Martin Dietzel

Model for kinetic effects on calcium isotope fractionation (δ44Ca) in inorganic aragonite and cultured planktonic foraminifera

The calcium isotope ratios (δ44Ca = [(44Ca/40Ca)sample/(44Ca/40Ca)standard −1] · 1000) of Orbulina universa and of inorganically precipitated aragonite are positively correlated to temperature. The slopes of 0.019 and 0.015‰ °C−1, respectively, are a factor of 13 and 16 times smaller than the previously determined fractionation from a second foraminifera, Globigerinoides sacculifer, having a slope of about 0.24‰ °C−1. The observation that δ44Ca is positively correlated to temperature is opposite in sign to the oxygen isotopic fractionation (δ18O) in calcium carbonate (CaCO3). These observations are explained by a model which considers that Ca2+-ions forming ionic bonds are affected by kinetic fractionation only, whereas covalently bound atoms like oxygen are affected by kinetic and equilibrium fractionation. From thermodynamic consideration of kinetic isotope fractionation, it can be shown that the slope of the enrichment factor α(T) is mass-dependent. However, for O. universa and the inorganic precipitates, the calculated mass of about 520 ± 60 and 640 ± 70 amu (atomic mass units) is not compatible with the expected ion mass for 40Ca and 44Ca. To reconcile this discrepancy, we propose that Ca diffusion and δ44Ca isotope fractionation at liquid/solid transitions involves Ca2+-aquocomplexes (Ca[H2O]n2+ · mH2O) rather than pure Ca2+-ion diffusion. From our measurements we calculate that such a hypothesized Ca2+-aquocomplex correlates to a hydration number of up to 25 water molecules (490 amu). For O. universa we propose that their biologically mediated Ca isotope fractionation resembles fractionation during inorganic precipitation of CaCO3 in seawater. To explain the different Ca isotope fractionation in O. universa and in G. sacculifer, we suggest that the latter species actively dehydrates the Ca2+-aquocomplex before calcification takes place. The very different temperature response of Ca isotopes in the two species suggests that the use of δ44Ca as a temperature proxy will require careful study of species effects.

Dissolution of silicates and the stability of polysilicic acid

Abstract In several experiments albite, orthoclase, diopside, muscovite, sepiolite, magadiite, kenyaite, natrosilite, δ-Na 2 Si 2 O 5, sodium metasilicate, kanemite, quartz, opal, amorphous silica, and silica gel were dissolved in sulfuric aqueous solutions at pH 3. The experimental results show that most of the solutions contain silicic acid as both polymeric and monomeric species. Polysilicic acid is measured as high- and as low-molecular-weight silica. The polymers may amount to ∼50 mol.% of total dissolved silica. The proportion of polymers varies with reaction time of the dissolution experiment. As a function of time, polymers decompose into monomeric species, which is the stable silica species at experimental conditions. Therefore, the transfer of polysilicic acid into the solution represents a transition stage during the dissolution of silicates. The pH, the temperature, and the kind of the dissolved components of most natural waters provide high depolymerisation capacities. Polysilicic acid decomposes to monomer within a few hours or days in river and seawater. Therefore, it is not surprising that in most natural waters silicic acid consists of monomeric species. However, in natural environments with acid solutions rich in bivalent cations, polysilicic acid may exist as a metastable component over several months.

THERMODYNAMICS OF THE SOLAR CORONA AND EVOLUTION OF THE SOLAR MAGNETIC FIELD AS INFERRED FROM THE TOTAL SOLAR ECLIPSE OBSERVATIONS OF 2010 JULY 11

