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Lithium isotope geochemistry of pore waters from ocean drilling program Sites 918 and 919, Irminger Basin
Geochimica et Cosmochimica Acta, Vol. 62, No. 14, pp. 2437–2450, 1998 Copyright © 1998 Elsevier Science Ltd Printed in the USA. All rights reserved 0016-7037/98 $19.00 1 .00
Pergamon
PII S0016-7037(98)00178-1
Lithium isotope geochemistry of pore waters from Ocean Drilling Program Sites 918 and 919, Irminger Basin LIBO ZHANG, LUI-HEUNG CHAN,*,1 and JORIS M. GIESKES2 1
Department of Geology and Geophysics, Louisiana State University, Baton Rouge, Louisiana 70803, USA 2 Scripps Institution of Oceanography, University of California, La Jolla, California 92093-0215, USA (Received November 12, 1997; accepted in revised form May 8, 1998)
Abstract—The distribution of Li isotopes in pore waters to a depth of 1157 m below seafloor is presented for ODP Sites 918 and 919 in the Irminger Basin, offshore Greenland. Lithium isotope data are accompanied by strontium isotope ratios to decipher diagenetic reactions in the sediments which are characterized by the pervasive presence of volcanic material, as well as by very high accumulation rates in the upper section. The lowering of the 87Sr/86Sr ratio below contemporaneous seawater values indicates several zones of volcanic material alteration. The Li isotope profiles are complex suggesting a variety of exchange reactions with the solid phases. These include cation exchange with NH1 4 and mobilization from sediments at depth, in addition to the alteration of volcanic matter. Lithium isotopes are, therefore, a sensitive indicator of sediment-water interaction. d6Li values of pore waters at these two sites vary between 242 and 225 ‰. At shallow depths (,100 mbsf), rapid decreases in the Li concentration, accompanied by a shift to heavier isotopic compositions, indicate uptake of Li into alteration products. A positive anomaly of d6Li observed at both sites is coincident with the NH1 4 maximum produced by organic matter decomposition and may be related to ion exchange of Li from the sediments by NH1 4 . In the lower sediment column at Site 918, dissolved Li increases with depth and is characterized by enrichment of 6Li. The Li isotopic compositions of both the waters and the solid phase suggest that the enrichment of Li in deep interstitial waters is a result of release from pelagic sediments. The significance of sediment diagenesis and adsorption as sinks of oceanic Li is evaluated. The maximum diffusive flux into the sediment due to volcanic matter alteration can be no more than 5% of the combined inputs from rivers and submarine hydrothermal solutions. Adsorption on to sediments can only account for 5–10% of the total inputs from rivers and submarine hot springs. Copyright © 1998 Elsevier Science Ltd concentrations show a rapid decrease by 10 mM within the top 10 m followed by an increase to approximately six times greater than the seawater value in the deeper section (Gieskes et al., 1998). In this paper, we report detailed Li isotope profiles in the interstitial waters at ODP Sites 918/919 and discuss the exchange processes affecting pore water Li distribution. The Li isotope study is accompanied by strontium isotope measurements which provide additional constraints on the sources and sinks of Li. This study is the first systematic attempt to understand the Li isotope geochemisry in sediment pore waters and aims at gaining an understanding of the Li exchange between the fluid and solid phases as well as the role of sediments in the marine budget of Li. The Li budget in the oceans has not been completely resolved. River flux, hydrothermal input, and fluid expulsion at convergent margins are three identified sources of Li in the oceans (Stoffyn-Egli and Mackenzie, 1984; Martin et al., 1991; You et al., 1995). The only well-defined sink in the oceans is low temperature alteration of oceanic crust, the magnitude of which is much smaller than the total input. Since marine sediments are enriched in Li in comparison to freshwater sediments (Holland, 1984), they are thought to be an important sink for oceanic Li. Lithium adsorption onto marine sediments (Seyfried et al., 1984) and incorporation into authigenic clay minerals (Stoffyn-Egli and Mackenzie, 1984) have been proposed as important removal processes of Li in the oceans. The rapid decrease in the uppermost layer at ODP 918/919 (Gieskes et al., 1998) suggests that Li has been removed from the pore
INTRODUCTION
Recently, Li and Li isotopes have been used to understand many geological processes in the oceans such as mid-ocean ridge hydrothermal activity, alteration of the oceanic crust, and fluid expulsion at convergent margins (Chan et al., 1992, 1993, 1994a; You et al., 1995). Relatively little, however, is known about the Li isotope geochemistry in the interstitial waters of marine sediments, even though they occupy roughly half of the marine sediment volume and constitute approximately 5% of the Earth’s hydrosphere (Lawrence, 1988). A previous Li isotope study of pore waters was carried out at Ocean Drilling Program (ODP) Site 808, Nankai Trough (You et al., 1995). At this site, a Li isotope anomaly at the de´collement zone provides strong evidence for the advection of fluids from deeper parts of the accretionary prism. The isotopic measurements were made by pooling samples from adjacent depths thus sacrificing depth resolution. This was due to sample size requirements for Li isotope analysis (;3.5 mg) by the most reliable mass spectrometric method (Chan, 1987). However, a recent improvement in the analytical sensitivity (You and Chan, 1996) has enabled Li isotopic determination of much smaller natural samples, thus permitting high resolution studies of pore waters. Our present study focuses on ODP Sites 918 and 919, Irminger Basin, off the east coast of Greenland (Fig. 1). At Site 918, pore water Li
*Author to whom correspondence ([email protected]).
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Fig. 1. Location map of Sites 918 and 919 (from Shipboard Scientific Party, 1994).
waters at shallow depths. These holes thus provide a suitable setting to study removal mechanisms of Li in the oceans other than the alteration of oceanic crust. Lithium concentrations greater than the seawater value have been observed in deep interstitial waters of many DSDP/ODP cores. The enrichment is thought to have originated from different processes such as transformation of biogenic opal to opal A, injection of hydrothermal fluids, and from reactions in the underlying oceanic or continental crust (Gieskes et al., 1982; Elderfield et al., 1990; Martin et al., 1991). The concentration increase in the deeper part of Hole 918 provides an opportunity to study the source of enrichment in the interstitial waters. Because of the distinct isotopic compositions of potential Li sources in the sediments (Chan et al., 1994b), the study of pore water Li isotopic composition at Site 918 may provide constraints on the processes controlling the distribution of Li in interstitial waters and thereby enhancing our understanding of its geochemical behavior in the marine sedimentary environment. 2. GEOLOGICAL SETTING
Samples were retrieved from Sites 918 and 919, ODP Leg 152, located on the upper continental rise of southeast Green-
land in the western part of the Irminger Basin (Fig. 1). The initial goal of drilling was to understand rifting of the continental lithosphere and formation of the oceanic lithosphere (Shipboard Scientific Party, 1994). Deep multichannel seismic profiling reveals that Site 918 is located on the transition zone between continental and oceanic crust with mantle derived flow basalts placed on top of thinned continental crust. Site 919 is situated away from the transition zone, probably on newly formed oceanic crust (Shipboard Scientific Party, 1994). The volcanic basement at Site 918 is covered by approximately 1200 m of sediments. The sedimentary section in Site 918 can be divided into five lithologic units: dark gray quartz silt with both volcaniclastic and continentally derived components (0 – 600.0 mbsf), nanofossil chalk and volcaniclastic and continental components (600 – 806.5 mbsf), nanofossil chalk and quartz-rich turbiditic sand (806.5–1108.2 mbsf), nanofossil chalk and volcanic silt (1108.2–1157.9 mbsf), and glauconitic sandy silt with interbedded calcareous sand (1157.9 –1189.4 mbsf; Shipboard Scientific Party, 1994). Volcanic material is present throughout the sediment column. Site 919 only extends to a shallow depth (;134 mbsf). The sediments are composed predominantly of volcaniclastic and continentally derived silty clays that are generally finer-grained than the equivalent sediments at Site 918 (Shipboard Scientific Party, 1994).
Lithium isotope geochemistry of pore waters 3. ANALYTICAL METHODS Pore water was extracted from 5 to 10 cm long, whole-round sediment samples by using the standard procedure of Manheim and Sayles (1974). The pore water was then filtered through an in-line filtration apparatus. Lithium and strontium concentrations were measured by atomic absorption spectrophotometry (Shipboard Scientific Party, 1994). Sediment samples were dissolved by digestion in a mixture of HF/HClO4, following the procedure described in Chan et al. (1992). The Li concentrations of the sediments were determined by flame emission spectrophotometry using standard additions. Strontium was separated from the pore water samples by a standard ion chromatography technique, using AG50Wx8 (BioRad) cation exchange resin. The eluted strontium fraction (;1 mg Sr) was mixed with 20 mL of 0.025 M H3PO4 and evaporated to dryness. The residue was redissolved in ;3 mL of 1 M HNO3, and loaded on a prebaked Re filament. Isotopic composition measurements were made on a 90° sector thermal ionization mass spectrometer (Finnigan MAT 262) using double Re filaments. Data were acquired using simultaneous collection with multi-Faraday cups. One hundred ratios were collected in ten blocks and the in-run precision (2sm) is better than 60.00001 or ;0.0014%. The 87Sr/86Sr ratios were normalized to 86Sr/88Sr 5 0.1194. Time series measurements of NBS 987 yielded an average value of 87Sr/86Sr 5 0.710262 6 7 (n 5 12, 2sm). Average 87Sr/86Sr ratio from replicate analyses of modern seawater was 0.709177 6 5 (n 5 7). Lithium isotope measurements were carried out following the procedure of You and Chan (1996). Lithium in pore water was eluted from the cation exchange column with 0.5 M HCl. The eluant was evaporated to dryness and redissolved in subboiling water. The solution was irradiated with ultraviolet light overnight to destroy any organic materials present and subsequently evaporated to a small drop. 0.04 mL 0.025 M H3PO4 was added and allowed to react on a hot plate for 5 h. Lithium was thus converted to phosphate that served as the ion source for the measurement. Lithium phosphate was loaded on a prebaked Re filament with a current of 1A. The current was then raised to 2A to expel the residual phosphoric acid. After introduction to the mass spectrometer, the ionization current was slowly raised to 2.1 A, and the evaporation current to 0.6 – 0.7 A. The 6Li/7Li ratio was measured directly on Li1 ions using a peak jumping mode while maintaining the ionization filament temperature at 1300 –1350°C. The sample size for the measurement was approximately 100 ng and a blank was determined to be 0.2 ng. Isotopic compositions are expressed as d6Li relative to a NBS standard L-SVEC (Flesch et al., 1973).
d6Li 5
F
(6Li/7Li)sample 21 (6Li/7Li)std.
