Downwind control of oceanic air by land: the land wake and its sensitivity to CO₂

We define the 'land wake' W_x of some variable x as the deviation from the zonal mean value of x of air over the ocean ($[x_{ocean}]$) at a given latitude (equation (1)).

$$\begin{equation} W_{x} = x - [x_{ocean}] \end{equation} \tag{ 1 }$$

W_x can be calculated globally or for a single ocean basin, at any atmospheric level. We focus on the near-surface wake averaged from 1000–900 hPa (the precise shape of the wake varies depending on the vertical levels selected), downwind of mid-latitude North America, where $[x_\mathrm{ocean}]$ is averaged longitudinally over the Atlantic ocean, defined here to be the ocean region between 100^∘ W–0^∘ W and 20^∘ N–70^∘ N. $W_x\lt0$ indicates x at that location is smaller than the average oceanic value of x for that latitude, while $W_x\gt0$ indicates x is larger than the average oceanic value of x for that latitude. In this study, we focus on the wakes of relative humidity ($W_\mathrm{RH}$), specific humidity ($W_\mathrm{Q}$), and temperature ($W_\mathrm{T}$).

We analyse the land wake using a combination of reanalysis data and climate model output. We use the ERA5 reanalysis global dataset at 0.25^∘ resolution, which spans from 1979–2019 (Hersbach et al 2020). Specifically, we use the ERA5 variables for: relative humidity ($\mathrm{RH}$), specific humidity (q), temperature (T), and winds (U, V).

We also evaluate the wake in Earth system models under low and high atmospheric CO₂ concentrations. Increased atmospheric CO₂ has both radiative and biological impacts. These effects are isolated in a pair of experiments from the Coupled Climate-Carbon Cycle Model Intercomparision Project (C4MIP Friedlingstein et al 2006, Jones et al 2016). In '1pctCO2-bgc', only the biogeochemical (bgc) components of the models experience CO₂ increasing at 1% per year from pre-industrial values (280 ppm), while the atmosphere experiences constant pre-industrial CO₂. In '1pctCO2-rad', only the atmospheric radiation experiences increasing CO₂ (at 1% per year), while the biogeochemical components of the model experience fixed pre-industrial CO₂. We also consider simulations from the '1pctCO2' experiment from the Coupled Model Intercomparison Project v. 6 (Eyring et al 2016), where both radiation and biogeochemistry experience the increase in CO₂. Simulations are run for 140 years, at which point CO₂ concentrations reach 4x pre-industrial levels (figure S1). We refer to the increase in CO₂ experienced by the biogeochemistry as bgcCO₂, the increase in CO₂ experienced by the atmospheric radiation as radCO₂, and use fullCO₂ to refer to simulations where both biogeochemistry and radiation experience increased CO₂.

Both the terrestrial and oceanic carbon cycles experience the increase in bgcCO₂, but impacts on surface fluxes are only driven by the response of terrestrial vegetation to increased CO₂. The bgcCO₂ simulations thus provide a clear demonstration of changes in land driving changing conditions over the ocean in the land wake. At higher CO₂ concentrations, plants can grow more (CO₂ fertilization; Field et al 1995, Sellers et al 1996, Medlyn et al 2001, Morison 2001), altering leaf area and land albedo, in turn leading to changes in surface temperatures and evapotranspiration (Donohue et al 2013, Swann et al 2016, Zarakas et al 2020a). Atmospheric CO₂ concentrations directly impact the stomatal conductance of vegetation. In global Earth system models, the equations governing stomatal conductance generate less conductance (thus less evapotranspiration, with all else held equal) with increased atmospheric CO₂ concentrations (Ball et al 1987, Medlyn et al 2011), changing plants' water use efficiency (Eamus 1991, Cheng et al 2014).

The effect of the biogoechemical response to increased CO₂ could be obtained by considering results from high-low CO₂ values in the bgcCO₂ simulations, or by subtracting the radCO₂ simulation from the fullCO₂ simulation. The first gives the effect of increasing bgcCO₂ in a low-CO₂ climate, while the latter gives the effect of increasing bgcCO₂ in a high-CO₂ climate; we focus on the first approach in this study, but show that the two methods give generally qualitatively similar responses of the land wake.

