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date: 15 October 2018

Noble Gases

Summary and Keywords

Although the second most abundant element in the cosmos is helium, noble gases are also called rare gases. The reason is that they are not abundant on terrestrial planets like the Earth, which is characterized by orders of magnitude depletion of—particularly light—noble gases when compared to the cosmic element abundance pattern. Indeed, geochemical depletion and enrichment processes mean that noble gases are highly versatile tracers of planetary formation and evolution. When our solar system formed—or even before—small grains and first condensates incorporated small amounts of noble gases from the surrounding gas of solar composition, resulting in depletion of light He and Ne relative to heavy Ar, Kr, and Xe, leading to the “planetary type” abundance pattern. Further noble gas depletion occurred during flash heating of mm- to cm-sized objects (chondrules and calcium, aluminum-rich inclusions), and subsequently during heating—and occasionally differentiation—on small planetesimals, which were precursors of planets. Some of these objects are present today in the asteroid belt and are the source of many meteorites. Many primitive meteorites contain very small (micron to sub-micron size) rare grains that are older than our Solar System and condensed billions of years ago in in the atmospheres of different stars, for example, Red Giant stars. These grains are characterized by nucleosynthetic anomalies, in particular the noble gases, such as so-called s-process xenon.

While planetesimals acquired a depleted noble gas component strongly fractionated in favor of heavy noble gases, the Sun and also gas giants like Jupiter attracted a much larger amount of gas from the protosolar nebula by gravitational capture. This resulted in a cosmic or “solar type” abundance pattern, containing the full complement of light noble gases. In contrast, terrestrial planets accreted from planetesimals with only minor contributions from the gaseous component of the protosolar nebula, which accounts for their high degree of depletion and essentially “planetary” elemental abundance pattern. The strong depletion in noble gases facilitates their application as noble gas geo- and cosmochronometers; chronological applications are based on being able to determine noble gas isotopes formed by radioactive decay processes, for example, 40Ar by 40K decay, 129Xe by 129I decay, or fission Xe from 238U or 244Pu decay. Particularly ingrowth of radiogenic xenon is only possible due to the depletion of primordial nuclides, which allows insight into the chronology of fractionation of lithophile parent nuclides and atmophile noble gas daughters. Applied to large-scale planetary reservoirs, this helps to elucidate the timing of mantle degassing and evolution of planetary atmospheres. Applied to individual rocks and minerals, it allows radioisotope chronology using short-lived (e.g., 129I–129Xe) or long-lived (e.g., 40K–40Ar) systems.

The dominance of 40Ar in the terrestrial atmosphere allowed von Weizsäcker to conclude that most of the terrestrial atmosphere originated by degassing of the solid Earth, which is an ongoing process today at mid-ocean ridges, as indicated by outgassing of primordial helium from newly forming ocean crust. Mantle degassing was much more massive in the past, with most of the terrestrial atmosphere probably formed during the first few 100 million years of Earth’s history, in response to major evolutionary processes of accretion, terrestrial core formation, and the terminal accretion stage of a giant impact that formed our Moon. During accretion, solar noble gases were added to the mantle, presumably by solar wind irradiation of the small planetesimals and dust accreting to form the Earth. While the Moon-forming impact likely dissipated a major fraction of the primordial atmosphere, today’s atmosphere originated by addition of a late veneer of asteroidal and possibly cometary material combined with a decreasing rate of mantle degassing over time. As other atmophile elements behave similarly to noble gases, they also trace the origin of major volatiles on Earth, for example, water, nitrogen, and carbon.

Keywords: noble gases, isotopes, atmosphere, mantle, radiogenic isotopes, solar, planetary, helium, argon, neon, xenon

Historical Notes

Discovery of Noble Gases, Radioactivity and the Age of the Earth

Compared to other stable elements, the noble gases were discovered late in the history of chemistry, due to their chemical inertness and scarcity. In 1784, Henry Cavendish experimentally separated nitrogen and oxygen from air and discovered a rather small gaseous residue. It was considered an experimental error. During the 1868 solar eclipse, a previously unknown line in the spectrum of the solar chromosphere was identified and named helium. The discovery of the noble gases was completed in the late 19th century by Rayleigh and Ramsey (Rayleigh, 1894; Rayleigh & Ramsay, 1895; Ramsay, 1904), almost simultaneously with the discovery of radioactivity by Henri Becquerel. The latter phenomenon was soon realized to have potential for measuring the age of geologic samples, and the first isotopic age reported involved a noble gas: Rutherford (1906) measured the U-content of ore minerals and the helium-alpha particles produced by U-decay. Rutherford obtained an age of the Earth of a few hundred million years, which at that time was in contrast to other estimates that assumed a considerably younger age. A competing method of calculating the age of the Earth at that time was based on the recognition that the Earth was hot in its interior and that the temperature profile was a result of cooling from an initial hot state. Isaac Newton had previously estimated that an iron sphere the size of the Earth would need 50,000 years to cool. In the 19th century, the ideas could be refined, as the temperature increase and gradient in the terrestrial crust was known (typically 1/28°C increase per meter), and the physical theory of heat conduction had been established. In 1862, William Thomson, later Lord Kelvin, assumed the Earth was initially 3,870°C—at that time the presumed melting point of rock—and lost heat with the typical heat conductivities of sand, sandstone, or basaltic rock. To achieve the typical crustal temperature gradient, Thomson had to assume an age of the Earth of 100 million years in his calculations. In 1893, King used experimentally determined melting temperatures of diabase (a crystalline basaltic rock) and excluded models above this temperature, because high degrees of partial melts in the Earth’s interior were not compatible with the stability of the terrestrial crust against the tidal forces exerted by the Moon. His refined value was 24 million years, a value later confirmed by Kelvin. These relatively young age values of the Earth were questioned by geologists, who had, however, only qualitative arguments based on, for example, sedimentation processes. A quantitative measure of geological time only became available after the discovery of radioactivity and the use of natural radioactive elements, for example, isotopes of uranium, thorium, rubidium, and potassium. The helium isotope (4He) used by Rutherford for measuring the first U-He age of rocks and minerals is emitted as an alpha particle. These particles were later used by Rutherford to conduct his famous irradiation experiment of gold foils, which gave rise to his model of largely empty atoms with electrons orbiting a spatially extremely confined nucleus. Soon after Rutherford’s first study of U-He dating, Boltwood (1907) developed the U-Pb method, which measures Pb, another daughter element arising by radioactive decay of U and Th. After several refinements of isotope mass spectrometry, Patterson in 1953 used both extraterrestrial and terrestrial samples to compute an age of the Earth of 4.55 Ga, which within uncertainty is indistinguishable from the value accepted today.

