ARTICLE
Received 29 Sep 2016 | Accepted 21 Apr 2017 | Published 1 Jun 2017
Tracing the oxygen isotope composition of the
upper Earth’s atmosphere using cosmic spherules
Andreas Pack1, Andres Ho¨weling1,2, Dominik C. Hezel3, Maren T. Stefanak1, Anne-Katrin Beck1,
Stefan T.M. Peters1, Sukanya Sengupta1, Daniel Herwartz3 & Luigi Folco4
Molten I-type cosmic spherules formed by heating, oxidation and melting of extraterrestrial
Fe,Ni metal alloys. The entire oxygen in these spherules sources from the atmosphere.
Therefore, I-type cosmic spherules are suitable tracers for the isotopic composition of
the upper atmosphere at altitudes between 80 and 115 km. Here we present data on
I-type cosmic spherules collected in Antarctica. Their composition is compared with
the composition of tropospheric O2. Our data suggest that the Earth’s atmospheric O2 is
isotopically homogenous up to the thermosphere. This makes fossil I-type micrometeorites
ideal proxies for ancient atmospheric CO2 levels.
DOI: 10.1038/ncomms15702 OPEN
1 Universita¨t Go¨ttingen, Geowissenschaftliches Zentrum, Goldschmidtstrae 1, 37077 Go¨ttingen, Germany. 2 Karlsruher Institut fu¨r Technologie, Institut fu¨r
Angewandte Materialien - Werkstoffprozesstechnik, Hermann-von-Helmholtz-Platz 1, 76344 Eggenstein-Leopoldshafen, Germany. 3 Universita¨t Ko¨ln, Institut
fu¨r Geologie und Mineralogie, Greinstrae 4-6, 50939 Ko¨ln, Germany. 4 Universita´ di Pisa, Dipartimento di Scienze della Terra, Via Santa Maria 53, 56126
Pisa, Italy. Correspondence and requests for materials should be addressed to A.P. (email: apack@uni-goettingen.de).
NATURE COMMUNICATIONS | 8:15702 | DOI: 10.1038/ncomms15702 |www.nature.com/naturecommunications 1
F
ree molecular oxygen (O2) is released by photosynthesis into
the atmosphere and is essential for all breathing animals.
With exception of data for the last 800,000 years from air
inclusions in polar ice, little direct information is available about
concentration and isotope composition of ancient atmospheric
O2. This is due to the limited interaction between air molecular
oxygen and the lithosphere.
Among the rare rocky materials that contain atmospheric
oxygen1–3 there are particular types of micrometeorites
(microscopic extraterrestrial dust particles) called cosmic
spherules4. Roughly 10 tons of small extraterrestrial particles
are deposited onto the Earth’s surface per day5. The particles
collide with the Earth’s atmosphere at velocities of 11–70 km s 1
(ref. 6) and are visible as shooting stars when they are decelerated
and at altitudes up to B80–115 km7,8. A portion of these
extraterrestrial particles totally melts during the atmospheric
entry and is termed cosmic spherules. Cosmic spherules that are
composed of Fe,Ni oxides are termed ‘I-type cosmic spherules’
(in the following, we use the short version ‘I-type spherules’9–11.
These I-type spherules formed by oxidation of extraterrestrial
Fe,Ni metal alloys, which are ubiquitous components of
meteorites.
Because oxygen in I-type spherules originates entirely from
the atmosphere, they are excellent probes for the isotopic
composition of upper atmospheric oxygen. The isotopic compo-
sition of atmospheric oxygen, in turn, is a proxy for the global
primary production (GPP) and atmospheric CO2 levels1,2,12–14. It
is not clear, however, if the atmospheric oxygen is isotopically
homogenous up to the meso- and thermosphere, where cosmic
spherules interact with air.
The stable isotope composition of tropospheric O2 (99.8% 16O,
0.04% 17O, 0.2% 18O) is controlled by the steady state between
photosynthesis and respiration (mass-dependent Dole effect15),
evapotranspiration and mass-independent fractionation in the
stratosphere14,16. For the composition of the modern troposphere
values of 23.4rd18Or24.2% and  0.566rD’17Or 0.430%
have been reported in the literature14,17–20 (for definitions, see
Methods) with little variations up to 61 km (ref. 17). The high
d18O of tropospheric O2 is caused by the Dole effect, whereas the
low D’17O value reflects mass-independent fractionation effects in
