Spatial differences in the monthly baseline CO2 since 1992
from Mauna Loa (mlo, 19.5∘ N, 155.6∘ W, 3379 m), Cape
Grim (cgo, 40.7∘ S, 144.7∘ E, 94 m), and South Pole (spo,
90∘ S, 2810 m) are examined for consistency between four
monitoring networks. For each site pair, a composite based on the average of
NOAA, CSIRO, and two independent Scripps Institution of Oceanography (SIO) analysis methods is presented. Averages
of the monthly standard deviations are 0.25, 0.23, and 0.16 ppm for mlo–cgo,
mlo–spo, and cgo–spo respectively. This high degree of consistency and
near-monthly temporal differentiation (compared to CO2 growth rates)
provide an opportunity to use the composite differences for verification of
global carbon cycle model simulations.
Interhemispheric CO2 variation is predominantly imparted by the mlo
data. The peaks and dips of the seasonal variation in interhemispheric
difference act largely independently. The peaks mainly occur in May, near
the peak of Northern Hemisphere (NH) terrestrial photosynthesis/respiration
cycle. February–April is when interhemispheric exchange via eddy processes
dominates, with increasing contributions from mean transport via the Hadley
circulation into boreal summer (May–July). The dips occur in September, when
the CO2 partial pressure difference is near zero. The cross-equatorial
flux variation is large and sufficient to significantly influence short-term
Northern Hemisphere growth rate variations. However, surface–air terrestrial
flux anomalies would need to be up to an order of magnitude larger than
found to explain the peak and dip CO2 difference variations.
Features throughout the composite CO2 difference records are
inconsistent in timing and amplitude with air–surface fluxes but are largely
consistent with interhemispheric transport variations. These include greater
variability prior to 2010 compared to the remarkable stability in annual
CO2 interhemispheric difference in the 5-year relatively El
Niño-quiet period 2010–2014 (despite a strong La Niña in 2011), and
the 2017 recovery in the CO2 interhemispheric gradient from the
unprecedented El Niño event in 2015–2016.
Introduction
Atmospheric CO2 measurements are normally introduced into global carbon
budgets as a “global growth rate … based on the average of multiple
stations selected from the marine boundary layer sites with well-mixed
background air …, after fitting each station with a smoothed curve as a
function of time, and averaging by latitude band …” (Le
Quéré et al., 2018). This approach encourages sampling at multiple
locations to seek atmospheric confirmation of national/continental emission
changes. Particularly in the Northern Hemisphere (NH), with more complicated
geography and atmospheric circulation, the influence of continental
emissions on marine boundary layer air can vary widely between sites.
A clearer indication of the global impact of regional emissions comes from
sites demonstrating maximum spatial representation. In this case, global
significance of biogeochemical CO2 exchanges between the surface will
be informed by their impact on validated baseline data with the least
continental influence. Such baseline data are more directly relevant to
changes in global ocean acidification and climate change, but places
heightened demands on sampling criteria and calibration.
Sites selected to maximize spatial representation in their respective
hemispheres, Mauna Loa (mlo, 19.5∘ N, 155.6∘ W, 3379 m)
and South Pole (spo, 90∘ S, 2810 m), also have the longest-term
(multi-decadal) coherent trace gas monitoring data, based on flask sampling
(Supplement S1). At these sites, and at Cape Grim (cgo, 40.7∘ S,
144.7∘ E, 94 m) since 1991, co-sampled baseline air has been
analysed at three different laboratories, using four different methodologies
summarized in Sect. 2.
To account for any persisting artefacts in the co-sampled data, we examine,
for each method, inter-site differences in the published monthly baseline
data from the three sites. The standard deviation in the average of the
co-sampled differences provides a practical uncertainty estimate. A key
advantage compared to the growth rate approach is that assumptions inherent
in the growth rate smoothing (where for example 22-month smoothing is used
to separate interannual and seasonal variations) are avoided so that in this
study near-monthly effective time resolution is achieved.
The inter-site difference approach was used by Francey and Frederiksen (2016; FF16) to conclude that the suppression of the normal eddy component of
interhemispheric (IH) CO2 exchange in February–April 2010 contributed to the
unprecedented 0.8 ppm step in the IH difference between 2009 and 2010. The
dynamical anomaly was associated with a moderately strong El Niño
leading to a NH build-up of CO2 in 2010. FF16 supplementary
information demonstrated a failure of atmospheric transport models of the
carbon cycle to simulate the step.
The 2015–2016 El Niño was stronger and has also been associated with
unprecedented behaviour in the global carbon cycle (elsewhere attributed to
the terrestrial biosphere anomalies, e.g. Yue et al., 2017). However,
Frederiksen and Francey (2018; FF18), argued that the unprecedented strength
in the Hadley circulation increased IH exchange (reduced IH CO2
difference) late in 2016, overwhelming the earlier reduced eddy exchange
linked to the strong 2015–2016 El Niño. They also indicated dynamical
contributions to IH CO2 during both El Niño and La Niña periods
(e.g. FF16 Fig. 5, and FF18 Sect. 6.2, on multi-species IH differences).
While El Niño–Southern Oscillation (ENSO) events are expected to impact on surface biology, it is also
clear that they also influence atmospheric IH CO2 fluxes. The timing of
the dynamical events suggests an alternate explanation for the CO2
behaviour discussed by Yue et al., if it can be demonstrated that IH
CO2 fluxes at the time exceed their postulated air–surface terrestrial
fluxes.
The scope of this paper includes (a) reduction of measurement uncertainties
in IH CO2 difference using a three decade composite of published
CO2 measurement results (distinguished by maximum spatial
representation and by well-documented sampling and measurement quality), and
(b) demonstration of the potential uses of the composite CO2 record by
comparing anomalies in the magnitude and phasing of composite IH CO2
variations with those in air–surface exchange model outputs, as well as in
dynamics indices representing atmospheric IH exchange.
Background information on flask networks
A historic overview of CO2 IH difference data is provided in Supplement S1.
By 1958 Charles David Keeling had identified mlo and spo as optimum sites to obtain
background CO2 in the respective hemispheres and by the 1970s was
obtaining a regular monthly supply of air admitted to 5 L evacuated glass
flasks from both sites (SIO1: Keeling et al., 2001). Since 1992, there
have been CO2 measurements as a by-product of a global network focussed on
O2/N2 ratios in baseline air (SIO2: Keeling and Schertz,
1992); this program uses 5 L glass flasks flushed and filled to ambient
pressure, with cryogenically dried air. While there is commonality regarding
calibration, in the context of spatial differences the Scripps Institution of Oceanography (SIO) networks can be
considered independent.
NOAA began sampling from all three sites, mlo, spo, and cgo (as part of a much
larger network), from 1984, using a variety of flask and filling methods.