We report on the first multi-wavelength coronal observations, taken simultaneously in white light, Hα 656.3 nm, Feix 435.9 nm, Fex 637.4 nm, Fexi 789.2 nm, Fexiii 1074.7 nm, Fexiv 530.3 nm, and Nixv 670.2 nm, during the total solar eclipse of 2010 July 11 from the atoll of Tatakoto in French Polynesia. The data enabled temperature differentiations as low as 0.2 × 10 6 K. The first-ever images of the corona in Feix and Nixv showed that there was very little plasma below 5 × 10 5 K and above 2.5 × 10 6 K. The suite of multi-wavelength observations also showed that open field lines have an electron temperature near 1×10 6 K, while the hottest, 2×10 6 K, plasma resides in intricate loops forming the bulges of streamers, also known as cavities, as discovered in our previous eclipse observations. The eclipse images also revealed unusual coronal structures, in the form of ripples and streaks, produced by the passage of coronal mass ejections and eruptive prominences prior to totality, which could be identified with distinct temperatures for the first time. These trails were most prominent at 10 6 K. Simultaneous Fex 17.4 nm observations from Proba2/SWAP provided the first opportunity to compare Fex emission at 637.4 nm with its extreme-ultraviolet (EUV) counterpart. This comparison demonstrated the unique diagnostic capabilities of the coronal forbidden lines for exploring the evolution of the coronal magnetic field and the thermodynamics of the coronal plasma, in comparison with their EUV counterparts in the distance range of 1–3 R� . These diagnostics are currently missing from present space-borne and ground-based observatories.

Heterotrophic prokaryotic production in ultraoligotrophic alpine karst aquifers and ecological implications.

Spring waters from alpine karst aquifers are important drinking water resources. To investigate in situ heterotrophic prokaryotic production and its controlling factors, two different alpine karst springs were studied over two annual cycles. Heterotrophic production in spring water, as determined by [(3)H]leucine incorporation, was extremely low ranging from 0.06 to 6.83 pmol C L(-1) h(-1) (DKAS1, dolomitic-karst-spring) and from 0.50 to 75.6 pmol C L(-1) h(-1) (LKAS2, limestone-karst-spring). Microautoradiography combined with catalyzed reporter deposition-FISH showed that only about 7% of the picoplankton community took up [(3)H]leucine, resulting in generation times of 3-684 days. Principal component analysis, applying hydrological, chemical and biological parameters demonstrated that planktonic heterotrophic production in LKAS2 was governed by the respective hydrological conditions, whereas variations in DKAS1 changed seemingly independent from discharge. Measurements in sediments recovered from LKAS2, DKAS1 and similar alpine karst aquifers (n=12) revealed a 10(6)-fold higher heterotrophic production (average 19 micromol C dm(-3) h(-1)) with significantly lower generation times as compared with the planktonic fraction, highlighting the potential of surface-associated communities to add to self-purification processes. Estimates of the microbially mediated CO(2) in this compartment indicated a possible contribution to karstification.

Chemical and 13C/12C- and 18O/16O-isotope evolution of alkaline drainage waters and the precipitation of calcite

Carbon and O isotope ratios demonstrate that calcite precipitating from alkaline tunnel drainage waters has formed by two distinct chemical mechanisms. Calcium-carbonate-sulfate ground water dissolves portlandite preferentially. This leads to an increase of Ca2, pH and the CO32-/HCO3− ratio so that calcite is precipitated. The C content of this calcite is a mixture of C from soil-CO2 and the dissolved limestone resulting in δ13C ≈ − 13‰ (PDB). The O isotope composition (δ18O ≈ 24‰, (SMOW) corresponds to thequilibrium fractionation between calcite and water. The precipitation continues as atmospheric CO2 is absorbed by the alkaline solutions. In this type of calcite the C isotope composition (δ13C ≈ −25‰, PDB) is controlled by the individual reaction rate constants for the hydroxylation of 13CO2 and 12CO2. The O isotope composition (δ18O ≈ 11‰, SMOW) is a mixture of the compositions of atmospheric CO2 and OH−. The isotope compositions may be used in order to evaluate the proportions of the two different mechanisms of calcite precipitation. Other mechanisms which cause chemical changes of the ground water are the formation of ettringite from initially unhydrated C3A-phases and the formation of brucite. Precipitation and dissolution rates for the various solids range from 10−3 to 10−4 kg/d/m2.

