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OPES Geochemistry





Lectured by : Zeqin Li Ph.D.geochemistry





Department Environmental Science & Engineering

Chengdu University of Technology

Chengdu, Sichuan Province, 610059

P. R. China





Phone: +86-28-84077542(O), 84078078(H)

Fax: +86-28-84077066

E-mail: lzq@cdut.edu.cn; zeqinlee@yahoo.com

Content



Principles of isotopic fractionation



Hydrogen and oxygen isotope geochemistry



Carbon isotope geochemistry



Sulfur isotope geochemistry



Isotope application

Review

LE ISOTOPES



Isotopes can be divided into 2 kinds:

 radioactive (unstable)

U-Pb, Rb-Sr, Th-Pb, K-Ar, Sm-Nb

 Nonradioactive(stable)

H, O, C, S, N, He, Li, B

Review

1 Principles of isotopic fractionation

Different isotopes of an element

Differences in mass result in

Minor differences in physical-chemical behaviour

Therefore, Stable Isotopic Fractionation due to differences

in isotopic mass by

1) Physical processes

2) Chemical reactions Isotopic Fractionation

3) Biological processes

Explain:

Molecules with the lighter isotope react faster.

why ? The molecules’ vibrational frequency is higher,

bond strength of the element is weaker ),

Review

differences in bond strengths affect kinetics of

reactions as the lighter isotope forms weaker bonds

than heavier isotope. Molecules with the lighter

isotope react faster, and are concentrated

preferentially into reaction products



Unidirectional reactions (e.g. open systems where

products escape) are effective at fractionating

isotopes



α goes to 1 with increasing T because small

Vibrational as

differences in bond strengths at high T are not frequency

depends on the temperature

of T.

effective in causing fractionation as at lowthe system.

Fractionation factor - As it is difficult to calculate

an equilibrium constant (K) for iso

differentiation of complex



 ab=Ra/Rb molecules









Ra is the ratio of heavy isotope and light isotope in A phase.

Rb is the ratio of heavy isotope and light isotope in B phase.



Example:

O isotopic fractionation between Liquid Water and Vapour Water

R l =[ 18O/16O]H2Ol R v =[ 18O/16O]H2Ov



 lv＝ R l /R v

Example:

For water (l) in EQ with vapour (v), the fractionation of O

isotopes in water molecules is:



 l

v = Rl/Rv = 1.0092 at 25oC



for H in water:

 lv = Rl/Rv = 1.074 at 25oC

General Rules for αab

T:  is T dependent, eg.  = f (T). Isotope fractionation factors

are greater at lower temperatures (vibration frequency gos to up

limit)

Reduced- oxidised: Light isotopes are enriched in reduced

species and heavy isotopes are enriched in oxidised species. S

Biogenic: Light isotopes are enriched in biogenic compounds, S

&C

the isotopeδ to express

Isotope δ Value a sample’s isotope composition





The data in Table 17.1 indicate that 18O/16O ratio of average terrestrial is

0.200/99.762 = 0.00200 and that D/H ration is 0.00015. Because it is

difficult to measure small differences in such small ratios, the isotope

composition of H, O, C and S are expressed as the per mille difference of the

isotope ratios of a sample(spl) and a standard(std).



δ spl = (Rspl - Rstd)/Rstd x 1000‰



e.g for O: δ18Ospl = [Rspl-Rstd/R std] x 103 ‰



δ18Ospl = [(18O/16O)spl - (18O/16O)std/(18O/16O) std] x 103 ‰



e.g for S: δ34Sspl = [Rspl-Rstd/R std] x 103 ‰



δ34Sspl = [(34S/32S)spl-(34S/32S)std/(34S/32S) std] x 103

 values express the deviation of Rspl from Rstd in

‰,

Again: δ spl = (Rspl - Rstd)/Rstd x 1000‰

δ18Ospl = [(18O/16O)spl - (18O/16O)std/(18O/16O) std] x 103 ‰





Positive δ values indicate that the sample is enriched

in heavy isotopes relative to standard or depleted in

light.

Negative δ values indicate that the sample is depleted

in heavy isotopes relative to standard or enriched

in light.

For δwe need

Stable Isotope Standards the isotope standards



 H, O: Standard Mean Ocean Water (SMOW) by Craig

(1961).



SMOW: D/H=0.0001558; 18O/16O=0.0020052





 C, O: Belemnite from the Peedee Formation (PDB) in

South Carolina (Craig, 1957)

PDB: 13C/12C=0.0112372





 S: Troilite from the Canon Diablo meteorite (CDT) for

sulfur isotopes (Jensen and Nakai, 1963).



CDT: 34S/32S=0.0450045



 Delta values for the standards?

Table 2-15 and Table 2-16

Summary of Isotope δ Value



We use the isotopeδvalue to express a sample’s

isotope composition

The isotope standards are used for reporting

isotopic compositions;

The isotope compositions of the standards are 0‰.