the stratosphere. The higher the atmospheric CO2 levels, the
lower the D’17O values12,14; a relation that was used as paleo-CO2
barometer1,2,13.
No experimental oxygen isotope data are available for the
upper atmosphere at altitudes 461 km. To obtain information
about the D’17O heterogeneity of the atmosphere, we measured
the oxygen (d17O, d18O) and iron (d56Fe, d57Fe; see Methods for
definition) isotope composition of I-type spherules from the
Transantarctic Mountains that have ages o2Ma (ref. 21). For
this time interval, atmospheric CO2 levels as well as D’17O of O2
did not deviate much from the modern level22,23 and data can be
used to test whether the atmosphere is isotopically homogenous.
Because of the B1,000 years residence time of atmospheric O2
(ref. 14), no effect on the man-made increase in CO2 is yet visible
in decreasing D’17O.
The oxygen isotope composition of I-type spherules is
controlled by the composition of the oxidizing species
(for example, atmospheric O2), the fractionation during oxidation
of the Fe,Ni alloys, and the fractionation during atmospheric
evaporation.
Cosmic spherules have higher d18O values (up to 56%; ref. 24)
than any terrestrial material reported so far. The high d18O led
Clayton et al.4 to propose a heavy oxygen isotope reservoir in the
upper atmosphere. Davis et al.25, however, showed that I-type
spherules are also enriched in heavy iron isotopes with extreme
d56Fe values of up to 45%. They concluded that high d18O values
are due to evaporation and do not reflect the isotope composition
of the upper atmosphere; a conclusion that was supported by
further measurements24,26–29. From the oxygen and iron isotope
composition, Engrand et al.24 modeled evaporative mass losses
for I-type spherules of 54–85%.
Because evaporative fractionation is strictly mass-dependent,
I-type spherules still provide unique information about the
mass-independent anomaly in D’17O of their upper mesospheric
oxygen source. The reconstruction of variations in atmospheric
D’17O from fossil cosmic spherules4,30–33 would be a new
paleo-CO2 proxy. The only published D’17O data on I-type
spherules by Clayton et al.4 and Engrand et al.24, however, have
intrinsic uncertainties that are too large (0.1 to41%) to provide
reliable information on the composition of the upper atmosphere.
We present new high-precision oxygen isotope data of
tropospheric O2 and compare these data with new high-precision
oxygen and iron isotope data from Antarctic I-type spherules.
These data are combined with results of oxidation and
evaporation experiments to test if the Earth atmosphere is
isotopically homogenous and if isotope ratios of fossil cosmic
spherules are suitable paleo-CO2 proxies.
Results
Oxygen isotope composition of tropospheric air. The mean
composition of air oxygen from our study (series B;
Supplementary Table 1) is d18O¼ 24.15±0.05% with
D’17O¼  0.469±0.007%. An earlier protocol (series A;
Supplementary Table 1) gave identical D’17O, but slightly lower
d18O. The datum from this study is within the range reported in
the literature and agrees with D’17O values of Thiemens et al.17
and Young et al.14. The measured data of Young et al.14 have
been corrected relative to the San Carlos olivine value reported by
Pack et al.34. The corrected measured datum of
D’17O¼  0.467±0.005% for air oxygen14 is then identical to
our measured datum of  0.469±0.007% (Supplementary
Table 1) and agrees with the model datum of  0.469%
presented by Young et al.14 The non-application of
VSMOW2-SLAP2 scaling34 to our air data shifts the D’17O
down to B 0.50%, which would be closer to the values of
Barkan and Luz18 and Kaiser and Abe20. For this study, we adopt
d18O¼ 24.15% and D’17O¼  0.47% for the troposphere.
Oxygen and iron isotope composition of cosmic spherules. The
isotope composition of the I-type spherules and the oxidation
experiment run products are listed in Supplementary Table 2. The
d18O of the spherules ranges from 36 to 42%. The corresponding
D’17O ranges from  0.72 to  0.62%. The d18O and the D’17O
