Zinc Isotopes - Geochemistry - Lecture Notes, Study notes for Geochemistry. Annamalai University
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Zinc Isotopes - Geochemistry - Lecture Notes, Study notes for Geochemistry. Annamalai University

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Following are the key concepts discussed in these Notes : Zinc Isotopes, Concentrations, Correlation, Indian Oceans, Isotopically Uniform, Manganese Nodules, Apparent Positive, Fractionation, Riverine, Principal
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IsotopeGeochemistry Chapter11

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al. (2010). In systems involving water, therefore, speciation may introduce significant Zn isotope frac- tionations. Data on zinc isotopes in seawater is very limited and the Zn isotopic budget of seawater has yet to be worked out. δ66Zn ≈ +0.3‰ in a single sample of English Channel water, while in the upper 400 m of North Pacific seawater, δ66Zn varying from -0.15 to +0.15‰ and correlating negatively with δ65Cu (Bermin et al., 2006). Although Zn concentrations are strongly depleted in surface water due to biologi- cal uptake, that study found no apparent correlation between δ66Zn and Zn concentration. Zn in eight deep water samples from the Pacific and Indian Oceans appears to be slightly heavy compared to sur- face water and somewhat more isotopically uniform, with δ66Zn varying from +0.22‰ to +0.64‰ and averaging δ66Zn = 0.45±0.14‰ (Bermin, 2006). Manganese nodules have δ66Zn that range from +0.53‰ to +1.16‰ and average 0.90‰ (Maréchal et al., 2000). Thus there is an apparent positive fractionation of Zn isotopes in Mn nodule formation. The. John et al. (2008) estimate that the riverine and atmos- pheric inputs to seawater are relatively light 66Zn ≈ 0.1 to 0.3‰; however, the principal source of Zn in seawater is hydrothermal fluids, which, as we found above, are isotopically variable, although perhaps somewhat heavier than the riverine source on average. Pichat et al., (2003) found that δ66Zn in sediments from ODP core 849 from the eastern equatorial Pa- cific, whose age ranged from 3 ka to 174ka, showed a general decrease with time, varying from +0.31 to +1.34‰. Superimposed on this decrease was a weak periodic variation with periodicities of 35.3 and 21.2 ka. The material in the core was primarily biogenic: carbonate tests of cocolithophorids and fora- minifera (and indeed, the analysis was restricted to the carbonate fraction of the sediment). The 21.2 ka periodicity corresponds closely to the mean periodicity of the precession periodicity of the Mi- lankovitch variations (Chapter 10). Pichat et al thus interpreted these variations as climatic effects of oceanic circulation (specifically equatorial upwelling) and biological productivity. Zinc is a remarkably important element in biology. It is the second most abundant transition metal, following iron, in organisms. It is found in all classes of enzymes as well as in transcription factors (pro- teins that bind to DNA sequences, effectively turning genes on and off). It is important in protein syn- thesis and membrane activity, conversion of CO2 to bicarbonate ion and visa versa (in both plants and animals), brain activity, leaf formation, digestion, and so on. Nevertheless data on Zn isotope fractiona- tion in organisms is fairly limited at present (this will undoubtedly change in the future). In hydro- ponic‡ growth experiments with tomatoes, lettuce, and rice, Weiss et al. (2005) found that roots of these plants preferentially took up isotopically heavy Zn from the growing solutions and that the shoots of the plants were isotopically light compared to the roots (by up to 0.5‰). Simple growth experiments of Moynier et al. (2009) with lentils (beans) and bamboo confirmed the preferential transport of isotopi- cally light Zn up the plant and into the leaves – suggesting the process is at least partially controlled by diffusion. Fractionations observed in bamboo were up to 1‰. John et al. (2007) found that experimen- tally grown marine diatoms preferentially took up isotopically light Zn and that the fractionation de- pended on Zn concentrations. At low concentrations, the fractionation was about -0.2‰ but at higher concentrations it was -0.8‰. The difference likely reflects two biochemical pathways being used to transport Zn into cell interiors: a high-affinity mechanism associated with a small isotopic fractionation predominates at low concentrations but becomes saturated at higher concentrations, where the low- affinity mechanism becomes dominant. In contrast, Andersen et al., (2011) found that Zn in cleaned dia- toms tests from sediment core tops from the Southern Ocean was isotopically heavy, with δ66Zn rang- ing from +0.7‰ to +1.5‰. δ66Zn correlated negatively with Zn concentrations in the tests and biogenic silica accumulation rates in the sediment. Assuming that Zn concentrations in the tests correlates with Zn concentrations in surface seawater in which the diatoms grew, this suggests surface seawater be- comes isotopically heavier as it becomes progressively Zn-depleted by biological uptake. Because autot- rophs such as plants and diatoms preferentially take up Zn2+ rather than organically complexed Zn,

