Breeden et al. (2021) showed that the mass of stratospheric air entering the lower troposphere over western North America is 3 times larger during the jet's spring transition phase as opposed to its late-winter or summer phases. This peak in mass transport is associated with enhanced synoptic-scale wave activity in the upper troposphere, tropopause folds that reach deeper into the troposphere, and a deeper planetary boundary layer. Because the seasonal evolution of the North Pacific jet impacts STT over western North America, in all of our analyses we consider changes in all fields as a function of the three phases of the seasonal transition of the North Pacific jet, as they are defined in Breeden et al. (2021). Therefore, the differences in transport arising from timing of the jet transition are inherently taken into account.

Figure 1 shows the seasonal evolution of the North Pacific jet in the preindustrial control (EXP1) and in the RCP8.5 experiment (EXP2). The jet is separated into winter, spring, and summer phases using the principal component time series associated with the first empirical orthogonal function (EOF) of the daily 200 hPa zonal winds averaged over the North Pacific region (100–280∘ E and 10–70∘ N). The zonal wind anomalies used for the EOF analysis are calculated with respect to the February to June years 11–60 average, rather than a daily climatology, in order to deliberately preserve the seasonal cycle that emerges as the first EOF. The associated principal component time series (PC1), calculated by projecting the gridded zonal wind for either the preindustrial control (EXP1) (Fig. 1d) or the RCP8.5 experiment (EXP2) (Fig. 1e) at each time step onto each experiment's EOF1, are smoothed with a 5 d running mean.

The winter jet is present when PC1 > 1 standard deviation (σ), during which the Pacific jet is strong and narrow (Fig. 1a). The spring jet is present when PC1 < 0.5σ and > -0.5σ, at which point the subtropical jet weakens and shifts north, and the secondary subtropical jet maximum extends between Hawaii and western North America (Fig. 1b). The summer jet is present when PC1 < -1σ (Fig. 1c). The jet weakens substantially and remains shifted poleward, and the secondary jet maximum over North America weakens. The structure of winter, spring, and summer jets (Fig. 1a–c) compares well with that from 1958–2017 Japanese Reanalysis-55 data (cf. Fig. 2, Breeden et al., 2021), as does the timing of the phase changes (Fig. 1d–g; cf. Figs. 1 and 3; Breeden et al., 2021).

The RCP8.5 North Pacific jet exhibits increases in variability compared to the preindustrial control during much of spring and summer (Fig. 1e). Recomputing Fig. 1e using EXP3, which includes RCP8.5 SSTs alone, confirms that the changing jet variability is associated with the SSTs (not shown). Despite these changes in variability, there is no statistically significant change when the spring transition begins (Fig. 1f) or ends (Fig. 1g). The start date for the preindustrial control (EXP1) is 31 March with a σ of ±13 d, and the end date is 11 May ±8 d. For the full RCP8.5 experiment (EXP2), the start date is 1 April with a σ of ±12 d, and the end date is 13 May ±13 d. Consistent with Fig. 1g, the enhanced jet variability due to RCP8.5 conditions manifests as a broader distribution of end dates. With no robust change in the timing of the spring transition, the calendar dates corresponding to the late-winter, spring, and summer jet phases are similar amongst the experiments. Therefore, in all subsequent figures, anomalies are calculated by binning each individual experiment's data according to that experiment's late-winter, spring, and summer days; time-averaging the data within each bin; and then differencing between the jet phase (e.g., late winter) bins from two different experiments (e.g., EXP2 minus EXP1). This approach would not be possible if, for instance, the annually averaged late-winter end date from EXP2 was 10 d after that from the EXP1. Similar results to those shown in Figs. 2–6 can be obtained by comparing like months (e.g., February–March, April–May) from two different experiments (not shown). However, we choose to show our results according to jet phase so that the STT inherently associated with each phase is accounted for.

Note that while no changes in the timing of the spring transition are found in these simulations, spring transition timing is heavily influenced by the El Niño–Southern Oscillation (ENSO; Breeden et al., 2021). Interannual SST fluctuations (which may arise, for instance, due to ENSO) are excluded from our experiments; hence, our results cannot comprehensively establish how RCP8.5 forcing modifies the timing of the spring transition.