G
3 1000
(1)
where (6Li/7Li )std. was determined to be 0.08261 6 0.00003 (1s) by the same method. Replicate analyses of seawater (n 5 6) produced a mean value for d6Li of 231.4 6 1.3‰ (You and Chan 1996). To understand the adsorption/desorption properties of Li on sediments, we have carried out laboratory experiments in a separate study. The method is described here briefly as relevant to this study and details will be given in another publication. Five identical mixtures of clay mineral (kaolinite or vermiculite) and Gulf of Mexico surface water with a water/sediment weight ratio of 20:1 were placed in poly tubes and allowed to equilibrate for 4900 h. The reaction in each vessel was terminated at a predetermined time, and the supernatant was filtered and analyzed. For this series of experimental fluids, the concentrations were determined by flame emission spectrophotometry with a precision of 62% and the isotopic analyses were as described above. The isotopic fractionation factors (a) for vermiculite and kaolinite were determined using a Raleigh fractionation model: R/R0 5 (C/C0)a21 6
7
(2) 6
where Ro is the initial Li/ Li ratio of Gulf water and R is the Li/7Li ratio of the solution after adsorption. Co and C are the corresponding Li concentrations. a was determined from the slope of the linear regression line when ln(R/Ro) was plotted against ln(C/Co) for different reaction times. In addition, Mississippi River suspended sediment was filtered from a large-volume water sample collected during high discharge and allowed to react with Gulf of Mexico water. Concentration
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and isotopic composition were determined only at the end of reaction after 4900 h. Lithium concentration was determined by mass spectrometry isotope dilution with a precision better than 0.5%. The isotopic fractionation factor for the Mississippi river sediment was calculated using both the Raleigh fractionation model (Eqn. 2) and an equilibrium model. The desorbability of Li from marine sediments was measured by reacting three Louisiana shelf sediments with 1M ammonium acetate. 4. RESULTS
Lithium and strontium isotope data of the pore waters from Sites 918 and 919 are presented in Table 1. At Site 919, the 87 Sr/86Sr ratio of pore water decreases from the seawater value (0.70918) to 0.70909 within the top 11 m. The ratio then decreases gradually to 0.70895 at 105 mbsf and subsequently rises slightly to 0.70901 at 134 mbsf. All 87Sr/86Sr ratios in this section are lower than the contemporaneous seawater values (Fig. 3). At Site 918, the87Sr/86Sr profile at the shallow depths is similar to that at Site 919 although the 87Sr/86Sr ratio at 112 mbsf rises to the contemporaneous seawater value (Fig. 2). At the depth interval of 110 –300 mbsf, 87Sr/86Sr ratios are again below the contemporaneous seawater values. The lowest 87Sr/ 86 Sr ratio (0.70819) occurs at 575 mbsf, corresponding to a Ca maximum and a Mg minimum (see Fig. 6; Gieskes et al., 1998). In the sand-rich layers (800 –1100 mbsf) and below, 87Sr/86Sr ratios indicate a slightly more radiogenic composition than the contemporaneous seawater values. In summary, there are three zones (0 –100 mbsf, 140 –300 mbsf, and 550 – 800 mbsf) at Site 918 and one zone (0 –134 mbsf) at Site 919 that have 87Sr/86Sr ratios lower than the contemporaneous seawater value. The downcore variations of pore water Li concentration and d6Li for Hole 918 are shown in Fig. 4a,b together with the present-day seawater values. The bottom of Hole 918 is of Eocene age (;50 Ma; Shipboard Scientific Party, 1994). Based on the Li/Ca ratio in foraminiferal shells, the Li concentration of seawater has not varied significantly over the last 40 Ma (Delaney and Boyle, 1986). In contrast, Hoefs and Sywall (1997) observed a Li isotope variation of about 30‰ in foraminifera over the past 60 Ma. In their study, relatively large analytical errors (5‰) were reported, and the possibility of diagenetic effects was acknowledged. There are other concerns regarding the separation procedures such as incomplete recovery of Li that could lead to isotopic fractionation. For these reasons, the variation of oceanic Li isotopic composition with time still needs to be confirmed. In the discussion of the present pore water data, we will use the present-day seawater values as a reference point to infer exchange processes. At Site 918, the pore water Li concentration decreases to 15 mM within the top 10 mbsf and reaches half of the present-day seawater value at 139 mbsf (Fig. 4a). At greater depths (.140 mbsf) dissolved Li concentrations increase gradually downcore. The Li concentration profile below 140 mbsf resembles a diffusion controlled curve and the concentration reaches a maximum (163 mM) at 765 mbsf. In the sand-rich layer (800 – 1100 mbsf), Li concentrations are lower than the maximum value observed at 765 mbsf. Correlated with the decrease in Li concentration at the shallow depths, the Li isotopic composition becomes heavier, reaching a minimum d6Li value of 241.4‰ at 55 mbsf, about 9‰ heavier than that of the present-day seawater (232.3‰, Chan and Edmond, 1988; Fig. 4b). Below 55 mbsf, the d6Li value increases to a maximum of 228‰,
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corresponding to the NH1 4 maximum from 200 to 300 mbsf (see Fig. 6). A second maximum occurs at 765 mbsf with an isotopic composition of ;7‰ lighter than the seawater value. The lightest isotopic composition observed, other than in the bottom most sample, coincides with the maximum Li concentration implying a 6Li-rich source at this depth ( ;765 mbsf). In the deeper interstitial waters, the d6Li profile indicates a heavier isotopic composition (236.4‰) in the sand-rich layer and a lighter isotopic composition (224.7‰) at the sediment and crust interface. The Li concentration profile at Site 919 is similar to that of Site 918 at shallow depths and the d6Li data are almost identical (Fig. 5a,b), showing a rapid decrease to a minimum isotopic composition of 241.9‰. Below 70 mbsf, the Li isotopic composition returns to a seawater-like value without any significant change in Li concentration. The increase in d6Li again corresponds to the rise of NH1 4 concentration (Fig. 7). Some sediment samples from Hole 918 were analyzed for Li content and isotopic composition for comparison with the pore waters. The data for the solids are given in Table 2. Concentrations are low (13.6 –19.1 ppm) in the shallow, quartz-rich silt above 232 mbsf (Fig. 8a). The concentration increases to a maximum of 52.7 ppm at 575 mbsf and is lower in the silt layer with nanofossils below (600 – 800 mbsf). Li content decreases to 33 ppm in the sand rich layer and to 32 ppm in the volcaniclastic silt at the bottom of the hole. The isotopic compositions, shown in Fig. 8b, fall mostly in the range of 28 to 212‰ with the exception of relatively light isotopic values of 24.9‰ at 660 mbsf and 23.2‰ at 1157 mbsf where volcanogenic components are prevalent.
The results of the adsorption and desorption experiments are given in Table 3. The Gulf of Mexico surface water used for the adsorption experiment had an initial concentration of 24.4 mM and d6Li of 231.5‰. The concentration decreased with time while isotopic composition became heavier. After approximately 4900 h of reaction, 34, 43, and 23% of dissolved Li was removed by vermiculite, kaolinite, and Mississippi River sediment. Their isotopic fractionation factors (a) were determined to be 1.029 6 0.005, 1.021 6 0.005, and 1.021 6 0.007 to 1.024 6 0.008, respectively. 5. DISCUSSION
5.1. Volcanic Alteration Zones at Sites 918 and 919 Studies of 87Sr/86Sr ratios of interstitial waters have shown a strong correlation between alteration of volcanic material in the sediment column and lowering of 87Sr/86Sr below the contemporaneous seawater values, especially in volcanic material rich sediments (Elderfield and Gieskes, 1982; Gieskes et al., 1986, 1987). Since volcanic material in the sediment column is the only source that has 87Sr/86Sr ratios below the contemporaneous seawater values, the decrease of 87Sr/86Sr ratios is probably caused by isotopic exchange with volcanic material. 87Sr/86Sr ratios of pore waters are lower than the contemporaneous seawater values in three different zones (,100 mbsf, 140 –300 mbsf, and 500 – 800 mbsf) at Site 918 and in the entire pore water column (,134 mbsf) of Site 919. The amount of lowering of 87Sr/86Sr ratios largely reflects the degree of alteration of volcanic material within sediments, although it can be compli-
Lithium isotope geochemistry of pore waters
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Fig. 2. Depth distribution of 87Sr/86Sr ratio in pore waters at Site 918. The solid line represents contemporaneous seawater values at the corresponding depths. Also shown are lithological units and age boundaries of the sedimentary column (Shipboard Scientific Party, 1994).
cated by the dissolution of carbonates. Volcanic materials are present throughout the sediment columns at Sites 918 and 919 (Shipboard Scientific Party, 1994) as evidenced by high Ti/Al ratios of bulk sediments (Gieskes et al., 1998). The lowest 87 Sr/86Sr ratio occurs at 575 mbsf and coincides with the Ca and Mg extrema (Fig. 6) which are caused by the breakdown of plagioclase in the volcanic matter and incorporation of Mg into alteration clays. 5.2. Lithium Depletion at Depths above 100 mbsf At shallow depths (,100 mbsf) at Sites 918 and 919, the Li concentration of the pore waters decreases rapidly to about half of the seawater value (Figs. 4a and 5a). 87Sr/86Sr ratio data (Figs. 2 and 3) indicate alteration of volcanic material. Lithium is known to be incorporated in the alteration phases during low temperature alteration of basalt (,150°C; Seyfried et al., 1984; Stoffyn-Egli and Mackenzie, 1984). Calcium concentrations, decreasing to a shallow minimum (Figs. 6 and 7; Gieskes et al., in press) suggest precipitation of calcite. Both precipitation of calcite and alteration of volcanic material could remove Li from the pore waters. However, based on the calculation using a typical Li/Ca molar ratio of 20 3 1026 in marine carbonates, Li concentrations of pore waters should only decrease at most by 0.2 mM due to precipitation of calcite. Hence, precipitation
of calcite cannot have a significant influence on the Li concentrations and isotopic compositions of the pore waters at Sites 918 and 919. The rapid decrease of pore water Li in the upper section of the two sites is accompanied by a shift to heavier d6Li values relative to present-day seawater value (Figs. 4 and 5). The uptake of Li into alteration clays favors the lighter isotope leaving the pore water isotopically heavier relative to the initial seawater composition. This is consistent with observations on weathered seafloor basalts (Chan et al., 1992). The isotopic fractionation factor during this process may be estimated by a Raleigh fractionation model (Eq. 2). A plot of ln(R/Ro) vs. ln(C/Co) for the upper 60 m of Site 919 is shown in Fig. 9, assuming present-day seawater concentration and isotopic ratio for the initial pore fluid. This is reasonable since the Pleistocene-Pliocene boundary lies at about 120 mbsf (Shipboard Scientific Party, 1994) and the oceanic Li composition probably has not changed during the Holocene-Pleistocene (Delaney and Boyle, 1986; Hoefs and Sywall, 1997). The fractionation factor thus determined is 1.020. The isotopic fractionation factor for submarine basalt alteration at sea bottom temperatures in the North Atlantic has been determined to be 1.019 (Chan et al., 1992). The temperature gradient at Site 918 is about 5.6°C/100 m (Shipboard Scientific Party, 1994). Thus the temperatures of
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Fig. 3. Depth distribution of 87Sr/86Sr ratio in pore waters at Site 919. The solid line represents contemporaneous seawater values at the corresponding depths.
