ION MICROPROBE MASS SPECTROMETRY: TECHNIQUES AND APPLICATIONS IN COSMOCHEMISTRY, GEOCHEMISTRY, AND GEOCHRONOLOGY

Trevor R. Ireland ______

I.!Introduction...... 2 2.!Sputtering ...... 3 A. Ion Production...... 3 B. Ionization Modeling ...... 10 C. Secondary Ionization Theory...... 12 3.!Hardware ...... 13 A. General Features...... 13 B. CAMECA...... 13 C. SHRIMP...... 19 D. VG Isolab 120...... 24 4. Methodology ...... 24 A. Isotopic analysis...... 24 B. Quantitative Analysis...... 46 ______Advances in Analytical Geochemistry Volume 2, pages 1-118. Copyright © 1995 by JAI Press Inc. All rights of reproduction in any form reserved. ISBN: 1-55938-785-8

1 2 TREVOR R. IRELAND

V.!Applications ...... 54 A. in Cosmochemistry...... 54 B. in Geochemistry ...... 82 C. in Geochronology...... 93 VI.!Prospects and Conclusions...... 105 Acknowledgments...... 107 References ...... 107

I.!INTRODUCTION

An ion microprobe is basically a mass spectrometer with a specialized source incorporating a finely focused primary ion beam to generate secondary ions from the target. The primary ions have energies of the order of 10 keV and so collisions between the primary ions and the surface physically erode, or “sputter”, the sample ejecting particles from the surface. Of these particles, a small fraction are ionized and can be electrostatically removed to the mass spectrometer where they are separated according to mass and the relative ion intensities can be measured. Ion microprobes have found a niche in materials research, particularly in the semiconductor industry where localized analysis of ppm to sub ppb concentrations are required. The commercially produced CAMECA ion microscopes have been particularly important in this area and over 200 instruments are in operation worldwide. However, only a small fraction of these instruments are used for research in the earth science field, and then mainly for research in cosmochemistry where isotopic and chemical abundance anomalies are at their largest. Of prime importance in the geological sciences has been the SHRIMP (Sensitive High mass- Resolution Ion MicroProbe) at the Australian National University that was designed for isotopic analysis of chemically complex targets. The main application of this instrument has been in situ U-Pb dating of zircon although a wide variety of applications have benefited from its high sensitivity. Until recently SHRIMP was the only instrument capable of undertaking routine U-Pb dating of zircon but now commercially produced machines are being marketed with these capabilities. The fundamentals for ion microprobe analysis were developed over thirty years ago with the goal of characterizing any sample for its chemical and isotopic abundances. In this same time frame, the electron microscope and microprobe were also being developed. The electron probe uses a focused electron beam to generate X-rays in the target and since the principles governing photon emission from a surface are well understood, quantitative analyses were readily forthcoming and the electron probe quickly became a necessary piece of apparatus in materials research laboratories worldwide. Development of the ion microprobe was much slower, because the principals that describe secondary ionization are not well understood and, despite a great deal of effort, no generally acceptable formulation has been forthcoming. However, the salient point involved in applying the ion microprobe to problems in geochemistry is that the primary ion beam does produce secondary Ion Microprobe Mass Spectrometry 3 ions that reflect in some way the isotopic and chemical characteristics of the sample. Accurately determining that relationship is the cornerstone of all successful applications. The purpose of this paper is to highlight the successful application of ion microprobe mass spectrometry to a variety of topics principally in the earth and planetary sciences. It is not a review of the development of the ion microscope and microprobe, previous reviews of ion microprobe mass spectrometry (Shimizu and Hart, 1982a; Reed, 1984; Benninghoven et al., 1987) have documented most of these aspects. Nor is it a review of the status of ion production models. While a complete model would be of great benefit to practitioners, it is not a requirement for successful application of the instruments as will be detailed in later sections. Moreover, the physics is complex and a detailed discussion is beyond the scope of this paper and probably beyond the requirements of the interested but non-expert to whom this work is addressed. Detailed discussions of sputtering models are given by Williams (1979; 1982), Harrison (1983) and Benninghoven et al. (1987). In the following sections, a brief overview of the sputtering process will be given followed by descriptions of the most recent generations of ion microprobes in use today. General aspects of analysis with the ion microprobe will then be discussed followed by specific examples from cosmochemistry, geochemistry, and geochronology.

II.!SPUTTERING A. Ion Production

16 – A O ion traveling at 350 km s-1 smashes into a wall; 10-10 s later it's all over! The damage is rather localized and extends only down through several atomic layers and only for several atomic radii around (Figure 1). For a 1 nA primary beam there are some 6 ¥ 109 impacts per second. An impact results in some 10 to 100 atoms and molecular fragments being ejected and in this way a sample is continually eroded. Felicitously, a small fraction of the emitted particles are ionized and can be electrostatically removed from the sputtering region to be analyzed in a mass spectrometer. While the basic scheme involved can be appreciated in macroscopic terms, the physical principles governing the ionization probability of individual particles are not readily quantifiable because of the diverse parameters that are needed to fully describe the characteristics of the ejected particles. The pioneering statistical theories of Thompson (1968) and Sigmund (1969) involved the partitioning of energy in collision cascades and were able to predict bulk properties of the sputtered material such as the sputtering yield, ejected atom energy, and secondary atom angular distributions as a function of primary ion energy. In order for a particle to be ejected it must obtain sufficient kinetic energy to overcome the surface binding energy. All sputtered atoms originate from the outer few atomic layers with the 4 TREVOR R. IRELAND

Figure 1. Sputtering of a solid surface by particles with energies of the order of 10 keV results in the disruption of the lattice and ejection of atoms and molecular fragments by the transfer of energy back towards the surface. These glancing collisions result in relatively low emission energies of the order of 10 eV although some particles may acquire energies of over 1000 eV. Most particles are neutral but a small fraction are ionized and this forms the basis of secondary ion mass spectrometry. majority coming from the surface itself (Williams, 1983). Since the atom must obtain momentum in the opposite direction to the incoming ion, most of the energy is transferred in glancing collisions and hence emitted ions have rather low energies peaking at around a few eV although a small fraction have energies extending up to the keV range (Metson et al., 1984). The highest energy fraction is atomic ions with complex molecular ions having increasingly narrower energy spreads in proportion to the number of constituent atoms. The vast majority of particles that are ejected are neutrals; only a small fraction (up to approximately 0.1%) are ionized and it appears that the ionized particles have an energy distribution similar to the neutral particles. If all particles emitted from Ion Microprobe Mass Spectrometry 5 the surface were ionized, a satisfactory model describing the relationship between sample chemistry and the secondary ion mass spectrum would be readily obtained. However, with only a small degree of ionization, there is scope for wide variations in ionization probabilities of a given species with changes in the sputtering environment. For example, the presence of oxygen, either as the primary ion species or simply leaked into the vicinity of the sputtering region, causes a large increase in the emission of positive ions. Similarly, sputtering with a Cs+ ion beam results in enhanced emission of negative ions. However, some elements essentially do not ionize, such as the inert gases He, Ne, Ar, Kr, and Xe, and also N. While the ion probe can operate in any of four configurations, incorporating positive or negative primary or secondary ions, the optimum configuration is to have opposite polarities of primary and secondary ions. The most common primary beam species are Cs+ and O– since these highly electropositive and electronegative species (respectively) have a chemical effect on the sputtering region enhancing ion emission many orders of magnitude over other source species such as Ar+. Since ion microprobe analysis relies on the transport of charged particles, care must be taken to avoid charge buildup, particularly at the sputtering site. The two most common analytical configurations are negative primary - positive secondary and positive primary - negative secondary. In the first mode, negative charge build up can be adequately dispersed with a conductive coating such as C or Au. However, with positive ion bombardment rapid charging is apparent. This is due to the combined addition of the positive ions from the primary beam and the extraction of both negative secondary ions and electrons. Even a relatively thick gold coat is not sufficient to allow electrons access to the sputtering site to neutralize the positive charge. This presents one of the major problems in the analysis of light stable isotopes of elements such as H, C, and O that are best analyzed with a Cs+ primary beam and negative secondary ions. There have been two main solutions to this problem, but neither has proved to be completely satisfactory. The first involves reducing the sample to 10-15 µm fragments and embedding them in gold foil. This removes some of the benefits of in situ analysis but still allows the analysis of reasonably small intact samples. The second method involves neutralizing the sputtering site by focusing an electron beam into the analysis area. CAMECA ion microprobes now include an electron gun for this purpose and have been successful in obtaining stable intense O– beams from oxides and silicates. However, the reproducibility of O isotope ratios for either method is only reliable to around the 2 ‰ level at best at present and further work is required to constrain the isotopic mass fractionation to a level which will allow precise and accurate analyses of insulators. The secondary ion mass spectrum from the ion microprobe is complicated, not only because of the chemical diversity in geological samples, but because molecular fragments of all types can be produced. The molecules are best described as fragments because they can have no stoichiometry in terms of valence that generally governs the configuration of molecular compounds. Instead, any algebraic combination of the components in the sample can be produced. Even for the simplest 6 TREVOR R. IRELAND

Figure 2. Negative secondary ion spectrum from a graphite substrate sputtered with an O+ primary beam (from McKeegan, 1987). Even from the simplest targets, the spectrum is complicated by the presence of a plethora of molecular species. The species can not be regarded as having stoichiometric affinities but rather should be regarded as simple algebraic combinations of the target atoms. samples, the spectra can be extremely complicated. For example, the secondary ion mass spectrum of graphite sputtered with an O– beam in Figure 2 shows combinations of C isotopes as well as molecules composed of C and O. For a geological sample the situation is even more chaotic. The data collected from an ion microprobe generally consists of ratioing peak heights of different isotopes of an element for isotopic analysis, or isotopes of different elements for a chemical analysis. Before data can be obtained any interferences under the peaks of interest must either be removed or their abundances quantified. Since the ion probe is a mass spectrometer the primary method of peak separation is according to mass. The first ion microprobes were constructed with low resolution spectrometers so that basically only unit masses could be separated. With the observation of complex molecular interferences it was soon realized that much higher mass resolutions would be essential for fully utilizing the capabilities of the ion probe. The regular variation of mass deficits of individual isotopes means that molecular interferences can be readily separated from atomic species for both low and high mass regions of the mass spectrum. For intermediate masses the mass deficit curve turns around and much higher mass resolutions are required. The masses are dispersed along the focal plane of the magnet with their separations dependent on the characteristics of the magnet as well as the width of the Ion Microprobe Mass Spectrometry 7

Figure 3. Schematic of peak shape showing relationship with slit widths. The 50 % peak-height width is proportional to the collector slit width and the width of the rise on the peak is proportional to the source slit width. Mass resolution is commonly defined as M/DM at the 10 % peak height. entrance and exit slits of the mass spectrometer. Since the characteristic dispersion of the magnetic analyzer is a fixed parameter, variation in mass resolution can only be accommodated through the opening or closing of the slits; the narrower the slits, the higher the mass resolution. The generalized peak shape is trapezoidal with the half peak width being equal to the collector slit width, and the slope dependent on the source slit width (Figure 3). The mass resolution is the mass of the peak divided by the base width of that peak (M/DM). There is some variability in defining mass resolution because of the different levels at which the base width is measured. In detail it is important to designate the height used to define the base width since a base width at the 1% level will be wider than that at the 10 % level and therefore the 10 % level will define a higher mass resolution than the 1 % level. A common convention is to use the 10 % base width and this will be used in this paper. An extension to the concept of mass resolution is the abundance sensitivity of a mass analyzer. Basically this can be viewed as the degree to which an adjacent mass interferes with a peak and is therefore related to peak tailing caused by, for example, gas scattering. A high abundance sensitivity requires as low a contribution as possible and this is an important parameter in the measurement of low abundance isotopes such as 10Be, 26Al, 234U, and 230Th. The abundance sensitivity can be defined for any mass, but a common convention is to measure the contribution at mass 237 from the 238U signal. While conventional mass spectrometers have mass resolutions of the order of 500!R, enabling separation of individual masses up to 500 amu, ion microprobes 8 TREVOR R. IRELAND

Figure 4. High mass resolution spectra of peaks in the region of the Ti isotopes from Fahey (1988). All singly charged molecular isobars, including hydrides, are resolved, and peak overlap occurs with only atomic species and doubly charged atomic species. While the 48Ca+ and 48Ti+ peaks are not fully resolved, 48Ca+ contributes less than 0.1 ‰ to the 48Ti+ signal when the latter is centered through the exit slit. However, this is not the case for mass 50 where there is no separation of 50Cr+ and 50V+ from 50Ti+ and the intensity of 50Ti+ must be corrected for the contributions of these isobaric interferences by monitoring 51V+ and 52Cr+. generally operate at much higher mass resolutions on the order of 3,000 to 10,000 R. The regular variation of the mass deficits of the chemical elements with mass enables the separation of all molecular isobaric interferences at mass resolution of the order of 10,000 R for elements of mass less than around 50 amu. Figure 4 shows the mass spectrum in the region of the titanium isotopes in a meteoritic hibonite at a mass resolution of ≈10,000 R. Molecular interferences are well resolved, for example 46TiH+ and 30Si16O+ are well separated from 47Ti+. In general hydrides require larger resolving powers than oxides but atomic isobaric interferences generally require much higher mass resolution for separation. In the case of the Ti isotopes only 48Ca is partially resolved from 48Ti, the other atomic isobaric interferences, 46Ca, 50V, and 50Cr are not resolved. For isotopic masses heavier than around 50 amu, molecular interferences become increasingly similar in mass to the atomic species and hence more difficult to resolve by mass separation. An alternative method for discriminating against molecular ions is the energy filtering technique. This method relies on the differences in energy distribution of Ion Microprobe Mass Spectrometry 9

+ + + Figure 5. Energy spectrum showing Si , SiO , and SiO2 sputtered from zircon as measured on SHRIMP I. The molecular species have increasingly narrow energy bandwidths as the complexity (number of constituent atoms increase). This forms the basis of the energy filtering technique which can used to discriminate against complex molecular interferences. atomic and molecular ions. Atomic species have a broad energy spread whereas molecular species have increasingly narrow energy bands depending on the number of atoms in the molecule. Figure 5 shows an energy spectrum for Si+, SiO+ and + + SiO2 sputtered from zircon. It can be seen that at an energy offset of ≥50 V, SiO2 + is excluded from collection. However, while SiO2 has fallen in intensity by 5 orders of magnitude, the atomic species has dropped by 2 orders of magnitude so interference discrimination is at the expense of intensity of the atomic species as well. A redeeming benefit of using this method is that low mass resolution can be used which relaxes the requirements of the magnetic field controller particularly over the large mass ranges required for chemical analyses. Even if an isobaric interference cannot be removed by high mass resolution or energy filtering, a correction can be applied to the peak of interest provided another isotope of the interfering element can be measured. For example, the intensities of 50V+ and 50Cr+ under the 50Ti+ peak can be estimated by monitoring 51V+ and 52Cr+ respectively. In both of these cases, the monitor isotope is much larger than the interfering isotope and is free of isobaric interferences at 10,000 R and so an accurate correction can be made. Of course, when the monitor isotope is smaller than the interference then there can be a substantial error propagated to the corrected ratio. Apart from interferences which are the result of coincident masses, interferences can also result from different charge states. For ion microprobe analysis the only 10 TREVOR R. IRELAND readily detected species are doubly charged atoms of elements with high ionization efficiencies, for example Mg, Ca, and Sr. Doubly charged Ca ions can interfere with Mg+, and Sr2+ can interfere with Ca+. In the latter case three Sr isotopes can interfere with Ca, 84Sr2+ with 42Ca+, 86Sr2+ with 43Ca+ and 88Sr2+ with 44Ca+ with quite different mass resolution requirements. The doubly charged species have one attribute that makes them relatively easy to detect - doubly charged odd-number isotopes occur at half masses and are therefore free of atomic interferences. The Sr2+ problem can therefore be readily monitored since 87Sr2+ occurs at mass 43.5 and hence is free of atomic isobaric interferences from Ca, although the intensity of 87Rb+ must also be estimated as well (from 85Rb+).

B. Ionization Modeling

Even though all isobaric interferences may be removed from the isotopes of an element, the measured isotopic abundances generally differ in detail from those measured in terrestrial materials by conventional mass spectrometry. This variation is due to mass dependent fractionation since the variation in the abundances is dependent on mass and can be modeled by a discrete function. Slodzian et al. (1980) observed that mass fractionation always enhances the abundance of the lighter isotope, the magnitude of the fractionation being proportional to the mass difference, and is a function of secondary ion energy. Shimizu and Hart (1982b) noted that for a given element, the mass fractionation is also matrix dependent in that it varies according to the chemistry and structure of the target. They also found that fractionation was a function of the spot alignment with the secondary ion extraction axis, and thus instrumental parameters could be critical in reproducibility of results. Shapiro et al. (1985) carried out a computer simulation of Ar sputtering of a Cu matrix and noted an enhancement of the lighter isotope in the ejected material relative to the isotopic composition of the target. Because the lighter 63Cu atoms transfer energy more efficiently to other light atoms, a small preference then arises for ejection of the light isotope. Moreover Shapiro et al. found that the 63Cu atoms carried a larger proportion of the energy back towards the sample surface than expected by proportion which further enhanced the discrimination. The fractionation also showed an angular dependence with the lightest component being ejected normal to the surface. The computer simulations therefore effectively show that isotopic mass fractionation can be expected from the nature of the collisions within the target. In chemical analysis, where isotopes of two different elements are ratioed, the variations with analytical conditions can be even more extreme. This is the most problematic aspect of ion microprobe measurements and has severely restricted its general application to quantitative chemical analysis. In modeling the sputtering process both theoretical aspects, involving the interaction between the sample and primary ion beam, and observational data, which are the manifestation of the sputtering process, are utilized. Theoretical models of the sputtering process can be Ion Microprobe Mass Spectrometry 11 constructed using complex computer simulations but the critical dependence of ionization efficiencies of both isotopes and elements with measurement conditions must be incorporated into any successful model of the sputtering phenomenon. However, it should be noted that most models have been formulated to take only a limited number of observations into account and therefore their success is only gauged on a limited number of parameters. Of particular importance in the development of ionization models for ion microprobe mass spectrometry is the local thermodynamic equilibrium (LTE) model of Andersen and Hinthorne (1973) who assumed the sputtering region was a dense plasma. They used a simple thermionic emission model to explain the effects of Cs and O on the electronic work function of the sample and argued that the work function affects only the absolute ion yield and not the relative ion yield of the individual elements. A sample could be characterized by measuring two elements of known concentration, and the concentrations of other elements could then be determined from the measured ionic ratios. This model produced results that were accurate to within a factor of 2 (Figure 6).

Figure 6. The local thermodynamic equilibrium model of Andersen and Hinthorne (1973) was able to predict the intensities of atomic species within a factor of two by normalizing to the intensity of only two peaks. However, in detail the model can not be supported according to the nature of the physical interactions at the sputtering site. No fully quantitative model exists that can be used to estimate elemental abundances from ionic intensities to the required accuracy levels. 12 TREVOR R. IRELAND

While Okuyama and Fujimoto (1986) found that the primary ion currents are sufficient to melt submicron Cr needles, thermal equilibrium models are problematic in that equilibrium conditions require on the order of 100 collisions and yet emission takes place after only several collisions (Harrison, 1983). Even though the physical processes do not directly support the LTE model, its success in determining elemental concentrations has led to the general use of the Saha-type equations in other quantitative models (Werner, 1980). A successful general theory of sputtering must be consistent with the form of the Saha equations although an equilibrium process is not necessarily implied. Other models have been developed from the observation of mass dependent mass fractionation. Slodzian et al. (1980) produced a qualitative model of ionization which incorporated the chemical bonding of the target and was dependent on relative velocities of the atoms in the matrix. Ionization was assumed to take place during the last collision that ejected the atom from the surface and hence matrix effects are readily appreciated. Shimizu and Hart (1982b) attempted to model fractionation data from pure metals with the quantum-mechanical based theory of Schroeer et al. (1973). This model assumes that ionization takes place after sputtering above the sample surface, and the observed fractionations were consistent with the formulation. Gnaser and Hutcheon (1988) examined the fractionation produced in a wide variety of polycomponent substrates for elements ranging in mass from Li to Ti and found that the isotopic mass fractionation is inversely proportional to the emission velocity of the ion but approaches a constant value for all samples at low velocities. These observations were largely consistent with the bond-breaking model initially proposed by Slodzian (1975) which addresses secondary ion emission from substrates with some ionic character. This model is an adaptation of the Landau - Zener - Stueckelberg model for ion-pair dissociation in the gas phase and in this case the ionization probability for an ejected particle depends on the distance between the ionized and neutral potential energy surfaces for that particle. By analogy with molecular dissociation, interactions out to ion - surface distances of 5 to 10 Å are important.

C. Secondary Ionization Theory

A large number of features must be accommodated into a general theory of secondary ionization including (1) the increased ionization probability for sputtering with O or Cs, (2) the partitioning of energy between the atomic and different molecules of a given element, (3) the matrix dependence of the ion yield which also indicates interaction between the surface and the ionization site. Beyond this are the isotopic characteristics of a given element, for example, the isotopic mass fractionation and its matrix dependence. As yet only components of a generalized theory are available. However a complete understanding is not specifically required for the application of the ion microprobe to problems in analytical geochemistry. Rather the characteristics of the secondary ions can be used to design recipes for Ion Microprobe Mass Spectrometry 13 isotopic and elemental abundance measurements. In this regard, applications in the earth sciences are based almost exclusively on relative measurements between a standard of known composition and identical mineralogy to the unknown. Then the relative ionization probability of standard and unknown can be assumed to be similar. A number of different techniques can be used to achieve the same end, notably the use of high mass resolution or energy filtering to remove isobaric interferences. In some part the different approaches are used to advantage on different ion microprobes, yet to a satisfactory degree, the results can be quite consistent.

III.!HARDWARE A. General Features

Benninghoven et al. (1987) have documented the historical development of ion microprobes and described the first generation machines. The first ion microprobes were recognized as potentially the ultimate weapon of the geoscientist (Lovering, 1975) but the difficulties in obtaining useful chemical data and isotopic ratios at low mass resolution and low sensitivity proved insurmountable to a large extent (see for example Williams et al., 1983). The latest generation of ion microprobes give high sensitivity at high mass resolution allowing high-precision (permil) measurements to be made with minimal corrections for isobaric interferences. It is with these machines that the future of secondary ion mass spectrometry in the geosciences lies. There are a large number of ion microprobes in use throughout the world but only a few are currently involved primarily with research in the geosciences. The CAMECA ims-3f and derivative 4f and 5f are currently the most widely used but these instruments do not possess the high sensitivity capabilities of the SHRIMP or CAMECA 1270. The benefits of the larger magnetic analyzer are readily apparent in U-Pb isotopic analyses and on this basis the latest generation of commercial ion microprobes are much larger and capable of high sensitivity at high mass resolution. In the following sections the capabilities and features of the CAMECA and SHRIMP ion microprobes will be described in some detail. It should be emphasized that while comparisons between the instruments are sometimes given, this article should not be used as a recommendation of one over the other. Designs are changing rapidly and additional features will almost certainly be available by the time of publication of this article. More limited information is given for the VG ISOLAB 120 since details of this instrument have not been forthcoming.

