Sediments, Diagenesis, andSedimentary RocksCitationsPlease use the following example for citations:Li Y.-H. and Schoonmaker J.E. (2003) Chemical composition and mineralogy of marine sediments, pp. 1-35. In Sediments,Diagenesis, and Sedimentary Rocks (ed. F.T. Mackenzie) Vol. 7 Treatise on Geochemistry (eds. H.D. Holland andK.K. Turekian), Elsevier -Pergamon, Oxford.Cover photo: Paleozoic section overlying Precambrian at the Grand Canyon: looking eastward toward Hermit`s Rest,Grand Canyon, USA. (Photograph provided by William Graustein)Sediments, Diagenesis, andSedimentary RocksEdited byF. T. MackenzieUniversity of Hawaii, HI, USATREATISE ON GEOCHEMISTRYVolume 7Executive EditorsH. D. HollandHarvard University, Cambridge, MA, USAandK. K. TurekianYale University, New Haven, CT, USA2005AMSTERDAM - BOSTON - HEIDELBERG - LONDON - NEW YORK - OXFORDPARIS - SAN DIEGO - SAN FRANCISCO - SINGAPORE - SYDNEY - TOKYOELSEVIER B.V.Radarweg 29P.O. Box 211, 1000 AE AmsterdamThe NetherlandsELSEVIER Inc.525 B Street, Suite 1900San Diego, CA 92101-4495USAELSEVIER LtdThe Boulevard, Langford LaneKidlington, Oxford OX5 1GBUKELSEVIER Ltd84 Theobalds RoadLondon WC1X 8RRUK 2005 Elsevier Ltd. All rights reserved.This work is protected under copyright by Elsevier Ltd., and the following terms and conditions apply to its use:PhotocopyingSingle photocopies of single chapters may be made for personal use as allowed by national copyright laws. Permission of thePublisher and payment of a fee is required for all other photocopying, including multiple or systematic copying, copying foradvertising or promotional purposes, resale, and all forms of document delivery. Special rates are available for educationalinstitutions that wish to make photocopies for non-proft educational classroom use.Permissions may be sought directly from Elsevier`s Rights Department in Oxford, UK: phone (+44) 1865 843830, fax (+44)1865 853333, e-mail: permissions@elsevier.com. Requests may also be completed on-line via the Elsevier homepage(http://www.elsevier.com/locate/permissions).In the USA, users may clear permissions and make payments through the Copyright Clearance Center, Inc., 222 RosewoodDrive, Danvers, MA 01923, USA; phone: (+1) (978) 7508400, fax: (+1) (978) 7504744, and in the UK through theCopyright Licensing Agency Rapid Clearance Service (CLARCS), 90 Tottenham Court Road, London W1P 0LP, UK;phone: (+44) 20 7631 5555; fax: (+44) 20 7631 5500. Other countries may have a local reprographic rights agency forpayments.Derivative WorksTables of contents may be reproduced for internal circulation, but permission of the Publisher is required for external resaleor distribution of such material. Permission of the Publisher is required for all other derivative works, including compilationsand translations.Electronic Storage or UsagePermission of the Publisher is required to store or use electronically any material contained in this work, including anychapter or part of a chapter.Except as outlined above, no part of this work may be reproduced, stored in a retrieval system or transmitted in any form orby any means, electronic, mechanical, photocopying, recording or otherwise, without prior written permission of thePublisher.Address permissions requests to: Elsevier`s Rights Department, at the fax and e-mail addresses noted above.NoticeNo responsibility is assumed by the Publisher for any injury and/or damage to persons or property as a matter of productsliability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained inthe material herein. Because of rapid advances in the medical sciences, in particular, independent verifcation of diagnosesand drug dosages should be made.First edition 2005Library of Congress Cataloging in Publication DataA catalog record is available from the Library of Congress.British Library Cataloguing in Publication DataA catalogue record is available from the British Library.ISBN: 0-08-044849-6 (Paperback)The following chapters are US Government works in the public domain and not subject to copyright:Sulfur-rich SedimentsCoal Formation and Geochemistry

W The paper used in this publication meets the requirements of ANSI/NISO Z39.48-1992 (Permanence of Paper).Printed in ItalyDEDICATEDTOROBERT GARRELS(1916~1988)Photograph provided by Cynthia GarrelsThis Page Intentionally Left BlankContentsExecutive Editors` Foreword ixContributors to Volume 7 xiiiVolume Editor`s Introduction xv7.01 Chemical Composition and Mineralogy of Marine Sediments 1Y.-H. LI and J. E. SCHOONMAKER7.02 The Recycling of Biogenic Material at the Seafoor 37W. R. MARTIN and F. L. SAYLES7.03 Formation and Diagenesis of Carbonate Sediments 67J. W. MORSE7.04 The Diagenesis of Biogenic Silica: Chemical Transformations Occurring inthe Water Column, Seabed, and Crust 87D. J. DEMASTER7.05 Formation and Geochemistry of Precambrian Cherts 99E. C. PERRY, Jr. and L. LEFTICARIU7.06 Geochemistry of Fine-grained Sediments and Sedimentary Rocks 115B. B. SAGEMAN and T. W. LYONS7.07 Late Diagenesis and Mass Transfer in Sandstone-Shale Sequences 159K. L. MILLIKEN7.08 Coal Formation and Geochemistry 191W. H. OREM and R. B. FINKELMAN7.09 Formation and Geochemistry of Oil and Gas 223R. P. PHILP7.10 Sulfur-rich Sediments 257M. B. GOLDHABER7.11 Manganiferous Sediments, Rocks, and Ores 289J. B. MAYNARD7.12 Green Clay Minerals 309B. VELDE7.13 Chronometry of Sediments and Sedimentary Rocks 325W. B. N. BERRY7.14 The Geochemistry of Mass Extinction 351L. R. KUMP7.15 Evolution of Sedimentary Rocks 369J. VEIZER and F. T. MACKENZIESubject Index 409viiThis Page Intentionally Left BlankExecutive Editors' ForewordH. D. HollandHarvard University, Cambridge, MA, USA&h?K. K. TurekianYale University, New Haven, CT, USAGeochemistry has deep roots. Its beginnings canbe traced back to antiquity, but many of thediscoveries that are basic to the science were madebetween 1800 and 1910. The periodic table ofelements was assembled, radioactivity was dis-covered, and the thermodynamics of hetero-geneous systems was developed. The solarspectrum was used to determine the compositionof the Sun. This information, together withchemical analyses of meteorites, provided anentry to a larger view of the universe.During the frst half of the twentieth century,a large number of scientists used a variety ofmethods to determine the major-element com-position of the Earth`s crust, and the geochemis-tries of many of the minor elements were defnedby V. M. Goldschmidt and his associates usingthe then new technique of emission spectrography.V. I. Vernadsky founded biogeochemistry. Thecrystal structures of most minerals were deter-mined by X-ray diffraction techniques. Isotopegeochemistry was born, and age determinationsbased on radiometric techniques began to defnethe absolute geologic timescale. The intensescientifc efforts during World War II yieldednew analytical tools and a group of people whotrained a new generation of geochemists at anumber of universities. But the feld grew slowly.In the 1950s, a few journals were able to report allof the important developments in trace-elementgeochemistry, isotopic geochronometry, theexploration of paleoclimatology and biogeochem-istry with light stable isotopes, and studies ofphase equilibria. At the meetings of the AmericanGeophysical Union, geochemical sessions werefew, none were concurrent, and they all rangedacross the entire feld.Since then the developments in instrumentationand the increases in computing power have beenspectacular. The education of geochemists hasbeen broadened beyond the old, rather narrowlydefned areas. Atmospheric and marine geochem-istry have become integrated into solid Earthgeochemistry; cosmochemistry and biogeochem-istry have contributed greatly to our understandingof the history of our planet. The study of Earth hasevolved into Earth System Science, whoseprogress since the 1940s has been truly dramatic.Major ocean expeditions have shown how andhow fast the oceans mix; they have demonstratedthe connections between the biologic pump,marine biology, physical oceanography, andmarine sedimentation. The discovery of hydro-thermal vents has shown how oceanography isrelated to economic geology. It has revealedformerly unknown oceanic biotas, and has clari-fed the factors that today control, and in the pasthave controlled the composition of seawater.Seafoor spreading, continental drift and platetectonics have permeated geochemistry. We fnallyunderstand the fate of sediments and oceanic crustin subduction zones, their burial and theirixexhumation. New experimental techniques attemperatures and pressures of the deep Earthinterior have clarifed the three-dimensional struc-ture of the mantle and the generation of magmas.Moon rocks, the treasure trove of photographsof the planets and their moons, and the successfulsearch for planets in other solar systems have allrevolutionized our understanding of Earth and theuniverse in which we are embedded.Geochemistry has also been propelled into thearena of local, regional, and global anthropogenicproblems. The discovery of the ozone hole cameas a great, unpleasant surprise, an object lessonfor optimists and a source of major new insightsinto the photochemistry and dynamics of theatmosphere. The rise of the CO2 content of theatmosphere due to the burning of fossil fuels anddeforestation has been and will continue to be atthe center of the global change controversy, andwill yield new insights into the coupling ofatmospheric chemistry to the biosphere, thecrust, and the oceans.The rush of scientifc progress in geochemistrysince World War II has been matched by organ-izational innovations. The frst issue of Geochi-mica