Ice core records of the major atmospheric greenhouse
gases (CO2, CH4, N2O) and their isotopologues covering recent
centuries provide evidence of biogeochemical variations during the
Late Holocene and pre-industrial periods and over the transition to the
industrial period. These records come from a number of ice core and firn air
sites and have been measured in several laboratories around the world and
show common features but also unresolved differences. Here we present
revised records, including new measurements, performed at the CSIRO Ice Core
Extraction LABoratory (ICELAB) on air samples from ice obtained at the high-accumulation site of Law Dome (East Antarctica). We are motivated by the
increasing use of the records by the scientific community and by recent
data-handling developments at CSIRO ICELAB. A number of cores and firn air
samples have been collected at Law Dome to provide high-resolution records
overlapping recent, direct atmospheric observations. The records have been
updated through a dynamic link to the calibration scales used in the Global
Atmospheric Sampling LABoratory (GASLAB) at CSIRO, which are periodically
revised with information from the latest calibration experiments. The
gas-age scales have been revised based on new ice-age scales and the
information derived from a new version of the CSIRO firn diffusion model.
Additionally, the records have been revised with new, rule-based selection
criteria and updated corrections for biases associated with the extraction
procedure and the effects of gravity and diffusion in the firn. All
measurements carried out in ICELAB–GASLAB over the last 25 years are now
managed through a database (the ICElab dataBASE or ICEBASE), which provides
consistent data management, automatic corrections and selection of
measurements, and a web-based user interface for data extraction. We present
the new records, discuss their strengths and limitations, and summarise their
main features. The records reveal changes in the carbon cycle and
atmospheric chemistry over the last 2 millennia, including the major
changes of the anthropogenic era and the smaller, mainly natural variations
beforehand. They provide the historical data to calibrate and test the next
inter-comparison of models used to predict future climate change (Coupled
Model Inter-comparison Project – phase 6, CMIP6). The datasets described in
this paper, including spline fits, are available at https://doi.org/10.25919/5bfe29ff807fb (Rubino et al., 2019).
Introduction
The three well-mixed (long-lived) atmospheric greenhouse gases (GHGs) that
contribute the most to current global warming are CO2, CH4, and
N2O. Their concentrations have been increasing since the beginning of
the industrial period, causing most of the current ∼1∘C
temperature increase above the average global temperature in
the period 1861–1880 (Stocker et al., 2013). The temperature increase limit
of 2 ∘C set by the Paris Agreement for 2100 requires substantial
reduction of GHG emissions in the next decades and, consequently,
significant reductions in the rates of GHG concentration increases.
Predicting how GHG concentrations will vary in the future requires a clear
understanding of the biogeochemical processes responsible for their
variations. However, models of future long-term climate changes predict a
large range in GHG concentrations for a given scenario of emissions
(Friedlingstein et al., 2014), and one of the key uncertainties is
associated with feedbacks in the coupled carbon–climate system (Arora et
al., 2013). Climate modellers have analysed and compared results from
state-of-the-art climate model simulations to gain insights into the
processes of climate variability, change, and feedbacks through the Coupled
Model Inter-comparison Project (CMIP). In CMIP, records of GHGs can be used
as either forcing or a diagnostic (Graven et al., 2017; Meinshausen et
al., 2017). However, real-time records of GHGs started in a period when
anthropogenic forcing was already very significant, and the atmosphere and
the Earth system were in strong disequilibrium, and therefore do not provide
a balanced state for model spin-up. Additionally, temperature and CO2
have both increased almost continuously through the 20th century, making it
difficult to separate the impacts of CO2 on carbon sinks from the
impacts of temperature increase on these sinks. Furthermore, real-time
records are often too short to draw strong conclusions on multi-decadal
variability. To provide a balanced system for model spin-up, and evaluate
the ability of models to capture observed variability on multi-decadal and
longer timescales, a branch of CMIP (“Historical Simulations”) starts
in 1850 CE (Eyring et al., 2016), while another branch (the “Paleoclimate
Modelling Intercomparison Project”, PMIP) focuses on paleo-climate
simulations (Schmidt et al., 2014). Yet, policymakers need short-term
predictions of global warming (next decades to century), and the
Intergovernmental Panel on Climate Change has very recently provided a
special report on the impacts of global warming of 1.5 ∘C above
pre-industrial levels. The last millennium is a very suitable period to
support these types of investigations since the Earth system was much closer
to its current state than previous periods of glacial–interglacial transition.
Ice cores are exceptional archives of factors influencing past climate
change because they contain a large range of substances, including water (H2O)
in the ice itself as well as ionic species, organic molecules,
and atmospheric gases sealed in bubbles (Barbante et al., 2010). They can
span from polar (Antarctica and Greenland mostly) to tropical (high
altitude) sites (Thompson et al., 2013) and extend several hundred thousand
years back in time (Higgins et al., 2015; Wolff et al., 2010). Ice cores
from different locations have different accumulation rates and temperatures,
which translate into differences in time resolution, the age of the deepest
layers, and archive suitability. Focusing on the last 2 millennia, multiple
ice core records of GHG concentration and isotopic composition are available:
CO2 from EDML (EPICA Dronning Maud Land, Antarctica) and the South Pole
(Siegenthaler et al., 2005), Law Dome, East Antarctica (Etheridge et al.;
1996; MacFarling Meure et al., 2006; Rubino et al., 2013), DML (Dronning
Maud Land; Rubino et al., 2016), and WAIS (West Antarctic Ice Sheet; Ahn et
al., 2012);
δ13C-CO2 from Law Dome (Francey et al., 1999; Rubino et
al., 2013), WAIS (Bauska et al., 2015), and DML (Rubino et al., 2016);
CH4 from NEEM (Rhodes et al., 2013) and from GISP2 (Mitchell et al.,
2013) in Greenland, Law Dome (Etheridge et al., 1998; MacFarling Meure et
al., 2006) and WAIS (Mitchell et al., 2011) in Antarctica;
N2O from EUROCORE and GRIP in Greenland (Flückiger et al., 1999),
Dome C (Flückiger et al., 2002), and Law Dome (MacFarling Meure et al.,
2006) in Antarctica.
There are other records focusing on periods other than the last centuries
but also covering the whole or part of the industrial and the pre-industrial
periods (i.e. for N2O: H15 by Machida et al., 1995; Styx glacier by Ryu
et al., 2018; Talos Dome by Schilt et al., 2010). We have decided not to
include them in our comparison because their temporal resolution (Schilt et
al., 2010) and/or coverage (Machida et al., 1995; Ryu et al., 2018) limits
their value for comparison with the records focusing on the last centuries.
There are real differences between records of the same GHG from different
sites caused by atmospheric features, such as the inter-hemispheric gradient
(north–south or Greenland vs. Antarctica). The inter-hemispheric gradient is
different from one GHG to another, depending on the balance between, and the
distribution of, sources and sinks for that specific GHG in the two
hemispheres, as well as on the atmospheric circulation and the atmospheric
lifetimes of the gases. There are also differences which do not reflect
atmospheric changes, due, for example, to the characteristics of the sites
where the ice is sampled. Ice core site characteristics influence the
measured gas records due to the gaseous diffusion through the uppermost
layers of porous, compacting snow, the “firn” (Schwander et al., 1993).
Together, diffusion in firn and gradual bubble close-off result in a
smoothed representation of the atmospheric history in ice core gas records.
