Three-isotope Method Research Articles

Iron isotope fractionation between aqueous ferrous iron and goethite

The equilibrium Fe isotope fractionation factor between aqueous Fe(II) and goethite has been experimentally measured to be −1.05±0.08‰ in 56Fe/54Fe (2σ) at 22°C, using the three-isotope method. Experiments were done using two sizes of goethite (81×11nm and 590×42nm), and the experimental products were subjected to serial extraction using acid partial dissolution techniques to determine if surface Fe(III) atoms have different isotopic properties than the bulk goethite. These experiments indicate that the interaction of Fe(II)aq and goethite is dynamic and results in complete or near-complete Fe isotope exchange over 30days, involving at least four components: Fe(II)aq, goethite, sorbed Fe(II), and Fe(III)surface. The equilibrium fractionation factor between Fe(II)aq and Fe(II)sorb is the same for both sizes of goethite, at Δ56FeFe(II)aq–Fe(II)sorb=−1.24±0.14‰; this fractionation factor is significantly different than the results of previous studies on Fe(II) sorption to goethite. The proportion of the Fe(III)surface component is greatest in the experiments that used the smallest goethite, and the Fe(III)surface–Fe(II)aq fractionation is estimated to be at least +2.1‰. The high Fe(III)surface–Fe(II)aq fractionation may exert a significant influence on the Fe isotope compositions of aqueous Fe(II) in natural systems that contain nanoparticulate goethite, including those involving bacterial iron reduction. These results demonstrate that the isotopic properties of nano-scale minerals may be distinct from micron-scale or larger minerals, as is the case for other thermodynamic properties of nanoparticles.

Equilibrium high-temperature Fe isotope fractionation between fayalite and magnetite: An experimental calibration

The iron isotopic fractionation factor between magnetite and fayalite was measured at high temperature in a piston–cylinder apparatus. For the first time, the three-isotope method [Matsuhisa, J., Goldsmith, J. R., and Clayton, R. N., 1978. Mechanisms of hydrothermal crystallisation of quartz at 250 °C and 15 kbar. Geochim. Cosmochim. Acta 42, 173–182.] has been used to determine equilibrium fractionation between two minerals directly. The experimentally-determined temperature-dependent iron isotope fractionation between fayalite and magnetite is described by 10 3ln α 57 Mag–Fa ≅ Δ 57Fe Mag–Fa = 0.30 (+/− 0.024) × 10 6 / T 2. The equation predicts measurable fractionation at magmatic temperatures. Our work bears on the magnitude of Fe isotope fractionation attending differentiation in magmatic systems and provides a new isotope thermometer for co-existing fayalite and magnetite.

Measurement of the radii of gyration of ribosomal components in situ by neutron scattering

The radii of gyration of 16S rRNA, 30S proteins and $4 have been measured in situ in the 30S ribosomal subunit from Escherichia coli by the three-isotope method (TIM) [Pavlov & Serdyuk (1987). J. Appl. Cryst. 20, 105-110]. The values obtained were 70.0 (25), 66.7 (24) and 23.5 (10),~, respectively. The data for ribosomal protein and RNA in this system are consistent with earlier findings, verifying that the TIM is reliable. The estimate obtained for $4 is one of the first high-quality estimates of the radius of gyration obtained for a ribosomal protein in situ.

18O/ 16O ratio determination of framework oxygen of apophyllite and wairakite by the preferential isotopic exchange of their water of crystallization

A new analytical method has been developed for the determination of the oxygen isotopic compositions of the Si 8O 20 framework of apophyllite and the Al 2Si 4O 12 framework of wairakite. The method consists of the preferential oxygen isotopic exchange of the water of crystallization with isotopically different external water, followed by the determination of the bulk isotopic composition of the minerals. The δ 180-value of the silicate framework can be obtained by subtracting the δ 18-O-value of the water of crystallization from that of the bulk mineral in proportion to the oxygen atomic fraction of the water of crystallization. Optimum conditions of the exchange for apophyllite have been found to be ∼ 160° C for ∼ 10 days during which the water of crystallization is almost completely exchanged with the external water, whereas the framework oxygen remains unchanged. For wairakite analysis, a graphical method utilizing a three-isotope plot ( 17O/ 16 vs. 18O/ 160) has been used. The advantage of the three-isotope method is that the δ 18O-value of the silicate framework can be graphically estimated with an accuracy better than ± 2‰, even if the complete exchange of external and crystallization waters is not attained.

