11  Rare gas geochemistry

 

The elements known as the rare, inert or noble gases possess unique properties which make them important in isotope geology. The low abundance of these rare gases allows them to sensitively record several types of nuclear process, even including rare nuclear-fission reactions. (In contrast, the relatively larger abundance of other fission product nuclides such as the ‘rare’ earths swamps fissiogenic production). Another property of these gases is their inertness, which allows unique insights into the Earth’s interior because of their lack of interaction with other materials. Finally, as isotopic tracers, rare gases can give information about the degassing history of the mantle, the formation of the atmosphere, and mixing relationships between different mantle reservoirs.

 

 

11.1     Helium

 

Helium has two isotopes, 4He and 3He. The former was recognised by Rutherford (1906) to be the " decay product of actinide elements, and hence comprised the first radiometric dating method. However, the great diffusivity of helium made the method very susceptible to thermal disturbance, and it has therefore been abandoned in all but the most specialised applications (e.g. Wernicke and Lippolt, 1993).

 

            Non-radiogenic 3He was first discovered in nature by Alvarez and Cornog (1939). Alvarez and Cornog estimated (using a cyclotron) that atmospheric helium had a 3He/4He ratio ten times greater than natural oil-well gases from the Earth’s crust. Aldrich and Nier (1948) confirmed this observation by mass spectrometric measurements, and determined atmospheric and well-gas 3He/4He ratios of ca. 1.2 H 10!6 and 1 H 10!7 respectively. They concluded that there must be independent sources of the two isotopes, one of which could be primordial.

 

 

11.1.1  Mass spectrometry

 

Mass spectrometric analysis of helium is broadly similar to argon isotope analysis in K)Ar dating (section 10.1.1). However, in helium isotope analysis there are no ‘extra’ isotopes available to allow accurate correction for atmospheric contamination. Therefore it is critical to minimise the extent of this contamination during helium extraction and analysis. Uncertainties in the atmospheric ‘blank’ may contribute the principal error in helium isotope analysis, especially for rock samples. Well-gas samples, being larger, are less susceptible to atmospheric contamination during analysis, but may have come from an open system in the natural environment. In the case of rock analysis, absorbed atmospheric helium is usually driven off by overnight heating at 200)300 oC. The sample gas may then be extracted by melting the rock or by crushing under vacuum. A combination of both techniques (e.g. Kurz and Jenkins, 1981) provides an extra check against the possibility of atmospheric contamination, both in the laboratory and the environment.

 

            Two steps are necessary in order to reduce blank levels in the mass spectrometer for all rare gas analyses. One is to polish all internal surfaces of a metal instrument to minimise the absorption of gases onto the walls of the vacuum system. Another is to reduce the internal surface area of the instrument as much as possible, for example by boring the flight tube out of a solid piece of steel rather then using welded pipe. A low internal volume also yields better sensitivity for very small samples.

 

            All rare gas analyses are performed in the static gas mode (i.e. with vacuum pumps isolated). As a result, hydrogen tends to build up in the instrument, so that its molecular ions HD+ and H3+ cause isobaric interferences onto 3He+. Therefore, the vacuum system in some older machines contains a small titanium ‘getter’, designed to absorb H2 released inside the instrument (Clarke et al., 1969). Nevertheless the peak composed of HD and H3 may still be much larger than that of 3He, so it is essential to separate them by mass. This can be done by making use of the 0.006 atomic mass unit difference between 3He and the other two species (Fig. 11.1) which results from their different nuclear binding energies. In order to achieve this separation at mass 3, a resolution of one mass unit in 600 is necessary, which can be achieved with an instrument of ca. 25 cm radius (Clarke et al., 1969; Kurz and Jenkins, 1981).

Fig. 11.1. Scan of peaks in the region of mass 3 during helium isotope analysis, showing the separation of molecular interference using high spectral resolution. Masses are quoted relative to 12C = 12.000. After Lupton and Craig (1975).

 

            In order to measure the very large difference in intensity between 3He and 4He signals, it is most convenient to measure the former on a multiplier detector and the latter by Faraday detector. These can only be used in the static collection mode if a branched flight tube is available, because of the extreme divergence of the mass-3 and mass-4 ion beams (Lupton and Craig, 1975). Alternatively, peak switching is performed by changing the accelerating potential or magnetic field (e.g. Clarke et al., 1969; Poreda and Farley, 1992).

 

 

11.1.2  Helium production in nature

 

In order to determine whether primordial helium is an important constituent in the Earth, it is necessary to determine the 3He/4He ratio of primordial solar system helium, and also the production ratio in nuclear and cosmogenic processes. A good indication of the composition of primordial helium is provided by the 3He/4He ratio of (2 ) 4) H 10!4 measured in gas-rich carbonaceous chondrites (Pepin and Signer, 1965). These meteorites have such high primordial gas contents that their composition is not significantly perturbed by ‘cosmogenic’ helium (a product of cosmic-ray spallation effects). In contrast, most 3He in iron meteorites is cosmogenic.

 

            Early calculations of the nuclear 3He/4He production ratio in igneous rocks were made by Morrison and Pine (1955). Radiogenic production of 4He is obvious, since the " particle is synonymous with a 4He nucleus. However, ‘nucleogenic’ 3He can also be generated by neutron bombardment of light atoms. Radioactive decay of uranium generates a neutron flux in rocks by two mechanisms. Spontaneous fission is a minor source, but by far the dominant source of neutrons is the collision of " particles with nuclei of light elements. Some of these neutrons reach epithermal energies, where they can induce the (n, ") reaction on lithium. The tritium thus produced decays to 3He:

 

            6Li    +   n   6     3H     +    "

 

            3H       6     3He    +    $     (t1/2 = 12 yr).

 

            Kunz and Schintlmeister (1965) calculated that 3He generation by this reaction is at least three orders of magnitude more efficient than all other neutron-induced reactions. Given the uranium (plus thorium) and lithium content of a rock, the 3He/4He yield can be calculated (Gerling et al., 1971). The results are consistent with the range of (1 ) 3) H 10!8 measured empirically in old granites. The calculations were also confirmed by experimental irradiation of ultrabasic rocks in a reactor (Tolstikhin et al., 1974). Some other possible sources of 3He (via tritium) are:

 

            238U   6   fission products   +  (2 H 10!4)  3H,

 

            7Li   +   "    6   8Be    +    3H,

 

            7Li   +   (    6   4He    +    3H.

 

However, Mamyrin and Tolstikhin (1984) calculated total 3He/4He production ratios of 8 H 10!12, < 7 H 10!9 and ca. 10!13, respectively, for these reactions, making them insignificant compared with the main (n, ") reaction. It was concluded from these observations and calculations that no nuclear process is capable of generating 3He/4He ratios significantly greater than 10!8 in normal rocks. However, uranium ores generate lower ratios, while Li-rich minerals generate abnormally high ratios.

 

            Another mineral in which high 3He/4He ratios have been observed is diamond. Values up to 3 H 10!4 were interpreted by Ozima and Zashu (1983) as indicative of primordial mantle reservoirs, but have been attributed by later workers to either nucleogenic or cosmogenic 3He production. For example, Lal et al. (1987) attributed high 3He/4He ratios in alluvial diamonds from Zaire to cosmogenic production while the diamonds were exposed at the surface.

 

            On the other hand, Kurz et al. (1987) and Zadnik et al. (1987) measured 3He/4He ratios as high as 1.4 H 10!3 in diamonds mined directly from kimberlite pipes at depths of ca. 26 and 200 m respectively. Since cosmic rays cannot penetrate to such depths, these helium signatures were attributed to nucleogenic production. Evidence for this interpretation came from the observation of isotopic variability within individual diamonds, and the determination of 3He/4He ratios higher than solar in the latter study. In both cases, 3He production was attributed to the (n, ") reaction on lithium. For this process to occur, the diamond and its inclusions must be irradiated by neutrons from outside the crystal, so that radiogenic 4He production in the diamond itself is suppressed.

 

            In situ cosmogenic helium production in terrestrial rocks was proposed by Jeffrey and Hagan (1969), but was not identified unambiguously until work by Kurz (1986a) and Craig and Poreda (1986). In a detailed helium isotope study of sub-aerial lavas from Haleakala volcano, Kurz discovered very high 3He/4He ratios released by step heating of some near-surface 0.5 ) 0.8 Myr-old alkali basalts. Low-temperature gas releases from samples within 0.5 m of the weathered surface gave 3He/4He ratios over 10!3 (Fig. 11.2). These values are even higher than those from primordial meteoritic or solar-wind helium.