Fig. 4. Lithium concentration (a) and isotopic composition (b) profiles at Site 918. The solid line in the respective diagrams represents the present-day seawater Li concentration and isotopic composition (Li concentration data are from Gieskes et al., 1998).
Lithium isotope geochemistry of pore waters
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Fig. 5. Lithium concentration (a) and isotopic composition (b) profiles at Site 919. The solid line in the respective diagrams represents the present-day seawater Li concentration and isotopic composition; Li concentration data are from Gieskes et al. (1998).
the pore waters in the upper 50 m section of Hole 919 are not significantly different from the sea bottom temperature, and the isotopic fractionation factor obtained at this site is consistent with low temperature alteration of volcanic material. Lithium depletion in interstitial waters observed at many ODP sites has been linked to the presence of altered volcanic material in the sediment column (Stoffyn-Egli and Mackenzie, 1984 and references therein). The Li isotopic data in this study support the suggestion that low Li concentrations in the interstitial waters are due to alteration of volcanic material. It has been speculated that diagenesis of volcanic material in sediments may be an important sink for Li in the oceans (StoffynEgli and Mackenzie, 1984) and that diffusion may carry dissolved Li in the oceans into sediments due to the concentration gradient caused by alteration of volcanic material. Using the concentration gradient (10 mM/10 m) at Site 919 and a diffusion coefficient of 1.48 3 102 cm2/year for Li at 0°C (Li and Gregory, 1974) and a porosity of 50%, the diffusive flux of Li into the sediments is estimated to be 3.7 3 1024 mmol/cm2. year (flux 5 (porosity)2 x diffusion coefficient x concentration gradient, Stoffyn-Egli and Mackenzie, 1984). The diffusive flux into the sediments in the entire ocean due to this diagenetic process is difficult to estimate because the content of volcanic matter and consequently the concentration gradient across the sediment-seawater interface would vary greatly geographically. At some locations the concentration gradient in the surface sediment is reversed, such as those observed in the Guaymas Basin (DSDP 477, Gieskes et al., 1982) and at the Peru margin (Martin et al., 1991). As an absolute upper limit, we assume that the diffusive flux estimated for Irminger Basin occurs over the entire ocean floor, yielding 1.3 3 109 mol/yr. This maxi-
mum flux may be compared to other Li fluxes in the ocean. The current estimates of the input and output fluxes are summarized in Table 4. It should be noted that the latest estimate of the river flux by Huh et al. (in press) is based on the flow-weighted mean concentration of the world’s major rivers and is considered the most accurate. It can be seen that Li removed by volcanic matter alteration in the sediments can be no more than 5% of the total inputs, mainly from rivers and submarine hydrothermal fluids (Stoffyn-Egli and Mackenzie, 1984 and Huh et al., in press). For this reason, post depositional alteration of volcanogenic components in the sediments is probably a minor sink of oceanic Li. 5.3. Source of Isotopically Light Lithium at 230 –350 mbsf Although 87Sr/86Sr and d18O data indicate that alteration of volcanic material occurs between 232 and 350 mbsf at Site 918 (Gieskes et al., 1998), d6Li shows an increase to 228‰, suggesting that other processes besides alteration of volcanic material take place in this depth interval. The reversal of d6Li to lighter isotopic values corresponds to the NH1 4 maximum in the zero SO22 zone and the zone of methanogenesis (Fig. 6; 4 Shipboard Scientific Party, 1994). A similar correlation is also observed at Site 919 (Fig. 7). Because the NH1 4 maximum is related to microbial decomposition of organic matter, the light Li isotopic composition might be partially derived from the organic matter. However, due to the relatively low concentration of Li in organic matter (0.5–1 ppm in dry organic matter; Heier and Billings, 1970) and the low organic content (;0.5%) of the sediments (Shipboard Scientific Party, 1994), the decom-
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Fig. 6. Vertical distribution of Ca, Mg, sulfate, and NH1 4 in pore waters at Site 918 (from Gieskes et al., 1998).
position of the organic matter cannot be a direct cause for the increase of d6Li from a mass balance point of view. Cation exchange of NH1 4 with clay minerals may better account for the light Li isotopic composition at the NH1 4 maximum. In some ODP sites, more than half of the NH1 4 produced in pore waters undergoes ion exchange with cations in clay minerals (Gieskes, 1983, and references therein). Hence, it is highly possible that NH1 4 may expel Li from exchangeable positions in clay minerals. A small Sr concentra-
tion maximum and more radiogenic strontium isotopic composition are also observed at ;300 mbsf. Since pore water Li and Sr concentrations are in the mM range and NH1 4 production is in the mM range, the exchange of NH1 4 with cations in clay minerals can potentially influence the Li and strontium isotopic compositions of pore waters. If the exchange of Li in clay minerals with NH1 4 is to account for the d6Li increase, the amount of Li in exchangeable positions of clay minerals has to be large enough to cause a shift of
Lithium isotope geochemistry of pore waters
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Fig. 7. Vertical distribution of Ca, Mg, sulfate, and NH1 4 in pore waters at Site 919 (from Gieskes et al., 1998.
about 10‰ in d6Li. The isotopic fractionation factors of the two clay minerals and natural river sediments are indistinguishable within analytical error (Table 3). If we assume an average isotopic fractionation factor of 1.025, absorbed Li from and in equilibrium with seawater would have an isotopic composition of 28‰. At equilibrium, approximately 1 mg Li was adsorbed per gram of clay mineral in the adsorption experiment (Table 3). Given a unit volume (cm3) of sediment with a porosity of 50% and a pore water Li concentration of 13 mM (at 140 mbsf) as the concentration before the onset of desorption at Site 918,
the total amount of Li in the pore water in this volume would be about 45 ng. By using 2.7 g/cm3 for the density of the bulk sediment and 40 –50% for the percentage of clay minerals (Shipboard Scientific Party, 1994), the amount of the adsorbed Li in clay minerals is on the order of 500 ng. If approximately 4% of the adsorbed Li that has a light isotopic composition 6 (28‰) were desorbed into pore water by NH1 4 , d Li in the pore water would effectively shift from 239‰ to a lighter isotopic composition by 10‰. On the basis of this calculation, there should be an increase of Li concentration to 19 mM at
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maximum exchange with NH1 4 . However, the actual concentration reaches 27 mM at the d6Li maximum (228.7‰) at 276 mbsf. The Li enrichment due to ion exchange with NH1 4 may, in part, be masked by the upward diffusion of Li from below (Fig. 4a). In contrast to the concentration increase with depth at Site 918, Li concentration remains essentially constant below 50 mbsf at Site 919 as NH1 4 rises (Fig. 5a). The lack of an explicit concentration increase at Site 919 may be the result of two competing processes: alteration of volcanic material and desorption. 87Sr/86Sr data indicate that volcanic alteration does occur throughout Hole 919. If we take the conditions at 50 mbsf as those before desorption, then 56 ng dissolved Li is present per cm3 of wet sediment (50% porosity). Assuming that d6Li of Li released from clay minerals by NH1 4 is 28‰ and the isotopic fractionation factor for low temperature alteration of volcanic material is 1.019, 112 ng Li that is desorbed by NH1 4 and later taken up by volcanic material is needed to shift the pore water 6Li from 242‰ to 232‰ as observed (Fig. 5b). This calculation demonstrates that the net result of these two
competing processes could cause the observed shift in d6Li in the pore water without any net change in dissolved Li concentration. The possibility of Li desorption by NH1 4 has been verified by the reaction of Louisiana shelf sediments with 1M ammonium acetate. After treatment with the solution, 0.5– 0.6 mg of Li was desorbed per gram of sediment (Table 3). We further note that the exchangeable Li is small compared to the Li present in the bulk sediment and is comparable to the amount adsorbed as
Fig. 8. Lithium concentration and isotopic composition of sediments from Site 918.
Lithium isotope geochemistry of pore waters
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Fig. 9. A plot of 6Li/7Li ratio vs. Li concentration ratio on the logarithmic scale for waters from the upper 50 m in Site 919 according to the Raleigh fractionation model. The slope of the linear regression line provides an estimation of isotopic fractionation factor for the alteration process at this site.
shown in the adsorption experiments. Even so, the exchange of NH1 4 with Li in clay minerals may exert significant control on the Li isotopic composition in sediment interstitial waters. As a corollary, we estimate the amount of oceanic Li that may be removed by adsorption onto sediments. If we assume adsorption of 1 mg Li per gram of sediment as shown in our experiments, and given a suspended load of 1.5 3 1016 g/yr for the world’s rivers (Martin and Meybeck, 1979) total Li removed from the oceans by adsorption on the suspended sediments would be about 2.2 3 109 moles Li/yr which is 5–10% of the total input estimated by Stoffyn-Egli and Mackenzie (1984) and Huh et al. (in press; Table 4). 5.4. Lithium Enrichment in Deeper Interstitial Waters From 500 to 800 mbsf, the Li concentration and the d6Li steadily increase with depth at Site 918, reaching a maximum concentration (163 mM) and the lightest isotopic composition (226‰) at 765 mbsf. Enrichment of Li in deep pore waters has been reported at many ODP Sites and several explanations have been proposed based on the correlation of Li concentration data with other geochemical data. At Site 474, the enrichment was interpreted as a result of the transformation of biogenic opal to opal-A (Gieskes et al., 1982). The injection of Li-rich hydro-
thermal solutions from the underlying basalt has been proposed to explain pore water Li enrichment in deep sea sediment at the Guaymas Basin spreading center (Gieskes et al., 1982). In contrast, at the Peru margin, the enrichment has been attributed to reactions with continental and oceanic crusts as indicated by strontium isotopic data (Elderfield et al., 1990; Martin et al., 1991). At Site 918, based on Ca, Mg, and 87Sr/86Sr data, the reactions in underlying basaltic basement are unlikely to have an important influence on the pore water chemistry. The maximum concentration and d6Li at 765 mbsf suggest a Li source within the sediment column. Characterization of the concentration and isotopic composition of various types of marine sediments is essential to identify the source of enrichment in the interstitial waters. In a recent survey of Li isotopes in marine sediments (Chan et al., 1994b), carbonates and siliceous oozes were found to be low (;1 ppm) in Li while hemipelagic and pelagic clays have 45– 62 ppm with d6Li values ranging from 210 to 214‰, distinct from oceanic crust (; 25‰; Chan et al., 1992). The Li concentration profile of the solids at Site 918 displays a similar shape as the dissolved Li profile (Figs. 4a and 8a). The silty sediments above 300 mbsf are quartz rich (Shipboard Scientific Party, 1994) and have relatively low Li contents.
