to increased thermohaline flow. This would be expected to result in greater heat distribution and more equable climates,rather than the episodic coolings observed. The principal restriction to the oceanic thermohaline flywheel, which tends to warm the world, has been expansion of Antarctic glaciers to sea level and then the spread of sea ice. Again, this could be a result of climatic cooling, rather than producing glaciation. An alternative, more Gaian view of grassland expansion and climatic cooling is also plausible (Retallack
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Chapter20
1 9 9 8 ~ )The . common thread to each of the observed episodes of grassland expansion and climatic cooling at roughly 33.20 and 7Ma is drawdown of atmospheric CO,, H,O and CH, from an earlier Cenozoic greenhouse. Grassland soilsthemselves may have played a role in this atmospheric greenhouse drawdown by rapid early successional colonization of nutrient-rich, new surfaces (Bestland 2000), and by creating soils of 6ne soil structure with high internal surface area (Retallack 1 9 9 8 ~ ) . Both processes promote weathering, which consumes atmospheric carbon dioxide as carbonic acid and exports it to the ocean as bicarbonate. Grassland soils also have high organic carbon contents (up to l O w t % ) to much greater depth (typically 1min tallgrassland) than in thedrywoodlands anddesertsthat theyreplaced (Fig. 20.13).Grassland soilscould thus have served as an important new sink for carbon dioxide both by weathering and organic matter accumulation. Grasslands also have high albedo and low evapotranspiration, thus lowering the water vapor content of air above them more than woodlands and forests. Suchdrying would also diminish the area of swamps as a source of CH,. Erosion and burial of moist, organic peds of grassland soils was a longterm sink for potential greenhouse gases. By this view, climatic drying, cooling, and drawdown of CO,, H,O and CH, would be consequences of the coevolution of grasses and grazers. Such coevolution was already well establishedby middle Miocene time (14 Ma) at Fort Ternan, Kenya, where fossil grasses are richly invested in phytoliths, grasses have low stomata1densities expected in open vegetation, antelope teeth have strongly striated microwear created by phytoliths, and antelope have the wide muzzles of grazers (Retallack 19926 Cerling et al. 1997b). Such adaptations reflect mammalian grazers’ ability to cope with coarse gritty graze and grasses’ ability to withstand trampling and cropping better than other plants. Grassland communities can be envisaged as a novel biologicalforce,driven by grass-grazer coevolution, that thrived in earliest Oligocene disturbance and then made great territorial gains in both the early and late Miocene. The question before us then is whether grasslands were born in crisis of global change as a result of mountain uplift or oceanic current diversion, or whether grasslands expanded their range by virtue of coevolution that had global geochemical consequences. There are not yet clear answers to such questions, but ma-
terials for addressing them can now be better identified, and these include paleosols.
Evolutionaryprocesses Compared with other kinds of paleosols, those formed under grasslands are unusually fossiliferouswith bones, snails, phytoliths and stony fruits. This is because the calcareous composition of these soils is especially favorable for preserving such hard parts (Retallack 1998a) and because they are such highly productive ecosystems (Bell 1982; McNaughton 1985a). Grassland paleosols and their fossils present a unique opportunity to study evolutionary processes both from the point of view of adaptations and from the point of view of the environments that selected for them.
Speciation Evolution of new species occurs on time scales that are too long to be revealed by study of modern populations and their ecology. The fossil record, in contrast, is replete with long-ranging fossil speciesthat can be grouped into complex genealogiesthought to have been produced by evolution. Yet the actual process of evolution is difficult to detect because it operates on time scales shorter than the million year resolution of many geological sequences. Thus, it is difficult to judge even simple questions about evolutionary change, such as how fast it is. Two extreme positions can be imagined (Gould & Eldredge 1977). On the one hand, it could be that evolution proceeds by slow modfications of morphology that are imperceptible from one population to the next, but that amount to distinctive differences over time. This phyletic gradualism can be contrasted with punctuated equilibrium, by which evolution occurs in spurts of geneticchangeover ageologicallyshortperiodof timethat is nevertheless long by human standards. Species produced in this way then persist for millions of years virtually unchanged. Whether either of these views or some intermediate one represents the more general tempo of evolution can be studiedusing the fossilrecord, provided there is some way of estimating the time elapsed between fossil samples. In nonmarine vertebrate-bearing sequences paleosols may fill this role because the time scales of soil formation are intermediate between those of ecology and geology.
Grasses in dry continental interiors From this perspective, the Oligocene sequence of paleosols in Badlands National Park, South Dakota, USA, is incomplete in a variety of ways. Not only are many of its paleosols well developed,but this aggrading sequence of paleosols is also punctuated by several episodes of erosion that may have lasted several million years (see Fig. 13.14).The vertebrate fossils of this region are not an adequate test of the evolutionary models of punctuated equilibrium or phyletic gradualism. Fossil mammals found there do seem to be stable in shape and size within each depositional unit and this could be interpreted as evidence for evolutionary stasis following rapid evolutionary origin (Prothero & Whittlesey 1998). However, it could also have been the result of a period of phyletic gradualism during the long periods of time unrecorded. Evidencefrom paleosols for the temporal resolution of this sequence compromises its relevance for fine-scaleevolutionary studies. A more promising study area has been identified in the paleosol sequence of the Early Eocene (Wasatchian) Willwood Formation of Wyoming (Bown&Kraus 198 1, 1983: Bown&Beard 1990).Eachsample of vertebrates used here for evolutionary studies has been collected from as many as four paleosols (Gingerich 19 80).which together may represent 50kyr or so of soil formation. This resolution can be improved by sampling paleosol by paleosol (Winkler 1983). Existing data on mammalian evolution from this sequence is suggestive of both phyletic gradualism and punctuated equilibrium for different features and evolutionary lineages. Further studies of this kind are needed to determine which is the more usual evolutionary pattern and under what circumstances each occurs. Natural selection In studying mammalian evolution from the fossil record, it is natural to focus on what was selected: the fossilmammals. In such studies, it is common to assume that their various features such as high-crowned teeth and elongate limbs were optimally designed for a particularpurpose. In this case the purpose would be eating and avoidingbeing eaten in open grasslands. It is this assumption that allows reconstruction of former vegetation from presumed adaptive features of mammalian fossils. This assumption of optimal adaptation has been lampooned (by Gould & Lewontin 1979) as the Pan-
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glossian paradigm, after the character in Voltaire’s novel Cundide who declaimed through a series of misfortunes that ‘all is for the best in this best of all possible worlds’. What is needed to avoid such false optimism is independent evidence of paleoenvironment against which the degree of adaptation or maladaptation of species or communities can be assessed. In some cases, paleosols can play that role. A Panglossian view of Oligocenemammalian faunas of Badlands National Park, South Dakota, USA, is that they lived in forest, because they resemble modern forest faunas in a variety of adaptive features (Van Valkenburgh 1985; Janis & Wilhelm 1993). From the new perspective offered by paleosols, Late Oligocene faunas there can now be seen as maladapted to grasslands compared with Miocene faunas of Fort Ternan in Kenya or of Nebraska, USA, or with modern faunas of these same regions. Nevertheless, these Late Oligocene faunas persisted for millions of years in dry open rangeland despite their apparent inadequacy. Similarly, an antiquated technology such as hand copying of manuscripts can remain viable in the context of its own times. It is only in hindsight of printing or photocopying that it seems archaic. If vegetation and climate are not selecting for nearconstant adaptive grades in mammalian populations over geologicaltime scales, then what is?One possibility is that evolution is constrained by difficultiesin altering highly co-ordinated biological systems for reproduction, growth and development.However, this also seems unlikely considering evidence for occasional rapid evolutionof individualspecies (Gingerich 1980)and faunal overturn of mammalian communities (Prothero & Whittlesey 1998). Could it be then that evolutionary novelties are shouldered aside by the web of ecological interactions in natural mammalian ecosystems? One way of approaching this question is to compare geological stability with ecological integration of fossil mammalian assemblages,as in the following section. Coevolution Assemblages of organisms that live together often are called communities and their various component organisms labeled as consumers, producers or similar terms that reflect their role in maintaining the community. An extreme version of this marketplace view of
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Chapter 20
ecology is that communities can be regarded as ‘superorganisms’ in which particular species are as indispensable to the function of the community as organ systems of a single creature (Clements 1928). Extending this concept, extreme versions of the Gaia hypothesis view the Earth itself as a ‘superorganism’. If communities really are so highly integrated, then presumably their evolution has more to do with biological forces such as competition between species and individuals (Van Valen 1973; Dawkms & Krebs 1979), independent of all but the most profound environmental perturbations (Jablonski & Sepkoski 1996). A contrasting, equally extreme and more Ereban view is that communities are only accidental associations of species whose abundance is determined by gradients in environmental factors, such as moisture, rather than by interaction with associated organisms (Whittaker 19 78).By this view, associations of organisms are difficult to definein both space and time.Their evolution and migration is a direct reaction to environmental perturbations such as climatic change (Bernabo & Webb 1977). Grassland ecosystems,as understood at present from the fossil record of mammals, grasses and paleosols, appear to have evolved from something a little l i e an accidental association to something a little more like a superorganism. It is remarkable that similar kinds of mammals persisted in South Dakota across the Eocene-Oligoceneboundary (33 Ma), despite profound changes in vegetation from woodland to rangeland
(Prothero & Whittlesey 1998), although both hypsodonty and cursoriality of some lineages increased slightly at this time (Bakker 1983; Janis & Wilhelm 1993). In this case, mammalian assemblages were remarkably stable in the face of aprofound environmental change. Greater interaction between mammals and vegetation is evident from the hypsodont and cursorial immigrant Miocene grassland ecosystem of East Africa, where this mammalian fauna also forms a recognizable dynasty (Pickford 1986). Similarly,in North America a more hypsodont and cursorial mammalian community established itself as a recognizable dynasty with the early Miocene spread of short sod grassland, and was severely tested by late Miocene spread of tall grassland. Mammalian communities following the late Miocene extinctions were much less diverse and retained mainly large, highly hypsodont and cursorial taxa such as the familiar monodactyl horses of today (MacFadden et aJ. 1999). Climatic and environmental change were extreme during Pleistocene ice ages with surprisingly little effect on mammalian community evolution (Barnosky 1994). In the long term, mammalian communities appear to have become more robust in the face of external environmental change, as a result of evolutionary improvements in the design of individual species and their web of interactions. Whether or not these tentative conclusions withstand future scrutiny, the fossil record of mammals and of vegetation revealed by paleosols can be seen as a fertile testing ground for basic ideas about evolution.
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Glossary
A horizon surface horizon of a soil, commonly including organic matter. AA or AAS atomic absorption spectrometry. accretionary lapilli small stones of volcanic tuff with concentric internal layering. acetate peel method of preparing large thin sections of carbonate rocks by means of acid etching and overlayingof acetone-softenedclear acetate sheets. acidic with a low pH (<7), and abundance of hydronium ion (H') in solution. adventitious roots roots attached to the stem rather than to the base of the plant. aerenchyma spongy,thin-walled, loosely meshed tissue, found within peg, propandkneerootsof swampplants. aerophore thin-walled area of epidermis,found in peg, prop andknee rootsof swampplants. agglomeroplasmicsoil microfabric in which there is an incomplete or local fine-grainedmatrix to skeleton grains. aggrotubule tubular feature of soilfilled with pellet-lie clasts made of clay and clastic grains. aging upwards sequence sedimentary successionincluding paleosolswhose degree of developmentis generally stronger for paleosolshigher within the succession than for those near the base. AlbaqualfAlfisolthat shows gley features and a sandy nearsurface horizon above the clayey subsurface horizon. albic horizon light-colored soil horizon characterized by less organic matter, less sesquioxides(Fe,O, and Al,O,), or lessclay than the underlying horizon: itslightcolor isdue largely to quartz and feldspar. albite feldspar mineral rich in sodium (NaAlSi,O,). AlbollMollisol with an albic horizon. A l k l fertileforest soil, with subsurface argillic horizon. alginitekind of liptinite (exinite)coal maceral, formed from unicellular algae,particularly botryococcoidalgae. alkalielementssodium (Na)and potassium (K). alkalinewithahighpH(>7),soalowactivityof hydronium (H+)ions in solution. alkaline earth elements calcium (Ca)and magnesium (Mg). alkaliation soil-formingprocess of accumulation of sodium and other alkaline elements in clays. allophane poorly orderedhydrous aluminum silicate.
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alluvial fan large conical landform deposited by streams entering intermontane basins fromnarrow mountainvalleys. alluvium sedimentary deposits of rivers. alumina aluminum oxide (A1203). alveolar-septa1structure micromorphologicalstructure of calcareous soils,with thin micritic compartments filled with sparry calcite,thought to form around fungi associated withroots. amphibole group of minerals including hornblende and tremolite. amphibolitefacies metamorphic assemblageformed at high temperatures (>SOOT),characterized by minerals such as hornblende, andalusite and sillimanite. anaerobic decay a metabolic reaction of organisms in which organic matter is broken down in the absence of oxygen into simpler compounds, including carbon dioxide. anatase brown, dark blue or black, high-relief mineral (TiO,). andesite silica-saturated calcakaline volcanicrock, often porphyritic with large crystals and groundmass commonly plagioclase. andalusite metamorphic mineral (AI,SiO,). andic withpropertiesli those of an Andisol. Andisol Inceptisol-lie soilsformed on volcanicash, with low bulk density,high porosity and great fertility. andisoliation soil-formingprocess of low-density,highporosity, fertile, noncrystalline weathering products from glassy volcanic shards. anorthite feldspar mineral rich in calcium (CaAI,Si,O,). anthoecium (plural anthoecia) flower and subtending bracts of grasses. anthracite grade or rank of compact, shiny coal, formed under burialconditionscorrespondingtoincipient metamorphism. anthropicepipedon surface horizon of a soil altered by human activity,such as plowing. anthrosoliation soil-formingprocesses of human disturbance. Ap horizon A horizon disrupted by plowing or other comparable disturbance. AqualfAKsolwith gley features. Aquand Andisolwith gley features. Aquent Entisolwith gley features.
