Iron in the Dead Sea

Earth and Planetary Science Letters, 71 (1984) 405-414

405

Elsevier Science Publishers B.V., Amsterdam - Printed in The Netherlands

[3]

Iron in the Dead Sea A. Nishri and M. Stiller Isotope Department, The Weizmann Institute of Science, Rehovot (Israel)

Received October 18, 1982 Revised version received November 29, 1984

The distributions of dissolved and of particulate iron in the Dead Sea during the period which preceeded its overturn and thereafter (1977-1980) are reported. During 1977-1978, the vertical profiles of the iron phases revealed facets of the mixing pattern: the progressive deepening of the pycnocline, restricted mixing within the upper water mass and penetration of surface waters into the deepest layer. The inventories of particulate iron suggest resuspension of bottom sediments in November 1978 and after the overturn the gradual disappearance from the water column of iron sulfides and iron oxy-hydroxides. Fluxes of iron from and to the lake in the undisturbed meromictic Dead Sea have been estimated: it appears that diffusion of divalent iron from bottom sediments was the major source for the standing crop of iron in the lower water mass. Low settling velocities of solid particles in the dense and viscous Dead Sea is one of the causes for the relatively large concentrations of particulate iron. The rate constant for oxidation of divalent iron in Dead Sea sediment interstitial waters is larger by two orders of magnitude than in other natural waters.

1. Introduction

The Dead Sea is a hypersaline chloride lake. Its water column was permanently stratified, i.e. meromictic, until February 1979 when a complete overturn was observed [1]. The stratified Dead Sea contained two distinct water bodies separated by a transition zone; an oxic upper water mass (UWM) and a denser anoxic lower water mass (LWM). During the seventies the density of the UWM increased gradually approaching that of the LWM. Until 1976 [2] the location of the permanent pycnocline separating the top of the LWM from the UWM was found to be at about the same depth (80-100 m) as described by Neev and Emery [3] for the early sixties. During the winter of 1976/77 an erosion took place at the top of the anoxic layer and the density profile in March 1977 revealed that a new, deeper, transition zone between 100 and 140 m had been formed. Futher erosion during the winter of 1977/78 and during 1978 finally led to a complete 0012-821X/84/$03.00

© 1984 Elsevier Science Publishers B.V.

overturn of the water column in the winter of 1978/79. The anoxic conditions which prevailed in the LWM were indicated by a strong odor of H2S. However, the measured concentrations of H2S reported for about the same period differ largely: Neev and Emery [3] found about 4 × 10 -4 mol 1 t in 1959/60 whereas in 1963-65 Nissembaum [4] reported about 2 × 10 -5 mol 1-1. Neev and Emery [3] also reported that ropes (on which sediment traps were hung) submerged in the LWM, blackened supposedly by coating with iron sulphides. It seems therefore that the presence of dissolved divalent iron in the anoxic LWM could be presumed. The steady state distribution of iron species which had probably prevailed until the mid seventies has been affected by the gradual erosions from the top of the anoxic LWM which took place during 1976-1978. In this study we present the behaviour of iron in the Dead Sea during stages of pycnocline deepening, at overturn and shortly thereafter. Also,

406

relying on the iron profile of March 1977, as most closely representing the undisturbed stratified Dead Sea, we attempt to estimate the iron fluxes which maintained the pre-1976 steady state distribution of iron species in the Dead Sea.

