Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
Phil. Trans. R. Soc. A (2008) 366, 4183–4203
doi:10.1098/rsta.2008.0150
Published online 30 September 2008
What CO2 well gases tell us about the origin of
noble gases in the mantle and their
relationship to the atmosphere
B Y C HRIS J. B ALLENTINE *
AND
G REG H OLLAND
School of Earth, Environmental and Atmospheric Sciences,
University of Manchester, Oxford Road, Manchester M13 9PL, UK
Study of commercially produced volcanic CO2 gas associated with the Colorado Plateau,
USA, has revealed substantial new information about the noble gas isotopic composition
and elemental abundance pattern of the mantle. Combined with published data from midocean ridge basalts, it is now clear that the convecting mantle has a maximum 20Ne/ 22Ne
isotopic composition, indistinguishable from that attributed to solar wind-implanted
(SWI) neon in meteorites. This is distinct from the higher 20Ne/ 22Ne isotopic value
expected for solar nebula gases. The non-radiogenic xenon isotopic composition of the well
gases shows that 20 per cent of the mantle Xe is ‘solar-like’ in origin, but cannot resolve
the small isotopic difference between the trapped meteorite ‘Q’-component and solar Xe.
The mantle primordial 20Ne/132Xe is approximately 1400 and is comparable with the
upper end of that observed in meteorites. Previous work using the terrestrial 129I–129Xe
mass balance demands that almost 99 per cent of the Xe (and therefore other noble gases)
has been lost from the accreting solids and that Pu–I closure age models have shown this to
have occurred in the first ca 100 Ma of the Earth’s history. The highest concentrations of
Q-Xe and solar wind-implanted (SWI)-Ne measured in meteorites allow for this loss and
these high-abundance samples have a Ne/Xe ratio range compatible with the ‘recycled-aircorrected’ terrestrial mantle. These observations do not support models in which the
terrestrial mantle acquired its volatiles from the primary capture of solar nebula gases
and, in turn, strongly suggest that the primary terrestrial atmosphere, before isotopic
fractionation, is most probably derived from degassed trapped volatiles in accreting material.
By contrast, the non-radiogenic argon, krypton and 80 per cent of the xenon in the
convecting mantle have the same isotopic composition and elemental abundance pattern
as that found in seawater with a small sedimentary Kr and Xe admix. These mantle heavy
noble gases are dominated by recycling of air dissolved in seawater back into the mantle.
Numerical simulations suggest that plumes sampling the core–mantle boundary would be
enriched in seawater-derived noble gases compared with the convecting mantle, and
therefore have substantially lower 40Ar/ 36Ar. This is compatible with observation. The
subduction process is not a complete barrier to volatile return to the mantle.
Keywords: carbon dioxide; rare gases; inert gases; convection
* Author for correspondence ([email protected]).
One contribution of 14 to a Discussion Meeting Issue ‘Origin and differentiation of the Earth: past
to present’.
4183
This journal is q 2008 The Royal Society
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
4184
C. J. Ballentine and G. Holland
1. Introduction
The mechanism, timing and origin of the formation of the Earth’s atmosphere
and oceans remain enigmatic. While this is ultimately responsible for creating
the conditions for life and the tectonic character of our own planet, small
differences in planetary volatile acquisition, in the case of Venus and Mars,
have resulted in totally different planet surfaces. In particular, the origin of
the inner planetary exospheres has variably been attributed to volatiles
trapped in accreting material, those from a residual solar nebula or those
deposited as a late cometary veneer (e.g. Dauphas et al. 2000; Dauphas 2003).
It has been clear for some time from both the major gas species and its inert
or ‘noble’ gas isotopic composition that the Earth’s atmosphere today is not a
simple relic of primary nebula gas capture (Brown 1949; Suess 1949), nor
indeed of degassed infalling accretionary or later cometary material. The noble
gas isotopic composition and elemental abundance pattern of the modern
atmosphere alone require significant primary atmosphere processing and loss
(e.g. Porcelli & Pepin 2004). Nevertheless, identifying the primary atmosphere
origin and composition is critical in being able to develop a quantitative
understanding of the mechanism and process of early planetary atmosphere
modification. The mechanism and magnitude of planetary gas loss itself
provides, in turn, stringent mass-balance constraints on the accretionary origin
of the volatiles.
One of the most important developments in our understanding of planetary
accretion from the early solar nebula is modelling work that suggests that
multiple Mars-sized bodies impacting on an accreting planet are a natural part of
the accretionary process (Wetherill 1994; Morbidelli 2002). In the more detailed
giant impactor simulations that indicate the extent of planetary-scale
vaporization, melting and material dispersal before re-accretion (Canup 2004,
2008), it is perhaps a question of whether or not we would expect to find any
volatiles trapped within the solid planets at all. However, we do, and an
important additional clue to planetary volatile origin lies in the isotopic and
elemental composition of the noble gases degassing from the solid Earth today.
We briefly review our current understanding of noble gases in the mantle
together with recent advances made since the comprehensive reviews by Graham
(2002) Hilton et al. (2002), Pepin & Porcelli (2002), Porcelli & Ballentine (2002)
and Porcelli & Pepin (2004). In the light of recent results, we discuss what
noble gases tell us about the planetary accretionary process, the subsequent
evolution of the Earth’s interior and, indeed, what relationship they have with
the present-day atmosphere.
2. Helium: the case for an accretionary volatile reservoir in the
terrestrial mantle
One of the most important observations is that of deep-Earth 3He degassing into the
oceans from mid-ocean ridges (Clarke et al. 1969). Unlike any other element except
hydrogen, helium has a residence time in the atmosphere of approximately 106 years
(Torgersen 1989) before being lost to space. This results in the atmosphere having a
very low He concentration and enables helium enriched in 3He relative to 4He
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
4185
What noble gases tell us about the mantle
East Pacific Rise
Juan de Fuca
Central Indian and Carlsberg Ridges
South West Indian Ridge
North Atlantic
South Atlantic
Yellowstone and Columbia River
Afar
Cameroon line continental
Cameroon line oceanic
St Helena, Mangaia and Tubuai
Reunion
Samoa
Galapagos
Iceland
Hawaii
104
40 30
MORB
20
10 8
6
4
2
continental
HIMU
oceanic
105
4He/ 3He
Figure 1. Compilation of He isotope data showing the low 4He/ 3He ratio and isotopic heterogeneity
of some ocean island systems compared with the relatively well-mixed isotopic composition of midocean ridges (MORB, mid-ocean ridge basalts; HIMU, high mu mantle). Adapted from Barfod
et al. (1999).
degassing from mid-ocean ridges to be readily identified. This also means that,
unlike any other element, helium cannot be recycled from the atmosphere back into
the mantle; the isotopic anomalies observed at mid-ocean ridges are a result of the
release of 3He trapped during accretion and modified only by radiogenic 4He
produced by U and Th decay in the mantle.
An equally important observation is the small 3He/4He ratio variance in midocean ridge basalts (MORB; figure 1). Away from oceanic plateaux and ocean
islands, these samples preserve, across the entire approximately 65 000 km length
of the mid-ocean ridge systems, a uniform 3He/4He isotopic composition (8.75G
2.14R a, where R a is the atmospheric ratio of 1.40!10K6; Graham 2002). This
points to a terrestrial mantle reservoir supplying the mid-ocean ridges that is
relatively homogeneous with respect to its 3He/(UCTh) composition and
provides a critical insight into the volatile–solid evolution of the mantle system
(e.g. Kurz et al. 1982; Allègre et al. 1983; O’Nions & Oxburgh 1983). Isotopic
evidence for the existence of a mantle reservoir separate and distinct from the
convecting mantle is provided by high 3He /4He ocean island volcanism (Kurz
et al. 1982; figure 1). Fluid dynamics require the plume to be initiated from a
boundary layer in the mantle, with the most recent seismic tomographic imaging
of many ocean island plume systems placing this reservoir near to or at the core–
mantle boundary (Montelli et al. 2004, 2006). Nevertheless, the source,
preservation and character of the mantle high 3He/4He reservoir remain a
current topic of debate.
Although the isotopic composition of He on accretion has some model
dependence (pre-deuterium burningZ120R a and solar windZ330R a; Mahaffy
et al. 1998; Podosek et al. 2000), owing to the small amount of 4He compared with
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
4186
C. J. Ballentine and G. Holland
current terrestrial values and model uncertainties, the actual endmember used
makes little difference to the mantle evolution models discussed. For a closed
system and known UCTh content, it is then simple to calculate the mantle
helium concentration that would produce a particular 3He/4He, and this provides
a critical insight into the source or process responsible for introducing volatiles
into the deep planet and its subsequent evolution.
