Marine Geochemistry || The Solid Phase of Marine Sediments

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1 The Solid Phase of Marine SedimentsDIETER K. FÜTTERER

1.1 Introduction

The oceans of the world represent a natural de-pository for the dissolved and particulate prod-ucts of continental weathering. After its input, thedissolved material consolidates by means of bio-logical and geochemical processes and is depos-ited on the ocean floor along with the particulatematter from weathered rock. The ocean floor de-posits therefore embody the history of the conti-nents, the oceans and their pertaining watermasses. They therefore provide the key for under-standing Earth’s history, especially valuable forthe reconstruction of past environmental condi-tions of continents and oceans. In particular, thequalitative and quantitative composition of thesedimentary components reflect the conditions oftheir own formation. This situation may be more orless clear depending on preservation of primarysediment composition, but the processes of earlydiagenesis do alter the original sediment composi-tion, and hence they alter or even wipe out theprimary environmental signal. Hence, only anentire understanding of nature and sequence ofprocesses in the course of sediment formation andits diagenetic alteration will enable us to infer theinitial environmental signal from the alteredcomposition of the sediments.

Looking at the sea-floor sediments from ageochemical point of view, the function of par-ticles, or rather the sediment body as a whole, i.e.the solid phase, can be quite differentlyconceived and will vary with the perspective ofthe investigator. The “classical” approach –simply applying studies conducted on thecontinents to the oceans – usually commenceswith a geological-sedimentological investigation,whereafter the mineral composition is recorded indetail. Both methods lead to a more or less overallgeochemical description of the entire system.

Another, more modern approach conceives theocean sediments as part of a global system inwhich the sediments themselves represent avariable component between original rock sourceand deposition. In such a rather process-relatedand globalized concept of the ocean as a system,sediments attain special importance. First, theyconstitute the environment, a solid framework forthe geochemical reactions during early diagenesisthat occur in the pore space between the particlesin the water-sediment boundary layer. Next to theaqueous phase, however, they are simultaneouslystarting material and reaction product, andprocure, together with the porous interspaces, amore or less passive environment in whichreactions take place during sediment formation.

1.2 Sources and Components ofMarine Sediments

Ocean sediments are heterogeneous with regardto their composition and also display a consider-able degree of geographical variation. Due to theorigin and formation of the components varioussediment types can be distinguished: Lithoge-nous sediments which are transported and dis-persed into the ocean as detrital particles, eitheras terrigenous particles – which is most fre-quently the case – or as volcanogenic particleshaving only local importance; biogenous sedi-ments which are directly produced by organismsor are formed by accumulation of skeletal frag-ments; hydrogenous or authigenic sedimentswhich precipitate directly out of solution as newformations, or are formed de novo when the par-ticles come into contact with the solution; finally,cosmogenic sediments which are only of second-ary importance and will therefore not be consid-ered in the following.

1 The Solid Phase of Marine Sediments

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1.2.1 Lithogenous Sediments

The main sources of lithogenous sediments are ul-timately continental rocks which have been bro-ken up, crushed and dissolved by means of physi-cal and chemical weathering, exposure to frost andheat, the effects of water and ice, and biologicalactivity. The nature of the parent rock and theprevalent climatic conditions determine the inten-sity at which weathering takes place. Informationabout these processes can be stored within theremnant particulate weathered material, the terri-genous detritus, which is transported by variousroutes to the oceans, such as rivers, glaciers andicebergs, or wind. Volcanic activity also contrib-utes to lithogenous sediment formation, however,to a lesser extent; volcanism is especially effectiveon the active boundaries of the lithosphericplates, the mid-ocean spreading ridges and thesubduction zones.

The major proportion of weathered material istransported from the continents into the oceansby rivers as dissolved or suspension load, i.e. inthe form of solid particulate material. Dependingon the intensity of turbulent flow suspension loadgenerally consists of particles smaller than 30 mi-crons, finer grained than coarse silt. As the min-eral composition depends on the type of parent

rock and the weathering conditions of the catch-ment area, it will accordingly vary with each riversystem under study. Furthermore, the mineralcomposition is strongly determined by the grain-size distribution of the suspension load. This canbe seen, for example, very clearly in the suspen-sion load transported by the Amazon River (Fig.1.1) which silt fraction (> 4 - 63 µm) predominantlyconsists of quartz and feldspars, whereas mica,kaolinite, and smectite predominate in the clayfraction.

It is not easy to quantify the amount of sus-pension load and traction load annually dis-charged by rivers into the oceans on a worldwidescale. In a conservative estimative approachwhich included 20 of the probably largest rivers,Milliman and Meade (1983) extrapolated thisamount to comprise approximately 13·109 tons.Recent estimations (Milliman and Syvitski 1992)which included smaller rivers flowing directly intothe ocean hold that an annual discharge of ap-proximately 20·109 tons might even exist.

Under the certainly not very realistic assump-tion of an even distribution over a surface area of362·106 km2 which covers the global ocean floor,this amount is equivalent to an accumulation rateof 55.2 tons km-2 y-1, or the deposition of an approxi-mately 35 mm-thick sediment layer every 1000 years.

Fig. 1.1 Grain-size distribution of mineral phases transported by the Amazon River (after Gibbs 1977).

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1.2 Sources and Components of Marine Sediments

Most of the sediment transported to the coast-line by the rivers today is deposited on protectedcoastal zones, in large estuaries, and on theshelves; only a rather small proportion of the sedi-ment is transported beyond the shelf edge andreaches the bottom of the deep sea. The geo-graphical distribution of the particulate dischargevaries greatly worldwide, depending on the geo-graphical distribution of the respective rivers,amount and concentration of the suspended mate-rial. According to Milliman and Syvitski (1992), theamount of suspension load is essentially a func-tion of the surface area and the relief of thecatchment region, and only secondarily does itdepend on the climate and the water mass of therivers. Apart from these influences, others like hu-man activity, climate, and geological conditionsare the essential factors for river systems insoutheast Asia.

The southeast Asian rivers of China, BanglaDesh, India, and Pakistan that drain the highmountain region of the Himalayan, and the riversof the western Pacific islands (Fig. 1.2), transportjust about one half of the global suspension loaddischarged to the ocean annually. This must natu-rally also exert an effect on the sedimentation

rates in the adjoining oceanic region of the Indo-Pacific.

Sediment transport by icebergs which calvefrom glaciers and inland ice into the ocean at polarand subpolar latitudes is an important process forthe discharge and dispersal of weathered coarsegrained terrigenous material over vast distances.Due to the prevailing frost weathering in nivalclimate regions, the sedimentary material which isentrained by and transported by the ice is hardlyaltered chemically. Owing to the passive transportvia glaciers the particles are hardly rounded andhardly sorted in fractions, instead, they comprisethe whole spectrum of possible grain sizes, frommeter thick boulders down to the clay-sizefraction.

As they drift with the oceanic currents, meltingicebergs are able to disperse weathered terrig-enous material over the oceans. In the southernhemisphere, icebergs drift from Antarctica north to40°S. In the Arctic, the iceberg-mediated transportis limited to the Atlantic Ocean; here, icebergsdrift southwards to 45°N, which is about thelatitude of Newfoundland. Coarse componentsreleased in the process of disintegration and melt-ing leave behind “ice rafted detritus” (IRD), or

Fig. 1.2 Magnitude of annual particulate sediment discharge of the world’s major rivers. The huge amount of sedimentdischarge in southeast Asia and the western Pacific islands is due to high relief, catchment, precipitation and humanactivity (Hillier 1995).


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“drop stones”, which represent characteristic sig-nals in the sediments and are of extremely highimportance in paleoclimate reconstructions.