Glossary Aquept Inceptisolwith gley features. Aquert Vertisol with gley features. aquic showing gley features. Aquod Spodosolwith gley features. AquollMollisolwith gley features. Aquox Oxisolwith gley features. Aquult Ultisol with gley features. aragonite mineral with the same chemical composition as calcite (CaCO,). but a differentcrystal structure, primarily found in shellsof mollusks. Archean era of geologicalt i e before 2 500Ma. Arent Entisol with surface layers mixed by plowing or other human activity. Argid Aridisolwith argillic or natric horizons. argillan cutan consisting of clay argillasepicsoil plasmic microfabricof mainly clay, lacking highly birefringent streaks when viewed in thin section under crossednicols. argillic horizons soil horizon of clay enrichment, recognized in the field by orientatedclay films that coat mineral grains, small channels or ped surfaces. Compared with eluvial horizons, argillichorizons have 3%more clay if eluvial clay is 10-15%, 12%moreclay if eluvialhorizon clay is 1 5 4 0 % or , 8%more if eluvial horizon clay is 40-60%,or 8%moreheclayif totalclayismore than 60%. argilluviation soil-formingprocess of washing down of clay deeper into the soil within cracks. ArgiustollUstoll with a clayeysubsurface horizon. Aridisol desert soil, usually thin profiles,commonly with calcareous nodules or salt crystals within l m of the surface. arthropod phylum of jointed-leggedanimals, including insects, spiders and crafish. asepic soil plasmic microfabric lacking highly birefringent streaks (plasma separations) when viewed under crossed nicols. atomic absorption spectrometry method of chemical analysis using light absorption wavelengths of flaming aerosol of sample. axial plane cleavagecleavage orientated parallel to the axis of the fold of folded metasedimentary rocks. B horizon subsurface horizon of soil, often enriched in clay or carbonate. backfillstructure sinuous or simplycurved layering within the filling sediment of a burrow produced by movement of an animal. badlands erosional landform of deeply gullied sedimentary rocks too unstable and in too dry aclimate to support a cover of vegetation. basalticrocksrockslike basalt, whicharerichinironandmag-
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nesium, h e grained and generally produced as volcanic flows. basesaturation percentage of the cation exchange capacityrelated to bases. bases principal cations of soil solutions (Ca2+,Mg2+.Na', K+). basket podzol locallythickened sandy subsurface (E or eluvial) horizon under an individual tree. bauxite highly weathered material rich in aluminum, and poor in humus, silica and bases, consisting mainly of gibbsite or similar minerals. Bc horizon B horizon with concretions or nodules. BghorizonB horizonwithstronggleying. bicarbonate common anion (HC0,-) in soil solution. billet small sawn slab of rock used to prepare a petrographic thin section. bmasepic sepic plasmic fabric with a network of highly birefringent streaks in two preferreddirections. biofunction mathematical relationship between a measured soilfeature andsome measure of soil biota. biogeochemicalcycling soil-formingprocesses of biologicalenhancement of base cations and other soil fertility. biosequenceset of soilsformed under similar climate, topographic setting, parent material and time, but differentvegetation or other organisms. biotite sheet silicate mineral of igneous and metamorphic rocks [K,(Mg,Fe2+),~(Fe3+,A1,Ti),,,(Sih3A1,_,0,,) (OH,F),l. bioturbatedmixed and moved by the burrowing, rooting and other activities of organisms. bmfringence interference colors of minerals when viewed in a microscope under crossed nicols. bunessite poorly crystallme mineral of dark iron and manganese oxides. bituminous coal grade or rank of black, dull to shiny banded coal, formedunder burial conditions close to those of the oil window. Bk horizon B horizon with accumulation of carbonates. usually calcite nodules. blocky peds a form of ped that is polygonal and nearly equant inshape. Bn horizon B horizon with accumulation of sodium. Bo horizon B horizon with residual accumulation of sesquioxides. boehmite aluminum hydroxide [y-AIO(OH)]. bog general term for wetland vegetation, often used for vegetation of mosses. BoralfAl6solof cool to coldclimates. Boroll Mollisol of cool to cold climates. botryococcoidalgae unicellular aquatic algae with thick waxy cell walls (Botryococcus).
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Glossary
brachiopod phylum of marine invertebrates with two-valved shells and lophophore. brakeland vegetation of well-drainedsoils formed of rooted herbaceous plants but no grasses. brecciacoarse-grainedrock with angular clasts. breunneritecarbonate mineral [(MgFe)CO,]. Bs horizon B horizon with illuvial accumulation of sesquioxides. Bt horizon B horizon withilhvial accumulation of clay. bulb plant storage organ of swollen stem base and leaves,as in onions. bulk density ameasure of mass in a given volume (grams per cubic centimeter or g cm-’): usually indicated by symbol p. burrow tunnel or other excavation of a soil animal. butte isolated hill or mountain with flat top and steep sides, usually having a summit widthcomparable with its height. Bw horizon withcolored or structural B horizon. By horizonB horizon with accumulation of gypsum. Bz horizonB horizon with accumulation of salts. C horizonsubsurface soil horizon, excluding bedrock, slightly more weathered than the material from which the soil formed or is presumed to have formed.Lacks properties of A andB horizons, but includes weathering as shown by mineral oxidation, accumulation of silica,carbonates or more soluble salts, and gleying. calcareous consisting largely of calcite. calciasepicsoil plasmic microfabricdominated by a mixture of clay and clay-sizedcarbonate, and lacking highly birefriigent streaks when viewed in thin section under crossed nicols. calcic horizon subsurface soil horizon enriched in calcite or dolomite in the form of coatings, wisps or nodules, and at least 1 5cm thick with at least 5% more carbonate than underlying horizons. Calcid Aridisol with calcic or petrocalcic horizon less than 1m deep. calcificationsoil-formingprocess of the accumulation of carbonate, usually as nodules in subsurface horizons. calcite carbonate mineral (CaCO,). calcrete rockcemented with calciumcarbonate. caliche pedogenic calcium carbonate nodules or layers. calorificvalue heat liberated by combustion of coal, measured in British thermal units (BTU) or kcal g-’ . cambic horizon subsurface soil horizon with at least enough pedogenicalteration to eradicate some rock structure,form some soil structure, and remove or redistribute primary carbonate. Its color has higher chroma or redder hue than does the color of theunderlying horizons. CambidAridisolwith weakly developedsubsurface (cambic) horizon.
Cambrian period of geological time about 5 4 4 4 9 5 Ma. cannel coal coal with dull waxy luster, black color, no development of cleat, little bedding and a characteristic conchoidal fracture. capillary action tendency of fluids to rise in a small cylinder as aresult of surface tension. carbonaceous with abundant organic matter. carbonatecommonanion(C0,2-)insoilsolution andcomponent of carbonate minerals such as calcite and siderite and skeletonssuch as mollusk shells. carbonatite igneous rock composedprimarily of carbonate minerals. Carboniferousperiodof geologicaltime about 3 54-290 Ma. carboxyl common radical of organic acids (COOH+). caries texture soil microfabricin which grains are deeply embayed because of local dissolution or hydrolysis. carnivore animal that eats meat. carr wetland vegetation of trees with alkaline groundwater. cathode ray streamof electronsreleased fromacathode ray tube. cathodoluminescence method observations of mineral luminescence under cathode rays, especially useful for differentiating carbonate cements. cation exchange capacity measure of a soil’sexchangeable cations (mainlyH+.A13+, Ca2+,Mg’+, Na+.K+), usually by displacement with acetate at pH 7 and titration for its abundance. celadonite bright green, illite-likeclay mineral [(K.Ca,Na)-, ,(Fe’+,AI,Mg,Fe2+),(Si, 3A11~,702,J(OH)4]. cement fine-grained binding substance that holds together rocks and 6rm parts of soil, typicallysilica, hematite or calcium carbonate. chalcedony microcrystallie to acicular quartz (SiO,). chaparral 6re-prone shrubland, a term used mainly in California. charcoal charred wood. chelate a chemical compound capable of transporting elements or compounds by means of a particularly favorable site of attachment (from Greekchelaforclaw). chlorite clay-likemineral ~~Mg,Fe2+,Fe3+,Mn,Al~1,[~Si,Al~~O~l~l~OH~~~~. chondritekindof meteorite containing chondrules. chondrule smail (silt to granule size), round mineral grain, usually olivine or pyroxene,found in meteorites. chroma purity of color,or degree to which a color is not masked by murkiness or lightness in theMunsel1system of color classification. Chromudert Vertisolof humid seasonally dry climates, which has some horizons that are not dark gray to black. chronofunction mathematical relationship between a soil feature and the time over which it developed.
Glossary chronosequence set of soilsformed under similar climate,vegetation, topographic position and parent material but over varying lengths of time. clarain coal with silky luster and bright and dull laminae alternating on a scaleof amillimeter or less. clasticdike crack in soil or sediment filled with contrasting material. ‘SiIan’in terminology of Brewer. clay skin coating of clay along cracks or grains withm a soil. kgillan’ in terminology of Brewer. cleat system of fractures outlining small (few millimeters) cubic fragments of coal. cleavageplanes of weakness in a metamorphic rock. climofunctionmathematical relationship between a soil feature and ameasure of climate. cliosequence set of soilsformed under similar vegetation, topographic setting. parent material and time, but varying climate. clinobiasepic sepicplasmic fabric with a network of highly birefringentstreaks in two preferred directions and at a low angle. clinometer device for measuring angular deviation from horizontal. clinozoisitea metamorphic mineral [Ca,AI,O.AIOH(Si,O,)(SiO,)]. coal black carbon-rich rock formed by burial alteration of peat. coal ball calcareous or pyritic nodule of uncompacted coal found within a coal seam, much sought by paleobotanists because of well-preservedpermineralied fossil plants they contain. coal measures sequences of sedimentary depositsincluding layers of coal. coal seam natural bed or layer of coal. coaliication process forming coal from peat by expulsion of volatile materials and enrichment in carbon as a result of heat and pressure of deep burial. collinitevitrinite coal maceral derivedfrom structureless, decayed wood. colloid material too h e grained to be visible under the optical microscope:includes iron oxides and clay minerals. compressionform of fossilpreservation in which the organic remains of the fossil are crushed and coaliliedbetween bedding planes. concretion glaebulewith concentric internal fabric, usually becauseof periodic additionof material. Theseare hard, locallycemented lumps of material with onion-skin internal layering. cone ellipsoidalstructure of helically arranged reproductive organs of plants, foundin conifers,horsetails and clubmosses. conglomeratetest a test to determine whether themagnetic
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direction recorded by rocks is an original paleomagnetic signalor a later overprint, by recording the random directions of magnetization within clasts of a conglomerate. conodont fossil tooth of extinct marine animals related to primitivefish. coprolitefossil feces. corestone spheroidallyweathered remnant of parent material, least weathered toward the center, and usually within the C horizon of deep weathering profiles. corm thickened root-bearing base of a plant stem. cornstone pedogenic calcium carbonate nodules of paleosols. cortexintermediatezone of soft,parenchymatous cellsbetween the central woody conducting strand and outer epidermisof the roots and stems of plants. corundum igneous and metamorphic alumina mineral (a~ 2 0 , ) .
cover slip thin glass cover glued on top of a petrographic thin section. Cretaceousperiod of geologicaltime about 142-65 Ma. cristobalitehigh-temperature polymorph of quartz (SiO,), commonly as white octahedra and groundmass of acidic volcanicrocks. crossbedding sedimentary Iayering that is inclined to regional layering,commonly as a result of the formation of dunes with slip-facesat an angle to the ground surface. Cryand Andisol with ice deformation features. Cryert Vertisol with ice deformation features. Cryid Aridisolwith ice deformation features. Cryod Spodosolwith ice deformation features. cryoturbation soil-forming process of heaving and cracking as a result of ground ice. crystal chamber irregular nodular masses of crystals. crystal sheet planar aggregates of crystals. crystal tube tubular aggregates of crystals. crystallaria single crystals or groups of crystals in soils. crystallinitydegree of perfection of crystal structure, free of defectsor other less regular arrangement of chemical constituents. crystallite small crystal, beyond the size resolvable by optical microscopy. crystic a soil microfabricdominated by crystals. cuirasse indurated hardpan or crust exposed at the surface, usually lateritic. cumulic horizon soil horizon that shows bedding or other evidence that it is accumuIating in a sedimentary fashion at the surface of the soil. cutan surface within a soil modified by enrichment, bleaching, coating or other alteration: formed at the surface of aped, channel, grain or other feature of the soil. cuticle tough coating of hydroxy fatty acids and waxes that covers the leaves and other aerial parts of land plants.
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Glossary
cutinite liptinite (exinite)coal maceral formed from plant cuticles. Cv horizon C horizon with plithite. decalcificationsoil-buildingprocessof leaching out carbonate from surface horizons. decarboxylationchemicalreaction removing carboxyl radical from organic matter. decompositiondecay, or breaking down of organic matter to simplercompounds such as carbon dioxide by decomposer microbes. dehydration chemical reaction involvingloss of water or hydroxyl. depth function a plot with depth from the surface of the profile as vertical axis (y) and measured soil feature as horizontal axis (x), a common graphical presentation of chemical and petrographic data on soils and paleosols. desert pavement surface layer of stones, s o m e t i e s as closely interlocking as a cobblestonestreet. desert rose sand cemented into the shape of a radiating group of crystals of the cementing material, usually gypsum. desert scrub vegetation of widely scattered thorny shrubs and succulents such as cactus, with patches of bare ground. deviatoricaligned in unpredictable and random directions. Devonian period of geologicaltime about 41 7-3 54 Ma. diagenesis alteration of sediments after burial but before metamorphism: includessoil formation. diasporearuminurn-hydroxidemineral [AlO(OH)]. diffractogramplot of the X-ray reflections produced for identification of minerals from anX-ray diffractometer. diffusioncutan cutan formed by concentration at the surface of a material that becomes less prominent away from the surface: for example, the strong oxidation of the margins but not interior of peds. dioritegray plutonicrock intermediate in composition between granitic andmafic rocks, usually with common hornblende and quartz. displacivefabric a soil microfabric in which one mineral (usually calcite)fillscavitiesopened by the expansion or rotation of large clods of soil or the cracking out of clods or grains. dolomitecarbonate mineral [CaMg(CO,),]. dolostone rocksimilar to limestone,but unreactive to acid and composed of dolomite. domed-columnarped ped shaped like a vertically orientated prism, usually as thickas the wholeB horizon and with hemispherical upper surface; common in salt-affected soils. drab-haloed root traces root traces that are surrounded by soil of a gray and also bluish or greenish color,unlike yellow to red soil or paleosol away from the root trace. driRwood wood that has been transported by floating on water.
durain hard, compact and dull coal. duricrust hard cemented horizon of soil or deep weathering profile; includes laterite, bauxite, calcrete andsilcrete. Durid Aridisol with duripan at depth of less than 1m. duripan subsurface soil horizon cemented firmly by clay and silica. Dystrochrept Ochrept soilthat is very low in weatherable bases.
E horizon soil horizon underlying 0 or A horizon, characterized by less organic matter, less sesquioxides(Fe,O, and A1,0,), or less clay than the underlying horizon. Its light color is due largely to quartz and feldspar.Also known as an eluvial or albic horizon. ecosystemthe complex of a biological community and its environment functioning as aunit in nature. EDX energy dispersiveX-ray spectrometry. Ehelectrodepotential (usuallyinmillivolts):for soilsameasure of the degree of oxidation.Oxidized soilshave high positiveEh; reduced soilshave low negative Eh. eluvial horizon soil horizon characterized by less organic matter, less sesquioxides(Fe,O, andAl,O,), or less clay than the underlying horizon. Its light color is due largely to quartz and feldspar. endocarp interior woody part of a seed or fruit coat of plants, also known as a pit or stone, as in cherries and peaches. endolithic microrelief system of surface cavitiesformed by microbes living within and on rocks. endoskeleton internal skeleton,as in mammals. energy dispersiveX-ray spectrometry method of chemical analysis using X-rays emitted from sample in scanning electron microscope. Entisolvery weakly developedsoils, usually with abundant sedimentary,igneous or metamorphic relicts from their parent material. Eoceneepoch of geologicalt i e about 5 5-34Ma. epidermisouter covering of cells of plant or animal. epsomitesalt mineral (MgSO,. 7H,O). estuary that part of a river mouth that is influenced by marine tides. Eutrochrept Ochrept soil rich in bases. evaporite sedimentary rock that forms by the accumulation of salts during the evaporation of water, includes rock salt and gypsum. exinite old term for coal maceral liptinite. exoskeletonexternal skeleton,as in arthropods. extinction angle angular differencebetween the orientation of a crystal and the point at which it can no longer transmit polarizedlight as viewed in petrographic thin section under a microscope, useful for identifyingminerals such as feldspars.
