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Sixty years of radiocarbon dioxide measurements at Wellington, New Zealand: 1954–2014

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We present 60 years of Δ¹⁴CO2 measurements from Wellington, New Zealand (41° S, 175° E). The record has been extended and fully revised. New measurements have been used to evaluate the existing record and to replace original measurements where warranted. This is the earliest direct atmospheric Δ¹⁴CO2 record and records the rise of the ¹⁴C bomb spike and the subsequent decline in Δ¹⁴CO2 as bomb ¹⁴C moved throughout the carbon cycle and increasing fossil fuel CO2 emissions further decreased atmospheric Δ¹⁴CO2. The initially large seasonal cycle in the 1960s reduces in amplitude and eventually reverses in phase, resulting in a small seasonal cycle of about 2 ‰ in the 2000s. The seasonal cycle at Wellington is dominated by the seasonality of cross-tropopause transport and differs slightly from that at Cape Grim, Australia, which is influenced by anthropogenic sources in winter. Δ¹⁴CO2 at Cape Grim and Wellington show very similar trends, with significant differences only during periods of known measurement uncertainty. In contrast, similar clean-air sites in the Northern Hemisphere show a higher and earlier bomb ¹⁴C peak, consistent with a 1.4-year interhemispheric exchange time. From the 1970s until the early 2000s, the Northern and Southern Hemisphere Δ¹⁴CO2 were quite similar, apparently due to the balance of ¹⁴C-free fossil fuel CO2 emissions in the north and ¹⁴C-depleted ocean upwelling in the south. The Southern Hemisphere sites have shown a consistent and marked elevation above the Northern Hemisphere sites since the early 2000s, which is most likely due to reduced upwelling of ¹⁴C-depleted and carbon-rich deep waters in the Southern Ocean, although an underestimate of fossil fuel CO2 emissions or changes in biospheric exchange are also possible explanations. This developing Δ¹⁴CO2 interhemispheric gradient is consistent with recent studies that indicate a reinvigorated Southern Ocean carbon sink since the mid-2000s and suggests that the upwelling of deep waters plays an important role in this change.
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Atmos. Chem. Phys., 17, 14771–14784, 2017
https://doi.org/10.5194/acp-17-14771-2017
© Author(s) 2017. This work is distributed under
the Creative Commons Attribution 3.0 License.
Sixty years of radiocarbon dioxide measurements at Wellington,
New Zealand: 1954–2014
Jocelyn C. Turnbull1,2, Sara E. Mikaloff Fletcher3, India Ansell1, Gordon W. Brailsford3, Rowena C. Moss3,
Margaret W. Norris1, and Kay Steinkamp3
1GNS Science, Rafter Radiocarbon Laboratory, Lower Hutt, New Zealand
2CIRES, University of Colorado at Boulder, Boulder, Colorado, USA
3National Institute of Water and Atmospheric Research (NIWA), Wellington, New Zealand
Correspondence: Jocelyn C. Turnbull (j.turnbull@gns.cri.nz)
Received: 9 December 2016 – Discussion started: 20 December 2016
Revised: 12 October 2017 – Accepted: 1 November 2017 – Published: 12 December 2017
Abstract. We present 60 years of 114CO2measurements
from Wellington, New Zealand (41S, 175E). The record
has been extended and fully revised. New measurements
have been used to evaluate the existing record and to re-
place original measurements where warranted. This is the
earliest direct atmospheric 114CO2record and records the
rise of the 14C “bomb spike” and the subsequent decline in
114CO2as bomb 14 C moved throughout the carbon cycle
and increasing fossil fuel CO2emissions further decreased
atmospheric 114CO2. The initially large seasonal cycle in
the 1960s reduces in amplitude and eventually reverses in
phase, resulting in a small seasonal cycle of about 2 ‰ in
the 2000s. The seasonal cycle at Wellington is dominated
by the seasonality of cross-tropopause transport and differs
slightly from that at Cape Grim, Australia, which is influ-
enced by anthropogenic sources in winter. 114CO2at Cape
Grim and Wellington show very similar trends, with signifi-
cant differences only during periods of known measurement
uncertainty. In contrast, similar clean-air sites in the Northern
Hemisphere show a higher and earlier bomb 14C peak, con-
sistent with a 1.4-year interhemispheric exchange time. From
the 1970s until the early 2000s, the Northern and Southern
Hemisphere 114CO2were quite similar, apparently due to
the balance of 14C-free fossil fuel CO2emissions in the north
and 14C-depleted ocean upwelling in the south. The South-
ern Hemisphere sites have shown a consistent and marked el-
evation above the Northern Hemisphere sites since the early
2000s, which is most likely due to reduced upwelling of 14C-
depleted and carbon-rich deep waters in the Southern Ocean,
although an underestimate of fossil fuel CO2emissions or
changes in biospheric exchange are also possible explana-
tions. This developing 114CO2interhemispheric gradient is
consistent with recent studies that indicate a reinvigorated
Southern Ocean carbon sink since the mid-2000s and sug-
gests that the upwelling of deep waters plays an important
role in this change.
1 Introduction
Measurements of radiocarbon in atmospheric carbon diox-
ide (114CO2)have long been used as a key to understanding
the global carbon cycle. The first atmospheric 114CO2mea-
surements were begun at Wellington, New Zealand in 1954
(Rafter, 1955; Rafter and Fergusson, 1959), aiming to better
understand carbon exchange processes (Otago Daily Times,
1957). Northern Hemisphere 114CO2measurements began
a few years later in 1962, in Norway (Nydal and Løvseth,
1983), and 1959, in Austria (Levin et al., 1985).
14C is a cosmogenic nuclide produced naturally in the up-
per atmosphere through neutron spallation; it reacts rapidly
to form 14CO and then oxidizes to 14 CO2over a period of
1–2 months, after which it moves throughout the global car-
bon cycle. Natural 14C production is roughly balanced by ra-
dioactive decay, which mostly occurs in the carbon-rich and
slowly overturning ocean carbon reservoir and to a lesser
extent in the faster cycling terrestrial carbon reservoir. The
perturbations to 114CO2from atmospheric nuclear weapons
testing in the mid-20th century and additions of 14C-free CO2
Published by Copernicus Publications on behalf of the European Geosciences Union.
14772 J. C. Turnbull et al.: Sixty years of radiocarbon dioxide measurements
from fossil fuel burning have both provided tools to investi-
gate CO2sources and sinks.
The penetration of bomb-14C into the oceans has been
used to understand ocean carbon uptake processes (Oeschger
et al., 1975; Broecker et al., 1985; Key et al., 2004; Naegler
et al., 2006; Sweeney et al., 2007). Terrestrial biosphere car-
bon residence times and exchange processes have also been
widely investigated using bomb-14C (e.g. Trumbore, 2000;
Naegler and Levin, 2009). Stratospheric residence times,
cross-tropopause transport and interhemispheric exchange
can also be examined with atmospheric 114CO2observa-
tions (Kjellström et al., 2000; Kanu et al., 2015).
The Suess effect, the decrease in atmospheric 114CO2due
to the addition of 14C-free fossil fuel CO2, was first identi-
fied in 1955 (Suess, 1955). It has subsequently been refined
(Meijer et al., 1996; Levin et al., 2003) and used to investi-
gate fossil fuel CO2additions on various scales (e.g. Turnbull
et al., 2009a, 2015; Djuricin et al., 2010; Miller et al., 2012;
Lopez et al., 2013).
