当前位置：文档库 › 3 Sedimentary Geochemistry

3 Sedimentary Geochemistry

Chapter3

Sedimentary Geochemistry How Sediments are Produced Knut Bj?rlykke

The composition and physical properties of sedimen-tary rocks are to a large extent controlled by chemical processes during weathering,transport and also during burial(diagenesis).We can not avoid studying chemi-cal processes if we want to understand the physical properties of sedimentary rocks.Sediment transport and distribution of sedimentary facies is strongly in?u-enced by the sediment composition such as the content of sand/clay ratio and the clay mineralogy.The primary composition is the starting point for the diagenetic processes during burial.

We will now consider some simple chemical and mineralogical concepts that are relevant to sedimen-tological processes.

Clastic sediments are derived from source rocks that have been disintegrated by erosion and weather-ing.The source rock may be igneous,metamorphic or sedimentary.The compositions of clastic sediments are therefore the product of the rock types within the drainage basin(provenance),of climate and relief.The dissolved portion?ows out into the sea or lakes,where it is precipitated as biological or chemical sediments. Weathering and abrasion of the grains continues dur-ing transport and sediments may be deposited and eroded several times before they are?nally stored in a sedimentary basin.

After deposition sediments are also being subjected to mineral dissolution and precipitation of new mine-rals as a part of the diagenetic processes.For the most part we are concerned with reactions between minerals K.Bj?rlykke( )

Department of Geosciences,University of Oslo,Oslo,Norway e-mail:knut.bjorlykke@geo.uio.no and water at relatively low temperatures.At tempera-tures above200–250?C these processes are referred to as metamorphism which is principally similar in that unstable minerals dissolve and minerals which are thermodynamically more stable at certain temperatures and pressures precipitate.

At low temperatures,however,unstable minerals and also amorphous phases may be preserved for a long time and there may be many metastable phases.

Many of the reactions associated with the dissolu-tion and precipitation of minerals proceed so slowly that only after an extremely long period can they achieve a degree of equilibrium.

Reactions will always be controlled by thermody-namics and will be driven towards more stable phases. The kinetic reaction rate is controlled by temperature.

Silicate reactions are very slow at low temperature and this makes it very dif?cult to study them in the laboratory.

Biological processes often accompany the purely chemical processes,adding to the complexity.Bacteria have been found to play an important role in both the weathering and precipitation of minerals.Their chief contribution is to increase reaction rates,particularly during weathering.

In this chapter we shall examine the processes between water and sediments from a simple physical-chemical viewpoint.A detailed treatment of sediment geochemistry is however beyond the scope of this book.

K.Bj?rlykke(ed.),Petroleum Geoscience:From Sedimentary Environments to Rock Physics, DOI10.1007/978-3-642-02332-3_3,?Springer-Verlag Berlin Heidelberg2010

K.Bj?rlykke

105°

Fig.3.1The strong dipole of water molecules causes them to be attracted to cations which thereby become hydrated.Small cations will be most strongly hydrated and less likely to be adsorbed on a clay mineral with a negative charge

Water (H 2O)consists of one oxygen atom linked to two hydrogen atoms,with the H-O-H bonds forming an angle of 105?(Fig.3.1).The distance between the O and the H atoms is 0.96?,and between the hydrogen atoms 1.51?.Water molecules therefore have a strong dipole with a negative charge on the opposite side from the hydrogen atoms (Fig.3.1).This is why water has a relatively high boiling point and high viscosity,and why it is a good solvent for polar substances.Another consequence of this molecular structure is that water has a high surface tension,important for enabling par-ticles and organisms to be transported on its surface.The capillary forces which cause water to be drawn up through ?ne-grained soils are also a result of this high surface tension.

A number of concepts are particularly useful for describing and explaining geochemical processes:1.Ionic potential 2.Redox potential Eh 3.pH

4.Hydration of ions in water

5.Distribution coef?cients

Isotopes

3.1Ionic Potential

Ionic potential is a term introduced by V .M.Goldschmidt to explain the distribution of elements in sediments and aqueous systems.It must not be

confused with ionisation potential.Recent authors have proposed the term “hydropotential”for the con-cept,to avoid confusion.

Ionic potential (I.P.)may be de?ned as the ratio between the charge (valency)Z and the ionic radius R :

IP =

Z R

The ionic potential is an expression of the charge on the surface of an ion,i.e.its capacity for adsorbing ions.Small ions carrying a large charge have a high ionic potential while large ions with a small charge have a low ionic potential (see Fig.3.2).

Ions with low ionic potential are unable to break the bonds in the water molecule and therefore remain in solution as hydrated cations (e.g.Na +,K +).This means that the ion is surrounded by water molecules with their negative dipole towards the cation (Fig.3.1).

This is because the O–H bond is stronger than the bond which the cation forms with oxygen (M–O bond-ing,M =metal);this is particularly true of alkali metal ions (Group I)and most alkaline earth elements (Group II,I.P.<3).Metals with an ionic potential only slightly lower than that required to form M–O bonds,namely Mg 2+,Fe 2+,Mn 2+,Li +and Na +,will be the most strongly hydrated.The hydration strongly affects the chemical properties of the ion and its capacity to be adsorbed or enter into the crystal structure of a mineral.Since the ions are surrounded by water molecules,we can use the expression “hydrated radius”to describe the space occupied by the ion and its water molecules within a crystal structure (Fig.3.3).

If the M–O bond is approximately equal in strength to the O–H bond (I.P.3–12),the metal ion replaces one of the hydrogen atoms to form very low solu-bility compounds of the type M(OH)n (see Fig.3.2).Examples of these so-called hydroxides that we com-monly encounter in sedimentary rocks as a result of weathering are Fe(OH)3,Al(OH)3and Mn(OH)4.These hydroxides have very low solubility.

Ions with high ionic potential (>12)form an M–O bond that is stronger than the H–O bond,giving soluble anion complexes such as SO 4??,CO 3??,PO 43?and releasing both of the H +ions into solution.

This approach can be used to explain the behaviour for elements on both sides of the Periodic Table (electropositive and electronegative)which form ionic bonds.The elements in the middle,however,have a greater tendency to form covalent bonds in which the

3Sedimentary Geochemistry

2.0

1.5

1.0

0.5

a nion g .I o n i c r a d i u s

Charge (valency)

Fig.3.2Ionic radius and charge (valence)for some

geochem-

ically important

elements.Ions

with low

ionic potential

are soluble

as cations

(e.g Na +,K +)while ions with intermediate ionic potentials will bond with OH ?groups and have very low

solubility,forming hydrolysates (e.g Al(OH)3),Fe (OH)3).High ionic potentials make soluble cation complexes like CO ??3

and SO ??

.The ratio between these parameters -the ionic potential -can be used to explain their behaviour in nature

Li +, r = 0.6 ? R = 3.8 ?Na +, r = 0.95 ?R = 3.6 ?

K +, r = 1.33 ?

R = 3.3 ?

Rb +, r = 1.48 ?R = 3.2 ?

Cs +, r = 1.69 ?R = 3.2 ?

Mg ++, r = 0.65 ?R = 4.2 ? Ca ++, r = 1.0 ?R = 4.0 ?

Sr ++, r = 1.13 ?

R = 4.0 ?

Ba ++, r = 1.43 ?R = 3.0 ?

