ION MICROPROBE MASS SPECTROMETRYpeople.rses.anu.edu.au/ireland_t/All_Publications... · Ion Microprobe Mass Spectrometry 3 ions that reflect in some way the isotopic and chemical

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ION MICROPROBE MASSSPECTROMETRY:TECHNIQUES AND APPLICATIONS INCOSMOCHEMISTRY, GEOCHEMISTRY, ANDGEOCHRONOLOGY

Trevor R. Ireland_______________________________________

I.!Introduction......................................................................................................22.!Sputtering ........................................................................................................3

A. Ion Production.....................................................................................3B. Ionization Modeling ............................................................................10C. Secondary Ionization Theory...............................................................12

3.!Hardware .........................................................................................................13A. General Features..................................................................................13B. CAMECA............................................................................................13C. SHRIMP.............................................................................................19D. VG Isolab 120.....................................................................................24

4. Methodology ...................................................................................................24A. Isotopic analysis..................................................................................24B. Quantitative Analysis...........................................................................46

__________________Advances in Analytical GeochemistryVolume 2, pages 1-118.Copyright © 1995 by JAI Press Inc.All rights of reproduction in any form reserved.ISBN: 1-55938-785-8

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V.!Applications ....................................................................................................54A. in Cosmochemistry.............................................................................54B. in Geochemistry ..................................................................................82C. in Geochronology................................................................................93

VI.!Prospects and Conclusions............................................................................105Acknowledgments................................................................................................107References ...........................................................................................................107

I.!INTRODUCTION

An ion microprobe is basically a mass spectrometer with a specialized sourceincorporating a finely focused primary ion beam to generate secondary ions fromthe target. The primary ions have energies of the order of 10 keV and so collisionsbetween the primary ions and the surface physically erode, or “sputter”, the sampleejecting particles from the surface. Of these particles, a small fraction are ionizedand can be electrostatically removed to the mass spectrometer where they areseparated according to mass and the relative ion intensities can be measured.

Ion microprobes have found a niche in materials research, particularly in thesemiconductor industry where localized analysis of ppm to sub ppb concentrationsare required. The commercially produced CAMECA ion microscopes have beenparticularly important in this area and over 200 instruments are in operationworldwide. However, only a small fraction of these instruments are used for researchin the earth science field, and then mainly for research in cosmochemistry whereisotopic and chemical abundance anomalies are at their largest. Of primeimportance in the geological sciences has been the SHRIMP (Sensitive High mass-Resolution Ion MicroProbe) at the Australian National University that was designedfor isotopic analysis of chemically complex targets. The main application of thisinstrument has been in situ U-Pb dating of zircon although a wide variety ofapplications have benefited from its high sensitivity. Until recently SHRIMP was theonly instrument capable of undertaking routine U-Pb dating of zircon but nowcommercially produced machines are being marketed with these capabilities.

The fundamentals for ion microprobe analysis were developed over thirty yearsago with the goal of characterizing any sample for its chemical and isotopicabundances. In this same time frame, the electron microscope and microprobe werealso being developed. The electron probe uses a focused electron beam to generateX-rays in the target and since the principles governing photon emission from asurface are well understood, quantitative analyses were readily forthcoming and theelectron probe quickly became a necessary piece of apparatus in materials researchlaboratories worldwide. Development of the ion microprobe was much slower,because the principals that describe secondary ionization are not well understoodand, despite a great deal of effort, no generally acceptable formulation has beenforthcoming. However, the salient point involved in applying the ion microprobe toproblems 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 successfulapplications.

The purpose of this paper is to highlight the successful application of ionmicroprobe mass spectrometry to a variety of topics principally in the earth andplanetary sciences. It is not a review of the development of the ion microscope andmicroprobe, previous reviews of ion microprobe mass spectrometry (Shimizu andHart, 1982a; Reed, 1984; Benninghoven et al., 1987) have documented most ofthese aspects. Nor is it a review of the status of ion production models. While acomplete model would be of great benefit to practitioners, it is not a requirement forsuccessful application of the instruments as will be detailed in later sections.Moreover, the physics is complex and a detailed discussion is beyond the scope ofthis paper and probably beyond the requirements of the interested but non-expert towhom this work is addressed. Detailed discussions of sputtering models are given byWilliams (1979; 1982), Harrison (1983) and Benninghoven et al. (1987).

In the following sections, a brief overview of the sputtering process will be givenfollowed by descriptions of the most recent generations of ion microprobes in usetoday. General aspects of analysis with the ion microprobe will then be discussedfollowed by specific examples from cosmochemistry, geochemistry, andgeochronology.

II.!SPUTTERINGA. Ion Production

A 16O– 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 layersand only for several atomic radii around (Figure 1). For a 1 nA primary beam thereare some 6 ¥ 109 impacts per second. An impact results in some 10 to 100 atomsand molecular fragments being ejected and in this way a sample is continuallyeroded. Felicitously, a small fraction of the emitted particles are ionized and can beelectrostatically removed from the sputtering region to be analyzed in a massspectrometer.

While the basic scheme involved can be appreciated in macroscopic terms, thephysical principles governing the ionization probability of individual particles arenot readily quantifiable because of the diverse parameters that are needed to fullydescribe the characteristics of the ejected particles. The pioneering statisticaltheories of Thompson (1968) and Sigmund (1969) involved the partitioning ofenergy in collision cascades and were able to predict bulk properties of the sputteredmaterial such as the sputtering yield, ejected atom energy, and secondary atomangular distributions as a function of primary ion energy. In order for a particle tobe ejected it must obtain sufficient kinetic energy to overcome the surface bindingenergy. All sputtered atoms originate from the outer few atomic layers with the

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Figure 1. Sputtering of a solid surface by particles with energies of the order of 10 keVresults in the disruption of the lattice and ejection of atoms and molecular fragments bythe transfer of energy back towards the surface. These glancing collisions result inrelatively low emission energies of the order of 10 eV although some particles mayacquire energies of over 1000 eV. Most particles are neutral but a small fraction areionized and this forms the basis of secondary ion mass spectrometry.

majority coming from the surface itself (Williams, 1983). Since the atom mustobtain momentum in the opposite direction to the incoming ion, most of the energyis transferred in glancing collisions and hence emitted ions have rather low energiespeaking at around a few eV although a small fraction have energies extending up tothe keV range (Metson et al., 1984). The highest energy fraction is atomic ions withcomplex molecular ions having increasingly narrower energy spreads in proportionto 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 havean energy distribution similar to the neutral particles. If all particles emitted from


the surface were ionized, a satisfactory model describing the relationship betweensample 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 inionization probabilities of a given species with changes in the sputteringenvironment. For example, the presence of oxygen, either as the primary ionspecies or simply leaked into the vicinity of the sputtering region, causes a largeincrease in the emission of positive ions. Similarly, sputtering with a Cs+ ion beamresults in enhanced emission of negative ions. However, some elements essentiallydo not ionize, such as the inert gases He, Ne, Ar, Kr, and Xe, and also N. While theion probe can operate in any of four configurations, incorporating positive ornegative primary or secondary ions, the optimum configuration is to have oppositepolarities of primary and secondary ions. The most common primary beam speciesare Cs+ and O– since these highly electropositive and electronegative species(respectively) have a chemical effect on the sputtering region enhancing ionemission many orders of magnitude over other source species such as Ar+.

Since ion microprobe analysis relies on the transport of charged particles, caremust be taken to avoid charge buildup, particularly at the sputtering site. The twomost common analytical configurations are negative primary - positive secondaryand positive primary - negative secondary. In the first mode, negative charge buildup 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 tothe combined addition of the positive ions from the primary beam and theextraction of both negative secondary ions and electrons. Even a relatively thickgold coat is not sufficient to allow electrons access to the sputtering site to neutralizethe positive charge. This presents one of the major problems in the analysis of lightstable 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 tothis problem, but neither has proved to be completely satisfactory. The first involvesreducing the sample to 10-15 µm fragments and embedding them in gold foil. Thisremoves some of the benefits of in situ analysis but still allows the analysis ofreasonably small intact samples. The second method involves neutralizing thesputtering site by focusing an electron beam into the analysis area. CAMECA ionmicroprobes now include an electron gun for this purpose and have been successfulin obtaining stable intense O– beams from oxides and silicates. However, thereproducibility 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 massfractionation to a level which will allow precise and accurate analyses of insulators.

The secondary ion mass spectrum from the ion microprobe is complicated, notonly because of the chemical diversity in geological samples, but because molecularfragments of all types can be produced. The molecules are best described asfragments because they can have no stoichiometry in terms of valence that generallygoverns the configuration of molecular compounds. Instead, any algebraiccombination of the components in the sample can be produced. Even for the simplest

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Figure 2. Negative secondary ion spectrum from a graphite substrate sputtered with anO+ primary beam (from McKeegan, 1987). Even from the simplest targets, thespectrum is complicated by the presence of a plethora of molecular species. Thespecies can not be regarded as having stoichiometric affinities but rather should beregarded as simple algebraic combinations of the target atoms.

samples, the spectra can be extremely complicated. For example, the secondary ionmass spectrum of graphite sputtered with an O– beam in Figure 2 showscombinations of C isotopes as well as molecules composed of C and O. For ageological sample the situation is even more chaotic.

The data collected from an ion microprobe generally consists of ratioing peakheights of different isotopes of an element for isotopic analysis, or isotopes ofdifferent elements for a chemical analysis. Before data can be obtained anyinterferences under the peaks of interest must either be removed or their abundancesquantified.

Since the ion probe is a mass spectrometer the primary method of peakseparation is according to mass. The first ion microprobes were constructed withlow resolution spectrometers so that basically only unit masses could be separated.With the observation of complex molecular interferences it was soon realized thatmuch higher mass resolutions would be essential for fully utilizing the capabilities ofthe ion probe. The regular variation of mass deficits of individual isotopes meansthat molecular interferences can be readily separated from atomic species for bothlow and high mass regions of the mass spectrum. For intermediate masses the massdeficit curve turns around and much higher mass resolutions are required.

The masses are dispersed along the focal plane of the magnet with theirseparations dependent on the characteristics of the magnet as well as the width of the


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 onthe peak is proportional to the source slit width. Mass resolution is commonly definedas M/DM at the 10 % peak height.

entrance and exit slits of the mass spectrometer. Since the characteristic dispersionof the magnetic analyzer is a fixed parameter, variation in mass resolution can onlybe 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 thehalf peak width being equal to the collector slit width, and the slope dependent onthe source slit width (Figure 3). The mass resolution is the mass of the peak dividedby the base width of that peak (M/DM). There is some variability in defining massresolution because of the different levels at which the base width is measured. Indetail it is important to designate the height used to define the base width since abase width at the 1% level will be wider than that at the 10 % level and therefore the10 % level will define a higher mass resolution than the 1 % level. A commonconvention 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 amass analyzer. Basically this can be viewed as the degree to which an adjacent massinterferes 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 aspossible and this is an important parameter in the measurement of low abundanceisotopes such as 10Be, 26Al, 234U, and 230Th. The abundance sensitivity can be definedfor any mass, but a common convention is to measure the contribution at mass 237from the 238U signal.

While conventional mass spectrometers have mass resolutions of the order of500!R, enabling separation of individual masses up to 500 amu, ion microprobes

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Figure 4. High mass resolution spectra of peaks in the region of the Ti isotopes fromFahey (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 than0.1 ‰ to the 48Ti+ signal when the latter is centered through the exit slit. However, thisis 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 isobaricinterferences by monitoring 51V+ and 52Cr+.

generally operate at much higher mass resolutions on the order of 3,000 to 10,000R. The regular variation of the mass deficits of the chemical elements with massenables the separation of all molecular isobaric interferences at mass resolution ofthe order of 10,000 R for elements of mass less than around 50 amu. Figure 4shows the mass spectrum in the region of the titanium isotopes in a meteoritichibonite at a mass resolution of ≈10,000 R. Molecular interferences are wellresolved, for example 46TiH+ and 30Si16O+ are well separated from 47Ti+. In generalhydrides require larger resolving powers than oxides but atomic isobaricinterferences generally require much higher mass resolution for separation. In thecase of the Ti isotopes only 48Ca is partially resolved from 48Ti, the other atomicisobaric interferences, 46Ca, 50V, and 50Cr are not resolved. For isotopic massesheavier than around 50 amu, molecular interferences become increasingly similar inmass to the atomic species and hence more difficult to resolve by mass separation.

An alternative method for discriminating against molecular ions is the energyfiltering technique. This method relies on the differences in energy distribution of


Figure 5. Energy spectrum showing Si+, SiO+, and SiO2+ sputtered from zircon as

measured on SHRIMP I. The molecular species have increasingly narrow energybandwidths as the complexity (number of constituent atoms increase). This forms thebasis of the energy filtering technique which can used to discriminate against complexmolecular interferences.

atomic and molecular ions. Atomic species have a broad energy spread whereasmolecular species have increasingly narrow energy bands depending on the numberof atoms in the molecule. Figure 5 shows an energy spectrum for Si+, SiO+ andSiO2+ 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 ordersof magnitude, the atomic species has dropped by 2 orders of magnitude sointerference discrimination is at the expense of intensity of the atomic species aswell. A redeeming benefit of using this method is that low mass resolution can beused which relaxes the requirements of the magnetic field controller particularlyover the large mass ranges required for chemical analyses.

Even if an isobaric interference cannot be removed by high mass resolution orenergy filtering, a correction can be applied to the peak of interest provided anotherisotope of the interfering element can be measured. For example, the intensities of50V+ 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 theinterfering isotope and is free of isobaric interferences at 10,000 R and so anaccurate correction can be made. Of course, when the monitor isotope is smallerthan the interference then there can be a substantial error propagated to thecorrected ratio.

Apart from interferences which are the result of coincident masses, interferencescan also result from different charge states. For ion microprobe analysis the only

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readily detected species are doubly charged atoms of elements with high ionizationefficiencies, for example Mg, Ca, and Sr. Doubly charged Ca ions can interfere withMg+, and Sr2+ can interfere with Ca+. In the latter case three Sr isotopes caninterfere with Ca, 84Sr2+ with 42Ca+, 86Sr2+ with 43Ca+ and 88Sr2+ with 44Ca+ with quitedifferent mass resolution requirements. The doubly charged species have oneattribute that makes them relatively easy to detect - doubly charged odd-numberisotopes occur at half masses and are therefore free of atomic interferences. TheSr2+ problem can therefore be readily monitored since 87Sr2+ occurs at mass 43.5and hence is free of atomic isobaric interferences from Ca, although the intensity of87Rb+ must also be estimated as well (from 85Rb+).

B. Ionization Modeling

Even though all isobaric interferences may be removed from the isotopes of anelement, the measured isotopic abundances generally differ in detail from thosemeasured in terrestrial materials by conventional mass spectrometry. This variationis due to mass dependent fractionation since the variation in the abundances isdependent on mass and can be modeled by a discrete function. Slodzian et al.(1980) observed that mass fractionation always enhances the abundance of thelighter isotope, the magnitude of the fractionation being proportional to the massdifference, 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 inthat it varies according to the chemistry and structure of the target. They alsofound that fractionation was a function of the spot alignment with the secondary ionextraction axis, and thus instrumental parameters could be critical in reproducibilityof results. Shapiro et al. (1985) carried out a computer simulation of Ar sputtering of aCu matrix and noted an enhancement of the lighter isotope in the ejected materialrelative to the isotopic composition of the target. Because the lighter 63Cu atomstransfer energy more efficiently to other light atoms, a small preference then arisesfor ejection of the light isotope. Moreover Shapiro et al. found that the 63Cu atomscarried a larger proportion of the energy back towards the sample surface thanexpected by proportion which further enhanced the discrimination. Thefractionation also showed an angular dependence with the lightest component beingejected normal to the surface. The computer simulations therefore effectively showthat isotopic mass fractionation can be expected from the nature of the collisionswithin the target.

In chemical analysis, where isotopes of two different elements are ratioed, thevariations with analytical conditions can be even more extreme. This is the mostproblematic aspect of ion microprobe measurements and has severely restricted itsgeneral application to quantitative chemical analysis. In modeling the sputteringprocess both theoretical aspects, involving the interaction between the sample andprimary ion beam, and observational data, which are the manifestation of thesputtering process, are utilized. Theoretical models of the sputtering process can be


constructed using complex computer simulations but the critical dependence ofionization efficiencies of both isotopes and elements with measurement conditionsmust be incorporated into any successful model of the sputtering phenomenon.However, it should be noted that most models have been formulated to take only alimited number of observations into account and therefore their success is onlygauged on a limited number of parameters.

Of particular importance in the development of ionization models for ionmicroprobe mass spectrometry is the local thermodynamic equilibrium (LTE)model of Andersen and Hinthorne (1973) who assumed the sputtering region was adense plasma. They used a simple thermionic emission model to explain the effectsof Cs and O on the electronic work function of the sample and argued that the workfunction affects only the absolute ion yield and not the relative ion yield of theindividual elements. A sample could be characterized by measuring two elements ofknown concentration, and the concentrations of other elements could then bedetermined from the measured ionic ratios. This model produced results that wereaccurate 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 bynormalizing to the intensity of only two peaks. However, in detail the model can not besupported according to the nature of the physical interactions at the sputtering site.No fully quantitative model exists that can be used to estimate elemental abundancesfrom ionic intensities to the required accuracy levels.


While Okuyama and Fujimoto (1986) found that the primary ion currents aresufficient to melt submicron Cr needles, thermal equilibrium models are problematicin that equilibrium conditions require on the order of 100 collisions and yetemission takes place after only several collisions (Harrison, 1983). Even though thephysical processes do not directly support the LTE model, its success in determiningelemental concentrations has led to the general use of the Saha-type equations inother quantitative models (Werner, 1980). A successful general theory of sputteringmust be consistent with the form of the Saha equations although an equilibriumprocess is not necessarily implied.

Other models have been developed from the observation of mass dependentmass fractionation. Slodzian et al. (1980) produced a qualitative model of ionizationwhich incorporated the chemical bonding of the target and was dependent onrelative velocities of the atoms in the matrix. Ionization was assumed to take placeduring the last collision that ejected the atom from the surface and hence matrixeffects are readily appreciated. Shimizu and Hart (1982b) attempted to modelfractionation data from pure metals with the quantum-mechanical based theory ofSchroeer et al. (1973). This model assumes that ionization takes place aftersputtering above the sample surface, and the observed fractionations were consistentwith the formulation. Gnaser and Hutcheon (1988) examined the fractionationproduced in a wide variety of polycomponent substrates for elements ranging inmass from Li to Ti and found that the isotopic mass fractionation is inverselyproportional to the emission velocity of the ion but approaches a constant value forall samples at low velocities. These observations were largely consistent with thebond-breaking model initially proposed by Slodzian (1975) which addressessecondary ion emission from substrates with some ionic character. This model is anadaptation of the Landau - Zener - Stueckelberg model for ion-pair dissociation inthe gas phase and in this case the ionization probability for an ejected particledepends on the distance between the ionized and neutral potential energy surfacesfor 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 ofsecondary ionization including (1) the increased ionization probability forsputtering with O or Cs, (2) the partitioning of energy between the atomic anddifferent molecules of a given element, (3) the matrix dependence of the ion yieldwhich also indicates interaction between the surface and the ionization site. Beyondthis are the isotopic characteristics of a given element, for example, the isotopic massfractionation and its matrix dependence. As yet only components of a generalizedtheory are available. However a complete understanding is not specifically requiredfor 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


isotopic and elemental abundance measurements. In this regard, applications in theearth sciences are based almost exclusively on relative measurements between astandard of known composition and identical mineralogy to the unknown. Then therelative ionization probability of standard and unknown can be assumed to besimilar.

A number of different techniques can be used to achieve the same end, notablythe use of high mass resolution or energy filtering to remove isobaric interferences.In some part the different approaches are used to advantage on different ionmicroprobes, yet to a satisfactory degree, the results can be quite consistent.

III.!HARDWAREA. General Features

Benninghoven et al. (1987) have documented the historical development of ionmicroprobes and described the first generation machines. The first ion microprobeswere recognized as potentially the ultimate weapon of the geoscientist (Lovering,1975) but the difficulties in obtaining useful chemical data and isotopic ratios at lowmass resolution and low sensitivity proved insurmountable to a large extent (see forexample Williams et al., 1983). The latest generation of ion microprobes give highsensitivity at high mass resolution allowing high-precision (permil) measurements tobe made with minimal corrections for isobaric interferences. It is with thesemachines 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 butonly a few are currently involved primarily with research in the geosciences. TheCAMECA ims-3f and derivative 4f and 5f are currently the most widely used butthese instruments do not possess the high sensitivity capabilities of the SHRIMP orCAMECA 1270. The benefits of the larger magnetic analyzer are readily apparent inU-Pb isotopic analyses and on this basis the latest generation of commercial ionmicroprobes are much larger and capable of high sensitivity at high mass resolution.In the following sections the capabilities and features of the CAMECA and SHRIMPion microprobes will be described in some detail. It should be emphasized thatwhile comparisons between the instruments are sometimes given, this article shouldnot be used as a recommendation of one over the other. Designs are changingrapidly and additional features will almost certainly be available by the time ofpublication of this article. More limited information is given for the VG ISOLAB120 since details of this instrument have not been forthcoming.

B. CAMECAThe 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


Figure 7. Schematic diagram of the CAMECA ims-3f.

elements and to obtain elemental abundance information with depth profiles. Thesecond generation machine, the 3f, did find its way into some research institutes foranalysis of geological materials with the promise of in situ analysis of isotopic ratiosand high sensitivity trace element abundance measurements. However, it was soonevident that there were shortfalls with the standard instrument. Specifically, the“off-the-shelf” 3f was incapable of isotopic analysis to the precision levels requiredfor useful terrestrial applications. The imaging capabilities of the CAMECA series isone of their most important attributes but in order to retain image informationcompromises must be made in terms of transmission. However, with perseverance,the CA M E C A ims-3f has turned out to be a highly capable and productiveinstrument for some specific applications. This is no better exemplified than by themodified CAMECA ims-3f at Washington University in St Louis, specific details ofwhich 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; aschematic diagram is shown in Figure 7. Two ion sources are available at the headof the primary column: a duoplasmatron for ionization of source gases on onechannel and a Cs gun on the other. The Cs gun produces Cs+ ions by thermally


ionizing Cs at a tungsten fritt whereas the duoplasmatron ionizes the source gas froma cold-hollow-cathode plasma discharge. Either positive or negative ions can beextracted 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 tothe magnetic prism. The magnetic prism also serves to remove contaminants fromthe primary ion beam such as H and N-bearing species which might otherwiseproduce unwanted interferences in the secondary ion mass spectrum.