Radioisotope Dating by Refined Noble Gas Mass Spectrometry

Noble gases are depleted on Earth by more than seven orders of magnitude compared to solar abundances, leading to their absolute abundances being very low (Ozima & Podosek, 2002). For example, a typical value for the concentration of Xe in a terrestrial rock is 10−11 cm3/g, equivalent to 0.00006 ppb. The reason is that they are extremely volatile and are depleted in the inner solar system. Nevertheless, using modern analytical techniques, such low xenon concentrations can be adequately measured. A one-gram sample provides sufficient Xe, allowing 5% uncertainty in absolute abundance and 1% or less uncertainty in relative abundances of the major isotopes. Such precision is rendered possible by the low natural background and by the intrinsic inertness of the noble gases; they can be measured using mass spectrometry in the absence of interfering molecular ions because noble gases can be purified by removing most reactive species using hot getters (e.g., Ti, Zr-Al metal alloys) in the extraction line.

The relative scarcity of the primordial noble gases is also an advantage because it enables the sometimes relatively small excess produced by radioactive decay on specific isotopes to be precisely determined. For example, the maximum excess of 129Xe in mantle-derived rocks is on the order of 10% and would hardly be detectable if primordial Xe was one order of magnitude more abundant. Another example is K-Ar dating, whereby 90% of 40K decays to 40Ca; however, 96.9% of Ca is comprised of 40Ca, so that in situ decay-produced 40Ca is virtually unrecognizable except in some unusual cases where the K/Ca is very high and the samples are billions of years old. In contrast, the 10% of the 40K decays that produce 40Ar overwhelm any Ar initially present and make the K-Ar method one of the most widely used isotope chronometers.

In the development of K-Ar dating (Schaeffer & Zähringer, 1966; McDougall & Harrison, 1988), an early important contribution was by Campbell and Wood (1906), who demonstrated that an alkali metal alloy containing potassium was radioactive. However, it was originally believed that a contaminating heavy element was the source of radioactivity, an idea unsuccessfully pursued until the late 1920s. The K-isotopes with masses 39 and 41 were discovered by Aston (1921) using his recent invention of the mass spectrograph. However, it became apparent that neither of these isotopes was likely to be radioactive. Klemperer (1935) and Newman and Walke (1935) considered 40K to be the radioactive isotope, decaying to 40Ca by β‎- decay. Shortly thereafter, Newman and Walke (1935) raised the possibility that 40Ar could be a second decay product. Using improved mass spectrometric techniques, Nier (1935) then confirmed the existence of 40K and determined its isotopic abundance to 39K/40K=8,600 ± 10%. Smythe and Hemmendinger (1937) separated the three naturally occurring K isotopes using mass spectrometry and demonstrated unequivocally the radioactive nature of 40K. Von Weizsäcker (1937) also argued for the dual decay mode of 40K into 40Ar and 40Ca, based on the observed overabundance of 40Ar in the terrestrial atmosphere. Realizing that 40Ar must be derived from the solid Earth by degassing was the first step leading to the use of noble gases as geochemical tracers of global evolutionary processes. Von Weizsäcker (1937) also suggested the correct decay mechanism to be electron capture, as the emission of positrons had not been observed prior to this (Sitte, 1935), an important advance taking into account that electron capture had not been proved until then for any other element. Aldrich and Nier (1948) measured argon released from ancient K-bearing minerals and confirmed the “excess” of 40Ar over 36Ar, that is, argon with a 40Ar/36Ar ratio higher than the atmospheric value of 296 (introduced by Nier) They noted that technical improvements might facilitate geochronological applications. Further important steps in developing K-Ar dating were made by the group of W. Gentner (Smits & Gentner, 1950; Gentner et al., 1953a, 1953b) and by Wasserburg and Hayden (1955). Today the 40Ar-39Ar dating technique (Merrihue & Turner, 1966; McDougall & Harrison, 1988) has largely replaced conventional K-Ar dating (Schaeffer & Zähringer, 1966).

Early Advances in Noble Gas Geo- and Cosmochemistry

Major developments in noble gas geo- and cosmochemistry were achieved in the 20th century, stimulated by continuous improvements in noble gas mass spectrometric techniques and instrumentation. Von Weizsäcker concluded in 1937 that most of the 40Ar in the Earth’s atmosphere originated by 40K decay in the solid earth, and Brown (1952) and Turekian (1964) extended this by recognizing that Earth’s atmosphere is secondary and was formed by degassing of the solid Earth. Clarke et al. (1975) and Lupton and Craig (1975) found evidence for ongoing degassing processes from the release of primordial 3He at mid-ocean ridges. From the late 1960s onward, several researchers measured argon with high 40Ar/36Ar ratios in submarine basalt glasses erupting at mid-ocean ridges, and recognized that the primordial isotope 36Ar is much more depleted in the upper mantle than 40Ar. This is interpreted to indicate that degassing occurred mainly early in Earth history, before most of the 40K decayed to 40Ar (Dalrymple & Moore, 1968; Fisher, 1971; Fanale, 1971; Ozima, 1975). An even more stringent constraint for early massive mantle degassing comes from the excess of 129Xe related to decay of 129I, which was first detected by Reynolds (1960) in meteorites and Butler et al. (1963) in terrestrial well gases. 129I is a short-lived nuclide with only 16 Ma half-life that was present in the early solar system and has been used for isotopic dating of early Solar System processes (e.g., Swindle & Podosek, 1988; Swindle, 2002a). The detection of excess of 129Xe in mantle-derived rocks (Staudacher et al., 1982) implies that major degassing should have occurred before 129I had fully decayed (within a few half-lives), more precise evaluations yield a value of ca. 100 Ma after formation of the Earth (e.g., Allègre et al., 1986).

The concept of degassing of the solid Earth has led to the view that the terrestrial atmosphere originated very early in Earth’s history by degassing of the mantle. One implication of early degassing is that, aside from radiogenic ingrowth, noble gas isotope ratios should be very similar in the mantle and atmosphere, particularly ratios involving isotopes not affected by radiogenic ingrowth, including 36Ar/38Ar and 20Ne/22Ne. Therefore it was both surprising and challenging when Honda et al. (1991) and Hiyagon et al. (1992) discovered that terrestrial mantle neon had a solar isotopic composition different from the planetary isotopic composition of the atmosphere. This solar signature was initially interpreted to represent an unfractionated solar gas with 20Ne/22Ne=13.8 and gave rise to models involving the gravitational attraction of solar nebula gas dissolved in a terrestrial magma ocean, followed by hydrodynamic escape and elemental and isotopic fractionation of the proto-atmosphere (Sasaki & Nakazawa, 1990; Pepin, 1991). However, it was shown later that the actual Ne endmember was 20Ne/22Ne=12.5 (Trieloff et al., 2000, 2002; Ballentine et al., 2005), identified as the solar wind implanted component known as Ne-B present in meteorites (Black & Pepin, 1969; Black, 1972). This requires that most solar gases were acquired from the solar wind after solar nebula dissipation, not before.

Many advances in geochemistry are intrinsically coupled to complementary developments in cosmochemistry. For example, the age of the Earth could only be derived precisely by Patterson (1953), because meteorites and the Earth were considered to have formed almost simultaneously. This concept had already been advocated in the 18th century by Kant and Laplace, who assumed that all solar system bodies were formed by a collapsing cloud of interstellar gas and dust, forming a protoplanetary disc due to conservation of angular momentum, causing today’s structure of our planetary systems, with all major bodies orbiting the sun in one direction and roughly a single plane.