Fig.3.3Ionic radius (in ?ngstrom units)of hydrated and non-hydrated (“naked”)ions of alkali metals and alkaline-earth metals.The smaller ions have higher ionic potentials and form stronger bonds with water molecules so that they become hydrated.This hydration effect is reduced with increasing temperature

strength of the M-O bond is not merely a function of the valency and radius,and the picture becomes far more complex.The concept of ionic potential is never-theless still useful;we see that during weathering,elements with low ionic potential remain in solution along with the anionic complexes of metals and non-metals with high ionic potential.This is re?ected in the composition of seawater.The hydrolysates,on the other hand,become enriched on land as insolu-ble residues or through weathering (Al 3+,Fe 3+,Mn 4+,Ti 4+,etc.).Note also that Fe ++and Mn ++which occur in reducing environments have lower ionic potential and are much more soluble that Fe 3+and Mn 4+.

The most soluble ions remain in the seawater until they are precipitated as salt when seawater is concentrated during evaporation.In addition to the chlorides (i.e.NaCl,KCl),these are mainly salts of cations with low ionic potential,and of anions with high ionic potential,e.g.CaSO 4.2H 2O,Na 2CO 3and carbonates such as CaCO 3(calcite),CaMg(CO 3)2(dolomite)and MgCO 3(magnesite).The principle of ionic hydration and the size of the ionic radius are capable of explaining a whole range of geochemical phenomena.Among the Group I ele-ments of the Periodic Table,we know that Li +and Na +are enriched in seawater.This because the strong hydration prevents absorption on clay minerals which usually have a negative surface charge.K +,Rb +and Cs +,on the other hand,have larger ionic radii and con-sequently are less strongly hydrated.This leaves them with a more effective positive surface charge which facilitates their adsorption onto clay minerals,etc.

90K.Bj?rlykke

This is demonstrated in nature during weathering

and transport.While similar amounts of potassium and

sodium are dissolved during weathering of basement

rocks,the potassium concentration in the sea is much

lower(K/Na ratio of only1:30).This is because K+is

more effectively removed by adsorption because it is

less protected by hydration.The same is true to an even

greater extent for Rb+and Cs+,which are adsorbed

even more readily.These ions therefore have a rela-

tively short residence time in seawater,between being

delivered by rivers and then removed by accumulating

sediment.

With regard to Group2elements,Mg++for example

will be more strongly hydrated than Ca++because it is

a smaller ion.As a result,Mg++has a greater tendency

to remain in solution in seawater.However,despite the

fact that the Mg/Ca ratio in seawater is5,it is calcium

carbonate which is the?rst to form through chemi-

cal and biological precipitation.Dolomite or magnesite

do not precipitate directly from seawater and this is in

part due to the strong hydration of Mg++.Normally,if

we had naked(unhydrated)ions,MgCO3and FeCO3

would be more stable than CaCO3because Mg++and

Fe++have greater ionic potentials and stronger bonding

to the CO2?3ion.However with increasing tempera-

ture the hydration declines because the bonds with the

dipole of the water molecules become weaker.Mg++is

then more likely to be incorporated into the carbonate

mineral structures.Therefore during diagenetic pro-

cesses at80–100?C,magnesium carbonates precipitate

more readily even if the Mg++/Ca++and Fe++/Ca++ ratios are low.Even if Mg is preferred in the carbon-

ate structure and also in the clay minerals,very little

magnesium is usually available in the deeper parts of

sedimentary basins except in the presence of evaporites

with Mg salts.

3.2Redox Potentials(Eh)

Oxidation potential(E)is an expression of the ten-

dency of an element to be oxidised,i.e.to give up

electrons so it is left with a more positive charge.

This potential can be measured by recording the

potential difference(positive or negative)which arises

when an element functions as one electrode in a

galvanic element.The other electrode is a standard

one,normally hydrogen.The oxidation potential of the reaction H2=2H++2e(electrons)is de?ned as E0=0.0V at1atm and H+concentration of1mol/l at 20?C.Different conventions have been used to assign plus and minus values.In geochemical literature, metals with a higher reducing potential than hydro-gen are assigned negative values,e.g.Na=Na++ e?=?2.71V,while strongly oxidising elements are given a positive sign,e.g.2F?=F2+2e=2.87V.

A list of redox potentials shows which elements will act as oxidising agents,and which will be reducing agents.Reactions which result in a negative oxidation potential(E)will proceed spontaneously,while those which have positive voltage will be dependent on the addition of energy from an outside source.We can predict whether a redox reaction will occur by using Nernst’s Law(see chemistry textbooks).

3.3pH

The ionisation product for water is

H+

OH?

= 10?14.The concentration of H+in neutral water will be10?7.pH is de?ned as the negative logarithm of the hydrogen ion concentration,and is therefore7for neu-tral water(at25?C).However,the ionisation constant (product)varies with temperature,e.g.at125?C the ionisation constant for water is[H+]·[OH?]=10?12. In other words,neutral water then has a pH of6.It is important to remember this when considering the pH of hot springs or in deep wells,for example oil wells.

In nature the pH of surface water mostly lies between4and9.Rainwater is frequently slightly acid due to dissolved CO2,which gives an acid reaction: H2O+CO2=H2CO3(carbonic acid)

H2CO3=H++HCO?3=2H++CO2?3 Humic acids may give the water in lakes and rivers a low pH.Sulphur pollution from burning oil and coal gives SO2,which is oxidised in water to sulphuric acid: 2SO2+O2+2H2O=2H2SO4

In areas with calcareous rocks or soils this sulphuric acid is immediately neutralised and the water becomes basic,as is the case across much of Europe.By con-trast,in areas with acidic granitic rocks as in the south of Norway and large areas of Sweden,the rock does

3Sedimentary Geochemistry91

not have suf?cient buffer capacity to counteract acid rain or acidic water produced by vegetation(due to humic acids).Organic material also contains a certain amount of sulphur,and drainage of bogs,or drought, can produce an acidic reaction.This is because H2S from organic material is oxidised to sulphate when the water table is lowered,allowing oxygen to penetrate deeper in these organic deposits.

The water near the surface of large lakes and the sea can have a high pH because CO2is consumed due to high organic production(photosynthesis).If the organic material decomposes(oxidises)on its way to the bottom,CO2is released again,causing the pH to decrease with depth since the solubility of the CO2 increases with the increasing pressure.

CO2is also less soluble in the warm surface water than in the colder water at greater depth.

Seawater is a buffered solution,with a typical pH close to8,though this varies somewhat with tempera-ture,pressure and the degree of biological activity.

Eh and pH are important parameters for describing natural geochemical environments,and the diagram obtained by combining these two parameters is partic-ularly useful.

The lower limit for Eh in natural environments is de?ned by the line Eh=?0.059pH,because other-wise we would have free oxygen,and the upper limit corresponds to Eh=1.22?0.059pH,beyond which free oxygen would be released from the water.If we also set pH limits at4and9in natural envi-ronments,we can divide the latter into four main categories:

1.Oxidising and acidic

2.Oxidising and basic

3.Reducing and acidic

4.Reducing and basic

Variations of pH and Eh are the major factors involved in chemical precipitation mechanisms in sedi-mentary environments where there is not strong evap-oration(evaporite environments).