is within the range reported by Clayton et al.4 and
Engrand et al.24 The d56Fe values are high for all spherules,
ranging from 22 to 32% (Supplementary Table 2).
Discussion
The interaction between cosmic Fe,Ni metal and the Earth
atmosphere during deceleration is considered to proceed in two
consecutive steps. The first step is the atmospheric heating and
oxidation of the infalling Fe,Ni metal alloy (fractionation in
oxygen isotopes only). The second step is the melting and
evaporation of the Fe,Ni oxides (fractionation in both, oxygen
and iron isotopes).
Information about the oxygen isotope fractionation that is
associated with the oxidation step is obtained from experiments
(this study; see Methods) and from iron meteorite fusion crust
data4,35,36. The products of the high-T metal oxidation
experiments have d18O values that are 4% lower than air
oxygen (Supplementary Table 2). Clayton et al.4 reported d18O
ARTICLE NATURE COMMUNICATIONS | DOI: 10.1038/ncomms15702
2 NATURE COMMUNICATIONS | 8:15702 | DOI: 10.1038/ncomms15702 | www.nature.com/naturecommunications
values of iron meteorite fusion crusts that wereB9% lower than
air oxygen. They concluded that the atmospheric oxidation of
Fe,Ni alloys is associated with kinetic fractionation and formation
of isotopically light oxides with d18OFe,Ni oxideod18Oair, which we
also suggest as the cause of the low d18O of the experimental
Fe,Ni oxides. Kinetic fractionation should be associated with a
low yO value (for definition, see Methods). The best-fit line
through air, experimental run products and iron meteorite fusion
crust4 has a slope of yO¼ 0.506±0.003 (Fig. 1). This low y value
clearly supports that kinetic fractionation is the cause for the
difference in d18OFe,Ni oxides and d18Oair O2 (ref. 37). Molecular
diffusion of O2 would give a slope of yO¼ 0.508. Iron meteorite
fusion crust data35 suggest that iron isotopes do not fractionate
during oxidation.
For the experiments and iron meteorite fusion crusts,
atmospheric oxygen is the oxidant. Above the ozone layer,
however, a considerable fraction of molecular oxygen is steadily
dissociated into atomic oxygen (for example,38). Atomic oxygen
is a hazard for low Earth orbit space flights due to its highly
corrosive nature. Because I-type spherules are oxidized at high
altitudes, atomic oxygen may have contributed to the oxidation.
However, Clayton et al.4 and Genge10 stated that no
discrimination between atomic and molecular oxygen is likely
during oxidation upon atmospheric entry because the collision
energy between infalling meteoroids and air particles is
higher than the O2 bond strength. This implies that I-type
spherules sample the bulk upper atmosphere oxygen (O and O2).
Our experiments and the fusion crust literature data4 show that
aoxidationo1 (for 18O/16O; see Equation 5), but also reveal
considerable variation. For I-type spherules, we assume that
0.9428raoxidationr1. The lower limit is given by pure Graham’s
law39 fractionation with atomic oxygen being the moving species.
The second process affecting the isotopic composition of I-type
spherules is evaporation25. Because iron isotopes are not affected
by oxidation35 but only by evaporation, d56Fe can be used as
monitor for the degree of evaporation f (refs 24,25,40;
Equation 1). The d56Fe of the infalling metal is assumed to be
0±1% relative to the IRMM-014 standard material41.
d56Fe ¼ 103ð1 f Þa
56=54
evaporation  1  1 ð1Þ
Wang et al.40 determined an aevaporation¼ 0.9820 (56Fe/54Fe). This
value in combination with d56Fe allows calculation of f for the
I-type spherules (Equation 1). The d56Fe values of the spherules
from this study (Supplementary Table 2) indicate evaporative
mass loss of 70rfr82%, which is well in the range of
54rfr85% reported by Engrand et al.24.
The experimental relation between d18O and d56Fe during
evaporation from Wang et al.40 (Fig. 2) now allows
reconstructing the d18O of the I-type spherules before
0.0
–0.1
–0.2
Iron meteorite
fusion crusts4
Oxidation
experiments
(this study)
Air O2
–0.3
–0.4
–0.5
–0.6
10 15 20 25
 = 0.506 ± 0.003
 = 0.5305
Δ'
17
 