The first measurements on meteorites already indicated the existence of two fundamental noble gas abundance patterns in the solar system; one was called “solar” and was represented by meteoritic regolith breccias that experienced solar wind ion implantation, the other was “planetary” and contained fractionated noble gases with an excess in heavy isotopes (Gerling & Levskii, 1956; Pepin, 1967; Black & Pepin, 1969; Mazor et al., 1970; Black, 1972). The planetary component was later found to be frequently bound to carbonaceous phases revealed by acid etching experiments (e.g., Alaerts et al., 1979): Most planetary gases are hosted by a phase surviving HF/HCl dissolution of most silicates. Revealed by further physical and chemical treatments, the acid residues were found to contain presolar grains (Lewis et al., 1975, 1987, 1990; Zinner, 1998) hosting some exotic nucleosynthetic noble gas components, for example s-process Xe. However, most planetary gases were released upon etching the HF/HCl residue (typically 1% of the original mass) with HNO3, which removed negligible further mass, but more than 90% of planetary xenon. Lacking knowledge of the physical or mineralogical nature of this phase, it was termed Q (“Quintessence”), which is still in use today, as it persistently eludes identification.

State of the Art

Noble Gas Isotopes and Their Provenances

Noble gas isotopes may be differentiated into two types: those that are nearly pure primordial isotopes which retain their original solar system abundances, and those isotopes which experience increase by radioactive decay or nuclear reactions. The extent to which radiogenic contributions are substantial or recognizable depends on the ratio of radioactive parent nuclides and primordial isotopes (e.g., K/Ar, U/Xe, U/He) which may vary in specific geochemical or cosmochemical environments. Ingrowth by nuclear reactions is also distinguished in terminology: noble gases from radioactive decay are termed “radiogenic,” those from secondary nuclear reactions “nucleogenic,” and those formed by interaction with cosmic ray particles “cosmogenic.”

This article will focus on noble gas acquisition during planetary formation and differentiation, with a primary focus on Earth. While radiogenic noble gases play a key role in understanding history and differentiation of the mantle-atmosphere system, neither the full width of geochronological achievements of Ar-Ar or (U-Th)/He dating nor the spectrum of applications in crustal, ocean, and surface processes will be represented. The reader is directed to respective reviews (McDougall & Harrison, 1988; Kelley, 2002; Farley, 2002; Niedermann, 2002; Ballentine et al., 2002; Ballentine & Burnard, 2002; Kipfer et al., 2002; Schlosser & Winckler, 2002).

Noble Gas Isotopes at the Origin of the Solar System—Interstellar Gas and Presolar Grains

Our solar system formed by collapse of an interstellar cloud of gas and dust about 4.57 Ga ago (e.g., Trieloff & Palme, 2006), dated by the oldest objects found in meteorites, the calcium aluminum rich inclusions (Amelin et al., 2002; Trieloff, 2009). Approximately 99% was gas (mainly hydrogen and helium), with only ~1% solid “condensable” elements in the form of mostly submicrometer-sized dust grains (Greenberg, 1984), which either originated in the stellar winds of asymptotic giant branch (AGB) stars and supernovae explosions (Anders & Zinner, 1993; Zinner, 1998) or formed directly in the interstellar medium (Zhukovska et al., 2008; Altobelli et al., 2016). The abundances of the elements in the collapsing cloud represent average cosmic abundances produced by stellar nucleosynthesis. Stellar nucleosynthesis produces heavy elements by nuclear fusion processes. In general, heavy elements with a high proton number are less abundant, with exceptions given by the stability of atomic nuclei, for example, nuclei with even numbers of protons or neutrons being more stable and more abundant than odd-numbered nuclei. By synthesizing nuclei lighter than Fe, nuclear reactions release net energy, keeping a star stable. Elements heavier than Fe can only be synthesized during supernova explosions. As galactic chemical evolution proceeds with time, there is a general increase of heavy elements or—astronomically speaking—increasing “metallicity.” The most significant change is in the ratio of heavy elements to hydrogen, and only to a minor extent the abundance ratios within condensable elements.

Table 1. Noble Gas Isotope Ratios in Extraterrestrial and Terrestrial Reservoirs

Ratio

Solar Gasa

Solar Windb

Solar Wind Implantedb

Planetary Qc

Planetary P3 Nanodiamondsd

Planetary -CV chondritese

Planetary Terrestrial airf

MORBg

OIB (Hawaii)g

4He/3He

6,000±2000

2,150 ±10

2,150 ±10

8,130±130

7,400±500

715 000

89 000±11 000

20 300±3 000

20Ne/22Ne

13.36

13.777±0.01

12.52±0.18/12.6

10.67/10.11

8.91±0.06

9.80

12.5±0.2

12.5±0.3

21Ne/22Ne

0.03236

0.03289±0.00007

0.03118±0.00048

0.0294±0.0010

0.029±0.001

0.029

0.0595±0.0003

0.0362±0.0003

36Ar/38Ar

5.37

5.47±0.03

5.38±0.12/5.29

5.34±0.02

5.26±0.03

5.32

~5.32

5.33±0.01

40Ar/36Ar

«1

«1

«1

«1

«1

296

32 400±4 200

8 000 ±1 000

129Xe/130Xe

6.306

6.436

6.557

6.496

7.79±0.25

6.75±0.06

136Xe/130Xe

1.819

1.954

1.948

2.176

2.616±0.100

2.258±0.020

3He/22Ne

1.8

4.8

4.8

0.15

1.01

(0.7)

(4.46x10−7)

7.4

7.0

36Ar/22Ne

0.34

0.32

0.32

253

45

10.3

18.8

9.7

7.0

84Kr/22Ne

1.39×10-4

1.30×10−4

1.30×10−4

2.7

0.30

0.069

0.386

130Xe/22Ne

1.03×10-6

0.97×10−6

0.97×10−6

0.540

0.018

0.0094

0.00211

0.0051

0.0049

Notes: (a) Solar gas neon and argon (outer convection zone) values by Heber et al. (2012), helium by Mahaffy et al. (1998, 2000). Elemental composition calculated from Lodders et al. (2009): 4He/132Xe=1.745×109; 20Ne/132Xe=2.13 ×106; 36Ar/132Xe =5.45 ×104; 84Kr/132Xe =22.1, and 4He/3He=6,000 (Mahaffy et al., 1998, 2000), as solar gas did not contain a 3He component by early solar deuterium burning, so had a higher 4He/3He ratio when compared to present-day solar wind of 2,150±10 (Heber et al., 2012).

(b) Present-day He, Ne, Ar solar wind values by Heber et al. (2012)—Ott (2014) also summarizes other studies (Meshik et al., 2007, 2014; Pepin et al., 2012). Xe by Meshik et al. (2014). Solar wind implanted isotopic composition by Heber et al. (2012) for He. Ne by Black (1972) and Moreira and Charnoz (2016), Ar by Black (1972) and Raquin and Moreira (2009). Elemental composition calculated from Lodders et al. (2009); see footnotea.

(c) Busemann et al. (2000), Wieler et al. (1992), and Ott (2014).

(d) Huss and Lewis (1994a, b), and Ott (2014).

