11.4 Xenon

Xenon is a heavy rare gas with nine stable isotopes. Reynolds (1960) first demonstrated variations in the meteoritic abundance of ¹²⁹Xe, produced from the extinct nuclide ¹²⁹I (section 15.3.1). In addition, the four heaviest isotopes are fission products of both ²³⁸U (Wetherill, 1953), and the extinct nuclide ²⁴⁴Pu (Kuroda, 1960). It is convenient to ratio xenon isotope abundances against ¹³⁰Xe, which is non-radiogenic and is also shielded from spallation production. Hence, xenon isotope data are usually presented on a plot of either ¹³⁴Xe/¹³⁰Xe or ¹³⁶Xe/¹³⁰Xe (uranogenic and plutogenic xenon) against ¹²⁹Xe/¹³⁰Xe (iodogenic xenon), as shown in Fig. 11.35.

11.4.1 Iodogenic xenon

Because Xe is a tracer for two extinct nuclides, meteorite ‘xenology’ is a powerful tool for studying the condensation of the solar system (Reynolds, 1963). However, ‘terrestrial xenology’ is also a powerful tool for understanding terrestrial differentiation (Staudacher and Allegre, 1982). The first evidence for excess ¹²⁹Xe in the Earth (relative to atmospheric xenon) was found in CO₂ well gases from Harding County, New Mexico (Butler et al., 1963). This evidence has such far-reaching implications that numerous subsequent studies have been devoted to Harding County well gases (e.g. Hennecke and Manuel, 1975; Phinney et al., 1978; Staudacher, 1987).

Studies of xenon isotope data from granitic rocks (Butler et al., 1963) showed that unlike the heavy xenon isotopes, ¹²⁹Xe is not generated in significant amounts (relative to non-radiogenic xenon) by fission or neutron activation reactions (Fig. 11.35). Therefore Butler et al. concluded that the excess ¹²⁹Xe in mantle-derived gases must have been due to decay, soon after the formation of the Earth, of extinct ¹²⁹I. The presence of this extinct nuclide (t_1/2 = 16 Myr) in the Earth demonstrates that its accretion must have occurred within a few million years of meteorites, which also commonly display ¹²⁹Xe anomalies (section 15.3.1).

Fig. 11.35. Plot of fissiogenic ¹³⁴Xe/¹³⁰Xe against iodogenic ¹²⁹Xe/¹³⁰Xe for granites and well gases analysed prior to 1978. Modified after Staudacher and Allegre (1982).

Further advances in terrestrial xenology required the well-gas data to be put into the wider perspective of major terrestrial reservoirs. Technical developments, allowing the xenon analysis of submarine glasses, made this possible. In early work, Staudacher and Allegre (1982) found correlated enrichments of ¹²⁹Xe and other heavy xenon isotopes in MORB glasses. This was confirmed by Staudacher (1987), who showed that the most ¹²⁹Xe-enriched well-gas analyses lay on the same correlation line as MORB (Fig. 11.36a). Well-gas and MORB compositions were also coherent on a plot of ⁴⁰Ar/³⁶Ar against ¹²⁹Xe/¹³⁰Xe (Fig. 11.36b). This provides strong evidence that the Harding County well gases sample an upper mantle reservoir similar to the depleted MORB source.

Fig. 11.36. Plots of (a) ¹³⁶Xe/¹³⁰Xe, and (b) ⁴⁰Ar/³⁶Ar against ¹²⁹Xe/¹³⁰Xe for MORB glasses, compared to other terrestrial components. After Staudacher (1987).

Staudacher and Allegre (1982) argued that the excess ¹²⁹Xe in MORB (relative to a Bulk Earth with atmospheric xenon) could be explained by degassing of the upper mantle very early in Earth history. Primordial rare gases including ¹³⁰Xe would also have been lost to the atmosphere, while ¹²⁹I still remained in the mantle. The high I/Xe ratios generated by this process would allow the small amount of ¹²⁹I remaining in the upper mantle to generate measurable excesses in ¹²⁹Xe after its decay. Hence, upper mantle degassing must have occurred much earlier than lithophile depletion (by crust formation), although these two processes seem to affect roughly similar fractions of the mantle. These data will be discussed further in section 11.4.2.

