II. S ENSITIVITY OF THE T ROPICAL P RECIPITATION R ESPONSE TO
2.6 S UMMARY AND D ISCUSSION
to shift the ITCZ in the equilibrium (Chiang and Bitz, 2005; Deser et al. 2015). However, the amplified seasonality of Arctic sea-ice under global warming (Haine and Martine 2017) would likely have no impact on the tropical precipitation, given that the mixed layer depth is 70 m in the northern extratropics (e.g., de Boyer Montégut et al. 2004). By contrast, a multi- decadal AMOC variability should effectively induce a corresponding meridional ITCZ variability (e.g., Delworth et al. 2017). Our results highlight the important role of decadal- and-longer extratropical climate variability in shaping the tropical climate, a potential precursor of tropical precipitation. For instance, it is suggested that the AMOC has predictability on decadal time scales (Tulloch and Marshall 2012) even though the amplitude of AMOC variability is usually underestimated by current climate models (Yan et al. 2018).
We also show that the transient response to extratropical climate variability depends on cloud radiative responses. The cloud radiative effects are known to affect the equilibrium tropical response to extratropical forcing (Kang et al. 2008; Zhang et al. 2010), and here we show that the transient behavior is also modulated by the cloud response. The cloud radiative response outside the deep tropics serves as a positive feedback, thereby accelerating the equatorward propagation of extratropical signal. By contrast, the cloud radiative effect in the deep tropics serves as a negative feedback, offsetting the tropical response. As a result, the locked cloud experiment exhibits a slower propagation of extratropical signal into the tropics and a more amplified ITCZ shift response compared to the interactive cloud experiment.
However, the negative CRE in the tropics is model-dependent, as pointed out in Kang et al.
(2014). This suggests that uncertainties in cloud modeling can create uncertainties in transient regional responses to climatic perturbations. For example, large uncertainties in cloud radiative effects may give rise to an uncertainty in how extratropical climate variability, such
translates into the tropics. Particularly, we think this dependency gives insights into the notorious precipitation bias in the tropics, double-ITCZ. Recent studies consistently beg the question of how much the Southern Ocean bias (or any extratropical bias) contributes to the tropical bias. An eight fully coupled climate models consistently present propagated trails through the eastern boundaries of the Pacific Ocean (Kang et al. 2019), where even state-of- the-art climate models cannot estimate the strength of stratocumulus feedback in the pathway.
If the stratocumulus feedback acts like that of our idealized configuration, the strength of stratocumulus feedback would highly control the overall mutuality between the southern extratropics and tropical climate, eventually able to induce our huge problem, the double- ITCZ bias. Hence, this dependency on the cloud response suggests that our results should be tested in other GCMs.
Acknowledgements
We thank Doyeon Kim, Hanjun Kim, Kiwoong Park, Yen-Ting Hwang, Dargan Frierson, Malte Stuecker, Young-Min Yang, and In-Sik Kang for insightful comments. Ken Takahashi provided solid and thoughtful suggestion from an earlier analysis. Comments from four reviewers, including Aaron Donohoe and Spencer Hill, greatly improved the paper. Y. S. was funded Global Ph.D. fellowship (2017H1A2A1044044) by the National Research Foundation of Korea (NRF).
Chapter 3
The Tropical Response to Periodic Extratropical Thermal Forcings: A three-step sequential mechanism
Abstract
In this chapter, we reveal the fundamental mechanism that governs the extratropics-to-tropics teleconnection. When extratropical forcing grants sufficient time scale, the atmospheric transient eddies diffusively transport surplus energy in to the midlatitude, perturbing SSTs outside the forcing region, as demonstrated by a one-dimensional energy balance model with a fixed diffusivity. As the transient eddies weaken in the subtropics, a further equatorward advection is accomplished by the Hadley circulation. The essential role of Hadley cell advection in connecting the subtropical signal to the equatorial region is supported by an idealized thermodynamical-advective model. Dynamic transition between eddy and mean- circulation explains the threshold behavior of extratropics-to-tropics teleconnection. The newly-revealed mechanism suggests climatological wind properties as a key feature which determines zonal tropical response.