(e) Elemental compositions of CV chondrites by Matsuda et al. (1993). Isotopic compositions are a mixture of planetary and solar noble gases and hence not listed. Helium in CV chondrites is affected by diffusional loss (Mazor et al., 1970) and hence 3He/22Ne only a lower limit.

(f) Ozima and Podosek (2002). In the terrestrial atmosphere, He is lost to space within short time scales, so 3He/22Ne ratios are not suitable for meaningful comparisons.

(g) Evaluations of primordial isotope ratios (OIB) by Trieloff et al. (2000, 2002). Radiogenic isotope composition by Trieloff and Kunz (2005). MORB and OIB elemental ratios derived by radiogenic excess ratios, and are consistent with Trieloff et al. (2002) values of 36Ar/22Ne=9.2±3, 84Kr/22Ne=0.33±0.05, 130Xe/22Ne=0.0055±00023, derived by averaging data at 20Ne/22Ne=12.5.

During solar system formation, the sun gravitationally acquired the main mass (>99%) of our solar system. Hence, solar elemental abundances (e.g., Lodders et al., 2009) largely correspond to cosmic abundances, including the volatile gaseous compounds and noble gases. For small planetary bodies or tiny dust grains, it is more difficult to acquire noble gases, although not impossible. If temperatures are sufficiently low, gases may be adsorbed on grain surfaces, particularly the heavier ones (Fanale & Cannon, 1972; Bernatowicz & Podosek, 1986). Another possibility is as ice clathrates (Sill & Wilkening, 1978; Lodders, 2004). A third mechanism could be ion implantation, for example ions from stellar winds—including our sun—or in the interstellar medium (e.g., Bernatowicz & Fahey, 1986; Ott, 2014).

Ion implantation may be mass selective or not, depending on the efficiency of ionization. At low ionization levels, heavy noble gases with lower ionization energy could be preferentially accelerated and implanted more efficiently than the lower mass elements helium or neon.

It is important to note that noble gases were already present in solids that accreted to the solar system; presolar grains including SiC and graphite are carriers of s-process nuclides, a signature of stellar nucleosynthesis that involves slow neutron capture at low neutron densities, related to asymptotic giant branch (AGB) stars (Amari et al., 1995; Amari, 2006). Nanodiamonds (it is uncertain what fraction is presolar) may contain solar-like or r- and p- process nuclides (Huss & Lewis, 1994a, 1994b; Dai et al., 2002; Stroud et al., 2011). Finally, phase Q, the ill-defined phase only characterized by acid etching properties but somehow related to the aforementioned carbonaceous grains, was considered to have acquired noble gases during residence in the solar nebula (Wieler et al., 1992; Ozima et al., 1998; Busemann et al., 2000). The important point to highlight is that these phases are a source of noble gases present in accreting solids in the early solar system (provided they did not reside in hot regions of the solar nebula) before being incorporated into larger planetesimals and planets. To emphasize their importance further, in a typical planetary component of a CI chondrite, more than 60% of the 22Ne is from nanodiamonds, more than 20% from SiC and graphite, and slightly more than 10% from phase Q. Due to the different elemental abundances, the proportions for Xe are different, but presolar grains are similarly important: more than 90% of the xenon is from phase Q, and slightly more than 5% from nanodiamonds (e.g., Ott, 2014).

Planetary and Solar Abundance Patterns

The planets show a clear dichotomy between the gas and ice giant planets Jupiter, Saturn, Uranus, and Neptune, which have a high proportion of gaseous compounds, and the solid rocky terrestrial planets Mercury, Venus, Earth, and Mars. The depletion of gaseous compounds in inner solar system bodies is twofold. First, the terrestrial planets accreted inside the “snow line,” where water (and more volatile species like methane and ammonia) existed only in gaseous but not icy form and hence could not be accreted as solids. Second, the relatively low gravitational fields associated with the lower mass terrestrial planets means they were unable to capture and retain gases by gravitational forces. Nevertheless, the terrestrial planets have more or less dense atmospheres. For example, Venus and Mars have CO2-dominated atmospheres with 93 and 0.006 bar surface pressure, respectively. The abundance of chemically reactive components in planetary atmospheres is subject to complex chemical interactions; for example, most of the carbon dioxide on Earth is sequestered by the terrestrial carbon cycle and bound in carbonates. The chemically inert noble gases are only influenced by physical processes (e.g., ion implantation, adsorption, solubility in liquid phases like water or magma) or nuclear processes like radioactive decay.

Noble GasesClick to view larger

Figure 1. Noble gas abundances in planetary atmospheres normalized to cosmic element abundances (Ozima & Podosek, 2002; Lodders et al., 2009; Anders & Grevesse, 1989; Pepin, 1991). Units are atoms per Si atoms—for planets, atmospheric noble gas concentrations are divided by planetary Si. In the case of argon, the less abundant primordial argon isotope 36Ar is used rather than 40Ar, which stems mostly from 40K-decay. The low terrestrial 4He value is not relevant, as helium is lost on geologically short time scales from Earth’s atmosphere.

When noble gas abundances in planetary atmospheres are compared to relative the solar composition, as shown in Figure 1 (Ozima & Podosek, 2002; Swindle, 2002b; Lodders et al., 2009; Anders & Owen, 1977), two important observations can be made. First, noble gases are depleted by up to 13 orders of magnitude, and second, the light noble gases are more depleted than heavy noble gas isotopes. Patterns of Mars, Earth, and Venus are somewhat similar to CI chondrites (carbonaceous chondrites of Ivuna type), which contain noble gases bound in solid phases, acquired before or during solar nebula residence (Figure 1). Chondrites are undifferentiated, primitive solar system rocks that originate from small asteroid-sized bodies that accreted in the solar nebula (Trieloff & Palme, 2006; Trieloff, 2009). They escaped severe heating and did not undergo differentiation to form metallic cores and silicate mantles. The carbonaceous chondrites in particular contain appreciable amounts of water (up to 18%, partly in OH-bearing minerals) and carbon (up to 3.5%). Hence their noble gas inventory can be considered pristine. The noble gas abundance pattern of CI chondrites, which were derived from small asteroid-sized parent bodies and therefore unaffected by gravitational capture of gases, is consistent with a mass selective process like ion implantation or adsorption of noble gases having solar or cosmic abundance patterns (Ott, 2014).

The similar noble gas abundance patterns of the terrestrial planets (Figure 1) suggest that their volatile inventory was part of the accreting planetesimals, with the atmosphere forming by degassing of the incoming solids. Subsequent loss of noble gases, most likely initiated by energetic giant impacts during terminal accretion, explains the generally stronger depletion of rare gases in terrestrial planets relative to carbonaceous chondrites.