The smoothing process depends on the depth of the firn layer and on how
quickly bubbles close off and trap air during firn-to-ice transition. Ice
core sites in Greenland generally have higher accumulation rates and
temperatures than in Antarctica. Consequently, GHG records from many
Antarctic sites are usually a more smoothed representation of the
atmospheric history. Unfortunately, there is no reliable CO2 record
available from Greenland because there is evidence of in situ production of
CO2 (Anklin et al., 1995; Barnola et al., 1995). The most likely
explanation for this is a high level of impurities in Greenland ice reacting
with acidity and/or hydrogen peroxide (Jenk et al., 2012; Tschumi and
Stauffer, 2000). Law Dome, Antarctica, provides the best time-resolved ice
core records due to the very high accumulation rate at this site (Etheridge
et al., 1996; Goodwin, 1990), even more so than Greenland. Also, records
from multiple Law Dome sites show no evidence of in situ production because they
agree with records from colder sites in Antarctica (Rubino et al., 2016;
Siegenthaler et al., 2005) and compare closely with each other, with air
extracted from the firn, and with modern atmospheric records (Rubino et al., 2013).
However, there are unexplained differences between records of the same GHG,
particularly for CO2. For example, while the CO2 records of the South
Pole and EDML over the last centuries are consistent with Law Dome when
their broader age smoothing is taken into account (Rubino et al., 2016;
Siegenthaler et al., 2005), the WAIS CO2 record is on average 3 ppm
higher than the Law Dome CO2 record (Ahn et al., 2012). The similarity
between the high-frequency variations in the CH4 records from Law Dome
and WAIS (Mitchell et al., 2011) suggests that the two sites (Law Dome and
WAIS) introduce similar smoothing of the atmospheric signals. However, the
Law Dome CO2 minimum measured around 1610 CE does not have a
corresponding feature at WAIS (Ahn et al., 2012). Considering that a
comparison between the two laboratories where Law Dome and WAIS samples were
measured has shown no significant offset (Ahn et al., 2012), the differences
between the WAIS and the Law Dome CO2 records could be explained by a
small effect of in situ production at WAIS. Additionally, there is a significant
difference in the mean pre-industrial level of δ13C-CO2
measured at WAIS and Law Dome (Bauska et al., 2015; Rubino et al., 2016).
These differences need to be resolved with inter-calibration campaigns
between different laboratories, using ice cores from different sites
(including new high-accumulation cores) and accurate modelling of gas age
(both the mean value and spread) at each site.
To provide the most consistent datasets possible for the past centuries, we
have previously compared the Law Dome records of CO2, CH4,
N2O, and δ13C-CO2 to firn and modern atmospheric
measurements (MacFarling Meure et al., 2006; Rubino et al., 2013). The
consistency between these measurements is evidence of our ability to extend
current atmospheric records back in time using ice and firn. However,
because of the emissions during the Industrial Revolution, our measurements
of modern and old (pre-industrial) air samples lie in different
concentration ranges and the calibrations used for measurements of modern
air samples are, therefore, in a concentration range rather different from
that used for measurements of old air samples. The measurements performed in
ICELAB–GASLAB at CSIRO have the advantage of being calibrated across the
range of concentrations of old and modern air sample measurements. Also, ice
core gas extractions and analyses are technically challenging, and different
people at CSIRO ICELAB have produced those measurements over almost 3
decades. Thus, it is possible that the extraction and analysis procedures have
introduced different biases over time, influencing the measurements by
variable amounts. However, except for minor developments over time
(Etheridge et al., 1996; Francey et al., 1999; MacFarling Meure et al.,
2006; Rubino et al., 2013, 2016) the equipment used for extraction and
analysis has not fundamentally changed.
In this study, we describe the procedure recently developed at CSIRO
ICELAB–GASLAB to perform calibration-scale updates and data selection and
correction automatically and in a consistent way for all measurements made
over the last 25 years. In the Supplement, we provide a detailed explanation
of the database recently created to store, process, and extract the
information about the samples analysed, the measurements performed, and the
results obtained. We present updated records of CO2, CH4,
N2O,
and δ13C-CO2 measured in ice and firn air from Law Dome
(Rubino et al., 2019). After merging them with other relevant records, they
will be used to run models participating at CMIP6 (Graven et al., 2017;
Meinshausen et al., 2017). We discuss the strengths and limitations of the
Law Dome GHG records and compare our records with other records from
different sites to show similarities and unresolved discrepancies. Finally,
we discuss the main features of those records, their implications for
biogeochemical, atmospheric, and climatic studies, and possible future lines of research.
MethodsLaw Dome
The ice cores used in this study, referred to as DE08, DE08-2, DSS, and
DSS0506, were drilled at Law Dome, East Antarctica (Fig. 1). Law Dome is a
relatively small (∼150 km diameter and 1390 m high) ice sheet
on the coast of Wilkes Land. It receives large and regular snowfall mainly
from the east, and the surface rarely melts in the colder central regions.
The ice flow is mainly independent of the flow of the main East Antarctic
ice sheet because of the drainage around Law Dome by two glacier systems
(the Totten and the Vanderford). Reworking of the accumulated snow is
insufficient to erase annual layers as high wind speeds are relatively
infrequent. The resulting annual layering is thick and regular and preserved
for much of the ice thickness (van Ommen et al., 2005).
Map of the Law Dome region, slightly modified from Smith et al. (2000),
showing the location of the drilling sites DE08, DE08-2, DSS, DSS0506, and DSSW20K
discussed in the text. Dotted lines are accumulation isopleths (kg m-2 yr-1)
and unbroken lines are elevation contours (m a.s.l.). The inset shows the
location of the region in Antarctica.
DE08 and DE08-2 were drilled in 1987 and 1993, respectively, only 300 m
apart and 16 km east of the summit of Law Dome (66∘44′ S,
112∘50′ E, 1390 m a.m.s.l. – above mean annual sea level), and have an
accumulation rate of approximately 1100 kg m-2 yr-1 (equivalent to
1.4 m of ice per year). DSS (Dome Summit South) was drilled between 1988
and 1993, 4.6 km south-southwest of the summit, and has an accumulation rate of
about 600 kg m-2 yr-1 (Etheridge et al., 1996; Goodwin, 1990;
van Ommen et al., 2005). In January–February 1993, air was sampled from the firn
layer at DE08-2, providing air with mean ages back to 1976 CE (Etheridge
et al., 1996). Another firn campaign was carried out at DSSW20K
(accumulation rate of approximately 150 kg m-2 yr-1), 20 km west
of DSS in December 1997 (Sturrock et al., 2002), which provided air dating
back to about 1940 CE (Trudinger et al., 2002b). DSS0506 was thermally
drilled in a dry hole (Burn-Nunes et al., 2011) during the 2005/2006 austral
summer near the Law Dome summit (66∘46′ S, 112∘48′ E,
1370 m a.s.l.). The site has an accumulation rate of about 600 kg m-2 yr-1
and a mean annual temperature of -22∘C.
ICELAB extraction
Measurement of the composition of air in ice core bubbles requires an
extraction step to release the air from ice. The dry extraction technique
used at ICELAB has been described in detail in previous publications
(Etheridge et al., 1996; MacFarling Meure et al., 2006), with recent minor
alterations to optimise extraction and measurement of
δ13C-CO2 analyses (Rubino et al., 2013). Briefly, after ice sample
selection and preparation (removing the outer 5–20 mm with a band saw),
typically 0.7–1.3 kg of ice is placed in a polyethylene bag (Layflat,
USP®) and cooled down to -80∘C in a chest freezer
for at least 24 h prior to extraction. The ice is then placed inside a
perforated inner cylinder (“cheese grater”) fixed inside an internally
electropolished stainless steel container, which is then evacuated to less
than 10-2 Pa and maintained at that pressure for at least 25 min. The
ice is then grated by mechanically shaking the container for 10 min,
which releases the trapped air. The process yields on average 70 mL (range
50–90 mL) of air, estimated from the pressure in the extraction line (whose
volume has been previously calibrated). The air is passed through a water
vapour trap (∼-100∘C) and then cryogenically
collected in an electropolished and preconditioned stainless steel trap at
around 20 K (-253∘C). The sample trap is warmed in a water bath
at room temperature (∼25∘C) for 5 min to vaporise
and mix the gases before being transported into the instrument laboratory.