The full atmospheric 14C budget has been investigated us-
ing long-term 114CO2records in conjunction with atmo-
spheric transport models (Caldiera et al., 1998; Randerson
et al., 2002; Naegler et al., 2006; Turnbull et al., 2009b;
Levin et al., 2010). These have shown changing controls on
114CO2through time. Prior to nuclear weapons testing, nat-
ural cosmogenic production added 14C to the upper atmo-
sphere, which reacted to CO2and moved throughout the at-
mosphere and the carbon cycle. The short carbon residence
time in the biosphere meant that biospheric exchange pro-
cesses only had a small influence on 114CO2, whereas the
ocean exerted a stronger influence due to radioactive decay
during its much longer (and temporally varying) turnover
time. The addition of bomb 14C in the 1950s and 1960s al-
most doubled the atmospheric 14C content. This meant that
both the ocean and biosphere were very 14C-poor relative to
the atmosphere in the two decades following the atmospheric
test ban treaty. As the bomb-14C was distributed throughout
the carbon cycle, this impact weakened, and by the 1990s,
the additions of fossil fuel CO2became the largest contribu-
tor to the 114CO2trend (Randerson et al., 2002; Turnbull et
al., 2007; Levin et al., 2010; Graven et al., 2012).
The long-term 114CO2records have been crucial in all of
these findings, and the Wellington 114CO2record is of spe-
cial importance, being the oldest direct atmospheric trace gas
record, even predating the CO2mole fraction record started
at Mauna Loa in 1958 (Keeling, 1961; Keeling and Whorf,
2005). It is the only Southern Hemisphere record recording
the bomb spike. Several short Southern Hemisphere records
do exist (Manning et al., 1990; Meijer et al., 2006; Graven et
al., 2012; Hua et al., 2013), and some longer records began
in the 1980s (Levin et al., 2010). Over the more than 60 years
of measurement, there have necessarily been changes in how
the Wellington samples are collected and measured. There
are no comparable records during the first 30 years of mea-
surement, so that the data quality has not been indepen-
dently evaluated. Comparison with other records since the
mid-1980s has suggested that there may be biases in some
parts of the Wellington record (Currie et al., 2011).
Here we present a revised and extended Wellington at-
mospheric 14CO2record, spanning 60 years from December
1954 to December 2014. We detail the different sampling,
preparation and measurement techniques used through the
record, compare it with new tree ring measurements, discuss
revisions to the previously published data and provide a final
dataset with an accompanying smooth curve fit.
In the results and discussion, we revisit the key findings
that the Wellington 14CO2record has provided over the years
and expand them with new findings based on the most recent
part of the record. The most recent publication of this dataset
included data to 2005 (Currie et al., 2011) and showed pe-
riods of variability and a seasonal cycle at Wellington that
differ markedly from the independent Cape Grim, Tasma-
nia, 14CO2record at a similar southern latitude (Levin et al.,
2010). Here we add complementary new data to investigate
these differences, fill gaps and extend the record to the near-
present.
2 Methods
Over 60 years of measurement, a number of different sam-
ple collection, preparation, measurement and reporting meth-
ods have been used. In this section, we give an overview of
the various methods and changes through time, and they are
summarized in Table 1. Full details of the sampling methods
used through time are provided in the Supplement, compil-
ing methodological information documented in previous re-
ports on the Wellington record (Rafter and Fergusson, 1959;
Manning et al., 1990; Currie et al., 2011) along with meth-
ods newly applied in this new extension and refinement of
the dataset.
2.1 Sampling sites
Samples from 15 December 1954 to 5 June 1987 were col-
lected at Makara (Lowe, 1974), on the south-west coast of the
North Island of New Zealand (MAK; 41.25S, 174.69E;
300 m above sea level). Samples since 8 July 1988 have been
collected at Baring Head (Brailsford et al., 2012) on the
south coast of the lower North Island and 23 km south-east
of Makara (BHD; 41.41S, 174.87E; 80 m a.s.l.; Fig. 1).
We also discuss tree ring samples collected from Eastbourne,
12 km north of Baring Head on Wellington Harbour.
2.2 Collection methods
2.2.1 NaOH absorption
The primary collection method is static absorption of CO2
into nominally CO2-free 0.5 or 1 M sodium hydroxide
(NaOH) solution, which is left exposed to air at the sam-
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J. C. Turnbull et al.: Sixty years of radiocarbon dioxide measurements 14773
Table 1. Wellington 14CO2measurement methods through time. Gas counting samples are identified by NZ numbers, AMS samples by NZA
numbers. NZ and NZA numbers do not overlap. Sites are Makara (MAK) and Baring Head (BHD). Collection and measurement methods
are described in detail in the text.
Sampling date Sample ID Site Collection Measurement
range NZ/NZA method method
1954–1986 0–7500 MAK tray gas counting
1987–1994 7500–8400 BHD tray gas counting
1995–2004 8400–30 000 BHD bottle AMS ENTandem 13C14 C
2005–2009 30 000–34 000 BHD bottle AMS ENTandem 12C13 C14C
2010–2011 34 000–50 000 BHD bottle AMS XCAMS
2012–present 50 000– BHD bottle AMS XCAMS/RG20
20m
NaOHsamples
Tree
BaringHead
Eastbourne
Makara
Wellington
CGO
NWR
VER
JFJ
20 km
100m
Wellington
Porirua
HuttValley
(a)
(b) (c)
Figure 1. Sampling locations. (a) Makara (1954–1986) and Baring
Head (1987–present) air sampling sites, the location of the East-
bourne tree samples, and the urbanized areas of Wellington, Porirua
and the Hutt Valley. (b) World location showing Wellington and
other sampling sites discussed in the text. (c) Close-up of the Bar-
ing Head site showing the relative positions of the air (NaOH) and
tree sampling locations.
pling site providing an integrated sample over a period of
2 weeks (Sect. S3.1 in the Supplement; Rafter, 1955).
From 1954–1995, 2 L NaOH solution was exposed to air
in a large (450 cm2surface area) Pyrex®tray. Since 1995,
wide-mouth high-density polyethylene (HDPE) bottles con-
taining 200 mL NaOH solution were left open inside a
Stevenson meteorological screen; the depth of the solution in
the bottles remained the same as that in the previously used
trays. No significant difference has been observed between
the two methods (Currie et al., 2011). A few early (1954–
1970) samples were collected using different vessels, using
air pumped through the NaOH (vs. passive absorption) or by
replacing NaOH with barium hydroxide (Rafter, 1955; Man-
ning et al., 1990). CO2is extracted from the NaOH solution
by acidification followed by cryogenic distillation (Rafter
and Fergusson, 1959; Currie et al., 2011). Static NaOH ab-
sorption necessarily fractionates relative to CO2in the at-
mosphere. Typical δ13C values are 15 to 25 ‰ for these
samples, and this is corrected for in the data analysis.
2.2.2 Whole-air flasks
In this study, we use whole-air flask samples collected at
Baring Head to supplement and/or replace NaOH samples.
Flasks of whole air are collected by flushing ambient air
through the flask for several minutes; they are then filled to
slightly over ambient pressure. Most flasks were collected
during southerly, clean-air conditions (Stephens et al., 2013).
CO2is extracted cryogenically (Turnbull et al., 2015b). For
whole-air samples collected from 1984 to 1993, the extracted
CO2was archived until 2012. We evaluated the quality of this
archived CO2using two methods. Tubes with major leakage
were readily detected by air present in the tube and were dis-
carded. δ13C from all the remaining samples was in agree-
ment with δ13C measured from separate flasks collected at
Baring Head and measured for δ13C by the Scripps Institu-
tion of Oceanography close to the time of collection (http:
//scrippsco2.ucsd.edu/data/atmospheric_co2/nzd, Keeling et
al., 2001). Whole-air samples collected since 2013 are anal-
ysed for δ13C and other trace gases and isotopes at NIWA
(National Institute of Water and Atmospheric Research, New
Zealand) (Ferretti et al., 2000), and for the 14CO2measure-
ment, CO2is extracted from whole air at the Rafter Radio-
carbon Laboratory (Turnbull et al., 2015b).