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Peak concentrations occur between 600 and 800 mbsf. At ;765 mbsf, where the maximum dissolved Li is found, the Li content in the solid phase is 47 ppm. Given a porosity of 50%, and a sediment density of 2.7 g/cm3, the Li content of the pore water represents only 1% of the Li inventory in the solid. Therefore, the similarity between solid and pore water Li profiles suggests that the Li concentration in the deeper interstitial water is largely controlled by the associated sediment. Experimental studies have shown that Li can be released from marine sediments at temperatures as low as 50 – 60°C (Chan et al., 1994a; You, 1994). In the experimental system with a water/sediment ratio of 3:1, Li released to the fluid reached 40 mM at 60°C after reaction for several days and is approximately 5– 6‰ heavier than that in the initial sediment. At Site 918, the geothermal gradient is about 5.6°C/100 m (Shipboard Scientific Party, 1994) and the temperature at the concentration maximum (;765 mbsf) is about ;45°C. With a much lower water/ sediment ratio (;1:3) and considerably longer reaction time, it is possible that Li can attain the observed concentration of 163 mM at this depth. That the enrichment originates from the pelagic sediments is also consistent with the Li isotopic composition in the pore waters. d6Li in pore water correlates well with 1/Li (r2 5 0.869) from 500 mbsf to 800 mbsf (Fig. 10). This linear correlation suggests addition of Li released from the sediments to the pore water of evolved seawater composition. If we assume the water at 515 mbsf is the evolved seawater endmember (with 57 mM Li and 237.7‰) that interacts with the deeper sediments, it can be calculated from the mixing relationship that Li released from the sediment has d6Li of
about 220‰. Available data show that Li isotopic compositions of the sediments at 918 mostly vary between 28 and 212‰ with two exceptions at 660 and 1157 mbsf where the isotopic compositions are very light (24.9 and 23.2‰; Fig. 8b). These depths are marked by volcanogenic components (Shipboard Scientific Party, 1994) which are isotopically light. Extraction of the sediment-bound Li with attendant isotopic fractionation may account for the enrichment of relatively light Li in the fluids between 550 and 800 mbsf. Although 87Sr/86Sr data indicate strong alteration of volcanic material in this section (a process that takes up Li from solution), the Li distribution appears to be dominantly controlled by the input from the sediments. The interval between 806 and 1100 mbsf consists of turbidite sands interbedded with nanofossil chalk (Shipboard Scientific Party, 1994). In this layer Li concentration decreases and d6Li shifts to 236.6‰. Chloride and oxygen isotopic data indicate advection of trapped meteoric water through the sand bed (Gieskes et al., 1998). However, the advecting fresh water probably has low Li content and would not significantly affect the Li isotopic composition of pore water in this layer. Hence, the d6Li value may largely reflect that of diluted pore water without significant input from the sediments at this horizon. The deepest water sample (1157 mbsf) has a relatively light isotopic composition (224.8‰). The concentration of pore water at this level (91 mM) is considerably higher than the seawater value. Thus the isotopic composition is probably not controlled by Li uptake into volcanic material, neither can hydrothermal fluids from the basement be the source of Li enrichment because their d6Li values are between 26 and 211‰ (Chan et al., 1993). The 87Sr/86Sr ratios below 925 mbsf are also higher than the contemporaneous seawater values which suggests alteration of more radiogenic material in the sediment column and possibly diffusion from the maximum at 925 mbsf (Gieskes et al., 1998). Similarly, the relatively light isotopic composition of Li in the deepest pore water may be attributed to interaction with the sediment at this depth. 6. SUMMARY AND IMPLICATIONS
At shallow depths (,100 mbsf), pore water Li concentrations are mainly controlled by alteration of volcanic material within the marine sediments. d6Li data support that Li has been taken up by altered volcanic material. This study confirms that alteration of volcanic material is not restricted to the exposed oceanic crust but also occurs in the volcanogenic components of marine sediments. It is estimated that the diffusive flux into the sediments as a result of this process may at most account for 5% of the sink for the dissolved Li in the global ocean. Relatively light isotopic compositions at Sites 918 and 919 are observed to correspond with the NH1 4 maximum . It is proposed that NH1 4 generated by bacterial decomposition of organic matter may expel adsorbed Li from clay minerals by ion exchange and significantly influence the isotopic composition of pore water. This suggestion is supported by results of adsorption and desorption experiments. Total Li removed from the ocean water by adsorption on to the suspended sediments is approximately 5–10% of the total input. In the deep sediment column (500 – 800 mbsf) at Site 918, Li isotopic data suggest that the enrichment of pore water Li is
Lithium isotope geochemistry of pore waters
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Fig. 10. d6Li vs. 1/Li plot for pore waters between 500 and 800 mbsf at Site 918. The linear relationship indicates addition of Li derived from sediments to the pore waters.
most likely derived from pelagic sediments. This is consistent with experimental studies which have shown that Li is easily mobilized from marine sediments at elevated temperatures. Our study suggests that Li enrichment in deep interstitial waters, observed in many ODP cores, may be explained by release of Li from sediments at elevated temperatures and low water/ sediment ratio. In conclusion, this study demonstrates that Li isotopic compositions of pore water can help elucidate a variety of processes in the sediment column and are a useful indicator of sediment-water interaction. It is also found that incorporation of Li into authigenic clay minerals from volcanic matter alteration in the sediment column and Li adsorption on sediments are relatively minor sinks of oceanic Li.
Acknowledgments—We are grateful to the Ocean Drilling Program for providing pore water and sediment samples from Leg 152 for this study. We thank Eric De Carlo and an anonymous reviewer for their valuable comments and suggestions. This work was performed with the support of National Science Foundation grant OCE 9314708 to L. H. C. and the support from the US Science Advisory Committee to J. M. G.
REFERENCES Chan L. H. (1987) Lithium isotope analysis by thermal ionization mass spectrometry of lithium tetraborate. Anal. Chem. 59, 2662–2665. Chan L. H. and Edmond J. M. (1988) Variation of lithium isotope composition in the marine environment: A preliminary report. Geochim. Cosmochim. Acta 52, 1711–1717. Chan L. H., Edmond J. M., Thompson G., and Gillis K. (1992) Lithium isotopic composition of submarine basalts: Implications for the lithium cycle in the oceans. Earth Planet. Sci. Lett. 108, 151–160. Chan L. H., Edmond J. M., and Thompson G. (1993) A lithium isotope study of hot springs and metabasalts from mid-ocean ridge hydrothermal systems. J. Geophys. Res. 98, 9653–9659.
Chan L. H., Gieskes J. M., You C. F., and Edmond J. M. (1994a) Lithium isotope geochemistry of sediments and hydrothermal fluids of the Guaymas Basin, Gulf of California. Geochim. Cosmochim. Acta 58, 4443– 4454. Chan L. H.., Zhang L., and Hein J. R. (1994b) Lithium isotope characteristics of marine sediments. EOS 75, 314 (abstr.). Delaney M. L. and Boyle E. A. (1986) Lithium in foraminiferal shells implications for high-temperature hydrothermal circulation fluxes and oceanic crustal generation rates. Earth Planet. Sci. Lett.. 80, 91–105. Demaster D. J. (1981) The supply and accumulation of silica in the marine environment. Geochim. Cosmochim. Acta 45, 1715–1732. Elderfield H. and Gieskes J. M. (1982) Sr isotopes in interstitial waters of marine sediments from Deep Sea Drilling Project cores. Nature 300, 493– 497. Elderfield H., Kastner M., and Martin J. B. (1990) Composition and sources of fluids in sediments of the Peru subduction zone. J. Geophys. Res. 95, 8819 – 8828. Flesch G. D., Anderson A. R. Jr., and Svec H. J. (1973) A secondary isotopic standard for lithium determinations. Intl. J. Mass Spectrom. Ion Phys. 12, 265–272. Gieskes J. M. (1983) The chemistry of interstitial waters of deep sea sediments: Interpretation of deep sea drilling data. In Chemical Oceanography (ed. J. P. Riley and R. Chester), Vol. 8, pp. 221–269. Academic Press. Gieskes J. M., Elderfied H., Lawrence J. R., Johnson J., Meyers B., and Campbell A. (1982) Geochemistry of interstitial waters and sediments, Leg 64, Gulf of California. Init. Repts. DSDP 64, 675– 694. Gieskes J. M., Elderfield H., and Palmer M. R. (1986) Strontium and its isotopic composition in interstitial waters of marine carbonate sediments. Earth. Planet. Sci. Lett. 77, 229 –235. Gieskes J. M., Lawrence J. R., Perry E. A., Grady S. J., and Elderfield H. (1987) Chemistry of interstitial waters and sediments in the Norwegian-Greenland Sea, DSDP Leg 38. Chem. Geol. 63, 143– 155. Gieskes J. M., Schrag D., Chan L. H., Zhang L., and Murray J. (1998) Geochemistry of interstitial Waters. Proc. ODP. Sci. Results 152, 293–305. Heier K. S. and Billings G. K. (1970) Lithium. In Handbook of
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Geochemistry (ed. K. H. Wedepohl), Vol. II-1, pp. 3-H-1. SpringerVerlag. Hoefs J. and Sywall M. (1997) Lithium isotope composition of Quaternary and Tertiary biogene carbonates and a global lithium isotope balance. Geochim. Cosmochim. Acta 61, 2679 –2690. Holland H. D. (1984) The Chemical Evolution of the Atmosphere and the Oceans. Princeton Univ. Press. Huh Y., Chan L. H., Zhang L., and Edmond J. M. (in press) Lithium and its isotopes in major world rivers: Implications for weathering and the oceanic budget. Geochim. Cosmochim. Acta . Kastner M., Elderfield H. and Martin J. B. (1991) Fluids in convergent margins: What do we know about their composition, origin, role in diagenesis and importance for oceanic chemical fluxes? Phil. Trans. Roy. Soc. London Ser. A 335, 243–259. Lawrence J. R. (1989) The stable isotope geochemistry of deep-sea pore water. In Handbook of Environmental Isotopes (ed. P. Fritz and J. C. Fontes), Vol. 3, pp. 317–356. Elsevier. Li Y. H. and Gregory S. (1974) Diffusion of ions in seawater and in deep sea sediments. Geochim. Cosmochim. Acta 38, 703–714. Manheim F. T. and Sayles F. L. (1974) Composition and origin of interstitial waters of marine sediments based on deep sea drill cores. In The Sea (ed. E. D. Goldberg), Vol. 5, pp. 527–568. Wiley Interscience. Martin J. B., Kastner M., and Elderfied H. (1991) Lithium: Sources in pore fluids of Peru slope sediments and implications for oceanic fluxes. Mar. Geol. 102, 281–292.