B. CAMECA The first CAMECA model, the ims-300 was designed primarily as an ion microscope for use in the materials industry to obtain images of the distribution of 14 TREVOR R. IRELAND

Figure 7. Schematic diagram of the CAMECA ims-3f. elements and to obtain elemental abundance information with depth profiles. The second generation machine, the 3f, did find its way into some research institutes for analysis of geological materials with the promise of in situ analysis of isotopic ratios and high sensitivity trace element abundance measurements. However, it was soon evident that there were shortfalls with the standard instrument. Specifically, the “off-the-shelf” 3f was incapable of isotopic analysis to the precision levels required for useful terrestrial applications. The imaging capabilities of the CAMECA series is one of their most important attributes but in order to retain image information compromises must be made in terms of transmission. However, with perseverance, the CAMECA ims-3f has turned out to be a highly capable and productive instrument for some specific applications. This is no better exemplified than by the modified CAMECA ims-3f at Washington University in St Louis, specific details of which are contained in unpublished PhD theses by McKeegan (1987) and Fahey (1988). Lepareur (1980) has described in detail the design of the CAMECA ims-3f; a schematic diagram is shown in Figure 7. Two ion sources are available at the head of the primary column: a duoplasmatron for ionization of source gases on one channel and a Cs gun on the other. The Cs gun produces Cs+ ions by thermally Ion Microprobe Mass Spectrometry 15 ionizing Cs at a tungsten fritt whereas the duoplasmatron ionizes the source gas from a cold-hollow-cathode plasma discharge. Either positive or negative ions can be extracted from the duoplasmatron depending on the polarity of the extraction cone; the potential of the extraction cone is generally operated in the range of 10-15 kV. An ion beam from either source can be selected by changing the current applied to the magnetic prism. The magnetic prism also serves to remove contaminants from the primary ion beam such as H and N-bearing species which might otherwise produce unwanted interferences in the secondary ion mass spectrum. The primary column ion-optical array consists of three electrostatic einzel lenses with deflection plates to focus and steer the beam to the sample surface. An octupole stigmator lens array is used to remove aberrations and change the shape of the beam at the sample surface. The primary beam is incident on the sample at an angle of 60˚. The sample surface is held at 4.5kV with respect to ground, the polarity dependent on whether positive or negative ions are being analyzed. The secondary ions are accelerated through a field gradient of 9 ¥ 105 V/m towards the grounded extraction plate, which acts like a divergent lens, forming a virtual image of the sample surface and a virtual crossover. The crossover reflects the radial energy distribution of the secondary ions with the distance from the optical axis being proportional to Esin2q where E is the kinetic energy of the ion and q is the angle of emission relative to the sample normal. The transfer lens system in the CAMECA 3f acts to transport the image of the crossover and sample surface into a field-free region, and to form real, magnified images of crossover and sample surface in the plane of the contrast aperture and the field aperture respectively. A continuous range of magnifications can be achieved through changing the potentials applied to the transfer lenses. However, as the magnification of the sample surface increases, the magnification of the crossover decreases and the spatial resolution of the sample image degrades, and so in practice a compromise must be attained depending on the application involved. For high sensitivity work e.g. trace element analysis, a 75 µm imaged field provides a smaller crossover and hence higher sensitivity. For isotopic analyses, a 150 µm imaged field produces less aberrations of the sample surface which facilitates high mass resolution measurements. The entrance slit to the mass spectrometer is located in the focal plane of the crossover. The contrast aperture restricts the diameter of the crossover thereby reducing spherical aberrations and improving the spatial resolution of the ion image of the sample (at the expense of sensitivity). The entrance slit is closed down to produce a vertical line image of the crossover (Figure 8) and hence for high mass resolution a large fraction of the total secondary ion beam is excluded. The contrast aperture consists of a sliding plate with different diameter holes that can be moved relative to the beam in both directions perpendicular to the beam path. The field aperture is positioned near the chromatic focal plane of the electrostatic analyzer which is coincident with the position of the real image of the sample surface. Again this aperture consists of a sliding plate with different hole diameters and allows 16 TREVOR R. IRELAND

Figure 8. CAMECA ims-3f entrance-slit images of 29Si+ and 28SiH+ following closure of the entrance slit to form the vertical line images; from Fahey (1988). The two species are separated by 0.0083 amu. transmission of ions from a selected area on the sample i.e. it allows for masking of unwanted regions of the sample surface. The mass analyzer is a double focusing mass spectrometer consisting of an electrostatic analyzer with turning radius of 17.3 cm and a magnetic prism with turning radius 12.7 cm. For a given charge to mass ratio (q/m) a real image of the crossover is focused at the exit slit which is also focused in terms of energy. In addition, the mass analyzer is designed to transport images through an electrostatic lens that couples the electrostatic analyzer (ESA) and magnet. The ESA consists of two electrodes cut from concentric spheres with inner and outer plates held at potentials such that the central beam path is at 0 V with respect to ground and the electric field between the plates accelerates ions of 4.5 kV (nominally) along the central beam path. The ESA forms an image of the crossover that is radially dispersed in energy; the energy slit at this image point allows for the selection of an energy window that is variable up to 130 eV wide. The energy slit is therefore useful for reducing chromatic aberrations for high mass resolution measurements and also for defining an energy bandwidth used in the energy filtering technique. Following the magnetic prism, a set of deflectors and stigmators can be used to align the image of the crossover with the image of the sample and also align the image of the entrance slit with the exit slit. The exit slit is closed down to mask unwanted species from the detectors. Two projector lenses can be operated to transform the virtual image produced in the magnetic prism to a real image that can Ion Microprobe Mass Spectrometry 17 be displayed on a fluorescent screen, or an electrostatic lens can be used to deflect the beam into a Faraday cup or electron multiplier for quantitative analyses. The CAMECA ims-4f has the same basic configuration as the 3f with several additional components. A normal incidence electron beam is designed to produce a cloud of low energy electrons just above the sample surface in order to compensate for sample charging during positive ion bombardment (usually Cs+) and negative ion extraction (including electrons). The charge compensation automatically adjusts to the sample charging with sufficient electrons extracted to neutralize the charge and excess electrons reflected from the surface. In this method electrons are accelerated from a filament and focused with a pair of quadrupole lenses and the resulting beam is introduced to the secondary ion extraction array with an electrostatic prism. The presence of the prism necessarily also affects the secondary ion trajectories and so a set of compensating deflectors are included in the secondary ion optical array to compensate for the perturbations. This system has been shown to work insofar as stable negative secondary-ion beams are produced, however it has not proved so successful in isotopic measurements. In particular, oxygen isotope fractionation is critically dependent on the position of the electron beam within the spot and variations in O-isotope ratios of several percent are evident (E. Zinner, pers. commun.). However, recent reports from CAMECA suggest that this problem may be solved by focusing the electron beam according to the image of the crossover. Full charge compensation is achieved with an evenly illuminated image and at this point reproducible O-isotopic mass fractionation can be measured. However, a full report of this procedure with measurements is not available as yet. The CAMECA ims-4f also includes a dynamic transfer system which is used in the raster imaging mode to minimize aberrations caused by secondary ion extraction away from the optimal central axis. An electrostatic deflection system is linked to the raster generator to recenter the trajectory of any extracted ions with the field aperture. In this way it appears to the secondary ion extraction system that the secondary ions are always coming from the focal point of the extraction system and hence relatively large areas can be imaged while maintaining high mass resolution. The CAMECA ims-5f is the latest version in the "small" CAMECA series. The basic configuration is the same but it includes a separate primary column for a liquid metal source that is used for producing secondary electron and secondary ion images. The 5f also has a dedicated programmable interface for controlling all lens settings so that tuning conditions for different applications can be recorded and recalled. The latest ion microprobe to be marketed by CAMECA is the ims-1270 (Figure 9) and the first instrument of this type has been delivered to the University of California at Los Angeles. This instrument incorporates a completely redesigned mass analyzer that has the imaging capabilities of the 3f-4f-5f series as well as a high transmission mode whereby beam transport (i.e. sensitivity) can be maximized at the expense of the imaging. The interrelationships between these different 18 TREVOR R. IRELAND

Figure 9. Schematic diagram of the CAMECA ims-1270. components and operational conditions make for an extremely complex set of lens configurations and so all lens control is through the dedicated interface developed for the ims-5f. The 1270 mass analyzer has large ESA (RESA = 585 mm) and magnet (Rm = 585 mm) turning radii with a magnification from unity to 5 times and a mass dispersion relationship of

h = 1.214 ¥ 106 DM/M where h is the perpendicular distance in micrometers at the collector between adjacent masses of M and M+DM. The primary column and source housing of the 1270 analyzer are similar to the 3f-4f-5f series and include a Cs source and a duoplasmatron. The sample interlock on the UCLA machine has been modified so that there are two stages of pumping to restrict contamination of the source chamber and maintain ultrahigh vacuum (≤5¥10-10 Torr); three samples can be loaded into the intermediate vacuum chamber at any time. The secondary ion extraction system and transfer lenses are similar to the earlier series as well. Besides the large turning radii of the ESA and magnet, there are a number of additional lens components to control the function of the mass analyzer. In the ion microscope mode, image quality is paramount and so lenses are operated to achieve the lowest possible spherical and chromatic Ion Microprobe Mass Spectrometry 19 aberrations in the imaging plane of the channel plate. Two circular coupling lenses are situated between the ESA and magnet to ensure optimal transfer of ions. This mode can also be utilized in an ion microprobe mode, but to achieve the highest possible sensitivity at high mass resolution a beam matching system operating in the XY mode allows for maximum transmission of ions at the expense of image information. In this mode, two slit einzel lenses provide a 1:5 magnification of the field aperture in the magnet while the entance slit image is magnified by a factor of 5; in this way the b aberration of the magnet is reduced. Slit einzel lenses are also present before and after the magnet to compensate for aberrations produced in the electrostatic peak-switching mode. In addition to these lens systems, sextupole lenses are positioned before and after the ESA and magnet to allow for maximum control of the transmission characteristics. Projection lenses are situated after the exit slit to transfer real images of the crossovers to an imaging plane containing the channel plate detector. The UCLA ims-1270 is currently undergoing specification testing and preliminary tests indicate that the Pb sensitivity of this machine is at least as good as SHRIMP I under similar operating conditions achieving 10 cps/ppm Pb/nA in zircon (Schumacher et al., 1993).

C. SHRIMP

The Sensitive High mass Resolution Ion MicroProbe, SHRIMP, was designed and constructed at the Australian National University with the purpose of analyzing geological materials with sufficiently high mass resolution to eliminate most major isobaric interferences while maintaining the highest possible sensitivity (Clement et al., 1977). This is achieved by a physically large mass analyzer (magnet turning radius of 100 cm) which gives high mass dispersion and therefore allows for high mass resolution operation with wide slits. In addition the secondary ion beam is matched in terms of its phase space characteristics with the acceptance of the magnetic analyzer to maximize the secondary ion beam transmission. This operation is at the expense of the ion optical image of the surface and so SHRIMP cannot act as a direct imaging ion microscope as can the CAMECA ion microscopes. A second ion microprobe of similar design (SHRIMP II) has now been produced as a commercial prototype. A schematic of SHRIMP I is shown in Figure 10. The SHRIMP I ion source is a cold-hollow-cathode duoplasmatron which lies on the primary beam central ray path. The essential elements of the ion-optical array are a Wien velocity filter, deflection plates and three electrostatic einzel lenses. The Wien filter consists of a crossed electromagnetic field that selects ions according to their velocities. It has 16 - been found that O2 at mass 32 is the most effective primary species on SHRIMP I since it has the same intensity as 16O– but the larger mass of the molecular species – creates a larger secondary ion signal. The use of O2 may also lead to a higher oxygen concentration at the sputtering site, hence producing a larger ionization yield. The 20 TREVOR R. IRELAND

Figure 10. (a) Schematic diagram of SHRIMP with (b) expanded view of source chamber and primary column.

SHRIMP einzel lenses generally run in a configuration which has been called Kohler illumination and relies on the insertion of an aperture between EL1 and EL2 at the focal point of EL2 (Figure 11). The aperture acts as the object for EL2 and the spot size is a demagnified image of that aperture alone. Kohler illumination has the major advantage that irregularities in the brightness of the source and aberrations upstream from the aperture do not cause irregularities in the final spot because the object for EL2 is in its focal plane and hence the rays emerge as a parallel beam. Several craters produced by SHRIMP I in Kohler illumination are shown in Figure 12. The spot size does not vary with primary beam current as is the case for critical illumination where the spot is a demagnified image of the extraction aperture. An einzel lens is situated after the primary extraction cone and allows the current delivered to the kohler aperture to be continuously varied. Drawbacks for this method are that the apertures are fixed in size and a continuously variable spot size is not available, and that the apertures burn out in a relatively short period of time (approximately 1 week continuous operation) and must be replaced regularly. Four of these apertures are held in a sliding plate that can be removed and the apertures changed during the routine servicing of the duoplasmatron. The primary beam has an incidence angle of 45˚ with respect to the sample surface yielding Ion Microprobe Mass Spectrometry 21

Figure 11. Primary column focusing modes on SHRIMP. Critical illumination relies on the demagnification of the source aperture by the two Einzel lenses EL1 and EL2; the greater the demagnification the smaller the spot and the lower the beam intensity. In Kohler illumination, the Kohler Einzel lens is activated to transfer the maximum beam intensity to the Kohler aperture which acts as the source for the final lens. In this configuration, the spot diameter is a function of the Kohler aperture diameter and aberration contributions are limited to those from the final lens only. Primary beam intensity can be controlled by the strength of the Kohler lens and does not change the spot diameter. distinctly elliptical craters. It should be noted that this method of Kohler illumina- tion can also be achieved with the CAMECA primary column. The sample is electrically isolated allowing the primary-beam flux to be measured at the sample. A 450 V bias, with respect to the first extraction electrode, then accelerates the secondary ions towards an intermediate electrode and the extraction aperture through a relatively low electric field. The extraction potential for positive secondary ions is approximately 10 kV. The secondary-ion optical array consists of the extraction system, the intermediate lens, and the phase-space- matching system. The extraction system is a simple pair of electrode tube lenses that produces a non-magnified image in the back focal plane of the intermediate einzel, which then in turn transfers the beam with unity magnification to the phase-space- matching system. This system is made up of three slit einzel lenses, two active in the XY plane and one in the XZ plane which are capable of steering and focusing the beam so as to match the acceptance of the mass analyzer. Whereas in the CAMECA 3f the crossover is masked to yield a line image, in SHRIMP I the beam is focused to a line image at the entrance slit thereby minimizing beam loss. A fraction of the secondary ion beam is monitored at an electrically isolated and suppressed aperture 22 TREVOR R. IRELAND

Figure 12. Ion probe crater produced by Kohler illumination. This method facilitates the production of steep sided evenly illuminated craters. Target is a sulfide with considerable S-isotopic heterogeneity as indicated by labelled d34S values (photo from C. S. Eldridge). between the intermediate einzel lens and the matching system. The aperture is approximately 3 mm in diameter and allows most (≥90 %) of the secondary ions through, but the remainder is collected as being representative of the total secondary beam. During data collection, the counts on each isotope can be divided by the counts on the secondary beam monitor during the same time interval. This procedure approximates to a double-collecting mode of operation and can result in significant improvements in precision by reducing the effects of instability in the primary beam. The double-focussing mass analyzer is based on a mass spectrometer design by Matsuda (1974) and the main ion-optical elements are a cylindrical 85˚ ESA, a quadrupole, and a 72.5˚ sector magnet with non-normal entry. The physical dimensions of the mass analyzer have been made as large as possible (Rm = 100 cm; RESA = 127.2 cm) in order to achieve high mass resolution while maintaining relatively wide slits to allow high sensitivity. The magnification of the mass analyzer is 0.4 and the mass dispersion relationship is

h = 7.73 ¥ 105 (DM/M) Ion Microprobe Mass Spectrometry 23 where h is the perpendicular distance in micrometers at the collector between adjacent masses of M and M+DM. An intermediate image point exists within the ESA which allows definition of energy limits within a bandwidth of 150 eV in 25 keV. Adjustable slits are also present at the entrance and exit of the ESA to limit beam divergence; the acceptance of the ESA is limited to <0.01 radians. A retractable Faraday cup, the post-ESA monitor, is positioned at the exit of the ESA and, in conjunction with the divergence slits, allows accurate alignment of the secondary beam along the central ray path of the ESA. During initial set-up routines, the post-ESA monitor can be used to optimize the secondary beam to the magnet without the necessity of a peak being accurately aligned through the collector slit. An optimum focal point exists in the plane of the mass analyzer for both angular and energy refocusing. Rotation of the collector slit and collector cradle, and in-line translation of the cradle are available to optimize the focal position. The focal point of the SHRIMP I mass analyzer has been found to be field-dependent with, for example, the optimum focus of 46Ti being 6 mm nearer the magnet than that for 50Ti (Ireland, 1986). A compromise position optimizing the peak shape of 48Ti produced significant degrading of the peak shapes of 46Ti and 50Ti, and ratios collected under these conditions showed an instrumental bias of approximately - 0.1% in the fractionation-corrected 46Ti/49Ti and 50Ti/49Ti ratios. The exact cause of this condition is not clear but in order to remedy its effects, the entire collector assembly is under computer control and a separate position can be assigned for each mass. SHRIMP I was originally fitted with a single collector assembly capable of receiving ions in a Faraday cup or electron multiplier. A sliding base-plate on the cup allowed the ion beam to pass to the ion counter. The single collector has since been replaced with a multiple collector housing that is designed to enable operation under a single-collecting mode or multiple-collection of up to eight masses. The individual ion beams can be centered on their respective collector slits by pre-slit deflection-plates, while post-slit deflection-plates optimize transmission to the electron multipliers. A Faraday cup is available on the central ray and can be switched in and out of the beam path. The collector slit widths for all but the central position are fixed; there is a collector-slit bar on the central ray that has three set slit- widths. For routine analysis of zircons, SHRIMP I operates in the single collector mode. The new commercial SHRIMP II uses the same mass analyzer configuration as SHRIMP I but there are substantial modifications particularly in the primary column - source chamber designs. The sample interlock can hold up to four samples and two samples can be in the source chamber at any time. The source elements of the primary column were manufactured by Cambridge Mass Spectrometry Ltd. The duoplasmatron operates with air cooling and a fixed magnetic field and is designed to optimize the ion transmission by producing a low divergence beam. The primary mode of operation of the primary column is kohler illumination and the demagni- 24 TREVOR R. IRELAND fication of the final lens element is ≈7 and so a final spot size of the order of 5 µm is produced for a 30 µm Kohler aperture. The secondary ion extraction system for SHRIMP II is similar to SHRIMP I and consists of an intermediate extraction electrode and transfer system to produce an image of the crossover at the focal point of the beam matching system. The beam matching system consists of three quadrupole lenses which have lower aberration characteristics than the slit einzels in use with SHRIMP I. The mass analyzer is based on the same design parameters as SHRIMP I although the the sensitivity of Pb+ in zircons has been measured at over 20 c/s/ppm!Pb/nA under routine analytical conditions which is superior to the 5-10 c/s/ppm!Pb/nA obtained with SHRIMP I (I. Williams, pers. commun.).

D. VG Isolab 120

The third of the large commercial instruments is the VG Isolab 120. A commercial prototype has been installed at Cambridge University, England and is now under operation although a detailed description of the machine and analytical capabilities is not available as yet. The configuration is similar to SHRIMP with a few notable differences. The source can operate with a thermal ionization source as well as the primary column used for ion microprobe analysis and ports are available for laser photoionization of sputtered or thermally produced neutrals. The mass analyzer is double focusing, with a focal plane normal to the ion optic axis, and features a 96 cm radius electrostatic analyzer with a 70˚ angle followed by a 60 cm radius magnet with an effective dispersion length of 120 cm. An energy band pass can be defined through slits at the intermediate energy focus between the ESA and magnet. The multiple collector has a microchannel plate as well as adjustable Faraday collectors. Very high abundance sensitivity (up to order of 10-11) is provided by a second electrostatic analyzer following the magnetic sector with a second collector assmbly consisting of a Daly/photomultiplier combination.

IV. METHODOLOGY A. Isotopic analysis

The ion microprobe is essentially a mass spectrometer and so the data it generates are treated in much the same manner as data from a conventional thermal ionization or gas mass spectrometer. Ion microprobes are generally fitted with both a Faraday cup and electron multiplier detection systems, however, the intensity of secondary ion beams is generally insufficient to warrant measurement by a Faraday cup, and isotopic measurements are typically made with the multipliers which have much lower noise levels, although they are limited to some extent by dead time uncertain- ties. Deadtime is that period following the triggering of the system by an incoming ion when any subsequent incoming ion will not retrigger the system For electron Ion Microprobe Mass Spectrometry 25 multipliers this time is of the order of 5 ns but the total deadtime of the counting system as a whole is typically around 20 ns. The total deadtime is affected by the response time of the discriminator and scalar used in the counting system , as well as the gain of the multiplier. In addition, Zinner et al. (1986a) have shown that pulse charge distributions (and hence deadtimes) in electron multipliers are a strong function of ion type (atomic vs. molecular species), element, polarity, and energy. Small but noticeable effects were also observed in the detection efficiencies for isotopes of Ca and Ti which introduced a small mass fractionation effect. The precision of a pulse-counted measurement is limited by the total number of counts, N, on the smallest peak to 1/√N. For example, to achieve 1 ‰ (10-3) precision in the 13C/ 12C ratio, 106 counts of 13C must be collected. Therefore the larger the number of total ions the higher the precision and the more intense the secondary signal the shorter the time period that is needed to obtain the required number of ions. However, in isotopic measurements at high count rates, the uncertainty in the dead time can ultimately be the limiting factor for the accuracy of a measurement. For a retriggerable system, the measured count rate, cmeas, is related to the actual count rate, ctrue, by

-tctrue cmeas = ctruee where t is the counting system dead time. The dead time correction can therefore be approximated by

c c †exp c exp c exp c exp c true ª meas ( meast ⋅ ( meast ⋅ ( meast ⋅ ( meast)))) For accurate isotopic measurements, the dead time must be accurately known and independent of count rate as is illustrated in Figure 13. Hayes and Schoeller (1977) have addressed in some detail the limits in precision† and accuracy attainable by pulse counting. When ratios differ significantly from unity, the uncertainty in the dead time determination, and not simply the magnitude of the dead time, becomes the limiting factor. Specifically,

2 2 -1 2 2 uN = (sN/N) ≈ (ƒt) + ƒ sr where sN is the coefficient of variation of N, the number of counts, ƒ is the frequency of events, t is the collection time, and sr is the standard deviation of r, the deadtime. Ignoring the uncertainty in the dead time, the relative standard deviation is simply 1/√ƒt, the total number of counts. However, when sr is non zero, the second term contributes to the relative standard deviation in a way that is dependent on the count rate. In this way Hayes and Schoeller (1977) found that the maximum count rate Fmax for any measurement is given by

Fmax = (uN)reqd/sr = sN/Nsr 26 TREVOR R. IRELAND

Figure 13. The counting system dead time is typically around 20 ns for an electron multiplier based system and should be independent of count rate as illustrated in this figure from Fahey et al. (1987b). where (uN)reqd is the required precision of the analysis. Hence for a 1‰ measurement of 13C/ 12C and a 2 ns error in the dead time, the maximum count rate is 5 ¥ 105 c/s and the minimum time for the measurement is 565 seconds, or around 10 minutes. In order to obtain a factor of ten improvement in precision, the maximum count rate must be decreased by a factor of 10 but the minimum measurement time is dependent on the inverse third power of sr and so the counting time must be increased by a factor of 1000. For ion microprobe analysis, it is often not possible to sustain analyses over such a period of time since the sample may be consumed or the crater become too deep for reliable analysis. For isotopic analysis the magnet is cyclically stepped through the peaks of interest. The counts are measured on the top of each peak and a number of such cycles are combined to give the primary ratio information for the analysis. It is important to combine a number of cycles in order to address the degree of temporal change in peak heights during analysis. The causes of temporal drift may be primary beam fluctuations or of more concern, change in the sputtered composition of the sample whether it be related to real compositional changes or related to a change in the sputtering/ionization conditions. The stability of the primary beam is an important factor in ion probe analysis since any noise from this source is propagated through to the secondary ion beam. Ion Microprobe Mass Spectrometry 27

Short-term noise simply expresses itself as an additional random contribution to the secondary-beam noise but longer-term drifts can cause more problematic variations which are of the same type as compositional variations. Such secular drifts must be corrected by a method of interpolation. The denominator peak is usually chosen to monitor the fluctuations in the medium to long-term signals and a typical means of temporal correction is simply to assume that there is a linear change in signal between one denominator peak measurement and the next. Clearly this is applicable so long as the fluctuations are of significantly longer time period than the measurement cycle. Correction of medium term fluctuations is quite difficult since at some stage the real variations must be separated from the actual beam noise. Such fluctuations might be addressed by attempting to fit a polynomial to the denominator peak, or by using two peaks to monitor beam fluctuations. Quite often medium term fluctuations are symptomatic of some hardware disorder and time is often better invested in trying to solve the problem than making elaborate algebraic corrections to the data set. When isobaric interferences cannot be resolved they can be stripped from the peak of interest by monitoring another isotope of the interference. This correction can be made from the final time-interpolated ratios, but if the magnitude of the correction changes through the analysis then an error must be propagated through to the mean ratio measurement that reflects the change. An alternative method is to strip the interference during each cycle so that the random error associated with each measurement is propagated directly to the stripped peak in each cycle. Then if the sample composition changes, the interference is stripped from the same period of isotope measurement and the mean isotopic ratio can be calculated without propagating an error due to fluctuation of the interference. However, systematic errors might still arise because of uncertainties in the isotopic ratios of the interference element. Such uncertainties may be due to the magnitude of isotopic mass fractionation, or in the case of meteoritic samples, could be due to uncertainty in the actual isotopic composition. Such errors are minimized if a large peak can be used to monitor a smaller interference on another element. Isotopic mass fractionation is one of the fundamental properties of isotopic ratios produced by secondary ionization. Isotopic mass fractionation is also an important parameter in systems subject to physicochemical reactions in nature, particularly for light elements, and so the instrumental mass fractionation must be removed in order to ascertain the intrinsic fractionation component. Instrumental mass fractionation is known to be a function of a number of machine parameters such as secondary ion energy and the axial position of the secondary ion extraction, as well as being matrix dependent (Shimizu and Hart, 1982b). The most straightforward approach to monitoring instrumental isotopic mass fractionation is to measure terrestrial standards of the same mineralogy as the unknowns, and whose compositions can be analyzed by conventional mass spectrometric techniques. This degree of mass fractionation can then be removed from the unknowns leaving the intrinsic fractionation. However, the uncertainty in the measurement of the terrestrial 28 TREVOR R. IRELAND standards must also be propagated onto the unknowns, during the course of the analyses of the unknowns. The convention adopted in this paper is that the total (=intrinsic+ instrumental) mass fractionation of element A (in permil per amu, ‰/amu) relative to the terrestrial ratio is given by

j j i j i D A = [( A/ A)meas/( A/ A)terr – 1] ¥ 1000/(j-i) (‰/amu) where iA and jA are two isotopes of element A, and the subscripts meas and terr refer to the measured and terrestrial ratios respectively; a positive mass fractionation is defined as heavy isotope enrichment. Typically the terrestrial ratio for the fractionation normalization is taken from conventional thermal ionization analyses in order to facilitate comparisons. The intrinsic mass fractionation in an analysis of element A can be expressed as the difference between the fractionation measured in the unknown with that measured from a standard i.e.

j j FA = D Ameas – D Astandard (‰/amu) This procedure is an external calibration of isotopic mass fractionation since it depends on the measurement of another reference material to calibrate the magnitude of the effect. For an element with only two isotopes this is the only method of removing the instrumental mass fractionation effects. In general the reproducibility of the measurements on both standards and unknowns is the limiting factor and only relatively large effects (≥ 1 ‰) can be resolved. Where there are three (or more) isotopes of a given element, an internal mass fractionation correction can also be applied by using one of the isotopic ratios to remove the mass fractionation from the other ratio(s). This is carried out by adopting a mass fractionation law that describes the relative fractionation of the isotope ratios according to their masses and expressing the corrected isotopic abundances relative to the terrestrial ratio. In this case the normalization is not dependent on the measurement of a standard but rather requires only that the mass fractionation obeys a predetermined law whose form is rather arbitrary since it need only describe the dependence of the measured ratios with isotopic mass. The precision of such fractionation-corrected ratios is limited only by the analysis time, the longer the analysis the higher the precision. Consider again element A except we now include a third isotope, kA. For three isotopes we can form two ratios, for example jA/iA and kA/iA. For the moment let us consider the algebraically simplest case of mass fractionation, that is it is linearly dependent on isotopic mass. In this case we will use jA/iA to monitor the mass fractionation and to derive a correction for kA/iA so that we can express kA/iA relative to the terrestrial ratio. As above we have DkA and DjA and we remove the effects of the linear mass fractionation from DkA to leave the residual dkA which is given by: Ion Microprobe Mass Spectrometry 29

dkA = DkAmeas – (k-i)/(j-i)DjAmeas (‰) In this scheme, terrestrial standards should all have residuals within error of zero since we are simply testing the form of the mass fractionation law. A number of mass fractionation laws have been used in both conventional and ion probe mass spectrometry although they all have a similar form. In the power law, the fractionation between two adjacent isotopes is a and is the same for anyadjacent isotopes within a region of interest. If we have the terrestrial j i k j abundances ( A/ A)0 and ( A/ A)0, and j = i+1 and k = j+1, then the measured ratios of the adjacent isotopes (jA/iA) and (kA/jA) are

(jA/iA) = (jA/iA)0 (1+a) and

k j k j ( A/ A) = ( A/ A)0 (1+a) and therefore by multiplying

k i k i 2 ( A/ A) = ( A/ A)0 (1+a) In practice, (jA/iA) can be measured to obtain a and hence allow a power law correction on (kA/iA). It can be seen that if we expand the power law formulation assuming a « 1, we obtain

(kA/iA) = (kA/iA)0 (1+2a) which is the linear law approximation. Russell et al. (1978) found that a power law could not produce a satisfactory fit to the fractionated isotopic abundances of Ca and so developed the exponential mass fractionation law which has the form

j i j i b ( A/ A)/( A/ A)0 = [mj/mi] i j where mi and mj are the isotopic masses of A and A. During thermal evaporation, it can be shown that the degree of fractionation is proportional to the square root of the masses of the evaporating species such that

(jA/iA)/(jA/iA)0 = [initial jA/final jA](√mj/mi– 1) where initial jA and final jA refer to the amount of isotope jA in the reservoir initially and after evaporation respectively. This is the Rayleigh law. Esat (1984) has shown that these laws all have the same basic form. If we define i k j k the ratios Rik and Rjk, where Rik = A/ A and Rjk = A/ A, then (m -m )/m -m ) power law: Rik = (Rjk) i k j k (√m -√m )/√m -√m ) Rayleigh law: Rik = (Rjk) i k j k 30 TREVOR R. IRELAND

log(m /m )/log(m /m ) exponential law: Rik = (Rjk) i k j k from which the generalized form of the fractionation laws can be seen to be

g Rik = [Rjk]

For example, the g values for the Mg isotopes are gpower = 1.996, gRayleigh = 1.976, and gexponential = 1.957. Therefore the different laws can be seen to be simply imparting different degrees of curvature to the mass fractionation function for any given element. In practice, the real functional relationship must be determined for each element on each mass spectrometer.