The primary column ion-optical array consists of three electrostatic einzellenses with deflection plates to focus and steer the beam to the sample surface. Anoctupole stigmator lens array is used to remove aberrations and change the shape ofthe beam at the sample surface. The primary beam is incident on the sample at anangle of 60˚. The sample surface is held at 4.5kV with respect to ground, thepolarity dependent on whether positive or negative ions are being analyzed. Thesecondary ions are accelerated through a field gradient of 9 ¥ 105 V/m towards thegrounded extraction plate, which acts like a divergent lens, forming a virtual imageof the sample surface and a virtual crossover. The crossover reflects the radialenergy distribution of the secondary ions with the distance from the optical axisbeing proportional to Esin2q where E is the kinetic energy of the ion and q is theangle of emission relative to the sample normal.

The transfer lens system in the CAMECA 3f acts to transport the image of thecrossover and sample surface into a field-free region, and to form real, magnifiedimages of crossover and sample surface in the plane of the contrast aperture and thefield aperture respectively. A continuous range of magnifications can be achievedthrough changing the potentials applied to the transfer lenses. However, as themagnification of the sample surface increases, the magnification of the crossoverdecreases and the spatial resolution of the sample image degrades, and so in practicea compromise must be attained depending on the application involved. For highsensitivity work e.g. trace element analysis, a 75 µm imaged field provides a smallercrossover and hence higher sensitivity. For isotopic analyses, a 150 µm imaged fieldproduces less aberrations of the sample surface which facilitates high mass resolutionmeasurements.

The entrance slit to the mass spectrometer is located in the focal plane of thecrossover. The contrast aperture restricts the diameter of the crossover therebyreducing spherical aberrations and improving the spatial resolution of the ion imageof the sample (at the expense of sensitivity). The entrance slit is closed down toproduce a vertical line image of the crossover (Figure 8) and hence for high massresolution a large fraction of the total secondary ion beam is excluded. The contrastaperture consists of a sliding plate with different diameter holes that can be movedrelative to the beam in both directions perpendicular to the beam path. The fieldaperture is positioned near the chromatic focal plane of the electrostatic analyzerwhich is coincident with the position of the real image of the sample surface. Againthis aperture consists of a sliding plate with different hole diameters and allows


Figure 8. CAMECA ims-3f entrance-slit images of 29Si+ and 28SiH+ following closure ofthe entrance slit to form the vertical line images; from Fahey (1988). The two speciesare separated by 0.0083 amu.

transmission of ions from a selected area on the sample i.e. it allows for masking ofunwanted regions of the sample surface.

The mass analyzer is a double focusing mass spectrometer consisting of anelectrostatic analyzer with turning radius of 17.3 cm and a magnetic prism withturning radius 12.7 cm. For a given charge to mass ratio (q/m) a real image of thecrossover is focused at the exit slit which is also focused in terms of energy. Inaddition, the mass analyzer is designed to transport images through an electrostaticlens that couples the electrostatic analyzer (ESA) and magnet. The ESA consists oftwo electrodes cut from concentric spheres with inner and outer plates held atpotentials such that the central beam path is at 0 V with respect to ground and theelectric field between the plates accelerates ions of 4.5 kV (nominally) along thecentral beam path. The ESA forms an image of the crossover that is radiallydispersed in energy; the energy slit at this image point allows for the selection of anenergy window that is variable up to 130 eV wide. The energy slit is thereforeuseful for reducing chromatic aberrations for high mass resolution measurementsand 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 toalign the image of the crossover with the image of the sample and also align theimage of the entrance slit with the exit slit. The exit slit is closed down to maskunwanted species from the detectors. Two projector lenses can be operated totransform the virtual image produced in the magnetic prism to a real image that can


be displayed on a fluorescent screen, or an electrostatic lens can be used to deflectthe beam into a Faraday cup or electron multiplier for quantitative analyses.

The CAMECA ims-4f has the same basic configuration as the 3f with severaladditional components. A normal incidence electron beam is designed to produce acloud of low energy electrons just above the sample surface in order to compensatefor sample charging during positive ion bombardment (usually Cs+) and negativeion extraction (including electrons). The charge compensation automatically adjuststo the sample charging with sufficient electrons extracted to neutralize the chargeand excess electrons reflected from the surface. In this method electrons areaccelerated from a filament and focused with a pair of quadrupole lenses and theresulting beam is introduced to the secondary ion extraction array with anelectrostatic prism. The presence of the prism necessarily also affects the secondaryion trajectories and so a set of compensating deflectors are included in thesecondary ion optical array to compensate for the perturbations. This system hasbeen 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 electronbeam within the spot and variations in O-isotope ratios of several percent are evident(E. Zinner, pers. commun.). However, recent reports from CAMECA suggest thatthis problem may be solved by focusing the electron beam according to the imageof the crossover. Full charge compensation is achieved with an evenly illuminatedimage 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 inthe raster imaging mode to minimize aberrations caused by secondary ion extractionaway from the optimal central axis. An electrostatic deflection system is linked tothe raster generator to recenter the trajectory of any extracted ions with the fieldaperture. In this way it appears to the secondary ion extraction system that thesecondary ions are always coming from the focal point of the extraction system andhence relatively large areas can be imaged while maintaining high mass resolution.

The CAMECA ims-5f is the latest version in the "small" CAMECA series. Thebasic configuration is the same but it includes a separate primary column for aliquid metal source that is used for producing secondary electron and secondary ionimages. The 5f also has a dedicated programmable interface for controlling all lenssettings so that tuning conditions for different applications can be recorded andrecalled.

The latest ion microprobe to be marketed by CAMECA is the ims-1270 (Figure9) and the first instrument of this type has been delivered to the University ofCalifornia at Los Angeles. This instrument incorporates a completely redesignedmass analyzer that has the imaging capabilities of the 3f-4f-5f series as well as a hightransmission mode whereby beam transport (i.e. sensitivity) can be maximized at theexpense of the imaging. The interrelationships between these different


Figure 9. Schematic diagram of the CAMECA ims-1270.

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

h = 1.214 ¥ 106 DM/Mwhere h is the perpendicular distance in micrometers at the collector betweenadjacent masses of M and M+DM.

The primary column and source housing of the 1270 analyzer are similar to the3f-4f-5f series and include a Cs source and a duoplasmatron. The sample interlockon the UCLA machine has been modified so that there are two stages of pumping torestrict contamination of the source chamber and maintain ultrahigh vacuum(≤5¥10-10 Torr); three samples can be loaded into the intermediate vacuum chamberat any time. The secondary ion extraction system and transfer lenses are similar tothe 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 themass analyzer. In the ion microscope mode, image quality is paramount and solenses are operated to achieve the lowest possible spherical and chromatic


aberrations in the imaging plane of the channel plate. Two circular coupling lensesare situated between the ESA and magnet to ensure optimal transfer of ions. Thismode can also be utilized in an ion microprobe mode, but to achieve the highestpossible sensitivity at high mass resolution a beam matching system operating in theXY mode allows for maximum transmission of ions at the expense of imageinformation. In this mode, two slit einzel lenses provide a 1:5 magnification of thefield aperture in the magnet while the entance slit image is magnified by a factor of5; in this way the b aberration of the magnet is reduced. Slit einzel lenses are alsopresent before and after the magnet to compensate for aberrations produced in theelectrostatic peak-switching mode. In addition to these lens systems, sextupolelenses are positioned before and after the ESA and magnet to allow for maximumcontrol of the transmission characteristics. Projection lenses are situated after theexit slit to transfer real images of the crossovers to an imaging plane containing thechannel plate detector.

The UCLA ims-1270 is currently undergoing specification testing andpreliminary tests indicate that the Pb sensitivity of this machine is at least as good asSHRIMP 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 designedand constructed at the Australian National University with the purpose of analyzinggeological materials with sufficiently high mass resolution to eliminate most majorisobaric interferences while maintaining the highest possible sensitivity (Clement etal., 1977). This is achieved by a physically large mass analyzer (magnet turningradius of 100 cm) which gives high mass dispersion and therefore allows for highmass resolution operation with wide slits. In addition the secondary ion beam ismatched in terms of its phase space characteristics with the acceptance of themagnetic analyzer to maximize the secondary ion beam transmission. Thisoperation is at the expense of the ion optical image of the surface and so SHRIMPcannot 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 asa commercial prototype.

A schematic of SHRIMP I is shown in Figure 10. The SHRIMP I ion source is acold-hollow-cathode duoplasmatron which lies on the primary beam central raypath. 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 acrossed electromagnetic field that selects ions according to their velocities. It hasbeen found that 16O2

- at mass 32 is the most effective primary species on SHRIMP Isince it has the same intensity as 16O– but the larger mass of the molecular speciescreates a larger secondary ion signal. The use of O2– may also lead to a higher oxygenconcentration at the sputtering site, hence producing a larger ionization yield. The


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

SHRIMP einzel lenses generally run in a configuration which has been called Kohlerillumination and relies on the insertion of an aperture between EL1 and EL2 at thefocal point of EL2 (Figure 11). The aperture acts as the object for EL2 and the spotsize is a demagnified image of that aperture alone. Kohler illumination has themajor advantage that irregularities in the brightness of the source and aberrationsupstream from the aperture do not cause irregularities in the final spot because theobject 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 Figure12. The spot size does not vary with primary beam current as is the case for criticalillumination where the spot is a demagnified image of the extraction aperture. Aneinzel lens is situated after the primary extraction cone and allows the currentdelivered to the kohler aperture to be continuously varied. Drawbacks for thismethod are that the apertures are fixed in size and a continuously variable spot sizeis 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. Fourof these apertures are held in a sliding plate that can be removed and the apertureschanged during the routine servicing of the duoplasmatron. The primary beam hasan incidence angle of 45˚ with respect to the sample surface yielding


Figure 11. Primary column focusing modes on SHRIMP. Critical illumination relies onthe demagnification of the source aperture by the two Einzel lenses EL1 and EL2; thegreater the demagnification the smaller the spot and the lower the beam intensity. InKohler illumination, the Kohler Einzel lens is activated to transfer the maximum beamintensity to the Kohler aperture which acts as the source for the final lens. In thisconfiguration, the spot diameter is a function of the Kohler aperture diameter andaberration contributions are limited to those from the final lens only. Primary beamintensity can be controlled by the strength of the Kohler lens and does not change thespot 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 bemeasured at the sample. A 450 V bias, with respect to the first extraction electrode,then accelerates the secondary ions towards an intermediate electrode and theextraction aperture through a relatively low electric field. The extraction potentialfor positive secondary ions is approximately 10 kV. The secondary-ion opticalarray 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 thatproduces 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 theXY plane and one in the XZ plane which are capable of steering and focusing thebeam so as to match the acceptance of the mass analyzer. Whereas in the CAMECA3f the crossover is masked to yield a line image, in SHRIMP I the beam is focused toa line image at the entrance slit thereby minimizing beam loss. A fraction of thesecondary ion beam is monitored at an electrically isolated and suppressed aperture


Figure 12. Ion probe crater produced by Kohler illumination. This method facilitatesthe production of steep sided evenly illuminated craters. Target is a sulfide withconsiderable S-isotopic heterogeneity as indicated by labelled d34S values (photo fromC. S. Eldridge).

between the intermediate einzel lens and the matching system. The aperture isapproximately 3 mm in diameter and allows most (≥90 %) of the secondary ionsthrough, but the remainder is collected as being representative of the total secondarybeam. During data collection, the counts on each isotope can be divided by thecounts on the secondary beam monitor during the same time interval. Thisprocedure approximates to a double-collecting mode of operation and can result insignificant improvements in precision by reducing the effects of instability in theprimary beam.

The double-focussing mass analyzer is based on a mass spectrometer design byMatsuda (1974) and the main ion-optical elements are a cylindrical 85˚ ESA, aquadrupole, and a 72.5˚ sector magnet with non-normal entry. The physicaldimensions 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 maintainingrelatively wide slits to allow high sensitivity. The magnification of the mass analyzeris 0.4 and the mass dispersion relationship is

h = 7.73 ¥ 105 (DM/M)


where h is the perpendicular distance in micrometers at the collector betweenadjacent masses of M and M+DM.

An intermediate image point exists within the ESA which allows definition ofenergy limits within a bandwidth of 150 eV in 25 keV. Adjustable slits are alsopresent at the entrance and exit of the ESA to limit beam divergence; the acceptanceof the ESA is limited to <0.01 radians. A retractable Faraday cup, the post-ESAmonitor, is positioned at the exit of the ESA and, in conjunction with the divergenceslits, allows accurate alignment of the secondary beam along the central ray path ofthe ESA. During initial set-up routines, the post-ESA monitor can be used tooptimize the secondary beam to the magnet without the necessity of a peak beingaccurately aligned through the collector slit.

An optimum focal point exists in the plane of the mass analyzer for bothangular 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. Thefocal point of the SHRIMP I mass analyzer has been found to be field-dependentwith, for example, the optimum focus of 46Ti being 6 mm nearer the magnet thanthat for 50Ti (Ireland, 1986). A compromise position optimizing the peak shape of48Ti produced significant degrading of the peak shapes of 46Ti and 50Ti, and ratioscollected under these conditions showed an instrumental bias of approximately -0.1% in the fractionation-corrected 46Ti/49Ti and 50Ti/49Ti ratios. The exact cause ofthis condition is not clear but in order to remedy its effects, the entire collectorassembly is under computer control and a separate position can be assigned for eachmass.

SHRIMP I was originally fitted with a single collector assembly capable ofreceiving ions in a Faraday cup or electron multiplier. A sliding base-plate on thecup allowed the ion beam to pass to the ion counter. The single collector has sincebeen replaced with a multiple collector housing that is designed to enable operationunder a single-collecting mode or multiple-collection of up to eight masses. Theindividual ion beams can be centered on their respective collector slits by pre-slitdeflection-plates, while post-slit deflection-plates optimize transmission to theelectron multipliers. A Faraday cup is available on the central ray and can beswitched in and out of the beam path. The collector slit widths for all but the centralposition 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 collectormode.

The new commercial SHRIMP II uses the same mass analyzer configuration asSHRIMP I but there are substantial modifications particularly in the primary column- source chamber designs. The sample interlock can hold up to four samples andtwo samples can be in the source chamber at any time. The source elements of theprimary column were manufactured by Cambridge Mass Spectrometry Ltd. Theduoplasmatron operates with air cooling and a fixed magnetic field and is designedto optimize the ion transmission by producing a low divergence beam. The primarymode of operation of the primary column is kohler illumination and the demagni-


fication of the final lens element is ≈7 and so a final spot size of the order of 5 µm isproduced for a 30 µm Kohler aperture.

The secondary ion extraction system for SHRIMP II is similar to SHRIMP I andconsists of an intermediate extraction electrode and transfer system to produce animage of the crossover at the focal point of the beam matching system. The beammatching system consists of three quadrupole lenses which have lower aberrationcharacteristics than the slit einzels in use with SHRIMP I. The mass analyzer is basedon the same design parameters as SHRIMP I although the the sensitivity of Pb+ inzircons has been measured at over 20 c/s/ppm!Pb/nA under routine analyticalconditions 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. Acommercial prototype has been installed at Cambridge University, England and isnow under operation although a detailed description of the machine and analyticalcapabilities is not available as yet. The configuration is similar to SHRIMP with afew notable differences. The source can operate with a thermal ionization source aswell as the primary column used for ion microprobe analysis and ports are availablefor laser photoionization of sputtered or thermally produced neutrals. The massanalyzer is double focusing, with a focal plane normal to the ion optic axis, andfeatures a 96 cm radius electrostatic analyzer with a 70˚ angle followed by a 60 cmradius magnet with an effective dispersion length of 120 cm. An energy band passcan be defined through slits at the intermediate energy focus between the ESA andmagnet. The multiple collector has a microchannel plate as well as adjustableFaraday collectors. Very high abundance sensitivity (up to order of 10-11) isprovided by a second electrostatic analyzer following the magnetic sector with asecond collector assmbly consisting of a Daly/photomultiplier combination.

IV. METHODOLOGYA. Isotopic analysis

The ion microprobe is essentially a mass spectrometer and so the data it generatesare treated in much the same manner as data from a conventional thermal ionizationor gas mass spectrometer. Ion microprobes are generally fitted with both a Faradaycup and electron multiplier detection systems, however, the intensity of secondaryion beams is generally insufficient to warrant measurement by a Faraday cup, andisotopic measurements are typically made with the multipliers which have muchlower 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 incomingion when any subsequent incoming ion will not retrigger the system For electron


multipliers this time is of the order of 5 ns but the total deadtime of the countingsystem as a whole is typically around 20 ns. The total deadtime is affected by theresponse time of the discriminator and scalar used in the counting system , as well asthe gain of the multiplier. In addition, Zinner et al. (1986a) have shown that pulsecharge distributions (and hence deadtimes) in electron multipliers are a strongfunction of ion type (atomic vs. molecular species), element, polarity, and energy.Small but noticeable effects were also observed in the detection efficiencies forisotopes of Ca and Ti which introduced a small mass fractionation effect.

The precision of a pulse-counted measurement is limited by the total number ofcounts, 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 thelarger the number of total ions the higher the precision and the more intense thesecondary signal the shorter the time period that is needed to obtain the requirednumber of ions. However, in isotopic measurements at high count rates, theuncertainty in the dead time can ultimately be the limiting factor for the accuracy ofa measurement.

For a retriggerable system, the measured count rate, cmeas, is related to theactual count rate, ctrue, by

†

cmeas = ctruee-tctrue

where t is the counting system dead time. The dead time correction can thereforebe approximated by

†

ctrue ª cmeas exp cmeast ⋅ exp cmeast ⋅ exp cmeast ⋅ exp cmeast( )( )( )( )For accurate isotopic measurements, the dead time must be accurately known andindependent of count rate as is illustrated in Figure 13.

Hayes and Schoeller (1977) have addressed in some detail the limits inprecision and accuracy attainable by pulse counting. When ratios differsignificantly from unity, the uncertainty in the dead time determination, and notsimply the magnitude of the dead time, becomes the limiting factor. Specifically,

uN2 = (sN/N)2 ≈ (ƒt)-1 + ƒ2sr2

where sN is the coefficient of variation of N, the number of counts, ƒ is thefrequency of events, t is the collection time, and sr is the standard deviation of r, thedeadtime. Ignoring the uncertainty in the dead time, the relative standard deviationis simply 1/√ƒt, the total number of counts. However, when sr is non zero, thesecond term contributes to the relative standard deviation in a way that is dependenton the count rate. In this way Hayes and Schoeller (1977) found that the maximumcount rate Fmax for any measurement is given by

Fmax = (uN)reqd/sr = sN/Nsr


Figure 13. The counting system dead time is typically around 20 ns for an electronmultiplier based system and should be independent of count rate as illustrated in thisfigure 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 is5 ¥ 105 c/s and the minimum time for the measurement is 565 seconds, or around 10minutes. In order to obtain a factor of ten improvement in precision, the maximumcount rate must be decreased by a factor of 10 but the minimum measurement timeis dependent on the inverse third power of sr and so the counting time must beincreased by a factor of 1000. For ion microprobe analysis, it is often not possibleto sustain analyses over such a period of time since the sample may be consumed orthe crater become too deep for reliable analysis.

For isotopic analysis the magnet is cyclically stepped through the peaks ofinterest. The counts are measured on the top of each peak and a number of suchcycles are combined to give the primary ratio information for the analysis. It isimportant to combine a number of cycles in order to address the degree of temporalchange in peak heights during analysis. The causes of temporal drift may beprimary beam fluctuations or of more concern, change in the sputtered compositionof the sample whether it be related to real compositional changes or related to achange in the sputtering/ionization conditions.

The stability of the primary beam is an important factor in ion probe analysissince any noise from this source is propagated through to the secondary ion beam.


Short-term noise simply expresses itself as an additional random contribution to thesecondary-beam noise but longer-term drifts can cause more problematic variationswhich are of the same type as compositional variations. Such secular drifts must becorrected by a method of interpolation. The denominator peak is usually chosen tomonitor the fluctuations in the medium to long-term signals and a typical means oftemporal correction is simply to assume that there is a linear change in signalbetween one denominator peak measurement and the next. Clearly this is applicableso long as the fluctuations are of significantly longer time period than themeasurement cycle. Correction of medium term fluctuations is quite difficult sinceat some stage the real variations must be separated from the actual beam noise. Suchfluctuations might be addressed by attempting to fit a polynomial to thedenominator peak, or by using two peaks to monitor beam fluctuations. Quite oftenmedium term fluctuations are symptomatic of some hardware disorder and time isoften better invested in trying to solve the problem than making elaborate algebraiccorrections to the data set.

When isobaric interferences cannot be resolved they can be stripped from thepeak of interest by monitoring another isotope of the interference. This correctioncan be made from the final time-interpolated ratios, but if the magnitude of thecorrection changes through the analysis then an error must be propagated throughto the mean ratio measurement that reflects the change. An alternative method is tostrip the interference during each cycle so that the random error associated with eachmeasurement is propagated directly to the stripped peak in each cycle. Then if thesample composition changes, the interference is stripped from the same period ofisotope measurement and the mean isotopic ratio can be calculated withoutpropagating an error due to fluctuation of the interference. However, systematicerrors might still arise because of uncertainties in the isotopic ratios of theinterference element. Such uncertainties may be due to the magnitude of isotopicmass fractionation, or in the case of meteoritic samples, could be due to uncertaintyin the actual isotopic composition. Such errors are minimized if a large peak can beused to monitor a smaller interference on another element.

Isotopic mass fractionation is one of the fundamental properties of isotopicratios produced by secondary ionization. Isotopic mass fractionation is also animportant parameter in systems subject to physicochemical reactions in nature,particularly for light elements, and so the instrumental mass fractionation must beremoved in order to ascertain the intrinsic fractionation component. Instrumentalmass fractionation is known to be a function of a number of machine parameterssuch as secondary ion energy and the axial position of the secondary ion extraction,as well as being matrix dependent (Shimizu and Hart, 1982b). The moststraightforward approach to monitoring instrumental isotopic mass fractionation isto measure terrestrial standards of the same mineralogy as the unknowns, and whosecompositions can be analyzed by conventional mass spectrometric techniques. Thisdegree of mass fractionation can then be removed from the unknowns leaving theintrinsic fractionation. However, the uncertainty in the measurement of the terrestrial


standards must also be propagated onto the unknowns, during the course of theanalyses 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 terrestrialratio is given by

DjA = [(jA/iA)meas/(jA/iA)terr – 1] ¥ 1000/(j-i) (‰/amu)

where iA and jA are two isotopes of element A, and the subscripts meas and terrrefer to the measured and terrestrial ratios respectively; a positive mass fractionationis defined as heavy isotope enrichment. Typically the terrestrial ratio for thefractionation normalization is taken from conventional thermal ionization analysesin order to facilitate comparisons. The intrinsic mass fractionation in an analysis ofelement A can be expressed as the difference between the fractionation measured inthe unknown with that measured from a standard i.e.