et Cosmochimica Acta appeared in June1950. The Geochemical Society was founded in1955 and adopted Geochimica et CosmochimicaActa as its offcial publication in 1957. TheInternational Association of Geochemistry andCosmochemistry was founded in 1966, and itsjournal, Applied Geochemistry, began publicationin 1986. Chemical Geology became the journal ofthe European Association for Geochemistry.The Goldschmidt Conferences were inauguratedin 1991 and have become large internationalmeetings. Geochemistry has become a major forcein the Geological Society of America and in theAmerican Geophysical Union. Needless to say,medals and other awards now recognize outstand-ing achievements in geochemistry in a numberof scientifc societies.During the phenomenal growth of the sciencesince the end of World War II an admirable num-ber of books on various aspects of geochemistrywere published. Of these only three attempted tocover the whole feld. The excellent Geochemistryby K. Rankama and Th.G. Sahama was publishedin 1950. V. M. Goldschmidt`s book with the sametitle was started by the author in the 1940s. Sadly,his health suffered during the German occupationof his native Norway, and he died in Englandbefore the book was completed. Alex Muir andseveral of Goldschmidt`s friends wrote the miss-ing chapters of this classic volume, which wasfnally published in 1954.Between 1969 and 1978 K. H. Wedepohltogether with a board of editors (C. W. Correns,D. M. Shaw, K. K. Turekian and J. Zeman) anda large number of individual authors assembledthe Handbook of Geochemistry. This and the othertwo major works on geochemistry begin withintegrating chapters followed by chapters devotedto the geochemistry of one or a small group ofelements. All three are now out of date, becausemajor innovations in instrumentation and theexpansion of the number of practitioners in thefeld have produced valuable sets of high-qualitydata, which have led to many new insights intofundamental geochemical problems.At the Goldschmidt Conference at Harvard in1999, Elsevier proposed to the Executive Editorsthat it was time to prepare a new, reasonably com-prehensive, integrated summary of geochemistry.We decided to approach our task somewhat dif-ferently from our predecessors. We divided geo-chemistry into nine parts. As shown below, eachpart was assigned a volume, and a distinguishededitor was chosen for each volume. A tenthvolume was reserved for a comprehensive index:(i) Meteorites, Comets, and Planets: AndrewM. Davis(ii) Geochemistry of the Mantle and Core:Richard Carlson(iii) The Earth's Crust: Roberta L. Rudnick(iv) Atmospheric Geochemistry: RalphF. Keeling(v) Freshwater Geochemistry, Weathering,and Soils: James I. Drever(vi) The Oceans and Marine Geochemistry:Harry Elderfeld(vii) Sediments, Diagenesis, and SedimentaryRocks: Fred T. Mackenzie(viii) Biogeochemistry: WilliamH. Schlesinger(ix) Environmental Geochemistry: BarbaraSherwood Lollar(x) IndexesThe editor of each volume was asked toassemble a group of authors to write a series ofchapters that together summarize the part of thefeld covered by the volume. The volume editorsand chapter authors joined the team enthusiasti-cally. Altogether there are 155 chapters and 9introductory essays in the Treatise. Naming thework proved to be somewhat problematic. It isclearly not meant to be an encyclopedia. The titlesComprehensive Geochemistry and Handbook ofGeochemistry were fnally abandoned in favor ofTreatise on Geochemistry.The major features of the Treatise were shapedat a meeting in Edinburgh during a conference onEarth System Processes sponsored by the Geo-logical Society of America and the GeologicalSociety of London in June 2001. The fact that theTreatise is being published in 2003 is due to agreat deal of hard work on the part of the editors,the authors, Mabel Peterson (the ManagingEditor), Angela Greenwell (the former Head ofMajor Reference Works), Diana Calvert (Devel-opmental Editor, Major Reference Works),Executive Editors' Foreword xBob Donaldson (Developmental Manager),Jerome Michalczyk and Rob Webb (ProductionEditors), and Friso Veenstra (Senior PublishingEditor). We extend our warm thanks to all ofthem. May their efforts be rewarded by adistinguished journey for the Treatise.Finally, we would like to express our thanks toJ. Laurence Kulp, our advisor as graduate studentsat Columbia University. He introduced us to theexcitement of doing science and convinced usthat all of the sciences are really subdivisions ofgeochemistry.Executive Editors' Foreword xiThis Page Intentionally Left BlankContributors to VoIume 7W. B. N. BerryUniversity of California, Berkeley, CA, USAD. J. DeMasterNorth Carolina State University, Raleigh, NC, USAR. B. FinkelmanUS Geological Survey, Reston, VA, USAM. B. GoldhaberUS Geological Survey, Denver, CO, USAL. R. KumpThe Pennsylvania State University, PA, USAL. LefticariuNorthern Illinois University, DeKalb, IL, USAY.-H. LiUniversity of Hawaii, Honolulu, HI, USAT. W. LyonsUniversity of Missouri, Columbia, MO, USAF. T. MackenzieUniversity of Hawaii, Honolulu, HI, USAW. R. MartinWoods Hole Oceanographic Institution, MA, USAJ. B. MaynardUniversity of Cincinnati, OH, USAK. L. MillikenThe University of Texas at Austin, TX, USAJ. W. MorseTexas A&M University, College Station, TX, USAW. H. OremUS Geological Survey, Reston, VA, USAE. C. Perry Jr.Northern Illinois University, DeKalb, IL, USAR. P. PhilpUniversity of Oklahoma, Norman, OK, USAxiiiB. B. SagemanNorthwestern University, Evanston, IL, USAF. L. SaylesWoods Hole Oceanographic Institution, MA, USAJ. E. SchoonmakerUniversity of Hawaii, Honolulu, HI, USAJ. VeizerRuhr University, Bochum, Germany and University of Ottawa, ON, CanadaB. VeldeEcole Normale Superieure, Paris, FranceContributors to Volume 7 xivVolume Editor's ntroductionF. T. MackenzieUniversity of Hawaii, HI, USAThis volume is dedicated to Robert M. Garrelsin recognition of his contributions to our under-standing of the fundamental geochemical process-es governing the formation and transformationof sedimentary rocks through geologic time andtheir use in deducing the evolution of the Earth`sexogenic system. In retrospect, Bob`s researchcareer can probably be divided into two mainphases. The frst dealt with theoretical andexperimental aspects of geochemistry, which ledto the publication of his book Mineral Equilibriaat Low Temperature and Pressure in 1960 andlater to the book`s successor Solutions, Minerals,and Equilibria co-authored with Charles Christ in1965. These books have become classics in thefeld of aqueous and sedimentary geochemistry, asthey were one of the frst efforts to apply chemicalthermodynamics to geological processes with anemphasis on the construction of stability diagramsfrom thermodynamic data. The second phase ofBob`s career led him to consider the importanceof the chemical cycling of sediments and theprocesses involved: weathering, erosion, trans-port, sedimentation, burial, diagenesis, and upliftrestarting the sedimentary cycle. This recyclingconcept had lain dormant for 150 years before thepublication of Evolution of Sedimentary Rockswith Fred Mackenzie in 1971. The papers in thisvolume largely address this phase of Bob`s career.Biogeochemical processes and the productsof those processes, the sedimentary rocks, arediscussed in this volume as is the use of variousgeochemical and isotopic proxies found insedimentary rocks to constrain the chemicalevolution of the Earth surface environment. Notall sedimentary rock facies are given a separatechapter because of space constraints, but thegeochemistry of most major facies is considered,particularly in terms of the evolution of sedimen-tary rocks and their usefulness in decipheringEarth history. The volume begins with fourchapters (Chapters 7.01, 7.02, 7.03, and 7.04)dedicated primarily to the geochemistry of major,modern marine, sediment types and their diagen-esis. Sediments at their origin have measurableproperties such as mineralogy, chemistry, isotopiccomposition of organic and inorganic phases,biological makeup, and texture. The diageneticprocesses of compaction, cementation, and solu-tion and reprecipitation leading to mineralstabilization can modify all of these properties.Thus, as sediments pass through the diageneticfence (Chave, 1960), they lose informationnecessary for environmental interpretation.Chapter 7.05 continues the theme of Chapter 7.04on the diagenesis of silica and discussesthe product of that diagenesis, chert, in the contextof Earth`s early temperature history. This isfollowed by chapters dealing with geochemistryand mass transfer in siliciclastic rocks and thegeochemistry of the organic-rich materials in coal,oil, and gas and the important element sulfur.Two minor but environmentally important sedi-ment components are then considered in the chap-ters on manganiferous sediments and the greenclay minerals. The fnal three chapters (Chapters7.13-7.15) are devoted to the chronological andhistorical aspects of sediments and sedimentaryrocks and the use of biogeochemical and isotopicproxies within them in interpreting the history ofEarth surface environments.In Chapter 7.01 Telu Li and Jane Schoonmakerreview the chemical composition of modernmarine sediments, building on the classicalworks of Tizzard et al. (1885), Murray andMenard (1891) and of more recent authorsxv(e.g., Arrhenius, 1963). Li and Schoonmakeraddress the mineralogy and geochemistry ofpelagic sediments in general and consider severalspecifc sediment types, particularly in terms oftheir hydrogenous components. Sediment elemen-tal enrichment factors and factor analysis are usedto determine the major