2.2.3 Tree rings
When trees photosynthesize, they faithfully record the 114C
of ambient CO2in their cellulose, the structural compo-
nent of wood. Annual tree rings therefore provide a sum-
mertime (approximately September–April in the Southern
Hemisphere) daytime average 114CO2. Photosynthetic up-
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14774 J. C. Turnbull et al.: Sixty years of radiocarbon dioxide measurements
take varies during the daylight hours depending on factors
including growth period, sunlight and temperature (Bozhi-
nova et al., 2013), resulting in a somewhat different effective
sampling pattern than the 1–2-week NaOH solution collec-
tions. We show in Sect. 3.5.1 that at the Wellington location
this difference is negligible. Note that we assign the mean
age of each ring as 1 January of the year in which growth
finished (i.e. the mean age of a ring growing from Septem-
ber to April), whereas dendrochronologists assign the “ring
year” as the year in which ring growth started (i.e. the previ-
ous year).
We collected cores from three trees close to the Baring
Head site. A pine (Pinus radiata) located 10 m from the
Baring Head sampling station (Fig. 1) yielded rings back to
1986 (Norris, 2015). A longer record was obtained from two
New Zealand kauri (Agathis australis) specimens planted
in 1919 and 1920, located 20 m from one another in East-
bourne, 12 km from Baring Head (Fig. 1). Kauri is a long-
lived high-density softwood species that has been widely
used in dendrochronology and radiocarbon calibration stud-
ies (e.g. Hogg et al., 2013).
Annual rings were counted from each core. Shifting the
Eastbourne record by 1 year in either direction moves the 14C
bomb spike maximum out of phase with the NaOH-based
Wellington 114CO2record (Supplement Fig. S1), confirm-
ing that the ring counts are correct. For the Baring Head pine,
rings go back to only 1986, and we verify them by comparing
them with the Eastbourne record. They show an insignificant
mean difference of 0.4 ±0.8 ‰ (Fig. S1).
In practice, it is difficult to ensure that one annual ring
is sampled without losing any material from that ring, and
no wood from surrounding rings is included. To evaluate the
potential bias from this source, we measured replicate sam-
ples from different cores from the same tree (Baring Head) or
two different trees (Eastbourne). For samples collected since
1985, all these replicates are consistent within their assigned
uncertainties (Fig. S2). However, for three replicates from
Eastbourne in 1963, 1965 and 1971, we see large differences
of 9.2, 44.5 and 4.9 ‰, which we attribute to small differ-
ences in the sampling of the rings that were magnified by the
rapid change in 114C of up to 200 ‰ yr1during this period.
Thus, the tree ring 114C values during this period should be
treated with caution.
Cellulose was isolated from whole tree rings by first re-
moving labile organics with solvent washes and then oxidiz-
ing the resultant material to isolate the cellulose from other
materials (Norris, 2015; Hua et al., 2000). The cellulose was
combusted and the CO2purified following standard methods
in the Rafter Radiocarbon Laboratory (Baisden et al., 2013).
2.3 14C measurement
Static NaOH samples were measured by conventional decay
counting on the CO2gas from 1954 to 1995 (Manning et
al., 1990; Currie et al., 2011), and these samples are identi-
fied by their unique “NZ” numbers. All measurements made
since 1995, including recent measurements of flask samples
collected in the 1980s and 1990s, were reduced to graphite,
measured by accelerator mass spectrometry (AMS), and are
identified by their unique “NZA” numbers. The LG1 graphi-
tization system was used from 1995 to 2011 (NZA < 50 000;
Lowe and Judd, 1987), and replaced with the RG20 graphite
system in 2011 (NZA >50 000; Turnbull et al., 2015b). Sam-
ples measured by AMS were stored for up to 3 years be-
tween sample collection and extraction, graphitization and
measurement.
For samples collected from 1995 to 2010, an EN Tandem
AMS was used for measurement (NZA < 35 000; Zondervan
and Sparks, 1996). Until 2005 (NZA < 30000, including all
previously reported Wellington 14 CO2data), only 13C and
14C were measured on the EN Tandem system, so the nor-
malization correction for isotopic fractionation (Stuiver and
Polach, 1977) was performed using an offline isotope ratio
mass spectrometer δ13C value. The data reported from 2005
onwards (NZA > 30000) show a reduction in scatter reflect-
ing the addition of online 12C measurement in the EN Tan-
dem system in 2005. This allows direct online correction for
isotopic fractionation that may occur during sample prepa-
ration and in the AMS system (Zondervan et al., 2015) and
results in improved long-term repeatability. Fractionation in
the AMS system may vary in sign depending on the par-
ticular conditions, but incomplete graphitization biases the
graphite towards lighter isotopes, which, if undiagnosed, will
bias 114C high. The LG1 graphitization system used during
this period did not directly evaluate whether graphitization
was complete, so it is possible or even likely that there was
a high bias in the 1995–2005 measurements. This is further
discussed in Sect. 3.5.3.
For all EN Tandem samples, a single large aliquot of ex-
tracted CO2was split into four separately graphitized and
measured targets and the results of all four were averaged.
We have revisited the multi-target averaging, applying a con-
sistent criterion to exclude outliers and using a weighted
mean of the retained measurements (Supplement). This re-
sults in differences of up to 5 ‰ relative to the values re-
ported by Currie et al. (2011) and is discussed in more detail
in the Supplement.
In 2010, the EN Tandem was replaced with a Na-
tional Electrostatics Corporation AMS, dubbed XCAMS
(NZA > 34 000). XCAMS measures all three carbon iso-
topes, such that the normalization correction is performed us-
ing the AMS-measured 13C values (Zondervan et al., 2015).
XCAMS measurements are made on single graphite targets
measured to high precision of typically 1.8 ‰ (Turnbull et
al., 2015b).
2.4 Results format
NaOH samples are collected over a period of typically
2 weeks and sometimes much longer. We report the date of
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J. C. Turnbull et al.: Sixty years of radiocarbon dioxide measurements 14775
collection as the average of the start and end dates. In cases
where the end date was not recorded, we use the start date.
For a few samples, the sampling dates were not recorded or
are ambiguous, and those results have been excluded from
the reported dataset.
Results are reported here as F14C (Reimer et al., 2004) and
114C (Turnbull et al., 2007). F14C is corrected for isotopic
fractionation and blank corrected. We calculated F14C from
the original measurement data recorded in our databases and
updated a handful of records where transcription errors were
found. 114C is derived from F14C and corrected for radioac-
tive decay since the time of collection; this is slightly dif-
ferent from 114C as defined by Stuiver and Polach (1977),
which is corrected to the date of measurement. 114C has
been recalculated using the date of collection for all results,
resulting in changes of a few tenths of per mil in most 114C
values relative to those reported by Currie et al. (2011) and
Manning et al. (1990). Uncertainties are reported based on
the counting statistical uncertainty, and for AMS measure-
ments, we add an additional error term, determined from
the long-term repeatability of secondary standard materi-
als (Turnbull et al., 2015b; Sect. S5.3). Samples for which
changes have been made relative to the previously published
results are indicated by the quality flag provided in the Sup-
plement dataset. Where more than one measurement was
made for a given date, we report the weighted mean (Bev-
ington and Robinson, 2003) of all measurements.
2.5 Smooth curve fits
In addition to the raw measured 114CO2values, we cal-
culate a smooth curve fit and deseasonalized trend from
the Wellington 114C and F14 C datasets. The deseasonalized
trend may be more useful than the raw data for the aging of
recent materials (e.g. Reimer et al., 2004; Hua et al., 2013).