To quantify the contributions of the residual advection, two-way isentropic mixing, and production and loss to the total O3S response, we calculate the terms in the transformed Eulerian mean (TEM) continuity equation for zonal mean tracer transport given by Andrews et al. (1987, Eq. 9.4.13) and discussed by Abalos et al. (2013). Daily data, time-averaged from the 6-hourly fields, are used to calculate each term. These terms are shown in Eq. (1): 1∂χ‾∂t+v∗‾∂χ‾∂y+w∗‾∂χ‾∂z=P-L+e-z/H∇⋅M, where overbars denote zonal averages, χ denotes the ozone concentration in parts per billion, P denotes chemical production and L chemical loss, H is the scale height equal to 7 km, y and x are the meridional and zonal Cartesian coordinates, z is log-pressure height, ∇ is the divergence operator, and M is the two-way isentropic mixing vector with meridional and vertical components given by Eqs. (2) and (3): 2My=-e-zHv′χ′-v′T′‾S∂χ‾∂z3Mz=-e-zHw′χ′+v′T′‾S∂χ‾∂y, where primes denote deviations from the zonal average, v and w are the meridional and vertical velocities, S equals (H⋅N2)/R in which N2 is the Brunt–Väisälä frequency, and R is the gas constant equal to 287 m2 s-2 K-1. The residual circulation velocities (v∗‾, w∗‾) are given by Eqs. (4) and (5): 4v∗‾=v‾-1ρ0∂∂zρ0v′θ′‾∂θ/∂z5w∗‾=w‾+1acosφ∂∂φcos⁡φv′θ′‾∂θ/∂z, where ρ0 is log-pressure density, θ is potential temperature, and a is Earth's radius.

To better understand how climate change may influence the amount of stratospheric ozone making it into the lower free troposphere over western North America, Fig. 2 shows the 700 hPa O3S responses to full RCP8.5 forcing, the change due to SSTs alone, and the change due to GHGs alone for the late-winter, spring, and summer North Pacific jet phases. In the preindustrial control climatology, lower-tropospheric O3S increases from low to high latitudes regardless of season, and mixing ratios are largest over the western North Pacific during the jet's spring phase, mimicking the observed seasonal maximum in deep STT over this region (Fig. 2 black lines; Škerlak et al., 2014; Breeden et al., 2021).

RCP8.5 forcing increases lower-tropospheric O3S over most of the longitudinal domain shown and over much of the hemisphere (not shown) during all three seasons (Fig. 2a–c). The RCP8.5 response is largest in late winter, during which there is up to a 50 % increase in O3S over the North Pacific and a 37 % increase over western North America (25–45∘ N, 235–260∘ E; Fig. 2a box). Although the responses are smaller in absolute magnitude during spring and summer compared to winter, they still imply roughly a 10 %–30 % change relative to climatology. Notably in spring, the largest increases are centered over western North America (Fig. 2b).

The SSTs alone (Fig. 2d–f) increase O3S by approximately 15 % over the eastern North Pacific during the jet's late-winter phase (Fig. 2d), explaining a portion of the aforementioned 50 % increase in late winter over this region due to full RCP8.5 forcing (cf. Fig. 2a). Over the low-latitude eastern North Pacific, close to Baja California, Mexico, the SSTs alone promote large increases in O3S during the jet's winter and spring phases relative to preindustrial climate (Fig. 2d–e). Conversely at high latitudes, O3S does not change during the late-winter phase in response to SSTs alone, and it decreases during both the jet's spring and summer phases. In summary, the SSTs alone can explain a portion of the full RCP8.5 response but clearly not the bulk of it.

The response to GHGs alone accounts for the majority of the full RCP8.5 700 hPa O3S response (Fig. 2g–i). Larger O3S increases develop during the jet's late-winter and spring phases compared to summer. Both SSTs alone and GHGs alone increase O3S over the eastern North Pacific and western North America during the jet's late-winter phase but have competing effects on O3S during the jet's spring and summer phases. To better understand the future changes in free tropospheric O3S and the relative roles of SST and GHG changes, the next sections consider in more detail how the North Pacific jet and the lower-stratospheric ozone reservoir respond to climate change.

