Variability in a four-network composite of atmospheric CO2 differences between three primary baseline sites
- CSIRO Oceans and Atmosphere, Aspendale, Victoria, Australia
Correspondence: Roger J. Francey (firstname.lastname@example.org)
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.
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.
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.
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.
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.
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.
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.
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.
5.1 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.)
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.
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.
5.2 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.
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 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.
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 ppm variation in the detrended annual record.
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.
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.
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.
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.
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.
The authors declare that they have no conflicts of interest.
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.
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.
This paper was edited by Christoph Zellweger and reviewed by two anonymous referees.
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