Isotopic Analysis by Ion Imaging

The ion microscope mode offers the possibility of taking images of different isotopes and through digital processing, the signals can be ratioed and the isotopic composition of the area in question can be determined. Clearly for a large area there may be no particular benefit, but when the samples are small, or heterogeneous over a small spatial scale, ion imaging can have significant benefits. For example, Nittler et al. (1993a) mapped the O isotopic compositions of a large number of corundum grains by imaging in 16O and 18O. In one image some 5-15 grains could be analyzed in about 6 minutes to a precision on the order of ±40 ‰ (1s). This

Figure 14. Ion images of silicon isotopes 28Si and 30Si from a mount containing interstellar SiC grains of 3-5 µm size range. Such ion images allow the rapid identification of the highly exotic but rare grains X. These grains are characterized by an extreme overabundance of 28Si relative to the minor isotopes compared to the normal values. The exposure times of the images are such that equal intensity in 28Si and 30Si equates with a normal composition. The X grain clearly has a lower intensity in 30Si relative to 28Si. (Photomicrographs provided by L. Nittler). Ion Microprobe Mass Spectrometry 31 allowed the clear identification of a highly anomalous grain with d18O of ≈-200 ‰, which could then be analyzed to higher precision, including the measurement of d17O, in a normal peak-jumping mode. Nittler et al. (1993b) have also used this technique to identify highly anomalous but rare SiC grains X. Over a period of three days around 1250 grains were analyzed and 9 grains X located. Successive Si ion images of a region containing an X grain are shown in Figure 14. The exposure times are such that a normal composition should result in equal intensities on each photograph. The intensity of the 30Si+ image is clearly lower than that of the 28Si+ image indicating its anomalous isotopic composition enriched in 28Si. Isotopic analysis by ion imaging will of course offer only limited precision but even moderate levels can be adequate for highly anomalous meteoritic samples.

Recipes for isotopic analysis The following section gives a brief outline of the problems that are particular to the isotopic analysis of some commonly measured elements. This is meant to be an outline only and for more detailed explanations the reader is referred to the primary source(s).

Hydrogen. The isotopic measurement of hydrogen is difficult for two reasons: the D abundance is very low (0.015 % of normal H) and the fractional mass difference between D and H is the most extreme case possible (Deloule et al., 1992). Isobaric interferences are no problem and a mass resolution of about 1000 R is + + required to separate D from H2 . The terrestrial D/H ratio of Standard Mean Ocean Water (SMOW) is 0.00015576 (Hagemann et al., 1970). – The first measurements by Hinton et al. (1983) were made with an O beam with hydrogen isotopes measured as H+ and D+. Since they were attempting to measure large effects, count times were limited to give precisions of the order of 50 ‰. Zinner et al. (1983) used a Cs+ primary beam and collected negative secondary – ions. They found that this technique has the advantage of producing far less H2 and this species was less than 0.5 % of the D– signal for all samples. They also found that the isotopic mass fractionation of the hydrogen isotopes was far less than the case for positive ions. In order to minimize the effects of sample charging with the Cs beam, the samples were pressed into a gold foil. During analysis, the energy distribution of the secondary ions was monitored and the sample charging was compensated by offsetting the accelerating voltage. Terrestrial standards were reproducible to around the ±30 ‰ level. Deloule et al. (1991a) were interested in measuring H isotopes in terrestrial samples thus high precision and accuracy is imperative. They used an O– primary + + + beam and measured H and D at a mass resolution of 1300 R to separate H2 from D+. They took care to remove moisture from the sample surface by baking it in the ion probe at 120 ˚C and used a liquid nitrogen cold trap to fix residual water in the + + vacuum. Measurements commenced when the H2 /H ratio was lower than 8 ¥ 10-4. Rather than simply comparing D/H ratios of standard and unknown, 32 TREVOR R. IRELAND

Deloule et al. (1991) found that the instrumental fractionation could be further calibrated by the measurement of Si, Ca, Ti, and Mn ion intensities on the same materials. The error of the best-fit calibration is around 7 ‰ and the dD can be measured to a precision of around ± 10 ‰. An interesting aspect of this work was the observation that the crystallographic orientation of mica may be an important factor in the reproducibility of the D/H ratios. In order to minimize these effects, Deloule et al. (1991a) crushed the micas first and potted them in epoxy. The corollary to this observation is that for successful in situ measurements it may be necessary to document the crystallographic orientation and apply a correction.

Boron. Boron has two stable isotopes, 10B and 11B with a normal 10B/ 11B ratio of 0.24726 as determined by De Bièvre and Debus (1969). Chaussidon et al. (1990) measured boron isotopic compositions with an O– primary beam and B+ secondary ions at a mass resolution of ≈2000 R to eliminate 10BH+ from 11B+. Instrumental mass fractionation for B was monitored in a similar fashion to D/H measurements in that mass fractionation was found to vary linearly with the product of mass/charge and ion emissivity of the octahedral cations (Fe+Mg+Mn+Ti+Li) within a range of - 65 ‰ for Li-rich tourmalines to -60 ‰ for Mg-Fe-rich tourmalines. The reproducibility of these measurements was better than 1 ‰.

Carbon Carbon has two stable isotopes, 12C and 13C, with a 13C/ 12C of 0.011237 in Pee Dee belemnite (Craig, 1957). The main isobaric interference is 12CH+ interfering with 13C+ which requires a mass resolution of ≈3500 R. Carbon can be analyzed as either negative or positive secondaries (with Cs+ and O– as the primary beam species respectively), although C– has several advantages: it is more efficiently ionized, i.e. sensitivity is higher, and the instrumental mass fractionation is less. McKeegan et al. (1985) analyzed C isotopic compositions as C– with Cs+ bombardment. They measured a variety of terrestrial standards of different mineralogical composition in order to address the possible effects of matrix effects on instrumental mass fractionation. While the absolute mass fractionation was quite large (45-50 ‰/amu), within measurement errors it was found that there was no difference between graphite and carbonate in terms of mass fractionation. There was a slight effect for kerogen, but McKeegan et al. (1985) concluded that matrix- dependent fractionation differences between the graphite standard and any reasonable C phases did not exceed 15 ‰. Harte and Otter (1992) also used a primary Cs+, secondary C– configuration to analyze C isotopes in diamonds. They found a similar range in instrumental fractionation from -56 to -34 ‰/amu and the fractionation was also found to vary as a function of age of the electron multiplier. The typical precision of an individual analysis for their work was around 0.6 ‰/amu for a restricted area of the standard. Over larger areas, variations of up to ± 1.5 ‰/amu were apparent which was interpreted as being due to heterogeneity in the standard. The standard had been Ion Microprobe Mass Spectrometry 33 analyzed previously by conventional mass spectrometry and was thought to be heterogeneous in d13C at the ± 1 ‰ level.

Nitrogen. Nitrogen has two stable isotopes 14N and 15N with a 14N/ 15N ratio of 272.0 (Junk and Svec, 1958). Nitrogen is a difficult element to analyze by secondary ion mass spectrometry. It does not form N– secondaries and yields N+ ions a thousand times less productively than Si (Zinner et al., 1987). However, Zinner et al. (1987) noted that nitrogen in the presence of carbon formed a very intense and stable CN– beam. For the acid residue samples Zinner et al. (1987) were analyzing, BO– was a ubiquitous contaminant and required a mass resolution of 6000 R to eliminate molecular interferences from the 12C14N– and 12C15N– peaks. The CN– signal depends on the bonding between C and N as well as the absolute concentrations of C and N and so no quantitative information on concentrations is possible unless the standards are identical to the unknowns. Zinner et al. (1987) were able to reproduce the 15 d Nair in 1-hydroxybenzotriazole to within 2 ‰, in synthetic SiC crystals, and in graphite on to which air was blown from a controlled vacuum leak.

Oxygen Oxygen has three stable isotopes, 16O, 17O, and 18O. The 18O/ 16O ratio of SMOW is 0.0020052 (Baertschi, 1976), but the absolute abundance of 17O is not well known. 17 16 McKeegan (1987) used ion microprobe measurements to derive a value of O/ O = 0.00038309 when normalized to the 18O/ 16O of Baertschi (1976). This is a difficult element to analyze because of the large difference between the abundance of 16O and the minor isotopes 17O and 18O and because of the large 16OH– interference on 17O–. In meteoritic work it is essential to measure all three O isotopes whereas in terrestrial problems measurement of the 18O/ 16O will suffice because the only parameter of interest is mass dependent fractionation. Since oxygen is a strongly electronegative element it is best analyzed by sputtering with Cs+ and measuring O–, however sample charging must be overcome in order to obtain reproducible and reliable data. McKeegan (1987) first described the mounting of 10-15 µm particles on gold foil which allows charge to dissipate rapidly. This technique has also been used by Fahey et al. (1987a), Virag et al. (1991) and Ireland et al. (1992). For inclusions larger than the required 10-15 µm, they must first be crushed, transferred to the gold foil, and then pressed in with a quartz plate. A primary beam of 14.5 kV Cs+ ions is used to sputter the samples and O isotopes are measured as 4.5 kV O– secondary ions. The primary beam is defocused such that the entire particle is sputtered and for hibonite and spinel particles between 8 and 15!µm a stable 16O– count rate of 500,000 c/s can be maintained for measurement. The major isobaric interference for O-isotopic measurements is 16OH– inter-fering with 17O–. The hydroxide species is mainly sputtered from the surrounding gold mount and results not only in an increased OH– signal, but also degrades the OH– peak shape since these ions originate from a larger source area than the O– ions. 34 TREVOR R. IRELAND

The OH– is apparently due to surface water migrating on the gold; water is a residual to the vacuum and the 16OH–/17O– falls with time after sample insertion and hence pumping time. A field aperture at the image point after the source slit is used to mask the surrounding gold from the sample and to exclude as much of the 16OH– from the gold mount as possible. In doing so it is important to defocus the Cs+ beam to a diameter that is larger than the diameter of the masked region on the gold mount in order to sputter-remove the migrating water before it reaches the area from which secondary ions are accepted. A mass resolution of 6500 R is sufficient to separate 16OH– from 17O–, and possible tail contributions of 16OH– are monitored by measuring the tail on the low mass side of 16O– at a mass offset of 16(m – m ) where m and m are 17 16OH 17O 16OH 17O the masses of 16OH and 17O respectively. Under these conditions the 16OH–/17O– is typically 2, corresponding to a correction of around 0.8 ‰ on 17O, after one day of pumping in the sample chamber. Analyses of the Burma-spinel standard are interspersed with the unknowns and are used to normalize the O-isotopic composition for instrumentally-induced isotopic mass fractionation. The instrumental mass fractionation measured by Ireland et al. (1992) was within a range of –3 to –20!‰/amu. The use of the Burma-spinel standard for the normalization of O-isotopic analyses of other oxide minerals, such as hibonite, has been examined by McKeegan (1987). He found no systematic differences between the instrumental mass-fractionation measured in terrestrial samples of hibonite and spinel that had previously been measured by conventional mass-spectrometric techniques and therefore a single standard could be used in the O-isotopic measurements. The O-isotopic compositions measured from the meteoritic samples are referenced to the O-isotopic composition of Baertschi (1976) and normalized to the mean oxygen composition measured from the Burma-spinel standard in the same 18 17 analytical session; the Burma spinel has d OSMOW of +22.1 ‰ and d OSMOW of 18 +11.6 ‰. While individual analyses have a precision in d OSMOW of around 3 ‰, the scatter of measurements on individual Burma-spinel grains in a session is typically around 8!‰. This variation, which limits the overall precision of the method, is mass-dependent fractionation since the Burma spinel data scatter along a slope 1/2 mass fractionation line on an oxygen three-isotope plot while deviations from this line are consistent within individual measurement errors. The reason for the variation in isotopic mass-fractionation is unclear, but is probably related to the irregular geometry of the grains producing variations in the isotopic-density distribution at the source slit of the mass spectrometer. In meteoritic work, it has been found that the 16O excess is an important parameter since large anomalies are found in the O-isotopic compositions of meteoritic materials. The anomaly is attributed to the 16O abundance since the 17O/ 18O ratio is generally close to terrestrial. The 16O excess is calculated according to Clayton and Mayeda (1983), Ion Microprobe Mass Spectrometry 35

0. 52 1 16O = d18O - d17O excess (1 - 0.52) (1 - 0. 52) For the graphical representation of the O-isotopic measurements individual measurement errors on the meteoritic samples are combined with the variation in instrumental mass fractionation as determined by the spread of the Burma spinel 1 data points along the slope /2 mass fractionation line (Virag et al., 1991). This 1 spread is expressed by an error ellipse whose main axis has a slope of /2 and whose axes are the standard deviations of the standards along the mass fractionation line and normal to it. Folding of this ellipse with the error ellipse representing individual measurement errors (in d17O and d18O) yields a final error ellipse whose axes are generally neither aligned with the coordinate axes nor aligned with the fractionation line and whose orientation is different for different data points. The method for oxygen isotope analysis described above is somewhat unsatisfactory because in situ analyses of insulators in polished thin sections are not possible. Material must first be removed from the thin section and then mounted in gold. The alternative is to use a charge compensation gun which neutralizes the charge build up at the surface of the sample. Lorin et al. (1989; see also Lorin, 1990) used such a charge compensation scheme to measure O-isotopic compositions of Allende refractory inclusions. This method has high sensitivity, up to 1 % of the O atoms are ionized and collected, and mass discrimination is modest, around 3.5 ‰/amu. However, the primary beam current used was less than 0.01 nA and so the sputtering rate is very low and extremely long counting times are required to achieve the cited analytical precision of 0.6 ‰ and 1.2 ‰ in 18O/ 16O and 17O/ 16O respectively. Hervig and Steele (1992; see also Hervig, 1992) describe a technique which uses charge compensation and extreme energy filtering (only ions with energies greater than 300 eV ion are collected) to discriminate against molecular interferences and reduce matrix effects. Precision for this technique is around ±2 ‰ for the ≈1.5 hour analysis times. The unsatisfactory aspect of this technique is the high loss of ion transmission with energy filtering at such high energies and therefore the loss of sensitivity. The problems of sample charging under Cs bombardment can be overcome by analyzing conductors and O isotopic compositions can be readily measured to high precision. Read et al. (1990) used a VG Isolab 54 ion microprobe (Lyon and Turner, 1992) with a multiple collector to analyze oxygen-isotopic compositions in polished sections stainless steel from the LDEF satellite. They obtained 16O– signals of ≈5 ¥ 106 c/s and a precision on the 17O/ 16O ratio of 0.1 % in thirty minutes with reproducibility at the same level. Valley and Graham (1991, 1992) have had success in measuring conductive Fe, Ti oxides such as magnetite (Fe3O4) and ilmenite (FeTiO3) on a CAMECA ims-4f. The 10 kV Cs+ primary beam is defocused to 30-40 µm with currents of 0.9 - 2.0 nA. Polished thin sections were gold coated but the conductivity of the gold was supplemented by a thin line of Ag colloid paint connecting the grains to the sample 36 TREVOR R. IRELAND holder. Secondary O– ions were accelerated through 4.5kV and the mass resolution of the spectrometer was 3000 R at the 10 % level. A field aperture was used to mask all but the central 8 µm of the primary beam spot. The 16O– count rate was limited to 500,000 c/s to minimize problems with dead time uncertainty and ≈106 counts of 18O were collected resulting in a theoretical uncertainty of 1 ‰. Reproducibility at this level appears to have been achieved since all 2000 individual 18O/ 16O ratios fit a Gaussian curve with the predicted precision. The data were normalized to SMOW by 18 measuring internal standards which were assigned the d OSMOW measured by conventional mass spectrometry. This method can be used for the quite homogeneous magnetites of this study but could prove unsatisfactory if quite heterogeneous samples are to be measured.

Magnesium. Magnesium has three stable isotopes, 24Mg, 25Mg, and 26Mg. The main isobaric interference is from hydrides, in particular the 24MgH+ interference with 25Mg+ which requires a mass resolution of 3500 R; at this level all other molecular interferences are resolved. The most commonly cited terrestrial Mg isotopic ratios are 25Mg/24Mg of 0.12663, and 26Mg/24Mg of 0.13932 (Catanzaro et al., 1966). The actual Mg isotopic composition has been a subject of some discussion since different laboratories report different 26Mg/24Mg ratios when normalized to the common value of 0.12663 for 25Mg/24Mg. This is mainly apparent in thermal ionization data which have much higher precision [see for example Esat (1984)] but even for ion probe data some differences do exist. Huneke et al. (1983) used a 26Mg/24Mg value of 0.139805 in agreement with Schramm et al. (1970), McKeegan et al. (1985) measured a value within 0.4 ‰ of the Catanzaro et al. (1966) value, as did Clayton et al. (1984), and Ireland et al. (1986) obtained a value of 0.139432. However, the absolute values are not important in this type of isotopic analysis since we are concerned with differences in isotopic composition rather than absolute ratios. Therefore the important aspect is to thoroughly document the operation of the specific instrument through measurement of standards until the behavior is sufficiently documented to allow reliable analysis of unknowns. The 25Mg/24Mg ratio is most commonly used to internally normalize isotopic mass fractionation since the main application of the Mg isotopic system is in the search for 26Mg excesses due to 26Al decay in meteoritic materials. Ion probe data have most commonly been normalized with a linear law since the largest effects in 26Mg are in meteoritic inclusions with high Al/Mg and therefore the data are insensitive to the scheme used. In order to document the veracity their analyses, Huneke et al. (1983) analyzed a series of glasses that had been gravimetrically doped with 25Mg. These glasses were subsequently analyzed by McKeegan et al. (1985) and the data from both laboratories is in excellent agreement with the gravimetrically calculated ratios. These laboratories also measured Mg isotopic fractionations in standard anorthite grains and, despite using very similar machines (CAMECA ims-3f) the range in mass fractionation was quite different, –6 to –17 ‰/amu for Huneke et al. (1983) Ion Microprobe Mass Spectrometry 37 and -4 to -10 ‰/amu for McKeegan et al. (1985). Ireland et al. (1986) documented well-resolved matrix dependent mass fractionation with Mg from spinel being some 4 ‰/amu lighter than hibonite, olivine, pyroxene, and kaersutite which all had fractionation values within ≈1 ‰/amu of each other. The documentation of the instrumental mass fractionation is therefore an important parameter to be obtained if reliable high-precision estimates of intrinsic mass fractionation are to be made. The CAMECA ims-3f has been used extensively for Mg-isotopic analysis and the experimental procedures have been described by Huneke et al. (1983) and Fahey et al. (1987b); similar procedures are used on the SHRIMP ion microprobe (Ireland et al., 1986). For these instruments, the preferred method of operation is to use high mass resolution (≈3500!R ) to resolve all molecular interferences including doubly charged Ca and Ti, as well as hydrides. Measurements are made by cyclically peak-hopping through the Mg isotopes, and 27Al if desired. The problem with measuring the 27Al abundance is that for high Al/Mg phases the Al+ signal can overload the electron multiplier. There are several options in this case. The secondary ion intensity can be reduced so that 27Al does not overload the multiplier; in this case the count rates on the Mg isotopes will also be reduced resulting in poorer precision per unit time of measurement. A Faraday cup can be employed to measure the 27Al+ intensity; in this case the high count rates of the Mg isotopes can be preserved, but the Faraday cup must be switched in and out of the beam line reliably and the electrometer must be calibrated accurately. Alternatively, the 27Al+/24Mg+ measurement can be made separately to the Mg isotopic measurement; however in this case information regarding correlated effects in 26Mg/24Mg with 27Al/24Mg may be compromised. Ireland et al. (1986) found that for minerals with more than 1-2 % MgO, there was sufficient Mg+ signal to use the Faraday cup for the Mg isotopic measurements as well. However, while there is sufficient intensity for the measurements, the method for automatic peak centering was susceptible to an interaction between the pre-slit deflection plates used to generate the deflection offset and the Faraday cup. The effect of this malaise is to offset the 24Mg+ peak but the offset is dependent on the absolute signal strength of 24Mg+; the lower the signal strength the larger the offset. The problem is clearly apparent in analyses of hibonite with less than 2 wt % MgO but has only a small effect on olivine (≈60 wt % MgO). Despite the problems found in this study, measurements using Faraday cups should be made where possible because of the benefits of no dead time correction. It is likely that such applications might become more common when the large high-sensitivity microprobes are analyzing major elements in a particular mineral phase. Mg isotopic compositions have been reported from a large number of other ion microprobe laboratories with different instruments particularly in the verification of the excess 26Mg correlated with 26Al in Allende refractory inclusions. However, the different practitioners used quite a variety of techniques. Bradley et al. (1978) used two different ARL (Applied Research Laboratory) ion microprobes to analyze anorthites from two Allende inclusions. The ARL ion microprobes were operated 38 TREVOR R. IRELAND at low mass resolution necessitating corrections for the presence of 48Ca2+ of up to 5.4 %, and for scattering of 27Al+ and 23Na+. Despite these difficulties, Bradley et al. 26 (1978) were able to show that the Allende inclusions had Mg excesses of up to 50 27 24 26 27 % at Al/ Mg ratios of 1000 corresponding to an initial ( Al/ Al)0 abundance of 5.5 ¥ 10-5 which is consistent with the original analyses of Lee et al. (1977). Shimizu et al. (1978) used an energy filtering technique on a CAMECA ims-300 to measure Mg isotopes in Leoville refractory inclusions and found that their data were also largely consistent with the 5 ¥ 10-5 initial abundance of 26Al, but hibonite in one Leoville inclusion had a large excess of 26Mg (160 ‰) at a low 27Al/24Mg of 22 corresponding to a 26Al/27Al of around 10-3. However, melilite from this same inclusion does not 26 27 show the extreme ( Al/ Al)0, but is consistent with the canonical value (Lorin et al., 1977). Macdougall and Phinney (1979) also used the ARL ion probe to show that excesses in 26Mg were present in CM meteorite hibonites. However, not only did they find 26Mg excesses, but they also found large variations in the 25Mg/24Mg ratio of up to 10 % which they interpreted as due to heterogeneities in the intrinsic mass fractionation.

Silicon. Silicon has a very similar isotopic system compared to Mg with three stable isotopes 28Si, 29Si, and 30Si with terrestrial ratios 29Si/28Si of 0.050633 and 30Si/28Si of 0.0336214 (Barnes et al., 1975). Again the main isobaric interferences to be eliminated are hydrides which are resolved at the same level (≈3500 R). There have been far fewer studies of Si isotopic compositions by ion microprobe, probably because of the lack of a suitable radiogenic precursor and therefore anomalies (in terms of the meteoritic work) would be limited to mass fractionation and nuclear anomalies, the effects of which have been found to be generally small in most samples. However, the discovery of presolar SiC grains has shown that large Si isotopic variations do exist. – McKeegan et al. (1985) used an O beam to sputter Si+ secondary ions and found that the Si was significantly mass fractionated by -31 to -37 ‰/amu. After a linear correction for mass fractionation based on the 29Si/28Si ratio, McKeegan et al. 30 28 (1985) found that there mean Si/ Si was 2.8 ‰ lower than the Barnes et al. (1975) value. This is not due to the choice of mass fractionation law since the exponential law was found to drive the corrected 30Si/28Si value even further below the Barnes et al. (1975) value. McKeegan et al. (1985) therefore chose to normalize to their data to a 30Si/28Si value of 0.03357. Zinner et al. (1987) used Cs+ to sputter Si– secondary ions so that C and Si isotopic analyses could be made on the same grains under the same analytical conditions. They found that the yield of Si– (Cs+ primary beam) per unit mass of sample was less than for Si+ (O– primary beam) but the ion yield per unit current of primary beam is substantially higher. The intrinsic mass fractionation of Si– was far less than that observed for Si+ increasing the precision achievable by the external normalization procedure. Since the isotopic anomalies in the SiC grains have turned out to be so large, an external mass fractionation correction is adequate and data Ion Microprobe Mass Spectrometry 39 are reported as deviations from the terrestrial ratios, d29Si and d30Si. Zinner et al. 30 28 29 28 (1987) also examined the normal Si/ Si for Si– after normalizing with the Si/ Si and found that the data were best described by an exponential mass fractionation law with a 30Si/28Si of 0.033474 ± 0.000013 quite different from the Barnes et al. (1975) value of 0.0336214.