RCP8.5 conditions accelerate, narrow, and elongate the late-winter North Pacific jet towards western North America at 200 hPa (Fig. 3a). This change is robust to varying severities of climate change (RCP4.5, Harvey et al., 2020; RCP6.0, Akritidis et al., 2019; and RCP8.5, Matsumura et al., 2021). Contrary to what takes place during the late-winter period, the subtropical jet shifts equatorward during the jet's spring and summer phases (Fig. 3b–c). At lower latitudes, westerly anomalies form over the subtropical eastern Pacific and central America, where there is a climatological minimum in the 200 hPa zonal wind (Fig. 3a–c). This response is present during all three jet phases and strengthens from late winter through summer (Fig. 3a–c).

The full RCP8.5 200 hPa zonal wind response is dominated by the contribution from the SSTs alone (Fig. 3d–f), with the GHGs alone (Fig. 3g–i) playing a comparatively minor role. The strong influence of the SSTs on the wind response arises in part because the SST forcing is associated with almost all of the ∼ 9–11 K warming of the tropical upper troposphere and the amplified Arctic surface warming (Fig. S1) and so dominates the influence on meridional temperature gradients and associated circulation changes that drive heat transport. Another consideration is that the zonal asymmetries in the pattern of SSTs prescribed in the experiments, particularly those over the tropical Pacific resembling El Niño (Fig. S2), may elicit teleconnections (e.g., Pacific North America (PNA) teleconnection pattern) that modify the upper-tropospheric circulation. In general, the impact of the GHGs alone on the 200 hPa winds is small, although the GHGs alone do have a large (compared to climatology) effect on the zonal wind over western North America during the jet's summer phase (Fig. 3i), illustrating that purely chemical changes in the stratosphere are capable of having significant dynamical impacts.

Figure 4 shows how RCP8.5 conditions modify 200 hPa O3S, allowing us to see both tropospheric and stratospheric ozone changes; at 200 hPa, the stratosphere is poleward of the anomalous thermal tropopause (blue lines), which can be compared with the preindustrial thermal tropopause (cyan lines) in each season. The 200 hPa O3S equatorward of the tropopause has already been transported into the troposphere and can be lost due to dry deposition, photolysis, and chemical loss or transported back to the stratosphere by reversible mixing processes.

In the preindustrial control (EXP1), O3S maxima and minima are co-located with the troughs and ridges of the climatological stationary wave (Fig. 4a–c). This is particularly clear in late winter, during which O3S mixing ratios exceed 600 ppb over the wave-1-scale trough of the climatological stationary wave, the Aleutian Low (Fig. 4a). O3S mixing ratios are, on the other hand, reduced over the climatological Alaskan ridge. Slightly out of view in Fig. 4a is a climatological wave-2-scale trough that resides over the Baffin Bay and Greenland; an O3S maximum is found over this region as well (Fig. 4a). As suggested by Reed (1950; see also Schoeberl and Kreuger, 1983; Salby and Callaghan, 1993), horizontal advection and vertical motion associated with waves act to concentrate ozone in troughs and reduce it over ridges. The climatological stationary wave influences the 200 hPa composition of O3S in this way.

Full RCP8.5 conditions increase lower-stratospheric O3S over much of the hemisphere during all seasons (Fig. 4d–f). The largest regional increase is a doubling of O3S over the North Pacific during the jet's late-winter phase (Fig. 4a, d). This regional O3S increase is co-located with the trough of an anomalous tropical–extratropical planetary-scale wave, whose signature is apparent in the zonal wind response (Fig. 3) and the stationary wave response (Fig. 4, black contours). As the amplitude of this wave diminishes during the spring and summer phases, so does the lower-stratospheric O3S maximum (Fig. 4e–f). The RCP8.5 O3S response is mostly contained in the lower-stratospheric (i.e., poleward of the tropopause) trough during the jet's late-winter phase, but in the absence of strong meridional potential vorticity gradients such as the high-latitude polar stratospheric westerlies (Manney et al., 1994; Salby and Callaghan, 2007) or the subtropical jet stream (Bönisch et al., 2009), which serve as transport barriers, the O3S response “smears out” during spring and summer, becoming more evenly distributed around the 200 hPa thermal tropopause (Fig. 4e–f).

