A kinetic model for iron diagenesis in a lake sediment

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GEOCHEMISTRY OF THE EARTH'S SURFACE AND OF MINERAL FORMATION 2nd INTERNATIONAL SYMPOSIUM, July, 2-8, 1990, Alx en Provence,

France.

A K I N E T I C M O D E L F O R I R O N DIAGENESIS IN A L A K E S E D I M E N T

WERSIN P. *, HOHENER P. *, GIOVANOLI R. ** and STUMM W. * • Institute for Aquatic Sciences (EAWAG), Swiss Federal Institute of Technology (ETH), Switzerland. • * Laboratory for Electron Microscopy, University of Bern, Switzerland.

INTRODUCTION We have studied the iron distribution in a calcium-rich anoxic lake sediment in terms of interstitial water and solid matter analysis. Lake Greifen (situated northwest of Zurich, Switzerland) is highly eutrophic due to anthropogenicaily released nuwitions. For interstitial water analysis we used the in situ dialysis method (Hesslein, 1976), for sediment core analysis we used X-ray fluorescence spectroscopy (XRF) and electron microscope techniques.

presence of iron(II) minerals such as siderite or vivianite. Figure 1 illustrates the distribution profile of both solid and dissolved iron and the concentration range of thermodynamically predicted dissolved Fe (II). Since roughly half of the total iron iron is constituted of reactive Fe(III) we imply the importance of slow reaction kinetics which control this distribution. Solid and dissolved iron (log M)

OBSERVED DISTRIBUTION OF DISSOLVED AND PARTICULATE IRON

10-8

10 -7

10 ~o 10 -5 10 -4

10 -3

10 -2

10 -1

10 0

-10

The depth profile for dissolved iron in the anoxic sediment exhibits a non linear increase in concentration with depth. This behavior of iron is in contrast with the distribution profile of other reduced compounds such as Mn(II) and sulfide, which show an immediate increase in concentration at the sediment-water interface from where a decrease to constant values is noted throughout the anoxic sediment. Equilibrium calculations show that the interstitial waters are clearly supersaturated with regard to a variety of Fe(II) mineral phases, e.g. pyrite, vivianite, siderite, iron monosulfides, whereas they are highly undersaturated with regard to the whole range of iron(III) (hydr)oxides. On the other hand XRF analysis shows that the solid iron concentrations are constant throughout the studied profile. Microscopic analysis shows that roughly half of the iron is tied up as detrital iron silicates whereas the remainder mainly builds up X-ray amorphous iron(III) hydroxides. Furthermore, we could not find any evidence for the

Age ~r~ss. Fe

total Fe" 1989

0

depth

-1975

(cm) 10"

1950 20-

Fe2~ -1925

30

Equilibria f predicted range for Fe2 !iiiti~:~i

1900

40

Figure 1 : Measured dissolved and total iron concentration profiles in comparison to calculated equilibrium situation. (*total iron concentrations calculated assuming constant porosity of 0.9 and solid density of 2.5 g cm-3).

GEOCHEMISTRY O F THE EARTH'S SURFACE AND OF MINERAL FORMATION 2nd INTERNATIONAL SYMPOSIUM, July, 2-8, 1990, Alx en Provence, France.

KINETIC MODELING The extent of kinetic control for iron transformation reactions can be visualized by a diagenetic model for dissolved iron (II) (Berner, 1980). Such a model expresses the effects of chemical reactions and transport (molecular diffusion, advection) on depth and time of dissolved iron in the form of a differential equation : ~OC

5 I 6C) ~C = ~'z ~Ds'~"z - Oc°'~"z + OR(z)

where C : Dissolved iron concentration Ds : Bulk sediment diffusion coefficient ~ : Porosity co : Velocity of sediment burial R(z) : Rate of chemical transformation z : Depth (cm) The subscript iw represents interstitial water, and s sediment. Analytical solutions of this equation can be obtained if steady state is assumed (Billen, 1982). Constant transport rates were calculated given (1) the porosity ( ~ ) (measured), (2) the velocity of sedimentation (co) which is derived from the sedimentation rate (Wan et al., 1987), (3) diffusion coefficient (Ds) which can be calculated from the molecular diffusivity of Fe 2+ (Li and Gregory, 1974; Ullman and Aller, 1982). Given the transport rates, the rates of chemical reactions (i.e. dissolution, precipitation) were then derived by assuming two extreme cases in order to have an upper and lower boundary for these reactions. In the first case the reductive dissolution of Fe(III) is assumed to be constant with depth whereas the precipitation of Fe(II) would follow first order kinetics. In the second case it is assumed that reductive dissolution decreases exponentially (first order) simultaneously to the observed exponential decrease of organic matter. In this case Fe(II) precipitation is completely inhibited. From solutions of these two cases upper and lower limits of rates of iron transformation reactions can be inferred. We derived a maximum dissolution rate of 0.07 lamol g-1 (dry sediment) yr-1. Thus, at a

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depth of 25 cm complete dissolution of Fe(III) hydroxides would take about 550 years. This calculated slow dissolution reaction is in agreement with the observed presence of Fe(III) hydroxides and in addition is supported by findings from dissolution experiments of iron oxides in the presence of organic ligands (Sulzberger et al., 1989). Furthermore, the diagenetic model allows estimates for precipitation rates of Fe(II) minerals. An upper limit for a rate constant of a first order reaction would be 0.04 yr -1. If we assume that only one mineral e.g. siderite would precipitate this would yield a maximum precipitation rate of 3 mg g-l (dry matter) y r 1 at a depth of 25 cm, which is only a very small fraction of the total iron in the sediment. This calculation agrees with the fact that neither siderite nor vivianite could be detected by microscopic techniques. It also confirms the observation that high supersaturations of Fe 2+ with respect to siderite and vivianite are built up in the interstitial waters of the sediments of Lake Greifen. REFERENCES BERNER R.A. (1980) Early diagenesis. A theoretical approach. Princeton University Press, 241 p. BILLEN G. (1982) Modelling the processes of organic matter degradation and nutrients recycling in sedimentary systems. In: NEDWELL DB & BROWN CM (eds.) : Sediment Microbiology. Academic Press, London, pp 15-52. HESSLEIN R.H. (1976) An in situ sampler for close interval pore water studies. Limnol. Ocenogr.21,912-915. LI Y.H. and GREGORY S. (1974) Diffusion of ions in sea

water and in deep-seasediments.Geochim.Cosmochim.Acta 38, 703-714. SULZBERGER B., SUTER D., SIFFERT C., BANWART S. and STUMMW. (1989) Dissolutionof Fe(III) (hydr)oxides in natural waters; laboratory assessment on the kinetics controlledby surfacespeciation. MarineChem. (in press). ULLMAN W.J. and ALLER R.C. (1982) Diffusion coefficients in nearshore marine sediments. Limnol.

Oceanogr. 27, 552-556. WAN G.J., SANTSCHI P.H., STURM M., FARRENKOTHEN K., LUECK A., WERTH E. and SCHULER C. (1987) Natural (210pb, 7Be) and fallout (137Cs, 239,240pu, 9°Sr) radionucfidesas geochemical tracers of sedimentation in Greifensee, Switzerland. Chem. Geol. 63, 181-196.