In certain regions, the transport and the distri-bution carried out by sea ice are important pro-cesses. This is especially true for the Arctic Oceanwhere specific processes in the shallow coastalareas of the Eurasian shelf induce the ice, in thecourse of its formation, to incorporate sedimentmaterial from the ocean floor and the watercolumn. The Transpolar Drift distributes the sedi-ment material across the Arctic Ocean all the wayto the North Atlantic. Glacio-marine sedimentationcovers one-fifth of present day’s ocean floor(Lisitzin 1996).

Terrigenous material can be carried from thecontinents to the oceans in the form of mineraldust over great distances measuring hundreds toup to thousands of kilometers. This is accom-plished by eolian transport. Wind, in contrast toice and water, only carries particles of finer grainsize, such as the silt and clay fraction. A grain sizeof approximately 80 µm is assumed to mark thehighest degree of coarseness transportable bywind. Along wind trajectories, coarser grains suchas fine sand and particles which as to their sizesare characteristic of continental loess soil (20-50µm) usually fall out in the coastal areas, whereas

the finer grains come to settle much farther away.The relevant sources for eolian dust transportare the semi-arid and arid regions, like the Sahelzone and the Sahara desert, the Central Asiandeserts and the Chinese loess regions (Pye1987). According to recent estimations (Prospero1996), a total rate of approximately 1-2·109 tons y-

1 dust is introduced into the atmosphere, ofwhich about 0.91·109 tons y-1 is deposited intothe oceans. This amount is, relative to the entireterrigenous amount of weathered material, notvery significant; yet, it contributes considerablyto sediment formation because the eolian trans-port of dust concentrates on few specific regions(Fig. 1.3). Dust from the Sahara contributes tosedimentation on the Antilles island Barbados ata rate of 0.6 mm y-1, confirming that it is not at alljustified to consider its contribution in buildingup deep-sea sediments in the tropical and sub-tropical zones of the North Atlantic as negligible.Similar conditions are to be found in the north-western Pacific and the Indian Ocean where greatamounts of dust are introduced into the oceanfrom the Central Asian deserts and the Arabicdesert.

According to rough estimates made by Lisitzin(1996), about 84 % of terrigenous sediment inputinto the ocean is effected by fluvial transport,

DESERTS SEMI-ARIDREGIONS

DUST TRANSPORTDIRECTIONS ANDDISTANCES

Fig. 1.3 The world’s major desert areas and semi-arid regions and potential long-distance eolian dust trajectories andoceanic depocenters (Hillier 1995).

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somewhat more than 7 % by eolian transport, andless than 7 % is due to the activity of icebergs.

A distinctly less, but still not insignificant pro-portion of lithogenous sediment is formed byvolcanic activity which is quite often coupled withprocesses of active subduction at the continentalplate boundaries. A large proportion of pyroclasticfragments becomes wind-dispersed over largeareas, whereafter they are usually retraced inoceanic sediments as finely distributed volcanicglass. Yet, the formation of single, distinct, cm-thick tephra layers might also occur in the deepsea where they represent genuine isochronousmarkers which can be used for correlationpurposes and the time calibration of stratigraphicunits. Layers of ash deposits in the eastern Medi-terranean are prominent examples indicative of theeruption of the volcanic island Ischia in prehis-toric times of more than 25,000 years ago, andSantorin about 3500 years ago.

Locally, there may be a frequent occurrence oftephra layers and significant concentrations offinely dispersed volcaniclastic material in deep-sea sediments especially in the proximal zones ofvolcanic activity, like in marginal zones of themodern Pacific Ocean. In a recent evaluation of te-phra input into the Pacific Ocean sediments basedon DSDP and ODP data Straub and Schmincke(1998) estimate that the minimum proportion ofvolcanic tephra corresponds to 23 vol.% of the ex-isting Pacific oceanic sediments.

Lithogenous detrital components of marinesediments, despite all regional variability, includeonly few basic minerals (Table 1.1). With the ex-ception of quartz, complete weathering, particu-larly the chemical weathering of metamorphic andigneous rock, leads to the formation of clayminerals. Consequently, this group represents,apart from the remaining quartz, the most impor-tant mineral constituent in sediments; clay miner-als make up nearly 50 % of the entire terrigenoussediment. To a lesser degree, terrigenous detrituscontains unweathered minerals, like feldspars.Furthermore, there are mica, non-biogenous cal-cite, dolomite in low quantities, as well as acces-sory heavy minerals, for instance, amphibole, py-roxene, apatite, disthene, garnet, rutile, anatase,zirconium, tourmaline, but they altogether seldomcomprise more than 1 % of the sediment. Basically,each mineral found in continental rock - apart fromtheir usually extreme low concentrations - mayalso be found in the oceanic sediments. Thepercentage in which the various minerals are

present in a sediment markedly depends on thegrain-size distribution.

The clay minerals are of special importance in-asmuch as they not only constitute the largestproportion of fine-grained and non-biogenoussediment, but they also have the special geo-chemical property of absorbing and easily givingoff ions, a property which affords more detailedobservation. Clay minerals result foremost fromthe weathering of primary, rock forming aluminoussilicates, like feldspar, hornblende and pyroxene,or even volcanic glass. Kaolinite, chlorite, illite,and smectite which represent the four mostimportant groups of clay minerals are formedpartly under very different conditions of weather-ing. Consequently, the analysis of their qualitativeand quantitative distribution will enable us todraw essential conclusions on origin and trans-port, weathering and hydrolysis, and therefore onclimate conditions of the rock’s source region(Biscay 1965; Chamley 1989). The extremely fine-granular structure of clay minerals, which is likelyto produce an active surface of 30 m2g-1 sediment,as well as their ability to absorb ions internallywithin the crystal structure, or bind themsuperficially by means of reversible adsorption, aswell as their capacity to temporarily bind largeramounts of water, all these properties arefundamental for us to consider clay sediments asa very active and effectively working “geochemi-cal factory”.

Clay minerals constitute a large part of the fam-ily of phyllosilicates. Their crystal structure ischaracterized by alternation of flat, parallel sheets,or layers of extreme thinness. For this reason clayminerals are called layer silicates. Two basic typesof layers, or sheets make up any given claymineral. One type of layer consists of tetrahedralsheets in which one silicon atom is surrounded byfour oxygen atoms in tetrahedral configuration.The second type of layer is composed ofoctahedron sheets in which aluminum or magne-sium is surrounded by hydroxyl groups and oxy-gen in a 6-fold coordinated arrangement (Fig. 1.4).Depending on the clay mineral under study, thereis still enough space for other cations possessinga larger ionic radius, like potassium, sodium,calcium, or iron to fit in the gaps between theoctahedrons and tetrahedrons. Some clay minerals- the so-called expanding or swelling clays - havea special property which allows them to incorpo-rate hydrated cations into their structure. Thisprocess is reversible; the water changes the



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volume of the clay particles significantly as itgoes into or out of the clay structure. All in all,hydration can vary the volume of a clay particleby 95 %.

Kaolinite is the most important clay mineral ofthe two-layer group, also referred to as the 7-Ång-strom clay minerals (Fig. 1.4), which consist ofinterlinked tetrahedron-octahedron units. Illiteand the smectites which have the capacity to bindwater by swelling belong to the group of three-layered minerals, also referred to as 10-Ångstromclay minerals. They are made of a combination oftwo tetrahedrally and one octahedrally coordi-

nated sheets. Four-layer clay minerals, also knownas 14-Ångstrom clay minerals, arise whenever a fur-ther autonomous octahedral layer emerges betweenthe three-layered assemblies. This group comprisesthe chlorites and an array of various composites.