O,W,

DISSOLVED Fe(l]) (/zmole kg-i) 400 800 1200

I0

2. Sample collection and analysis "BATHO" ~ / ~ . ~ " / Samples for this investigation were collected during 12 cruises, between March 1977 and May 1980, at a deep basin station, located at 105°N, 195.5°E (local coordinates). Water samples were collected in 5-liter PVC Niskin bottles. Dissolved divalent iron was "fixed" on board as follows: the brine was allowed to flow directly from the Niskin sampler into a 30-ml glass "reaction bottle" which already contained 0.5 ml of 0.1% bathophenantroline disulfonic acid disodium solution. In the presence of Fe(II) a pink color appears instantaneously. We assume that there was no substantial oxidation of divalent iron during transfer of the sample from the Niskin bottle into the "reaction bottle", as the "half time" for oxidation in Dead Sea pore waters is about 13.5 hours (see also next sections). The amount of dissolved divalent iron was determined spectrophotometrically within 3 days after sampling (assuming that the bathophenantroline reagent does not react with particulate iron during this period), by measuring the absorbance of the bathophenantroline-divalent iron complex at 535 ffm with a Cary No. 14 model Spectrophotometer, in 5-cm light-path cells. The absorbance which is due to the natural turbidity of the Dead Sea brines was measured for each sample (at 535 #m) and subtracted from the absorbance of the bathophenantroline complex. The limit of detection for this procedure is about 0.4 /~mol kg 1, the precision better than 5% and the color of the complex was found to be stable for about one week. Comparison of Fe(II) analysis of filtered pore waters by the above method of complex formation and subsequent spectrophotometric determination and by immediate acidification with nitric acid and subsequent analysis by atomic absorbtion spectrophotometry (AAS) proved to be satisfactory as shown in Fig. 1.

E "II-IX. hi tm

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Fig. 1. Comparison of Fe(II) concentrations in pore waters of core GF measured both by AAS and by spectrophotometric determination with bathophenantoline disulfonic disodium salt. Core G F was taken from the deep basin in April 1978. The concentration of Fe(II) in waters overlying the sediment (O.W.) in core GF (about 1 cm) is marked by an arrow.

Separation of particulate iron by immediate filtration of the brines was not possible on board. Thus, we prefered to first allow complete oxidation of dissolved Fe(II) to particulate iron and then perform and analysis of the total iron content of the sample. Therefore our procedure for the analysis of particulate iron was as following: a separate l-liter aliquot taken from the Niskin sampler, was stored in polyethylene bottles for one week or more, during which complete oxidation of dissolved Fe(II) present in the brine was found to have been accomplished. Then the brines were vacuum filtered through pre-weighed 0.45 /~m, 47mm Millipore HA membranes. After filtration, the pre-weighed membrane filters were flushed with double distilled water to remove entrapped Dead Sea brines, dried in a vacuum over for 48 hours at 25°C, and stored in a desiccator at room temperature until being reweighed in order to obtain the

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Density (g cm ~) Fig. 2. Vertical profiles of dissolved Fe(ll) ( x ) and of particulate iron, Fe* (O) from February 1978 to May 1980. In March 1977 only total iron, Fe T (A) has been measured. The density profiles at the respective sampling dates [2] are also shown by solid lines.

408 weight of suspended matter. Unfortunately, this conventional technique was found to be inadequate because efficient flushing of the brines from the filters was not feasible. By immersing the filters in 8 M H N O 3 and heating on a plate (in covered beakers) for about half an hour, leaching of iron phases was obtained along with complete dissolution of the filter membranes. The iron content, Fe T, was measured by AAS. For each set of samples, blanks were run in parallel. The particulate iron, Fe*, was calculated by subtraction of dissolved Fe(II) from the total iron content, F e r , measured by AAS. The reproducibility was about _+7%.

3. Vertical profiles of iron in the Dead Sea The vertical profiles of dissolved Fe(II) and of particulate iron, Fe* measured during 1978-1980 are shown in Fig. 2, along with the density profiles at the respective sampling dates [2]. Fe(II) was not measured in March 1977 and therefore only the total iron, Fe T profile, is presented in Fig. 2. The dissolved Fe(II) concentrations in the LWM, during February-November 1978, were about 4.7/~mol kg 1 (ranging from about 2 ~mol kg -1 to 9.5 ~mol kg 1; the largest concentrations of Fe(II) were detected at 275 m, 6.9 and 9.5/~mol kg a in August and November 1978, respectively. The presence of dissolved Fe(II) in the LWM indicates the prevailance of anoxic conditions until the very last stages of meromixis, i.e., in spite of the small density gradients between the UWM and the LWM, the LWM was still isolated. The total disappearance of Fe(II) from the water column coincided with the turnover observed in February 1979. Ever since, dissolved Fe(II) has not been detected in the Dead Sea. Before the overturn, the concentration of Fe* in the LWM ranged from about 8.8 to 14.5 ~mol kg-1, while in the UWM it ranged between 3.6 to 7.3 /~mol k g - k After February 1979, the vertical profiles of particulate iron became more evenly distributed within the entire water column with concentrations gradually decreasing from about 8 ~mol kg i in March 1979 to 4.4 gmol kg -1 in May 1980.