Martin J. M. and Meybeck M. (1979) Elemental mass-balance of material carried by major world rivers. Mar. Chem. 7, 173–206. Milliman J. D. (1993) Production and accumulation of calcium carbonate in the ocean: Budget of a nonsteady state. Global Biogeochem. Cycles, 7, 927–957. Seyfried W. E. Jr., Janecky D. R., and Mottl M. J. (1984) Alteration of the oceanic crust: Implications for geochemical cycles of lithium and boron. Geochim. Cosmochim. Acta 48, 557–569. Shipboard Scientific Party (1994) Site 918 and Site 919. Proc. ODP, Init. Rept. 152, 177–272. Stoffyn-Egli P. and Mackenzie F. T. (1984) Mass balance of dissolved lithium in the oceans. Geochim. Cosmochim. Acta 48, 859 – 872. Thompson G. (1983) Hydrothermal fluxes in the ocean. In Chemical Oceanography (ed. J. P. Riley and R. Chester) Vol. 8, pp. 272–337. Academic Press. You C. F. (1994) Lithium, beryllium, and boron isotope geochemistry: Implication for fluid processes in convergent margins. Ph.D. thesis, Univ. California. You C. F., Chan L. H., Spivack A. J., and Gieskes J. M. (1995) Lithium, boron, and their isotopes in sediments and pore waters of Ocean Drilling Program Site 808, Nankai Trough: Implications for fluid expulsion in accretionary prisms. Geology 23, 37– 40. You C. F. and Chan L. H. (1996) Precise determination of lithium isotopic composition in low concentration natural samples. Geochim. Cosmochim. Acta 60, 909 –915.
Pergamon
PII S0016-7037(98)00178-1
Lithium isotope geochemistry of pore waters from Ocean Drilling Program Sites 918 and 919, Irminger Basin LIBO ZHANG, LUI-HEUNG CHAN,*,1 and JORIS M. GIESKES2 1
Department of Geology and Geophysics, Louisiana State University, Baton Rouge, Louisiana 70803, USA 2 Scripps Institution of Oceanography, University of California, La Jolla, California 92093-0215, USA (Received November 12, 1997; accepted in revised form May 8, 1998)
Abstract—The distribution of Li isotopes in pore waters to a depth of 1157 m below seafloor is presented for ODP Sites 918 and 919 in the Irminger Basin, offshore Greenland. Lithium isotope data are accompanied by strontium isotope ratios to decipher diagenetic reactions in the sediments which are characterized by the pervasive presence of volcanic material, as well as by very high accumulation rates in the upper section. The lowering of the 87Sr/86Sr ratio below contemporaneous seawater values indicates several zones of volcanic material alteration. The Li isotope profiles are complex suggesting a variety of exchange reactions with the solid phases. These include cation exchange with NH1 4 and mobilization from sediments at depth, in addition to the alteration of volcanic matter. Lithium isotopes are, therefore, a sensitive indicator of sediment-water interaction. d6Li values of pore waters at these two sites vary between 242 and 225 ‰. At shallow depths (,100 mbsf), rapid decreases in the Li concentration, accompanied by a shift to heavier isotopic compositions, indicate uptake of Li into alteration products. A positive anomaly of d6Li observed at both sites is coincident with the NH1 4 maximum produced by organic matter decomposition and may be related to ion exchange of Li from the sediments by NH1 4 . In the lower sediment column at Site 918, dissolved Li increases with depth and is characterized by enrichment of 6Li. The Li isotopic compositions of both the waters and the solid phase suggest that the enrichment of Li in deep interstitial waters is a result of release from pelagic sediments. The significance of sediment diagenesis and adsorption as sinks of oceanic Li is evaluated. The maximum diffusive flux into the sediment due to volcanic matter alteration can be no more than 5% of the combined inputs from rivers and submarine hydrothermal solutions. Adsorption on to sediments can only account for 5–10% of the total inputs from rivers and submarine hot springs. Copyright © 1998 Elsevier Science Ltd concentrations show a rapid decrease by 10 mM within the top 10 m followed by an increase to approximately six times greater than the seawater value in the deeper section (Gieskes et al., 1998). In this paper, we report detailed Li isotope profiles in the interstitial waters at ODP Sites 918/919 and discuss the exchange processes affecting pore water Li distribution. The Li isotope study is accompanied by strontium isotope measurements which provide additional constraints on the sources and sinks of Li. This study is the first systematic attempt to understand the Li isotope geochemisry in sediment pore waters and aims at gaining an understanding of the Li exchange between the fluid and solid phases as well as the role of sediments in the marine budget of Li. The Li budget in the oceans has not been completely resolved. River flux, hydrothermal input, and fluid expulsion at convergent margins are three identified sources of Li in the oceans (Stoffyn-Egli and Mackenzie, 1984; Martin et al., 1991; You et al., 1995). The only well-defined sink in the oceans is low temperature alteration of oceanic crust, the magnitude of which is much smaller than the total input. Since marine sediments are enriched in Li in comparison to freshwater sediments (Holland, 1984), they are thought to be an important sink for oceanic Li. Lithium adsorption onto marine sediments (Seyfried et al., 1984) and incorporation into authigenic clay minerals (Stoffyn-Egli and Mackenzie, 1984) have been proposed as important removal processes of Li in the oceans. The rapid decrease in the uppermost layer at ODP 918/919 (Gieskes et al., 1998) suggests that Li has been removed from the pore
INTRODUCTION
Recently, Li and Li isotopes have been used to understand many geological processes in the oceans such as mid-ocean ridge hydrothermal activity, alteration of the oceanic crust, and fluid expulsion at convergent margins (Chan et al., 1992, 1993, 1994a; You et al., 1995). Relatively little, however, is known about the Li isotope geochemistry in the interstitial waters of marine sediments, even though they occupy roughly half of the marine sediment volume and constitute approximately 5% of the Earth’s hydrosphere (Lawrence, 1988). A previous Li isotope study of pore waters was carried out at Ocean Drilling Program (ODP) Site 808, Nankai Trough (You et al., 1995). At this site, a Li isotope anomaly at the de´collement zone provides strong evidence for the advection of fluids from deeper parts of the accretionary prism. The isotopic measurements were made by pooling samples from adjacent depths thus sacrificing depth resolution. This was due to sample size requirements for Li isotope analysis (;3.5 mg) by the most reliable mass spectrometric method (Chan, 1987). However, a recent improvement in the analytical sensitivity (You and Chan, 1996) has enabled Li isotopic determination of much smaller natural samples, thus permitting high resolution studies of pore waters. Our present study focuses on ODP Sites 918 and 919, Irminger Basin, off the east coast of Greenland (Fig. 1). At Site 918, pore water Li
*Author to whom correspondence ([email protected]).
should
be
addressed 2437
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L. Zhang, L.-H. Chan, and J. M. Gieskes
Fig. 1. Location map of Sites 918 and 919 (from Shipboard Scientific Party, 1994).
waters at shallow depths. These holes thus provide a suitable setting to study removal mechanisms of Li in the oceans other than the alteration of oceanic crust. Lithium concentrations greater than the seawater value have been observed in deep interstitial waters of many DSDP/ODP cores. The enrichment is thought to have originated from different processes such as transformation of biogenic opal to opal A, injection of hydrothermal fluids, and from reactions in the underlying oceanic or continental crust (Gieskes et al., 1982; Elderfield et al., 1990; Martin et al., 1991). The concentration increase in the deeper part of Hole 918 provides an opportunity to study the source of enrichment in the interstitial waters. Because of the distinct isotopic compositions of potential Li sources in the sediments (Chan et al., 1994b), the study of pore water Li isotopic composition at Site 918 may provide constraints on the processes controlling the distribution of Li in interstitial waters and thereby enhancing our understanding of its geochemical behavior in the marine sedimentary environment. 2. GEOLOGICAL SETTING
Samples were retrieved from Sites 918 and 919, ODP Leg 152, located on the upper continental rise of southeast Green-
land in the western part of the Irminger Basin (Fig. 1). The initial goal of drilling was to understand rifting of the continental lithosphere and formation of the oceanic lithosphere (Shipboard Scientific Party, 1994). Deep multichannel seismic profiling reveals that Site 918 is located on the transition zone between continental and oceanic crust with mantle derived flow basalts placed on top of thinned continental crust. Site 919 is situated away from the transition zone, probably on newly formed oceanic crust (Shipboard Scientific Party, 1994). The volcanic basement at Site 918 is covered by approximately 1200 m of sediments. The sedimentary section in Site 918 can be divided into five lithologic units: dark gray quartz silt with both volcaniclastic and continentally derived components (0 – 600.0 mbsf), nanofossil chalk and volcaniclastic and continental components (600 – 806.5 mbsf), nanofossil chalk and quartz-rich turbiditic sand (806.5–1108.2 mbsf), nanofossil chalk and volcanic silt (1108.2–1157.9 mbsf), and glauconitic sandy silt with interbedded calcareous sand (1157.9 –1189.4 mbsf; Shipboard Scientific Party, 1994). Volcanic material is present throughout the sediment column. Site 919 only extends to a shallow depth (;134 mbsf). The sediments are composed predominantly of volcaniclastic and continentally derived silty clays that are generally finer-grained than the equivalent sediments at Site 918 (Shipboard Scientific Party, 1994).