Sulfur. Sulfur has four stable isotopes, 32S, 33S , 34S, and 36S with abundances of approximately 95, 0.75, 4.2, and 0.02 % respectively. In terms of terrestrial fractionation processes, it is customary to only measure the 34S/ 32S ratio and reference it to the Canyon Diablo standard ratio of 0.0450045 (Thode et al., 1961). Ion microprobe S-isotopic measurements have been made in a variety of modes, including O+ primary and S– secondary, O– - S+, and O– - S–. Pimminger et al. (1984) used O+ primary ions and analyzed and S– secondaries in galena. No charge compensation is necessary because galena is an electrically conductive phase, however, if this technique is to be applied to any other sulfides some form of charge compensation will probably be necessary. Eldridge et al. (1987) have described the techniques used for the analysis of S+ secondary ions from O– primary ions on the SHRIMP; similar procedures have been used by Chaussidon et al. (1989) on a CAMECA ims-3f. A mass resolution of 16 + ≈2000 R is sufficient to eliminate the S hydride interferences and O2 but doubly charged species such as 64Zn2+ require 4500 R for complete exclusion. However, even in sphalerite (ZnS) the presence of Zn2+ could not be detected and so it is probably safe to use the lower mass resolution. Secondary ion intensity is a strong function of matrix as is the instrumental fractionation which ranges from –15 ‰/amu for galena to –60 ‰/amu for barite. Eldridge et al. (1987) found that the different fractionation factors between the minerals were largely a function of the relative bond strengths with the weak S-metal bonds producing less fractionation than the stronger S-O bonds. The precisions obtainable ranged from 1 ‰/amu for galena to 2 ‰/amu for barite and were largely limited by counting statistics. Macfarlane and Shimizu (1991) used an unusual analytical configuration of O– primary ions and S– secondary ions. This is an unusual configuration because the greatest sensitivity is generally obtained with primary and secondary ions of opposite sign, and because the power supplies need to be banked to achieve the required potential gradients for primary and secondary ions of the same sign. Macfarlane and Shimizu (1991) operated in this mode to ensure exclusion of Zn – interferences which do not form negatively charged secondary ions. However, O2 was still a problem and so an energy offset of ≈40 eV was used to reduce the interference to ≈0.017 %. This analytical procedure usually yielded in-run precisions of ≈0.6 ‰. Fractionation is also a function of energy offset; the variations in fractionation measured at 0 eV were substantially reduced at higher energy offsets although long term variations were still observed. At an energy offset of 30 eV, the instrumental fractionation ranged from -29 ‰/amu for galena, to -47 and -48 ‰/amu for troilite and pyrite. 40 TREVOR R. IRELAND

Calcium and Titanium. Calcium and titanium have six and five stable isotopes respectively, namely 40Ca , 42Ca, 43Ca, 44Ca, 46Ca, 48Ca, and 46Ti, 47Ti, 48Ti, 49Ti, 50Ti. The terrestrial abundance ratios for Ca are most commonly normalized to 40Ca/44Ca = 47.153, with 42Ca/44Ca = 0.31221, 43Ca/44Ca = 0.06486, 46Ca/44Ca = 0.00153, and 48Ca/44Ca = 0.088727 (Niederer and Papanastassiou, 1984). For Ti, compositions are generally normalized to a 46Ti/48Ti of 0.108548, with 47Ti/48Ti = 0.099315, 49Ti/48Ti = 0.074463, and 50Ti/48Ti = 0.072418 (Niederer et al., 1981). Calcium and titanium have two common isobars at masses 46 and 48. The mass resolution required to separate 46Ca from 46Ti is ≈43,000 R and so is essentially unresolveable with the present techniques. However, 48Ca and 48Ti can be separated at a mass resolution of 10,500 R. Titanium also has isobaric interferences from 50Cr and 50V which require mass resolutions of 21,000 R and 42,000 R respectively for full separation. These isobaric interferences can be monitored and the contributions at mass 50 stripped through the signals of the major isotopes 51V (99.75 %) and 52Cr (83.79 %). The main application of Ca and Ti-isotopic analyses has been in analyzing meteoritic inclusions for isotopic anomalies. The largest anomalies have been found in the mineral hibonite (CaAl12O19, with Ti and Mg substitution for Al), therefore the analytical techniques have been designed primarily for measurements of this mineral (Zinner et al., 1986; Fahey et al., 1987; Ireland et al., 1992). All isotopes of Ca, except for 46Ca (masses 40, 42, 43, 44, 48), and all isotopes of Ti (masses 46 through 50) can be measured. Contributions from 46Ca, 50V, and 50Cr are stripped from their respective Ti isobars. Mass 43.5 is checked for the presence of 87Sr2+ which might indicate the presence of Sr2+ interferences at masses 42, 43, and 44. The possible presence of Zr2+ can also be checked but the effects are likely to be small since 48Ca+ is resolved from 96Zr2+ under these analytical conditions and the 92Zr2+, 94Zr2+, and 96Zr2+ count rates would cause negligible effects on 46Ti+, 47Ti+, and 48Ti+. Ca and Ti isotopes can be measured in a single cycle provided the 48Ti+ signal is higher than the 48Ca+ signal, such as in the majority of hibonites and perovskites (Ireland, 1990). The 48Ca+ is excluded from the 48Ti+ measurement at 10,500 R, but the converse is not true since the 48Ti+ signal is up to 100 times the height of 48Ca+. In order to monitor the magnitude of any 48Ti+ tail under 48Ca+, the high-mass edge 40 + -3 of the Ca peak is measured at a distance Dm40 = 40/48 ¥ 4.58 ¥ 10 from the center of 40Ca+ which gives the fractional proportion of the tail. This technique works as long as 48Ca+/48Ti+ is larger than approximately 0.01, for smaller ratios tailing from the 48Ti+ peak becomes unacceptably large. On the other hand, if the 48Ca+ signal is comparable in intensity to or larger than the 48Ti+ signal, as in the case of the HAL-type hibonites analyzed by Ireland et al. (1992), the peaks must be well resolved from each other to enable 48Ca+ and 48Ti+ to be centered separately. Isotopic mass-fractionation is monitored using 40Ca/44Ca and 46Ti/48Ti ratios. The instrumental mass fractionations for these two elements are quite different, for 40 46 example Ireland (1990) used mean values of D Ca = –0.6 ‰/amu and D Ti = Ion Microprobe Mass Spectrometry 41

–14.5 ‰/amu to correct the meteoritic data. Fahey et al. (1987b) found that the exponential mass fractionation law provided the best fit to the Ti isotopic data and has also been used for the Ca isotopes. However, Ireland et al. (1992) used a Rayleigh law for the HAL-type hibonites since it was argued that the mass fractionation of these hibonites was intrinsic and due to distillation, and the instrumental mass fractionation is a constant component.

Hafnium. Hafnium has six stable isotopes 174Hf (0.14 %), 176Hf (5.2 %), 177Hf (18.6 %), 178Hf (27.1 %), 179Hf (13.7 %), and 180Hf (35.2 %). The main interest in Hf 176 176 10 isotopes is through the radiogenic decay of Lu to Hf (t1/2 = 3.7 ¥ 10 yrs); Hf isotopic compositions can therefore be used as a chronometer or a radiogenic isotope tracer. Kinny et al. (1991) measured Hf isotopic compositions in zircons in order to compare the Hf isotopic systematics with U-Pb systematics. Hf is very similar geochemically to Zr and so is present at a reasonably high level in zircon (typically 1-2 wt% HfO2). The techniques used by Kinny et al. (1991) illustrate some of the difficulties encountered when analyzing isotopic compositions in the REE region of the periodic table. The total range in 176Hf/177Hf over geologic time is about 1 % and therefore high precision is required to obtain useful geologic resolution. Hafnium has atomic isobaric interferences from 176Yb and 176Lu (requiring mass resolutions of 35,000 and 90,000 R for separation) and monoxide interferences from the middle REEs which require ≈8,000 R. However, Kinny et al. (1991) noted + + + + + + that the HfO /Hf is about 4, the YbO /Yb is about 0.5, and the REEO2 /REEO is negligible and therefore more Hf signal is available from the oxide species with an additional advantage of reduced isobaric interferences. Major molecular interferences composed of atoms of Zr, Y, and O require a mass resolution of ≈ 2000 R leaving contributions only from monoxide peaks and hydroxide peaks to be addressed. An iterative correction procedure was applied to the data that included stripping the 176YbO and 176LuO interferences from the 176HfO peak, removing the contributions from 17O and 18O, determining the isotopic mass fractionation based on the 178Hf/180Hf ratio, and correcting all peaks for hydroxide interferences as measured at mass 197 (180Hf 16OH+). The most important interferences for the determination of 176Hf/177Hf are the 176YbO (1-10 % of signal at mass 192) and 176LuO (<1 %). In order to minimize corrections, low REE areas of zircon grains were selected and a liquid nitrogen cold trap was used to inhibit the formation of hydroxide species. Kinny et al. (1991) found that the REEOH/REEO ratios were not identical to the HfOH/HfO but were related by a constant factor (2.5) derived by repeated measurements on a suite of cogenetic zircons with different REE concentrations. A standard zircon was analyzed for its REE and Hf concentrations so that a correction for in situ decay of 176Lu could be made, and analyzed isotopically for an 176 177 internal isotopic standard. The Sri Lanka 7 zircon standard has a ( Hf/ Hf)init of 0.28183 ± 15 or eHf of -30 ± 1 as measured by thermal ionization. The ion probe 42 TREVOR R. IRELAND

data are in reasonable agreement with this value having eHf of -23 ± 5. The precision for the ion probe data equates to ≈ ± 0.5 ‰ (2s) and can provide an age resolution of the order of a few hundred million years.

Uranium-Thorium-Lead. The U-Th-series decay schemes are widely used in geochronology. The two main isotopes of U, 235U (0.72 %) and 238U (99.27 %) decay to form 207Pb and 206Pb with half lives of 7.0 ¥ 108 yrs and 4.5 ¥ 109 yrs respectively, while 232Th decays to 208Pb with a half life of 1.4 ¥ 1010 yrs. Lead has one stable isotope, 204Pb, that is non-radiogenic. In order to produce concordia diagrams, the standard method of portraying U-Pb data, both the Pb isotopic composition and the U/Pb ratio must be known. In ion microprobe analysis this involves two types of determinations; measurement of Pb isotope ratios and measurement of an interelement ratio, U/Pb. The main application of the U-Th-Pb measurements on the ion microprobe has been zircon dating (Williams, 1992), although procedures for perovskite have also been described (Ireland et al., 1990) and it is likely that other minerals, such as monazite, will be found to be suitable as well. In zircon analysis, the mass spectrum around the Pb peaks has a variety of Zr and Hf oxides and silicides, as well as REE molecules. The molecules can be completely resolved at around 6500 R and only hydrides are not resolved. The measurement of isotopic ratios of Pb at trace levels and at high mass resolution is reliant on high sensitivity; the SHRIMP I ion microprobe operates with typical sensitivities of around 5-10 c/s/ppm Pb/nA. The techniques used for U-Th-Pb isotopic measurements on SHRIMP were first described by Compston et al. (1984). An O– primary beam, preferably mass filtered, is used to sputter Pb+ secondary ions. The measurement cycle includes the 90 16 + 238 + 232 16 + 238 16 + Pb isotopes, Zr2 O as a reference for concentrations, U , Th O and U O . An important aspect of this technique is the normalization of the Pb+/U+ ratios to a standard zircon. This is required for the determination of the interelement ratios, but also allows a check on the measurement of Pb isotopic ratios for the presence of isotopic mass fractionation and hydrides. Compston et al. (1984) took the Pb isotopic ratios as measured. The ratios were not corrected for isotopic mass fractionation or the presence of hydrides because these effects were probably small (in terms of the lunar zircons being analyzed with 207Pb/206Pb ratios around 0.53) and also act in opposite directions, hydrides to increase the measured 207Pb/206Pb (from the contribution of 206PbH+ under 207Pb+), and mass fractionation to decrease 207Pb/206Pb (since in general the lighter species is preferentially ionized). However, for younger zircons these effects, especially the hydride contamination, should be taken into account. The combined effects of hydride and mass fractionation can be examined by comparing the measured 207Pb/206Pb ratio of the standard with the conventionally measured values. A correction can be applied to the unknowns on this basis assuming that the unknowns are affected in the same way as the standard. In general, such a Ion Microprobe Mass Spectrometry 43 correction has been found to be small and a correction is only necessary after recent insertion of the sample when there is a significant residual water level in the vacuum. If the sample is left to pump overnight before analysis, then the measured ratios are usually within error of the conventionally measured values. This indicates that the instrumental fractionation of Pb from zircon is also low. The ThO+/UO+ ratios as measured are linearly proportional to the 232Th/238U in zircon with a fixed discrimination as determined by concurrent 208Pb/206Pb measurements. For a closed system,

l t Ê 208Pbˆ 232Th È (e 232 -1)˘ Á 206 ˜ = 238 Í ˙ Pb U el238t 1 Ë ¯ rad ÎÍ ( - )˚˙ 232 238 -11 -1 where l232 and l238 denote the decay constants of Th and U (4.9475 ¥ 10 a and 1.55125 ¥ 10-10 a-1 respectively). The discrimination constant is 1.11, i.e. † 232Th Ê ThO+ ˆ 238 = 1.11Á + ˜ U Ë UO ¯ and the error on this factor is around 1 % based on the error of a linear correlation determined from measurements on two Sri Lanka zircons with different Th/U. The Pb/U calibration is derived from the observation by Andersen and Hinthorne (1972) that the Pb/U ratio should be related to the concurrent UO+/U+ measurement based on their local thermodynamic equilibrium model. Compston et al. (1984) emphasized that the use of this correlation in their normalization procedure is purely an empirical treatment and is not in itself used as support for the local thermodynamic equilibrium model. It should also be noted that the UO+/U+ in any analysis does not appear to be controllable but is a somewhat random function of the day to day operation of the ion probe. Ratios of 206Pb+/U+ are measured in the standard and unknown and the actual Pb/U of the unknown is derived from the known Pb/U value for the standard with the relationship + + (Pb U ) (Pb U) unk = unk Pb+ U + (Pb U ) ( )std std In Compston et al. (1984) a linear regression was found to be adequate to describe the relationship between 206Pb+/U+ and UO+/U+ which could be used to normalize the measured Pb+/U+ to a single value of UO+/U+ i.e. Pb+ U + = 0.0764 x - 2. 77 ( )std ( ) where x is UO+/U+. Subsequently, Williams and Claesson (1987) found a degree of curvature in the standard calibration that could be described by a quadratic calibration, such that, 44 TREVOR R. IRELAND

Pb+ U + = 0.0048x 2 + 0. 0265x - 0.0825 ( )std Most recently a power law fit, albeit very similar to the quadratic formulation, has been found to fit better the accumulated standard data and has a form

Pb+ U + = 0. 0055x2 The actual form of the standard calibration curve is only particularly important when the mean UO+/U+ of the standards and the unknowns differ. If the means are the same then a poor functional fit has the main effect of increasing the error on the Pb/U calibration based on the reproducibility of the standards. Figure 15 illustrates the fit of the three schemes to a typical set of analyses of the SL13 standard zircon. All curves give an acceptable calibration to the data and the discrepancies are only apparent at large deviations from the centroid. Within a given day, the scatter of the data about the calibration curve is found to be around 2 %. This scatter is almost certainly an instrumental effect since repeated

Figure 15. U-Pb calibration curves are used to normalize out variations in sputtered ion emission. All curves give a good functional relationship to the data and the differences are only apparent when large extrapolations are required. For most data sets however, the centroids of the standard and unknowns are similar and the exact form of the curve is not a critical parameter. Ion Microprobe Mass Spectrometry 45 conventional analyses of the SL13 standard zircon have yielded the same age (572.2 ± 0.2 Ma) within error. The SL3 standard first used by Compston et al. (1984) was found to be actually a heterogeneous mix of baddeleyite and silica-rich glass (McLaren et al., 1993) and so use of this standard has been abandoned in favor of SL13. Approximate concentrations can be calculated using a similar formulation + + + + between Zr2O /U and UO /U and the unknowns are normalized to the 220 ppm U concentration of the SL13 standard. The absolute error on these concentration measurements is around 20 % since this is the variation observed in individual fragments of SL13. Despite having low initial Pb concentrations, a correction for this Pb has to be applied to the measured Pb isotopic composition in order to obtain age information. In general this has a small effect on old zircons, but for young zircons it can have a large effect, particularly with regards to the amount of radiogenic 207Pb in the analysis. Common Pb can be estimated in three different ways. Since 204Pb is nonradiogenic, its abundance is fixed and reflects the initial Pb in the system. Therefore for a given initial Pb composition, the contribution of this Pb can be subtracted from the analysis. We can define the common Pb content as the fractional amount of initial 206Pb in the analysis, i.e.

206 206 ¶ = Pbinit Pbtot then for the 204Pb/206Pb correction method, ( 204Pb 206Pb) ¶ = tot 204Pb 206Pb ( )init The isotopic composition of the common Pb can be estimated from coexisting common Pb-rich minerals such as feldspar, estimated from the common Pb growth curves, or an average composition can be used, such as Broken Hill Pb, when the common Pb is largely surface related (and the contribution is small). This is the most reliable method for common Pb correction in old zircons, but for young zircons the uncertainties in the small amount of common Pb become too large for an accurate correction, particularly for the low abundance 207Pb isotope. The second estimate is based on the 208Pb/206Pb ratio. The radiogenic 208Pb/206Pb can be estimated from the ThO+/U+ ratio for an assumed formation age. The actual formation age assumed has only a small effect on the correction since the common 208Pb/206Pb ratio has changed only from 0.28 to 0.32 over geological time. Then the fraction of common 206Pb can be estimated with ( 208Pb 206Pb) - ( 208Pb 206Pb) f = tot rad 208Pb 206Pb - 208Pb 206Pb ( )init ( )rad For well-behaved populations of young zircons, this method can give very reliable data however, it can be unreliable for grains that have lost Pb during geological disturbances, or for grains in which Th and U have moved independently. 46 TREVOR R. IRELAND

A third method for estimating common Pb is based on an assumption of concordance. This method is used mainly for young zircons when insufficient counts of 204Pb are available for a precise common Pb correction. The 207Pb correction assumes concordance between the 206Pb/238U age and the 207Pb/206Pb age and calculates a fraction of common Pb on that basis, ( 207Pb 206Pb) - ( 207Pb 206Pb) f = tot rad 207Pb 206Pb - 207Pb 206Pb ( )init ( )rad This method is useful for obtaining mean 206Pb/238U ages without propagating a correlated error from uncertainty in the common Pb correction.

B. Quantitative Analysis

The secondary ion mass spectrum can also be used to quantify the abundances of major and trace elements in a target; this type of application is generally referred to as quantitative analysis. Ion microprobe mass spectrometry has great benefits for elemental abundance measurements such as inherently low background levels and reasonably high ionization efficiencies for most rock-forming elements. However, the ionization processes responsible for the benefits of the ion microprobe also constitute one of its major problems, namely the relationship between ionic intensities of different elements in the secondary ion spectrum and their actual concentrations in the sample is not necessarily a predictable function. A great deal of effort has been put into formulating a general theory of ionization which would allow all elements to be normalized according to the ion to element ratios of only a few elements. For example, in the local thermodynamic equilibrium model of Andersen and Hinthorne (1973) the relationship between two adjacent charge states of a single element could be used to characterize an analysis and derive quantitative analyses of all other elements within a factor of two (see Figure 6). However, it is difficult to justify the physical properties attributed to the sputtered region from this model and even a factor of two is basically not good enough for modern geochemical analysis. There are two main problems in elemental abundance measurements. The first is the presence of complex isobaric interferences particularly around the region of the rare earth element (REE) spectrum, and the second is the matrix dependence of the ionization efficiency. As outlined above, isobaric interferences can be discriminated against by either high mass resolution or energy filtering. For most ion microprobes of the CAMECA 3f-4f-5f series, the loss of intensity associated with high mass resolution restricts measurement precision and also calls for extremely good magnet control, a situation that is exacerbated by the practical limitations of centering on very low intensity ion beams. Therefore, while the energy filtering technique drastically reduces the secondary ion signal from a given species, the mass analyzer can be operated at lower mass resolution and hence the magnet control Ion Microprobe Mass Spectrometry 47

requirements are not as stringent. Another potential benefit of the energy filtering technique is that ionic ratios do not appear to be as matrix sensitive as in the high mass-resolution approach, that is the ions which leave the sample surface with low energies tend to have been most greatly affected by the matrix. For high mass resolution analysis, it is imperative to have a suite of standards with different elemental abundances and the same mineralogy as the unknown. However, this does not mean that the energy filtering technique can be used without standards. It might be possible to use a single standard for a variety of minerals once the relationship between the matrix effects of the standard and unknown mineral matrices have been established, but the calibration of any technique through the analysis of well- documented standards is imperative.

Energy-filtering technique The energy filtering technique is used to discriminate against complex molecular interferences by selecting a high-energy window where such interferences are reduced by orders of magnitude relative to the atomic species (Figure 16). The technique was proposed by Shimizu (1978) and was used to analyze a series of major and trace elements, ranging in mass from 23Na to 59Co, in plagioclase from different magmatic environments. Shimizu et al. (1978) extended the range

Figure 16. Low mass resolution scan of the peaks in the mass region of the REE. The unfilled bars indicate the signal that is obtained without energy filtering while the filled bars are the signal with the filtered secondary ion beam; from Zinner and Crozaz (1986). 48 TREVOR R. IRELAND

Figure 17 Energy distributions for (a) 44Ca+with 1 eV window and 30 eV window and (b) 40 31 16 159 mass 159 isobars Ca2 P O3+ and Tb+. Typical operating conditions for energy filtering of REE are a 100 V offset from the 10 % low energy edge of the distribution. Graph “b” illustrates the rapid decrease in the signal of the complex molecular species compared to the gradual decrease from the atomic ion species. Figure from Zinner and Crozaz (1986). of analyses using this technique to Pb isotopic compositions in galena, Mg isotopic compositions in a refractory inclusion from the Leoville meteorite, and analyses of major and trace elements (including REE) in various terrestrial minerals. They showed that at high energies (150 eV offset) the isotopic compositions of a variety of elements approached their terrestrial compositions indicating the virtual elimination of complex molecular interferences. In general they found linear correlations between ionic and atomic ratios however, they also showed that this is not necessarily the case with a systematic variation of Ca+/Si+ as a function of Fe concentration in Ca-rich pyroxenes. Clearly it is necessary to not only analyze standards of the same mineral composition as the unknowns, but also a range of compositions that brackets the unknown in terms of chemistry. For the REE measurements Shimizu et al. (1978) assumed constant ionization efficiencies for all the REE and plotted the data as CI-normalized and La-normalized patterns. They recognized the shortcomings of this method and proposed that further work was necessary to determine the relative ionization efficiencies of the different REE as well as to analyze well-documented standards so as to obtain absolute concentrations. Zinner and Crozaz (1986) used the energy filtering technique to develop a method specifically for measuring REE concentrations but other elements could also be included in the same analysis. Their experimental procedures involved monitoring the energy offset during the analysis by measuring the 10 % lower Ion Microprobe Mass Spectrometry 49 energy edge of the energy distribution and maintaining a 100 V offset from that position (Figure 17). In this way, differences in ion energy due to charge build-up do not affect the energy range of the ions being analyzed. Rather than analyzing a single isotope for each REE, the determination is based on the measurement of the mass spectrum between masses 133 and 191 and its deconvolution into element and element oxide species. Counts from the 59 measured masses form an overdetermined system for the unknown intensities of 33 M+ and MO+ species and the solution is obtained by minimizing the least squares of the difference between the counts at a given mass and the various contributions obtainable at that mass (divided by the uncertainty on the counts at that mass). The minimum solution gives a goodness-of-fit parameter (c2) of the match between the measured spectrum and the intensities calculated for each of the components. A low c2 indicates that there are no additional interferences and the data are consistent within counting statistics; a high c2 indicates that either there are additional interferences to be taken into account, or the errors on the analyses are larger than those derived from counting statistics. The latter might commonly be the case when samples with relatively high count rates are deconvolved and then 1/√N is an underestimate of the error (for example primary ion beam noise can not be disregarded and isotopic mass fractionation might also be significant). Other problems arise in the deconvolution of the heavy REE Yb, Lu, and Gd. The similarity in the isotopic patterns of GdO and Yb means that small changes in the counts of different isotopes can substantially effect the resultant abundance of Yb. In order to better constrain the deconvolution, GdO is subtracted by assuming a GdO/Gd ratio measured in terrestrial standards. Similarly 159TbO interferes with 175Lu and since Tb is monoisotopic and 97 % of Lu is 175Lu it is impossible to separate the signal at mass 175 into TbO and Lu. Again a contribution based on TbO/Tb is removed from mass 175. While Gd has seven isotopes, 152Gd (0.2 %) and 154Gd (2.2 %) are too small for accurate concentration assessments, 155Gd and 157Gd have contributions from 139LaO (99.9%) and 141PrO (100%), and 156Gd, 158Gd, and 160Gd cannot be easily deconvolved from CeO and NdO. In order to improve the precision of the Gd measurement, a fixed pattern of MO+/M+ for the elements La, Ce, Pr, Nd, and Sm is used with an overall multiplication factor applied to this pattern in the least squares fit. + The relative ion signals of the individual elements, [REEi ], are normalized to the ion signal of a major element in the sample (e.g. Ca+ or Si+) whose concentration [CaO] or [SiO2] can be measured by electron probe and multiplying by sensitivity factors Fi derived by analyzing suitable standards. Therefore, the concentration of each REE, [REEi], is given by + REEi [REEi ] = Fi [CaO] Ca + for the case of normalization to the CaO concentration of the matrix. 50 TREVOR R. IRELAND

Figure 18. Ion yields of REE relative to Ca+ from four different mineral matrices. While the patterns are similar, there are significant differences in the absolute yields from the different minerals. Data from Zinner and Crozaz (1986), Fahey et al. (1987c), and Ireland et al. (1991).

Zinner and Crozaz (1986) found that the ion yields for the REE relative to 44Ca+ in a terrestrial apatite standard range by over a factor of two from 0.53 (Lu) to 1.23 (Eu) and hence the sensitivity factors range from 5.81 to 2.20 respectively. The ion yields from different matrices follow a similar pattern although the absolute sensitivities are different (Figure 18). The uncertainty in an analysis can be divided into the precision (usually defined by counting statistics) and the accuracy (which is a function of the suitability of the sensitivity factors and the determination of the major element abundance for the analyzed spot). The use of the sensitivity factors Fi inherently assumes that the secondary ion intensity is linearly proportional to the concentration. Such a relationship has been demonstrated by a number of experiments and is one of the advantages of this technique. However, such a relationship cannot be assumed since in some cases a non-linear working curve has been advocated (see below). The relative precisions of REE concentrations can be better than 10 % depending on the concentrations of the REE, although the absolute concentrations may not be known to such a high level. In geochemical modeling, where the absolute concentrations are important in constraining various parameters such as parent magma compositions, it is important to try and assess the external errors applicable to the analyses. However, in determining variations in REE concentrations on a spot by spot basis the relative abundance is the important parameter and the absolute calibration will only produce a systematic shift on all concentrations.

Ion Microprobe Mass Spectrometry 51

Figure 19. REE abundance patterns of a zircon as measured by SHRIMP I and the CAMECA ims-3f at Washington University under the same conditions with the same set of sensitivity factors gives very similar results despite the different ion extraction and beam transport conditions in the two mass spectrometers.

It is interesting to note that the same sensitivity factors (relative to Si+) could be used on two quite different machines with different secondary-ion extraction systems, viz. the CAMECA ims-3f at Washington University and SHRIMP at the ANU. Ireland and Wlotzka (1992) presented REE patterns of a zircon measured on these two instruments using the same sensitivity factors and very similar results were obtained (Figure 19). Overall the energy filtering technique is a robust method for the determination of a wide variety of trace elements.

Working curve calibrations Whether high mass-resolution or energy filtering is used to discriminate against isobaric interferences some form of calibration is required which relates the ionic intensities to the atomic concentrations. In its simplest form an empirical working curve can simply show the relationship between measured ionic ratios and the atomic ratio in a mineral phase for which several standards are available with different concentrations. While the functional relationship must be related in some way to the physical mechanisms involved in sputtering, it cannot be stressed too greatly that the working curves are empirical and their purpose is solely as a calibration tool. For the most basic calibration, one standard might be used and a linear function passing through the origin could be assumed. This is basically the formulation used by Zinner and Crozaz (1986) with the sensitivity factors; these factors simply apply a single factor for any given ratio. Maas et al. (1992) have also used a simple linear calibration with high mass-resolving power analyses to derive REE abundances in Archean zircons by ratioing the counts obtained on an 52 TREVOR R. IRELAND isotope of a given element with the counts of the same peak obtained from a standard zircon. However work from other laboratories suggests that the linear calibration is not always appropriate. Hinton (1990) notes that while Si+ is the most obvious species to use for normalization, the relative intensity of Si+ is strongly affected by changes in analytical conditions. Furthermore, Botazzi et al. (1991) found that the Si-normalized intensity ratios had to be adjusted to take into account relative REE intensity enhancements due to compositional matrix effects between felsic and mafic/ultramafic samples. The normalization was based on the relationship

REEi REEc = b l + a Sii Sic where the best value for the exponent l was found to be 1.3. However, Botazzi et al. (1991) used a relatively low energy offset (80 ± 25 eV) and so residual matrix effects should be more pronounced than in the case of the data from Zinner and Crozaz (1986), who used an energy window of 100 ± 30 eV. Ray and Hart (1982) demonstrated good linear correlations (coefficients better than 0.98) for a number of elements including Na, Mg, Al, Ca, Ti, Sc, Ba, Rb, Zr, and Fe sputtered from natural clinopyroxene and synthetic glass. While linear correlations were obtained for both clinopyroxene and glass for a given element, the slopes of the linear correlations were different with ions preferentially emitted from the crystal matrix. This is a matrix effect and is probably a function of the energy offset as well since unpublished data from SHRIMP (T. Ireland, W. McDonough, R. Rudnick) indicate that a glass standard can provide suitable calibration for clinopyroxene and garnet at the relatively high energy offsets used in the SHRIMP analyses. On the other hand, Shimizu et al. (1978) used a similar offset in their clinopyroxene analyses and found a correlation in Ca+/Si+ with the Ca/Fe ratio of the sample. It is clear that the analysis of well-calibrated standards is necessary in order to constrain the correct formulation of the working curve. Only in this way can problems associated with matrix effects be addressed. The exact formulation of the curve is not important provided it can be demonstrated that it produces accurate and reproducible results for a wide variety of target compositions.