The location of the depth zone in which large concentration gradients of Fe(II) and of Fe* were observed are compared in Table 1 with the depths of the pycnocline (the one separating the top of the fossil LWM from overlying waters). Although the March 1977 and April 1978 profiles present limited information ,1, several features can be observed in Table 1: (1) The transition layer indicated by the Fe(II) concentration gradient seems to be, with one exception in February 1978, at an identical depth as the pycnocline layer which separates the deep LWM from overlying waters. In fact, during 1978, the on-site detection of the Fe(II) transition zone served as a practical tracer for the stratified condition of the Dead Sea and its resolution compared well with subsequent shore measurements of other tracers. (2) The progressive erosion of the meromictic structure, between August 1978 and November 1978 is indicated by the successive deepening of the Fe(II) devoid zone. (3) The distinct increase of particulate iron is always observed at a relatively shallower depth than the distinct increase of dissolved Fe(II). Following a stage of deepening of the pycnocline, local oxidation of Fe(II) formerly present in the LWM is occurring. However, this local oxidation of Fe(II) can account only for a part of the increase in Fe*. The rest of it is attributed to the relatively larger amounts of Fe* contained within the deep water which mix with the overlying layers. This mixing is restricted to a certain adjacent layer of overlying waters and not to the entire UWM. The above suggested concept of restricted (incomplete) mixing within the UWM is also supported by the shape of the tritium and of the 228Ra profiles in the UWM and in the transition zone [5,6]. A characteristic feature observed in all four Fe(II) profiles is a decrease in its concentration at the bottom of the LWM (Fig. 2). It is suggested that this may have been caused by surface waters ,1 The first because no Fe(II) determinations have been performed and the second because technical difficultiesarose on board and prevented adequate sampling within the transition zone.

409 TABLE 1 Comparison of the depths with large concentration gradients for dissolved iron, particulate iron and density Sampling date

Depth range of layer with large gradient (m) Fe*

Fe(II)

density a

Mar. 29, 1977 Feb. 20-21,1978 Apr. 17,1978 Aug. 16,1978 Nov. 10, 1978 Feb. 8,1979

n.d. 120-140 180-225 100 150 150-175

n.d. 160-180 180-225 170-200 200-225 n.d.

100-140 170-180 170-180 170-180 - 190-200 none

Remarks

Fe T transition zone at 130 170m poor resolution, no sampling between 180 and 225 m

overturn

a Pycnocline between the U W M and the top of the fossil L W M [2].

which penetrate to this depth. Supporting evidence for the penetration of surface waters into the deepest layers in 1978 (at 300 m and below) is also provided by the presence of tritium with activities of about 1-1.5 TU in these layers while in the rest of the LWM, tritium activities of less than 1 TU have been measured [5]. Therefore, we believe that the relatively lower Fe(II) concentrations in the deepest layers can be explained by mixing (and subsequent oxidation) with surface waters which contain tritium and oxygen. We estimated the amount of oxygen needed to oxidize the "missing Fe(II)" in the deepest layers for each profile separately, by assuming that oxygen-loaded surface waters (44/~mol 02 1 t; [7]) reached the bottom without losing any oxygen on its way. It is thus calculated that the deepest layers, from 280 m to bottom, were "diluted" by 1.2% in February 1978, 1.9% in April 1978, 1.8% in August 1978 and 3.5% in November 1978, by intrusion of surface waters. This is a minimal estimate since part of the oxygen could have been consumed for oxidation of other species, such as organic matter, HzS and also some of the "lost" Fe(II) might have been replenished by upward diffusion of Fe(II) from the pore waters. By interpreting the tritium content of the deepest water layers Steinhorn [2] obtained similar "intrusion" estimates. 4. The rate of oxidation of Fe(II) in the Dead Sea pore waters