The solubility of many elements is highest in the reduced state and they are precipitated by oxidation. This is particularly characteristic of iron and man-ganese,whereas others such as uranium and vanadium are least soluble in the reduced state.3.3.1Distribution Coef?cients

When a mineral crystallises out of solution,the com-position of the mineral will be a function of the compo-sition of the solution and the temperature and pressure. Trace elements which are incorporated in the mineral structure are particularly sensitive to variations of these factors.With constant temperature and pressure,the concentration of an element within a mineral which is being precipitated,is proportional to its concentration in the solution.The ratio between the concentration of an element in the mineral and its concentration in the solution(water)is called the distribution coef?cient.

A number of elements substitute for Ca++in the calcite lattice:Mn++,Fe++and Zn++have distribution coef?cients(k)<1.This means that they will be cap-tured,so that the mineral becomes enriched in these elements relative to the solution.

Mn++/Ca++(mineral)=k·Mn++/Ca++(solution) k here is about17,that is to say the manganese con-centration in the calcite is17times greater than in the solution.

At low temperatures(25?C)Mg++,Sr++,Ba++and Na+have distribution coef?cients<1.This means that the mineral phase will contain proportionately less of these elements than the aqueous phase.For Sr++, k is about0.1(0.05–0.14)in calcite,such that the Sr content in calcite is relatively low.The Sr con-tent in aragonite is considerably higher because the Sr++ion,which is larger than the Ca++ion,is more easily accommodated within the lattice.By analysing trace elements in minerals like calcite we can infer something about the environment when the minerals precipitated.Limestones with a high content of stron-tium may have had much primary aragonite which was replaced by calcite.Calcite containing signi?cant amounts of iron must have precipitated under reducing conditions because only Fe++would be admitted into the calcite structure.

3.4Isotopes

A number of elements occur in nature as differ-ent isotopes:the atomic number(protons)is constant but there are different numbers of neutrons.They

92K.Bj?rlykke

therefore have the same chemical properties although their masses are slightly different.Isotopes which are radioactive(unstable)break down at a speci?c rate characteristic for the isotope species(the disintegration constant).By analysing the reaction products formed in the minerals they can be dated.The87Rb?87Sr and the40K?40Ar methods are the ones most com-monly used in determining the age of rocks.The ratios between lead isotopes can also be employed because of the235U?207Pb,238U?206Pb and238Th?208Pb reactions.

Dating sedimentary rocks is a complicated proce-dure and the results are often dif?cult to interpret. The main problem is that clastic sediments are com-prised of fragments and minerals which have been eroded from older rocks and the measured radio-metric age may be strongly in?uenced by the age of these source rocks.Separating the newly formed (authigenic)mineral to be dated,can be particularly challenging.

The fact that isotopes have different masses causes fractionation to take place through both chemical and biological processes.The simplest example is water, H2O,which contains two oxygen isotopes and two hydrogen isotopes.The oxygen isotopes are fraction-ated through evaporation,with more H162O evaporating than H182O.This is because the18O isotope has greater mass and a phase change from?uid to vapour therefore requires more energy.H162O has higher vapour pressure than H182O.This is the reason why rainwater and ice contain less18O than seawater.

Isotope fractionation is a function of temperature, however,and is much more effective with evaporation at low temperatures than at high ones.The explana-tion for this is that at high temperatures the energy of the molecules are so great that the difference in mass between18O and16O is of less consequence.At low temperatures the isotopic separation evaporation is much more selective so that the water evaporated is more enriched in16O.When water vapour condenses to rainwater,molecules with18O are most stable.Rain and snow becomes enriched in the heavier isotope (18O),so that the water vapour remaining in the air becomes more enriched in16O.Most of the evapora-tion takes place at low latitudes and the water vapour in the air has a progressively lower18O-content towards higher latitudes as the air cools and it rains.The con-centration of oxygen isotopes is expressed in relation to a standard:

δ18O=

18O/16O

sample/18O/16O std?1

·1000

This standard may be the average composition of seawater,called SMOW(Standard Mean Ocean Water).Another commonly used standard is PDB(Pee Dee Belemnite),which is the composition of calcite in a Cretaceous belemnite.The calcite(CaCO3)was precipitated in the sea and its composition was in equilibrium with the seawater at normal temperatures (15–20?C).There is more18O in calcite than in the water(positive fractionation),but with higher temper-atures the less effective fractionation of oxygen lowers theδ18O values.The relationship between the two standards is:

δ18O SMOW=1.031·δ18O PDB+30.8

PDB values are preferred for carbonate minerals while the SMOW scale is mainly used for water sam-ples and silicate minerals.

Hydrogen has two stable isotopes,1H and2H(deu-terium),and an unstable one,3H(tritium),which has a half-life of12years.The hydrogen isotopes are even more strongly fractionated than oxygen isotopes during evaporation.Water molecules with deuterium (heavy water)have lower vapour pressure that water moles with hydrogen.

In meteoric water there is a linear relation between the deuterium/hydrogen ratio(D/H)and theδ18O.

The isotopic composition of seawater has varied through geological time,though not so much during the last200–300million years.During glacial periods, seawater acquires more positiveδ18O values because the water bound as ice has more negativeδ18O values. Rainwater(meteoric water)has normalδ18O values from–2to–15.The values become more negative towards higher latitudes,and near the poles one can measureδ18O values of about–50andδD(2H)values close to–350(see Fig.3.4).Minerals that form in sea-water show decreased18O/16O ratios with increased ambient temperature during formation.Theδ18O/16O ratio in carbonate secreting marine organisms,for example,is thus a function of both temperature and salinity.The seawater changes itsδ18O values by around1–1.5‰.Isotopes can thus provide important proxy evidence for palaeoclimate studies.

Cold freshwater gives strongly negativeδ18O val-ues,whereas evaporites are enriched in18O isotopes

3Sedimentary Geochemistry 93

δ1800

δD 0

–300

–200

–100

Fig.3.4Ratio between the isotopic composition of seawater and freshwater.Evaporites will deviate from the mixing line between these endmembers

(positive δ18O).Shallow marine carbonates that are diagenically modi?ed by freshwater,give lower δ18O values than marine carbonates deposited in deeper water.

Stable oxygen isotope analyses were ?rst used by Urey,in 1951,to demonstrate past temperature changes in seawater.By taking samples through a cross-section of a belemnite it was possible to regis-ter annual variations in seawater temperature from 150million years ago (Fig.3.5).

The precipitation of newly formed (authigenic)min-erals gives an oxygen isotope composition which is a function of the composition of the porewater in which the mineral is precipitated,and the tempera-ture.If the porewater isotope composition is known,the temperature (T )can be calculated,and vice versa.The calcite precipitation formula is:

T =16.9?4.38(18O carb ?18O water )+

0.1(18O carb ?18O water )2

Here the values for calcite are given in PDB and for water in SMOW.We see that if the δ18O value for cal-cite is 0(PDB)and seawater has 0(SMOW),the tem-perature is 16.9?C,which may have been a typical sea temperature when the standards were precipitated.).The above formula can be expressed graphically,enabling the temperature to be read off a curve as a function of the isotopic composition of the calcite,

which is the assumed composition of the porewater during precipitation (Fig.3.6).Similar calculations can be done for other precipitated minerals,for example for quartz using the δ18O fractionation as a function of the temperature for quartz.

