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In the early fifties, before Richard Feynman famously seeded the concept of nanoscience in his 1959 talk “there’s plenty of room at the bottom” [1], and well before the concept of nanotechnology became popular in the late 80’s, a significant research effort was already underway into the fundamental nanoscience associated with high-field effects at surfaces and the resulting emission of ions and electrons [2]. Born from this work, in 1955, field ion microscopy (FIM) became the first true atomic scale microscopy technique, allowing us to ‘see atoms’ for the very first time [3]. The technique, invented by Erwin Müller in 1951 employed a specimen shaped into a sharp point, enabling it to act as a point projection field ion emitter. The specimen was cooled to 78K in the presence of He gas. This gas was adsorbed and subsequently field ionized and detected, with the distribution of detected atoms showing the arrangement of the specimen atoms at the surface of the tip. Sixty years on, this seminal work by Erwin Müller has spurred important and wide-ranging research, including many significant discoveries and inventions [4]. Progressive field evaporated of surface atoms can be detected [5] and their positions reconstructed to create high-resolution 3D atom maps in a technique known as atom probe tomography [6], which has become an established microscopy technique. It’s use in materials characterisation has led to ground-breaking research including the first 3D images of segregation to dislocations [7], understanding the growth of nanowires [8], determining the kinetics of elemental steps of catalytic surface reactions [9], revealing precipitation pathways in important engineering alloys [10] and confirmation of the dating of the oldest minerals on earth [11], to name just a few examples. Other contributions from field-emission science include the development of the liquid metal ion source that now forms the basis of focused ion beam instruments [12], field electron emission from new forms of emitter [13] along with the sustained development of theory around high-field effects at surfaces [14]. It is timely that we recognize these exceptional contributions. The International Field Emission Society (IFES) originally grew from pioneering research on high-field nanoscience, and supports the development and application of techniques and instruments based on these effects. It has hosted symposia since 1952 occurring every one to two years. In 2016, this conference, “Atom Probe Tomography & Microscopy (55th IFES)” will be held in Gyeongju, South Korea (June 12-17). At the event, the Steering Committee of the IFES (see note at end of this article for a list of members) is proud to award an inaugural round of “Fellows of the International Field Emission Society”, elected in recognition of eminence in the field of field emission, field ionization, and related phenomena. These individuals have been nominated and elected by their peers for outstanding research that has pushed the frontiers of knowledge in the field. Many have also undertaken distinguished service to the IFES. Those to be honored as IFES fellows in 2016 are listed below: Hans-Olof Andren, Chalmers University of Technology:  For development of atom probe techniques, and for his use of atom probe instruments as materials science tools to study the detailed microstructure of primarily metallic materials. Didier Blavette, Université de Rouen:  For unique contributions to atom probe field ion microscopy spanning the fundamental physics of the technique, instrumentation, and cutting-edge materials characterization. Alfred Cerezo, University of Oxford:  For development of the position sensitive atom probe, which opened new dimensions and perspectives in both material science and instrumentation. Paul Cutler, The Pennsylvania State University:  For working on theory of field electron and ion emission over more than 50 years, developing quantum mechanical models to explain and predict the behavior of field electron emitters. Richard Forbes, University of Surrey:  For his many contributions to the growth of the theory and understanding of field electron and ion emission as well as his contributions to the society. Georgiy Fursey, St Petersburg University of Telecommunications:  For wide-ranging, outstanding contributions to field electron emission science and technology, particularly explosive emission and emission from semiconductors. Robert Gomer, University of Chicago:  For outstanding contributions to science, especially areas of field electron and ion emission and their application to problems in surface chemistry, and for public service. Kazuhiro Hono, National Institute for Materials Science:  For key contributions to the growth of atom probe, developments in instrumentation, and broad utilization of the technique to impact the study of magnetic materials and precipitation hardening. Gary Kellogg, Retired:  For fundamental technical contributions to laser-pulsed atom probe instrumentation and numerous aspects of surface and materials science, and for extraordinary service to the nanoscience community. Thomas Kelly, Cameca Inc.:  For revolutionizing atom probe technology with the invention of the LEAP, and for service to the IFES community as President of the society. Hans-Juergen Kreuzer, Dalhousie University:  Published more than 325 papers, 8 books, and 6 patents in the area of physics and chemistry of high electric fields. Norbert Kruse, Washington State University:  For sustained contributions towards understanding chemical physics at materials surfaces and outstanding service to the high field nanoscience and atom probe communities. Allan Melmed, Retired:  One of the most distinguished scientists of the IFES community, with a lifetime experience in field emission since his PhD thesis with the late EW Müller. Michael Miller, Retired:  For seminal contributions in the development and application of atom probe tomography as demonstrated by his 600+ publications, service to the community, and impactful collaborations with numerous international scientists and engineers in their development and use of atom probe tomography. Marwan Mousa, Mu'tah University:  For outstanding contributions to field emission science and for service to the society including organization of the 45th IFES. Osamu Nishikawa, Kanazawa Institute of Technology:  For outstanding contributions to atom probe becoming a mainstream scientific instrument in hundreds of laboratories around the world. John Panitz, University New Mexico:  As one of the inventors of the atom probe technique, John Panitz’ contributions and vision for the technique enabled its large acceptance in the international realm of materials characterization. Simon Ringer, The University of Sydney:  For outstanding research in atom probe science, sustained IFES community service, including as Vice President and conference organiser and his role in training a new generation of field emission scientists. Guido Schmitz, University of Stuttgart:  For his contribution to understanding diffusion and other atomic scale metallurgical processes studied using atom probe tomography. David Seidman, Northwestern University:  Having advised more than 120 individuals and with 450+ publications, David Seidman's materials research based on APT and technique developments has laid a solid groundwork for atom probe groups worldwide. George Smith, University of Oxford:  For more than 45 years of contributions and commitment to the field of atom probe field ion microscopy. Krystyna Stiller, Chalmers University of Technology:  For fruitful use and development of atom probe techniques contributing to understanding of radiation damage, phase transformations, interfacial segregation and high temperature oxidation, and for promoting atom probe techniques. Lynwood Swanson, FEI:  For outstanding scientific contributions to characterisation and development of field electron/ion emitters, and technical and managerial leadership of FEI Company in commercially developing these emitters and related instruments. Tien Tsong, Academia Sinica:  For observation of the interaction between adsorbates on metal surfaces and for seminal research involving the use of a laser to promote thermal field evaporation. The Steering Committee of the IFES currently consists of: [1] Feynman RP. There's Plenty of Room at the Bottom. Engineering and Science 1960:22-36. [5] Cerezo A, Godfrey TJ, Smith GDW. Application of a position-sensitive detector to atom probe microanalysis. Review of Scientific Instruments 1988;59:862-6. [8] Perea DE, Hemesath ER, Schwalbach EJ, Lensch-Falk JL, Voorhees PW, Lauhon LJ. Direct measurement of dopant distribution in an individual vapour-liquid-solid nanowire. Nature Nanotechnology 2009;4:315-9. [9] Kruse N, Abend G, Block JH. The kinetics of adsorption and thermal desorption of NO on stepped Pt single crystal surfaces. The Journal of Chemical Physics 1988;88:1307-12. [10] Ringer SP, Hono K. Microstructural evolution and age hardening in aluminium alloys: atom probe field-ion microscopy and transmission electron microscopy studies. Materials Characterization 2000;44:101-31. [11] Valley JW, Cavosie AJ, Ushikubo T, Reinhard DA, Lawrence DF, Larson DJ, et al. Hadean age for a post-magma-ocean zircon confirmed by atom-probe tomography. Nature Geoscience 2014;7:219-23. [13] Li Z, Xu N, Kreuzer HJ. Coherent field emission image of graphene predicted with a microscopic theory. Physical Review B - Condensed Matter and Materials Physics 2012;85. [14] Forbes RG, Edgcombe CJ, Valdrè U. Some comments on models for field enhancement. Ultramicroscopy 2003;95:57-65.