 

            In contrast, step heating of samples from a similar stratigraphic horizon that were buried under ca. 160 m of younger flows yielded MORB-like helium (3He/4He = 1.2 x 10!5). Helium released by crushing of phenocrysts also gave a MORB signature, for both the buried and surface samples. Therefore, Kurz argued that crushing released magmatic helium from vesicles, but step heating of old surface samples released dispersed cosmogenic helium from the rock matrix. Young surface samples such as the 1790 flow on Haleakala do not show these effects, ruling out anthropogenic bomb tritium as the source of the 3He.

Fig. 11.2. Step heating helium isotope analysis of a surface sample of Haleakala lava showing a large cosmogenic component, especially  in the low-temperature release steps. Crushed vesicles yield the ‘true’ mantle value. After Kurz (1986a).

 

            Kurz (1986b) went on to examine cosmogenic 3He production as a function of depth below the surface of a lava flow. Spallation reactions caused by cosmogenic neutrons are the dominant source of 3He at the surface, but neutron fluxes are attenuated exponentially downwards. Nevertheless, 3He abundances showed less attenuation with depth than expected. This was attributed to production by cosmic-ray muons, which have a greater penetration depth than neutrons. Muon capture by nuclei causes neutron emission, which in turn produces 3He via the (n,") reaction on lithium. The depth dependences of different production routes for 3He are summarised in Fig. 11.3 (Lal, 1987).

Fig. 11.3. Calculated production rates for 3He by different processes as a function of depth in a rock surface. Depths are expressed as kg / cm2, which is approximately equal to 1/3 H depth in m. After Lal (1987).

 

            Cosmogenic isotopes represent a useful tool for determining exposure ages of rock surfaces (section 14.6). However, the great diffusivity of helium may be a problem in using 3He in such studies. For example, Cerling (1989) showed that helium was often not quantitatively retained in quartz, the most widely used material in surface exposure dating. On the other hand, 21Ne displays cosmogenic production with an attenuation depth similar to that of 3He (Sarda et al., 1993). The lower diffusivity of neon may therefore make it more widely useful in surface-exposure dating (section 11.2.1).

 

 

11.1.3  Terrestrial primordial helium

 

The first accurate determinations of the atmospheric 3He/4He ratio were made by Mamyrin et al. (1970) and Clarke et al. (1976), yielding ratios of 1.40 H 10!6 and 1.38 H 10!6. Because atmospheric helium is universally used as a mass spectrometric standard, it is convenient to express 3He/4He ratios in unknown samples relative to the atmospheric ratio in the form Runknown/Rair (R/RA). However, because cosmogenic 3He production in the atmosphere is difficult to quantify accurately, it is not possible to prove the existence of a primordial helium source in the Earth simply by the fact that the atmosphere is two orders of magnitude richer in 3He than radioactive production in rocks.

 

            Stronger evidence of a primordial helium signature in the Earth was provided by Clarke et al. (1969), when they discovered that deep water from the Pacific ocean was enriched in 3He by up to 20% relative to the atmosphere. However, Sheldon and Kern (1972) and Lupton and Craig (1975) hypothesised that this could conceivably be due to a temporary weakening of the Earth’s magnetic field, during which the atmospheric 3He/4He ratio was elevated by greater cosmic ray penetration.

 

            Convincing evidence of primordial helium in the Earth was first provided by Mamyrin et al. (1969), who found 3He/4He ratios ten times higher than atmospheric values in thermal fluids from the Kurile Islands. Subsequently, 3He/4He ratios as high as 20 times atmospheric were found in hot springs from Iceland (Mamyrin et al., 1972). Even higher R/RA values have been found in oceanic volcanic rocks: up to 26 in sub-glacial basaltic glasses from Iceland and up to 32 in basaltic glass from Loihi Seamount off Hawaii (Kurz et al., 1982).  These values are regarded as reliable indicators of primordial helium because they come from rocks which have been shielded from in situ cosmogenic production, either by water or by ice.

 

            The highest R/RA value for any plume source (38) was found in crushed phenocrysts from an olivine basalt in the neovolcanic zone of NW Iceland (Hilton et al., 1999). These rocks are not shielded from cosmogenic production, but it was argued that cosmogenic contamination of the data could be excluded. This was based on the observation that the residues from crushing (which should contain any dispersed cosmogenic component) had less elevated 3He/4He ratios than the gases released by crushing. The helium signature of the Iceland plume can also be seen spreading out over a large area of the North Atlantic as the plume head contaminates the asthenospheric upper mantle (Fig. 11.4).

Fig. 11.4. Plot of helium isotope ratios along the Mid Atlantic Ridge, expresses as deviations from the atmospheric value (R/RA).  The primordial 3He signature of the Iceland plume ( x ) is elevated relative to MORB ( ! ). Shaded fields display mixing of sources. After Kurz et al. (1985).

 

 

11.1.4  The ‘two reservoir’ model

 

In contrast to the variable helium isotope signatures in OIB, Craig and Lupton (1976) found a relatively narrow range of R/RA ratios around 9 in MORB glasses from various ocean basins. Subsequent data have confirmed this narrow helium isotope range in MORB, relative to the large variations in plumes (e.g. Fig. 11.4). The intermediate helium isotope composition of MORB, between those of atmospheric and plume sources, can be explained by partial outgassing of primordial helium from the upper mantle, followed by radiogenic helium production. This caused the upper mantle to develop a lower 3He/4He composition than the un-degassed lower mantle, where radiogenic production is swamped by primordial helium.

 

This partial degassing or ‘two reservoir’ model for the mantle was originally proposed to explain argon isotope systematics (Hart et al., 1979), and was applied to helium by Kaneoka and Takaoka (1980). Unfortunately Kaneoka and Takaoka based their case on rare gas compositions in phenocrysts from Haleakala volcano, Hawaii, which were subsequently shown to be contaminated with atmospheric argon and cosmogenic helium (Fisher, 1983; Kurz, 1986a). However, more recent data from OIB samples have confirmed the elevated R/RA value of the plume source, as noted above. Hence the two-reservoir model for mantle helium has been widely accepted.

 

            An early test of the two-reservoir model was made by comparing helium and heat fluxes from the Earth (O’Nions and Oxburgh, 1983). These fluxes must be related because the decay of uranium and thorium produces both radiogenic helium (alpha particles) and also radioactive heating. Taking account of the small amount of heat also derived from 40K decay, O’Nions and Oxburgh (1983) calculated that 1012 atoms of 4He would be generated in the mantle per joule of heat production. They then calculated the concentration of U necessary to generate the observed helium and heat fluxes. The results were somewhat surprising, because the amount of uranium required to generate 88% of the Earth’s oceanic helium flux can produce only 3% of the oceanic heat flow.

 

            The logical source for some of the remaining heat flux is crystallisation of the inner core, which releases heat through the outer core and mantle by convection. However, this convection must operate in such a way that the reservoir of primordial 3He in the Earth’s interior is not completely exhausted. Therefore, O’Nions and Oxburgh proposed that a boundary layer inhibits upward transport of helium from the ‘primordial’ reservoir much more effectively than it inhibits the transport of heat. They envisaged this boundary layer at 700 km depth, separating the upper and lower mantle. This would imply that the whole of the lower mantle is a kind of primordial helium reservoir. Other workers have preferred the core)mantle boundary, although it is not clear whether the core could represent the repository of primordial Earth helium. This question will be discussed further below.

 

            Another challenge for the two reservoir model is to explain the respective concentrations of helium and other rare gases in the two reservoirs. Thus, if OIB come from the un-degassed source, we would expect them to contain more helium than MORB glasses from the degassed upper mantle. However, OIB glasses actually have ten times less 3He than MORB glasses (Fisher, 1985). This observation has sometimes been called the ‘helium paradox’ (e.g. Hilton et al., 2000). However, although this evidence is problematical, it is not definitive, due to the poorly constrained behaviour of rare gases during the melting process. For example, the dynamics of mantle convection and melt segregation under ridges must be different from those of plumes; so that ridge magmas collect helium from a greater volume of mantle during the melting process (section 13.3.6). Hence, most workers have taken the isotopic evidence in favour of the two-reservoir model for helium as definitive, and over-riding any problems involving rare gas abundances. The case for the heavy rare gases will be discussed later.

 

            The existence of a primordial helium reservoir in the Earth was challenged more recently by Anderson (1993), who attributed this signature to the subduction of cosmic (interplanetary) dust particles. These particles were found to accumulate in ocean-floor sediments by Merrihue (1964). Cosmic dust has 3He/4He ratios similar to gas-rich meteorites (ca. 3 H 10!4), but unlike meteorites, these particles can fall to Earth without burning up in the atmosphere (Nier and Schlutter, 1990). Hence, ocean-floor sediments develop a ‘primordial’ helium isotope signature (Fig. 11.5a).

Fig. 11.5. Histograms of a) 3He/4He and b) 3He/20Ne in cosmic dust particles (stipple) and ocean-floor sediments (white) compared with the rare gas composition of MORB (hatched) and OIB (black). The Solar wind composition is shown for reference. Data from Allegre et al. (1993).