components and miner-alogical phases in which the elements of marinesediments reside. The major components ofmarine pelagic sediments are shown to be(i) aluminosilicates, derived mainly from theweathering of shales and eolian and riverinetransport of the weathered products to the ocean;and (ii) ferromanganese oxides, carbonate fuor-apatite, zeolites, biogenic carbonate, silica, andbarite, derived from the products of organicproductivity and chemical reactions involvingelements directly derived from the general back-ground of elements in seawater or resulting fromdiagenetic reactions in the oxic, suboxic, andanaerobic realms of early diagenesis (Chester,1990).In Chapter 7.02, Bill Martin and Fred Saylesconsider the early diagenetic reactions occurringin sediments, including those involving organicmatter degradation, calcium carbonate dissolu-tion, and silica cycling, and the role of bioturba-tion in mixing particles downward from thesediment -water interface. The authors primarilyuse the wealth of solute concentration data frompore-water vertical profles that they themselves(e.g., Sayles, 1979, 1981; Martin et al., 2000), aswell as other investigators (e.g., Archer et al.,1989; Bender et al., 1989), have obtained duringthe past several decades to interpret diageneticprocesses. Mass balances provided by the authorsshow that only a small fraction of the organicmatter and the opal that falls to the seafoorsurvives early diagenesis; thus, only a smallportion accumulates in the sediment and isavailable for downcore interpretations of accumu-lation rates. The incorporation of aluminum intoopal and the hypothesized reprecipitation ofH4SiO4 from dissolving opal into neoformedminerals highlight the signifcant infuence of thesilica cycle on the cycling behavior of the majorconstituents of pore water and the likelihoodthat reverse weathering (e.g., Mackenzie andGarrels, 1966; Michalopoulos and Aller, 1995;Mackenzie and Kump, 1995) is an importantprocess affecting ocean chemistry. For calciumcarbonate, the role of dissolution in sedimentsabove the calcite lysocline driven by acidsgenerated by oxic respiration is highlighted,which also brings into question present interpret-ations of the temporal record of CaCO3 accumu-lation rates. About half of the O2 consumption inthe ocean below 1,000 m occurs in deep-seasediments, and sedimentary denitrifcation isshown by Martin and Sayles to be a dominantterm in the marine fxed nitrogen budget. As anaside, these authors point out that pore-waterstudies probably began with Murray and Irvine`s(1885) work.In Chapter 7.03, John Morse discusses thesources and early diagenesis of deep-sea andshoal-water (shallow-water) carbonate sedi-ments. One of the long-standing controversialareas of carbonate geochemistry has been therelationship between calcium carbonate accumu-lation in deep-sea sediments and the saturationstate of the overlying waters. Hypotheses consid-ered have ranged from a nearly thermodynamicocean where the calcite and aragonite compen-sation depths are at the calcite and aragonitesaturation horizons (e.g., Li et al., 1969 ) to astrongly kinetically controlled ocean systemwhere major differences exist between the car-bonate compensation depths, lysoclines, andsaturation depths (e.g., Morse and Berner, 1972 ).Morse points out that if recent calculations relatedto these features are correct, the long-cherishedidea of a tight coupling between seawaterchemistry and carbonate depositional faciesrequires re-examination. That such a couplingmay not exist has been suggested by Millimanet al. (1999), who evaluated several lines ofevidence and concluded that considerable dissol-ution of calcium carbonate, perhaps as much as60-80%, occurs in the upper 500-1,000 m of theocean, well above the chemical lysocline. The roleof organic matter in this water-column dissolutionprocess and in microbially driven organic oxi-dation processes in both shoal-water and deep-seasediments (as discussed by Morse) is now wellrecognized. Morse also considers what is knownabout early dolomite formation. He points out thatmost modern dolomite is forming from relativelyhigh-ionic-strength seawater solutions in whichthe Mg2| to Ca2| solution ratios are considerablyhigher than in normal seawater. However, he alsorecognizes that factors other than the Mg2|/Ca2|ratio may be involved in dolomite formation.Sulfate depletion and alkalinity increase (increasein saturation state) in pore waters, resulting fromanaerobic oxidation of organic matter by bacteriausing SO42ras a substrate, as well as directparticipation by bacteria, may play a role in thedolomite formation process. Dolomite is not foundin abundance in modern sediments but is animportant component of older sedimentary rocks,and its origin has important implications forinterpreting the history of seawater and theatmosphere through geologic time (see Chapters6.21 and 7.15).Except for the sponges of the latest Precam-brian Ediacaran fauna, most of the organisms thatmake their skeletons of amorphous silica are ofCambrian age or younger. The appearance ofradiolarians and diatoms in the PhanerozoicVolume Editor's Introduction xvichanged the manner of deposition of amorphoussilica on the seafoor from a primarily abioticprocess to one in which biochemical processesplay a strong role, and led to a shift in the locus ofsilica deposition from the shallow shelf to the deepsea. Prior to the evolution and spread of siliceousbiota in the Cambrian, the oceans containedsignifcant concentrations of dissolved silica(Garrels and Perry, Jr., 1974; Siever, 1992),reaching saturation levels with respect to amor-phous silica in Precambrian aquatic systems. Assiliceous organisms became more prevalent in thePhanerozoic, the surface waters of the oceanbecame depleted in dissolved silica, and theaverage concentration of dissolved silica inseawater fell from a level of about 1,000 6M to100 6M. In the context of these signifcant eventsin biotic evolution and seawater chemistry, DaveDeMaster in Chapter 7.04 and Ed Perry andLiliana Lefticariu in Chapter 7.05 discuss thediagenesis of biogenic silica (chemical transform-ations occurring in the water column, seabed, andcrust), and the formation and geochemistry ofPrecambrian cherts, respectively. DeMasterexamines the chemical and structural transform-ations that take place following the death ofsiliceous organisms in the oceans and that lead tochanges in the solubility and dissolution kineticsof the original biogenic amorphous silica phase.He documents the stabilization pathways of thetransformation of biogenic silica from amorphousopal-A to opal-CT and fnally in many cases tomicrocrystalline quartz (the rock chert) as thebiogenic silica is buried beneath meters to kilo-meters of sediment and is exposed to highertemperatures and pressures (e.g., Williams andCrerar, 1985). DeMaster also recognizes theimportance of the tie between the aluminum andsilicon cycles and reverse weathering diageneticprocesses as a sink for elements in the ocean.Perry and Lefticariu highlight the fact thatcherts precipitated from aqueous solution havethe potential to preserve a record of the oxygenisotopic composition of contemporaneous sea-water or lake water and hence the temperatureof the environmental medium. The problems,of course, are that the oxygen isotopic signal ofcherts is a function of both temperature andsolution composition and that the siliceous preci-pitates undergo diagenesis and may encountertemperatures and pressures of metamorphismin their passage through the sedimentary cycleleading to isotopic exchange. Recognizing theseproblems, Perry and Lefticariu, nevertheless,suggest that Earth surface temperatures derivedfrom the %18O signatures of 1,900-2,400-Myr-oldcherts could have been as high as 64 ?C and thoseobtained from Archean, 3,400-3,700-Myr-oldcherts suggest temperatures as high as 90 ?C.Such evidence of high early Earth surfacetemperatures, although controversial, can beused in conjunction with other evidence to suggestthat one of the earliest progenitors of life was ananaerobic thermophilic microorganism evolvingin a hydrothermal or volcanic environment (e.g.,Kasting and Chang, 1992).Chapters 7.06 and 7.07 focus mainly on thegeochemistry and sedimentology of detrital sedi-ments and sedimentary rocks (siliciclastics),These sedimentary materials preserve records ofmultiple biogeochemical proxies that reveal theprocesses and conditions of sediment formation,transport, deposition, burial, and diagenesis andprovide information regarding the evolution ofEarth`s surface environment. The proxies includesedimentary mineralogical suites and their distri-bution in time and space, major-, minor-, andtrace-element concentration data, the stable iso-topic composition of carbon, oxygen, sulfur, andnitrogen, radiogenic isotopes such as 40K, and theabundance and isotopic composition of biomarkercompounds. Brad Sageman and Tim Lyons dealmainly with the fne-grained, mixed siliciclastic-biogenic sedimentary facies, commonly referredto as hemipelagic facies. Based on the results ofcase studies of modern Black Sea and Cariacobasin sediments, Cretaceous hemipelagic sedi-ments, Devonian black-to-gray shale transitionalzones, and Precambrian shales and argillites, theseauthors propose a unifed Earth-systems modelfor the biogeochemical analysis of hemipelagicrocks. Important linkages in the model includethose between detrital fuxes and bulk sedimen-tation, productivity and nutrients, productivity and%13C records, molybdenum and iron, and controlson redox conditions, and elemental ratios andaccumulation rates. This work demonstrates anabundance