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  • View in gallery

    Downstream development during the ET of Supertyphoon Nuri (2014) at (a) 0600 UTC 5 Nov, (b) 1800 UTC 6 Nov, (c) 1800 UTC 7 Nov, and (d) 0000 UTC 9 Nov 2014. Panels shows IR Gridsat clouds [brightness temperature in °C as in color bar; Knapp et al. (2011)], dynamical tropopause [2 PVU on the 330-K isentropic surface (1 PVU = 106 K kg−1 m2 s−1); red contour], and wind speed maxima highlighting jet streaks (wind speed on the 330-K isentropic surface as semitransparent shading in yellow for 55 and orange for 65 m s−1). Dynamical tropopause and wind speed are taken from ERA-Interim. The TC symbol indicates the position of the transitioning cyclone (encircled during extratropical stage) and the “L” the position of the downstream cyclone.

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    Overview of synoptic features and processes involved in Northern Hemispheric ET. Labels indicate relevant processes, starting with the section in which they are discussed. Transitioning cyclone presented by the black and white pictogram. The dark red line indicates axis of the undulating midlatitude jet stream separating stratospheric high PV air (dark gray; poleward) and tropospheric low PV air (light gray; equatorward), with the dashed line indicating an alternate configuration. The orange ellipse denotes the jet streak. The purple, semitransparent area signifies the forecast uncertainty for the downstream flow. The downstream cyclone is indicated by the “L” symbol and its associated fronts. The airplane symbol represents observation systems used for ET reconnaissance.

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    PV anomalies involved in ET and their respective contribution to RWP modification near ET. The (a) 2D and (b) 3D schematics of the different flow features associated with ET. 1) Cyclonic circulation (thin orange arrows) associated with the cyclonic PV tower [orange TC symbol in (a), gray column in (b)] and with low-level cyclonic PV anomalies (small gray clouds and circulation) at the developing warm front. 2) Anticyclonic circulation associated with the anticyclonic PV anomaly of the upper-tropospheric outflow [white/gray cloud symbol in (a),(b) and thin blue arrows]. 3) Upper-tropospheric divergent outflow associated with latent heat release below (thin green arrows). The advective contribution of these flow features to the amplification of the downstream ridge and trough are given by the bold arrows in (a). The red contour and shading in (a),(b) are similar to Fig. 2. The lower layer in (b) exemplifies the developing warm front with colder air masses toward the pole. (c) Contributions to ridge amplification from the advection of potential temperature on the dynamic tropopause (a surrogate for upper-level PV advection) by the several flow features, integrated over a ridge for a 72-h forecast of an idealized scenario [in K km2 s−1; colors as in (a),(b)]. [Figure 8 from Riemer and Jones (2010), with modifications.]

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    Jet streak formation during ET from an energetics perspective. (a) Schematic representation, showing midlatitude jet (black line), developing Ke maxima (jet streak; gray ellipses), baroclinic conversion of Ke (clouds), ageostrophic geopotential flux (orange arrow), and its divergence (blue ellipses) and convergence (red ellipses). (b),(c) TC-relative composite of Ke budget for western North Pacific ETs, based on ERA-Interim for 1980–2010 [after Quinting and Jones (2016), their Figs. 12a,b]: vertically integrated Ke (shaded in 105 J m−2), 200-hPa geopotential (contours every 200 m2 s−2; thick black contour illustrates 11 800 m2 s−2), and (b) ageostrophic geopotential flux (vectors, reference vector in 106 W m−1; divergence as colored contours every 8 W m−2, divergence in blue, 0 W m−2 omitted) and (c) vertically integrated baroclinic conversion of Ke (red contours every 8 W m−2). The black box approximates the area that is captured by (b),(c). Composites are shown relative to the mean TC position.

  • View in gallery

    Lagrangian trajectories for the ET of Typhoon Jangmi showing ridge building and jet streak formation at (a) 0000 UTC 30 Sep and (b) 1200 UTC 1 Oct 2008. Shown are the 1.5-PVU surface (blue shading), 320-K surface of equivalent potential temperature (transparent gray shading), representing 3D baroclinic zone; |V| = 60 m s−1 (green shading), highlighting upper-level midlatitude jet; potential temperature at 990 hPa (shading at bottom; brown colors > 300 K, green ≈ 290 K) and geopotential height at 990 hPa (black contours; every 25 dam). Paths of representative trajectories [(a) 1200 UTC 28 Sep–1200 UTC 30 Sep 2008 and (b) 1200 UTC 30 Sep–1200 UTC 2 Oct 2008] colored by PV of air parcel moving along trajectory. Anticyclonic PV air (PV < 0.6 PVU) in gray shades and cyclonic PV values (PV > 0.6 PVU) in red shades (see legend in bottom right). [Figure 5 from Grams et al. (2013a).]