Carbon has two stable isotopes (12C ?98.9%and 13C ?1.1%).During photosynthesis a greater propor-tion of 12CO 2than 13CO 2forms organic compounds,because 12CO 2has a smaller https://www.wendangku.net/doc/ba8774721.html,anic material is therefore enriched in 12C relative to atmospheric CO 2and HCO ?3in seawater.The isotopic composition of carbon is expressed as δ13C values:

δ13C =[13C /12C(sample)/13C /12C(std)?1]·1000All samples are compared against a standard of marine calcite,the PDB belemnite,which by de?ni-tion has δ13C =0‰PDB.The isotopic composition of dissolved carbon (CO 2)has been relatively constant during the last 300–400million years,but limestones can nevertheless be dated and correlated using differ-ences due to variation in the composition of seawater.Towards the end of the Precambrian the composition of seawater seems to have been more variable,and there this type of correlation is particularly valuable since there are no fossils.In large massive limestones the isotope composition does not change signi?cantly during diagenesis,because the volume is so great.Atmospheric CO 2has δ13C =–7‰.Land plants have an average δ13C value of –24(–15to –30‰),and marine organisms have a similar range of values.Freshwater containing CO 2released by the breakdown of organic matter,and groundwater ?ltered through a soil pro?le,will take up CO 2with negative δ13C values from roots and organic material.

Bacterial fermentation of organic material (2CH 2O =CH 4+CO 2)forms gas (methane)which is very strongly enriched in 12C (δ13C =–55to –90‰)and CO 2which is positive (δ13C =+15).

Thermal breakdown (thermal decarboxylation)of organic matter produces δ13C values of –10to –25.The strontium isotope ratio (87Sr /86Sr)in seawa-ter has varied considerably through geological time (Fig.3.7).This is because there are two radically dif-ferent sources of strontium in seawater.Continental weathering supplies much 87Sr to the sea since granitic rocks contains relatively high concentrations of rubid-ium which can decay to 87Sr.Dissolution of basalt at

94K.Bj?rlykke

202.0

016018

19t °c

181716150

0.2

0.4

0.6

0.8 1.0

1.2

1.4cm

Radius

Fig.3.5Analyses of oxygen isotopes in a Jurassic belemnite from the centre to the outermost layer.Colder water during the winter is recorded by lower 18O/16O ratios.We can see that the belemnite lived for 4.5years and died in the spring.(From Urey et al.1951)

the mid-oceanic ridges will supply strontium with a relatively low 87Sr /86Sr ratio because basalt contains a little potassium and also rubidium.

When there is rapid sea?oor spreading a great deal of water passes through the mid-oceanic ridges,so that the seawater receives much Sr with a low 87Sr /86S ratio.During such periods,for example in the Jurassic –Cretaceous,the creation of new warm sea-?oor will lead to a transgression onto the continents.This reduces the gradients and hence transporting capacity of rivers,limiting the supply of clastic mate-rial to the ocean.

Since the Jurassic,the 87Sr /86Sr ratio has risen almost continuously,and by analysing marine calcitic

3Sedimentary Geochemistry

δ180 Calcite (PDB)

204080120

Late Calcite

Cement ( type II )

T e m p e r a t u r e o C

10060140160

Late Calcite Cement (type I)

Measured range in calcites ( type I & II )

δ1

Fig.3.6Relation between the isotopic composition of porewater and carbonate cement,as a function of temperature (from Saigal and Bj?rlykke 1987)0.709

0.708

0.707

0100200300400500600 m.yr.

Sr/86Sr Fig.3.787Sr/86Sr ratio in seawater from the Cambrian to present.Based on

MacArthur et al.(2001).This ratio re?ects the relative

contribution form weathering of continental rocks with high contents of 87Sr and exchange with basaltic rocks with low contents of 87Sr on the oceanic spreading ridges

fossils such as foraminifera,one can obtain rather accurate age determinations.This applies particularly to the Tertiary period,when the rise in the 87Sr /86Sr ratio was particularly rapid (Fig.3.7).

The isotopic composition of clastic sediments can also be used for stratigraphical correlation.Then it can be more useful to employ isotopes which do not go into solution and react with water,but retain the orig-inal age of the rocks from which they were eroded.In the North Sea and on Haltenbanken the ratio between the rare earth elements samarium and neodymium

(147Sm /143Nd)was used to correlate reservoir rocks both in-?eld and regionally.

3.5Clay Minerals

A number of minerals are referred to as clay min-erals because they predominantly occur in the ?nest grain-size fraction (clay fraction)of sediments and sedimentary rocks.However,this is not an accurate

96K.Bj?rlykke

de?nition,because the clay fraction contains many other minerals than those we call clay minerals,and because the clay minerals themselves are often larger than4μm(0.004mm).By“clay minerals”we usually mean sheet silicates which consist chie?y of oxygen, silicon,aluminium,magnesium,iron and water(H2O, OH–).Clay minerals in sedimentary basins are partly derived from sheet silicate minerals occurring in meta-morphic and eruptive rocks(e.g.biotite,muscovite and chlorite),but during weathering and transport these clastic minerals are typically altered from their initial composition in the parent rock.

Mica(muscovite and biotite)lose some potassium which is replaced by water(H2O,H3O+)to form illite (hydro mica).Clay minerals are also formed through weathering reactions,for example by the breakdown of feldspar and mica.Clay minerals which are formed by the breakdown of other minerals within the sediment, are called authigenic.

Sheet silicates have a structure consisting of sheets of alternating layers of SiO4tetrahedra and octahe-dra.In the tetrahedral layers,silicon or aluminium atoms are surrounded by four oxygen atoms.In the octahedral layers the cation is surrounded by six oxy-gen or hydroxyl ions.Both bi and trivalent ions can act as cations in the octahedral layer.In sheet sili-cates with trivalent ions(e.g.Al3+)only two of the three positions in the octahedral layer are occupied, and such minerals are therefore called dioctahedral. With bivalent ions(Mg++,Fe++)all three positions must be?lled to achieve a balance between the posi-tive and negative charges,so these minerals are called trioctohedral.

The main method of identifying clay minerals is X-ray diffraction(XRD),by which the thickness of the sheet silicates is determined using X-rays which are diffracted according to Bragg’s Law:nλ=2d sin?. Hereλis the wavelength of the X-ray,?the angle of incidence and d the thickness of the re?ecting silicate layers;d is thus a function of angle?.

Sheet silicates may also be identi?ed by means of differential thermal analysis(DTA),which records characteristic exothermal or endothermal reactions.

Figure3.8shows the structure of some of the main clay minerals.Illite consists of sheets with two layers of tetrahedra and one of octohedra,bonded together by potassium.This ionic bonding is relatively weak so the mineral cleaves easily along this plane.The bonds within the tetrahedral and octahedral layers are more covalent and stronger.The potassium content in mica corresponding to the formula of mica is about 9%K2O,while illite has a greater or lesser de?cit of potassium.Smectite(montmorillonite)has the same structure except that most of the potassium is replaced by water(H3O+),other cations or organic compounds (e.g.glycol).There are strong indications that smectite consists of small particles of10?,plus water.

Illite is most likely comprised of several layers of these small10?particles stacked on top of one another.