At the Earth's surface, coexisting compounds of H,

O, C, & S have different isotopic composition

(value ).



What is the relationship of  to  ? (also R)

For phases a and b coexisting in EQ in natural systems,

The relationship between δ values and the isotopic

fractionation factor () is given by:

Equation (2) is usually determine

δa - δb =Δa b≈ 103lna b (1) experimentally

or from natural systems

For C, N, O & S in EQ





103ln  ab = (A x 106/T2) + B ( TºK) (2)

(A & B are Constants)



103lnαab ≈Δab ≈ (A x 106/T2) + B

y =A x+B (Let: y= Δab ; x = 106/T2)

linearity Why? Ref. the text

Superscript p304-305, in detail

subscript

Δab = δa – δb





Δab = 103lnαab





Figure : Temperature dependence of sphalerite-

galena sulphur isotope fractionation factor

Summary for Principles of isotopic fractionation



Isotopic fractionation factor －α



δvalues



δa - δb =Δab ≈ 103lnαab



103lnαab = (A x 106/T2) + B

(A & B are Constants)



103lnαab ≈Δab ≈ (A x 106/T2) + B

2 Oxygen and hydrogen isotopes

 Oxygen has 3 isotopes: 16O, 17O and 18O and their

abundances are 99.756%, 0.039%, 0.205% respectively.

Because 17O is less abundant and closer in mass to the

dominant oxygen isotope 16O, Most measurements

of oxygen isotopes are concerned with the ratio 18O/16O.



 Hydrogen has 2 isotopes: 1H and D and their

abundances are 99.985 % and 0.015%, respectively.



Variations in isotopic composition of O and H caused

mainly by:

1) Evaporation and condensation

2) Water-mineral interaction

3) Mixing of different type of water

O and H isotopes in the hydrosphere

Characters:



1) δ18O & δD of seawater is 0 ‰

2) δ18O & δD for meteoric waters vary



because of fractionation during evaporation

and condensation .



Rayleigh fractionation:



Evaporation: vapour depleted in D & 18O

Condensation: liquid enriched in D & 18O

Figure 2

Figure 17.2 Example of Rayleigh fractionation



Figure 2

Atmospheric vapour has δ values more negative than

atmospheric precipitation



δ values of atmospheric vapour and precipitation become

more negative from coast to inland because atmospheric

vapour and precipitation are derived from the evaporation of

continental water which is already depleted with respect to

heavy isotope



δ values is lower in polar regions than equatorial regions

Figure δD and δ 18O

δ values of vapour & precipitation correlate with latitude

because fractionation increases with decreasing T

Since H and O occur together in water molecules and

since both experience the same sequence of events during the

migration of air masses, the δD and δ18O values of meteoric

water are strongly correlated and satisfy an empirical equation

know as the meteoric-water-line (MWL):



δD=8 δ18O + 10 (Craig, 1961) Figure 4 (MWL”)









But, each local area has its own meteoric water line (LMWL),

which differs from the global average.

See Figure (“MWL”)









Site (north and east of Yucca Mountain) has higher δ

Figure 1 Plot of δD vs δ18O for ground waters of the Ash Meadows and Alkali

Flat/Furnace Creek (Alk./Furn.) flow systems.

surface waters and the meteoric water line (“MWL”) of Craig

(1961).

“LMWL” is the local meteoric water line.. Analytical uncertainty is

±0.2‰ for δ18O and ±2.5‰ for δD.

Isotope reequilibration of wate/rock

Figure5

(minerals) interaction



Underground water types

Meteoric water: derived from rain, snow, rivers, lakes, sea, percolates

through pore spaces in rocks and displaces interstitial water



Connate water: deposited with sediments, and is out of contact with

atmosphere from the time of deposition.



Seawater trapped in pore spaces for marine sediments



Diagenetic water: released from solid phases as a result of mineral

reactions during diagenesis;

e.g., gypsum to anhydrite, smectite to illite.

Formation water exists in layers of sedimentary rock prior to drilling,

regardless of origin



Brine: water with salinity exceeding that of seawater

Oxygen isotope composition of minerals (SMOW)

Oxygen isotope composition of minerals



Table 17.2 Isotope fractionation Factors for H and O between

Clay Minerals--Water at Eath-Surface Temperature





Mineral Hydrogen Oxygen

Montmorillonite 0.94 1.027



Kaolinite 0.97 1.027



Glauconite 0.93 1.026



Illite -- 1.0236



Gibbsite 0.984 1.018

Figure 17.3, 17.4

Fractionation factor (mw) of Mineral –



water is the function of T







Table 2-15 and Table 2-16

Mixing with the water with

different H & O isotope reservoir



Magmatic water



Metamorphic water



Water derived from the meteoric-water



hydrothermal water (W/R interaction & waters mixing )



Figure5, Figure 6

3 Carbon Isotopes





Carbon has 3 isotopes:

12C 98.9%, 13C 1.1%, negligible 14C

Carbon is the principal element in the biosphere, but it occurs also in the

lithosphere, hydrosphere, and atmosphere of the Earth.