From around 1992 the current system of Peltier-dried air in pressurized 2.5 L flasks (Tans et al., 1992; Conway et al., 1994; Dlugokencky et al., 2014)
was phased in. NOAA has maintained the World Meteorological Organization
(WMO) Central CO2 Calibration Laboratory since 1996 (a role previously
carried out by SIO). The NOAA atmospheric sampling is generally more
frequent (typically 8–10 flasks per month) than is the case for the SIO or
CSIRO programs (except for the CSIRO cgo program); however, the size and
sampling frequency in the NOAA network amplifies calibration challenges due
to shorter lifetimes of reference and calibration standards.
Both NOAA and SIO use non-dispersive infrared analysers (NDIR) for CO2
measurement (CSIRO flask sampling at cgo, spo, and mlo in the early 1980s
used NDIR for analysis of chemically dried air, pressurized into 5 L glass
flasks). However, analyses here are restricted to CSIRO's measurements from
1992 using chemically dried, pressurized air in 0.5 L glass flasks, but with
retention of 5 L flasks at spo (Francey et al., 1996). Gas chromatography
with flame ionization detection (GC/FID) was introduced to measure CO2
in flasks, a technique providing a more linear response than NDIR
(Supplement S2). Hourly radon measurements at Cape Grim (Chambers et al.
2016) were introduced around this time. Air mass history is further informed
by a decade of vertical profiling (Langenfelds et al., 2003; Pak et al.,
1996), back trajectory analysis, and other tracers (e.g. Dunse et al.,
2001), demonstrating that selected cgo data can achieve a degree of spatial
representation matching, or sometime exceeding, that at the more remote
high-altitude sites at mlo and spo.
Apart from longevity, the flask records offer other advantages over in situ
monitoring, but are more susceptible to some unfavourable factors as are
discussed in Supplement S3. The challenges of maintaining high quality over
decades in any one monitoring program are many. They include external
factors, acknowledged but not pursued here, such as high turnover of skilled
staff particularly at remote air sampling sites or changes in institutional
strategic and economic priorities. The latter are well described by Keeling (1998), with CSIRO sharing similar institutional experiences.
Monthly CO2 differences (in ppm) from NOAA for sites mlo,
cgo, and spo and networks SIO1, SIO2, and CSIRO. Five-month running means are
highlighted. 1996–2016 mean and standard deviation (bracketed) values are
included.
Network intercomparison
NOAA, which has the most extensive global network (and since 1996 has also
operated the WMO CO2 Central Calibration Laboratory) is selected as the
reference for an initial inter-network comparison. For each of the three
baseline sites, Fig. 1 shows systematic behaviour in the SIO1, SIO2, and
CSIRO monthly CO2 differences from NOAA. Five-month running means aid
discussion. The CO2 mixing ratios used here are referred to in the
commonly used units of parts per million (ppm) rather than the more strictly
correct term of μ mole of CO2 per mole of dry air. Note that data
independently flagged for sampling or measurement anomalies are rejected by
individual laboratories prior to publication as monthly averages. Typically,
a small number of gross outliers in individual flask data (e.g. in
flask-pair differences) are also rejected prior to publication.
In Fig. 1, there is clear evidence of systematic differences in mean
offsets, seasonality, and between sites within one network. In the context
of interhemispheric exchange, the typical 0.5 ppm range of variation remains
relatively small compared to the 7–10 ppm maximum CO2 interhemispheric
difference (IH ΔCO2). Net IH exchange is proportional to IH
partial pressure difference.
Between 1991 and 1993, there is a marked inconsistency between NOAA mlo–spo and
mlo–cgo, particularly in seasonal amplitude; CSIRO has comparable
measurements that are more consistent (Supplement S4). This is a reason for
caution when interpreting the data in this period.
In the post-1996 statistics the SIO1 offsets from NOAA behave similarly for
mlo and spo. This is not the case for SIO2, which has similar offsets at mlo
(-0.18 ppm) and cgo (-0.19 ppm), but not at spo (-0.04 ppm), or for CSIRO
(mlo: -0.08, cgo: +0.01, spo: +0.13 ppm).
CSIRO records at cgo exhibit the smallest offset and scatter relative to
NOAA (±0.08 ppm) while SIO2 mlo data exhibit the largest scatter
(±0.37 ppm).
Remnant seasonality is still evident in the CSIRO cgo differences from NOAA.
While a small effect, the CSIRO GC/FID near-linear response for CO2
means results are not so sensitive to differences between sample and
reference CO2. This advantage is reinforced in the CSIRO SH data since
reference gases use recent SH baseline air. This is generally not the case
for non-linear NDIR measurement and particularly in the NH if relatively
short-lived reference gases sourced in the NH have a less-than-optimum match
with ambient CO2 from a site.
Three-month averaged CO2 differences between sites for each
of the four sampling networks: for years spanning 1990 to 2020.
While Fig. 1 reveals some un-resolved systematic differences between data
sets, Fig. 2 emphasizes that they are generally small compared to the IH
partial pressure differences that are a pre-requisite for IH net exchange.
Data from each method are presented as 3-month seasonal averages in order to
minimize potential influences related to network sample frequency (by
ensuring an adequate number of individual flask samples per period). In addition, the particular 3-month seasonal selection distinguishes periods of
distinct relatively stable partial pressure differences between hemispheres
and the selected seasons also distinguish eddy and mean IH transport
mechanisms (FF18).
Figure 2 demonstrates the considerable coherence between data sets.
For the most part, and particularly in the August–October season when IH CO2
difference (IH ΔCO2) is at a minimum, there is a high level of
consistency in the year-to-year variation in seasonal spatial differences
from each network.
There are relatively few examples of one record differing markedly from the
others; when it occurs, it is often for reasons evident in Fig. 1. For
example, in Fig. 2 NOAA cgo–spo appears low in 1992–1993; CSIRO mlo–spo
shows negative outliers in May–July 2009 and November–January 2002, but not for
mlo–cgo. SIO2 outliers in 2002 and 2006 exhibit similar characteristics;
positive outliers, e.g. SIO2 from November to January 2016, suggest a cgo problem. In
February–April 2005 NOAA data indicate a possible mlo problem; however, this is
also when the “volatility” of the records (and in IH transport) is large, so
it is conceivable that different flask sampling numbers and times could
contribute to lower values by both SIO and CSIRO.
The largest IH ΔCO2 variability is recorded in February–April and in
May–July, both seasons having near-equally large IH differences. The large
seasonality in NH CO2 is widely linked to the photosynthesis/respiration in NH forests. February–April is also when IH exchange by eddy
processes is most influential (FF16), whereas mean transport via the Hadley
circulation is the main dynamical influence in May–July (FF18).
Systematic differences due to sampling and measurement methodology can
possibly arise from factors such as the linearity of instrument response,
flask storage effects or undetected entrainment of laboratory air. Records
with the sparsest sample density (e.g. at spo and particularly in CSIRO spo
data) may be more susceptible to undetected anomalies. Closer inspection of
individual flask metadata, or of the less extensive in situ monitoring, may resolve
some of these infrequent anomalies, but for the present, composite averaging
of the flask data is relied on to moderate their influence.