‡ Hydroponics is a method of growing plants from a solution without soil.

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Zn, and because there is isotopic fractionation between species in solution, the degree to which frac- tionations in biological systems reflect fractionations between dissolved Zn species rather than bio- chemical processes will require further study. Zn isotope ratios in animals is remarkably variable, from -0.4 to 0.8‰; human blood appears to be more uniform and isotopically heavy (δ66Zn +0.3 to +0.4‰) (Cloquet et al., 2008; Marèchal et al., 1999).

11.6 ISOTOPES OF BORON AND LITHIUM Although there are a few earlier works in the literature, there was little interest in the isotopic compo- sition of boron and lithium until about 25 years ago. This are perhaps several reasons for this: because both elements have low abundances in the Earth, they have only one valance state in nature, neither B nor Li form gaseous species that can be analyzed in the gas-source mass spectrometers conventionally used for analysis of other stable isotopes, and the fractionation produced in thermal ionization mass spectrometers, would exceed the natural ones. Since the development of new analytical techniques that overcame this latter problem in the 1980’s, the fields of boron and lithium isotope geochemistry have developed rapidly. Though both lithium and boron can occur as stoichiometric components of minerals, these minerals have limited occurrence, and these elements generally substitute for other elements in silicates. Boron is relatively abundant in seawater, with a concentration of 4.5 ppm. Lithium is somewhat less abun- dant, with a concentration of 0.17 ppm. Both are “conservative” species in seawater, which is to say they are always present in a constant ratio to salinity. In silicate rocks, the concentration of boron ranges from a few tenths of a ppm or less in fresh basalts and peridotites to several tens of ppm in clays. Lithium concentrations in these materials typically range from a few ppm to a few tens of ppm. Boron has two isotopes: 10B and 11B whose abundances are 19.9% and 80.1%, respectively (Coursey et al., 2011). The 11B/10B is reported as per mil variations, δ11B, from the NIST SRM 951 standard. Li also has two isotopes: 6Li and 7Li whose abundances are 7.59% and 92.41%, respectively (Coursey et al., 2011). The 7Li/6Li ratio is reported as per mil variation, δ7Li, from the NIST SRM 8545 Li2CO3 (L-SVEC) standard (Table 11.1). Prior to 1996, Li isotope ratios were often reported as δ6Li, i.e., deviations from the 6Li/7Li ratio. However, the standard used was the same, so that for variations less than about 10‰, δ7Li ≈ -δ6Li; at higher deviations, a more exact conversion is necessary, e.g., -38.5‰ δ6Li = 40‰ δ7Li. The analytical precision for most of the Li isotope data now in the literature is about 1‰, but recent advances, particularly the use of multiple-collector inductively coupled plasma mass spectrometers, has reduced uncertainty to as little as 0.2‰.