O
 (‰
)
103
δ18OVSMOW2103 × In( + 1) (‰)
Figure 1 | Triple oxygen isotope fractionation during metal oxidation. Plot
of D’17O versus d’18O of the run products of the high-T oxidation
experiments (solid red squares) along with air oxygen (this study) and iron
meteorite oxide fusion crust (solid blue circles;4; the lowest values were not
considered as it was likely affected by exchange with low-d18O water).
Shown are values with 1s error bars. The oxidation is associated with a
kinetic isotope fractionation with slope yO¼0.506±0.003 (solid line;
dashed lines indicate the 1s s.d. uncertainty interval). The high-T
equilibrium slope y¼0.5305 (dotted horizontal line) is shown for
comparison.
10
3 
 
In
(
10
3
+
 1
) (
‰
)
δ1
8 O
st
ar
tin
g 
m
at
er
ia
l
sa
m
pl
e
103  In(
103
+ 1) (‰)δ
56Festarting material
sample
50
40
Experiments40
Eva
por
atio
n tr
end
(slop
e = 1
.18 ±
 0.05
)
30
20
10
0
0 10 20 30 40
Figure 2 | Oxygen and iron isotope fractionation during evaporation.
Plot of the linearized d18O versus d56Fe of evaporation experiments
by Wang et al.40. The iron isotope fractionation is a function of fraction
f evaporated (Equation 1). The close coupling between d18O and d56O
allows prediction of the pre-evaporation d18O of the I-type spherules.
The 1s-errors (s.d.) were estimated on base of the scatter of the data. The
regression line (solid) is shown with its respective 1s-error envelope
(dashed lines).
35
0.8
0.7
0.6
0.2
0.4
0
30
32 –1
37–2 7m–01
Forbidden
f e
va
po
ra
te
d
32–2
25
20
15
10
5
0
–10 0 10 20 30 40 50 60
δ5
6 F
e 
(‰
)
δ18O (‰)
δ18O of Fe,Ni oxide
before evaporation