Acknowledging that the 1995–2005 period is variable and
possibly biased in the Wellington record (Sect. 4.3), we also
provide in the Supplement an alternative midlatitude South-
ern Hemisphere smooth curve fit and deseasonalized trend in
which the Wellington data for 1995–2005 has been removed
and replaced with the Cape Grim, Tasmania, data for that pe-
riod (Levin et al., 2010).
Curve fitting is particularly challenging for the 114CO2
record since (a) there are data gaps and inconsistent sam-
pling frequency, (b) the growth rate and trend vary dramati-
cally, and (c) the seasonal cycle changes both in magnitude
and phase (Sect. 5.2). We chose to use the CCGCRV fitting
procedure (Thoning et al., 1989; www.esrl.noaa.gov/gmd/
ccgg/mbl/crvfit/), which can readily handle the data gaps,
inconsistent sampling frequency and rapid changes in the
trend. To address the changing seasonal cycle, we make sep-
arate fits to the record for five time periods: 1954–1965,
1966–1979, 1980–1989, 1990–2004 and 2005–2014. These
divisions were chosen based on major changes in the raw
observational growth rate, seasonal cycle and data quality
(Sect. S6). For each time period, we use CCGCRV with one
linear and two harmonic terms, and fit residuals are added
back using a low-pass filter with an 80-day cut-off in the fre-
quency domain. At each transition, we overlapped a 2-year
period and linearly interpolated the two fits across that 2-
year period to smooth the transitions caused by end effects.
The 1σuncertainty in the smoothed curve and deseasonal-
ized trend were determined using a Monte Carlo technique.
Further details of the fitting procedure and choice of time pe-
riod cut-offs are provided in the Supplement.
The mean difference between the fitted curve and the mea-
sured 114CO2values is 3.8 ‰, consistent with the typical
measurement uncertainty for the full dataset. Further, the
residuals are highest for the early period (1954–1970) at 6 ‰,
consistent with the larger measurement errors at that time of
6 ‰. The residuals improve as the measurement errors re-
duce, such that since 2005, the mean residual is 2 ‰, consis-
tent with the reported 2 ‰ uncertainties. The exception is the
1995–2005 period where the mean residual difference of 5 ‰
is substantially higher than the mean reported uncertainty of
2.5 ‰, reflecting the apparent larger scatter during this pe-
riod (Sect. 4.3).
3 Data validation
3.1 Tree ring comparison
Over the more than 60 years of the Wellington 114 CO2
record, there have necessarily been many changes in method-
ology, and the tree rings provide a way to validate the full
record, albeit with lower resolution. Due to the possible sam-
pling biases in the tree rings (Sect. 3.2.3.), we do not include
them in the final updated record but use them to validate the
existing measurements.
During the rapid 114CO2change in the early 1960s, there
are some differences between the kauri tree ring and Welling-
ton 114CO2records (Fig. 2). The 1963 and 1964 tree ring
samples are slightly lower than the concurrent 114CO2sam-
ples. The peak 114CO2measurement in the tree rings is
30 ‰ lower than the smoothed 114 CO2record and 100 ‰
lower than the two highest 114CO2measurements in 1965.
These differences are likely due to small errors in the sam-
pling of the rings, which will be most apparent during periods
of rapid change.
Prior to 1960 and from the peak of the bomb spike in
1965 until 1990, there is remarkable agreement between the
tree rings and Wellington 114CO2record, with the variabil-
ity replicated in both records. Since 2005, there is excellent
agreement across all the different records. Some differences
are observed in 1990–1993 and 1995–2005, which we dis-
cuss in the following sections.
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14776 J. C. Turnbull et al.: Sixty years of radiocarbon dioxide measurements
Figure 2. Wellington 14 CO2record showing all collection and measurement methods for the full record (a) and zoomed in for the period
since 1980 (b). Tree rings (green) and outliers (grey pluses) are excluded from the reported final dataset. The black line is the smooth curve
fit to the final dataset.
3.2 1990–1993 anomaly
An anomaly in the gas counting measurements between 1990
and 1993 has previously been noted (Figs. 2, 3) as a devia-
tion from the Cape Grim 114CO2record (CGO; 40.68S,
144.68E; 94 m a.s.l.; Levin et al., 2010) during the same
period. Cape Grim is at similar latitude, and observes a mix-
ture of air from the midlatitude Southern Ocean sector and
mainland Australia (Ziehn et al., 2014; Law et al., 2010). The
Wellington and Cape Grim records overlap during almost all
other periods (Fig. 3).
We use archived CO2from flask samples to evaluate this
period of deviation. First, the recent flask samples collected
since 2013 (n=12) agree very well with the NaOH static
samples from the same period (Fig. 2), indicating that de-
spite the difference in sampling period for the two methods,
flask samples reflect the 114CO2observed in the longer-term
NaOH static samples. We then selected a subset of archived
1984–1992 extracted CO2samples for measurement, mostly
from southerly wind conditions but including a few from
other wind conditions. These flask 114CO2measurements
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J. C. Turnbull et al.: Sixty years of radiocarbon dioxide measurements 14777
do not exhibit the anomaly seen in the NaOH static samples
(Fig. 2), implying that the deviation observed in the origi-
nal NaOH static samples may be a consequence of sampling,
storage or measurement errors. Annual tree rings from both
the kauri and pine follow the flask measurements for this pe-
riod (Fig. 2), confirming that the NaOH static samples are
anomalous.
The 1990–1993 period was characterized by major
changes in New Zealand science, both in the organizational
structure and personnel. Although we are unable to ex-
actly reconstruct events at that time, we hypothesize that
the NaOH solution was prepared slightly differently, per-
haps omitting the barium chloride precipitation step for these
samples. This would result in contaminating CO2absorbed
on the NaOH before the solution was prepared. Since at-
mospheric 114CO2is declining, this would result in higher
114CO2observed in these samples than in the ambient air.
Another possibility is that there were known issues with the
background contamination in the proportional counters dur-
ing this period that could result in a high-bias 114CO2. In
any case, these values are anomalous, and we remove the
original NaOH static sample measurements between 1990
and 1993 and replace them with the new flask measurements
for the same period.
3.3 1995–2005 variability
As already discussed in Sect. 3.3, the measurement method
was changed from gas counting to AMS for samples col-
lected in 1995 and thereafter. During the first 10 years of
AMS measurements, the record is much noisier than during
any other period (Fig. 2). Until 2005, offline δ13C measure-
ments on the evolved CO2were used in the normalization
correction. In 2005, online 12C measurement was added to
the AMS system, allowing online AMS measurement of the
δ13C value and accounting for any fractionation during sam-
ple preparation and AMS measurement (Zondervan et al.,
2015; see also Sect. 3.3). This substantially improved the
measurement accuracy, and the noise in the 114CO2record
immediately reduced as can be seen in the lower panel of
Fig. 2. Therefore, we suspect that the variability and appar-
ent high bias in the 1995–2005 period of the 114CO2record
is due to measurement uncertainty and bias rather than atmo-
spheric variability.
The remaining NaOH solution for all samples collected
since 1995 has been archived, and typically only every sec-
ond sample collected was measured, with the remainder
archived without extraction. In 2011–2016, we revisited the
1995–2005 period, remeasuring some samples that had pre-
viously been measured and some that had never been mea-
sured for a total of 52 new analyses.
The new measurements for this time period do show re-
duced scatter over the original analyses, particularly for the
period from 1998 to 2001, when the original analyses ap-
pear anomalously low, and in 2002–2003, when the original
Figure 3. Comparison of the final Wellington and Cape Grim (Levin
et al., 2010) 114CO2records. Wellington tree ring measurements
are also shown.
analyses appear anomalously high. Yet there remain a num-
ber of both low and high outliers in the new measurements.