Depolymerization of soluble silicate in dilute aqueous solutions

Experiments with diluted solutions of a customary water glas show that the rate of depolymerization depends not only upon pH and the SiO2 concentration, but also varies systematically as a function of the type and concentration of an additional electrolyte. Increasing cation activities of metal chlorides are causing a decrease of the rate constant in the order 1) Na+, K+, 2) Mn2+, Mg2+, Ca2+, Sr2+, 3) Zn2+, Ni2+, Ce3+, Cu2+. With respect to anions of sodium salts the rate constants are increasing with increasing activities in the order NO3−, HCO3−, Cl−, SO42−, whereas HPO42− causes a decrease. The results permit to identify those components of water which are most responsible for a change of the depolymerization rate and may be used to evalute the properties of a water glass as a possible anticorrosive agent for water supply systems.

Interaction of polysilicic and monosilicic acid with mineral surfaces

Interaction of polysilicic and monosilicic acid was studied via adsorption experiments with lepidocrocite, hematite, feroxyhyte, goethite, akaganeite, magnetite, ferrihydrite, and gibbsite. The kinetics of monosilicic acid adsorption follows a first order reaction. At equilibrium monosilicic acid adsorption may be described by surface complexation with an adsorption maximum at pH 9.8. If polysilicic acid is adsorbed to the surface, one part is bound to the surface within a relatively short time. The other part decomposes to monomer in the solution. The polymeric silica at the surface is stabilised at pH < 6. Thus the present results show that polymerization of silica at the mineral surface has to be considered only in acidic solutions.

THE RATE AND MECHANISM OF DEEP-SEA GLAUCONITE FORMATION AT THE IVORY COAST–GHANA MARGINAL RIDGE

The environmental conditions and reaction paths of shallow-water glauconitization (<500 m water depth, ~15°C) close to the sediment-seawater interface are generally considered to be well understood. In contrast, the key factors controlling deep-sea glauconite formation are still poorly constrained. In the present study, green grains formed in the recent deep-sea environment of the ODP Site 959, Ivory Coast-Ghana Marginal Ridge, (~2100 m water depth, 3-6°C) were investigated by X-ray diffraction and electron microscopic methods in order to determine the rate and mechanism of glauconitization.Green clay authigenesis at Hole 959C occurred mainly in the tests of calcareous foraminifera which provided post-depositional conditions ideal for glauconitization. Within this organic-rich microenvironment, Fe-smectite developed <10 ky after deposition of the sediments by precipitation from precursor gels containing Fe, Mg, Al, and silica. This gel formation was supported by microbial activity and cation supply from the interstitial solution by diffusion. At a later stage of early marine diagenesis (900 ky), the Fe-smectites reacted to form mixed-layer glauconite-smectite. Further down (~2500 ky), almost pure glauconite with no compositional gaps between the Fe-smectite and glauconite end members formed. This burial-related Fe-smectite-to-glauconite reaction indicates that the glauconitization process was controlled mainly by the chemistry of the interstitial solutions. The composition of the interstitial solution depends heavily on micro-environmental changes related to early diagenetic oxidation of biodegradable (marine) organic matter, microbial sulfate reduction, silicate mineral alteration, carbonate dissolution, and Fe redox reactions. The availability of Fe is suggested as the probable limiting factor for glauconitization, explaining the various states of green-grain maturity within the samples, and this cation may be the most important rate-determining element.The rate of glauconite formation at ODP Site 959 is given by %GlSed = 22.6·log(ageSed) + 1.6 (R2 = 0.97) where %GlSed is the state of glauconitization in the sediment and ageSed is the sediment age (in ky). This glauconitization rate depends mainly on continuous cation supply (in particular Fe) and is about five times less than that in shallow-shelf regions, suggesting significantly slower reaction at the lower temperature of deep-sea environments.