kin
eti
c
ox
ida
tio
n
=
 0.
94
28

kin
eti
c
ox
ida
tio
n
=
 1
Figure 3 | Model for the triple oxygen isotope fractionation during
oxidation. Plot of measured d56Fe versus d18O of spherules for the
reconstruction of the pre-evaporative d18O (32-01: black solid circle; 32-02:
blue solid square; 37-02: red solid triangle; 7m-01: rotated green solid
triangle). The right axis displays the fraction f that evaporated (70–82%).
The slopes of the lines (32-01: black; 32-02: blue; 37-02: red; 7m-01:
green) displayed were taken from the experiments by Wang et al.40) (Fig. 2.
The grey shaded ‘forbidden’ areas outline cases for unreasonable kinetic
fractionation factors a associated with atmospheric oxidation of Fe,Ni
metal; all studied spherules fall outside the forbidden areas. The star marks
the d18O of modern air oxygen.
NATURE COMMUNICATIONS | DOI: 10.1038/ncomms15702 ARTICLE
NATURE COMMUNICATIONS | 8:15702 | DOI: 10.1038/ncomms15702 |www.nature.com/naturecommunications 3
evaporation (f¼ 0; Equation 2).
ln
d18Obefore evaporation
103
þ 1
 
¼ ln d
18Ospherule
103
þ 1
 
 1:18  ln d
56Fespherule
103
þ 1
 
ð2Þ
The difference in d18O between the spherules before evaporation
and air oxygen gives the degree of fractionation during the oxidation
step. We obtained values between  22 and  12% during
oxidation of the infalling Fe,Ni alloys (Fig. 3). The greater magni-
tudes of fractionation compared with our experiments ( 4%) and
iron meteorite fusion crust (B  9%; ref. 4) is attributed to the
much shorter heating time of I-type spherules10 compared with iron
meteorites and experiments. Shorter heating and oxidation time
prevents equilibration between oxides and air oxygen.
For the reconstruction of the D’17O of the upper mesospheric
oxygen, we need the yO values for the oxidation (0.506±0.003;
Fig. 1) and evaporation. Wang et al.40 measured yO¼ 0.5096 for
their evaporation experiments. As in case of oxidation, the low yO
value is indicative of kinetic fractionation during evaporation37. The
observed yFe values during evaporation (yFe¼ 0.671; ref. 40) and for
the I-type spherules from this study (yFe¼ 0.674) are significantly
lower than the high-T equilibrium value of yFe¼ 0.687 (see ref. 37
for details) and support kinetic fractionation.
The results of the calculation are listed in Supplementary
Table 3 and illustrated in Fig. 4. The gas that oxidized the studied
Antarctic I-type spherules had  0.510rD’17Or 0.420%
(mean  0.460±0.020%; this study;14,17 which is in excellent
agreement with measured D’17O¼  0.469% for the modern
troposphere (Fig. 4).
Our data from oxygen and iron isotope analyses of Antarctic
I-type spherules are consistent with an oxygen source with a
D’17O similar to that of modern tropo- and stratospheric
molecular oxygen within±0.02%. No oxygen reservoir with a
markedly different D’17O participated in the oxidation of I-type
spherules, suggesting that the Earth atmosphere is isotopically
homogenous up to the mesosphere in B70–80 km (Fig. 5).
Our results imply that the oxygen isotope composition (D’17O)
of the bulk atmosphere can be reconstructed from combined
oxygen and iron isotope data of I-type cosmic spherules.
This has an important implication for the reconstruction of
past atmospheric CO2 levels. Blunier et al.42 showed that the
D’17O of atmospheric molecular oxygen, indeed, varies with CO2
partial pressures. This is predicted from experiments12 and mass
balance modeling14. The oxygen and iron isotope composition
of unaltered fossil I-type cosmic spherules4,30–33 will thus provide
information on the D’17O of the ancient atmosphere and
past CO2 levels. The resolution of the calculated D’17OO2
isB0.07% (single I-type cosmic spherule; see Fig. 4), which
translates (at modern GPP) to an uncertainty in the CO2 mixing
ratio ofB200 p.p.m.14. Lower than modern GPP levels would
lead to an even higher resolution of calculated CO2 levels.
The reconstruction of CO2 levels based on 17O in I-type spherules
is therefore considerably more precise than CO2 reconstruction
from 17O of sulfate1 and reaches far more back into Earth
history than 17O from air inclusions in ice cores42 and fossil
mammal bioapatite2,13. However, to use D’17O of air oxygen as
paleo-CO2-barometer the GPP at that time needs to be known14.
This may limit the usability of the new proxy. The apparent
disadvantage, however, can be turned into a fortune. If the
atmospheric CO2 concentration is known from other,
independent proxies43, D’17O of atmospheric molecular oxygen
103
δ18OVSMOW2103 × In( + 1) (‰)
0.4a b
c d
0.2
32–01
37–02
7m–01
32–02
0.0
–0.2
–0.4
–0.6
–0.8
–10 0 10 20 30 40 50
Δ'
17
 
O
0.
53
05
 
(‰
)
103
δ18OVSMOW2103 × In( + 1) (‰)
0.4
0.2
0.0
–0.2
–0.4
–0.6
–0.8
–10 0 10 20 30 40 50
Δ'
17
 
O
0.
53
05
 
(‰
)
103
δ18OVSMOW2103 × In( + 1) (‰)
0.4
0.2
0.0
–0.2
–0.4
–0.6
–0.8
–10 0 10 20 30 40 50
Δ'
17
 
O
0.
53
05
 
(‰
)
103
δ18OVSMOW2103 × In( + 1) (‰)
0.4
0.2
0.0
–0.2
–0.4
–0.6
–0.8
–10 0 10 20 30 40 50
Δ'
17
 