3.1 Introduction
This chapter investigates the sequential mechanism of extratropics-to-tropics teleconnection.
The previous chapter shows that the temporal condition sufficiently controls the extratropics- to-tropics teleconnection, mainly accompanied with changes of lower boundaries. Therefore, a mechanism that governs the SST propagation would explain the sensitivity of extratropics- to-tropics teleconnection. We utilize the advantages of periodic time-varying forcing that allows objective identification of multiple processes based on lead-lag phase comparison.
Also, we build and adapt simplified models as fundamental tools to describe outlined atmospheric dynamics.
While most modeling studies examine the equilibrium tropical response to a time- invariant extratropical forcing, some studies have evaluated the transient features of extratropics-to-tropics teleconnection. Chiang and Bitz (2005) proposed the wind- evaporation-SST (WES) feedback as the dominant mechanism for the equatorward propagation of the high-latitude signal. This would be quite similar the role of WES feedback on the evolution of Pacific Meridional Mode (PMM) (e.g., Amaya 2019). However, suppressing the WES feedback had little impact on the tropical response to extratropical forcing in different climate models (Mahajan et al. 2011; Kang et al. 2012; Kang et al. 2014).
Their modification is made in the bulk aerodynamic formula of surface turbulent fluxes over the ocean. The parameterizations of surface turbulent heat fluxes is governed by following equations:
SHF = {4;|}U X(>*(?Y− ?) and LHF = {4;|}U X(@)(AY− A) (3.1) with ;|U the drag coefficient, {4 air density, AY the surface saturation specific humidity, A the specific humidity of the atmospheric bottom level, >* the specific heat capacity of moist
air, @) the latent heat of vaporization of water, and }X( the wind speed at 10 m reference height. The }X( is nudged or prescribed to the climatology, resulting in breaking the WES feedback. Nonetheless, if the energetic imbalance is stably imposed in the extratropics, atmosphere tends to compensate the interhemispheric energy contrast, resulting in ITCZ shift even without surface wind speed changes. We emphasize that the WES feedback would take sufficient role in the extratropics-to-tropics teleconnection for controlling magnitude.
However, it is obvious that the WES feedback is neither essential nor fundamental mechanism for the tropical responses to the extratropical changes. If it is essential, such as ocean dynamics for ENSO, the tropics would not response to the extratropical forcing.
The energetic theory has been developed over the past two dacades to understand the meridional position of ITCZ (e.g., Kang 2020), which is a main consequence of the extratropics-to-tropics teleconnection. The main idea is that an energetic imbalance in one hemisphere necessitates an anomalous cross-equatorial energy transport into the hemisphere with net energy loss. This is accomplished with baroclinic eddies spreading high-latitude energetic imbalance toward the tropics, followed by shifted ascending branch of tropical general circulation, leading to net energy transport following the upper branch to the hemisphere. Although the energetic framework is proven useful for interpreting the ITCZ response to radiative perturbations, it is diagnostic expression of cause-and-effect relationship rather than prognostic illustration of propagation process, providing little information about the teleconnection itself, repeatedly pointed in previous studies (e.g., Chapter 4 in Chiang and Friedman 2012; Emerging Challenges in Kang 2020). The energetic constraint itself hinted the possible roles of transient eddies in the extratropics and the Hadley circulation in the tropics based on their relative importance in transporting the moist static
in an experiment where energy flux perturbations are abruptly introduced, it is often difficult to disentangle the sequential order of multiple dynamical processes at work without a large number of ensemble members. Therefore, despite its implication for real world, a mechanistic theory of the teleconnection from extratropical cooling to the tropical ITCZ shift and Hadley circulation is yet to be fully understood.
In the previous chapter, we show a temporal requirement of the extratropics-to- tropics teleconnection by considering the transient facet of the energetic framework. The energetic imbalance in extratropics is almost completely resolved by local processes (i.e., local atmospheric dynamics and radiative response) when it fluctuates with high-frequency.