Oxygen isotope fractionation between zoisite and water

Oxygen isotope fractionations between zoisite and water have been studied at 400–700°C, PH 2 O = 13.4 kbar, using the three-isotope method described by Matsuhisa et al. (1978) and Matthews et al. (1983a). The zoisite-waier exchange reaction takes place extremely slowly and consequently direct-exchange calibration of equilibrium 18O 16O fractionation factors was possible only at 600 and 700°C. Fractionation factors at 400–600°C were determined from samples hydrothermally crystallized from a glass of the anhydrous zoisite composition. At 600°C, both exchange procedures gave identical fractionations within experimental error. Scanning electron microscope studies showed that the zoisite-water exchange reaction occurs largely by solution-precipitation mass-transfer mechanisms. The slow kinetics of zoisite-water exchange may be typical of hydrous silicates, since additional experiments on tremolite-water and chlorite-water exchange also showed very low rates. When the zoisite-water fractionation factors determined in this study are combined with the quartz and albite-water data of Matsuhisa et al. (1979) and the calcite-water data of O'Nell et al. (1969), mineral-pair fractionations are obtained for which the coefficients “ A” in the equation 1000 In α = A × 10 6 T −2 are: Ab Cc Zo Q 0.50 0.50 1.56 Ab 0.00 1.06 Cc 1.06 Fractionation data on coexisting quartz-zoisite (epidote) ± calcite in metamorphic rocks yield temperatures which do not appear to represent peak metamorphic conditions and possibly reflect the complexity of the chemistry of the epidotes and/or the complexity of zoisite and epidote parageneses.

Oxygen isotope fractionations involving pyroxenes: The calibration of mineral-pair geothermometers

Oxygen isotope fractionations between wollastonite, diopside, jadeite, hedenbergite and water have been experimentally studied at high pressures (1<- PH 2O ≥ 24 kbar) and temperatures (400/ dg ≤ T <- 800/ dgC) using the three-isotope method (Matsuhisa et al., 1978). Initial 18 O 16 O fractionations were made close to equilibrium and initial 17O 16O ratios were well removed from equilibrium, allowing accurate determinations of the equilibrium 18 O 16 O fractionations and of the extent of isotopic exchange. Scanning electron microscope and rate studies show that the wollastonite-water and diopside-water exchange reactions occur largely by solution-precipitation (Ostwald Ripening) mechanisms. Equilibrium 18 O 16 O fractionations between water and the minerals wollastonite, diopside, and hedenbergite are in close agreement with one another, whereas significantly more positive fractionations are found for jadeite-water. These isotopic substitution effects can be ascribed to replacement of SiOM bonds (M is a divalent metal cation in octahedral coordination) by higher frequency SiOAl bonds. The fractionations determined in this study can be combined with quartz- and feldspar-water data of Matsuhisa et al. (1979) and revised magnetite-water data of O'NEIL (1963), to provide a coherent set of mineral-pair fractionations satisfactorily represented by straight lines through the origin on a conventional graph of In /ga versus T −2 . Mineral-water data, on the other hand, cannot readily be fitted to the simple relationship suggested by Bottinga and Javoy (1973). Coefficients “ A” for the mineral-pair fractionations 1000 ln α = A × 10 6 T −2 are: Ab Jd An Di Wo Mt Q 0.50 1.09 1.59 2.08 2.20 6.11 Ab 0.59 1.09 1.58 1.70 5.61 Jd 0.50 0.99 1.11 5.02 An 0.49 0.61 4.52 Di 0.12 4.03 Wo 3.91 Geologically reasonable temperatures are found for appropriate natural assemblages.