In an atmosphere of glycol vapour,smectite will swell from14to17?,while illite is unable to expand because there are numerous layers bonded together with K+or other cations,for example NH+.Smectite has a very high ion-exchange capacity and to some extent can exchange ions in the octahedral layer.The stability of smectite declines in aqueous solutions with high K+/H+

Na+/H+

ratio and with increasing tem-perature,and it converts to illite.Vermiculite has a structure reminiscent of the smectites,and also under-goes ion exchange and thus charge de?cit in the tetra-hedral layer,so that the bonding between each layer is too strong for much swelling to occur.Vermiculites are mostly trioctahedral,containing mostly Mg or Fe in the octahedral layer.

Glauconite is a green mineral which forms on the seabed.It is a potassium and iron bearing silicate somewhat similar to illite and contains both di-and trivalent iron.It is therefore formed right on the redox boundary,and during periods with little or no clastic sedimentation this can result in relatively pure beds of glauconite.

Kaolinite consists just of a tetrahedral layer and an octahedral layer and is very stable at low temperatures. There are no positions in the structure where exchange can precede easily,which gives kaolinite a much lower ion exchange capacity than smectite.At higher tem-peratures kaolinite becomes unstable and will convert to illite if K-feldspar or other sources of potassium are available(at130?C)or pyrophyllite(Al3Si4O10(OH)2) at higher temperatures.

Kaolinite is part of the kaolin mineral group,which includes dickite which tends to form at slightly higher temperatures(100?C).

Chlorite is a mineral which consists of two tetra-hedral layers and two octahedral layers,totalling 14?.The octahedral layer is?lled with Mg++and Fe++.Magnesium-rich chlorites are typical of high

3Sedimentary Geochemistry

{

Tetrahedron

Octahedron

Si 4+, Al 3+

O – – ,OH –

O – – ,OH –Tetrahedral layer

(SI,Al) (O, OH)4

Octahedral Al 3+

/Mg ++

, Fe ++

layer

Al 3+,Fe 3+,Fe ++,Mg ++

K Al 2 AlSi 3O 10 (OH)2Muscovite

(Dioctahedral)

K (Mg, Fe) AlSi 3 O 10 (OH)2Biotite

(Trioctahedral)

Illite

(K 1–x) (Al, Mg, Fe)2–3 AlSi 3 O 10(OH)2

Tetrahedral layer (SI,Al) (O, OH)4Octahedral layer (Al, Mg)

n. H 3O + (Exchangeable cations x)(Glycol)

Smectite × (Al 2–x , Mg x ) Si 4O 10(OH)2(Montmorillonite)

(Si 3, Al) (O, OH)4(Al 1)

Kaolinite – Al 2SiO 2O 5 (OH)4

(Si 3, Al)(Mg, Fe, Al)

Octahedral layer (Me, Fe, Al)(Biotite layer)

Chlorite – (Mg, Fe, Al)6 (Si,Al)4 O 10 (OH)8

14 ?

7 ?

14–17 ?

10 ?

K +

Fig.3.8Simpli?ed illustration of the main groups of clay min-erals.Their physical properties can be explained by their crystal

structure.The chemical bonds between SiO 4?4and O

2?in the tetrahedral structure are very strong.In the octahedral layers the

bonds are weaker because Mg ++is surrounded (co-ordinated)with 6oxygen.In the illite (mica)structure potassium is co-ordinated with 12oxygen molecules,resulting in weak bonds that produce a strong cleavage

temperature metamorphic rocks,while iron-rich ones may form authigenically in sediments near the sea?oor or at shallow depth.Chamosite is an iron-rich chlorite mineral which often forms close to the sediment surface in reduc-ing conditions.Together with siderite,chamosite is

98K.Bj?rlykke

an important mineral in sedimentary ironstones which, especially in England,earlier were used as iron ore.

Clay minerals have a number of properties which distinguish them from most other minerals.Because of their very large speci?c surface area they have a great capacity for adsorbing ions,which is increased by the fact that clay minerals have negatively charged edges due to broken bonds.In water with a low electrolyte content,clay minerals will therefore repel each other.If cations are added,clay minerals will then accumulate a layer of positive ions(a double layer),and repulsion between negatively charged clay minerals declines as the strength of the electrolyte increases.Van der Waal’s forces will therefore cause?occulation more easily in saltwater,where repulsion due to the negative charge is reduced.This is why clays transported by rivers?oc-culate into larger particles which sink more rapidly to the sea?oor.

A colloid solution of clay in water is called a sol, which can be regarded as a Newtonian?uid.Floccu-lated clay is called a gel and has thixotropic properties. This means that the shear strength decreases with increasing deformation so that it changes from a gel to a sol which consists of dispersed colloidal particles in water(hydrosol).After a time without deforma-tion,it regains its strength.This is typical for smectitic clays.

Sediments which increase their volume when they are deformed,are called dilatant.When the original packing of the grains is destroyed,the new packing may be less effective and the volume then increases. Walking on a beach the deformation of the sand causes increased porosity so that water is sucked into the sand at the surface.

Norwegian clays deposited in the sea when the ice sheet retreated about10,000years ago,consist of crushed rock fragments and have approximately the same composition as the parent rock.During deposi-tion these clays acquired a markedly porous structure, with the minerals stacked like a card house.Saltwater helped to hold this structure together because it neu-tralised the negative charges.When the clays are ele-vated above the sea by isostatic recovery of the land, they are exposed to meteoric water.Even if the clay has a low permeability,freshwater will slowly seep through it and remove the saline water.This reduces the strength of the clay’s structure and hence its sta-bility and the clay becomes“quick”.This is caused by overpressure and weak effective stress between the grains and hence little friction.The card house structure may then collapse,releasing the interstitial water to produce a low viscosity clay slurry that will ?ow even down very gentle gradients.The addition of salt(NaCl or KCl)binds the clay particles so that the shear strength is increased,and this is employed when the ground has to be stabilised for buildings and construction.

3.6Weathering

The composition of clastic rocks in sedimentary sequences depends to a large extent on the supply of sediments from source areas undergoing weathering and erosion.The physical properties of sedimentary rocks are controlled by the primary sediment compo-sition and changes during burial(diagenesis).We shall here brie?y look at the processes producing sediments.

Mechanical weathering is the physical breakdown of rocks into smaller pieces which can then be trans-ported as clastic sediment.

Chemical weathering involves the dissolution of minerals and rocks and precipitation of new minerals which are more stable at low temperatures and high water contents.Parts of the parent rock will then be carried away in aqueous solution by the groundwater and rivers into the ocean.

Erosion is the combined result of the disintegration of rock and the removal of the products.

As we shall see,biological processes are not only important in connection with chemical weathering but also with respect to mechanical weathering.Both mechanical and chemical weathering,which are essen-tially land surface processes,are due to the chemical instability of rocks that were formed or modi?ed under other conditions(at greater depths,higher pressures or temperatures,or in different chemical environments). They are no longer stable when exposed to the atmo-sphere,water and biological activity.

3.6.1Mechanical Weathering

Igneous and metamorphic rocks which have been formed at many km depth at higher temperatures and pressures are not stable when exposed at the surface.

3Sedimentary Geochemistry99

When uplifted and unloaded the rocks expand, mostly in the vertical direction,producing horizontal fractures(sheeting)parallel to the land surface.This is because as the rocks near the surface are unloaded by the reduction in overlying rock,they can expand vertically but not horizontally.In this way the ver-tical stresses become less than the horizontal ones, and joints develop normal to the lowest stresses.We see this most clearly in granites,which are homo-geneous,while expansion in metamorphic and sedi-mentary rocks occurs along bedding surfaces,along tectonically weak zones with crushing,or along frac-tures that were formed at great depth.Joints opened by stress release in turn provide pathways for ground-water to circulate,increasing the surface area of rock exposed to chemical weathering.