Composite records of baseline station spatial differences
For each of mlo–cgo, mlo–spo, and cgo–spo monthly CO2 differences, Table 1 shows the number of months between 1992 and 2017 contributing to a
composite value, arranged in columns indicating the number of contributing
networks; e.g. 266 of 31 months have four networks contributing to mlo–spo,
while 279 months have three contributing networks at mlo–cgo.
The percentage of missing months for each network, and scatter in the
composite differences for different historic periods, are tabulated in
Supplement S1.
Number of months of data available for composite differences at the
baseline sites.
The monthly composite CO2 differences are shown in Fig. 3 (and
tabulated in Supplement S5). The small error bars represent the ensemble
standard deviation (the one exception is for cgo–spo in February 2009, with only
the NOAA network contributing. It is arbitrarily assigned 100 %
uncertainty and appears as an outlier in Fig. 3c). The seasonality at mlo,
generally attributed to the NH forest photosynthesis/respiration cycle, is
the dominant variation in IH ΔCO2. The composite uncertainties
are small compared to seasonal amplitudes, especially for the IH
differences. Average standard deviations of mlo–cgo, mlo–spo, and cgo–spo are
0.25, 0.23, and 0.16 ppm respectively. Systematic year-to-year variability is
well defined and is reflected similarly in both IH records and is consistent
with mlo driving most of the seasonal variation.
Variations that exceed the ensemble monthly standard deviations include the
following.
The overall increase in IH difference, generally attributed mainly to
increasing NH fossil fuel CO2 emissions, is indicated by a linear
regression through the mlo–cgo values (with slope 0.056±0.021 ppm yr-1; mlo–spo gives 0.062±0.021 ppm yr-1). The slope of
such regressions is much higher for the April–May data (0.087±0.011 ppm yr-1) than for September–October data (0.049±0.011 ppm yr-1).
From 1992 to 2017, most minima occur in September; of 26 minima, 24 occur in September and
2 in October (1992 and 1995). Of the 26 maxima, 20 are in May, and 6 in April
(1997, 1999, 2000, 2004, 2005, and 2016).
Scatter in the amplitude of seasonal maxima (boreal winter/spring) is
smaller before 1999. The step-like behaviour in April–May from 2009 to 2010
remains the major anomaly.
In contrast, the minima (in boreal summer/autumn) exhibit greater scatter
before 2011, replaced afterwards by a smooth decline to a marked 2016
minimum, then sudden reset in 2017.
Unusually low boreal summer/autumn IH minima also occur in 1993–1994. Apart
from being a period when measurement and calibration methods were
consolidating (as discussed in the next section) the most significant volcanic
influence (Pinatubo) is potentially an influence at this time.
A question arises as to how well mlo data represents the NH. Of more
relevance to this study is how well the mlo samples represent air that is
transferred into the Southern Hemisphere. Flask samples are collected at mlo
above 3 km altitude in downslope winds, close to the upper troposphere
regions where the IH transfer processes defined in FF18 occur (see Fig. 5
below), circumstances not shared by other NH surface monitoring sites.
Unlike in typical growth rate analyses, the peak and dip values are largely
independent. This is visually explored in Fig. 3 using plotting software,
with spline polylines linking peaks (solid) and dips (dashed) months of IH
ΔCO2. Trace gas mixing within extratropical (ET) hemispheres is
typically estimated at 1–2 months or less, and interhemispheric exchange
times are estimated at 6–12 months or more (e.g. Bowman and Cowan, 1997; Jacob,
1999). Monthly changes in the peak and dip IH ΔCO2 largely
reflect flux changes in or out of the extratropical northern troposphere
close to that month. The following sections seek similarities with possible
causal forcing processes.
Processes influencing CO2 IH difference variations
Global carbon cycle models generally attribute short-term variations in
atmospheric CO2 to exchanges with the terrestrial biosphere (Le
Quéré et al., 2018; Rödenbeck et al., 2018; Yue et al., 2017)
and implicitly assume model atmospheric transport is correct on all time
frames. While the models have demonstrated an impressive ability to predict
mid-to-high-latitude CO2 variations influenced by weather, it is less
clear that short-term variations in IH exchange (of a magnitude sufficient
to influence hemispheric growth rates) have been adequately captured.
Air–surface fluxes influencing IH ΔCO2
The relative magnitude and timing of monthly variations of IH ΔCO2
are compared to those in the terrestrial biosphere, wildfires, and fossil
fuel (possible contributions from air–sea exchange are discussed below in
relation to Fig. 8.)
Composite station difference data showing the network ensemble
average and standard deviation of monthly CO2 for (a) mlo–spo, (b) mlo–cgo, and (c) cgo–spo (on a doubly expanded scale). Linear regressions
through the IH records are black-dotted lines. Spline polylines visually
link peaks (blue, solid) and dips (red, dotted) of the seasonal IH
differences. Grey-shaded and rounded panels indicate El Niño periods with
strongly anomalous equatorial zonal winds.
The primary determinant of the well-defined seasonality in IH ΔCO2 in Fig. 3 is widely attributed to the temperature-moderated
photosynthesis/respiration cycle of NH forests. Monthly dynamic vegetation
model (DVM) estimates of terrestrial net biosphere production (NBP) in three
latitude bands 90 to 30∘ N, 30∘ N to 30∘ S, and 30 to 90∘ S, over the 1992–2016
period are obtained using the Community Atmosphere Biosphere Land Exchange
(CABLE) model (Kowalczyk et al., 2006; Haverd et al., 2018). In addition,
ET NBP from an ensemble of 16 land surface models (shown in
Fig. 2 of Bastos et al., 2018) are considered. Because of the small SH
contribution, the ensemble ET values are most comparable to CABLE NH NBP.
Note: We do not discuss air–surface fluxes derived from CO2 data that
are less spatially representative and/or rely on atmospheric transport
modelling. The latter introduce additional model degrees of freedom and
potentially overestimate terrestrial variability if the variability in
atmospheric IH transport is not adequately captured.
NBP signs are reversed and are described as terrestrial-to-air carbon
fluxes. Global wildfire emissions from the Global Fire Emissions Database
(Randerson et al., 2018, GFED4.1) from 1997 to 2015 are classified as NH, EQ
and EQ/SH. Seasonal anthropogenic emission anomalies are calculated as
differences from the detrended 2000 to 2016 monthly data of Oda et al. (2018). For each data set, anomalies (in PgC month-1) in seasonal
behaviour for each latitude band were determined by subtracting the mean
seasonality from the monthly values.
Comparison of the timing and amplitude of terrestrial emission
anomalies (i.e. mean seasonality subtracted) with variations of the peaks
and dips in Fig. 3. (a) shows seasonal anomalies in CABLE emissions (dark
green) and in 16-DGVM TRENDY ET emissions (Bastos et al., 2018;
light green) and (b) shows GFED4.1 wildfire seasonal anomalies, for NH
(green), EQ (pink), and SH (blue, SH/EQ for GFED4.1). In (c) the largest
anomalies (CABLE NH, CABLE EQ, and GFED4.1 EQ) on the left axis are compared
to the ppm variation in peaks (red) and dips (blue) on separate right axes.