Lithium isotope geochemistry of pore waters 3. ANALYTICAL METHODS Pore water was extracted from 5 to 10 cm long, whole-round sediment samples by using the standard procedure of Manheim and Sayles (1974). The pore water was then filtered through an in-line filtration apparatus. Lithium and strontium concentrations were measured by atomic absorption spectrophotometry (Shipboard Scientific Party, 1994). Sediment samples were dissolved by digestion in a mixture of HF/HClO4, following the procedure described in Chan et al. (1992). The Li concentrations of the sediments were determined by flame emission spectrophotometry using standard additions. Strontium was separated from the pore water samples by a standard ion chromatography technique, using AG50Wx8 (BioRad) cation exchange resin. The eluted strontium fraction (;1 mg Sr) was mixed with 20 mL of 0.025 M H3PO4 and evaporated to dryness. The residue was redissolved in ;3 mL of 1 M HNO3, and loaded on a prebaked Re filament. Isotopic composition measurements were made on a 90° sector thermal ionization mass spectrometer (Finnigan MAT 262) using double Re filaments. Data were acquired using simultaneous collection with multi-Faraday cups. One hundred ratios were collected in ten blocks and the in-run precision (2sm) is better than 60.00001 or ;0.0014%. The 87Sr/86Sr ratios were normalized to 86Sr/88Sr 5 0.1194. Time series measurements of NBS 987 yielded an average value of 87Sr/86Sr 5 0.710262 6 7 (n 5 12, 2sm). Average 87Sr/86Sr ratio from replicate analyses of modern seawater was 0.709177 6 5 (n 5 7). Lithium isotope measurements were carried out following the procedure of You and Chan (1996). Lithium in pore water was eluted from the cation exchange column with 0.5 M HCl. The eluant was evaporated to dryness and redissolved in subboiling water. The solution was irradiated with ultraviolet light overnight to destroy any organic materials present and subsequently evaporated to a small drop. 0.04 mL 0.025 M H3PO4 was added and allowed to react on a hot plate for 5 h. Lithium was thus converted to phosphate that served as the ion source for the measurement. Lithium phosphate was loaded on a prebaked Re filament with a current of 1A. The current was then raised to 2A to expel the residual phosphoric acid. After introduction to the mass spectrometer, the ionization current was slowly raised to 2.1 A, and the evaporation current to 0.6 – 0.7 A. The 6Li/7Li ratio was measured directly on Li1 ions using a peak jumping mode while maintaining the ionization filament temperature at 1300 –1350°C. The sample size for the measurement was approximately 100 ng and a blank was determined to be 0.2 ng. Isotopic compositions are expressed as d6Li relative to a NBS standard L-SVEC (Flesch et al., 1973).
d6Li 5
F
(6Li/7Li)sample 21 (6Li/7Li)std.
G
3 1000
(1)
where (6Li/7Li )std. was determined to be 0.08261 6 0.00003 (1s) by the same method. Replicate analyses of seawater (n 5 6) produced a mean value for d6Li of 231.4 6 1.3‰ (You and Chan 1996). To understand the adsorption/desorption properties of Li on sediments, we have carried out laboratory experiments in a separate study. The method is described here briefly as relevant to this study and details will be given in another publication. Five identical mixtures of clay mineral (kaolinite or vermiculite) and Gulf of Mexico surface water with a water/sediment weight ratio of 20:1 were placed in poly tubes and allowed to equilibrate for 4900 h. The reaction in each vessel was terminated at a predetermined time, and the supernatant was filtered and analyzed. For this series of experimental fluids, the concentrations were determined by flame emission spectrophotometry with a precision of 62% and the isotopic analyses were as described above. The isotopic fractionation factors (a) for vermiculite and kaolinite were determined using a Raleigh fractionation model: R/R0 5 (C/C0)a21 6
7
(2) 6
where Ro is the initial Li/ Li ratio of Gulf water and R is the Li/7Li ratio of the solution after adsorption. Co and C are the corresponding Li concentrations. a was determined from the slope of the linear regression line when ln(R/Ro) was plotted against ln(C/Co) for different reaction times. In addition, Mississippi River suspended sediment was filtered from a large-volume water sample collected during high discharge and allowed to react with Gulf of Mexico water. Concentration
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and isotopic composition were determined only at the end of reaction after 4900 h. Lithium concentration was determined by mass spectrometry isotope dilution with a precision better than 0.5%. The isotopic fractionation factor for the Mississippi river sediment was calculated using both the Raleigh fractionation model (Eqn. 2) and an equilibrium model. The desorbability of Li from marine sediments was measured by reacting three Louisiana shelf sediments with 1M ammonium acetate. 4. RESULTS
Lithium and strontium isotope data of the pore waters from Sites 918 and 919 are presented in Table 1. At Site 919, the 87 Sr/86Sr ratio of pore water decreases from the seawater value (0.70918) to 0.70909 within the top 11 m. The ratio then decreases gradually to 0.70895 at 105 mbsf and subsequently rises slightly to 0.70901 at 134 mbsf. All 87Sr/86Sr ratios in this section are lower than the contemporaneous seawater values (Fig. 3). At Site 918, the87Sr/86Sr profile at the shallow depths is similar to that at Site 919 although the 87Sr/86Sr ratio at 112 mbsf rises to the contemporaneous seawater value (Fig. 2). At the depth interval of 110 –300 mbsf, 87Sr/86Sr ratios are again below the contemporaneous seawater values. The lowest 87Sr/ 86 Sr ratio (0.70819) occurs at 575 mbsf, corresponding to a Ca maximum and a Mg minimum (see Fig. 6; Gieskes et al., 1998). In the sand-rich layers (800 –1100 mbsf) and below, 87Sr/86Sr ratios indicate a slightly more radiogenic composition than the contemporaneous seawater values. In summary, there are three zones (0 –100 mbsf, 140 –300 mbsf, and 550 – 800 mbsf) at Site 918 and one zone (0 –134 mbsf) at Site 919 that have 87Sr/86Sr ratios lower than the contemporaneous seawater value. The downcore variations of pore water Li concentration and d6Li for Hole 918 are shown in Fig. 4a,b together with the present-day seawater values. The bottom of Hole 918 is of Eocene age (;50 Ma; Shipboard Scientific Party, 1994). Based on the Li/Ca ratio in foraminiferal shells, the Li concentration of seawater has not varied significantly over the last 40 Ma (Delaney and Boyle, 1986). In contrast, Hoefs and Sywall (1997) observed a Li isotope variation of about 30‰ in foraminifera over the past 60 Ma. In their study, relatively large analytical errors (5‰) were reported, and the possibility of diagenetic effects was acknowledged. There are other concerns regarding the separation procedures such as incomplete recovery of Li that could lead to isotopic fractionation. For these reasons, the variation of oceanic Li isotopic composition with time still needs to be confirmed. In the discussion of the present pore water data, we will use the present-day seawater values as a reference point to infer exchange processes. At Site 918, the pore water Li concentration decreases to 15 mM within the top 10 mbsf and reaches half of the present-day seawater value at 139 mbsf (Fig. 4a). At greater depths (.140 mbsf) dissolved Li concentrations increase gradually downcore. The Li concentration profile below 140 mbsf resembles a diffusion controlled curve and the concentration reaches a maximum (163 mM) at 765 mbsf. In the sand-rich layer (800 – 1100 mbsf), Li concentrations are lower than the maximum value observed at 765 mbsf. Correlated with the decrease in Li concentration at the shallow depths, the Li isotopic composition becomes heavier, reaching a minimum d6Li value of 241.4‰ at 55 mbsf, about 9‰ heavier than that of the present-day seawater (232.3‰, Chan and Edmond, 1988; Fig. 4b). Below 55 mbsf, the d6Li value increases to a maximum of 228‰,
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L. Zhang, L.-H. Chan, and J. M. Gieskes
corresponding to the NH1 4 maximum from 200 to 300 mbsf (see Fig. 6). A second maximum occurs at 765 mbsf with an isotopic composition of ;7‰ lighter than the seawater value. The lightest isotopic composition observed, other than in the bottom most sample, coincides with the maximum Li concentration implying a 6Li-rich source at this depth ( ;765 mbsf). In the deeper interstitial waters, the d6Li profile indicates a heavier isotopic composition (236.4‰) in the sand-rich layer and a lighter isotopic composition (224.7‰) at the sediment and crust interface. The Li concentration profile at Site 919 is similar to that of Site 918 at shallow depths and the d6Li data are almost identical (Fig. 5a,b), showing a rapid decrease to a minimum isotopic composition of 241.9‰. Below 70 mbsf, the Li isotopic composition returns to a seawater-like value without any significant change in Li concentration. The increase in d6Li again corresponds to the rise of NH1 4 concentration (Fig. 7). Some sediment samples from Hole 918 were analyzed for Li content and isotopic composition for comparison with the pore waters. The data for the solids are given in Table 2. Concentrations are low (13.6 –19.1 ppm) in the shallow, quartz-rich silt above 232 mbsf (Fig. 8a). The concentration increases to a maximum of 52.7 ppm at 575 mbsf and is lower in the silt layer with nanofossils below (600 – 800 mbsf). Li content decreases to 33 ppm in the sand rich layer and to 32 ppm in the volcaniclastic silt at the bottom of the hole. The isotopic compositions, shown in Fig. 8b, fall mostly in the range of 28 to 212‰ with the exception of relatively light isotopic values of 24.9‰ at 660 mbsf and 23.2‰ at 1157 mbsf where volcanogenic components are prevalent.