We use the nine models participating in C4MIP that provided temperature, specific humidity, and RH data for the 1pctCO2-bgc and 1pctCO2-rad C4MIP experiments and were available for download from the LLNL ESGF node (https://esgf-node.llnl.gov/projects/cmip6/) on 11 August 2021 (table 1). These models were BCC-CSM2-MR, CanESM5, CNRM-ESM2-1, ACCESS-ESM1-5, IPSL-CM6A-LR, MIROC-ES2L, UKESM1-0-LL, MPI-ESM1-2-LR, and CESM2. Cloud cover data was available for all models except CanESM5, so low cloud and downwelling solar radiation are evaluated for the remaining eight models only; snow cover and leaf area changes are shown for all models except CanESM5 and BCC-CESM2-MR as neither model had both variables available. Where possible, the r1i1p1f1 ensemble member was used, otherwise the r1i1p1f2 ensemble member was used; only a single ensemble member per model was considered. All output was re-gridded to a common 1^∘ resolution horizontal grid. For each field considered, we took the difference of the average of years 121–140 (high CO₂) and subtracted the average of years 1–20 (low CO₂).

To compare changes in the $\mathrm{RH}$ wake in the C4MIP simulations with increased bgcCO₂ and radCO₂, we define the 'magnitude' M of the wake as the area-weighted sum of the spatial footprint of the Atlantic $\mathrm{RH}$ wake, in this case, anywhere RH $\lt0$, multiplied by the intensity of the wake $W_\mathrm{RH}$, where $W_\mathrm{RH}$ is how negative $\mathrm{RH}$ is at each location:

$$\begin{align} M_\mathrm{RH,Atlantic} = \frac{\int_{105~W}^{25~W} \int_{25~N}^{65~N} \big( \cos{\phi} W_\mathrm{RH} \big)|_{W_\mathrm{RH}\lt0} d\phi d\lambda}{\int_{105~W}^{25~W} \int_{25~N}^{65~N} \big( \cos{\phi} \big)|_{W_\mathrm{RH}\lt0} d\phi d\lambda}, \end{align} \tag{ 2 }$$

where φ is the latitude. $M_\mathrm{RH,Atlantic}$ is calculated using $W_\mathrm{RH}$ in the domain 25–65^∘ N and 105–25^∘ W, only where $W_\mathrm{RH}$ is negative.

We take the difference of the average monthly $M_\mathrm{RH,Atlantic}$ of years 121–140 (high bgcCO₂ and radCO₂) minus the average monthly $M_\mathrm{RH,Atlantic}$ of years 1–20 (low bgcCO₂ and radCO₂) for the 1pctCO2-bgc and 1pctCO2-rad simulations, respectively. We similarly calculate the change in continental $\mathrm{RH}$ (averaged over land areas from 1000 to 90 hPa, between 25^∘–65^∘ N and 125^∘–55^∘ W) for each month of the year, then take seasonal averages. The wake area changes slightly with increasing bgcCO₂ and radCO₂ (table S1), but most of the change in the wake magnitude comes from changes within the climatological wake footprint.

To evaluate the individual contributions of T and q to the change in $W_\mathrm{RH}$, we define an additional metric, the change in the wake 'intensity' (equation (3)). $I_\mathrm{RH}$ is calculated in the box 25^∘–65^∘ N and 105^∘–25^∘ W, only where $\Delta W_\mathrm{RH} \lt 0$ (see figure 3) for each individual model. Within that footprint, a similar approach is used as with the calculation of the magnitude $M_\mathrm{RH}$, i.e. taking the area-weighted sum of the wake values in that region (but only where the wake changes, i.e. where $\Delta W_\mathrm{RH} \lt 0$).

$$\begin{equation} I_\mathrm{RH} = \frac{\int_{105~W}^{25~W} \int_{25~N}^{65~N} \big( \cos{\phi} W_\mathrm{RH} \big)|_{\Delta W_\mathrm{RH}\lt0} d\phi d\lambda}{\int_{105~W}^{25~W} \int_{25~N}^{65~N} \big( \cos{\phi} \big)|_{\Delta W_\mathrm{RH}\lt0} d\phi d\lambda}, \end{equation} \tag{ 3 }$$

To separate the impact of T and q on $W_\mathrm{RH}$, $\Delta W_\mathrm{RH}$, and $I_\mathrm{RH}$, we calculate $W_\mathrm{RH}$ in three distinct ways. First, we calculate $W_\mathrm{RH}$ in the standard manner (equation (1)). Second, we calculate what the $\mathrm{RH}$ field would be in the high CO₂ case if we hold q fixed to the low CO₂ value, but increase T to the high CO₂ value; we then calculate the wake of that new $\mathrm{RH}$ value that only accounts for the change in T. Third, we calculate what the $\mathrm{ RH}$ field and $W_\mathrm{RH}$ would look like with high CO₂ if only q is increased to high CO₂, with T held to the low CO₂ value. $I_\mathrm{RH}$ can then be calculated within the $\Delta W_\mathrm{RH}$ footprint for low CO₂ (using $W_\mathrm{RH}$ at low CO₂), for $\mathrm{RH}$ associated with low CO₂ q and high CO₂ T, for $\mathrm{RH}$ associated with low CO₂ T and high CO₂ Q, and for $\mathrm{RH}$ at high CO₂. These terms do not add linearly as $\mathrm{RH}$ is not a linear function of T and q and we approximate the effect of Δ T and Δ q on $\Delta\mathrm{RH}$ using monthly mean T and q values. The actual change in $\mathrm{RH}$ is qualitatively similar but numerically differs slightly from the change in $\mathrm{RH}$ calculated from monthly mean T and q together (not shown).