  • View in gallery

    Idealized scenario for steering flow topology during ET. Midtropospheric geopotential to illustrate the midlatitude wave pattern and winds in the frame of reference moving with this pattern (at 620 hPa; shaded and arrows, respectively) and different tracks of transitioning cyclones from sensitivity experiments (thin black lines). The bifurcation points associated with the upstream trough and the developing downstream ridge, deciding upon the development of a “no ET,” an “NW pattern,” or an “NE pattern” scenario, are marked by the dot and the cross, respectively. Dashed contours depict the streamlines that emanate from the bifurcation points. [Figure 11 from Riemer and Jones (2014).]

  • View in gallery

    Three-dimensional schematic depiction of the preconditioning stage with a PRE during western North Pacific ET anticyclonic PV air in the upper-level outflow of a TC and associated PRE as blue shading in the upper panel, and jet streak as green shading and 200-hPa waveguide as red contour separating high PV air (>3 PVU; orange shading) from lower PV air (<3 PVU; unshaded). Midlevel baroclinic zone as blue tilted surface. Trajectories of rapidly ascending air parcels as blue–red–blue lines, reflecting the diabatic PV modification of the parcels from low to high to low PVU, respectively. Mean sea level pressure (gray contours; every 8 hPa) and equivalent potential temperature (violet contours; 320 and 330 K) are indicated in the lower panel. [Figure 11a from Grams and Archambault (2016).]

  • View in gallery

    Conceptual model of the key synoptic-scale features during the occurrence of a PRE. Shown is the 200-hPa geopotential height (contours), the upper-tropospheric jet streak (gray shading), the midlatitude baroclinic zone [surface frontal structure with 925-hPa streamlines associated with warm (red) and cold (blue) air advection], and lower-tropospheric moisture flux along the eastern side of the recurving TC and the western flank of a subtropical high. The formation of the PRE is indicated by the yellow–green ellipse. [Fig. 20a of Moore et al. (2013).]

  • View in gallery

    Downstream development during the ET of Typhoon Nabi (2005). Vertically integrated Ke (shaded; 105 J m−2) and 500-hPa heights (light gray contours; 60-m intervals) and (a) vertically integrated divergence (dashed) and convergence (solid) of the ageostrophic geopotential flux (contours; W m−2); (b) vertically integrated total Ke flux vectors (advection + dispersion; reference vector in lower right; 105 W m−1); and (c) vertically integrated baroclinic conversion (contours; W m−2). [Figures 3a–c from Harr and Dea (2009).]

  • View in gallery

    Interaction between a transitioning cyclone and the midlatitude flow expressed as advection of low PV air by the upper-level divergent outflow. (a) Idealized representation of ridge amplification and jet streak intensification. Vectors represent the upper-tropospheric divergent outflow associated with the transitioning cyclone. Shading denotes anticyclonic PV advection by the divergent wind (Archambault et al. 2013, their Fig. 4). Composite analyses of objectively defined (b) strong and (c) weak interactions at the time of maximum interaction. The 500-hPa ascent (green; every 2 × 10−3 hPa s−1, negative values only), total column precipitable water (shaded according to grayscale; mm), 200-hPa PV (blue; every 1 PVU), irrotational wind (vectors, > 2 m s−1; purple vectors, > 8 m s−1), negative PV advection by the irrotational wind (dashed red; every 2 PVU day−1 starting at −2 PVU day−1), and total wind speed (shaded according to color bar; m s−1). The star denotes point of maximum interaction. The TC symbol denotes composite TC position. Downstream development of (d) strong and (e) weak interactions 36 h after time of maximum interaction as represented by 250-hPa meridional wind anomalies (shaded; m s−1; enclosed by black contours where significant at the 99% confidence level), PV (blue; every 1 PVU), and irrotational wind (vectors; >2 m s−1). [Figures 8a,b, 5d, and 6d from Archambault et al. (2015).]

  • View in gallery

    Recurvature-relative composites of enhanced RWP frequency anomaly (shaded in %) (a) relative to June–November climatology for western North Pacific transitioning cyclones and (b) relative to December–April climatology for south Indian Ocean transitioning cyclones. Statistical significance at 95% confidence level hatched, mean track given by black line. Mean and range of recurvature longitudes indicated by white star and black bar, respectively. Data smoothed with a Gaussian filter. [Figures 3a and 5a from Quinting and Jones (2016).]