While the planetary noble gas abundance pattern is present in chondrites and planetary atmospheres, the solar-like abundance pattern is not restricted to the Sun, but is also found in Jupiter’s atmosphere (e.g., Mahaffy et al., 1998). Jupiter formed a core equivalent in size to about several Earth masses when the solar nebula was still extant, and was able to attract solar gases by its gravitational forces. However, gravitation is not the only way to collect solar gases. The sun emits a steady flow of matter, the solar wind. It resembles the bulk composition of the sun and consists of ions that are driven outward by the solar magnetic field. These ions have sufficient energy to be implanted into the outer few microns of grain surfaces exposed to space. The highest solar wind concentrations occur in small dust particles due to their high surface-to-volume ratio. Indeed, some chondrites contain a mixture of solar wind implanted and planetary noble gases, with the light noble gases helium and neon dominated by solar wind implanted gases and the heavier noble gases predominantly of planetary origin (e.g., Matsuda et al., 1993).

In addition to their differing noble gas abundance patterns, planetary and solar noble gases also differ in the isotopic compositions of their constituent noble gases. For example, as shown in Figure 2, the 20Ne/22Ne ratio of solar wind has a value of 13.8 (e.g., Benkert et al., 1993; Grimberg et al., 2006; Heber et al., 2009, 2012; Pepin et al., 2012), whereas planetary neon is depleted in the light isotope 20Ne with a lower ratio of 8.2 (e.g., Black & Pepin, 1969; Black, 1972).

The Early Earth: Acquisition of Solar Noble Gases

The Earth’s mantle was also found to contain a mixture of solar type He and Ne, and planetary type Ar, Kr, and Xe (Honda et al., 1991; Moreira et al., 1998; Kunz, 1999; Trieloff et al., 2000, 2002; Ballentine et al., 2005; Holland et al., 2009; Caracausi et al., 2016), while the terrestrial atmosphere seems to be a mixture of solar and planetary neon (Marty, 1989, 2012; Vogt et al., 2017, see Figure 2). Particularly important constraints on terrestrial accretion come from the two possibilities, if solar neon in Earth’s mantle was acquired as solar gas neon dissolved from a massive proto-atmosphere within the solar nebula (Sasaki & Nakazawa, 1990; Honda et al., 1991; Pepin, 1991; Hiyagon et al., 1992; Porcelli et al., 2001) or as solar wind by ion implantation into small accreting grains or bodies after nebular dissipation (Sasaki, 1991; Trieloff et al., 2000, 2002; Hopp & Trieloff, 2005; Ballentine et al., 2005; Moreira & Charnoz, 2016; Péron et al., 2017; Jaupart et al., 2017), with the latter scenario becoming increasingly accepted as the origin for most of the solar gases in the mantle.

The acquisition of solar wind implanted gases requires matter to be exposed to solar radiation, particularly the finer fraction of nebula matter. The initial solar nebula comprising gas and dust existed for 3–10 Ma (Haisch et al., 2001; Pfalzner et al., 2014). The early nebula was largely self-shielding, but nevertheless offered regions close to the surface of the protoplanetary disc where a small fraction of solids could have been irradiated (e.g., Sasaki, 1991). After dissipation of the solar nebula, terrestrial accretion continued for another ~100 Ma before the Earth gained its final mass. This time interval is largely derived from numerical models (e.g., Morbidelli et al., 2012) and terrestrial core formation ages using 182Hf-182W dating (Kleine et al., 2002, 2009; Halliday & Wood, 2009; Rubie et al., 2011, 2015). Large planetesimals were accreted by the Earth as well as debris dust from collisions between smaller and larger bodies populating the early solar system. This collisional equilibrium caused a similar size distribution as today, where the influx of extraterrestrial matter to Earth is dominated by two size regimes (Grün et al., 1985; Love & Brownlee, 1993; Bland & Artemieva, 2006; Cremonese et al., 2012; Vogt et al., 2017): (i) a continuous flux of very small, less than mm-sized interplanetary dust particles and micrometeorites, and (ii) very rarely occurring impacts of km-sized asteroidal bodies. Of course, when the Earth grew in its infant stage, total accretion rates were much higher.

Noble GasesClick to view larger

Figure 2. Ne three isotope plot showing solar, planetary, and terrestrial compositions. Solar wind 20Ne/22Ne ratio is slightly higher than the bulk composition of the sun and its outer convection zone (13.36), while solar Ne-B implanted in mineral surfaces is lower (12.5). The explanation for this is that Ne-B implanted into a mineral surface exposed in space is enriched in the lighter isotope 20Ne, which is more shallowly implanted due to its lower kinetic energy, and preferentially lost by grain surface sputtering. Planetary neon (Ne-A) found in meteorites is also depleted in the light isotopes, and mass fractionated along the mass fractionation line. MORB and OIB data (references in text) are shown as density distributions. Both Earth’s upper MORB mantle and OIB sources fed by deep mantle plumes (e.g., Loihi, Hawaii) are Ne-B like in their 20Ne/22Ne ratios, with various additions by nuclear reactions reflected by increases in the 21Ne/22Ne ratio. Furthermore, MORB and OIB samples are locally contaminated by ubiquitous atmospheric neon, which is not indigenous to the mantle sources, and causing mixing lines to terrestrial atmospheric composition (Air). While solar neon in Earth’s mantle reflects the first accretion stage, when solar gases were incorporated into the solid earth, the atmosphere is likely a mixture of two components, which is solar neon from mantle degassing and later accretion of planetary neon. Planetary Ne-A is a complex mixture of neon from phase Q, P3 and HL (nanodiamonds) and presolar grains (SiC and graphite)—see Table 1 for isotopic compositions.

Accreting irradiated material may have been directly incorporated into Earth’s interior, in particular micrometer-sized dust, which is decelerated more gently in a proto-atmosphere without severe atmospheric entry heating (Flynn, 1989; Love & Brownlee, 1991; Farley et al., 1997; Engrand et al., 2005; Füri et al., 2013), and therefore more likely to have retained solar wind implanted noble gases in grain boundaries. This dust component will have resembled the solar gas–rich interplanetary dust particles (Nier & Schlutter, 1990; Pepin et al., 2000; Schwarz et al., 2005) and micrometeorites (Engrand & Maurette, 1998; Maurette et al., 2000; Osawa et al., 2000, 2003a, 2003b, 2010; Osawa & Nagao, 2002; Baecker, 2014) that are found in modern marine sediments, collected in the stratosphere and in Antarctic ice (Baecker, 2014). During this stage of terrestrial accretion, the influx of helium and neon is dominated by solar wind gases, because planetary precursors forming in the terrestrial planet region carried very minor amounts of intrinsic volatiles or planetary noble gases, unlike outer solar system bodies.

Nevertheless, the accretionary flux may have contained up to a few percent of water-bearing planetesimals from beyond the snow line. In that case, impact degassing of water from incoming planetesimals would form a steam atmosphere, causing an efficient greenhouse effect. Loss of accretionary impact heat would be prevented, and temperatures would rise until a magma ocean developed (Matsui & Abe, 1986; Zahnle et al., 1988; Abe, 1993; Elkins-Tanton, 2008, 2012; Schaefer & Fegley, 2010; Tucker & Mukhopadhyay, 2014; Fegley et al., 2016; de Vries et al., 2016). Hot temperatures would have released solar wind gases from the dust particles, and these would have become part of the proto-atmosphere. In this environment, solar gases could enter the magma by solubility-controlled dissolution (Jambon et al., 1986; Hayashi et al., 1979; Mizuno et al., 1980; Sasaki & Nakazawa, 1990; Harper & Jacobsen, 1996; Porcelli et al., 2001; Pepin & Porcelli, 2002; Paonita, 2005). Equilibrium between atmosphere and the magma ocean would be maintained until the accretionary flux and impact heating ceased. In the case of the Earth, the main phase of accretion likely terminated with a giant impact that formed our Moon.