The Earth’s convecting mantle is degassing and also has a changing UCTh
content due to continental crust extraction. The present-day UCTh concentration in the MORB-source or ‘convecting’ mantle, however, is reasonably well
known (e.g. Jochum et al. 1983). An estimate of the 3He concentration in this
mantle reservoir can be made from the present-day 3He flux into the ocean
(Lupton & Craig 1975) and the average oceanic crust production rate (Parsons
1981). The estimated 3He concentration is too low to have preserved the
observed 3He/4He ratio for the given UCTh concentration and cannot be a
simple closed system. Three approaches have been adopted to explore how the
He isotope system in the mantle works: (i) analytically (e.g. Tolstikhin 1975;
Staudacher & Allègre 1982; Anderson 1998; Porcelli & Elliot 2008), (ii) through
numerical simulations of mantle convection and degassing (e.g. Van Keken &
Ballentine 1998, 1999; Tackley 2000), or (iii) by considering a steady-state He
flux through the convecting mantle from a deep volatile-rich mantle reservoir
(Kellogg & Wasserburg 1990; O’Nions & Tolstikhin 1994; Porcelli & Wasserburg
1995). We first discuss the steady-state models.
A widely accepted explanation for the He observations has been a steady-state
flux of high 3He/4He material from beneath the 670 km discontinuity in the
mantle into the convecting upper mantle. High 3He/4He in plumes provides
evidence for a reservoir capable of supplying this flux (Kellogg & Wasserburg
1990; O’Nions & Tolstikhin 1994; Porcelli & Wasserburg 1995). A combination
of seismic tomography (Grand 1987; Van der Hilst et al. 1997) and numerical
models (Van Keken & Ballentine 1998, 1999) demonstrated that the 670 km
mantle phase change does not prevent efficient mixing, and that the mantle
cannot preserve a compositional boundary at this depth. This does not rule out a
flux of volatile-rich and high 3He/4He material from a yet deeper geochemical
reservoir within the Earth (e.g. Kellogg et al. 1999; Porcelli & Halliday 2001;
Tolstikhin & Hofmann 2005), but makes any assumption of steady state in a
much larger convecting system less certain. Numerical simulations do not need to
make an assumption of steady state and are equally capable of balancing such a
deep 3He flux into a convecting mantle system with the observed 3He/4He and
3
He concentration (e.g. Ballentine et al. 2007).
Alternative explanations to a deep 3He flux require higher than currently
accepted 3He concentrations in the convecting mantle or faster degassing rates in
the past to have only recently brought the He concentration to current ‘low’
values. A consequence of the former would require that the present-day 3He flux
into the oceans is a factor of 3.5 times lower than the average over the time
required to produce the ocean crust and is called the ‘zero-paradox’ reference
concentration (Ballentine et al. 2002). While there is considerable uncertainty in
the variance of the past 3He degassing flux (Van Keken et al. 2001), there is
no evidence to support such a high zero-paradox convecting mantle 3He
concentration (e.g. Saal et al. 2002).
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
What noble gases tell us about the mantle
4187
A recent parametrized analytical study by Porcelli & Elliot (2008) assumes an
exponentially decreasing degassing rate over time and a UCTh decrease in the
convecting mantle linked to different crustal growth rates. This work explores
the evolution of 3He/4He in the mantle and provides a conservative estimate of
the highest 3He/4He produced by this reservoir over time. One analytical
approach has been to explore the concept of the high 3He/4He mantle reservoir
originating from a depletion in UCTh, thereby preserving the ambient 3He/4He
at the time of radio-element depletion without requiring a reservoir with a high
3
He concentration (e.g. Anderson 1998). The study by Porcelli & Elliot (2008) in
a similar model context does not provide any support for the required ancient
high 3He/4He convecting mantle required by the ‘low UCTh, not high 3He’
branch of enquiry.
The high-helium high-degassing model explored by Porcelli & Elliot (2008) also
raises an interesting problem. In order to match the present-day convecting
mantle, 3He/4He (8R a) and He concentration (3.2!10K11 cm3(STP) 3He gK1),
a mantle initial 3He of 4.1!10K9 cm3(STP) 3He gK1 is required. In the last 4 Ga,
the model then requires the whole convecting mantle to have lost approximately
99 per cent of its volatiles to the atmosphere. 40Ar mass balance between the
mantle and the atmosphere, discussed in §5, provides evidence of only
approximately 50 per cent mantle outgassing efficiency (Hart et al. 1979), and is
reasonably well reproduced by numerical models of mantle convection
(Van Keken & Ballentine 1998, 1999) and more recent models that attempt to
take into account changes in past heat flow (Ballentine et al. 2007). The same fluid
dynamical models only outgas up to 70 per cent of their original 3He. To effect the
extreme convection required for such efficient outgassing will require a substantial
advance in our understanding of the mantle rheology of the early Earth. Until then,
we argue that the strongest explanation for the observed mantle 3He/4He, low [3He]
and [UCTh] remains a flux of high 3He/4He material from a gas-rich reservoir
deeper in the mantle system (Kurz et al. 1982; Allègre et al 1983; O’Nions & Oxburgh
1983; Kellogg & Wasserburg 1990; Tolstikhin & O’Nions 1994; Porcelli &
Wasserburg 1995).
3. Neon: a trapped meteorite origin for terrestrial mantle volatiles?
Key information on the source composition of the volatiles in the mantle is
provided by the Ne isotopic system. Some of the earliest terrestrial neon isotopic
anomalies, different from atmospheric values, were observed in natural gases and
radioactive minerals (Wetherill 1954; Sharif-Zade et al. 1972, 1973; Verchovsky &
Shukolyukov 1976). These allowed the identification of the different Ne isotope
production mechanisms, often referred to as ‘Wetherill’ reactions (Wetherill
1954). A combination of laboratory observation (Hunemore 1989) and modelling
work (Yatsevich & Honda 1997) provides a good understanding of the rates of
production as a function of radio-element, target element and respective spatial
distribution (Ballentine et al. 2002). In the terrestrial mantle, production of
20
Ne and 22Ne is negligible, while nucleogenic 21Ne addition is readily observed.
The initial observation of 20Ne/ 22Ne values above atmospheric values in both
CO2-rich well gas (Phinney et al. 1978) and MORB (Sarda et al. 1988) was then
due to either mass fractionation of an atmospheric component or an original
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
4188
C. J. Ballentine and G. Holland
14.0
solar Ne
13.5
13.0
MORB
20Ne/ 22Ne
12.5
12.0
SWI-Ne
air-MORB mixing
11.5
11.0
10.5
10.0
9.5
radiogenic
(crust)
air
9.0
0.02
0.04
0.06
0.08
0.10
21Ne / 22Ne
Figure 2. Resolution of the mantle Ne isotopic composition from Ne isotopes in CO2 well gas
(Ballentine et al. 2005; Holland & Ballentine 2006). Groundwater that has accumulated crustal
21
Ne contains dissolved air Ne with a natural radiogenic 21Ne isotope ‘spike’. When magmatic CO2
contacts the groundwater, mixing between the groundwater and the magmatic fluid produces a
unique mixing line. Intersection of this mixing line with the well-defined MORB air–mantle
mixing line provides an unambiguous resolution of the convecting mantle 20Ne/ 22Ne value
(Ballentine et al. 2005), unobtainable with either dataset alone. SWI, solar wind-implanted. Data
adapted from Holland & Ballentine (2006).
accretionary signature. Honda et al. (1991) were able to show that the resolved
3
He/22Ne, assuming an accretionary origin, was similar to solar values. This
interpretation was reinforced by Patterson et al. (1990), who successfully argued
against any fractionation of a shallow air contaminant. It is worth noting at this
point that recent work (Holland & Ballentine 2006) rules out a recycled-air
component significantly contributing to the mantle Ne isotope system; this is
discussed in more detail later in this paper.
The Ne isotopic composition observed in mantle-derived samples is then a
three-component mixture of an atmospheric component most probably
introduced during the eruption process or even during sample handling (Patterson
et al. 1990; Ballentine & Barfod 2000), nucleogenic 21Ne and the original
accretionary 20Ne/22Ne isotope signature (figure 2).
Identification of the accretionary origin of Ne is particularly important because,
alone of all of the noble gases, there is a clear and simple isotopic difference
between Ne trapped within meteorites and solar Ne. Solar Ne has a
20
Ne/22NeZ13.8 (Benkert et al. 1993; Grimberg et al. 2006), while the maximum
observed trapped in meteorites is 20Ne/22NeZ12.5 (Black 1972; Wieler
2002; Trieloff et al. 2000, 2002). Models of terrestrial volatile acquisition invoking a major role for solar nebula gases are predicted to produce a mantle
20
Ne/22New13.8. Solar nebula models include, for example, gravitational capture
of a massive early atmosphere and magma ocean equilibration (Mizuno et al. 1980;
Porcelli et al. 2001; Zahnle et al. 2007). By contrast, models of mantle volatile
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
What noble gases tell us about the mantle
4189
capture that require solar wind irradiation of accreting material (Podosek
et al. 2000; Trieloff et al. 2000, 2002; Ballentine et al. 2005; Tolstikhin & Hofmann
2005) would predict a lower 20Ne/ 22Ne value. But it is critical to distinguish
between a solar Ne isotopic composition that has been reduced by eruptionrelated air contamination (we rule out atmospheric Ne recycling later) and
a systematic mantle 20Ne/22Ne upper limit (Ballentine et al. 2001; Trieloff
et al. 2001).