News Article | August 22, 2016

The thermodynamic process of lithophile element incorporation in iron involves the solubility of mantle components in the metal phase (equation (1)), rather than redox exchange as in the case of siderophile element partitioning. The magnesium concentration in the metal ranges between 0.2 mol% and 1 mol% in our experiments. The equilibrium constant of the dissolution reaction given in equation (1) (reprinted here for convenience): (where X is the mole fraction of Mg in the metal and X the mole fraction of MgO in the silicate) and its logarithm is proportional to the change in Gibbs free energy of the reaction defined by equation (1): where the parameters a, b and c correspond to the changes in entropy, enthalpy and volume in the reaction in equation (1), respectively. These parameters were fitted to the data using linear regression, and c was found to be statistically irrelevant (no pressure dependence), yielding equation (2) (reprinted here for convenience): Similarly, the aluminium concentration of in the metal ranges from 0 mol% (below the detection limit, explaining two fewer points for the Al plot in Extended Data Fig. 2) to 1.1 mol%. The equilibrium constant of the dissolution reaction (where X is the mole fractions of Al in the metal and the mole fraction of AlO in the silicate) Its logarithm is proportional to the change in Gibbs free energy of the reaction in equation (6) and can be written in the same form as equation (5); fitting to the data using linear regression shows that c is once again statistically irrelevant, and we find where the numbers in parentheses are the standard errors of the parameters. Equations (2), (4), (7) and (8) allow us to calculate the Mg and Al concentrations in molten iron as a function of temperature and silicate composition. An important case is that of the equilibrium value in the core at the core–mantle boundary (CMB). As shown above, MgO dissolution in iron has no pressure dependence. This means that MgO exsolves in the coldest part of the core, which is the CMB. The equilibrium value at the CMB is therefore the MgO saturation value; if the MgO concentration in the core is above saturation, then MgO will be exsolved until it reaches that value. Figure 1b shows the equilibrium value of MgO concentration in the core as a function of CMB temperature, for a core buffered by (that is, in local equilibrium with) a pyrolitic magma ocean (50 mol% MgO in the mantle). The silicate glasses were produced in an aerodynamic levitation laser furnace. The starting mixes were made by grinding and mixing from pure oxide (SiO , MgO, Fe O , Al O ) and carbonate (CaCO ) components, pressing them into pellets, and then fusing them at constant oxygen fugacity at 1,900–2,100 °C for 5 min in a laser furnace using a 120-W CO laser. The fused samples were quenched to glasses, and analysed for recrystallization, homogeneity and composition on a Zeiss Auriga field-emission scanning electron microscope (IPGP, Paris). The glass beads were thinned down to 20-μm-thick double-parallel thin sections and were processed using a femtosecond laser machining platform to cut disks of identical size for loading in the diamond-anvil cell. Spherical iron balls 1–3 μm in size were flattened between two such silicate disks, and constituted the layered starting sample. Pressure was measured from the frequency shift of the first-order Raman mode in diamond, measured on the anvil tips. Temperature was measured every second, simultaneously from both sides, by spectroradiometry. Electronic laser shutdown operates in about 2–4 μs, and temperature quench occurs in approximately 10 μs (owing to thermal diffusion in the sample) ensuring an ultrafast quench of the sample. After decompression, a thin section (20 μm × 10 μm wide, 1–3-μm thick) was extracted from the centre of the laser-heated spot using a Zeiss Auriga crossbeam focused-ion-beam microscope (IPGP, Paris). The sample was imaged and then transferred to a TEM copper grid, and the metal and silicate phases were analysed using a Cameca SX-Five electron microprobe (CAMPARIS, Paris) with five large-area analysers. Metal and silicate phases of the run products are large enough (>5 μm) to perform reliable analyses with an electron probe micro-analyser (EPMA) on focused ion beam (FIB) thin sections. Metal and silicate phases were analysed using Cameca SX100 and Cameca SX FIVE (CamParis, UPMC–IPGP) electron probe micro-analysers. X-ray intensities were reduced using the CITZAF correction routine. Operating conditions were 15-kV accelerating voltage, and 10–20-nA beam current and counting times of 10–20 s on peak and background for major elements and 20–40 s for trace elements (including Mg and Al in the metallic phases). Pure Fe metal was used as standard for metal. Fe O , SiO , MgO and Al O were used as standards to measure solubility of oxygen, silicon, magnesium and aluminium in metal. Diopside glass (Si), wollastonite (Ca), orthoclase (K), anorthite (Al), albite (Na), rutile (Ti) and pure oxides (Fe O , MgO, SiO , CaO and Al O ) were used as standards for the silicate. We verified that the geometry of the metal and silicate phases was identical from both sides of the FIB sections, so that the EPMA analyses only a single phase. The EPMA uses a beam size of 1–2 μm, which is large enough to integrate the small quench features of metal and silicate phases (<200 nm) and to determine their bulk compositions. When a few small metallic blobs were present in the silicate (500 nm to 2 μm in diameter), special care was taken to avoid them during analysis of the silicates. The core of Earth formed in the first approximately 50 million years24, 25 of the Solar System, by an iterative addition of material to the proto-Earth. The accreting material, consisting of mixtures of iron-rich metals and silicates similar to those found in extra-terrestrial bodies (such as chondrite parent bodies, HEDs and angrites), impacted the growing planet. The heat generated by those impacts maintained the outermost portion of the planet in a molten state known as a magma ocean26. At temperatures below the solvus of iron and silicate, the two phases un-mix and the metal (twice denser) segregates towards the centre and forms the core. Along with the segregating metal, the siderophile elements are stripped to the core, among which are light elements such as Si and O. The depletion of siderophile elements from the mantle has been widely used to constrain the pressure–temperature–composition path of core formation, and has shown that the core formed in a deep magma ocean27, 28. As the planet accretes, the magma ocean grows deeper; recent models11 show that the concentrations of Ni, Co, Cr and V in the mantle satisfy terrestrial observables for a final magma ocean depth of between 1,000 km and 1,700 km, corresponding to final pressures of between 40 GPa and 75 GPa and final temperatures of between 3,000 K and 4,180 K, respectively. We ran a series of traditional multistage core-formation models11 where the planet was accreted to its present mass in increments of 0.1% of Earth’s mass, without giant impacts. At each stage, the planet grows and the pressure and temperature of equilibration increase accordingly. The concentrations of Ni, Co, V, Cr, O, Si and Mg in the core were calculated iteratively during the 1,000 steps of the accretion process. The simulations were run for a variety of redox paths (ranging from very reduced to very oxidized), several geotherms (between the solidus and the liquidus of peridotite), and for all possible magma ocean depths, ranging from 0% (magma lake) to 100% (fully molten Earth) of the mantle. We forward-propagated all uncertainties on the thermodynamic parameters governing the partitioning equations using Monte Carlo simulation. Most models (very deep or very shallow) do not satisfy, within uncertainties, the observed geochemical abundances of Ni, Co, V and Cr in the mantle and therefore are not relevant. We selected only the models that do reproduce the geochemical abundances of Ni, Co, V and Cr in the present-day mantle, and found that the maximum MgO concentration in the core at the end of accretion is 0.8 wt%. In the Moon-forming giant-impact scenario12, the impactor is typically thought of as a Mars-sized planetary embryo, but the masses used in models range from 2.5% to 20% of Earth’s mass19, 20. With such a size, the impactor is a differentiated object with a core and mantle (as opposed to small undifferentiated bodies) and, hence, it does not fully equilibrate with the entire magma ocean, but rather partially equilibrates29 with a small portion9, 21 of that magma ocean. The impactor and the magma ocean (in the impact zone) reach tremendous temperatures during the impact, as shown by smoothed-particle hydrodynamic simulations19, 20. Even though the temperatures from those simulations can be inaccurate because of intrinsic inaccuracies in the equations of state that they are based on, the minimum temperature19 for the impactor core is 8,000 K and that of the magma ocean in the impacted area is 7,000 K. Therefore, the system consisting of the impactor core and the surrounding silicate mantle is necessarily always hotter than 7,000 K, and turns into a single miscible metal-silicate phase. We calculated the composition of Earth’s core after the giant impact in two steps. First, we modelled the pre-giant-impact accretionary phase. The Earth was partially accreted, as described in the previous paragraph, until it reached 80% to 99% of Earth’s mass, leaving the planet in the state it was in before the giant impact. We considered only the models that reproduce the present-day geochemical abundances of Ni, Co, V and Cr in the mantle. Then the final accretion event took place, consisting of the giant impact bringing in the remaining 1% to 20% of Earth’s mass. We calculated the composition of the hybridized impactor core (HIC) as a function of its size (Fig. 2a) by considering the fact that, as opposed to small accretionary building blocks, the core of the giant impactor does not fully equilibrate with the entire magma ocean; instead, it partially equilibrates29 with a small portion9, 21 of the magma ocean (see Methods section ‘Partial core equilibration and turbulent fragmentation and mixing’ below). It is clear from Fig. 2a that the bigger the impactor, the smaller the relative mass of magma ocean it interacts and equilibrates with and, consequently, the less mantle components (Mg, O, Si) the HIC contains. The net effect on Earth’s core, once the HIC is added, is mitigated as shown in Fig. 2b; it is the result of the balance between larger HICs being less enriched in mantle component, but contributing more mass to the whole core. The composition of the HIC was calculated by taking into account two main parameters that are usually neglected in traditional core-formation models9, 10, 11, 28, 30. First, the degree of partial equilibration—that is, the fraction of the core that equilibrates with the mantle—has been constrained by geochemical modelling, from the combined analysis of the Hf–W and U–Pb isotopic systems, and shown to be at least25, 29, 31 40%. We used this conservative lower bound, meaning that 60% of the impactor core merges with Earth’s core without equilibration (and therefore with no compositional effect), whereas the other half equilibrates in the magma ocean before merging with the core. Second, the impactor core only ‘sees’ a portion9 of the magma ocean, with the fraction involved in the equilibration estimated from fragmentation and turbulent mixing scaling laws21; these laws show that the ratio of equilibrated silicate to equilibrated metal (dilution ratio Δ) in the magma ocean is given by where ρ and ρ are the densities of silicate and metal, and δ is the ratio of impactor to Earth mass. The energy release depends on how the HIC mixes with Earth’s core, as shown by the dependence on ρ(r) in equation (3). Even though simulations20 and energetic arguments32 suggest that the HIC should thoroughly mix with Earth’s core, we investigated two extreme models of mixing: (i) full mixing of the HIC with Earth’s core producing a homogeneous core and (ii) full layering where the HIC sits atop Earth’s core. In the mixed case, the HIC is diluted in the bulk of Earth’s core and therefore the Si and O content delivered by the impactor are below the saturation limit of those elements11, 30, 33 (Fig. 2b); those concentrations are under-saturated with respect to the overlying conditions imposed by the magma ocean at the CMB, and there is no chemical drive to force those components out of the system. In that case, we considered that MgO is the only phase to exsolve so that the associated energy release is a conservative lower bound. In the layered case, the HIC is concentrated atop the proto-core, and all three mantle components (MgO, SiO and FeO) are highly concentrated in the layer and over-saturated with respect to CMB conditions prevailing atop that layer. In that case, all of those components would exsolve and remix with the overlying magma ocean. In our energy calculations, we fixed the present-day CMB temperature to 4,100 K. Lower temperatures imply a lower saturation level in the core, and mean that more MgO exsolves and more energy is produced, and vice versa. The final density and radius of the core are the present-day values (10.6 g cm−3 and 3,485 km, respectively). We considered a uniform core of density ρ and radius R; it subsequently undergoes un-mixing into an inner (dense) region with density ρ and radius R (the present-day values given above), and an outer buoyant layer with density ρ . The volume fraction of the outer layer is f, which we take to be . We may write where equation (10) is correct to first order in f. In practice, we specify ρ and ρ (4.8 g cm−3 for MgO) and calculate ρ and R for a given value of f, with the current core boundary taken to be R . The gravitational energy E of the core in either state may be derived using equation (3), and the change in energy ΔE in going from the uniform to the unmixed state can be available to do work (for example, to drive a dynamo). Making use of equations (3), (9) and (10), it may be shown that, to first order in f: For f = 20%, equation (11) overestimates the full calculation (plotted in the figures) by about 5%; the discrepancy is smaller with smaller f, and equation (11) can be used, to a good approximation, to estimate the amount of energy released by mantle-component exsolution from the core. This equation shows the correct limiting behaviour in the cases of f = 0 and ρ  = ρ . In this case we take the mass fraction of the Earth’s core added by the HIC to be f . With a HIC density of ρ and a present-day total core mass of M , the radius of the base of the impactor layer R before un-mixing of this layer is given by The HIC layer then undergoes un-mixing into two components: ‘mantle components’ (ρ , 5.6 g cm−3) and ‘core material’ (ρ , 10.6 g cm−3). The HIC density ρ may then be derived using where R is the radius of the base of the light-element layer after un-mixing. To make the total core mass correct, the density of the pre-impact core, ρ , is also calculated. Once ρ , R and R have been calculated, the energy change due to un-mixing within the layer can be calculated using successive applications of equation (11), as before. Using a CMB temperature evolution model, we can estimate the MgO exsolution rate and, hence, an exsolution power, as a function of time. A typical CMB temperature evolution is shown in Extended Data Fig. 4a, along with the associated MgO content of the core (Extended Data Fig. 4b) obtained by rewriting the MgO equilibrium curve (Fig. 1b) as a function of time. The time derivatives are the cooling rate of the core and its MgO exsolution rate as a function of time, and are plotted in Extended Data Fig. 4c, d, respectively. Very early in Earth’s history, the core was so hot that the equilibrium MgO concentration at the CMB (Extended Data Fig. 4b) is higher than the MgO content of the core, and no exsolution occurs. The reverse reaction—that is, the potential for MgO to be dissolved from the mantle to the core—is limited; it is prone to affect only a thin layer below the CMB that is enriched in MgO, that becomes light and stably stratified, and that is therefore unable to recycle and affect the entire core. As the core cools, exsolution starts once the temperature at the CMB reaches a critical value corresponding to an MgO equilibrium concentration equal to that in the core. This is shown in Extended Data Fig. 5, and is highlighted for two models: the Mars-size impact19 leaving behind a core containing 2.9 wt% MgO and a small fast-spinning impact20 producing a core containing 2.1 wt% MgO (see Fig. 2b and Extended Data Fig. 3). The power produced by MgO exsolution is linked to the exsolution rate, and can be estimated from the energy release (Fig. 3 and Extended Data Fig. 8) to be between 5.5 TW wt% Gyr−1 and 7 TW wt% Gyr−1. This estimate allows us to translate an exsolution rate (Extended Data Fig. 4d) into exsolution power, as shown in Extended Data Figs 5b and 8. What is noteworthy is that the initial MgO core content does not directly affect exsolution power. The latter is a function of only the exsolution rate, which is itself a function of core cooling rate. Initial MgO content sets only the onset of exsolution, as shown in Extended Data Fig. 5. Of course, higher MgO contents in the core entail an earlier onset of exsolution, a longer duration for buoyancy-driven exsolution power and, hence, much higher total exsolution energies, as shown in Fig. 3. This dichotomy could be mitigated had we self-consistently included MgO exsolution in the thermal evolution model of the core. MgO exsolution power markedly drops with the onset of inner-core growth, as a consequence of the drop in core cooling rate. At the present day, MgO exsolution should still produce about 1 TW of power, much lower than the approximately 3 TW produced by inner-core growth and driving the geodynamo. However, before inner-core growth, exsolution power is always higher than about 3 TW, demonstrating that MgO exsolution can conceivably drive a geodynamo as early as around 1 Gyr after core formation, and until the onset of inner-core growth. Assuming an entirely bottom-driven present-day dynamo, corresponding to a CMB heat flow exactly at the adiabatic value (Q ) of 15 TW (refs 34,35), the convective power sustaining the geomagnetic field P = εQ is 3 TW, where ε = 0.2 is the thermodynamic efficiency of latent heat and light-element release at the inner-core boundary22. Power-based scaling laws of the magnetic intensity36 then predict an internal magnetic field of about 1–4 mT, the higher estimate being in agreement with the observation of magnetic Alfvén waves in the core37 coupled to length-of-day variations at periods close to 6 years (ref. 38). Dynamo strength increases as buoyancy flux increases39, 40, so the MgO exsolution mechanism represents a potent driver of an early geodynamo7. Although a giant impact might cause thermal stratification in the core6, 41, the stabilizing thermal buoyancy will be completely overwhelmed by the compositional buoyancy associated with MgO exsolution.