Samples are analysed on gas chromatographs (GCs) for CO2, CH4, CO,
H2, and N2O within 24 h after extraction, and on the isotope ratio
mass spectrometer (IRMS) for δ13C and δ18O within 12 h.
To estimate the uncertainty and any possible bias introduced by the
extraction procedure (called the blank correction), test samples are run
together with the real ice samples. The test samples can be either reference
air samples of known composition processed with no ice present (named
“blanks”) or reference air samples injected over ice grown with no visible
bubbles in it and grated as for an actual ice core sample (the so-called
bubble-free ice, or BFI). BFI is grown in ICELAB by keeping a container
filled with deionised water in thermal equilibrium, in order to grow ice as
slowly as possible from the bottom to the top of the container. The
container features Plexiglas sidewalls that are electrically heated. The
water exchanges heat only through the metallic base and freezes from the
bottom to the top. If the process is slow enough, the produced ice is free
of visible bubbles. The results of the tests performed on ICELAB BFI, as
well as on other externally grown BFI, have been extensively described by
Rubino et al. (2013).
GASLAB analysis
Each extracted air sample is analysed for trace gas concentrations (defined
as mole fractions in parts per million (ppm) or parts per billion (ppb) in
dry air) using several GCs in GASLAB. A Series 400 CARLE/EG & G (Tulsa,
Oklahoma, USA) GC equipped with a flame ionisation detector is used to measure
CH4 and CO2 (the latter converted, after column separation, to
CH4 using a nickel catalyst at 400 ∘C). A Trace
Analytical RGA3 (Menlo Park, California, USA) GC, equipped with a mercuric
oxide reduction gas detector, is used to measure CO and H2, which
reduces HgO to gaseous Hg for detection by UV absorption. N2O is
measured on a Shimadzu GC-8AIE (Kyoto, Japan) equipped with an electron
capture detector. In normal GASLAB operation, air samples (including those
sampled from firn) in low-pressure flasks and high-pressure cylinders are
injected and analysed on the GCs using automated inlet systems to ensure
reproducibility and minimum sample consumption (Francey et al., 2003).
Because of the limited amount of air available, a semi-automated procedure
is used to inject the small volume of ice core air into the GC inlet
systems. Approximately 15–20 mL is used to measure CO2, CH4,
CO,
and H2, and 12–15 mL is used to measure N2O. The remaining air
(typically 40 mL) is used for δ13C and δ18O
measurements. The volumes indicated are total volume used for flushing gas
transfer lines as well as for analysis.
The δ13C and δ18O of the CO2 in the residual
air are measured using the MAT252 (Finnigan MAT GmbH, Bremen) IRMS located
in GASLAB. Low-pressure, large-volume whole-air samples from flasks
(atmospheric or firn air samples) and the small-volume, high-pressure, whole-air samples from ICELAB are introduced into the IRMS through a common inlet
(multiport) equipped with an all stainless steel mass flow controller
(Brooks 5850) to ensure constant mass flow conditions for all samples. The
IRMS uses two cryogenic traps (MT Box C, Finnigan) to retain water vapour
and other condensable gases and to extract CO2 (plus N2O) from
air. The residual CO2 (and N2O) is injected into the MAT252 ion
source via a dedicated micro-volume and crimped capillary. Nitrous oxide
(N2O) has identical molecular masses to CO2 and interferes with
the isotopic analyses. To remove this interference, a correction is made to
the IRMS output in GASLAB using the relative ionisation efficiency of
N2O and CO2, the isotopic composition of N2O, and the measured
N2O and CO2 concentrations, as described in detail by Allison and
Francey (2007). High-precision isotopic ratios are determined by alternating
sample CO2 and reference CO2 injected via matched crimped
capillaries. The carbon isotopic ratio of the sample (sa) is expressed
relative to the reference (ref) following Eq. (1):
δ13C=13C/12Csa13C/12Cref-1×1000.
When comparing measurements performed more than 20 years apart, rigorous
traceability in the propagation of calibration scales becomes an important
factor. This is obtained with a long-term, continuous comparison of standard
cylinders for both GC (Francey et al., 2003) and IRMS (Allison and Francey,
2007) analyses. The current calibration scales used are WMO X2004A for
CH4, WMO X2007 for CO2, NOAA 2006A for N2O, and CSIRO2005 for
CO2 isotopes (CO2-in-air scale, which is linked to the
VPDB CO2 scale).
ICELAB database
A new database allows storage, selection, correction, updating and
extraction of the data produced in ICELAB–GASLAB. It allows results to be
dynamically updated if changes in analytical methods or calibration scales
are implemented, keeping ice, firn, and atmospheric measurements consistent
with each other. Data are stored in tables where the information associated
with each specific sample is linked via a Universal Analysis Number (UAN)
that acts as the index for combining all information. The structure of the
database and its tables are described in detail in the Supplement. The
database includes procedures, which automatically perform sample selection,
correct results and estimate uncertainty, and provide a routine for data
extraction (see Supplement for details).
Results and discussionsDevelopment of the Law Dome GHG records, internal consistency, unexplained discrepancies, and limitations
The first pre-industrial record of CO2 from Law Dome covering the whole
last millennium was published by Etheridge et al. (1996, reported here with
red squares in Figs. 2a and 3a). It was one of the first ice core records to
show the overlap with firn (Fig. 3a) and contemporary atmospheric
measurements. The overlap of ice core and contemporary atmospheric
measurements is one of the main advantages of the Law Dome ice core sites,
due to their high snow accumulation rates and the resultant relatively quick
bubble close-off time and recently enclosed air. This feature has been
described extensively in previous papers (Etheridge et al., 1996; MacFarling Meure
et al., 2006; Rubino et al., 2013) and used, together with the overlap
between different cores, to demonstrate our confidence in extending
contemporary GHG concentration measurements back in time. Based on replicate
analyses of test samples (blanks and BFI) over time periods of several
months, on ice samples within an annual layer, and the overlap mentioned
above, Etheridge et al. (1996) estimated that the uncertainty of the
CO2 measurements was 1.2 ppm (1σ). The major biogeochemical
events discussed in Etheridge et al. (1996) were the LIA (Little Ice Age)
CO2 decline between 1550 and 1750 CE with the subsequent recovery from
the LIA perturbation between 1750 and 1800 CE in the pre-industrial period,
and the 1940s stabilisation of atmospheric CO2 concentration
(Fig. 3a),
which ended just before the Mauna Loa and South Pole atmospheric records began.
Published pre-industrial period (1–1900 CE) GHG records from Law Dome
ice extracted and measured at CSIRO ICELAB–GASLAB.
(a)CO2,
(b)CH4, (c)δ13C-CO2, and
(d)N2O.
Published and unpublished industrial period (1750–2000) GHG records
from Law Dome ice and firn, extracted and measured at CSIRO ICELAB–GASLAB.
(a)CO2, (b)CH4,
(c)δ13C-CO2,
and (d)N2O.