FA = DjAmeas – DjAstandard (‰/amu)This procedure is an external calibration of isotopic mass fractionation since itdepends on the measurement of another reference material to calibrate themagnitude of the effect. For an element with only two isotopes this is the onlymethod of removing the instrumental mass fractionation effects. In general thereproducibility of the measurements on both standards and unknowns is the limitingfactor and only relatively large effects (≥ 1 ‰) can be resolved.

Where there are three (or more) isotopes of a given element, an internal massfractionation correction can also be applied by using one of the isotopic ratios toremove the mass fractionation from the other ratio(s). This is carried out byadopting a mass fractionation law that describes the relative fractionation of theisotope ratios according to their masses and expressing the corrected isotopicabundances relative to the terrestrial ratio. In this case the normalization is notdependent on the measurement of a standard but rather requires only that the massfractionation obeys a predetermined law whose form is rather arbitrary since it needonly describe the dependence of the measured ratios with isotopic mass. Theprecision 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 threeisotopes we can form two ratios, for example jA/iA and kA/iA. For the moment let usconsider the algebraically simplest case of mass fractionation, that is it is linearlydependent on isotopic mass. In this case we will use jA/iA to monitor the massfractionation and to derive a correction for kA/iA so that we can express kA/iArelative to the terrestrial ratio. As above we have DkA and D jA and we remove theeffects of the linear mass fractionation from DkA to leave the residual dkA which isgiven by:


dkA = DkAmeas – (k-i)/(j-i)DjAmeas (‰)In this scheme, terrestrial standards should all have residuals within error of zerosince we are simply testing the form of the mass fractionation law.

A number of mass fractionation laws have been used in both conventional andion probe mass spectrometry although they all have a similar form. In the powerlaw, the fractionation between two adjacent isotopes is a and is the same fora n y adjacent isotopes within a region of interest. If we have the terrestrialabundances (jA/iA)0 and (kA/jA)0, and j = i+1 and k = j+1, then the measured ratiosof the adjacent isotopes (jA/iA) and (kA/jA) are

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

(kA/jA) = (kA/jA)0 (1+a)and therefore by multiplying

(kA/iA) = (kA/iA)0 (1+a)2

In practice, (jA/iA) can be measured to obtain a and hence allow a power lawcorrection on (kA/iA). It can be seen that if we expand the power law formulationassuming 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 fitto the fractionated isotopic abundances of Ca and so developed the exponential massfractionation law which has the form

(jA/iA)/(jA/iA)0 = [mj/mi]b

where mi and mj are the isotopic masses of iA and jA.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 initiallyand after evaporation respectively. This is the Rayleigh law.

Esat (1984) has shown that these laws all have the same basic form. If we definethe ratios Rik and Rjk, where Rik = iA/kA and Rjk = jA/kA, then

power law: Rik = (Rjk) (mi-mk)/mj-mk)

Rayleigh law: Rik = (Rjk) (√mi-√mk)/√mj-√mk)


exponential law: Rik = (Rjk) log(mi/mk)/log(mj/mk)

from which the generalized form of the fractionation laws can be seen to be

Rik = [Rjk]g

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 simplyimparting different degrees of curvature to the mass fractionation function for anygiven element. In practice, the real functional relationship must be determined foreach element on each mass spectrometer.

Isotopic Analysis by Ion Imaging

The ion microscope mode offers the possibility of taking images of differentisotopes and through digital processing, the signals can be ratioed and the isotopiccomposition of the area in question can be determined. Clearly for a large areathere may be no particular benefit, but when the samples are small, or heterogeneousover 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 ofcorundum grains by imaging in 16O and 18O. In one image some 5-15 grains couldbe 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 containinginterstellar SiC grains of 3-5 µm size range. Such ion images allow the rapididentification of the highly exotic but rare grains X. These grains are characterized byan extreme overabundance of 28Si relative to the minor isotopes compared to thenormal values. The exposure times of the images are such that equal intensity in 28Siand 30Si equates with a normal composition. The X grain clearly has a lower intensity in30Si relative to 28Si. (Photomicrographs provided by L. Nittler).


allowed the clear identification of a highly anomalous grain with d18O of ≈-200 ‰,which could then be analyzed to higher precision, including the measurement ofd17O, in a normal peak-jumping mode. Nittler et al. (1993b) have also used thistechnique to identify highly anomalous but rare SiC grains X. Over a period ofthree days around 1250 grains were analyzed and 9 grains X located. Successive Siion images of a region containing an X grain are shown in Figure 14. The exposuretimes are such that a normal composition should result in equal intensities on eachphotograph. The intensity of the 30Si+ image is clearly lower than that of the 28Si+image indicating its anomalous isotopic composition enriched in 28Si. Isotopicanalysis by ion imaging will of course offer only limited precision but evenmoderate levels can be adequate for highly anomalous meteoritic samples.

Recipes for isotopic analysisThe 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 anoutline only and for more detailed explanations the reader is referred to the primarysource(s).

Hydrogen. The isotopic measurement of hydrogen is difficult for tworeasons: the D abundance is very low (0.015 % of normal H) and the fractional massdifference 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 isrequired to separate D+ from H2+. The terrestrial D/H ratio of Standard Mean OceanWater (SMOW) is 0.00015576 (Hagemann et al., 1970).

The first measurements by Hinton et al. (1983) were made with an O– beamwith hydrogen isotopes measured as H+ and D+. Since they were attempting tomeasure 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 secondaryions. 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 alsofound that the isotopic mass fractionation of the hydrogen isotopes was far less thanthe case for positive ions. In order to minimize the effects of sample charging withthe Cs beam, the samples were pressed into a gold foil. During analysis, the energydistribution of the secondary ions was monitored and the sample charging wascompensated by offsetting the accelerating voltage. Terrestrial standards werereproducible to around the ±30 ‰ level.

Deloule et al. (1991a) were interested in measuring H isotopes in terrestrialsamples thus high precision and accuracy is imperative. They used an O– primarybeam and measured H+ and D+ at a mass resolution of 1300 R to separate H2+ fromD+. They took care to remove moisture from the sample surface by baking it in theion probe at 120 ˚C and used a liquid nitrogen cold trap to fix residual water in thevacuum. Measurements commenced when the H2+/H+ ratio was lower than 8 ¥ 10-4. Rather than simply comparing D/H ratios of standard and unknown,


Deloule et al. (1991) found that the instrumental fractionation could be furthercalibrated by the measurement of Si, Ca, Ti, and Mn ion intensities on the samematerials. The error of the best-fit calibration is around 7 ‰ and the dD can bemeasured to a precision of around ± 10 ‰. An interesting aspect of this work wasthe observation that the crystallographic orientation of mica may be an importantfactor 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. Thecorollary to this observation is that for successful in situ measurements it may benecessary to document the crystallographic orientation and apply a correction.

Boron. Boron has two stable isotopes, 10B and 11B with a normal 10B/11B ratio of0.24726 as determined by De Bièvre and Debus (1969). Chaussidon et al. (1990)measured boron isotopic compositions with an O– primary beam and B+ secondaryions at a mass resolution of ≈2000 R to eliminate 10BH+ from 11B+. Instrumentalmass fractionation for B was monitored in a similar fashion to D/H measurements inthat mass fractionation was found to vary linearly with the product of mass/chargeand 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. Thereproducibility of these measurements was better than 1 ‰.

Carbon Carbon has two stable isotopes, 12C and 13C, with a 13C/12C of 0.011237 inPee Dee belemnite (Craig, 1957). The main isobaric interference is 12CH+

interfering with 13C+ which requires a mass resolution of ≈3500 R. Carbon can beanalyzed as either negative or positive secondaries (with Cs+ and O– as the primarybeam species respectively), although C– has several advantages: it is more efficientlyionized, 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 differentmineralogical composition in order to address the possible effects of matrix effectson instrumental mass fractionation. While the absolute mass fractionation was quitelarge (45-50 ‰/amu), within measurement errors it was found that there was nodifference between graphite and carbonate in terms of mass fractionation. Therewas a slight effect for kerogen, but McKeegan et al. (1985) concluded that matrix-dependent fractionation differences between the graphite standard and anyreasonable C phases did not exceed 15 ‰.

Harte and Otter (1992) also used a primary Cs+, secondary C– configuration toanalyze C isotopes in diamonds. They found a similar range in instrumentalfractionation from -56 to -34 ‰/amu and the fractionation was also found to vary asa function of age of the electron multiplier. The typical precision of an individualanalysis 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 wasinterpreted as being due to heterogeneity in the standard. The standard had been


analyzed previously by conventional mass spectrometry and was thought to beheterogeneous 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 ionmass spectrometry. It does not form N– secondaries and yields N+ ions a thousandtimes 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 aubiquitous contaminant and required a mass resolution of 6000 R to eliminatemolecular interferences from the 12C14N– and 12C15N– peaks. The CN– signal dependson the bonding between C and N as well as the absolute concentrations of C and Nand so no quantitative information on concentrations is possible unless the standardsare identical to the unknowns. Zinner et al. (1987) were able to reproduce thed15Nair in 1-hydroxybenzotriazole to within 2 ‰, in synthetic SiC crystals, and ingraphite on to which air was blown from a controlled vacuum leak.

OxygenOxygen 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.McKeegan (1987) used ion microprobe measurements to derive a value of 17O/16O =0.00038309 when normalized to the 18O/16O of Baertschi (1976). This is a difficultelement to analyze because of the large difference between the abundance of 16Oand the minor isotopes 17O and 18O and because of the large 16OH– interference on17O–. In meteoritic work it is essential to measure all three O isotopes whereas interrestrial problems measurement of the 18O/16O will suffice because the onlyparameter of interest is mass dependent fractionation. Since oxygen is a stronglyelectronegative element it is best analyzed by sputtering with Cs+ and measuring O–,however sample charging must be overcome in order to obtain reproducible andreliable data.

McKeegan (1987) first described the mounting of 10-15 µm particles on goldfoil which allows charge to dissipate rapidly. This technique has also been used byFahey et al. (1987a), Virag et al. (1991) and Ireland et al. (1992). For inclusionslarger than the required 10-15 µm, they must first be crushed, transferred to the goldfoil, and then pressed in with a quartz plate. A primary beam of 14.5 kV Cs+ ions isused to sputter the samples and O isotopes are measured as 4.5 kV O– secondaryions. The primary beam is defocused such that the entire particle is sputtered andfor hibonite and spinel particles between 8 and 15!µm a stable 16O– count rate of500,000 c/s can be maintained for measurement.

The major isobaric interference for O-isotopic measurements is 16OH– inter-feringwith 17O–. The hydroxide species is mainly sputtered from the surrounding goldmount 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.


The OH– is apparently due to surface water migrating on the gold; water is a residualto the vacuum and the 16OH–/17O– falls with time after sample insertion and hencepumping time. A field aperture at the image point after the source slit is used tomask 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 goldmount in order to sputter-remove the migrating water before it reaches the area fromwhich secondary ions are accepted.

A mass resolution of 6500 R is sufficient to separate 16OH– from 17O–, andpossible tail contributions of 16OH– are monitored by measuring the tail on the lowmass side of 16O– at a mass offset of 16

17(m16OH – m17O) where m16OH and m17O arethe masses of 16OH and 17O respectively. Under these conditions the 16OH–/17O– istypically 2, corresponding to a correction of around 0.8 ‰ on 17O, after one day ofpumping in the sample chamber.

Analyses of the Burma-spinel standard are interspersed with the unknowns andare used to normalize the O-isotopic composition for instrumentally-inducedisotopic mass fractionation. The instrumental mass fractionation measured byIreland et al. (1992) was within a range of –3 to –20!‰/amu. The use of theBurma-spinel standard for the normalization of O-isotopic analyses of other oxideminerals, such as hibonite, has been examined by McKeegan (1987). He found nosystematic differences between the instrumental mass-fractionation measured interrestrial samples of hibonite and spinel that had previously been measured byconventional mass-spectrometric techniques and therefore a single standard could beused in the O-isotopic measurements.

The O-isotopic compositions measured from the meteoritic samples arereferenced to the O-isotopic composition of Baertschi (1976) and normalized to themean oxygen composition measured from the Burma-spinel standard in the sameanalytical session; the Burma spinel has d18OSMOW of +22.1 ‰ and d17OSMOW of+11.6 ‰. While individual analyses have a precision in d18OSMOW of around 3 ‰,the scatter of measurements on individual Burma-spinel grains in a session istypically around 8!‰. This variation, which limits the overall precision of themethod, is mass-dependent fractionation since the Burma spinel data scatter along aslope 1/2 mass fractionation line on an oxygen three-isotope plot while deviationsfrom this line are consistent within individual measurement errors. The reason forthe variation in isotopic mass-fractionation is unclear, but is probably related to theirregular geometry of the grains producing variations in the isotopic-densitydistribution at the source slit of the mass spectrometer.

In meteoritic work, it has been found that the 16O excess is an importantparameter since large anomalies are found in the O-isotopic compositions ofmeteoritic materials. The anomaly is attributed to the 16O abundance since the17O/18O ratio is generally close to terrestrial. The 16O excess is calculated according toClayton and Mayeda (1983),


16Oexcess =0. 52

1 - 0.52( )d18O -

11 - 0. 52( )

d17O

For the graphical representation of the O-isotopic measurements individualmeasurement errors on the meteoritic samples are combined with the variation ininstrumental mass fractionation as determined by the spread of the Burma spineldata points along the slope 1/2 mass fractionation line (Virag et al., 1991). Thisspread is expressed by an error ellipse whose main axis has a slope of 1/2 and whoseaxes are the standard deviations of the standards along the mass fractionation lineand normal to it. Folding of this ellipse with the error ellipse representing individualmeasurement errors (in d17O and d18O) yields a final error ellipse whose axes aregenerally neither aligned with the coordinate axes nor aligned with the fractionationline and whose orientation is different for different data points.

The method for oxygen isotope analysis described above is somewhatunsatisfactory because in situ analyses of insulators in polished thin sections are notpossible. Material must first be removed from the thin section and then mounted ingold. The alternative is to use a charge compensation gun which neutralizes thecharge 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 compositionsof Allende refractory inclusions. This method has high sensitivity, up to 1 % of theO 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 thesputtering rate is very low and extremely long counting times are required to achievethe 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 usescharge compensation and extreme energy filtering (only ions with energies greaterthan 300 eV ion are collected) to discriminate against molecular interferences andreduce matrix effects. Precision for this technique is around ±2 ‰ for the ≈1.5 houranalysis times. The unsatisfactory aspect of this technique is the high loss of iontransmission with energy filtering at such high energies and therefore the loss ofsensitivity.

The problems of sample charging under Cs bombardment can be overcome byanalyzing conductors and O isotopic compositions can be readily measured to highprecision. Read et al. (1990) used a VG Isolab 54 ion microprobe (Lyon andTurner, 1992) with a multiple collector to analyze oxygen-isotopic compositions inpolished sections stainless steel from the LDEF satellite. They obtained 16O– signalsof ≈5 ¥ 106 c/s and a precision on the 17O/16O ratio of 0.1 % in thirty minutes withreproducibility 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 - 2.0nA. Polished thin sections were gold coated but the conductivity of the gold wassupplemented by a thin line of Ag colloid paint connecting the grains to the sample


holder. Secondary O– ions were accelerated through 4.5kV and the mass resolutionof the spectrometer was 3000 R at the 10 % level. A field aperture was used to maskall but the central 8 µm of the primary beam spot. The 16O– count rate was limited to500,000 c/s to minimize problems with dead time uncertainty and ≈106 counts of 18Owere collected resulting in a theoretical uncertainty of 1 ‰. Reproducibility at thislevel appears to have been achieved since all 2000 individual 18O/16O ratios fit aGaussian curve with the predicted precision. The data were normalized to SMOW bymeasuring internal standards which were assigned the d18OSMOW measured byconventional mass spectrometry. This method can be used for the quitehomogeneous magnetites of this study but could prove unsatisfactory if quiteheterogeneous samples are to be measured.

Magnesium. Magnesium has three stable isotopes, 24Mg, 25Mg, and 26Mg. Themain isobaric interference is from hydrides, in particular the 24MgH+ interferencewith 25Mg+ which requires a mass resolution of 3500 R; at this level all othermolecular interferences are resolved. The most commonly cited terrestrial Mgisotopic ratios are 25Mg/24Mg of 0.12663, and 26Mg/24Mg of 0.13932 (Catanzaro etal., 1966). The actual Mg isotopic composition has been a subject of somediscussion since different laboratories report different 26Mg/24Mg ratios whennormalized to the common value of 0.12663 for 25Mg/24Mg. This is mainly apparentin 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 etal. (1966) value, as did Clayton et al. (1984), and Ireland et al. (1986) obtained avalue of 0.139432. However, the absolute values are not important in this type ofisotopic analysis since we are concerned with differences in isotopic compositionrather than absolute ratios. Therefore the important aspect is to thoroughlydocument the operation of the specific instrument through measurement ofstandards until the behavior is sufficiently documented to allow reliable analysis ofunknowns.

The 25Mg/24Mg ratio is most commonly used to internally normalize isotopicmass fractionation since the main application of the Mg isotopic system is in thesearch for 26Mg excesses due to 26Al decay in meteoritic materials. Ion probe datahave most commonly been normalized with a linear law since the largest effects in26Mg are in meteoritic inclusions with high Al/Mg and therefore the data areinsensitive to the scheme used.

In order to document the veracity their analyses, Huneke et al. (1983) analyzeda series of glasses that had been gravimetrically doped with 25Mg. These glasses weresubsequently analyzed by McKeegan et al. (1985) and the data from bothlaboratories is in excellent agreement with the gravimetrically calculated ratios.These laboratories also measured Mg isotopic fractionations in standard anorthitegrains and, despite using very similar machines (CAMECA ims-3f) the range in massfractionation was quite different, –6 to –17 ‰/amu for Huneke et a l . (1983)


and -4 to -10 ‰/amu for McKeegan et al. (1985). Ireland et al. (1986) documentedwell-resolved matrix dependent mass fractionation with Mg from spinel being some4 ‰/amu lighter than hibonite, olivine, pyroxene, and kaersutite which all hadfractionation values within ≈1 ‰/amu of each other. The documentation of theinstrumental mass fractionation is therefore an important parameter to be obtained ifreliable high-precision estimates of intrinsic mass fractionation are to be made.

The CAMECA ims-3f has been used extensively for Mg-isotopic analysis andthe experimental procedures have been described by Huneke et al. (1983) andFahey 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 touse high mass resolution (≈3500!R ) to resolve all molecular interferences includingdoubly charged Ca and Ti, as well as hydrides. Measurements are made bycyclically peak-hopping through the Mg isotopes, and 27Al if desired. The problemwith measuring the 27Al abundance is that for high Al/Mg phases the Al+ signal canoverload the electron multiplier. There are several options in this case. Thesecondary 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 inpoorer precision per unit time of measurement. A Faraday cup can be employed tomeasure the 27Al+ intensity; in this case the high count rates of the Mg isotopes canbe preserved, but the Faraday cup must be switched in and out of the beam linereliably and the electrometer must be calibrated accurately. Alternatively, the27Al+/24Mg+ measurement can be made separately to the Mg isotopic measurement;however in this case information regarding correlated effects in 26Mg/24Mg with27Al/24Mg may be compromised.

Ireland et al. (1986) found that for minerals with more than 1-2 % MgO, therewas sufficient Mg+ signal to use the Faraday cup for the Mg isotopic measurementsas well. However, while there is sufficient intensity for the measurements, themethod for automatic peak centering was susceptible to an interaction between thepre-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 onthe absolute signal strength of 24Mg+; the lower the signal strength the larger theoffset. 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 problemsfound in this study, measurements using Faraday cups should be made wherepossible because of the benefits of no dead time correction. It is likely that suchapplications might become more common when the large high-sensitivitymicroprobes are analyzing major elements in a particular mineral phase.

Mg isotopic compositions have been reported from a large number of other ionmicroprobe laboratories with different instruments particularly in the verification ofthe excess 26Mg correlated with 26Al in Allende refractory inclusions. However, thedifferent practitioners used quite a variety of techniques. Bradley et al. (1978) usedtwo different ARL (Applied Research Laboratory) ion microprobes to analyzeanorthites from two Allende inclusions. The ARL ion microprobes were operated


at low mass resolution necessitating corrections for the presence of 48Ca2+ of up to5.4 %, and for scattering of 27Al+ and 23Na+. Despite these difficulties, Bradley et al.(1978) were able to show that the Allende inclusions had 26Mg excesses of up to 50% at 27Al/24Mg ratios of 1000 corresponding to an initial (26Al/27Al)0 abundance of 5.5¥ 10-5 which is consistent with the original analyses of Lee et al. (1977). Shimizu etal. (1978) used an energy filtering technique on a CAMECA ims-300 to measure Mgisotopes in Leoville refractory inclusions and found that their data were also largelyconsistent with the 5 ¥ 10-5 initial abundance of 26Al, but hibonite in one Leovilleinclusion had a large excess of 26Mg (160 ‰) at a low 27Al/24Mg of 22 correspondingto a 26Al/27Al of around 10-3. However, melilite from this same inclusion does notshow the extreme (26Al/27Al)0, but is consistent with the canonical value (Lorin et al.,1977). Macdougall and Phinney (1979) also used the ARL ion probe to show thatexcesses in 26Mg were present in CM meteorite hibonites. However, not only didthey find 26Mg excesses, but they also found large variations in the 25Mg/24Mg ratioof up to 10 % which they interpreted as due to heterogeneities in the intrinsic massfractionation.

Silicon. Silicon has a very similar isotopic system compared to Mg with threestable isotopes 28Si, 29Si, and 30Si with terrestrial ratios 29Si/28Si of 0.050633 and30Si/28Si of 0.0336214 (Barnes et al., 1975). Again the main isobaric interferences tobe eliminated are hydrides which are resolved at the same level (≈3500 R). Therehave been far fewer studies of Si isotopic compositions by ion microprobe, probablybecause of the lack of a suitable radiogenic precursor and therefore anomalies (interms of the meteoritic work) would be limited to mass fractionation and nuclearanomalies, the effects of which have been found to be generally small in mostsamples. However, the discovery of presolar SiC grains has shown that large Siisotopic variations do exist.

McKeegan et al. (1985) used an O– beam to sputter Si+ secondary ions andfound that the Si was significantly mass fractionated by -31 to -37 ‰/amu. After alinear correction for mass fractionation based on the 29Si/28Si ratio, McKeegan et al.(1985) found that there mean 30Si/28Si was 2.8 ‰ lower than the Barnes et al. (1975)value. This is not due to the choice of mass fractionation law since the exponentiallaw was found to drive the corrected 30Si/28Si value even further below the Barnes etal. (1975) value. McKeegan et al. (1985) therefore chose to normalize to their datato a 30Si/28Si value of 0.03357.