The kinetics of oxidation of Fe(II) were found to be described by: - d F e ( I I ) / d t = kpO2a~m-Fe(II)

where k is the rate constant (1: mol 2 atm-1 min 1) and pO: is the partial pressure of oxygen [8,9]; hence, the rate of oxidation is strongly pH dependent. In the Dead Sea, the oxidation rate of Fe(II) is expected to be rather low since its pH ranges between 5.9 and 6.25. We measured the rate of oxidation of Fe(II) in filtered interstitial waters, extracted from sediments of the deep basin, which had an initial concentration of dissolved Fe(II) of 539 /~mol kg-1 and a pH of 6.0. Assuming the above kinetics to hold for our experiment, we obtained at constant p O 2 and pH, k aZiq pO 2 = 8.1 × 10 4 min-1 (half-life of 820 minutes). Then, k was calculated by assuming that one may refer to the pH measured in Dead Sea waters by a combined glass calomel electrode as to the pH measured in solutions of much lower ionic strength and by using a value of 0.75 [10] for the activity of H 2 0 , aH20, in the Dead Sea. The magnitude of k was found to be 7.2 × 1013 12 mo1-2 atm 1 m i n - l . This figure is surprisingly close to the k value measured by Singer and Stumm [9] for HCO 3 solutions 8.0 x 1013 12 mo1-2 atm 1 min-a and is about two orders of magnitude larger than the k's measured in natural waters [11,12]. The rather large k value obtained for the Dead Sea pore waters might be attributed to the speciation of Fe(II) in this highly saline interstitial waters (which is yet not known), whereas the very low rate observed for the oxidation of Fe(II) in these waters is attributed to its dependency on a~)H-. By assuming that the k measured on Dead Sea pore waters also represents the rate constant for oxidation in overlying brines it is probable that only negligible amounts of Fe(II) might have been

410 TABLE 2 The inventory of iron in the Dead Sea during 1977-1980 Date of sampling

Dissolved iron, Fe(II) (ton/lake)

Particulate iron, Fe* (ton/lake)

Total iron, FeT (ton/lake)

Fe(II)/Fe*

Mar. 29,1977 Feb. 20 21,1978 Apr. 17, 1978 Aug. 16, 1978 Nov. 10, 1978 Feb. 8, 1979 Mar. 25,1979 Jun. 28,1979 Oct. 2, 1979 Dec. 12, 1979 Mar. 5, 1980 May 26, 1980

n.m. 17,860 13,700 16,020 14,000 n.d. n.d. n.d. n.d. n.d. n.d. n.d.

n.m. 70,500 68,700 52,900 88,600 86,300 81,500 61,300 65,300 53,900 46,100 39,200

98,500 88,400 82,400 68,900 102,600 86,300 81,500 61,300 65,300 53,900 46,100 39,200

0.44 0.37 0.38 0.56 0.38

a

(140 (160 (200 (170 (200

m bottom) m-bottom) m-bottom) m bottom) m bottom)

n.m. = not measured; n.d. = not detected, below detection limit. The ratio for March 1977 has been estimated as follows: (a) the shape of the pycnocline and 21°Pb [13] indicate that Fe(ll) was probably present below 140 m: if its concentration was about 4.7/Lmol kg ] as in the fossil LWM in February 1978, then 20,800 ton Fe(II) were present in the LWM of March 1977; (b) above 140 m, all the measured Fe T is assumed to be Fe*, i.e. 30,000 ton; (c) Fe* from 140 m to bottom was calculated by subtracting the above 20,800 ton Fe(ll) and 30,00 ton Fe* from the total inventory of 98,500 and is 47,700 tons.

lost by oxidation during the short time lapse between sampling and fixation on board. On the other hand, the oxidation rate is fast enough to have caused disappearance of dissolved Fe(II) (below our detection limit) from the Dead Sea within about 3 days, when complete turnover took place.

5. The inventory of iron in the Dead Sea The inventories of iron were calculated by assuming a volume of 146 km 3 at a lake level of - 4 0 2 m MSL. In fact the lake level decreased from - 4 0 2 . 2 m in February 1978 to - 4 0 3 . 5 m in December 1979, and rose again to about - 4 0 2 m in May 1980, involving a maximal change in volume of about 1.1 km 3 (which is negligible compared to the precision of our inventories). As shown by Table 2, large variations in the inventories of particulate iron are observed during 1977-1980. The total amount of allochthonous material annually brought to the Dead Sea by the Jordan and by floods is about 2 × 10 5 to 3 × 10 5 tons (Stiller, unpublished). With concentrations of about 1.5% Fe [7] in the detrital material the yearly input of