Until recently, most work has simply relied on the uppermost observed
values. For example, Ozima & Zashu (1991) concluded from diamonds that the
mantle upper limit was 20Ne/22Ne!12.0, and therefore included a significant
meteorite-trapped ‘planetary’ component in their model. Farley & Poreda
(1993) noted that the mantle value appeared to ‘top out’ at 20Ne/22NeZ12.5
but, owing to limited data, did not place any significance on this value. More
recently, variable 40Ar/36Ar for a constant maximum 20Ne/22NeZ12.5 from
both Iceland and Loihi (Hawaiian seamount) and newer high-quality MORB
data (Moreira et al. 1998) were used to argue for a SWI origin for the Ne in
the mantle (Trieloff et al. 2000). It was not possible, however, to rule out a
ubiquitous air component mixing with solar wind with a higher 20Ne/22Ne
(Ballentine et al. 2001; Trieloff et al. 2001). Interpretation was further complicated
by reports of 20Ne/22NeO12.5, but these either had large error bars, or were nonreproducible or both (e.g. Sarda et al. 1988; Harrison et al. 1999). The only
indication of reproducible 20Ne/22NeO12.5 is in aliquots of gas released from
inclusions from one Kola plume sample (Yokochi & Marty 2004).
4. Neon: evidence from CO2-rich well gases
A major breakthrough occurred by transferring focus from ocean island and midocean ridge basalts to continental natural gases. In particular, CO2-rich natural
gases in New Mexico were used to identify the first terrestrial isotopic anomalies
(129Xe/130Xe) that could be ascribed to the presence of an extinct radionuclide
(129I) within the Earth (Butler et al. 1963). From their particular Xe isotopic
composition, these gases were soon identified as also being remarkably similar to
those found in MORB (Staudacher 1987), although with the addition of more
radiogenic noble gases from the crust.
It has been known for a long time that crustal radiogenic noble gases
accumulate in groundwater (Torgersen & Clarke 1985). Groundwater 4He
ages are now routinely used in ancient groundwater studies (e.g. Kipfer et al.
2002), while the accumulation of 4He in natural gases that have equilibrated
with the groundwater is well documented (Ballentine & Sherwood Lollar 2002;
Zhou & Ballentine 2006). The latter is particularly important when a magmatic
gas passes through or contacts an ‘old’ groundwater. Instead of picking up
simple air ‘contamination’, the air component has been ‘spiked’ with a natural
crustal radiogenic/nucleogenic addition. In the case of the Ne isotopes,
this means that mixing between the premixed dissolved air and crustal
component and the magmatic endmember will define a mixing line with a
different starting composition and gradient from those of simple air–mantle
mixing. With the principal assumption that the convecting mantle has supplied
the magmatic gases, intersection of the (crustCair)–mantle mixing line with
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
4190
C. J. Ballentine and G. Holland
the air–MORB mixing line then uniquely defines the mantle upper limit
20
Ne/22Ne. This value, 20Ne/22NeZ12.5 (Ballentine et al. 2005; Holland &
Ballentine 2006; figure 2), is indistinguishable from the upper limit found from
SWI meteorites. Models requiring simple solar nebula gas input into the early
planet need a rethink.
5. Argon: a note on mantle degassing
The main thrust of this paper is to review the advances made in understanding
the accretionary origin of volatiles in the solid Earth. Nevertheless, the
evolution and magnitude of degassing of the planet play an important role in
developing an understanding of the interaction between the deep-Earth volatile
reservoir and the atmosphere. Foremost among the tools we have is 40Ar.
Comparing the solar 40Ar/36ArZ3!10K4 (Anders & Grevesse 1989) with the
terrestrial atmosphere value of 295 clearly shows the important input of 40Ar
into the atmosphere derived from the decay of 40K from the solid Earth. If we
take bulk silicate Earth (BSE) estimates of the K concentration based on the
K/Uw12 500 ratio found in basalts (Jochum et al. 1983), then some 45–50 per
cent of the planet’s 40Ar is now found in the Earth’s atmosphere (e.g. Hart
et al. 1979; Allègre et al. 1996). This reflects not only the ability of the Earth
over 4.5 Ga to concentrate the lithophile elements (including K) in the
continental crust, but also the efficiency of the Earth to degas the 40Ar
produced within it.
We can estimate the MORB-source mantle 40Ar concentration from the
mantle 3He/40Ar defined from well gases and MORB (Moreira et al. 1998;
Ballentine et al. 2005) and the 3He mantle concentration derived above. If we
take the whole mantle to have this 40Ar concentration, the total 40Ar in the
atmosphere and the mantle combined is significantly less than the total
amount of 40Ar estimated to have been produced since the Earth’s accretion.
Hart et al. (1979) concluded that this could be explained by a hidden 40Ar-rich
reservoir in the deep mantle. It is often assumed that the 3He- and 40Ar-rich
reservoirs are one and the same, although there is no a priori reason for
this assumption.
We note that if the convecting mantle had a 40Ar concentration 3.5 times higher
(the zero-paradox noble gas concentration), there would be no need for a
hidden 40Ar reservoir (Ballentine et al. 2002). This is exactly the same as the 3He
and 4He concentration discrepancy discussed previously and illustrates a strong
coherence between radiogenic 4He and 40Ar. Arguments to explain the ‘40Ar
paradox’ by, for example, low BSE K/U (e.g. Lassiter 2004) or high compatibility
of Ar ( Watson et al. 2007) do not account for this coherence and are
weaker accordingly.
6. Non-radiogenic argon: no evidence of a solar component
There is no significant production of either 36Ar or 38Ar in the terrestrial mantle
(Ballentine & Burnard 2002). These isotope species then represent either an air
component that has been recycled into the mantle, an accretionary signature,
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
4191
What noble gases tell us about the mantle
0.1950
38Ar/ 36Ar
0.1900
air
0.1850
0.1800
range of
solar values
0.1750
0.1700
0
5
10
40Ar/ 36Ar
15
20
25
(×103)
Figure 3. CO2 well-gas 38Ar/ 36Ar versus 40Ar/36Ar. Data adapted from Holland & Ballentine
(2006). Bravo Dome (filled square) and Sheep Mountain (open square) samples with resolvable
mantle-derived noble gases show no deviation from air 38Ar/ 36Ar (dashed line), irrespective of
40
Ar/ 36Ar. Gases dominated by a crustal radiogenic signature (open diamond, McElmo Dome)
with high 40Ar/36Ar are expected to have air values of 38Ar/ 36Ar and show that these values are
not an experimental artefact caused by high partial pressures of 40Ar at high 40Ar/36Ar. The
mantle has 38Ar/36Ar values irresolvable from air. The data are presented in table 1.
or a hybrid of the two. Although deviations from air 38Ar/36Ar in MORB have
been reported (Niedermann et al. 1997; Pepin 1998), these were not reproduced
in later detailed studies of the same MORB samples (Trieloff et al. 2002).
We present here 38Ar/36Ar data from the mantle volatile-rich well-gas samples
detailed in Holland & Ballentine (2006). One-sigma error envelopes of 38Ar/36Ar
ratios in all the samples are within error of air (figure 3; table 1), even in well-gas
samples containing as much as 50 per cent mantle argon (Holland & Ballentine
2006). A York fit to the Bravo Dome data shows that 38Ar/36ArZ0.1882G0.0003
at 40Ar/36ArZ22 000 (figure 3). This is indistinguishable from modern air.
Solar 38Ar/36Ar ratios are distinct from air with estimates ranging from 0.172
to 0.18 (Becker et al. 1998; Pepin et al. 1999) compared with 0.1880 for air. Other
possible accretionary components such as ‘Q’ have 38Ar/36ArZ0.1873 (Busemann
et al. 2000), almost indistinguishable from air, while fractionated solar Ar can
have 38Ar/36Ar ratios greater than air (38Ar/36ArZ0.205; Benkert et al. 1993)
due to implantation fractionation analogous to that observed in Ne isotopes
(Grimberg et al. 2006).
While it is clear that simple consideration of only 38Ar/36Ar cannot in itself
enable resolution of a potential Q-component or hybrid solar/implanted solar Ar
contribution to the mantle 38Ar/36Ar, the simplest explanation is that this Ar
isotope pair is dominated by air that has been recycled into the mantle (Holland &
Ballentine 2006). We develop this argument further in the discussion of mantle
Kr and Xe.