Zinner et al. (1987) used Cs+ to sputter Si– secondary ions so that C and Siisotopic analyses could be made on the same grains under the same analyticalconditions. They found that the yield of Si– (Cs+ primary beam) per unit mass ofsample was less than for Si+ (O– primary beam) but the ion yield per unit current ofprimary beam is substantially higher. The intrinsic mass fractionation of Si– was farless than that observed for Si+ increasing the precision achievable by the externalnormalization procedure. Since the isotopic anomalies in the SiC grains have turnedout to be so large, an external mass fractionation correction is adequate and data


are reported as deviations from the terrestrial ratios, d29Si and d30Si. Zinner et al.(1987) also examined the normal 30Si/28Si for Si– after normalizing with the 29Si/28Siand found that the data were best described by an exponential mass fractionation lawwith 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 ofapproximately 95, 0.75, 4.2, and 0.02 % respectively. In terms of terrestrialfractionation processes, it is customary to only measure the 34S/32S ratio and referenceit to the Canyon Diablo standard ratio of 0.0450045 (Thode et al., 1961). Ionmicroprobe 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– secondariesin galena. No charge compensation is necessary because galena is an electricallyconductive phase, however, if this technique is to be applied to any other sulfidessome 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 beenused by Chaussidon et al. (1989) on a CAMECA ims-3f. A mass resolution of≈2000 R is sufficient to eliminate the S hydride interferences and 16O2+ but doublycharged 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 isprobably safe to use the lower mass resolution. Secondary ion intensity is a strongfunction 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 thedifferent fractionation factors between the minerals were largely a function of therelative bond strengths with the weak S-metal bonds producing less fractionationthan the stronger S-O bonds. The precisions obtainable ranged from 1 ‰/amu forgalena 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 thegreatest sensitivity is generally obtained with primary and secondary ions ofopposite sign, and because the power supplies need to be banked to achieve therequired potential gradients for primary and secondary ions of the same sign.Macfarlane and Shimizu (1991) operated in this mode to ensure exclusion of Zninterferences 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 theinterference to ≈0.017 %. This analytical procedure usually yielded in-runprecisions of ≈0.6 ‰. Fractionation is also a function of energy offset; thevariations in fractionation measured at 0 eV were substantially reduced at higherenergy offsets although long term variations were still observed. At an energy offsetof 30 eV, the instrumental fractionation ranged from -29 ‰/amu for galena, to -47and -48 ‰/amu for troilite and pyrite.


Calcium and Titanium. Calcium and titanium have six and five stable isotopesrespectively, namely 40Ca , 42Ca, 43Ca, 44Ca, 46Ca, 48Ca, and 46Ti, 47Ti, 48Ti, 49Ti, 50Ti. Theterrestrial 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, and48Ca/44Ca = 0.088727 (Niederer and Papanastassiou, 1984). For Ti, compositions aregenerally 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 titaniumhave two common isobars at masses 46 and 48. The mass resolution required toseparate 46Ca from 46Ti is ≈43,000 R and so is essentially unresolveable with thepresent techniques. However, 48Ca and 48Ti can be separated at a mass resolution of10,500 R. Titanium also has isobaric interferences from 50Cr and 50V which requiremass resolutions of 21,000 R and 42,000 R respectively for full separation. Theseisobaric interferences can be monitored and the contributions at mass 50 strippedthrough 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 analyzingmeteoritic inclusions for isotopic anomalies. The largest anomalies have been foundin the mineral hibonite (CaAl12O19, with Ti and Mg substitution for Al), therefore theanalytical 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 through50) can be measured. Contributions from 46Ca, 50V, and 50Cr are stripped from theirrespective Ti isobars. Mass 43.5 is checked for the presence of 87Sr2+ which mightindicate the presence of Sr2+ interferences at masses 42, 43, and 44. The possiblepresence 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+, and96Zr2+ 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 ishigher 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, butthe 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 edgeof the 40Ca+ peak is measured at a distance Dm40 = 40/48 ¥ 4.58 ¥ 10-3 from thecenter of 40Ca+ which gives the fractional proportion of the tail. This techniqueworks as long as 48Ca+/48Ti+ is larger than approximately 0.01, for smaller ratiostailing from the 48Ti+ peak becomes unacceptably large. On the other hand, if the48Ca+ signal is comparable in intensity to or larger than the 48Ti+ signal, as in the caseof the HAL-type hibonites analyzed by Ireland et al. (1992), the peaks must be wellresolved 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, forexample Ireland (1990) used mean values of D40Ca = –0.6 ‰/amu and D 46Ti =


–14.5 ‰/amu to correct the meteoritic data. Fahey et al. (1987b) found that theexponential mass fractionation law provided the best fit to the Ti isotopic data andhas also been used for the Ca isotopes. However, Ireland et al. (1992) used aRayleigh law for the HAL-type hibonites since it was argued that the massfractionation of these hibonites was intrinsic and due to distillation, and theinstrumental 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 Hfisotopes is through the radiogenic decay of 176Lu to 176Hf (t1/2 = 3.7 ¥ 1010 yrs); Hfisotopic compositions can therefore be used as a chronometer or a radiogenicisotope tracer.

Kinny et al. (1991) measured Hf isotopic compositions in zircons in order tocompare the Hf isotopic systematics with U-Pb systematics. Hf is very similargeochemically to Zr and so is present at a reasonably high level in zircon (typically1-2 wt% HfO2). The techniques used by Kinny et al. (1991) illustrate some of thedifficulties encountered when analyzing isotopic compositions in the REE region ofthe periodic table. The total range in 176Hf/177Hf over geologic time is about 1 % andtherefore high precision is required to obtain useful geologic resolution.

Hafnium has atomic isobaric interferences from 176Yb and 176Lu (requiring massresolutions of 35,000 and 90,000 R for separation) and monoxide interferencesfrom the middle REEs which require ≈8,000 R. However, Kinny et al. (1991) notedthat the HfO+/Hf+ is about 4, the YbO+/Yb+ is about 0.5, and the REEO2+/REEO+ isnegligible and therefore more Hf signal is available from the oxide species with anadditional advantage of reduced isobaric interferences. Major molecularinterferences 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 beaddressed. An iterative correction procedure was applied to the data that includedstripping the 176YbO and 176LuO interferences from the 176HfO peak, removing thecontributions from 17O and 18O, determining the isotopic mass fractionation basedon the 178Hf/180Hf ratio, and correcting all peaks for hydroxide interferences asmeasured at mass 197 (180Hf 16OH+). The most important interferences for thedetermination 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 wereselected and a liquid nitrogen cold trap was used to inhibit the formation ofhydroxide species. Kinny et al. (1991) found that the REEOH/REEO ratios werenot identical to the HfOH/HfO but were related by a constant factor (2.5) derived byrepeated measurements on a suite of cogenetic zircons with different REEconcentrations.

A standard zircon was analyzed for its REE and Hf concentrations so that acorrection for in situ decay of 176Lu could be made, and analyzed isotopically for aninternal isotopic standard. The Sri Lanka 7 zircon standard has a (176Hf/177Hf)init of0.28183 ± 15 or eHf of -30 ± 1 as measured by thermal ionization. The ion probe


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

Uranium-Thorium-Lead. The U-Th-series decay schemes are widely used ingeochronology. 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 yrsrespectively, while 232Th decays to 208Pb with a half life of 1.4 ¥ 1010 yrs. Lead hasone stable isotope, 204Pb, that is non-radiogenic. In order to produce concordiadiagrams, the standard method of portraying U-Pb data, both the Pb isotopiccomposition and the U/Pb ratio must be known. In ion microprobe analysis thisinvolves two types of determinations; measurement of Pb isotope ratios andmeasurement of an interelement ratio, U/Pb. The main application of the U-Th-Pbmeasurements 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 aswell.

In zircon analysis, the mass spectrum around the Pb peaks has a variety of Zrand Hf oxides and silicides, as well as REE molecules. The molecules can becompletely resolved at around 6500 R and only hydrides are not resolved. Themeasurement of isotopic ratios of Pb at trace levels and at high mass resolution isreliant on high sensitivity; the SHRIMP I ion microprobe operates with typicalsensitivities of around 5-10 c/s/ppm Pb/nA.

The techniques used for U-Th-Pb isotopic measurements on SHRIMP were firstdescribed by Compston et al. (1984). An O– primary beam, preferably massfiltered, is used to sputter Pb+ secondary ions. The measurement cycle includes thePb isotopes, 90Zr2

16O+ as a reference for concentrations, 238U+, 232Th16O+ and 238U16O+.An important aspect of this technique is the normalization of the Pb+/U+ ratios to astandard 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 ofisotopic mass fractionation and hydrides.

Compston et al. (1984) took the Pb isotopic ratios as measured. The ratios werenot corrected for isotopic mass fractionation or the presence of hydrides becausethese effects were probably small (in terms of the lunar zircons being analyzed with207Pb/206Pb ratios around 0.53) and also act in opposite directions, hydrides toincrease the measured 207Pb/206Pb (from the contribution of 206PbH+ under 207Pb+), andmass fractionation to decrease 207Pb/206Pb (since in general the lighter species ispreferentially ionized). However, for younger zircons these effects, especially thehydride contamination, should be taken into account. The combined effects ofhydride and mass fractionation can be examined by comparing the measured207Pb/206Pb ratio of the standard with the conventionally measured values. Acorrection can be applied to the unknowns on this basis assuming that the unknownsare affected in the same way as the standard. In general, such a


correction has been found to be small and a correction is only necessary after recentinsertion 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 areusually within error of the conventionally measured values. This indicates that theinstrumental fractionation of Pb from zircon is also low.

The ThO+/UO+ ratios as measured are linearly proportional to the 232Th/238U inzircon with a fixed discrimination as determined by concurrent 208Pb/206Pbmeasurements. For a closed system,

†

208Pb206Pb

Ê

Ë Á

ˆ

¯ ˜

rad

=232Th238U

el232t -1( )el238t -1( )

È

Î Í Í

˘

˚ ˙ ˙

where l232 and l238 denote the decay constants of 232Th and 238U (4.9475 ¥ 10-11 a-1

and 1.55125 ¥ 10-10 a-1 respectively). The discrimination constant is 1.11, i.e.232Th238U

= 1.11 ThO+

UO+

Ê

Ë Á

ˆ

¯ ˜

and the error on this factor is around 1 % based on the error of a linear correlationdetermined from measurements on two Sri Lanka zircons with different Th/U.

The Pb/U calibration is derived from the observation by Andersen andHinthorne (1972) that the Pb/U ratio should be related to the concurrent UO+/U+

measurement based on their local thermodynamic equilibrium model. Compston etal. (1984) emphasized that the use of this correlation in their normalizationprocedure is purely an empirical treatment and is not in itself used as support for thelocal thermodynamic equilibrium model. It should also be noted that the UO+/U+ inany analysis does not appear to be controllable but is a somewhat random functionof the day to day operation of the ion probe.

Ratios of 206Pb+/U+ are measured in the standard and unknown and the actualPb/U of the unknown is derived from the known Pb/U value for the standard with therelationship

Pb+ U +( )unk

Pb+ U +( )std

=Pb U( )unkPb U( )std

In Compston et al. (1984) a linear regression was found to be adequate to describethe relationship between 206Pb+/U+ and UO+/U+ which could be used to normalize themeasured Pb+/U+ to a single value of UO+/U+ i.e.

Pb+ U +( )std= 0.0764 x - 2. 77( )

where x is UO+/U+. Subsequently, Williams and Claesson (1987) found a degree ofcurvature in the standard calibration that could be described by a quadraticcalibration, such that,


Pb+ U +( )std= 0.0048x 2 + 0. 0265x - 0.0825

Most recently a power law fit, albeit very similar to the quadratic formulation, hasbeen 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 importantwhen the mean UO+/U+ of the standards and the unknowns differ. If the means arethe same then a poor functional fit has the main effect of increasing the error on thePb/U calibration based on the reproducibility of the standards. Figure 15 illustratesthe 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 onlyapparent at large deviations from the centroid.

Within a given day, the scatter of the data about the calibration curve is found tobe 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 sputteredion emission. All curves give a good functional relationship to the data and thedifferences are only apparent when large extrapolations are required. For most datasets however, the centroids of the standard and unknowns are similar and the exactform of the curve is not a critical parameter.


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) wasfound 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 ofSL13.

Approximate concentrations can be calculated using a similar formulationbetween Zr2O+/U+ and UO+/U+ and the unknowns are normalized to the 220 ppm Uconcentration of the SL13 standard. The absolute error on these concentrationmeasurements is around 20 % since this is the variation observed in individualfragments of SL13.

Despite having low initial Pb concentrations, a correction for this Pb has to beapplied 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 alarge effect, particularly with regards to the amount of radiogenic 207Pb in theanalysis. Common Pb can be estimated in three different ways. Since 204Pb isnonradiogenic, 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 besubtracted from the analysis. We can define the common Pb content as thefractional amount of initial 206Pb in the analysis, i.e.

¶ = 206Pbinit206Pbtot

then for the 204Pb/206Pb correction method,

¶ =

204Pb 206Pb( )tot204Pb 206Pb( )init

The isotopic composition of the common Pb can be estimated from coexistingcommon Pb-rich minerals such as feldspar, estimated from the common Pb growthcurves, or an average composition can be used, such as Broken Hill Pb, when thecommon Pb is largely surface related (and the contribution is small). This is themost reliable method for common Pb correction in old zircons, but for youngzircons the uncertainties in the small amount of common Pb become too large foran accurate correction, particularly for the low abundance 207Pb isotope.

The second estimate is based on the 208Pb/206Pb ratio. The radiogenic 208Pb/206Pbcan be estimated from the ThO+/U+ ratio for an assumed formation age. The actualformation age assumed has only a small effect on the correction since the common208Pb/206Pb ratio has changed only from 0.28 to 0.32 over geological time. Then thefraction of common 206Pb can be estimated with

f =

208Pb 206Pb( )tot- 208Pb 206Pb( )rad

208Pb 206Pb( )init- 208Pb 206Pb( )rad

For well-behaved populations of young zircons, this method can give very reliabledata however, it can be unreliable for grains that have lost Pb during geologicaldisturbances, or for grains in which Th and U have moved independently.


A third method for estimating common Pb is based on an assumption ofconcordance. This method is used mainly for young zircons when insufficientcounts of 204Pb are available for a precise common Pb correction. The 207Pbcorrection assumes concordance between the 206Pb/238U age and the 207Pb/206Pb ageand calculates a fraction of common Pb on that basis,

f =

207Pb 206Pb( )tot- 207Pb 206Pb( )rad

207Pb 206Pb( )init- 207Pb 206Pb( )rad

This method is useful for obtaining mean 206Pb/238U ages without propagating acorrelated error from uncertainty in the common Pb correction.

B. Quantitative Analysis

The secondary ion mass spectrum can also be used to quantify the abundancesof major and trace elements in a target; this type of application is generally referredto as quantitative analysis. Ion microprobe mass spectrometry has great benefits forelemental abundance measurements such as inherently low background levels andreasonably high ionization efficiencies for most rock-forming elements. However,the ionization processes responsible for the benefits of the ion microprobe alsoconstitute one of its major problems, namely the relationship between ionicintensities of different elements in the secondary ion spectrum and their actualconcentrations in the sample is not necessarily a predictable function. A great dealof effort has been put into formulating a general theory of ionization which wouldallow all elements to be normalized according to the ion to element ratios of only afew elements. For example, in the local thermodynamic equilibrium model ofAndersen and Hinthorne (1973) the relationship between two adjacent charge statesof a single element could be used to characterize an analysis and derive quantitativeanalyses of all other elements within a factor of two (see Figure 6). However, it isdifficult to justify the physical properties attributed to the sputtered region from thismodel and even a factor of two is basically not good enough for moderngeochemical analysis.

There are two main problems in elemental abundance measurements. The firstis the presence of complex isobaric interferences particularly around the region ofthe rare earth element (REE) spectrum, and the second is the matrix dependence ofthe ionization efficiency. As outlined above, isobaric interferences can bediscriminated against by either high mass resolution or energy filtering. For mostion microprobes of the CAMECA 3f-4f-5f series, the loss of intensity associated withhigh mass resolution restricts measurement precision and also calls for extremelygood magnet control, a situation that is exacerbated by the practical limitations ofcentering on very low intensity ion beams. Therefore, while the energy filteringtechnique drastically reduces the secondary ion signal from a given species, the massanalyzer can be operated at lower mass resolution and hence the magnet control


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

Energy-filtering techniqueThe energy filtering technique is used to discriminate against complex

molecular interferences by selecting a high-energy window where such interferencesare reduced by orders of magnitude relative to the atomic species (Figure 16). Thetechnique was proposed by Shimizu (1978) and was used to analyze a series ofmajor and trace elements, ranging in mass from 23Na to 59Co, in plagioclase fromdifferent 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 thefilled bars are the signal with the filtered secondary ion beam; from Zinner and Crozaz(1986).


Figure 17 Energy distributions for (a) 44Ca+with 1 eV window and 30 eV window and (b)mass 159 isobars 40Ca2

31P16O3+ and 159Tb+. Typical operating conditions for energyfiltering 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 speciescompared to the gradual decrease from the atomic ion species. Figure from Zinnerand Crozaz (1986).

of analyses using this technique to Pb isotopic compositions in galena, Mg isotopiccompositions in a refractory inclusion from the Leoville meteorite, and analyses ofmajor and trace elements (including REE) in various terrestrial minerals. Theyshowed that at high energies (150 eV offset) the isotopic compositions of a varietyof elements approached their terrestrial compositions indicating the virtualelimination of complex molecular interferences. In general they found linearcorrelations between ionic and atomic ratios however, they also showed that this isnot necessarily the case with a systematic variation of Ca+/Si+ as a function of Feconcentration in Ca-rich pyroxenes. Clearly it is necessary to not only analyzestandards of the same mineral composition as the unknowns, but also a range ofcompositions that brackets the unknown in terms of chemistry. For the REEmeasurements Shimizu et al. (1978) assumed constant ionization efficiencies for allthe REE and plotted the data as CI-normalized and La-normalized patterns. Theyrecognized the shortcomings of this method and proposed that further work wasnecessary to determine the relative ionization efficiencies of the different REE aswell as to analyze well-documented standards so as to obtain absoluteconcentrations.

Zinner and Crozaz (1986) used the energy filtering technique to develop amethod specifically for measuring REE concentrations but other elements could alsobe included in the same analysis. Their experimental procedures involvedmonitoring the energy offset during the analysis by measuring the 10 % lower


energy edge of the energy distribution and maintaining a 100 V offset from thatposition (Figure 17). In this way, differences in ion energy due to charge build-updo not affect the energy range of the ions being analyzed. Rather than analyzing asingle isotope for each REE, the determination is based on the measurement of themass spectrum between masses 133 and 191 and its deconvolution into element andelement oxide species. Counts from the 59 measured masses form anoverdetermined system for the unknown intensities of 33 M+ and MO+ species andthe solution is obtained by minimizing the least squares of the difference betweenthe 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 solutiongives a goodness-of-fit parameter (c2) of the match between the measured spectrumand the intensities calculated for each of the components. A low c2 indicates thatthere are no additional interferences and the data are consistent within countingstatistics; a high c2 indicates that either there are additional interferences to be takeninto account, or the errors on the analyses are larger than those derived fromcounting statistics. The latter might commonly be the case when samples withrelatively high count rates are deconvolved and then 1/√N is an underestimate of theerror (for example primary ion beam noise can not be disregarded and isotopicmass 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 inthe counts of different isotopes can substantially effect the resultant abundance ofYb. In order to better constrain the deconvolution, GdO is subtracted by assuminga GdO/Gd ratio measured in terrestrial standards. Similarly 159TbO interferes with175Lu and since Tb is monoisotopic and 97 % of Lu is 175Lu it is impossible toseparate the signal at mass 175 into TbO and Lu. Again a contribution based onTbO/Tb is removed from mass 175. While Gd has seven isotopes, 152Gd (0.2 %) and154Gd (2.2 %) are too small for accurate concentration assessments, 155Gd and 157Gdhave contributions from 139LaO (99.9%) and 141PrO (100%), and 156Gd, 158Gd, and160Gd cannot be easily deconvolved from CeO and NdO. In order to improve theprecision 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 patternin the least squares fit.

The relative ion signals of the individual elements, [REEi+], are normalized tothe ion signal of a major element in the sample (e.g. Ca+ or Si+) whoseconcentration [CaO] or [SiO2] can be measured by electron probe and multiplyingby sensitivity factors Fi derived by analyzing suitable standards. Therefore, theconcentration of each REE, [REEi], is given by

REEi[ ] = FiREEi

+

Ca+ CaO[ ]

for the case of normalization to the CaO concentration of the matrix.


Figure 18. Ion yields of REE relative to Ca+ from four different mineral matrices. Whilethe patterns are similar, there are significant differences in the absolute yields from thedifferent minerals. Data from Zinner and Crozaz (1986), Fahey et al. (1987c), andIreland 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 ionyields from different matrices follow a similar pattern although the absolutesensitivities are different (Figure 18). The uncertainty in an analysis can be dividedinto the precision (usually defined by counting statistics) and the accuracy (which isa function of the suitability of the sensitivity factors and the determination of themajor element abundance for the analyzed spot).

The use of the sensitivity factors Fi inherently assumes that the secondary ionintensity is linearly proportional to the concentration. Such a relationship has beendemonstrated by a number of experiments and is one of the advantages of thistechnique. However, such a relationship cannot be assumed since in some cases anon-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 concentrationsmay not be known to such a high level. In geochemical modeling, where theabsolute concentrations are important in constraining various parameters such asparent magma compositions, it is important to try and assess the external errorsapplicable to the analyses. However, in determining variations in REE concentrationson a spot by spot basis the relative abundance is the important parameter and theabsolute calibration will only produce a systematic shift on all concentrations.


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

It is interesting to note that the same sensitivity factors (relative to Si+) could beused on two quite different machines with different secondary-ion extractionsystems, viz. the CAMECA ims-3f at Washington University and SHRIMP at theANU. Ireland and Wlotzka (1992) presented REE patterns of a zircon measured onthese two instruments using the same sensitivity factors and very similar results wereobtained (Figure 19). Overall the energy filtering technique is a robust method forthe 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 theionic intensities to the atomic concentrations. In its simplest form an empiricalworking curve can simply show the relationship between measured ionic ratios andthe atomic ratio in a mineral phase for which several standards are available withdifferent concentrations. While the functional relationship must be related in someway to the physical mechanisms involved in sputtering, it cannot be stressed toogreatly that the working curves are empirical and their purpose is solely as acalibration tool. For the most basic calibration, one standard might be used and alinear function passing through the origin could be assumed. This is basically theformulation used by Zinner and Crozaz (1986) with the sensitivity factors; thesefactors simply apply a single factor for any given ratio. Maas et al. (1992) have alsoused a simple linear calibration with high mass-resolving power analyses to deriveREE abundances in Archean zircons by ratioing the counts obtained on an


isotope of a given element with the counts of the same peak obtained from astandard zircon. However work from other laboratories suggests that the linearcalibration is not always appropriate. Hinton (1990) notes that while Si+ is the mostobvious species to use for normalization, the relative intensity of Si+ is stronglyaffected by changes in analytical conditions. Furthermore, Botazzi et al. (1991)found that the Si-normalized intensity ratios had to be adjusted to take into accountrelative REE intensity enhancements due to compositional matrix effects betweenfelsic and mafic/ultramafic samples. The normalization was based on therelationship

REEiSii

= bREEcSic

l + a

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 matrixeffects should be more pronounced than in the case of the data from Zinner andCrozaz (1986), who used an energy window of 100 ± 30 eV.