Glossary faciesan informal rock unit, usually designated by features thought to be significantfor interpreting sedimentary paleoenvironment. factor-function approach study of environmental control in the expressionof soilfeatures: can beused also to infer paleoenvironmental conditions from paleosols. falsetrunkplantstructureoutwardlysimiIar toatreetrunk, but composedmainly of adventitiousroots and leaf bases, as in tree ferns and palms. fecal pellet small ovoid to spherical feces produced by small animals. feldspar group of aluminosilicateminerals including microcline and plagioclase. fen wetland vegetation of grasses and other herbs with alkaline groundwater. ferrallitiiation soil-formingprocess of enrichment to dominance of sesquioxidesof iron and aluminum. ferran cutan consisting of oxides or hydroxidesof iron. ferric iron iron in the Fe3+valencestate,usually within red or yellow minerals or compounds. ferrous iron iron in the Fe2+valence state, usually within gray to green minerals or compounds. ferruginous rhuoconcretion rhnoconcretion cemented by goethite, hematite or other iron hydroxidesor oxides. fibric peat peat with abundant recognizableplant material, not completelydecayed. FibristHistosolconsisting largely of plant remains so little decomposed that their botanical origin can be determined. fibrous roots numerous fine (usually<2 mmdiameter) roots radiating from the base of a piant, as in palms and grasses. fining upwardssequence sedimentary layer that varies toward smaller grain size from the bottom to the top. fie-prone shrubland closely spaced woody shrubs, less than 2 m tall, adapted to frequent burning. fishtailtwin crystallographic twin in which the paired crystals divergeat a low angle like the tail of a Esh. flocculationaggregation into a coherent mass from fine suspendedparticles, as can happen to clay particles in turbid water with changes in salinity. floodplainfrequentlyflooded low-lyingregion flanking large rivers. fluorescenceemission of electromagnetic radiation usually as a response to absorption of another form of radiation. Fluvent Entisol soilformed on silt and clay with conspicuous relict bedding. fluvial relating to rivers. flysch sedimentary rock sequences of sandy graded beds interbedded with gray shale, formed indeep sea. Folist kind of Histosol,freely drained, consisting primarily of organic horizons derived from leaf litter,twigs and branches
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resting onrock, gravelor boulders, the interstices of which are 6Ued with organic material. footslopelower convexpart of a hill slope, between steeper backslope and more leveltoeslope. foraminifermarine microorganisms that secrete or manufacture a small shell. forb herbaceous plant other than grasses, found typicaIlyin rangeland. forest vegetation of closely spaced trees more than 8 m tall. fragipan dense, subsurface hardpan of clay, often mottled, prismatic and siliceous. framboidminute (10pm to 1mm diameter) round spherules or groups of spherules of pyrite, usually produced by anaerobic bacteria. freeze-thaw banding structure of cracks and silty seams, usually in abrickworkpattern, produced by frost insoils. friableeasy to break into constituent grains: oppositeof cemented or indurated. fusain soft, powdery,friable, soot-likecoal, which is dirty to the touch. fusinite kind of inertinite coal maceral. formed from fossilized charcoal. ganister silicified quartz sandstone, containing root traces and underlying a coal seam. These are in most cases paleosol Eor eluvial horizons. Geliilsoilswithpermafrostorothergroundicewithin 1mof the surface. geosol soil stratigraphic unit in the North American stratigraphic code. geothermal gradient variation in temperature of the Earth with depth into the crust. gibbsite aluminum hydroxide mineral [Al(OH),J. gilgaimicrorelief soil surface of ridges and swales or potholes produced by the shrinking and swelling of Vertisols. glacialpavement rock surface scoured and scratched by a glacier. glaebule segregations of materials distinct from other parts of the soil, including nodules, concretions and septaria. gleization process of gley formation. gley soil that is bluegray or green-gray colored, strongly mottled or with abundant iron-manganese nodules, usually as a result of waterlogging, but sometimes produced also by burial. glossic features locally penetrating tongues or tubes of lightcolored sandy material deep within a soil profile. glycolation experimental treatment of clays by exposure to open container of ethylene glycol at 80°C for 4h, used to study expansion of smectite clays. gneiss foliatedmetamorphic rock with alternating bands of light and dark minerals.
340
Glossary
goethiteyellow to brown iron hydroxidemineral [a-FeO(OH)]. gradation processes that produce regolith on a planetary body: includespedogenic.sedimentary,volcanic and other sur6cia1processes acting on bedrock. graded bed sedimentary bed that varies in grain size from the bottom to the top of the bed. Normally graded beds are finer grained to the top, and inversely graded beds are coarser grained toward the top. graniticrocksrocksliiegranite,whicharerichinquartz and feldspar,usually coarsegrained and formed from the coolingof molten rockdeep within theEarth’scrust. granotubule tubular feature of soil filled with clastic grains and little clay. granular microfabricsoil microfabric in which skeletongrains are touching with little or no fine-grainedmatrix in the interstices. granular peda form of ped or clod that is small and rounded. graphite metamorphosed coal, consisting largelyof carbon. grassland vegetation of mainly grasses. ‘GreenClay’kind of paleosol found in Precambrian rocks, with unusual combination of indications of deep weathering (corestones,cracks) and gley colors (green-gray with chemically reduced iron minerals). greenschist facies of regional metamorphism characterized by extensive recrystallizationand metamorphic chlorite, usually under temperatures of about 3 50-500°C. groundwater calcrete calcrete formed by precipitation of carbonate cement from groundwater. groundwater gley gley features formed by the ponding of groundwater from an elevated water table. groveland vegetation of clumps of trees separated by open grassland. guano excrement of birds and bats. Gypsid AridisoI with gypsicor petrogypsic horizon at depth of less than 1m. gypsumevaporite mineral (CaS0,.2H20). hackly having the appearance of something chopped or cut up roughly. halite salt mineral of desert soils (NaCl). halloysite hydrated form of kaolinite. Haplaquept typicalAquept, lacking distinguishing features of other kinds of Aquepts. Haplohumult typical Humult ,lacking distinguishing features of other kinds of Humults. Haplorthod typical Orthod, lacking distinguishing features of otherkindsof Orthod. Haploxerandtypical Xerand. lacking an especiallydark surface horizon. HapludalftypicalUdalf, lacking distinguishing features of other kinds of Udalfs.
Hapludolltypical Udoll, lacking distinguishing features of other kinds of Udolls. Hapludult typical Udult, lacking distinguishing features of other kinds of Udults. Haplustalftypical Ustalf, lacking distinguishing features of other kinds of Ustalfs. Haplustoll typical Ustoll, lacking distinguishing features of other kindsof Ustolls. Harden index quantitative measure of the degree of soil development, calculated by addition of scores for a variety of soil features thought to vary with time of formation. hardground indurated cemented sea floor. hematite red iron oxide mineral (Fe,O,). hemic peat peat in which some but not all of the plant material is so decayed as to be unrecognizable. Hemist Histosol in which organic matter is decomposed to the extent that as much as two-thirds of the plant material is unidentifiable. herbivore animal that eats plants. Histel Gelisolswith a peaty surface horizon. histicepipedon peat: a soil surface horizon with at least 18% organic matter if the mineral fraction contains more than 60%clay or 12%organic matter if the mineral fraction has no clay, for adepth of 20 cm. Histosolsoil with peaty surface, which must be at least 40 cm thick if composed mainly of woody material, 80 cm if mainly moss peat. Holoceneepoch of geologicaltime from 10Ka to present. hoodoo erosional pinnacle of badlands. in which a pillar of sediment has been protected from erosion by a capstone. hornblende aluminosilicate mineral of igneous and metamorphicrocks [(Na.K),,-,Ca, (Mg,Fe2+.Fe3+,Al), Si,~,.,AI,~,,O,,(OH),l. hue color. such as red, yellow or green, independent of murkiness or lightness of the color,in the MunselIsystem of color classification. Humod Spodosolwith a surface horizon rich in organic matter. Humult Ultisol with a surface horizon rich in organic matter. hydrocarbon compound mainly of hydrogen and carbon, including methane and paraffin. hydrolysiscommon weathering reaction in soil solutions. converting aluminosilicate minerals to clay and cations in solution. hydromorphismformation of gley features by waterlogging (broadlysynonymous with gleization). hydronium hydrogen ion (H+). hydrothermal related to groundwater of elevatedtemperature, commonly associatedwithvolcanic activity. hydroxide compound including hydroxyl. hydroxylchernicalanion (OH-).
Glossary ice wedge cast vertical wedge-like disruption of a soil Ned with sediment after melting of ice that created the crack. ichnogenus formal taxonomic category for a specific kind of trace fossil,similar to a genus of biologicalclassification. ichnospeciesas ichnogenus, but similar to biologicalspecies. ICP inductively coupled plasma emission spectrometry. illitepotassium-rich clay mineral
LK I . 5- 1.0A14(si6. 5 - 7 . 1.5~ ~1.0°2doH)d ~ illitization common process during deep burial that converts smectite clays to illite. illuvial horizon soil horizon enriched by illuviation of a component, such as clay. illuviation soil-buildingprocess of enrichment of clay washed in from higher horizons. illuviation cutan cutan formed by washing down of material from higher within a soil. ilmeniteopaque heavy mineral (FeTiO,). impressionform of fossil preservation in which only an outline of the fossil remains with none of the original organic material. INAA instrumental neutron activation analysis. I n c e p t i l soil with some weathering and incipient development of a variety of differentkinds of horizons, but none well enough developed that the soil could be identifledwith another order. incisor front tooth of mammals. greatly elongatedin rodents. inductivelycoupled plasma emissionspectrometry method of chemical analysis using wavelengthsof light emitted from atomized sample introduced into an argon plasma. indurated firm. hard, cemented. inertinite component of coal that is light gray to white inreflected light and nearly opaquein transmitted light, but not nearly so bright or opaque as pyrite. insepicsepicplasmicfabric with small isolated patches of highly birefringent plasma. intercalary crystal single crystal embedded in soil matrix. interstratied clays clays with crystal layers of differingchemical composition. intertextic soil microfabric in which skeleton grains are more prominent than 6ne-grained matrix, which forms intergranular braces and fills local pockets. intrabasaltic within a sequence of basalts. inundulic soil plasmic microfabric similar to undulic, but cloudy, with large irregular isotropicpatches that appear dark when viewed under crossed nicols. isoticsoil plasmic microfabric that is dark when viewed under crossednicolsbecause either isotropic,likeopal, or opaque, like hematite. isotope alternative forms of a chemical element that differ slightly in mass and sometimesalso in other properties.
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Isotopes of carbon include the radiogenic isotope 14C used forcarbondatiing, andthestableisotopes 13Cand12C. isotropicmineral a mineral whose crystal structure shows no preferred orientations: appears black in thin section viewed under crossed nicols. isotubule tubular feature of soil filled with mixed clay and clastic grains without any preferred orientation. Jacobstaffgraduatedsurveypoleusedfortakingangulardifferencesin elevation and horizontat differencesin distance with clinometer. jarosite powdery,yellow mineral smelling of rotten eggs [KFe,(OH,)(SO,),], forms by oxidativeweathering of pyrite. Jurassicperiod of geologicaltime about 206-142 Ma.
K feldspar potassium feldspar (KAlSi,O,) including microcline, orthoclase and sanidine. K horizon subsurface soil horizon so impregnated with carbonate that its morphology is determined by the carbonate. Authigenic carbonate coats or engulfs all primary grains in a continuous medium and makes up 50%or more by volume of the horizon. The uppermost part of the horizon commonlyis laminated. kandic horizon subsurface soil horizon similar to argillichorizonin clay enrichment, but clays are kaolinitic and there are very few weatherable minerals remaining. kaoliite base-poor clay mineral [AI,(Si,O,,)(OH),]. karst landform formed from the weathering and dissolution of limestone, often rugged with caves and pinnacles. karst bauxite kind of bauxite formed within depressionsof karst landscapes. kerogen particulate organic matter lacking regular chemical structure and insolublein organic solvents and mineral acids, present in sedimentary rocks. kerogen cracking conversion of kerogen to oil or to natural gas under elevated temperatures and pressures of deep burial. kieserite salt mineral (MgSO,.H,O). Kjeldahl titration wet chemical method for determining nitrogencontent of soilsandrocks. knee root root that extends upwards out of the ground and then bends over and rejoins the main root system, found in bald cypressand other swamp trees. krotovina tubular feature in a soil Ned with material from a higher soilhorizon, usually animal burrows. Kiibler index measure of clay crystallinity using the width of X-ray diffractometerpeaks. kunkar pedogenic calcium carbonate nodules or layers. labileminerals minerals such as olivine and pyroxene that are relatively easily weathered.
342
Glossary
lacustrine relating to lakes. larval cell chamber used for rearing young in the complex nests of social insects such as termites, bees and ants. laterite highly weathered material rich in iron, and poor in humus, silica and bases. lattiipic sepicplasmic fabric with a network of highly birefringent streaks in two preferred directions that are at aright angle. lentil ped soil clodsthat are shaped l i e an elongate parallelogram, usually with slickensidedfaces and characteristic of Vertisols. lessivagesoil-buildingprocess of washing down of clay into subsurface cracks. lichen plant-lie organism formed by the symbioticassociation of fungi and algae. ligament tough tissue that connects bones or supports internalorgans of animals. liecalciumoxide (CaO). liestone rock formed mainly of calcite. h e a r gilgaigilgaimicrorelief of elongate ridges and swales, usually running downslope. liptiiite component of coal that includes avariety of oily, waxy and resinous plant materials, dark gray to black in reflected light, and yellowtoreddishbrownin transmittedlight. lithic sandstone sandstone whose clasts are mainly rock fragments. lithofunction mathematical relationship between soil features and parent material of the soil. lithorelictrock fragment in a soil remaining fromits parent material. lithosequence set of soilsformed under similar climate. vegetation, topographic setting and time, but varying parent material. litter accumulation of leaves,wood and other decaying organic matter on the surface of soil. liiviation soil-formingprocess of hydrolyticleaching of base cations (Ca2+,Mg2+,Na+,KC). loessdeposits of wind-blown glacial dust. lumen (plural lumina) hollow cavities found within roots of swamp plants. lycopsid group of spore-bearing plants with narrow, linear, helically arranged leaves,including clubmosses. maceral microscopicgrains of coal, derivedfrom alteration of plant Gagments during coalication. and best examined in polished sectionsunder areflectedlight microscope. macronutrients elements needed in large amounts for the nourishment of plants, including hydrogen (H),carbon (C), nitrogen (N), oxygen (0).magnesium (Mg),phosphorus (P), sulfur (S),potassium (K) andcalcium (Ca). maghemite spinelmineral (yFe,O,).
magnesia magnesium oxide (MgO). magnetite an iron-rich mineral (Fe2+Fe3+,0,). mangal vegetation consisting of mangroves. mangancutan consisting of oxides or hydroxidesof manganese. mangrove tree capable of living within the intertidal zone of the ocean. maquis fire-proneshrubland of the Mediterranean region. marsh wetlandvegetation of grasses and other herbs with acidicto neutral groundwater. masepic sepicplasmic fabric with highly bireftiingentstreaks forming an extensivecriss-crossingnetwork. matorral fire-prone shrubland and heath of South America. megaspore large (usually more than 60pm diameter) spores of some kinds of ferns, lycopsids and similar plants. melanic epipedon surface horizon darkened with he-grained organic matter. melanization soil-formingprocess of darkening mineral soils by accumulation of he-grained organic matter. mesa hill with a broad flat top. metagranotubule tubular feature in a soilfilled with sandy material from a higher soilhorizon. metamorphism alteration of rocks during deep burial and heating, generally to more than 200°C or greater than 7 km, whichever comes first. metasediment metamorphosed sedimentary rock. metatubule tubular feature of soil Wed with material different from soil matrix and derived from some other soil horizon. micrinite kind of inertinite coal maceral, formed from granular degraded material. micrite very h e grained sediment of calcite and clay minerals. micritiiation soil-formingprocess whereby coarsely crystalline calcite or other materials are converted tomicrite. microarthropod microscopic arthropod, including springtails and mites. microbe microscopic organism. microbial earth vegetation of microbes living on and within a friablesoil. microbial rockland vegetation of microbesliving on the surface and within the surface weathering rinds of rocks. microcline potassium-rich feldspar mineral [K(AlSi,O,)], commonlywith tartan twinning. microfabricarrangement of constituents visibleby microscopy. middle lamella seam between adjacent cellulose cellwalls in wood. ministromatolite small (oftenmicroscopic)domed structure withinternal lamination, formed by the growth of microbial colonies. Mioceneepoch of geologicalt i e about 2 4 5 Ma.