The SSTs alone are almost solely responsible for the development of the anomalous planetary wave and are therefore a key reason why there are zonal asymmetries in the lower-stratospheric ozone reservoir (Fig. 4g–i). Similar effects of large-scale planetary wave trains on lower-stratospheric ozone have been noted in relation to ENSO (Zhang et al., 2015; Albers et al., 2022). The SST forcing considered in this study displays SST warming globally but contains some zonal asymmetries, one of them being an El Niño-like eastern tropical Pacific warming (Fig. S2). This zonal asymmetry may explain why the planetary wave response to the SSTs alone during late winter (Fig. 4g) resembles the PNA wave train known to develop with El Niño (albeit the Canadian ridge in Fig. 4g is displaced east relative to the PNA Canadian ridge). Note though that there is large inter-model and inter-generational (CMIP5 vs. CMIP6) spread in how ENSO responds to climate change (Beobide-Arsuga et al., 2021; Cai et al., 2022), suggesting that this planetary wave response could vary amongst climate models should it in fact be related to the El Niño-like warming superimposed on the global SST increase (Fig. S2).

Contrary to the SSTs alone, the GHGs alone have little effect on the planetary-scale eddies and elicit more zonally symmetric O3S responses (Fig. 4j–l). The lower-stratospheric O3S response to the GHGs alone develops largely due to net chemical production of stratospheric ozone, likely associated with the large methane increase in RCP8.5, which enhances O3 mixing ratios in the extratropical stratosphere (Morgenstern et al., 2018), and changes in transport associated with the BDC's deep branch.

To further clarify how the lower-stratospheric reservoir responds to RCP8.5 conditions, Fig. 5 shows latitude–pressure transects of O3S anomalies and isentropes averaged between 235 and 260∘ E (over western North America; the same longitudinal bounds used for the box in Fig. 2a). Climatologically, extratropical lower-stratospheric O3S mixing ratios are larger during winter and spring (Fig. 5b), following from transport by the BDC's deep branch (Ray et al., 1999; Hegglin and Shepherd, 2007; Bönisch et al., 2009; Butchart, 2014; Konopka et al., 2015; Ploeger and Birner, 2016; Albers et al., 2018). During summer in climatology, enhanced isentropic mixing between the tropical and extratropical lowermost stratosphere (Hegglin and Shepherd, 2007; Abalos et al., 2013) and rising tropopause heights (Schoeberl et al., 2004) act to flush ozone out of the lowermost stratosphere.

During every jet phase, RCP8.5 conditions reduce O3S in the low-latitude stratosphere while promoting accumulation of O3S at high latitudes (Fig. 5d–f). Some of this O3S accumulating in the extratropical lower stratosphere may enter the troposphere along the subtropical upper-tropospheric and lower-stratospheric isentropes (e.g., 360 K). Both the GHGs alone and SSTs alone play a role in making this happen. The upper-tropospheric warming induced by the SSTs alone depresses the isentropes (e.g., 360 K) to lower altitudes, enhancing the access of the troposphere to lower-stratospheric air (Fig. 5g–i), where wave breaking is able to transport the ozone into the subtropical and tropical upper troposphere (e.g., Waugh and Polvani, 2000; Albers et al., 2016, and references therein). The GHGs alone on the other hand mainly contribute by more broadly enhancing the extratropical lower-stratospheric O3S concentrations (Fig. 5j–l).

O3S is reduced near the extratropical tropopause in all seasons in response to RCP8.5 forcing (Fig. 5d–f). This is associated with the increased height of the tropopause (Abalos et al., 2017) resulting from the SSTs alone. Due to steep vertical gradients in tracers near the tropopause (e.g., Pan et al., 2004), taking the difference between an experiment with a lifted tropopause (EXP2 or EXP3) and an experiment without this feature (preindustrial control, EXP1) amounts to taking the difference between relatively O3S-depleted tropospheric air and O3S-rich stratospheric air; hence, the negative O3S anomalies develop near the tropopause (Fig. 5d–i). This negative O3S response can largely be removed by remapping the vertical axis of each data field used to make, for instance, Fig. 5d–f (zonally averaged RCP8.5 (EXP2) O3S and preindustrial (EXP1) O3S) to tropopause-relative coordinates (meters above or below the thermal tropopause), then taking the difference between these two modified data fields and remapping this set of anomalies (axes: tropopause-relative × latitude) to a log-pressure coordinate system (axes: pressure × latitude) (Abalos et al., 2017). Using annual cycles of thermal tropopause and O3S data, which should help to smooth out the large hourly/daily fluctuations in these fields near the tropopause, the aforementioned procedure was applied to a zonally averaged transect over the North Pacific (Fig. S3) and applied at all grid points at 200 hPa (Fig. S4) and 300 hPa (Fig. S5). While this tropopause-relative analysis does remove the majority of the negative O3S response associated with the tropopause lift, the strong O3S zonal asymmetries associated with the planetary wave response to the SSTs alone persist, namely the negative O3S response corresponding to the planetary wave's ridge near Alaska (cf. Figs. 4k, S5e). This analysis corroborates that the higher tropopause in RCP8.5 is largely responsible for the presence of the negative O3S response in the extratropical upper troposphere and lower stratosphere, however, not entirely, as we find that a portion of this negative O3S is associated with the anomalous planetary wave's zonally asymmetric effects on the upper-tropospheric and lower-stratospheric O3S distribution.