The axes scaling equates 1 PgC with 1 ppm (see text). To highlight seasonal
differences, (d) shows the standard deviation in the seasonal
anomalies for each month, including those in anthropogenic emissions (Oda et
al., 2018).
The major seasonal anomalies in NBP and wildfire emissions that potentially
influence IH ΔCO2 are shown in Fig. 4a and b. The
largest anomalous surface-to-air flux is the extreme equatorial emission
anomaly from equatorial wildfire in late 1997 (∼0.9 PgC over
3 months); it is not associated with unusual behaviour in the IH ΔCO2 records.
Despite mixing of CO2 within the ET Northern Hemisphere
being as rapid as 1–2 weeks (Jacob, 1999) compared to IH exchange times of
greater than 6 months (Bowman and Cowan, 1997), we see strong correlations
with transport for unlagged 3-month averages. And since IH ΔCO2
peaks re-occur within 1 month of the same time each year, close
correspondence in timing of terrestrial anomalies and the IH ΔCO2 peaks would be expected if NH terrestrial exchange was the main
determinant. This is not evident in Fig. 4. More importantly, the
amplitude range of terrestrial anomalies appears to be far too small to
account for the magnitude of the changes in the peaks and dips of IH ΔCO2.
Over the last 25 years the annual relationship between global (mainly NH)
fossil fuel combustion emissions and IH ΔCO2 has been 2.8 PgC ppm-1 (equivalent to the 0.36 ppm (PgC)-1 used by FF18). This is
applicable when northern fossil fuel emissions effectively mix globally. The
volume of the troposphere north of Mauna Loa is around 33 % of the global
troposphere, so that on the shorter time frame of within-hemisphere mixing,
only ∼0.92 PgC is required to change the NH background
CO2 by 1 ppm. In Fig. 4c we round this to 1 PgC =1 ppm for
simplicity.
The variability in the air–surface fluxes, relative to that in IH ΔCO2, is displayed in Fig. 4d, which plots the standard deviations
of residuals from the mean seasonality, for each month over the available
record. As for the peaks and dips, we assume a 1:1 relationship between ppm
and PgC month-1 in IH ΔCO2. The main variation in IH
ΔCO2 occurs in March–April, when variability in surface–air fluxes
is small but variability in eddy IH exchange is large (see below). A second
peak in IH ΔCO2 standard deviation occurs August–September, around the
time of the dips (but also when equatorial wildfires are more active
suggesting a possible contribution from the equatorial emissions at this
time).
Accepting the precision and near-hemispheric spatial representation of the
composite IH ΔCO2 records, these inconsistences with surface
emissions in both timing and magnitude suggest that there are other
short-term influences on IH ΔCO2 of greater magnitude than
air–surface exchange.
Wind indices reflecting CO2 IH transport
In contrast to the case for air–surface exchanges, there are a number of
prominent features in the composite IH ΔCO2 records that are
shared with behaviour in the dynamical indices of FF18. Interhemispheric
exchange of CO2 occurs mainly by eddy processes in the boreal
winter–spring and by mean convection and advection associated with the
Hadley circulation in the boreal summer–autumn (FF18 and references
therein). FF18 developed wind indices that characterize both types of IH
transport based on reanalysis data sets focussing on the National Center for
Environmental Prediction (NCEP) and National Center for Atmospheric Research
(NCAR) reanalysis (NRR) data (Kalnay et al., 1996). Eddy transport is
described by uduct, the average 300 hPa zonal velocity in the Pacific
westerly duct region (Frederiksen and Webster, 1988) of 5∘ N to
5∘ S, 140 to 170∘ W (FF16, FF18). Here we use that index
and two of the four indices for mean transport introduced in FF18. These are
ωP, the average 300 hPa vertical velocity in pressure
coordinates in the region 10 to 15∘ N, 120 to 240∘ E, and
vP the average 200 hPa meridional velocity in the region
5 to 10∘ N, 120 to 240∘ E. Figure 5 provides a schematic of
the geographical location of regions used by FF18, and time series of the
monthly values of wind indices are shown in Fig. 6.
Schematic of the boundaries and altitudes of regions used in FF18
to define wind indices that describe eddy IH transfer (uduct, westerlies positive) and mean transfer (uplift, negative ωP)
and north-to-south transfer (negative vP). The shaded area brackets the
austral summer extent of the Intertropical Convergence Zone in the south
(blue dash) and boreal summer extent in the north (red dash).
Monthly values of (a)uduct, (b)ωP, and (c)vP. The shading to zero indicates months of enhanced transport which
act to reduce the IH ΔCO2. Anomalous dynamical periods are
highlighted with grey-shaded rectangles.
The top panel in Fig. 6a shows a 3-decade time series of the uduct
index which characterizes cross-equatorial Rossby wave dispersion, Rossby
wave breaking, and corresponding increases in transient kinetic energy and
eddy transport in the near-equatorial upper troposphere (Webster and Holton,
1982, Frederiksen and Webster, 1988, Ortega et al., 2018). The large-scale
Rossby waves are generated by thermal anomalies and topographic features
including the Himalayan mountains from which they propagate south-eastward
and are able to penetrate into the SH when uduct is positive,
corresponding to an open Pacific westerly duct.
The ωP and vP indices in Fig. 6b and c describe the
strength of the mean transport by the Hadley cell in the Pacific region with
negative ωP corresponding to uplift and negative vP to
north-to-south transport.
Net interhemispheric trace gas exchange requires a partial pressure
difference between hemispheres. For CO2 the average seasonal cycle of
a 25-year mean partial pressure difference, represented here by monthly
baseline mlo–cgo, is shown in Fig. 7a (mlo–spo is not shown here since,
reflecting on data quality, it is effectively identical).
The monthly averages of dynamical factors governing CO2 IH
exchange over the last 25 years. (a) Detrended CO2 partial pressure
differences mlo–cgo (green), (b) Pacific eddy transport index uduct (dark blue), (c) Pacific Hadley transport indicated by uplift at
10–15∘ N (-ωP, light red), and (d) north-to-south
transport (-vP, dark red). On average, coincidence of shading in wind
indices and shaded months of IH ΔCO2 is a precondition for
increased IH mixing (reduced IH gradient). The more anomalous transport
years, 1998 (dots), 2010 (dashes), and 2016 (black line) are shown for each
wind index, and for mlo–cgo IH ΔCO2.
The positive mean IH ΔCO2 is largely due to fossil fuel
emissions. Months of positive (north–south) IH difference are shaded green
and only in September–October is there a small reverse gradient. Transport of CO2
from the Northern to the Southern Hemisphere occurs when green-shaded areas
in Fig. 7a coincide (on average) with blue-shaded areas (Fig. 7b, via
eddy transfer with index uduct) or with red-shaded areas (Fig. 7c
and d, via mean transport with indices ωP and vP).