A few years later, the same authors published the Law Dome pre-industrial
record of CH4 covering the last millennium (Etheridge et al., 1998, red
squares in Figs. 2b and 3b). The tight overlap for the first time between
ice, firn and contemporary atmospheric CH4 measurements that began more
than 20 years later than for CO2 confirmed that the ice core air
record is a faithful representation of the past atmospheric CH4
concentration (Fig. 3b). The estimated uncertainty was 5 ppb. The major
features discussed in Etheridge et al. (1998) were the LIA CH4 decline,
supporting a terrestrial origin for the synchronous CH4/CO2
decrease, and the rapid increase in CH4 growth rates after
1945 CE, which peaked in 1981 CE, just as atmospheric monitoring began. It was also
possible to determine the pre-industrial inter-hemispheric difference in
CH4 (24–58±10 ppb), based on comparison with CH4
measurements from Greenland (EUROCORE and GISP2), also made in
ICELAB–GASLAB, and supporting evidence from Blunier et al. (1993) and
Chappellaz et al. (1997). The variability over time of the CH4
pre-industrial inter-hemispheric gradient provides a constraint to quantify
variations in sources and sinks of CH4 (Mitchell et al., 2013). The
same is not possible for CO2 because of the above-mentioned in situ production
in Greenland ice.
To quantify variations in the sources and sinks of CO2, Francey et
al. (1999) measured its isotopic ratio (δ13C, red squares in
Figs. 2c and 3c) in Law Dome ice. This record provided a means to quantify the
relative CO2 uptake by the land and the ocean to the total atmospheric
CO2 change (Trudinger et al., 2002a), when the emissions from fossil
fuel and land use change were taken into account for the industrial period,
and assuming that, in the pre-industrial period, there was no significant
influence of anthropogenic activities on the atmospheric δ13C-CO2
(Pongratz and Caldeira, 2012; Stocker et al., 2011), and
also assuming no significant changes in inter-hemispheric CO2 exchange
times (Francey and Frederiksen, 2016; Frederiksen and Francey, 2018). The
δ13C-CO2 decrease measured by Francey et al. (1999) in the
last 2 centuries (Fig. 3c) is mainly due to 13C-depleted CO2
derived from fossil fuel CO2 emissions and is important evidence of
the prominent role of anthropogenic emissions in the industrial period
CO2 increase. Francey et al. (1999) also discussed the increase in
δ13C-CO2 during the LIA, supporting the interpretation of
a terrestrial origin for the synchronous CH4/CO2 decrease
(Trudinger et al., 1999), though with lower sampling resolution compared to
the CO2 in Etheridge et al. (1996). Francey et al. (1999) estimated
statistical and systematic δ13C-CO2 biases between 0.025
and 0.07 ‰ and uncertainties of up to ±0.05 ‰ but
found an unexplained discrepancy of up to 0.2 ‰ (Trudinger, 2000, Sect. 3.8) around 1900 CE from the
South Pole δ13C-CO2 firn record measured at NOAA INSTAAR
(National Oceanic and Atmospheric Administration Institute of Arctic and
Alpine Research, Boulder, Colorado).
The early Law Dome GHG records have been revised and extended over time as follows.
MacFarling Meure et al. (2006) extended the CO2 and CH4 records
back through the last 2 millennia (green diamonds in Fig. 2a and b) and
increased sample density in the industrial period (green diamonds in Fig. 3a
and b). They also confirmed the LIA CO2/CH4 decrease as well as
the 1940s CO2 stabilisation and produced a record of N2O (green
diamonds in Fig. 2d) which, in turn, overlaps with firn N2O
measurements (green diamonds in Fig. 3d). The authors interpreted the
N2O decrease of about 5 ppb during the LIA as additional evidence for
the terrestrial origin of the LIA GHG decrease. The measurement uncertainty
remained the same for CO2 (1.2 ppm) as for Etheridge et al. (1996) but
decreased slightly for CH4 (from 5 ppb in Etheridge et al., 1998, to
4 ppb in MacFarling Meure et al., 2006). The uncertainty of the N2O
measurements was 6.5 ppb. The authors also found an increase in N2O
concentration of about 10 ppb between 675 and 800 CE, which does not seem to
be related to any known climatic event.
Rubino et al. (2013) revised the δ13C-CO2 record (see
yellow triangles in Figs. 2c and 3c) by updating the calibration scale and
revisiting the corrections applied in Francey et al. (1999) for blank,
gravity, and diffusion effects, using the revised CSIRO firn model (Trudinger
et al., 2013) for the gravity and diffusion corrections (Trudinger et al.,
1997). In doing so, they resolved the 0.2 ‰ discrepancy
found between the Law Dome δ13C-CO2 record and the South
Pole δ13C-CO2 firn record (the South Pole firn records
have been reported in Fig. 5, but see Rubino et al., 2013, for more details).
They also increased sample density during the industrial period and applied
the new chronology available for Law Dome ice (Plummer et al., 2012), which
caused a shift of about 150 years for samples that are 2000 years old (see
difference between the ages of green diamonds and yellow triangles in Fig. 2a).
The age difference becomes negligible in the last millennium, as
evident by comparing red squares and yellow triangles in Fig. 2c.
Rubino et al. (2016) carried out additional CO2 and
δ13C-CO2 measurements (see blue circles in Figs. 2–3a and c) from
ice cores sampled at the Law Dome site of DSS0506 (Pedro et al., 2011). The
data extended back to 1700 CE effective air age and provided additional
evidence of consistent results between different ice cores and firn records
where they overlapped. However, the increasing CO2 trend measured in
DSS0506 between 1700 and 1850 CE does not tightly match that previously
attributed to recovery from the LIA (Etheridge et al., 1996).
It is also worth mentioning the results of two other studies performed using
Law Dome ice and firn, which were sampled but not measured using CSIRO GASLAB instruments.
To investigate changes in pre-industrial sources of CH4, Ferretti et
al. (2005) produced a record of δ13C-CH4 in Law Dome ice
covering the last 2000 years (not shown). They reported unexpected changes
of the global CH4 budget, mainly attributed to variations in biomass
burning emissions during the late pre-industrial Holocene (LPIH) through an
atmospheric box model (Lassey et al., 2000). The δ13C-CH4
record from Ferretti et al. (2005) has not been included in ICEBASE because
the air samples extracted in ICELAB were measured on a mass spectrometer not
maintained by CSIRO GASLAB. Therefore, the δ13C-CH4 data
are not on a CSIRO calibration scale and have not been included in ICEBASE.
Park et al. (2012) measured oxygen and intramolecular nitrogen isotopic
compositions of N2O (not shown) covering 1940 to 2005 in Law Dome firn
air and archived air samples from Cape Grim (Tasmania). In doing so, they
confirmed that the rise in atmospheric N2O levels is largely the result
of an increased reliance on nitrogen-based fertilisers. These isotopic
measurements are also not included in ICEBASE.
Revised records (100–1900 CE) of
(a)δ13C-CO2,
(b)CO2, (c)CH4, and (d)N2O
from Law Dome ice compared to published records from other sites:
WAIS δ13C-CO2 from Bauska et al. (2015), WAIS CO2
from Ahn et al. (2012), DML CO2 and δ13C-CO2
from Rubino et al. (2016), EDML and South Pole CO2 from Siegenthaler
et al. (2005), WAIS CH4 from Mitchell et al. (2011), NEEM CH4 from
Rhodes et al. (2013) plotted using the revised age scale as explained in the text, GISP2 CH4 from Mitchell et al. (2013),
EUROCORE and GRIP N2O from Flückiger et al. (1999), and EDC N2O from
Flückiger et al. (2002).
The new Law Dome GHG records and comparison with records from other sites
Figures 4 and 5 show the newly revised Law Dome GHG records (red circles).