11.6.1 Boron Isotopes In nature, boron has a valence of +3 and is almost always bound to oxygen or hydroxyl groups in ei- ther trigonal (e.g., BO3) or tetrahedral (e.g.,B(OH)4

– ) coordination (the only exception is boron bound to fluorine, e.g., BF3). Since the bond strengths and vibrational frequencies of trigonal and tetrahedral forms differ, we can expect that isotopic fractionation will occur between these two forms. This is con- firmed by experiments which show a roughly 20‰ fractionation between B(OH)3 and B(OH)4

– , with 11B preferentially found in the B(OH)3 form. In solution, the reaction between trigonal and tetrahedral bo- ron is quite fast, implying equilibrium can be expected and that the isotopic fractionation should reflect this equilibrium. In natural aqueous solutions boron occurs as both boric acid, B(OH)3, and the borate ion, B(OH)4

– , the dominant species being determined by pH. At pH of around 9 and above, B(OH)4

– , dominates, at lower pH B(OH)3 dominates. In seawater which has a pH in the range of 7.6 to 8.1, about 80-90% of boron will be in the B(OH)3 form. Most fresh waters are a little more acidic so B(OH)3 will be more dominant; only in highly alkaline solutions, such as saline lakes, will B(OH)4

– be dominant. The most common bo- ron mineral in the crust is tourmaline (Na(Mg,Fe,Li,Al)3Si6O18 (BO3)3(OH,F)4), in which boron is present in BO3 groups. In clays, boron appears to occur primarily as B(OH)4

– , most likely substituting for silica

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in tetrahedral layers. The coordination of boron in common igneous minerals is uncertain, possibly substituting for Si in tetrahedral sites. Boron is an incompatible element in igneous rocks and is very fluid-mobile. It is also readily adsorbed onto the surfaces of clays. There is an isotopic fractionation be- tween dissolved and adsorbed B of –20 to -30‰ (i.e, adsorbed B is 11B poor), depending on pH and temperature (Palmer et al., 1987). Figure 11.12 illustrates the variation in B isotopic composition in a variety of geologic materials. Spivack and Edmond (1987) found the δ11B of seawater was uniform within analytical error. This has more recently been confirmed by Foster et al. (2010) who determined a δ11B value of 39.61±0.04‰. Fresh mid-ocean ridge basalts have a mean δ11B of -4.25±2.1‰. Oceanic island basalts (OIB) have slightly lighter δ11B (e.g., Chaussidon and Jambon, 1994). Bulk chondrites have δ11B similar to MORB, which presumably is approximately the bulk silicate Earth value. However, meteoritic materials can have quite variable δ11B as a consequence of a variety of processes, including cosmogenic production and decay of 10Be both in the early solar system and subsequent exposure of the meteorites to cosmic rays. The average B iso- topic composition of the continental crust probably lies between -13‰ and -8‰ (Chaus- sidon and Abarède, 1992). Perhaps the most remarkable aspect of B iso- tope geochemistry is the very large fractiona- tion of B isotopes between the oceans and the silicate Earth. It was recognized very early that this difference reflected the fractionation that occurred during adsorption of boron on clays (e.g., Schwarcz, et al., 1969). However, as we noted above, this fractionation is only about 30‰ or less, whereas the difference between the continental crust and seawater is close to 50‰. Furthermore, the net effect of hydro- thermal exchange between the oceanic crust and seawater is to decrease the δ11B of seawater (Spivack and Edmond, 1987). The discrepancy reflects the fact that the ocean is not a simple equilibrium system, but rather a kinetically controlled open one. Since all processes oper- ating in the ocean appear to preferentially re- move 10B from the ocean, seawater is driven to an extremely 11B-rich composition. Ishikawa and Nakamura (1993) noted that ancient lime- stones and cherts have more negative δ11B than their modern equivalents, calcareous and sili- ceous oozes, and suggested that further frac- tionation occurred during diagenesis. Another large fractionation occurs during evaporation and precipitation. Rose-Koga et al. (2006) found that rain and snow is substantially lighter than seawater with d11B ranging from - 10 to +34‰. They estimate a seawater–vapor

Figure 11.12. Boron isotopic composition in crystalline rocks (MORB: mid-ocean ridge basalts; OIB: oceanic island basalts; BABB: back-arc basin basalts, IAV: is- land arc volcanics), sediments, groundwater, freshwa- ter, salt lakes, seawater, and mid-ocean ridge hydro- thermal solutions.