of biogeochemical information andproxies is present in the often-overlookedsedimentary hemipelagic facies that can be usedto interpret Earth surface conditions throughgeologic time.Chapter 7.07 discusses what is known about thechemical and physical processes that lead to thetransformation of detrital siliciclastic sedimentsinto rocks. These processes take place duringburial diagenesis and low-grade metamophismand certainly involve basin-scale mass transfer ofhelium, water, and petroleum, and signifcanttransport of carbon out of shales into the ocean-atmosphere system. Large-scale silica transferfrom shales to sandstones is also a persistentfeature of burial diagenetic sequences, as theprocesses of smectite dissolution and illite pre-cipitation and quartz pressure dissolution in shalesand quartz precipitation in sandstones actin concert, and silica solubility increases withincreasing burial and temperature. Calcium is lostfrom shales, as is carbonate, due to the dissolu-tion of CaCO3 and, although controversial, theVolume Editor's Introduction xviiprobable loss of calcium from smectite as it istransformed to illite (Garrels and Mackenzie,1974). Potassium progressively increases insome shale sequences undergoing burial diagen-esis but decreases in others. However, the globaleffect appears to be the transport of potassiumfrom other lithologies into shales (Garrels andMackenzie, 1972). The fate of magnesium andsodium is not well constrained, although Garrelsand Mackenzie (1971, 1974) argue that bothelements are lost from shales with diagenesis andtime, and that some magnesium is transferred vialate dolomitization processes to the carbonatereservoir with increasing age of the rock mass(see Chapters 6.21 and 7.15). Milliken concludesher chapter with emphasis on the conclusion thatmost precipitation reactions occurring during thediagenesis of sandstones and shales are acid-releasing reverse weathering reactions and caninfuence the crustal CO2 cycle (e.g., Kastner,1974; Kerrick et al., 1995).Chapters 7.08-7.10 deal with the formationand fate of materials derived from living organicmatter, its degradation products, and elementsgenerally associated with organic-rich sedi-ments deposited in aquatic environments. InChapter 7.08 Bill Orem and Bob Finkelmanpoint out that because coal is an organic rocklargely composed of an assemblage of amorphous,degraded plant material, an admixture of synge-netic, diagenetic, epigenetic, and detrital mineralgrains, and contains within its structure water,oils, and gases, it is one of the most complex andchallenging sedimentary rocks to analyze andunderstand. Coal is the most abundant of the fossilfuels and the second-largest source of energy inthe world. In 2002 it supplied 830% of theworld`s commercial energy (Mackenzie, 2003). Inthe USA more than half of the electricitygenerated is by coal-fred power plants.Coal formation begins during the process ofpeatifcation in which microorganisms preferen-tially degrade plant biomolecules such as carbo-hydrates and leave other materials such as ligninand waxes selectively preserved. In the earlystages of coal formation (coalifcation), numerouschemical changes to biomolecules occur such asloss of oxygen functionality and changes in lignincomposition but the overall structure of plantmaterials of low-rank coals remains relativelyintact. Later-stage coalifcation producing bitumi-nous and higher-rank coals involves more sub-stantial changes in the structure of the coal. Theseinclude the transformation of coal to a hard, black,lustrous organic rock depleted in both oxygen andhydrogen relative to the precursor material andwith aromatic structures that are more condensed.In contrast to coal formation, in which the mostimportant starting materials are biopolymersfound in vascular plants with some admixture ofresistant algae polymers as well as bacterialpolymers, kerogens are the insoluble organicmaterials that are the immediate precursors tothe formation of most oil and gas. Kerogen is avery heterogeneous and complex aggregate ofmacerals, discrete particles of insoluble organicmaterial identifable under the microscope andrepresenting residual detritus from various sourcesof organic matter (e.g., Tissot et al., 1974). Thetypes and maturity levels of kerogen play animportant role in determining the type of oil orgas generated. In 1992, oil provided 40%, and gasprovided 20% of the commercial global energyconsumption (Mackenzie, 2003).In Chapter 7.09, Paul Philp describes the earlysteps of oil and gas formation, starting with theprocess of photosynthesis, ultimately responsiblefor most oil, gas, and coal formation. Today only80.1% of global net primary production oforganic matter accumulates in sediments (Tissotand Welte, 1978; Mackenzie et al., 1993 ) and isavailable for the formation of fossil fuels; thus,these commercial deposits of the stored energy ofphotosynthesis actually represent but a very smallportion of the gross production of organic matterand of total organic matter present in thesedimentary mass. Philp discusses the nature ofthe insoluble organic materials, kerogen, and thesource, depositional environment, maturity, bio-degradation, age dating, and migration of thesoluble components of kerogen. It is nowgenerally accepted that the major mechanism forthe formation of most oil (and gas) is throughthermal degradation of kerogen: kerogen =bitumen =oil | gas | residue.In conclusion, Philp argues that the integrationof geochemical parameters, such as biomarkers,with sequence stratigraphy models will becomemore important in the future in petroleum explor-ation in terms of predicting how the processes ofoil and gas generation, expulsion from sourcerocks, migration to reservoir rocks and subsequenttrapping, and preservation control the volume,quality, and distribution of petroleum.In Chapter 7.10, Marty Goldhaber considers indetail the geochemistry of sulfur, one of theimportant elements commonly associated withorganic-rich sediments and a principal elementfound in living organisms. He points out thatnearly all marine sediments with more than a fewtenths of a percent of organic carbon and thosenonmarine sediments with signifcant concen-trations of SO42rin depositional waters containthe mineral pyrite (FeS2). Pyrite, along withsulfur-bearing organic compounds and othermineral phases of sulfur such as mackinawite,greigite, and elemental sulfur, forms at Earth sur-face temperatures and pressures primarily throughthe metabolic activities of sulfate-reducing micro-organisms. Goldhaber discusses the biochemistryVolume Editor's Introduction xviiiof bacterial sulfate reduction along with theecology of the bacteria responsible for the overallprocess, emphasizing the fact that in general theultimate products of sulfate reduction are depletedin 34S. He then considers the forms, mechanisms,diagenesis, abundance, and isotopic systematics ofsulfur in marine sediments. In the global context,the evolution of the cycling behavior of sulfurthrough geologic time is closely tied to the carboncycle and the evolution of atmospheric oxygen(e.g., Petsch and Berner, 1998). In addition, thevariations in the sulfur isotopic composition ofsedimentary sulfate minerals and pyrite provideconstraints on the isotopic composition of sea-water through geologic time and suggest that bothatmospheric O2 and oceanic SO42rwere very lowprior to 2.4 Ga. The presence of very signifcantmass-independent sulfur fractionations in sulfateand sulfde minerals older than 2.45 Ga (Farquharet al., 2000) supports this conclusion.Chapters 7.11 and 7.12 are concerned with theformation of manganiferous phases in sediments,rocks, and ores and the sedimentary green clayminerals. Manganese is the tenth most abundantelement in the crust and is very similar to iron inits chemical properties; however, the fact that ithas higher valence states distinguishes its beha-vior from iron and gives rise to a plethora of Mn-oxide minerals. In Chapter 7.11 Barry Maynarddiscusses the basic geochemical properties ofmanganese and its common minerals, its distri-bution in rocks and natural waters, and thecomposition of its signifcant accumulations. Headdresses the behavior of manganese especially inmid-ocean ridge vent systems and during sedi-mentation. Throughout the chapter, Maynardemphasizes the difference between the geochem-istry of manganese and iron, and points out thatunder reducing and mildly oxidizing conditions,manganese is exported from low-oxygen environ-ments, such as basalt -hydrothermal systems oreuxinic basins, and it accumulates in oxidizingenvironments of the shallow ocean or in low-productivity areas of the deep sea. Thus the Mn/Feratios of sediments and waters provide a clue tothe oxidation structure of ocean basins, soils, andgroundwater systems. The strong affnity ofmanganese for certain transition elements andthe rare earth elements is an important property. Inthe case of cerium and europium, these elementswhen preserved in managnese accumulations candefne the relative importance of hydrothermaland diagenetic processes in the sedimentaryrecord.In Chapter 7.12, Bruce Velde demonstrates thatthe basic green color of glauconite, celadonite,berthierine, verdine, chamosite, nontronite, andtalc is due to the presence of iron in the structuresof these minerals. This is not a startling revelationto mineralogists, but it is the key to understandingthe origin and stability of the phases in nature.Velde concludes that green clay mineral namescan refect compositional differences, mineralstructural differences, or differences in the geo-logical occurrence of these phases. This in itselfleads