allochthonous iron should not be more than about 4000 tons. Hence, seasonal variations of allochthonous inputs cannot account for the large variations, of more than 10,000 tons iron, between the inventories of two successive samplings. It follows that these variations are probably related to the successive deepenings of the pycnocline, to the turnover event and to the different chemical behavior of iron in the entirely oxic post-turnover Dead Sea. Comparison of the Fe* inventory of February 1978 with that of August 1978,2 indicates a decrease of 17,600 tons by August 1978. The loss of Fe* is most pronounced in the deeper layers (see Fig. 3). It should be noted that the pycnocline zone had deepened by about 40 m in winter of 1977/78, from 100 140 m to 170-180 m [2]. Within this zone about 3000 tons of Fe(II) have then been oxidized and also the production of FeS within this zone has ceased. We believe that the decrease in the Fe* inventory of August 1978, as compared to that of February 1978 is due to the ,2 The inventory of April 1978 is not discussed because it might be inaccurate. We did not have data points between 180 and 225 m and we used interpolated values for this zone.

411

presumed formation (during winter of 1977/78) and subsequent settling of iron oxyhydroxides and also to the gradual disappearance of iron sulfides from the newly oxidized zones. The disappearance in August 1978 of about 1800 tons from the February 1978 inventory of Fe(II) was obviously caused by the oxidation of Fe(II) in the 160-175 m layer. November 1978 is characterized by a sudden rise in the inventory of Fe* (Table 2). An increase in Fe* concentrations is observed all along the water column (Fig. 2), and is most pronounced in layers adjacent to the UWM-LWM interface (Fig. 3). In the 150-200 m layer, part of this increase can be attributed to the deepening of the pycnocline'and subsequent oxidation of Fe(II) within the layer 175-200 m. We speculate that the overall increase in Fe* in November 1978 is due to re-

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suspension of the top few millimeters of sediments, from bottom surfaces underlying the zone of pycnocline deepening. It could be that the top few millimeters of the sediment became enriched with iron oxyhydroxides formed during the 1976-1978 destratifications. In February 1979, there was no more Fe(II) in the entire water column and all of it has probably been transformed into iron oxyhydroxides. As the inventory of Fe* had already partly declined in comparison to that of November 1978, we suggest that complete destruction of the meromictic structure must have occurred a few weeks after the November 1978 cruise. (The February 1979 sampling can be regarded as a post-overturn cruise and not as "the overturn cruise".) However, complete homogenization of the water layers was observed only in March 1979 (Fig. 2). Since March 1979 the inventories of Fe* ( = FeT) gradually decreased (Table 2). After the overturn no more Fe* (iron sulfide and iron oxyhydroxide), was produced within the lake. One would therefore expect the post overturn inventories to lose all of the iron sulphides, iron oxyhydroxides and the material assumed to have been resuspended in November 1978. Hence, with the exception of the October 1979 inventory which may be inexact because of few data points, the general trend of diminishing inventories from March 1979 to May 1980 seems reasonable.

6. An estimate of the steady state fluxes of iron species in the Dead Sea before the destruction of meromixis

F A J A 0'1]~1g A J A ODIF'A 1978

1979

1980

Fig. 3. Mean concentrations of particulate iron (weighted means) in the water layers of the Dead Sea vs. time.

Until 1976, the location of the permanent pycnocline separating the top of the LWM from the UWM was at about the same depth (80-100 m) as in the early sixties. It is probable that until 1976 concentrations of elements controlled by redox reactions were at steady state in both water masses. During the winter of 1976/77 an erosion of the top of the anoxic layer took place. The density profile of March 1977 revealed a new deeper transition zone located between 100 and 140 m. Despite this recent erosion of the stable meromictic structure we had no better choice than