Glossary Missiiippianperiodof geologicaltiie about 3 54-323 Ma. moder humus organic matter consisting of a mix of recognizable plant material and other organic material completelydecayed:intermediate between mor and mull humus. moderately developedsoil or paleosol with surface rooted zone and obvious subsurface clayey, sesquioxidic,humic or calcareous or surface organic horizons, qualifyingas argillic. spodic or calcichorizons or Histosoland developedat least to the extent of nodules for calcic horizons. moisture equivalent percentage moisture in a soil at ‘field capacity’,which is the point at which water is no longer moving through or W i g soil pores, but bound in immobile films to grains and roots. molemassingramsof Avogadro’snumber(6.022~1023)of atoms or molecules of an element or compound, calculated by dividing weight per cent of analyzedelement or compound by its atomic weight. molecular weathering ratio ratio of chemical constituents in moles, calculated to understand changing chemical proportions as aresult of weathering. mollicepipedon soil surface horizon of grassland soils,with b e structure (usually granular peds),dark color (chroma of three or less, value darker than five when dry),contains at least l%organicmatter(O.S8%organiccarbon),andhasa base saturationof over 50%. Mollilgrasslandsoilwithamollicepipedonatleast 18cm thick. monsoonal climate climate of very marked seasonal rainfall, as in the Indian subcontinent. mor humus organic material consisting of little-decayedplant material, such as the dried pine needles common under conifer forest. moraine accumulation of earth and stones carried and depositedby glacier. mosepic sepicplasmic fabric with partly adjoining highlybirefringent streaks. mottle glaebulesof very irregular shape and diffuseboundaries, usually expressed as differentlycolored areas of soil. mucigel gelatinous zone within the rhizosphere. rich in bacteria and fungi. mudflow rapid downslopemovement of a slurry of mud and boulders after rain storms or volcaniceruptions. mukkara structure subsurface deformation of soil horizons, usually with the surface horizon festoonedbetween ridges of exposed cracked subsurface horizon, produced by shrinking and swellingof Vertiiols. mull humus organic matter decayed so that plant structure is no longer visible. as in the surface horizon of grassland soils or Mollisols.
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murram loose aggregate of pea-sized lateritic gravel used for road building. muscovite potassium-rich mica mineral [K,A14(Si6A1202,)(0H,F)41.
mylonitecrushed and shearedrock found within fault zones. natric horizon soil subsurface horizon more than 15% saturated with exchangeable sodium, commonly also with salts and prismatic or domed-columnar peds. Natrustoll Ustoll soil with abundant sodium or other indications of salt accumulation such as domed-columnar peds. natural gas mixture of methane and other gaseous hydrocarbonsfoundin the Earth’s crust. needle fiber calcite calcite crystals in the form of microscopic needles. commonly precipitated by soil fungi. neocutan altered area within a soil at a surface as well as some distance in from the surface.These are unusually thick cutans. neoferran neocutan of iron oxidesor hydroxides. neomorphim change in form of crystals, including recrystalhiation. nephelie feldspathoidmineral of igneous rocks nepheliite igneous rock with common nephelie. and iron- and magnesium-rich minerals, such as olivine and pyroxene, and little if any quartz. neutron activation analysis method of chemical analysis using radiation induced after neutron irradiation in a nuclear reactor. nicols polarizing light filterson a petrographic microscope. NitosolUltisol or Oxisolwith abundant slickensidedclay skins in the FA0 soil classfication. nodule glaebulewith anundifferentiated, massive internal fabric.These are local, usually hard, cemented lumps of soil material. nomogram graph that consists of lines marked in such a way that needed and unknown values can be obtained from measured or knownvalues. normative mineral composition estimate of the proportions of minerals present in a specimen calculated from thechemical composition of the specimen and ideal compositions of the minerals. nuram gilgai gilgai microrelief of large more or less circular holes. 0 horizon surface accumulation of organic material overlying mineral soil. Ochrept Inceptisol soilwith ochric and cambic horizons, and sometimesalso poorly developedcalcic horizons, fragipans or duripans.
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Glossary
ochric epipedon soil surface horizon too light in color and low in organic matter to be mollic or umbric. oil liquid hydrocarbon formed inthe Earth from burial alteration of organic matter. oil window stage in burial history conducive to creation of oil, typicallybetween burial temperatures of 50°C and 15OoC, or depths of some 1-5 km,depending on local geological age, tectonic history and geothermal gradient. Oligoceneepochof geologicalt i e about 34-24Ma. oligotrophicforest forest living inlow-nutrient soilsuch as a Spodosol. olivinemineral of igneousrocks [(Mg,Fe),(Si04)]. ombrotrophicpeatland elevated above water table so that nutrients are deliveredprimarilyby rain and dust. omnisepicsepicplasmic fabric dominated by highly birefringent, orientated plasma, with a ‘woven’appearance. omnivoreanimal that eats meat and plants. ooidsmall(1-2 mmdiameter) roundgrainwithinternalconcentric structure. opal silicamineral (SiO,.nH,O). opalite silica-cementedshale or claystone. opaque not pervious to light, and so black when viewed by transmitted light in a petrographic thin section. Ordovicianperiod of geologicaltime about 49 5 4 4 3 Ma. organan cutan consisting of organic matter. Orthel Gelisol lacking a peaty surface or obvious indications of cryoturbation. Orthent Entisol soil formed on erosional remnants such that hard bedrock, or relict soil material, is within 2 5 cm of the surface. Orthid Aridisolsoil lacking an argillic horizon (not recognized since 1997 revision of soiltaxonomy). orthoclase alkali felspar (KAISi,O,) common in granite, similar to quartz under the petrographic microscopebut with lower relief. Orthod Spodosolsoil with a subsurfacehorizon including both sesquioxidesand organic matter. orthotubule tubular feature of soilfilled with material of very similarfabric and compositionto soil matrix. ortstein a spodichorizon that is unusually heavily cemented, withmore than 50%by volume opaque cement. Ostwald ripening process of increasing of crystal size that minimizes freeenergy of crystal faces,thought to account for burial illitiiation. oxic horizon highly weathered subsurface horizoncharacterized by hydrated oxides of iron and aluminum, 1: 1lattice clays, and low cation exchange capacity.Few primary silicate minerals remain, with the exception of quartz, which is resistant to weathering. oxidation chemical reaction in whichelectrons are lost to valence of elements with multiple valence states, for exam-
ple Fe2+to Pel+,which is achieved commonly by means of oxygenas a n electron acceptor. oxidizinggroundwater water within soils and rocks that is rich in dissolved oxygen. Oxidizing groundwaters are rare, and largely found within activelyrecharged, sandy aquifers, because of oxygen scavengingby microbes and by minerals. Oxisol deeply weathered soilswith kaoliniticclays. quartz and few weatherable minerals. Paleocene epoch of geologicaltime about 6 5-5 5 Ma. paleochannel formerriver channel, usually marked by an elongate depositbroadly lenticular in cross-section,of cross-beddedsandstone or conglomerate. paleoecologystudy of the ecology of fossil organisms. paleomagnetiim former orientation of Earth’s magnetic field as recorded in magnetic minerals in rocks. paleopedologystudy of paleosols. paleosolsoilof alandscape of thepast: apast surfacematerial of a planet or similar body altered in place by biological, chemical or physical processes, or a combination of these. This spelling (rather than palaeosol or paleosoil)has been formally adopted by the INQUA-ISSSPaleopedology Commission. PaleudalfUdalf that is very strongly developed,usually over a long period of time (hundreds of thousands of years). Paleudult Udult that is very strongly developed,usually over a long period of time (hundreds of thousands of years). PaleustalfUstalf that isvery strongly developed,commonly with a laterally continuous subsurface calcareous horizon. pallidzone horizonof whiteclay and bleachedormottledrock beneath laterite in a deep weathering profile. paludization soil-formingprocess of surface-water ponding and accumulation of peat. palygorskitemineral of desert soils[(OH2)4MgSSiR0,,,.4H,0]. pampassynonym of grassland, used mainly in South America. papule glaebule of clay, ausefulnongenetic term if one is not sure whether they are clay galls of the original parent material, void fills or segregations of clay. paratubule tubular feature of soil filled with material different from soil matrix and u n l i anything else within the profile. parkland vegetation of woodland with numerous large grassy clearings. pednatural aggregate of soil; that is, stable lumps or clods of soil between roots, burrows, cracks or other planes of weakness. pedoderm soil stratigraphic unit in Australian stratigraphic code, equivalent to geosol of North American and internationalusage.
Glossary pedofacieslaterally contiguous bodies of sedimentary rock that differ in their contained laterally contiguous paleosols as a result of their distance (during formation) from areasof relatively high sediment accumulation. pedogenicformed in association with soil. pedolithsedimentary depositcomposed of clasts that are clearlyderived from soils,such as talus slopes of lateritic clasts below alaterite scarp. pedorelict soil structure that formed in adifferent soilthan the one in which it is found for example,acalcareous nodule within the gravellyparent material of a noncalcareous soil. pedotubule tubular features of soils,including roots and burrows. pedotype asoil orpaleosolmapping unit, basedon areference profile for dehition. peg root short, bluntly ending roots that protrude upward from theground, as inmangroves. pegmatitecoarsely crystalline rockmainly of quartz and feldspar,usually forming veins within granitic rocks. Pennsylvanian period of geologicaltime about 323-2 90 Ma. pentlandite sulfidemineral [(FeNi),S8]. perched water table level of water ponded in soil by an impermeable subsurface layer. peridotite plutonicrock formedmainly of olivine. periostracum outer, thin, brown organic layer covering the shellsof some mollusks,such assnails and freshwater clams. Permian period of geologicalt i e about 290-250Ma. permineraliiedin6Itratedwith cementing minerals, as in the silica that fills woody cell contents to create permineralized wood. Such wood is often incorrectly referred to as petrified, but the cellulosecell walls remain, and have not been turned to stone. Perox Oxisols of very wet climates. petrocalcichorizon subsurface soilhorizon cemented firmly with calcium carbonate. petroferrichorizon subsurface soilhorizon cemented firmly with iron oxides and hydroxides. petrography description of rocks, usually including study in thinsection. petrogypsic horizon subsurface soilhorizon cemented h l y with gypsum. pH negativelogarithm of the activity of the hydroniumion (H+):for soilsameasure of acidity.Acidic soils have a low pH (<7) andalkahesoilshavea highpH(>7),witha total observedrangeinnatureof 3-1 1,froma theoretical 1-14. phenocryst large crystal in a porphyriticrock. phonolitemafic volcanic rock with alkali feldspar and a feldspathoid such as nepheline, lacking quartz. phosphate phosphorus anion (PO,3-). phytokarst form of limestone weathering where the surface is
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blackened and made spongy by abundant microscopic algae andlichenslivingwithin andontherock. phytolith mineral particle made bya plant, such as opal bodies of grasses. pingo large conical structure, often appearing likea small volcaniccone, and produced by the growth of ground ice. pisolite spherical concretion usually about 2-1 5 mmin diameter. plagioclase group of feldspar minerals of igneous rocks, ranging in composition from albite [Na(AlSi,O,)] to anorthite [Ca(AlSi,O,)]. planetimal small planet-like world, hypothetical early precursor of planets. plasma he-grained material of soilmicrofabric,making up peds, including clay andiron oxides. plasmicfabric appearance of plasma under the petrographic microscope under crossednicols. platypedformof ped thatis thinbutwide,oftenformed by weathering of sedimentary layers. playalakedesert basin, rarelyinundated, usuallydryand coveredin salt and clay. Pleistoceneepoch of geologicalt i e about 1.8Ma to 1OKa. plinthite iron-rich, humus-poor clayey part of a soil, usually mottledred and yellow, and with the distinctiveproperty of hardening irreversiblyto an iron hardpan upon drying. This term is used for those kinds of laterites that are found within soilprofiles. Plioceneepoch of geologicaltime about 5-1.6 Ma. plutonicrockrock formed at great depth, usually coarsely crystallme. pneumatophore erect, upward-growing root with aerophores or aerenchyma in mangal and swamp plants. podzoliation soil-formingprocess of acidic leaching leading to a subsurface accumulation of sesquioxides,organic matter or combinations of these. point-counting systematic search and record of grain type or sue made to determine mineral compositionor grain size of paleosols. polsterlandvegetation of well-drainedsoils consisting only of rootlessplants such as bryophytes and other rootless organisms such as lichens. porcellanous with the appearance of china or porcelain. porosity per cent void space between grains and peds of soil; complement to solidity. porphyritic crystalline texture of large crystals isolated within ahe-grained matrix. porphyroskelic soil microfabricin which grains are dispersed within he-grained matrix, like phenocrysts in a porphyritic rock. potash potassiumoxide (K,O).
346
Glossary
prairie synonym of grassland, used mainly for vegetation in North America. Precambrian period of geologicaltime before about 544Ma. prehnite-purnpellyite facies assemblageof burial metamorphic rocks formed at temperatures of about 200-300°C and burialdepthsof about 6-12 km. pressure solution dissolutionof grains of rock by the pressures of deep burial focused at grain contacts. primary porosity proportion of void space between grains and peds in the original soil or sediment. primary productivitybiologicalproductivityfrom plants and other creatures with similar nonconsuming metabolism. productivity a measure of biologicalaccumulation of organic matter, measured as grams of carbon or of dry organic matter per square meter per year. progymnosperm extinct group of plants with gymnospermous anatomy but pteridophyticreproduction. Proterozoicperiod of geologicaltime about 544-2 500 Ma. ‘Protorendziia’ weakly developedsoil with mollic surface horizon on limestonebedrock. Psamment Entisolformed on sand, especiallyeolian dunes, with relict bedding. pseudoanticlieuparched bedding planes that are confined to a particular layer,and so thought to be due to local clay heave or crystallizationrather than regional foldingthat produces anticlines. pseudornorphmineral grain that has adopted the form of another mineral, usually as aresult of replacement of that mineral: for example, chalcedony pseudomorphs of gypsum. pseudomyceliumh e irregular filamentsof calcium carbonate in soil. ptygmatic folded backon itself inacomplex way. pumice light-colored vesicular volcanic rock with very low bulkdensity(itfloatsin water) and achemicalcomposition l i e that of rhyolite. pyrite common mineral of mangal and salt marsh soils (FeS, ). pyroxene mineral of igneous rocks [(Mg,Fe)Si,O,]. quartz common mineral of soils (SO,). quasicutan altered area within a soil that is thick and shows a relationship to a surface,but is not right at thesurface. Quasicutans form akind of halo following the outline of peds, grains and other features. quasimangan quasicutan of iron-manganese. Quaternary period of geologicaltime from 1.8Ma to present.
R horizon consolidatedor weathered bedrock underlying the soil. radiogenicprone to decay with the release of radioactivity. rainforest forest living in avery humidclimate. rangeland region of open vegetation, including wooded grassland, grassland and desert scrub. rareearthelementselementswithatomicnumbers5 7-71. also known as lanthanides, of which lanthanum (La), cerium (Ce),neodymium (Nd).samarium (Sm),europium (Eu),gadolinium (Gd),ytterbium (Yb) and lutetium (Lu)are commonly analyzedin rocks and soils. recrystallizationprocess of formingnew crystals without a change in chemical composition. redbedssediments or sedimentary rocks that arelargelyred in color. reddeningchangein color from brown (Munsell 10YR)to red (Munsell5R)that occurs during soil formation and during burial of soils. reductant chemicalcompound capable of inducing chemical reduction. reduction achemical reaction in which electrons are donated to change the valence of elements with multiple valence states, for example converting ferric iron (Pe3+)to ferrous iron (Fe2’). REE rare earthelements. regolith unconsolidated material at the surface of a planetary body, including sediment, soil, and weathered and hydrothermally altered materials. relict structures features persisting in soil from its parent material, including bedding, crystalline structure and schistosity. reliefdegreeto which the margins of a mineral grain stand out fromits surroundings as viewed in thin section under apetrographic microscope. RendollMollisolformed on limestone bedrock. replacivefabric fabric of soils in which one mineral is convertedincompletely and over an irregular front into another mineral. residuum material remaining after a long period of weathering. resiniteliptinite (exinite)coal maceral, formed from plant resins, including amber and waxes. resistatemineralsminerals such as quartz and zircon that are resistant to weathering and tend to persist as other minerals aredestroyed in soils. respiration metabolicprocess of organisms, whereby organic foodisconvertedintoenergyandcarbondioxide. rheotrophic peatland that receivesnutrients largely through groundwater. rhizoconcretion concretion that forms by episodiccementation of soil around a root.