Additional evidence for the xenon isotope evolution of the upper mantle has been obtained from the analysis of ‘coated’ diamonds (Ozima and Zashu, 1991). The coats of these diamonds contain relatively large rare gas contents, and are thus suited to isotopic analysis. Ozima and Zashu found xenon isotope ratios identical to the MORB correlation line (Fig. 11.37), suggesting that the same evolution processes gave rise to the mantle sources of MORB and diamonds. In the light of neon and argon isotope evidence (sections 11.2 and 11.3), the MORB array is best attributed to atmospheric contamination of magmas with radiogenic xenon. Similarly, Ozima and Zashu attributed the diamond array to mixing between a radiogenic xenon source and atmospheric xenon contamination (represented by the cores).

Fig. 11.37. Plot of ¹³⁶Xe/¹³⁰Xe against ¹²⁹Xe/¹³⁰Xe for coated diamonds, compared with the MORB correlation line of Staudacher (1987). ( " ) = cores; ( ! ) = coats. Dashed line shows result of uranogenic production from a 2 Byr-old source. After Ozima and Zashu (1991).

Analysis of xenon isotope compositions in glasses from Loihi Seamount and the Reykjanes Ridge gave results within error of atmospheric rare gases, but the origin of this signature was widely disputed. Allegre et al. (1983) and Hart et al. (1983) attributed it to a less-degassed mantle source, in which radiogenic ¹²⁹Xe and ⁴⁰Ar are swamped by large primordial rare gas contents. Alternatively, Ozima et al. (1985) proposed that atmospheric xenon is recycled into the deep mantle, whereas Patterson et al. (1990) argued that the xenon isotope ratios of Loihi glasses (together with their argon signatures) are due to atmospheric contamination, introduced directly into plume magmas from seawater.

Poreda and Farley (1992) suggested that the relatively shallow ocean depths at which Loihi basalts were erupted allowed much of the magmatic gas content to be lost, making them very susceptible to contamination by rare gases in seawater. To avoid these problems, Poreda and Farley analysed two suites of very volatile-rich harzburgite xenoliths from the Samoan hot-spot, which are more resistant to atmospheric contamination. The helium and neon isotope signatures of these rocks were somewhat more radiogenic than Loihi (section 11.2.3), but still well below the levels in MORB. Therefore, these samples clearly contain a large plume-derived component. Xenon isotope ratios were found to be elevated by up to 6% relative to the atmospheric point, providing the first evidence for a non-atmospheric xenon signature from the lower mantle.

Based on this evidence for non-atmospheric xenon in plumes, Porcelli and Wasserburg (1995a) extended the steady-state helium model of Kellogg and Wasserburg (1990) to xenon. They argued that the radiogenic signature of upper mantle xenon (relative to the atmosphere) is not a residue of early degassing of the Earth, as previously proposed, but is the result of input from the lower mantle, along with minor in situ production. The supply of rare gases from lower to upper mantle was attributed to mass transfer in plumes. A fraction of the plume is degassed at the hot-spot, but the bulk is mixed into the upper mantle. In this mass transfer model, rare gases are not fractionated from one another, so they are all thought to have the same upper mantle residence time of around 1.4 Byr.

Since the upper mantle is argued to be in steady state, it bears no memory of its early history. Hence, ¹²⁹Xe/¹³⁰Xe variations in the upper mantle (attributed to extinct iodine) are explained solely by mixing of lower mantle and atmospheric xenon. The atmosphere is at the unradiogenic end of the MORB array, and must therefore have been outgassed from the Earth before a substantial amount of the iodine budget had decayed. However, Porcelli and Wasserburg suggested that some of the atmosphere was probably contributed by late accretion of volatile-rich material after degassing of the deep earth. In contrast to this composition, the lower mantle is predicted to be at the radiogenic end of the MORB range.

Experimental confirmation that the xenon isotope composition of the plume source has a radiogenic signature resembling MORB came from new analyses of Loihi and Iceland samples (Trieloff et al., 2000). These were the first Loihi data to be clearly distinct from the atmospheric composition, and extended half way up the MORB array (Fig. 11.38). These data therefore confirm the similarities between rare gas signatures of the plume source and the upper mantle, providing strong support for the steady state model of Porcelli and Wasserburg (1995a). However, the evidence continues to support the two-reservoir model for terrestrial rare gases by maintaining the need for a primordial rare gas reservoir in the deep Earth, at the same time as a strongly degassed upper mantle. It remains unclear whether the primordial rare gas reservoir is located in the deep mantle or the core. However, Jephcoat (1998) suggested that heavy rare gases could actually be in the solid phase at lower mantle pressures.