Loss of the Primary Atmosphere During the Giant Impact Forming the Moon

The continuous growth of planetary bodies in the terrestrial planet–forming region of the nebula will have been characterized by increasingly large impacting bodies during the final stages of planetary accretion. The final stage of terrestrial accretion was marked by the giant impact of a proto-planet of the size of Mars, called Theia, or alternatively by multiple smaller impacts (Rufu et al., 2017). While most of the metallic core of the impacting proto-planet merged with the terrestrial core, large amounts of silicate material were launched into Earth’s orbit, where it condensed to form the Moon. This giant impact scenario can explain some dynamic peculiarities of the Earth-Moon system, the partial loss of (non-gaseous) volatile elements in the Earth and more so in the Moon, as well as the Moon’s small metallic core (Hartmann, 1986; Benz et al., 1987; Benz & Cameron, 1990; Canup & Righter, 2000; Canup & Asphaug, 2001; Canup, 2004).

The impact of Theia caused significant—possibly total—loss of the primary atmosphere, which was probably comprised of solar-wind helium and neon (Pepin, 1997; Canup & Righter, 2000; Genda & Abe, 2003, 2005; Zahnle et al., 2007). However, the remnants of this atmosphere were retained within Earth’s magma ocean and the subsequently solidifying mantle. Upon solidification, the convecting mantle degassed a major portion of its volatiles to give rise to a secondary atmosphere in the aftermath of the Theia impact.

Model calculations based on typical concentrations of solar neon in interplanetary dust particles and micrometeorites demonstrate that a few percent of irradiated matter are sufficient to account for the solar noble gas inventory of the Earth’s mantle and atmosphere, assuming broad constraints for both the possible depth of a magma ocean and various degrees of loss of the primary atmosphere (up to 100%) during the Theia impact (Vogt et al., 2017).

Mantle Degassing and Late Accretion of a Planetary Component

The secondary atmosphere formed by mantle degassing after the impact of Theia was compositionally and isotopically similar to the present mantle inventory of noble gases, a hybrid of solar-type helium and neon, and planetary-type heavy noble gases, with minor contributions of radiogenic noble gases accumulating over geological time.

Furthermore, the secondary atmosphere was supplemented by late-accreting volatile-rich objects containing planetary-type neon. As most of the planetesimals in the inner solar system were consumed by planetary growth or ejected from the solar system by close encounters, late-accreting planetesimals very likely originated from the volatile-rich regions of the asteroid belt or the Kuiper belt, where most small solar system bodies are still located. The isotopic signatures of reactive gaseous elements, including nitrogen and hydrogen, imply that most terrestrial volatiles in the atmosphere and the hydrosphere were delivered by carbonaceous chondrite-type planetesimals (Marty, 2012; Albarede et al., 2013; Halliday, 2013), and possibly a small contribution from comets, explaining the peculiar terrestrial xenon signature of the atmosphere (Owen, 2008; Marty et al., 2017). The isotopic signature of krypton is compatible with carbonaceous chondrites (Holland et al., 2009). The delivery of planetary volatiles also lowered the terrestrial atmospheric 20Ne/22Ne ratio from ~12.5 down to the present value of 9.8.

The term “late accretion” means the last few percent of the terrestrial mass that accreted in the interval between Moon formation and close to the end of the Hadean 4.0 Ga ago. Formation of the lunar ring basins appears to be restricted to a relatively short interval 3.8–3.9 Ga, usually referred to as the late or lunar heavy bombardment (LHB), as recorded by impact ages of lunar rocks returned by the Apollo missions (Tera et al., 1974; Wetherill, 1975; Turner, 1977; Chou, 1978; Ryder, 1990; Hartmann et al., 2000; Koeberl, 2004, 2006; Albarede, 2009; Albarede et al., 2013; Morbidelli & Wood, 2015). There is some evidence, however, that earlier, between 4.4 and 3.9 Ga ago, significant impact activity was ongoing on the moon (e.g., Nemchin et al., 2008), which would imply volatile delivery to Earth.

The Noble Gas State of Earth’s Mantle: Convecting Shallow Mantle Versus Deep Mantle Plume Reservoirs

Radiogenic noble gases produced by radioactive decay are present in the mantle (Graham, 2002); these include (Table 1): 21Ne from nuclear reactions during actinide decay, 40Ar from 40K, 129Xe from 129I, and Xe isotopes from spontaneous fission of 238U or 244Pu (Staudacher et al., 1982; Allègre et al., 1986; Sarda et al., 1985, 1988; Marty, 1989; Burnard et al., 1997; Moreira et al., 1998; Kunz et al., 1998; Caffee et al., 1999; Trieloff et al., 2000, 2002, 2003; Trieloff & Kunz, 2005). Ratios of all radiogenic/nucleogenic versus primordial nuclides in the mantle (40Ar/36Ar, 129Xe/130Xe, 21Ne/22Ne) are higher than in the atmosphere (Table 1). This is a direct consequence of mantle degassing, which allows accumulation of radiogenic or nucleogenic nuclides that stem from decay of lithophile elements in the solid Earth (Ozima & Zahnle, 1993).

The 129Xe excess resulting from the short-lived nuclide 129I (16 Ma half-life) indicates that most mantle degassing occurred early mostly within about the first 100 Ma of Earth’s history (Staudacher et al., 1982; Allègre et al., 1986). The most likely process was outgassing after the Theia impact or more generally in the aftermath of the terminal accretionary bombardment, after which the mantle solidified and previously dissolved gases were expelled (Elkins-Tanton, 2008; Tucker & Mukhopadhyay, 2014). Although degassing is ongoing today, mostly at mid-ocean ridges, degassed amounts—even calculated back over 4 Ga—are minor when compared to the early Earth (Turner, 1989).

The higher proportion of radiogenic isotopes is most prominent in the mantle reservoir sampled by mid-ocean ridge volcanism, the MORB mantle (Staudacher et al., 1982; Allègre et al., 1986; Sarda et al., 1985, 1988; Marty, 1989; Burnard et al., 1997; Moreira et al., 1998; Kunz et al., 1998; Trieloff et al., 2003; Trieloff & Kunz, 2005). MORB form by partial melting at 60–80 km depth in regions of the convecting upper mantle which are depleted in incompatible elements now present in the Earth’s crust (Hofmann, 1988; Taylor & McLennan, 1995). Excesses of radiogenic isotopes are systematically less pronounced for oceanic island basalt (OIB) mantle reservoirs, that is, deep mantle plumes sampled by the hot spot volcanism generating oceanic islands like Hawaii (Kurz et al., 1982, 1983; Honda et al., 1991; Hiyagon et al., 1992; Valbracht et al., 1997; Trieloff et al., 2000, 2002, 2003) or Iceland (Hilton et al., 1999; Harrison et al., 1999; Trieloff et al., 2000, 2002; Dixon et al., 2000; Moreira et al., 2001). For example, 4He/3He ratios are lower in OIB than in MORB, as shown in Table 1. This is commonly interpreted as reflecting the less degassed character of the OIB source.