Saka S.K.,University of Gottingen | Vogts A.,Leibniz Institute for Baltic Sea Research | Krohnert K.,University of Gottingen | Hillion F.,Cameca | And 2 more authors.
Nature communications | Year: 2014

The isotopic composition of different materials can be imaged by secondary ion mass spectrometry. In biology, this method is mainly used to study cellular metabolism and turnover, by pulsing the cells with marker molecules such as amino acids labelled with stable isotopes ((15)N, (13)C). The incorporation of the markers is then imaged with a lateral resolution that can surpass 100 nm. However, secondary ion mass spectrometry cannot identify specific subcellular structures like organelles, and needs to be correlated with a second technique, such as fluorescence imaging. Here, we present a method based on stimulated emission depletion microscopy that provides correlated optical and isotopic nanoscopy (COIN) images. We use this approach to study the protein turnover in different organelles from cultured hippocampal neurons. Correlated optical and isotopic nanoscopy can be applied to a variety of biological samples, and should therefore enable the investigation of the isotopic composition of many organelles and subcellular structures.

Valley J.W.,University of Wisconsin - Madison | Cavosie A.J.,University of Wisconsin - Madison | Cavosie A.J.,University of Puerto Rico at Mayaguez | Ushikubo T.,University of Wisconsin - Madison | And 8 more authors.
Nature Geoscience | Year: 2014

The only physical evidence from the earliest phases of Earth's evolution comes from zircons, ancient mineral grains that can be dated using the U-Th-Pb geochronometer. Oxygen isotope ratios from such zircons have been used to infer when the hydrosphere and conditions habitable to life were established. Chemical homogenization of Earth's crust and the existence of a magma ocean have not been dated directly, but must have occurred earlier. However, the accuracy of the U-Pb zircon ages can plausibly be biased by poorly understood processes of intracrystalline Pb mobility. Here we use atom-probe tomography to identify and map individual atoms in the oldest concordant grain from Earth, a 4.4-Gyr-old Hadean zircon with a high-temperature overgrowth that formed about 1 Gyr after the mineral's core. Isolated nanoclusters, measuring about 10 nm and spaced 10-50 nm apart, are enriched in incompatible elements including radiogenic Pb with unusually high 207 Pb/ 206 Pb ratios. We demonstrate that the length scales of these clusters make U-Pb age biasing impossible, and that they formed during the later reheating event. Our tomography data thereby confirm that any mixing event of the silicate Earth must have occurred before 4.4 Gyr ago, consistent with magma ocean formation by an early moon-forming impact about 4.5 Gyr ago. © 2014 Macmillan Publishers Limited.

Gyngard F.,University of Washington | Gyngard F.,Carnegie Institution of Washington | Zinner E.,University of Washington | Nittler L.R.,Carnegie Institution of Washington | And 3 more authors.
Astrophysical Journal | Year: 2010

We report new O isotopic data on 41 presolar oxide grains, 38 MgAl 2O4 (spinel) and 3 Al2O3 from the CM2 meteorite Murray, identified with a recently developed automated measurement system for NanoSIMS. We have also obtained Mg-Al isotopic results on 29 of the same grains (26 spinel and 3 Al2O3). The majority of the grains have O isotopic compositions typical of most presolar oxides, fall well into the four previously defined groups, and are most likely condensates from either red giant branch or asymptotic giant branch stars. We have also discovered several grains with more unusual O and Mg compositions suggesting formation in extreme astrophysical environments, such as novae and supernovae (SNe). One of these grains has massive enrichments in 17O, 25Mg, and 26Mg, which are isotopic signatures indicative of condensation from nova ejecta. Two grains of SN origin were also discovered: one has a large 18O/16O ratio typical of Group 4 presolar oxides; another grain is substantially enriched in 16O, and also contains radiogenic 44Ca from the decay of 44Ti, a likely condensate from material originating in the O-rich inner zones of a Type II SN. In addition, several Group 2 presolar spinel grains also have large 25Mg and 26Mg isotopic anomalies that are difficult to explain by standard nucleosynthesis in low-mass stars. Auger elemental spectral analyses were performed on the grains and qualitatively suggest that presolar spinel may not have higher-than-stoichiometric Al/Mg ratios, in contrast to SIMS results obtained here and reported previously. © 2010. The American Astronomical Society. All rights reserved.

An achromatic magnetic mass spectrometer, for example of the SIMS type with double focusing, comprises means for canceling the four aberrations of the second order, and means for canceling the off-axis achromatism and for modulating the dispersion in mass.