In areas that experience freeze-thaw cycles,frost weathering becomes very important.When water freezes in cracks in rock,it expands by9%and can generate very high stresses,further widening cracks near the surface.The surfaces of exposed rocks are also subjected to daily temperature?uctuations which cause greater expansion of the outer layers relative to the rest of the rock.Desert regions in particular expe-rience very wide daily temperature ranges,though the importance of this process for mechanical weathering has been questioned.The roots of plants and moss can also contribute to mechanical weathering as they grow into fractures,take up water and expand.

3.6.2Biological Weathering

Rocks are a source of nutrients for plants,and plants are capable of dissolving and breaking down the major rock-forming minerals.Moss,which consists of algae and fungi living in symbiosis,produces organic com-pounds that can slowly dissolve silicate minerals.Even in the earliest stages of weathering,we see that fungus hyphae penetrate into microscopic cracks.

Plant roots produce CO2which helps to lower the pH and dissolve minerals such as feldspar and mica,thus freeing an important plant nutrient,potas-sium.Plants also produce humic acids,which likewise strongly in?uence the solubility of silicate minerals, and also affect the stability of clay minerals.The pro-duction of humic acids is perhaps the major factor in?uencing the rate of weathering.In heavily vegetated areas,such as rain forest in the tropics,the weathering rate is exceptionally high because so much humic acid is produced.

Bacteria and fungi,which are found in almost all soil types,are active in breaking down minerals. Animals also contribute to weathering,and certain marine organisms such as mussels are able to bore into solid rock(see Chap.8).Microbiology has become a key area of research in the quest to understand how minerals are dissolved and precipitated.

3.6.3Chemical Weathering

There is no sharp demarcation between biological and chemical weathering,because we?nd biological activ-ity in almost all soils and rocks near the surface. The chemical environment in water at the surface of the earth is very much affected by local biological activity,and in most cases it is biological processes that cause weathering to continue after rainwater has been neutralised through reaction with minerals.We will therefore use the term“weathering”here for both chemical and biological processes.

3.6.4Weathering Pro?les(Soil Pro?les) Both chemical and biological weathering are to a large extent controlled by climate.The crucial factor is the ratio between precipitation and evaporation in an area. In areas where precipitation far exceeds evaporation, podsol pro?les develop in which there is a net trans-port of ions down through the soil pro?le as minerals are dissolved.In other words,we get weathering due to the fact that rainwater is slightly acidic(on account of its CO2and H2SO4content)and contains oxygen. Rainwater is initially undersaturated with respect to all minerals.Some minerals are only very slightly soluble, others more soluble in this slightly acidic,oxidising water.Dissolved ions are transported down to the water table,but ferrous iron liberated from iron-bearing min-erals will be oxidised and precipitated as ferric iron (Fe(OH)3).Vegetation at the top of the soil pro?le produces CO2from roots and organic compounds,par-ticularly humic acid,which will increase the solubility of silicate minerals.Similarly,aluminium derived from

100K.Bj?rlykke

Podsol profile Layer A

Leaching (grey)

Precipitation

of Fe(OH)3 (red)

Layer B

Capillary water

Water table

Phreatic (ground) water

Layer C

Podsol soil profile. Downwards transport.

Brown-Earth

profile

Little leaching.

Precipitation of

carbonates

(caliche)

Water table

Phreatic

Layer C

or a tion

Evaporitic soil

Precipitation

of salts and

carbonates

Water table

phreatic

(ground) water

Layer C

or a tion

u noff

Fig.3.9Simpli?ed representation of soil pro?les as a function of rainfall (precipitation)and evaporation

a solution of feldspar and mica,for example,pre-cipitates as Al(OH)3but is less noticeable because aluminium hydroxide is white.The uppermost part of the soil pro?le,where dissolution due to under-saturated rainwater and organic acids dominates,is called the A-horizon.Some of the dissolved salts and particularly iron hydroxide is precipitated in the B-horizon below(Fig.3.9).These may develop into a layer of solid rock(hard-pan)cemented with iron and aluminium oxides and hydroxides.

Where precipitation is approximately equal to eva-poration,there is less leaching within the soil pro?le. At a certain depth(about0.5–1m)carbonate will be precipitated and form an indurated layer(calcrete) which may be eroded and from conglomerates.

The organic content is greater in the B-horizon which is brown due to less oxidation of organic mat-ter,hence the term brown-earth pro?les.If evaporation is greater than precipitation there will be a net upward transport of porewater,causing dissolved salts from the groundwater to be precipitated high in the soil pro?le.

3.6.5What Factors Control Weathering

Rate and Products?

Because weathering is the most important sediment-producing process,we are interested in understanding how the rate of weathering is related to rock type,pre-cipitation,temperature,vegetation,relief etc.We also try to establish correlations between weathering prod-ucts,particularly clay minerals,and these factors.By studying sediments from older geological periods,we can learn something about weathering conditions at those times.Weathering products will also bear the stamp of the rocks undergoing weathering.The stabil-ity of a mineral during weathering is largely a function of the strength of the bonds holding the cations in the crystal lattice.Potassium(K+)in mica is held by weak bonds(low ionic potential)which are responsible for the pronounced cleavage.In biotite,the Mg++and Fe++ in the octahedral layer will also be weakly bonded. During weathering cations like K+,Na+,Ca++,Mg++ and Fe++can be attacked by protons(H+)which will replace them and send them into solution.Chain sil-icates like hornblendes and pyroxenes will also be relatively unstable and rapidly weather.In feldspars the alkali ions are dissolved so that the whole mineral disintegrates.Stability is lowest in calcium-rich plagio-clase,while pure albite(sodium feldspar),orthoclase and microcline(potassium feldspar)are more stable. The breakdown of these silicate minerals will pri-marily liberate alkali cations.Silicon and aluminium have very low solubility and form new silicate min-erals,largely clay minerals,though some silicic acid (H4SiO4)goes into solution.

3Sedimentary Geochemistry 101

1.2K(Na)Al 2AlSi 3O 10(OH)2(muscovite)+2H +

+3H 2O =3Al 2Si 2O 5(OH)4(kaolinite)+2K +(Na +)

2.2K(Mg,Fe)3AlSi 3O 10(OH)2(biotite)+12H ++2e +O 2=

Al 2Si 2O 5(OH)4(kaolinite)+4SiO 2(in solution)+Fe 2O 3+4Mg +++6H 2O +2K +

3.2K(Na)AlSi 3O 8(feldspar)+2H ++9H 2O =

Al 2Si 2O 5(OH)4(kaolinite)+4H 4SiO 4+K +(Na +)We see that potassium has been replaced by hydro-gen ions in the new silicate minerals.The same applies to sodium in albite.The equations show that the reac-tions are driven to the right by low K +/H +and Na +/H +ratios.