The fraction of the mantle representing the more strongly degassed MORB source is considered to be ca. 70–90% of the mantle volume, clearly more voluminous than the less degassed OIB source (Kimura et al., 2017). The important point is that there seems to be a less degassed mantle reservoir in the Earth that escaped convective mixing with the rest of the mantle. Some models explain the particular trace element and isotope signatures of OIB reservoirs by adding old, deeply subducted oceanic lithosphere recycled through the mantle (Hofmann & White, 1982; Zindler & Hart, 1986; Hofmann, 1997), but this is difficult to reconcile with a less degassed and hence primordial signature. Alternatively, some models explain the less degassed OIB signature as being derived from convectively isolated mantle domains (Becker et al., 1999; Kellogg et al., 1999) or even the core (Porcelli & Halliday, 2001; Trieloff & Kunz, 2005), as liquid metal during core formation has a certain affinity to incorporate noble gases (Sudo et al., 1994). On the other hand, a significant fraction—particularly of the heavy primordial noble gases—is considered to be due to subduction of atmospheric noble gases (Holland & Ballentine, 2006; Trieloff & Kunz, 2005; Parai & Mukhopadhyay, 2015).

Noble Gases in the Continental Crust and the Subcontinental Lithospheric Mantle

The continental crust is a geochemical reservoir that is the product of ongoing crust-mantle differentiation (Hofmann, 1988; Taylor & McLennan, 1995). Parts of the Earth’s continental crust date back more than 4 Ga, while oceanic crust is rarely older than 200 Ma. The continental crust hosts incompatible elements, and among the most incompatible are those with radiogenic parent nuclides of noble gas isotopes: K, U, and Th. Hence, crustal reservoirs are characterized by highly radiogenic noble gases, enriched in 21Ne, 40Ar, and Kr and U isotopes from U fission. In contrast to the noble gas signature of the convecting MORB mantle, the continental crust is much more heterogeneous (e.g., from the age of local crustal compartments), and so are the noble gas signatures of crustal rocks, whether dominated by volcanics or sediments (Ozima & Podosek, 2002; Ballentine & Burnard, 2002; Ballentine et al., 2002; Kennedy et al., 1990; Drescher et al., 1998; Pinti et al., 2001; Pujol et al., 2009; Hopp et al., 2016). Besides radiogenic additions, atmospheric contributions are significant, because atmospheric noble gases are dissolved in water, so fluid rock interaction can introduce fractionated air.

The subcontinental lithospheric mantle seems to represent a somewhat decoupled reservoir from the convecting mantle, partly old and metasomatized. It has a noble gas isotopic signature which is slightly more radiogenic than the convecting MORB mantle, but nevertheless quite similar. For example, its helium and neon isotopic signature is only slightly more radiogenic (Dunai & Baur, 1995; Hopp et al., 2004; Buikin et al., 2005; Gautheron et al., 2005; Hopp et al., 2007a, 2007b; Hopp & Ionov, 2011), and also influenced by local plume-type contributions, for example carbonatites of the Kola peninsula (Yokochi & Marty, 2004, 2005), Central European volcanism (Buikin et al., 2005), or the Red Sea rift shoulders (Hopp et al., 2004). Diamonds are special windows to the deep subcontinental lithospheric mantle and reflect the full range between OIB, MORB, SCLM, and crustal noble gas isotopic signatures (Honda et al., 1987, 2004, 2011; Ozima & Zashu, 1987, 1991).

Important basic isotopic signatures have been inferred from the subcontinental lithospheric mantle and generalized to the whole mantle; for example, a small meteoritic contribution within a subduction related air-like component was found for Kr and Xe isotopes in CO2-dominated well and spring gases (Holland & Ballentine, 2006; Caracausi et al., 2016). In this context it is also important to discuss pros and cons of the possibility of recycling noble gases back into the mantle (e.g., Staudacher & Allègre, 1988; Trieloff et al., 2000; Trieloff & Kunz, 2005; Holland & Ballentine, 2006).

Open Questions and Future Directions

Although the noble gas isotopic state of Earth’s mantle is increasingly well understood (particularly for helium, which is hardly susceptible to atmospheric contamination), important findings remain to be confirmed. Most of the solar helium and neon in Earth’s mantle seem to originate by solar wind implantation, but it remains to be shown what fraction of the Earth accreted within the solar nebula (Hayashi et al., 1979; Jaupart et al., 2017) and if massive loss of the proto-atmosphere caused elemental or isotopic fractionation (Hunten et al., 1987; Pepin, 1991). For xenon with isotopes being generated by short-lived (129I, 244Pu) and long-lived (238U) precursors, it is important to deconvolve radiogenic contributions by iodogenic, uranogenic, and plutonogenic xenon and relate this to plausible degassing histories of mantle reservoirs (Kunz et al., 1998; Mukhopadhyay, 2012; Caracausi et al., 2016), particularly for oceanic island mantle plumes, for which radiogenic xenon excess has been demonstrated in only a few cases (Poreda & Farley, 1992; Trieloff et al., 2000, 2002; Hopp & Trieloff, 2005; Mukhopadhyay, 2012). The reservoir feeding noble gases in mantle plumes is still mysterious, for example if it is sourced by deep unmixed mantle reservoirs or the core. For krypton, the only anomalies measured were on well gas of subcontinental mantle origin (Holland et al., 2009) which needs to be verified for other mantle samples.

Furthermore, the relationship between atmospheric, mantle, and solar or meteoritic (or cometary) xenon—both primordial and radiogenic—is still unsolved, for example the nature of the mass-independent fractionation process generating atmospheric xenon (e.g., Pujol et al., 2009). Very recently, Marty et al. (2017) obtained the first xenon measurements on a comet (67P/Churyumov-Gerasimenko, visited by the Rosetta mission) and found a unique isotopic composition highly enriched in light xenon isotopes, unlike solar, chondritic, or terrestrial xenon, which could serve as a missing link and explain minor contributions to the terrestrial atmosphere. Clearly more measurements—both in situ and sample return—are needed for cometary samples, but also (more precise) measurements for the atmospheres of terrestrial planets, in order to elucidate the genetic relationship between the Earth and extraterrestrial bodies.

The nature and carriers of extraterrestrial components, the ill-defined phase Q and its origin, certainly deserve further interest and intense studies. Phase Q and the P3 component associated with nanodiamonds contain rather non-exotic noble gases and are found within meteoritic acid residues that also contain interstellar SiC and graphite-carrying nucleosynthetic anomalies. Interstellar grains are likely a twofold population: a certain fraction consists of grains that condensed in specific stellar atmospheres with characteristic isotopic anomalies of the nucleosynthetic stage of the parent star, while another fraction results from grain growth in the interstellar medium. Such grains have been hypothesized from theoretical considerations (Zhukovska et al., 2008) and have recently been detected by the dust detector on board the Cassini spacecraft (Altobelli et al., 2016). These grains would contain a mixture of all nucleosynthetic components dispersed in the interstellar medium, and therefore represent an average isotopic abundance similar to the bulk solar system. Nanodiamonds for which an interstellar origin was suggested, for example due to the presence of nanophase carbon glass (Stroud et al., 2011), could be such candidates. However, for phase Q too such an origin could be envisaged—amorphization by ion radiation and simultaneous implantation of noble gas ions could happen at variable temperatures and densities when grains are cycling between the hot, diffuse warm and cold interstellar medium.