The results of the adsorption and desorption experiments are given in Table 3. The Gulf of Mexico surface water used for the adsorption experiment had an initial concentration of 24.4 mM and d6Li of 231.5‰. The concentration decreased with time while isotopic composition became heavier. After approximately 4900 h of reaction, 34, 43, and 23% of dissolved Li was removed by vermiculite, kaolinite, and Mississippi River sediment. Their isotopic fractionation factors (a) were determined to be 1.029 6 0.005, 1.021 6 0.005, and 1.021 6 0.007 to 1.024 6 0.008, respectively. 5. DISCUSSION
5.1. Volcanic Alteration Zones at Sites 918 and 919 Studies of 87Sr/86Sr ratios of interstitial waters have shown a strong correlation between alteration of volcanic material in the sediment column and lowering of 87Sr/86Sr below the contemporaneous seawater values, especially in volcanic material rich sediments (Elderfield and Gieskes, 1982; Gieskes et al., 1986, 1987). Since volcanic material in the sediment column is the only source that has 87Sr/86Sr ratios below the contemporaneous seawater values, the decrease of 87Sr/86Sr ratios is probably caused by isotopic exchange with volcanic material. 87Sr/86Sr ratios of pore waters are lower than the contemporaneous seawater values in three different zones (,100 mbsf, 140 –300 mbsf, and 500 – 800 mbsf) at Site 918 and in the entire pore water column (,134 mbsf) of Site 919. The amount of lowering of 87Sr/86Sr ratios largely reflects the degree of alteration of volcanic material within sediments, although it can be compli-
Lithium isotope geochemistry of pore waters
2441
Fig. 2. Depth distribution of 87Sr/86Sr ratio in pore waters at Site 918. The solid line represents contemporaneous seawater values at the corresponding depths. Also shown are lithological units and age boundaries of the sedimentary column (Shipboard Scientific Party, 1994).
cated by the dissolution of carbonates. Volcanic materials are present throughout the sediment columns at Sites 918 and 919 (Shipboard Scientific Party, 1994) as evidenced by high Ti/Al ratios of bulk sediments (Gieskes et al., 1998). The lowest 87 Sr/86Sr ratio occurs at 575 mbsf and coincides with the Ca and Mg extrema (Fig. 6) which are caused by the breakdown of plagioclase in the volcanic matter and incorporation of Mg into alteration clays. 5.2. Lithium Depletion at Depths above 100 mbsf At shallow depths (,100 mbsf) at Sites 918 and 919, the Li concentration of the pore waters decreases rapidly to about half of the seawater value (Figs. 4a and 5a). 87Sr/86Sr ratio data (Figs. 2 and 3) indicate alteration of volcanic material. Lithium is known to be incorporated in the alteration phases during low temperature alteration of basalt (,150°C; Seyfried et al., 1984; Stoffyn-Egli and Mackenzie, 1984). Calcium concentrations, decreasing to a shallow minimum (Figs. 6 and 7; Gieskes et al., in press) suggest precipitation of calcite. Both precipitation of calcite and alteration of volcanic material could remove Li from the pore waters. However, based on the calculation using a typical Li/Ca molar ratio of 20 3 1026 in marine carbonates, Li concentrations of pore waters should only decrease at most by 0.2 mM due to precipitation of calcite. Hence, precipitation
of calcite cannot have a significant influence on the Li concentrations and isotopic compositions of the pore waters at Sites 918 and 919. The rapid decrease of pore water Li in the upper section of the two sites is accompanied by a shift to heavier d6Li values relative to present-day seawater value (Figs. 4 and 5). The uptake of Li into alteration clays favors the lighter isotope leaving the pore water isotopically heavier relative to the initial seawater composition. This is consistent with observations on weathered seafloor basalts (Chan et al., 1992). The isotopic fractionation factor during this process may be estimated by a Raleigh fractionation model (Eq. 2). A plot of ln(R/Ro) vs. ln(C/Co) for the upper 60 m of Site 919 is shown in Fig. 9, assuming present-day seawater concentration and isotopic ratio for the initial pore fluid. This is reasonable since the Pleistocene-Pliocene boundary lies at about 120 mbsf (Shipboard Scientific Party, 1994) and the oceanic Li composition probably has not changed during the Holocene-Pleistocene (Delaney and Boyle, 1986; Hoefs and Sywall, 1997). The fractionation factor thus determined is 1.020. The isotopic fractionation factor for submarine basalt alteration at sea bottom temperatures in the North Atlantic has been determined to be 1.019 (Chan et al., 1992). The temperature gradient at Site 918 is about 5.6°C/100 m (Shipboard Scientific Party, 1994). Thus the temperatures of
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L. Zhang, L.-H. Chan, and J. M. Gieskes
Fig. 3. Depth distribution of 87Sr/86Sr ratio in pore waters at Site 919. The solid line represents contemporaneous seawater values at the corresponding depths.
Fig. 4. Lithium concentration (a) and isotopic composition (b) profiles at Site 918. The solid line in the respective diagrams represents the present-day seawater Li concentration and isotopic composition (Li concentration data are from Gieskes et al., 1998).
Lithium isotope geochemistry of pore waters
2443
Fig. 5. Lithium concentration (a) and isotopic composition (b) profiles at Site 919. The solid line in the respective diagrams represents the present-day seawater Li concentration and isotopic composition; Li concentration data are from Gieskes et al. (1998).
the pore waters in the upper 50 m section of Hole 919 are not significantly different from the sea bottom temperature, and the isotopic fractionation factor obtained at this site is consistent with low temperature alteration of volcanic material. Lithium depletion in interstitial waters observed at many ODP sites has been linked to the presence of altered volcanic material in the sediment column (Stoffyn-Egli and Mackenzie, 1984 and references therein). The Li isotopic data in this study support the suggestion that low Li concentrations in the interstitial waters are due to alteration of volcanic material. It has been speculated that diagenesis of volcanic material in sediments may be an important sink for Li in the oceans (StoffynEgli and Mackenzie, 1984) and that diffusion may carry dissolved Li in the oceans into sediments due to the concentration gradient caused by alteration of volcanic material. Using the concentration gradient (10 mM/10 m) at Site 919 and a diffusion coefficient of 1.48 3 102 cm2/year for Li at 0°C (Li and Gregory, 1974) and a porosity of 50%, the diffusive flux of Li into the sediments is estimated to be 3.7 3 1024 mmol/cm2. year (flux 5 (porosity)2 x diffusion coefficient x concentration gradient, Stoffyn-Egli and Mackenzie, 1984). The diffusive flux into the sediments in the entire ocean due to this diagenetic process is difficult to estimate because the content of volcanic matter and consequently the concentration gradient across the sediment-seawater interface would vary greatly geographically. At some locations the concentration gradient in the surface sediment is reversed, such as those observed in the Guaymas Basin (DSDP 477, Gieskes et al., 1982) and at the Peru margin (Martin et al., 1991). As an absolute upper limit, we assume that the diffusive flux estimated for Irminger Basin occurs over the entire ocean floor, yielding 1.3 3 109 mol/yr. This maxi-
mum flux may be compared to other Li fluxes in the ocean. The current estimates of the input and output fluxes are summarized in Table 4. It should be noted that the latest estimate of the river flux by Huh et al. (in press) is based on the flow-weighted mean concentration of the world’s major rivers and is considered the most accurate. It can be seen that Li removed by volcanic matter alteration in the sediments can be no more than 5% of the total inputs, mainly from rivers and submarine hydrothermal fluids (Stoffyn-Egli and Mackenzie, 1984 and Huh et al., in press). For this reason, post depositional alteration of volcanogenic components in the sediments is probably a minor sink of oceanic Li. 5.3. Source of Isotopically Light Lithium at 230 –350 mbsf Although 87Sr/86Sr and d18O data indicate that alteration of volcanic material occurs between 232 and 350 mbsf at Site 918 (Gieskes et al., 1998), d6Li shows an increase to 228‰, suggesting that other processes besides alteration of volcanic material take place in this depth interval. The reversal of d6Li to lighter isotopic values corresponds to the NH1 4 maximum in the zero SO22 zone and the zone of methanogenesis (Fig. 6; 4 Shipboard Scientific Party, 1994). A similar correlation is also observed at Site 919 (Fig. 7). Because the NH1 4 maximum is related to microbial decomposition of organic matter, the light Li isotopic composition might be partially derived from the organic matter. However, due to the relatively low concentration of Li in organic matter (0.5–1 ppm in dry organic matter; Heier and Billings, 1970) and the low organic content (;0.5%) of the sediments (Shipboard Scientific Party, 1994), the decom-
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L. Zhang, L.-H. Chan, and J. M. Gieskes
Fig. 6. Vertical distribution of Ca, Mg, sulfate, and NH1 4 in pore waters at Site 918 (from Gieskes et al., 1998).
position of the organic matter cannot be a direct cause for the increase of d6Li from a mass balance point of view. Cation exchange of NH1 4 with clay minerals may better account for the light Li isotopic composition at the NH1 4 maximum. In some ODP sites, more than half of the NH1 4 produced in pore waters undergoes ion exchange with cations in clay minerals (Gieskes, 1983, and references therein). Hence, it is highly possible that NH1 4 may expel Li from exchangeable positions in clay minerals. A small Sr concentra-
tion maximum and more radiogenic strontium isotopic composition are also observed at ;300 mbsf. Since pore water Li and Sr concentrations are in the mM range and NH1 4 production is in the mM range, the exchange of NH1 4 with cations in clay minerals can potentially influence the Li and strontium isotopic compositions of pore waters. If the exchange of Li in clay minerals with NH1 4 is to account for the d6Li increase, the amount of Li in exchangeable positions of clay minerals has to be large enough to cause a shift of
Lithium isotope geochemistry of pore waters
2445
Fig. 7. Vertical distribution of Ca, Mg, sulfate, and NH1 4 in pore waters at Site 919 (from Gieskes et al., 1998.