412

to refer to the iron profile of March 1977 as the one which most closely represents the steady state iron distribution, which prevailed in the undisturbed, permanently stratified, Dead Sea. It is assumed that before the start of pycnocline deepenings, i.e., before winter of 1976/77, particulate iron within the UWM was mostly of terrigenous origin. On the other hand, in the LWM, particulate iron was essentially iron sulfide produced by interaction between dissolved divalent iron and hydrogen sulfide. In March 1977, the measured particulate iron inventory of 30,000 tons in the UWM within the depths 0-140 m (see footnote, Table 2) was actually composed of 3 components: (a) approximately 6350 tons of Fe(II) were formerly present within the 100-140 m depth range (obtained by nmltiplying the average Fe(II) concentration of 4.7 /~mol kg-1 Fe(II) by 19.86 km 3, the volume between 100 and 140 m depth) and have been oxidized during winter 1976/77; (b) if between 100 and 140 m the ratio Fe(II)/Fe*, (Fe* present as FeS), was about 0.44, the same as the ratio Fe(II)/Fe* in March 1977 below 140 m (Table 2), then about 14,400 tons Fe* were present, supposedly as iron sulfides. (By taking the above mentioned ratio of 0.44, it is implicitly assumed that all Fe* below 140 m is iron sulfide and Fe* of terrigeneous origin is thus disregarded. In fact, the true ratio Fe(II)/(Fe* present as FeS) should be somewhat larger.) As the March 1977 sampling was performed shortly after the pycnocline deepening of winter 1976/77, we assume that the entire amounts of components (a) and (b) were still present above 140 m. (c) The remaining 9200 tons represent the standing crop of particulate iron of terrigenous origin in the UWM, in March 1977. It follows that its mean residence time in the upper 140 m was about 2.3 years, since the maximum annual input of terrigenous iron is approximately 4000 tons. A mean residence time of 2.3 years implies a settling velocity of allochthonous particles, v, of about 44 m yr-~ (calculated by v = h/~'p, where h is the mean depth to 140 m and ~'p is the mean residence time with respect to settling of particles). This rather low settling velocity is not surprising; due to the large density (1.23 g cm 3) and viscosity (about 3.1 centipoise at 22°C, [7]) of the Dead Sea, particles

settle in this environment with velocities which are about 5 times lower than in a fresh water lake. It should be mentioned that a similar settling velocity, of about 50 m yr -1 was derived [13] for 21°Po-loaded particles. It seems reasonable to assume that in March 1977 the distribution of iron species within the LWM, below 140 m, was still at a steady state. The fluxes of particulate iron into and out of the LWM and of dissolved Fe(II) into the anoxic LWM required to maintain the iron distribution of March 1977 and also relevant to the former, pre1976, undisturbed steady state will now be estimated. At steady-state, the fluxes of iron into the LWM should balance the fluxes of iron leaving the LWM, as described schematically by: F [Fe(II)] + F[Fe*]uwM = F[Fe*]LWM + F [Fe(II)]

L W M ~ UWM

F[Fe(II)] denotes the upward flux of divalent iron from the sediments into the LWM which is unknown. F[Fe*]vwM denotes the flux of particulate iron settling from the UWM to the LWM, F[Fe*]LwM is the flux of particulate iron from the LWM to the sediments and F[Fe(II)]LwM~vw M is the upward diffusion of dissolved Fe(II) across the pycnocline. The eddy diffusion coefficients across the pycnocline of the meromictic Dead Sea are expected to have been very small in view of the large and stable density gradients (about 1 × 10 -5 g cm 4) across this layer. The upward diffusion of Fe(II) depends upon the eddy diffusion coefficent and the dissolved Fe(II) gradient across the pycnocline zone. To estimate K, the eddy diffusion coefficient in the pycnocline zone, we have used the empirical relationship [14] between the eddy diffusion coefficient and the Brunt-V~is~ilh f r e q u e n c y N 2 ( N 2 = - g lip ~p/~Z; it is a measure for the intensity of stratification), K = 0.178 exp ( - 5 5 . 8 N). For N 2 = 8 . 8 6 × 1 0 -3 s 2, K = 9.4 × 10 4 cm 2 S I. With a K of about 10 -3 cm 2 s -1 and a gradient of about 4.7 #mol Fe(II) kg 1 across 30 m ( = 1.1 × 10 - 4 #g cm-4), an upward flux of about 17 ton Fe(II) yr -a (through an interface of about 500 km 2) is obtained. In comparison with the other iron fluxes (see below) this flux is negligible.