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fractionation is +25.5‰ Spivack and Edmond (1987) investigated the exchange of boron between seawater and oceanic crust. Boron is readily incorporated into the alteration products of basalt, so that even slightly altered basalts show a dramatic increase in B concentration and an increase in δ11B, with altered oceanic crust having δ11B in the range of 0 to +25‰. Smith et al. (1995) estimated that average altered oceanic crust contains 5 ppm B and δ11B of +3.4‰. During high temperature interaction between seawater and oceanic crust, Spivack and Edmond (1987) concluded that boron was quantitatively extracted from the oceanic crust by hydrothermal fluids. The δ11B of these fluids is slightly lower than that of seawater. They inferred that the B in these fluids is a simple mixture of seawater- and basalt-derived B and that little or no iso- topic fractionation was involved. Analysis of hydrothermal altered basalts recovered by the Ocean Drilling Project generally confirm these inferences, as they are boron-poor and have δ11B close to 0 (Ishikawa and Nakamura, 1992). Island arc volcanics (IAV) have distinctly more positive δ11B than either MORB or OIB. This may in part reflect the incorporation of subducted marine sediment and altered oceanic crust into the sources of island arc magmas (e.g., Palmer, 1991), but a more important factor is likely to be fractionation of B isotopes during dehydration of the subducting slab, in which 11B is strongly partitioned into the fluid phase (You et al., 1995). This process begins at quite shallow depths in the subduction zone, in the forearc region, where fluids released from the overlying slab serpentinize the forearc mantle and enrich it in 11B, producing forearc serpentinites with δ11B as high as +30‰. By the time the slab reaches the sub-arc magmagenesis region, fluids released are less 11B-rich, such that island arc volcanics typically have δ11B of around +5‰. As Figure 11.12 shows, that the distribution of δ11B in island arc volcanics is distinctly bimodal, with modes at +5‰ and +16‰. All samples with δ11B > 12 are from the South Sandwich arc. Tonarinni et al. (2011) interpret these very high values as a consequence of subduction erosion of the forearc region; i.e., they suggest the strongly 11B serpentinites of the forearc are being car- ried into the magmagenesis zone. The differences in δ11B between oceanic island basalts (OIB) and MORB is perhaps more problematic. Though no experimental or theoretical studies have been carried out, it seems unlikely that significant fractionation of boron isotopes will occur during melting of the mantle, both because the temperatures are high, and because the atomic environment of B in silicate melts is probably similar to that in silicate solids. Thus, as we found was the case for O isotopes, B isotope fractionation probably occurs only at the surface of the Earth, and the difference between OIB and MORB must somehow reflect surface processes. Chaussidon and Marty (1995) argued that the boron isotopic composition of the mantle is that of OIB (-10‰) and that the higher δ11B of MORB reflects contamination of MORB magmas by al- tered oceanic crust. This seems unlikely for several reasons. First, although there are still relatively few data available, MORB appear to be relatively homogeneous in their boron isotopic composition. This means assimilation would have to be systematic and pervasive and that all MORB magmas would have to assimilate roughly the same amount of material. Both of these seem highly improbable. Second, there is little or no other evidence for pervasive assimilation of oceanic crust by MORB magmas. Third, oceanic island basalts have an opportunity to assimilate not only altered oceanic crust, but also overly- ing sediment. Yet, according to the Chaussidon and Marty (1995) hypothesis, they are not systemati- cally contaminated. Although they are not systematically contaminated, there is evidence of occasional assimilation of oceanic crust and/or sediment by oceanic island basalt magmas from both B and Os iso- tope geochemistry. This may explain some of the lower δ11B values in OIB seen in Figure 11.04 (Chaus- sidon and Jambon, 1994). The alternative explanation for the lower δ11B in OIB is that they contain a component of material re- cycled into the mantle, through subduction, from the surface of the Earth. The idea that mantle plumes, and the OIB magmas they produce, contain material recycled from the surface of the Earth has been suggested on other grounds. As we noted above significant fractionation of B isotopes occurs during sediment dehydration during subduction (You, 1994). The fluid produced will be enriched in 11B, leav-