Apart from these types of clay minerals, thereis a relatively large number of clay minerals thatpossess a mixed-layered structure made up of acomposite of different basic structures. The resultis a sheet by sheet chemical mixture on the scaleof the crystallite. The most frequent mixed-layerstructure consists of a substitution of illite andsmectite layers.

relativeimportance idealized composition

Quarz +++ SiO2

Calcite + CaCO3

Dolomite + (Ca,Mg)CO3

Feldspars Plagioclase ++ (Na,Ca)[Al(Si,Al)Si2O8] Orthoclase ++ K[AlSi3O8]Muscovite ++ KAl2[(AlSi3)O10](OH)2

C l a y m i n e r a l sKaolinite +++ Al2Si2O5(OH)4Mica Group e.g. Illite +++ K0.8-0.9(Al,Fe,Mg)2(Si,Al)4O10(OH)2Chlorite Groupe e.g. Chlorite s.s. +++ (Mg3-yAl1Fey )Mg3(Si4-xAl)O10(OH)8Smektite Groupe e.g. Montmorillonite +++ Na0.33(Al1.67Mg0.33)Si4O10(OH)2•nH2O

H e a v y m i n e r a l s , e. g.Amphiboles e.g. Hornblende + Ca2(Mg,Fe)4Al[Si7,Al22](OH)2Pyroxene e.g. Augite + (Ca,Na)(Mg,Fe,Al)[(Si,Al)2O6]Magnetite - Fe3O4

Ilmenite - FeTiO3

Rutile - TiO2

Zircon - ZrO2

Tourmaline - (Na,Ca)(Mg,Fe,Al,Li)3Al6(BO3)3Si6O18(OH)4Garnet e.g. Grossular - Ca3Al2(SiO4)3

Table 1.1 Mineralogy and relative importance of main lithogeneous sediment components.

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Kaolinite is a regularly structured di-octahedraltwo-layer mineral and preferentially developsunder warm and humid conditions, by chemicalweathering of feldspars in tropical soil. Gooddrainage is essential to assure the removal of cat-ions released during hydrolysis. Abundance anddistribution of kaolinite reflects soil-forming pro-cesses in the area of its origin which is optimal inlateritic weathering in the tropics. Owing to thefact that the occurrence of this mineral in oceansediments is distinctly latitude-dependent, it is of-ten referred to as the “mineral of low latitudes”.

Illite is a three-layered mineral of the micagroup and not really a specified mineral, insteadthe term illite refers to a group of mica-like miner-als in the clay fraction; as such, it belongs to themost frequently encountered type of clay miner-als. Illites are formed as detrital clay minerals byfragmentation in physical weathering. Chemicalweathering (soil formation) in which potassium isreleased from muscovite also leads to theformation of illites. For this reason, illite is oftenreferred to as incomplete mica or hydromica. Thedistribution of illites clearly reflects its terrestrialand detrital origin which is also corroborated byK/Ar-age determinations made on illites obtainedfrom recent sediments. There are as yet no indica-tions as to in-situ formations of illites in marineenvironments.

Similar to the illites are the cation-rich, ex-pandable, three-layered minerals of the smectitegroup. The smectites, a product of weathering andpedogenic formation in temperate and sub-aridzones, hold an intermediate position with regardto their global distribution. However, smectites areoften considered as indicative of volcanic envi-ronments, in fact smectite formation due to low-temperature chemical alteration of volcanic rocksis even a quite typical finding. Similarly, finelydispersed particles of volcanic glass may trans-form into smectite after a sufficiently longexposure to seawater. Smectites may also arisefrom muscovite after the release of potassium andits substitution by other cations. There is conse-quently no distinct pattern of smectite distributiondiscernible in the oceans.

The generally higher occurrence of smectiteconcentrations in the southern hemisphere can beexplained with the relatively higher input of volca-nic detritus. This is especially the case in theSouthern Pacific.

Chlorite predominately displays a trioctahedralstructure and is composed of a series of threelayers resembling mica with an interlayered sheetof brucite (hydroxide interlayer). Chlorite is mainlyreleased from altered magmatic rocks and frommetamorphic rocks of the green schist facies as aresult of physical weathering. It therefore charac-


7 A Kaolinite

10 A SmectiteIllite

14 A Chlorite

Exchangeable cations nH2O

Oxygens Hydroxyls Aluminium, iron, magnesium

and Silicon, occasionaly aluminium

tetrahedra

octahedra

hydroxyoctahedra

tetrahedra

tetrahedra

tetrahedra

tetrahedra

octahedra

octahedra

Fig. 1.4 Schematic diagram of clay mineral types: Left: According to the combination of tetrahedral- and octahedral-coordinated sheets; Right: Diagrammatic sketch of the structure of smectite (after Hillier 1995 and Grim 1968).


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teristically depends on the type of parent rock.On account of its iron content, chlorite is proneto chemical weathering. Chlorite distinctly dis-plays a distribution pattern of latitudinal zona-tion and due to its abundance in polar regions itis considered as the “mineral of high latitudes”(Griffin et al. 1968). The grain size of chlorite min-erals is – similar to illite – not limited to the clayfraction (< 4 µm), but in addition encompassesthe entire silt fraction (4 - 63 µm) as well.

1.2.2 Biogenous Sediments

Biogenous sediments generally refer to bioclasticsediments, hence sediments which are built ofremnants and fragments of shells and tests pro-duced by organisms – calcareous, siliceous or

phosphatic particles. In a broader sense, bio-genous sediments comprise all solid materialformed in the biosphere, i.e. all the hard parts in-clusive of the organic substance, the causto-bioliths. The organic substance will be treatedmore comprehensively in Chapter 4, therefore theywill not be discussed here.

The amount of carbonate which is deposited inthe oceans today is almost exclusively of bio-genous origin. The long controversy whetherchemical precipitation of lime occurs directly inthe shallow waters of the tropical seas, such asthe banks of the Bahamas and in the Persian Gulf(Fig. 1.5), during the formation of calcareousooids and oozes of acicular aragonite, has beensettled in preference of the concept of biomin-eralization (Fabricius 1977). The tiny aragonite

Fig. 1.5 SEM photographs of calcareous sediments composed of aragonite needles which are probably of biogenic origin.Upper left: slightly etched section of a calcareous ooid from the Bahamas showing subconcentric laminae of primary ooidcoatings; upper right: close-up showing ooid laminae formed by small acicular aragonite needles. Lower left: silt-sizedparticle of aragonite mud from the Persian Gulf; lower right: close-up showing details of acicular aragonite needles measuring upto 10 µm in length, scalebar 5 µm.

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needles (few micrometers) within the more or lessconcentric layers of calcareous ooids have beenconsidered for a long time as primary precipitatesfrom seawater. However, further investigations in-cluding the distribution of stable isotopes dis-tinctly evidenced their biological origin as prod-ucts of calcification by unicellular algae. Yet, onepart of the acicular aragonite ooze might still origi-nate from the mechanical disruption of shells andskeletal elements.

Although marine plants and animals are numer-ous and diverse, only relatively few groups pro-duce hard parts capable of contributing to theformation of sediments, and only very few groupsoccur in an abundance relevant for sediment for-mation (Table 1.2). Relevant for sediment forma-tion are only carbonate minerals in the form of ara-gonite, Mg-calcite and calcite, as well as biogenic

opal in the form of amorphous SiO2·nH2O. Thesulfates of strontium and barium as well asvarious compounds of iron, manganese, and alu-minum are of secondary importance, yet they areof geochemical interest, e.g. as tracers for the re-construction of past environmental conditions.For example, the phosphatic particles formed byvarious organisms, such as teeth, bone, and shellsof crustaceans, are major components of phos-phorite rocks which permit us to draw conclusionsabout nutrient cycles in the ocean.