Fig. 11.38. Comparison of new data from Loihi ( " ) and Iceland ( ! ) with the location of the MORB array ( + ) on the xenon 3-isotope plot. Modified after Trieloff et al. (2000).

11.4.2 Fissiogenic xenon

Variations in fissiogenic xenon are more complex than those of iodogenic xenon because there are components from fission of both extinct plutonium and extant uranium. The early work by Butler et al. (1963) and Hennecke and Manuel (1975) suggested that some fissiogenic xenon in Harding County well gases was plutogenic. However, more precise analysis by Phinney et al. (1978) showed that the excess abundances of ¹³¹Xe, ¹³²Xe and ¹³⁴Xe relative to ¹³⁶Xe in well gas are a better match to the calculated isotope production from spontaneous fission of ²³⁸U rather than ²⁴⁴Pu (Fig. 11.39).

Fig. 11.39. Plot of excess abundances of xenon isotopes in well gas relative to the atmosphere, ratioed against excess ¹³⁶Xe. The data ( ! ) are compared with modelled production of fissiogenic xenon from ²³⁸U and ²⁴⁴Pu. After Phinney et al. (1978).

The heavy xenon isotope signatures in diamonds can also be used to test for plutogenic or uranogenic production, as demonstrated above for well gases. Ozima and Zashu (1991) showed that production from ²³⁸U yields the best fit to the xenon data on diamond coats, which in turn places a tight constraint on the time of formation of these coats. Uranogenic xenon (unlike iodogenic and plutogenic xenon) grows in the mantle over Earth history. Therefore, if the rare gases in diamonds have been isolated from the silicate upper mantle for the past few billion years, they should have developed less ¹³⁶Xe than young MORB samples (dashed line in Fig. 11.34). The fact that MORB, diamond coats, and well gases all lie on the same array (if not a coincidence) suggests that all of these samples were formed relatively recently in geological time. This presents a problem, since other studies have indicated ancient formation ages for diamonds (section 4.2.1). However, it is possible that old diamond cores were overgrown by younger coats shortly before or during kimberlite magmatism (ca. 100 Myr ago).

Ozima et al. (1985) showed that fissiongenic xenon data can be evaluated in a more quantitative manner on a plot in which two different fissiogenic xenon isotopes are ratioed against ¹³⁰Xe. Either ¹³¹Xe, ¹³²Xe or ¹³⁴Xe can be compared with ¹³⁶Xe, but the latter two isotopes are usually compared because they have the best analytical precision. These isotopes are plotted in Fig. 11.40, from a re-evaluation of published data by Tolstikhin and O’Nions (1996). The errors on the MORB data available at the time this plot was drawn were too large to reliably resolve these components, but it appeared that MORB xenon resembles well gas xenon in having a uranogenic signature. More recently, xenon isotope analysis of carbonatites from Brazil and Canada gave a signature that was distinctly uranogenic (Sasada et al., 1997). However, new analyses of MORB by Kunz et al. (1998) suggested a mixture of uranogenic and plutogenic xenon.

Fig. 11.40. Plot of two fissiongenic xenon isotopes, ratioed to ¹³⁰Xe, to compare data for well gas ( ! ) and MORB ( " ) with alternative uranogenic and plutogenic production routes. The * ¹³⁶Xe diagram above shows the 95% confidence envelope for the best-fit regression of MORB data (dashed lines), relative to the alternative production routes. After Tolstikhin and O’Nions (1996).

The evolution of fissiogenic and iodogenic xenon in the steady-state model of Porcelli and Wasserburg (1995a) is shown in Fig. 11.41. In response to iodine and plutonium decay, it is thought that lower mantle xenon evolved from an initial solar or planetary composition to the hypothetical composition of present day lower mantle ( " plume source). In the steady state model, xenon escaping from the lower mantle is supplemented by uranium decay in the upper mantle, and then mixed with subducted atmospheric xenon, to reach the upper mantle composition (! Depleted mantle). The MORB xenon array is then attributed to mixing between D and the atmosphere, due to contamination of magmas at the sea floor.

Plume sources should define a similar array of gentler slope, between the atmosphere and P ( " plume source ). However, this prediction cannot yet be tested reliably, due to the difficulty of obtaining definitive non-atmospheric xenon signatures from hot spots. The available data for plume sources (Fig. 11.38) do not appear to support the model, since all data appear to be colinear with the MORB array. However, two processes may act to homogenise the plume and MORB signatures: firstly ongoing contamination of the upper mantle by plume material; and secondly, contamination of ascending plumes by entrainment of upper mantle xenon. It should also be borne in mind that many plume sources (e.g. Samoa) contain recycled lithospheric material, almost certainly including subducted atmospheric xenon.