The degree of weathering depends on how under-saturated the water is with respect to the minerals comprising the rock,and on the volume of water ?ow-ing through the rock.If the reaction products in the solution,K +,Na +and silica (H 4SiO 4),are not removed by water ?ow,the reactions will cease.This is why weathering is always found to begin along cracks where water can penetrate (Fig.3.10).The vertical and horizontal joints that often develop in response to pressure release when previously deeply buried rock is exposed at the land surface provide the initial path-ways.As the weathering process spreads outward from joints,blocks of unweathered rock are gradually iso-lated.They have rounded corners and may become entirely round (spheroidal weathering)(Fig.3.11a,b,c).In desert areas,where there is little rainfall,weathering proceeds much more slowly.Illite and montmorillonite may be formed under higher K +/H +and Na +/H +ratios than kaolinite,and they are frequently formed where there is less water percolation and the removal of potassium or sodium is slower.

There is often also a high silica content in the water in desert areas due to frequent silica algae

(diatom)blooms and because silica is concentrated by water evaporation.This helps enhance the stability of smectite.

Granites subjected to weathering over a very long period often develop a special topography.Fractures and fault zones weather fastest and form valleys where the groundwater collects,which further accelerates the weathering.The more massive granitic areas will stand out in the terrain,and because precipitation runs swiftly down into the depressions between the ele-vated portions,the topographic difference will become more and more pronounced.Granites surrounded by sedimentary rocks will,because of their high content of feldspar,normally weather faster than the sediments which contain more quartz and other stable minerals.Particularly if the sediments consist of quartzites and shales,the granite will form a depression in the terrain.The weathering products from granites will normally be quartz grains which form sand grains the same size as the quartz crystals in the granite,and clay con-sisting of kaolinite,and possibly also some illite and smectite formed from feldspars and micas.We have at the outset a bimodal grain-size distribution with sand and clay,but very little silt.

Basic rocks (e.g.gabbro)will weather far more rapidly than granite because basic plagioclase (Ca-feldspar),pyroxenes and hornblende are very unstable and dissolve faster than silica-rich (acid)minerals.During progressive weathering sodium,potassium,magnesisum and calcium will be removed by the groundwater but some potassium may be adsorbed on clay minerals.The weathering residue will be enriched in elements with low solubility such as Ti,Al,Si and Mn (Fig.3.12).In a normal,oxidising weathering environment,all the iron will precipitate out again as iron oxide (Fe(OH)3)while the magnesium will tend to remain in solution.In areas with high rainfall the concentration of ions like K +,Na +and silica will be

Weathering = Rock + H 2O + H +

Weathering product + Ions in solution

Exfoliation cracks

Granite Rainwater

Fig.3.10Weathering of granites along extensional fractures

102K.

Bj?rlykke

Fig.3.11(a,b)Weathered granite(North of Kampala,Uganda)showing different stages of weathering.(c)Concentric–spheroidal weathering in a basic intrusive rock,Weathering is much faster in the absence of quartz

diluted and kaolinite will precipitate.Where porewa-ter circulation is slower we may get a higher build-up of Mg++,Ca++and silica concentrations in the water, so that smectite(montmorillonite)or chlorite preci-pitates.Smectite requires porewater with a relatively high silica concentration(Fig.3.13)and is therefore often found in sediments derived from volcanic rocks that contain glass or soluble silicate minerals.Biogenic sources of silica(diatoms,radiolaria)will also increase the silica concentration in porewater because amor-phous silica is much more soluble than quartz. In desert environments evaporation of water after rainfalls will concentrate silica in the porewater and make smectite stable.Figure3.12shows analyses of rocks at various stages of transformation due to weath-ering.Weathering proceeds particularly rapidly in amphibolites:Na+,Mg++and Ca++are quickly leached out,while A13+,Fe3+and Ti4+become enriched.K+is

3Sedimentary Geochemistry

103

%200180160140120100908070605040302010J1

CaO MgO Na 2O K 2O

SiO 2Al 2O 3

Fe 2O 3

TiO 2

Fig.3.12Chemical analyses of changes in the chemical com-position of samples representing progressive weathering of an amphibolite compared to an unweathered sample.The stable elements (Ti,Fe and Al)are enriched while there is a strong depletion of Na,Mg and Ca.Potassium is depleted but is adsorbed on clay minerals in the soil

8.0

6.0

0.0

4.0

2.0

Log aSiO 2 (aq)

Log

aK +/aH +

Fig.3.13Activity diagram showing the stability of some min-erals as a function of silica and the K +/H +ratio (after Aagaard and Helgeson 1982).Weathering reactions are characterised by

reduction in both these two parameters towards the lower left of the diagram

104K.Bj?rlykke

however to large extent adsorbed on clay minerals and

much more Na+than K+is therefore supplied to the

oceans with the rivers.

After the alkali cations have been dissolved out of

the silicates and kaolinite has been formed,extremely

slow leaching of quartz commences.When the concen-

tration of silica in the porewater is suf?ciently low(see

Fig.3.13),kaolinite will be unstable and be replaced

by gibbsite Al(OH)3or(A12O3?3H2O).Since gibbsite cannot form as long as the porewater is in equilibrium

with quartz,all the quartz must have dissolved?rst or

become encapsulated(e.g.in a layer of iron oxides).

Clearly,gibbsite will form far more rapidly during the

weathering of basic rocks than of granites,since the

initial silica content is considerably lower.It takes a

very long time to dissolve all the quartz in a granite.

The solubility of quartz at surface temperatures and

pH7–8is about5ppm,increasing at higher pH val-

ues.Alkaline(basic)water can therefore increase the

solution rate of quartz.The end product of the weath-

ering process is laterite,which consists of gibbsite and

iron oxides or hydroxides.Under atmospheric(oxidis-

ing)conditions with a neutral pH,aluminum hydroxide

and iron oxides may for practical purposes be regarded

as insoluble.

At low pH values,e.g.under the in?uence of humic

acids in humid tropical climates,aluminum is more

soluble than iron and selective leaching of A13+will

produce iron-rich laterites.Aluminum hydroxide can

be dissolved at low(acidic)or high(basic)pH val-

ues and may reprecipitate as aluminum oxide,which

has a lower iron content and is thus a higher grade

aluminium source(bauxite).

3.7Distribution of Clay Minerals

and other Authigenic Minerals

as a Function of Erosion

and Weathering

3.7.1What Determines the Type of Clay

Minerals We Find in Sediments

and Sedimentary Rocks?

When rocks are subjected to erosion and weather-

ing,clastic minerals are broken down and perhaps

somewhat altered,with respect to the minerals in the parent rocks.We can also get new minerals formed in the source rock itself,and precipitation of new minerals through weathering.

The more rapidly erosion and transport take place compared to the rate of weathering,the closer the com-position of sediments is to the source rock.Glacial sediments represent one end of the scale in terms of sediment composition.Because glaciations are char-acterised by very high rates of erosion and low tem-peratures,chemical weathering will be very weak,and Quaternary sediments–including clays–will have a composition which essentially represents the aver-age of the rocks which have been eroded.Sediments deposited in fault-controlled basins(e.g.rift basins) have a short transport distance between the site of erosion and deposition and the sediments have lit-tle time to weather on the way.Clastic chlorite and biotite break down relatively rapidly during weather-ing and are therefore likely to be preserved in addition to feldspar in rift basins.These unstable minerals are good indicators of rapid erosion and/or cold climates, since otherwise they are unlikely to survive.