These are present in both the samples that were remeasured
and in those for which this was the first extraction of the sam-
ple. This suggests that a subset of the archived sample bot-
tles were either contaminated at the time of collection or that
some bottles were insufficiently sealed, causing contamina-
tion with more recent CO2during storage. Comparison with
the tree ring measurements and with the Cape Grim record
(Levin et al., 2010) suggests that the measurements during
this period may, on average, be biased high as well as having
additional scatter (Fig. 3). Nonetheless, in the absence of bet-
ter data, we retain both the original and remeasured NaOH
sample results in the full Wellington record, with a special
flag to allow users to easily remove the questionable results
if they prefer. We also provide a smoothed fit that excludes
these data (Sect. 3.6).
4 Results and discussion
4.1 Variability in the Wellington record through time
The Wellington 114CO2record begins in December 1954, at
a roughly pre-industrial 114CO2level of 20 ‰ (Fig. 2).
From 1955, 114CO2increased rapidly, near-doubling to
700 ‰ in 1965 at Wellington, due to the production of 14C
during atmospheric nuclear weapons tests. Nuclear tests in
the early 1950s contributed to the rise; then a hiatus in test-
ing in the late 1950s led to a plateau in Wellington 114CO2
before a series of very large atmospheric tests in the early
1960s led to further increases (Rafter and Ferguson, 1959;
Manning et al., 1990).
Most atmospheric nuclear weapons testing ceased in 1963,
and the Wellington 114CO2record peaks in 1965 and then
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begins to decline, at first rapidly at 30 ‰ yr1in the 1970s
and gradually slowing to 5 ‰ yr1after 2005. The initial
rapid decline has been attributed primarily to the uptake of
the excess radiocarbon into the oceans and, to a lesser extent,
to uptake into the terrestrial biosphere (Naegler et al., 2006;
Randerson et al., 2002; Manning et al., 1990; Stuiver and
Quay, 1981). The short residence time of carbon in the bio-
sphere means that from the 1980s, the terrestrial biosphere
changed from a 14C sink to a 14 C source as the bomb pulse
was re-released (Randerson et al., 2002; Levin et al., 2010).
Natural cosmogenic production of 14C damps the rate of
decline since the bomb peak by 5 ‰ yr1in 114CO2; this
may vary with the solar cycle, but there is no known long-
term trend in this component of the signal (Turnbull et al.,
2009b; Naegler et al., 2006). There is also a small positive
contribution from the nuclear industry, which emits 14 C to
the atmosphere, and this has increased from 0 in the 1950s
to 0.5–1 ‰ yr1in the last decade (Turnbull et al., 2009b;
Levin et al., 2010; Graven and Gruber, 2011).
The Suess effect, the decrease in atmospheric 114CO2due
to the addition of 14C-free fossil fuel CO2to the atmosphere
(Suess, 1955; Tans et al., 1979; Levin et al., 2003), was first
recognized in 1955 and has played a role throughout the
record. Although the magnitude of fossil fuel CO2emissions
has grown through time, when convolved with the declining
atmospheric 114CO2history, the impact on 114CO2stayed
roughly constant at 10 ‰ yr1from the 1970s to the mid-
2000s (Randerson et al., 2002; Levin et al., 2010; Graven et
al., 2012). Yet the continued increase in fossil fuel CO2emis-
sions has slightly increased the impact of fossil fuel CO2in
the last few years, to about 12 ‰ yr1in 2014 (using an-
nual global fossil fuel CO2estimates from CDIAC (Carbon
Dioxide Information Analysis Center); Boden et al., 2017).
Since the 1990s, the Suess effect has been the largest driver
of the ongoing negative growth rate (Turnbull et al., 2009b;
Levin et al., 2010).
The most recent part of the Wellington 114CO2record
from 2005–2014 is reported here for the first time. It shows
a continuing downward trend in 114CO2of 5 ‰ yr1and
a slight slowing in the negative trend relative to the 1990–
2004 period, which had a trend of 5.8 yr1. This slight
slowing in the downward 114CO2trend is the opposite of
what might be expected due to the Suess effect alone. Pos-
sible explanations are a slowing of the rate of uptake of 14C
into the oceans, an increase in the return rate of bomb 14C to
the atmosphere from the biosphere and a long-term increase
in 14C production.
4.2 Seasonal variability in the Wellington record
We determine the changing seasonal cycle from smooth
curve fits to five separate periods of the record (1954–1965,
1966–1979, 1980–1989, 1990–2004, 2005–2014; Fig. 4a).
This subdivision is necessary to allow the seasonal cycle to
vary through time since the CCGCRV curve fitting routine
Figure 4. Detrended seasonal cycle in the Wellington 114 CO2
record. (a) BHD monthly detrended seasonal cycle averaged over
four time periods as described in the text and the CGO (Levin et al.,
2010) detrended seasonal cycle. Error bars are the standard devia-
tion of all years averaged. Points for each time period are slightly
offset for clarity. (b) Full seasonal cycle record determined sepa-
rately for each time period shown in (a) plus 1954–1965 (black)
and detrended observations without any smoothing (grey).
assigns a single set of harmonics to the time period fitted (see
Sect. 3.5). The choice of time periods is discussed in Sect. S6.
We also created detrended, fitted 114CO2seasonal cycles by
subtracting the deseasonalized trend from the observations.
Comparison with the detrended, fitted seasonal cycle deter-
mined from the smooth curve fits (Fig. 4b) shows that the
smooth curve fit, as might be expected, does not capture the
largest deviations from the trend seen in the observations but
represents the changing seasonal cycle quite well.
The 1966–1979 period shows a strong seasonal cycle
(Fig. 4) with a consistent phase and an amplitude that varies
from a maximum in 1966 of 30 ‰ gradually declining to 3 ‰
in 1979, with a mean amplitude of about 6 ‰. This is primar-
ily attributed to seasonally varying stratosphere–troposphere
exchange bringing bomb 14C into the troposphere (Manning
et al., 1990; Randerson et al., 2002). Manning et al. (1990)
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were unable to simulate the correct phasing of the seasonal
cycle, apparently because their model distributed bomb 14C
production throughout both the northern and southern strato-
sphere. In fact, the majority of the bomb 14C was produced in
the Northern Hemisphere stratosphere (Enting et al., 1982).
Randerson et al. (2002) were able to match the amplitude
of the Wellington seasonal cycle during this time period, al-
though their model was out of phase with the observations
by about 1.5 months. They attribute the seasonal cycle dur-
ing this period mostly to the seasonality in Northern Hemi-
sphere stratosphere–troposphere exchange with a phase lag
caused by cross-equator exchange into the Southern Hemi-
sphere. The seasonal cycle kept the same phase but gradu-
ally decreased in amplitude until the late 1970s, attributed
to the declining disequilibrium between the stratosphere and
troposphere as the bomb 14C moved throughout the carbon
reservoirs.
Between 1978 and 1980 the seasonal cycle weakened and
then reversed during the 1980s, with a maximum in win-
ter (June–August) and an amplitude of about 2 ‰. The de-
trended observations show that this change in phase is not an
artefact of the fitting method (Fig. 4b). This result is compa-
rable to that obtained by Manning et al. (1990) and Currie et
al. (2011), who both used a seasonal trend loss (STL) pro-
cedure to determine the seasonal cycle from the same data.
This is consistent with the seasonality in atmospheric trans-
port convolving with a change in sign of the terrestrial bio-
sphere contribution as the bomb 14C pulse began to return to
the atmosphere from the biosphere (Randerson et al., 2002).