 

            The rare gases in cosmic dust particles are encapsulated in magnetite grains, which are relatively resistant to thermal degassing (Matsuda et al., 1990). Therefore, the cosmic helium in ocean-floor sediments might survive the subduction process and be transported into the deep mantle. In contrast, atmospheric rare gases trapped in ocean-floor sediments are very susceptible to thermal degassing. Staudacher and Allegre (1988) argued that subduction-related volcanism is at least 98% efficient in scavenging these atmospheric gases from subducted sediments before they can reach the deep mantle.

 

            Because cosmic dust might survive the ‘subduction barrier’ against atmospheric rare gases, it has the potential to deliver helium with a primordial signature into the deep mantle. This possibility was been recognised by several workers (e.g. Allegre et al., 1993), but Anderson (1993) took the model a step further by attributing most of the primordial helium signal in mantle hot-spots to subducted cosmic dust. However, Allegre et al. (1993) used neon isotope data to place upper limits on the amount of cosmic 3He which can enter plume sources. They noted that the 3He/20Ne ratio in cosmic dust is one to two orders of magnitude lower than 3He/20Ne in the upper mantle (Fig. 11.5b). Furthermore, helium has a much greater diffusivity than neon, which would promote its preferential degassing from grains of cosmic dust during subduction (Hiyagon, 1994). Therefore it appears that subduction of cosmic dust cannot contribute more than a small fraction of the mantle 3He budget without causing excessive enrichment of 20Ne in submarine glasses.

 

            Although helium isotope ratios provide the best evidence for a primordial gas reservoir in the Earth, this single isotope ratio cannot provide enough degrees of freedom to constrain the complex mixing processes expected to occur in the mantle. Hence, various attempts have been made to compare helium isotope signatures with other isotope ratios in oceanic volcanics, in order to provide extra constraints on mantle processes.

 

            One such approach is the comparison of helium and strontium isotope data (Kurz et al., 1982; Lupton, 1983). MORB samples define a restricted range of compositions on a plot of helium isotope ratio against 87Sr/86Sr, but ocean islands are widely scattered (Fig. 11.6). While Loihi defines the most primordial helium composition, some ocean islands such as Tristan, Gough and the Azores have 3He/4He ratios lower (more radiogenic) than MORB. Similar low ratios have subsequently been found in the HIMU islands of the SW Pacific (Hanyu and Kaneoka, 1997). These low 3He/4He ratios require a component of radiogenic helium from a long-lived U- or Th-rich source, which can most easily be satisfied by the recycling of oceanic crust and sediments into the mantle, as inferred from lithophile isotope data (section 6.5). Finally, a possible third type of plume source is exemplified by the data distribution from Samoa in Fig. 11.6. This was originally attributed to mixing of primordial and recycled helium in the deep mantle, but more recent work suggests a mixing process in the shallow mantle as the cause (section 11.1.5).

Fig. 11.6. Plot of 3He/4He against Sr isotope ratio to show mixing between the MORB reservoir and primordial and recycled plume sources. After Lupton (1983).

 

            The MORB field in Fig. 11.6 breaks into two lobes with geographical boundaries. The main field trends slightly towards the primordial source, while the Mid Atlantic Ridge between 33 and 50 oN defines a subsidiary field with more radiogenic helium, consistent with contamination by the nearby Azores plume. This evidence suggests that plumes break into trains of blobs which locally contaminate the MORB source. According to the nature of the plume material contaminating any given ridge segment, the MORB array can trend towards either the primordial or the recycled type of plume.

 

            Models which involve large-scale recycling of crustal sources back into the mantle imply that the upper mantle has a short residence time for many elements, and would therefore have reached a steady state condition at the present day (e.g. sections 6.3.3 and 13.3.7). In view of the ease with which helium can escape from any system at elevated temperatures, it represents the ultimate incompatible element, and should therefore have the shortest residence time. Since it was argued above that negligible helium is subducted, input to the upper mantle must be restricted to primordial helium escaping from the lower mantle, plus in situ production of radiogenic helium from U-series isotopes.

 

            Kellogg and Wasserburg (1990) assumed a steady state between supply and degassing in order to determine the residence time of helium in the upper mantle. They argued that ridges are the principal sites where helium escapes from the upper mantle (whereas hot spots dominate in outgassing the lower mantle). Hence, they used a simple calculation to estimate the residence time of helium in the upper mantle:

 

tau = mass of upper mantle / rate of outgassing                         

 

Based on a depth of 670 km, the upper mantle has a mass of 1 H 1027 g. Also, assuming ocean floor production at 3.5 km3 / yr and a melting depth of 60 km, the rate of mantle outgassing is estimated at 7 H 1017 g / yr. Hence, Kellogg and Wasserburg calculated a helium residence time of 1.4 Byr in the upper mantle. On the other hand, O’Nions and Tolstikhin (1994) estimated a somewhat shorter residence time of 1.1 Byr, based on an upper mantle mass of 1.1 H 1027 g and an outgassing rate of 1 H 1018 g / yr (corresponding to a melting depth of 90 km). These relatively short residence times suggest that the upper mantle has been completely outgassed of primordial helium. However, in view of the extreme volatility of helium, they also represent a minimum for the upper mantle residence times of lithophile elements (section 6.3.3).

 

            Contamination of the MORB source by OIB sources can occur at various scales, from plumes to isolated blobs and sheets. To evaluate this process, Allegre et al. (1995) examined the dispersion of helium isotope data in MORB as a function of the spreading rates of various ridges. They found that several ridges defined a strong inverse correlation between isotopic dispersion and spreading rate (Fig. 11.7). This led them to suggest that the MORB source has a stirring time for helium about four times shorter than the mean residence time of helium in this source. Other workers have examined isotopic variations of particular ridge segments in more detail, and shown how these can be explained by local contamination by plums or plumes. For example, Graham et al. (1996) observed correlated He–Pb and He–Sr isotope systematics in the South Atlantic.

Fig. 11.7. Plot of the standard deviation of helium isotope ratios for different ridges against the reciprocal of spreading rate, showing a good correlation for several ridges. The South Pacific displays more isotope heterogeneity than expected. After Allegre et al. (1995).

 

            The relatively coherent account of the two reservoir model given above has been threatened more recently by increasing geophysical evidence for single layer mantle convection (section 6.2.3). This militates against traditional box models which make the lower mantle the source of primordial helium signatures. Several alternative approaches have been proposed to deal with this problem.

 

            An attractive way of preserving the two reservoir model in a mantle with single layer convection is to invoke increasing mantle viscosity with depth. It is argued that this could cause large lumps of the lower mantle to be preserved intact, without being streaked out and homogenised by convection. This model was tested by two-dimensional numerical modelling of one-layer convection in such a mantle (van Keken and Ballentine, 1998; 1999). However, these workers argued that models which were realistically close to the real Earth in terms of viscosity and phase transformations could not preserve lower mantle domains large enough to retain a significant primordial helium reservoir. Opinion about the effect of three dimensional modelling is divided: van Keken and Ballentine (1999) argued that this would cause more rapid homogenisation, whereas other workers (e.g. Schmalzl et al., 1995) suggested that it would cause less rapid homogenisation.

 

            A second approach to preserving the two reservoir model is to place the primordial reservoir in the core. Work by Matsuda et al. (1993) suggested that the core would have only a limited helium budget, but on the other hand, osmium isotope data (section 8.3.5) support this model. A recent review of the evidence (Porcelli and Halliday, 2001) is equivocal.

 

            A third approach to this problem is to at least partially dismantle the two reservoir model. For example, studies by Coltice and Ricard (1999), Anderson (2001) and Seta et al. (2001) suggested that the ‘primordial’ or ‘relatively un-degassed’ helium reservoir does not exist. Instead, they argued that all of the mantle is equally degassed, but some parts (e.g. the MORB reservoir) are more enriched in radiogenic helium due to uranium recycling via enriched OIB plumes. Anderson again took the most extreme position, suggesting that OIB are actually derived from a shallow mantle reservoir. In contrast, other workers still require a lower mantle plume source, with relatively high 3He/4He ratio which must be preserved against convective homogenisation with the more radiogenic upper mantle. However, Seta et al. argued that such a distinction is easier to preserve than a primordial helium reservoir, because it only needs to be preserved for ca. 2 Byr, whereas a primordial reservoir must be preserved for 4.5 Byr. This argument is illustrated in more detail in Fig. 11.8.

Fig. 11.8. Plot of helium isotope ratios against time to show how uniform degassing, plus excess production of radiogenic helium in the MORB reservoir (due to U recycling) could give rise to similar present day signatures as the traditional variable degassing model (dashed lines). After Seta et al. (2001).