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ing the sediment and oceanic depleted in 11B. Thus the effect of subduction zone processes will be to lower the δ11B of oceanic crust and sediment carried into the deep mantle. One of the more interesting applications of boron isotopes has been determining the paleo-pH of the oceans. Boron is readily incorporated into carbonates, with modern marine carbonates having B con- centrations in the range of 15-60 ppm. In modern foraminifera, δ11B is roughly 20‰ lighter than the seawater in which they grow. This fractionation seems to result from the kinetics of B co-precipitation in CaCO3, in which incorporation of B in carbonate is preceded by surface adsorption of B(OH)4

– (Ven- gosh et al., 1991; Heming and Hanson, 1992). We noted above that boron is present in seawater both as B(OH)3, and B(OH)4

– . Since the reaction be-

Figure 11.13. Top graph shows the variation pH of surface seawater during Tertiary time as in- ferred from δ11B in shells of planktonic foraminifera in ODP cores. Bottom graph shows the concen- tration of atmospheric CO2 calculated from seawater pH. Adapted from Pearson and Palmer (2000).

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tween the two may be written as: B(OH)3+H2O ®   B(OH)4

- +H+ 11.4 The equilibrium constant for this reaction is:

pKapp = ln B(OH)4 -

B(OH)3 − pH 11.5

The relative abundance of these two species is thus pH dependent. Furthermore, we can easily show that the isotopic composition of these two species must vary if the isotopic composition of seawater is constant. From mass balance we have: δ11BSW = δ11B3ƒ + δ11B4(1-ƒ) 11.6 where ƒ is the fraction of B(OH)3 (and 1 - ƒ is therefore the fraction of B(OH)4

− ), δ11B3 is the isotopic com- position of B(OH)3, and δ11B4 is the isotopic composition of B(OH)4

− . If the isotopic composition of the two species are related by a constant fractionation factor, ∆3-4, we can write 11.6 as: δ11BSW = δ11B3ƒ + δ11B4-δ11B4ƒ = δ11B4 - ∆3-4ƒ 11.7 Solving for δ11B4, we have: δ11B4 = 11BSW + ∆3-4ƒ 11.8 Thus assuming a constant fractionation factor and isotopic composition of seawater, the δ11B of the two B species will depend only on ƒ, which, as we can see in equation 11.5, will depend on pH. If the mechanism for incorporation of B in carbonate suggested by Vengosh et al. (1991) is correct, the fractionation of 11B/10B between calcite and seawater will be pH dependent. There is still some debate as to as to the exact mechanism of boron incorporation in carbonate, in particular whether only borate ion or both boric acid and borate ion can be incorporated. Regardless of the exact mechanism, the iso- topic composition of boron in carbonate precipitated from seawater has been shown to be a strong function of pH (Sanyal et al., 1996), allowing in principle the reconstruction of paleo-seawater pH from carbonates. There are a some additional factors that must be considered: (1) different carbonate- secreting species may fractionate B isotopes slightly differently, perhaps because they alter the pH of their micro-environment, or perhaps because B(OH)3 is also utilized to varying degrees, (2) the frac- tionation factor is temperature dependent, and (3) the B isotopic composition of seawater may vary with time. Nevertheless, if care is exercised to account for “vital” effects and variation in the isotopic composition and temperature of seawater, the B isotopic composition of marine biogenic carbonate pre- served in sediment should reflect the pH of the water from which they were precipitated. The pH of seawater, in turn, is largely controlled by the carbonate equilibrium, and depends there- fore on the partial pressure of CO2 in the atmosphere. Thus if the pH of ancient seawater can be deter- mined, it should be possible to estimate pCO2 of the ancient atmosphere. Given the concern about the re- lation of future climate to future pCO2, it is obviously interesting to know how these factors related in the past. Pearson and Palmer (2000) measured δ11B in foraminiferal carbonate extracted from Ocean Drilling Program (ODP) cores and from this calculated pH based on equations 11.5 through 11.8. To minimize the effect of temperature on the fractionation factor, they restricted their study to cores from regions that were tropical at the time of deposition. To minimize vital effects, they used only 1 species of planktonic foraminifera for the Neogene period, G. trilobus (also known as G. sacculifer), which is thought to incorporate B with no vital effect. For the Paleogene, they used 6 species where the vital ef- fect was arguably minimal. They argue that changes in the B isotopic composition of seawater should be slow to occur since the residence time of B in seawater is roughly 20 million years. Nevertheless, they account for a small variation, roughly 1.7‰, in seawater δ11B over Tertiary time. The results sug- gest dramatically lower seawater pH and dramatically higher pCO2 in the Paleogene. The apparent variation in pCO2 is qualitatively consistent with what is known about Tertiary climate change – namely