A mass analysis device with wide angular acceptance, notably of the mass spectrometer or atom probe microscope type, includes means for receiving a sample, means for extracting ions from the surface of the sample, and a reflectron producing a torroidal electrostatic field whose equipotential lines are defined by a first curvature in a first direction and a first center of curvature, and a second curvature in a second direction perpendicular to the first direction and a second center of curvature, the sample being positioned close to the first center of curvature.

Agency: European Commission | Branch: H2020 | Program: ECSEL-RIA | Phase: ECSEL-06-2015 | Award Amount: 23.11M | Year: 2016

The objective of the 3DAM project is to develop a new generation of metrology and characterization tools and methodologies enabling the development of the next semiconductor technology nodes. As nano-electronics technology is moving beyond the boundaries of (strained) silicon in planar or finFETs, new 3D device architectures and new materials bring major metrology and characterization challenges which cannot be met by pushing the present techniques to their limits. 3DAM will be a path-finding project which supports and complements several existing and future ECSEL pilot-line projects and is linked to the MASP area 7.1 (subsection More Moore). Innovative demonstrators and methodologies will be built and evaluated within the themes of metrology and characterization of 3D device architectures and new materials, across the full IC manufacturing cycle from Front to Back-End-Of-Line. 3D structural metrology and defect analysis techniques will be developed and correlated to address challenges around 3D CD, strain and crystal defects at the nm scale. 3D compositional analysis and electrical properties will be investigated with special attention to interfaces, alloys and 2D materials. The project will develop new workflows combining different technologies for more reliable and faster results; fit for use in future semiconductor processes. The consortium includes major European semiconductor equipment companies in the area of metrology and characterization. The link to future needs of the industry, as well as critical evaluation of concepts and demonstrators, is ensured by the participation of IMEC and LETI. The project will directly increase the competitiveness of the strong Europe-based semiconductor Equipment industry. Closely connected European IC manufacturers will benefit by accelerated R&D and process ramp-up. The project will generate technologies essential for future semiconductor processes and for the applications enabled by the new technology nodes.