Other techniques The isolated specimen technique (Metson et al., 1984) is another form of the energy filtering technique and relies on sample charging to achieve the energy selection of the secondary ions. However, the build up of charge on the sample surface degrades the primary beam spot and the extraction conditions cannot be controlled to the same degree as in the conventional energy filtering technique. Ion microprobes have been used extensively in materials science research and in particular in the semiconductor industry to measure concentrations of contaminants and dopants. However in this field the preferred method of analysis is by depth- Ion Microprobe Mass Spectrometry 53

Figure 20. (a) Depth profiles of an amorphous Si:H/Cr layer on crystalline silica. The isotopic species monitored are 1H, 16O, 29Si, 50Cr, and 52Cr. Images (b) and (c) correspond to the distribution of 52Cr at the depths (sputter time) labelled in (a). Figures adapted from Herion et al. (1989) reprinted with permission from John Wiley Ltd.). profiling. This method makes use of the sequential manner in which the sample is sputtered away and ion intensity versus time is monitored as the sample is sputtered. In this way, information from the third dimension can be obtained (hence the name) and potentially a three dimensional reconstruction of a sample could be made (Figure 20). Quantitative analyses can also be accomplished by implanting ions of the element under examination. A known dose of the element is implanted into the sample and hence the depth integrated signal from the ion probe over this implantation depth is proportional to the concentration. Once the implanted layer is sputtered away, the signal is proportional to the intrinsic concentration of the element in the sample. The known concentration of the dopant can then be used to normalize the intrinsic concentration of the sample. The main difficulty in this approach is that the sputtering probably does not occur in a layer-by-layer fashion but represents some integrated mixture of layers because of heterogeneities in the primary beam and knock-on effects of surface atoms to deeper levels. Furthermore, the absolute concentrations of the dopants are probably known to no better than approximately 25 % and so the accuracy of a given analysis can be no better than this. However, 54 TREVOR R. IRELAND while the ultimate precisions achievable by this technique may not be high, the concentrations of contaminants and dopants can be extremely low and so even large errors can give useful results. Useful reviews and details of the depth profiling technique are given by Wilson et al. (1989) and Zinner (1980). In the earth sciences depth profiling has had more limited application but is of particular use in measuring diffusion profiles, for example Bancroft et al. (1987) assessed the stability of synthetic and natural titanites through depth profiles after leaching the samples in pure D2O, Giletti and Yund (1984) studied diffusional oxygen exchange between quartz and water by determining 18O depth profiles, and Ryerson and McKeegan (1993) have measured oxygen diffusion profiles in meteoritic minerals.

V.!APPLICATIONS

Successful applications of the ion microprobe are becoming increasingly abundant as are the variety of problems to which ion microprobes can be applied. It is beyond the scope of this paper to present all applications of ion microprobes that are currently in use. The applications that will be highlighted in this review are those which concern applications in the earth sciences, particularly those in cosmochemistry and geochemistry. The subdivision of relevant work into effective subsections is rather difficult. A natural subdivision occurs between terrestrial and extraterrestrial applications and chemical and isotopic work because the problems are quite different in each field. In cosmochemical applications there can be large variations in isotopic and chemical abundances, whereas in terrestrial samples the effects are generally more subtle. For this reason alone, there have been far more papers written concerning the analysis of extraterrestrial materials than terrestrial and hence the discussion is somewhat weighted in favor of these (perhaps) rather esoteric studies. One of the greatest advantages of the ion probe is that a large variety of analyses can be performed on the same sample. This has been especially notable in the field of cosmochemistry where isotopic abundances and elemental abundances are measured for grains less than 10 µm in diameter. For the extraterrestrial studies, rather than simply progressing through the periodic table examining the different elements, it is more useful to structure the discussion around the samples. In the case of terrestrial studies however, there are far fewer cases of analyzing different systems in the same sample and so for these studies the applications are more structured around the techniques.

A. in Cosmochemistry

By far the greatest number of publications dealing with ion microprobe analysis of “geologic” materials has been concerned with formation and evolution of extraterrestrial samples. The reason for this is clear; isotopic and chemical anoma- lies can be extreme. This is exemplified by the use of logarithmic plots for both Ion Microprobe Mass Spectrometry 55

chemical and isotopic data from some presolar materials. Furthermore, these samples are often small and the maximum sensitivity per unit weight of sample is required to enable the greatest number of complementary analyses to be made on individual grains. For some of the recent research into interstellar grains, a single grain of 5 µm in diameter has been analyzed for Si, C, N, Mg, Ca, and Ti isotopic compositions. This work would have been impossible without the recent advances made in ion microprobe mass spectrometry.

Introduction to Chemical and Isotopic Systematics It has been recognized for over thirty-five years now that nucleosynthesis in stars is responsible for the production of the elements from carbon to the heaviest nuclides (Burbidge et al., 1957). It is also clear that the isotopic abundances present in the solar system cannot have been produced in a single nucleosynthetic site but instead record a history from a variety of stellar sources. However, prior to 1973 the canonical view of the solar nebula was of a hot turbulent body in which heating was so pervasive as to cause widespread volatilization of all elements thereby homogenizing any initial isotopic heterogeneities that were present. This view was changed irrevocably when Clayton et al. (1973) measured correlated effects in the 17O/ 16Oand 18O/ 16O ratios from refractory inclusions in carbonaceous chondrites that were interpreted to represent a reservoir of 16O-rich material in the solar nebula. With the finding of isotopic anomalies in the most common of rock forming elements, the likelihood was that isotopic effects in other elements could be found provided you looked in the right place at the right element with sufficiently high analytical precision. And so thermal ionization mass spectrometers were tuned up to yield the best possible precision and it was not long before anomalies were found in the abundance of 26Mg (Gray and Compston, 1974; Lee and Papanastassiou, 1974), which could be attributed to the in situ decay of 26Al with the finding that excess 26Mg was correlated with Al/Mg (Lee et al., 1977). Isotopic anomalies in Ti were then found at around the 1 ‰ level in all inclusions that were analyzed (Heydegger et al., 1979; Niederer et al., 1981; Niemeyer and Lugmair, 1981). The effects were predominantly in the abundance of 50Ti indicating that the isotopic abundances of the Fe group elements could be affected by the admixture of a neutron-rich nucleosynthetic product. This interpretation was supported by the discovery of effects in the isotopic compositions of Ca, Cr, Fe, Ni, and Zn (Jungck et al., 1984; Birck and Allègre, 1984; Birck and Lugmair, 1988; Völkening and Papanastassiou, 1989; Völkening and Papanastassiou, 1990). The fact that these isotopic anomalies were found in refractory inclusions was not believed to be coincidental. The mineralogy of the Allende inclusions - primar- ily melilite, spinel, and Ti-rich pyroxene, is consistent with their formation as condensates from a cooling gas of solar composition (Grossman, 1972). The sequence basically follows the sequential condensation of the refractory oxides of Al, Ca, and Ti followed by Mg and Si. The trace element compositions of these objects also indicate a high temperature origin with enrichments in the refractory 56 TREVOR R. IRELAND lithophile elements to around 20 ¥ chondrite abundances and fractionation of the REE according to their relative volatilities. In this regard, there are two main types of refractory inclusion: those showing fractionations in the most volatile of the REE, and those showing fractionation in the most refractory (Fegley and Ireland, 1991). These fractionations can yield patterns either enriched or depleted relative to solar abundances. One particular type of pattern, that with depleted ultrarefractory REE has been used to argue for a condensation origin of these inclusions since the pattern can only be produced in a condensate that has been separated from a more refractory residue (Boynton, 1975). The correlation of the presence of isotopic anomalies and the refractory chemical compositions of the inclusions is suggestive of their early formation before the initial isotopic heterogeneities of the solar system were homogenized through mixing and high temperature processing. If this were the case, then larger anomalies might be expected in more refractory types of inclusion consisting of corundum or hibonite. Corundum- bearing inclusions are rather rare but hibonite-bearing inclusions are quite common constituents of the CM2 class of meteorites. These meteorites are generally much finer grained than the Allende CV3 type and the refractory inclusions are also much smaller being typically less than 500 µm compared with up to 2 cm inclusions in Allende. With such small objects, it is difficult to carry out conventional mass spectrometric analyses but they are ideally suited to ion microprobe analysis.

Refractory Inclusion. The first isotopic measurements to be performed on the ion microprobe were of Mg since it had already been demonstrated that large effects from 26Al decay might be expected. Mg isotopic measurements were made on a wide variety of machines following quite different experimental procedures (for further details see Methodology section). These analyses showed that while a large number of refractory inclusions do preserve the effects of 26Al decay at the 26 27 canonical level of ( Al/ Al)0 of 5¥10-5 e.g. (Hutcheon et al., 1978; Hutcheon, 1982; Stegmann and Begemann, 1981; Huneke et al., 1983). However a large number do not have excess 26Mg consistent with "live" 26Al, in particular hibonite-bearing inclusions (e.g. Macdougall and Phinney, 1979; Hutcheon et al., 1980; Bar- Matthews et al., 1982; Fahey et al., 1987b; Ireland, 1988; 1990). A simple chronological interpretation would place those inclusions without excess 26Mg at a formation time after those preserving 26Mg at the canonical level. However, the lack of effects in a lot of the hibonite grains was perplexing since the CM inclusions were thought to be amongst the earliest condensates based on the presence of hibonite, and occasionally corundum. This requires that our understanding of the formation of refractory inclusions in a simple thermal progression is at fault or that 26Al was not homogeneously distributed throughout the solar nebula in space and/or time. The formation history for hibonite became even more puzzling with analyses of Ca and Ti. Ca is of course an essential element in hibonite at 8.5 wt % CaO while the Ti concentration ranges from under 1 up to 9 wt % TiO2. Conventional isotopic Ion Microprobe Mass Spectrometry 57 analyses of Allende inclusions had shown that small anomalies (generally excesses) were present in the abundances of 48Ca and 50Ti. The first ion-microprobe Ti isotopic analyses relied on peak stripping 48Ca away from 48Ti since the transmission loss on the AEI IM-20 at high mass resolution is prohibitive (Hutcheon et al., 1983). These analyses showed that some Murchison hibonites may have 50Ti deficits as large as 15 ‰ and Hinton et al. (1987) used the same technique to measure a 70 ‰ deficit in the Murchison inclusion BB-5. However, with such large anomalies apparent in 50Ti, there is a high degree of likelihood that 48Ca is also anomalous and therefore a correction cannot be made for the presence of 48Ca under 48Ti based on normal isotopic ratios. For this reason Ca and Ti isotopic measurements are best made at high mass resolution (≈10,000 R) where the peaks of 48Ca and 48Ti are separated (Fahey et al., 1985; Ireland et al., 1985). In contrast to the large deficits observed by Hutcheon et al. (1983) and Hinton et al. (1987), Fahey et al. (1985) measured a 100 ‰ excess of 50Ti in two Murray hibonites and smaller excesses in two others, while Ireland et al. (1985) measured a range of compositions from 40 ‰ deficits to 10 ‰ excesses. Initially, the three laboratories reporting Ti isotopic measurements adopted different normalization schemes but as it became apparent that effects were also present in 49Ti, a 46Ti/48Ti normalization scheme was adopted by all. For most meteoritic hibonites, the largest effects are in 50Ti with smaller anomalies sometimes resolved in 49Ti while d 47Ti is generally close to normal (Figure 21). Zinner et al. (1986b) reported the Ca-isotopic compositions in a suite of hibonites that had a range of d50Ti from -40 ‰ to +100 ‰. The Ca isotopic compositions were normalized to the 40Ca/44Ca ratio and no anomalies were observed for d42Ca and d43Ca but large effects were evident in the abundance of 48Ca with anomalies ranging from –46 to +56 ‰ (Figure 21). The sign of d48Ca is nearly always the same as that for d50Ti (Figure 22) indicating that the anomalies are associated. The main exception to this correspondence is HAL which has a small deficit in 48Ca and a 15 ‰ excess of 50Ti (Fahey et al. 1987b). The veracity of the ion probe Ca isotopic compositions has been demonstrated by a conventional thermal ionization measurement of BB-5 which has a d48Ca of –56.1 ± 3.7 ‰ by ion microprobe (Fahey et al., 1987a) and –58.6 ± 1.0 ‰ by the conventional method (Podosek and Brannon, 1988) . Fahey et al. (1987a) measured O-isotopic compositions in a suite of inclusions including those analyzed by Fahey et al. (1985), Zinner et al. (1986b), and BB-5 which was analyzed by Hinton et al. (1987). The O-isotopic compositions of the hibonite grains all showed enrichments in 16O relative to the normal O-isotopic composition with the 7 grains all lying within error of the Allende mixing line defined by conventional analyses of Allende inclusions (Figure 23). The maximum 16O enrichment was shown by BB-5 at ≈ 70 ‰, whereas the maximum shown by the conventional data from Allende inclusions is around 50 ‰. The interesting aspect of the O-isotopic anomalies is that there is no correlation with the Ca and Ti isotopic compositions. Despite individual grains having a wide range in 48Ca and 58 TREVOR R. IRELAND

Figure 21. The Ca and Ti isotopic compositions of Murchison hibonites BB-5 and 13-13 show the largest 48Ca and 50Ti anomalies. BB-5 is depleted in 48Ca and 50Ti whereas 13-13 shows large enrichments. Data from Fahey et al. (1987a) and Ireland (1990).

50Ti anomalies from large deficits to large enrichments, they all show enrichments of 16O between 4 and 7%. This suggests that while the Ca and Ti have a common nucleosynthetic origin, the O-isotopic anomalies are probably related to another source. Besides isotopic measurements, trace-element compositions could also be measured from these grains. Neutron activation analyses had suggested that the patterns were not very different from the larger Allende inclusions (Ekambaram et al., 1984a) although the INAA analyses often had quite large uncertainties because of the small size of the samples. The benefit of ion probe analysis is that all REE elements can be measured as well as other refractory lithophile elements on the same discrete areas that have been analyzed for isotopic compositions. Analyses can also be made on a relatively short time scale (2 hours) so that a large data base of isotopic and chemical data on hibonite-bearing inclusions was rapidly built up. Around 100 individual grains have now been analyzed by ion microprobe (Ireland, 1990 and references therein) and have allowed detailed studies of the correlations between isotopic systems of different elements as well as chemical and morphological effects. The correlation of Ca and Ti isotopic anomalies is almost certainly a nucleosyn- thetic effect, but other correlations between morphological, chemical, and isotopic Ion Microprobe Mass Spectrometry 59

Figure 22. The 48Ca and 50Ti isotopic anomalies in hibonites are generally correlated in that samples with enrichments (or deficits) in 48Ca are accompanied by enrichments (or deficits) in 50Ti. For comparison, the largest anomalies measured by thermal ionization mass spectrometry are represented by the FUN inclusions, EK-1-4-1, and C1 as indicated in the inset. Figure adapted from Ireland (1990). characteristics have not been so readily forthcoming. The problem with the systematics of the hibonite-bearing inclusions was that the correlations were not quantitative but rather indicated associations of features. Hutcheon et al. (1986) first recognized that inclusions consisting solely of hibonite rarely had 26Mg* at a level of 5¥10-5¥ 27Al whereas hibonite associated with spinel or melilite generally did have 26Mg* at the canonical level. Ireland (1988) noted that the hibonite platy crystal fragments (PLACs) were not only characterized by low levels of 26Mg* (Figure 24), but were also the most common carriers of highly anomalous 50Ti. On the other hand, the spinel - hibonite inclusions (SHIBs) generally had small 50Ti effects but had 26Mg* at a level of 5¥10-5¥27Al (Figure 25). Ireland et al. (1988) showed that the PLACs were dominated by REE patterns showing depletions in Eu and Yb whereas the SHIBs often had patterns showing fractionations in the ultrarefractory REE (Figure 25). Clayton et al. (1988) noted that radiogenic 26Mg* had not been found in any of the inclusions with 50Ti anomalies larger than 10 ‰. Ireland (1990) also found that inclusions with ultrarefractory-depleted patterns all had radiogenic 26Mg* unless they also had large 50Ti isotopic anomalies. These observations clearly establish a link between the preservation of isotopic anomalies 60 TREVOR R. IRELAND

Figure 23. Ion microprobe measurements O isotopic compositions of seven hibonites from the Murchison and Murray carbonaceous chondrites. Despite large variations in their Ti isotopic compositions, ranging from enrichments to depletions, all of the hibonites are enriched in 16O to a degree not represented in conventional analyses of the larger Allende inclusions. Also shown are conventional analyses of a Murchison spinel fraction, FUN inclusions EK1-4-1, C1, and HAL, and terrestrial Burma spinel, New York spinel, and Madagascar hibonite which have been used as ion microprobe standards. Figure adapted from Fahey et al. (1987a). and the physicochemical processes responsible for the formation and processing of the precursors of refractory inclusions. However, while such a link can be demonstrated, converting these associations into a formation mechanism is not so readily forthcoming. The nature of the carriers of the isotopically anomalous material, and the timing and location of the high temperature event(s) responsible for the trace-element fractionations are still matters of conjecture. However, a link with a discrete formation mechanism can be proposed for another set of refractory objects, the fractionated and unknown nuclear (FUN) inclusions (Wasserburg et al., 1977). In these inclusions, which are not markedly different mineralogically or morphologically to other Allende refractory inclusions, isotopic anomalies in every element analyzed thus far are accompanied by mass dependent fractionation effects with enrichments in the heavy isotopes. This heavy isotope enrichment is a characteristic feature of residues left after partial evaporation (Clayton et al., 1985; Davis et al., 1990). One of the FUN inclusions, HAL, has been extensively studied conventionally and by ion microprobe for its isotopic and chemical characteristics (Hinton and Bischoff, 1984; Hinton et al., 1988; Fahey et al., 1987b). Three other hibonite inclusions are similar to HAL, viz. DH-H1 from Ion Microprobe Mass Spectrometry 61

Figure 24. Mg isotopic compositions of two types of hibonite-bearing inclusions from CM meteorites. PLACs are PLAty hibonite crystal fragments and are characterized by high Al/Mg and generally low 26Al/27Al of less than 1 ¥ 10-5 while SHIBs are spinel- hibonite inclusions with generally low Al/Mg and 26Al/27Al around 5 ¥ 10-5. Figure adapted from Ireland (1990). the Dhajala H3 chondrite (Hinton and Bischoff, 1984; Hinton et al., 1988) and 7- 404 and 7-971 from Murchison (Ireland et al., 1988). Ireland et al. (1992) analyzed O, Mg, Ca, Ti-isotopic compositions in these inclusions and found that O, Ca, and Ti were substantially mass fractionated in favor of the heavy isotopes but only 7-971 had mass fractionated Mg (Ireland and Compston, 1987). This is probably due to the almost total evaporation of Mg from these inclusions. While Ca and Ti isotopic mass fractionations are not correlated, the Ti isotopic mass fractionation is inversely correlated with Ti concentration in the four inclusions (Figure 26) reinforcing the conclusion that distillation has played a major role in their formation. Corundum-bearing inclusions are relatively rare but are important since they are likely to be higher temperature objects than the hibonite-bearing inclusions and, by extension of the relationship between the large inclusions from Allende and the hibonite inclusions from Murchison, may have extremely large isotopic effects. The 62 TREVOR R. IRELAND

Figure 25. REE patterns of a representative PLAC and SHIB. PLACs typically show deficits in Eu, Yb and the less refractory of the trace elements (cf. Allende Group III in 35) whereas SHIBs show deficits in the ultra refractory REE Gd-Er and Lu as well as Sc, Y, Zr, and Hf. Data from Ireland et al. (1988). inclusion BB-5, which has the largest deficits of 48Ca and 50Ti and the largest excess of 16O, contains substantial amounts of corundum. Bar-Matthews et al. (1982) first described this inclusion and used the Chicago AEI IM-20 ion microprobe to analyze its Mg isotopic composition. Despite the extreme Al/24Mg ratios in the corundum, 26 26 27 the d Mg was only slightly above normal corresponding to a maximum ( Al/ Al)0 of only 1.5 ¥ 10-8. Hinton et al. (1988) analyzed BB-5 and another corundum- hibonite inclusion GR-1 for their trace element abundances and found evidence for small ultrarefractory phases (possibly composed of Zr oxide and hibonite respectively). A more systematic study of a large number of corundum grains from Murchison was carried out by Virag et al. (1991). Twenty-six corundum grains were identified from acid residues that had been produced to concentrate interstellar carbon-bear- ing grains and so the petrographic context of the grains had been lost. Virag et al. (1991) were able to divide the corundum grains into three groups on the basis of O Ion Microprobe Mass Spectrometry 63

Figure 26. The inverse correlation between Ti isotopic mass fractionation, FTi, and Ti concentration in HAL-type hibonites is suggestive of mass loss by evaporation. The two cases shown are (1) measured compositions and (2) TiO2 concentrations corrected for total mass loss. Figure adapted from Ireland et al. (1992). and Mg isotopic compositions and trace element abundances. Some corundum 26 27 -5 26 24 grains have ( Al/ Al)0 of 5 ¥ 10 which yields Mg/ Mg ratios up to 56 ¥ solar because of the high Al/Mg ratios of corundum. The O-isotopic compositions are anomalous but are still similar to those of other refractory inclusions in that the predominant effect is enrichment of 16O by up to 7 %. The corundum in these grains therefore appears to have a similar origin to the other phases identified in refractory inclusions. However, one corundum grain from the Orgueil CI chondrite has a 26Mg*/27Al of 9 ¥ 10-4 far in excess of the canonical value of 5 ¥ 10-5 and this might be the only oxide grain analyzed to date for which a presolar origin could be proposed (Huss et al., 1992). Nittler et al. (1993a) have also analyzed a possible presolar corundum grain from the Murchison meteorite which had the same 26Mg*/27Al of 9 ¥ 10-4 as the Orgueil grain analyzed by Huss et al. (1992). Nittler et al. (1993a) also measured the oxygen isotopic composition and found it to be highly anomalous with a d17O in excess of 1000 ‰ and a d18O of -200 ‰, clearly indicative of an interstellar origin. Another component within refractory inclusions which may have exotic origins are Fremdlinge, complex aggregates of metal grains, sulfides, phosphates, oxides, and silicates rich in refractory siderophile elements such as Pt, Ir, Os, Re, and Ru. 64 TREVOR R. IRELAND

Their origins are enigmatic since they contain highly reduced as well as highly oxidized phases and both ultrarefractory- and volatile-rich components. One proposal was that they were presolar in origin and should therefore have large isotopic anomalies, however Hutcheon et al. (1987) used a CAMECA ims-3f ion microprobe to show that there were no isotopic anomalies for Mg, Fe, Mo, Ru, and W and therefore it was likely that the Fremdlinge did form within the early solar nebula. Ion microprobe analysis also has a role to play in the analysis of the larger refractory inclusions as well. These objects appear to have been subjected to more than one thermal event and deciphering their histories is reliant on analyzing selected portions of inclusions. In particular the rims of these objects appear to have some genetic association with the cores, for example Boynton and Wark (1987) measured REE in the rim of an Efremovka refractory inclusion by INAA and argued that the five fold enrichment of the REE in the rim was due to the flash heating of the core. On the other hand, Fahey et al. (1987c) used a CAMECA ims-3f to measure Mg isotopes and REE abundances in the core and rim, and concluded that the rim of this inclusion could not have formed by flash heating but rather by condensation from another reservoir which had higher (26Al/27Al)0 than the reservoir from which the core formed.

Interplanetary Dust Particles. Of the hundreds of tons of extraterrestrial material that the Earth captures each day, only a small fraction reaches the ground in a macroscopic form that can be recovered and identified. A significant fraction of the material is in the 1 to 10 µm size range and are known as stratospheric or interplanetary dust particles (IDPs). The importance of this class of material is that it may represent material that is fundamentally different than that preserved in the larger meteorite bodies that reach the earth. Specifically, at least a fraction of this material may be derived from comets and therefore yield material that has escaped the high temperature processes at the beginning of the solar system. The collection of these particles is non-trivial since man-made material can contaminate the samples. In order to prevent contamination, particles are collected on “flags” which are mounted on high-flying air craft (>20 km) and deployed for the high altitude sections of the flight. These particles are caught in a silicone-oil medium and then are picked off the mount and washed in xylene. The particles are analyzed by as many non-destructive techniques as possible before being allowed into an ion microprobe for destructive analysis (McKeegan et al., 1985). Fragments are analyzed in the SEM and are classified as chondritic if all chondritic major elements are observable by EDX analysis. The mineral assemblages in these particles were thought to be different to previously analyzed meteoritic materials, and so isotopic measurements would be of great interest. The small size of these grains (often less than 10 µm) means that ion probe analysis is the only means of analyzing a number of isotopic systems and finding correlated isotopic effects. McKeegan et al. (1985) measured H and C (Cs+ Ion Microprobe Mass Spectrometry 65 primary beam, negative secondaries) and Mg and Si (O– primary beam, positive secondaries) isotopic compositions from a suite of eight chondritic particles. The D/H measurements of the individual grains showed that large isotopic anomalies were commonly present with 5 out of the 8 grains having enrichments over 500 ‰ with a maximum enrichment over 2000 ‰. However, different fragments of the same inclusion showed markedly different compositions and so even the extreme values measured were mixtures of even more exotic components. Three of the grains with heavy hydrogen were also analyzed for C and in comparison to the large D/H variations, the range in C isotopic compositions is more modest, ≈ 40 ‰. The C compositions of different fragments of the same inclusion appeared to be the same in contrast to the H isotopic heterogeneities. Three particles were also measured for Mg and Si isotopic compositions but no anomalies beyond analytical errors were found. McKeegan et al. (1985) also used the ion microscope mode of the CAMECA ims-3f to image the distribution of H, C, O, Si, and S in the particles (Figure 27). They found that there was a strong correlation in the distributions of C and H, and

Figure 27. Negative secondary ion images of squashed Mosquito, an IDP fragment pressed into gold foil. Also shown is an optical photomicrograph of the same region. dD in this particle ranges up to 2500 ‰. (Figure courtesy of K. McKeegan). 66 TREVOR R. IRELAND that the D enrichment was strongly correlated with 12C/ 16O, indicating the deuterium carrier was probably carbonaceous in nature. While these are extreme enrichments in deuterium, large enrichments have been measured conventionally from bulk samples of Renazzo and Semarkona, and ion probe measurements of individual particles of these meteorites have extreme dD values up to 3700 ‰ (Hinton et al., 1983; Zinner et al., 1983; McKeegan and Zinner, 1984). Therefore, despite the large deuterium enrichments in the IDPs, they are not markedly more anomalous than some meteorite matrices. McKeegan (1987) measured Mg and O isotopic compositions in five IDPs that had refractory bulk compositions; two contained corundum, one contained hibonite, and a further two grains were composed of a spinel-hibonite-perovskite-melilite assemblage. Four of the particles had 16O excesses similar to refractory inclusions from carbonaceous chondrites, while a fifth, one of the corundum particles, had normal oxygen and was probably a terrestrial contaminant. All four extraterrestrial particles had normal Mg isotopic compositions. At this stage, while there are mineralogical and petrographic distinctions between IDPs and other meteoritic classes of material, there appear to be no major isotopic distinctions.