Figure 7 also demonstrates that differences from the long-term mean in
transport indices (average for each month) vary between the significant El
Niño events in 1998, 2010, and 2016:
In 2010, the IH ΔCO2 exceeds the average between February and July (Fig. 7a) with reduced eddy transfer between February and April, associated with lower that
average uduct (Fig. 7b). Further, between June and September, there is weaker
ascent (Fig. 7c) and north-to-south upper tropospheric wind (Fig. 7d) in
the key regions defining ωP and vP. As noted in FF18,
the IH ΔCO2 eddy and mean transports reinforce to contribute to
the unprecedented 2009–2010 step in IH ΔCO2.
In 2016, the IH ΔCO2 is larger than average between February and June and
smaller than average between July and October (Fig. 7a). These results are again
consistent with the behaviour of the dynamical indices. There is reduced IH
ΔCO2 eddy transfer in the first half of the year (Fig. 7b)
but very strong mean transport in the second half of the year (Fig. 7c
and d) that accounts for the annual IH ΔCO2, as noted in FF18.
In 1998, the IH ΔCO2 exceeds the average from May to December and is
close to the mean annual cycle for the rest of the year. We note from Fig. 7 that the annual increase in IH ΔCO2, also shown in Fig. 2
of FF18, is largely induced by the June–August mean Hadley circulation.
It is suggestive that the relative variation in IH ΔCO2 February–May for the three big El Niño years matches that in uduct,
however it is puzzling that the largest uduct anomaly, in 1998, is when IH
ΔCO2 is closest to the mean behaviour. The fact that the mean
transport indices at this time of year are also consistently well below
their long-term average is also of note, since with uduct close to zero
and -ωP, -vP indicating descent and south-to-north
meridional winds, there is no obvious mechanism for IH exchange in this
season. Yet, over the 25 years, correlation of the April–May IH ΔCO2 peaks with -ωP-vP is significant, r≈
0.4. One possible explanation for these behaviours in the early part of the
boreal winter/austral summer may be found in changes in the volume of the
well-mixed portion of the Northern Hemisphere (see Discussion, Sect. 7).
Different responses of IH ΔCO2 to wind indices at different
ENSO events, and from non-ENSO periods, are discussed in Sect. 7.
As an aside, we also include a similar plot for the average SH cgo–spo
differences in Fig. 8. Despite some concerns about artefacts in spo data
(e.g. due to long flask-air storage times), all networks indicate that on
average spo baseline CO2 exceeds that at cgo in the austral summer
months. The minimum cgo–spo appears to precede inversion estimates of
Southern Ocean CO2 uptake south of 30∘ S (Lenton et al.,
2013). High-precision continuous CO2 monitoring across the Southern
Ocean (Stavert et al., 2019; Ann Stavert, personal communication, 2018) confirm small and
relatively smooth seasonal variation. The earlier November–December minimum
in the
CO2 difference coincides with a seasonal dip in fossil fuel emissions
(Oda et al., 2018) perhaps indicating an alternative explanation.
Composite 25-year average of monthly baseline cgo–spo CO2
(dark blue). Individual network values are shown in orange (NOAA), dark blue
(SIO2), and CSIRO (black). Estimates of sea–air CO2 flux seasonality are
shown in light blue.
Year-to-year variation in the composite records
The annual net impacts of the various potential influences on site IH
ΔCO2 (when typical terrestrial biosphere seasonal variations
are balanced) appear in Fig. 9. Uncertainty in annual values obtained from
combining composite standard deviations of normalized monthly values for the
cgo case, averaging ±0.08 ppm, compares to ∼±0.3 ppm variation in the detrended annual record.
Annual changes in the baseline CO2 difference between sites.
Interhemispheric differences mlo–cgo (orange, with dashed linear regression)
and mlo–spo (light blue) are plotted on the left axis. The peak magnitudes
of strong El Niños (brown, ONI index > 1) and strong La
Niñas (purple, ONI index < -1) are indicated. The cgo–spo annual
differences are plotted on a doubled right-hand scale. Annual fossil fuel
emissions from FF18, are shown on the top right axis.
Working through Fig. 9 from the left in order to highlight other
systematic features:
Except for 2016, every major El Niño event (as indicated by the magnitude of
the peak Oceanic Niño Index, when ONI > 1) corresponds to a
transition from a low to high IH difference. The ONI is a 3-month running mean
of the Nino3.4 index; similar treatment of Nino3 yields correlation
coefficients with Nino3.4 of 0.94 for both the annual maxima and minima from
1992 to 2018. However, the CO2 response is not proportional to ONI,
e.g. when
comparing 2009–2010 to 1997–1998, or most noticeably to 2015–2016 (the strongest
ONI but the smallest IH ΔCO2 step).
There is remarkable stability in IH ΔCO2 from 2010 to 2014
(despite the strong La Niña in 2011). After 2010, there are no
significantly positive ONI anomalies (El Niños), and the 5-year increase
of ∼0.1 ppm is lower than that generally attributed to the
increasing mean fossil fuel emissions (the 2010–2014 change in FF is 0.73 ppm yr-1, which at 0.36 ppm (PgC)-1 would result in a 0.26 ppm
increase). The FF with this scaling is shown at the top. There is markedly
less variability (the composite standard deviation of de-trended annual
means is 0.04 ppm) than any equivalent period over the previous 16 years
(0.31 ppm).
The 2009–2010 year-to-year change of ∼0.8 ppm (addressed in
FF16 using CSIRO data only) remains the major year-to-year change in the
annual records. The current composite data confirm the general FF16
conclusion.
The linear regression through the 25-year mlo–cgo annual data gives a slope
of 0.067±0.006 ppm yr-1 compared to that through monthly values
of 0.56±0.021 ppm yr-1 in Fig. 3, or through the peaks of
0.087±0.011 ppm yr-1 or the dips of 0.049±0.011 ppm yr-1. We interpret this as indicating the combined long-term influence
of both eddy and mean transport on the annual mean IH ΔCO2.
In 2017, the IH difference is close to the 3-decade trend, with the duct
open and Hadley strength returning to be close to its long-term mean.
Discussion
The composite monthly IH differences reveal variation from monthly to
decadal timescales that exceed measurement and sampling error (as indicated
by the composite standard deviations) thus requiring biogeochemical
explanation. This discussion focusses on the potential of IH transport
measured by wind indices to explain major features in IH ΔCO2
variation, with emphasis on periods and events when they are likely to be
the dominant influence on IH ΔCO2. It complements the more
general statistical analyses in FF18. In Fig. 6, decreasing uduct
acts to lessen eddy IH exchange and increase IH ΔCO2, while the
increasing Hadley circulation (decreasing vP and ωP)
decreases IH ΔCO2.
The fact that the magnitude of IH ΔCO2 response varies greatly
between the 1998, 2010, and 2016 El Niño events (with little or no eddy
transfer occurring in boreal winter/spring in these years) is consistent
with a quasi-decadal variation in the negative excursions of vP and
ωP in Fig. 6 (most obvious in ωP). In 1998 and
2010, the Hadley boreal summer/autumn indices are closer to zero, while 2016
registers an unprecedented negative excursion.