The seasonal variability of both tropical (Abalos et al., 2013) and extratropical (Albers et al., 2018) lower-stratospheric ozone tendencies is heavily influenced by upwelling and downwelling associated with BDC's residual mean meridional circulation component. This circulation is made up of a shallow and a deep branch. Transport associated with the shallow branch proceeds more horizontally, and the air masses enter the stratosphere closer to the subtropics, whereas transport associated with the deep branch is more vertical, and the air masses enter the stratosphere through the deep tropics and descend at high latitudes (Birner and Bönisch, 2011). To quantify the influence of RCP8.5 forcing on these physical processes, Fig. 6 shows the residual mass streamfunction response to RCP8.5 forcing in black contours and in shading the local changes in O3S tendencies as a result of transport by the residual mean meridional circulation terms in the TEM continuity equation (v∗‾∂χ‾∂y+w∗‾∂χ‾∂z). As in reanalysis (cf. Rosenlof, 1995), in the preindustrial control, the tropical upward mass flux peaks in amplitude during boreal winter when the residual mass streamfunction is strongest (Fig. 6a). As the zonal momentum budget changes in each hemisphere during spring and summer, the tropical upward mass flux shifts into the Northern Hemisphere, and the residual mass streamfunction weakens and shifts downward towards the troposphere (Fig. 6b, c). The negative O3S tendencies in the tropical lower stratosphere track the latitudinal shifting of the tropical upward mass flux over time. The positive O3S tendencies in the extratropical lower stratosphere associated with poleward transport of stratospheric ozone from its tropical source region peak in amplitude during winter when the BDC's deep branch is strongest and weaken thereafter.

RCP8.5 forcing strengthens the shallow branch of the BDC during all three seasons, reducing tropical stratospheric O3S tendencies (Fig. 6d–f). The SSTs alone (Fig. 6g–i) are primarily responsible for the acceleration of the residual mass streamfunction in the subtropical lower stratosphere (50 hPa/30∘ N) when compared against the GHGs alone (Fig. 6j–l), consistent with Oberländer et al. (2013) and Chrysanthou et al., (2020). The upper component of the Hadley circulation near 150 hPa and 15∘ N accelerates, as previously reported by Abalos et al. (2020). All models that they studied included this response. This feature acts cooperatively with the reinforced BDC shallow branch to increase O3S transport through the subtropical tropopause into the upper troposphere (200 hPa and 30∘ N), with the largest increase occurring during summer in response to the SSTs alone (Fig. 6i). The GHGs alone accelerate the deep branch well above 30 hPa during winter (Fig. 6j); the deep branch's high-latitude downwelling increases lower-stratospheric O3S during spring (Fig. 6k) and then disappears by summer (Fig. 6i).

Another aspect of the BDC is two-way isentropic mixing, which climatologically increases subtropical O3S tendencies above and south of the subtropical jet, while reducing extratropical O3S tendencies throughout the stratosphere (Fig. 7a–c). In the tropical lower stratosphere (∼ 80 hPa), tendencies peak during summer in present-day analyses (Abalos et al., 2013) and in the preindustrial control climatology (Fig. 7c). RCP8.5 forcing generally reinforces the climatological two-way isentropic mixing in the stratosphere during each season, increasing subtropical tendencies and reducing extratropical tendencies (Fig. 7d–f). Additionally, enhanced cross-tropopause mixing by eddies increases upper-tropospheric O3S tendencies from 30–60∘ N, with stronger signals during summer than winter. These anomalies are primarily associated with the SSTs alone (Fig. 7g–i). Hardly any part of the two-way isentropic mixing responses to GHGs alone are statistically significant (Fig. 7j–l).