to much confusion in the literature concern-ing the origin and diagenesis of the green clayminerals. Regardless, the mineralogy and chem-istry of the green clay minerals provide valuableclues to the temperature and solution compositionof the environment in which they formed. Forexample, both glauconite and berthierine-verdineminerals are low-temperature phases that form inshallow-water sediments at low sedimentationrates. The formation of glauconite probablyrequires the more oxidizing conditions oforganic-poor sediments rather than the lowerredox conditions for berthierine formation inorganic-rich sediments. The original glauconiteprecursor evolves toward a potassic, ferric claymineral, whereas that of berthierine evolvestoward an alkali-free, ferrous mineral. Thus, thetwo minerals can be used to infer their environ-mental conditions of deposition and diagenesis.Many of Velde`s conclusions are based on workstarted in the 1980s (e.g., Velde, 1985).The fnal three chapters (Chapter 7.13-7.15) ofthe volume concentrate on the longer-termhistorical aspects of sediments and sedimentaryrocks, their cycling behavior, and the biogeo-chemical and isotopic proxies found within themthat permit the interpretation of the evolution ofEarth`s surface environment through geologictime. Interpretations of this nature frst require atemporal framework in which to place informationand events, and this is the subject of Bill Berry`schapter (Chapter 7.13). Berry points out that it waseconomically imperative to develop a chronome-try of sedimentary rocks in order to fnd andrecover larger volumes of nature`s resources ofcoal, ores, building stone, sand, gravel, andeventually petroleum. The basic principle ofchronometry based on the fossil record isembodied in the Table of Strata that WilliamSmith dictated to his ecclesiastic associates inJune 1799 (see Berry, 1987) and still holds truetoday, that of the succession of fauna and forafound in the rock record. Berry`s description of thedevelopment of chronometry based on the fossilrecord is essentially that of the early developmentof the science of geology itself. Chorology, thescience that deals with the geographical distri-bution of living organisms, is discussed in termsof the constraints it imposes on chronometry.As Berry points out, chronometry cannot providethe absolute ages of rocks. Fossil-based chrono-metry has been enhanced through studies ofradiochronometry and magnetobiochronometry,as well as orbital chronometry. The chronometryof sedimentary rocks supplies the necessaryVolume Editor's Introduction xixframework for interpretation of events in Earthhistory and the longer-term secular evolution andcycling of the sediment-ocean-atmosphere-biotasystem, the subject matter of the next twochapters.In Chapter 7.14, Lee Kump points out that thecourse of biological evolution is inextricablylinked to that of the environment in a complicatednetwork of feedback at all time and space scales.Perturbations of the environment and disruptionswithin have biological consequences and viceversa. Kump considers mainly the elemental andcarbon, sulfur, strontium, and oxygen isotopicevidence for the environmental nature of the bigfve extinctions of the Phanerozoic (Sepkoski,1993): the Late Ordovician (8440 Ma), the LateDevonian (8367 Ma), the Permian-Triassic(8251 Ma), the Triassic-Jurassic (8200 Ma),and the Cretaceous -Tertiary (865 Ma). Hisconclusion is that there are no universal geochem-ical precursors or responses to extinction events inthe Phanerozoic. However, the development ofwidespread ocean anoxia appears to be associatedwith all three Paleozoic mass extinctions. Suchanoxic conditions can certainly lead to massmortality of extant organisms, and potentialrapid turnover of anoxic ocean waters couldproduce elevated concentrations of CO2 and H2Sin surface waters resulting in the death of aerobicorganisms. Rapid ocean turnover could be theresult of an oceanic impactor during the LateDevonian and Permian-Triassic events. How-ever, a purely terrestrial origin of the LateOrdovician extinction event is probable due toa glacio-eustatic sea-level fall causing a loss ofshallow-water marine habitats. In contrast, itappears that both the Triassic-Jurassic andCretaceous-Tertiary mass extinctions occurredduring periods of well-oxygenated oceans. Anasteroid or cometary impact appears to be theproximal cause of these events. At least for theK-T event, the consequent release of hugequantities of CO2 and SO2 to the atmospheremay have changed the climate and caused massmortality of organisms.In the fnal chapter (Chapter 7.15), Veizer andMackenzie show that the progressive increasewith decreasing age of the observed extantthickness and areal extent of Phanerozoic sedi-mentary strata and the secular variations in therelative proportions of lithological types and theirchemistry and mineralogy are a function of bothevolution and recycling of the sedimentary mass.They also conclude that the global sedimentarymass of 2.7 t 1024g is largely cannibalistic andhas a half-mass age of about 600 Myr, resulting intotal sediment deposition and destruction equi-valent to about fve sedimentary masses overthe course of geologic time (e.g., Garrels andMackenzie, 1971). The differential recycling ofvarious components of the sedimentary rock massis primarily controlled by the probability ofpreservation of different tectonic settings inwhich the sediments are found and the recyclingrates of the tectonic realms (e.g., Veizer, 1988 ).Because of differential recycling, the preservationof the sedimentary mass is inherently biased. Forexample, the half-life of the post-Devoniansedimentary mass is about 130 Ma, whereas thehalf-life of post-Permian carbonates is only886 Ma; thus, the carbonates are recycled at arate 81.5 times faster than the total sedimentarymass. This difference is due to the fact that duringpost-Permian time, the locus of carbonate depo-sition has progressively shifted to the deep sea,where post-Permian carbonates are more suscep-tible to destruction because of subduction. Theabsence of pre-Permian deep-sea carbonatesimplies that continental carbonates are the onlyrecord of carbonate deposition and chemical,mineralogical, and isotopic data for most ofgeologic time. Using the full range of miner-alogical, chemical, and strontium, osmium, sulfur,carbon, and oxygen isotopic evidence, Veizer andMackenzie also consider the chemical evolutionof the ocean-atmosphere system from the anoxicconditions of the early Precambrian to the modernoxic global environment (see also Chapter 6.21).REFERENCESArcher D., Emerson S., and Smith C. R. (1989) Dissolution ofcalcite in deep-sea sediments: pH and O2 microelectroderesults. Geochim. Cosmochim. Acta 53, 2831-2845.Arrhenius G. (1963) Pelagic sediments. In The Sea, Ideas andObservations on Progress in the Study of the Seas, The EarthBeneath the Sea, History (ed. M. N. Hill). IntersciencePublishers, New York, vol. 3, pp. 655-727.Bender M., Jahnke R., Weiss R., Martin W., Heggie D.,Orchardo J., and Sowers T. (1989) Organic carbon oxidationand benthic nitrogen and silica dynamics in San ClementeBasin, a continental borderland site. Geochim. Cosmochim.Acta 53, 685-697.Berry W. B. N. (1987). Growth of the Prehistoric Time Scale(revised edn.). Blackwell, Oxford.Chave K. E. (1960) Carbonate skeletons to limestones:problems. Trans. NY Acad. Sci. 23, 14-24.Chester R. (1990) Marine Geochemistry. Unwin Hyman,London, 698p.Farquhar J., Huiming B., and Thiemens M. H. (2000)Atmospheric infuence of Earth`s earliest sulfur cycle.Science 289, 756-758.Garrels F. T. and Mackenzie F. T. (1971) Evolution ofSedimentary Rocks. WW Norton, New York, 397p.Garrels R. M. and Mackenzie F. T. (1972) Aquantitative modelfor the sedimentary rock cycle. Mar. Chem. 1, 27-41.Garrels R. M. and Mackenzie F. T. (1974) Chemical history ofthe oceans deduced from post-depositional changes insedimentary rocks. In Studies in Paleo-Oceanography.Society of Economic Paleontologists and MineralogistsSpecial Publication No. 20 (ed. W. H. Hay). Society ofEconomic Paleontologists and Mineralogists, Tulsa, OK,pp. 193-204.Garrels R. M. and Perry E. A., Jr. (1974) Cycling of carbon,sulfur, and oxygen through geologic time. In The Sea (ed.E. D. Goldberg). Wiley, New York, vol. 5, pp. 303-357.Volume Editor's Introduction xxKasting J. F. and Chang S. (1992) Formation of the earth andthe origin of life. In The Proterozoic Biosphere.An Interdisciplinary Study (eds. J. W. Schopf and C.Klein). Cambridge University Press, New York, pp. 9-12.Kastner M. (1974) The contribution of authigenic feldspars tothe geochemical balance of the alkali metals. Geochim.Cosmochim. Acta 38, 650-653.Kerrick D. M., McKibben M. A., Seward T. M., and CaldeiraK. (1995) Convective hydrothermal CO2 emission from highheat fow regions. Chem. Geol. 121, 17-27.Li Y.-H., Takahashi T., and Broecker W. S. (1969) Degree ofsaturation of CaCO3 in the oceans. J. Geophys. Res. 74,5507-5525.Mackenzie F. T. (2003) Our Changing Planet. An Introductionto Earth System Science and Global Environmental Change.Prentice Hall, New York.Mackenzie F. T. and Garrels R. M. (1966) Chemical massbalance between rivers and oceans. Am. J. Sci. 284, 507-525.Mackenzie F. T. and Kump L. R. (1995) Reverse weathering,clay mineral formation, and oceanic element cycles. Science270, 586-587.Mackenzie F. T., Ver L. M., Sabine C., and Lane M. (1993)C, N, P, S global biogeochemical cycles and modeling ofglobal change. In Interactions of C, N, P and S Biogeo-chemical Cycles and Global Change (eds. R. Wollast,F. T. Mackenzie, and L. Chou). Springer, Berlin, pp. 1-61.Martin W. R., McNichol A. P., and McCorkle D. C. (2000) Theradiocarbon age of calcite dissolving at the sea foor:estimates from pore water data. Geochim. Cosmochim. Acta64, 1391-1404.Michalopoulos P. and Aller R. C. (1995) Rapid clay mineralformation in Amazon Delta sediments: reverse weatheringand oceanic elemental cycles. Science 270, 614-617.Milliman J. D., Troy P. J., Balch W. M., Adams A. K., LiY.