Following the rule-based selection described in the Supplement,
there are 299 ice core measurements for CO2, 307 for CH4,
147 for N2O (compared to 212, 228, and 103, respectively, in MacFarling Meure
et al., 2006), and 86 for δ13C-CO2 (compared to 58 in
Francey et al., 1999, and 69 in Rubino et al., 2013). All of the major
features described in previous publications are retained. However, the
differences mentioned above can potentially influence the biogeochemical and
climatic interpretation of these records. Given that the Law Dome GHG
records are a major source of information for models used to predict the
future behaviour of the Earth system (Graven et al., 2017; Köhler et
al., 2017b; Meinshausen et al., 2017), in the following paragraphs we
provide an explanation of the main reasons for these differences.
Changes to the calibration scale result in small, mostly negligible, differences.
All records (except the δ13C-CO2) start in 154 CE
(effective age for CO2) rather than 0 CE. This causes a revision of air
ages towards more recent times for all events recorded in the second-to-last
millennium (e.g. the 10 ppb increase in N2O between 675 and 800 CE
discussed in MacFarling Meure et al. (2006) is now dated 701–822 CE,
Fig. 4d) but a less than 2-year change in dating after about 1000 CE.
Each data point has an uncertainty, which is independently calculated
based on the weighting and flagging systems described in the Supplementary
Material. The uncertainty does not include any additional uncertainty
associated with inter-core variability. For example, based on comparisons
between samples of the same ages, the discrepancy found between DSS0506 and
other Law Dome cores in the period 1700–1850 CE suggests that the inter-core
variability can potentially add a random, extra uncertainty of up to 5 ppm for CO2.
Further research is needed to precisely quantify the inter-core variability.
Revised records (1750–2010 CE) of
(a)δ13C-CO2,
(b)CO2, (c)CH4, and (d)N2O
from Law Dome ice and firn compared to the South Pole firn records of δ13C-CO2,
CO2, and N2O (measured at NOAA-INSTAAR) and to published records from other sites:
WAIS δ13C-CO2 from Bauska et al. (2015), WAIS CO2
from Ahn et al. (2012), DML CO2 and δ13C-CO2
from Rubino et al. (2016), EDML and South Pole CO2 from Siegenthaler
et al. (2005), WAIS CH4 from Mitchell et al. (2011), NEEM CH4 from
Rhodes et al. (2013), GISP2 CH4 from Mitchell et al. (2013), and
EUROCORE N2O from Flückiger et al. (1999).
The following list compares the new Law Dome records with records from other
sites and discusses the main differences.
There is good agreement between the revised CO2/δ13C-CO2
Law Dome records and the CO2/δ13C-CO2 records from DML produced in ICELAB–GASLAB (blue
triangles in Figs. 4–5a and b). Once the different gas-age distributions of
the ice cores are taken into account, the two records are in very good
agreement, with a difference of less than 2 ppm for CO2 and differences
within error bars for δ13C-CO2 (Rubino et al., 2016).
Given that both records have been produced at CSIRO ICELAB–GASLAB, we can
exclude any possible inter-laboratory offset.
There is also acceptable agreement (random differences up to 4 ppm)
between the CO2 Law Dome record and the CO2 records from EDML and
the South Pole (Siegenthaler et al., 2005, white and green squares in Fig. 4b).
Considering that the records have been produced in different laboratories
(CSIRO ICELAB–GASLAB and University of Bern), it is possible that part of
the difference is explained by an inter-laboratory offset.
There is a systematic difference (3 ppm on average) between the Law Dome
CO2 record and the WAIS CO2 record in the pre-industrial (Ahn et
al., 2012, see grey squares in Fig. 4b). Though small, the difference is of
concern because it is systematic throughout the whole record. The two
laboratories (CSIRO ICELAB–GASLAB and Oregon State University) that produced
these records have also run a comparison experiment to quantify the
contribution of a possible inter-laboratory offset to the total discrepancy
(Ahn et al., 2012). The good agreement (measurements from the two
laboratories were consistent within the 1σ analytical uncertainty)
found by the inter-comparison experiment suggests that the discrepancy is
mostly due to an inter-core difference.
There is an increase in this difference to >8 ppm between the
Law Dome CO2 dip around 1610 CE and the WAIS CO2 decrease during
the LIA (Ahn et al., 2012, compare grey squares and red circles in Fig. 4b).
WAIS is considered a high-accumulation site and should retain the same
events as those recorded at Law Dome. The difference is even more surprising
when the tight agreement between the Law Dome CH4 record and the WAIS
CH4 record (Mitchell et al., 2011) around this time is considered
(compare red circles and grey squares in Fig. 4c). The consistency between
the Law Dome and the WAIS CH4 record rules out dating issues or large
differences in smoothing of the atmospheric signals between the two sites.
This suggests a chemical origin (in situ production) of the CO2
discrepancy, which is more likely to occur for CO2 than for CH4.
This interpretation is supported by additional evidence of 6 ppm discrepancy
(Köhler et al., 2017b) during the Last Glacial Maximum, last termination,
and Early Holocene between the EDC CO2 record (Monnin et al., 2001,
2004) and the WAIS CO2 record (Marcott et al., 2014).
There is a difference of up to 0.15 ‰ (compare grey
squares and red circles in Fig. 4a) between the Law Dome
δ13C-CO2 record (Rubino et al., 2013) and the WAIS
δ13C-CO2 record (Bauska et al., 2015). This difference is most
likely due to an inter-laboratory offset but may also indicate a
contribution from the CO2 discrepancy to its δ13C, or a
combination of the two. It is not possible to quantify the inter-laboratory
offset without running an inter-comparison study. However, it is possible to
calculate a maximum effect of the in situ CO2 production on the
δ13C measured at WAIS, assuming that the 3 ppm extra CO2 measured
in WAIS (compared to an average pre-industrial CO2 concentration of 280 ppm measured at Law Dome) is totally due to the in situ production and that it all
comes from carbonate carbon (δ13C= 0 ‰)
because organic carbon with δ13C=-27 ‰ would
make the WAIS δ13C-CO2 more negative than the Law
Dome δ13C-CO2. Under this assumption, we calculate a
possible shift of 0.07 ‰ through an isotope mass balance
(=[0-(-6.55)]⋅3/283). This can only explain part of the discrepancy but
can go up to 0.14 ‰ if an extra amount of 6 ppm is
assumed (Ahn et al., 2012; Köhler et al., 2017b). The Bauska et al. (2015)
record agrees within uncertainties with the Francey et al. (1999)
dataset. However, Rubino et al. (2013) is the only record to show
consistency with all firn records and direct atmospheric measurements (see
Figs. 3c and 5a). This would suggest that the Rubino et al. (2013) study is
currently the most accurate record and should be used to set a
pre-industrial baseline. However, no definite conclusion can be drawn until
a thorough inter-comparison study is carried out between the labs that have
produced the WAIS and the Law Dome δ13C-CO2 datasets
(Oregon State University-University of Colorado-Institute of Arctic and
Alpine Research, INSTAAR and CSIRO). It is important to resolve the
difference between the Law Dome and the WAIS δ13C-CO2
records in order to establish a pre-industrial baseline and, thus, a
pre-industrial-to-industrial δ13C-CO2 difference. Setting
a pre-industrial baseline could have consequences on the Last Glacial
Maximum-to-pre-industrial δ13C-CO2 difference as well
(Schmitt et al., 2012). These values are useful for biogeochemical
interpretation (Broecker and McGee, 2013; Krakauer et al., 2006).