 

            Although this model can be represented numerically, this does not necessarily mean that it is a realistic earth model. Indeed it has several problems. Firstly, the new helium data from Iceland (R/RA = 38) suggest that some plume sources are less degassed than previously thought, placing tighter limits on the amount of degassing possible from this source. Furthermore, evidence for a relatively un-degassed mantle source is supported by new neon data (section 11.2.2). Finally, new data from some plume sites with low R/RA values suggest that some of these may be due to shallow mixing with radiogenic sources. Therefore, the average R/RA value of the plume source is higher than proposed by Anderson. It is concluded that a genuine primordial helium source is probably still present in the Earth. To the present author, the evidence now favours the core.

 

 

11.1.5  Crustal helium

 

Of the continental helium flux, 99% is radiogenic, and can be sustained by a U equivalent concentration of 6 ppm in the upper 8 km of the crust. This can also explain 50% of the continental heat flux. Hence, the other 50% of continental heat flow must be sub-continental, whereas less than 1% (primordial plus radiogenic) of the continental helium flux comes from the mantle. In this case it is clear that the continental crust is a boundary layer. Mantle-derived heat can be carried across it conductively, but mantle-derived helium only leaks through the crust in certain discrete areas. These are normally areas of active magmatism.

 

            Well-gas studies demonstrate the local nature of mantle helium transport through the crust. Oxburgh et al. (1986) showed that sedimentary basins which result from crustal loading, such as the Alpine Molasse basin, yield helium with very low R/RA values around 0.05, whereas sedimentary basins formed by extensional tectonics, such as the Rhine Graben and the Pannonian basin of Hungary, may yield helium with much higher R/RA values around unity (Fig. 11.9a). The huge ‘Panhandle’ gas field in the southern USA is particularly interesting. It is one of the world’s largest gas fields, and has helium contents of up to 2%. In the south, the reservoir is draped over uplifted Proterozoic)Paleozoic basement, and in this region R/RA values as low as 0.06 have been measured (Fig. 11.9b). In contrast, the northern part of the reservoir is in an area of recent igneous activity. Here, R/RA values up to 0.2 have been measured, corresponding to 2% MORB-type helium (Oxburgh et al., 1986).

Fig. 11.9. Histograms showing variation in R/RA values (on a log scale) in different types of sedimentary basin. After Oxburgh et al. (1986).

 

            More recently, helium analysis of geothermal fluids was used by Hilton et al. (1993b) and Hoke et al. (1994) to probe the width of the mantle melting zone behind the Andean subduction zone. Both studies revealed high R/RA values (indicative of a significant fraction of mantle helium) in the magmatic zone centred on the Altiplano. In contrast, R/RA values below 0.5 were found in the trench zone in front of the magmatic arc, and behind the Eastern Cordillera (Fig. 11.10). Hilton et al. found good agreement between helium analyses of geothermal fluids and phenocrysts from nearby volcanoes. However, because of the high altitude of these volcanoes, special rapid-crushing procedures were necessary to minimise contamination from a large in situ cosmogenic component. Hoke et al. attributed high R/RA values in the central section of the Andes to thinned lithosphere resulting from subduction erosion. The decrease in R/RA values behind the Eastern Cordillera probably marks the transition from hot thin lithosphere above the subduction zone to the thick cold lithosphere of the Brazilian shield.

Fig. 11.10. Variation of R/RA values across the Central Andes, compared with a 1/1 cross section showing proposed thinning of the lithosphere within the mantle wedge. ( " ) = gas sample; ( ! ) = water sample. After Hoke et al. (1994).

 

            It is now widely understood that heavy rare gases in OIB magmas are very susceptible to shallow-level contamination in magmatic systems (sections 11.2 – 11.4). However, recent work shows that helium isotope signatures in OIB may also be susceptible to such contamination processes. These processes can potentially be mistaken for the signature of deep mantle sources in OIB.

 

            For example, Hilton et al. (1993a) found a strong correlation between R/RA value and petrology in submarine volcanic glasses from the Lau back-arc basin, situated behind the Tongan arc. Basaltic samples from the centre of the basin had relatively high helium contents (up to 10 micro cc/g), and normal MORB-like R/RA values of 8 (Fig. 11.11). However, more differentiated glasses from just behind the magmatic arc had much lower helium contents (< 0.2 micro cc/g), and R/RA values as low as unity. Given this correlation between 3He/4He ratio, helium content, and petrology, Hilton et al. attributed the lower R/RA values in differentiated glasses to shallow level contamination, probably due to crustal assimilation by magmas which had been largely degassed of mantle helium.

Fig. 11.11. Plot of helium isotope composition (in vesicles) against the silica content of Lau Basin submarine volcanics, showing an inverse correlation. After Hilton et al. (1993a). Fig. 3

 

            Hilton et al. (1995) observed similar evidence in phenocryst phases from two different lava series on Heard Island, in the Kerguelen Archipelago. Phenocrysts from the Laurens Peninsula series had a plume-type helium signature with R/RA of 16 – 18. However, phenocrysts from the Big Ben series had MORB-like R/RA values in helium-rich samples, but lower R/RA values in helium-poor samples (Fig. 11.12a). In the latter sample suite, isotopic disequilibrium between olivine and cpx phenocrysts was observed. Therefore, Hilton et al. suggested that contamination with radiogenic helium occurred immediately before eruption, probably during phenocryst growth in shallow magma chambers.

Fig. 11.12. Plot of helium isotope ratio against abundance; showing, a) depressed R/RA values in helium-poor samples from Big Ben volcano, Heard Island; b) similar features in other OIB. After Hilton et al. (1995).

 

            Several other hot-spots with R/RA values lower than MORB and elevated Sr isotope ratios (up to 0.705) have been attributed to crustal recycling into the OIB source reservoir (Fig. 11.6). However, Hilton et al. observed that low R/RA values in these islands were also associated with low total helium contents (Fig. 11.12b). Hence, they suggested that the radiogenic helium signatures in these islands also result from contamination within the oceanic lithosphere, rather than sediment recycling into the deep mantle.

 

            A different scenario is seen in Samoan basalts, which have a unique combination of elevated R/RA and radiogenic Sr (Fig. 11.6). Peridotite xenoliths in these lavas have very radiogenic Sr isotope signatures, which are attributed to metasomatism by an EMII mantle component (section 6.5.2). However, helium isotope analysis of fluid inclusions from the xenoliths revealed high R/RA values of around 12 (Farley, 1995a). This was unexpected, since a recycled sediment component should have radiogenic helium with low R/RA values. However, the fluids also had C/3He ratios of ca. 3 H 109 which were typical of mantle values, and distinct from ratios of over 1011 seen in sediments (see below). The combined evidence suggests that the volatile component of the metasomatic fluid was derived from the deep mantle and only recently mixed with a volatile-poor melt of subducted sediment. Because subducted sediment accumulates radiogenic helium from uranium decay, the high R/RA value of the metasomatic fluid places limits on the mantle residence time of the sediment since subduction. Based on binary strontium–helium mixing calculations, Farley estimated a residence time of only 10 Myr, suggesting that the sediments were incorporated into the plume from the nearby Tongan trench.

 

 

11.1.6  Helium and volatiles

 

The comparison of helium isotope data with other rare gas tracers will be discussed below. However, helium isotope compositions can also be used to place important constraints on the interpretation of other volatile species, the most important of which are carbon dioxide and methane.

 

            Carbon fluxes in the Earth are difficult to constrain because of the reactivity of this element. However, Marty and Jambon (1987) argued that if the abundance of carbon could be tied to 3He in major mantle products such as MORB, helium fluxes might be usable as a measure of the carbon flux in a variety of environments. They collated C/3He data for MORB from various ocean basins, and found a relatively narrow range with an average C/3He ratio of ca. 2 H 109.

 

            In a parallel study, Jambon and Zimmerman (1987) showed that the C/3He ratios measured by heating MORB glass and by crushing of vesicles were similar (Fig. 11.13), suggesting that the measured ratios are indicative of the basaltic magma itself, and are not severely fractionated relative to one another during eruption. This is attributed to the similar solubilities of helium and carbon dioxide in basaltic magma. Taken together, these pieces of evidence suggest that the measured ratio is typical of the C/3He flux from the upper mantle on a world-wide scale.

Fig. 11.13. Plot of C/3He ratio against glass vesicularity for MORB samples. ( " ) = heated glass; ( ! ) = crushed vesicles. After Marty and Jambon (1987).