Glossary rhizoid elongatedepidermal cell that functions as a root, as in mosses, liverwortsand primitive landplants. rhizome a root-likestructure of plants that lies along or within the ground, but that is really a stem, as revealed anatomically and in its pattern of branching and budding. rhuomorphcorm-like and root-like stems at base of plant in quillworts (Isoetes)and extinct tree lycopsids (Stigmaria). rhmsphere area of influence of aroot in the soil. rhyolite silica-richvolcanicrock. ripplemark sedimentary structure of small-scaleundulations of a bedding plane: miniature dune-like forms produced by wind or water currents. rock varnish thin crust of red to black iron and manganese oxides and clay formed on the surface of rocks in deserts, lakes and streams. largely as a result of microbial activity. root branching subterranean structure of plants, oftenwith some woody internal thickening. root hair elongate cell erect on the surface of roots, most common a short distance behind the growing tip of roots. root trace tubular cavity or irregular disturbance left in soils and paleosolsby roots, recognized by irregular tubular shape, tapering downwards, branching downward or outward from a center, and (for deeply buried paleosols) concertina-like shape as a result of compaction of surrounding sediment around the main lateral rootlets. rootlet side branch fromroots. rutile red to brown, high-relief,heavy mineral (TiO,). sabkhacoastalmudilat withsalt crusts. saccharoidal l i e sugar crystals. salichorizons subsurface soil horizon with accumulations of salt. Salid Aridisolwith a salic horizon at depth of less than 1m. salinization soil-formingprocess of salt accumulation by evaporation. salinized affectedby salt accumulation. salt marsh wetland vegetation of grasses and other herbs with saline groundwater, usually within the intertidal zone of bays. sand crystal sand cemented into the shape of a crystal of cementing material, usually gypsum. sand wedge vertically orientated wedge-like disruption of soil 6lled with vertically banded sediment, from permafrost ice. sandstone sedimentary rock formed of sand-sizegrains. sanidme alkali feldspar (KAlSi,O,) common as clear tabular crystals involcanic rocks. sapricpeatpeatin whichorganicmatter hasdecayed tosuch an extent that little if any of the original plant components are recognizable.
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Saprist Histosol soil consisting mainly of decomposed and unrecognizableplant matter. saproliteless alteredlower portion of weathering profiles, showing some weathering, but relict bedding, schistosityor crystalline structure remaining from parent material. savanna commonly used as a synonym of wooded grassland, but also widely taken to include grassy woodland. scanning electron microscopemicroscopethat uses electron beams to create images. schist metamorphic rock of clay grain size showing pronounced schistosity. schistositydegree of development of planes of hsility and foliation characteristic of he-grained metamorphic rocks. sclerotiarounded. woody bodies, with interlaced, elongate, hollows produced by theresting stages of fungi. sclerotmitekind of inertinite coal maceral, formed from sclerotia. scoriadarkvesiculatedvolcanic rock of basaltic to andesitic composition. secondary porosity a system of tubes and vesicles developedin rock during deep burial associated with maturation of buried organic matter. secondary productivity biological productivity from animals and other consumers. SEM scanning electron microscopy. semifusinitekind of inertinite coal maceral, formed from partly degraded charcoal. sepicplasmic fabric appearance of the he-grained part of a soil in petrographic thin sections viewed under crossed nicols as wisps or streaks of highly orientated and highly birefringent clay in a less organizeddark matrii: a characteristic microfabricof soils. sepiolitemineralof desert soils [(OH,),Mg,Si, L03,,.8H20]. septarium (plural septaria) glaebule with a complex system of internal cracks, as aresult of shrinkage or hydraulic fracturing. sericitehe-grained white mica, which may be muscovite, illite or mixtures of these, but difEcultto determine optically:common in metamorphosed paleosols. serpentine mineral of metamorphic rocks [Mg,(Si,O,)(OH),]. sesquan cutan consisting of sesquioxidesof iron and aluminum. sesquioxidesalumina (A1,0,) and ferric iron (Fe,O,). shipworm bivalved mollusks capable of boring into hard substrates such as rock and wood. shrublandvegetation of low-growing woody shrubs, such as sagebrush and saltbush. siderite carbonate mineraI of waterlogged and organic-rich soils (FeCO,). silan cutanconsisting of silica.
348
Glossary
silasepicsoil microfabricdominated by silt and sand grains, and lacking highly birefriigent streaks when viewed under crossednicols. silcretea silica-cementedrock, commonly associatedwith weathering profiles. silicasilicondioxide(Si0,). siliceouscomposedmainly of silica. silicificationsoil-formingprocess of cementation with silica, suchas opal, chalcedonyandquartz. silicifiedcemented by silica. sillianite metamorphic mineral (Al,SiO,). Silurian period of geological time about 443417 Ma. sinkerdeeplypenetratinglargeroot,which takes advantage of deep groundwater. sinter silica-rich, often vuggy, depositsformed around volcanic hot springs. skeletan cutan consisting of clastic grains such as quartz or feldspar. skeleton grainsclastic grains such as quartz and feldspar within the soil microfabric. skeletonizedleafplant leaf decayed in such a way that cuticle and soft tissues have been removed to reveal the woody vascular traces or veins. skelsepicsepic plasmic fabric with highly birefringent plasma associated with the outer surface of skeleton grains. slickensidessmooth to striated surfacesof soil or rockproduced by shearing withinsoils. during burial crushing of soil peds and during faulting. slidesmall glass pane used for supporting thin section of rock or soil. smectite a base-rich clay mineral {('/,Ca,Na),,7(A1.Mg,Fe),~,[(Si.Al),0,,,l(OH)I.nH,O}. soda sodium oxide (Na,O). soil creep downslope movement of hillside soils,as revealed by bending into the surface of near-vertical veins and bedding planes. soilhorizons gradational changes in texture or mineral content down into parent material of a soil or paleosol from the truncated land surface. soilstructure three-dimensional features characteristic of soils. soilscapethat part of the landscape consisting of soil. solidityper cent solid grains or peds in a soil: complement to porosity. solodization soil-formingprocess of accumulation of soda and other alkalis in subsurface clays, but with acidic surface horizon, giving bleached columnar peds. solonizationsoil-formingprocess of accumulation of soda and other alkalis insubsurface, giving domed-columnar peds. soluan cutan consisting of soluble salts such as gypsum or calcite.
solum altered upper part of a weathering profile,including the various named soil horizons. sparry calcite calcite crystals large enough to be discernible under an optical microscope. sphaerosideritesmall (usually 1-2 mm diameter) spherules of radiating crystals of siderite. spherical microped microscopicsphere-shaped soil clod. typical of tropical soils and produced as oral and fecal pellets of termites. spherulite spherical aggregate of radiating crystals. spiculesmall pointed mineral body made by an animal as part of its skeletalsupport, such as the opal bodies of freshwater sponges. spodic horizon subsurface soil horizon formed by concentration of organic matter and sesquioxidesthat have been translocateddownward from anE horizon. Spodosol acidic sandy soil with B horizons enriched in organic matter, iron and aluminum or combinations of these, but not clay. sporinite liptinite (exinite)coal maceral, formed from spores and pollen. standard deviation measure of the variation of a set of data points around a mean value or a fitted curve, defined as the square root of the variance; indicated by symbolofor population ands for sample. standard error measure of the variation of a set of data points around a mean value or fittedcurve, defined as the square root of variance afterdivision by thenumber of datapoints. stele central conducting strand of tracheids or xylem found within roots and stems of plants. steppe open grassland, used mainly for dry Asian grasslands. stilt root thick adventitious roots arising from the trunk and passing down through the air for a short distance before entering the ground, as in mangroves. stoichiometry measurement of proportions of components for conservation of matter in chemical equations and formulae. stone l i e layer of pebbles or other large rock fragments confined to a narrow horizon, commonly an erosional plane, and conspicuous as the only large clasts in an otherwise fine-grainedsoil. strain volume gain or loss during soil formation and burial. stratovolcano steep volcanic cone constructed by successive layersof ashandlava. stress cutan cutan formed by differentialforceswithin the soil such as shearing as a result of swelling and shrinking induced by wetting anddrying. The surface of stress cutans is commonlyslickensided. striotubule tubular feature of soilfilled with mixed clay and clastic grains withcurved internallayering. strongly developedsoilor pdeosolwith especiallythick
Glossary (2-3 m), red, clayey or humic subsurface (B) horizons or surface organic horizons (coal or lignites)or especiallywelldevelopedsoil structure or calcic horizons as a continuous layer. subcutanic featuresmodificationsof soil material that show a relationship to a surface, but do not occur only at that surface. Sulfaquent Aquent with common sulfur minerals such as pyrite or jarosite, commonly formedunder salt marsh and mangal vegetation. Sulfaqueptkind of Aquept with common sulfur minerals such as pyrite or jarosite, commonly formedunder salt marsh and mangal vegetation. sulfate common anion (Soh2-) in soils,found in minerals such as gypsum. surface-water gley gley features formed by the pondmg of water by impermeablesoil layers above drier subsoil. swale local elongate depressionon the landscape, typically from abandoned flood channels. swampwetland vegetation of trees with acidic to neutral groundwater. syeniteplutonic rock with little or no quartz, mainly alkali feldspar,plagioclase and hornblende.
taproot single, thick, verticalroot, likethatof acarrot. taxonomic uniformitarianism assumption that soil types of the past formed under similar environmental conditions to taxonomicallysimilar soils of the present. taxonomy classification:for soils,used to distinguish the soil classificationof the US Soil Conservation Service(originally entitled Soil Taxonomy). teliitekind of vitrinitecoal maceral, formed from wood fragments with some crushed, cellular structure remaining. tepee structure inverted V-shaped local disruption of layering within a bed, as aresult of action of roots or clay swelling. terrace flat geomorphologicalsurface representing the erosional remnant of a former land surface. Tertiaryperiod of geologicaltime about 65-1.8 Ma. thin section transparent, thin (10-30 pn)slice of rock or soil mounted between glass covers,used for microscopic examination. till deposit of unweathered boulders, sand and clay dumped by a glacier. topofunction mathematical relationship between soil features and topographic setting of soil. toposequence set of soilsformed under similar climate, vegetation, parent material and time, but varying topographic setting. torbanite boghead coal: similar to cannelcoal, but more waxy, less brittle and browner in color. Torrand Andisol soil of very dry climate, with little clay or col-
349
loidal material, and abundant salts and carbonate at shallow levels within the profile. Torrert Vertisol soil of very dry climate. Torrox Oxisol soil of very dry climates. tower karst a form of karst characterized by large steep-sided pinnacles. trace fossil fossilized evidenceof the activity of an organism, such as a fossil footprint or burrow. trachytevolcanic intrusive and flow rock close to saturation with silica, including common alkali feldspar. transpiration evaporativeloss of water from leaves of plants. transportedmass fraction proportion of the mass of achemical element lost or gained during soil formation or burial. tremolite amphibole mineral, common in metamorphosed limestones {Ca2(Mg.Fe2+),[Si,0,,I(OH,F),J. Triassicperiod of geologicaltime about 2 50-206 Ma. trimasepic sepic plasmic fabric with a network of highly birefringent streaks in three preferred directions. Tropept Inceptisolsoilfound in intertropical regions. tuber potato-Iie underground storage organ of plants. tuffcompacted pyroclasticdeposit of volcanic ash, dust or lapilli. Turbel Gelisol with ice wedges or other indications of substantial cryoturbation. twin adjoining crystals of the same mineral in a differentcrystallographic orientation. UdalfAllkolsoilformed in a humid climate, usually noncalcareous. Udand Andisol soil formedin a humid climate, usually noncalcareous. Udert Vertisolsoilformed in a humid climate, usually noncalcareous. Udifluvent Entisol soil discerniblydecalcifiedand with clear relict bedding remaining from clayey alluvial parent material. Udoll Mollisol soil formed in a humid climate, usually noncalcareous. Udult Ultisol soil formed in a humid climate, deeply weathered and noncalcareous. Ultisol acidic,deeply weathered forest soil. with clayey, ferric, aluminous or humic subsurface horizon. ultramafic rocks rocks likeperidotite that are even more rich in iron and magnesium and the minerals that contain these elements than basalt, usually formed frommoltenrock deep within the Earth’s crust. umbric epipedon soilsurface horizon similar to mollic except for platy to massive structure and base saturation less than 50%,generally associatedwith forest vegetation. unconformity major temporal break in the accumulation of a sedimentary rock sequence, as indicated by fossil evidenceof
350
Glossary
age or by deformation of underlying layers before deposition of overlying layers. underclay clayey paleosol beneath a coal seam. undulic soil plasmic microfabricthat is almost but not quite isotropic,sovery dark whenviewed in thin section under crossed nicols. Ustalf Al6sol soil of dry summer-wet climates. Ustand Andisol soil of dry summer-wet climates. Ustert Vertisol soil of dry summer-wet climates. Ustochrept Ochrept soil formedin a dry climate, usually with carbonate nodules. Ustoll Mollisol soil of dry summer-wet climates. Ustox Oxisol soil of seasonally dry climates. UstropeptTropept soil formed in a dry climate. valley calcrete form of groundwater calcrete formed by precipitation from the water table nearstreams. value thedegreeof lightness of a color in theMunsel1system of color classification. variance a measure of the variation of a set of data points around a mean value or a fitted curve, defined as the sum of squares of the deviations from the mean dividedby the number of data points: indicated by symbol o2for population and s2for sample. vascular trace bundle of tracheidcells that form veins of leaves and conducting strandsof plants. vein anarrow crack through rock, commonly filled with minerals such as quartz or calcite. vermicular wormy, full of elongate cavities. vermiculiteswellingclay mineral {(Mg,Ca),,,,~,,.,(Mg,Fe3+,A1),,,,[(Si,Al)XOLOl(OH),.nH,O}. vertic showing some properties of Vertisols,such as slickensides and deepcracks. Vertisol thick, very clayey, slickensidedsoil,often with internal deformation of horizons. vertisolizationsoil-formingprocess of cracking and heaving of soil clays as aresult of drying and wetting. vertiiation soil-forming process of cracking and heaving of soilclays as aresult of drying and wetting, synonymous with vertisoliation. very strongly developedsoil or paleosolwith unusually thick (>3 m) subsurface (B) horizons or surface horizons (coal or lignites):such a degree of developmentis found mainly at major geologicalunconformities. very weakly developedsoilor paleosolwith little evidence of soil development apart from root traces and abundant sedimentary, metamorphic or igneous textures remaining from parent material. vesicular full of small, near-spherical cavities. vitrain bright coal: brilliant, black, nonlaminated coal, clean to the touch and breaking withconchoidal fracture.
Vitrand Andisol soil rich in glassy volcanic shards. vitrinite component of coal, grayish and yellowish white, with low relief under thereflected-light microscope,andlight orange to dark red under the transmitted-light microscope, derived from coalification of wood fragments. vitrinite reflectancemeasure of the shininess of the coal maceralvitrinite. the percentage of light reflected from the maceral. vivianite blue to white mineralof marsh soils
[Fe,(P04),.8H,01.
void small open spaces within soil microfabric, generally crushed out of buriedpaleosols. volatilematter easily moved materials: for coal, volatile material includes water, sulfur and nitrogen. volcanic ash small particles of volcanic rock, crystals and glass that settle out through the atmosphere, lakes or oceans after volcanic eruptions. volcaniclasticformed from particles of volcanicrock. vosepicsepicplasmic fabricwith highly birefringent plasma associated with the walls of voids: may be difficult to recognize in paleosolsbecause their voids are usually crushed during burial. vug small unfilled cavity in rock or soil. Wakley-Black method wet chemical titration method for the determination of abundance (weight percent) of soil organic carbon. water potential negative water pressure maintained within tracheids that transports water through a plant. waterloggedsoil soil that is saturated with water. weakly developed soil or paleosolwith a surface rooted zone (A horizon), as well as incipient subsurface clayey, calcareous, sesquioxidicor humic or surface organic horizons, but none of these developedto the extent that they would qualify as argillic. spodicor calcic horizons or histic epipedons. weathering rind thin outer zone of weathering found on rock and mineral grains within a soil; usually adiffusion sesquan in Brewer terminology. Weaver index measure of clay crystallinity using height of Xray diffractometer peak. Weber index measure of clay crystallinity using width of X-ray diffractometer peak of clay compared with that of quartz. wetland part of the landscape that is waterlogged or inundated for a substantial part of the year. wooded grassland trees giving 1 M 0 %cover,isolated and scattered among grasses. woodedshrubland trees giving 1 0 4 0 %cover,isolated and scattered among shrubs such as sagebrush. woodlandvegetationof closely spacedtrees 2-l0mtall.