From equation (1), changes in the wake W_x could occur either because of a change in the spatial pattern of x, or because of a change in the zonal mean of x averaged over the Atlantic basin. In the case of increased bgcCO₂, the changes in the $\mathrm{RH}$ wake are driven by spatial changes in $\mathrm{RH}$ off the North American east coast, with very little change in the zonal mean value of $\mathrm{RH}$ over the Atlantic ocean. This is not surprising, as bgcCO₂ primarily impacts surface energy and water fluxes over the continents; this signal is then advected over the ocean and thus is strongest in the wake region, with very little change in the average Atlantic $\mathrm{RH}$ at each latitude (figure S2).

Increasing radCO₂ has a strong impact on the atmospheric radiative budget and surface temperatures over both land and ocean regions, causing the zonal mean Atlantic values of T and q to increase, but zonally averaged RH over the ocean remains roughly the same (O'Gorman and Muller 2010). Regionally however, $\mathrm{RH}$ and the $\mathrm{RH}$ wake become drier with increasing radCO₂ off the North American east coast, as a result of continental warming (figure S3).

Western boundary currents, such as the Gulf Stream, cause a tongue of anomalously warm water extending into cooler waters at higher latitudes along the east coast of mid-latitude continents, releasing heat and moisture to the atmosphere above it. In the context of our definition of the wake, this would raise the overall zonally averaged $\mathrm{RH}$, T, and q of the ocean basin. As such, air that is near the coast but has not yet reached the eastern boundary current could look anomalously cold or dry simply because it has not yet come into contact the warmer waters. The C4MIP simulations analyzed here show that the $\mathrm{RH}$, T, and q of the wake region change in response to increased bgcCO₂ and radCO₂—that is, even if the wake only existed as an artifact of the eastern boundary currents, our analysis still shows the sensitivity of oceanic air in this region to terrestrially-driven changes in continental air. Changes in the location and strength of the Gulf Stream, e.g. with changing climate, could alter both the shape and strength of the continental wake by modifying wind patterns and ocean-atmosphere exchange of heat and moisture. For example, a warmer gulf stream could potentially reduce the continental wake as increased ocean evaporation could more quickly make continental air resemble ocean air.

Analysis was conducted using the Python programming language, heavily leveraging the NumPy (Harris et al 2020) and xarray (Hoyer and Hamman 2017) packages.

In reanalysis data, $W_\mathrm{RH}$ downwind of mid-latitude North America is negative in both winter (December–February, DJF) and summer (June–August, JJA), becoming positive as air moves further from shore (figures 1(a) and (b)); that is, air near the continent has a lower RH than air farther from land. The shape of the wake changes with season, but its longest dimension is generally aligned parallel to the time-mean low-level wind with a peak length of at least 1000 km. In winter, $W_\mathrm{RH}$ extends from the northern coast of the Gulf of Mexico along the eastern seaboard of the USA to around 50^∘ N, with an additional region of anomalously low $\mathrm{RH}$ around Greenland (figure 1(a)). In summer, the wake lies closer to shore along the east coast of the USA, but extends further north along the east coast of Canada and Hudson Bay, while the Greenland wake retreats towards the coast of that land mass (figure 1(b)). Eastward winds off the continent are stronger in winter, and the winter wake correspondingly extends farther from shore between 30^∘–45^∘ N. In contrast, the summer wake is more intense and occupies a larger fraction of the coastline of North America.