 

            O’Nions and Oxburgh (1988) took these deductions further by arguing that the oceanic upper mantle flux of C/3He could also be applied to mantle-derived volatile fluxes through the continental lithosphere. They examined as an example the Pannonian basin of Hungary, which has an R/RA value as high as 6, suggesting that mantle-derived helium comprises up to 90% of the total helium flux in parts of the basin. This is attributed to its extensional tectonic setting. Measurements of helium abundance for a major aquifer in the Pannonian basin were used by Martel et al. (1989) to estimate a mantle 3He flux of 8 H 104 atoms/m2/s for the basin as a whole, which is greater than the globally averaged oceanic flux. Hence, O’Nions and Oxburgh concluded that if extensional zones such as the Pannonian basin make up a significant fraction of the lithosphere, they must also make a major contribution to the global carbon inventory of the crust.

 

 

11.1.7  Helium and interplanetary dust

 

It is well established that high 3He/4He ratios in ocean floor sediments reflect the accumulation of inter-planetary dust particles (IDPs). However, the question of temporal variability in the IDP flux has only recently been examined (Takayanagi and Ozima, 1987). These authors studied 3He variability in a 10 m pelagic clay core from the Central Pacific and a 150 m nano-fossil ooze core from the South Atlantic. The former spanned 0 – 3 Myr, while the latter, with generally higher sedimentation rates, spanned 0 – 40 Myr. Sedimentation rates were determined in both cases by paleomagnetism, supplemented in the 3 Myr-old core with 10Be data (section 14.3.4). The observed range of 3He/4He ratios was attributed to mixing of 0.1 – 1 ppm of IDPs with terrestrial sediment (Fig. 11.14). However, the 3He content of IDPs is ten orders of magnitude higher than that of terrestrial sediment, so the IDP fraction totally dominates the 3He budget of the samples.

Fig. 11.14. Plot of 3He abundance against isotope ratio in pelagic clays from the Central Pacific, compared with mixing lines between cosmic dust and terrestrial sediments. After Takayanagi and Ozima (1987).

 

            In both cores studied by Takayanagi and Ozima, 3He contents were inversely correlated with sedimentation rate. The 3He deposition flux was therefore determined by multiplying the 3He content by the sediment mass accumulation rate (mass is used because ocean floor sediments undergo compaction after deposition). The results suggested flux variations over time, but did not display any overall trend. The average 3He flux over the past 40 Myr was estimated as 1.5 (" 1) H 10!15 cc/cm2/yr (at STP).

 

            Generally similar results were obtained by Farley (1995b) on a 22 m core of pelagic clay from the central North Pacific, spanning the past 72 Myr. During the Quaternary, the sedimentation rate was high, yielding a 3He flux of about 1.1 H 10!15 cc/cm2/yr, in good agreement with Takayanagi and Ozima (1987). However, in the deeper part of the core, the calculated 3He flux was lower, averaging about 0.5 H 10!15 cc/cm2/yr. It is not clear whether this represents a real variation in the interplanetary dust flux over time, or a reduction in the retentivity of 3He with depth. In a detailed study, Mkopadhyay et al. (2001) observed no 3He peak at the K–T boundary, indicating that the extra-terrestrial signals from iridium and helium are decoupled (Fig. 11.15). This was attributed to impact-induced vaporisation and outgassing of the K–T bolide. In contrast, Farley et al. (1998) discovered a spike of 3He in the late Eocene (35–36 Myr ago) which did correlate with the iridium signal (Fig. 11.15). This was attributed to a comet shower.

Fig. 11.15. Record of 3He abundances in Cretaceous to Tertiary ages sediments, showing a 3He peak in the late Eocene, but no peak at the K–T boundary. After Mkopadhyay et al. (2001).

 

            A more controversial question concerns the evidence for variation of the 3He flux during the glacial cycles of the Quaternary period. A detailed study in this time range was performed by Marcantonio et al. (1995) on a 4 m core of carbonate-rich sediment from the Central Pacific, spanning the last 200 kyr. After correcting for dilution by biogenic carbonate, their 3He/4He data lay on the same mixing line as that observed by Takayanagi and Ozima between terrigenous and IDP components. However, Marcantonio also determined initial excess 230Th activities on the same samples. Normalisation of 3He with respect to 230Th can remove the effects of variable sediment dilution, because 230Th is constantly produced in seawater from 234U, and is rapidly transported to the ocean floor by adsorption onto sinking particulate matter (section 12.3.3). When plotted against ages from oxygen isotope stratigraphy, 3He and 230Th showed strong co-variation, with peak signals during interglacial periods (Fig. 11.16). These peaks were attributed, not to variations of the IDP flux, but to intensified dissolution of carbonate during interglacial periods. Hence, from the ratio of 230Th activity to 3He content, an average 3He deposition flux of 0.96 H 10!15 cc/cm2/yr was determined for the past 200 kyr.

Fig. 11.16. Isotope stratigraphy of a carbonate-rich sediment from the Central Pacific, showing strong co-variation between 3He abundance ( ! ) and excess initial 230Th activity ( " ). After Marcantonio et al. (1995).

 

            Farley and Patterson (1995) performed a similar study of Quaternary 3He variation on a 9 m core of foram–nano-fossil ooze from the flank of the Mid Atlantic Ridge, spanning the period 250 – 450 kyr BP. 3He contents were inversely correlated with * 18O variations, which were interpreted as monitors of glacial–interglacial cycles. Similar results were found by Patterson and Farley (1998) in a Pacific Ocean core.

 

            The 100 kyr cycle of 3He abundance variations observed in these studies, if interpreted as a proxy of IDP flux variations, ties in with a proposal by Muller and MacDonald (1995, 1997) that Quaternary climate variations might be due to a 100 kyr cycle of variations in the Earth’s orbital inclination, causing periodic encounters with a cloud of IDPs which could partially block out solar radiation. Hence, Farley and Patterson speculated that the helium isotope data might be recording a causal relationship between IDP accumulation and climate. However, according to this model, interglacial periods should be characterised by the lowest 3He flux, whereas Farley and Patterson found the opposite relationship. Therefore, an alternative interpretation of the data (Marcantonio et al., 1995) is that climatically induced variations in sedimentation rate caused apparent variations in the 3He flux which could not be adequately corrected with the available age data for the core.

 

            Curiously enough, new modelling of the orbits of IDPs and the Earth (Kortenkamp and Dermott (1998) confirmed the 100 kyr periodicity of the IDP flux predicted by Muller and MacDonald, but attributed this periodicity to variations in the eccentricity of the Earth’s orbit. However, the predicted IDP flux variations were still anti-correlated with the ocean floor 3He record, and their magnitude was also judged to be too small to cause climatic cycles. 

 

            Confirmation that Quaternary 3He variations in ocean floor sediments are due to variations in sedimentation rate rather than the IDP flux was provided in new work (Marcantonio et al. (1996, 1999) on cores from the Atlantic and Indian Oceans respectively. In both studies 3He abundances were again correlated with 230Th data, which Marcantonio et al. attributed to variations in the amount of sediment focussing during different climatic periods. After using 230Th data to correct for these effects, constant 3He accumulation rates were found, with magnitudes of 0.8, 1.2 and 1.1 H 10!15 cc / cm2 /yr for the equatorial Pacific , Atlantic and Indian oceans respectively. Hence, it is concluded that the 3He flux has been essentially uniform and constant over the past 200 kyr. 

 

 

References

 

Aldrich, L. T. and Nier, A. O. (1948). The occurrence of He3 in natural sources of helium. Phys. Rev. 74, 1590)4.

 

Allegre, C. J., Moreira, M. and Staudacher, T. (1995). 4He/3He dispersion and mantle convection. Geophys. Res. Lett. 22, 2325–8.

 

Allegre, C. J., Sarda, P. and Staudacher, T. (1993). Speculations about the cosmic origin of He and Ne in the interior of the Earth. Earth Planet. Sci. Lett. 117, 229)33.

 

Allegre, C. J., Staudacher, T. and Sarda, P. (1986). Rare gas systematics: formation of the atmosphere, evolution and structure of the Earth’s mantle. Earth Planet. Sci. Lett. 81, 127–50.

 

Allegre, C. J., Staudacher, T., Sarda, P. and Kurz, M. (1983). Constraints on evolution of Earth’s mantle from rare gas systematics. Nature 303, 762)6.

 

Alvarez, L. W. and Cornog, R. (1939). Helium and hydrogen of mass 3. Phys. Rev. 56, 613.

 

Anderson, D. L. (1993). Helium-3 from the mantle: primordial signal or cosmic dust? Science 261, 170)6.

 

Anderson, D. L. (2001). A statistical test of the two reservoir model for helium isotopes. Earth Planet. Sci. Lett. 193, 77–82.

 

Burnard, P., Graham, D. and Turner, G. (1997). Vesicle-specific noble gas analyses of "popping rock": implications for primordial noble gases in Earth. Science 276, 568–71.

 

Butler, W. A., Jeffery, P. M., Reynolds, J. H. and Wasserburg, G. J. (1963). Isotopic variations in terrestrial xenon. J. Geophys. Res. 68, 3283)91.