Interstellar Grains. While refractory inclusions in carbonaceous chondrites contain isotopically anomalous material, it is likely that they acquired their present form during high temperature processes in the early solar nebula and are not, themselves, interstellar grains. Unfortunately, it is likely that the interstellar precursors to the refractory inclusions are probably oxides and therefore they are swamped by an overbearing proportion of material that has been processed within the solar nebula. Recently, two needles in the oxide haystack have been identified; Huss et al. (1992; 1993) and Nittler et al. (1993a) discovered corundum grains with extreme excesses of 26Mg and 17O. However, it turns out that nature has not been so unkind as to bury all interstellar grains. E. Anders and coworkers isolated to levels of increasing purity the carriers of the exotic noble gas components observed in bulk meteorites. These carriers were highly resistant to a variety of acid treatments and along with separation by size, density, and flotation properties, the noble gases were traced to a series of C-bearing phases. The first phase identified was Cd, which was mineralogically a form of diamond (Lewis et al., 1987). This was the carrier of a component called Xe-HL, so named because the Xe isotopic composition shows large enrichments in both the heavy and light isotopes of Xe. These diamonds are exceedingly small - 2 nm on average - and their chemistry is dominated by the dangling bonds on the surface (Bernatowicz et al., 1990). In tracing a component called Ne-E(H), - nearly mono- isotopic 22Ne, and Xe-S, enriched in 128Xe, 130Xe, and 132Xe - a residue containing abundant SiC was produced (Bernatowicz et al., 1987). The SiC grains cover a large size range from submicron to over 20 µm in diameter. In following Ne-E(L), again essentially monoisotopic 22Ne but in a low density fraction, interstellar graphite Ion Microprobe Mass Spectrometry 67 was identified (Amari et al., 1990). The graphite occurs as 1-4 µm grains with a range of morphologies from euhedral platy grains to compact spherules and spherulitic aggregates. A fourth interstellar phase was subsequently identified as 7- 21 nm inclusions in isotopically anomalous graphite spherules and were identified as TiC based on their electron diffraction patterns (Bernatowicz et al., 1991). The fact that these first interstellar grains are carbon-bearing is no coincidence. The early solar system was an oxidizing environment and the stable phases in the solar nebula are oxides and silicates. Diamond, graphite and SiC are unstable in solar nebula conditions and are stable condensates only when C/O is greater than 0.83 (Larimer and Bartholomay, 1979) whereas the solar ratio is only 0.42 (Anders and Grevesse, 1989). Therefore interstellar carbon phases are undiluted in the early solar nebula by locally produced material. While they are not chemically stable, these phases are highly resistant to physical, chemical, and thermal processes and so they can survive through processes in the interstellar medium, the solar nebula, as well as the chemical treatment for purifying them (Zinner et al., 1987). The origin of these grains must therefore be in stars with high C/O such as C-rich red giants, however the exact nature of the nucleosynthetic environment is perhaps best constrained by their isotopic signatures. These interstellar grains do not comprise a large fraction of meteorites; their abundances range from 400 ppm for diamond, 6-9 ppm for SiC and <2 ppm for graphite (Amari et al., 1990); TiC comprises only a few tens of ppm of the graphite spherules in which it was observed (Bernatowicz et al., 1991). These materials are therefore very precious and the maximum information should be extracted from them. In this regard ion microprobe analysis has been an essential tool in determining the isotopic systematics of these grains particularly where heterogeneous populations are present. The first isotopic measurements of Cd were carried out by stepped pyrolysis of bulk samples. Somewhat surprisingly, the average d13C of –38 ‰ in interstellar diamond is within the terrestrial range (Swart et al., 1983) although the minimum d15N is anomalous at –330 ‰ (Lewis et al., 1983) and the N composition was observed to be highly heterogeneous between individual steps of the pyrolysis. Since the diamonds are so small, it is possible that the C values represent an average for which differences might be resolved on the scale of ion microprobe analysis. However, it should be noted that while an ion microprobe analysis consumes are far smaller quantity of material, it is still a bulk analysis in terms of the size of the diamond grains and therefore it too will represent an average composition. Virag et al. (1989) reported C and N isotopic compositions from six Cd samples from the Allende (4) and Murray (2) carbonaceous chondrites, and dD from two of the Allende separates. The samples were loaded as 20 µm agglomerates onto gold foil and all elements were measured as negative ions. The d13C of the four Allende samples is the same within error (–32 to –36 ‰) and are also consistent with the bulk measurements previously reported. However, one of the Cd residues from Murray, CN, appears to be substantially lighter at –40.2 ± 2.8 ‰ 2s. The source of 68 TREVOR R. IRELAND this deviation is unclear but hints at some form of C-isotopic heterogeneity between the diamond samples. All residues are depleted in 15N but there are substantial variations in d15N between the four Allende samples and the two from Murray although not as great as the differences observed by stepped combustion. The average of all six samples agrees well with the average d15N of –164 ‰ measured by Lewis et al. (1983) from Allende Cd residue CC but the ion probe analyses fall far short of the minimum d15N of –330 ‰. It is possible that individual diamond grains are heterogeneous and the stepped combustion technique can separate the different N components based on the thermal release characteristics. In this regard, the high H concentrations (10-40 at%) of the Cd samples may be important since this shows that preponderance of surface bonds on the diamonds and the chances are quite high for contamination of the N and H isotopic compositions either in the interstellar medium or in the beaker during the acid digestion.

Figure 28. C- and N- isotopic ratios of individual SiC grains from Murchison separate KJH (average size 4.6 µm). The thin lines dividing the plots into quadrants depict solar isotopic ratios. Note the extreme range of the isotopic compositions and the use of a log scale. Most grains fall into the left upper quadrant, grains in the right upper quadrant with 12C/13C > 150 are named grains Y (they are also distinguished by unusual Si-isotopic compositions), those in the right lower quadrant grains X (they also deviate in their Si-isotopic ratios from the bulk of the SiC grains). (Figure adapted from Hoppe et al., 1993). Ion Microprobe Mass Spectrometry 69

SiC has probably been the most intriguing of the interstellar grains studied so far. This is in part due to the relatively large size of these grains which enables different isotopic analyses to be made on individual grains. Correlations can then be drawn between the different isotopic compositions which can then be examined in the light of theoretical nucleosynthetic calculations. The first analyses of Si, C, and N isotopic compositions in SiC (Zinner et al., 1987) exceeded the largest anomalies in these elements by factors of up to 50. While agglomerates of a fine-grain-size fraction were relatively homogeneous, analyses of clusters of coarse (>2 µm) SiC scattered widely in their isotopic compositions. The Si in both samples was highly anomalous lying off the terrestrial mass fractionation line indicating that the compositions cannot be derived from terrestrial silicon by mass fractionation. Most of the compositions lay on the 29Si enriched side of the terrestrial mass fractionation line. The C-N isotopic compositions were characterized by enrichments in 13C and depletions in 15N. While the majority of grains still occupy the heavy C, light N quadrant, there are representatives now in all four quadrants (Figure 28). These compositions represent diverse nucleosynthetic conditions and it is highly unlikely that all grains were

Figure 29. Si isotopic compositions of individual SiC grains from Murchison separates KJH (average size 4.6 µm) and KJG (3.0 µm). Most grains plot close to a line with slope 1.34 that does not go through the solar composition. Grains Y plot on a separate line with slope 0.35 (broken line) indicating that they belong to a separate population; and grains X have Si isotopic compositions substantially enriched in 28Si. (Figure adapted from Hoppe et al., 1993). 70 TREVOR R. IRELAND

Figure 30. C- and N- isotopic compositions of individual graphite grains from Murchison separate LFC (1-4 µm diameter). Only the compact grains and dense clusters show large C-isotopic anomalies but N is rather indeterminate because of N contamination from nearby nitrogen-rich aggregates. The dominant nucleosynthetic processes in each quadrant are indicated as is the composition of micro diamonds (Cd). Figure adapted from Amari et al., 1990). produced in a single star. Zinner et al. (1989) argued that the Si isotopic compositions would not be affected in the nucleosynthetic conditions responsible for the C and N compositions and so individual compositions represented different stars. However, an alternate viewpoint has been put forward by Stone et al. (1991). They found that a texturally distinctive type of SiC showed correlated 29Si and 30Si effects that were consistent with two component mixing possibly within the same star. However, it is clear from the diversity of compositions now measured that a number of stellar sources are required (Figure 29). The isotopic compositions of C, N, and Si have been used to distinguish distinct subsets of grains which probably reflect their origins in distinct nucleosynthetic environments. After ion probe analyses of around 100 grains Zinner et al. (1991b) noted that two grains had completely different characteristics to the rest. The majority have heavy C and light N and heavy Si while the two extremely exotic grains (termed grains X) have light C, heavy N, and light Si. After measuring nearly 700 SiC grains, Amari et al.(1992a) discovered 3 further X grains. Another type of SiC identified by Amari et al. (1992a) were grains Y which lie to the right of the main SiC trend on a Si 3-isotope plot and also have light C. Ion Microprobe Mass Spectrometry 71

Figure 31. Mg isotopic compositions in SiC and graphite often show large excesses of 26Mg and yield 26Al/27Al ratios ranging up to 0.2, 4000 times the canonical maximum value of 5 ¥ 10-5 for refractory inclusions. (Figure adapted from Zinner et al., 1991a).

Graphite C isotopic compositions deviate from terrestrial even further than those of the SiC grains with enrichments and depletions of 13C by up to a factor of twenty (Amari et al., 1990). However, these effects are only in one morphological type of graphite, the rounded grains. While the C compositions are highly variable, N isotopic compositions appear to be more uniform (at least in terms of a log ratio scale; Figure 30) and are slightly heavy (to a factor of two). A number of other minor elements have also been analyzed for their isotopic compositions in SiC and graphite grains. A startling find was that large amounts of extinct 26Al (i.e. 26Mg) were present in both SiC and graphite (Zinner et al., 1991a). Initial Mg in these grains is very low and 26Mg/24Mg ratios ranged up to nearly 1000 (7000 ¥ solar). The resulting 26Al/27Al ratios range up to 0.06 in graphite and 0.2 in SiC of type X (Figure 31), well beyond the canonical 5¥10-5 value for the early solar system found in refractory oxide inclusions. The Al abundance in these grains appears to be correlated with the N abundance suggesting that aluminium nitride may be the condensed phase, however no discrete particles of Al-nitride within the graphite spherules have been observed as yet. Titanium is also often present in sufficient concentrations to allow ion probe isotopic measurements. Ti concentrations of SiC grains analyzed by Ireland et al. (1991b) range from ≈30 ppm to several thousand ppm. While the Ti concentrations 72 TREVOR R. IRELAND

Figure 32. Most SiC grains have V-shaped isotopic patterns with excesses of all isotopes relative to 48Ti compared to solar ratios. Note however that in one grain in (a) the pattern is inverted. Grains X have large excesses of 49Ti which may be due to the decay of 49V. (Figure adapted from Amari et al.,1992). are low, so too are the concentrations of the elements which have isobaric contributions, Ca, V, and Cr, and so the analyses are generally limited by counting statistics. The Ti isotopic compositions of SiC are distinct from those in refractory- oxide inclusions by showing large anomalies in all ratios not just 50Ti (Ireland et al., 1991b). The characteristic feature of the majority of the grains is a V-shaped pattern with enrichments in 46Ti and 50Ti of up to 300 ‰ (Figure 32). This pattern is characteristic of nucleosynthesis in an environment with slow addition of neutrons (s-process) whereby the isotopic abundances are controlled by the neutron capture cross sections. A different signature is present in grains X which have large 49Ti and 44Ca excesses (Amari et al., 1992b) (Figure 32). These two isotopes have radiogenic precursors in 44Ti and 49V which are both p-process (i.e. they lie to the proton-rich side of the main stream of stable nuclides) and they have the longest half-lives (47 years and 331 days respectively) of the p-process radionuclides in this region of the chart of the nuclides. These two nuclides are produced in supernova ejecta, however, not all of the grains X show 44Ca excesses and the 26Al/27Al ratios inferred for these grains are too high for supernova production models. The isotopic compositions of some heavier nuclides have also been measured by ion microprobe. Zinner et al. (1991c) measured Ba and Nd isotopic compositions Ion Microprobe Mass Spectrometry 73 in various fractions of SiC and showed the s-process nature of their compositions. These compositions have also been determined by conventional mass spectrometry with good agreement between the results (Ott and Begemann, 1990; Prombo et al., 1992; Richter et al., 1992). The isotopic data coming out of these interstellar grains is of paramount importance in understanding the nuclear reactions that are going on in their source stars. Gallino et al. (1990) have already attempted to model the measured isotopic compositions of Kr and predict the isotopic compositions of a wide variety of elements. Now high precision isotopic ratios are available for a number of elements from an extra-solar source with which we can compare theoretical and measured values.

Short-livednuclides in theearly solar system. Short-lived nuclides are important for models of the evolution of the solar system since they can act as chronometers for thermal events and heat sources for melting planetessimals. In this paper we are concerned with the role of ion probe analysis in the detection of these nuclides; a more general recent review of extinct radionuclides is given by Podosek and Swindle (1988). The prime advantage of the ion microprobe in these studies is being able to select a very small volume which has a very high parent daughter ratio so that the maximum effects in the isotopic composition of the daughter can be measured. Excesses of a given radionuclide in a single phase of a refractory inclusion can only give an indication that the isotope was alive, an alternative possibility is that the phase inherited the daughter during a later thermal event. However, the finding of excess 26Mg which is correlated with 27Al/24Mg in different phases gives the strongest possible evidence that 26Al was alive in refractory inclusions. With the finding of live 26Al in refractory inclusions, it was proposed that 26Al could be a suitable heat source for melting planetessimals. With an abundance of 5¥10-5¥27Al, there would be 26 sufficient heat generated from Al decay (t1/2 = 0.7 Ma) to melt a small planetary body . However, 26Mg excesses at this level have only been found in refractory inclusions from carbonaceous chondrites. The only “non-refractory” assemblage in which excess 26Mg has been found is a olivine-pyroxene clast containing plagioclase from the Semarkona ordinary chondrite (Hutcheon and Hutchison, 26 27 -6 1989). For this assemblage a ( Al/ Al)0 of 7.7 ± 2.1 ¥ 10 (Figure 33) is sufficient to produce incipient melting in well-insulated bodies of chondritic composition. However, excess 26Mg has never been found in a differentiated meteorite and the role of 26Al in planetary heating remains open. The other potential use of short-lived radionuclides is for short term chronome- ters of early solar system events. With a half-life of 0.7 Ma, the apparent 26Al abundance can potentially yield very high age resolution provided it can be shown that it was uniformly distributed throughout the solar system. The current data from refractory inclusions indicates that 26Al was not homogeneously distributed and 74 TREVOR R. IRELAND

Figure 33. Al-Mg diagram for coexisting olivine, pyroxene, and anorthite in Semarkona chondrule CC-1. The linear correlation between excess 26Mg and Al/Mg is consistent with the in situ decay of 26Al with an initial 26Al/27Al of 7.7±2.1 ¥ 10-6. The dashed line represents the canonical 26Al/27Al of 5 ¥ 10-5 found in coarse-grained Allende inclusions. (Figure adapted from Hutcheon and Hutchison, 1989). therefore its use as a chronometer may be limited. However, there may be some potential use of 26Al in determining the duration of events in the solar nebula by analyzing related objects. In this case it may be possible to argue for a uniform 26 distribution of Al. Podosek et al. (1991) have used ion microprobe Al-Mg isotopic data and conventional Rb-Sr data in an attempt to see if the systems are correlated within a suite of coarse grained refractory inclusions. The inclusions analyzed ranged from relatively pristine samples of basically igneous droplets, to highly altered inclusions whose textures appear to be metamorphic in origin. However, both Al-Mg and Rb-Sr systematics appear to have been affected in all inclusions but nonetheless the comparison of both isotopic systems did not reveal any chronological inconsistencies in terms of a heterogeneous distribution of 26Al. 53 53 Excesses of Cr due to the decay of Mn (t1/2 = 3.7 Ma) have also been found in chondritic meteorites. The analysis of Mn-bearing sulfides in enstatite chondrites by ion microprobe has been particularly fruitful in the search for excess 53Mn since Cr is incompatible in the sulfide phases and 55Mn/52Cr ratios of up to 4 ¥ 106 have been measured in sphalerite from EL3 chondrites (El Goresy et al., 1992) (Figure 34). The 53Mn/55Mn ratios are quite variable within individual meteorites as well as between meteorites ranging from < 10-7 to near 10-5 (Zinner et al., 1991d; El Goresy et al., 1992) compared with the abundance in refractory inclusions of around 4 ¥ 10-5 (Birck and Allègre, 1985; Birck and Allègre, 1988). Hutcheon and 53 Olsen (1991) measured excess Cr in differentiated meteorites including iron meteorites and determined 53Mn/55Mn ratios of up to 2 ¥ 10-5, which is close to the Ion Microprobe Mass Spectrometry 75

Figure 34. 53Mn - 53Cr systematics of Mn-bearing sulfides from MAC88136. The high Mn/Cr in these sulfides (up to 4 ¥ 106) in these phases makes them ideal candidates in the search for 53Cr excesses due to the decay of 53Mn. While excesses of 53Mn are clearly detected, there appears to have been remobilization of Cr after 53Mn decay since the 53Mn/55Mn ratio in the sphalerites ranges from 5.4 ¥ 10-8 to 1.6 ¥ 10-7. Alabandites show even higher 53Mn/55Mn ratios up to 4.5 ¥ 10-7. Figure adapted from El Goresy et al. (1992). values from Allende refractory inclusions, down to 8 ¥ 10-7. 53Mn was clearly alive in a wide variety of early solar system objects but its use as a chronometer is still being evaluated. 60 60 Attempts have also been made to find Ni excesses from the decay of Fe (t1/2 = 1.5 Ma) by ion microprobe. Hyman et al. (1988) measured Ni isotopic compositions in individual Orgueil magnetite grains that had 56Fe+/62Ni+ ratios of up to 1.4 ¥ 105 but could find no excesses attributable to 60Fe decay; the upper limit for 60Fe/56Fe was ≈10-4. Evidence for live 60Fe at a level of 1.6 ¥ 10-6 ¥ 56Fe has been found recently in the Chervony Kut eucrite by Shukolyukov and Lugmair (1992) who used conventional thermal ionization mass spectrometry. An important aspect of this work is that the 60Fe is alive in a differentiated meteorite and at this abundance level, 60Fe could be a very important heat source in planetary differentiation. Hutcheon et al. (1984) measured the K-isotopic composition in two Allende 41 refractory inclusions in an attempt to find evidence for live Ca (t1/2 = 0.1 Ma). This is a difficult measurement because of the presence of large isobaric interferences 76 TREVOR R. IRELAND at mass 41 (e.g. 25Mg16O and 40CaH). They measured excesses at 41K of up to 1400 ‰ at 40Ca/39K ratios up to 1.6 ¥ 107 yielding a 41Ca/40Ca ratio of around 5 ¥ 10-8. Ireland (1991) analyzed meteoritic zircons in an attempt to find excess 182W 182 which would be indicative of the presence of Hf (t1/2 = 9 Ma) in the early solar system. This radionuclide is a key indicator to the last contribution of r-process material to the solar system. However, the measurement of W isotopic compositions in zircons is difficult because of the presence of molecular interferences from the REE elements. The limiting interference for these measurements appears to have 180 + 182 been HfH2 and therefore the measurements of the W signal represent an upper limit to its abundance. No effects were resolved relative to terrestrial standards in two zircons with 180Hf+/184W+ of around 8 ¥ 104 yielding an upper limit of around 5¥10-5 for the 182Hf/180Hf. This ratio corresponds to a free-decay interval of no less than 120 Ma for a production ratio of 182Hf/180Hf ≈ 0.5 which is consistent with estimates based on the 244Pu abundance.

Chemistry of Solar System Materials The chemical systematics of early solar system materials are affected by a number of high temperature processes but the effects can be viewed as being related to either volatility-related (solid-gas or liquid-gas fractionation) or igneous (solid- melt fractionation) mechanisms. Volatility fractionations, particularly in the refractory elements, are particularly well preserved in refractory inclusions of carbonaceous chondrites while igneous fractionations can be found in a large range of meteorites and meteoritic components. The trace-element signatures tell us a great deal about the processes that were responsible for the formation of these objects.

Volatility Fractionations in Refractory Inclusions. The major element chemistry of refractory inclusions is dominated by the refractory oxides of Mg, Al, Si, Ca, and Ti, but it is in the refractory trace elements that the most diagnostic signatures of high temperature processes are found. The REE patterns of terrestrial minerals are largely orderly functions of ionic radius but in refractory inclusions large variations are apparent that are strongly correlated with volatility. Allende inclusions were classified into six groups according to their REE patterns [for a summary see Mason and Taylor (1982)]. Inclusions belonging to Groups I, III, V, and VI show an overall enrichment of the REE by a factor of 10-20 relative to CI-chondrite abundances and are characterized by the relative abundances of Eu and Yb. Variations in Eu abundances in terrestrial materials are usually caused by the tendency of Eu to exist in the divalent oxidation state, but the operative parameter under solar nebula conditions is that Eu and Yb are the most volatile of the REE. Group V inclusions show no anomalies, Group I has a positive Eu anomaly, Group VI has positive Eu and Yb anomalies and Group III has negative Eu and Yb anomalies (Figure 35a). Group II patterns have completely different characteristics. The patterns of this group have depletions in the most volatile elements (Eu and Yb) as well as large depletions in the most refractory elements Ion Microprobe Mass Spectrometry 77

Figure 35. Allende refractory inclusion classification based on REE abundance patterns. Allende inclusions are large and can be analyzed by conventional methods. Groups I, III, V, and VI have relatively flat patterns at 20 ¥ CI abundances with the abundances of the relatively volatile REE Eu and Yb being the distinguishing characteristics. Group II shows a depletion in the most refractory elements, Gd-Er and Lu, as well as the most volatile, Eu and Yb and must be obtained from the fractional condensation of a gas from which an ultra refractory component was removed. The Group IV pattern is obtained from olivine-rich chondrules.

(Gd, Tb, Dy, Ho, Er, Lu), so that the characteristic features are a flat light REE pattern and a Tm anomaly (Figure 35b). This pattern was attributed to condensation of the inclusion after the removal of a small amount of the most refractory material from the gas phase (Boynton, 1975). The ultrarefractory pattern complementary to the Group II pattern has never been found in Allende. It was first identified in inclusion MH-115 which was one of the first inclusions analyzed from the Murchison carbonaceous chondrite (Boynton et al., 1980). This inclusion contained hibonite and so it was not surprising that a more refractory inclusion composed of hibonite should have a more refractory trace element pattern. However, most Murchison inclusions have patterns that are similar 78 TREVOR R. IRELAND

Figure 36. Ion microprobe analysis of smaller CM meteorite inclusions revealed a large range of patterns, three of which are shown here. (a) The ultra refractory-enriched pattern is complementary to the Allende Group II pattern. (b) This pattern shows the effects of a high local oxygen fugacity in which Ce and Pr become as volatile as Eu and Yb. (c) This pattern shows deficits of Ce and Yb but Eu is not anomalous. (Data from Ireland et al.,1988; Ireland, 1990). to those observed in Allende inclusions. Initially, a lot of painstaking work analyzing small amounts of refractory inclusions were carried out by neutron activation analyses, for example Ekambaram et al. (1984b). However, ion microprobe analysis of these inclusions has now become rather routine and has the benefit of allowing analyses of very small inclusions which can also be analyzed for their isotopic abundances. Around one hundred inclusions have now been analyzed for their trace-element characteristics by ion microprobe. Most of these have been hibonite-bearing inclusions from Murchison but inclusions from other meteorites have also been studied (Fegley and Ireland, 1991). The vast majority of refractory inclusions from Murchison have patterns that are similar to Groups II and III whereas the larger Allende inclusions are dominated by Groups I and V. The chemical characteristics are quite strongly correlated with the morphological characteristics with the Group III patterns preserved in PLAC crystal fragments, while Group II patterns are Ion Microprobe Mass Spectrometry 79 commonly preserved in spinel-hibonite SHIB inclusions (Ireland et al., 1988). However there are distinct variations of the Group II and III patterns as well as a variety of patterns that are not seen in Allende (Figure 36). Fegley and Ireland (1991) compiled an inventory of over 280 inclusions from 19 meteorites for which REE patterns had been determined and found that the Allende group classification scheme was inadequate. They proposed an alternative classification scheme which was not reliant on adding groups or having modified patterns from the existing scheme. They noted that the patterns can basically be described in terms of three processes: fractionation affecting the ultrarefractory elements Gd-Er and Lu, fractionation of the relatively volatile REE Ce, Eu, and Yb, and an overall fractionation that affects the slope but does not produce anomalies (igneous fractionation). From the tabulation of patterns, it was found that fractionated patterns (in terms of the ultrarefractory elements) comprised nearly half of all patterns measured with most being depleted (37%) relative to enriched (7%). As well as being invaluable in analyzing small inclusions, ion microprobe analysis has also enabled the crystallization history of larger inclusions to be examined. The large Type B Allende inclusions have the textural characteristics of being produced by solidification of molten droplets. However, they have some unusual features such as thick melilite-rich rims, and spinel-free islands which are difficult to rationalize in a traditional fractional crystallization model. However, Beckett et al. (1990) produced synthetically zoned melilite crystals and were able to show that the distribution coefficients were a function of the composition of the crystallizing melilite, for example Be is incompatible in gehlinitic melilites, but compatible in åkermanitic melilites (Figure 37). Similar results were obtained by Kuehner et al. (1989). The compositions measured from meteoritic melilite-rich inclusions (Beckett et al., 1990; MacPherson et al., 1989) are broadly consistent with sequential crystallization from a single melt, but the spinel-free islands are probably trapped xenoliths which were partially assimilated (MacPherson et al., 1989). PLAC hibonites also have a roll-off in the heavy REE abundances which is suggestive of igneous fractionation in their formation, however the phase in which the heavy REE were accommodated has been lost (Ireland et al., 1988). The other phase may have been a glass since this is observed in microspherules from Murchison and ALH85085 but of the three Murchison hibonite-glass microspherules analyzed by Ireland et al. (1991a), only one shows the partitioning expected between hibonite and silicate melt (Drake and Boynton, 1988). The most unusual features are shown by spherule 7-753 for which both hibonite and glass have ultrarefractory-depleted patterns and the differences between the compositions appear to be related to the presence of a Gd-rich ultrarefractory phase as has been observed in some Murchison inclusions such as GR-1 (Hinton et al., 1988), and 13- 60 (Ireland, 1990). HAL and DH-H1 also show substantial igneous fractionation effects and while the heavy REE enriched phase has been lost from DH-H1, the fractionation of the REE in HAL has been successfully modeled by Hinton et al. 80 TREVOR R. IRELAND

Figure 37. Distribution coefficients for Be melilite/liquid as a function of åkermanite content, XAk. Beryllium has the unusual property of being incompatible in melilite with XAk less than 0.5 and compatible at higher XAk. (Figure adapted from Beckett et al., 1990).