As expected, the Law Dome and WAIS CH4 concentrations are lower than the
NEEM high-resolution CH4 record (Rhodes et al., 2013, the white
triangles
in Fig. 4c show the median CH4 concentrations for 5-year time slices,
after data outliers have been removed) by an amount which is consistent with
an inter-hemispheric CH4 difference of 40–60 ppb (Mitchell et al.,
2013). Interestingly, the LIA CH4 decrease measured at NEEM appears to
start before the CH4 decrease measured at Law Dome and WAIS. The age scale
of the NEEM CH4 record published in Rhodes et al. (2013) (Fig. 4c) has been
revised with the updated ice-age scale published in Sigl et al. (2015) and
the new estimate of Δage provided by Buizert et al. (2014).
Mitchell et al. (2013) have synchronised the GISP2 CH4 record
with the WAIS CH4 record to investigate changes of the interpolar
difference in the pre-industrial based on the reasoning that “the
multidecadal events observed in both ice core records must have occurred
simultaneously since the durations of the events were much larger than the
atmospheric mixing time (∼1 year)” (Mitchell et al., 2013).
The NEEM CH4 record has not been synchronised with the others, and
there are multiple possible reasons, including age scale issues, different
smoothing of the atmospheric signals at the different sites, and inadequate
sampling resolution, to explain the discrepancy found between the NEEM and
the GISP2, Law Dome, and WAIS CH4 records during the LIA. A thorough
investigation is out of the scope of this paper, but, in the future, this
discrepancy should be resolved to obtain a precise synchronisation of all
ice core records available over the LIA.
The N2O records from Greenland (Flückiger et al., 1999, EUROCORE
and GRIP, grey and white squares in Fig. 4d, respectively) and from EDC
(Flückiger et al., 2002, green squares in Fig. 4d) show higher scatter
than the Law Dome N2O record. All records need higher sampling
resolution to investigate changes of atmospheric N2O concentration over
the last centuries with more confidence. A new N2O record from a
high-resolution site is required to explore the real variations in N2O
in the pre-industrial period (Ryu et al., 2018).
The LIA and the 1610 CE CO2 minimum in DSS (Law Dome)
In an attempt to produce δ13C data around the Law Dome 1610 CE
CO2 minimum (sparsely sampled by Francey et al., 1999) and confirm
the interpretation of its terrestrial origin (Rubino et al., 2016; Trudinger
et al., 1999), in 2012 we measured 18 samples from DSS, the only core from Law
Dome covering the whole LPIH. The results both for CO2 and for
δ13C are significantly lower than the spline fit to the revised records
from Law Dome (results not shown). Conversely, the CH4
concentration measured is very consistent with the spline fit to the revised
CH4 record. In the past, abnormally low CO2 concentration was
interpreted as a sign of post coring melting (PCM), since CO2 is much more
soluble than CH4. With PCM, the N2O concentration is usually low
as well. However, in 7 of the 18 DSS samples that provided enough air to
measure N2O, the N2O concentration was, on average, not
significantly lower than the spline fit to the revised Law Dome record. This
argues against the hypothesis of PCM. Another possibility is the effect of
clathrate formation, which could alter CO2 and δ13C, but
issues due to clathrate can be ruled out because none of the ice cores in
this study reached depths or temperatures sufficient for clathrate formation,
which can affect the extraction and measurement of enclosed gases and
13C-CO2 in particular (e.g. Schaefer et al., 2011). We do not have
a definite explanation for the low CO2 and δ13C measured,
but the results suggest that the DSS core may no longer be a reliable
archive for CO2. This experiment was conducted while we were
measuring the DSS0506 CO2 samples published in Rubino et al. (2016).
During that survey, we also found a similar behaviour for 6 of the
34 DSS0506 samples measured. Since the two cores – DSS0506 and DSS – were
stored in the same freezer in Hobart (Tasmania, Australia), with DSS0506
stored for a shorter period, we suggest the low CO2 may be a recent
effect of storage (see Supplement to Rubino et al., 2016).
The LIA, and particularly the 1610 CE CO2 event, is important for our
understanding of the carbon cycle dynamics and the carbon–climate system in
the past. It has been used to derive the CO2 sensitivity to temperature
(Cox and Jones, 2008; Rubino et al., 2016), it is the most prominent
biogeochemical event during the LPIH, and it has even been suggested as the
beginning of the new geologic era called the Anthropocene (Lewis and Maslin,
2015). Therefore, it is of fundamental importance that we understand the
amplitude of the minimum as recorded by the ice and the timing and likely
size of the original atmospheric decrease before smoothing during firn
diffusion and enclosure into bubbles.
Biogeochemical and climatic interpretation of the Law Dome GHG records
The Law Dome GHG records have been used for biogeochemical and climatic
interpretation of changes in CO2 (Joos et al., 1999; Joos and Bruno,
1998; Rubino et al., 2013; Trudinger et al., 2002a), CH4 (Ferretti et
al., 2005; Ghosh et al., 2015), and N2O (Park et al., 2012) over the
past decades to millennia. They are also used as reference atmospheric GHG
records for model simulations of the carbon–climate system of the LPIH
(Graven et al., 2017; Köhler et al., 2017b; Meinshausen et al., 2017).
Here we present an overview of the insight obtained through interpretation
of the Law Dome GHG records and provide some perspective on the challenges
we will have to face to obtain a deeper understanding of the carbon–climate
system during the LPIH and the industrial period.
The biogeochemical interpretation of GHG variations depends on quantifying
their sources and sinks. The concentration of GHGs in the atmosphere is the
net result of the processes releasing GHGs to the atmosphere (sources) and
processes removing GHGs from the atmosphere (sinks). The atmospheric
circulation then redistributes GHGs assuming consistency in the reasonably
well-known patterns of air movement. Measuring the atmospheric concentration
of GHGs provides one constraint on the net sources vs. sinks. However,
generally, multiple sources and sinks act simultaneously. Therefore,
multiple constraints are necessary to partition the contribution of each
source or sink. Measuring the isotopic composition of each GHG provides an
additional constraint (albeit usually with additional complexity), but there
are also other independent constraints, such as the inter-hemispheric
difference or evidence coming from other species, that help quantify the
contribution of sources and sinks. Additionally, the net GHG emission
strength is reflected in the rate of change in atmospheric concentration.
Thus, ice core records that track rapid changes (i.e. high resolution) are
best to infer budgets and hence biogeochemical information before direct
atmospheric observations. The Law Dome records provide the highest
resolution among existing Antarctic ice cores. There have been recent
studies looking into the effects of firn microstructure (including density
layers) on bubble trapping (Burr et al., 2018; Fourteau et al., 2017;
Gregory et al., 2014; Mitchell et al., 2015). Improved understanding of
these processes, how they affect smoothing of atmospheric GHG signals, and
their incorporation into numerical models may lead to a more accurate
quantification of the relationship between ice core GHG measurements at
different sites and the original atmospheric variations.
The two major reservoirs of CO2 that can change atmospheric
concentrations over years to millennia are the terrestrial biosphere (land)
and the oceans. Fossil fuel and land use emissions have added to these over
recent centuries. Measurements of δ13C-CO2 have been used
to quantify the contribution of land and ocean to the atmospheric CO2
variations measured (Joos et al., 1999; Joos and Bruno, 1998; Trudinger et
al., 2002a). For example, the interpretation of CO2 and
δ13C-CO2 variations through a double deconvolution (Fig. 6a) has
identified the terrestrial biosphere as the main contributor to the LIA
CO2 decline (Rubino et al., 2013, 2016; Trudinger et al., 2002a). This
agrees with the findings of Bauska et al. (2015), who used the WAIS CO2
and δ13C-CO2 records to suggest that changes in
terrestrial organic carbon stores best explain the observed multi-decadal
variations in the δ13C-CO2 and in CO2 concentrations
from 755 to 1850 CE. This agreement of interpretation from the DSS and WAIS
records shows that what matters for the biogeochemical interpretation is the
change in concentration over time, rather than the absolute concentration
measured in different ice cores. The above studies assume consistency in the
inter-hemispheric transport of Northern Hemisphere terrestrial emissions
over multiple decades. Preliminary examination suggests Southern Hemisphere
δ13C-CO2 records over the last decades are more
susceptible to multi-year changes in the ratio of eddy to mean advective
inter-hemispheric transport (Francey and Frederiksen, 2016; Frederiksen and
Francey, 2018) than is the case for CO2 concentration, as a result of
differences in isotopic equilibration that occur for the two transport modes.