Fig. 11.41. Conceptual model showing the evolution of fissiogenic xenon in the Earth, subdivided into plutogenic and uranogenic components. ( " ) = hypothetical compositions. After Porcelli and Wasserburg (1995a).

Further uncertainties in fissiogenic Xe systematics arise from the findings of Meshik et al. (1995), who suggested that variations in the relative abundances of heavy xenon isotopes could be due to open-system conditions during the beta decay of uranium fission products. High temperature conditions could allow preferential escape of some short-lived intermediates, affecting the isotopic composition of xenon isotopes, which are the final decay products. In the light of all the above problems, fissiogenic xenon systematics remain in considerable doubt.

11.4.3 Solar xenon

The non-radiogenic light isotopes of xenon are potentially very useful for unravelling the mysteries of terrestrial xenology. These isotopes should be usable in the same way as the non-radiogenic neon isotopes to assess the relationship between atmospheric xenon and primordial xenon in the Earth. However, because of the low abundances of the light isotopes, analytical errors have prevented their application until fairly recently.

Well gas ¹²⁸Xe data presented by Caffee et al. (1988) provided the first evidence for a correlation between iodogenic ¹²⁹Xe/¹³⁰Xe and non-radiogenic ¹²⁸Xe/¹³⁰Xe ratios. This implies that the atmosphere and the (upper) mantle reservoir sampled by well gases exhibit relative fractionation of xenon isotopes, in a manner analogous to neon. These systematics were investigated in more detail by Jacobsen and Harper (1996), Tolstikhin and O’Nions (1996) and Caffee et al. (1999), as illustrated in Fig. 11.42. By showing that the Earth’s mantle is slightly enriched in ¹²⁸Xe relative to the atmosphere, the data support the model of Pepin (1991), who argued that the atmosphere has undergone mass fractionation of xenon relative to an original solar composition.

Fig. 11.42. Plot of non-radiogenic ¹²⁸Xe/¹³⁰Xe against iodogenic ¹²⁹Xe/¹³⁰Xe, showing the evolution of mantle and atmospheric xenon and their mixing in well gases. Modified after Jacobsen and Harper (1996).

It is generally agreed that the compositions of solar xenon, the present-day Earth atmosphere, and well gases in Fig. 11.42 can be related by mass fractionation and by iodine decay, but there is some disagreement about the order of events. This author proposes to follow the ‘classical’ view that the order of events was 1): accretion of the Earth, 2): early degassing, 3): atmospheric fractionation, and 4): iodine decay in the deep Earth.

Ozima and Igarashi (1989) suggested that mass fractionation occurred during accretion (and therefore before degassing). However, the distinct solar composition of the Earth’s mantle seems to rule out this model. Similarly, Caffee et al. (1999) suggested that substantial decay of iodine occurred before accretion. However, this would require extreme variation in the I/Xe ratios of accreted components in order to create large inherited xenon isotope heterogeneity between the mantle and atmosphere. This also seems unlikely. Therefore, the best explanation (as shown in Fig. 11.41) is that the atmosphere was degassed after only minor iodine decay in the Earth, and then underwent mass fractionation.

The analysis of additional non-radiogenic light xenon isotopes by Caffee et al. (1999) provided further support for the ‘solar xenon’ model by revealing a correlation between ¹²⁶Xe/¹³⁰Xe and ¹²⁸Xe/¹³⁰Xe (Fig. 11.43). This array, formed entirely by non-radiogenic xenon isotopes, trends from the atmosphere composition towards the solar wind composition. This confirms that the Earth’s mantle has a solar xenon signature, and continues to suggest that the atmosphere underwent mass fractionation of xenon, probably from a combination of giant impact and solar radiation (section 11.2.2).

Fig. 11.43. Plot of non-radiogenic xenon isotope systematics, showing that well gases lie on a mixing line between atmospheric and solar xenon. After Caffee et al. (1999).

In conclusion, much progress has been made in understanding terrestrial xenology. The overall pattern of early Earth degassing, the evolution of the plume source, and upper mantle buffering now constitute a fairly robust model. However, many details remain to be worked out.

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