Glossary Xeralf M s o l soil of dry winter-wet climates. Xerand Andisol soil of dry winter-wetclimates. Xerert Vertisol soil of dry winter-wetclimates. xeric of dry winter-wet climates. Xeroll Mollisol soil of dry winter-wet climates. Xerult Ultiol soil of dry winter-wetclimates. X-ray diilkactometermachine used to identifyminerals by the angles at which their crystal facesreflect a focused beam of X-rays. X-ray fluorescencespectrometry method of chemical analysis
351
using wavelengthsof secondary radiation induced by bombardment of sample withx-rays. XRF X-ray fluorescencespectrometry. xylem woody tissue of tree trunks, roots and veins of leaves, consistmgof tracheidcells. zeolitegroup of water-solubleminerals of low relief under the petrographic microscope
{(Na,,K2,Ca.Ba)[(AI,Si)0,1,.xH20}.
zircon highly birefringentheavy mineral [Zr(SiO,)].
Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
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39 3
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Soils of the Past: An introduction to paleopedology Edited by: Gregory J. Retallack Copyright 02001 Gregory J. Retallack
Index
Page numbers in bold refer to figures and tables. Definitionsare also to be found in the Glossary. acarines 140 AcastaGneiss 209 accumulation of paleosol sequences
194-203
achondrites 221 acid consumption, oxidation, chemical weathering 46 acidification British Isles,Lake District,effectsof soil use 3 30-3 and cation exchangecapacity (CEC) 5 5 chemical weathering 42-5 Eh-pHstability 156 LakeDistrict,effectsof soil use 3 30-3 and soil structure 29 Acrudox 290 Adirondacks 177 aeration, soilstructure 29 aerophores 15 afforestation 280-99 s~walsoforcstsoils: forests agglutinates 211,212 agric horizon 71 air circulation patterns 111 albic horizon 71 AldanShield 240
Alfisol(fertileforestsoil)75.147.177,178. 179,208,282,285
profile, grain-size and molecular weathering ratios 44 algae charophytes 152,263 see also microbes alkalinesoils 42-5 allophane 175 alluvial-volcaniclasticsediments 11 alluvium 171,172,177 architecture. time scales 196-8 clay distribution 189 alpinefellfield 121.152 aluminium Al-hydroxyions 42 chemistry 175 Ambitisporitrs 2645 amphibians 144-5,291 anaerobic processes 248 bacterial decay 19 andesite 179
weatheringrinds. timescales 193 andesitic arcs 231 Andic Ustochrepts 1 5 8 Andisol(volcanicashsoil) 73,179,208,
287
andisolization 60 angiosperms 135-6 afforestation 292-5 early 292 animals land 263-79 origins 273 annelids 137,183 anorthosite. lunar 210 Antarctica AztecSiltstoue 282,283.289 Graphite Peak, Permian-Triassic boundary
295-7
organic matter carbon isotopiccomposition 126,127 microbial 152 Rosemarypaleosol 282,283,284 temperate rainforest soil 148 Weller Coal Measures, permafrost structures 119-20,122,123 anthropic epipedon 70-1 anthrosolization 60 ants 143-4 apes 321-33 aphids 141 Aqualfs 178 Aquods 148,332 Aquolls 177 arachnids 140 Archaeopteris 280.281.287 Archean greenstonebelts 231 paleosols 241-2 terraues 239 argillaus 165 seasonality indicators 120-2 argillic horizons 70,90 clayey subsurface (Bt)development 171,
188-9
thickness, chronofunctions 190 argillic minerals, chemistry 175 argillic regimes 186 argilluviation 60 argondating 183 Aridisol(desertsoi1) h5.74.150,164.
177-8,208,289
profile, grain-size and molecular weathering
ratios 44 Arizona desert 1 51 MescalLimestone 257,259 PetrifiedForest National Park I63 arthropods 13844 ashbeds 167 ashsoils 175 Athabasca Basin, Canada 240 atmosphere change 230,279 indicators of greenhouse effect 125-7 oxygen abundance 238 oxygenation 231-9 Paleozoic 279 reducing 246-52 atmosphericoxidation, Precambrian soils 2 37-9 augite 175 Australia Ediacaran fossils 263 Gelisols 286 Mt Isa. LochnessFormation 260 MtRoepaleosol 257.259 mukkara 39,120.124.181 Narryer 209 PilbaraRegion 209 Queensland.Australia early life 25 7 weathering profiles, stratigraphic relationship 8 5 Toorong Granodiorite 96 Warrawoona Supergroup 240 Austrulopithecus 317-19,324-8 AztecSiltstone. Antarctica 282.283,
289 bacterial decay see microbes Badlands National Park 158 Andisols 287 hone preservation 157 BruleFormation. Poleslide Member 199 EoceneOligocenepedotypes 8&1,15 7 ,
158,194-203
fluvialdeposition 198 Oligocenegrasslands 305-6 ScenicMernber 199 SharpsFormation 180,202 time scales, accumulatiou of paleosol sequences 19G-203 unconformities 162 vegetation 81
395
396
Index
banded iron formations 229 Baragwanathia 263-4 Barberton. South Africa 209.240 Barremian 292 BanassCoal, Siberia 272,276 basalt, weatheringrinds, timescales 193 basaltic rocks 179 basket podzol 1 35 bauxites 1 8 5 , 2 3 9 4 0 bauxitic paleosols. relict 9 bees 144,295 beetles 142 beidellite 1 71 berylliumdating 183 Besselia 264 Betula 331-2 BhuraSeries 1 6 9 , 1 7 0 biogeochemicalcycling 60 biological weathering bioturbation 56-7 humitication 50-3 nutrient consumption 53-6 biotite 175,178 bioturbation 56-7 and nodules 8 7-8 seealsoburrows birds 145 birnessite 47-8.132 blackphytokarst 114,134 Bloomsburg Formation 23 3 , 26 9 , 2 7 0 , 2 7 1 , 273 bog 152 productivity,stature and biomass of vegetation 5 9 'bogpeople' 154-5 boxcanyons 161 brakelands 150 British Isles, LakeDistrict acidification 330-3 pollendiagram 333 BruleFormation 1 9 6 , 1 9 7 PoleslideMember 1 9 9 , 2 0 2 bryophytes 1 3 4 - 5 2 6 4 bugs 141 Buntley and Westin color scores 194 burial and alteration processes 8 7-1 01 ageing 87-9 cementation and neomorphism. differentiation 100 cementationof primary porosity 9 1-3 coalificdtionof peat 97-8 common patterns of alteration 100-1 compaction by overburden 93-5 compaction equation 94 correlation of thermal maturation indices 99 decompositionof organic matter 89 gleization 9 0 graincontact 94 illitizationof smectite 39,95-7 kerogen maturation andcracking 98-9 metamorphism 100 neomorphismof carbonate 99
reddening of iron oxides and hydroxides 90-1 zeolitizationand celadonitization of volcanicrocks 97 burial compaction equation 9 4 burrows examples 146 liverwort-millipede polsterlands 2 66-9 mammals 145 stone lines 163 toads 144-5 water table indicators 165 see also bioturbation; pedotubules butterfliesandmoths 1 4 1 C (saprolite)horizon 171 cactusdesert 151 caddisflies 141 caecilians 145 calcareous nodules 288 calcareous soils 42 calcichorizon 72,151,165 Bk 114-15 timescales 185-8 climaticdrying 116 nodules 114 precipitation indicators 114-15 . 1 1 5 calcic regimes 186 Calcids 244 see also calcichorizon calcification,formation of soil 62 calcite cementation 91-2 neomorphism 9 9 , 1 0 0 pedogenicand groundwater 92 calcium chemistry 175 in soils, rocks and organisms 2 5 0 calcretes 240 precipitationindicators 112-14,240 calichenodules 180,241 California LakeTahoe 1 7 4 , 1 9 3 MercedRiveralluvium 1 8 9 , 1 9 4 Pleistocene 1 9 2 Quaternaryunit stratigraphy and sedimentary formations 83 CaliforniaCoast Ranges 1 95 Callixylon 280.281 Calluna 331-2 cambic horizon 71 Canada AthabascaBasin 240 karst, SteeprockGroup 241 PitzFormation 289 pre-Huronian paleosols 232-5 redpaleosols.NovaScotia 266,275 ThelonBasin 2 4 s 1 Canadian Shield,relict desert 241 carbon isotopiccomposition 54,123-5.150.183. 227,258-9.305 seealsocharcoal: organic matter
carbon cycling bacteria 1 31 carbon dioxide and grasslands 31 3 greenhouse atmosphere indicators 125-7 oxygen demand ratio 237-8 carbonaceouschondrites 2 2 1 , 2 2 2 4 carbonate accumulation,time scales 187 caries texture 186 fieldscale 24 neomorphism 99 carbonatecements 91-3 carbonate fixation 152 carbonate-rich soils 1 7 carbonate/calcareousnodules 12,29-30 carbonatite 73 volcanic flows 174 carbonatite-nephelinitecindercones 9 0 Carboniferous 276-9 cam 153 catena 1 6 6 7 0 defined 160 downslopevariation 1 6 1 lateralvariation in paleosols 166-7 cationexchange capacity (CEC) 54-5 and soil acidity 5 5 Catskills 177 celadonite 73 celadonitization of volcanic rocks 9 7 cementation andneomorphism, differentiation 100 of primary porosity, paleosol alteration after burial 91-3 Cenomanian 292 centipedes 1 3 9 4 0 ChadronFormation 1 9 6 , 1 9 8 , 2 0 2 ChamberlainPassFormation 195,200, 202 chaparral 149 charcoal distinguishing fromcoalifiedwood 40,47 seasonality indicator 123 wildfire 154.299 seealsocarbon charophytes 1 5 2 , 2 6 3 Chasmatapsis 265 chelicerates 140 CheltenhamFormation 2 8 9 , 2 9 0 chemical composition calculation 4 3 parentmaterial 175-6 precipitationindicators 115-16 chemical weathering acidification 42-5 chemicalreactions 41.45 formulae relating volume, thickness and chemical changes 46 hydration 49 hydrolysis 42 oxidation 41,45-50 acid and oxygen consumption 46 salinization 49-50.62
Index chert chemistry 175 RhynieChert 269 chlorite 178,238 chondrites carbonaceous 222-4 enstatite 221 chronosequences 183-5 cicadas 141 classification of climate 109-1 2 classificationof paleosols 64 classification of soils 63-76 cautionarynote 76 FA0 world map 6 3 4 , 6 6 7 peds 26 US soiltaxonomy 6 4 7 6 climate,vegetation and profile 69 diagnostic horizons and properties 67-72 classificationof vegetation 111 clastic dikes 93 clay minerals bright (mosepichasepic)clay fabric 32-5. 189 bright (sepicplasmic)clay fabric 189 chemistry 175 ascrystalgenes 2 5 3 distribution, alluvium 189 andearlylife 252-3 precipitation indicators 116-1 7 selfish soil. originsof life 2 54-6 stability series 4 3 synthesisof monomers 2 52 total profile. chronofunctions 190 clay swelling mukkaraandgilgai 120,163 physical weathering 39 cliffs 161 climate 108-27 change, oceans 3 13 classification 109-12 Koppen's categories 110 Koppen'sclassi6cation 109 defined 108 greenhouse atmosphereindicators 1 25-7 precipitation indicators 112-18 seasonalityindicators 120-5 temperatureindicators 118-20 clinoptilolite 73 CLORPT acronym 106 coal Antarctica, permafrost structures 119-20, 122,123 BanassCoa1,Siberia 272,276 earliest 276 precipitation indicators 112 rock types of coal measures 5 2 coal balls 153 coalification of peat, alteration after burial 9 7-8 coalifiedwood, distinguishingfrom charcoal 40,47 coastalcliffs 161
collembolans 141 color scores Buntley and Westin 194 Hurst 194 MunsellChart 23 Colorado 187 compaction 181 by overburden 93-5 Conataassemblages 157.158 concretions glaebules 30 seasonalityindicators 120-2 continental crust crustalplate tectonics 23 1 differentiation 2 3 9 4 2 continental emergence 241-2 Cooksonia 264,2 75 Cooksonia-trigonotarbid brakelands 2 69 Coprtnisphaera 29 1 coprolites 145,155 corestones 173 cornstone.defined 12.20 cradleknoll 136,147 Cretaceous reconstruction 294 Cretaceous-Tertiary boundary 295,297-8 crickets 141 cross-cutting,minerals 88 crustaceans 138-9 crustalplate tectonics 23 1 crustal rocks differentiation 2 3 9 4 2 eight elements 41 cryoturbation 60 crystallinity, parent material 175 crystals 31 cumulic horizons 19,164 cutans 28-9,33,165 terminology 2 8 cyanobacteria 131,134.257 cypress.bald 153,191 definitions 7-9 dehydration 41 Dendropithecus 32 1 Denisonpaleosols 233,234.258 desert pavement, precipitationindicators 1 18 desertscrub 151 desert soilseeAridiso1 desert varnish 132,258 desertification,Greece, human impact 328-30 Devonian 2 76-9 DhokPathanFormation 1 6 6 , 1 6 7 , 1 6 8 , 1 6 9 diagenesis defined 87 iron hydroxide minerals 194 pressurekemperature 88 Dinosaur National Monument 293 dinosaurs 145,208.292 Cretaceous-Tertiary boundary 295, 297-8 effecton afforestation 292 diorite 179
397
disconformities 10-11.161-2 dissolution 41,49 DNA 246-8.251 dolostone,karstiication 113,166,177-8 drabpaleosols 234 drab-haloedroottraces 17-19,90,149 drywoodland 149 duricrusts 10,289-90 Durids 244 seealsosilcretes duripans 1 7 Dystrochrepts 148.260.332 Earth early landscapes 2 2 7 4 5 surfaceenvironments 229 earlylifeonland 246-62 magnetic field reversal 166-7 origins 22 5-6 earthworms 137,183,264 ecosystemevolution 275-6 ecosystemtraces 145-53 bog, marsh and fen 152 brakeland 150 desert scrub 1 51 dry woodland 149 early successionalvegetation 147 fire-prone shrubland 149 forest and woodland 147-8 heath 148-9 microbialearthandrockland 151-2 oligotrophic forest 148 open grassland 149-50 polsterland 150-1 rainforest 148 saltmarsh 152-3 shrubland 150 swamp,carr and mangal 153 taigaandkrummholz 152 tundraandalpine fellfield 152 woodedgrassland andshrubland 149 see also headingsabove ecotypes, productivity, stature and biomass of vegetation 5 9 Ediacaranfossils 263 Egypt 177 Eh (electrode potential) 45-7 Eh-pH stability 156 elements,traceelements 260 Elginerpeton 29 1 enchytraeids 13 7 endocarps 154 energy, solar 249-50 Entisol(incipientsoi1) 72-3,147,155, 177-8,208,285-6 Eoastrion 2 5 6 , 2 5 7 Eocene-Oligocene, Badlands National Park, SouthDakota 80-1.157.194203 Eohostirnella 263,269 eolian dunesandstones 241 eolianpaleosols 182 Eospermatopteris 280,287 epipedons.soilhorizons 65-71.150
398
Index