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
4192
C. J. Ballentine and G. Holland
Table 1. Ar isotopic ratios from Bravo Dome, Sheep Mountain and McElmo Dome (all data 1s).
sample
40
Ar/36Ar
BD11-B
BD20-B
WBD04-B
BD19-B
BD15-B
BD02-B
BD07-B
BD13-B
BD05-B
BD04-B
WBD01-B
WBD03-B
19733
21401
22548
9888
13472
4197
11205
7655
5288
8612
18496
16792
300
392
730
152
224
69
172
117
81
133
298
268
0.1878
0.1886
0.1884
0.1887
0.1886
0.1880
0.1890
0.1878
0.1883
0.1884
0.1879
0.1881
0.00037
0.00089
0.00079
0.00056
0.00055
0.00046
0.00047
0.00047
0.00044
0.00065
0.00050
0.00099
Sheep 5-9-a
Sheep 8-2-p
Sheep 2-10-0
Sheep 9-26
Sheep 2-9-h
Sheep 3-15-B
Sheep 4-13
Sheep 4-26-E
Sheep 3-23-D
Sheep 7-35-L
Sheep 2-35-C
Sheep 1-15-C
Sheep 4-14-M
Sheep 3-4-0
Sheep 5-15-0
Sleep 4-4-P
17506
14100
15548
17026
4348
8680
17617
22033
16836
16857
16420
5102
17663
4469
8562
15802
229
182
1433
230
126
114
269
292
219
221
226
71
264
59
114
209
0.1893
0.1881
0.1885
0.1875
0.1874
0.1884
0.1890
0.1893
0.1893
0.1877
0.1876
0.1887
0.1880
0.1897
0.1892
0.1881
0.00065
0.00023
0.00040
0.00043
0.00060
0.00034
0.00072
0.00109
0.00044
0.00039
0.00063
0.00089
0.00092
0.00076
0.00079
0.00069
McElmo
McElmo
McElmo
McElmo
McElmo
16825
14297
16420
8471
14425
157
156
181
78
114
0.1877
0.1883
0.1891
0.1883
0.1896
0.00053
0.00115
0.00113
0.00032
0.00076
Moquil
He2
Sc9
Doe Canyon
Yb2
error
38
Ar/36Ar
error
7. Krypton: both hidden and ignored
The krypton isotopic composition of mantle gases remains largely unreported,
with most analyses providing isotopic compositions little different from modern
air. Unlike the unambiguous identification of the isotopic composition of mantle
Ne, similar to mantle 38Ar/ 36Ar, it has been impossible to distinguish mantle Kr
from atmosphere contamination. The mantle and atmosphere Kr isotopic
composition is uniformly depleted in light isotopes by approximately 0.8
per cent/amu from solar. The Q-component in carbonaceous chondrites is
depleted in light isotopes by approximately 1.1 per cent/amu compared with solar
(e.g. Pepin & Porcelli 2002). Again, it is impossible to use simple consideration
of mantle Kr isotopes to distinguish between a recycled-air and a possible
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
What noble gases tell us about the mantle
4193
35
ocean
sediment
130Xe/ 36Ar
(×10 – 4)
30
25
20
15
mantle (well gas and MORB)
10
seawater
5
air
0
0.02 0.04 0.06 0.08 0.10 0.12 0.14 0.16 0.18 0.20
84Kr/ 36Ar
Figure 4. The 130Xe/ 36Ar versus 84Kr/36Ar for air, seawater, MORB 2PD43 popping rock at
20
Ne/22NeZ12.5 (Moreira et al. 1998), Bravo Dome well-gas ‘mantle’ (Holland & Ballentine 2006)
and ocean sediments (Matsuda & Nagao 1986). The mantle components in Bravo Dome and
popping rock are identical and suggest minimal elemental fractionation in these complementary
sample types. The mantle 130Xe/ 84Kr/ 36Ar composition is identical to seawater with a small ocean
sediment contribution.
Q-component origin for Kr in the Earth’s mantle. However, the mantle 84Kr/36Ar
in both the MORB-source mantle and well-gas mantle are remarkably similar to
the elemental abundance (and isotopic composition) of seawater (figure 4). This is
used to support a recycled air-dissolved-in-seawater origin for these species in the
mantle (Holland & Ballentine 2006).
8. Xenon: massive early planetary volatile loss and fractionation
The full detail of the Xe isotope system is reviewed in several papers (Pepin 2000;
Pepin & Porcelli 2002, 2006; Porcelli & Ballentine 2002). A key observation is
that the terrestrial atmosphere is depleted in the light isotopes of Xe by
approximately 4.2 per cent/amu compared with solar and is significantly more
fractionated than Kr. The terrestrial atmosphere in addition contains 129I
contributions to 129Xe, and 244Pu and UCTh contributions to 131,132,134,136Xe.
Nevertheless, if the 244Pu and UCTh additions to the atmosphere are subtracted, the resulting non-radiogenic atmosphere composition cannot be simply
related to the fractionation of solar Xe, being underabundant in the two heaviest
isotopes of Xe. This has led to the use of ‘U–Xe’ as the non-radiogenic precursor
Xe isotopic composition of the Earth (Pepin 2000). The non-radiogenic
fractionated U–Xe, called NEA-Xe, provides the main reference for resolving
subsequent addition of U-, Pu- and I-derived isotopes of Xe to the atmosphere
(Pepin & Porcelli 2006).
The timing of this volatile loss can be accessed by consideration of model
‘closure’ ages for both the I–Xe and Pu–Xe systems separately (e.g. Wetherill
1975) or by combining the two decay systems (Pepin & Phinney 1976). There are
variants of this approach (e.g. Pepin & Porcelli 2006), but all converge on a closure
age for terrestrial volatile loss at ca 80 Ma after Solar System formation. It should
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
4194
C. J. Ballentine and G. Holland
2.7
radiogenic
(crust)
MORB
2.6
136Xe/ 130Xe
2.5
2.4
air-MORB mixing
2.3
2.2
air
2.1
6.4
6.6
6.8
7.0
7.2
7.4
7.6
7.8
8.0
8.2
129Xe / 130Xe
Figure 5. The intersection of the Bravo Dome Xe array with the MORB array defines the Xe
mantle composition. In an analogous way to Ne isotope systematics, the Xe data from the Bravo
Dome system show a two-component air–crust mixture that intersects with the MORB air–mantle
mixing line to define the mantle Xe isotopic endmember to be 129Xe/130XeZ7.90G0.14. The
MORB–air mixing line is defined by popping rock data (Moreira et al. 1998). Vertical and
horizontal dashed lines are the intersection of the MORB array with the Bravo Dome array and
indicate the convecting mantle values.
be noted that the 129Xe*/136Xe* (where the asterisk designates 129I-derived Xe and
244
Pu-derived Xe) in the convecting mantle (MORB-source) and atmosphere are
identical—it is not possible to resolve closure of one reservoir from the other
(Pepin & Porcelli 2006). Critically for this work, from I–Pu systematics, it can
be calculated that the silicate Earth has lost almost 99 per cent of its Xe and
therefore other noble gases during the accretionary process (Porcelli et al. 2001).
9. Non-radiogenic xenon: an accretionary component in the mantle
mixed with air-derived Xe
The non-radiogenic Xe isotopes, 124,126,128,130Xe, in the convecting mantle
(MORB-source) are dominated by Xe with an isotopic composition very similar
to air. An excess in these isotopes was first shown by Caffee et al. (1999) in
mantle-rich CO2 well gases, with samples lying on the air–solar mixing line. The
most solar-rich gas, from Bravo Dome in New Mexico, had a 10 per cent excess
above air value. Since solar and Q-Xe are so similar in composition, it is not
possible from these samples to distinguish isotopically between sources.
Holland & Ballentine (2006), collecting and analysing a full suite of samples,
were able to investigate the Xe isotope system at Bravo Dome in more detail. In
particular, we were able to define, independent of Ne, the mantle endmember to
be 129Xe/130XeZ7.90G0.14 (figure 5), similar to the highest observed value in
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
4195
What noble gases tell us about the mantle
10 4
terrestrial mantle
20Ne/ 132Xe
10 3
10 2
10
1
1
10 2
10
20 Ne
10 3
10 4
(×10 – 8 cm3 g–1)
Figure 6. Carbonaceous chondrite 20Ne/132Xe versus 20Ne. The range of Ne concentration observed in
carbonaceous chondrites is sufficient to provide Ne in the gas-rich mantle source reservoir (vertical dotdashed line; see text). Also, the Ne/Xe ratio observed in the highest-concentration meteorites is similar
to accretionary Ne/Xe in the terrestrial mantle (horizontal dot-dashed line). Carbonaceous chondrite
(CC) data (diamonds, CC type 1; squares, CC type 2; triangles, CC type 3; circles, CC type 4) from
Mazor et al. (1970) and Scherer & Schultz (2000).