about 10‰ in d6Li. The isotopic fractionation factors of the two clay minerals and natural river sediments are indistinguishable within analytical error (Table 3). If we assume an average isotopic fractionation factor of 1.025, absorbed Li from and in equilibrium with seawater would have an isotopic composition of 28‰. At equilibrium, approximately 1 mg Li was adsorbed per gram of clay mineral in the adsorption experiment (Table 3). Given a unit volume (cm3) of sediment with a porosity of 50% and a pore water Li concentration of 13 mM (at 140 mbsf) as the concentration before the onset of desorption at Site 918,
the total amount of Li in the pore water in this volume would be about 45 ng. By using 2.7 g/cm3 for the density of the bulk sediment and 40 –50% for the percentage of clay minerals (Shipboard Scientific Party, 1994), the amount of the adsorbed Li in clay minerals is on the order of 500 ng. If approximately 4% of the adsorbed Li that has a light isotopic composition 6 (28‰) were desorbed into pore water by NH1 4 , d Li in the pore water would effectively shift from 239‰ to a lighter isotopic composition by 10‰. On the basis of this calculation, there should be an increase of Li concentration to 19 mM at
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L. Zhang, L.-H. Chan, and J. M. Gieskes
maximum exchange with NH1 4 . However, the actual concentration reaches 27 mM at the d6Li maximum (228.7‰) at 276 mbsf. The Li enrichment due to ion exchange with NH1 4 may, in part, be masked by the upward diffusion of Li from below (Fig. 4a). In contrast to the concentration increase with depth at Site 918, Li concentration remains essentially constant below 50 mbsf at Site 919 as NH1 4 rises (Fig. 5a). The lack of an explicit concentration increase at Site 919 may be the result of two competing processes: alteration of volcanic material and desorption. 87Sr/86Sr data indicate that volcanic alteration does occur throughout Hole 919. If we take the conditions at 50 mbsf as those before desorption, then 56 ng dissolved Li is present per cm3 of wet sediment (50% porosity). Assuming that d6Li of Li released from clay minerals by NH1 4 is 28‰ and the isotopic fractionation factor for low temperature alteration of volcanic material is 1.019, 112 ng Li that is desorbed by NH1 4 and later taken up by volcanic material is needed to shift the pore water 6Li from 242‰ to 232‰ as observed (Fig. 5b). This calculation demonstrates that the net result of these two
competing processes could cause the observed shift in d6Li in the pore water without any net change in dissolved Li concentration. The possibility of Li desorption by NH1 4 has been verified by the reaction of Louisiana shelf sediments with 1M ammonium acetate. After treatment with the solution, 0.5– 0.6 mg of Li was desorbed per gram of sediment (Table 3). We further note that the exchangeable Li is small compared to the Li present in the bulk sediment and is comparable to the amount adsorbed as
Fig. 8. Lithium concentration and isotopic composition of sediments from Site 918.
Lithium isotope geochemistry of pore waters
2447
Fig. 9. A plot of 6Li/7Li ratio vs. Li concentration ratio on the logarithmic scale for waters from the upper 50 m in Site 919 according to the Raleigh fractionation model. The slope of the linear regression line provides an estimation of isotopic fractionation factor for the alteration process at this site.
shown in the adsorption experiments. Even so, the exchange of NH1 4 with Li in clay minerals may exert significant control on the Li isotopic composition in sediment interstitial waters. As a corollary, we estimate the amount of oceanic Li that may be removed by adsorption onto sediments. If we assume adsorption of 1 mg Li per gram of sediment as shown in our experiments, and given a suspended load of 1.5 3 1016 g/yr for the world’s rivers (Martin and Meybeck, 1979) total Li removed from the oceans by adsorption on the suspended sediments would be about 2.2 3 109 moles Li/yr which is 5–10% of the total input estimated by Stoffyn-Egli and Mackenzie (1984) and Huh et al. (in press; Table 4). 5.4. Lithium Enrichment in Deeper Interstitial Waters From 500 to 800 mbsf, the Li concentration and the d6Li steadily increase with depth at Site 918, reaching a maximum concentration (163 mM) and the lightest isotopic composition (226‰) at 765 mbsf. Enrichment of Li in deep pore waters has been reported at many ODP Sites and several explanations have been proposed based on the correlation of Li concentration data with other geochemical data. At Site 474, the enrichment was interpreted as a result of the transformation of biogenic opal to opal-A (Gieskes et al., 1982). The injection of Li-rich hydro-
thermal solutions from the underlying basalt has been proposed to explain pore water Li enrichment in deep sea sediment at the Guaymas Basin spreading center (Gieskes et al., 1982). In contrast, at the Peru margin, the enrichment has been attributed to reactions with continental and oceanic crusts as indicated by strontium isotopic data (Elderfield et al., 1990; Martin et al., 1991). At Site 918, based on Ca, Mg, and 87Sr/86Sr data, the reactions in underlying basaltic basement are unlikely to have an important influence on the pore water chemistry. The maximum concentration and d6Li at 765 mbsf suggest a Li source within the sediment column. Characterization of the concentration and isotopic composition of various types of marine sediments is essential to identify the source of enrichment in the interstitial waters. In a recent survey of Li isotopes in marine sediments (Chan et al., 1994b), carbonates and siliceous oozes were found to be low (;1 ppm) in Li while hemipelagic and pelagic clays have 45– 62 ppm with d6Li values ranging from 210 to 214‰, distinct from oceanic crust (; 25‰; Chan et al., 1992). The Li concentration profile of the solids at Site 918 displays a similar shape as the dissolved Li profile (Figs. 4a and 8a). The silty sediments above 300 mbsf are quartz rich (Shipboard Scientific Party, 1994) and have relatively low Li contents.
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L. Zhang, L.-H. Chan, and J. M. Gieskes
Peak concentrations occur between 600 and 800 mbsf. At ;765 mbsf, where the maximum dissolved Li is found, the Li content in the solid phase is 47 ppm. Given a porosity of 50%, and a sediment density of 2.7 g/cm3, the Li content of the pore water represents only 1% of the Li inventory in the solid. Therefore, the similarity between solid and pore water Li profiles suggests that the Li concentration in the deeper interstitial water is largely controlled by the associated sediment. Experimental studies have shown that Li can be released from marine sediments at temperatures as low as 50 – 60°C (Chan et al., 1994a; You, 1994). In the experimental system with a water/sediment ratio of 3:1, Li released to the fluid reached 40 mM at 60°C after reaction for several days and is approximately 5– 6‰ heavier than that in the initial sediment. At Site 918, the geothermal gradient is about 5.6°C/100 m (Shipboard Scientific Party, 1994) and the temperature at the concentration maximum (;765 mbsf) is about ;45°C. With a much lower water/ sediment ratio (;1:3) and considerably longer reaction time, it is possible that Li can attain the observed concentration of 163 mM at this depth. That the enrichment originates from the pelagic sediments is also consistent with the Li isotopic composition in the pore waters. d6Li in pore water correlates well with 1/Li (r2 5 0.869) from 500 mbsf to 800 mbsf (Fig. 10). This linear correlation suggests addition of Li released from the sediments to the pore water of evolved seawater composition. If we assume the water at 515 mbsf is the evolved seawater endmember (with 57 mM Li and 237.7‰) that interacts with the deeper sediments, it can be calculated from the mixing relationship that Li released from the sediment has d6Li of
about 220‰. Available data show that Li isotopic compositions of the sediments at 918 mostly vary between 28 and 212‰ with two exceptions at 660 and 1157 mbsf where the isotopic compositions are very light (24.9 and 23.2‰; Fig. 8b). These depths are marked by volcanogenic components (Shipboard Scientific Party, 1994) which are isotopically light. Extraction of the sediment-bound Li with attendant isotopic fractionation may account for the enrichment of relatively light Li in the fluids between 550 and 800 mbsf. Although 87Sr/86Sr data indicate strong alteration of volcanic material in this section (a process that takes up Li from solution), the Li distribution appears to be dominantly controlled by the input from the sediments. The interval between 806 and 1100 mbsf consists of turbidite sands interbedded with nanofossil chalk (Shipboard Scientific Party, 1994). In this layer Li concentration decreases and d6Li shifts to 236.6‰. Chloride and oxygen isotopic data indicate advection of trapped meteoric water through the sand bed (Gieskes et al., 1998). However, the advecting fresh water probably has low Li content and would not significantly affect the Li isotopic composition of pore water in this layer. Hence, the d6Li value may largely reflect that of diluted pore water without significant input from the sediments at this horizon. The deepest water sample (1157 mbsf) has a relatively light isotopic composition (224.8‰). The concentration of pore water at this level (91 mM) is considerably higher than the seawater value. Thus the isotopic composition is probably not controlled by Li uptake into volcanic material, neither can hydrothermal fluids from the basement be the source of Li enrichment because their d6Li values are between 26 and 211‰ (Chan et al., 1993). The 87Sr/86Sr ratios below 925 mbsf are also higher than the contemporaneous seawater values which suggests alteration of more radiogenic material in the sediment column and possibly diffusion from the maximum at 925 mbsf (Gieskes et al., 1998). Similarly, the relatively light isotopic composition of Li in the deepest pore water may be attributed to interaction with the sediment at this depth. 6. SUMMARY AND IMPLICATIONS
At shallow depths (,100 mbsf), pore water Li concentrations are mainly controlled by alteration of volcanic material within the marine sediments. d6Li data support that Li has been taken up by altered volcanic material. This study confirms that alteration of volcanic material is not restricted to the exposed oceanic crust but also occurs in the volcanogenic components of marine sediments. It is estimated that the diffusive flux into the sediments as a result of this process may at most account for 5% of the sink for the dissolved Li in the global ocean. Relatively light isotopic compositions at Sites 918 and 919 are observed to correspond with the NH1 4 maximum . It is proposed that NH1 4 generated by bacterial decomposition of organic matter may expel adsorbed Li from clay minerals by ion exchange and significantly influence the isotopic composition of pore water. This suggestion is supported by results of adsorption and desorption experiments. Total Li removed from the ocean water by adsorption on to the suspended sediments is approximately 5–10% of the total input. In the deep sediment column (500 – 800 mbsf) at Site 918, Li isotopic data suggest that the enrichment of pore water Li is
Lithium isotope geochemistry of pore waters
2449
Fig. 10. d6Li vs. 1/Li plot for pore waters between 500 and 800 mbsf at Site 918. The linear relationship indicates addition of Li derived from sediments to the pore waters.
most likely derived from pelagic sediments. This is consistent with experimental studies which have shown that Li is easily mobilized from marine sediments at elevated temperatures. Our study suggests that Li enrichment in deep interstitial waters, observed in many ODP cores, may be explained by release of Li from sediments at elevated temperatures and low water/ sediment ratio. In conclusion, this study demonstrates that Li isotopic compositions of pore water can help elucidate a variety of processes in the sediment column and are a useful indicator of sediment-water interaction. It is also found that incorporation of Li into authigenic clay minerals from volcanic matter alteration in the sediment column and Li adsorption on sediments are relatively minor sinks of oceanic Li.
Acknowledgments—We are grateful to the Ocean Drilling Program for providing pore water and sediment samples from Leg 152 for this study. We thank Eric De Carlo and an anonymous reviewer for their valuable comments and suggestions. This work was performed with the support of National Science Foundation grant OCE 9314708 to L. H. C. and the support from the US Science Advisory Committee to J. M. G.
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