Biogeochemical and climatic interpretation of the Law Dome GHG records:
(a) atmospheric CO2 fluxes from (negative values on the
y axis) and to (positive values on the y axis) the terrestrial biosphere (land:
green line) and the ocean (blue line), resulting from the double deconvolution
of CO2 and δ13C-CO2 (Rubino et al., 2016).
(b) Flux of atmospheric CH4 from biomass burning (Ferretti
et al., 2005). (c) Temperature variations in different continents in
the Northern Hemisphere (Asia: yellow line; Europe: blue line; North America:
red line) from Pages2k (2013). (d) Palmer Drought Severity Index (PSDI)
of different continents in the Northern Hemisphere (Asia: yellow line; Europe:
blue line; North America: red line) from Cook et al. (2009, 2010, 2015).
(e) Atmospheric CO2 fluxes from different continents in the
Northern Hemisphere (Asia: yellow line; Europe: blue line; North America: red
line) due to pre-industrial land use change from Pongratz et al. (2012).
An additional constraint for the biogeochemical interpretation of the LIA
CO2 decrease has recently come from the new interpretation (Rubino et
al., 2016) of the record of carbonyl sulfide (COS; Aydin et al., 2008) from
the SPRESSO ice core (South Pole Remote Earth Science and Seismological
Observatory) and modelling of its increase over the LIA (Rubino et al.,
2016). Rubino et al. (2016) showed that the simultaneous COS increase during
the LIA confirms that the LIA CO2 decline was caused by net terrestrial
uptake due to cooling (heterotrophic respiration declining more than gross
primary production, due to its higher sensitivity to temperature changes),
though a very recent paper estimating the amount of carbon taken up by land
use change following the colonisation of the Americas by the Europeans (Koch
et al., 2019) provides a different view. Nonetheless, the multi-species
approach used by Rubino et al. (2016, e.g. using trends of CO2,
δ13C-CO2, and COS) can provide multiple constraints to help
understand the biogeochemical processes behind atmospheric CO2
variations over the recent past.
The 1940s CO2 plateau is a prominent feature in the industrial part of
the Law Dome CO2 record and occurs at a time of continued fossil fuel
emissions. Taking into consideration the smoothing effects of firn diffusion
and bubble trapping on the rate of change in potential atmospheric signals,
an uptake of around 2–3 GtC yr-1 between 1942 and 1949 would be required to explain
the observed plateau. The Law Dome δ13C-CO2 measurements
suggest that the oceans were responsible for at least two-thirds of this
uptake (Rubino et al., 2013; Trudinger, 2000, Sect. 6.4; Trudinger et al.,
2002a). Bastos et al. (2016) used the latest estimates of fossil fuel and
land-use change emissions, ocean uptake reconstructions, and terrestrial
models but were not able to explain the plateau, although they did not
consider decadal variability in the ocean carbon sink that may have been
important. Better understanding of the 1940s feature is needed to quantify
how variable ocean and land carbon exchange can be on multi-year to decadal
timescales and how this variability relates to climate variability.
Improved understanding is expected to come from more high-precision,
high-time-resolution ice core measurements to confirm
δ13C-CO2 variation through the industrial period, additional
constraints such as COS, better understanding of the smoothing effects on
trapped air, and improved modelling of other influences on atmospheric
δ13C-CO2 such as the effect of climate on isotopic
discrimination (Randerson et al., 2002; Scholze et al., 2003).
There are multiple natural sources–sinks of CH4: geological, wetlands,
wildfires, termites, and ocean sediments are the main sources, while
oxidation by tropospheric species (OH), oxidation by stratospheric species
(OH, Cl, and O(1-D)), oxidation in soils and reactive chlorine in the marine
boundary layer are the main sinks. Also for CH4, measurements of its
isotopic composition (δ13C-CH4 and δD-CH4)
have helped quantify the contribution of land vs. ocean to the measured
atmospheric CH4 variations (Ferretti et al., 2005; Mischler et al.,
2009; Sapart et al., 2012). However, because there are more distinct
CH4 source types than isotopic tracers, and more spatially distributed
sources than can be resolved by the geographically restricted suite of ice
cores (despite the fact that the ice cores of Antarctica and Greenland are
both known to provide reliable CH4 records), a unique solution for the
history of CH4 sources is not possible. Nonetheless, the full suite of
isotopic tracers and bipolar ice core data provides important boundary
conditions for testing hypothetical CH4 source–sink histories, allowing
elimination of large classes of scenarios. By measuring
δ13C-CH4 in Law Dome ice, Ferretti et al. (2005) provided evidence
of a remarkable decrease in pyrogenic CH4 during the last millennium
(Fig. 6b). This interpretation was confirmed by Mischler et al. (2009) and
Sapart et al. (2012), who measured both δ13C-CH4 and
δD-CH4 in ice cores from Antarctica and Greenland, respectively.
They also found an increasing agricultural source of CH4 throughout the
last millennium, with most of the change occurring between the 1500s and the 1600s,
supporting the hypothesis of a pre-industrial anthropogenic influence on
atmospheric CH4. Additional source information is provided by
measurements of carbon monoxide (CO) and its isotopes (δ13C-CO
and δ18O-CO; Wang et al., 2010), through evidence of variations
in biomass burning, and 14CH4, which identifies CH4 emissions
from fossil sources. However, the measurement of 14CH4 is limited
so far to large air samples extracted from firn (Lassey et al., 2007a, b)
or from large ice samples collected where glacial-age ice outcrops at
the surface (Petrenko et al., 2009, 2017).
There are also multiple natural sources and sinks of N2O,
both on land and in the ocean. The main sources are microbiological
processes, especially in tropical soils, while the main sinks are
photochemical reactions in the atmosphere. To the best of our knowledge,
there are only two attempts to use δ15N-N2O and
δ18O-N2O records to better constrain the land vs. ocean sources of
N2O over the last century (Park et al., 2012; Prokopiou et al., 2017),
and these used the larger volumes of air available in firn. Only one study (Prokopiou et al., 2018) has
extended the investigation to the last millennia, but there is an analysis
covering the last deglaciation by Schilt et al. (2014). As already mentioned,
the pre-industrial inter-hemispheric N2O difference is also poorly
constrained. Thus, there is room for vast improvement to understand the N
cycle from measurements of N2O concentration and its isotopes in ice
cores (Schilt et al., 2014). However, there is a risk of in situ production of
N2O, especially in Greenland ice (Flückiger et al., 2002). More
records, with high sampling resolution, from both hemispheres are needed to
confirm the features found in Law Dome (MacFarling Meure et al., 2006) and
understand the causes of the changes in N2O concentration over time
(Ryu et al., 2018).
The climatic interpretation of the Law Dome GHG records has generally been
carried out by comparing the timing of GHG variations and temperature
changes and testing hypotheses of mechanistic relationships between the two
with coupled carbon cycle–climate models.