The compensation is hardly related to the tropical climate, hence with muffled ITCZ shift.
The ITCZ response happens when the energetic imbalance occurs in low latitudes that tropical general circulation compensates. And, with sufficient time scale, the energetic imbalance emerges in the tropics with SST responses. To understand the dependency, we set out to understand how the SST propagates into the tropics, the sequential processes of extratropics-to-tropics teleconnection. For this purpose, our experiment setup with periodical forcing is particularly helpful as that allows us to objectively differentiate the sequential order of different physical processes based on the phase lag between the extratropical forcing and the response of any variable of interest.
The chapter is organized as follows. In section 3.2, we again outline overall strategies to examine the sequential mechanism of extratropics-to-tropics teleconnection. Here, we mainly explain idealized and theoretical models that we adapt to underpin how atmospheric dynamics govern the propagation. In section 3.3, the propagation mechanism is suggested.
An energy balance model suggests that diffusive eddies are in charge of an equatorward SST propagation in the mid-latitudes. As the eddy activities weaken in the subtropics, the Hadley
circulation takes over to advect moist static energy into the deep tropics, as demonstrated by an idealized thermodynamical-advective model. Then we additionally discuss the intermediate links between governing dynamics offers conclusions in section 3.4.
3.2 Methodology
In this chapter, we mainly analyze the series of the experiment that impose time-varying surface heat flux into extratropics, summarized in Table 2.2.
3.2.1 Energy Balance Model
To delineate sequential mechanism of extratropics-to-tropics teleconnection, we firstly conduct an Energy Balance Model (EBM). As indicated by their name, the EBM try to estimate the changes in the climate system from the energetic balance of the Earth. The EBM was introduced by both Budkyo (1969) and Sellers (1969), continually employed in many different climate studies: sea-ice and climate (e.g., Wagner and Eiesnman 2015), extratropical changes and energy transport (e.g., Hwang and Frierson 2010), storm tracks (e.g., Mbengue and Schneider 2018), and Hadley cell and tropical hydrologic cycle (e.g., Siler et al. 2018). Particularly, Kang et al. (2009) employ the EBM based on the diffusive nature of eddy fluxes, reproducing energetic transport to the extratropical forcing. Similarly, we use a one-dimensional EBM to predict the mid-latitude response. We first remind the atmospheric and oceanic energy is governed by the following equations, respectively:
$
$%〈>&? + @'A〉 = 9!"# − .:; + ∇ ⋅ :#, (3.2)
;V$%$ ?Y = .:; + b, (3.3)
Where the net downward top-of-atmosphere (TOA) radiation 9!"#, the net downward surface energy flux .:;, the angle bracket denoting the vertical average, imposed surface flux b, ;6= 2 × 10Z J K-1 m-2 is the specified heat capacity of 50 m oceanic mixed layer,
and ?2 is the surface temperature.
The meridional MSE flux at any given latitude (∇ ⋅ :#) is proportional to the surface MSE gradient (see e.g., Hwang and Frierson 2010; Wagner and Eisenman 2015), hence meridional energy flux is parametrized to −&[0Ä∇+Å2, where the surface MSE is a function of surface temperature Å2 = >*?2+ @)ℋA2∗ with the specific heat of dry air >* = 1004 J K−1 kg−1, the heat of vaporization @) = 2.5 × 10\ J kg-1, the surface pressure J2 = 1 × 10] Pa the surface relative humidity ℋ (fixed as 76 %), and the saturation specific humidity A2∗, calculated using ?2 in the Clausius–Clapeyron equation. The diffusion coefficient Ä = 1.2 × 10\ m s-1 is tuned to best fit the atmospheric energy flux :# from the reference experiment. Combining Eq (3.2) and (3.3) with ignoring the atmospheric energy tendency gives a governing equation of EBM:
;6$A$%0= . − @^_`+ @^_a+ b −&[0Ä∇+Å2, (3.4) where . is the net incoming shortwave radiation at TOA, @^_` is the clear-sky outgoing longwave radiation (OLR), @^_a is the longwave CRE (OLR in clear sky minus OLR in all sky). The clear-sky OLR is parametrized by the least squares regression of ?2 at each latitude from the reference experiment. As we prescribe ., @^_a, and b from the corresponding experiment, Eq. (3.4) can be solved for ?2. Figure 3.1a shows temporal evolution of atmospheric energy transport predicted by EBM, reaching steady-state condition.