News Article | January 6, 2016

The starting material for experiments to determine the melting-phase relations of carbonated MORB (ATCM1) replicates basalts from the IODP 1256D from the Eastern Pacific Rise20 (the reported composition of IODP 1256D basalts is the average of all analyses presented in table T17 of ref. 20) with an added 2.5 wt% CO (Extended Data Table 1). This material was formed by mixing high-purity SiO , TiO , Al O , FeO, MnO, MgO, Ca (PO ) and CaCO , which were fired overnight at temperatures of 400–1,000 °C, of appropriate weights in an agate mortar under ethanol. This mixture was decarbonated and fused into a crystal-free glass in a one-atmosphere tube furnace by incrementally increasing the temperature from 400 to 1,500 °C before drop quenching into water. Subsequently weighed amounts of CaCO , Na CO and K CO were ground into the glass, introducing the alkali and CO components. After creation, the starting material was stored at 120 °C to avoid absorption of atmospheric water. Starting material ATCM2 replicates the near-solidus melt composition measured in melting experiments at 20.7 GPa and 1,400/1,480 °C. This was created by grinding natural magnesite and synthetic siderite with high-purity CaCO , Na CO , K CO , SiO , TiO , Al O and Ca (PO ) . Synthetic siderite was created in a cold-seal pressure vessel experiment run at 2 kbar and 375 °C for 7 days. A double Au capsule design containing iron (II) oxalate dehydrate in the inner and a 1:1 mixture of CaCO and SiO in the outer capsule produced a pale beige powder confirmed as siderite using Raman spectroscopy. The material for a sandwich experiment, to ensure near-solidus melt compositions were accurately determined at 20.7 GPa, was formed of a 3:1 mixture of ATCM1:ATCM2. The transition-zone peridotite mineral assemblage in reaction experiments was synthesized at 20.7 GPa and 1,600 °C for 8 h from a mixture of KR4003 natural peridotite31 with an added 2.5 wt% Fe metal. In reaction runs the recovered synthetic peridotite was loaded in a second capsule, surrounded by the ATCM2 near-solidus melt composition. Additional reaction-type experiments were performed on ground mixtures of peridotite and melt compositions. In these experiments PM1 pyrolite32 was used as the peridotite component and mixed with ATCM2 melt in 9:1, 7:3 and 1:1 weight ratios in Fe capsules. A single mixed experiment was performed in a Au capsule and used a starting mix of PM1:Fe:ATCM2 in 16:1:4 molar ratio. High-pressure experiments were performed using a combination of end-loaded piston cylinder (3 GPa) and Walker-type multi anvil (5–21 GPa) experiments at the University of Bristol. Piston cylinder experiments employed a NaCl-pyrex assembly with a straight graphite furnace and Al O inner parts. Temperature was measured using type D thermocouple wires contained in an alumina sleeve and positioned immediately adjacent to the Au Pd sample capsule that contained the powdered starting material. We assume that the temperature gradient across the entire capsule (<2 mm) was smaller than 20 °C (refs 33, 34). The hot piston-in technique was used with a friction correction of 3% applied to the theoretical oil pressure to achieve the desired run conditions35. Multi-anvil experiments were performed using Toshiba F-grade tungsten carbide cubes bearing 11, 8 or 4 mm truncated corners in combination with a pre-fabricated Cr-doped MgO octahedron of 18, 14 or 10 mm edge length, respectively. The relationship between oil-reservoir and sample pressure for each cell was calibrated at room and high temperature (1,200 °C) by detecting appropriate room temperature phase transitions of Bi, ZnTe and GaAs and bracketing transformations of SiO (quartz-coesite and coesite-stishovite), Mg SiO (α-β and β-γ) and CaGeO (garnet-perovskite). Calibrations are estimated to be accurate within ±1 GPa. In all experiments, desired run pressure was achieved using a slow, Eurotherm controlled, pressure ramp of ≤50 tonnes per hour. Experiments were heated after high pressure was reached with high temperatures generated using stepped graphite (18/11 cell) or straight LaCrO furnaces (14/8 and 10/4 cells) and monitored with type C thermocouple wires. Two 10/4 experiments, performed during a period of repeated LaCrO heater failures, used rolled 40-μm-thick Re furnaces. Temperature was quenched by turning off the furnace power before a slow decompression ramp (half the rate of experiment compression) to ambient conditions. Samples were contained in Au capsules unless temperatures exceeded its thermal stability, in which case Au Pd or Au Pd capsules were used. Run durations all exceeded 600 min and are reported in Extended Data Tables 2 and 3. Temperature uncertainties were believed to be less than ±20, 30 or 50 °C for 18/11, 14/8 and 10/4 cells respectively36, 37. Recovered samples were mounted longitudinally in epoxy, polished under oil and repeatedly re-impregnated with a low viscosity epoxy (Buelher EpoHeat) to preserve soft and water-soluble alkali carbonate components present in run products. Polished and carbon-coated run products were imaged in backscatter electron mode (BSE) using a Hitachi S-3500N scanning electron microscope (SEM) with an EDAX Genesis energy dispersive spectrometer to identify stable phases and observe product textures. Subsequently, wavelength dispersive spectroscopy (WDS) was performed using the Cameca SX100 Electron Microprobe or the Field Emission Gun Jeol JXA8530F Hyperprobe at the University of Bristol to achieve high-precision chemical analyses of run products. Analyses were performed using an accelerating voltage of 15 or 12 kV on the respective instruments, with a beam current of 10 nA. Calibrations were performed during each session using a range of natural mineral and metal standards and were verified by analysing secondary standards (as described previously6). Silicate phases were measured using a focused electron beam whereas carbonates and melts were analysed using an incident beam defocused up to a maximum size of 10 μm. Count times for Na and K were limited to 10 s on peak and 5 s on positive and negative background positions. Peak count times for other elements were 20–40 s. Additional analyses of the calcium perovskite phases grown during reaction experiments, measuring only SiO and MgO content, were made using the Jeol instrument at 5 kV and 10 nA to ensure reported MgO contents were not influenced by secondary fluorescence from surrounding material. The identity of experimental-produced minerals was determined using Raman spectroscopy as a fingerprint technique. Spectra were collected using a Thermo Scientific DXRxi Raman microscope equipped with an excitation laser of either 455 or 532 nm. Studies that investigate the alteration of oceanic crust have demonstrated that carbon incorporation does not simply occur by the addition of a single carbonate species to MORB9. It instead appears to occur by a complex amalgamation of hydrocarbon and graphite deposition related to hydrothermal fluxing above magma chambers at the mid-ocean ridge8 and underwater weathering9, 38, 39, 40 where seawater-derived CO reacts with leached crustal cations, often in veins. It is believed that the quantity of biotic organic carbon in the crustal assemblage is negligible compared with abiotic organic compounds and inorganic carbonates8. These processes result in a layered crustal assemblage that, in the uppermost few hundred metres can contain up to a maximum of 4 wt% CO in rare cases9, 39 but more commonly <2 wt% CO (refs 8, 9, 39). Beneath 500 m depth the carbon content drops to between 100 and 5,000 p.p.m. CO throughout the remainder of the 7-km-thick basaltic section8, and is mostly organic hydrocarbon species. The upper 300 m are regularly altered and can be generally thought to have compositions similar to the altered MORB rocks analysed previously41. Deeper portions of the MORB crust retain their pristine MORB compositions. It is therefore apparent that carbonated eclogite bulk compositions used in previous studies, where at least 4.4 wt% CO was added to an eclogite by addition of ~10 wt% carbonate minerals, may not be good analogues of naturally subducting crustal sections. The compositions of these starting materials from previous studies19, 42, 43, 44, 45, 46 can be found in Extended Data Table 1. We do not include the composition of the starting material used by refs 47 or 48 as these studies were conducted in simplified chemical systems so are not directly comparable with these natural system compositions. However, as some of the previous studies rightly identify and discuss, the composition of deeply subducted MORB is unlikely to be the same as that entering the subduction system. One process widely believed to alter the composition of downwelling MORB is sub-arc slab dehydration. Pressure (P)–temperature (T) paths of subducted slabs26 can be compared with experimental studies of hydrous, carbonated and H O-CO -bearing eclogite compositions12, 24, 42, 43, 49 and thermodynamic models11, 50 to conclude that slabs experience dehydration at sub-arc conditions (that is, 1–5 GPa) but will generally not reach high enough temperatures to undergo melting. Therefore, they will by and large retain their carbon components although some fraction may be lost by dissolution into aqueous fluids51, 52. It is believed that sub-arc dehydration is capable of removing SiO from the subducting assemblage, and previous