The Wellington 114CO2seasonal cycle declined in the
1990s, and the larger variability in the observations between
1995 and 2005 makes it difficult to discern a seasonal cycle
during that period. Since 2005, the more precise measure-
ments allow us to detect a small seasonal cycle with an am-
plitude of about 2 ‰ (Fig. 4). We compare the seasonal cycle
at Wellington from 2005 to 2015 with the seasonal cycle at
Cape Grim, Australia, from 1995 to 2010. There is no sig-
nificant difference in the seasonal cycle at either site if we
select only the overlapping time period of 2005–2010. Both
sites show a similar magnitude seasonal cycle during this
period, and Cape Grim shows a maximum in March–April
that has been attributed primarily to the seasonality of atmo-
spheric transport of Northern Hemisphere fossil fuel emis-
sions to the southern troposphere (Levin et al., 2010). This
maximum at Cape Grim coincides with a seasonal maximum
in the Wellington record. However, Wellington 114CO2ex-
hibits a second maximum in the austral spring (October) that
is not apparent at Cape Grim.
Recent work has shown that during the winter, the Cape
Grim station is influenced by air coming off the Australian
mainland, including the city of Melbourne (Ziehn et al.,
2014), which would act to reduce 114CO2at Cape Grim
relative to Southern Ocean clean air. This shift is shown to
be the result of seasonal variations in atmospheric transport.
The 2-week integrated sampling used for 114CO2at both
Cape Grim and Baring Head means that in contrast to other
species, 114CO2measurements cannot be screened to re-
move these pollution events.
In contrast, the Baring Head location near Wellington does
not show significant seasonal variation in atmospheric trans-
port (Steinkamp et al., 2017) and Baring Head is less likely
than Cape Grim to be influenced by anthropogenic emis-
sions in any season. Air is typically from the ocean, and
the local geography means that the urban emission plume
from Wellington and its northern suburbs of Lower Hutt very
rarely passes over Baring Head (Fig. 1), and the typically
high wind speeds further reduce the influence of the local
urban area (Stephens et al., 2013). During the austral au-
tumn, there is some land influence from the Christchurch re-
gion in the South Island, but emissions from Christchurch
are much smaller than the Melbourne emissions influenc-
ing Cape Grim: State of Victoria fossil fuel CO2emissions
for 2013 were 23 MtC, whereas Wellington and Christchurch
each emitted 0.4 MtC of fossil fuel CO2in 2013 (Boden et
al., 2017; AECOM, 2016; Australian Government, 2016).
Although broad-scale flow from the west is common, the
local topography means that local air flow is almost always
either southerly or northerly (Stephens et al., 2013), but dur-
ing rare (< 5 % of the time) westerly wind events, fossil fuel
emissions from Wellington do appear to cause enhancements
of up to 2 ppm in CO2(Stephens et al., 2013), which would
decrease 114CO2by 1 ‰ during such an event. Yet there is
no evidence of seasonality in the infrequent westerly events.
Northerly conditions bring a terrestrial biosphere influence
that elevates CO2by about 1 ppm (Stephens et al., 2013),
which could result in a maximum increase in 114CO2of
0.2 ‰ relative to background conditions, but there is no ev-
idence that this influence is seasonally variable either. Thus,
although there are some local influences on the Baring Head
114CO2, none of these appear to be seasonally dependent
and instead, the observed Baring Head 114CO2maximum
in spring in the recent part of the record may be explained by
the seasonal maximum in cross-tropopause exchange bring-
ing 14C-enriched air at this time of year.
4.3 Comparison with other atmospheric 114CO2
records
We compare the Wellington 114CO2record with several
other 114CO2records, located as indicated in Fig. 1. First,
we compare it with measurements from Cape Grim, Aus-
tralia (CGO; 40.68S, 144.68E; 94 m a.s.l.). Cape Grim
is at similar latitude to Wellington and also frequently re-
ceives air from the Southern Ocean (Levin et al., 2010).
Samples are collected by a similar method to the Welling-
ton record using NaOH absorption and are measured by
gas counting to 2 ‰ precision. Next, we compare them
with midlatitude high-altitude clean-air sites in the Northern
Hemisphere. The Vermunt, Austria (VER, 47.07N, 9.57E,
1800 m a.s.l.), record began in 1958, only a few years after
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Table 2. 114CO2gradients between sites, determined as the mean
of the monthly differences for each time period. Errors are the stan-
dard deviation of the monthly differences.
Site difference Time period 114CO2difference
(‰)
BHD-CGO 1986–1990 1.8 ±2.5
BHD-CGO 2005–2013 1.3 ±3.4
BHD-JFJ 1986–1990 0.8 ±3.9
BHD-JFJ 2005–2013 4.8 ±2.7
BHD-NWR 2005–2013 6.9 ±2.5
the Wellington record began, and in the 1980s the site was
moved to Jungfraujoch, Switzerland (JFJ, 46.55N, 7.98E,
3450 m a.s.l.); these measurements are made in the same
manner and by the same laboratory as the Cape Grim record
(Levin et al., 2013). We also consider the Niwot Ridge, USA
114CO2record (NWR, 40.05N, 105.59W, 3523 m a.s.l.),
which began in 2003 (Turnbull et al., 2007; Lehman et al.,
2013). Niwot Ridge is also a midlatitude, high-altitude site,
but samples are collected as whole air in flasks and mea-
sured by AMS in a similar manner to that described for the
Wellington flask samples. Thus, we are comparing two inde-
pendent Southern Hemisphere records with two independent
Northern Hemisphere records, with the two hemispheres tied
together by the common measurement laboratory used for
Cape Grim and Jungfraujoch. Results from all records are
compared in Fig. 5.
The Wellington and Cape Grim records are generally con-
sistent with one another (Fig. 3), with the exception of the
1995–2005 period, when the Wellington record is slightly
higher, apparently due to bias in the Wellington record (dis-
cussed in Sect. 3.5.3.). Differences between the sites are
smaller than the measurement uncertainty for all other pe-
riods (Table 2). This implies that 114CO2is homogeneous
across Southern Hemisphere clean-air sites within the same
latitude band, at least since the 1980s when the two records
overlap. Similarly, the high-altitude, midlatitude Northern
Hemisphere sites are consistent with one another, although
there are some differences in seasonal cycles in recent years
(Turnbull et al., 2009b).
The bomb spike maximum is higher and earlier in the
Northern Hemisphere records (Fig. 5), consistent with the
production of most bomb 14C in the Northern Hemisphere
stratosphere. We make a new, simple estimate of the inter-
hemispheric exchange time during the 1963–1965 period us-
ing the difference in the timing of the Northern and Southern
Hemisphere bomb peaks. The first maximum of the bomb
peak was in July 1963 in the Northern Hemisphere and Jan-
uary 1965 in the Southern Hemisphere, a 1.4-year offset, im-
plying a 1.4-year exchange time. This is consistent with other
more detailed interhemispheric exchange time estimates that
Figure 5. Comparison of Wellington and other atmospheric
114CO2records (Levin et al., 2010; Turnbull et al., 2007; Lehman
et al., 2013).
have been determined from long-term measurements of SF6
of 1.3 to 1.4 years (Geller at el., 1997; Patra et al., 2011).
Northern Hemisphere 114CO2remains higher than South-
ern Hemisphere 114CO2by about 20 ‰ until 1972. Al-
though most nuclear weapons testing ceased in 1963, a few
smaller tests continued in the late 1960s, contributing to this
continued interhemispheric offset (Enting, 1982). The inter-
hemispheric gradient disappeared within about 1.5 years af-
ter atmospheric testing essentially stopped in 1970. Except
for periods of noisy data from Vermunt in the late 1970s and
Wellington in 1995–2005, there are only small (< 2‰) inter-
hemispheric gradients from 1972 until 2002 (Fig. 5, Table 2).