-H., and Mackenzie F. T. (1999) Biologically mediateddissolution of calcium carbonate above the chemicallysocline? Deep-Sea Res. 46, 1653-1669.Morse J. W. and Berner R. A. (1972) Dissolution kinetics ofcalcium carbonate in seawater: II. A kinetic origin for thelysocline. Am. J. Sci. 274, 638-647.Murray J. and Irvine R. (1885) On the chemical changes whichtake place in the composition of seawater associated withblue muds on the foor of the ocean. Trans. Roy. Soc.Edinburgh 37, 481-507.Murray J. and Menard A. (1891) Deep sea deposits. InChallenger Reports. Longmans, London, 525p.Petsch S. T. and Berner R. A. (1998) Coupling the geochemicalcycles of C, P, Fe, and S: the effect on atmospheric O2 andthe isotopic records of carbon and sulfur. Am. J. Sci. 298,246-262.Sayles F. L. (1979) The composition and diagenesis ofinterstitial solutions: I. Fluxes across the seawater-sedimentinterface in the Atlantic Ocean. Geochim. Cosmochim. Acta43, 527-545.Sayles F. L. (1981) The composition and diagenesis ofinterstitial solutions: II. Fluxes and diagenesis at thesediment-water interface in the high latitude North andSouth Atlantic. Geochim. Cosmochim. Acta 45, 1061-1086.Sepkoski J., Jr. (1993) Ten years in the library: newdata confrmpaleontological patterns. Paleobiology 19, 43-51.Siever R. (1992) The silica cycle in the Precambrian. GeochimCosmochim. Acta 56, 3265-3272.Tissot B. and Welte D. H. (1978) Petroleum Formation andOccurrence. Springer, Berlin.Tissot B., Durand B., Espitalie J., and Combaz A. (1974)Infuence of nature and diagenesis of organic matter information of petroleum. Am. Assoc. Petrol. Geol. Bull. 58,499-506.Tizzard T. H., Moseley H. N., Buchanan M. A., and Murray J.(1885) Report of the Scientifc Results of the Voyage of H. M.S. Challenger during the Years 1873-1876, vol. 1, pt. 1.Her Majesty`s Stationery Offce, London, 509p.Veizer J. (1988) The evolving exogenic cycle. In ChemicalCycles in the Evolution of the Earth (eds. C. B. Gregor,R. M. Garrels, F. T. Mackenzie, and J. B. Maynard). Wiley,New York, pp. 175-220.Velde B. (1985) Clay Minerals. A Physio-chemical Expla-nation of Their Occurrence. Elsevier, Amsterdam, 427p.Williams L. A. and Crerar D. A. (1985) Silica diagenesis:II. General mechanisms. J. Sedim. Petrol. 55, 312-321.Volume Editor's Introduction xxiThis Page Intentionally Left Blank7.01Chemical Composition andMineralogy of Marine SedimentsY.-H. Li and J. E. SchoonmakerUniversity of Hawaii, Honolulu, HI, USA7.01.1 INTRODUCTION 17.01.2 PELAGIC SEDIMENTS 47.01.2.1 Equatorial Pacic 77.01.2.2 South Pacic 87.01.2.3 Central Indian Basin 107.01.3 FERROMANGANESE NODULES AND CRUSTS 117.01.3.1 Equatorial Pacic Nodules 157.01.3.2 Seamount Ferromanganese Crusts from the Central Pacic 177.01.4 METALLIFEROUS RIDGE AND BASAL SEDIMENTS 187.01.4.1 Metalliferous Ridge Sediments 237.01.4.2 Metalliferous Basal Sediments 247.01.5 MARINE PHOSPHORITES 247.01.6 CONCLUSIONS 28REFERENCES 307.01.1 INTRODUCTIONThe earliest reports on the composition of deep-sea sediments resulted from the ChallengerExpedition (18731876) (e.g., Tizzard et al.,1885; Murray and Renard, 1891). Many reviewpapers on marine sediment composition havesubsequently been published, including the onesby Revelle (1944), El Wakeel and Riley (1961),Arrhenius (1963), Goldberg (1963), Chester andAston (1976), Glasby (1977), Bischoff and Piper(1979), Baturin (1982, 1988), Notholt and Jarvis(1990), Nicholson et al. (1997), Glenn et al.(2000), and Li (2000). The constituents of amarine sediment are often classied according totheir origin (Table 1; after Goldberg, 1963). Thedetrital component is made up of cosmogenousand lithogenous materials. Cosmic spherulescontain particles of FeNi that are formed byablation of iron meteorites as they pass throughEarths atmosphere, as well as fragments ofsilicate chondrules (Arrhenius, 1963). Lithogen-ous constituents of marine sediments are theminerals derived from weathering of rock on landor on the seaoor, or from the volcanic eruptions(Goldberg, 1963; see review in Windom (1976)).The biogenous component is made up of the testsof planktic and benthic organisms, as well asbiogenic apatite (see review in Berger (1976)).The hydrogenous fraction of marine sedimentencompasses phases formed by inorganic precipi-tation from seawater. Eldereld (1976) and Piperand Heath (1989) provide comprehensive reviewsof hydrogenous material in marine sediments.In this chapter, we present a review of thecomposition of the marine sediments, rst addres-sing pelagic sediments in general, and thenconsidering several specic types of sediment,each dominated by hydrogenous components.In each section, a review of mineralogy isfollowed by an examination of the geochemistryusing enrichment factors and factor analysis.Table 2 summarizes the average compositionsof the Earths upper crust, shale, and marinesediments, including pelagic sediments, ferro-manganese nodules and crusts, metalliferous1ridge and basal sediments, and marine phosphor-ites. Also given in Table 2 are the compositions ofseawater and the hydrothermal vent solution fromthe Hanging Garden vent on the mid-Pacic rise at218 N (Li (2000) and references therein; additionaldata sources are given in the footnote to Table 2).One convenient way to compare a given sampleto a chosen reference material is the so-calledenrichment factor. The enrichment factor (Eij) isdened as the concentration ratio of a givenelement i and the normalizing element j(Xi=Xj) inthe given sample divided by the same ratio in thereference material, i.e.,Eij = (Xi=Xj)sample(Xi=Xj)referenceAn Eij value of greater than one representsenrichment of element i in the sample as comparedto the reference; whereas a value less than onemeans depletion. In order to avoid possibleconfusion, sample and reference names can beadded after Eij: As shown in Figure 1(a), theaverage shale composition is very similar to that ofthe upper continental crust (EiAl = 1:0 ^0:3):The obvious exceptions are Li, and volatileelements B, C, N, S, Se, Te, Br, I, As, Cd, In, Sb,Hg, and Bi, which are enriched in the shale incomparison to the upper continental crust whenaluminum is chosen as the normalizing element(Li, 2000). These excess volatile elements camefrom the interior of the Earth by magmaticdegassing processes during the Earths earlyhistory (Li, 1972). The depletion of Ca, Sr, andNa in the shale relative to the Earths upper crust isbalanced by the increased presence of thoseelements found in carbonate rocks, evaporates,and seawater. The average compositions of shaleand marine pelagic clay are similar (Figure 1(b);within a factor of 2). The obvious exceptions are B,Na, P, Mn, Co, Ni, Cu, Mo, Pd, Te, Ba, W, Os, Ir,Pt, Tl, and Pb, which are enriched in the pelagicclay. As will be shown later, many of theseelements are associated with manganese oxidephases in pelagic clay. Shale materials arecontinuously transported to the oceans via riversas suspended particles and via air as aerosols.One disadvantage in dealing with an averagecomposition of any rock type is that importantinformation on the variability of original data,interrelationships among measured elements, andchemical kinship or uniqueness of individualsamples is lost during the process of averaging.Therefore, in the following sections, the statisticaltechnique of factor analysis (Davis, 1973) isapplied to original data in order to nd theunderlying interrelationship among elements (oldvariables), and kinship or variability among a setof given samples. Useful outputs from the factoranalysis are the means, the standard deviationsof variables, correlation coefcient matrices,eigenvalues, factor loadings, and factor scores.The extracted new factor 1 (F1), or new variable 1,from a factor analysis is the best linearTable 1 Mineral constituents of marine sediments classied by origin.Cosmogenous BiogenousSpherules Calcite CaCO3; (Ca12xMgx)CO3Iron FeNi Aragonite CaCO3Olivine (Mg,Fe)2SiO4 Opal SiO2nH2OPyroxene (Mg,Fe)2Si2O6 Francolite Ca102x2yNaxMgy(PO4)62z(CO3)zF0.4zF2Barite BaSO4Celestite SrSO4Lithogenous HydrogenousQuartz SiO2 FeMn oxides/ See Table 3Plagioclase (Ca,Na)(Al,Si)AlSi2O8 oxyhydroxidesClay minerals Francolite Ca102x2yNaxMgy(PO4)62z(CO3)zF0.4zF2Illite KxAl2(Si42xAlx)O10(OH)2 Barite BaSO4Chlorite (Mg,Fe)5(Al,Fe)2Si3O10(OH)8 Celestite SrSO4Kaolinite Al2Si2O5(OH)4 Montmorillonite (Na,K)x(Al22xRx)Si4O10(OH)2Smectite (M)x2y(R3+22yR2+y )(Si42xAlx) Nontronite (Na,K)xFe2(AlxSi42x)O10(OH)2O10(OH)2 Glauconite K0.85(Fe,Al)1.34(Mg,Fe)0.66Volcanic glass (Si3.76Al0.24)O10(OH)2Amphiboles Ca2(Mg,Fe)5Si8O22(OH)2 ZeolitesPyroxene (Mg,Fe)2Si2O6 phillipsite K2.8Na1.6Al4.4Si11.6O3210H2OOlivine (Mg,Fe)2SiO4 clinoptilolite K2.3Na0.8Al3.1Si14.9O3612H2OGeothite FeOOHPalygorskite (OH2)4Mg5Si8O20(OH)24H2OSepiolite (OH2)4Mg8Si12O30(OH)24H2OAfter Goldberg (1986).Chemical Composition and Mineralogy of Marine Sediments 2Table 2 Average compositions of the Earths upper continental crust, shale, marine sediments (all in unitsof ppm, noted otherwise), along with seawater and hydrothermal vent solution from the East Pacic Rise (both inunits of 1029g L2l).Element Z Uppercrust(1)Shale(1)Pelagicclay(1)FeMnnodule(1)FeMncrust(1)Basalsed.(2)Ridgesed.