carbonated MORB compositions were therefore designed to be considerably silica undersaturated (relative to fresh/altered MORB)19, 43, 44, 45. While studies53, 54, 55, 56 do indicate that SiO can become soluble in H O at high pressures, they infer that the solubility of silica in hydrous fluids only exceeds ~1 wt% at T > 900 °C at 1 GPa (higher T at higher P). In contrast, slab dehydration occurs on all prograde slab paths at T < 850 °C. Additionally, the composition of quenched hydrous fluids coexisting with MORB at 4 GPa and 800 °C (ref. 57) indicate that a maximum of ~12 wt% SiO can dissolve in the fluid. Given that there should be considerably less than 10 wt% H O (more likely << 5 wt% H O) in subducting assemblages, this suggests a maximum SiO loss in subducting MORB lithologies of ~0.6–1.2 wt%. The compositions used in previous studies have SiO depletions ranging from 3 wt% up to, more commonly, 6–10 wt% SiO relative to MORB. We further investigated the effect of oceanic crust alteration and sub-arc dehydration on the composition of subducted MORB rocks by compiling a data set of altered MORB41 and exhumed blueschist, greenschist and eclogite facies rocks from exhumed terrains worldwide to compare them with fresh MORB21, our starting material and previous starting materials. We then assess the relevance of our starting material based on the composition of natural MORB rocks, rather than using models of the subduction process that contain few observable constraints. Results of this comparison are plotted in Extended Data Fig. 1. This analysis confirms that relative to fresh MORB, altered MORB and exhumed crustal rocks are somewhat depleted in SiO , up to a maximum of 6 wt% SiO in the most extreme case, but more commonly 0–3 wt% SiO . Thus, many previous starting materials are too silica undersaturated to be good analogues of subducting MORB. Furthermore, this analysis reveals that altered and exhumed MORB are not enriched in CaO compared with fresh MORB, if anything they actually contain lower CaO on average. In contrast, all previous starting materials are enriched in CaO compared with fresh MORB. This is because most previous studies introduced the carbon component to their experiment by adding ~10 wt% calcite to an eclogite-base composition. We note that SLEC1 (ref. 43) was not created in this manner, but instead this composition falls far from the MORB field as the authors used an eclogite xenolith erupted by a Hawaiian volcano as a base material. By plotting the position of the maj–cpx join, defined by the composition of our experimental phases plotted in Extended Data Fig. 5, onto Extended Data Fig. 1a, we demonstrate that our bulk composition (ATCM1), ALL-MORB21, the vast majority of the fresh MORB field, altered41 and exhumed MORB samples fall on the CaO-poor side of this join, that is, on the Mg+Fe-rich side. Therefore, magnesite will be the stable carbonate phase in these compositions at high pressure (above dolomite breakdown). In contrast, all previous bulk compositions plot on the Ca-rich side of this join, or are very depleted in SiO , and therefore fall in a different phase field to the overwhelming majority of subducted MORB. This difference causes a considerable difference in the phase relations of our starting material relative to those used in previous studies. We acknowledge that no single bulk composition can be a perfect analogue for the entire range of subducting MORB compositions, however, ATCM1 is a good proxy for sections of the MORB crust between ~300 m and 7 km depth that have unaltered major element compositions and low CO contents. Additionally, ATCM1 remains a better analogue for the uppermost portions of the MORB crust than starting materials employed in previous studies because its CO content is within the range of natural rocks while it is also not oversaturated in CaO or over depleted in SiO . This is despite it falling towards the SiO -rich end of the compositional spectrum of subducting MORB rocks. Recent experiments have suggested that carbonate in eclogitic assemblages may be reduced to elemental carbon, either graphite or diamond, at depths shallower than 250 km (ref. 58). However, subducting slab geotherms are much colder than the experimental conditions investigated by this study, and additionally they are believed to contain considerable ferric iron that is further increased during de-serpentinization10. Indeed, several observations of carbonate inclusions in sub-lithospheric diamonds6, 7, 59 require that slab carbon remains oxidized and mobile until diamond formation, far deeper than 250 km. Given the numerous observations from natural diamond samples, the general uncertainty in the mantle’s fO structure and the lack of any conclusive experimental evidence that subducting carbon becomes reduced before reaching the transition zone we posit that nearly all subducting carbon is stable as carbonate throughout the upper mantle in subducting MORB assemblages. Extended Data Table 2 presents the run conditions, durations and phase proportions in all carbonated MORB melting experiments, which are also summarized in Extended Data Fig. 2. Phase and melt compositions are presented in the Supplementary Tables 1–4. Phase proportions are calculated by mass balance calculations that use the mean composition of each phase as well as the reported 1σ uncertainty in this mean as inputs. We note that the 1σ uncertainty for some oxides in garnet and clinopyroxene minerals occasionally exceeds 1 wt%, although it is normally much smaller than this. These large uncertainties are a function of the small crystal sizes present in some runs, and not a function of sluggish reaction kinetics. Phase proportion calculations were run in a Monte Carlo loop of 10,000 calculation cycles where a varying random error was added to each oxide in each mineral phase during each iteration. Overall the distribution of varying random errors for each oxide form a Gaussian distribution with standard deviation equal to the reported 1σ uncertainty of measurements. The reported proportions are the numerical mean of all calculation cycles and the r2 value reports the average squared sum of residuals. Low r2 values indicate that chemical equilibrium is likely to have been achieved and that mineral and melt compositions have been accurately determined. Representative BSE images of the polished experiments are shown in Extended Data Fig. 3. Garnets in experiments at all pressures contain abundant SiO inclusions. In subsolidus experiments the number of inclusions increases and the definition of mineral boundaries deteriorates, which makes accurate analysis of garnet compositions increasingly challenging. In supersolidus runs, garnet minerals adjacent, or near to, carbonatite melt pools have well defined edges and contain fewer inclusions. However, far from quenched melts the textures of garnets remain small and pervasively filled with inclusions, indicating the influence of melt fluxing on mineral growth. With increasing pressure, garnets become increasingly majoritic, with increasing quantities of octahedral silicon. Clinopyroxene was observed in all subsolidus experiments, as euhedral crystals that are often spatially associated with the carbon-bearing phase. Cpx abundance falls with increasing pressure and their compositions becoming increasingly dominated by sodic components (jadeite, aegerine and NaMg Si O ) at high pressure (Extended Data Fig. 5). Cpx only disappears from the stable phase assemblage in supersolidus experiments at 20.7 GPa. SiO is observed in all runs and are small, often elongated tabular-shaped crystals. An oxide, either TiO at low pressure or an Fe-Ti oxide above 13 GPa (as described previously24) are observed in all subsolidus runs. The carbon-bearing phase in subsolidus experiments changes with increasing pressure. At 3 GPa CO , marked by the presence of voids in the polished sample, is stable. This converts to dolomite at 7.9 GPa, consistent with the position of the reaction 2cs + dol = cpx + CO (ref. 22). Beyond ~9 GPa dolomite becomes unstable and breaks down into magnesite + aragonite23. Therefore, because the ATCM1 bulk composition lies on the Mg+Fe2+-rich side of the garnet–cpx join (Extended Data Figs 1a and 5), magnesite replaces dolomite as the carbon host in the experimental phase assemblage. This differs from experiments in previous studies, where aragonite was dominant because bulk compositions fall on the opposite side of the garnet–cpx join. It is clear from the ternary diagrams (Extended Data Fig. 5) that while the tie-line between garnet and cpx remains, magnesite and aragonite cannot coexist in a MORB bulk composition. Finally, at pressures above 15 GPa, Na-carbonate becomes stable in the subsolidus phase assemblage. This is chemographically explained by the rotation of the garnet–cpx tie-line with increasing pressure (EDF5). Its appearance can also be justified as a necessary host of sodium at increasing pressure, since aside from clinopyroxene there is no other Na-rich phase stable on the Mg+Fe side of the maj–cpx join. The appearance of silicate melt, containing dissolved CO (estimated by difference), defines the solidus at 3 GPa. This may initially appear to contradict the results of some previous studies, which find carbonatite melts are produced near the solidus of carbonated eclogite at pressures lower than 7 GPa (refs 43, 45, 46). However, this is easily explained by the differences