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While the $\mathrm{RH}$ wake is apparent in both winter and summer, it is driven by different mechanisms seasonally. In winter, cold, dry air flows from the continent over ocean, producing low temperature (T) and low specific humidity (q) offshore (figures 2(a) and (c)). However, because saturation vapor pressure decreases rapidly as temperature drops, the low temperatures in more northern regions prevent anomalously low $\mathrm{RH}$ downwind of the continent. The winter wakes in T and q extend along nearly the entire coastline of North America, but their interaction limits the $\mathrm{RH}$ wake to latitudes south of Nova Scotia (figure 1(a)). In summer, time-mean winds along eastern North America flow more parallel to the coast due to the summertime anticyclone over the Atlantic (figure 1(b)). At more southern latitudes, q is anomalously positive due to moisture transport from the low-latitude ocean, while at more northern latitudes q is anomalously negative where there is flow off the continent (figure 2(b)). Close to the continent, air is warmer than average ocean air (figure 2(d)); as in winter, the shape of $W_\mathrm{RH}$ results from the strong response of saturation vapor pressure to temperature, yielding low $\mathrm{RH}$ along the whole east coast of North America despite anomalously high q south of 40^∘ N (figures 1(b) and 2(b)).

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Mixing between continental and maritime regions produced by shorter timescale motions and vertical circulations is likely also a factor in determining the wake. For example, along the northern coast of the Gulf of Mexico, monthly mean winds are climatologically onshore, yet a distinct band of low RH air hugs the coast. The difference in seasonal drivers of $W_\mathrm{RH}$ illustrate the potential for long-term changes to land, such as those caused by agriculture, fire, drought, and rising atmospheric CO₂, to impact ocean regions downwind of continents through multiple mechanisms.

Temperature increases due to climate change are occurring more quickly over land than over most ocean regions (IPCC 2007), and the existence of the land wake provides a pathway for land's response to climate change to influence ocean regions. Earth System Models (ESMs) participating in C4MIP produce a qualitatively similar $W_\mathrm{RH}$ to that seen in reanalysis (cf top and bottom rows of figure 1), with the ERA5 wake having a slightly larger area than the models (table S1).

Increases in bgcCO₂ and radCO₂ each separately strengthen the land wake (figure 3). The spatially integrated magnitude of the RH wake ($M_\mathrm{RH,Atlantic}$, see section 2) intensifies in response to both increased bgcCO₂ and radCO₂, with both driving similar magnitudes of change in $M_\mathrm{RH,Atlantic}$ during summer (figures 4(a) and (b)). This intensification of the wake accompanies a large reduction of $\mathrm{RH}$ over continental North America during spring, summer, and autumn (figure 4), although the spring and autumn changes are less pronounced in the radCO₂ experiment. There is a strong correlation across models ($r^2 = 0.72$) between the change in $\mathrm{RH}$ over North America and the change in the $\mathrm{RH}$ wake magnitude associated with increased bgcCO₂, and a more modest correlation ($r^2 = 0.57$) for radCO₂. These results support the hypothesis that changes in continental air induced by the radiative and biogeophysical effects of CO₂ drive changes in the wake.

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When radiative and biogeochemical effects of CO₂ are both considered (using the fullCO₂ simulations), the radCO₂ effects are seen to dominate the spatial pattern of change in the wake found in the fullCO₂ simulation (figure 3). However, the effects of bgcCO₂ contribute substantial changes to the wake magnitude (figure 4(c)). We would not expect the wake changes in the bgcCO₂ and radCO₂ simulations to sum perfectly linearly to the fullCO₂ wake change because the results of the ΔbgcCO₂ simulations show the effect of bgcCO₂ in a low-CO₂ background climate, while the difference between the fullCO₂ and radCO₂ simulations show the effect of bgcCO₂ in a high-CO₂ background climate. However, the full−rad wake response is qualitatively similar to the bgcCO₂ wake response in most regions, except in the Hudson Bay and Davis Straight areas (figures S4–S6). These differences are not the focus of our study, but we suspect they result from large differences in base-state Arctic climate with low vs. high CO₂.

We now focus on how the response of terrestrial evapotranspiration to increased bgcCO₂ impacts the wake. The summer response across the C4MIP simulations to increased bgcCO₂ is a widespread reduction in terrestrial evapotranspiration (figure S7). The terrestrial biogeophysical effects of increased CO₂ produce a wake that is drier (i.e. $W_\mathrm{RH}$ is anomalously low) over the coastal ocean in both summer and winter with increased bgcCO₂ (figures 3(a) and (b)). Note that though the RH wake becomes drier, specific humidity in the summer wake actually increases because of circulation changes, discussed below. While the horizontal extent of this change in the wake is smaller in summer, the peak intensity (defined by equation (3)) of this change is larger during that season. Although the definition of W (equation (1)) encompasses the whole Atlantic, the change in $W_\mathrm{RH}$ is primarily the result of changes near North America, with little influence from the zonal-mean change (figures S2, S7; see section 2 for further discussion).