(3)Phosphorite(1)Seawater(1)Ventsolution(1)Ag 47 0.06 0.07 0.11 0.09 0.18 6.2h 2 2.5 4,000Al (%) 13 7.83 8.8 8.4 2.7 0.41 2.73 0.5 0.91 300 120,000As 33 1.6 13 20 140 230 145 23 1,700 35,000Au (ppb) 79 2.3 2.5 2 2 250w 16c 1.4 0.03B 5 12 100 230 300 123 500 16 4.5E-6 6.0E-6Ba 56 570 580 2,300 2,300 1,000 6,230 6,000hd 350 1.5E-4 1.5E-6Be 4 3.2 3 2.6 2.5 6.7 2.6 0.21 120Bi 83 0.054 0.43 0.53 7 29w 0.17 0.06 0.004Br 35 2.1 20 21 58 7E-7 6.9E-7C (%) 6 0.023 1.2 0.45 0.1 0.12 2.1 3E-7 7.1E-7Ca (%) 20 3.15 1.6 1 2.3 2.2 1.47 31.4 5E-8 4.8E-8Cd 48 0.1 0.3 0.42 10 3 0.4 4 18 76 20,000Ce 58 58 82 101 530 900 34 8.4 104 1.6 1,640Cl 17 150 180 300 1.9E-10 1.8E-10Co 27 17 19 74 2,700 8,400 82 105 7 1.2 13,000Cr 24 69 90 90 35 9.1 15 d 55 125 252.6Cs 55 3.7 5 6 1 310 28,000Cu 29 39 45 250 4,500 380 790 730 75 210 2.8E-6Dy 66 3.5 4.7 7.4 31 20.7 7.3 19.2 1.5 69Er 68 2 3 4.1 18 24 12.9 5.6 23.3 1.3 35Eu 63 1.1 1.2 1.85 9 8.1 5.4 1.5 6.5 0.21 275F 9 700 740 1,300 200 466 31,000 1.3E-6 0.14E-6Fe (%) 26 4.17 4.72 6.5 12.5 12.3 20 18 0.77 250 1.39E-8Ga 31 18 19 20 10 6.8 4 1.7Gd 64 3.9 5.1 8.3 32 39 22.6 6 12.8 1.3 92Ge 32 1.5 1.6 1.6 0.8 3.3 4.3Hf 72 4 5 4.1 8 1.6 3.4Hg 80 0.08 0.18 0.1 0.15 0.4bf 0.06 0.42Ho 67 0.74 1.1 1.5 7 8.2 4.7 4.2 0.45I 53 0.5 19 28 400 24 58,000In 49 0.05 0.1 0.08 0.25 0.1Ir (ppb) 77 0.05 0.05 0.4 7 10v 0.8c 0.0015K (%) 19 2.56 2.66 2.5 0.7 0.38 1.15 0.42 3.9E-8 9.5E-8La 57 30 43 42 157 190 98 29 133 5.6Li 3 23 66 57 80 125 5 1.8E-5 94E-5Lu 71 0.32 0.42 0.55 1.8 3.9 2.2 0.88 2.7 0.32Mg (%) 12 1.64 1.5 2.1 1.6 0.88 2.08 0.18 1.3E-9 ,0Mn (%) 25 0.077 0.085 0.67 18.6 20.4 6.1 6 0.12 72 4.9E-7Mo 42 1.6 2.6 27 400 370 30 9 10,300N 7 20 1,000 600 200 100 4.20E-5Na (%) 11 2.54 0.59 2.8 1.7 1.5 2.56 0.45 1.8E-10 1.0E-10Nb 41 15 11 14 50 5.1 10 10Nd 60 26 33 43 158 150 87 23 98 4.2 500Ni 28 55 50 230 6,600 3,900 460 430 53 530Os (ppb) 76 0.05 0.05 0.14 2 0.8v 0.0017P 15 860 700 1,500 2,500 3,900 9,000b 138,000 65,000 18,000Pb 82 17 20 80 900 1,400 100 152h 50 2.7 75,000Pd (ppb) 46 1 (1) 6 6 1.1 21c 0.07Pr 59 6.6 9.8 10 36 34 19.3 21 0.87Pt (ppb) 78 1 (1) 5 200 350 0.1Rb 37 110 140 110 17 16 0.12E-6 2.8E-6Re (ppb) 75 0.4 0.4 0.3 1 8Rh (ppb) 45 0.4 13 14Ru (ppb) 44 0.2 8 0.005S 16 530 2,400 2,000 4,700 7,200 8.98E-8 1.3E-7Sb 51 0.2 1.5 1 40 17d 7 150Sc 21 14 13 19 10 11 0.86(continued)Introduction 3combination of the old variables to account for thelargest fraction of the total variance in the wholedata set. The factor 2 (F2) is the next best linearcombination to account for the residual variance,and so on. Only those factors with eigenvaluesgreater than one are extracted here. Therefore,the number of extracted factors (new variables)is much smaller than the number of originalvariables (elements). One may visualize theextracted factor loadings as the correlationcoefcients between old variables and newfactors, and the extracted factor scores as thenew concentrations of new factors in eachsample. Factor loading ranges between 21 and-1, and factor scores are allowed to have negativeconcentrations. Any correlation coefcient (g)or factor loading value between 0.7 and 1 repre-sents in this context a strong correlation, between0.5 and 0.69 a moderate correlation, between 0.3and 0.49 a weak correlation, and between 0 and0.29, no correlation at all. Similarly, a negativecorrelation coefcient represents anti- or inversecorrelation.7.01.2 PELAGIC SEDIMENTSPelagic sediments are generally dened as thosedeposited from dilute suspensions of detritalmaterial that are distributed throughout deep-ocean water (Arrhenius, 1963). They are charac-terized by low accumulation rates of terrigenousmaterial, and presence of relatively highpercentages of authigenic minerals, cosmogenicmaterial, and biogenic planktonic debris. Becauseof their slow depositional rates (,mm Kyr21),pelagic sediments tend to undergo high degrees ofoxidation. The two major types of pelagicsediment are pelagic clay and biogenous oozes.Pelagic clays accumulate at abyssal depths (belowthe compensation depths for carbonate minerals),in oligotrophic regions of the ocean, far fromsources of terrigenous turbidites (e.g., Leinen,1989). They consist of ne-grained (generally,3 mm; Horn et al. (1970)) terrigenous material,largely eolian in origin, and generally contain afew percent of authigenic minerals. In shallowerregions of the oceans, or under areas of highproductivity, calcareous and/or siliceous testsaccumulate to form biogenous oozes (Figure 2).The dominant lithogenous components of pela-gic sediments are quartz, plagioclase, and clayminerals (Table 1). The primary clay minerals areillite, chlorite, kaolinite, and smectite. Relativeabundance of these constituents depends onproximity to sources; on an average, the clay-sizefraction of North Pacic pelagic clay contains3040% illite, 1015% chlorite, 1015% quartz,1015% plagioclase, 1015% kaolinite, and05% smectite (Leinen (1989) and referencestherein). Submarine weathering of basaltic rock,and marine volcanic activity, are sources ofplagioclase, amphiboles, pyroxenes, and olivines(Windom, 1976). Windom (1976) notes a numberof other lithogenous phases that are generallypresent in marine sediments in minor amounts,Table 2 (continued).Element Z Uppercrust(1)Shale(1)Pelagicclay(1)FeMnnodule(1)FeMncrust(1)Basalsed.(2)Ridgesed.(3)Phosphorite(1)Seawater(1)Ventsolution(1)Se 34 0.14 0.6 0.2 0.6 2.6 4.6 145 4,800Si (%) 14 30 27.5 25 7.7 2.2 10.8 6.1 5.6 2.5E-6 4.5E-8Sm 62 4.5 6.2 8.35 35 30 18.6 5 20 0.84 137Sn 50 3.3 3 4 2 0.6 3 0.6Sr 38 350 170 180 830 1,200 351 750 7.8E-6 5.8E-6Ta 73 1.5 1.3 1 10 2.1 2.5Tb 65 0.6 0.84 1.42 5.4 5 3.2 0.21Te 52 0.003 0.07 1 10 0.07Th 90 11 12 13 30 2.4 6.5 0.05 0.3Ti 22 3,300 4,600 4,600 6,700 7,700 240 640 10Tl 81 0.53 0.7 1.8 150 4.8 34h 14Tm 69 0.32 0.44 0.57 2.3 3.6 1.2 0.25U 92 2.8 2.7 2.6 5 4.2 22f 120 3,200 ,0V 23 140 130 120 500 500 450 100 2,150W 74 1.3 1.8 4 100 100Y 39 22 26 40 150 190 128 260 13Yb 70 2 2.8 3.82 20 24 13 5.7 13 1.5 33Zn 30 67 95 170 1,200 540 470 380 200 320 6.9E-6Zr 40 170 160 150 560 225 70 17Data sources: (1) Li (2000) and references therein; (2) Cronan (1976); (3) Bostrom and Peterson (1969); b = Bostrom (1973); bf = Bostrom andFisher (1969); c = Crocket et al. (1973); d = Dymond et al. (1973); f = Fisher and Bostrom (1969); h = Horowitz (1970); hd = Heath andDymond (1977); v = Vonderhaar et al. (2000); w = Wen et al. (1997); REE = Jarvis (1985) for basal sediments, and Bender et al. (1971) for ridgesediments. Pd and Pt values for shale are an educated guess based on the upper crust values.Chemical Composition and Mineralogy of Marine Sediments 4particularly in coastal regions near their sources.Most lithogenous constituents undergo little or notransformation during deposition in the oceanbasins (Windom, 1976). Eldereld (1976) dis-cusses the cation exchange and sorption reactionsthat clay minerals undergo in seawater. In additionto its detrital source, the clay mineral smectiteforms authigenically in the ocean basins. Theformation of smectite at low temperature via thehydrogenous and hydrothermal means is discussedbelow.Biogenous oozes are either calcareous orsiliceous. Calcareous oozes are predominantlythe calcitic tests of coccolithophores and/orforaminifera, or the aragonitic tests of pteropods.The solubility of CaCO3 increases with decreasingtemperature and increasing pressure, and thus withincreasing depth in the oceans. Aragonite is ,1.45times more soluble than calcite (Morse andMackenzie, 1990), so aragonitic oozes are con-ned to shallower depths than the calcitic oozes.The compensation depth for each mineral isdened as the depth at which the rates ofdeposition and dissolution of that mineral areequal, so the content of that mineral drops to0 wt.% (e.g., Morse and Mackenzie, 1990; Pinet,2000). Compensation depths are dependent on anumber of factors and vary between, and across,ocean basins (e.g., Berger et al., 1976; Biscayeet al., 1976; Kolla et al., 1976; see discussion inMorse and Mackenzie, 1990). The aragonitecompensation depth (ACD) generally averagesaround 3,000 m (Berger, 1978), whereas thecalcite compensation depth (CCD) ranges from5,0006,000 m in the Atlantic (Biscaye et al.,1976) to 4,0005,000 m in the Pacic (Bergeret al., 1976). Calcareous oozes therefore com-monly occur on ridge crests and other topographichighs.Although calcite and aragonite are both madeup of calcium carbonate, the different structures ofthe two minerals accommodate different elemen-tal substitutions (see review in Mucci and Morse(1990)). Strontium and, to a lesser extent,magnesium and sodium can substitute for calciumin skeletal aragonite (Speer, 1983). MagnesianEAl(shale/uppercrust)iEAl(pelagicclay/shale)iCNBLiSBrIBiHgThUTlAuPbPt OsTaRe IrWHfEuBaCs CeSnAgPdNb SrZrYRbZnGeGaNaFeCuCo MnCr TiVScKCaClAlSiPMgFBeTeSbCdInMoAsSe1021011001010 10 20 30 40 50 60 70 80 90Atomic number (Z)1021011001010 10 20 30 40 50 60 70 80 90Atomic number (Z)s blockp blockd blockf (La)f (Ac)s blockp blockd blockf (La)f (Ac)UThBiTlPbIrPtOsWLuHf ReTaAuHgEuBaTeCeLaCsIPdAgSnCdSbInNbYZrSrRbAsZnGaGeMoMnCuNiCoNaPSc FeTiCrVSeCaKSSiAlMgFBLiNBeC(a)(b)Figure 1 Enrichment factors of elements (EiAl) in: (a) average shale as compared with the Earths upper continentalcrust, and in (b) marine pelagic clay as compared with average shale. Data are from Table 2.Pelagic Sediments 5calcites contain up to 30 mol.