O
0.
53
05
 
(‰
)
Δ'17 O
air,proxy = –0.43 ± 0.08 ‰
Δ'17 O
air,proxy = –0.52 ± 0.06 ‰
Δ'17 O
air,proxy = –0.46 ± 0.07 ‰
Δ'17 O
air,proxy = –0.42 ± 0.07 ‰
Figure 4 | Reconstruction of the D’17O of upper atmospheric oxygen. Plot of D’17O versus linearized d18O showing the result of the model calculations
(including MonteCarlo error calculation (N¼ 500) grey dots) for I-type spherules 32–01 (a), 32–02 (b), 37–02 (c) and 7m-01 (d). The measured
composition of the I-type spherules is displayed (yellow circle). The oxygen isotope composition is controlled by oxidation (red arrow) and evaporation
(blue arrow). The resultant D’17O (at given d18O¼ 24%; this study) is displayed (yellow square) along with the composition of modern tropospheric O2
(filled upside–down triangle). Given are 1s-error bars.
ARTICLE NATURE COMMUNICATIONS | DOI: 10.1038/ncomms15702
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in combination with mass balance modeling14 turns into a proxy
for the GPP. There is little doubt that I-type cosmic spherules
were deposited during the entire geological history of the Earth.
The question is whether sufficiently large and unaltered fossil
I-type spherules can actually be recovered from sediments. The
recent find of unaltered 2.7 Ga old I-type cosmic spherules33 is
very promising in this respect.
Methods
Sampling and experiments. We studied a total of 21 aliquots of four samples for
oxygen isotopes and one aliquot of the four samples each for iron isotopes. The
samples are part of the Transantarctic Mountain collection44. The sample sizes
range between 400 and 550 mm with masses between 160 and 370mg. The samples
were inspected for weathering products by optical microscopy. No weathering
products (for example, brownish ferrihydride or goethite) were observed.
The spherule densities were determined prior to crushing. Their diameters and
masses were measured and from this their densities calculated (Supplementary
Table 4). The I-type spherule densities vary between 4.3 and 4.8 g cm 3
(Supplementary Table 4).
The density data confirm that the I-type spherules are composed of Fe,Ni oxides
with little or no remaining Fe,Ni metal. The studied samples fall within the density
range (5.0±0.5 g cm 3) observed by Feng et al.45). Wu¨stite ([Fe,Ni]O0.94) has a
density of 5.7 and magnetite ([Fe,Ni]2FeO4) of 5.2 g cm 3. In contrast, iron metal
has a density of 7.9 g cm 3. The apparent lower density of the spherules compared
to wu¨stite and magnetite is explained by B20 vol.% pore space.
The samples are all spherical due to melting during their atmospheric entry.
The bulk elemental composition and the mineralogy of the studied spherules were
not determined. Electron microprobe analysis of I-type spherules from the same
collection yielded 91±5 wt.% FeO, 2.8±0.5 wt.% NiO, and MgO, Al2O3, ando0.5
wt.% SiO2 (ref. 44). This composition is similar to results of Engrand et al.24 and
Herzog et al.28 who report values of 92–93 wt.% FeO and 4–5 wt.% NiO.
For isotope analysis, spherules were wrapped in Al foil and gently crushed in a
steel mortar. We obtained 21 aliquots (4–8 per spherule) with masses of 20–50 mg.
As magnetite is the dominant phase in I-type spherules24,46, we used terrestrial
magnetite along with NBS-28 quartz for tests. For NBS-28, we adopted a
d18O¼ 9.65% and D’17O¼  0.054% (D’17O from47, with revision from34).
I-type spherules form by oxidation of Fe,Ni alloys at high temperatures during
their atmospheric entry. We conducted three metal oxidation experiments at the
University of Go¨ttingen to study the oxygen isotope fractionation associated with
high temperature oxidation of metal in air. A powdered iron base alloy with 7.5
wt.% Ni and 0.6 wt.% Co was used as an analogue material for I-type spherules.
Between 0.8 and 1.1mg powder was placed on a ceramic plate in the hot zone of a
Gero HTRV vertical gas-mixing furnace. The furnace was flushed with
300mlmin 1 air. Oxidation occurred between 1,510 and 1,590 C for 30min.
Air samples were taken at the Go¨ttingen University North Campus outside the
Geoscience Building (513302300 N 95604600 E). The air was taken from the balcony
on the 4th floor using a 5ml syringe yieldingB1ml standard temperature and
pressure O2 gas.
Oxygen isotope analyses. Variations in stable oxygen isotope ratios of a sample
are expressed in form of the d notation relative to the ratios in VSMOW2 water
(Equation 3) with i standing for masses 17 and 18:
diOsample ¼
iO
16Osample
iO
16O VSMOW2
 1 ð3Þ
Deviations from an otherwise close correlation between d17O and d18O are
expressed in form of the D’17O notation (Equation 4). We choose a reference line
with slope 0.5305 and zero intercept. Deviations of D’17O from zero can be caused
by both, non-mass-dependent and mass-dependent processes.
D
017Osample ¼ lnðd17Osample þ 1Þ 0:5305  lnðd18Osample þ 1Þ ð4Þ
The fractionation between two reservoirs (A, B) is expressed in form of the
fractionation factor a (Equation 5). The reservoirs could be two phases in
equilibrium or products (B) and educts (A) of a kinetic process.
ai=16AB ¼
diOA þ 1000
diOB þ 1000 ð5Þ
The i in Equation 5 stands for isotopes with masses 17 and 18. For mass-dependent
processes, the relation between a17/16 and a18/16 is linked through the triple oxygen
isotope fractionation exponent yO (Equation 6).
a
17O=16O
AB ¼ ða
18O=16O
AB ÞyO ð6Þ
For oxygen yO varies between 0.5000 and 0.5305 (refs 37,48,49). Only for very
small a values, y values may fall outside the 0.5–0.5305 range50. Such effects are
neglected here. As a rule, low y values are associated with kinetic effects, whereas
higher y values are associated with equilibrium fractionation processes.
The triple oxygen isotope ratios of I-type spherules and the high-T oxidation
experiment run products were analysed at the University of Go¨ttingen on O2
extracted by infrared laser fluorination51, following the protocol described in
Pack et al.2. In brief, sample O2 was liberated by laser fluorination (F2) and
analysed in continuous flow mode in a Thermo MAT253 gas source mass
spectrometer. NBS-28 quartz was used for normalisation relative to VSMOW2
scale (d17O¼ 5.04%, d18O¼ 9.65%, D’17O¼  0.054%; using the revised
calibration of San Carlos olivine from34). The total masses of the spheres were
160–370mg and thus suitable for measurement of multiple aliquots of a single
spherule. For the oxidation experiments, B0.6–1mg aliquots were analysed.
Oxygen from air was extracted using the same line that was described by
Pack et al.34 for their water (including VSMOW2 and SLAP2) and silicate analyses
(San Carlos olivine) (Fig. 6).
For each extraction, 5ml aliquots of air standard temperature and pressure
(STP) were injected through a liquid nitrogen cooled glass U-trap (for removal of
moisture and CO2; ‘trap 7’, Fig. 6). The dry, CO2 free mixture of Ar, N2 and O2 was
transferred to ‘trap 2’ that was filled with 5Å molecular sieve pellets. In an early
protocol (S01–S05; Supplementary Table 1), Ar was separated from O2 at
 100 C, which resulted in very long trapping times. In an improved protocol
(B01–BP2; Supplementary Table 1), separation of Ar was performed using the cryo
unit of the Hewlett-Packard 5890 gas chromatograph at  80 C. After Ar had
passed through the gas chromatograph (monitored using a Pfeiffer Prisma
80 Thermosphere
This study
Troposphere
Mesosphere
Stratosphere
Stratosphere
Ozone layer
Mass–dependent
fractionationMa
ss
–i
nd
ep
en
de
nt
 c
on
tri
bu
tio
n 
= 
f(p
co
2,
G
PP
)
At
m
os
ph
er
ic
 O
2
(D
ole
 ef
fec
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nd
 e
va
po
tra
n
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ira
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Ph
ot
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(km
)
20
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–0.8 –0.6 –0.4 –0.2 0.0
R
oc
ke
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at
a1
7
Δ'17 O0.5305 (‰)
Figure 5 | Isotope profile through the Earth atmosphere. Diagram
showing the D’17O profile of the modern atmosphere. The data up to
61 km are measurements on air samples (solid circles:17; solid
triangles:14,18–20,54). These data agree well with the mesospheric I-type
spherule proxy data from this study (solid red squares with 1s-error bars).
NATURE COMMUNICATIONS | DOI: 10.1038/ncomms15702 ARTICLE
NATURE COMMUNICATIONS | 8:15702 | DOI: 10.1038/ncomms15702 |www.nature.com/naturecommunications 5
quadrupole mass spectrometer at the end of the He capillary; ‘He out’ in Fig. 6),
temperature was raised to  30 C for the elution of O2 and separation of N2. The
improved protocol is similar to the protocol described by Young et al.14. The
purified O2 was analysed forB60–90min in dual inlet mode.
To test accuracy and precision of laser F2 in combination with continuous flow
mass spectrometry of small samples, a set of experiments were performed on