(1988) based on partitioning between the core hibonite and perovskite that occurs in the rim layers. Enstatite chondrites, the reduced analog of chondritic meteorites, may also preserve primitive components. These meteorites formed under extremely reducing conditions such that elements which we normally regard as lithophile become siderophile. Calcium is an important example here since the major rare earth element carrier in these meteorites is oldhamite, CaS. Oldhamite reacts with atmospheric water and so is difficult to handle but can be preserved in thin section and analyzed in situ by ion microprobe. Lundberg et al. (1989; 1991) documented four distinct REE patterns in oldhamite from the Qingzhen (EH3) meteorite (A-D in Figure 38) and found that these patterns were associated with oldhamite from different petrographic contexts. The fifth type, E, has ≈500 ¥ CI abundaces for the REE excluding Eu and has not been found in Qingzhen. Oldhamite is possibly a reduced high-temperature condensate that is analogous to hibonite and perovskite in the more-oxidized chondrites, and so these patterns may reflect nebular formation of oldhamite before incorporation into the meteorite.

Igneous Fractionations in Differentiated Bodies. While a lot of REE meas- urements have been directed at primitive components in carbonaceous chondrites, some attention has also been paid to REE inventories in more evolved meteorite types. It is not always easy to identify the trace-element carriers in a rock particu- larly when whole rock analyses show large heterogeneities. This is a common problem in meteorite studies since available material is precious and heterogeneity Ion Microprobe Mass Spectrometry 81

Figure 38. The major REE carrier in enstatite chondrites is oldhamite, CaS, which shows a variety of abundance patterns generally with anomalies only in Eu and Yb. (Figure adapted from Lundberg et al., 1991). can simply be a result of the small sample size that is available for analysis. For rare meteorites, ion microprobe analysis can be used to analyze individual components and the bulk composition reconstructed from the modal abundances. Floss et al. (1990) confirmed that oldhamite was the major REE carrier in the Bishopville aubrite. These meteorites are relatively rare and are the achondritic equivalent to the enstatite chondrites. They are probably of igneous origin, although Floss et al. (1990) found that individual oldhamite inclusions had quite different REE patterns suggesting oldhamite may be a relict phase even in these differentiated meteorites. In this case bulk REE patterns would tell us little about the igneous processes involved in the formation of these rocks. Lundberg et al. (1988; 1990) analyzed the REE carriers in two shergottite meteorites to examine the implications for magma generation and evolution (purportedly) of Mars. Most of the REE reside in whitlockite whose modal concentration is only ≈1 %. REE patterns of individual phases can be used to model the magma composition and the agreement of these model compositions for pyroxene, plagioclase, and whitlockite for ALHA77005 suggests that crystallization took place in a closed system. The mineral REE data are consistent with the petrographic observations that plagioclase and whitlockite crystallized late and ALHA77005 crystallized low-Ca pyroxene before high-Ca pyroxene, whereas nakhlites and chassignites have higher CaO sources and magmas derived from them crystallized high-Ca pyroxene first. Getting closer to home, but still in the extraterrestrial line, analyses of lunar materials by ion microprobe have been extremely useful because of their ability to 82 TREVOR R. IRELAND analyze small volumes. This is important, not only curatorally, but also because of the brecciated nature of the most primitive lunar rocks and the presence of glass microspherules on the lunar surface. These glasses contain much higher siderophile and volatile element abundances than mare basalts and appear to originate from a source region deeper than the devolatilized source of the mare basalts. The primitive nature of these glasses is therefore of great interest since there characteristics might better reflect the bulk chemistry of the Moon for these elements and facilitate comparisons between elemental abundances in the Earth and Moon. Delano and coworkers [e.g. Delano (1986)] have established the presence of 25 chemically distinct types of lunar glass and ion microprobe studies (e.g. Shearer et al., 1990) of the rare earth elements and several other refractory trace elements have identified at least seven types and also indicate heterogeneous source regions at depth in the Moon.

B. in Geochemistry

While terrestrial materials are far more abundant than meteorites and returned lunar samples, far less attention has been paid to them in terms of ion microprobe analysis. In part this is related to the relatively large terrestrial sample sizes available which means conventional analyses can be made without significantly depleting the total inventory of that sample. However, it is also related to the much smaller effects, both chemical and isotopic, that need to be resolved in terrestrial work and hence high precision is of greater importance. While the ion probe offers the best in terms of sensitivity per unit volume, the volumes available to conventional analysis are generally not so restricted and the extreme precision necessary can be achieved. However, when the samples are small, or are shown to be complex and the assumption of sample homogeneity cannot be sustained, then the ion probe can again become a useful tool.

Isotope Geochemistry Applications in the field of isotope geochemistry can be divided into two main fields concerning stable and radiogenic isotopes. Stable isotope studies are based on the physicochemical isotopic mass fractionation of the relatively light elements H, C, O, and S and are used as indicators of the mechanisms of mineral formation, or as reflectors of the formation environment. Radiogenic isotope systems can be used as chronometers or as tracers of petrogenetic processes. In this section, the discussion will be concerned with radiogenic isotopes as tracers and ion microprobe applications in geochronology will be discussed later.

Stable Isotopes. We have seen that extraterrestrial dD values can range up to several thousand permil, however the range for most terrestrial samples is around –200 to + 20 ‰ and therefore higher precision is required to differentiate the range of values. Deloule et al. (1991b) analyzed amphiboles occurring as rare disseminated grains in peridotites from four petrologically distinct settings to a Ion Microprobe Mass Spectrometry 83

precision of ≈10‰. For two of the samples, Massif Central, France and Salt Lake Crater, Hawaii, dD SMOW variations were measured both within single crystals and among crystals from the same sample. Based on the diffusion of H in amphibole, Deloule et al. (1991b) concluded that the dD of the fluids was heterogeneous and the interval between exchange and volcanism was very short (less than a few months). For the Lherz peridotite, France, and Nunivak Island, Alaska, the dD values are within the range shown by uncontaminated mantle minerals. However, the Massif Central xenoliths have a higher dD which may have involved contamination of the lithosphere with subducted sea-water-altered oceanic crust. The Hawaiian peridotites have very light H which may be due to the influence of a deep mantle plume. Chaussidon and Albarède (1990) measured B isotopic compositions in tourmalines from a variety of magmatic and sedimentary rocks. Reproducibility of the measurements on tourmaline standards is cited as being better than ± 1 ‰. The d11B values in the unknown tourmalines range from –30 ‰ to +17 ‰ with individual variations in single tourmalines being less than 5 ‰ despite chemical zoning in some crystals. There was no correlation of d11B with age of the source rocks, and the large variations are related in part to the bulk chemistry of the tourmaline grains. Carbon isotopes were measured in a heterogeneous Bultfontein diamond by Wilding et al. (1990) with a Cs+ beam. Since diamond is electrically conductive sample charging is probably minimal. The precision of each analysis was around ± 0.5 ‰ and the data are believed to be accurate to 1 ‰. They found that the d13C values of the diamond ranged from -5.8 to -10.9 ‰ and showed a consistent correlation with the cathodoluminescent zones indicating either a change in the precipitation conditions or a change in the source C isotopic composition. Harte and Otter (1992) also measured C isotopes in diamonds and found that the K3 diamond alone covered the entire range of diamonds measured by conventional techniques of -9 to -2 ‰. In this and four other diamonds there was no clear correlation with cathodoluminescent zones in contrast to the findings of Wilding et al. (1990). Terrestrial oxygen-isotopic compositions, in terms of the fractionation of 18O/16O, range from around -55 ‰ for Antarctic meteoric water to around +35 ‰ in low temperature minerals. However, the range for lithospheric materials is only around 35 ‰, which is less than 20 ‰/amu fractionation. Therefore, in order to assess even gross O-isotopic heterogeneities in terrestrial rocks, precisions of the order of a few permil are required. So far this has only been achieved for conductive samples such as magnetites and ilmenites. Valley and Graham (1991) found that individual magnetite grains from the Adirondack Mountains, New York, 18 were largely homogeneous with d OSMOW of +8.9 ± 1.0 ‰, but analyses from close to the rims were depleted in 18O. Depth profiles of the rims showed d18O to be 9 ‰ lower than the cores. The difference in d18O between calcite and magnetite (9.3 ‰) yields an apparent equilibration temperature of 525 ˚C, over 200 ˚C below the temperature of regional metamorphism. From their data, Valley and Graham 84 TREVOR R. IRELAND

(1991) concluded that the Adirondacks had experienced a two stage cooling history, either slow cooling followed by rapid cooling at high P(H2O) or slow cooling through the blocking temperature of calcite and magnetite followed by exchange with late low temperature fluids. Sulfur has been the most analyzed of any of the stable isotope systems because of the relatively high isotopic ratio of 34S to 32S (≈0.045) and the large fractionations that are apparent in natural rocks. The total range in d34S in sedimentary sulfides is from –70‰ to +70 ‰ (Ohmoto and Rye, 1979) which is largely due to bacterial fractionation of S. The original approach for S isotopic compositions was to interpret sulfides with d34S around 0 ‰ as magmatic and sulfides with variable d34S values (≥ ± 10 ‰) as sedimentary (Ohmoto and Rye, 1979). However, much more variability is apparent in magmatic systems, due to, for example, kinetic effects, and such interpretations based on S isotopes alone cannot be definitive. S isotopes can also be used for geothermometry, based on the S isotopic fractionation between two coexisting S-bearing minerals. Ion microprobe applications in S-isotope geochemistry have been particularly important in ore studies where large variations in S-isotopic compositions can be found even within the same crystals. The ion microprobe allows measurements of these variations and can be used to pin point the petrographic carrier of the anomalous S. Deloule et al. (1986) combined S- and Pb-isotopic measurements to study the microstratigraphy of galena crystals from Mississippi Valley-type deposits and found that there were frequent changes in the sources of the brines which circulate through the ore deposit localities. The S-isotopic compositions ranged from -6.3 to -12.0 ‰ for the Picher mine and from +11.8 to +20.8 ‰ for the Buick mine. Eldridge and coworkers (Eldridge et al., 1988; 1989; 1993; McKibben and Eldridge, 1989) have analyzed a variety of sediment-hosted massive sulfide deposits including the McArthur River HYC and Mount Isa lead zinc deposits, Australia, the Rammelsberg deposits, Germany, and the Salton Sea geothermal system, California. In all of these cases, more isotopic variation was found in a single thin section than had been reported from the entire bodies by conventional analysis. Of particular importance was the analysis of sulfide from the H.Y.C. deposit McArthur R., since the minerals are noted for their fine grain size (<200 µm in diameter) and lack of metamorphic overprint. The relationships between the different sulfide minerals had remained enigmatic because of the difficulties in obtaining pure mineral separates. Furthermore Eldridge et al. (1993) found that there were two generations of pyrite, py1 a euhedral form and py2 a framboidal form, which would have been mixed in a conventional analysis anyway. Both forms of pyrite are extremely heterogeneous with d34S in py1 ranging from -15 to + 12 ‰ and py2 ranging from -1 to +45 ‰. In comparison, the base metal sulfides range from -5 to +14 ‰ (Figure 39). Macfarlane and Shimizu (1991) analyzed sulfides from vein and stratabound ores from the Hualgayoc district of northern Peru and found the galena from the veins Ion Microprobe Mass Spectrometry 85

Figure 39. S-isotopic compositions in sulfides from the HYC deposit at McArthur River. Minerals shown are chalcopyrite (Ccp), sphalerite (Sp), galena (Gn), and two forms of pyrite (py1, py2). Each of the pyrite types has a skewed isotopic distribution typical of biogenic pyrite formed in a system with limited sulfate supply and the means of py1 and py2 differ by 15 ‰. Typical error for these analyses is ± 2 ‰. (Figure adapted from Eldridge et al., 1993). is about on average 5 ‰ lighter than galena from the stratabound assemblage. The shift in d34S was interpreted as a result of oxidation of the ore fluid during the later stages of deposition. Chaussidon et al. (1989) measured sulfide inclusions in rocks of mantle origin. A range of d34S from -4.9 to +8 ‰ was found with the most heterogeneity being found in sulfide inclusions with low Ni content from pyroxenites, ocean island basalts, and kimberlites, whereas sulfides with high Ni content, mostly peridotitic, had a more limited range from -3.2 to +3.6 ‰. Eldridge et al. (1991) have analyzed sulfide inclusions in diamonds and found a 25‰ range in d34S (from -11 to +14 ‰) which is the largest reported for mantle samples. However the mean d34S is 0.4 ‰ and such heterogeneity could have gone unnoticed in a bulk analysis. The largest variations were preserved in inclusions with Ni contents less than 8 wt% which is a proposed maximum for sulfides of eclogitic affinity whereas the peridotitic inclusions show little deviation from 0 ‰. These results suggest that the sulfides in the diamonds have sampled crustal sulfur possibly in the form of recycled sediments. 86 TREVOR R. IRELAND

Radiogenic Isotopes as Tracers. Lead isotopes are one of the most important tracers of crustal evolution. The Pb isotopic ratios are affected by the ratio of 238U/ 204Pb (µ) and 232Th/204Pb (k) in the source region and are also a function of residence time in a reservoir. Additionally, Pb isotopes can be used to fingerprint ore bodies and the Pb-isotopic compositions of the various economic horizons can be used to identify or at least restrict possible sources. One of the classic studies in secondary ion mass spectrometry is the measurement of differences in Pb-isotopic compositions in different growth zones in a galena from the Buick mine, southeast Missouri (Hart et al., 1981). Variations in the Pb isotopes of up to 4 % in 207Pb/206Pb were found which was more than the total range previously reported from the whole southeast Missouri ore district. Deloule et al. (1986) also measured Pb isotopic compositions in the samples from the Mississippi Valley-type ores they analyzed for S-isotopic compositions. The variability noted in the S was also reflected in the Pb compositions. For the Picher deposit, the Pb and S appear to have been derived from the same source whereas in the Buick mine different sources are required with S being extracted earlier than Pb in the hydrothermal circulation. As well as analyzing S-isotopic compositions of sulfides in diamonds, Eldridge et al. (1991) measured the Pb isotopic compositions of the same inclusions. The Pb data are consistent with the S data in that the peridotitic samples have Pb compositions on the growth curve with average crustal µ at around 2 Gyr ago (Figure 40). The eclogitic samples on the other hand, formed in a reservoir with much higher µ (for an average age of 1 Gyr, µ ≈ 300) further reinforcing the possible contribution of crustal material in the diamond formation region. The possibility of ion probe analysis of strontium isotopic compositions has been examined by Exley (1983) who used a modified AEI-IM20 ion microprobe at low mass resolution. Carbonates were analyzed because of their very low concentrations of 87Rb+ which cannot be resolved from 87Sr+. However, molecular + + interferences predominantly from Ca2 and CaMgO had to be peak stripped from 86 + 87 + 88 + 40 + 40 25 16 + Sr , Sr and Sr by monitoring Ca2 and Ca Mg O respectively. A precision in 87Sr/86Sr of ≈1 ‰ could be obtained for Sr concentrations of greater than 5000 ppm. Ion microprobe analyses were compared with conventional solid source data for a suite of standards with good agreement between the two techniques, and two applications were presented regarding isotopic variation across a calcite vein in a hydrothermally altered basalt and the analysis of 35 µm calcite grains in a lherzolite. Exley and Jones (1983) examined Sr isotopic compositions in kimberlitic carbonates and were able to distinguish between primary carbonates with 87Sr/86Sr of ≈0.705 from secondary calcite with 87Sr/86Sr greater than 0.710. Hafnium isotopic compositions in zircons have been measured by Kinny et al. (1991) and can be used to obtain model Hf ages (from a chondritic reservoir) for protoliths. In this way Hf can be used in the same way as Nd for bulk rock samples. Zircon is an ideal target for Hf isotopic studies since it contains around 1 wt % HfO2 and the Hf model ages can be compared with U-Pb ages of the same areas. Kinny Ion Microprobe Mass Spectrometry 87

Figure 40. Pb isotopic compositions of sulfide inclusions in diamonds. The inclusions are classified according to their Ni content into eclogitic and peridotitic suites. Many of the peridotitic inclusions fall on the terrestrial growth curve indicating a formation age of 2.0 Ga with average terrestrial µ. Some of the eclogitic inclusions lie off the curve towards the origin indicating growth in a high µ environment such as the crust. This is supported by S-isotopic compositions for which only the eclogitic inclusions show significant variability. Figure adapted from Eldridge et al. (1991). et al. (1991) found that the Hf model ages of Mt. Narryer zircons were consistent with their 4.2 Ga U-Pb age, and concordant Hf and U-Pb ages were also obtained from a Mt. Narryer anorthosite (3.73 Ga) and White Cloud gneiss (2.72 Ga), but for the younger Pacoima Canyon pegmatite (1.17 Ga) and Jwaneng kimberlite zircons (0.24 Ga) eHf was significantly elevated suggesting incorporation of old Hf into the source regions (Figure 41). On the other hand, a Sri Lanka zircon was highly depleted (eHf of -23 at 570 Ma) indicating that the concordant U-Pb ages of these zircons have been reset or that they were derived from a very unradiogenic source. These zircons probably grew during high-grade metamorphism of a late Archean protolith. Kinny et al. (1991) also analyzed zircons from the Watersmeet tonalitic gneiss for which conventional analysis had suggested a correlation between U-Pb discordance and Hf model age. However, the ion probe data showed no such correlation and all cores had the same Hf isotopic composition despite the almost complete loss of Pb from some grains. However, 2.7 Ga rims do have younger Hf model ages suggesting the addition of younger Hf during rim formation.

Trace-element Geochemistry There have been a wide variety of applications in terrestrial trace-element geochemistry. As well as the more typical applications involving the measurement 88 TREVOR R. IRELAND

Figure 41. 176Hf evolution diagram showing results of ion microprobe zircon analyses (open symbols) with conventional thermal ionization data (solid symbols). MAINZ refers to the mean 176Hf/177Hf value of the Sri Lanka zircon analyzed by thermal ionization data. Most of the zircons are consistent with the trend of the conventional data but the Sri Lanka zircon probably grew metamorphically from a late Archean protolith. (Figure adapted from Kinny et al., 1991). of trace element abundances in and across individual mineral grains, there have been some highly innovative techniques ranging from cracking fluid inclusions and analyzing cation ratios (Diamond et al., 1990) and even CO2 concentrations (Pan et al., 1991), to analyzing REE concentrations in fish teeth and conodonts (Grandjean and Albarède, 1989). In the following section some selected applications from the recent literature are described which cover some of the more intensive research areas now being examined.

Diamond Inclusions. Besides sulfide inclusions for which S and Pb isotopic data have been obtained, a number of silicate phases are present in diamonds. Of major interest is whether these silicates are related to the diamond formation region or whether they represent more diverse provenances as is the case for the sulfides. Shimizu and Richardson (1987) measured REE, Ti, Zr, and Sc concentrations in individual crystals of sub-calcic peridotite suite garnets. All inclusions display a general enrichment of light REE relative to heavy, and the peridotite-suite garnets are depleted in Ti compared to those from garnet lherzolites. These characteristics are contrary to those predicted for garnet in association with olivine, orthopyroxene Ion Microprobe Mass Spectrometry 89 and diamond, and in particular the partitioning between garnet and liquid should produce a light REE depleted pattern. Such patterns have been measured from eclogitic garnet inclusions from Monastery diamonds which have a majorite component in solid solution (Moore et al., 1991). Navon et al. (1988) analyzed microinclusions in diamonds from Zaire and Botswana and found that they differed in composition from the larger eclogitic and peridotitic inclusions. They measured the major rock-forming elements as well as trace elements by ion probe with each analysis representing the average of a large number of inclusions since the inclusions were smaller than the probe diameter. The elemental concentrations were highly variable, probably due to the number density of the inclusions, and the submicron grains were found to resemble potassic magmas in their compositions. A different aspect of diamond genesis was examined by Phinney (1988). Large 3He/4He ratios have been measured in diamonds and lie in the range observed only in meteorites and this suggests that terrestrial diamonds might have a primordial component. However, 3He can also be produced by (n,a) reactions on 6Li but estimates of the importance of this reaction were restricted because of the absence of Li concentration data in diamonds. Phinney (1988) measured Li abundances in five terrestrial diamonds by implanting 7Li at a known dose and measuring the abundance by depth profiling. The highest concentration was 2.8 ppb with all the others having concentrations under 1 ppb; the analytical uncertainty of the measurements was around 25 %. In order to explain the 3He abundance, a Li concentration of around 30 ppm is required and so it appears that 6Li(n,a)3He does not make a significant contribution to the 3He abundance.

Mantle xenoliths Ultramafic xenoliths in basaltic magmas are the best samples available of Earth’s mantle. They are composed predominantly of olivine, orthopyroxene, and clinopyroxene, (dunites, harzburgites, pyroxenites, etc) with spinel (spinel lherzolites) or garnet (garnet lherzolites). The amount of available sample is small but is generally sufficient for conventional analyses. However, painstaking mineral separation is required and minerals in these xenoliths can be heterogeneous on a grain by grain scale and could also be internally zoned or the minerals can be locally altered. Shimizu and Allègre (1978) first analyzed garnet lherzolite nodules from kimberlites for major and trace elements including Sc, Ti, V, Cr, Mn, Co, Sr, and Zr. They found that the garnet lherzolites could be classified into three groups, one of which could be close in composition to primitive mantle. Salters and Shimizu (1988) found that clinopyroxene from some harzburgites and peridotites are depleted in the high field strength elements Ti, Zr, Hf, Nb, and Ta relative to the REE. They argued that the clinopyroxenes from spinel lherzolites were an adequate representation of the bulk inclusion because of the similarity of the patterns in clinopyroxenes and the bulk inclusion from which they were derived. These depletions were present in xenoliths from continental regions as well as oceanic 90 TREVOR R. IRELAND

Figure 42. Chondrite-normalized REE, Ti, and Zr concentrations for orthopyroxene (opx) and clinopyroxene (cpx) from mantle xenolith ER-N1/4. The cpx and opx show complementary anomalies in Ti and Zr while the calculated bulk rock pattern (bulk) shows a smooth variation. (Figure adapted from Rampone et al., 1991). regions and therefore appeared to be of world-wide occurrence in Earth’s upper mantle. However, Rampone et al. (1991) also used an ion microprobe to show that orthopyroxene had complementary Ti and Zr anomalies to the clinopyroxene and the recalculated bulk rocks show only negligible anomalies (Figure 42). Therefore, the complete inventory of any inclusion should be examined before the existence of an anomaly in the bulk rock can be established.

Glass inclusions. One of the principal aims of petrology is to use the measured compositions of coexisting minerals to reconstruct the parent magma from which they were derived. Potentially an equally useful approach may be to analyze the small glass inclusions that are commonly present in magmatic rocks. These inclusions are typically small (<100 µm) and so ion microprobe analysis is probably the most effective means of analysis. Glass inclusions can be found in a wide variety of compositions from felsic to mafic. Clinopyroxenes from abyssal peridotites have been found to have strongly light REE depleted patterns ([Ce/Yb]n = 0.002 - 0.05), depletions of Ti (300 - 1600 ppm) and Zr (0.1 - 10 ppm) and strongly fractionated Ti/Zr (250 - 4000) (Johnson et al., Ion Microprobe Mass Spectrometry 91

Figure 43. Ion microprobe analyses of clinopyroxenes in abyssal peridotites show evidence of near fractional melting of oceanic mantle and therefore require the presence of ultra-depleted melts. Such a composition has been discovered in a melt inclusion in olivine from a typical N-MORB and would be in equilibrium with a cpx even more depleted than the range previously found (shaded area). (Figure adapted from Sobolev and Shimizu, 1991).

1990). These compositions demonstrate that the peridotites are the result of variable degrees of fractional melting in the garnet and spinel peridotite fields, and that basalts might evolve from aggregation of these fractional melts. Support for this interpretation has recently been found in the form of melt inclusions in olivine from typical NMORB (Sobolev and Shimizu, 1991) (Figure 43). Surprisingly, the major element chemistry of the melt is consistent with only a low percentage of melting suggesting a very efficient mechanism for the fractionation of the incompatible elements.

Trace elements in zircon. The discovery of 4.2 Ga zircons in Western Australia has opened up a new era in Earth’s history. However, the zircons are preserved in a much younger quartzite and appear to be the sole representatives of the earliest crust; no rocks per se have been discovered that are the source of the 4.2 Ga zircons. Therefore all information concerning the source rocks and formation conditions must be obtained from the zircons. In this regard the trace element chemistry of the zircons may be the most important indicator of the nature of the source rocks. Hinton and Upton (1991) have analyzed large zircons from syenite and alkali basalt xenoliths and describe the basic features of zircon chemistry: they are 92 TREVOR R. IRELAND systematically enriched in the heavy REE relative to the light REE, with a distinctive positive Ce anomaly which reflects the more favorable incorporation of Ce4+ (relative to Ce3+) into the zircon structure. This feature appears to be ubiquitous in terrestrial zircons but Ireland and Wlotzka (1992) found that neither meteoritic zircons nor a lunar zircon assemblage had Ce anomalies reflecting the reduced nature of their parent bodies. However, Hinton and Meyer (1991) described an unusually oxidized lunar granite which does show the Ce anomaly. Snyder et al. (1993) also measured two zircons from lunar glasses and found, like Ireland and Wlotzka (1992) that there was no Ce anomaly. It would appear that the Moon is significantly more reduced the the Earth but that under certain conditions the ƒO2 may rise to a level where Ce4+ is stabilized, particularly in late-stage residual liquids such as those responsible for the assemblage analyzed by Hinton and Meyer (1991). In this regard the Ce4+/Ce3+ in zircon could become a useful oxygen barometer.

Figure 44. REE abundance patterns of detrital zircons from Mt Narryer, western Australia. The pre-4.0 Ga grains (a) show similar patterns to the 3.3 - 3.75 Ga grains suggesting a similar petrogenetic environment for all these grains. (Figure adapted from Maas et al., 1992). Ion Microprobe Mass Spectrometry 93

Maas et al. (1992) studied a variety of Archean zircons as well as the 4.2 Ga zircons from Mt. Narryer and Jack Hills. The old zircons could not be distinguished from the younger on the basis of the REE patterns (Figure 44). All zircons are characterized by the HREE enrichment, positive Ce anomalies and smaller negative Eu anomalies. Since most of the younger zircons are thought to be derived from a mature continental source dominated by K-rich granitic rocks, it is likely that the old zircons have a similar origin. To a large extent, it appears that the trace element chemistry of zircon is not a sensitive indicator of source rock chemistry. The trace element characteristics are reflecting the conditions under which zircon crystallizes which might be similar in different rock types despite initial differences in chemistry. This is because zircon is forming at a very late stage in the crystallization sequence and the local conditions may be of more importance than the bulk chemistry. However, as noted by Maas et al. (1992), there is only a limited geochemical data base with which to compare zircons from different petrogenetic environments and further work might elucidate some of the observed variations in the 4.2 Ga zircon suite.