Large amounts of carbonate sediment accumu-late on the relatively small surface of the shallowshelf seas – as compared to entire oceans surface– of the tropical and subtropical warm water re-gions, primarily by few lime-secreting benthicmacrofossil groups. Scleritic corals, living in sym-biosis with algae, and encrusting red algae consti-


Aragonite Aragonite Mg-Calcite Calcite Calcite + Opal divers+ Calcite Mg-Calcite

Plankton Pteropods x Radiolarians x celestite Foraminifera (x) x x x Coccolithophores x Dinoflagellates x organic Silicoflagellates x Diatoms x

Benthos Chlorophyta x Rhodophyta x x (x) Phaeophyta x Sponges x x x celestite Scleratinian corals x Octocorals x x Bryozoens x x Brachiopods x phosphate Gastropods x x Pelecypods x x x Decapods x phosphate Ostracods (x) x Barnacles (x) x Annelid worms x x x (x) phosphate Echinoderms x phosphate Ascidians x

Table 1.2 Major groups of marine organisms contributing to biogenic sediment formation and mineralogy of skeletalhard parts. Foraminifera and diatoms (underlined) are important groups of both plankton and benthos. x = common,(x) = rare (mainly after Flügel 1978, 2004 and Milliman 1974).


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tute the major proportion of the massivestructures of coral reefs. Together with calcifyinggreen algae, foraminifera, and mollusks, these or-ganisms participate in a highly productive ecosys-tem. Here, coarse-grained calcareous sands andgravel are essentially composed of various bio-clasts attributable to the reef structure, of lime-se-creting algae, mollusks, echinoderms and largeforaminifera.

Fine-grained calcareous mud is produced bygreen algae and benthic foraminifera as well as bythe mechanical abrasion of shells of themacrobenthos. Considerable amounts of sedimentare formed by bioerosion, through the action ofboring, grazing and browsing, and predating or-ganisms. Not all the details have been elucidatedas to which measure the chemical and biologicaldecomposition of the organic matter in biogenichard materials might lead to the formation of pri-mary skeletal chrystalites on the micrometer scale,and consequently contribute to the fine-grainedcalcareous mud formation.

It is obvious that the various calcareous-shelled groups, especially of those organisms whosecrete aragonite and Mg-calcite, contribute sig-nificantly to the sediment formation in the shallowseas, whereas greater deposits of biogenic opalare rather absent in the shallow shelf seas. Theisostatically over-deepened shelf region ofAntarctica, where locally a significant accumula-tion of siliceous sponge oozes occurs, however,makes a remarkable exception. The relatively lowopal concentration in recent shelf deposits doesnot result from an eventual dilution with terri-genous material. The reason is rather that recenttropical shallow waters have low silicate concen-trations, from which it follows that diatoms andsponges are only capable of forming slightlysilicified skeletons that quickly remineralize inmarkedly silicate-deficient waters.

With an increasing distance from the coastalareas, out toward the open ocean, the relevance ofplanktonic shells and tests in the formation ofsediments increases as well (Fig. 1.6). Planktoniclime-secreting algae and silica-secreting algae,coccolithophorids and diatoms that dwell as pri-mary producers in the photic zone which thick-ness measures approximately 100 m, as well as thecalcareous foraminifers and siliceous radiolarians,and silicoflagellates (Table 1.2), are the producersof the by far most widespread and essential deep-sea sediments: the calcareous and siliceousbiogenic oozes. Apart from the groups mentioned,

planktonic mollusks, the aragonite-shelledpteropods, and some calcareous cysts formingdinoflagellates, also contribute to a considerabledegree to sediment formation.

1.2.3 Hydrogenous Sediments

Hydrogenous sediments may be widely distrib-uted, but as to their recent quantity they are rela-tively insignificant. They will be briefly mentionedin this context merely for reasons of beingcomplete. According to Elderfield (1976), hydrog-enous sediments can be subdivided into “precipi-tates”, primary inorganic components which haveprecipitated directly from seawater, like sodiumchloride, and “halmyrolysates”, secondarycomponents which are the reaction products ofsediment particles with seawater, formed subse-quent to in-situ weathering, but prior to dia-genesis. Of these, manganese nodules give an ex-ample. In the scope of this book, these compo-nents are not conceived as being part of the “pri-mary” solid phase sediment, but as “secondary”authigenic formations which only emerge in thecourse of diagenesis, as for instance some clayminerals like glauconite, zeolite, hydroxides of ironand manganese etc. In the subsequent Chapters11 and 13 some aspects of these new formationswill be more thoroughly discussed.

The distinction between detrital and newlyformed, authigenic clay minerals is basically diffi-cult to make on account of the small grain size andtheir amalgamation with quite similar detritalmaterial. Yet it has been ascertained that the by farlargest clay mineral proportion – probably morethan 90 % – located in recent to subrecent sedi-ments is of detrital origin (Chamley 1989; Hillier1995). There are essentially three ways forsmectites to be formed, which demand specificconditions as they are confined to local areas. Al-terations produced in volcanic material is one way,especially by means of hydration of basaltic andvolcanic glasses. This process is referred to aspalagonitization. The probably best studiedsmectite formation consists in the vents of hydro-thermal solutions and their admixture with seawa-ter at the mid-oceanic mountain ranges. Shouldauthigenic clay minerals form merely in recent sur-face sediments in very small amounts, their fre-quency during diagenesis (burial diagenesis), willdemonstrate a distinct elevation. However, thisaspect will not be considered any further beyondthis point.

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1.3 Classification of MarineSediments

As yet, there is no general classification schemeapplicable to marine sediments that combines allthe essential characteristics pertaining to a sedi-ment. A large number of different schemes hasbeen proposed in the literature which focus eitheron origin, grain-size distribution, chemical andmineralogical features of the sediment compo-nents, or the facial development of the sediments– all depending on the specific problem under

study. The advantages and disadvantages of thevarious schemes will not be of any concern here.In the following, an attempt will be made to give acomprehensive concise summary, based on thecombination of the various concepts.

Murray and Renard (1891) early introduced abasically simple concept which in its essentialsdifferentiated according to the area of sedimentdeposition as well as to sediment sources. It dis-tinguished (i) “shallow-water deposits from low-water mark to 100 fathoms”, (ii) “terrigenous de-posits in deep and shallow water close to land”,i.e. the combined terrigenous deposits from the

Fig. 1.6 SEM photographs of the important sediment-forming planktonic organisms in various stages of decay. Above:Siliceous centered diatoms. Middle: Siliceous radiolarian. Below: Calcareous coccolithophorid of single placoliths =coccoliths and lutitic abrasion of coccoliths. Identical scale of 5 µm.

1.3 Classification of Marine Sediments


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deep sea and the shelf seas, and (iii) “pelagic de-posits in deep water removed from land”. Thisscheme is very much appropriate to provide the ba-sic framework for a simple classification scheme onthe basis of terrigenous sediments, inclusive of the“shallow-water deposits” in the sense of Murrayand Renard (1891), and the deep-sea sediments.The latter usually are subdivided into hemipelagicsediments and pelagic sediments.

1.3.1 Terrigenous Sediments

Terrigenous sediments, i.e. clastics consisting ofmaterial eroded from the land surface, are not onlyunderstood as nearshore shallow-water depositson the shelf seas, but also comprise the deltaicforeset beds of continental margins, slump depos-its at continental slopes produced by gravitytransport, and the terrigenous-detrital shelf sedi-ments redistributed into the deep sea by the activ-ity of debris flows and turbidity currents.

The sand, silt, and clay containing shelf sedi-ments primarily consist of terrigenous siliciclasticcomponents transported downstream by rivers;they also contain various amounts of autochtho-nous biogenic shell material. Depending on theavailability of terrigenous discharge, biogenic car-bonate sedimentation might predominate in broadshelf regions. A great variety of grain sizes istypical for the sediments on the continental shelf,with very coarse sand or gravel accumulating inhigh-energy environments and very fine-grainedmaterial accumulating in low-energy environ-ments. Coarse material in the terrigenous sedi-ments of the deep sea is restricted to debris flowdeposits on and near the continental margins andthe proximal depocenters of episodic turbidites.