MORB (Kunz et al. 1998). This dataset then showed that, even for the most
mantle-rich sample, approximately 50 per cent of the Xe was due to shallow air
contamination. For the first time, we were then able to define the convecting
mantle as a mixture of 20 per cent solar or Q-derived Xe, but dominated (80%)
by a component with an isotopic composition identical to air. The mantle
accretionary/primitive 20Ne/132Xe ratio is approximately 1400. This is clearly
not a solar ratio (solar 20Ne/130Xex500 000–1 000 000; Wieler 2002), is similar to
the values observed in meteorites (figure 6) and compatible with a model in
which the source of mantle Xe is also a trapped accretionary component.
The well-gas sample suite of Holland & Ballentine (2006) is able to
unambiguously resolve, for the first time, the mantle endmember components for
all noble gas isotopes. It is now possible to compare the elemental abundance pattern
of the non-radiogenic heavy noble gases in the mantle, 36Ar/84Kr/(80%)130Xe, with
those in MORB and potential sources of these species, the atmosphere and oceans.
The well-gas and least degassed MORB sample, ‘popping rock’ 2PD43 (Moreira
et al. 1998), have indistinguishable mantle elemental abundance patterns. In the
simplest case, one might expect the degassing residue (MORB) and gas released
from a magma (well gases) to fractionate in opposite directions. The similar value in
both sample suites strongly suggests that both represent a minimally fractionated
mantle system. This system is identical to a seawater composition admixed with
sedimentary-derived Kr and Xe (figure 4). Since Ne (and He) has such a low
solubility in seawater, this source would have no effect on the primitive Ne (or He)
isotopic composition of the mantle. If the noble gases are not decoupled from the
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
4196
C. J. Ballentine and G. Holland
seawater in which they are dissolved during the subduction process, an ongoing
area of research, recycled seawater accounts for approximately 50 per cent of the
water now in the convecting mantle and is compatible with constraints imposed
by deuterium isotope considerations (Holland & Ballentine 2006).
We add one note of caution. High-precision well-gas and MORB Xe isotope data
have been compared and shown to be different in their plutogenic fission Xe
spectrum (Pepin & Porcelli 2006). These authors have attributed this difference to
a different closure age of distinct reservoirs in the mantle. The well-gas samples
are, however, otherwise almost identical to MORB popping rock in noble gas
elemental and isotopic compositions, including 129Xe/130Xe. It is difficult to
understand how differential closure and therefore differential volatile loss would
then balance with the noble gas input over time from U, Th and K decay, and
possibly air recycling, in different reservoirs yet still provide for such similar noble
gas compositions. While the Xe-fission isotopic difference between MORB and
well gases remains unresolved, this does not affect the conclusions above drawn
from the non-radiogenic Xe isotopes, 124,126,128,130Xe.
10. Support for subducted ‘air’ noble gases in the terrestrial mantle
The concept of seawater recycling into the mantle has been controversial, with
authors, respectively, arguing for a possible air component in Ar and Xe (Porcelli &
Wasserburg 1995) or a subduction barrier effectively preventing any surface return
to the mantle (Staudacher & Allègre 1988). While observations of air-derived
components in MORB have been used to support a recycling model (Bach &
Niedermann 1998; Sarda et al. 1999), shallow air contamination of the samples has
been a significant complicating factor (Ballentine & Barfod 2000).
Serpentine inclusions within olivine from the Higashi-akaishi peridotite body,
Japan, provide a new insight into the subduction process. These have been
exhumed from the depths of approximately 120 km and originate through
dewatering of the subducting slab and subsequent hydration reactions with the
host mantle-wedge olivine (Mizukami & Wallis 2005). These inclusions contain
atmosphere-derived noble gases in seawater proportions and halogen ratios
indistinguishable from marine pore fluids. First results indicate that subduction
of fluids with a seawater (marine pore fluid) noble gas signature occurs to depths
of at least 120 km (Sumino et al. 2007).
The global impact of seawater noble gas subduction can be investigated by
adapting the numerical modelling approach of Van Keken & Ballentine (1998,
1999). In these models, a seawater noble gas flux into the mantle can be balanced
with the model 40Ar production by K decay to reproduce the observed convecting
mantle (MORB-source) 40Ar/ 36Ar. These models produce a plume-forming
region at the core–mantle boundary with a 40Ar/ 36Ar ratio three to four times
lower than the average convecting mantle (Ballentine et al. 2007). This compares
well with observation, where we now know the convecting mantle (MORBsource) 40Ar/ 36ArZ42 450G4420 (Holland & Ballentine 2006), which is between
5.1 and 9.3 times higher than plume maximum values of 40Ar/ 36ArZ4530
(Iceland—Trieloff et al. 2000), 8000 ( Juan Fernandez—Farley & Craig 1994),
8300 (Loihi—Trieloff et al. 2000) and 7600 (Reunion—Trieloff et al. 2002).
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
What noble gases tell us about the mantle
4197
11. Mass balance for a mantle-‘trapped’ noble gas meteorite source
Using a 3He concentration in the MORB-source region of 4.4!10K11cm3 (STP)
3
He gK1 (Porcelli & Ballentine 2002), convecting mantle He/Ne and He/Xe
ratios at 20Ne/ 22NeZ12.5 (Moreira et al. 1998) and knowledge that even
these samples are 80 per cent atmospheric Xe (Holland & Ballentine 2006),
we obtain a primordial Ne concentration of 1.1!10K10 cm3(STP) 20Ne gK1 and
a primordial Xe concentration of 7.8!10K14 cm3(STP) 132Xe gK1 in the
convecting mantle.
We have already established earlier in this paper that we require a gas-rich
mantle reservoir to flux 3He into the convecting mantle. The concentration of
3
He in this mantle reservoir can be estimated if we assume a closed system, BSE
UCTh, an initial 3He/4He of 120R a and a present-day 3He/ 4He similar to the
maximum observed in plumes of 50R a, which is approximately 100 times the 3He
concentration of the convecting mantle itself (e.g. Porcelli & Wasserburg 1995).
Scaling Ne and Xe concentrations to 3He then gives us gas-rich mantle source
reservoir concentrations of 1.1!10K8 cm3(STP) 20Ne gK1 and 7.8!10K12
cm3(STP) 132Xe gK1.
These values compare with the highest concentrations observed in bulk
meteorite samples of approximately 1!10K5 cm3(STP) 20Ne gK1 (figure 6) and
approximately 1!10K8 cm3(STP) 132Xe gK1 (compiled from Mazor et al. (1970)
and Scherer & Schultz (2000)). As discussed earlier, Xe isotopes require that we
must allow for at least a 100-fold loss of noble gases during the accretionary
process (Porcelli et al. 2001). This reduces the maximum noble gas concentrations
that could have been present on the Earth at the end of accretion to 1!10K7
cm3(STP) 20Ne gK1 and 1!10K10 cm3(STP) 132Xe gK1. These are still
approximately an order of magnitude higher than the gas-rich mantle source
reservoir concentrations calculated above.
While a necessarily crude calculation, we do not need to resort to special
circumstances or to only the most gas-rich material to form the source of
the volatile-rich terrestrial mantle reservoir—this is well within the range of
observed concentrations in bulk meteorite samples, even allowing for 99 per cent
volatile loss.
12. Synthesis and conclusions
(i) The terrestrial atmosphere is unique in its noble gas isotopic and
elemental composition within the Solar System.
(ii) This isotopic signature dominates the mantle Ar/Kr/(80%)Xe nonradiogenic isotopes but in seawater proportions, allowing for a small
ocean sediment contribution.
(iii) Ocean island volcanic systems have significantly lower maximum observed
40
Ar/36Ar values at 20Ne/ 22NeZ12.5 than the convecting mantle.
(iv) This last observation is consistent with numerical models that investigate recycled atmosphere-derived noble gases into the convecting
mantle and show the mantle source region of plumes to have lower
40
Ar/ 36Ar than the convecting mantle.
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
4198
C. J. Ballentine and G. Holland
It would be an extreme coincidence if the unique isotopic composition and
abundance pattern of the atmosphere, modified by the respective noble gas
solubility in water, were to be found in the mantle by any process other than
subduction of air dissolved in seawater into the mantle system. The dominant
source of the non-radiogenic heavy noble gases (Ar, Kr and (80%)Xe) in the
convecting mantle is due to the subduction of air dissolved in seawater-derived
pore fluids. These species do not then represent a primary or accretionary
reservoir within the mantle that later degassed to form the atmosphere. Further
observations of the mantle strongly suggest that we have evidence for a primitive
reservoir within the Earth, unfractionated since accretion.
(v) The helium isotopic composition of the convecting mantle is too high for
a closed system model to simply match the observed 3He and UCTh
concentrations.
(vi) A 3He flux into the convecting mantle from a deep volatile-rich mantle
reservoir provides one explanation for the high convecting mantle 3He/ 4He.
(vii) The Ne isotopic composition of the convecting reservoir is 20Ne/ 22NeZ12.5,
and is indistinguishable from that expected for a SWI accretionary source.