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that a long-term cooling trend began in the early to middle Eocene. In contrast to the Paleogene, Figure 11.13 shows that the Neogene is characterized by atmospheric pCO2 near or slightly below modern pre-industrial levels. This aspect of the result is somewhat surprising, given the evidence for cooling through the Neogene, particularly over the last 10- 15 Ma (e.g., Figure 10.36). If, as Figure 11.13 suggests, atmospheric pCO2 has been nearly constant through the Neogene, some factor other than the greenhouse effect must account for the cooling over this time. On a more limited time scale, Hönisch and Hemming (2005) investigated δ11B over the last two glacial cycles (0- 140 and 300-420 ka). In this study, they controlled for temperature by analyzing the Mg/Ca ratio of the carbonate shells, which is know to be strongly temperature dependent. Their calculated

pH values ranged from 8.11 to 8.32, which in turn correspond to a pCO2range of ~180 to ~325 ppm. These calculated pCO2 values are in good agreement with CO2 concentrations measured in bubbles in the Vostok ice core. Hönisch et al. (2009) extended the boron isotope-based pCO2 record through the Pleis- tocene. The were particularly interested in the mid-Pleistocene time when the dominant orbital forcing frequency switched from 40,000 years to 100,000 years. They found that whereas over the last 400,000 years pCO2 has varied from 180 and 300 ppmv in glacial and interglacial periods, respectively, the variation had been 210 and 280 ppmv in the early Pleistocene. The calculated average pCO2 of the early and late Pleistocene are 248 and 241 ppmv, respectively, and statistically indistinguishable. This a de- crease in pCO2 does not seem to be responsible for the more severe glaciations of the late Pleistocene. Pearson et al. (2009), working with foraminifera from a well-preserved Paleogene section in Tanzania determined pCO2 over the Oligocene-Eocene transition, a time of great interest because it was then that Antarctic glaciation began. Some had suspected that a drop in pCO2 below 750 ppmv may have been responsible for the cooling that led to formation of the ice sheet. Their δ11B results suggest that pCO2 did dip below this level near this boundary, but recovered in the early Oligocene before declining again (Figure 11.14). Although still somewhat controversial, paleo-seawater pH reconstruction based on boron isotopes has become better grounded in the last few years. Some of the controversy centered around the lack of

Figure 11.14. a, pCO2 (± uncertainty from δ11B measurements at 95% confidence; blue symbols calculated assuming varying [CO3-2], red symbols are for constant [CO3-2]). The grey band is the assumed threshold for Antarctic glaciation. b, Deep-sea oxygen isotopes from DSDP site 522 (crosses) and ODP site 744 (diamonds). c, δ11B proxy measurements compared to alkenone proxy estimates (green symbols with green shaded band for maximum and minimum).