  • View in gallery

    (a) Illustration of omega block and high-impact weather downstream of Typhoon Nuri (2014) after Bosart et al. (2015). (b) Downstream impact of Typhoon Choi-Wan (2009), based on NWP experiments where the storm has been removed from initial conditions (Keller and Grams 2015). Black items represent midlatitude flow features in the presence of ET, and red items the evolution if ET influences were not present: 300-hPa geopotential height contour indicates upper-level waveguide (950 dam at 0000 UTC 22 Sep 2009). Arrows indicate shift of high-impact weather (precipitation, sunny and hot conditions, cold conditions) with symbol size representing magnitude.

  • View in gallery

    Vertically integrated initial-time dry total energy singular vector sensitivities (shaded; values in color bar: J kg−1) with 500-hPa streamlines for TC Shanshan from 0000 UTC 15 Sep 2006. [Figure 3c from Reynolds et al. (2009).]

  • View in gallery

    Increase in standard deviation of the 500-hPa geopotential height (in dam) in the Australian, the Canadian, the ECMWF, and the TIGGE (Swinbank et al. 2016) multimodel EPS for the ET of Hurricane Ike. Forecast initialized 0000 UTC 10 Sep 2008. TC position in ensemble members is marked by the black dots, best track position at ET time by the red dot. [Figure 1 from Keller et al. (2011).]

  • View in gallery

    (a) Schematic of the shift and amplitude pattern of ensemble forecast uncertainty derived from the first two EOFs (thin solid and dashed lines) of potential temperature on the dynamical tropopause in ensemble members. The thick black line represents the strong potential temperature gradient on the dynamic tropopause in the midlatitudes. (b) Synoptic patterns (shape of ridge) that result from the contribution to the variability patterns. (c) EOF 1 (left; contours) and 2 (right; contours) for potential temperature at 2 PVU (shaded in K) in an ECMWF ensemble forecast for Typhoon Maemi (2003). Values indicate percentage of total uncertainty captured by the respective EOF. [Figures 9, 10a, and 10b from Anwender et al. (2008).]

  • View in gallery

    Normalized ensemble spread of 500-hPa geopotential height as a function of forecast hour for NOAA’s second-generation global reforecasts initialized at recurvature time. Data cover all western North Pacific tropical cyclones from 1985 to 2013. Solid line shows the mean, dashed line shows the median, and the shaded region shows the 25th–75th percentile range of the distribution. Statistically significant values of the mean are shown as the thicker line. Hurricane symbol marks the time of recurvature, circle marks the median time of ET, and the thin vertical line marks the peak spread. [Figure 3a from Aiyyer (2015).]

  • View in gallery

    Forecast degradation due to data denied in (a) extratropical sensitive regions (SVout) and (b) vicinity of the storm (ETout), expressed as root-mean-square difference total energy. Box-and-whisker plot of the percentage impact over Europe (35°–75°N, 10°–30°E) for all denial cases. The 25 and the 75 quantile, median, and most extreme outliers are indicated by the box edges, red line, and whiskers, respectively. Vertical dashed lines separate ET cases, and vertical dotted lines indicate ET times. [Adapted from Fig. 5 of Anwender et al. (2012).]

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The Extratropical Transition of Tropical Cyclones. Part II: Interaction with the Midlatitude Flow, Downstream Impacts, and Implications for Predictability

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  • 1 Deutscher Wetterdienst, Offenbach, Germany
  • | 2 World Meteorological Organization, Geneva, Switzerland
  • | 3 Institute for Atmospheric and Climate Science, ETH Zürich, Zürich, Switzerland
  • | 4 Institute of Meteorology and Climate Research (IMK-TRO), Karlsruhe Institute of Technology, Karlsruhe, Germany
  • | 5 Johannes Gutenberg-Universität Mainz, Mainz, Germany
  • | 6 NOAA/Geophysical Fluid Dynamics Laboratory, Princeton, New Jersey
  • | 7 Department of Atmospheric and Environmental Sciences, University at Albany, State University of New York, Albany, New York
  • | 8 Naval Research Laboratory, Monterey, California
  • | 9 The Pennsylvania State University, University Park, Pennsylvania
  • | 10 The University of Arizona, Tucson, Arizona
  • | 11 RiskPulse, Madison, Wisconsin
  • | 12 Naval Postgraduate School, Monterey, California
  • | 13 Meteorological College, Kashiwa, Chiba, Japan
  • | 14 Numerical Weather Prediction Research Section, Environment and Climate Change Canada, Dorval, Québec, Canada
  • | 15 School of Earth, Atmosphere and Environment, and ARC Centre of Excellence for Climate System Science, Monash University, Clayton, Victoria, Australia
  • | 16 University of New South Wales, Canberra, Australian Capital Territory, Australia
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Abstract