% MgCO3 (e.g.,Tribble et al., 1995). Magnesian calcites areprecipitated by a wide variety of shallow watermarine organisms, but the increase in solubilityaccompanying magnesium substitution limitsaccumulation of these phases to shallow depthsin the oceans (e.g., Bischoff et al., 1987;Mackenzie et al., 1983). Calcite also may containminor substituents such as Sr2-, Na-, Mn2-, andSO422ions.Siliceous oozes are accumulations of opalinesilica (opal-A, an amorphous phase of high watercontent and porosity) in the tests of diatoms,radiolarians, and/or silicoagellates. Opal-A solu-bility at 25 8C is 60130 ppm SiO2(aq) (e.g.,Williams et al., 1985), and solubility increases withincreasing temperature and pressure (Walther andHelgeson, 1977). Adsorption of aluminumand ironon the surfaces of siliceous tests decreases theirsolubility (Iler, 1955; Lewin, 1961). Opal-A is ametastable phase that with burial eventuallyrecrystallizes to quartz, often with another meta-stable intermediary phase, opal-CT (e.g., Heinet al., 1978; Williams et al., 1985; Williams andCrerar, 1985). Opal-CT structurally resembles aninter-layering of the two silica phases, cristobaliteand tridymite. Siliceous oozes are usually found onthe oor of the Southern Ocean around Antarctica,along the equatorial oceans (especially in thePacic), and around the northernedge of the PacicOcean (Broecker and Peng, 1982; Pinet, 2000).The hydrogenous fraction of pelagic sedimentmay contain a wide variety of authigenic minerals(Table 1). The discussion here will be limited to themore abundant phases not considered in sub-sequent sections, or in separate chapters in thisvolume (e.g., glauconite, see Chapter 7.12). Bariteis the only sulfate mineral present in abundance inpelagic sediment (e.g., Piper and Heath, 1989).Barite of hydrothermal origin is associated withmid-ocean ridge deposits (e.g., Edmond et al.,1980). Crystals of nonhydrothermal barite havebeen recovered by seawater ltration (Dehairset al., 1980), and barite is commonly found atconcentrations of 15% in sediments that are alsorich in organic matter (Piper and Heath, 1989).Church (1979) suggested that the microenviron-ment of pore water might be suitable forprecipitation of barite. Most of the worlds oceansare undersaturated with respect to barite. Thesolution model for barite, thanks to Monnin et al.(1999), shows that an equilibrium with pure bariteFigure 2 Distribution of the principal types of marine sediment on the seaoor. Calcareous oozes are found inrelatively shallow oceanic regions, such as along ridge crests. Siliceous oozes are found predominantly under theequatorial regions of high productivity and in the high-latitude oceans. Deep-sea or pelagic clay occupies the abyssaldepths of the ocean basins, far from terrigenous sources of river-borne or glacial sediment (source Davies andGorsline, 1976) (reproduced by permission of Elsevier from Chemical Oceanography, 1976, 5).Chemical Composition and Mineralogy of Marine Sediments 6is reached or slightly exceeded only in the coldsurface waters of the Southern Ocean, at inter-mediate depths in the Pacic, and in the deepwaters of the Bay of Bengal. It is possible that thesolubility of barite decreases upon substitution ofstrontium for barium stabilizing the phase(Church, 1979). Barite is secreted by certainmarine plankton, and it is possible that much ofmarine barite is biogenic in origin (Dehairs et al.,1980, 1990; Bishop, 1988; Bertram and Cowen,1997).Phillipsite is the most abundant zeolite in thesurface sediments of the Pacic (Boles, 1977;Kastner and Stonecipher, 1978). Although it maybe locally abundant (.50 wt.% on a carbonate-free basis, Bonatti (1963)), its etched surface, andthe pattern of its decreasing abundance with theburial depth, suggest that it is a metastable phaseunder deep-sea conditions (Kastner, 1979). Theprimary mechanism of formation is thought to bealteration of basaltic glass, but it may also formby reaction of biogenic silica and dissolved Al3-(Arrhenius, 1963). Phillipsite is commonly foundin association with authigenic smectite, andthe combined formation of the two mineralsmay be represented asbasalt - K- - Na- - SiO2 - H2O - H-!phillipsite - smectite - Ca2- - Fe2O3(Piper and Heath, 1989).The present-day sediments of the Atlantic do notcontain a substantial amount of zeolite (Eldereld,1976). Clinoptilolite has been found, however, inassociation with sepiolite, quartz, and mont-morillonite (Hathaway and Sachs, 1965; Bonattiand Joensuu, 1968). The clayminerals palygorskiteand sepiolite are usually minor constituents ofmarine sediments (Hathaway, 1979), and may bedetrital (Weaver and Beck, 1977) or authigenic.Their hydrogenous occurrences are usually in basalsediment sections exposed to uids of elevatedtemperatures (Bonatti and Joensuu, 1968; Churchand Velde, 1979; see below).Geochemically, pelagic sediments can beconsidered as mixtures of the major constituents:pelagic clay, calcium carbonate, opaline silica,and apatite. The following subsections providesome examples to show how the compositions ofpelagic sediments change in different environ-ments and to elucidate the factors controllingobserved compositional changes.7.01.2.1 Equatorial PacicResults of factor analysis (varimax; used in thischapter throughout) on chemical composition datafor the surface siliceous oozes from the Wahinesurvey area (8820/ N, 1538 W; 34 samples; Calvertet al. (1978)) are shown in Figures 3(a) (factorloadings) and (b) (factor scores). The results forthe surface pelagic sediments from the widerequatorial Pacic (288 N138 S, 1178 W1758 E;17 samples; Calvert and Price (1977)) are shownin Figures 3(c) and (d). One unusual sampleAMPH-85PG from the equatorial Pacic, whichcontains micronodules, is excluded from thefactor analysis. A group of elements in Figures3(a) and (c) is enclosed by dotted ellipse,whenever the correlation coefcients are greaterthan 0.49 among all pairs within a dotted ellipse(based on the correlation coefcient matrix). Anytwo elements connected by a solid line also have acorrelation coefcient higher than 0.49. The samerepresentation applies to all other similar guresin this chapter.In Figure 3(a) (F1 loading versus F2 loading), F1is characterized by Al, Ti, K, and Rb, which areassociated with terrigenous shale in siliceous oozesof the Wahine survey area. Silicon is only partlyassociated with this group due to additionalindependent input of siliceous radiolarian shells.F2 is represented by Mn, Ni, Cu, Zn, Mo, Ba, andSr. These elements are preferentially incorporatedinto MnO2 phases. Factor 3 (F3) (not plotted here)represents a carbonate-uorapatite (CFA) phase(Ca, P, and Y). Mixing of these four end-members(F1 to F3 and biogenic silica) in differentproportions forms the observed sediments in theWahine survey area. Rare earth elements (REEs),except cerium, are shown to be closely associatedwith CFA(Ca, P, and Y; Eldereld et al. (1981)) inthe sediments from the Wahine survey area.As shown in Figure 3(b) (factor scores of F1 versusF2), most samples are chemically similar, exceptthat two samples have high F2 group elements andone sample has high F1 group elements. Theelements Fe, Co, As, Pb, Ce, and Mg are notincluded in the nal factor analysis, because theconcentrations of these elements correlate onlyweakly, or not at all, with one another and with anyother element. The implication is that theseelements are more or less evenly distributedamong different phases of samples, and do notconcentrate in any particular phase.Factor analysis of the equatorial Pacic sedi-ments (Figure 3(c)) indicates four major com-ponents: F1 (aluminosilicates: Al, Si, Ti, Th, Zr,K, Rb, Fe, As, and Mg), negative F1 (carbonates:Ca and Sr), F2 (manganese oxides: Mn, Ni, Co,Mo, Pd, Cu, and Zn), and F3 (CFA: P and Y).Most samples are chemically similar (Figure 3(d)),except that four samples from topographic highsare carbonate rich and one pelagic clay sample ismanganese rich.In the equatorial Pacic (8840/ N10810/ N,173850/ W1758 W), several 67 m long coresand many surface sediment samples at waterdepths of ,6,000 m (below the carbonatecompensation depth) were collected during thePelagic Sediments 7GH80-5 cruise by the Geological Survey of Japan.Samples were analyzed for major elements(Sugisaki and Yamamoto, 1984) and traceelements (Mita and Nakao, 1984). They classiedthe samples into three major groups: siliceous clay(radiolarian-rich) on the core top, pelagic clay inmid-core, and zeolitic clay (zeolite .5%) below.Factor analysis of 101 carbonate-free samples(excluding the four samples containing Al2O3 lessthan 6.1%, caused by a large dilution effect ofbiogenic silica) shows four major components(Figures 4(a) and (b)): F1 (aluminosilicates: Al,Fe, and Ti), negative F1 (CFA: Ca and P), F2(manganese oxides: Mn, Ni, Co, Cu, Zn, and Mo),and F3 (zeolite - biogenic silica: K, Na, and Si).According to the correlation coefcient matrix, K,Na, and Si are also moderately correlated to Al.According to Figures 4(c) and (d), pelagic claysamples are mostly characterized by the presenceof high aluminosilicate (F1), siliceous clay by lowMnO2 (F2) and moderately high silica (F3), andzeolitic clay by high MnO2 (F2) and high CFA(negative F1). The xy-plots of the originalconcentration data of some selected key elementsfor F1F4 (Figures 5(a) (f)) conrm thesecharacterizations. Notice the similarity betweenFigures 4(c) (F1versus F2) and5(a) (Al versus Mn),and Figures 4(d) (F1 versus F3) and 5(b)(Al versus Si). Al and Si are highly correlatedfor pelagic clay (Figure 5(b); open circles with adotted dilution line). In contrast, silica concen-tration is high in siliceous and zeolitic claysrelative to pelagic clay at a given aluminumconcentration. Similarly, K and Na are stronglycorrelated with Al for pelagic and siliceous clays(Figure 5(d); open circles and crosses). However,K and Na concentrations can be very high inzeolitic clay. Data points in Figure 5(e) fall on thedotted line with a CaO/P2O5 slope of 1.32(in weight) for uora