Erehus 261.275-6 cmsional gullies, unconformities 162. 198 erosional planes 162-3 Ethiopia AfarTriangle 326 HadarFormation 3 2 5 , 3 2 6 Pliocene woodlands 324 Europe.soils, history 33 1 eurypterids 264-5 Eutrandepts 180 evaporite minerals p d k O S O k 164 precipitation indicators 117-1 8 evolution 29 1-9.314-1 6 angiosperms 292-5 coevolution 315-16 crisis and recovery 2 9 5-8 dinosaurs 29 2 massextinction 297-8 early land invertebrates 291-2 effectof afforestation 291-9 grasses 314-16 naturalselection 3 15 soil invertebrates 291 speciation 3 14-1 5 evolution of continental ecosystems 275-9 blanketing bogs 2 76 filtering air 2 79 tamingstreams 276-8 extinctsoils 208 extinctions 295-8 FaciesLaw, Walther’s I 6 7 FA0 world map, classificationof soils 6 3 4 . 66-7 feces see coprolites feldspar 175 fen 152 seealso bog; marsh ferallitization.formation of soil 60,62 fermentation 248 field capacity,defined 38 Finland 237 lire charcoal seasonality indicator 40, 47. 12 3 wildfire 154,299 and grasslands 3 1 3 humanuse 147 succession 150 fire heating, physical weathering 40 fire-proneshrubland 149.1 50 fish 144 lungfishburrows 144.291 tlag trees 152 flatworms 136 Fluvents 321 fluvial-colian paleosols 182 fogloma 150 foliated rocks 174 folisticcpipedon 70 footprints 129.145
forestsoils 2824,282-91 diversification 28 5-9 1 origin offorestecosystems 285 similaritieswithmodern world 290-1 seealso Altisol: Spodosol;Ultisol forests afforestation 280-99 earlyMiocenc.Kenya,SW 321-3 ecosystem traces 147-8 oligotrophic forest 1 4 8 productivity,stature and biomass of vegetation 5 9 rainforest 148 temperate rainforest 1 4 8 see also afrorestation: woodlands formation of soil 8-9.37-62,60,171-82 base line assumptions 179-82 fresh parent material 180 onestableconstituent 180-1 quantifying soil and paleosol development 182 uniform parent material 180 volumechange 181 biological weathering 50-7 humitication 50-3 rock typesof coalmeasures 52 chemicalweathering 91-50 common processes 5 7-62 development stages 186.187 models 105-7 physical weathering 3 7-40 productivity,stature and biomassof vegetation 5 9 timescales 9 , 1 8 5 see also calcification;ferallitization: gleization:lessivage: parent material: podzolization:salinization formulae. relating oxidation. volume, thickness andchemical changes 46 fossils 155-9 coprolites 145,155 microfossils 256-7 preservation in paleosols 15 3-9 chemistry 154-5 time 1 5 5 trace fossils 129-30.2 5 7-8 fragipans 17, 72 framboids 46,132 France, GresdeChaunoy 240 friability. induration scale I 7 3 fulgurites. defined 14 fungi 1 3 3 4 in root traces, calcite fibers 1 33 fusinite (charcoal).seasonality indicator 123 Gaia 261.275-6 galls 136 ganisters.defined 12 . 2 0 garrigue 149 Gelisol(permafrostsoi1) 75,208,286 geochemistry 5 geological range of paleosols 208
geological time scale 184 geomorphologic indicators 160-4 Georgia. Okefenokee Swamp 191 geosol 83-6 defined 83.84 lateral variants 8 4 naming 86 for stratigraphic mapping 84-6 gibbsite 175. I 7 8 gilgai Australia 39 seasonality indicators 1 2 0 gilgai microrelief 39,163 glacial features,seasonality indicators 121 glaciation periglacial soils 241 Precambrian 243 glaebules 25.29-3 1 description of glaebules and mottles 3 1 nodulesand concretions 30 glass andesitic 175 lunar 21 1-12 volcanic 175, 179,181 gleization formationof soil 59-60 paleosol alteration after burial 90, 149 Cleskaclay 1 5 7 glossicfeatures 90 gneiss 162.166.209 foliation 1 7 3 goethite 165.166 ironcements 9 3 gradation, defined 209 graincoatings 1 7 3 grainsize parent material 174-5 weathering 192 granite 1 7 1 , 1 7 2 granitic rocks 178-9 chemistry 1 7 5 , 2 3 1 grasses 3 0 ( t 1 6 grasslands dry continental interiors 30Lb-16 early grassland soils 303-5 evolutionaryproccsses 314-1 b coevolution 3 1 5-1 6 naturalselcction 31 5 speciation 314-1 5 human impact on landscapes. Kenya and Tanzania. Pliocene mosaic 32 5-7 Kenya. Middle Mioceneshort sod grasslands 306-8,309-10,323 opengrassland 149-50 Oregon. Late Miocenetall sod grasslands 308 origins of grasslands 3 12-14 plant formations 147 productivity,stature and biomass of vegetation 5 9 roots 1 6
Index soil seeMollisol South Dakota. Oligocene rangelands 305-6 timerecord 308-12 woodedgrassland and shrubland 149 Greatplains. grasslands 312 Greece, desertification 32 8-3 0 Green Clays 179,209.235.237 irondepletion 243 Green RiverFormation, Wyoming 163 greenhouse effect atmospheric indicators 12 5-7 Precambrian atmosphere 241 Greenland, G6thaab 209 Gresde Chaunoy, France 240 grus 179 gypsic horizon 72.118 gypsum cements 9 3 pseudomorphs 227 Hadar Formation, Ethiopia 326 Hapludalf 332 Hapludult 20 Haplustalf 283, 321 Haplustoll 20 heath 148-9 heather 148-9,331-2 heathland. productivity,stature and biomass of vegetation 5 9 hematite 165.166.178 ironcements 9 3 hemipterans 141 hillslope development, time scales 194-6 histic epipedon 70,113,189 histic regimes 186 history of paleopedology 3-6 soil classification 63 Histosol(peatysoi1) 70.73-4.155.179.208, 287-8 defined 190 geological history 288 oldest 276 precipitation indicators 112 Hokalampi paleosol.Finland 237 Holdridge'sclassificationof vegetation 111 Homo spp. 3 18-20.32 7-8 horizons see soil horizons hornblende 178, I 79 hornfcls 171 horses 157, 320 horseshoecrabs 140 Houston Black Clay 178 howardites 224-5 human impact on landscapes 3 17-33 early human ecology 324-8 effectsofsoiluse 328-33 British Isles. Lake District, acidification 3 30-3 Greece.desertification 328-30 humanorigins 320-4 early human evolution 324
Ethiopia, Pliocene woodlands 324 Kenya,South-Western early Mioceneforests 32 1-3 Mid-Miocenegrassland mosaic 323 landuse 328 Pakistan, Northern, Late Miocene woodlands 3 2 3 4 humification 5Q-3 coalmacerals and their origin 53 rock types of coal measures 5 2 hummockand swale gilgai microrelief 39, 163 Huronian paleosols,Canada 232-5 Hurstcolor scores 194 hydration 45 chemical weathering 4 9 hydrolysis,chemical weathering 4 2 , 4 5 hydromorphism 60 hydronium ion. and pH 42-5 hydrothermal alteration 234 hymenopterans 1 4 3 4 hypersthene 1 7 5 , 1 9 2 icewedges 40 illitization of smectite. paleosol alteration aftc:r burial 39,95-7 illuvial horizon 70.90.165 imogolite 175 Inceptisol(youngsoi1) 73,147,149,208, 286-7 Green Clays 179, 2 0 9 , 2 35 induration, parent material 1 7 3 4 IndusRiver, Pakistan 1 6 7 , 1 6 8 . 1 6 9 , 170 inertinite macerals 5 1 , 53 insects 140-4 springtails 140-1 intermediate redox (lowEh) soils 47-8 Iowa 160,161,189 iron banded formations 229 carbon dioxide-oxygen demand ratio 237-8 cements 93 chemistry 175, 1 7 6 drab-haloedroottraces 18.90 ferallitization.formation of soil 60, 62 inmeteorites 220-5 neoferrans 28-9 Staca pedotype 235 weatheringratios 48-9 seealsogoethite: hematite: pyrite: siderite iron carbonate see siderite iron oxides and hydroxides burial reddening 90-1 chemistry 1 75 diagenesis 194 and early life 2 5 3 ironoxyhydrates 17.48 Mossbauer spectroscopy 49 ironsilicates 238 isopterans 143
399
jarosite 4 6 joints 173-4 JuniataFormation 240.266-7,270,271, 273,274,275 funiperus 331 Jurassic paleosols 293
Kukabekia 2 5 6 . 2 5 7 KalaSeries 1 6 9 , 1 7 0 Kampecuris 265 kandic horizon 70 Kansas, Cretaceous reconstruction 294 kaolinite 175,178.278 oxygenisotopic composition, temperature indicator 1 1 9 . 1 2 0 karst 49.16h. 241 phytokarst, black 114.134 precipitation indicators 112-7 4 SouthAfrica 241 SteeprockGroup 241 karstbauxites 239-40 Kenya Amhoseli 1 5 7 carbonatite tuff 1 7 4 Early Mioceneforests 32 1-3 Kanapoi, Koobi Fora 326 LakeMagadi 163 LakeTurkana 167,326 Middle Miocene short sod grasslands 306-8.323 kerogen maturation and cracking,paleosol alteration after burial 98-9 KhakistariSeries 1 6 9 , 1 7 0 Koppen's classification of climates 109 Krakatau. volcanism, colonization 129 krotovinas 1 9 , 9 0 . 9 3 krummholz 152 lacewings 1 4 1 lacustrinesediments 1 9 1 lakes 157-8 LalSeries 1 6 9 , 1 7 0 land invertebrates 1 4 1 effect of afforestation 291-2 landorganisms 2 72-5 1andscapes.early 7 - 1 2 , 2 2 7 4 5 atmosphericoxygenation 2 3 1-39 calculating frompaleosols 23 7-9 Canada, Ontario, Pre-Huronian (2450 Mafpaleosols 232-5 Scotland, North Western, PreTorridonian (810Ma)paleosols 235-7 blanketingsoils 7-9 calibratingcontinental emergencefrom paleosols 241-2 differentiation of continental crust 23942 human impact 3 17-3 3 effectsof soil use, Greece 328-30 majorunconformities 10-1 I Precambrian scenery 242-5
400 Index landscapes,early (continued) Quaternarypaleosols 9-10 sedimentary and volcanicsequences 11-12 topographic relief indicators 39,160-70 laterites 72 and Oxisols 240 paleosolunconformities 1&11 lateritic paleosols,relict 9 , l O leaflitter 50-1,152,154,157 lepidoptera 141 Lepidosigillaria 2 8 7 lessivage,formationof soil 61-2. 177 levees 191 Lewisiangneiss 162,166 lichens 1 3 4 Liesegang banding 173 life, early 246-62 ecology and environments, early Paleozoic 273-5 ecosystemtraces 145-53 metabolic processes of organisms 248 multicellular,inpaleosols 265-6 organismtraces 1 2 9 4 5 origins 248-52 selfishsoil 254-6 paleosoltraces 153-9,256-60 uniqueness andevolution 260-62 limestone 177,178 seealsokarst liptinitemacerals 51,53 liverwort-millipede polsterlands 266-9,2 72 liverworts andmosses 134-5 liwiviation 6 0 Llandovery 272 loess 150,177 Palouse Loess 15 1 Louisiana 1 9 1 lungfishburrows 144,291 Lycopodium 269 macerals 5 1 origin 5 3 McKenzie river, Oregon, soil series and vegetation 79 mafic minerals 175 magnesium chemistry 175 in soils, rocks andorganisms 2 5 0 magnetic reversal 166-7 mallee 149 mammals 145 andgrasses 302-3 mangal 153 productivity, stature and biomassof vegetation 5 9 mangalpaleosol 1 4 mapping andnamingpaleosols 77-86 deeply weathered rocks 86 paleoenvironmental studies 7 7-82 stratigraphicstudies 83-6 maquis 149
marcasite 153,165 Marianas 231 marinefossils 264-5 marinesediments 177-8 marls 178 marsh 152,269-72 productivity,stature and biomass of vegetation 5 9 time record 2 72 see also bog; fen Martian soils 2 16-20 comparison of planetary bodies 2 10 composition 21 7-19 development 219-20 massextinctions 295-8 mass transport function 181.182 mattoral 149 melanic epipedon 70 melanization 60 mesosiderites 224-5 metabolic activity 246-52 metamorphism, paleosol alteration after burial 100 meteorites 220-5 carbonaceous chondrites 2 2 2 4 mesosideritesand howardites 224-5 types 2 2 1 methane 238 greenhouse atmosphere indicators 12 5-7 methanogenic bacteria 1 3 1 mica 178 microbes 13&3 anaerobic bacterial decay 18-20,130 earlylife 246-52 mineralcycles 1 31 ministromatolites 132 noduleformation 132 organizationof E. coli 2 4 7 rockvarnish 132 and soil stability 13&1 microbial earth, productivity.statureand biomassof vegetation 5 9 microbialrockland 151-2 microcline 1 7 5 , 1 7 6 microfabrics 32-6 asepic 34 brightclay fabric 32-5,189 bright (mosepichnasepic) 32-5,189 grain 3 5 insepic plasmic 39 omnisepicplasmic 39 sepicplasmic 32,166,189 terms for plasmic microfabric 33 termsfor soilmicrofabric 3 5 water table indicators 166 microfossils 256-7 micropeds.and termite nests, temperature indicators 118-19 millipedes 139-40.154.264-5 mineral grains. weathering 192 minerals availability 53-6
composition,parent material 1 75 cross-cutting 88 crustalrocks 4 1 in soils, rocks and organisms 2 5 0 traceelements 260 Minnesota 2 5 7 Miocene C,photosynthesis 150 DhokPathanFormation 1 6 6 , 1 6 7 , 1 6 8 , 169 Mississippi Delta plain 191 modelsof soil formation 105-7 types 106 moder humus 5 1 molecular weathering ratios 42-3.44,45, 48-9 mollic epipedon 6 5 Mollisol(grasslandsoi1) 74,147,149,150, 177-8,208 distribution 3 11 productivity,carbon storage. calcic horizon andstructure 57 profile, grain-size and molecular weathering ratios 44 mollusks 136-7 montmorillonite 2 5 2 Moon soils 209-1 3 comparison of planetary bodies 2 1 0 soilcomposition 21 1-12 soildevelopment 212-13 morhumus 51 MorrisonPormation 292,293 mosses 264 mottles 31 mukkara 181 Australia 3 9 , 1 2 0 , 1 2 4 , 1 8 1 seasonality indicators 120 SouthAfrica 124 mullhumus 51 Munsellcolorscores 23.194 muscovite 175,178 mycorrhua 1 33 4 myriapods 1 3 9 , 2 6 4 , 2 6 5 natric horizon 7 1 Natrustoll 326 natural gas, kerogen maturation and cracking 98-9 natural selection 3 15 nematodes 1 36 neoferrans 28-9 neomorphism. and cementation, differentiation 100 nephelinite 73 neuropterans 1 4 1 New Mexico, carbonate accumulation 188 New Zealand. andesiticglass 175 NileRiver 177 nitrogen cycling bacteria 131 nodules and bioturbation 87-8 calcichorizon 114,186-7,288
Index caliche 180 formation by microbes 132 glaebules 30 oxygenavailability 165-6 nomenclature. soil taxonomy 65 nomenclature, stratigraphy 77 NovaScotia, redpaleosols 266,275 nucleicacids 246-8.251 nutrients early availability 249-52 soils 53-6 in soils. rocks andorganisms 2 5 0 Ohorizon 152 oakforest 191 oceanic islandarcs 239 oceans, climatechange 3 13 Ochrepts 244 ochric epipedon 70 OlduvaiGorge,Tanzania 205.327 Oligocene 80-1.157.305-6 oligotrophicforest 148 olivine 175.179 OneontaFormation 271 onychophorans 138 opal,plant 93.302 opal bodies 154 opaline silica 175 opengrassland 149-50 Ordovician 270-1.272.276-9 Oregon 116.199.200 Late Miocene tallsodgrasslands 308 Painted Hills, calcic horizon and climatic drying 116 RattlesnakeTufT 3 11 soil series and vegetation,McKenzie and WillametteRivers 79 oreodons 157 Orestovia 272 organicmatter C by Walkley-Black titration 89 carbon isotopiccomposition 1 2 6 drab-haloedroottraces 17-20,90 gleization 90 paleosol alteration after burial 89-90 organisms 128-59 originsof life 248- 52 dating 260 land animals 2 73 land plants 2 72-3 selfishsoil 254-6 Orthents 244 orthopterans 141 overburden.compaction,paleosol alteration afterburial 93-5 oxichorizon 71 oxic micropeds. temperature indicators 118-19 oxidation chemicalweathering 41.45-50.45 Green Clays 2 3 7 oxidized (highEh) soils 48