Ray and Hart (1982) demonstrated good linear correlations (coefficients betterthan 0.98) for a number of elements including Na, Mg, Al, Ca, Ti, Sc, Ba, Rb, Zr, andFe sputtered from natural clinopyroxene and synthetic glass. While linearcorrelations were obtained for both clinopyroxene and glass for a given element, theslopes of the linear correlations were different with ions preferentially emitted fromthe crystal matrix. This is a matrix effect and is probably a function of the energyoffset as well since unpublished data from SHRIMP (T. Ireland, W. McDonough, R.Rudnick) indicate that a glass standard can provide suitable calibration forclinopyroxene and garnet at the relatively high energy offsets used in the SHRIMPanalyses. On the other hand, Shimizu et al. (1978) used a similar offset in theirclinopyroxene analyses and found a correlation in Ca+/Si+ with the Ca/Fe ratio of thesample.

It is clear that the analysis of well-calibrated standards is necessary in order toconstrain the correct formulation of the working curve. Only in this way canproblems associated with matrix effects be addressed. The exact formulation of thecurve is not important provided it can be demonstrated that it produces accurate andreproducible results for a wide variety of target compositions.

Other techniquesThe isolated specimen technique (Metson et al., 1984) is another form of the

energy filtering technique and relies on sample charging to achieve the energyselection of the secondary ions. However, the build up of charge on the samplesurface degrades the primary beam spot and the extraction conditions cannot becontrolled to the same degree as in the conventional energy filtering technique.

Ion microprobes have been used extensively in materials science research and inparticular in the semiconductor industry to measure concentrations of contaminantsand dopants. However in this field the preferred method of analysis is by depth-


Figure 20. (a) Depth profiles of an amorphous Si:H/Cr layer on crystalline silica. Theisotopic 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). Figuresadapted 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 issputtered 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 ofthe element under examination. A known dose of the element is implanted into thesample and hence the depth integrated signal from the ion probe over this implantationdepth is proportional to the concentration. Once the implanted layer is sputteredaway, the signal is proportional to the intrinsic concentration of the element in thesample. The known concentration of the dopant can then be used to normalize theintrinsic concentration of the sample. The main difficulty in this approach is that thesputtering probably does not occur in a layer-by-layer fashion but represents someintegrated mixture of layers because of heterogeneities in the primary beam andknock-on effects of surface atoms to deeper levels. Furthermore, the absoluteconcentrations of the dopants are probably known to no better than approximately25 % and so the accuracy of a given analysis can be no better than this. However,


while the ultimate precisions achievable by this technique may not be high, theconcentrations of contaminants and dopants can be extremely low and so even largeerrors can give useful results.

Useful reviews and details of the depth profiling technique are given by Wilsonet al. (1989) and Zinner (1980). In the earth sciences depth profiling has had morelimited application but is of particular use in measuring diffusion profiles, forexample Bancroft et al. (1987) assessed the stability of synthetic and natural titanitesthrough depth profiles after leaching the samples in pure D2O, Giletti and Yund(1984) studied diffusional oxygen exchange between quartz and water bydetermining 18O depth profiles, and Ryerson and McKeegan (1993) have measuredoxygen diffusion profiles in meteoritic minerals.

V.!APPLICATIONS

Successful applications of the ion microprobe are becoming increasingly abundantas are the variety of problems to which ion microprobes can be applied. It isbeyond the scope of this paper to present all applications of ion microprobes thatare currently in use. The applications that will be highlighted in this review are thosewhich concern applications in the earth sciences, particularly those incosmochemistry and geochemistry. The subdivision of relevant work into effectivesubsections is rather difficult. A natural subdivision occurs between terrestrial andextraterrestrial applications and chemical and isotopic work because the problemsare quite different in each field. In cosmochemical applications there can be largevariations in isotopic and chemical abundances, whereas in terrestrial samples theeffects are generally more subtle. For this reason alone, there have been far morepapers written concerning the analysis of extraterrestrial materials than terrestrial andhence the discussion is somewhat weighted in favor of these (perhaps) rather esotericstudies. One of the greatest advantages of the ion probe is that a large variety ofanalyses can be performed on the same sample. This has been especially notable inthe field of cosmochemistry where isotopic abundances and elemental abundancesare measured for grains less than 10 µm in diameter. For the extraterrestrial studies,rather than simply progressing through the periodic table examining the differentelements, it is more useful to structure the discussion around the samples. In thecase of terrestrial studies however, there are far fewer cases of analyzing differentsystems in the same sample and so for these studies the applications are morestructured around the techniques.

A. in Cosmochemistry

By far the greatest number of publications dealing with ion microprobe analysisof “geologic” materials has been concerned with formation and evolution ofextraterrestrial 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


chemical and isotopic data from some presolar materials. Furthermore, thesesamples are often small and the maximum sensitivity per unit weight of sample isrequired to enable the greatest number of complementary analyses to be made onindividual grains. For some of the recent research into interstellar grains, a singlegrain of 5 µm in diameter has been analyzed for Si, C, N, Mg, Ca, and Ti isotopiccompositions. This work would have been impossible without the recent advancesmade in ion microprobe mass spectrometry.

Introduction to Chemical and Isotopic SystematicsIt 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 heaviestnuclides (Burbidge et al., 1957). It is also clear that the isotopic abundances presentin the solar system cannot have been produced in a single nucleosynthetic site butinstead record a history from a variety of stellar sources. However, prior to 1973 thecanonical view of the solar nebula was of a hot turbulent body in which heating wasso pervasive as to cause widespread volatilization of all elements therebyhomogenizing any initial isotopic heterogeneities that were present. This view waschanged irrevocably when Clayton et al. (1973) measured correlated effects in the17O/16Oand 18O/16O ratios from refractory inclusions in carbonaceous chondrites thatwere 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 formingelements, the likelihood was that isotopic effects in other elements could be foundprovided you looked in the right place at the right element with sufficiently highanalytical precision. And so thermal ionization mass spectrometers were tuned up toyield the best possible precision and it was not long before anomalies were found inthe 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 excess26Mg was correlated with Al/Mg (Lee et al., 1977). Isotopic anomalies in Ti werethen found at around the 1 ‰ level in all inclusions that were analyzed (Heydeggeret al., 1979; Niederer et al., 1981; Niemeyer and Lugmair, 1981). The effects werepredominantly in the abundance of 50Ti indicating that the isotopic abundances ofthe Fe group elements could be affected by the admixture of a neutron-richnucleosynthetic product. This interpretation was supported by the discovery ofeffects 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 wasnot believed to be coincidental. The mineralogy of the Allende inclusions - primar-ily melilite, spinel, and Ti-rich pyroxene, is consistent with their formation ascondensates from a cooling gas of solar composition (Grossman, 1972). Thesequence basically follows the sequential condensation of the refractory oxides ofAl, Ca, and Ti followed by Mg and Si. The trace element compositions of theseobjects also indicate a high temperature origin with enrichments in the refractory


lithophile elements to around 20 ¥ chondrite abundances and fractionation of theREE according to their relative volatilities. In this regard, there are two main typesof 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 solarabundances. One particular type of pattern, that with depleted ultrarefractory REEhas been used to argue for a condensation origin of these inclusions since thepattern can only be produced in a condensate that has been separated from a morerefractory residue (Boynton, 1975). The correlation of the presence of isotopicanomalies and the refractory chemical compositions of the inclusions is suggestiveof their early formation before the initial isotopic heterogeneities of the solar systemwere homogenized through mixing and high temperature processing.

If this were the case, then larger anomalies might be expected in morerefractory types of inclusion consisting of corundum or hibonite. Corundum-bearing inclusions are rather rare but hibonite-bearing inclusions are quite commonconstituents of the CM2 class of meteorites. These meteorites are generally muchfiner grained than the Allende CV3 type and the refractory inclusions are also muchsmaller being typically less than 500 µm compared with up to 2 cm inclusions inAllende. With such small objects, it is difficult to carry out conventional massspectrometric analyses but they are ideally suited to ion microprobe analysis.

Refractory Inclusion. The first isotopic measurements to be performed on the ionmicroprobe were of Mg since it had already been demonstrated that large effectsfrom 26Al decay might be expected. Mg isotopic measurements were made on awide variety of machines following quite different experimental procedures (forfurther details see Methodology section). These analyses showed that while a largenumber of refractory inclusions do preserve the effects of 26Al decay at thecanonical level of (26Al/27Al)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 donot have excess 26Mg consistent with "live" 26Al, in particular hibonite-bearinginclusions (e.g. Macdougall and Phinney, 1979; Hutcheon et al., 1980; Bar-Matthews et al., 1982; Fahey et al., 1987b; Ireland, 1988; 1990). A simplechronological interpretation would place those inclusions without excess 26Mg ata formation time after those preserving 26Mg at the canonical level. However, thelack of effects in a lot of the hibonite grains was perplexing since the CM inclusionswere thought to be amongst the earliest condensates based on the presence ofhibonite, and occasionally corundum. This requires that our understanding of theformation of refractory inclusions in a simple thermal progression is at fault or that26Al was not homogeneously distributed throughout the solar nebula in space and/ortime.

The formation history for hibonite became even more puzzling with analyses ofCa and Ti. Ca is of course an essential element in hibonite at 8.5 wt % CaO while theTi concentration ranges from under 1 up to 9 wt % TiO2. Conventional isotopic


analyses of Allende inclusions had shown that small anomalies (generally excesses)were present in the abundances of 48Ca and 50Ti. The first ion-microprobe Tiisotopic analyses relied on peak stripping 48Ca away from 48Ti since the transmissionloss 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 largeas 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 anomaliesapparent in 50Ti, there is a high degree of likelihood that 48Ca is also anomalous andtherefore a correction cannot be made for the presence of 48Ca under 48Ti based onnormal isotopic ratios.

For this reason Ca and Ti isotopic measurements are best made at high massresolution (≈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 etal. (1983) and Hinton et al. (1987), Fahey et al. (1985) measured a 100 ‰ excessof 50Ti in two Murray hibonites and smaller excesses in two others, while Ireland etal. (1985) measured a range of compositions from 40 ‰ deficits to 10 ‰ excesses.Initially, the three laboratories reporting Ti isotopic measurements adopted differentnormalization schemes but as it became apparent that effects were also present in49Ti, a 46Ti/48Ti normalization scheme was adopted by all. For most meteoritichibonites, the largest effects are in 50Ti with smaller anomalies sometimes resolved in49Ti while d 47Ti is generally close to normal (Figure 21).

Zinner et al. (1986b) reported the Ca-isotopic compositions in a suite ofhibonites that had a range of d50Ti from -40 ‰ to +100 ‰. The Ca isotopiccompositions were normalized to the 40Ca/44Ca ratio and no anomalies were observedfor d42Ca and d43Ca but large effects were evident in the abundance of 48Ca withanomalies ranging from –46 to +56 ‰ (Figure 21). The sign of d48Ca is nearlyalways the same as that for d50Ti (Figure 22) indicating that the anomalies areassociated. The main exception to this correspondence is HAL which has a smalldeficit in 48Ca and a 15 ‰ excess of 50Ti (Fahey et al. 1987b). The veracity of theion probe Ca isotopic compositions has been demonstrated by a conventionalthermal ionization measurement of BB-5 which has a d48Ca of –56.1 ± 3.7 ‰ by ionmicroprobe (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 inclusionsincluding those analyzed by Fahey et al. (1985), Zinner et al. (1986b), and BB-5which was analyzed by Hinton et al. (1987). The O-isotopic compositions of thehibonite grains all showed enrichments in 16O relative to the normal O-isotopiccomposition with the 7 grains all lying within error of the Allende mixing linedefined by conventional analyses of Allende inclusions (Figure 23). The maximum16O enrichment was shown by BB-5 at ≈ 70 ‰, whereas the maximum shown by theconventional data from Allende inclusions is around 50 ‰. The interesting aspectof the O-isotopic anomalies is that there is no correlation with the Ca and Ti isotopiccompositions. Despite individual grains having a wide range in 48Ca and


Figure 21. The Ca and Ti isotopic compositions of Murchison hibonites BB-5 and13-13 show the largest 48Ca and 50Ti anomalies. BB-5 is depleted in 48Ca and 50Tiwhereas 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 of16O between 4 and 7%. This suggests that while the Ca and Ti have a commonnucleosynthetic origin, the O-isotopic anomalies are probably related to anothersource.

Besides isotopic measurements, trace-element compositions could also bemeasured from these grains. Neutron activation analyses had suggested that thepatterns were not very different from the larger Allende inclusions (Ekambaram etal., 1984a) although the INAA analyses often had quite large uncertainties becauseof the small size of the samples. The benefit of ion probe analysis is that all REEelements can be measured as well as other refractory lithophile elements on the samediscrete areas that have been analyzed for isotopic compositions. Analyses can alsobe made on a relatively short time scale (2 hours) so that a large data base ofisotopic 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 correlationsbetween isotopic systems of different elements as well as chemical andmorphological effects.

The correlation of Ca and Ti isotopic anomalies is almost certainly a nucleosyn-thetic effect, but other correlations between morphological, chemical, and isotopic


Figure 22. The 48Ca and 50Ti isotopic anomalies in hibonites are generally correlated inthat samples with enrichments (or deficits) in 48Ca are accompanied by enrichments (ordeficits) in 50Ti. For comparison, the largest anomalies measured by thermal ionizationmass spectrometry are represented by the FUN inclusions, EK-1-4-1, and C1 asindicated in the inset. Figure adapted from Ireland (1990).

characteristics have not been so readily forthcoming. The problem with thesystematics of the hibonite-bearing inclusions was that the correlations were notquantitative but rather indicated associations of features. Hutcheon et al. (1986)first recognized that inclusions consisting solely of hibonite rarely had 26Mg* at alevel of 5¥10-5¥ 27Al whereas hibonite associated with spinel or melilite generally didhave 26Mg* at the canonical level. Ireland (1988) noted that the hibonite platycrystal fragments (PLACs) were not only characterized by low levels of 26Mg*

(Figure 24), but were also the most common carriers of highly anomalous 50Ti. Onthe other hand, the spinel - hibonite inclusions (SHIBs) generally had small 50Tieffects 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 Euand Yb whereas the SHIBs often had patterns showing fractionations in theultrarefractory 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 allhad radiogenic 26Mg* unless they also had large 50Ti isotopic anomalies. Theseobservations clearly establish a link between the preservation of isotopic anomalies


Figure 23. Ion microprobe measurements O isotopic compositions of seven hibonitesfrom the Murchison and Murray carbonaceous chondrites. Despite large variations intheir Ti isotopic compositions, ranging from enrichments to depletions, all of thehibonites are enriched in 16O to a degree not represented in conventional analyses ofthe larger Allende inclusions. Also shown are conventional analyses of a Murchisonspinel 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 microprobestandards. Figure adapted from Fahey et al. (1987a).

and the physicochemical processes responsible for the formation and processing ofthe precursors of refractory inclusions. However, while such a link can bedemonstrated, converting these associations into a formation mechanism is not soreadily forthcoming. The nature of the carriers of the isotopically anomalousmaterial, and the timing and location of the high temperature event(s) responsiblefor the trace-element fractionations are still matters of conjecture.

However, a link with a discrete formation mechanism can be proposed for anotherset of refractory objects, the fractionated and unknown nuclear (FUN) inclusions(Wasserburg et al., 1977). In these inclusions, which are not markedly differentmineralogically or morphologically to other Allende refractory inclusions, isotopicanomalies in every element analyzed thus far are accompanied by mass dependentfractionation effects with enrichments in the heavy isotopes. This heavy isotopeenrichment 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 beenextensively studied conventionally and by ion microprobe for its isotopic andchemical characteristics (Hinton and Bischoff, 1984; Hinton et al., 1988; Fahey etal., 1987b). Three other hibonite inclusions are similar to HAL, viz. DH-H1 from


Figure 24. Mg isotopic compositions of two types of hibonite-bearing inclusions fromCM meteorites. PLACs are PLAty hibonite crystal fragments and are characterized byhigh 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. Figureadapted 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 butonly 7-971 had mass fractionated Mg (Ireland and Compston, 1987). This isprobably due to the almost total evaporation of Mg from these inclusions. While Caand Ti isotopic mass fractionations are not correlated, the Ti isotopic massfractionation is inversely correlated with Ti concentration in the four inclusions(Figure 26) reinforcing the conclusion that distillation has played a major role intheir formation.

Corundum-bearing inclusions are relatively rare but are important since theyare likely to be higher temperature objects than the hibonite-bearing inclusions and,by extension of the relationship between the large inclusions from Allende and thehibonite inclusions from Murchison, may have extremely large isotopic effects. The


Figure 25. REE patterns of a representative PLAC and SHIB. PLACs typically showdeficits in Eu, Yb and the less refractory of the trace elements (cf. Allende Group III in35) whereas SHIBs show deficits in the ultra refractory REE Gd-Er and Lu as well asSc, 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 excessof 16O, contains substantial amounts of corundum. Bar-Matthews et al. (1982) firstdescribed this inclusion and used the Chicago AEI IM-20 ion microprobe to analyzeits Mg isotopic composition. Despite the extreme Al/24Mg ratios in the corundum,the d26Mg was only slightly above normal corresponding to a maximum (26Al/27Al)0of 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 forsmall ultrarefractory phases (possibly composed of Zr oxide and hiboniterespectively).

A more systematic study of a large number of corundum grains from Murchisonwas carried out by Virag et al. (1991). Twenty-six corundum grains were identifiedfrom 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


Figure 26. The inverse correlation between Ti isotopic mass fractionation, FTi, and Ticoncentration in HAL-type hibonites is suggestive of mass loss by evaporation. Thetwo cases shown are (1) measured compositions and (2) TiO2 concentrations correctedfor total mass loss. Figure adapted from Ireland et al. (1992).

and Mg isotopic compositions and trace element abundances. Some corundumgrains have (26Al/27Al)0 of 5 ¥ 10-5 which yields 26Mg/24Mg ratios up to 56 ¥ solarbecause of the high Al/Mg ratios of corundum. The O-isotopic compositions areanomalous but are still similar to those of other refractory inclusions in that thepredominant effect is enrichment of 16O by up to 7 %. The corundum in thesegrains therefore appears to have a similar origin to the other phases identified inrefractory inclusions. However, one corundum grain from the Orgueil CI chondritehas a 26Mg*/27Al of 9 ¥ 10-4 far in excess of the canonical value of 5 ¥ 10-5 and thismight be the only oxide grain analyzed to date for which a presolar origin could beproposed (Huss et al., 1992). Nittler et al. (1993a) have also analyzed a possiblepresolar corundum grain from the Murchison meteorite which had the same26Mg*/27Al of 9 ¥ 10-4 as the Orgueil grain analyzed by Huss et al. (1992). Nittler etal. (1993a) also measured the oxygen isotopic composition and found it to behighly anomalous with a d17O in excess of 1000 ‰ and a d18O of -200 ‰, clearlyindicative of an interstellar origin.

Another component within refractory inclusions which may have exotic originsare Fremdlinge, complex aggregates of metal grains, sulfides, phosphates, oxides,and silicates rich in refractory siderophile elements such as Pt, Ir, Os, Re, and Ru.


Their origins are enigmatic since they contain highly reduced as well as highlyoxidized phases and both ultrarefractory- and volatile-rich components. Oneproposal was that they were presolar in origin and should therefore have largeisotopic anomalies, however Hutcheon et al. (1987) used a CAMECA ims-3f ionmicroprobe to show that there were no isotopic anomalies for Mg, Fe, Mo, Ru, and Wand therefore it was likely that the Fremdlinge did form within the early solarnebula.

Ion microprobe analysis also has a role to play in the analysis of the largerrefractory inclusions as well. These objects appear to have been subjected to morethan one thermal event and deciphering their histories is reliant on analyzingselected portions of inclusions. In particular the rims of these objects appear to havesome genetic association with the cores, for example Boynton and Wark (1987)measured REE in the rim of an Efremovka refractory inclusion by INAA andargued that the five fold enrichment of the REE in the rim was due to the flashheating of the core. On the other hand, Fahey et al. (1987c) used a CAMECA ims-3fto measure Mg isotopes and REE abundances in the core and rim, and concludedthat the rim of this inclusion could not have formed by flash heating but rather bycondensation from another reservoir which had higher (26Al/27Al)0 than thereservoir from which the core formed.

Interplanetary Dust Particles. Of the hundreds of tons of extraterrestrialmaterial that the Earth captures each day, only a small fraction reaches the ground ina macroscopic form that can be recovered and identified. A significant fraction ofthe material is in the 1 to 10 µm size range and are known as stratospheric orinterplanetary dust particles (IDPs). The importance of this class of material is that itmay represent material that is fundamentally different than that preserved in thelarger meteorite bodies that reach the earth. Specifically, at least a fraction of thismaterial may be derived from comets and therefore yield material that has escapedthe high temperature processes at the beginning of the solar system.

The collection of these particles is non-trivial since man-made material cancontaminate the samples. In order to prevent contamination, particles are collectedon “flags” which are mounted on high-flying air craft (>20 km) and deployed forthe high altitude sections of the flight. These particles are caught in a silicone-oilmedium and then are picked off the mount and washed in xylene. The particles areanalyzed by as many non-destructive techniques as possible before being allowedinto an ion microprobe for destructive analysis (McKeegan et al., 1985). Fragmentsare analyzed in the SEM and are classified as chondritic if all chondritic majorelements are observable by EDX analysis.

The mineral assemblages in these particles were thought to be different topreviously analyzed meteoritic materials, and so isotopic measurements would be ofgreat interest. The small size of these grains (often less than 10 µm) means that ionprobe analysis is the only means of analyzing a number of isotopic systems andfinding correlated isotopic effects. McKeegan et al. (1985) measured H and C (Cs+


primary beam, negative secondaries) and Mg and Si (O– primary beam, positivesecondaries) isotopic compositions from a suite of eight chondritic particles. TheD/H measurements of the individual grains showed that large isotopic anomalieswere commonly present with 5 out of the 8 grains having enrichments over 500 ‰with a maximum enrichment over 2000 ‰. However, different fragments of thesame inclusion showed markedly different compositions and so even the extremevalues measured were mixtures of even more exotic components. Three of thegrains with heavy hydrogen were also analyzed for C and in comparison to the largeD/H variations, the range in C isotopic compositions is more modest, ≈ 40 ‰. The Ccompositions of different fragments of the same inclusion appeared to be the samein contrast to the H isotopic heterogeneities. Three particles were also measured forMg and Si isotopic compositions but no anomalies beyond analytical errors werefound.