The complication of IH ΔCO2 variations in the boreal
winter/austral summer when uduct, ωp and vP indices
indicate that little or no IH exchange occurs (and uduct closure tends
to increase mlo CO2) is at a time when the north-to-south seasonal
variation in the Intertropical Convergence Zone (ITCZ) is near maximum. If
NH peak terrestrial emissions (biospheric and industrial) at that time are
diluted into a larger volume of well-mixed NH air, it could offset the mlo
CO2 increase anticipated from uduct closure. This volume effect is
likely to be a second-order effect in non-El Niño years.
NH terrestrial biosphere emission anomalies in the 2010–2014 period (Fig. 4) are more variable than those in 2000–2005, the opposite of the relative
behaviour in IH ΔCO2 variability in Fig. 9. These emissions
are relatively small, and frequently occur after the larger IH ΔCO2 anomalies, all inconsistent with a significant contribution to the
composite IH differences; thus, they are considered second order. The small
2010–2014 trend (∼0.1 ppm compared to 0.26 ppm expected from
fossil fuel emissions), and the steadily decreasing westerly wind strength
in uduct over the period, should increase IH ΔCO2 over the
fossil fuel trend (FF16). The flattening trend is consistent with the IH
ΔCO2 flux due to IH mixing by the Hadley process overwhelming
the increases expected from fossil fuel combustion and from decreasing
uduct strength. There is a linear relationship between uduct and
equatorial upper troposphere transient kinetic energy shown in Fig. 6 of
Frederiksen and Webster (1988) and discussed in FF16 and FF18. Note that in
Fig. 6, there is no precedent for similar sustained opposing behaviour in
the two modes of IH transfer. The trend and lack of scatter in 2010–2014 IH
ΔCO2 can be understood by the IH ΔCO2 fluxes being
significantly larger than air–surface exchanges at the time.
The magnitude of the IH flux anomalies of up to ∼2 PgC month-1 exceed known air–surface fluxes in the NH and are of a
sufficient magnitude to significantly influence NH CO2 growth rate variability.
With increasing fossil fuel fluxes, the role of IH exchange on IH ΔCO2, and NH CO2 growth, is expected to become increasingly
important.
The previous inability of carbon cycle models to simulate the 2009–2010 step
(FF16, Supplement) suggests that there is inadequate
parameterization of IH ΔCO2 transfer, particularly by eddy
exchange, in some global carbon cycle models. If this is the case, then
studies that interpret CO2 behaviour during ENSO events as a guide to
terrestrial biosphere responses to climate (e.g. Rödenbeck et al., 2018)
will also be compromised. The ability to simulate the identified features of
the composite IH ΔCO2 (within the standard deviations) would
provide convincing independent confirmation of atmospheric transport
implementation.
Conclusions
Over the last 25 years there has been a high degree of agreement in the
measurement of monthly spatial differences in background CO2 levels by
three measurement laboratories using four different sampling methodologies
and sampling frequencies. Geographic isolation of sample collection sites
and consistent sophisticated background selection over the 25 years, as well
as coincident monitoring of a wide range of atmospheric species, excludes
local and regional influence on CO2 at mlo, spo, and cgo to an extent
not generally available at other surface monitoring sites.
The temporal variation in the composite IH ΔCO2 exhibit several
systematic features on monthly to multi-year time frames that are not
reflected in independent evidence of air–surface exchange but do correspond
to features in dynamical indices selected to represent both eddy and mean IH
exchange. The comparisons in this paper imply a major role for IH exchange
of CO2 in NH growth rate variations.
The evidence for a significant influence of atmospheric dynamics on the
CO2 IH gradient has relevance for global carbon cycle studies. It
implies that both eddy and mean transport processes, and volume effects,
need to be specifically included in transport model simulations, since the
balance between the two is constantly changing, particularly in El Niño
periods when eddy transport is reduced. It also means that El Niño
events may be a poor predictor of the carbon cycle behaviour in non-ENSO
years.
Global carbon cycle model simulations should be able to reproduce the major
features identified here in the composite IH records if the re-analyses
transport is correctly implemented. In attempting to simulate the composite
differences, one complication is model selection of a baseline that matches
the flask sampling criteria. While monthly baseline averages appear to
succeed in this respect, a more comprehensive treatment (outside the scope
of this study) based on individual flask measurements rather than monthly
averages, and on other trace gas observations (FF16, FF18), and in particular
radon (Chambers et al., 2016), could possibly improve this process.
Data availability
Monthly average NOAA/ESRL, SIO, and CSIRO CO2 data were obtained
respectively from
ftp://aftp.cmdl.noaa.gov/data/trace_gases/co2/flask/ (Dlugokencky et al., 2014),
https://scrippsco2.ucsd.edu/data/atmospheric_co2/sampling_stations.html (Keeling et al., 2005), and
ftp://pftp.csiro.au/pub/data/gaslab/ (CSIRO, 2018).
Ocean Nino Index data were obtained from https://origin.cpc.ncep.noaa.gov/products/analysis_monitoring/ensostuff/ONI_v5.php (NOAA, 2019).
Meteorological data are available from the NOAA/ESRL website at
http://www.esrl.noaa.gov/psd/ (Kalnay et al., 1996).
The supplement related to this article is available online at: https://doi.org/10.5194/acp-19-14741-2019-supplement.
Author contributions
RJF generated the composite records and their analyses, while JSF provided
information on atmospheric dynamics and the roles of transport mechanisms.
LPS and RLL contributed CO2 measurement quality assessments. All four
authors contributed to the written document.
Competing interests
The authors declare that they have no conflicts of interest.
Special issue statement
This article is part of the special issue “The 10th International Carbon Dioxide Conference (ICDC10) and the 19th WMO/IAEA Meeting on Carbon Dioxide, other Greenhouse Gases and Related Measurement Techniques (GGMT-2017) (AMT/ACP/BG/CP/ESD inter-journal SI)”. It is a result of the 10th International Carbon Dioxide Conference, Interlaken, Switzerland, 21–25 August 2017.
Acknowledgements
We thank Ralph Keeling of SIO and Ed Dlugokencky of NOAA/ESRL for approval
of our use of their CO2 data; in particular we thank Ralph Keeling for his
suggestion to include SIO data from the O2/N2 program, and also
Brad Hall from NOAA for information on the early NOAA data. The sustained
focus and innovation of CSIRO GASLAB personnel, plus skilled trace gas
sample collection by personnel at the Bureau of Meteorology Cape Grim
Baseline Atmospheric Program, and NOAA personnel at Mauna Loa and South Pole
stations underpin this work. From CSIRO, Vanessa Haverd was generous with
her time in providing regional groupings and updates of the CABLE DVM data;
Ying-Ping Wang also provided other DVM data for scrutiny; Paul Krummel and
Nada Derek advised on data processing and graphics, and Cathy Trudinger and
Rachel Law on the global carbon budget. The dynamics contributions were
prepared using data and software from the NOAA/ESRL Physical Sciences
Division website at http://www.esrl.noaa.gov/psd/ (last access: 27 November 2019). Helpful comments from two
anonymous referees are acknowledged.