The extratropical transition (ET) of tropical cyclones often has an important impact on the nature and predictability of the midlatitude flow. This review synthesizes the current understanding of the dynamical and physical processes that govern this impact and highlights the relationship of downstream development during ET to high-impact weather, with a focus on downstream regions. It updates a previous review from 2003 and identifies new and emerging challenges and future research needs. First, the mechanisms through which the transitioning cyclone impacts the midlatitude flow in its immediate vicinity are discussed. This “direct impact” manifests in the formation of a jet streak and the amplification of a ridge directly downstream of the cyclone. This initial flow modification triggers or amplifies a midlatitude Rossby wave packet, which disperses the impact of ET into downstream regions (downstream impact) and may contribute to the formation of high-impact weather. Details are provided concerning the impact of ET on forecast uncertainty in downstream regions and on the impact of observations on forecast skill. The sources and characteristics of the following key features and processes that may determine the manifestation of the impact of ET on the midlatitude flow are discussed: the upper-tropospheric divergent outflow, mainly associated with latent heat release in the troposphere below, and the phasing between the transitioning cyclone and the midlatitude wave pattern. Improving the representation of diabatic processes during ET in models and a climatological assessment of the ET’s impact on downstream high-impact weather are examples for future research directions.

Denotes content that is immediately available upon publication as open access.

© 2019 American Meteorological Society. For information regarding reuse of this content and general copyright information, consult the AMS Copyright Policy (www.ametsoc.org/PUBSReuseLicenses).

Corresponding author: Julia H. Keller, julia.keller@dwd.de

This article is included in the Predictability and Dynamics of Weather Systems in the Atlantic-European Sector (PANDOWAE) Special Collection.

This article is included in the Waves to Weather (W2W) Special Collection.

This article has a companion article which can be found at http://journals.ametsoc.org/doi/abs/10.1175/MWR-D-17-0027.1

Abstract

The extratropical transition (ET) of tropical cyclones often has an important impact on the nature and predictability of the midlatitude flow. This review synthesizes the current understanding of the dynamical and physical processes that govern this impact and highlights the relationship of downstream development during ET to high-impact weather, with a focus on downstream regions. It updates a previous review from 2003 and identifies new and emerging challenges and future research needs. First, the mechanisms through which the transitioning cyclone impacts the midlatitude flow in its immediate vicinity are discussed. This “direct impact” manifests in the formation of a jet streak and the amplification of a ridge directly downstream of the cyclone. This initial flow modification triggers or amplifies a midlatitude Rossby wave packet, which disperses the impact of ET into downstream regions (downstream impact) and may contribute to the formation of high-impact weather. Details are provided concerning the impact of ET on forecast uncertainty in downstream regions and on the impact of observations on forecast skill. The sources and characteristics of the following key features and processes that may determine the manifestation of the impact of ET on the midlatitude flow are discussed: the upper-tropospheric divergent outflow, mainly associated with latent heat release in the troposphere below, and the phasing between the transitioning cyclone and the midlatitude wave pattern. Improving the representation of diabatic processes during ET in models and a climatological assessment of the ET’s impact on downstream high-impact weather are examples for future research directions.

Denotes content that is immediately available upon publication as open access.

© 2019 American Meteorological Society. For information regarding reuse of this content and general copyright information, consult the AMS Copyright Policy (www.ametsoc.org/PUBSReuseLicenses).

Corresponding author: Julia H. Keller, julia.keller@dwd.de

This article is included in the Predictability and Dynamics of Weather Systems in the Atlantic-European Sector (PANDOWAE) Special Collection.

This article is included in the Waves to Weather (W2W) Special Collection.

This article has a companion article which can be found at http://journals.ametsoc.org/doi/abs/10.1175/MWR-D-17-0027.1

1. Introduction and motivation

Tropical cyclones (TCs) that move poleward often interact with the midlatitude flow, undergo profound structural changes, and transition into extratropical cyclones. This process is known as extratropical transition (ET; Sekioka 1956; Palmén 1958). In recent years, several ET cases were associated with extreme weather events, thus attracting the attention of the general public. Hurricane Sandy (2012) inflicted widespread damage and severe disruption of public life along the northeast U.S. coast as it underwent ET (Blake et al. 2013; Halverson and Rabenhorst 2013). Hurricane Gonzalo (2014), having undergone ET, tracked across Europe and brought flooding and extreme winds to the Balkans (Brown 2015; Feser et al. 2015). Extreme precipitation associated with Tropical Storm Etau (2015) during and after its ET over Japan flooded areas north and east of Tokyo (AIR Worldwide 2015; Kitabatake et al. 2017). In these examples, the high-impact weather was associated directly with the transitioning cyclone. Such impacts, along with the structural evolution of the cyclone during ET, are discussed in the first part of this review (Evans et al. 2017, hereafter Part I).