McKeegan et al. (1985) also used the ion microscope mode of the CAMECAims-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 fragmentpressed 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).


that the D enrichment was strongly correlated with 12C/16O, indicating the deuteriumcarrier was probably carbonaceous in nature.

While these are extreme enrichments in deuterium, large enrichments have beenmeasured conventionally from bulk samples of Renazzo and Semarkona, and ionprobe measurements of individual particles of these meteorites have extreme dDvalues up to 3700 ‰ (Hinton et al., 1983; Zinner et al., 1983; McKeegan andZinner, 1984). Therefore, despite the large deuterium enrichments in the IDPs, theyare not markedly more anomalous than some meteorite matrices.

McKeegan (1987) measured Mg and O isotopic compositions in five IDPs thathad refractory bulk compositions; two contained corundum, one contained hibonite,and a further two grains were composed of a spinel-hibonite-perovskite-meliliteassemblage. Four of the particles had 16O excesses similar to refractory inclusionsfrom carbonaceous chondrites, while a fifth, one of the corundum particles, hadnormal oxygen and was probably a terrestrial contaminant. All four extraterrestrialparticles had normal Mg isotopic compositions.

At this stage, while there are mineralogical and petrographic distinctionsbetween IDPs and other meteoritic classes of material, there appear to be no majorisotopic distinctions.

Interstellar Grains . While refractory inclusions in carbonaceous chondritescontain isotopically anomalous material, it is likely that they acquired their presentform during high temperature processes in the early solar nebula and are not,themselves, interstellar grains. Unfortunately, it is likely that the interstellarprecursors to the refractory inclusions are probably oxides and therefore they areswamped by an overbearing proportion of material that has been processed withinthe 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 withextreme excesses of 26Mg and 17O.

However, it turns out that nature has not been so unkind as to bury all interstellargrains. E. Anders and coworkers isolated to levels of increasing purity the carriersof the exotic noble gas components observed in bulk meteorites. These carriers werehighly 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-bearingphases. The first phase identified was Cd, which was mineralogically a form ofdiamond (Lewis et al., 1987). This was the carrier of a component called Xe-HL, sonamed because the Xe isotopic composition shows large enrichments in both theheavy and light isotopes of Xe. These diamonds are exceedingly small - 2 nm onaverage - 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 containingabundant SiC was produced (Bernatowicz et al., 1987). The SiC grains cover a largesize range from submicron to over 20 µm in diameter. In following Ne-E(L), againessentially monoisotopic 22Ne but in a low density fraction, interstellar graphite


was identified (Amari et al., 1990). The graphite occurs as 1-4 µm grains with arange of morphologies from euhedral platy grains to compact spherules andspherulitic aggregates. A fourth interstellar phase was subsequently identified as 7-21 nm inclusions in isotopically anomalous graphite spherules and were identified asTiC 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 thesolar nebula are oxides and silicates. Diamond, graphite and SiC are unstable insolar nebula conditions and are stable condensates only when C/O is greater than0.83 (Larimer and Bartholomay, 1979) whereas the solar ratio is only 0.42 (Andersand Grevesse, 1989). Therefore interstellar carbon phases are undiluted in the earlysolar nebula by locally produced material. While they are not chemically stable,these phases are highly resistant to physical, chemical, and thermal processes and sothey can survive through processes in the interstellar medium, the solar nebula, aswell as the chemical treatment for purifying them (Zinner et al., 1987). The originof 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 bestconstrained by their isotopic signatures.

These interstellar grains do not comprise a large fraction of meteorites; theirabundances range from 400 ppm for diamond, 6-9 ppm for SiC and <2 ppm forgraphite (Amari et al., 1990); TiC comprises only a few tens of ppm of the graphitespherules in which it was observed (Bernatowicz et al., 1991). These materials aretherefore very precious and the maximum information should be extracted fromthem. In this regard ion microprobe analysis has been an essential tool indetermining the isotopic systematics of these grains particularly whereheterogeneous populations are present.

The first isotopic measurements of Cd were carried out by stepped pyrolysis ofbulk samples. Somewhat surprisingly, the average d13C of –38 ‰ in interstellardiamond is within the terrestrial range (Swart et al., 1983) although the minimumd15N is anomalous at –330 ‰ (Lewis et al., 1983) and the N composition wasobserved 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 averagefor 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 farsmaller quantity of material, it is still a bulk analysis in terms of the size of thediamond grains and therefore it too will represent an average composition.

Virag et al. (1989) reported C and N isotopic compositions from six Cd samplesfrom the Allende (4) and Murray (2) carbonaceous chondrites, and dD from two ofthe Allende separates. The samples were loaded as 20 µm agglomerates onto goldfoil and all elements were measured as negative ions. The d13C of the four Allendesamples is the same within error (–32 to –36 ‰) and are also consistent with thebulk measurements previously reported. However, one of the Cd residues fromMurray, CN, appears to be substantially lighter at –40.2 ± 2.8 ‰ 2s. The source of


this deviation is unclear but hints at some form of C-isotopic heterogeneity betweenthe diamond samples. All residues are depleted in 15N but there are substantialvariations in d15N between the four Allende samples and the two from Murrayalthough not as great as the differences observed by stepped combustion. Theaverage of all six samples agrees well with the average d15N of –164 ‰ measured byLewis et al. (1983) from Allende Cd residue CC but the ion probe analyses fall farshort of the minimum d15N of –330 ‰. It is possible that individual diamond grainsare heterogeneous and the stepped combustion technique can separate the differentN components based on the thermal release characteristics. In this regard, the highH concentrations (10-40 at%) of the Cd samples may be important since this showsthat preponderance of surface bonds on the diamonds and the chances are quitehigh for contamination of the N and H isotopic compositions either in theinterstellar medium or in the beaker during the acid digestion.

Figure 28. C- and N- isotopic ratios of individual SiC grains from Murchison separateKJH (average size 4.6 µm). The thin lines dividing the plots into quadrants depict solarisotopic ratios. Note the extreme range of the isotopic compositions and the use of alog scale. Most grains fall into the left upper quadrant, grains in the right upperquadrant with 12C/13C > 150 are named grains Y (they are also distinguished byunusual Si-isotopic compositions), those in the right lower quadrant grains X (they alsodeviate in their Si-isotopic ratios from the bulk of the SiC grains). (Figure adapted fromHoppe et al., 1993).


SiC has probably been the most intriguing of the interstellar grains studied sofar. This is in part due to the relatively large size of these grains which enablesdifferent isotopic analyses to be made on individual grains. Correlations can then bedrawn between the different isotopic compositions which can then be examined inthe light of theoretical nucleosynthetic calculations. The first analyses of Si, C, andN isotopic compositions in SiC (Zinner et al., 1987) exceeded the largest anomaliesin these elements by factors of up to 50. While agglomerates of a fine-grain-sizefraction were relatively homogeneous, analyses of clusters of coarse (>2 µm) SiCscattered widely in their isotopic compositions. The Si in both samples was highlyanomalous lying off the terrestrial mass fractionation line indicating that thecompositions cannot be derived from terrestrial silicon by mass fractionation. Mostof the compositions lay on the 29Si enriched side of the terrestrial mass fractionationline. The C-N isotopic compositions were characterized by enrichments in 13C anddepletions in 15N.

While the majority of grains still occupy the heavy C, light N quadrant, there arerepresentatives now in all four quadrants (Figure 28). These compositions representdiverse nucleosynthetic conditions and it is highly unlikely that all grains were

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


Figure 30. C- and N- isotopic compositions of individual graphite grains from Murchisonseparate LFC (1-4 µm diameter). Only the compact grains and dense clusters showlarge C-isotopic anomalies but N is rather indeterminate because of N contaminationfrom nearby nitrogen-rich aggregates. The dominant nucleosynthetic processes ineach quadrant are indicated as is the composition of micro diamonds (Cd). Figureadapted from Amari et al., 1990).

produced in a single star. Zinner et al . (1989) argued that the Si isotopiccompositions would not be affected in the nucleosynthetic conditions responsiblefor the C and N compositions and so individual compositions represented differentstars. 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 30Sieffects that were consistent with two component mixing possibly within the samestar. However, it is clear from the diversity of compositions now measured that anumber 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 probablyreflect their origins in distinct nucleosynthetic environments. After ion probe analyses of around 100 grains Zinner et al. (1991b) noted thattwo grains had completely different characteristics to the rest. The majority haveheavy C and light N and heavy Si while the two extremely exotic grains (termedgrains X) have light C, heavy N, and light Si. After measuring nearly 700 SiCgrains, Amari et al.(1992a) discovered 3 further X grains. Another type of SiCidentified by Amari et al. (1992a) were grains Y which lie to the right of the mainSiC trend on a Si 3-isotope plot and also have light C.


Figure 31. Mg isotopic compositions in SiC and graphite often show large excesses of26Mg and yield 26Al/27Al ratios ranging up to 0.2, 4000 times the canonical maximumvalue of 5 ¥ 10-5 for refractory inclusions. (Figure adapted from Zinner et al., 1991a).

Graphite C isotopic compositions deviate from terrestrial even further than thoseof 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 ofgraphite, the rounded grains. While the C compositions are highly variable, Nisotopic compositions appear to be more uniform (at least in terms of a log ratioscale; Figure 30) and are slightly heavy (to a factor of two).

A number of other minor elements have also been analyzed for their isotopiccompositions in SiC and graphite grains. A startling find was that large amounts ofextinct 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 inSiC of type X (Figure 31), well beyond the canonical 5¥10-5 value for the early solarsystem found in refractory oxide inclusions. The Al abundance in these grainsappears to be correlated with the N abundance suggesting that aluminium nitridemay be the condensed phase, however no discrete particles of Al-nitride within thegraphite spherules have been observed as yet.

Titanium is also often present in sufficient concentrations to allow ion probeisotopic measurements. Ti concentrations of SiC grains analyzed by Ireland et al.(1991b) range from ≈30 ppm to several thousand ppm. While the Ti concentrations


Figure 32. Most SiC grains have V-shaped isotopic patterns with excesses of allisotopes 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 thedecay of 49V. (Figure adapted from Amari et al.,1992).

are low, so too are the concentrations of the elements which have isobariccontributions, Ca, V, and Cr, and so the analyses are generally limited by countingstatistics. 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-shapedpattern with enrichments in 46Ti and 50Ti of up to 300 ‰ (Figure 32). This pattern ischaracteristic of nucleosynthesis in an environment with slow addition of neutrons(s-process) whereby the isotopic abundances are controlled by the neutron capturecross sections. A different signature is present in grains X which have large 49Ti and44Ca excesses (Amari et al., 1992b) (Figure 32). These two isotopes have radiogenicprecursors in 44Ti and 49V which are both p-process (i.e. they lie to the proton-richside of the main stream of stable nuclides) and they have the longest half-lives (47years and 331 days respectively) of the p-process radionuclides in this region of thechart 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 thesegrains are too high for supernova production models.

The isotopic compositions of some heavier nuclides have also been measured byion microprobe. Zinner et al. (1991c) measured Ba and Nd isotopic compositions


in various fractions of SiC and showed the s-process nature of their compositions.These compositions have also been determined by conventional mass spectrometrywith 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 paramountimportance in understanding the nuclear reactions that are going on in their sourcestars. Gallino et al. (1990) have already attempted to model the measured isotopiccompositions of Kr and predict the isotopic compositions of a wide variety ofelements. Now high precision isotopic ratios are available for a number of elementsfrom an extra-solar source with which we can compare theoretical and measuredvalues.

Short-lived nuclides in the early solar system. Short-lived nuclides areimportant for models of the evolution of the solar system since they can act aschronometers for thermal events and heat sources for melting planetessimals. Inthis paper we are concerned with the role of ion probe analysis in the detection ofthese nuclides; a more general recent review of extinct radionuclides is given byPodosek and Swindle (1988). The prime advantage of the ion microprobe in thesestudies is being able to select a very small volume which has a very high parentdaughter ratio so that the maximum effects in the isotopic composition of thedaughter can be measured.

Excesses of a given radionuclide in a single phase of a refractory inclusion canonly give an indication that the isotope was alive, an alternative possibility is that thephase inherited the daughter during a later thermal event. However, the finding ofexcess 26Mg which is correlated with 27Al/24Mg in different phases gives the strongestpossible evidence that 26Al was alive in refractory inclusions. With the finding of live26Al in refractory inclusions, it was proposed that 26Al could be a suitable heat sourcefor melting planetessimals. With an abundance of 5¥10-5¥27Al, there would besufficient heat generated from 26Al decay (t1/2 = 0.7 Ma) to melt a small planetarybody .

However, 26Mg excesses at this level have only been found in refractoryinclusions from carbonaceous chondrites. The only “non-refractory” assemblagein which excess 26Mg has been found is a olivine-pyroxene clast containingplagioclase from the Semarkona ordinary chondrite (Hutcheon and Hutchison,1989). For this assemblage a (26Al/27Al)0 of 7.7 ± 2.1 ¥ 10-6 (Figure 33) is sufficientto produce incipient melting in well-insulated bodies of chondritic composition.However, excess 26Mg has never been found in a differentiated meteorite and the roleof 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 26Alabundance can potentially yield very high age resolution provided it can be shownthat it was uniformly distributed throughout the solar system. The current data fromrefractory inclusions indicates that 26Al was not homogeneously distributed and


Figure 33. Al-Mg diagram for coexisting olivine, pyroxene, and anorthite in Semarkonachondrule CC-1. The linear correlation between excess 26Mg and Al/Mg is consistentwith the in situ decay of 26Al with an initial 26Al/27Al of 7.7±2.1 ¥ 10-6. The dashed linerepresents the canonical 26Al/27Al of 5 ¥ 10-5 found in coarse-grained Allendeinclusions. (Figure adapted from Hutcheon and Hutchison, 1989).

therefore its use as a chronometer may be limited. However, there may be somepotential use of 26Al in determining the duration of events in the solar nebula byanalyzing related objects. In this case it may be possible to argue for a uniformdistribution of 26Al. Podosek et al. (1991) have used ion microprobe Al-Mgisotopic data and conventional Rb-Sr data in an attempt to see if the systems arecorrelated within a suite of coarse grained refractory inclusions. The inclusionsanalyzed ranged from relatively pristine samples of basically igneous droplets, tohighly altered inclusions whose textures appear to be metamorphic in origin.However, both Al-Mg and Rb-Sr systematics appear to have been affected in allinclusions but nonetheless the comparison of both isotopic systems did not revealany chronological inconsistencies in terms of a heterogeneous distribution of 26Al.

Excesses of 53Cr due to the decay of 53Mn (t1/2 = 3.7 Ma) have also been found inchondritic meteorites. The analysis of Mn-bearing sulfides in enstatite chondrites byion microprobe has been particularly fruitful in the search for excess 53Mn since Cris incompatible in the sulfide phases and 55Mn/52Cr ratios of up to 4 ¥ 106 have beenmeasured in sphalerite from EL3 chondrites (El Goresy et al., 1992) (Figure 34).The 53Mn/55Mn ratios are quite variable within individual meteorites as well asbetween meteorites ranging from < 10-7 to near 10- 5 (Zinner et al., 1991d; ElGoresy et al., 1992) compared with the abundance in refractory inclusions ofaround 4 ¥ 10-5 (Birck and Allègre, 1985; Birck and Allègre, 1988). Hutcheon andOlsen (1991) measured excess 53Cr in differentiated meteorites including ironmeteorites and determined 53Mn/55Mn ratios of up to 2 ¥ 10-5, which is close to the


Figure 34. 53Mn - 53Cr systematics of Mn-bearing sulfides from MAC88136. The highMn/Cr in these sulfides (up to 4 ¥ 106) in these phases makes them ideal candidates inthe search for 53Cr excesses due to the decay of 53Mn. While excesses of 53Mn areclearly detected, there appears to have been remobilization of Cr after 53Mn decaysince 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 fromEl Goresy et al. (1992).

values from Allende refractory inclusions, down to 8 ¥ 10-7. 53Mn was clearly alivein a wide variety of early solar system objects but its use as a chronometer is stillbeing evaluated.

Attempts have also been made to find 60Ni excesses from the decay of 60Fe (t1/2= 1.5 Ma) by ion microprobe. Hyman et a l . (1988) measured Ni isotopiccompositions in individual Orgueil magnetite grains that had 56Fe+/62Ni+ ratios of upto 1.4 ¥ 105 but could find no excesses attributable to 60Fe decay; the upper limit for60Fe/56Fe was ≈10-4. Evidence for live 60Fe at a level of 1.6 ¥ 10-6 ¥ 56Fe has beenfound recently in the Chervony Kut eucrite by Shukolyukov and Lugmair (1992)who used conventional thermal ionization mass spectrometry. An important aspectof this work is that the 60Fe is alive in a differentiated meteorite and at this abundancelevel, 60Fe could be a very important heat source in planetary differentiation.

Hutcheon et al. (1984) measured the K-isotopic composition in two Allenderefractory inclusions in an attempt to find evidence for live 41Ca (t1/2 = 0.1 Ma).This is a difficult measurement because of the presence of large isobaric interferences


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 182Wwhich would be indicative of the presence of 182Hf (t1/2 = 9 Ma) in the early solarsystem. This radionuclide is a key indicator to the last contribution of r-processmaterial to the solar system. However, the measurement of W isotopic compositionsin zircons is difficult because of the presence of molecular interferences from theREE elements. The limiting interference for these measurements appears to havebeen 180HfH2+ and therefore the measurements of the 182W signal represent an upperlimit to its abundance. No effects were resolved relative to terrestrial standards intwo zircons with 180Hf+/184W+ of around 8 ¥ 104 yielding an upper limit of around5¥10-5 for the 182Hf/180Hf. This ratio corresponds to a free-decay interval of no lessthan 120 Ma for a production ratio of 182Hf/180Hf ≈ 0.5 which is consistent withestimates based on the 244Pu abundance.

Chemistry of Solar System MaterialsThe chemical systematics of early solar system materials are affected by a

number of high temperature processes but the effects can be viewed as being relatedto either volatility-related (solid-gas or liquid-gas fractionation) or igneous (solid-melt fractionation) mechanisms. Volatility fractionations, particularly in therefractory elements, are particularly well preserved in refractory inclusions ofcarbonaceous chondrites while igneous fractionations can be found in a large rangeof meteorites and meteoritic components. The trace-element signatures tell us agreat deal about the processes that were responsible for the formation of theseobjects.

Volatility Fractionations in Refractory Inclusions. The major element chemistryof refractory inclusions is dominated by the refractory oxides of Mg, Al, Si, Ca, andTi, but it is in the refractory trace elements that the most diagnostic signatures ofhigh temperature processes are found. The REE patterns of terrestrial minerals arelargely orderly functions of ionic radius but in refractory inclusions large variationsare 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 toCI-chondrite abundances and are characterized by the relative abundances of Euand Yb. Variations in Eu abundances in terrestrial materials are usually caused bythe tendency of Eu to exist in the divalent oxidation state, but the operativeparameter under solar nebula conditions is that Eu and Yb are the most volatile ofthe REE. Group V inclusions show no anomalies, Group I has a positive Euanomaly, Group VI has positive Eu and Yb anomalies and Group III has negative Euand Yb anomalies (Figure 35a). Group II patterns have completely differentcharacteristics. The patterns of this group have depletions in the most volatileelements (Eu and Yb) as well as large depletions in the most refractory elements


Figure 35. Allende refractory inclusion classification based on REE abundancepatterns. 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 theabundances of the relatively volatile REE Eu and Yb being the distinguishingcharacteristics. Group II shows a depletion in the most refractory elements, Gd-Er andLu, as well as the most volatile, Eu and Yb and must be obtained from the fractionalcondensation of a gas from which an ultra refractory component was removed. TheGroup IV pattern is obtained from olivine-rich chondrules.

(Gd, Tb, Dy, Ho, Er, Lu), so that the characteristic features are a flat light REEpattern and a Tm anomaly (Figure 35b). This pattern was attributed to condensationof the inclusion after the removal of a small amount of the most refractory materialfrom the gas phase (Boynton, 1975).

The ultrarefractory pattern complementary to the Group II pattern has neverbeen found in Allende. It was first identified in inclusion MH-115 which was one ofthe first inclusions analyzed from the Murchison carbonaceous chondrite (Boyntonet al., 1980). This inclusion contained hibonite and so it was not surprising that amore refractory inclusion composed of hibonite should have a more refractory traceelement pattern. However, most Murchison inclusions have patterns that are similar


Figure 36. Ion microprobe analysis of smaller CM meteorite inclusions revealed a largerange of patterns, three of which are shown here. (a) The ultra refractory-enrichedpattern is complementary to the Allende Group II pattern. (b) This pattern shows theeffects of a high local oxygen fugacity in which Ce and Pr become as volatile as Euand Yb. (c) This pattern shows deficits of Ce and Yb but Eu is not anomalous. (Datafrom Ireland et al.,1988; Ireland, 1990).

to those observed in Allende inclusions. Initially, a lot of painstaking workanalyzing small amounts of refractory inclusions were carried out by neutronactivation analyses, for example Ekambaram et al. (1984b). However, ionmicroprobe analysis of these inclusions has now become rather routine and has thebenefit of allowing analyses of very small inclusions which can also be analyzed fortheir isotopic abundances.

Around one hundred inclusions have now been analyzed for their trace-elementcharacteristics by ion microprobe. Most of these have been hibonite-bearinginclusions from Murchison but inclusions from other meteorites have also beenstudied (Fegley and Ireland, 1991). The vast majority of refractory inclusions fromMurchison have patterns that are similar to Groups II and III whereas the largerAllende inclusions are dominated by Groups I and V. The chemical characteristicsare quite strongly correlated with the morphological characteristics with the GroupIII patterns preserved in PLAC crystal fragments, while Group II patterns are


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 avariety 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 whichREE patterns had been determined and found that the Allende group classificationscheme was inadequate. They proposed an alternative classification scheme whichwas not reliant on adding groups or having modified patterns from the existingscheme. They noted that the patterns can basically be described in terms of threeprocesses: fractionation affecting the ultrarefractory elements Gd-Er and Lu,fractionation of the relatively volatile REE Ce, Eu, and Yb, and an overallfractionation that affects the slope but does not produce anomalies (igneousfractionation). From the tabulation of patterns, it was found that fractionatedpatterns (in terms of the ultrarefractory elements) comprised nearly half of allpatterns measured with most being depleted (37%) relative to enriched (7%).