The Fe-Mg-saponite solid solution series - A hydrothermal synthesis study

Abstract The boundary conditions of saponite formation are generally considered to be well known, but significant gaps in our knowledge persist in respect to the influence of solution chemistry, temperature, and reaction time on the mineralogy, structure, stability, and chemical composition of laboratory-grown ferrous saponite. In the present study, ferrous saponite and Mgsaponite were synthesized in Teflon-lined, stainless steel autoclaves at 60, 120 and 180°C, alkaline pH, reducing conditions, and initial solutions with molar Si:Fe:Mg ratios of 4:0:2, 4:1:1, 4:1.5:0.5, 4:1.75:0.25, and 4:1.82:0.18. The experimental solutions were prepared by dissolution of sodium orthosilicate (Na4SiO4), iron(II)sulfate (FeSO4·6H2O) and magnesium chloride salts (MgCl2·6H2O with 40.005 mass% of K and Ca) in 50 mL ultrapure water that contained 0.05% sodium dithionite as the reducing agent. The precipitates obtained at two, five and seven days of reaction time were investigated by X-ray diffraction techniques, transmission electron microscopy analysis, infra-red spectroscopy, and thermo-analytical methods. The precipitates were composed mainly of trioctahedral ferrous saponite, with small admixtures of co-precipitated brucite, opal-CT, and 2-line ferrihydrite, and nontronite as the probable alteration product of ferrous saponite. The compositions of the obtained ferrous saponites were highly variable, (Na0.44-0.59K0.00-0.05Ca0.00-0.02) (Fe2+0.37-2.41Mg0.24-2.44Fe3+0.00-0.28 )S2.65-2.85 [(Fe3+0.00-0.37Si3.63-4.00)O10](OH)2, but show similarities with naturally occurring trioctahedral Fe and Mg end members, except for the Al content. This suggests that a complete solid solution may exist in the Fe-Mg-saponite series. A conceptual reaction sequence for the formation of ferrous saponite is developed based on the experimental solution and solid compositions. Initially, at pH ≥ 10.4, brucite-type octahedral template sheets are formed, where dissolved Si-O tetrahedra are condensed. Subsequent reorganization of the octahedra and tetrahedra via multiple dissolution-precipitation processes finally results in the formation of saponite structures, together with brucite and partly amorphous silica. The extent of Fe2+ incorporation in the octahedral template sheets via isomorphic substitution is suggested to stabilize the saponite structure, explaining (i) the abundance of saponite enriched in VIFe2+ at elevated Fe supply and (ii) the effect of structural Fe on controlling the net formation rates of ferrous saponite.