in CO and SiO content used in these studies. The higher CO and lower SiO contents of previous studies stabilize carbonate melt to lower temperatures relative to silicate melts. Indeed, we note that our results are consistent with those described previously42, 44 (the two previous studies with the least depleted SiO ), which also observed that near-solidus melts below 5 GPa were basaltic to dacitic silicate melts containing dissolved CO . The results of one paper19 are not entirely self-consistent, in that at some pressures between 3.5 and 5.5 GPa the authors observed silicate melts before carbonate melts (4.5 and 5 GPa), whereas this relationship is sometimes reversed (5 GPa in AuPd capsules) or both melts were observed together (3.5 GPa). The observation of two immiscible melts in previous studies probably reflects the maximum CO solubility in silicate melts. Since our bulk composition has less CO , akin to natural rocks, we do not observe liquid immiscibility. In all experiments above 7 GPa, near-solidus melt compositions are carbonatititc and essentially silica-free. This result is notably different from those described previously19, which reported that near-solidus melts were a mixture of silicate, carbonated silicate and carbonatite melts. We believe this contrast is caused by the interpretation of experimental run textures. Whereas ref. 19 identified regions of fine-grained material consisting of mixtures of stable phases from elsewhere in the capsule as quenched melts, we have not followed the same interpretation of these features. Although we do recognize similar features in some run products, we have interpreted these features as a consequence of poor crystal growth in regions far from the influence of melt fluxing. In all supersolidus experiments, we observed regions of carbonatite material (typically <1 wt% SiO ) that is fully segregated from surrounding silicate minerals and possesses a typical carbonate-melt quench texture (Extended Data Fig. 3). Silicate minerals in close proximity to these melt pools are larger than those elsewhere in the same experiment, have well-defined crystal boundaries and contain few inclusions. Therefore, we attribute the variable texture and regions of fine-grained material present in experiments to the location of melt within experiments, which has a tendency to segregate to isolated regions of capsules under influence of temperature gradients. Although melt segregation occurs in all supersolidus experiments, the efficiency of segregation and size of melt pools considerably increases with rising temperature above the solidus. Extended Data Figure 4 shows the highly systematic evolution of the melt compositions reported from our study with increasing pressure, strongly supporting our interpretations. Carbonatite melts are calcic, Ca number > 0.5 (Ca number = Ca/[Ca+Mg+Fe]), despite subsolidus carbonates being dominated by magnesite (Extended Data Fig. 4). Melts have high concentrations of TiO (typically 1–3.5 wt%), P O (0.4–1.5 wt%) and K O (0.3–1.5 wt%) and a variable Mg number (0.33–0.7 defined as Mg/[Mg+Fe]). The alkali content of melts, strongly dominated by Na O due to the bulk composition, increases with pressure (from 1 to ~15 wt% Na O at 7.9 and 20.7 GPa respectively; Extended Data Fig. 4). This increasing Na O content is driven by the decreasing compatibility of Na O in the residual mantle phase assemblages as the abundance of stable clinopyroxene falls. At 20.7 GPa the melt composition, as evidenced both by constant phase proportions and consistent melt/majorite compositions, remains constant over a temperature interval of ~350 °C above the solidus. It is only when temperature reaches 1,530–1,600 °C (runs #16 and #31) that the silica content of the melt begins to increase (to 8.7 wt%) and CO content falls as melts start to become silica-carbonatites. One experiment (#33) aimed to verify that measured low-degree melt compositions are accurate, and are not affected by analytical problems related to the small size of melt pools, was conducted at 20.7 GPa. In this experiment the abundance of carbonate melt was increased by adding a mix replicating the low degree melt composition ATCM2 to ATCM1 in a mass ratio of 1:3. If the composition of low-degree melts has been accurately determined in ‘normal’ experiments then this addition will have a negligible effect on phase relations or the compositions of the garnet, SiO or melt; it would simply increase the melt abundance. The result of this experiment has a similar texture to all other experiments, where carbonatite melt segregates to one end of the capsule and is adjacent to large, well-formed majoritic garnets. The far end of the capsule has a much smaller crystal size, crystals have ragged edges, garnets are full of inclusions and SiO is present along grain-boundaries and triple junctions (Extended Data Fig. 3h). Mineral and melt compositions, although not exactly identical, are similar to those measured in ‘normal’ experiments (to achieve identical compositions an iterative approach would be required that was not deemed to be necessary) thus confirming that near-solidus melt compositions have been accurately determined. The presence of fine-grained material away from segregated melt also acts to further confirm our hypothesis regarding the vital importance of melt presence for growing large crystals during experiments. Comparing our starting material and results with those of previous studies using ternary and quaternary projections (Extended Data Fig. 5) reveals that it is not possible for both magnesite and aragonite to coexist alongside majorite and clinopyroxene owing to stable mineral phase fields (see earlier). Thus, in Mg-Fe-dominated compositions, such as our starting material, magnesite is the stable carbonate at high-pressure subsolidus conditions. Whereas in Ca-dominated compositions aragonite will be the stable carbonate beyond the pressure of dolomite dissociation. Natural subducting MORB compositions, which contain, at most, a similar quantity of CO to our bulk composition11, almost all lie on the Ca-poor side of the majorite–clinopyroxene join (Extended Data Figs 1 and 5). In this situation, as our experiments demonstrate, cpx remains an important Na-host in MORB assemblages to high pressures alongside [Na,K] Ca CO structured carbonate. Ca-rich compositions containing subsolidus CaCO experience different phase relations because aragonite can dissolve considerable Na O and so is the sole Na-host in these compositions. We conclude that because the majority of natural MORB rocks fall on the Mg+Fe side of the maj–cpx join, like our bulk composition, that the phase relations determined in this study are applicable to the case of natural subduction. Therefore, the melting point depression we observe along the carbonated MORB solidus at uppermost transition zone pressures is generally applicable to subducted oceanic crust. Without the influence of slab-derived melts, the anhydrous transition zone peridotite assemblage at 20.7 GPa and 1,600 °C (experiment G168 and G176) is dominated by Na-poor majorite and wadsleyite (Mg number = 0.90) (Extended Data Fig. 6, Extended Data Table 3 and Supplementary Table 5a). Upon reaction with the near-solidus alkaline carbonatite defined during melting experiments, ATCM2, a clearly defined reaction zone is observed between this ambient peridotite assemblage and the infiltrating melt (Extended Data Fig. 6). The products of this reaction are garnet containing a notable Na X2+Si O majorite component, Ca(Si,Ti)O perovskite, ringwoodite, ferropericlase and diamond. All of these phases were identified using Raman spectroscopy (Extended Data Fig. 7) and their compositions are presented in Supplementary Table 5a. Raman spectroscopy alone, which was performed before any sample polishing using diamond-based products, confirms the creation of diamond during these reactions. We have not observed diamond using SEM techniques and believe that it resides as sub-micrometre-sized inclusions in the various reaction-product minerals where it is seen by spectroscopic methods. The experiments performed on intimately mixed powders of melt and pyrolite also form the same phase assemblages (Extended Data Table 3) and mineral compositions from those runs are also presented in Supplementary Table 5b, c. We observed the reaction products as new crystals floating in the residual carbonatite melt and/or nucleated on the relics of the peridotite assemblage, thus creating zoned minerals. We have demonstrated that the composition of majorite minerals crystallizing during the reactions lie between those expected for peridotitic and eclogitic minerals at a similar pressure and possibly explain intermediate-composition diamond-hosted majorites (Fig. 2). We suggest that the full range of intermediate inclusion compositions might be created by the gradual shift in phase compositions, from those we observe towards more peridotitic minerals as the melt composition reacts with increasing quantities of mantle material. Additionally we have shown that the compositions of calcium perovskite (Extended Data Fig. 8) and ferropericlase (Fig. 3) formed during the reactions are consistent with diamond-hosted minerals of those species. Further experiments, across the solidus ledge and into the uppermost lower mantle pressure range are required to test whether melt–mantle interactions account for all diamond-hosted inclusions.