As well as being invaluable in analyzing small inclusions, ion microprobeanalysis has also enabled the crystallization history of larger inclusions to beexamined. The large Type B Allende inclusions have the textural characteristics ofbeing produced by solidification of molten droplets. However, they have someunusual features such as thick melilite-rich rims, and spinel-free islands which aredifficult to rationalize in a traditional fractional crystallization model. However,Beckett et al. (1990) produced synthetically zoned melilite crystals and were able toshow that the distribution coefficients were a function of the composition of thecrystallizing melilite, for example Be is incompatible in gehlinitic melilites, butcompatible in åkermanitic melilites (Figure 37). Similar results were obtained byKuehner et al. (1989). The compositions measured from meteoritic melilite-richinclusions (Beckett et al., 1990; MacPherson et al., 1989) are broadly consistent withsequential crystallization from a single melt, but the spinel-free islands are probablytrapped xenoliths which were partially assimilated (MacPherson et al., 1989).

PLAC hibonites also have a roll-off in the heavy REE abundances which issuggestive of igneous fractionation in their formation, however the phase in whichthe heavy REE were accommodated has been lost (Ireland et al., 1988). The otherphase may have been a glass since this is observed in microspherules fromMurchison and ALH85085 but of the three Murchison hibonite-glassmicrospherules analyzed by Ireland et al. (1991a), only one shows the partitioningexpected between hibonite and silicate melt (Drake and Boynton, 1988). The mostunusual features are shown by spherule 7-753 for which both hibonite and glasshave ultrarefractory-depleted patterns and the differences between the compositionsappear to be related to the presence of a Gd-rich ultrarefractory phase as has beenobserved 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 fractionationeffects and while the heavy REE enriched phase has been lost from DH-H1, thefractionation of the REE in HAL has been successfully modeled by Hinton et al.


Figure 37. Distribution coefficients for Be melilite/liquid as a function of åkermanitecontent, XAk. Beryllium has the unusual property of being incompatible in melilite withXAk 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 inthe rim layers.

Enstatite chondrites, the reduced analog of chondritic meteorites, may alsopreserve primitive components. These meteorites formed under extremely reducingconditions such that elements which we normally regard as lithophile becomesiderophile. Calcium is an important example here since the major rare earthelement carrier in these meteorites is oldhamite, CaS. Oldhamite reacts withatmospheric water and so is difficult to handle but can be preserved in thin sectionand analyzed in situ by ion microprobe. Lundberg et al. (1989; 1991) documentedfour distinct REE patterns in oldhamite from the Qingzhen (EH3) meteorite (A-D inFigure 38) and found that these patterns were associated with oldhamite fromdifferent petrographic contexts. The fifth type, E, has ≈500 ¥ CI abundaces for theREE excluding Eu and has not been found in Qingzhen. Oldhamite is possibly areduced high-temperature condensate that is analogous to hibonite and perovskite inthe more-oxidized chondrites, and so these patterns may reflect nebular formationof 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 meteoritetypes. 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 commonproblem in meteorite studies since available material is precious and heterogeneity


Figure 38. The major REE carrier in enstatite chondrites is oldhamite, CaS, whichshows 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 raremeteorites, ion microprobe analysis can be used to analyze individual componentsand the bulk composition reconstructed from the modal abundances.

Floss et al. (1990) confirmed that oldhamite was the major REE carrier in theBishopville aubrite. These meteorites are relatively rare and are the achondriticequivalent to the enstatite chondrites. They are probably of igneous origin,although Floss et al. (1990) found that individual oldhamite inclusions had quitedifferent REE patterns suggesting oldhamite may be a relict phase even in thesedifferentiated meteorites. In this case bulk REE patterns would tell us little about theigneous processes involved in the formation of these rocks.

Lundberg et al. (1988; 1990) analyzed the REE carriers in two shergottitemeteorites to examine the implications for magma generation and evolution(purportedly) of Mars. Most of the REE reside in whitlockite whose modalconcentration is only ≈1 %. REE patterns of individual phases can be used to modelthe magma composition and the agreement of these model compositions forpyroxene, plagioclase, and whitlockite for ALHA77005 suggests that crystallizationtook place in a closed system. The mineral REE data are consistent with thepetrographic observations that plagioclase and whitlockite crystallized late andALHA77005 crystallized low-Ca pyroxene before high-Ca pyroxene, whereasnakhlites and chassignites have higher CaO sources and magmas derived from themcrystallized high-Ca pyroxene first.

Getting closer to home, but still in the extraterrestrial line, analyses of lunarmaterials by ion microprobe have been extremely useful because of their ability to


analyze small volumes. This is important, not only curatorally, but also because ofthe brecciated nature of the most primitive lunar rocks and the presence of glassmicrospherules on the lunar surface. These glasses contain much higher siderophileand volatile element abundances than mare basalts and appear to originate from asource region deeper than the devolatilized source of the mare basalts. Theprimitive nature of these glasses is therefore of great interest since therecharacteristics might better reflect the bulk chemistry of the Moon for theseelements and facilitate comparisons between elemental abundances in the Earth andMoon. Delano and coworkers [e.g. Delano (1986)] have established the presence of25 chemically distinct types of lunar glass and ion microprobe studies (e.g. Sheareret al., 1990) of the rare earth elements and several other refractory trace elementshave identified at least seven types and also indicate heterogeneous source regions atdepth in the Moon.

B. in Geochemistry

While terrestrial materials are far more abundant than meteorites and returned lunarsamples, 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 whichmeans conventional analyses can be made without significantly depleting the totalinventory of that sample. However, it is also related to the much smaller effects, bothchemical and isotopic, that need to be resolved in terrestrial work and hence highprecision is of greater importance. While the ion probe offers the best in terms ofsensitivity per unit volume, the volumes available to conventional analysis aregenerally not so restricted and the extreme precision necessary can be achieved.However, when the samples are small, or are shown to be complex and theassumption of sample homogeneity cannot be sustained, then the ion probe canagain become a useful tool.

Isotope GeochemistryApplications in the field of isotope geochemistry can be divided into two main

fields concerning stable and radiogenic isotopes. Stable isotope studies are based onthe 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 asreflectors of the formation environment. Radiogenic isotope systems can be used aschronometers or as tracers of petrogenetic processes. In this section, the discussionwill be concerned with radiogenic isotopes as tracers and ion microprobeapplications in geochronology will be discussed later.

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


precision of ≈10‰. For two of the samples, Massif Central, France and Salt LakeCrater, Hawaii, dDSMOW variations were measured both within single crystals andamong 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 andthe interval between exchange and volcanism was very short (less than a fewmonths). For the Lherz peridotite, France, and Nunivak Island, Alaska, the dD valuesare within the range shown by uncontaminated mantle minerals. However, theMassif Central xenoliths have a higher dD which may have involved contaminationof the lithosphere with subducted sea-water-altered oceanic crust. The Hawaiianperidotites have very light H which may be due to the influence of a deep mantleplume.

Chaussidon and Albarède (1990) measured B isotopic compositions intourmalines from a variety of magmatic and sedimentary rocks. Reproducibility ofthe measurements on tourmaline standards is cited as being better than ± 1 ‰. Thed11B values in the unknown tourmalines range from –30 ‰ to +17 ‰ with individualvariations in single tourmalines being less than 5 ‰ despite chemical zoning in somecrystals. There was no correlation of d11B with age of the source rocks, and the largevariations are related in part to the bulk chemistry of the tourmaline grains.

Carbon isotopes were measured in a heterogeneous Bultfontein diamond byWilding et al. (1990) with a Cs+ beam. Since diamond is electrically conductivesample 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 d13Cvalues of the diamond ranged from -5.8 to -10.9 ‰ and showed a consistentcorrelation with the cathodoluminescent zones indicating either a change in theprecipitation conditions or a change in the source C isotopic composition. Harteand Otter (1992) also measured C isotopes in diamonds and found that the K3diamond alone covered the entire range of diamonds measured by conventionaltechniques of -9 to -2 ‰. In this and four other diamonds there was no clearcorrelation with cathodoluminescent zones in contrast to the findings of Wilding etal. (1990).

Terrestrial oxygen-isotopic compositions, in terms of the fractionation of18O/16O, range from around -55 ‰ for Antarctic meteoric water to around +35 ‰in low temperature minerals. However, the range for lithospheric materials is onlyaround 35 ‰, which is less than 20 ‰/amu fractionation. Therefore, in order toassess even gross O-isotopic heterogeneities in terrestrial rocks, precisions of theorder of a few permil are required. So far this has only been achieved forconductive samples such as magnetites and ilmenites. Valley and Graham (1991)found that individual magnetite grains from the Adirondack Mountains, New York,were largely homogeneous with d18OSMOW of +8.9 ± 1.0 ‰, but analyses from closeto 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 thetemperature of regional metamorphism. From their data, Valley and Graham


(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 coolingthrough the blocking temperature of calcite and magnetite followed by exchangewith late low temperature fluids.

Sulfur has been the most analyzed of any of the stable isotope systems becauseof the relatively high isotopic ratio of 34S to 32S (≈0.045) and the large fractionationsthat are apparent in natural rocks. The total range in d34S in sedimentary sulfides isfrom –70‰ to +70 ‰ (Ohmoto and Rye, 1979) which is largely due to bacterialfractionation of S. The original approach for S isotopic compositions was tointerpret sulfides with d34S around 0 ‰ as magmatic and sulfides with variable d34Svalues (≥ ± 10 ‰) as sedimentary (Ohmoto and Rye, 1979). However, much morevariability is apparent in magmatic systems, due to, for example, kinetic effects, andsuch interpretations based on S isotopes alone cannot be definitive. S isotopes canalso be used for geothermometry, based on the S isotopic fractionation between twocoexisting S-bearing minerals.

Ion microprobe applications in S-isotope geochemistry have been particularlyimportant in ore studies where large variations in S-isotopic compositions can befound even within the same crystals. The ion microprobe allows measurements ofthese variations and can be used to pin point the petrographic carrier of theanomalous S.

Deloule et al. (1986) combined S- and Pb-isotopic measurements to study themicrostratigraphy of galena crystals from Mississippi Valley-type deposits andfound that there were frequent changes in the sources of the brines which circulatethrough 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 andEldridge, 1989) have analyzed a variety of sediment-hosted massive sulfide depositsincluding the McArthur River HYC and Mount Isa lead zinc deposits, Australia, theRammelsberg deposits, Germany, and the Salton Sea geothermal system, California.In all of these cases, more isotopic variation was found in a single thin section thanhad been reported from the entire bodies by conventional analysis. Of particularimportance was the analysis of sulfide from the H.Y.C. deposit McArthur R., sincethe minerals are noted for their fine grain size (<200 µm in diameter) and lack ofmetamorphic overprint. The relationships between the different sulfide minerals hadremained 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 aconventional analysis anyway. Both forms of pyrite are extremely heterogeneouswith 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 oresfrom the Hualgayoc district of northern Peru and found the galena from the veins


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 ofpyrite (py1, py2). Each of the pyrite types has a skewed isotopic distribution typical ofbiogenic pyrite formed in a system with limited sulfate supply and the means of py1and py2 differ by 15 ‰. Typical error for these analyses is ± 2 ‰. (Figure adaptedfrom Eldridge et al., 1993).

is about on average 5 ‰ lighter than galena from the stratabound assemblage. Theshift in d34S was interpreted as a result of oxidation of the ore fluid during the laterstages 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 beingfound in sulfide inclusions with low Ni content from pyroxenites, ocean islandbasalts, and kimberlites, whereas sulfides with high Ni content, mostly peridotitic, hada more limited range from -3.2 to +3.6 ‰.

Eldridge et al. (1991) have analyzed sulfide inclusions in diamonds and founda 25‰ range in d34S (from -11 to +14 ‰) which is the largest reported for mantlesamples. However the mean d34S is 0.4 ‰ and such heterogeneity could have goneunnoticed in a bulk analysis. The largest variations were preserved in inclusions withNi contents less than 8 wt% which is a proposed maximum for sulfides of eclogiticaffinity whereas the peridotitic inclusions show little deviation from 0 ‰. Theseresults suggest that the sulfides in the diamonds have sampled crustal sulfur possiblyin the form of recycled sediments.


Radiogenic Isotopes as Tracers. Lead isotopes are one of the most importanttracers of crustal evolution. The Pb isotopic ratios are affected by the ratio of238U/204Pb (µ) and 232Th/204Pb (k) in the source region and are also a function ofresidence time in a reservoir. Additionally, Pb isotopes can be used to fingerprintore bodies and the Pb-isotopic compositions of the various economic horizons canbe used to identify or at least restrict possible sources.

One of the classic studies in secondary ion mass spectrometry is themeasurement of differences in Pb-isotopic compositions in different growth zones ina galena from the Buick mine, southeast Missouri (Hart et al., 1981). Variations inthe Pb isotopes of up to 4 % in 207Pb/206Pb were found which was more than the totalrange previously reported from the whole southeast Missouri ore district. Deloule etal. (1986) also measured Pb isotopic compositions in the samples from theMississippi Valley-type ores they analyzed for S-isotopic compositions. Thevariability noted in the S was also reflected in the Pb compositions. For the Picherdeposit, the Pb and S appear to have been derived from the same source whereas inthe Buick mine different sources are required with S being extracted earlier than Pbin the hydrothermal circulation.

As well as analyzing S-isotopic compositions of sulfides in diamonds, Eldridgeet al. (1991) measured the Pb isotopic compositions of the same inclusions. The Pbdata are consistent with the S data in that the peridotitic samples have Pbcompositions 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 withmuch higher µ (for an average age of 1 Gyr, µ ≈ 300) further reinforcing thepossible contribution of crustal material in the diamond formation region.

The possibility of ion probe analysis of strontium isotopic compositions hasbeen examined by Exley (1983) who used a modified AEI-IM20 ion microprobe atlow mass resolution. Carbonates were analyzed because of their very lowconcentrations of 87Rb+ which cannot be resolved from 87Sr+. However, molecularinterferences predominantly from Ca2+ and CaMgO+ had to be peak stripped from86Sr+, 87Sr+ and 88Sr+ by monitoring 40Ca2+ and 40Ca25Mg16O+ respectively. Aprecision in 87Sr/86Sr of ≈1 ‰ could be obtained for Sr concentrations of greater than5000 ppm. Ion microprobe analyses were compared with conventional solid sourcedata for a suite of standards with good agreement between the two techniques, andtwo applications were presented regarding isotopic variation across a calcite vein in ahydrothermally altered basalt and the analysis of 35 µm calcite grains in a lherzolite.Exley and Jones (1983) examined Sr isotopic compositions in kimberliticcarbonates 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) forprotoliths. 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 % HfO2and the Hf model ages can be compared with U-Pb ages of the same areas. Kinny


Figure 40. Pb isotopic compositions of sulfide inclusions in diamonds. The inclusionsare classified according to their Ni content into eclogitic and peridotitic suites. Many ofthe peridotitic inclusions fall on the terrestrial growth curve indicating a formation age of2.0 Ga with average terrestrial µ. Some of the eclogitic inclusions lie off the curvetowards the origin indicating growth in a high µ environment such as the crust. This issupported by S-isotopic compositions for which only the eclogitic inclusions showsignificant variability. Figure adapted from Eldridge et al. (1991).

et al. (1991) found that the Hf model ages of Mt. Narryer zircons were consistentwith their 4.2 Ga U-Pb age, and concordant Hf and U-Pb ages were also obtainedfrom a Mt. Narryer anorthosite (3.73 Ga) and White Cloud gneiss (2.72 Ga), but forthe younger Pacoima Canyon pegmatite (1.17 Ga) and Jwaneng kimberlite zircons(0.24 Ga) eHf was significantly elevated suggesting incorporation of old Hf into thesource regions (Figure 41). On the other hand, a Sri Lanka zircon was highlydepleted (eHf of -23 at 570 Ma) indicating that the concordant U-Pb ages of thesezircons have been reset or that they were derived from a very unradiogenic source.These zircons probably grew during high-grade metamorphism of a late Archeanprotolith.

Kinny et al. (1991) also analyzed zircons from the Watersmeet tonalitic gneissfor which conventional analysis had suggested a correlation between U-Pbdiscordance and Hf model age. However, the ion probe data showed no suchcorrelation and all cores had the same Hf isotopic composition despite the almostcomplete loss of Pb from some grains. However, 2.7 Ga rims do have younger Hfmodel ages suggesting the addition of younger Hf during rim formation.

Trace-element GeochemistryThere have been a wide variety of applications in terrestrial trace-element

geochemistry. As well as the more typical applications involving the measurement


Figure 41. 176Hf evolution diagram showing results of ion microprobe zircon analyses(open symbols) with conventional thermal ionization data (solid symbols). MAINZ refersto the mean 176Hf/177Hf value of the Sri Lanka zircon analyzed by thermal ionizationdata. Most of the zircons are consistent with the trend of the conventional data butthe 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 beensome highly innovative techniques ranging from cracking fluid inclusions andanalyzing cation ratios (Diamond et al., 1990) and even CO2 concentrations (Pan etal., 1991), to analyzing REE concentrations in fish teeth and conodonts (Grandjeanand Albarède, 1989). In the following section some selected applications from therecent literature are described which cover some of the more intensive research areasnow being examined.

Diamond Inclusions. Besides sulfide inclusions for which S and Pb isotopic datahave been obtained, a number of silicate phases are present in diamonds. Of majorinterest is whether these silicates are related to the diamond formation region orwhether they represent more diverse provenances as is the case for the sulfides.

Shimizu and Richardson (1987) measured REE, Ti, Zr, and Sc concentrations inindividual crystals of sub-calcic peridotite suite garnets. All inclusions display ageneral enrichment of light REE relative to heavy, and the peridotite-suite garnetsare depleted in Ti compared to those from garnet lherzolites. These characteristicsare contrary to those predicted for garnet in association with olivine, orthopyroxene


and diamond, and in particular the partitioning between garnet and liquid shouldproduce a light REE depleted pattern. Such patterns have been measured fromeclogitic garnet inclusions from Monastery diamonds which have a majoritecomponent in solid solution (Moore et al., 1991).

Navon et al. (1988) analyzed microinclusions in diamonds from Zaire andBotswana and found that they differed in composition from the larger eclogitic andperidotitic inclusions. They measured the major rock-forming elements as well astrace elements by ion probe with each analysis representing the average of a largenumber of inclusions since the inclusions were smaller than the probe diameter. Theelemental concentrations were highly variable, probably due to the number densityof the inclusions, and the submicron grains were found to resemble potassic magmasin their compositions.

A different aspect of diamond genesis was examined by Phinney (1988). Large3He/4He ratios have been measured in diamonds and lie in the range observed only inmeteorites and this suggests that terrestrial diamonds might have a primordialcomponent. However, 3He can also be produced by (n,a) reactions on 6Li butestimates of the importance of this reaction were restricted because of the absence ofLi concentration data in diamonds. Phinney (1988) measured Li abundances in fiveterrestrial diamonds by implanting 7Li at a known dose and measuring theabundance by depth profiling. The highest concentration was 2.8 ppb with all theothers having concentrations under 1 ppb; the analytical uncertainty of themeasurements was around 25 %. In order to explain the 3He abundance, a Liconcentration of around 30 ppm is required and so it appears that 6Li(n,a)3He doesnot make a significant contribution to the 3He abundance.

Mantle xenolithsUltramafic xenoliths in basaltic magmas are the best samples available of

Earth’s mantle. They are composed predominantly of olivine, orthopyroxene, andclinopyroxene, (dunites, harzburgites, pyroxenites, etc) with spinel (spinellherzolites) or garnet (garnet lherzolites). The amount of available sample is smallbut is generally sufficient for conventional analyses. However, painstaking mineralseparation is required and minerals in these xenoliths can be heterogeneous on agrain by grain scale and could also be internally zoned or the minerals can belocally altered.

Shimizu and Allègre (1978) first analyzed garnet lherzolite nodules fromkimberlites 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 ofwhich could be close in composition to primitive mantle. Salters and Shimizu(1988) found that clinopyroxene from some harzburgites and peridotites aredepleted in the high field strength elements Ti, Zr, Hf, Nb, and Ta relative to theREE. They argued that the clinopyroxenes from spinel lherzolites were an adequaterepresentation of the bulk inclusion because of the similarity of the patterns inclinopyroxenes and the bulk inclusion from which they were derived. Thesedepletions were present in xenoliths from continental regions as well as oceanic


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 showcomplementary 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 uppermantle. However, Rampone et al. (1991) also used an ion microprobe to show thatorthopyroxene had complementary Ti and Zr anomalies to the clinopyroxene andthe recalculated bulk rocks show only negligible anomalies (Figure 42). Therefore,the complete inventory of any inclusion should be examined before the existence ofan anomaly in the bulk rock can be established.

Glass inclusions. One of the principal aims of petrology is to use themeasured compositions of coexisting minerals to reconstruct the parent magmafrom which they were derived. Potentially an equally useful approach may be toanalyze the small glass inclusions that are commonly present in magmatic rocks.These inclusions are typically small (<100 µm) and so ion microprobe analysis isprobably the most effective means of analysis. Glass inclusions can be found in awide variety of compositions from felsic to mafic.

Clinopyroxenes from abyssal peridotites have been found to have strongly lightREE 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.,


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

1990). These compositions demonstrate that the peridotites are the result of variabledegrees of fractional melting in the garnet and spinel peridotite fields, and thatbasalts might evolve from aggregation of these fractional melts. Support for thisinterpretation has recently been found in the form of melt inclusions in olivine fromtypical NMORB (Sobolev and Shimizu, 1991) (Figure 43). Surprisingly, the majorelement chemistry of the melt is consistent with only a low percentage of meltingsuggesting a very efficient mechanism for the fractionation of the incompatibleelements.

Trace elements in zircon. The discovery of 4.2 Ga zircons in Western Australiahas opened up a new era in Earth’s history. However, the zircons are preserved in amuch younger quartzite and appear to be the sole representatives of the earliestcrust; no rocks per se have been discovered that are the source of the 4.2 Gazircons. Therefore all information concerning the source rocks and formationconditions must be obtained from the zircons. In this regard the trace elementchemistry of the zircons may be the most important indicator of the nature of thesource rocks.

Hinton and Upton (1991) have analyzed large zircons from syenite and alkalibasalt xenoliths and describe the basic features of zircon chemistry: they are


systematically enriched in the heavy REE relative to the light REE, with a distinctivepositive Ce anomaly which reflects the more favorable incorporation of Ce4+

(relative to Ce3+) into the zircon structure. This feature appears to be ubiquitous interrestrial zircons but Ireland and Wlotzka (1992) found that neither meteoriticzircons nor a lunar zircon assemblage had Ce anomalies reflecting the reducednature of their parent bodies. However, Hinton and Meyer (1991) described anunusually oxidized lunar granite which does show the Ce anomaly. Snyder et al.(1993) also measured two zircons from lunar glasses and found, like Ireland andWlotzka (1992) that there was no Ce anomaly. It would appear that the Moon issignificantly more reduced the the Earth but that under certain conditions the ƒO2may rise to a level where Ce4+ is stabilized, particularly in late-stage residual liquidssuch 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, westernAustralia. The pre-4.0 Ga grains (a) show similar patterns to the 3.3 - 3.75 Ga grainssuggesting a similar petrogenetic environment for all these grains. (Figure adaptedfrom Maas et al., 1992).