Review statement
This paper was edited by Christoph Zellweger and reviewed by two anonymous referees.
ReferencesBastos, A., Freidlingstein, P., Sitch, S., Chi Chen, Mialon, A. Wigneron,
J-P., Arora, V. K.,, Briggs, P.R., Canadell, J. G., Ciais, P., Chevallie,
F., Lei Cheng, Delire, C., Haverd, V., Jain, A. K., Joos, F., Kato, E.,
Lienert, S., Lombardozzi, D., Melto, J. R.,Myneni, R., Nabel, J. E. M. S.,
Pongratz, J., Poulter, B., Rödenbeck, C., Séférian, R., Tian,
H., van Eck, C., Viovy, N., Vuichard, N., Walker, A. P., Wiltshire, A., Jia
Yang, Zaehle, S., Ning Zeng, and Zhu, D: Impact of the 2015/2016 El Niño
on the terrestrial carbon cycle constrained by bottom-up and top-down
approaches, Philos. T. Roy. Soc. B, 373, 20170304,
10.1098/rstb.2017.0304, 2018.
Bowman, K. P. and Cohen, P. J.: Interhemispheric exchange by seasonal
modulation of the Hadley Circulation, J. Atmos. Sci., 54, 2045–2059, 1997.Chambers, S. D., Williams, A. G., Conen, F., Griffiths, A. D., Reimann, S., Steinbacher, M., Krummel, P. B., Steele, L. P., van der Schoot, M. V.,
Galbally, I. E., Molloy, S. B., and Barnes J. E.: Towards a Universal
“Baseline” Characterisation of Air Masses for High- and Low-Altitude
Observing Stations Using Radon-222, Aerosol Air Qual. Res., 16,
885–899, 2016, 10.4209/aaqr.2015.06.0391, 2016.CSIRO: CSIRO Oceans and Atmosphere GASLAB data October 2018, Commonwealth Scientific and Industrial Research Organisation, available at: ftp://gaspublic:gaspublic@pftp.csiro.au/pub/data/gaslab/ (last access: 28 January 2019), 2018.
Conway, T. J., Tans, P. P., Waterman, L. S., Thoning, K. W., Kitzis, D. R.,
Masarie, K. A., and Zhang, N.: Evidence for interannual variability of the
carbon cycle from the National Oceanic and Atmospheric
Administration/Climate Monitoring and Diagnostics Laboratory Global Air
Sampling Network, J. Geophys. Res., 99, 22831–22855, 1994.Dlugokencky, E. J., Lang, P. M., Masarie, K. A., Crotwell, A. M., and Crotwell,
M. J.: Atmospheric Carbon Dioxide Dry Air Mole Fractions from the NOAA ESRL
Carbon Cycle Cooperative Global Air Sampling Network, 1968–2013, Version:
2014-06-27, available at: ftp://aftp.cmdl.noaa.gov/data/trace_gases/co2/flask/surface/ (last access: 28 January 2019), 2014.
Dunse, B. L., Steele, L. P., Fraser, P. J., and Wilson, S. R.: An analysis of
Melbourne pollution episodes observed at Cape Grim from 1995–1998, in:
Baseline Atmospheric Program (Australia) 1997–1998, edited by: Tindale, N. W., Derek, N., and Francey, R. J., Bureau of
Meteorology and CSIRO Atmospheric Research, Melbourne, Australia, 34–42,
2001.Francey, R. J. and Frederiksen, J. S.: The 2009–2010 step in atmospheric CO2 interhemispheric difference, Biogeosciences, 13, 873–885, 10.5194/bg-13-873-2016,
2016.
Francey, R. J., Steele, L. P., Langenfelds, R. L., Lucarelli, M., Allison, C.
E., Beardsmore, D. J. Coram, S. A., Derek, N., de Silva, F. R., Etheridge,
D. M., Fraser, P. J., Henry, R. J., Turner, B., and Welch, E. D.: Global
Atmospheric Sampling Laboratory (GASLAB): supporting and extending the Cape
Grim trace gas programs, in: Baseline Atmospheric Program (Australia) 1993, edited by: Francey, R. J., Dick, A. L., and Derek, N., Bureau
of Meteorology and CSIRO Division of Atmospheric Research, Melbourne,
8–29, 1996.Frederiksen, J. S. and Francey, R. J.: Unprecedented strength of Hadley circulation in 2015–2016 impacts on CO2 interhemispheric difference, Atmos. Chem. Phys., 18, 14837–14850, 10.5194/acp-18-14837-2018, 2018.
Frederiksen, J. S. and Webster, P. J.: Alternative theories of atmospheric
teleconnections and low-frequency fluctuations, Rev. Geophys., 26, 459–494,
1988.Haverd, V., Smith, B., Nieradzik, L., Briggs, P. R., Woodgate, W., Trudinger, C. M., Canadell, J. G., and Cuntz, M.: A new version of the CABLE land surface model (Subversion revision r4601) incorporating land use and land cover change, woody vegetation demography, and a novel optimisation-based approach to plant coordination of photosynthesis, Geosci. Model Dev., 11, 2995–3026, 10.5194/gmd-11-2995-2018, 2018.
Jacob, D. J.: Introduction to Atmospheric Chemistry, Princeton University
Press, 1999.Kalnay, E., Kanamitsu, M., Kistler, R., Collins, W., Deaven, D., Gandin, L.,
Iredell, M., Saha, S., White, G., Woollen, J., Zhu, Y., Leetmaa, A.,
Reynolds, R., Chelliah, M., Ebisuzaki, W., Higgins, W., Janowiak, J., Mo, K.
C., Ropelewski, C., Wang, J., Jenne, R., and Joseph, D.: The NCEP/NCAR
Reanalysis 40-year Project, B. Am. Meteorol. Soc., 77, 437–471, 1996 (data available at: http://www.esrl.noaa.gov/psd/timeseries/, last access: 28 January 2019).
Keeling, C. D.: Rewards and penalties of monitoring the earth, Annu. Rev. Energy Environ.,
23, 25–82, 1998.Keeling, C. D., Piper, S. C., Bacastow, R. B., Wahlen, M., Whorf, T. P.,
Heimann, M., and Meijer, H. A.: Exchanges of atmospheric CO2 and
13CO2 with the terrestrial biosphere and oceans from 1978 to 2000.