Oxisol (tropicaldeeplyweatheredsoil) 76,148, 178.185.208.240.289-90 oxygen chemistry 175 isotopic composition, temperature indicator 119.120 oxygen consumption,chemical weathering 46 oxygen/carbondioxidedemand ratio 237-8 oxygenation of atmosphere 231-9 PaintedHills.Oregon Big Basin Member 199 calcic horizon and climaticdrying 116 timescales 199,200 Pakistan DhokPathanFormation 166,167,168, 169 Miocenewoodlands 3 2 3 4 paleocatenae 166-70 paleoclimate 111-12 paleopedology,history 3-6 paleosols developmentstages 186.187 features 13-3 6 Paleozoic,globalchange 2 7 7 Palouse Loess 151 paludization 60 pampas 150 PanduSeries 1 6 9 , 1 7 0 parent material 171-82 general properties 172-6 chemicalcomposition 175-6 crystallinity 175 grainsize 174-5 induration 1 7 3 4 scaleforstrength 173 uniformity 173 types 176-9 weathering indications 171 seealso formation of soil Parka 2 6 3 4 peats 51 accumulation, time scales 189-91 bogandfen 152 coalification of peat, paleosol alteration after burial 97-8 andcoals precipitationindicators 112 thermal maturation indices 99 Sphagnumbog 152,191,331-2 swamp 153 see also Histosol pedalfers and pedocals, precipitation indicators 114 pedogenesisstudies 20 7-8 pedogenic and groundwater calcite 9 2 pedolith. defined 9 pedology,history 3-6 pedon,defined 8.78
401
pedotubules 25.31-2.33 and nodules 8 7-8 terms 33 see also bioturbation: burrows pedotypes, soil facies 82 peds 26-8,165.186 classification 26 lentil 289 Pennsylvania BloomsburgFormation 270,271,273 rhyniophytes 269 clay horizons 189 JuniataFormation 240,266-7,270,271, 273.274.275 SchuykillMemberUltisol 282 Pennsylvanian, CheltenhamFormation 289, 290 peridotite 179 periglacialfeatures,climatic inferences 1 2 1 periglacialsoils 241 permafrostsoil 75,286 permafrost structures 119-20,122 permeability,defined 3 8 Permian-Triassic boundary 295-7 Persian Gulf 163 PetrifiedForest National Park 163 petrocalcic horizons 163 petroferic contact 72 petrography 26 PH broadcategories 42 Eh-pHstability 156 hydroniumion 42-5 soilstructure 29 see also acidification phonolites 179 photooxidation 229 photosynthesis 246-7 C,andC,pathway 150,304 CAMplants 150,304 physical weathering clay swelling 39 distinguishing charcoal from coalied wood 40 fire heating 40 fluid flow 3 7-9 formation of soil 3 7 4 0 freezing 40 loosening 3 7 phytoliths 150,154 Pilaseries 169,170 pingos 4 0 , 1 2 1 Pinus 331-2 pisolitic laterites 240 PitzFormation, Canada 289 plaggen epipedon 70 plagioclase 175 plant formations, geological ranges 2 9 8 plants afforestation 280-99 angiosperms 135-6,292-5 grasses 30@16
402
Index
plants (continued) large plantsonland 263-79 origin of land plants 2 72-3 plant formations, defined 146-7 treestumps 136 vascular plants 1 35-6 seealsosperifictypes:vegetation plate tectonics 231 Pleistocene 326-7 hypersthene 192 loess 150 plinthite 72 Plio-Pleistocenegrassland mosaics, Kenya and Tanzania 167.326-7 pneumatophores 15 podzol 6 3 4 basket 135 podzolizdtion. formationof soil 60-1.136 PokegamaQuartrite 257 pollen 152,154,155,157 grasses 301 pollen diagram 333 polsterland 150-1,266-9,272 productivity,stature and biomass of vegetation 5 9 polypedon, defined 8 porosity,defined 38-9 potassium chemistry 175 metasomatism 234 in soils,rocks and organisms 2 5 0 potworms 137 prairie 150 prairie grasses, roots 1 6 Precambrian 1 7 6 , 1 7 9 , 2 2 7 4 5 Entisols 285-6 evolution of soils,atmosphere and continents 244 glaciation 243 ProntoandDenison paleosols 232 scenery 242-5 soils 227-31,242-5 calculating atmospheric oxidation 2 3 7-9 Stacaand Sheigrapaleosols 243 unconformities 10-11 Precambrian atmosphere greenhouseconditions 241 oxygen abundance 238 precipitation indicators 112-1 8 base/aluminaratio 117 chemical composition 115-16 clay minerals 116-1 7 deptbtocalcichorizon 114-15 desertpavement 118 evaporite minerals 1 17-1 8 histosols, peats and coals 112 karst 112-14 pedocalsand pedalfers 114 Proconsul 3 1 8 , 3 2 1 Pronto paleosols 2 33,234.2 58 Proterozoic 1 7 6 . 2 2 7 4 5
Prototaxites 264.275 Psamments 292 PsammosteusLimestone 2 78 Pseudobornea 280 pumice 175,179 pyrite 153,164,165,229 oxidation 46 pyroxene 179 pyroxenite 179 quart2 caries texture 1 65 structure 175, 17 6 Quartzipsamments 148 Quaternary paleosols 9-10.15 1 Quaternary units, and sedimentary formations, California 83 Quercus 331-2 Quinag 162 rainforest 148 rccyclingof rocks 231 redsoils 153,165,185.229 seealso Oxisol reddening. iron oxides and hydroxides 90-1, 165 reduced(1owEh)soils 45-7,156 regolith, defined 209 relict paleosols 9 Rendolls 178 reptiles 145.292 respiration 248 rhieoconcretions 17,149.150 rhizolite 1 7 RhynieChert 269 rhyniophytes 269-70 RNA 246-8,251 rockglaciers 4 0 , 1 2 1 rock record. timescales 198-9 rock types of coal measures 5 2 rockvarnish 132,258 Romanagriculture 332 root systems 15-1 7 root traces 13-20 calcite fibers 1 33 carbonaceous 153 drab-haloed 17-20,90.149 indicatorsof pastwatertable 1 6 4 . 1 6 5 patterns 15-1 7 rhizoconcretions 1 7 seasonality indicators 122-3 types 14-15 Rosemarypaleosol 2 8 2 , 2 8 3 , 2 8 4 rubidium, isotopiccomposition 88-9 Russia, AIdan Shield 240 sabhkas 163 salic horizon 72 Salids 244 salinization chemical weathering 45.49-50.62 formationof soil 49-50.62
salt crusts 1 6 3 4 saltmarsh 152-3.269-72 saltyclaysoils 244 Samnaclay 1 8 0 sandwedges 40 sandstones 1 7 8 chemistry 1 7 5 saprolite 180 defined 7 horizons 1 7 1 SarangSeries 1 6 9 , 1 7 0 saturation conductivity, salinization 49 savanna 149, 1 5 7 scaleinsects 1 4 1 schist 1 7 8 scoria 179 scorpions 140 Scotland 166.176 pre-Torridonian paleosols 2 3 2 , 2 3 5,236. 23 7 , 2 4 3 , 2 5 9 RhynieChert 269 unconformities 162 Scoyenia 266-8.291 seasonalityindicators 12&5 carbon isotopic composition 123-5 charcoal 1 2 3 climatic inferences from periglacial features 121 concretionsand argillans 120-2 mukaraandgilgai 1 2 0 patternsof roottraces 122-3 sedimentary sequences 11-12 disconformities 161-2 marine 177-8 Quaternary units 83 Selaginella 269 selfishsoil,early life 2 54-6 sequences 9 stratigraphy, model 203 time scales. Badlands National Park 194-203 sericite 176 serpentinite 1 7 9 shales 178 abundanceof clay minerals 95 illitization 95-7 Sheigra, Scotland,pre-Torridonian 232, 235.236.237.243.259 shrubland 1 5 0 fire-prone 150 wooded 149 sidereal periods 210 siderite 47-8,153,164.165.238.253 silcretes 240-1 silica, chemistry I 7 5 silica cement 9 2 silicification 60 silt, stability series 43 Silurian 272.2 76-9 Sivapithecus 3 17 SiwalikGroup 1 6 7 , 1 6 9 , 2 0 3 slickensides 288
Index Slioch 162 smectite 175.178.278 illitination 3 9 , 9 5 7 sodium chemistry 175 in soils, rocks and organisms 2 5 0 soil, textural classes 2 3 soil acidity see acidification soil analysis 24 soil climate seeclimate soil creep 16@1 soildensity 37 moisture equivalents and porosity 38 soil facies 82 soil horizons 19-24.70b2 A 188 boundaries 2 4 C(sapro1ite) 171 calcic (Bk) horizon development climate 114-15,165 timescales 185-8 clayey subsurface (Bt)horizon development 171.188-9 cumulic horizons 164 density 38 depths 21 describing 214.67-76 descriptiveshorthand 2 2 diagnostic horizons and properties 65-72 epipedons 67,70-1 field scale of acid reactions to approximate carbonate content 24 indicatorsof past water table 164-5 master and subordinate 22 0 152 past water table 164-5 petrocalcic 163 types 20-1 soil invertebrates, effect on afforestation 29 1 soil mapping and naming paleosols 77-8 6 soil nodules. indicators of water table 1 65-6 soil sediment, defined 9 soil structure 24-6 aeration 29 indicatorsof pastwatertable 165 structuralelements 2 5 soils and plant formations 146-7 solarenergy 249-50 variationonEarth 111-12 SolarSystem 209-23 comparisonof planetary bodies 2 1 0 comparison of soil formation 2 2 5 origins 225 solodization 60 solonization 60 solum. defined 7.8 solum thickness, chronofunctions 190 sombrichorixon 71 SonitaSeries 1 6 9 , 1 7 0 South Africa 209,237,240,247 Barberton 209,240 Black Reef Quartzite 258-9
karst 241 mukkara 1 2 4 PongolaSupergroup 258 Schagenpaleosol 2 5 9 WatervalOnderpaleosol 257-8,260. 288-9 South Dakota seeBadlandsNational Park spalling. fire heating 4 0 speciation 314-15 sphaerosiderite 47-8 Sphagnumhog 152.191.331-2 spiders and mites 140 spodichorizon 71 Spodosol(sandyforestsoi1)75.147-8.178. 208,284-5 springtails, wingless insects 140-1 Stacapedotype 235,237,243 steadystate 185 steppe 149,150 stone lines 173 burrows 163 stone spalling,fire heating 40 strain, andmass transportfunction 181.182 stratigraphy geosolsformapping 84-6 model 2 0 3 nomenclature (NACSN) 77 paleosols, asmarker horizons 83-6 Quaternary units and sedimentary formations, California 83 timescales 199-203 weathering profiles. Queensland. Australia 85 strength, induration scale 173 string bogs 40 stromatolites 2 2 8 lichen 134 ministromatolites 132 strontium, isotopic composition 88-9 succession, fire 150 successional vegetation 147 sulfur 229-31 isotope ratios 2 3 0 , 2 4 7 sulfurbacteria 131 sulfuric horizon 72 swamp 153 peat accumulation,time scales 189-9 1 Sweden 1 9 1 taiga 152 Tanzania OlduvaiGorge 2 0 5 , 3 2 7 Laetoli. Pliocenegrassland mosaics 325-6 tardigrades 138.264 temperature indicators oxygenisotopiccomposition 119 permafroststructures 119-20 spherical micropedsand termite nests 118-19 termite nests 143 temperatureindicators 118-19
403
termites 143 Termitichnus 29 1 TerraRossasoils 174 Tetrahedruletes 2 6 4 5 thecamoebans 133 Thelon Basin, Canada 240 therapsids 145 thermal maturation indices 99 coals, peats 9 9 tholeites 179 till 160 defined 176-7 tilth and bioturbation 56-7 microbialaction 132 timescales 183-203,184 accumulation of paleosolsequences 194-203 alluvialarchitecture 196-8 completeness of rockrecord 198-9 hillslopedevelopment 194-6 stratigraphy 199-203 calcic (Bk)horizon development, soil horizons 185-8 formation of soil 9 grasslands 308-12 marsh 272 paleosol development 185-94 clayey subsurface (Bt) horizon development 188-9 color scores 194 combinationsof features 1 9 3 4 mineralweathering 191-3 peat accumulation 189-9 1 stages 187 stages of carbonate accumulation in soils 187 weathering of mineral grains 1 9 2 weatheringrinds 193 preservationof fossils 155 titania 1 8 1 toad. spadefoot 144-5 topographic relief indicators 160-70 geomorphologicsetting 3 9 , 1 6 3 4 paleocatenae 166-70 water table 164-6 see also landscapes Torridonian pre- 2 3 2 unconformities 161-2 Torridonian sequences 166 tourmaline 1 8 1 trace elements 260 tracefossils 129-30.257-8 traces of ecosystems 145-53 tracheophyte-trigonotarbidbrakelands 269 treeferns 15 trichopterans 141 trigonotarbids 269 tropicaldeeply weathered soil seeOxisol tundra 1 2 1 , 1 5 2
T Ulmus 331-2 Ultisol(base-poorforestsoil)75-6.147.148. 177-8,208,282-5 profile,grain-size and molecular weathering ratios 44 ultramaftcrocks 179 ultravioletradiation 250,274 unconformities 1%: 1 erosionalgullies 162,198 past geomorphologicalsettings 1 C 1 1 , 161-2 uranium 229 USsoil taxonomy 64-7.68-9 Ustochrepts 158 Ustollic Eutrandepts 180 Ustropepts 3 2 1 Utah, Dinosaur National Monument 293 vegetation climate and profile, US soil taxonomy 69 Fncene-Oligocene,Badlands National Park, SouthDakota 81 Holdridge’sclassification 111 productivity, stature and biomass 59 succession 147 volcanism, colonization 1 2 9 see also plants velvetworms 137,264 Venusian soils 2 13-1 6 comparison of planetary bodies 2 10 soil composition 2 14-1 6 soildevelopment 2 16
Vertisol(swellingclaysoil) 74.178.179.208, 288-9 compaction 94 vertization 60 vesicles 25 vitrinitemacerals 51, 53 vitrinite reflectance 98 vivianite 46 volcanicash 179 volcanicglass 175,179,181 volcanic hot springs 2 5 0 4 volcanic rocks, zeolitization 9 7 volcanic sequences,paleosols 11-12 volcanism, Krakatau 1 2 9 volume, formulae relating thickness and chemicalchanges 4 6 volume change, formationof soil 181 vugs 25
Walkley-Black titration. organic matter 89 wallows 145 Walther’s FaciesLaw 167 wasps 144 waterbears 138 watertable indicators of past, soil horizons 164-5 topographicrelief indicators 164-6 water table indicators 164-6 weather,defined 108 weathering biological weathering 50-7 chemicalweathering 41-50 deep profiles 8 6
andgrainsize 192 hydrolytic 22 7 , 2 2 8 indications 171 physical weathering 3 7 4 0 stratigraphic relationship,Queensland. Australia 8 5 weathering index 43 weathering profites, CuraUe silcrete, Queensland.Australia 85 weathering ratios 42-3.94-5.48-9 weatheringrinds. timescales 193 weeds 332-3 WhiteRiverGroup 196 Willamette River, Oregon, soil series and vegetation 79 wind-throws 136 Wisconsinan 1 6 woodlands dry 149 Pakistan,N.Miocene 3 2 3 4 plant formations 147 productivity, stature and biomass of vegetation 59 wooded grassland and shrubland 147-9 Wyoming, Green River Formation 163 YellowMoundspedotype 195,196 young soil seeInceptisol zeolitiation. volcanic rocks, paleosol alteration after burial 9 7 zircon,age 209 zirconium 180