 

Caffee, M. W., Hudson, G. B., Velsko, C., Alexander, E. C., Huss, G. R. and Chivas, A. R. (1988). Non-atmospheric noble gases from CO2 well gases. Lunar Planet. Sci. 19, 154–5 (abs).

 

Caffee, M. W., Hudson, G. B., Velsko, C., Huss, G. R., Alexander, E. C. and Chivas, A. R. (1999). Primordial noble gases from Earth’s mantle: identification of a primitive volatile component. Science 285, 2115–18.

 

Cerling, T. E. (1989). Dating geomorphologic surfaces using cosmogenic 3He. Quaternary. Res. 33, 148)56.

 

Clarke, W. B., Beg, M. A. and Craig, H. (1969). Excess 3He in the sea: evidence for terrestrial primordial helium. Earth Planet. Sci. Lett. 6, 213)20.

 

Clarke, W. B., Jenkins, W. J. and Top, Z. (1976). Determination of tritium by mass-spectrometric measurement of 3He. Int. J. Appl. Rad. Isot. 27, 515)22.

 

Coltice, N. and Ricard, Y. (1999). Geochemical observations and one layer mantle convection. Earth Planet. Sci. Lett. 174, 125–37.

 

Craig, H. and Lupton, J. E. (1976). Primordial neon, helium, and hydrogen in oceanic basalts. Earth Planet. Sci. Lett. 31, 369)85.

 

Craig, H. and Poreda, R. J. (1986). Cosmogenic 3He in terrestrial rocks: the summit lavas of Maui. Proc. Nat. Acad. Sci. USA 83, 1970)4.

 

Damon, P. E. and Kulp, L. (1958). Excess helium and argon in beryl and other minerals. Amer. Miner. 43, 433)59.

 

Dixon, E. T., Honda, M., McDougall, I., Campbell, I. H. and Sigurdsson, I. (2000). Preservation of near-solar neon isotopic ratios in Icelandic basalts. Earth Planet. Sci. Lett. 180, 309–24.

 

Fanale, F. P. (1971). A case for catastrophic early degassing of the Earth. Chem. Geol. 8, 79)105.

 

Farley, K. A. (1995a). Rapid cycling of subducted sediments into the Samoan mantle plume. Geology 23, 531–4.

 

Farley, K. A. (1995b). Cenozoic variations in the flux of interplanetary dust recorded by 3He in a deep-sea sediment. Nature 376, 153–6.

 

Farley, K. A. and Craig, H. (1994). Atmospheric argon contamination of ocean island basalt olivine phenocrysts. Geochim. Cosmochim. Acta 58, 2509)17.

 

Farley, K. A., Montanari, A., Shoemaker, E. M. and Shoemaker, C. S. (1998). Geochemical evidence for a comet shower in the late Eocene. Science 280, 1250–3.

 

Farley, K. A. and Patterson, D. B. (1995). A 100-kyr periodicity in the flux of extraterrestrial 3He to the sea floor. Nature 378, 600–3.

 

Farley, K. A. and Poreda, R. J. (1993). Mantle neon and atmospheric contamination. Earth Planet. Sci. Lett. 114, 325)39.

 

Fisher, D. E. (1971). Incorporation of Ar in East Pacific basalts. Earth Planet. Sci. Lett. 12, 321)4.

 

Fisher, D. E. (1983). Rare gases from the undepleted mantle? Nature 305, 298)300.

 

Fisher, D. E. (1985). Noble gases from oceanic island basalts do not require an undepleted mantle source. Nature 316, 716)18.

 

Fisher, D. E. (1986). Rare gas abundances in MORB. Geochim. Cosmochim. Acta 50, 2531)41.

 

Gerling, E. K., Mamyrin, B. A., Tolstikhin, I. N. and Yakovleva, S. S. (1971). Isotope composition of helium in some rocks. Geokhimiya 10, 1209)17.

 

Graf, T., Kohl, C. P., Marti, K. and Nishiizumi, K. (1991). Cosmic-ray-produced neon in Antarctic rocks. Geophys. Res. Lett. 18, 203)6.

 

Graham, D. W., Castillo, P. R., Lupton, J. E. and Batiza, R. (1996). Correlated He and Sr isotope ratios in South Atlantic near-ridge seamounts and implications for mantle dynamics. Earth Planet. Sci. Lett. 144, 491–503.

 

Hanyu, T. and Kaneoka, I. (1997). The uniform and low 3He/4He ratios of HIMU basalts as evidence for their origin as recycled materials. Nature 390, 273–6.

 

Harrison, D., Burnard, P. and Turner, G. (1999). Noble gas behaviour and composition in the mantle: constraints from the Iceland Plume. Earth Planet. Sci. Lett. 171, 199–207.

 

Hart, R., Dymond, J. and Hogan, L. (1979). Preferential formation of the atmosphere)sialic crust system from the upper mantle. Nature 278, 156)9.

 

Hart, R., Dymond, J., Hogan, L. and Schilling, J. G. (1983). Mantle plume noble gas component in glassy basalts from Reykjanes Ridge. Nature 305, 403)7.

 

Hart, R., Hogan, L. and Dymond, J. (1985). The closed-system approximation for evolution of argon and helium in the mantle, crust and atmosphere. Chem. Geol. (Isot. Geosci. Sect.) 52, 45)73.

 

Hennecke, E. W. and Manuel, O. K. (1975). Noble gases in CO2 well gas, Harding County, New Mexico. Earth Planet. Sci. Lett. 27, 346)55.

 

Hilton, D. R., Barling, J. and Wheller, G. E. (1995). Effect of shallow-level contamination on the isotope systematics of ocean-island lavas. Nature 373, 330–3.

 

Hilton, D. R., Gronvold, K, Macpherson, C. G. and Castillo, P. R. (1999). Extreme 3He/4He ratios in northwest Iceland: constraining the common component in mantle plumes. Earth Planet. Sci. Lett. 173, 53–60.

 

Hilton, D. R., Hammerschmidt, K., Loock, G. and Friedrichsen, H. (1993a). Helium and argon isotope systematics of the central Lau Basin and Valu Fa Ridge: evidence of crust/mantle interactions in a back-arc basin. Geochim. Cosmochim. Acta 57, 2819–41.

 

Hilton, D. R., Hammerschmidt, K., Teufel, S. and Friedrichsen, H. (1993b). Helium isotope characteristics of Andean geothermal fluids and lavas. Earth Planet. Sci. Lett. 120, 265–82.

 

Hilton, D. R., Thirlwall, M. F., Taylor, R. N., Murton, B. J. and Nichols, A. (2000). Controls on magmatic degassing along the Reykjanes Ridge with implications for the helium paradox. Earth Planet. Sci. Lett. 183, 43–50.

 

Hiyagon, H. (1994). Retention of solar helium and neon in IDPs in deep sea sediment. Science 263, 1257)9.

 

Hiyagon, H., Ozima, M., Marty, B., Zashu, S. and Sakai, H. (1992). Noble gases in submarine glasses from mid-ocean ridges and Loihi seamount: constraints on the early history of the Earth. Geochim. Cosmochim. Acta 56, 1301)16.

 

Hoke, L., Hilton, D. R., Lamb, S. H., Hammerschmidt, K. and Friedrichsen, H. (1994). 3He evidence for a wide zone of active mantle melting beneath the Central Andes. Earth Planet. Sci. Lett. 128, 341–55.

 

Honda, M., Reynolds, J. H., Roedder, E. and Epstein, S. (1987). Noble gases in diamonds: occurrences of solar-like helium and neon. J. Geophys. Res. 92, 12507)21.

 

Jacobsen, S. B. and Harper, C. L. (1996). Accretion and early differentiation history of the Earth based on extinct radionuclides. In: Basu, A. and Hart, S. R. (Eds.), Earth Processes: Reading the Isotopic Code. Geophys. Monograph 95, pp. 47–74.

 

Jambon, A. and Zimmermann, J. L. (1987). Major volatiles from an Atlantic MORB glass: a size fraction analysis. Chem. Geol. 62, 177)89.

 

Jeffrey, P. M. and Hagan, P. J. (1969). Negative muons and the isotopic composition of the rare gases in the Earth’s atmosphere. Nature 223, 1253.

 

Jephcoat, A. P. (1998). Rare-gas solids in the Earth's deep interior. Nature 393, 355–8.

 

Honda, M., McDougall, I., Patterson, D. B., Doulgeris, A. and Clague, D. A. (1991). Possible solar noble-gas component in Hawaiian basalts. Nature 349, 149)51.

 

Kaneoka, I. and Takaoka, N. (1980). Rare gas isotopes in Hawaiian ultramafic nodules and volcanic rocks: constraints on genetic relationships. Science 208, 1366–8.

 

Kellog, L. H. and Wasserburg, G. J. (1990). The role of plumes in mantle helium fluxes. Earth Planet. Sci. Lett. 99, 276–89.