(viii) Twenty per cent of the convecting mantle non-radiogenic Xe is primordial
but its isotopic composition cannot distinguish between a solar nebula and
trapped (Q) meteorite source.
(ix) The primordial 20Ne/ 132Xe ratio in the mantle of approximately 1400
(figure 6) is within the range observed in the higher concentration
component of bulk meteorites and distinct from the solar ratio (approx.
500 000–1 000 000).
The presence of neon indistinguishable from meteorite SWI neon in the
terrestrial mantle together with meteorite proportions of primitive Ne/Xe
strongly supports a trapped accretionary noble gas source for these volatiles now
in a deep volatile-rich reservoir in the mantle. The mechanism and timing of
formation of this reservoir remain enigmatic. Possible explanations may be dry
subduction of a late solar wind-irradiated veneer (Tolstikhin & Hofmann 2005).
This would allow for bulk planetary degassing, volatile loss and fractionation of
the atmosphere to its current noble gas state independently of arrival.
Alternatively, the Earth may have retained some volatiles from the earliest
phase of accretion. Xe isotopes require extensive gas loss (more than 99%) during
accretion, with only a small percentage of the mantle preserving any accretionary
signature at all. In the light of the recent modelling work pointing to multiple
massive impacts in the formation of planets with Earth-like orbits (Morbidelli
2002), it is no surprise that, during the process of accretion, there has been a
substantial loss of volatile elements from the accreting material.
Nonetheless, after correction for atmospheric contamination and water
recycling, we have identified primordial noble gas signatures in the deep Earth
that strongly suggest they were acquired from meteorites rather than
gravitationally captured from the solar nebula. Trapped meteorite gases should
therefore provide the starting point for a more refined understanding of the
accretionary and post-accretionary loss processes that occurred to leave the
volatile abundances observed in the terrestrial planets today.
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
What noble gases tell us about the mantle
4199
We would like to thank Andrew Jephcoat and Alex Halliday for organizing the meeting at the
Royal Society. We thank Don Porcelli for a thorough and detailed review of this paper.
References
Allègre, C. J., Staudacher, T., Sarda, P. & Kurz, M. 1983 Constraints on evolution of Earth’s
mantle from rare gas systematics. Nature 303, 762–766. (doi:10.1038/303762a0)
Allègre, C. J., Hofmann, A. W. & O’Nions, R. K. 1996 The argon constraint on mantle structures.
Geophys. Res. Lett. 23, 3555–3557. (doi:10.1029/96GL03373)
Anders, E. & Grevesse, N. 1989 Abundances of the elements: meteoritic and solar. Geochim.
Cosmochim. Acta 53, 197–214. (doi:10.1016/0016-7037(89)90286-X)
Anderson, D. L. 1998 A model to explain the various paradoxes associated with mantle noble gas
geochemistry. Proc. Natl Acad. Sci. USA 95, 9087–9092. (doi:10.1073/pnas.95.16.9087)
Bach, W. & Niedermann, S. 1998 Atmospheric noble gases in volcanic glasses from the southern
Lau Basin: origin from the subducting slab? Earth Planet. Sci. Lett. 160, 297–309. (doi:10.1016/
S0012-821X(98)00091-0)
Ballentine, C. J. & Barfod, D. N. 2000 The origin of air-like noble gases in MORB and OIB. Earth
Planet. Sci. Lett. 180, 39–48. (doi:10.1016/S0012-821X(00)00161-8)
Ballentine, C. J. & Burnard, P. G. 2002 Production, release and transport of noble gases in the
continental crust. Rev. Miner. Geochem. 47, 481–538. (doi:10.2138/rmg.2002.47.12)
Ballentine, C. J. & Sherwood Lollar, B. 2002 Regional groundwater focusing of nitrogen and noble
gases into the Hugoton–Panhandle giant gas field, USA. Geochim. Cosmochim. Acta 66,
2483–2497. (doi:10.1016/S0016-7037(02)00850-5)
Ballentine, C. J., Porcelli, D. & Wieler, R. 2001 Technical comment on ‘Noble gases in mantle
plumes’. Science 291, 2269a. (doi:10.1126/science.291.5512.2269a)
Ballentine, C. J., van Keken, P. E., Porcelli, D. & Hauri, E. H. 2002 Numerical models,
geochemistry and the zero paradox noble-gas mantle. Phil. Trans. R. Soc. A 360, 2611–2631.
(doi:10.1098/rsta.2002.1083)
Ballentine, C. J., Marty, B., Sherwood Lollar, B. & Cassidy, M. 2005 Neon isotopes constrain
convection and volatile origin in the Earth’s mantle. Nature 433, 33–38. (doi:10.1038/
nature03182)
Ballentine, C. J., Brandenburg, J. P., van Keken, P. E. & Holland, G. 2007 Seawater recycling into
the deep mantle—and the source of He-3? Geochim. Cosmochim. Acta 71, A56–A56. (doi:10.
1016/j.gca.2006.07.037)
Barfod, D. N., Ballentine, C. J., Halliday, A. N. & Fitton, J. G. 1999 Noble gases in the Cameroon
line and the He, Ne, and Ar isotopic compositions of high mu (HIMU) mantle. J. Geophys. Res.
Solid Earth 104, 29 509–29 527. (doi:10.1029/1999JB900280)
Becker, R. H., Schlutter, D. J., Rider, P. E. & Pepin, R. O. 1998 An acid-etch study of the Kapoeta
achondrite: implications for the argon-36/argon-38 ratio in the solar wind. Meteorit. Planet. Sci.
33, 109–113.
Benkert, J.-P., Baur, H., Signer, P. & Wieler, R. 1993 He, Ne, and Ar from solar wind and solar
energetic particles in lunar ilmenites and pyroxenes. J. Geophys. Res. 98, 13 147–13 162.
(doi:10.1029/93JE01460)
Black, D. C. 1972 On the origins of trapped helium, neon and argon isotopic variations in
meteorites—I. Gas-rich meteorites, lunar soil and breccia. Geochim. Cosmochim. Acta 36,
347–375. (doi:10.1016/0016-7037(72)90028-2)
Brown, H. 1949 Rares gases and the formation of the Earth’s atmosphere. In The atmospheres of
the Earth and planets (ed. G. P. Kuiper), pp. 258–266. Chicago, IL: University of Chicago Press.
Busemann, H., Baur, H. & Wieler, R. 2000 Primordial noble gases in ‘phase Q’ in carbonaceous and
ordinary chondrites studied by closed system stepped etching. Meteorit. Planet. Sci. 35, 949–973.
Butler, W. A., Jeffery, P. M. & Reynolds, J. H. 1963 Isotopic variations in terrestrial xenon.
J. Geophys. Res. 68, 3283–3291. (doi:10.1029/JZ068i010p03283)
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
4200
C. J. Ballentine and G. Holland
Caffee, M. W., Hudson, G. U., Velsko, C., Huss, G. R., Alexander, E. C. & Chivas, A. R. 1999
Primordial noble gases from Earth’s mantle: identification of a primitive volatile component.
Science 285, 2115–2118. (doi:10.1126/science.285.5436.2115)
Canup, R. M. 2004 Simulations of a late lunar-forming impact. Icarus 168, 433–456. (doi:10.1016/
j.icarus.2003.09.028)
Canup, R. M. 2008 Accretion of the Earth. Phil. Trans. R. Soc. A 366, 4061–4075. (doi:10.1098/
rsta.2008.0101)
Clarke, W. B., Beg, M. A. & Craig, H. 1969 Excess 3He in the sea: evidence for terrestrial
primordial helium. Earth Planet. Sci. Lett. 6, 213–220. (doi:10.1016/0012-821X(69)90093-4)
Dauphas, N. 2003 The dual origin of the terrestrial atmosphere. Icarus 165, 326–339. (doi:10.1016/
S0019-1035(03)00198-2)
Dauphas, N., Robert, F. & Marty, B. 2000 The late asteroidal and cometary bombardment of
Earth as recorded in water deuterium to protium ratio. Icarus 148, 508–512. (doi:10.1006/icar.
2000.6489)
Farley, K. A. & Craig, H. 1994 Atmospheric argon contamination of ocean island basalt olivine
phenocrysts. Geochim. Cosmochim. Acta 58, 2509–2517. (doi:10.1016/0016-7037(94)90027-2)
Farley, K. A. & Poreda, R. J. 1993 Mantle neon and atmospheric contamination. Earth Planet. Sci.
Lett. 114, 325–339. (doi:10.1016/0012-821X(93)90034-7)
Graham, D. W. 2002 Noble gas isotope geochemistry of mid-ocean ridge and ocean island basalts:
characterization of mantle source reservoirs. Rev. Miner. Geochem. 47, 247–317. (doi:10.2138/
rmg.2002.47.8)
Grand, S. P. 1987 Tomographic inversion for shear velocity beneath the North American Plate.