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experimental determination of the fractionation factor between B(OH)3 and B(OH)4– (Pagani et al., 2005; Zeebe 2005). This has be partially resolved by Klochko et al. (2006) who experimentally determined the fractionation factor, α, to be 1.0272± 0.0006 and independent of temperature. This value differs from the theoretically estimated value of 1.0194, which had previously been used for paleo-oceanographic calculations. However, the theoretical is hampered by uncertainties in measured vibrational frequen- cies, and attempts to calculate it ab initio from molecular orbital theory result in uncertainties that are too large to be useful. In addition to this, there are additional uncertainties in the mechanism of incor- poration of boron into carbonate and the associated fractionation factor, and uncertainties in how δ11B in seawater has varied with time, and the requirement of high precision measurement of B isotope ra- tios. Foster (2008) has shown that paleo-CO2 determined from foraminifera in cores from the Caribbean agree well with CO2 in the Vostok ice core using the experimentally determined fractionation factor.

11.6.2 Li Isotopes Terrestrial lithium isotopic variation is dominated by the strong fractionation that occurs between minerals, particularly silicates, and water. Indeed, this was first demonstrated experimentally by Urey in the 1930’s. This fractionation in turn reflects the chemical behavior of Li. The ionic radius of Li1+ is small (78 pm) and Li readily substitutes for Mg2+, Fe2+, and Al3+ in crystal lattices, mainly in octahedral sites. In aqueous solution, it is tetrahedrally coordinated by 4 water molecules (the solvation shell) to which it is strongly bound, judging from the high solvation energy. These differences in atomic envi- ronment, differences in binding energies, the partly covalent nature of bonds, and the low mass of Li all lead to strong fractionation of Li isotopes. Modern study of Li isotope ratios began with the work of Chan and Edmond (1988). They found that the isotopic composition of seawater was uniform within analytical error with a δ7Li value of +33‰ (which has subsequently been revised to +31‰ based on more accurate techniques). Subsequent work suggests that δ7Li in seawater might vary by as much as 4‰, but the degree to which this variation re- flects analytical error and inter-laboratory biases remains unclear, as the residence time of Li in the ocean (1.5-3 Ma) is much longer than the mixing time. Fresh MORB have δ7Li of +3.8±1.3‰ (White and Klein, 2013), a range not much larger than that expected from analytical error alone. Oceanic island basalts (OIB) have on average higher δ7Li: +4.9±1.2‰. The highest δ7Li occurs on islands characterized by particularly radiogenic Pb (the so- called HIMU OIB group), such as on some islands of the Cook-Austral chain. Here, δ7Li may be as high as +8‰. This may reflect a recycled crustal component in their sources (e.g., Vlastelic et al., 2009). High precision analyses of whole carbonaceous and ordinary chondrites have a narrow range of values with a mean of 2.96±0.77‰; enstatite chondrites appear to be systematically lighter, with a mean of 1.69±0.73‰ (Pogge von Strandmann, et al., 2011). Individual components of meteorites are more vari- able. The mean δ7Li of fertile, unaltered and unmetasomatized peridotites is +3.5±0.5‰, which is pre- sumably the value of the bulk silicate Earth, is essentially indistinguishable from that of carbonaceous and ordinary chondrites (Pogge von Strandmann, et al., 2011). Other peridotites exhibit a larger range: -15‰ to +10‰. There appears to be little the relationship between δ7Li and other geochemical parameters in igneous rocks, such as MgO concentration, suggesting Li experiences little isotopic fractionation during frac- tional crystallization, and perhaps also partial melting (Tomascak et al., 1999; Teng et al., 2004). How- ever, Pogge von Stratdmann et al. (2011) found a correlation between δ7Li and δ26Mg in peridotites xenoliths, which they attribute to kinetic isotope fractionation occurring during transport from the mantle. Alpine eclogites can have distinctly light isotopic compositions. These rocks are thought to be fragments of basaltic oceanic crust deeply subducted and metamorphosed then subsequently rapidly exhumed during the Alpine orogeny. Their light isotopic compositions presumably reflect preferential partitioning of 7Li into the fluid produced as the rocks were dehydrated during metamorphism. δ7Li in