The EBM appropriately predicts climatological meridional energy transport (Fig. 3.1b).
Figure 3.1. (left) Time evolution of meridional energy transport predicted by EBM. Positive indicates northward transport. Unit is PW. (right) Climatological energy transport in AM2 (solid) and prediction of EBM (dashed).
3.2.2 Idealized Thermodynamical-Advective Model
We also adopt an idealized thermodynamical-advective model (Takahashi et al. 2007a,b) to demonstrate the importance of mean meridional advection of MSE anomaly in the lower troposphere for the equatorward propagation of the SST response from the subtropics (i.e., the Hadley circulation terminus). The simple model represents the coupled lower atmosphere-upper ocean system (Fig. 3.2). The atmospheric mixed layer (AML) is coupled to an ocean mixed layer (OML), interacting via sensible (SHF) and latent heat fluxes (LHF).
The AML and OML are governed by the following equations, respectively:
'$b
$c = SHF + LHF + ℛ + ℳ, (3.5)
;6$A+
$% = −SHF − LHF − ℛ, (3.6) where ' is the AML meridional wind, ℛ is the net radiative damping, ℳ is the damping
through entrainment from the free troposphere, and ?2 is the OML temperature. The moist static energy in the AML is Å ≡ N@)A + >*?P{4ℎ4, with the air density {4 = 1.1 kg m-3 and the AML height ℎ4 = 1 km. We specify the OML heat capacity to be ;6= 2 × 10Z J K-1 m-2 , as in our comprehensive model AM2. Note that we ignore the time tendency of Å in Eq. (3.5) as the AML is assumed to be in a quasi-equilibrium state associated with a small heat capacity.
Figure 3.2. Schematic of idealized thermodynamical-advective model
3.3 Propagation Mechanism
The SSTs outside the forcing region are only a function of net surface energy flux determined by the local air-sea interaction in a slab ocean model; thus, we first examine how the atmospheric moist static energy (MSE) propagates away from the forcing region. Figure 3.3a shows the lag time at which the correlation coefficient is maximized between the zonal-mean tropospheric MSE anomalies and the Qflux forcing in the standard experiment with ? = 10 years (E10). This time lag distribution in latitude and pressure illustrates the zonal-mean MSE propagation pattern as a smaller time lag indicates that the MSE response is established faster at a given location. The impact of extratropical forcing propagates upward and equatorward simultaneously. On the equatorward side of the forcing edge (40°N), the fastest propagation takes place near the surface, implying that the most preferential pathway for extratropical perturbations to reach the tropics is through the lower troposphere. This propagation pattern is largely shared by other experiments in which the extratropically forced teleconnection emerges (Fig. 3.4). The rapid lower tropospheric propagation can be further confirmed by a Hovmöller diagram of the zonal-mean MSE response at different vertical levels (Figs. 3.3b- d). The MSE response at 850 hPa propagates equatorward from the extratropical forcing region (Fig. 3.3d) whereas at higher levels the tropical response emerges earlier than the extratropical response, indicative of a poleward propagation (Figs. 3.3b-c). That is, the mid- to-upper tropospheric MSE in the subtropics responds only after the near-surface propagation reaches the deep tropics.
Figure 3.3. (a) The lag time for which the correlation coefficient is maximized between the tropospheric zonal-mean MSE response and the Qflux forcing as a function of latitude and height (shading; year) in E10. The time-mean meridional stream function (clockwise circulation in solid and counter-clockwise circulation in dashed; interval = 0.5×1011 kg s-1) and the equivalent potential temperature (red contour; interval = 10 K from 270 K) are from the reference experiment. The purple arrow illustrates the propagation pattern of zonal-mean MSE anomalies. The composite across the forcing cycle of the zonal mean MSE response at (b) 300 hPa, (c) 500 hPa, and (d) 850 hPa. The equatorial edge of the forcing region (φ) is indicated by a cyan line.