Partition Coefficients. The determination of partition coefficients of trace elements is an obvious area for ion microprobe analysis especially where grains are zoned or trace impurities exist which are significant repositories of the trace elements. The ion microprobe analyses also allow a check for internal equilibrium of an assemblage, Joliff et al. (1989) found that the apatite crystals from a pegmatite were certainly not in equilibrium and were probably symptomatic of rapid crystal growth or local heterogeneities of the melt. Sisson (1991) and Sisson and Bacon (1992) used ion microprobe analysis to avoid contamination of pyroxenes and garnets by trace element rich accessory phases in high-silica rhyolites.

Diffusion and Dissolution Studies. The spatial resolution of the ion microprobe spot is of the order of 10 µm. However, a depth profile has a resolution of the order of hundreds of angstroms because the primary beam is sequentially sputtering layers away. This has great application in surface studies since elemental profiles can be measured as a function depth yielding information on dissolution or diffusion. As an example, Muir et al. (1989; 1990) analyzed surfaces of feldspars after dissolution and found that 600-1200 Å thick layers were formed that were devoid of Na, Ca, and Al. The thickness of the layers is strongly dependent on the pH of the solution and the composition of the plagioclase.

C. in Geochronology

The dating of rocks and minerals relies on the precise measurement of radionu- clides and their decay products. Possible geochronometers are more limited on the ion microprobe than conventional methods since the ion microprobe relies on the 94 TREVOR R. IRELAND mass differences between nuclides to resolve isobars whereas these can be chemically separated for conventional analysis. This effectively rules out b-decay schemes, such as 87Rb-87Sr, because of the extremely small mass difference between these nuclides. Furthermore, another common dating scheme, 40K-40Ar, as well as suffering from the problems of mass resolution, is also not suitable because of the poor ionization efficiency of Ar. For b decay schemes, measurements can generally only be made of samples which have a high daughter to parent ratio and are therefore limited to radiogenic tracer applications such as measuring Sr in carbonates and Hf in zircons. The data can be used to estimate ages, but these are model ages dependent on initial composition estimates and elemental fractionations. The ion microprobe has been used to advantage in one b decay scheme viz. 187Re-187Os. The difficulty in this measurement is that Re and Os are difficult to ionize into positive ion species in conventional thermal ionization mass spectrometry. They can be ionized by sputtering quite efficiently and so the first measurements of this system were made by an ion microprobe (Luck et al., 1980). The samples were first chemically separated, spiked, and loaded onto Al discs for analysis. However, this method has been superseded by conventional mass spectrometry after the discovery that both Os and Re yield quite intense negatively ionized oxide beams by thermal ionization (Creaser et al., 1991). Since heavy isobars cannot be resolved, ion microprobe dating is reliant on a decay schemes for which the daughter differs from the parent by four mass units or more. The best example of this for ion microprobe analysis is the U-Th-Pb scheme which includes four decay chains, each having multiple a decays, for example 238U decays to 206Pb with 8a emissions. An example of a single a decay scheme is 147Sm- 143Nd. The use of these decay schemes requires a precise determination of the isotopic composition of the daughter, as well as a determination of the parent daughter ratio. Of these schemes, U-Pb is the most amenable to ion probe analysis because of the large variation in parent-daughter ratios and also the use of the 207Pb/206Pb ratio for age determination. Over geologic time 206Pb/238U and 207Pb/206Pb range from zero to 1.0 and 0.05 to 0.6 respectively. On the other hand, extreme parent daughter fractionations of Sm-Nd are not found and very high precision is required for useful geologic age resolution.

U-Th-Pb in zircon The U-Th-Pb decay scheme is ideally suited to geochronologic applications since quite extreme fractionations in the parent to daughter elemental ratios are found. Of particular interest has been the mineral zircon, which contains U and Th at significant concentrations but strongly excludes Pb, and is geologically resistant to alteration in thermal processes. This indestructibility has hindered some conventional applications because a zircon crystal can become a complex mix of different zones that formed at different times. The ion microprobe is ideally suited to analyzing these different zones and has been paramount in the deconvolution of Ion Microprobe Mass Spectrometry 95 complex igneous and metamorphic terranes (Williams, 1992). Furthermore zircons can be rapidly dated to determine provenance age distributions in sediments, and recent advances have allowed the dating of Phanerozoic rocks for time scale calibration. Early attempts at zircon dating relied on stripping interferences by calculating the contributions from other isobars. This technique can be unreliable if the interferences are not well characterized and some of the early SHRIMP work was concerned with rectifying some of the earlier attempts (Williams et al., 1983). SHRIMP is still the only ion microprobe for which both U-Pb and Pb-Pb ratios can be accurately obtained although the commercial production of more large mass analyzer probes should alleviate that situation. It would not be realistic to review all papers which have been published concerning ion probe dating of zircon and so only a selection of topics has been chosen as representative of the applications undertaken. First however, it is perhaps important to verify the abilities of SHRIMP with some comparisons with conventional isotope dilution zircon dating.

Evaluation of the SHRIMP Zircon-Dating Technique. Over the past 10 years or so, SHRIMP has been used to date zircons that range in age from the beginning of the solar system, to some of the youngest tectonically active regions of the Earth. For the simplest zircons, a comparison can be made between conventional dating techniques (using chemically separated aliquots, and zircons dated by SHRIMP. It should be noted that even for morphologically simple zircons, the U- Th-Pb systematics are not necessarily well-behaved since heterogeneous Pb loss, the presence of thin U-rich rims, and invisible cores can all play havoc with the anticipated results on a micron scale. In this section, six comparisons are made between conventional dating methods and ion probe zircon dates in samples ranging in age from 4560 Ma to 2 Ma. The oldest zircons, for which a conventional tie point is available, are not terrestrial but are found in differentiated meteorites. The howardite-eucrite- diogenite association consists of basalts and gabbros, often brecciated, that apparently formed on a relatively large asteroidal body referred to as the HED Parent Body. Small zircons (≤30 µm) have been found in several meteorites of this group and some of these have been analyzed by ion microprobe for their U-Th-Pb systematics. Ireland and Wlotzka (1992) presented data for zircons from the Vaca Muerta mesosiderite and for one zircon with approximately 50 ppm U a concordant 207Pb/206Pb age of 4563 ± 15 Ma was obtained (Figure 45) in good agreement with the canonical age of the solar system. Zircons from the Isua supracrustal belt have been analyzed extensively both by ion microprobe and conventionally. Compston et al. (1986) obtained a mean age of 3807 ± 2 Ma for a suite of Isua zircons, in good agreement with a conventional determination of 3813 ± 9 Ma by Baadsgaard et al. (1984). The conventional data show a degree of very early lead loss whereas most of the ion probe data are consistent with only a small degree of recent lead loss with the ancient Pb loss 96 TREVOR R. IRELAND

Figure 45. Concordia plot for Vaca Muerta zircons. VM-1 has only 0.6 ppm U compared to 50 ppm in VM-2 and hence the very large error on the VM-1 analysis. The mean of the two VM-2 analyses is concordant with a 207Pb/206Pb age of 4563 ± 15 Ma (2s) which agrees with the canonical age of the solar system. (Figure adapted from Ireland and Wlotzka, 1992). restricted to a few individual grains. This is exemplified in Figure 46a, which shows isotope dilution analyses of single crystals and ion probe analyses of grains from the same population. Zircons from a syenite from East Greenland are concordant with a 207Pb/206Pb age of 2701 ± 5 Ma as indicated by single crystal analyses of abraded grains. Ion probe analyses of these same abraded grains (analyzed prior to isotope dilution analyses) give the same age within error at 2698 ± 7 Ma. There is some dispersion in the 207Pb/206Pb ages tending to lower values suggesting some domains in the zircons have experienced a small degree of early lead loss (Figure 46b). For zircons younger than around 800 Ma, the 207Pb/206Pb ratio becomes a rather insensitive indicator of age as measured on SHRIMP I. For such zircons the 206Pb/238U ratio can be used and the U-Pb age is simply determined as an age relative to the SL13 standard zircon. Zircons dated by the 206Pb/238U ratio at 802 ± 10 Ma by MSID give the same age within error by ion microprobe (801 ± 5 Ma) (Figure 47a). Paleozoic zircons dated with the 206Pb/238U ratio at 330 ± 4 Ma by MSID also give an identical result by ion microprobe Figure 47b. Both MSID and ion probe results are concordant within rather large errors in the 207Pb/206Pb ratio because of the small sample size analyzed. Ion Microprobe Mass Spectrometry 97

Figure 46. Comparison of ion microprobe (open boxes) and isotope dilution ages (grey) for single zircons from (a) 3800 Ma Isukasia volcanics, and (b) a 2700 Ma syenite from East Greenland. The isotope dilution results from Isukasia indicate the sample has experienced ancient lead loss and the oldest ages are similar to the ages obtained by ion microprobe analysis. The ages shown for the syenite are derived from the weighted mean 207Pb/206Pb ratio. (Unpublished data from A.P. Nutman and C.M. Fanning.) 98 TREVOR R. IRELAND

Figure 47. Comparison of ion microprobe (open boxes) and isotope dilution ages (grey) for single zircons from (a) 800 Ma, and (b) 330 Ma source rocks. One abraded multi-grain sample is shown in black in (a). The isotope dilution results are from small- number multi grain samples from the U.S.G.S., Denver. (Unpublished data from (a) J. Claoué-Long and C. M. Fanning and (b) W. Compston, C. M. Fanning, and K. R. Ludwig.) Ion Microprobe Mass Spectrometry 99

Figure 48. Concordia plot of zircons from the Omara granodiorite showing U-Pb data uncorrected for common lead. The age of the granodiorite is interpreted as the intercept of the common Pb trajectory with the concordia which is 2.1 ± 0.1 Ma, and is in good agreement with the 40Ar/39Ar age of this rock. (Unpublished data from S. Baldwin.)

For very young zircons, the common lead contribution cannot be reliably constrained by the 204Pb/206Pb ratio. Even the 208Pb method of correction results in large uncertainties on the radiogenic 207Pb/206Pb ratio because of the low concentration of 207Pb and the low number of counts at this isotope. In this case, the uncorrected data can be used to extrapolate to the concordia to obtain an age estimate. This can be done by either assuming a fixed common Pb composition, or if there is sufficient dispersion in the data, the common Pb composition can be determined from the data. This approach is demonstrated with ≈ 2 Ma zircons from the Omara granodiorite from the D’Entrecasteaux Islands, a tectonically active area off Papua - New Guinea. The zircons have quite high U concentrations (500 - 1000 ppm) which yield data quite close to concordia (Figure 48). The age is estimated to be 2.1 ± 0.1 Ma. These data clearly demonstrate that the ion microprobe is capable of reliably dating extremely small amounts of Pb in zircon. This allows the complexities of natural zircons to be unraveled into distinct geological events for which good age control can be obtained. Well over a thousand samples have been dated on SHRIMP covering the full range of geologic time. In the following sections, only a few of the problems tackled by SHRIMP are outlined as representative of the power of the ion probe zircon-dating technique in solving geological problems. 100 TREVOR R. IRELAND

The Oldest Terrestrial Zircons. One of the most conspicuous applications of the SHRIMP zircon-dating program has been the search for the oldest zircons on Earth. Prior to the SHRIMP analyses, the oldest-known terrestrial material was obtained from the Isua supracrustal belt for which some units were found to be a little over 3.8 Ga. This still leaves some 750 Ma between the formation time of the meteorites, believed to be close to the formation time of the earth, and the preservation of felsic crust on Earth. Within the first few years of routine zircon analysis, zircons that significantly predated Isua were found in Archean quartzites from Western Australia (Froude et al., 1983; Compston and Pidgeon, 1986). These findings have been paramount in developing models for the formation of Earth’s earliest crust and have pushed back the oldest recognizable vestiges of Earth’s history to 4.28 Ga. The zircons were found as a small proportion (<2 % in Mt Narryer) of the total and so a large number of zircons had to be analyzed to ascertain firstly their presence and then their abundance. Schärer and Allègre (1985) attempted to verify the findings of Froude et al. (1983) with conventional single-grain analysis. After analyzing 32 grains, Schärer and Allègre had found no “old” zircons. As argued by Compston et al. (1985) this simply reflects the statistics of small numbers. At an abundance level of 0.02, there is roughly a 50 % chance of finding an old grain in 32 attempts. Schärer and Allègre may have inadvertently biased their population by choosing grains according to morphology. However, since there is no apparent correlation between morphology and age characteristics it is more likely that they were simply unlucky. Schärer and Allègre also claimed that the discordance pattern of the younger grains as deduced by ion microprobe analysis did not match their conventional data set. While the ion probe data set was consistent with three main populations at 3300, 3500, and 3750 Ma, the conventional data set also shows intermediary ages which are likely to be due to mixed ages between cores and rims of the more structured zircons. This is illustrated in Figure 49 which shows a number density plot of 275 ion probe analyses with the conventional data points superimposed (Kinny et al., 1990). The exchange between the ANU and Paris groups is an indication of the healthy skepticism to which the report of > 4 Ga zircons was first met. It should be noted that this was the first paper on zircon analysis to be published from the ANU ion probe group and so the technique had not had the chance to substantiate itself on less complicated and less controversial problems initially. Conventional analyses of > 4 Ga zircons have now been made and confirm the antiquity of these grains. Kober et al. (1989) analyzed 41 zircons from Jack Hills by stepwise evaporation and found that five of the grains had minimum crystallization ages between 4.07 and 4.17 Ga. Single-grain isotope dilution analyses also confirm the old lead compositions in these detrital grains but while concordant domains could be analyzed with the ion microprobe, conventional analyses have been often discordant (C. M. Fanning, personal commun.). Ion Microprobe Mass Spectrometry 101

Figure 49. Concordia plot of 275 detrital zircon grains from the Mt Narryer quartzite. The ion microprobe analyses are contoured according to number density. Conventional single grain and grain-fragment analyses are superimposed as crosses. In general there is an excellent agreement between the two techniques for the discordia trend of the majority of the grains. (Figure adapted from Kinny et al., 1990).

Intact remnants of the source of the 4.1-4.3 Ga zircons from western Australia have not been found to date with the oldest gneisses having formation ages no greater than ≈3730 Ma (Nutman et al., 1991). Pre 3.8 Ga detrital zircons have also been found in Archean quartzites from Greenland, the oldest being 3850 Ma (Nutman and Collerson, 1991), from China, 3850 Ma (Liu et al., 1992), and from the Beartooth Mountains, Montana where ages range up to 3960 Ma (Mueller et al., 1992). In contrast to western Australia, gneisses with similar formation ages have been identified in these cratonic areas with the 3.96 Ga Acasta gneiss being the oldest rock identified so far (Bowring et al., 1989).

Sedimentary Provenances. Ion microprobe dating of single zircons has an obvious application to determining sedimentary provenances. This technique has been used in conventional analysis as well, but the ion microprobe can produce a usable age (≤25 Ma error) in 15 minutes with no chemical separation required, which is far more time effective than undertaking conventional analysis. There are two types of data that result from the analysis of zircons in sedimentary rocks: first is the age of the individual provenances from which the zircons are derived; second is the relative abundances of the zircons from those provenances. The ages of the 102 TREVOR R. IRELAND

Figure 50. Zircon age histogram from a paragneiss from the West Coast of New Zealand. This pattern is also found in Ordovician sediments in the same region and in Ordovician sediments from eastern Australia and Western Antarctica indicating original contiguity of the terranes. (Data from Ireland, 1992). zircons allow the evaluation of possible sources for the zircons and hence the sediments; the ages are simply matched with possible source terranes. The second piece of information is more conjectural in its use. The relative proportions of the zircons may not in fact reflect the proportions of the detritus from different terranes for a number of reasons. Firstly, different rock types have different zircon concentrations and so a granite will bias the sample relative to an intermediate rock and mafic rocks will not even be represented. Secondly there can be a sampling bias during zircon separation (size, magnetic susceptibility), and then in the choice of grains for analysis. A large number of randomly selected grains from unfractionated separates reduces some of these problems but bias cannot be excluded. Perhaps of equal importance to the interpretation of the abundance peaks is the use of the zircon signature as a fingerprint of the rock. This might allow correlations of sedimentary units over large distances. For example, it is evident in south eastern Australia, western New Zealand and western Antarctica that Cambro- Ordovician sediments have identical detrital zircon patterns (Figure 50) which is strong evidence for a common provenance and original contiguity (Williams et al., 1990; Ireland, 1992). Surprisingly, the ages of the zircons in the sediments cannot be Ion Microprobe Mass Spectrometry 103

Figure 51. Concordia plot of zircon analyses from an orthogneiss at Mt Sones, Antarctica (Black et al., 1986). Four zircon ages are apparent from this one rock. Cores represent the original igneous protolith at 3.93 Ga, with discordant analyses in strongly zoned crystals (filled symbols) projecting towards a granulite event at 2.95 Ga represented by overgrowths (inset). Structureless grains grew at around 2.46 Ga and minor Pb loss occurred during an event at around 1.0 Ga. readily attributed to any source rocks in New Zealand or eastern Australia. However, the fingerprinting of other rocks of this age from around the Pacific margin of Antarctica may be useful in firstly correlating the rock units and secondly determining the provenance directions for the distinctive age components. On a slightly different note, Brimhall et al. (1992) used zircon morphology and age characteristics to show that soils develop by dissolution and collapse of the bedrock and by invasive transport of aerosol detritus from above. Euhedral zircons throughout the profile have the age of the unweathered granite beneath, but rounded zircons are found to have a wide provenance age. These foreign zircons are transported as aerosols and penetrate the soil horizon to a maximum depth of 2 m by movement down channels left by decayed roots and by bioturbation.

Complex Metamorphic Terranes. The ion microprobe is a particularly useful device for analyzing complexly zoned zircons from metamorphic terranes. Zircon is a highly persistent mineral once crystallized and so successive thermal episodes can deposit zircon resulting in several generations within the same crystal. This was 104 TREVOR R. IRELAND highlighted by Black et al. (1986) who found that four distinct generations of zircon growth had occurred in an orthogneiss from Mt Sones, Antarctica (Figure 51). These zones could be independently sampled and correlated with conventional isotope data from rocks within this terrane. Zircon cores have a mean age of 3927 ±10 Ma and this is believed to be the age of the protolith of the orthogneiss. After +31 granulite facies metamorphism dated at 2948-17Ma, extensive zircon recrystallization occurred but only a small amount of new zircon growth. This is approximately the age of extensive isotopic resetting in the Rb-Sr system corresponding to the D2 event marking the end of granulite facies metamorphism but is significantly younger than the D1 event marking the onset of the metamorphism at 3070 ± 34 Ma. Structureless pink zircons crystallized at 2479 ± +8 23 Ma which is within error of the D3 event determined by Rb-Sr at 2456-5 Ma (Black et al., 1983). Finally tectonism at around 1000 Ma caused a small degree of zircon Pb loss.

Calibration of the Geologic Time Scale. The high selectivity available in ion probe analysis has led to its application to geological time scale work. This is perhaps the most stringent type of analysis possible because of the poor age resolution afforded by the 207Pb/206Pb ratio and therefore the need for accurate and precise calibration of Pb/U ratios. The external reproducibility of repeat analyses of the SL13 standard appears to limit the precision of individual analyses to around 2 %. Therefore a large number of analyses are required of both standard and unknown in order to precisely calibrate the age. The benefits of ion probe analysis are the ability to detect and reject statistical outliers that may be due to inheritance or Pb loss; the inclusion of such domains in a conventional analysis could bias the result to higher or lower ages respectively. Zircons are typically separated from tuffaceous units with well-controlled stratigraphic relationships. Analyses of zircons from early Cambrian tuffs from Morocco, China, and Australia give identical results at 520-525 Ma indicating the Cambrian - Precambrian boundary is substantially younger than previously thought (Compston et al., 1992; Cooper et al., 1992). Compston and Williams (1992) measured zircons separated from British Ordovician stratotypes and found that the ages were clearly younger than the canonical time scale of Harland (1989). The ion probe ages generally agree well with conventional analyses of the same rocks (Tucker et al., 1990) although the former can be younger by up to ca. 10 Ma. Compston and Williams (1992) attributed this difference to a small degree of inheritance in the multigrain samples analyzed by Tucker et al. (1990) since ion probe ages agreed closely with high-precision K-Ar and Rb-Sr for other volcanic rocks. At the other end of the Paleozoic, Claoué-Long et al. (1991) have dated the Permian - Triassic boundary at 250 ± 3 Ma.

Other chronometers Probably no other system can be exploited in geochronology to the same degree on the ion microprobe as that of U-Th-Pb. The large variation in the parent daughter Ion Microprobe Mass Spectrometry 105 ratios as well as the two coupled decay systems are unique. There are several systems that may be of potential interest in terms of model ages. Hf isotopic compositions in zircons has been shown to be useful in identifying the crustal extraction age of the zircons thereby seeing through metamorphic events that may have reset the U-Pb system. Sr isotopes too might be used in a similar fashion as a radiogenic isotope tracer. With both these systems however, the parents and daughters are isobaric and require high mass resolution for them to be separated. Hence there is no foreseeable direct chronological application of these systems. In the Re-Os system, this has been sidestepped by chemically separating these elements prior to their analysis on the ion probe. However, there is again no real possibility of achieving in situ analysis as yet. Of the commonly used a-decay systems, Sm-Nd dating may prove to be possible but there will have to be substantial advances in ionization and transmission efficiency at high mass resolution for it to be worthwhile.

VI.!PROSPECTS AND CONCLUSIONS

The ion microprobe has indeed come of age. Its capabilities are now recognized although it is often viewed as simply a zircon-analysis device or a highly sensitive trace-element detector. Despite the quite dedicated application of individual ion probes, they are all capable of the same retinue of techniques. Basically what works on one ion probe will work on all. The limitation to this concept however is the transmission efficiency and hence sensitivity of the individual machines. For this reason, it is likely that the new “consumers” of these devices will require large high sensitivity machines capable of U-Pb analysis which can also be used for the lower sensitivity applications as well. There are ultimate limits to the sensitivity that can be achieved. The number of atoms in a 10 µm spot sputtered to a reasonable depth is finite. Of the atoms present only a certain percentage are of interest for the analysis and of these only a small fraction are ionized and will make it to the detector. The large ion probes are designed to collect and transfer as large a fraction of the ions from the sample into the mass spectrometer and only modest gains (factor of 3-10) will be forthcoming with better designs. If higher sensitivity is to be achieved then more ions must be produced per unit volume or the ions that are produced must be collected more efficiently. There has been a great deal of effort expended already in the application of lasers in the selective excitation of various species. It is still not clear whether this will result in a greater sensitivity per unit time than that already achieved in SHRIMP- style ion microprobes. The ionized fraction of sputtered material is generally taken to be around 0.1 %. Therefore there is, at most, a factor of 103 to be achieved if resonance ionization was completely efficient. The problem is that sufficient power is only available in pulsed lasers with only nanosecond duration. In order to achieve comparable ion throughput, the laser would need a repetition rate of the order of 1 106 TREVOR R. IRELAND

MHz (as opposed to the currently available repetition rates of 1kHz at a maximum). There is also the problem of measuring ion abundances in such ionic parcels and a charge collection mode would have to be employed rather than the ion counting schemes currently utilized. Even then, the isotopic abundances measured following resonance ionization can be highly variable and dependent on the exact experimental setup (Spiegel et al., 1991). Until high-powered continuous lasers are available and the fundamental mechanisms of ion production under laser bombardment can be understood, resonance ionization ion microprobes will be severely limited in their application. One of the main losses of ions in analysis is due to the collection cycle efficiency, that is, the time that is spent counting on one peak means that the other peak(s) is(are) not being counted. For equal counting time on two peaks the collection efficiency is 50 %. For simple systems, the obvious method to improve collection efficiency is to go to multiple collection, as has been done in conventional mass spectrometers. There are clearly major benefits to be had and it is a direction that commercial manufacturers are addressing. The main difficulties concern the low count rates on some peaks requiring ion counters to be used, and the physically small space into which these counters must fit. Both the CAMECA 1270 and SHRIMP II ion microprobes are to be fitted with multiple collectors based on ion counters but these have not been tested as yet. The limits on sensitivity also place limits on the ultimate spot size that can be effectively used. In rastered ion imaging there is a clear advantage in a small spot for looking at distributions of elements in the target. However, in high precision isotopic analysis, there is almost no prospect of using a 1 µm spot to achieve sub permil precision in any elemental system. There are simply not enough atoms and hence ions produced. For isotopic analysis the minimum number of ions required can be calculated according to Poisson counting statistics and from this a minimum volume to be consumed can be calculated. A reasonable minimum probe diameter for isotopic analysis will probably be in the 5-10 µm range. There are clearly some great difficulties in the next stages of development but future generations of these machines may be very different to what we have today and capable of a range of applications that were unforeseen. The main difficulty in the present generation of ion microprobes is in obtaining access. Large ion microprobes are expensive and are likely to be available only through regional facilities. Even then there is only a finite amount of time available on a single machine and it is likely that the smaller user will be unable to compete for time. In this regard, the availability of the ion microprobe is likely to be quite different to the electron probe. The price of electron probes was soon within reach of earth science departments with only quite modest budgets. It is unlikely that this will ever be the case for high sensitivity ion microprobes at least in their present configurations. The ion microprobe has proved to be a highly versatile machine. Despite the lack of a thorough understanding of the ionization processes involved, the ion probe can be used for a variety of types of quantitative analysis both in terms of isotope ratios Ion Microprobe Mass Spectrometry 107 and elemental abundances. These analyses are generally relative measurements in that a well-known standard must be analyzed in order to quantify the effects of selective ionization in the sputtering process. There are still a large number of applications that have not been fully tamed as yet. The most obvious case is that of oxygen isotopic analysis of silicates where despite a great deal of effort, the fractionation of the oxygen isotopes cannot be controlled sufficiently for use in terms of stable isotope geochemistry. It is, however, only a matter of time before a recipe can be developed that will enable routine measurements of this type. So too with other applications. Individual recipes must be developed, firstly by evaluating isobaric interferences and ion intensities and then determining the reproducibility of the measurements. In this way it is likely that a large retinue of new applications will be forthcoming in the near future.

ACKNOWLEDGMENTS I wish to thank Bill Compston for introducing me to ion microprobe measurement on SHRIMP and Ernst Zinner, Albert Fahey, Kevin McKeegan, and Bob Walker for my experiences on the CAMECA ims-3f at Washington University. Thanks also to Mark Harrison for my introduction to the CAMECA 1270. I have enjoyed discussions with these practitioners as well as Ian Williams, Peter Kinny, Derek Froude, Stewart Eldridge, Allen Nutman, and Mark Fanning, all of whom are in some way responsible for aspects of this work. Bill McDonough, Ghislaine Crozaz, Mark Harrison, and Emmanuel de Chambost are thanked for their comments, and Larry Nittler and Stewart Eldridge for photomicrographs.

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