Further Reading

Ballentine, C. J., Burgess, R., & Marty, B. (2002). Tracing fluid origin, transport and interaction in the crust. Reviews in Mineralogy and Geochemistry, 47, 539–614.Find this resource:

Ballentine, C. J., & Burnard, P. G. (2002). Production of noble gases in the continental crust. Reviews in Mineralogy and Geochemistry, 47, 481–538.Find this resource:

Farley, K. A. (2002). (U-Th)/He dating: Techniques, calibrations, and applications. Reviews in Mineralogy and Geochemistry, 47, 819–844.Find this resource:

Graham, D. (2002). Noble gas isotope geochemistry of mid-ocean ridge and ocean island basalts: Characterization of mantle source reservoirs. Reviews in Mineralogy and Geochemistry, 47, 247–317.Find this resource:

Hofmann, A. W. (1997). Mantle geochemistry: The message from oceanic volcanism. Nature, 385, 219–229.Find this resource:

Kelley, S. (2002). K-Ar and Ar-Ar dating. Reviews in Mineralogy and Geochemistry, 47, 785–818.Find this resource:

Kipfer, R., Aeschbach-Hertig, W., Peeters, F., & Stute, M. (2002). Noble gases in lakes and groundwaters. Reviews in Mineralogy and Geochemistry, 47, 615–700.Find this resource:

Lodders, K., Palme, H., & Gail, H.-P. (2009). Abundances of the elements in the solar system. In Springer materials: The Landolt Börnstein Database; New Series VI/4B (pp. 1–59). Berlin, Germany: Springer.Find this resource:

McDougall, I., & Harrison, T. M. (1988). Geochronology and thermochronology by the 40Ar/39Ar method. New York, NY: Oxford University Press.Find this resource:

Morbidelli, A., & Wood, B. J. (2015). Late accretion and the late veneer. In J. Badro & M. J. Walter (Eds.), The early Earth (pp. 71–82). New York, NY: John Wiley and Sons.Find this resource:

Morbidelli, A., Lunine, J. I., O’Brien, D. P., Raymond, S. N., & Walsh, K. J. (2012). Building terrestrial planets. Annual Review of Earth and Planetary Sciences, 40, 251–275.Find this resource:

Moreira, M. (2013). Noble gas constraints on the origin and evolution of Earth’s volatiles. Geochemical Perspectives, 2, 229–403.Find this resource:

Niedermann, S. (2002). Cosmic-ray produced noble gases in terrestrial rocks: Dating tools for surface processes. Reviews in Mineralogy and Geochemistry, 47, 731–784.Find this resource:

Ott, U. (2014). Planetary and pre-solar noble gases in meteorites. Chemie der Erde—Geochemistry, 74, 519–544.Find this resource:

Owen T. (2008). The contributions of comets to planets, atmospheres, and life: Insights from Cassini-Huygens, Galileo, Giotto, and inner planet missions. Space Science Reviews, 138, 301–316.Find this resource:

Ozima, M., & Podosek, F. A. (2002). Noble gas geochemistry (2nd ed.). Cambridge, UK: Cambridge University Press.Find this resource:

Pepin, R. O. (2006). Atmospheres on the terrestrial planets: Clues to origin and evolution. Earth and Planetary Science Letters, 252, 1–14.Find this resource:

Pepin, R. O., & Porcelli, D. (2002). Origin of noble gases in the terrestrial planets. Reviews in Mineralogy and Geochemistry, 47, 191–246.Find this resource:

Porcelli, D., & Ballentine, C. J. (2002). Models for distribution of terrestrial noble gases and evolution of the atmosphere. Reviews in Mineralogy and Geochemistry, 47, 411–480.Find this resource:

Porcelli, D., Ballentine, C. J., & Wieler, R. (2002). An overview of noble gas geochemisry and cosmochemistry. Reviews in Mineralogy and Geochemistry, 47, 1–19.Find this resource:

Porcelli, D., & Pepin, R. O. (2011). The origin of noble gases and major volatiles in the terrestrial planets. In H. D. Holland & K. K. Turekian (Eds.), Isotope geochemistry—From the treatise on geochemistry (pp. 493–520). London, UK: Elsevier Academic Press.Find this resource:

Schaefer, L., & Fegley, B., Jr. (2010). Chemistry of atmospheres formed during accretion of the Earth and other terrestrial planets. Icarus, 208, 438–448.Find this resource:

Schlosser, P., & Winckler, G. (2002). Noble gases in ocean waters and sediments. Reviews in Mineralogy and Geochemistry, 47, 701–730.Find this resource:

Swindle, T. D. (2002a). Noble gases in the Moon and meteorites: Radiogenic components and early volatile chronologies. Reviews in Mineralogy and Geochemistry, 47, 101–124.Find this resource:

Swindle, T. D. (2002b). Martian noble gases. Reviews in Mineralogy and Geochemistry, 47, 171–190.Find this resource:

Trieloff, M., & Kunz, J. (2005). Isotope systematics of noble gases in the Earth’s mantle: Possible sources of primordial isotopes and implications for mantle structure. Physics of the Earth and Planetary Interiors, 148, 13–38.Find this resource:

Trieloff, M., & Palme, H. (2006). The origin of solids in the early solar system. In H. Klahr & W. Brandner (Eds.), Planet formation: Theory, observations, and experiments (pp. 64–89). Cambridge, UK: Cambridge University Press.Find this resource:

Trieloff, M. (2009). Chronology of the solar system. In J. Trümper (Ed.), Landolt-Börnstein: Numerical data and functional relationships; New series Vol. VI/4 Astronomy, astrophysics, cosmology (pp. 599–612). Berlin, Germany: Springer.Find this resource:

Wieler, R. (2002). Noble gases in the Solar System. Reviews in Mineralogy and Geochemistry, 47, 21–70.Find this resource:

Zahnle, K., Schaefer, L., & Fegley, B. (2010). Earth’s earliest atmospheres. Cold Spring Harbor Perspectives in Biology, 2, 1–17.Find this resource:

Zhang, Y. (2014). Degassing history of Earth. In K. K. Turekian (Ed.), Treatise on geochemistry (2d ed., Vol. 6, pp. 37–69). Oxford, UK: Elsevier.Find this resource:

Zindler, A., & Hart, S. (1986). Chemical geodynamics. Annual Review of Earth and Planetary Sciences, 14, 493–571.Find this resource:

Zinner, E. (1998). Stellar nucleosynthesis and the isotopic composition of presolar grains from primitive meteorites. Annual Review of Earth and Planetary Sciences, 26, 147–188.Find this resource:

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