Figure 3.4. Same as Fig. 3.3a, but for the (a) E1, (b) E1_10m, (c) E1_10m_5A, (d) E3, (e) E5, and (f) E10_CEF experiments.
Figure 3.5. Time lag at which the lag correlation between the anomalous 900 hPa MSE at φ0 and each latitude attains the maximum value in E3 (gray), E5 (dark gray), E10 (black), E10_CEF (purple),
Since the near-surface propagation is the preferential pathway for extratropics-to- tropics teleconnection, we examine the lower atmospheric MSE propagation in more detail.
Figure 3.5 shows the time lag of the composite zonal-mean MSE response at 900 hPa, representing the phase line of lower tropospheric MSE propagation. Based on the propagation speed, estimated by the slope of the phase line, we identify three distinct regimes for the extratropics-to-tropics teleconnection: the mid-latitude eddy regime, the subtropics transition regime, and the Hadley cell regime.
3.3.1 Mid-Latitude Eddy Regime
Previous studies consistently suggest the importance of atmospheric eddy in transporting energy meridionally (e.g., Kang et al., 2009). We also decompose the atmospheric MSE flux into the mean and eddy components (see e.g., Hill et al. 2015; Xiang et al. 2018). While the two components act in the opposite direction, the meridional MSE flux in the mid-latitude is dominated by the eddy component in all experiments (Fig.3.6a). As the EBM is adequate to represent diffusive nature of eddy fluxes, we adopt it to support the role of eddy energy transport as a first step of extratropics-to-tropics teleconnection.
Figure 3.7 compares the temporal evolution of the MSE flux response !"! and the surface temperature response !#" in all standard experiments and its prediction by the EBM.
Even with the same diffusivity in the reference and perturbed experiments, both the magnitude and the phase of !"! between 30°N-40°N are well predicted in the extratropics (Fig. 3.7a). The EBM also captures the sensitivity of the magnitude of extratropical SST variability to the forcing period (Fig. 3.7c). That is, larger SST fluctuations are predicted for lower-frequency forcing, albeit weaker in magnitude by about 20 %. The overall success of EBM indicates that the eddy diffusivity is not changing but they are acting diffusively on an
altered MSE gradient. The agreement between AM2 and EBM implies that an equatorward propagation within the extratropics is primarily accomplished by the eddy energy transport.
In response to lower-frequency forcing, the eddies diffuse energy in one direction for a longer period, allowing for a larger SST response to develop (Fig. 3.7c).
Figure 3.6. Composite across the forcing cycle of the MSE flux response δFA (solid), and its decomposition into the mean (dashed) and eddy (dotted) components averaged over (a) 30°N-40°N and (b) 10°S-10°N for the E1 (light gray), E3 (gray), E5 (dark gray), and E10 (black) experiments. Units are PW.
Figure 3.7. Composite across the forcing cycle of the MSE flux response δFA (in PW) averaged over (a) 30°N-40°N and (b) 10°S-10°N, and (c) the latitudinal distribution of root-mean-square of δTS composite (in K) for the E1 (light gray), E3 (gray), E5 (dark gray), and E10 (black) experiments in the AM2.0 (solid) and EBM (dashed).
Overlaid in Fig. 2.3 is the eddy heat transport response composite at its peak level of 900 hPa. Anomalous eddy heat transport is equatorward during the period of extratropical warming, while it is poleward during the period of extratropical cooling. The eddy heat transport response decreases with the forcing period, by a factor of 3 from E1 to E10, associated with smaller meridional gradients in SST anomalies. Despite the weaker eddy heat transport response for lower-frequency cases, the SST response shows a clearer equatorward propagation into the deep tropics. This is because the duration of the anomalous eddy heat transport convergence with the same sign (which corresponds to half of forcing period)