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granites ranges from about -5 to +10‰, with an average of +1.7‰. Based on this and analysis of other crustal materials, Teng et al. (2009) estimate the average continental crust to have δ7Li = +1.2‰. δ7Li in oceanic crust altered by seawater at low temperatures takes up Li from solution and has high δ7Li compared to fresh basalt. In hydrothermal reactions, however, Li is extracted from basalt into the solution and hydrothermal fluids can Li con- centrations up to 50 times greater than sea- water. 7Li is extracted more efficiently than 6Li during this process, so hydrothermally al- tered basalt can have δ7Li as low as -2‰. Serpentinites (hydrothermally altered peri- dotite) can have even lower δ7Li. Because they extract Li from oceanic crust so com- pletely, hydrothermal solutions have Li iso- topic compositions intermediate between MORB despite this fractionation. The average δ7Li of island arc volcanics, +4.05±1.5‰ is only slightly higher than that of MORB. This is somewhat surprising since other isotopic evidence clearly demonstrates island arc magmas contain components de- rived from subducted oceanic crust and sediment. Furthermore, while δ7Li have been shown in some cases to correlate with chemi- cal and isotopic indicators of a subduction component, this is not always the case. It seems nevertheless likely that the sub- duction process has profoundly influenced the isotopic composition of the mantle over time. As a consequence of fractionation oc- curring during weathering, seawater is strongly enriched in 7Li. This enrichment is imprinted upon the oceanic crust as it reacts with seawater. When the oceanic crust is re- turned to the mantle during subduction, the mantle becomes progressively enriched in 7Li. The continental crust, on the other hand, becomes progressively depleted in 7Li over time. Elliot et al. (2004) calculate that this process has increased δ7Li in the mantle by 0.5 to 1‰ and decreased δ7Li of the continen- tal crust by 3‰ over geologic time. As is the case for boron, seawater repre- sents one extreme of the spectrum of isotopic compositions in the Earth (Figure 11.15). During mineral-water reactions, the heavier isotope, 7Li, is preferential partitioned into the solution. Weathering on the continents results in river water being isotopically heavy, +23.5‰ on average and the sus-

Figure 11.15. Li isotopic composition of terrestrial and meteoritic materials. Diamond represent average val- ues.

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pended load of rivers which have δ7Li ≈ +2 (Teng et al., 2004). The riverine flux is slightly smaller than the hydrothermal one, which has an average δ7Li of +8‰ and the total inputs to seawater are on aver- age about +15‰. Thus seawater is some 16 per mil heavier than average river water, so additional frac- tionation must occur marine environment. This includes adsorption on particles (although Li is less prone to absorption than most other metals), authigenic clay formation, and low temperature alteration of oceanic crust. Misra and Froelich (2012) estimate that authigenic clay formation accounts for about 70% of Li removal, with alteration of the oceanic crust accounting for the remainder and that the net fractionation factor for these processes is -16‰, consistent with other observations and experiments. δ7Li in shale range from -3 to +5‰. Marine carbonate sediments, which tend to be Li-poor, typically have higher δ7Li than non-carbonate sediment. Misra and Froehlich (2012) compiled lithium isotopic analyses of Cenozoic fora- minifera and have shown the δ7Li of seawater has in- increased by 9‰ over this period. Although not identical, it tracks the rise in 87Sr/86Sr and 187Os/188Os remarkably well (Figure 11.16). Misra and Froehlich attribute the change in all three ratios to changes in weathering of the continents due primarily to changes in tectonism over the Cenozoic, which has seen the rise of the Rocky Mountains, the Andes, the Himalayas, and the Alps. As they point out, low-lying terrains where removal of weathering products is transport- limited, especially those in the tropics, undergo congru- ent weathering Li isotope ra- tios that reflect their bed- rocks. Mountainous terrains with weathering undergo high weathering and denu- dation rates with incongru- ent weathering, and large rates of secondary clay for- mation and hence dissolved Li is 7Li-enriched. The sig- nificance of this is the impor- tant role that silicate weath- ering plays in drawing down atmospheric CO2. They note

Figure 11.16. δ7Li, 87Sr/856Sr, and 187Os/188Os in seawater over the Ceno- zoic. From Misra and Froehlich (2012).

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