The development and the hydrodynamic his-tory of terrigenous sediments is described in itsessentials by the grain-size distribution and thederived sediment characteristics. Therefore, aclassification on the basis of textural features ap-pears suitable to describe the terrigenous sedi-ments. The subdivision of sedimentary particlesaccording the Udden-Wentworth scale encom-passes four major categories: gravel (> 2 mm),sand (2 - 0.0625 mm), silt (0.0625 - 0.0039 mm) andclay (< 0.0039 mm), each further divided into anumber of subcategories (Table 1.3). Plotting thepercentages of these grain sizes in a ternary dia-gram results in a basically quite simple and clearclassification of the terrigenous sediments (Fig.1.7). Further subdivisions and classifications

within the various fields of the ternary diagram aremade quite differently and manifold, so that acommonly accepted standard nomenclature hasnot yet been established. The reason for this lies,to some extent, in the fact that variable amounts ofbiogenic components may also be present in thesediment, next to the prevalent terrigenous,siliclastic components. Accordingly, sedimentclassifications are likely to vary and certainly willbecome confusing as well; this is especially trueof mixed sediments. However, a certain degree ofstandardization has developed owing to the fre-quent and identical usage of terms descriptive ofsediment cores, as employed in the internationalOcean Drilling Program (ODP) (Mazullo and Gra-ham 1987).

Although very imprecise, the collective term“mud” is often used in literature to describe thetexture of fine-grained, mainly non-biogenicsediments which essentially consist of a mixtureof silt and clay. The reason for this is that thedifferentiation of the grain-size fractions, silt andclay, is not easy to manage and is also very time-consuming with regard to the applied methods.However, it is of high importance for the geneticinterpretation of sediments to acquire thisinformation. For example, any sediment mainlyconsisting of silt can be distinguished, such as adistal turbidite or contourite, from a hemipelagicsediment mainly consisting of clay. Theclassification of clastic sediments as proposedby Folk (1980) works with this distinction,providing a more precise definition of mud as aterm (Fig. 1.8).

Clay

Clay

Siltyclay

Sandy clay

Sand Silt

Clayey silt

Clayey sand

Siltysand

Sandy silt

Sand-silt-clay

Sand Silt

75

20 20

2075 75

Fig. 1.7 Ternary diagram of sand-silt-clay grain-size dis-tribution showing principal names for siliciclastic,terrigeneous sediments (from Shepard 1954).

13

It needs to be stated in this context, that, in orderto compare the quantitative reports of publishedgrain sizes, the international literature does not drawthe line between silt and clay at 0.0039 mm– as the U.S standard and the Udden-Wentworthscale does (Table 1.3), or the French AFNOR-norm, but sets the limit at 0.002 mm, a value veryoften found in German literature and complyingwith the DIN-standards. To add to further diver-sity, modern publications from Russia mark thesilt-clay transition at a grain size of 0.01 mm (e.g.Lisitzin 1996).

1.3.2 Deep-sea Sediments

The sediments in the deep sea consist of only fewbasic types which in their manifold combinationsare suited for the description of a varied facialpattern (Table 1.4). The characteristic pelagicdeep-sea sediment far from coastal areas is deep-sea red clay, an extremely fine-grained (median< 1 µm) red-brown clay sediment which covers theoceanic deep-sea basins below the CalciteCompensation Depth (CCD). More than 90 % iscomposed of clay minerals, other hydrogenousminerals, like zeolite, iron-manganese precipi-tates and volcanic debris. Such sediment com-position demonstrates an authigenic origin. The


Udden-Wenthworth Terminologymm phi values (φ )

1024 -10 Boulder512 -9256 -8128 -7 Cobbles64 -632 -516 -4 Pebble8 -34 -2

3,36 -1,752,83 -1,5 Granule2,38 -1,252,00 -1,0

1,68 -0,751,41 -0,5 Very coarse sand1,19 -0,251,00 0,00,84 0,250,71 0,50 Coarse sand0,59 0,750,50 1,000,42 1,250,35 1,50 Medium sand0,3 1,750,25 2,000,21 2,25

0,177 2,50 Fine sand0,149 2,750,125 3,000,105 3,250,088 3,50 Very fine sand0,074 3,750,0625 4,00

0,053 4,250,044 4,50 Coarse silt0,037 4,750,031 5,00

5,255,50 Medium silt5,75

0,0156 6,006,256,50 Fine silt6,75

0,0078 7,007,257,50 Very fine silt7,75

0,0039 8,00

0,00200 9,00,00098 10,0 Clay0,00049 11,0

Table 1.3 Grain-size scales and textural classificationfollowing the Udden-Wentworth US Standard. The phivalues (φ) according to Krumbein (1934) and Krumbeinand Graybill (1965); φ = - log2 d/d0 where d0 is the stan-dard grain diameter (i.e. 1 mm).

10

50

90 S

cS mS zS

sC sM sZ

C M ZSiltClay

Sand

% Sand

2:1 1:2

S, sand ; s, sandyZ, silt ; z, siltyM, mud ; m, muddyC, clay ; c, clayey

Fig. 1.8 Textural classification of clastic sediments(modified after Folk 1980).


14

small percentage of lithogenic minerals, such asquartz, feldspar and heavy minerals, confirms theexistence of terrigenous components which in partshould have originated from eolian transportprocesses. The biogenic oozes represent the mostfrequent type of deep-sea sediments; they mainlyconsist of shells and skeletal material from plank-tonic organisms living in the ocean where theydrizzle from higher photic zones down to theocean floor, like continuous rainfall, once theyhave died (Fig. 1.6). The fragments of the calcare-ous-shelled pteropods, foraminifera, and cocco-lithophorids constitute the calcareous oozes (pter-opod ooze, foraminiferal ooze, or nannofossilooze), whereas the siliceous radiolarians, silico-flagellates, and diatoms constitute the siliceousoozes (radiolarian ooze or diatomaceous ooze).

The hemipelagic sediments are basically madeof the same components as the deep-sea red clay

and the biogenic oozes, clay minerals and bio-genic particles respectively, but they also containan additional and sometimes dominating amountof terrigenous material, such as quartz, feldspars,detrital clay minerals, and some reworked bio-genous components from the shelves (Table 1.4,Fig. 1.9).

Most deep-sea sediments can be described ac-cording to their composition or origin as a three-component system consisting of

(i) biogenic carbonate,(ii) biogenic opal, and(iii) non-biogenic mineral constituents.

The latter group comprises the components of thedeep-sea red clay and the terrigenous silici-clastics. On the basis of experiences made in theDeep Sea Drilling Project, Dean et al. (1985) have

Table 1.4 Classification of deep-sea sediments according to Berger (1974).

I. (Eu-)pelagic deposits (oozes and clays)< 25 % of fraction > 5µm is of terrigenic, volcanogenic, and/or neritic originMedian grain size < 5µm (except in authigenic minerals and pelagic organisms)A. Pelagic clays. CaCO3 and siliceous fossils < 30 % 1. CaCO3 1 - 10 %. (Slightly) calcareous clay 2. CaCO3 10 - 30 %. Very calcareous (or marl) clay 3. Siliceous fossils 1 - 10 %. (Slightly) siliceous clay 4. Siliceous fossils 10 - 30 %. Very siliceous clayB. Oozes. CaCO3 or siliceous fossils > 30 % 1. CaCO3 > 30 %. < 2/3 CaCO3: marl ooze. > 2/3 CaCO3: chalk ooze 2. CaCO3 < 30 %. > 30 % siliceous fossils: diatom or radiolarian ooze

II. Hemipelagic deposits (muds)> 25 % of fraction > 5 µm is of terrigenic, volcanogenic, and/or neritic originMedian grain size >5 µm (except in authigenic minerals and pelagic organisms)A. Calcareous muds. CaCO3 > 30 % 1. < 2/3 CaCO3: marl mud. > 2/3 CaCO3: chalk mud 2. Skeletal CaCO3 > 30 %: foram ~, nanno ~, coquina ~B. Terrigenous muds, CaCO3 < 30 %. Quarz, feldspar, mica dominant Prefixes: quartzose, arkosic, micaceousC. Volcanogenic muds. CaCO3 < 30 %. Ash, palagonite, etc., dominant

III. Special pelagic and/or hemipelagic deposits 1. Carbonate-sapropelite cycles (Cretaceous). 2. Black (carbonaceous) clay and mud: sapropelites (e.g., Black Sea) 3. Silicified claystones and mudstones: chert (pre-Neogene) 4. Limestone (pre-Neogene)

15

Fig. 1.9 Hemipelagic sediment from Sierra Leone Rise, tropical North Atlantic. Left: Fine silt-size fraction composedof coccoliths (c), foraminiferal fragments (f) and of detrital quartz (q) and mica (m). Right: Coarse silt-sized fractionpredominately composed of foraminiferal fragments (f), some detrital quartz (q) and mica (m), scalebar 10 µm.