Maas et al. (1992) studied a variety of Archean zircons as well as the 4.2 Gazircons from Mt. Narryer and Jack Hills. The old zircons could not be distinguishedfrom the younger on the basis of the REE patterns (Figure 44). All zircons arecharacterized by the HREE enrichment, positive Ce anomalies and smaller negativeEu anomalies. Since most of the younger zircons are thought to be derived from amature continental source dominated by K-rich granitic rocks, it is likely that the oldzircons have a similar origin.

To a large extent, it appears that the trace element chemistry of zircon is not asensitive indicator of source rock chemistry. The trace element characteristics arereflecting the conditions under which zircon crystallizes which might be similar indifferent rock types despite initial differences in chemistry. This is because zircon isforming at a very late stage in the crystallization sequence and the local conditionsmay be of more importance than the bulk chemistry. However, as noted by Maas etal. (1992), there is only a limited geochemical data base with which to comparezircons from different petrogenetic environments and further work might elucidatesome of the observed variations in the 4.2 Ga zircon suite.

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

Diffusion and Dissolution Studies. The spatial resolution of the ion microprobespot is of the order of 10 µm. However, a depth profile has a resolution of the orderof hundreds of angstroms because the primary beam is sequentially sputtering layersaway. This has great application in surface studies since elemental profiles can bemeasured as a function depth yielding information on dissolution or diffusion.

As an example, Muir et al. (1989; 1990) analyzed surfaces of feldspars afterdissolution and found that 600-1200 Å thick layers were formed that were devoid ofNa, Ca, and Al. The thickness of the layers is strongly dependent on the pH of thesolution 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 theion microprobe than conventional methods since the ion microprobe relies on the


mass differences between nuclides to resolve isobars whereas these can be chemicallyseparated for conventional analysis. This effectively rules out b-decay schemes,such as 87Rb-87Sr, because of the extremely small mass difference between thesenuclides. Furthermore, another common dating scheme, 40K-40Ar, as well assuffering from the problems of mass resolution, is also not suitable because of thepoor ionization efficiency of Ar. For b decay schemes, measurements can generallyonly be made of samples which have a high daughter to parent ratio and aretherefore limited to radiogenic tracer applications such as measuring Sr incarbonates and Hf in zircons. The data can be used to estimate ages, but these aremodel 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 toionize into positive ion species in conventional thermal ionization massspectrometry. They can be ionized by sputtering quite efficiently and so the firstmeasurements 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 foranalysis. However, this method has been superseded by conventional massspectrometry after the discovery that both Os and Re yield quite intense negativelyionized oxide beams by thermal ionization (Creaser et al., 1991).

Since heavy isobars cannot be resolved, ion microprobe dating is reliant on adecay schemes for which the daughter differs from the parent by four mass units ormore. The best example of this for ion microprobe analysis is the U-Th-Pb schemewhich includes four decay chains, each having multiple a decays, for example 238Udecays 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 theisotopic composition of the daughter, as well as a determination of the parentdaughter ratio. Of these schemes, U-Pb is the most amenable to ion probe analysisbecause of the large variation in parent-daughter ratios and also the use of the207Pb/206Pb ratio for age determination. Over geologic time 206Pb/238U and 207Pb/206Pbrange from zero to 1.0 and 0.05 to 0.6 respectively. On the other hand, extremeparent daughter fractionations of Sm-Nd are not found and very high precision isrequired for useful geologic age resolution.

U-Th-Pb in zirconThe U-Th-Pb decay scheme is ideally suited to geochronologic applications

since quite extreme fractionations in the parent to daughter elemental ratios arefound. Of particular interest has been the mineral zircon, which contains U and That significant concentrations but strongly excludes Pb, and is geologically resistant toalteration in thermal processes. This indestructibility has hindered someconventional applications because a zircon crystal can become a complex mix ofdifferent zones that formed at different times. The ion microprobe is ideally suitedto analyzing these different zones and has been paramount in the deconvolution of


complex igneous and metamorphic terranes (Williams, 1992). Furthermore zirconscan be rapidly dated to determine provenance age distributions in sediments, andrecent advances have allowed the dating of Phanerozoic rocks for time scalecalibration.

Early attempts at zircon dating relied on stripping interferences by calculatingthe contributions from other isobars. This technique can be unreliable if theinterferences are not well characterized and some of the early SHRIMP work wasconcerned 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 canbe accurately obtained although the commercial production of more large massanalyzer probes should alleviate that situation. It would not be realistic to review allpapers which have been published concerning ion probe dating of zircon and soonly a selection of topics has been chosen as representative of the applicationsundertaken. First however, it is perhaps important to verify the abilities of SHRIMPwith some comparisons with conventional isotope dilution zircon dating.

Evaluation of the SHRIMP Zircon-Dating Technique. Over the past 10years or so, SHRIMP has been used to date zircons that range in age from thebeginning of the solar system, to some of the youngest tectonically active regions ofthe Earth. For the simplest zircons, a comparison can be made between conventionaldating techniques (using chemically separated aliquots, and zircons dated bySHRIMP. It should be noted that even for morphologically simple zircons, the U-Th-Pb systematics are not necessarily well-behaved since heterogeneous Pb loss, thepresence of thin U-rich rims, and invisible cores can all play havoc with theanticipated results on a micron scale. In this section, six comparisons are madebetween conventional dating methods and ion probe zircon dates in samples rangingin age from 4560 Ma to 2 Ma.

The oldest zircons, for which a conventional tie point is available, are notterrestrial but are found in differentiated meteorites. The howardite-eucrite-diogenite association consists of basalts and gabbros, often brecciated, thatapparently formed on a relatively large asteroidal body referred to as the HEDParent Body. Small zircons (≤30 µm) have been found in several meteorites of thisgroup and some of these have been analyzed by ion microprobe for their U-Th-Pbsystematics. Ireland and Wlotzka (1992) presented data for zircons from the VacaMuerta mesosiderite and for one zircon with approximately 50 ppm U a concordant207Pb/206Pb age of 4563 ± 15 Ma was obtained (Figure 45) in good agreement withthe canonical age of the solar system.

Zircons from the Isua supracrustal belt have been analyzed extensively both byion microprobe and conventionally. Compston et al. (1986) obtained a mean ageof 3807 ± 2 Ma for a suite of Isua zircons, in good agreement with a conventionaldetermination of 3813 ± 9 Ma by Baadsgaard et al. (1984). The conventional datashow a degree of very early lead loss whereas most of the ion probe data areconsistent with only a small degree of recent lead loss with the ancient Pb loss


Figure 45. Concordia plot for Vaca Muerta zircons. VM-1 has only 0.6 ppm Ucompared 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 adaptedfrom Ireland and Wlotzka, 1992).

restricted to a few individual grains. This is exemplified in Figure 46a, which showsisotope dilution analyses of single crystals and ion probe analyses of grains from thesame population.

Zircons from a syenite from East Greenland are concordant with a 207Pb/206Pbage of 2701 ± 5 Ma as indicated by single crystal analyses of abraded grains. Ionprobe analyses of these same abraded grains (analyzed prior to isotope dilutionanalyses) give the same age within error at 2698 ± 7 Ma. There is some dispersionin the 207Pb/206Pb ages tending to lower values suggesting some domains in thezircons have experienced a small degree of early lead loss (Figure 46b).

For zircons younger than around 800 Ma, the 207Pb/206Pb ratio becomes a ratherinsensitive indicator of age as measured on SHRIMP I. For such zircons the206Pb/238U ratio can be used and the U-Pb age is simply determined as an age relativeto the SL13 standard zircon. Zircons dated by the 206Pb/238U ratio at 802 ± 10 Ma byMSID 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 anidentical result by ion microprobe Figure 47b. Both MSID and ion probe results areconcordant within rather large errors in the 207Pb/206Pb ratio because of the smallsample size analyzed.


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 syenitefrom East Greenland. The isotope dilution results from Isukasia indicate the samplehas experienced ancient lead loss and the oldest ages are similar to the agesobtained by ion microprobe analysis. The ages shown for the syenite are derived fromthe weighted mean 207Pb/206Pb ratio. (Unpublished data from A.P. Nutman and C.M.Fanning.)


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 abradedmulti-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.)


Figure 48. Concordia plot of zircons from the Omara granodiorite showing U-Pb datauncorrected for common lead. The age of the granodiorite is interpreted as theintercept of the common Pb trajectory with the concordia which is 2.1 ± 0.1 Ma, and isin 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 reliablyconstrained by the 204Pb/206Pb ratio. Even the 208Pb method of correction results inlarge uncertainties on the radiogenic 207Pb/206Pb ratio because of the lowconcentration of 207Pb and the low number of counts at this isotope. In this case, theuncorrected data can be used to extrapolate to the concordia to obtain an ageestimate. This can be done by either assuming a fixed common Pb composition, orif there is sufficient dispersion in the data, the common Pb composition can bedetermined from the data. This approach is demonstrated with ≈ 2 Ma zircons fromthe Omara granodiorite from the D’Entrecasteaux Islands, a tectonically active areaoff Papua - New Guinea. The zircons have quite high U concentrations (500 - 1000ppm) which yield data quite close to concordia (Figure 48). The age is estimated tobe 2.1 ± 0.1 Ma.

These data clearly demonstrate that the ion microprobe is capable of reliablydating extremely small amounts of Pb in zircon. This allows the complexities ofnatural zircons to be unraveled into distinct geological events for which good agecontrol can be obtained. Well over a thousand samples have been dated onSHRIMP covering the full range of geologic time. In the following sections, only afew of the problems tackled by SHRIMP are outlined as representative of the powerof the ion probe zircon-dating technique in solving geological problems.


The Oldest Terrestrial Zircons. One of the most conspicuous applications ofthe SHRIMP zircon-dating program has been the search for the oldest zircons onEarth. Prior to the SHRIMP analyses, the oldest-known terrestrial material wasobtained from the Isua supracrustal belt for which some units were found to be alittle over 3.8 Ga. This still leaves some 750 Ma between the formation time of themeteorites, believed to be close to the formation time of the earth, and thepreservation of felsic crust on Earth. Within the first few years of routine zirconanalysis, zircons that significantly predated Isua were found in Archean quartzitesfrom Western Australia (Froude et al., 1983; Compston and Pidgeon, 1986). Thesefindings have been paramount in developing models for the formation of Earth’searliest crust and have pushed back the oldest recognizable vestiges of Earth’shistory to 4.28 Ga.

The zircons were found as a small proportion (<2 % in Mt Narryer) of the totaland so a large number of zircons had to be analyzed to ascertain firstly theirpresence and then their abundance. Schärer and Allègre (1985) attempted to verifythe findings of Froude et al. (1983) with conventional single-grain analysis. Afteranalyzing 32 grains, Schärer and Allègre had found no “old” zircons. As arguedby Compston et al. (1985) this simply reflects the statistics of small numbers. At anabundance level of 0.02, there is roughly a 50 % chance of finding an old grain in32 attempts. Schärer and Allègre may have inadvertently biased their population bychoosing grains according to morphology. However, since there is no apparentcorrelation between morphology and age characteristics it is more likely that theywere simply unlucky. Schärer and Allègre also claimed that the discordance patternof the younger grains as deduced by ion microprobe analysis did not match theirconventional data set. While the ion probe data set was consistent with three mainpopulations at 3300, 3500, and 3750 Ma, the conventional data set also showsintermediary ages which are likely to be due to mixed ages between cores and rimsof the more structured zircons. This is illustrated in Figure 49 which shows anumber density plot of 275 ion probe analyses with the conventional data pointssuperimposed (Kinny et al., 1990).

The exchange between the ANU and Paris groups is an indication of the healthyskepticism to which the report of > 4 Ga zircons was first met. It should be notedthat this was the first paper on zircon analysis to be published from the ANU ionprobe group and so the technique had not had the chance to substantiate itself onless complicated and less controversial problems initially.

Conventional analyses of > 4 Ga zircons have now been made and confirm theantiquity of these grains. Kober et al. (1989) analyzed 41 zircons from Jack Hillsby stepwise evaporation and found that five of the grains had minimumcrystallization ages between 4.07 and 4.17 Ga. Single-grain isotope dilutionanalyses also confirm the old lead compositions in these detrital grains but whileconcordant domains could be analyzed with the ion microprobe, conventionalanalyses have been often discordant (C. M. Fanning, personal commun.).


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 thediscordia 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 Australiahave not been found to date with the oldest gneisses having formation ages nogreater than ≈3730 Ma (Nutman et al., 1991). Pre 3.8 Ga detrital zircons have alsobeen found in Archean quartzites from Greenland, the oldest being 3850 Ma(Nutman and Collerson, 1991), from China, 3850 Ma (Liu et al., 1992), and fromthe Beartooth Mountains, Montana where ages range up to 3960 Ma (Mueller et al.,1992). In contrast to western Australia, gneisses with similar formation ages havebeen identified in these cratonic areas with the 3.96 Ga Acasta gneiss being theoldest rock identified so far (Bowring et al., 1989).

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


Figure 50. Zircon age histogram from a paragneiss from the West Coast of NewZealand. This pattern is also found in Ordovician sediments in the same region and inOrdovician sediments from eastern Australia and Western Antarctica indicating originalcontiguity of the terranes. (Data from Ireland, 1992).

zircons allow the evaluation of possible sources for the zircons and hence thesediments; the ages are simply matched with possible source terranes. The secondpiece of information is more conjectural in its use. The relative proportions of thezircons may not in fact reflect the proportions of the detritus from different terranesfor a number of reasons. Firstly, different rock types have different zirconconcentrations and so a granite will bias the sample relative to an intermediate rockand mafic rocks will not even be represented. Secondly there can be a sampling biasduring zircon separation (size, magnetic susceptibility), and then in the choice ofgrains for analysis. A large number of randomly selected grains fromunfractionated separates reduces some of these problems but bias cannot beexcluded.

Perhaps of equal importance to the interpretation of the abundance peaks is theuse of the zircon signature as a fingerprint of the rock. This might allowcorrelations of sedimentary units over large distances. For example, it is evident insouth eastern Australia, western New Zealand and western Antarctica that Cambro-Ordovician sediments have identical detrital zircon patterns (Figure 50) which isstrong 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


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 instrongly zoned crystals (filled symbols) projecting towards a granulite event at 2.95 Garepresented by overgrowths (inset). Structureless grains grew at around 2.46 Ga andminor 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 ofAntarctica may be useful in firstly correlating the rock units and secondlydetermining the provenance directions for the distinctive age components.

On a slightly different note, Brimhall et al. (1992) used zircon morphology andage characteristics to show that soils develop by dissolution and collapse of thebedrock and by invasive transport of aerosol detritus from above. Euhedral zirconsthroughout the profile have the age of the unweathered granite beneath, but roundedzircons are found to have a wide provenance age. These foreign zircons aretransported as aerosols and penetrate the soil horizon to a maximum depth of 2 mby movement down channels left by decayed roots and by bioturbation.

Complex Metamorphic Terranes. The ion microprobe is a particularly usefuldevice for analyzing complexly zoned zircons from metamorphic terranes. Zirconis a highly persistent mineral once crystallized and so successive thermal episodescan deposit zircon resulting in several generations within the same crystal. This was


highlighted by Black et al. (1986) who found that four distinct generations ofzircon growth had occurred in an orthogneiss from Mt Sones, Antarctica (Figure51). These zones could be independently sampled and correlated with conventionalisotope 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. Aftergranulite facies metamorphism dated at 2948-17

+31Ma, extensive zirconrecrystallization occurred but only a small amount of new zircon growth. This isapproximately the age of extensive isotopic resetting in the Rb-Sr systemcorresponding to the D2 event marking the end of granulite facies metamorphismbut is significantly younger than the D1 event marking the onset of themetamorphism at 3070 ± 34 Ma. Structureless pink zircons crystallized at 2479 ±23 Ma which is within error of the D3 event determined by Rb-Sr at 2456-5

+8 Ma(Black et al., 1983). Finally tectonism at around 1000 Ma caused a small degree ofzircon Pb loss.

Calibration of the Geologic Time Scale. The high selectivity available in ionprobe analysis has led to its application to geological time scale work. This isperhaps the most stringent type of analysis possible because of the poor ageresolution afforded by the 207Pb/206Pb ratio and therefore the need for accurate andprecise calibration of Pb/U ratios. The external reproducibility of repeat analyses ofthe SL13 standard appears to limit the precision of individual analyses to around 2%. Therefore a large number of analyses are required of both standard andunknown in order to precisely calibrate the age. The benefits of ion probe analysisare the ability to detect and reject statistical outliers that may be due to inheritance orPb loss; the inclusion of such domains in a conventional analysis could bias theresult to higher or lower ages respectively.

Zircons are typically separated from tuffaceous units with well-controlledstratigraphic relationships. Analyses of zircons from early Cambrian tuffs fromMorocco, China, and Australia give identical results at 520-525 Ma indicating theCambrian - 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 theages were clearly younger than the canonical time scale of Harland (1989). The ionprobe 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 ofinheritance in the multigrain samples analyzed by Tucker et al. (1990) since ionprobe ages agreed closely with high-precision K-Ar and Rb-Sr for other volcanicrocks. At the other end of the Paleozoic, Claoué-Long et al. (1991) have dated thePermian - Triassic boundary at 250 ± 3 Ma.

Other chronometersProbably 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


ratios as well as the two coupled decay systems are unique. There are severalsystems that may be of potential interest in terms of model ages. Hf isotopiccompositions in zircons has been shown to be useful in identifying the crustalextraction age of the zircons thereby seeing through metamorphic events that mayhave reset the U-Pb system. Sr isotopes too might be used in a similar fashion as aradiogenic isotope tracer. With both these systems however, the parents anddaughters are isobaric and require high mass resolution for them to be separated.Hence there is no foreseeable direct chronological application of these systems. Inthe Re-Os system, this has been sidestepped by chemically separating these elementsprior to their analysis on the ion probe. However, there is again no real possibilityof achieving in situ analysis as yet. Of the commonly used a-decay systems, Sm-Nddating may prove to be possible but there will have to be substantial advances inionization and transmission efficiency at high mass resolution for it to beworthwhile.

VI.!PROSPECTS AND CONCLUSIONS

The ion microprobe has indeed come of age. Its capabilities are now recognizedalthough it is often viewed as simply a zircon-analysis device or a highly sensitivetrace-element detector. Despite the quite dedicated application of individual ionprobes, they are all capable of the same retinue of techniques. Basically what workson one ion probe will work on all. The limitation to this concept however is thetransmission efficiency and hence sensitivity of the individual machines. For thisreason, it is likely that the new “consumers” of these devices will require large highsensitivity machines capable of U-Pb analysis which can also be used for the lowersensitivity applications as well.

There are ultimate limits to the sensitivity that can be achieved. The number ofatoms in a 10 µm spot sputtered to a reasonable depth is finite. Of the atoms presentonly a certain percentage are of interest for the analysis and of these only a smallfraction are ionized and will make it to the detector. The large ion probes aredesigned to collect and transfer as large a fraction of the ions from the sample intothe mass spectrometer and only modest gains (factor of 3-10) will be forthcomingwith better designs. If higher sensitivity is to be achieved then more ions must beproduced per unit volume or the ions that are produced must be collected moreefficiently.

There has been a great deal of effort expended already in the application of lasersin the selective excitation of various species. It is still not clear whether this willresult in a greater sensitivity per unit time than that already achieved in SHRIMP-style ion microprobes. The ionized fraction of sputtered material is generally takento be around 0.1 %. Therefore there is, at most, a factor of 103 to be achieved ifresonance ionization was completely efficient. The problem is that sufficient poweris only available in pulsed lasers with only nanosecond duration. In order to achievecomparable ion throughput, the laser would need a repetition rate of the order of 1


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 acharge collection mode would have to be employed rather than the ion countingschemes currently utilized. Even then, the isotopic abundances measured followingresonance ionization can be highly variable and dependent on the exactexperimental setup (Spiegel et al., 1991). Until high-powered continuous lasers areavailable and the fundamental mechanisms of ion production under laserbombardment can be understood, resonance ionization ion microprobes will beseverely limited in their application.

One of the main losses of ions in analysis is due to the collection cycleefficiency, that is, the time that is spent counting on one peak means that the otherpeak(s) is(are) not being counted. For equal counting time on two peaks thecollection efficiency is 50 %. For simple systems, the obvious method to improvecollection efficiency is to go to multiple collection, as has been done in conventionalmass spectrometers. There are clearly major benefits to be had and it is a directionthat commercial manufacturers are addressing. The main difficulties concern thelow count rates on some peaks requiring ion counters to be used, and the physicallysmall space into which these counters must fit. Both the CAMECA 1270 andSHRIMP II ion microprobes are to be fitted with multiple collectors based on ioncounters but these have not been tested as yet.

The limits on sensitivity also place limits on the ultimate spot size that can beeffectively used. In rastered ion imaging there is a clear advantage in a small spotfor looking at distributions of elements in the target. However, in high precisionisotopic analysis, there is almost no prospect of using a 1 µm spot to achieve subpermil precision in any elemental system. There are simply not enough atoms andhence ions produced. For isotopic analysis the minimum number of ions requiredcan be calculated according to Poisson counting statistics and from this a minimumvolume to be consumed can be calculated. A reasonable minimum probe diameterfor isotopic analysis will probably be in the 5-10 µm range.

There are clearly some great difficulties in the next stages of development butfuture generations of these machines may be very different to what we have todayand capable of a range of applications that were unforeseen. The main difficulty inthe present generation of ion microprobes is in obtaining access. Large ionmicroprobes are expensive and are likely to be available only through regionalfacilities. Even then there is only a finite amount of time available on a singlemachine and it is likely that the smaller user will be unable to compete for time. Inthis regard, the availability of the ion microprobe is likely to be quite different to theelectron probe. The price of electron probes was soon within reach of earth sciencedepartments with only quite modest budgets. It is unlikely that this will ever be thecase for high sensitivity ion microprobes at least in their present configurations.

The ion microprobe has proved to be a highly versatile machine. Despite the lackof a thorough understanding of the ionization processes involved, the ion probe canbe used for a variety of types of quantitative analysis both in terms of isotope ratios


and elemental abundances. These analyses are generally relative measurements inthat a well-known standard must be analyzed in order to quantify the effects ofselective ionization in the sputtering process. There are still a large number ofapplications that have not been fully tamed as yet. The most obvious case is that ofoxygen isotopic analysis of silicates where despite a great deal of effort, thefractionation of the oxygen isotopes cannot be controlled sufficiently for use interms of stable isotope geochemistry. It is, however, only a matter of time before arecipe can be developed that will enable routine measurements of this type. So toowith other applications. Individual recipes must be developed, firstly by evaluatingisobaric interferences and ion intensities and then determining the reproducibility ofthe measurements. In this way it is likely that a large retinue of new applications willbe forthcoming in the near future.

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

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