J. Geophys. Res. 92, 14 065–14 090. (doi:10.1029/JB092iB13p14065)
Grimberg, A., Baur, H., Bochsler, P., Buhler, F., Burnett, D. S., Hays, C. C., Heber, V. S.,
Jurewicz, A. J. G. & Wieler, R. 2006 Solar wind neon from genesis: implications for the lunar
noble gas record. Science 314, 1133–1135. (doi:10.1126/science.1133568)
Harrison, D., Burnard, P. & Turner, G. 1999 Noble gas behaviour and composition in the mantle:
constraints from the Iceland plume. Earth Planet. Sci. Lett. 171, 199–207. (doi:10.1016/S0012821X(99)00143-0)
Hart, R., Dymond, J. & Hogan, L. 1979 Preferential formation of the atmosphere–sialic crust
system from the upper mantle. Nature 278, 156–159. (doi:10.1038/278156a0)
Hilton, D. R., Fischer, T. P. & Marty, B. 2002 Noble gases and volatile recycling at subduction
zones. Rev. Miner. Geochem. 47, 319–370. (doi:10.2138/rmg.2002.47.9)
Holland, G. & Ballentine, C. J. 2006 Seawater subduction controls the heavy noble gas composition
of the mantle. Nature 441, 186–191. (doi:10.1038/nature04761)
Honda, M., McDougall, I., Patterson, D. B., Doulgeris, A. & Clague, D. A. 1991 Possible solar
noble gas component in Hawaiian basalts. Nature 349, 149–151. (doi:10.1038/349149a0)
Hünemohr H. 1989 Edelgase in U- und Th-reichen Mineralen und die Bistimmung der
21Ne-Dicktarget der 18(a,n)21Ne-Kernreaction im Bereich 4.0–8.8 MeV. PhD thesis,
Johannes-Gutenberg-Universität.
Jochum, K. P., Hofmann, A. W., Ito, E., Seufert, H. M. & White, W. M. 1983 K, U and Th in midocean ridge basalt glasses and heat-production, K/U and K/Rb in the mantle. Nature 306,
431–436. (doi:10.1038/306431a0)
Kellogg, L. H. & Wasserburg, G. J. 1990 The role of plumes in mantle helium fluxes. Earth Planet.
Sci. Lett. 99, 276–289. (doi:10.1016/0012-821X(90)90116-F)
Kellogg, L. H., Hager, B. H. & van der Hilst, R. D. 1999 Compositional stratification in the deep
mantle. Science 283, 1881–1884. (doi:10.1126/science.283.5409.1881)
Kipfer, R., Aeschbach-Hertig, W., Peeters, F. & Stute, M. 2002 Noble gases in lakes and ground
waters. Rev. Miner. Geochem. 47, 615–700. (doi:10.2138/rmg.2002.47.14)
Kurz, M. D., Jenkins, W. J. & Hart, S. R. 1982 Helium isotopic systematics of oceanic islands and
mantle heterogeneity. Nature 297, 43–47. (doi:10.1038/297043a0)
Kunz, J., Staudacher, T. & Allègre, C. J. 1998 Plutonium-fission xenon found in Earth’s mantle.
Science 280, 877–880. (doi:10.1126/science.280.5365.877)
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
What noble gases tell us about the mantle
4201
Lassiter, J. C. 2004 Role of recycled oceanic crust in the potassium and argon budget of the Earth:
toward a resolution of the ‘missing argon’ problem. Geochem. Geophys. Geosyst. 5, Q11012.
(doi:10.1029/2004GC000711)
Lupton, J. E. & Craig, H. 1975 Excess 3He in oceanic basalts, evidence for terrestrial primordial
helium. Earth Planet. Sci. Lett. 26, 133–139. (doi:10.1016/0012-821X(75)90080-1)
Mahaffy, P. R., Donahue, T. M., Atrey, S. K., Owen, T. C. & Niemann, H. B. 1998 Galileo probe
measurements of D/H and 3He/4He in Jupiter’s atmosphere. Space Sci. Rev. 84, 251–263.
(doi:10.1023/A:1005091806594)
Matsuda, J. & Nagao, K. 1986 Noble gas abundances in a deep-sea sediment core from eastern
equatorial Pacific. Geochem. J. 20, 71–80.
Mazor, E., Heymann, D. & Anders, E. 1970 Noble gases in carbonaceous chondrites. Geochim.
Cosmochim. Acta 34, 781–824. (doi:10.1016/0016-7037(70)90031-1)
Mizukami, T. & Wallis, S. R. 2005 Structural and petrological constraints on the tectonic evolution
of the garnet–lherzolite facies Higashi-akaishi peridotite body, Sanbagawa belt, SW Japan.
Tectonics 24, TC6012. (doi:10.1029/2004TC001733)
Mizuno, H., Nakazawa, K. & Hayashi, C. 1980 Dissolution of the primordial rare gases into the
molten Earth’s material. Earth Planet. Sci. Lett. 50, 202–210. (doi:10.1016/0012-821X(80)
90131-4)
Montelli, R., Nolet, G., Dahlen, F. A., Masters, G., Engdahl, E. R. & Hung, S. H. 2004 Finitefrequency tomography reveals a variety of plumes in the mantle. Science 303, 338–343. (doi:10.
1126/science.1092485)
Montelli, R., Nolet, G., Dahlen, F. A. & Masters, G. 2006 A catalogue of deep mantle plumes: new
results from finite-frequency tomography. Geochem. Geophys. Geosys. 7, Q11007. (doi:10.1029/
2006GC001248)
Morbidelli, A. 2002 Modern integrations of solar system dynamics. Annu. Rev. Earth Planet. Sci.
30, 89–112. (doi:10.1146/annurev.earth.30.091201.140243)
Moreira, M., Kunz, J. & Allègre, C. J. 1998 Rare gas systematics in popping rock: isotopic and
elemental compositions in the upper mantle. Science 279, 1178–1181. (doi:10.1126/science.279.
5354.1178)
Niedermann, S., Bach, W. & 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, 2697–2715. (doi:10.1016/S0016-7037(97)00102-6)
O’Nions, R. K. & Oxburgh, E. R. 1983 Heat and helium in the Earth. Nature 306, 429–431. (doi:10.
1038/306429a0)
O’Nions, R. K. & Tolstikhin, I. N. 1994 Behaviour and residence times of lithophile and rare gas
tracers in the upper mantle. Earth Planet Sci. Lett. 124, 131–138. (doi:10.1016/0012821X(94)00070-0)
Ozima, M. & 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. (doi:10.1016/0012-821X(91)90117-Z)
Parsons, B. 1981 The rates of plate creation and consumption. Geophys. J. R. Astron. Soc. 67,
437–448. (doi:10.1111/j.1365-246x.1981.tb02759.x)
Patterson, D. B., Honda, M. & McDougall, I. 1990 Atmospheric contamination—a possible source
for heavy noble gases in basalts from Loihi Seamount, Hawaii. Geophys. Res. Lett. 17, 705–708.
(doi:10.1029/GL017i006p00705)
Pepin, R. O. 1998 Isotopic evidence for a solar argon component in the Earth’s mantle. Nature 394,
664–667. (doi:10.1038/29272)
Pepin, R. O. 2000 On the isotopic composition of primordial xenon in terrestrial planet
atmospheres. Space Sci. Rev. 92, 371–395. (doi:10.1023/A:1005236405730)
Pepin, R. O. & Phinney, D. 1976 The formation interval of the Earth. Proc. 7th Lunar and
Planetary Science Conf., Lunar and Planetary Institute, Houston, TX, pp. 682–684.
Pepin, R. O. & Porcelli, D. 2002 Origin of noble gases in the terrestrial planets. Rev. Miner.
Geochem. 47, 191–246. (doi:10.2138/rmg.2002.47.7)
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
4202
C. J. Ballentine and G. Holland
Pepin, R. O. & Porcelli, D. 2006 Xenon isotope systematics, giant impacts, and mantle degassing
on the early Earth. Earth Planet. Sci. Lett. 250, 470–485. (doi:10.1016/j.epsl.2006.08.014)
Pepin, R. O., Becker, R. H. & Schlutter, D. J. 1999 Irradiation records in regolith materials, I:
isotopic compositions of solar-wind neon and argon in single lunar mineral grains. Geochim.
Cosmochim. Acta 63, 2145–2162. (doi:10.1016/S0016-7037(99)00002-2)
Phinney, D., Tennyson, J. & Frick, U. 1978 Xenon in CO2 well gas revisited. J. Geophys. Res. 83,
2313–2319. (doi:10.1029/JB083iB05p02313)
Podosek, F. A., Woolum, D. S., Cassen, P. & Nichols, R. H. 2000 Solar gases in the Earth by solar
wind irradiation? Geochim. Cosmochim. Acta 64, 804.