 

Kennedy, B. M., Hiyagon, H. and Reynolds, J. H. (1990). Crustal neon: a striking uniformity. Earth Planet. Sci. Lett. 98, 277)86.

 

Kortenkamp, S. J. and Dermott, S. F. (1998). a 100,000-year periodicity in the accretion rate of interplanetary dust. Science 280, 874–6.

 

Kunz, J. (1999). Is there solar argon in the Earth’s mantle? Nature 399, 649–50.

 

Kunz, J., Staudacher, T. and Allegre, C. J. (1998). Plutonium-fission xenon found in the Earth’s mantle. Science  280, 877–80.

 

Kunz, W. and Schintlmeister, I. (1965). Tabellen der Atomkerne, Teil II, Kernreaktionen. Akademie-Verlag, 1022 p.

 

Kuroda, P. K. (1960). Nuclear fission in the early history of the Earth. Nature 187, 36)8.

 

Kurz, M. D. (1986a). Cosmogenic helium in a terrestrial igneous rock. Nature 320, 435)9.

 

Kurz, M. D. (1986b). In-situ production of terrestrial cosmogenic helium and some applications to geochronology. Geochim. Cosmochim. Acta 50, 2855)62.

 

Kurz, M. D., Gurney, J. J., Jenkins, W. J. and Lott, D. E. (1987). Helium isotopic variability within single diamonds from the Orapa kimberlite pipe. Earth Planet. Sci. Lett. 86, 57)68.

 

Kurz, M. D. and Jenkins, W. J. (1981). The distribution of helium in oceanic basalt glasses. Earth Planet. Sci. Lett. 53, 41)54.

 

Kurz, M. D., Jenkins, W. J. and Hart, S. R. (1982). Helium isotopic systematics of oceanic islands and mantle heterogeneity. Nature 297, 43)6.

 

Kurz, M. D., Meyer, P. S. and Sigurdsson, H. (1985). Helium isotopic systematics within the neovolcanic zones of Iceland. Earth Planet. Sci. Lett. 74, 291)305.

 

Kyser, T. K. and Rison, W. (1982). Systematics of rare gas isotopes in basic lavas and ultramafic xenoliths. J. Geophys. Res. 87, 5611)30.

 

Lal, D. (1987). Production of 3He in terrestrial rocks. Chem. Geol. (Isot. Geosci. Sect.) 66, 89)98.

 

Lal, D., Nishiizumi, K., Klein, J., Middleton, R. and Craig, H. (1987). Cosmogenic 10Be in Zaire alluvial diamonds: implications to 3He excess in diamonds. Nature 328, 139)41.

 

Lupton, J. E., (1983). Terrestrial inert gases: isotope tracer studies and clues to primordial components in the mantle. Ann. Rev. Earth Planet. Sci. 11, 371)414.

 

Lupton, J. E. and Craig. H. (1975). Excess 3He in oceanic basalts: evidence for terrestrial primordial helium. Earth Planet. Sci. Lett. 26, 133)9.

 

Mamyrin, B. A., Anufriyev, G. S., Kamenskiy, I. L. and Tolstikhin, I. N. (1970). Determination of the composition of atmospheric helium. Geochem. Int. 7, 498)505.

 

Mamyrin, B. A. and Tolstikhin, I. N. (1984). Helium Isotopes in Nature. Elsevier, 273 p.

 

Mamyrin, B. A., Tolstikhin, I. N., Anufriyev, G. S. and Kamenskiy, I. L. (1969). Anomalous isotopic composition of helium in volcanic gases. Dokl. Akad. Nauka SSSR 184, 1197)9.

 

Mamyrin, B. A., Tolstikhin, I. N. Anufriyev, G. S. and Kamenskiy, I. L. (1972). Isotopic composition of helium in Icelandic hot springs. Geokhimiya 11, 1396.

 

Marcantonio, F., Anderson, R. F., Stute, M., Kumar, N., Schlosser, P. and Mix, A. (1996). Extraterrestrial 3He as a tracer of marine sediment transport and accumulation. Nature 383, 705–7.

 

Marcantonio, F., Kumar, N., Stute, M., Anderson, R. F., Seidl, M. A., Schlosser, P. and Mix, A. (1995). A comparative study of accumulation rates derived by He and Th isotope analysis of marine sediments. Earth Planet. Sci. Lett. 133, 549–55.

 

Marcantonio, F., Turekian, K. K., Higgins, S., Anderson, R. F., Stute, M. and Schlosser, P. (1999). The accretion rate of extraterrestrial 3He based on the oceanic 230Th flux and the relation to os isotope variation over the past 200,000 years in an Indian Ocean core. Earth Planet. Sci. Lett. 170, 157–68.

 

Martel, D. J., Deak, J., Dovenyi, P., Horvath, F., O’Nions, R. K., Oxburgh, E. R., Stegena, L. and Stute, M. (1989). Leakage of helium from the Pannonian basin. Nature 342, 908)12.

 

Marti, K. and Craig, H. (1987). Cosmic-ray-produced neon and helium in the summit lavas of Maui. Nature 325, 335)7.

 

Marty, B. (1989). Neon and xenon isotopes in MORB: implications for the Earth)atmosphere evolution. Earth Planet. Sci. Lett. 94, 45)56.

 

Marty, B. and Jambon, A. (1987). C/3He in volatile fluxes from the solid Earth: implications for carbon geodynamics. Earth Planet. Sci. Lett. 83, 16)26.

 

Marty, B., Tolstikhin, I., Kamensky, I. L., Nivin, V., Balaganskaya, E. and Zimmermann, J.-L. (1998). Plume-derived rare gases in 380 Ma carbonatites from the Kola region (Russia) and the argon isotopic composition in the deep mantle. Earth Planet. Sci. Lett. 164, 179–192.

 

Matsuda, J., Murota, M. and Nagao, K. (1990). He and Ne isotopic studies on the extraterrestrial material in deep-sea sediments. J. Geophys. Res. 95, 7111)17.

 

Matsuda, J., Sudo, M., Ozima, M., Ito, K., Ohtaka, O. and Ito, E. (1993). Noble gas partitioning between metal and silicate under high pressures. Science 259, 788)90.

 

Merrihue, C. (1964). Rare gas evidence for cosmogenic dust in modern Pacific red clay. Ann. N. Y. Acad. Sci. 119, 351)67.

 

Meshik, A. P., Shukolyukov, Y. A. and Je$berger, E. K. (1995). Chemically fractionated fission xenon (CFF-Xe) on the Earth and in meteorites. In: Busso, M. et al. (Eds), Nuclei in the Cosmos III, Amer. Inst. Phys. Conf. Proc. 327, 603–6.

 

Moreira, M., Breddam, K., Curtice, J. and Kurz, M. D. (2001). Solar neon in the Icelandic mantle: new evidence for an undegassed lower mantle. Earth Planet. Sci. Lett. 185, 15–23.

 

Moreira, M., Kunz, J. and Allegre, C. J. (1998). Rare gas systematics in popping rock: isotopic and elemental compositions in the upper mantle. Science 279, 1178–81.

 

Morrison, P. and Pine, J. (1955). Radiogenic origin of the helium isotopes in rocks. Ann. N. Y. Acad. Sci. 62, 69)92.

 

Mukhopadhyay, S., Farley, K. A. and Montanari, A. (2001). A 35 Myr record of helium in pelagic limestones from Italy: implications for interplanetary dust accretion from the early Maastrichtian to the middle Eocene. Geochim. Cosmochim. Acta 65, 653–69.

 

Muller, R. A. and Macdonald, G. J. (1995). Glacial cycles and orbital inclination. Nature 377, 107–8.

 

Muller, R. A. and Macdonald, G. J. (1997). Glacial cycles and astronomical forcing. Science 277, 215–18.

 

Nagao, K., Takaoka, N. and Matsubayashi, O. (1979). Isotopic anomalies of rare gases in the Nigorikawa geothermal area, Hokkaido, Japan. Earth Planet. Sci. Lett. 44, 82)90.

 

Niedermann, S., Bach, W. and Erzinger, J. (1997). Noble gas evidence for a lower mantle component in MORBs from the southern East Pacific Rise: decoupling of helium and neon isotope systematics. Geochim. Cosmochim. Acta 61, 2697–715.

 

Nier, A. O. and Schlutter, D. J. (1990). Helium and neon isotopes in stratospheric particles. Meteoritics 25, 263)7.

 

O’Nions, R. K. and Oxburgh, E. R. (1983). Heat and helium in the Earth. Nature 306, 429)36.

 

O’Nions, R. K. and Oxburgh, E. R. (1988). Helium volatile fluxes and the development of continental crust. Earth Planet. Sci. Lett. 90, 331)47.