News Article | April 11, 2016

The U.S. Naval Research Laboratory (NRL), Materials Science and Technology Division, has taken delivery of a state-of-the-art Cameca 4000X Si Local Electrode Atom Probe (LEAP), a high performance microscope that provides precise atom-by-atom dissection of a material volume, enabling true three-dimensional (3D) atomic-scale reconstructions of material microstructures. "Exact knowledge of where individual atoms are in a material is of tremendous benefit when engineering new materials," said Keith Knipling, NRL Materials Science and Technology Division. "We expect the LEAP to greatly enhance our capability to develop new materials, including the next generation of structural alloys for stronger ship hulls and more advanced turbine engines, new electronic materials for tomorrow's faster integrated circuits, and advanced solar cell and battery materials with improved power and energy efficiency." The LEAP works using the principle of field evaporation, whereby a strong electric field applied to a needle-like specimen is sufficient to cause removal of atoms by ionization. Atom evaporation is triggered either by a voltage or laser pulse applied to the sample. The resulting ions are accelerated away from the specimen and identified chemically by time-of-flight mass spectrometry and their positions are deduced from the coordinates of ion impacts on a position-sensitive detector. By repeating this sequence, the atoms are progressively removed from the tip, and a 3D image of the material can be reconstructed at the atomic scale. "Mapping the location of each chemical species in a material microstructure enables an unprecedented understanding of the true effects of alloying and material synthesis, which is essential for truly optimizing the properties of any material. We expect the LEAP to deliver new atomic-scale perspective and insights into a wide range of materials science investigations. For example, researchers at NRL are developing semiconductor materials with very dilute levels of added 'dopant' atoms for the purpose of tailoring their electrical properties," Knipling adds. "These doped materials form the building blocks of nearly all semiconductor electronic devices such as diodes, transistors, solar cells, LEDs, and integrated circuits. Most analytical techniques are incapable of measuring these small concentration levels, much less where the dopants segregate within the microstructure. With the LEAP, researchers can now answer these questions." Worldwide, there are only a handful of LEAP microscopes in use. NRL Materials Science and Technology Division possesses the only one in use by the U.S. Department of Defense (DoD). About the U.S. Naval Research Laboratory The U.S. Naval Research Laboratory provides the advanced scientific capabilities required to bolster our country's position of global naval leadership. The Laboratory, with a total complement of approximately 2,500 personnel, is located in southwest Washington, D.C., with other major sites at the Stennis Space Center, Miss., and Monterey, Calif. NRL has served the Navy and the nation for over 90 years and continues to advance research further than you can imagine. For more information, visit the NRL website or join the conversation on Twitter, Facebook, and YouTube.

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