Fig. 1.10 Classification of deep-sea sediments according to the main constituents, e.g. clay (non-biogenic), diatoms(siliceous biogenic), and nannofossils (calcareous biogenic); (modified from Dean et al. 1985).



16

developed a very detailed and purely descriptiveclassification scheme (Fig. 1.10):

• The most frequently occurring componentwith a percentage higher than 50 % deter-mines the designation of the sediment. Non-biogenic materials are specified accordinglyon the basis of the grain-size fractions: sand,silt or clay. Biogenic material is referred to as“ooze” and preceded by the most abundantbiogenic component: nannofossil ooze, fora-minifera ooze, diatom ooze, and radiolarianooze respectively.

• Each component measuring between 25-50 %is characterized by the following attributes:sandy, silty, clayey, or nannofossil, foramini-feral, diatomaceous, or radiolarian.

• Components with percentages between 10-25 %are referred to by adding the suffix “-bearing”,as in “clay-bearing”, “diatom-bearing”.

• Components with percentages below 10 % arenot expressed at all, but may be included byaddition of the suffix “rich”, as in “Corg-rich”.

The thus established, four-divided nomencla-ture of deep-sea sediments with threshold limits of10 %, 25 % and 50 % easily permits a quite de-tailed categorization of the sediment which isadaptable to generally rare, but locally frequentoccurrences of components, e.g. zeolite, eventu-ally important for a more complete description.

1.4 Global Patterns ofSediment Distribution

The overall distribution pattern of sediment typesin the world’s oceans depends on few elementaryfactors. The most important factor is the relativeamount with which one particle species contrib-utes to sediment formation. Particle preservationand eventual dilution with other sediment compo-nents will modify the basic pattern. The formationand dispersal of terrigenous constituents derivedfrom weathering processes on the continents, aswell as autochthonous oceanic-biogenic constitu-ents, both strongly depend on the prevalent cli-mate conditions, so that, in the oceans, a latitude-dependent and climate-related global pattern ofsediment distribution will be the ultimate result.

1.4.1 Distribution Patterns ofShelf Sediments

The particulate terrigenous weathering productsmainly transported from the continents by riversare not homogeneously distributed over the oceanfloor, but concentrate preferentially along the con-tinental margins, captured either on the shelf orthe continental slope (Fig. 1.11). Massive sedi-ment layers are built where continental inputs areparticularly high, and preferentially during glacialperiods when sea levels were low.

Approximately 70 % of the continental shelfsurface is covered with relict sediment, i.e. sedi-ment deposited during the last glacial period un-der conditions different from today’s, especially attimes when the sea level was comparatively low(Emery 1968). It has to be assumed that there is akind of textural equilibrium between these relictsediments and recent conditions. The fine-grainedconstituents of shelf sediments were elutedduring the rise of the sea level in the Holoceneand thereafter deposited, over the edge of theshelf onto the upper part of the continental slope,so that extended modern shelf surface areasbecame covered with sandy relict sediment(Milliman et al. 1972; Milliman and Summerhays1975).

According to Emery (1968), the sediment distri-bution on recent shelves displays a plain and dis-tinctly zonal pattern (Fig. 1.12):

Biogenic sediments with coarse-grained calcare-ous sediments predominate at lower latitudes,

Detrital sediments with riverine terrigenoussiliciclastic material at moderate latitudes, and

Glacial sediments of terrigenous origin trans-ported by ice are limited to high latitudes.

In detail, this well pronounced pattern may be-come strongly modified by local superimpositions.Coarse-grained biogenic carbonate sediments willbe found at moderate and high latitudes as well, atplaces where the riverine terrigenous inputs arevery low (Nelson 1988).

Today, as a result of the post-glacial high sealevels, most river-transported fine-grained materialis deposited in the estuaries and on the flat innershelves in the immediate proximity of rivermouths. Only a small proportion is transportedover the edge of the shelf onto the continentalslope. These processes account for the develop-ment of mud belts on the shelf, of which 5 types

17

of shelf mud accumulation can be distinguished(McCave 1972, 1985): “muddy coasts, nearshore,mid- and outer-shelf mud belts and mud blankets’(Fig. 1.13). Mud belts depend on the amount ofdischarged mud load, the prevalent tidal and/orcurrent system, or the distribution of the sus-

pended sediment. Muddy coasts will preferentiallyform near river mouths, whereas mid-shelf mudbelts are characteristic of regions where wave andtidal activity are relatively lower than on the inneror outer shelf.

Especially in delta regions where the supplyrates of terrigenous material are high – as in thetropics – even the entire shelf might become cov-ered with a consistent blanket of mud, althoughthe shelf represents a region of high energy con-version.

1.4.2 Distribution Patterns ofDeep-sea Sediments

The two most essential boundary conditions inpelagic sedimentation are the nutrient content inthe surface water which controls biogenic produc-tivity and by this biogenic particle production,and the position of the calcite compensationdepth (CCD) controlling the preservation ofcarbonate. The CCD, below which no calcite isfound, describes a level at which the dissolutionof biogenic carbonate is compensated for by itssupply rate. The depth of the CCD is generallysomewhere between 4 and 5 km below the surface,however, it varies rather strongly within the threegreat oceans due to differences in the water massand the rates of carbonate production.

Calcareous ooze and pelagic clays are the pre-dominant deep-sea sediments in offshore regions

Spreading Center Pacific Equatorial Bulge, 0.5 km >1 Km <1 Km

miles

km

0 2000

0 3000

north pole

GLACIAL

WATER-CONTRIBUTEDDETRITAL

AUTHIGENIC

BIOGENICequator

south pole

CONT

INEN

TS

CONT

INEN

TS

0°C

20°C

0°C

20°C

Fig. 1.11 Sediment thickness to acoustic basement in the world ocean (from Berger 1974).

Fig. 1.12 Latitudinal distribution of sedimentary faciesof the shallow marine environment of continentalshelves in an idealized ocean. Bold arrows = cold water;light arrows = warm water; grey arrows = upwelling water(modified from Reineck and Singh 1973).

1.4 Global Patterns of Sediment Distribution


18

(Table 1.5). The distribution of these sediments inthe three great oceans shows a considerable de-gree of variation (Fig. 1.14). The distribution pat-terns strongly depend on the water depth, i.e. theposition of the CCD. Calcareous ooze, primarilyconsisting of foraminiferal oozes and nanno-plankton oozes, covers vast stretches of the sea-

floor at water depths less than 3-4 km androughly retraces the contours of the mid-oceanicridges as well as other plateaus and islands,whereas pelagic clay covers the vast deep-seaplains in the form of deep-sea red clay. This par-ticular pattern is especially obvious in the Atlan-tic Ocean.

muddy coast

1 2

nearshore mud-belt

3 4

mid-shelfmud-belt

outer-shelf mud-belt

5

mud blanket (off delta)*

S H E L F E D G E

* or under advective mud stream

C O A S T

Fig. 1.13 Schematic representation of modern mud accumulation on continental shelves (modified from McCave(1972)).