 

O’Nions, R. K. and Tolstikhin, I. N. (1994). Behaviour and residence times of lithophile and rare gas tracers in the upper mantle. Earth Planet. Sci. Lett. 124, 131–8.

 

Oxburgh, E. R., O’Nions, R. K. and Hill, R. I. (1986). Helium isotopes in sedimentary basins. Nature 324, 632)5.

 

Ozima, M. and Igarashi, G. (1989). Terrestrial noble gases: constraints and implications on atmospheric evolution. In: Atreya, S. K., Pollack, J. B. and Matthews, M. (Eds.) Origin and Evolution of Planetary and Satellite Atmospheres. Univ. Arizona Press, pp. 306–27.

 

Ozima, M. and Igarashi, G. (2000). The primordial noble gases in the Earth: a key constraint on Earth evolution models. Earth Planet. Sci. Lett. 176, 219–32.

 

Ozima, M., Podosek, F. A. and Igarashi, G. (1985). Terrestrial xenon isotope constraints on the early history of the Earth. Nature 315, 471)4.

 

Ozima, M. and Zashu, S. (1983). Primitive helium in diamonds. Science 219, 1067)8.

 

Ozima, M. and Zashu, S. (1988). Solar-type Ne in Zaire cubic diamonds. Geochim. Cosmochim. Acta 52, 19)25.

 

Ozima, M. and Zashu, S. (1991). Noble gas state of the ancient mantle as deduced from noble gases in coated diamonds. Earth Planet. Sci. Lett. 105, 13)27.

 

Patterson, D. B. and Farley, K. A. (1998). Extraterrestrial 3He in seafloor sediments: evidence for correlated 100 kyr periodicity in the accretion rate of interplanetary dust, orbital parameters, and Quaternary climate. Geochim. Cosmochim. Acta 62, 3669–82.

 

Patterson, D. B., Honda, M. and McDougall, I. (1990). Atmospheric contamination: a possible source for heavy noble gases in basalts from Loihi Seamount, Hawaii. Geophys. Res. Lett. 17, 705)8.

 

Pepin, R. O. (1991). On the origin and early evolution of terrestrial planet atmospheres and meteoritic volatiles. Icarus 92, 2–79.

 

Pepin, R. O. (1997). Evolution of Earth’s noble gases: consequences of assuming hydrodynamic loss driven by giant impact. Icarus 126, 148–56.

 

Pepin, R. O. (1998). Isotopic evidence for a solar argon component in the Earth’s mantle. Nature 394, 664–7.

 

Pepin, R. O. and Signer, P. (1965). Primordial rare gases in meteorites. Science 149, 253)65.

 

Phinney, D., Tennyson, J. and Frick, U. (1978). Xenon in CO2 well gas revisited. J. Geophys. Res. 83, 2313)19.

 

Porcelli, D. and Halliday, A. N. (2001). The core as a possible source of mantle helium. Earth Planet. Sci. Lett. 192, 45–56.

 

Porcelli, D. and Wasserburg, G. J. (1995a). Mass transfer of xenon through a steady-state upper mantle. Geochim. Cosmochim. Acta 59, 1991–2007.

 

Porcelli, D. and Wasserburg, G. J. (1995b). Mass transfer of helium, neon, argon, and xenon through a steady-state upper mantle. Geochim. Cosmochim. Acta 59, 4921–37.

 

Poreda, R. J. and Farley, K. A. (1992). Rare gases in Samoan xenoliths. Earth Planet. Sci. Lett. 113, 129)44.

 

Reynolds, J. H. (1960). Determination of the age of the elements. Phys. Rev. Lett. 4, 8)10.

 

Reynolds, J. H. (1963). Xenology. J. Geophys. Res. 68, 2939)56.

 

Rutherford, E. (1906). The production of helium from radium and the transformation of matter. In: Rutherford, E., Radioactive Transformations. Yale Univ. Press, pp. 187)93.

 

Sarda, P., Staudacher, T. and Allegre, C. J. (1988). Neon isotopes in submarine basalts. Earth Planet. Sci. Lett. 91, 73)88.

 

Sarda, P., Staudacher, T., Allegre, C. J. and Lecomte, A. (1993). Cosmogenic neon and helium at Reunion: measurement of erosion rate. Earth Planet. Sci. Lett. 119, 405)17.

 

Sasada, T., Hiyagon, H., Bell, K. and Ebihara, M. (1997). Mantle-derived noble gases in carbonatites. Geochim. Cosmochim. Acta 61, 4219–28.

 

Schmalzl, J., Houseman, G. A. and Hansen, U. (1995). Mixing properties of 3-dimensional (3-D) stationary convection. Phys. Fluids 7, 1027–33.

 

Schwartzman, D. W. (1973). Argon degassing models of the Earth. Nature Phys. Sci. 245, 20)1.

 

Seta, A., Matsumoto, T. and Matsuda, J.-I. (2001). Concurrent evolution of 3He/4He ratio in the Earth’s mantle reservoirs for the first 2 Ga. Earth Planet. Sci. Lett. 188, 211–19.

 

Sheldon, W. R. and Kern, J. W. (1972). Atmospheric helium and geomagnetic field reversals. J. Geophys. Res. 77, 6194)201.

 

Staudacher, T. (1987). Upper mantle origin for Harding County well gases. Nature 325, 605)7.

 

Staudacher, T. and Allegre, C. J. (1982). Terrestrial xenology. Earth Planet. Sci. Lett. 60, 389)406.

 

Staudacher, T. and Allegre, C. J. (1988). Recycling of oceanic crust and sediments: the noble gas subduction barrier. Earth Planet. Sci. Lett. 89, 173)83.

 

Staudacher, T., Kurz, M. D. and Allegre, C. J. (1986). New noble-gas data on glass samples from Loihi Seamount and Hualalai and on dunite samples from Loihi and Reunion Island. Chem. Geol. 56, 193)205.

 

Staudacher, T., Sarda, P. and Allegre, C. J. (1990). Noble gas systematics of Reunion Island, Indian Ocean. Chem. Geol. 89, 1)17.

 

Staudacher, T., Sarda, P., Richardson, S. H., Allegre, C. J., Sagna, I. and Dmitriev, L. V. (1989). Noble gases in basalt glasses from a Mid-Atlantic Ridge topographic high at 14 oN: geodynamic consequences. Earth Planet. Sci. Lett. 96, 119)33.

 

Takayanagi, M. and Ozima, M. (1987). Temporal variation of 3He/4He ratio recorded in deep-sea sediment cores. J. Geophys. Res. 92, 12 531–8.

 

Tolstikhin, I. N., Mamyrin, B. A., Khabarin, L. V. and Erlikh, E. N. (1974). Isotopic composition of helium in ultrabasic xenoliths from volcanic rocks of Kamchatka. Earth Planet. Sci. Lett. 22, 75)84.

 

Tolstikhin, I. N. and O’Nions, R. K. (1996). Some comments on isotopic structure of terrestrial xenon. Chem. Geol. 129, 185–99.

 

Trieloff, M., Kunz, J., Clague, D. A., Harrison, D. and Allegre, C. J. (2000). The nature of pristine noble gases in mantle plumes. Science 288, 1036–8.

 

Turekian, K. K. (1964). Outgassing of argon and helium from the Earth. In: Brancazio, P. and Cameron, A. G. W. (Eds), The Origin and Evolution of Atmospheres and Oceans. Wiley, pp. 74)83.

 

Valbracht, P. J., Staudacher, T., Malahoff, A. and Allegre, C. J. (1997). Noble gas systematics of deep rift zone glasses from Loihi Seamount, Hawaii. Earth Planet. Sci. Lett. 150, 399–411.

 

van Keken, P. E. and Ballentine, C. J. (1998). Whole-mantle versus layered mantle convection and the role of a high-viscosity lower mantle in terrestrial volatile evolution. Earth Planet. Sci. Lett. 156, 19–32.

 

van Keken, P. E. and Ballentine, C. J. (1999). Dynamical models of mantle volatile evolution and the role of phase transitions and temperature-dependent rheology. J. Geophys. Res. 104, 7137–51.

 

Wernicke, R. S. and Lippolt, H. J. (1993). Botryoidal hematite from the Schwarzwald (Germany): heterogeneous uranium distributions and their bearing on the helium dating method. Earth Planet. Sci. Lett. 114, 287)300.

 

Wetherill, G. W. (1953). Spontaneous fission yields from uranium and thorium. Phys. Rev. 82, 907)12.

 

Wetherill, G. W. (1954). Variations in the isotopic abundances of neon and argon extracted from radioactive materials. Phys. Rev. 96, 679)83.

 

Zadnik, M. G., Smith, C. B., Ott, U. and Begemann, F. (1987). Crushing of a terrestrial diamond: 3He/4He higher than solar meteorites. Meteoritics 22, 540)1.