Fractionated by : organic & inorganic reactions



Many of these are not equilibrium processes. Values of  13C/12C

range from +4 ‰ (freshwater limestone) to –31 ‰ (petroleum).

Biogenic organic C is lighter than inorganic C, due to fractionation during

photosynthesis and during other biological processes

Table 2-15 and Table 2-16

Sources of C in groundwater:

1) dissolution of carbonate rocks (introduction of heavy

C relative to PDB)

2) oxidation of organic matter (introduction of light C in

CO2)

3) CO2 from soil gas (introduction of light C)





C may be lost from water by the precipitation of

carbonate minerals or loss of CO2 gas

4 Sulfur Isotopes



Sulfur has four stable isotopes 32S, 33S, 34S, 36S

32S (95.0 %) and 34S (4.2%) are the most abundant.







Variations in isotopic composition of S caused by:

1) anaerobic bacteria:

(a) reduction of SO42- to H2S, H2S is enriched in 32S/34S

(b) oxidation of S2- to SO42-, SO42- enriched in 34S/32S

2) isotope exchange reactions: 34S concentrates in species with

highest S oxidation state (-2, 0, +4, +6) or greatest bond strength

3) Redox reactions can produce large fractionations. They may

be inorganically or biologically mediated



Figure 20.1

S isotope composition in geological materials:



1) It is generally assumed that primordial earth sulfur

and the earth’s mantle also have a 32S/34S value of 22.22 (

34S/32S=0).

2) Sulfides from igneous rocks show a narrow range of

values, with most slightly enriched in 34S.

3) Other types of material, particularly sedimentary

sulfates and sulfides from hydrothermal ore deposits, show

wide variations in isotopic composition.



Such variations probably reflect the variety of processes

and geologic conditions involved in forming these sulfides.

Table 2-19 General Range of 32S/34S Ratios and of  34S/32S

Example for S fractionation Figure 28.4







α is the fractionation factor for S partitioned between SO2

and melt. (Fig. 28.4)

When α > 1, 34S is partitioned into SO2 and the melt becomes

depleted, δ34S is negative.

When α < 1, the melt is enriched in 34S, δ34S is positive





The sign of α depends on whether S in the melt is oxidized

(SO42-) or reduced (S2-)

However, S-isotope variations are complex and large ranges

in δ34S are not necessarily well understood.



Equilibrium fractionation is generally inorganic whereas

kinetic fractionation is generally bacterial.



Total S decreases with depth because SO4-2 is more soluble

than S-2, δ34S increases with depth for both SO4-2 and S-2



Near surface: SO4-2 is bacterially reduced to S-2, sulphate is

enriched in 34S.



With burial: more sulphate is reduced to sulphide but as

source sulphate was enriched in 34S, S-2 is enriched in 34S

relative to shallow sediments Figure 28.1

5 Isotope applicatoin



 Isotope geothermometry



δa - δb =Δab ≈ 103lnαab



Δba ≈ 103lnαba = (A x 106/T2) + B ( TºK)

for H, O & S



If EQ has been reached, S, O, and H isotopes can be used



for geothermometry

Δab = δa – δb







Δab ≈ 103lnαab









Δba ≈ 103lnαba = (A x 106/T2) + B ( TºK)

Figure : Temperature dependence of sphalerite-

galena sulphur isotope fractionation factor

δ34S values of pairs of coexisting sulphide minerals increase

with decreasing T (Table 17.5)



δ18O values of pairs of coexisting quartz-other minerals

increase with decreasing T (Fig. 2.19)

Global climate changes



- Global changes in δw: During the last few million years, the earth has

experienced dramatic changes in global climate between glacial and

interglacial periods, resulting in the waxing and waning of the large ice caps

on the North American and Eurasian continent. Since the light isotope 16O is

preferentially vaporized, freshwater and ice are isotopically lighter than

seawater . During ice ages, large amounts of freshwater were locked in the

polar ice caps, resulting in a drop in sea level of about 130m. Because the

130m of water that have been removed from the ocean were isotopically

light, the seawater remaining in ocean basins during

glacial ages was isotopically heavy, i.e. enriched in 18O compared to the

seawater present in the oceans during interglacial periods. This effect

produced heavy δw values during glacial periods reflected in the heavy

δ18O recorded by foraminifera during the “even-numbered” isotopic (or

glacial) stages, and light δw values during interglacial periods recorded

during the “oddnumbered” isotopic stages (Fig. Global changes ).

During ice ages, 130m

of water that have been

removed from the ocean

to form the ice sheet

which were isotopically

light, and the ocean is

enriched in heavy O and

H isotopes. During

interglacial periods. the

ice is back to the

ocean, which result in

it enriched in light

isotopes 16O and 1H .

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