Table 1.5 Relative areas of world oceans covered with pelagic sediments; area of deep-sea floor = 268.1·106 km2

(from Berger (1976)).

Sediments (%) Atlantic Pacific Indian World

Calcareous ooze 65,1 36,2 54,3 47,1Pteropod ooze 2,4 0,1 --- 0,6Diatom ooze 6,7 10,1 19,9 11,6Radiolarian ooze --- 4,6 0,5 2,6Red clays 25,8 49,1 25,3 38,1

Relative size of ocean (%) 23,0 53,4 23,6 100,0

19

Siliceous oozes, mostly consisting of diatomooze, form a conspicuous ring around Antarcticawhich clearly marks the zone of the Polar Front inthe Antarctic Circumpolar Current. A broad bandof radiolarian ooze covers the Pacific Oceanbelow the equatorial upwelling zone, whilst dia-tom ooze covers the oceanic margins of theNorthern Pacific. Terrigenous sediments, espe-cially in the form of mass flow deposits and tur-bidites cover vast stretches of the near-continen-tal zones in the North Atlantic, the northeasternPacific, and the broad deep-sea fans off the bigriver mouths in the northern Indian Ocean.Glacio-marine sediments are restricted to thecontinental margins of Antarctica and to the highlatitudes of the North Atlantic.

1.4.3 Distribution Patterns ofClay Minerals

Apart from the pelagic clays of the abyssal plainsand the terrigenous sediments deposited along

the continental margins, primarily or exclusivelyconsisting of clay minerals, various amounts oflithogenous clay minerals can be found in alltypes of ocean sediments (Table 1.6). The relativeproportion of the various clay minerals in the sedi-ments is a function of their original source, theirmode of transport into the area of deposition –either by eolian or volcanic transport, or by meansof water and ice – and finally the route oftransportation (Petschick et al. 1996). In a globalsurvey, it is easy to identify the particular interac-tions which climate, weathering on the continents,wind patterns, riverine transport, and oceaniccurrents have with regard to the relative distribu-tion of the relevant groups of clay minerals(kaolinite, illite, smectite, and chlorite).

The distribution of kaolinite in marine sedi-ments (Fig. 1.15) depends on the intensity ofchemical weathering at the site of the rock’s originand the essential patterns of eolian and fluvialtransport. Due to its concentration at equatorialand tropical latitudes, kaolinite is usually referred

Fig, 1.14 Distribution of dominant sediment types on the present-day deep-sea floor. The main sediment types aredeep-sea clay and calcareous oozes which patterns are predominately depth-controlled. (from Davies and Gorsline(1976)).



20

to as the “clay mineral of low latitudes” (Griffin etal. 1968).

Illite is the most frequent clay mineral to befound in ocean sediments (Fig. 1.16). It demon-strates a distinctly higher concentration in sedi-ments at mid-latitudes of the northern oceanswhich are surrounded by great land masses. Thisfollows particularly from its terrigenous origin andbecomes evident when the Northern Pacific iscompared with the Southern Pacific. The illite con-centration impressively reflects the percentageand distribution of particles which were intro-

duced into marine sediments by fluvial transport.The predominance of illites in the sediments of thePacific and Atlantic Oceans at moderate latitudes,below the trajectories of the jet-stream, indicatesthe great importance of the wind system in thetransport of fine-dispersed particulate matter.

The distribution pattern of smectite differsgreatly in the three oceans (Fig. 1.17), and alongwith some other factors may be explained as an ef-fect induced by dilution. Smectite is generallyconsidered as an indicator of a “volcanic regime”(Griffin et al. 1968). Thus, high smectite concentra-

Kaolinite Illite Smectite Chlorite

North Atlantic 20 56 16 10 Gulf of Mexico 12 25 45 18 Caribbian Sea 24 36 27 11South Atlanic 17 47 26 11North Pacific 8 40 35 18South Pacific 8 26 53 13Indian Ocean 16 30 47 10 Bay of Bengal 12 29 45 14 Arabian Sea 9 46 28 18

Table 1.6 Average relative concentrations [wt.%] of the principal clay mineral groups in the < 2 �m carbonate-freefraction in sediments from the major ocean basins (data from Windom 1976).

Fig. 1.15 Relative distribution of kaolinite in the world ocean, concentration in the carbonate-free < 2 µm size frac-tion (from Windom 1976).

21

tions are usually observed in sediments of theSouthern Pacific, in regions of high volcanic activ-ity, where the sedimentation rates are very lowdue to great distances from the shoreline, andwhere the dilution with other clay minerals is low

Fig. 1.16 Relative distribution of illite in the world ocean, concentration in the carbonate-free < 2 µm size fraction(from Windom 1976).

Fig. 1.17 Relative distribution of smectite in the world ocean, concentration in the carbonate-free < 2 µm sizefraction (from Windom 1976).


as well. The low smectite concentration in theNorth Atlantic results from terrigenous detritus in-puts which are rich in illites and chlorites.

The distribution of chlorite in deep-sea sedi-ments (Fig. 1.18) is essentially inversely related to


22

Fig. 1.18 Relative distribution of chlorite in the world ocean, concentration in the carbonate-free < 2 µm sizefraction (from Windom 1976).

Table 1.7 Sedimentation rates of red clay in variousdeep-sea basins of the world ocean (data from varioussources, e.g. Berger 1974, Gross 1987).

the pattern of kaolinite. Although chlorite is dis-tributed homogeneously over the oceans, its high-est concentration is measured in polar regions andtherefore is referred to as the “high latitudemineral” (Griffin et al. 1968).

1.4.4 Sedimentation Rates

As can be seen in Table 1.8, the sedimentationrates of typical types of deep-sea sediments showa strong geographical variability which is basedon the regionally unsteady import of terrigenousmaterial and a highly variable biogenic productiv-ity in the ocean.

Basically, it can be stated that the sedimenta-tion rate decreases with increasing distance from asediment source, may this either be a continent oran area of high biogenic productivity. The highestrates of terrigenous mud formation are recordedon the shelf off river mouth’s and on thecontinental slope, where sedimentation rates canamount up to several meters per one thousandyears. Distinctly lower values are observed at de-tritus-starved continental margins, for example ofAntarctica. The lowest sedimentation rates everrecorded lie between 1 and 3 mm ky-1. and areconnected to deep-sea red clay in the offshoredeep-sea basins (Table 1.7), especially in the cen-tral Pacific Ocean.

Calcareous biogenic oozes demonstrate in-termediate rates which frequently lie between 10and 40 mm ky-1. Their distribution pattern de-pends on the biogenic production and on the wa-ter depth, or the depth of the CCD. As yet, ratevalues between 2 and 10 mm ky-1. were con-sidered as normal for the sedimentation of sili-ceous oozes. Recent investigations in the regionof the Antarctic Circumpolar Current have re-vealed that a very high biogenic production, inconnection with lateral advection and “sedimentfocusing”, can even give rise to sedimentationrates of more than 750 mm per thousand years(Fig. 1.19).

Rate (mm ky-1)Mean Range

North Atlantic 1,8 0.5-6.2South Atlantic 1,9 0.2-7.5North Pacific 1,5 0.4-6.0South Pacific 0,45 0.3-0.6

23

Table 1.8 Typical sedimentation rates of recent and subrecent marine sediments (data from various sources, e.g.Berger 1974, Gross 1987).