I. Global aspects, SIO Reference Series, No. 01–06, Scripps Institution of
Oceanography, San Diego, 88 pp., 2001.Keeling, C. D., Piper, S. C., Bacastow, R. B., Wahlen, M., Whorf, T. P., Heimann, M., and Meijer, H. A.: Atmospheric CO2 and 13CO2 exchange with the terrestrial biosphere and oceans from 1978 to 2000: observations and carbon cycle implications, 83–113, in: A History of Atmospheric CO2 and its effects on Plants, Animals, and Ecosystems, edited by: Ehleringer, J. R., Cerling, T. E., Dearing, M. D., Springer Verlag, New York, 2005 (data available at: https://scrippsco2.ucsd.edu/data/atmospheric_co2/sampling_stations.html, last access: 28 January 2019).
Keeling, R. F. and Schertz, S. R.: Seasonal and interannual variations in
atmospheric oxygen and implications for the global carbon cycle, Nature,
358, 723–727, 1992.
Kowalczyk, E. A., Wang Y. P., Law, R. M., Davies, H. L. McGregor, J. L., and
Abramowitz, G.: The CSIRO Atmosphere Biosphere Land Exchange (CABLE) model
for use in climate models and as an offline model, National Library of
Australia Cataloguing-in-Publication, ISBN 1 921232 39 0, 2006.
Langenfelds, R. L., Francey, R. J., Steele, L. P., Dunse, B. L., Butler, T.
M., Spencer, D. A., Kivlinghon, L. M., and Meyer, C. P.: Flask sampling from
Cape Grim overflights, in: Baseline Atmospheric Program (Australia) 1999–2000, edited by: Tindale, N. W., Derek, N., and
Fraser, P. J., Bureau of Meteorology and CSIRO Atmospheric Research, Melbourne,
Australia, 73–75, 2003.Lenton, A., Tilbrook, B., Law, R. M., Bakker, D., Doney, S. C., Gruber, N., Ishii, M., Hoppema, M., Lovenduski, N. S., Matear, R. J., McNeil, B. I., Metzl, N., Mikaloff Fletcher, S. E., Monteiro, P. M. S., Rödenbeck, C., Sweeney, C., and Takahashi, T.: Sea–air CO2 fluxes in the Southern Ocean for the period 1990–2009, Biogeosciences, 10, 4037–4054, 10.5194/bg-10-4037-2013, 2013.Le Quéré, C., Andrew, R. M., Friedlingstein, P., Sitch, S., Pongratz, J., Manning, A. C., Korsbakken, J. I., Peters, G. P., Canadell, J. G., Jackson, R. B., Boden, T. A., Tans, P. P., Andrews, O. D., Arora, V. K., Bakker, D. C. E., Barbero, L., Becker, M., Betts, R. A., Bopp, L., Chevallier, F., Chini, L. P., Ciais, P., Cosca, C. E., Cross, J., Currie, K., Gasser, T., Harris, I., Hauck, J., Haverd, V., Houghton, R. A., Hunt, C. W., Hurtt, G., Ilyina, T., Jain, A. K., Kato, E., Kautz, M., Keeling, R. F., Klein Goldewijk, K., Körtzinger, A., Landschützer, P., Lefèvre, N., Lenton, A., Lienert, S., Lima, I., Lombardozzi, D., Metzl, N., Millero, F., Monteiro, P. M. S., Munro, D. R., Nabel, J. E. M. S., Nakaoka, S., Nojiri, Y., Padin, X. A., Peregon, A., Pfeil, B., Pierrot, D., Poulter, B., Rehder, G., Reimer, J., Rödenbeck, C., Schwinger, J., Séférian, R., Skjelvan, I., Stocker, B. D., Tian, H., Tilbrook, B., Tubiello, F. N., van der Laan-Luijkx, I. T., van der Werf, G. R., van Heuven, S., Viovy, N., Vuichard, N., Walker, A. P., Watson, A. J., Wiltshire, A. J., Zaehle, S., and Zhu, D.: Global Carbon Budget 2017, Earth Syst. Sci. Data, 10, 405–448, 10.5194/essd-10-405-2018, 2018.NOAA: Cold & Warm Episodes by Season, data from Climate Prediction Center, National Oceanic and Atmospheric Administration, data available at: https://origin.cpc.ncep.noaa.gov/products/analysis_monitoring/ensostuff/ONI_v5.php, last access: 27 November 2019.Oda, T., Maksyutov, S., and Andres, R. J.: The Open-source Data Inventory for Anthropogenic CO2, version 2016 (ODIAC2016): a global monthly fossil fuel CO2 gridded emissions data product for tracer transport simulations and surface flux inversions, Earth Syst. Sci. Data, 10, 87–107, 10.5194/essd-10-87-2018, 2018.Ortega, S., Webster, P. J., Toma, V., and Chang, H. R.: The effect of
potential vorticity fluxes on the circulation of the tropical upper
troposphere, Q. J. Roy. Meteorol. Soc., 144, 848–860, 10.1002/qj.3261, 2018.
Pak, B. C., Langenfelds, R. L., Francey, R. J., Steele, L. P., and Simmonds,
I.: A climatology of trace gases from the Cape Grim Overflights, 1992–1995,
in: Baseline Atmospheric Program (Australia), 1994–95, edited by: Francey, R. J., Dick, A. L., and Derek, N., Bureau of
Meteorology and CSIRO Division of Atmospheric Research, Melbourne,
Australia, 41–52, 1996.Randerson, J. T., van der Werf, G. R., Giglio, L., Collatz, G. J., and
Kasibhatla, P. S.: Global Fire Emissions Database, Version 4.1 (GFEDv4). ORNL
DAAC, Oak Ridge, Tennessee, USA,
10.3334/ORNLDAAC/1293, 2018.Rödenbeck, C., Zaehle, S., Keeling, R., and Heimann, M.: How does the terrestrial carbon exchange respond to inter-annual climatic variations? A quantification based on atmospheric CO2 data, Biogeosciences, 15, 2481–2498, 10.5194/bg-15-2481-2018, 2018.Stavert, A. R., Law, R. M., van der Schoot, M., Langenfelds, R. L., Spencer, D. A., Krummel, P. B., Chambers, S. D., Williams, A. G., Werczynski, S., Francey, R. J., and Howden, R. T.: The Macquarie Island (LoFlo2G) high-precision continuous atmospheric carbon dioxide record, Atmos. Meas. Tech., 12, 1103–1121, 10.5194/amt-12-1103-2019, 2019.
Tans, P. P., Conway, T. J., Dlugokencky, E. J., Thoning, K. W., Lang, P. M.,
Masarie, K. A., Novelli, P., and Waterman, L. S.: 2. Carbon Cycle Division,
2.1 Continuing Programs, Climate Monitoring and Diagnostics Laboratory, 20,
15–24, 1992.
Webster, P. J. and Holton, J. R.: Cross-equatorial response to mid-latitude
forcing in a zonally varying basic state, J. Atmos. Sci., 39, 722–733, 1982.Yue, C., Ciais, P., Bastos, A., Chevallier, F., Yin, Y., Rödenbeck, C., and Park, T.: Vegetation greenness and land carbon-flux anomalies associated with climate variations: a focus on the year 2015, Atmos. Chem. Phys., 17, 13903–13919, 10.5194/acp-17-13903-2017, 2017.