An important climatic event of the LPIH is the
Medieval Climate Anomaly (MCA, roughly 950–1250 CE), which showed higher
temperature in some regions (Goosse et al., 2005; Mann et al., 2009). During
the MCA, there are generally higher levels of GHGs, but the timings of
increase vary from one gas to another, with N2O showing a rise between
701 and 822 CE (well before the start of the MCA), CO2 increasing
between 950 and 1200 CE, and CH4 showing some variability superimposed
on a long-term increasing trend (Fig. 4b–d).
The main climatic event of the last 2 millennia is the LIA
(roughly 1400–1700 CE; Mann et al., 2008; Neukom et al., 2014; Pages2k,
2013). There is a clear decrease in all GHGs during the LIA, which, together
with the other evidence mentioned above, suggests that all processes
releasing GHG from land slowed down during the cold phase (Fig. 4b–d).
However, as for the MCA, the change in concentration is not simultaneous for
all GHGs. While both CO2 and N2O seem to decrease starting around
1550 CE (but N2O would need higher sampling resolution to confirm
this), CH4 has a later decline, starting around 1580 CE. Also, the
CH4 decrease seems to last for a shorter period of time, ending around
1610 CE, whereas the CO2 low is maintained longer, ending about
1750 CE. At the same time, there was a significant decrease in biomass burning
(Ferretti et al., 2005; Mischler et al., 2009; Sapart et al., 2012; Wang et
al., 2010), interpreted to be mostly a consequence of decreased fire
emissions. While the relationship between CO2 and temperature variation
has been used to infer the sensitivity of CO2 to temperature (Cox and
Jones, 2008; Rubino et al., 2016), confirming and quantifying a positive
feedback of terrestrial carbon with temperature (Rubino et al., 2016), it is
now time to investigate the regional contribution to the total CO2
change, as attempted by Bauska et al. (2015), and the contribution from
different processes within the terrestrial biosphere, such as net primary
production, heterotrophic respiration, and biomass burning. There are
regional (Mann et al., 2008), continental (Pages2k, 2013), and hemispheric
(Neukom et al., 2014) temperature reconstructions that can be used to drive
models describing the relationship between climate and carbon cycle to
quantify the contribution from each region to the total CO2 decrease
(Fig. 6c). There are also hydro-climatic reconstructions (Cook et al., 2009,
2010, 2015) providing evidence of dry and wet periods over the LPIH
(Fig. 6d), which can be used together with records of charcoal (Marlon et al.,
2013, 2016) and biomass burning tracers in ice cores (Grieman et al., 2017)
to quantify the contribution of declining biomass burning in each region to
the total CO2 decrease. The emissions from anthropogenic land use
change have also been quantified for each world region (Pongratz and
Caldeira, 2012; see Fig. 6e) and can be used to subtract the human
contribution from the total CO2 change, even though there is a debate
on the amount of land use change following the European colonisation of the
Americas (Koch et al., 2019). We suggest that the LIA provides a suitable
epoch to further study carbon cycle–climate feedbacks, for predictions of
the future carbon–climate system, and particularly to understand the role of
different regions of the world in changes of atmospheric chemistry and
biogeochemical fluxes and carbon pools. The consequences of the LIA climatic
changes on societal development are important for understanding why
different communities were more or less vulnerable, resilient, or even
adaptive (Degroot, 2018). Being able to quantify which regions have been
more vulnerable to past climate change, also in terms of the response of the
natural carbon cycle, could help plan future adaptation strategies.
Data availability
Data connected with this paper are available in the CSIRO
Data Access Portal (https://doi.org/10.25919/5bfe29ff807fb; Rubino et al.,
2019). Each species has been ordered by core and by age. In detail, for each
species and each record the following fields are available (Rubino et al., 2019):
sample ID,
ice age,
gas age,
value,
uncertainty.
For each species, the calculated spline fit with time steps of 1 year and
growth rate are also given. The spline fits attenuate variations with
periods of less than 20 years by 50 % for CO2 and CH4, 100 years
for N2O, and 50 years for δ13CO2. When using these
data please consider citing the original publications from which the data
underlying this compilation have been taken.
Conclusions
The records of GHG (CO2, CH4, N2O) concentrations and the
isotopic composition of CO2 (δ13C-CO2) from the Law
Dome ice cores are among the most important sources of information for
models trying to predict the future behaviour of biogeochemical cycles and
their influence on the climate system. The records of CO2, CH4,
N2O, and δ13C-CO2 are constantly being updated with
new measurements and revised for changes in calibration scales and
corrections for the effects of laboratory extractions and those of gravity
and diffusion in firn. This paper has provided an in-depth explanation of
the procedures used to extract and measure the samples in CSIRO ICELAB and
store, correct, and select the results obtained. Given the widespread use of
the datasets produced at CSIRO ICELAB–GASLAB, it is important to provide a
track record of the reasons for changes carried out over time and keep the
records open to the scientific community. This should help with their use in
modelling and avoid misinterpretation (for example, the mistake spotted by
Köhler et al., 2017a, in the paper from Machado and Froehner, 2017).
The GHG records from Law Dome have already provided significant insight into
biogeochemical and climatic feature over the last centuries. However, there
is room for deeper understanding, for example studying the influence of
different regions on variations in atmospheric chemistry. Also, there are
unresolved discrepancies (e.g. the LIA CO2 decrease) which need to be
resolved and Law Dome appears to provide the most suitable ice cores for
high-resolution investigation of atmospheric changes over the last
centuries. Considering that
all GHG records in ice cores are a smoothed representation of the real
atmospheric history,
DSS is the highest accumulation rate site ever sampled in Antarctica
recording the LIA CO2 event,
there is the risk that the WAIS core is affected by in situ production of
CO2, and
accurate CO2 records have not been derived from Greenland ice
cores,
we suggest that there is a need to sample a new, clean, and deep ice core
from Law Dome, to confirm or improve our knowledge of the atmospheric LIA
CO2 decrease and other rapid changes in atmospheric composition during
pre-industrial millennia.
Finally, there are open questions about the real size of past atmospheric
variations in some species, such as N2O and COS, and the reasons for
those variations, which, once resolved with new measurements, could provide
additional understanding and insights, useful to better predict the future
behaviour of biogeochemical cycles and their influence on the climate system.
The supplement related to this article is available online at: https://doi.org/10.5194/essd-11-473-2019-supplement.
Author contributions
MR conceived the database structure, performed measurements
between 2009 and 2013, and led the writing of the paper. DME supervised the process,
performed measurements, and collected samples. RH wrote the database queries
and procedures in the Microsoft SQL server and developed the web-based user
interface with Microsoft Visual Studio. DPT, CEA, RJF, RLL, LPS, and DAS carried
out measurements of samples and standards in GASLAB and performed calibration
scale updates. CMT carried out the numerical modelling, including the CSIRO firn
model, the double deconvolution, and the spline fit to the data. MAJC, TDVO, and
AMS collected, stored, and distributed samples and contributed to ice
dating.
Competing interests
The authors declare that they have no conflict of interest.
Acknowledgements
We acknowledge the long-term support provided to this work by CSIRO, the
Department of Energy and Environment, the Australian Antarctic Division and
Australian Antarctic Program, and the many people involved in collecting the
several ice cores and firn air samples from Law Dome since the 1980s. We
thank Rachael Rhodes and Michael Sigl for providing updated ice and gas-age
scales for NEEM. We also thank the two anonymous referees for their
comments, which have helped improve the original paper.
In particular, we acknowledge the enormous contribution of chief driller and
close friend Alan Elcheikh (1960–2019), who sadly passed away during the
production of this paper.
Review statement
This paper was edited by Thomas Blunier and reviewed by two anonymous referees.
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