Fig. 1.19 Distribution of major sediment facies across the frontal system of the Antarctic Circumpolar Current(ACC) between Africa and Antarctica. Numbers are typical sedimentation rates in mm ky-1. PF = Polar Front, SAF =Sub Antarctic Front, STF = Subtropical Front.

sandymud

foraminiferalmud

diatomooze

diatommud

diatomforaminiferal

ooze

foraminiferalnannofossil

mud

foraminiferalooze

sandymud

65 25 25

15

15

907505

AgulhasRidge

PF SAF STFAntarctica Africa

0

1

2

3

4

5

Dep

th [k

m]

0

1

2

3

4

5

CCD

2

deep-seamud

Weddell Abyssal Plain Cape Basin

MeteorRise

Atlanic-IndianRidge

AgulhasRidge

18

Fa cie s Are a Ave ra geS e dim e nta tion Ra te

(m m ky-1)

Terrigeneous m ud California Borderland 50 - 2,000Ceara A bys sal P lain 200A ntarc tic Continental M argin 30-65

Calcareous ooze North A tlantic (40-50 °N) 35-60North A tlantic ( 5-20 °N) 40-14E quatorial A tlantic 20-40Caribbean ~28

E quatorial Pac ific 5-18E as tern E quatorial Pac ific ~30E as t P ac ific Rise (0-20 °N) 20-40E as t P ac ific Rise (~ 30 °N) 3-10E as t P ac ific Rise (40-50 °N) 10-60

S iliceous ooze E quatorial Pac ific 2-5A ntarc tic , Indian S ec tor 2-10A ntarc tic , A tlantic S ec tor 25-750

Red c lay Northern North P ac ific (m uddy) 10-15Central North P ac ific 1-2Tropic al North Pac ific 0-1S outh P ac ific 0.3-0.6A ntarc tic , A tlantic S ec tor 1-2


24

1.5 Problems

Problem 1

Is there any correlation between the distributionof calcareous and siliceous sediments, terrigenoussediments and deep-sea clay?

Problem 2

Explain the distribution pattern of clay mineralskaolinite, illite, chlorite and smectite in the worldocean. Why should be illite more common in theNorth Atlantic than in the South Atlantic?

Problem 3

What are the reasons why sediments of theAtlantic Ocean are thicker along the continentalmargin than near the mid-Atlantic ridge?

Problem 4

What causes the variability of sedimentationrates?

Problem 5

Summarize and explain the main aspects of deepsea sediment classification.

References

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Berger, W.H. and Herguera, J.C., 1991. Reading the sedi-mentary record of the ocean’s productivity. Pimaryproductivity and biogeochemical cycles in the sea. In:Falkowski, P.G. and Woodhead, E.D. (eds) PlenumPress, New York, London, pp 455-486.

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Flügel, E., 2004. Microfacies of carbonate rocks:analysis, interpretation and application. SpringerVerlag, Berlin, Heidelberg, New York, 976 pp.

Folk, R.L., 1980. Petrology of sedimentary rocks.Hemphill. Publ. Co., Austin, 182 pp.

Garrels, R.M. and Mackenzie, F.T., 1971. Evolution ofsedimentary rocks. Norton & Co, New York, 397 pp.

Gibbs, R.J., 1977. Transport phase of transition metals inthe Amazon and Yukon rivers. Geological Society ofAmerica Bulletin, 88: 829-843.

Gingele, F., 1992. Zur klimaabhängigen Bildung biogenerund terrigener Sedimente und ihre Veränderungendurch die Frühdiagnese im zentralen und östlichenSüdatlantic. Berichte, Fachbereich Geowissenschaften,Universität Bremen, 85, 202 pp.

Velocity of sedimentation or deposition ofmaterial through time, i.e. sedimentation rate inthis chapter is given solely as millimeter per 1000years (mm/1000 y or mm ky-1). This describes thevertical thickness of sediment deposited in acertain period of time. It assumes more or lesslinear and constant processes – not regarding anyinconsistency hiatuses – between two datumlevels and is, therefore, often referred to as linearsedimentation rate. This term has to be dis-tinguished from mass accumulation rate whichrefers to a sediment mass deposited in a certainperiod of time within a defined area, usuallygiven as grams per centimeter square per 1000years (g cm-2 ky-1). For further reading see Brunsand Hass (1999).

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Goldberg, E.D. and Griffin, J.J., 1970. The sediments of thenorthern Indian Ocean. Deep-Sea Research, 17: 513-537.

Griffin, J.J., Windom, H. and Goldberg, E.D., 1968. Thedistribution of clay minerals in the World Ocean.Deep-Sea Research, 15: 433-459.

Grim, R.E., 1968. Clay Mineralogy, McGraw-Hill Publ.,NY, 596 pp.

Gross, M.G., 1987. Oceanography: A view of the Earth.Prentice Hall, Englewood Cliffs, NJ, 406 pp.

Hillier, S., 1995. Erosion, sedimentation and sedimentary ori-gin of clays. In: Velde, B. (ed) Origin and mineralogy ofclays. Springer Verlag, Berlin, Heidelberg, New York, pp162-219.

Krumbein, W.C., 1934. Size frequency distribution ofsediments. Journal of Sediment Petrology, 4: 65-77.

Krumbein, W.C. and Graybill, F.A., 1965. An introductionto statistical models in Geology. McGraw-Hill, NewYork, 328 pp.

Lisitzin, A.P., 1996. Oceanic sedimentation: Lithology andGeochemistry. Amer. Geophys. Union, Washington, D.C., 400 pp.

Mazullo, J. and Graham, A.G., 1987. Handbook forshipboard sedimentologists. Ocean Drilling Program,Technical Note, 8, 67 pp.

McCave, I.N., 1972. Transport and escape of fine-grainedsediment from shelf areas. In: Swift, D.J.P., Duane, D.B.and Pilkey, O.H. (eds), Shelf sediment transport, processand pattern. Dowden, Hutchison & Ross, Stroudsburg, pp225-248.

Milliman, J.D., Pilkey, O.H. and Ross, D.A., 1972.Sediments of the continental margins off the easternUnited States. Geological Society of America Bulletin,83: 1315-1334.

Milliman, J.D., 1974. Marine carbonates. Springer Verlag,Berlin, Heidelberg, New York, 375 pp.

Milliman, J.D. and Summerhays, C.P., 1975. Upper conti-nental margin sedimentation of Brasil. Contributionto Sedimentology, 4, Schweizerbart, Stuttgart, 175pp .

Milliman, J.D. and Meade, R.H., 1983. World-wide deliveryof river sediment to the ocean. The Journal of Geology,91: 1-21.

Milliman, J.D. and Syvitski, J.P.M., 1992. Geomorphic/tectonic control of sediment discharge to the ocean:The importance of small mountainous rivers. TheJournal of Geology, 100: 525-544.

Murray, J. and Renard, A.F., 1891. Deep sea deposits - Re-port on deep sea deposits based on specimens collectedduring the voyage of H.M.S. Challenger in the years1873-1876. „Challenger“ Reports, Eyre & Spottiswood,London; J. Menzies & Co, Edinburgh; Hodges, Figgis &Co, Dublin, 525 pp.

Nelson, C.S., 1988. Non-tropical shelf carbonates - Modernand Ancient. Sedimentary Geology, 60, 1-367 pp.

Petschick, R., Kuhn, G. and Gingele, F., 1996. Claymineral distribution in surface sediments of the SouthAtlantic: sources, transport, and relation tooceanography. Marine Geology, 130: 203-229.

Prospero, J.M., 1996. The atmospheric transport ofparticles in the ocean. In: Ittekkot, V., Schäfer, P.,

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Shepard, F., 1954. Nomenclature based on sand-silt-clayratios. Journal of Sediment Petrology, 24: 151-158.

Straub, S.M. and Schmincke, H.U., 1998. Evaluating thetephra input into Pacific Ocean sediments:distribution in space and time. GeologischeRundschau, 87: 461-476.

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