Porcelli, D. & Ballentine, C. J. 2002 Models for the distribution of terrestrial noble gases and
evolution of the atmosphere. Rev. Miner. Geochem. 47, 411–480. (doi:10.2138/rmg.2002.47.11)
Porcelli, D. & Elliot, T. 2008 The evolution of He isotopes in the convecting mantle and the
preservation of high 3He/4He ratios. Earth Planet. Sci. Lett. 269, 175–185. (doi:10.1016/j.epsl.
2008.02.002)
Porcelli, D. & Halliday, A. N. 2001 The core as a possible source of mantle helium. Earth Planet.
Sci. Lett. 192, 237–251. (doi:10.1016/S0012-821X(01)00493-9)
Porcelli, D. Pepin, R. O. The origin of noble gases and major volatiles in the terrestrial planets. In
The atmosphere (ed. R.F. Keeling). Treatise on geochemistry (eds H. D. Holland & K. K.
Turekian), vol. 4, pp. 319–342. Oxford, UK: Elsevier-Pergamon.
Porcelli, D. & Wasserburg, G. J. 1995 Mass transfer of helium, neon, argon, and xenon through a
steady-state upper mantle. Geochim. Cosmochim. Acta 59, 4921–4937. (doi:10.1016/00167037(95)00336-3)
Porcelli, D., Woolum, D. S. & Cassen, P. 2001 Deep Earth rare gases: initial inventories, capture
from the solar nebula, and losses during Moon formation. Earth Planet. Sci. Lett. 193, 237–251.
(doi:10.1016/S0012-821X(01)00493-9)
Saal, E. S., Hauri, E. H., Langmuir, C. H. & Perfit, M. R. 2002 Vapour undersaturation in
primitive mid-ocean-ridge basalt and the volatile content of Earth’s upper mantle. Nature 419,
451–455. (doi:10.1038/nature01073)
Sarda, P., Staudacher, T. & Allègre, C. J. 1988 Neon isotopes in submarine basalts. Earth Planet.
Sci. Lett. 91, 73–88. (doi:10.1016/0012-821X(88)90152-5)
Sarda, P., Moreira, M. & Staudacher, T. 1999 Argon–lead isotopic correlation in Mid-Atlantic
Ridge basalts. Science 283, 666–668. (doi:10.1126/science.283.5402.666)
Scherer, P. & Schultz, L. 2000 Noble gas record, collisional history, and pairing of CV, CO, CK and
other carbonaceous chondrites. Meteorit. Planet. Sci. 35, 145–153.
Sharif-Zade, V. B., Shukolyukov, Y. A., Gerling, E. K. & Ashkinadze, G. S. 1972 Neon isotopes in
radioactive minerals. Geochem. Int. 10, 199–207.
Shukolyukov, Y. A., Sharif-Zade, V. B. & Ashkinadze, G. S. 1973 Neon isotopes in natural gases.
Geochem. Int. 4, 346–354.
Staudacher, T. 1987 Upper mantle origin for Harding County well gases. Nature 325, 605–607.
(doi:10.1038/325605a0)
Staudacher, T. & Allègre, C. J. 1982 Terrestrial xenology. Earth Planet. Sci. Lett. 188, 211–219.
Staudacher, T. & Allègre, C. J. 1988 Recycling of oceanic crust and sediments: the noble gas
subduction barrier. Earth Planet. Sci. Lett. 89, 173–183. (doi:10.1016/0012-821X(88)90170-7)
Suess, H. E. 1949 The abundance of noble gases in the Earth and the cosmos. [In German.] J. Geol.
57, 600–607.
Sumino, H., Burgess, R., Mizukami, T., Wallis, S. R. & Ballentine, C. J. 2007 Subducted noble gas
and halogen preserved in wedge mantle peridotite from the Sanbagawa belt, SW Japan.
Geochim. Cosmochim. Acta 71, A985–A985.
Tackley, P. J. 2000 Mantle convection and plate tectonics: toward an integrated physical and
chemical theory. Science 288, 2002–2007. (doi:10.1126/science.288.5473.2002)
Tolstikhin, I. N. 1975 Helium isotopes in the Earth’s interior and in the atmosphere: a degassing
model of the Earth. Earth Planet. Sci. Lett. 26, 88–96. (doi:10.1016/0012-821X(75)90180-6)
Phil. Trans. R. Soc. A (2008)
Downloaded from http://rsta.royalsocietypublishing.org/ on June 16, 2017
What noble gases tell us about the mantle
4203
Tolstikhin, I. N. & Hofmann, A. W. 2005 Early crust on top of the Earth’s core. Phys. Earth
Planet. Inter. 148, 109–130. (doi:10.1016/j.pepi.2004.05.011)
Tolstikhin, I. N. & O’Nions, R. K. 1994 The Earth’s missing xenon: a combination of early
degassing and of rare gas loss from the atmosphere. Chem. Geol. 115, 1–6. (doi:10.1016/00092541(94)90142-2)
Torgersen, T. 1989 Terrestrial helium degassing fluxes and the atmospheric helium budget:
implications with respect to the degassing processes of continental crust. Chem. Geol. 79, 1–14.
Torgersen, T. & Clarke, W. B. 1985 Helium accumulation in groundwater I: an evaluation of
sources and the continental flux of crustal 4He in the Great Artesian Basin, Australia. Geochim.
Cosmochim. Acta 49, 1211–1218. (doi:10.1016/0016-7037(85)90011-0)
Trieloff, M., Kunz, J., Clague, D. A., Harrison, D. & Allègre, C. J. 2000 The nature of pristine
noble gases in mantle plumes. Science 288, 1036–1038. (doi:10.1126/science.288.5468.1036)
Trieloff, M., Kunz, J., Clague, D. A., Harrison, D. & Allègre, C. J. 2001 Reply to comment on noble
gases in mantle plumes. Science 291, 2269. (doi:10.1126/science.291.5512.2269a)
Trieloff, M., Kunz, J. & Allègre, C. J. 2002 Noble gas systematics of the Réunion mantle plume
source and the origin of primordial noble gases in Earth’s mantle. Earth Planet. Sci. Lett. 200,
297–313. (doi:10.1016/S0012-821X(02)00639-8)
Van der Hilst, R. D., Winiyantoro, S. & Engdahl, E. R. 1997 Evidence for deep mantle circulation
from global tomography. Nature 386, 578–584. (doi:10.1038/386578a0)
Van Keken, P. E. & 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. (doi:10.1016/S0012-821X(98)00023-5)
Van Keken, P. E. & Ballentine, C. J. 1999 Dynamical models of mantle volatile evolution and the
role of phase transitions and temperature-dependent rheology. J. Geophys. Res. Solid Earth 104,
7137–7151. (doi:10.1029/1999JB900003)
Van Keken, P. E., Ballentine, C. J. & Porcelli, D. 2001 A dynamical investigation of the heat and
helium imbalance. Earth Planet. Sci. Lett. 188, 421–434. (doi:10.1016/S0012-821X(01)00343-0)
Verchovsky, A. B. & Shukolyukov, Y. A. 1976 Isotopic composition of neon in the crustal rocks
and the origin of Ne21rad in natural gases. Geochem. Int. 13, 95–98.
Watson, B. E., Thomas, J. B. & Cherniak, D. J. 2007 40Ar retention in the terrestrial planets.
Nature 449, 299–304. (doi:10.1038/nature06144)
Wieler, R. 2002 Noble gases in the solar system. Rev. Miner. Geochem. 47, 21–70. (doi:10.2138/
rmg.2002.47.2)
Wetherill, G. W. 1954 Variations in the isotopic abundances of neon and argon extracted from
radioactive minerals. Phys. Rev. 96, 679–683. (doi:10.1103/PhysRev.96.679)
Wetherill, G. W. 1975 Radiometric chronology of the early solar system. Annu. Rev. Nucl. Sci. 25,
283–328. (doi:10.1146/annurev.ns.25.120175.001435)
Wetherill, G. W. 1994 Provenance of the terrestrial planets. Geochim. Cosmochim. Acta 58,
4513–4520. (doi:10.1016/0016-7037(94)90352-2)
Yatsevich, I. & Honda, M. 1997 Production of nucleogenic neon in the Earth from natural
radioactive decay. J. Geophys. Res. Solid Earth 102, 10 291–10 298. (doi:10.1029/97JB00395)
Yokochi, R. & Marty, B. 2004 A determination of the neon isotopic composition of the deep
mantle. Earth Planet. Sci. Lett. 225, 77–88. (doi:10.1016/j.epsl.2004.06.010)
Zahnle, K., Arndt, N., Cockell, C., Halliday, A., Nisbet, E., Selsis, F. & Sleep, N. H. 2007
Emergence of a habitable planet. Space Sci. Rev. 129, 35–78. (doi:10.1007/s11214-007-9225-z)
Zhou, Z. & Ballentine, C. J. 2006 He-4 dating of groundwaters associated with hydrocarbon
reservoirs. Chem. Geol. 226, 309–327. (doi:10.1016/j.chemgeo.2005.09.030)
Phil. Trans. R. Soc. A (2008)