magnetite (as analogues of I-type spherules) and NBS-28 quartz (Fig. 7).
Our tests on NBS-28 quartz and magnetite (Fig. 7) show that precise analyses of
d18O and D’17O are possible down to 10 mg samples. The blank intensities
corresponded to rB 0.005 mmol O2 (rB6% of the sample). The tests reveal
uncertainties in d18O of±0.6% for quartz and±1% for magnetite. The
uncertainties in D’17O were±0.03% for quartz and ±0.05% for magnetite.
The higher uncertainty observed for magnetite may be due to heterogeneity on
small scale (for eample,52). For our I-type spherule analyses, we adopt analytical
uncertainties of 1% for d18O and 0.06% for D’17O. We have not adopted the
VSMOW2-SLAP2 scaling for the continuous flow measurements since conditions
during the measurement are less controlled (for example, variable peak heights and
widths) and VSMOW2-SLAP2 correction for D’17O(ref. 34) would be smaller than
the uncertainty in D’17O. The agreement between the spherule D’17O from this
study and conventional F2 data from Clayton et al.4 confirm that the continuous
flow d17O data are on VSMOW2 scale.
Iron isotopes. Three iron isotopes have been analysed (54Fe, 56Fe, 57Fe). Isotope
ratios are expressed relative to the IRMM-014 standard in form of the d notation
with:
diFe ¼
iFe
54Fe sample
iFe
54Fe IRMM14
 1 ð7Þ
and i corresponding to masses 56 and 57. All d56Fe and d57Fe data are reported in
% variations. Variations in the three iron isotopes are coupled through the triple
iron isotope exponent yFe (Equation 8).
a56=54AB ¼ ða56=54ABÞyFe ð8Þ
The symbols A and B in Equation 8 can either stand for two phases that are in
equilibrium or for educts (B) and products (A) of a reaction with associated kinetic
fractionation. The high-T approximation for equilibrium iron isotope fractionation
is yFe¼ 0.6784 (ref. 37). As in case of oxygen, variations in yFe provide insights into
the fractionation process.
The iron isotope compositions were measured at the University of Bonn using a
Thermo Scientific Neptune MC-ICP-MS instrument and glassware for sample
introduction. Samples were measured 2–3 times, non-consecutively, during long
analytical sessions of around 8 hours. Analyses were carried out in high-resolution
mode with sufficient transmission to allow routine analyses of an 1 p.p.m. iron
sample solutions. Each sample measurement was bracketed by two analyses of an
IRMM-014 iron solution that was made up to closely match the iron concentration
of the sample. The external reproducibility of the data was typically±0.07%,
whereas the internal reproducibility was generally about a factor of two better. A
detailed description on sample preparation and mass spectrometry can be found in
Hezel et al.53 and Hezel et al.36.
Monte Carlo simulation. The errors in d18O and D’17O of the upper atmospheric
oxygen were estimated using a Monte Carlo approach. The composition was
computed 500 times. Each input parameter was varied independently for each
run (Supplementary Table 5). We used a normal distribution for the variation
within the respective error interval. The computation was performed using the
Mathematica software.
Data availability. The authors declare that the data supporting the findings of this
study are available within the paper and its supplementary information.
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Acknowledgements
The authors thank Reinhold Przybilla and Ingrid Reuber for assistance in the laboratory.
The collection and characterisation of the micrometeorites studied in this work was
supported by the Italian MIUR (grant PNRA16_00029 and PRIN2015_20158W4JZ7).
The project was financially supported by Deutsche Forschungsgemeinschaft (grant
PA909/11-1).
Author contributions
A.P. conceived and supervised the project. A.H., D.C.H., M.T.S., A.-K.B., S.T.M.P., S.S.
and D.H. designed the experiments, prepared the samples and conducted the analyses.
L.F. provided the samples. A.P. wrote the manuscript.
Additional information
Supplementary Information accompanies this paper at http://www.nature.com/
naturecommunications
Competing interests: The authors declare no competing financial interests.
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How to cite this article: Pack, A. et al. Tracing the oxygen isotope composition
of the upper Earth’s atmosphere using cosmic spherules. Nat. Commun. 8, 15702
doi: 10.1038/ncomms15702 (2017).
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