Atlas and data of solid-solution equilibria of marine evaporites

I Introduction.- I.1 Purpose of the Volume.- I.2 Use of the Volume.- I.2.1 Derivation of Phase Diagrams from Solubility Data.- I.2.2 Data and Diagrams for Solid-Solution Equilibria.- I.2.3 The Computer Program SALSYS.- I.3 Handling of the Volume.- I.3.1 General Procedure.- I.3.2 Examples.- I.3.2.1 Binary Systems.- I.3.2.2 Ternary Systems.- I.3.2.3 Quaternary Systems.- I.3.2.4 The Quinary System.- I.4 Synopsis of Marine Evaporites and Solubility Data.- I.4.1 Salt Deposits.- I.4.2 Significance of Solubility Data.- I.4.3 Data Bases.- I.4.4 Individual Articles.- I.4.5 Comments.- I.4.5.1 Binary Systems.- I.4.5.2 Ternary Systems.- I.4.5.3 Quaternary Systems.- I.4.5.4 The Quinary System.- II Derivation of Phase Diagrams from Solubility Data.- II. 1 The Phase Rule.- II.2 The Binary System AX-H2O.- II.3 The Ternary System AX-BX-H2O.- II.3.1 Phase Diagrams.- II.3.2 Isothermal Evaporation and Dissolution.- II.3.3 Formation of a Compound.- II.4 Quaternary Systems.- II.4.1 Systems with a Common Ion.- II.4.2 Reciprocal Salt Pairs.- II.4.2.1 Phase Diagrams.- II.4.2.2 Concentrations.- II.4.2.3 Polytherms.- II.5 The Quinary System AX-BX-CX-AY-BY-CY-H2O.- II.5.1 Isotherms.- II.5.2 Concentrations.- II.5.3 Projections of Isotherms.- II.5.4 Polytherms.- II.6 Numerical Examples.- II.6.1 System 2.1 Na2Cl2-H2O.- II.6.1.1 Isothermal Evaporation.- II.6.1.2 Dissolution and Precipitation.- II.6.2 System 2.6 MgSO4-H2O.- II.6.2.1 Dissolution of Kieserite.- II.6.2.2 Precipitation of Hexahydrite.- II.6.3 System 3.1 Na2Cl2-K2Cl2-H2O.- II.6.3.1 Projections.- II.6.3.2 Isothermal Evaporation.- II.6.3.3 Polythermal Precipitation of Sylvite.- II.6.4 System 3.3 K2Cl2-MgCl2-H2O.- II.6.4.1 Isothermal Evaporation.- II.6.4.2 Dissolution of Carnallite.- II.6.5 System 4.1 Na2Cl2-K2Cl2-MgCl2-H2O.- II.6.5.1 Calculation of Ion Fractions from Concentrations.- II.6.6 System 4.2 Na2SO4-K2SO4-MgSO4-H2O.- II.6.6.1 Isothermal Evaporation.- II.6.6.2 Dissolution of Glaserite and Picromerite.- II.6.7 System 4.3 Na2Cl2-K2Cl2-Na2SO4-K2SO4-H2O.- II.6.7.1 Calculation of Ion Fractions from Concentrations.- II.6.7.2 Combinations of Concentrations.- II.6.7.3 Isothermal Evaporation.- II.6.8 System 4.4 Na2Cl2-MgCl2-Na2SO4-MgSO4-H2O.- II.6.8.1 Isothermal Evaporation.- II.6.8.2 Generation of a Solution from Hydrated Solids.- II.6.9 System 5 Na2Cl2-K2Cl2-MgCl2-Na2SO4-K2SO4-MgSO4-H2O.- II.6.9.1 Calculation of Ion Fractions from Concentrations.- II.6.9.2 Crystallization and Transition Lines.- II.6.9.3 Evaporation of Solution W.- II.6.9.4 Evaporation of Solution M.- II.6.9.5 Incongruent Dissolution of Bloedite and Carnallite.- II.6.9.6 Generation of a Solution from Hydrated Solids.- III Data and Diagrams for Solid-Solution Equilibria.- III.1 Compositions and Concentrations.- III. 1.1 Minerals and Seawater.- III. 1.2 Concentration Units.- III.2 Binary Systems.- III.2.1 Na2Cl2-H2O.- III.2.2 K2Cl2-H2O.- III.2.3 MgCl2-H2O.- III.2.4 Na2SO4-H2O.- III.2.5 K2SO4-H2O.- III.2.6 MgSO4-H2O.- III.3 Ternary Systems.- III.3.1 Na2Cl2-K2Cl2-H2O.- III.3.2 Na2Cl2-MgCl2-H2O.- III.3.3 K2Cl2-MgCl2-H2O.- III.3.4 Na2SO4-K2SO4-H2O.- III.3.5 Na2SO4-MgSO4-H2O.- III.3.6 K2S04-MgSO4-H2O.- III.3.7 Na2Cl2-Na2SO4-H2O.- III.3.8 K2Cl2-K2SO4-H2O.- III.3.9 MgCl2-MgSO4-H2O.- III.4 Quaternary Systems.- III.4.1 Na2Cl2-K2Cl2-MgCl2-H2O.- IIL4.2 Na2SO4-K2SO4-MgSO4-H2O.- III.4.3 Na2Cl2-K2Cl2-Na2SO4-K2SO4-H2O.- III.4.4 Na2Cl2-MgCl2-Na2SO4-MgSO4-H2O.- III.4.5 K2Cl2-MgCl2-K2SO4-MgSO4-H2O.- III.5 Quinary System Na2Cl2-K2Cl2-MgCl2-Na2SO4-K2SO4-MgS04-H2O.- IV Subject Index.