In summer, continental $\mathrm{RH}$ decreases as a result of reduced continental evapotranspiration driven by the stomatal response to increased bgcCO₂, but the influence of these continental changes on the $\mathrm{RH}$ wake is more complex than a simple downwind transport of low-$\mathrm{RH}$ air. The reduced continental evapotranspiration causes lower continental q as well as higher continental T due to reduced latent cooling of land (see land regions in figures 3(b) and S7((b), (d)). Downwind transport of high-T air over the wake region drives an intensification of the $\mathrm{RH}$ wake, but this is partially opposed by an increase in q over much of the western Atlantic (figures S7(b) and (d)). Because continental q decreases with increased bgcCO₂, the downwind transport of lower q continental air cannot explain the summer increase in q in the wake region. Summer ocean evaporation also does not change substantially in the wake region (figure S8). Instead, the increase in summer q over the western Atlantic is due to a change in atmospheric circulation. The continental warming resulting from reduced evapotranspiration induces an anomalous low-level cyclone centered over the eastern edge of the continent, bringing warm, moist air from the Gulf of Mexico and Caribbean northward along the coast of North America (figure S9). The transport of moisture into the wake region by this anomalous circulation damps the temperature-induced intensification of the RH wake. The combination of the opposing effects of the T and q changes on RH produces a slight intensification of the RH wake with increased bgcCO₂ across all models (figure 5).

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The drivers of winter $W_\mathrm{RH}$ change are different than those in summer. In winter, there is little change in continental $\mathrm{RH}$, but $W_\mathrm{RH}$ still intensifies (figure 3). Rather than being driven directly by downwind transport of anomalous continental $\mathrm{RH}$, the reduction in winter $W_\mathrm{RH}$ is due to increased continental temperatures, which are caused by a darker land surface resulting from reduced snow cover and increased winter leaf area (figures S7 and S10). The darkened land surface absorbs more shortwave energy, leading to warming of the surface and overlying atmosphere in the winter.

The radiative effects of enhanced CO₂ intensify $W_\mathrm{RH}$ in both winter and summer, but with spatial patterns and magnitudes that differ greatly from those caused by the biogeophysical effects of CO₂ (figure 3). The intensification of $W_\mathrm{RH}$ with increased radCO₂ results from the radiatively driven increase in land surface temperatures (figure S11). In contrast to increased bgcCO₂, which has the largest impact on temperatures over land, increased radCO₂ drives widespread warming over both land and ocean, with stronger warming than bgcCO₂; the hotter atmosphere has a greater evaporative demand and causes increased terrestrial evapotranspiration in many regions (figure S11). This response of evapotranspiration is not large enough to fill the enhanced vapor pressure deficit that drives it, so the large increase in land temperatures produces a decrease in summer continental $\mathrm{RH}$ and an intensification of $W_\mathrm{RH}$. As in the bgcCO₂ case, q increases in the region where $W_\mathrm{RH}$ intensifies as a result of anomalous poleward winds over the western Atlantic, but for radCO₂ this poleward flow is part of an anomalous anticyclone centered over Florida and the Gulf of Mexico (figure S9). On its own, this increase in q would lead to a weaker $W_\mathrm{RH}$, while the increase in T would lead to a strong intensification of $W_\mathrm{RH}$ (figure 5). The final result is a modest intensification of the wake as a result of increased radCO₂.

While the peak magnitude of the wintertime $W_\mathrm{RH}$ change is more than twice as large for radCO₂ than for bgcCO₂, the wake change induced by bgcCO₂ occupies a larger spatial area (cf figures 3(a) and (c)). In summer, the radCO₂ wake changes are largest in Hudson Bay and the Gulf of Mexico, while for bgcCO₂ the changes occur largely along the eastern seaboard of Canada and the USA, suggesting that future plant changes driven by the biogeophysical effects of CO₂ will dominate changes in the wake from Florida to Newfoundland.

Increased CO₂ drives changes not only in coastal $\mathrm{RH}$, but in coastal cloud cover and radiation. With increased bgcCO₂, summer low cloud cover decreases along the full length of North America's east coast, from the Gulf of Mexico to Hudson Bay (figure 6(a)). The reduced cloud fraction occurs year-round, with some seasonal changes in structure; cloud loss in the Gulf of Mexico is largest in winter (not shown), while cloud loss in Hudson Bay is largest in summer. Co-located with the reduction in low cloud cover is an increase of roughly 3 W m⁻² in downwelling shortwave radiation (figure 6(b)). Coastal low cloud cover is also reduced as a result of increased radCO₂, though for that forcing the total change in downwelling shortwave radiation at the surface is complicated by the response of clouds and other absorbers to the radiative effects of CO₂ (not shown).

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