413

As dictated by the morphology of the lake bottom, the amount of particulate iron settling from the UWM and actually reaching the LWM could be only about 60% of the 4000 tons (upper limit); hence the annual supply of allochthonous iron, i.e. F[Fe*]uWM amounts to about 2400 tons yr -1 or less. F[Fe*]LWM, the flux of particulate iron settling from the LWM, is in fact equal to Fe[wM/~"p, i.e. the ratio between the standing crop of particulate iron in the LWM, FeLWM, and its mean residence time, rp, in the LWM. Stiller and Kaufman [13] have estimated a residence time, ~p of about 4.4 years for particulate 21°pb in the LWM (below 140 m) in March 1977 and suggested that particulate 2]°pb settling through the LWM was adsorbed mainly onto iron sulfide particles. The standing crop of particulate iron in March 1977 in the LWM was estimated to have been about 47,700 tons (see footnote, Table 2). It follows that the settling flux of particulate iron at the lake bottom, F[Fe*]LWM, was about 10,800 tons yr -1. In order to maintain this settling flux, the upward flux of dissolved Fe(II), from the sediment pore-water system, F[Fe(II)] (see equation (1)) should have been about 8400 tons yr -1. An upward flux of 8400 tons Fe(II) per year is equivalent to 5 . 6 x 1 0 5 / ~ g c m - 2 s - f o r 1 f t m o l c m 2 s -1 (the anoxic bottom area below 120 m is 495 km 2). The fluxes of iron deduced from the data of march 1977 are summarized in Fig. 4. The residence time (~a) for dissolved Fe(II) in the LWM was about 2.5 years (20,800 tons Fe(II)/8400 tons yr-1). An independent estimate of F[Fe(II)], through the sediment-water interface is performed by asINFLOW 4000

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(BELOW 140 M)

Fig. 4. Schematic representation of major iron fluxes (tons yr -1) in the Dead Sea, in March 1977.

suming that this flux is described by Fick's first law F[Fe(II)] = - dPDbAC/AZ.The concentration gradient across the sediment-water interface AC/AZ is about 0.14 /~mol Fe(II) cm -4 (Fig. 1) and the porosity q~, at the sediment-water interface is assumed to be close to 1.0 [15]. O b , the bulk sediment diffusion coefficient is given approximately by O b = ~bzOi [16], where D i is the diffusion coefficient at infinite dilution. In this particular case, as ff = 1, D b = D i = 6.59 × 1 0 . 6 c m 2 S - 1 [15] at 22°C (the temperature of the deep basin sediments). It follows that F[Fe(II)] = 0.92/~mol cm 2 s-1, and compares reasonably well with our previous estimate. It should be noted, however, that we have used the diffusion coefficient of Fe(II) at infinite dilution, because the effect upon it of the Dead Sea salinity is unknown; a decrease is expected with an increase of salinity [17]. In the above calculation, the possible transformation of F[Fe*]uwM into Fe(II) while settling through the LWM was disregarded; if this process actually happens it may only slightly reduce the residence time of dissolved Fe(II) within the LWM. Despite the inherent approximations involved in the above calculations we can safely assume that the distribution of iron species in the LWM of the meromictic Dead Sea could not have been maintained by external fluxes of terrigenous iron and that internal recycling between the sediments and the LWM played an important role. After the overturn the formation of iron sulfides within the deep waters probably ceased, as penetration of oxygen took place all through the water column. The gradually diminishing inventories of iron during 1979 and 1980 confirm this expectation. Also a thin, reddish-brownish layer of iron oxyhydroxides was observed at the sediment water interface of the deep basin sediments, in cores taken after February 1979. One may forecast that if oxic conditions will persist for several years in the entire lake, a new "steady state" inventory, supported exclusively by allochthonous inputs will be approached. We estimate that this new steady state inventory will be much smaller than the inventories of the meromictic Dead Sea, only about 15,000-20,000 tons of particulate iron. The physical properties of the Dead Sea waters (high density and viscosity) attenuate the settling

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velocities of solid particles. Therefore, the inventories of particulate iron in the stratified Dead Sea, as well as in the completely oxic, post-turnover Dead Sea are about 5 times larger than in any freshwater lake of comparable dimensions and input fluxes.

Acknowledgements We thank Prof. J.R. Gat and Dr. A. Nissenbaum for critically reading the manuscript. The technical assistance of Mrs. N. Bauman is acknowledged. This research was supported by grant No. 936 from the U.S.-Israel Binational Science Foundation (BSF), Jerusalem, Israel.

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