As previously noted by Levin et al. (2010), using a shorter
dataset, an interhemispheric gradient of 5–7 ‰ develops in
2002, with the Southern Hemisphere sites higher than the
Northern Hemisphere sites (Table 2). We choose 1986–1990
and 2005–2013 as time periods to compare, to avoid the pe-
riods where the Wellington record is noisy (1995–2005) and
where we substituted flask measurements from 1990 to 1993.
In 1986–1990, there is less than 2 ‰ difference between
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J. C. Turnbull et al.: Sixty years of radiocarbon dioxide measurements 14781
Wellington and either Cape Grim or Jungfraujoch. There is
also no difference between the Cape Grim and Jungfrau-
joch records during this time period. The Wellington and
Cape Grim records still agree within 2 ‰ after 2005, but
both Jungfraujoch and Niwot Ridge diverge from Welling-
ton, by 4.8 ±2.7 and 6.9 ±2.5 ‰, respectively; Jungfraujoch
and Niwot Ridge are not significantly different from one an-
other. This new interhemispheric gradient is robust, being
consistent amongst the sites measured by three different re-
search groups each with their own methods. It is not an arte-
fact of interlaboratory offsets, since Cape Grim and Jungfrau-
joch measurements are made by the same group using the
same sampling and measurement methods, and the Welling-
ton and Niwot Ridge measurements (measured by different
techniques) agree well with the other sites at similar latitude
(Cape Grim and Jungfraujoch, respectively). This develop-
ing gradient is also apparent in the larger sampling network
of Levin et al. (2010) and in a separate 114CO2sampling
network (Graven et al., 2012), although that dataset extends
only to 2007.
Graven et al. (2012) demonstrated that increasing (mostly
Northern Hemisphere) fossil fuel CO2emissions cannot ex-
plain this 114CO2interhemispheric gradient, and instead,
they postulated that 14C uptake into the Southern Ocean re-
duced over time. Levin et al. (2010) were able to roughly
replicate this interhemispheric gradient in their GRACE
model by tuning the terrestrial biosphere fluxes to match
the observed global average atmospheric CO2and 114CO2
records. Where the observations suggest the rapid devel-
opment of an interhemispheric gradient in the early 2000s
(Fig. 5), the GRACE model simulates a more gradual transi-
tion over a period of roughly two decades. Independent evi-
dence suggests that the Southern Ocean is more likely to be
responsible for this rapid shift in the atmospheric 114CO2
gradient. That is, an apparent reorganization of Southern
Ocean carbon exchange in the early 2000s (Landschützer et
al., 2015) is postulated to be associated with changes in the
upwelling of deep water (DeVries et al., 2017), to which at-
mospheric 114CO2is highly sensitive (Rodgers et al., 2011;
Graven et al., 2012). The observed 114CO2interhemispheric
gradient is consistent with these postulated changes in up-
welling. Other possible explanations for this new interhemi-
spheric 114CO2gradient are a substantial underreporting of
Northern Hemisphere fossil CO2emissions (e.g. Francey et
al., 2013) or changes in the land carbon sink (Wang et al.,
2013; Sitch et al., 2015), although this latter is less likely
since 114CO2is much less sensitive to biospheric fluxes than
to either ocean or fossil fuel fluxes (e.g. Levin et al., 2010;
Turnbull et al., 2009b). Given the limited spatial coverage of
the current 114CO2observing network, it is not possible to
robustly determine which of these processes causes the in-
terhemispheric gradient. This could be achieved with more
observations of the spatial and temporal variations in atmo-
spheric 114CO2.
5 Conclusions
The 60-year-long Wellington 114 CO2record was revised
and extended to 2014. Most revisions were minor, but we
particularly note that the earlier reported 1990–1993 mea-
surements have been entirely replaced with new measure-
ments. A second period form 1995–2005 has poorer data
quality than the rest of the record, and may also be biased
high by a few per mil. These data have been revised substan-
tially, and new measurements have been added to this period,
but we were unable to definitively identify or correct for bias,
so the data have been retained, albeit with caution. We fur-
ther validated the record by comparison with tree ring sam-
ples collected from the Baring Head sampling location and
from nearby Eastbourne, Wellington; both tree ring records
show excellent agreement with the original record, and in-
dicate that there are no other periods in which the original
measurements are problematic.
The Wellington 114CO2time series records the history
of atmospheric nuclear weapons testing and the subsequent
decline in 114CO2as the bomb 14 C moved throughout the
carbon cycle and 14C-free fossil fuel emissions further de-
creased 114CO2. The timing of the first appearance of the
bomb-14C peak at Wellington is consistent with other recent
estimates of interhemispheric exchange time at 1.4 years.
The seasonal cycle at Wellington evolves through the
record, apparently dominated by the seasonality of cross-
tropopause transport, which drives a changing seasonal cy-
cle through time. In the early post-bomb period, the sea-
sonally variable movement of bomb 14C from the northern
stratosphere through the northern troposphere to the south-
ern troposphere appears to be the dominant control on the
seasonal cycle at Wellington. The seasonal cycle reversed in
later years, possibly due to a change in sign of the terres-
trial biosphere 114C signal. In recent years, the seasonal cy-
cle has an amplitude of only 2 ‰, with a maximum in the
austral spring. Cape Grim exhibits a similar seasonal cycle
magnitude but appears to be very slightly influenced by a
terrestrial/anthropogenic signal during the austral winter that
is not apparent at Wellington.
During the 1980s and 1990s, 114CO2was similar at mid-
latitude clean-air sites in both hemispheres, but since the
early 2000s, the Northern Hemisphere 114CO2has dropped
below the Southern Hemisphere by 5–7 ‰. The control on
this changing interhemispheric gradient cannot be robustly
determined from the existing sparse 114CO2observations
but may be due to a change in Southern Ocean dynamics
reducing the upwelling of old, 14C-poor deep waters, con-
sistent with recent evidence of an increasing Southern Ocean
carbon sink. Alternative explanations are an underestimate of
Northern Hemisphere fossil fuel CO2emissions or a chang-
ing land carbon sink. This implies that ongoing and expanded
Southern Hemisphere 114CO2observations and modelling
may provide a fundamental constraint on our understanding
of Southern Ocean dynamics and exchange processes.
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14782 J. C. Turnbull et al.: Sixty years of radiocarbon dioxide measurements
Data availability. The datasets presented in this paper are in-
cluded as a Supplement. The datasets (including updates as they
are available) can be accessed through the World Data Cen-
tre for Greenhouse Gases (http://ds.data.jma.go.jp/gmd/wdcgg/) or
directly through GNS Science (https://gns.cri.nz/Home/Products/
Databases/Wellington-atmospheric-14CO2-record) or NIWA (ftp:
//ftp.niwa.co.nz/tropac/).
The Supplement related to this article is available
online at https://doi.org/10.5194/acp-17-14771-2017-
supplement.
Competing interests. The authors declare that they have no conflict
of interest.
Acknowledgements. A 60-year-long record takes more than a
handful of authors to produce. This work was possible only
because of the amazing foresight and scientific understanding of
Athol Rafter and Gordon Fergusson, who began this record in the
1950s. Their work was continued over the years by a number of
people, including Hugh Melhuish, Martin Manning, Dave Lowe,
Rodger Sparks, Charlie McGill, Max Burr and Graeme Lyon.
This work was funded by the Government of New Zealand as
GNS Science Global Change Through Time core funding and
NIWA Greenhouse Gases, Emissions, and Carbon Cycle Science
Programme core funding. The authors wish to acknowledge the
contribution of New Zealand eScience Infrastructure (NeSI) to
the results of this research. New Zealand’s national computer and
analytics services and team are supported by the NeSI and funded
jointly by NeSI’s collaborator institutions and through the Ministry
of Business, Innovation and Employment (http://www.nesi.org.nz).