An analysis of the distribution of clay minerals in modern sediments shows that we?nd clastic chlo-rite almost exclusively in high latitudes,except around islands of basic volcanic rocks(e.g.basalts).Clastic chlorites from metamorphic rocks or altered basic rocks are more magnesium-rich than authigenic chlo-rites,which are iron-rich(chamositic).In temperate areas with moderate to high precipitation,weathering proceeds relatively rapidly.In desert areas weathering proceeds very slowly because all weathering reactions require water.The type of weathering and the distri-bution of clay minerals are clearly related to latitude (Fig.3.14).

Smectite and illite are typical of deserts because the weathering is much slower when water is nearly absent.When feldspar and other unstable minerals are altered(weathered)gradually in a dry climate,alkali ions and alkaline earth ions such as K+,Na+,Ca2+and Mg2+will not be removed rapidly enough due to little percolation of fresh rainwater.

Rainwater will cause some dissolution of minerals and,during the periods of drying,the silica con-centration may be higher so that smectite becomes stable.

As a result we will usually get illite or smectite formed as authigenic minerals because they are stable in the presence of high K+/H+ratios in the porewater.

3Sedimentary Geochemistry

105

2700mm Evaporation Precipitation

Temperature

°C 252015105

Chlorite Illite and smectite Laterite Kaolinite

Smectite (montmorillonite)

24002100190015001200900600300

Fig.3.14Simpli?ed diagram showing the distribution of weathering and common clay minerals as a function of latitude and rainfall.Cold areas and deserts are characterised by little weathering and more mechanical erosion

Smectite (montmorillonite)is thus a common clay mineral in desert areas,and its ability to swell when wet renders sediments very plastic during ?oods.This expansion of smectite also lowers its permeability and may be the reason why water can ?ow over the surface for a long time before sinking into the ground.In addition,capillary forces will prevent rapid percolation of water through dry soil.On the ocean ?oor near desert regions we ?nd that illite and smec-tite are typical minerals,brought there by aeolian transport.

In tropical areas where precipitation is relatively high,the rate of weathering will be very rapid.This is not only because weathering processes accelerate with temperature,but also because vegetation produces large amounts of organic acids (humic acids)which are very effective in breaking down silicate minerals.Microbiological organisms such as fungi and bacteria also help in the breakdown process by producing CO 2which forms carbonic acid,H 2CO 3.

Gibbsite (A12O 3?3H 2O)and iron oxides (haematite,goethite,Fe 2O 3?3H 2O)are constituents of the lat-erite which we ?nd only in tropical areas with rapid weathering and slow erosion.Whereas iron oxides are also found at higher latitudes,gibbsite occurs almost exclusively in humid tropical areas.

Laterisation is a very slow process and takes millions of years,even in tropical regions with rapid weathering.It is therefore primarily in tropical areas that we ?nd bauxite for the aluminum indus-try.Iron-rich laterites may have an iron content of over 50%,and in some areas (e.g.India)have been exploited as iron https://www.wendangku.net/doc/ba8774721.html,terite forms a very hard cement-like crust over the weathering pro?le and is also virtually devoid of nutrients,so crop cultivation is impossible.In East Africa (especially Uganda),how-ever,erosion has incised through a layer of Tertiary lat-erites.While the laterite cover remains on ?at elevated surfaces,fresher,more fertile,rocks and weathering material is exposed in the valley sides.The vegetation in some tropical areas is more abundant,even if the soils are very poor in nutrients,because the vegetation recycles those nutrients which are available.If the veg-etation is removed and organic material is no longer produced,oxidation and the absence of humic acids (increased pH)will lead to precipitation of oxides and hydroxides which make the soil hard and uncultivable.V olcanic ash consisting of glass and unstable vol-canic mineral assemblages may alter to smectite on land or on the sea?oor.In deep sea sediments zeolites like phillipsite are common.

Areas with volcanic rocks,particularly amorphous material (volcanic glass),will often form zeolites.These require a high concentration of both silica and alkaline ions in the water,which is the situation when glass dissolves.Zeolites,particularly phillipsite,are formed authigenically on the Paci?c Ocean bed and are also found in lakes (e.g.in East Africa).

106K.Bj?rlykke

In summary we can say that the factors which deter-mine which types of clay minerals are“produced”in the various areas are:

1.The rocks which are eroded/weathered(source

rocks)

2.Rate of erosion

3.Temperature

4.Precipitation

5.Vegetation

6.Permeability of source rocks and sediments(perco-

lation of water).

Typical distribution of various minerals:

1.Chlorite and biotite–high latitudes(cold climate)

–rapid erosion.

2.Kaolinite–humid temperate and humid tropical

regions–good drainage.

3.Smectite(montmorillonite)–low precipitation or

poor drainage.Typical of desert environments,but also formed in impervious,e.g.basaltic,rocks in more humid environments.Typically formed from volcanic rocks.

4.Gibbsite–tropical humid climate–long weathering

period.

5.Zeolites–formed in areas with volcanic material

and restricted porewater circulation.Require a high concentration of silica and alkali ions.

3.8Geochemical Processes in the Ocean The ocean can be regarded as a reservoir of chemicals dissolved in water.It looks as though the composi-tion of seawater has not altered radically throughout the geological ages from the early Palaeozoic until the present day,although there have certainly been some variations.

The supply of elements to ocean water from rivers and by water circulation at the spreading ridges must be balanced by a removal of the same elements from the ocean water(Fig. 3.15).The annual addition of salts dissolved in river water is about2×109 tonnes/year.The?gure was probably less in the geo-logical past because vegetation was sparse or absent. The development of land plants that produce humic acids,which in turn produced more rapid weathering, has probably increased the supply of salts(since Devonian times).This trend was sometimes slowed by periods with higher sea level that caused widespread transgressions and converted huge tracts of coastal land areas into continental shelf(e.g.during the Upper Cretaceous)reducing the weathering and the supply of salts and nutrients to the ocean.

Some ions like potassium(K+)are adsorbed to clay minerals supplied by rivers(B in Fig.3.15).Sodium (Na+)is so strongly hydrated that it has a tendency to remain in solution,while potassium will be far less hydrated and can be more easily adsorbed onto clay minerals and rapidly removed from seawater.

Most of the elements which the rivers bring to the ocean are precipitated by organic processes. Organisms can build their own internal chemical envi-ronment,and use their energy to precipitate minerals which are not normally stable in seawater.Carbonate-secreting organisms,e.g.foraminifera,molluscs etc., will precipitate aragonite or calcite even when the water is cold and undersaturated with respect to these minerals.Diatoms are so effective in precipitating sil-ica(amorphous silicon dioxide,SiO2),that in most places the seawater near the surface in the photic zone(photosynthetic zone)is much more depleted with respect to silica than to https://www.wendangku.net/doc/ba8774721.html,anisms,when they die,will in most cases start to break down by oxidation of organic matter and by dissolution of the mineral skeletons.In reducing environments much of the organic matter will however be pre-served.In shallow tropical waters,like on a carbonate bank,the seawater may be saturated with respect to carbonate(calcite),but carbonate also accumu-lates in cold water like the Barents Sea because the rate of dissolution is slower than the rate of precipitation.

The more ef?cient organisms are at building skele-tons despite undersaturation,the more rapidly they will dissolve.Diatoms,for example,dissolve to the extent of99–99.9%before they have sunk to the seabed. Only a very small proportion is therefore preserved in sediments.

Photosynthesising organisms in the surface water use up CO2and produce oxygen and organic matter: CO2+H2O+nutrients=CH2O+O2