We thank Scott Lehman (University of Colorado) and Inge-
borg Levin (University of Heidelberg) for sharing their 114 CO2
datasets for comparison with the Wellington record.
Edited by: Neil Harris
Reviewed by: Samuel Hammer, John Miller, and
one anonymous referee
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Rapid urbanization and population growth drives increased air pollution across Sub-Saharan Africa with serious implications for human health, yet pollutant sources are poorly constrained. Here, we analyse fine particulate aerosol concentrations and radiocarbon composition of black carbon over a full annual cycle in Nairobi, Kenya. We find that particle concentrations exceed the World Health Organisation’s recommended safe limit throughout the year, with little seasonal variability in particle concentration or composition. Organics (49 ± 7%) and water-soluble inorganic ions, dominated by sulfates (13 ± 5%), constitute the largest contributors to the particle loadings. Unlike large cities on other continents, the fraction of black carbon in particles is high (15 ± 4%) suggesting black carbon is a prominent air pollutant in Nairobi. Radiocarbon-based source quantification indicates that fossil fuel combustion emissions are a dominant source of black carbon throughout the year (85 ± 3%). Taken together, this indicates that black carbon emissions from traffic are a key stressor for air quality in Nairobi.
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A basic assumption of climate change made by the United Nations Intergovernmental Panel on Climate Change (IPCC) is natural CO2 stayed constant after 1750 and human CO2 dominated the CO2 increase. IPCC's basic assumption requires human CO2 to stay in the atmosphere longer than natural CO2. But human CO2 and natural CO2 molecules are identical. So, human CO2 and natural CO2 must flow out of the atmosphere at the same rate, or e-time. The 14 CO2 e-time, derived from δ 14 C data, is 10.0 years, making the 12 CO2 e-time less than 10 years. The IPCC says the 12 CO2 e-time is about 4 years and IPCC's carbon cycle uses 3.5 years. A new physics carbon cycle model replicates IPCC's natural carbon cycle. Then, using IPCC's natural carbon cycle data, it calculates human carbon has added only 33 [24-48] ppmv to the atmosphere as of 2020, which means natural carbon has added 100 ppmv. The physics model calculates if human CO2 emissions had stopped at the end of 2020, the human CO2 level of 33 ppmv would fall to 10 ppmv in 2100. After the bomb tests, δ 14 C returned to its original balance level of zero even as 12 CO2 increased, which suggests a natural source dominates the 12 CO2 increase.
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The 14C-free fossil carbon added to atmospheric CO2 by combustion dilutes the atmospheric 14C/C ratio (Δ14C), potentially providing a means to verify fossil CO2 emissions calculated using economic inventories. However, sources of 14C from nuclear power generation and spent fuel reprocessing can counteract this dilution and may bias 14C/C-based estimates of fossil fuel-derived CO2 if these nuclear influences are not correctly accounted for. Previous studies have examined nuclear influences on local scales, but the potential for continental-scale influences on Δ14C has not yet been explored. We estimate annual 14C emissions from each nuclear site in the world and conduct an Eulerian transport modeling study to investigate the continental-scale, steady-state gradients of Δ14C caused by nuclear activities and fossil fuel combustion. Over large regions of Europe, North America and East Asia, nuclear enrichment may offset at least 20% of the fossil fuel dilution in Δ14C, corresponding to potential biases of more than −0.25 ppm in the CO2 attributed to fossil fuel emissions, larger than the bias from plant and soil respiration in some areas. Model grid cells including high 14C-release reactors or fuel reprocessing sites showed much larger nuclear enrichment, despite the coarse model resolution of 1.8°×1.8°. The recent growth of nuclear 14C emissions increased the potential nuclear bias over 1985–2005, suggesting that changing nuclear activities may complicate the use of Δ14C observations to identify trends in fossil fuel emissions. The magnitude of the potential nuclear bias is largely independent of the choice of reference station in the context of continental-scale Eulerian transport and inversion studies, but could potentially be reduced by an appropriate choice of reference station in the context of local-scale assessments.
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A regional atmospheric inversion method has been developed to determine the spatial and temporal distribution of CO2 sinks and sources across New Zealand for 2011–2013. This approach infers net air–sea and air–land CO2 fluxes from measurement records, using back-trajectory simulations from the Numerical Atmospheric dispersion Modelling Environment (NAME) Lagrangian dispersion model, driven by meteorology from the New Zealand Limited Area Model (NZLAM) weather prediction model. The inversion uses in situ measurements from two fixed sites, Baring Head on the southern tip of New Zealand's North Island (41.408° S, 174.871° E) and Lauder from the central South Island (45.038° S, 169.684° E), and ship board data from monthly cruises between Japan, New Zealand, and Australia. A range of scenarios is used to assess the sensitivity of the inversion method to underlying assumptions and to ensure robustness of the results. The results indicate a strong seasonal cycle in terrestrial land fluxes from the South Island of New Zealand, especially in western regions covered by indigenous forest, suggesting higher photosynthetic and respiratory activity than is evident in the current a priori land process model. On the annual scale, the terrestrial biosphere in New Zealand is estimated to be a net CO2 sink, removing 98 (±37) Tg CO2 yr⁻¹ from the atmosphere on average during 2011–2013. This sink is much larger than the reported 27 Tg CO2 yr⁻¹ from the national inventory for the same time period. The difference can be partially reconciled when factors related to forest and agricultural management and exports, fossil fuel emission estimates, hydrologic fluxes, and soil carbon change are considered, but some differences are likely to remain. Baseline uncertainty, model transport uncertainty, and limited sensitivity to the northern half of the North Island are the main contributors to flux uncertainty.
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The ocean is the largest sink for anthropogenic carbon dioxide (CO2), having absorbed roughly 40 per cent of CO2 emissions since the beginning of the industrial era. Recent data show that oceanic CO2 uptake rates have been growing over the past decade, reversing a trend of stagnant or declining carbon uptake during the 1990s. Here we show that ocean circulation variability is the primary driver of these changes in oceanic CO2 uptake over the past several decades. We use a global inverse model to quantify the mean ocean circulation during the 1980s, 1990s and 2000s, and then estimate the impact of decadal circulation changes on the oceanic CO2 sink using a carbon cycling model. We find that during the 1990s an enhanced upper-ocean overturning circulation drove increased outgassing of natural CO2, thus weakening the global CO2 sink. This trend reversed during the 2000s as the overturning circulation weakened. Continued weakening of the upper-ocean overturning is likely to strengthen the CO2 sink in the near future by trapping natural CO2 in the deep ocean, but ultimately may limit oceanic uptake of anthropogenic CO2.
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Two tracers, SF 6 with sources at the earth's surface and 14 CO 2 with a source in the lower stratosphere, are used to investigate the simulation of global scale transport in an atmospheric general circulation model. The simulated mixing ratios of SF 6 in the troposphere are generally in close agreement to observations revealing a realistic description of the large scale tropospheric transport. The interhemispheric exchange time for SF 6 is calculated to be 0.9 years, indicating a slightly too strong interhemispheric exchange. In the lowermost stratosphere the simulated vertical gradient of SF 6 is in good agreement with observations within the 1st 4 to 5 km above the tropopause indicating that the flux from the troposphere to the lowermost stratosphere is captured by the model. On the other hand, downward transport of 14 CO 2 from the stratosphere into the troposphere is found to be overestimated. From a comparison with observations it is concluded that it is the downward transport in the subtropics that is overestimated, at high latitudes the vertical gradients in the tropopause region are close to observations. Finally, the tracer tests show that the transport into the uppermost two levels, above 20 km, is underestimated as these levels serve as sponge layers and not as layers with a reasonable transport. Consequently, the tracer concentrations in that altitude interval are underestimated, by up to a factor 2. DOI: 10.1034/j.1600-0889.2000.00882.x