Extratropical transition may also lead to high-impact weather far downstream from the actual cyclone. A prominent example for such a “downstream impact” is provided by the ET of Supertyphoon Nuri (2014) in the western North Pacific, displayed in Fig. 1. At the onset of ET, Nuri moves poleward and starts to interact with the midlatitude flow (Fig. 1a). This results in the formation of a jet streak (Fig. 1b) and a poleward deflection of the jet near the transitioning cyclone in conjunction with the development of a ridge–trough couplet (Fig. 1b). At the same time, a region of enhanced moisture flux—a so-called atmospheric river (Zhu and Newell 1998)—forms ahead of the downstream trough. The ridge–trough couplet continues to amplify, a new cyclone develops farther downstream, and the next downstream ridge builds, which signifies the downstream propagation that arises from the initial local changes in the jet near the site of ET (Fig. 1c). Meanwhile, Nuri reintensifies into a strong extratropical cyclone and initiates cyclonic wave breaking over the western North Pacific (Fig. 1c). Subsequently, the upper-level wave pattern amplifies farther downstream, establishing a high-amplitude ridge–trough couplet over North America. A heat wave develops in the high-pressure conditions along the North American west coast, with highest values occurring along the coast of California and over Alaska. The atmospheric river in the western flank of the second downstream ridge (Fig. 1c) makes landfall in Alaska and British Columbia, resulting in heavy precipitation (Fig. 1d). A cold-air outbreak occurs over continental and eastern North America. Further amplification of this pattern eventually leads to a massive omega block over the west coast of North America and associated cold surges and heavy snowfall in the continental and eastern United States (Bosart et al. 2015). Nuri is just one example of the type of midlatitude flow modification due to ET. The processes acting during such a midlatitude flow modification and the associated implications on downstream extratropical regions are the subjects of this review.

Fig. 1.
Fig. 1.

Downstream development during the ET of Supertyphoon Nuri (2014) at (a) 0600 UTC 5 Nov, (b) 1800 UTC 6 Nov, (c) 1800 UTC 7 Nov, and (d) 0000 UTC 9 Nov 2014. Panels shows IR Gridsat clouds [brightness temperature in °C as in color bar; Knapp et al. (2011)], dynamical tropopause [2 PVU on the 330-K isentropic surface (1 PVU = 106 K kg−1 m2 s−1); red contour], and wind speed maxima highlighting jet streaks (wind speed on the 330-K isentropic surface as semitransparent shading in yellow for 55 and orange for 65 m s−1). Dynamical tropopause and wind speed are taken from ERA-Interim. The TC symbol indicates the position of the transitioning cyclone (encircled during extratropical stage) and the “L” the position of the downstream cyclone.

Citation: Monthly Weather Review 147, 4; 10.1175/MWR-D-17-0329.1

Together with Part I, this review describes developments in our understanding of ET since the first ET review by Jones et al. (2003, hereafter J2003). The review by J2003 was motivated by the challenges that ET typically poses to forecasters in terms of predicting the structural evolution of the transitioning cyclone itself, and the high-impact weather that might be associated with it, mostly in the immediate proximity of the storm. Since the publication of J2003, it has become increasingly apparent that a forecasting challenge is also present for the region downstream of ET because ET often leads to a basin-wide or even hemispheric reduction in the forecast skill of numerical weather prediction (NWP) models. J2003 reviewed the then-current insights into ET and highlighted the need for a better understanding of the physical and dynamic processes involved in ET and their representation in NWP models. Since then, the research community’s understanding of the interactions that occur between a transitioning cyclone and the midlatitude flow during ET has progressed considerably. The impact of ET on the midlatitude flow configuration and on predictability both near the transitioning cyclone and in downstream regions has now been quantified. These advancements motivate this second part of the updated review, which synthesizes our current understanding, and highlights open questions and current challenges, thus providing guidance for future research activities.

The structure of the paper largely follows the sequence of processes involved in downstream development during ET and is visualized in Fig. 2. The color of the labels in Fig. 2 indicates the section, whereas the index number refers to the subsection in which these aspects are discussed. Section 2 reviews the impact of ET on the midlatitude flow in the direct vicinity of the transitioning cyclone. The amplification of the downstream ridge, the formation of a jet streak, and the amplification of the downstream trough are discussed in section 2a because this material is key background information for the material that follows. Section 2b introduces aspects that arise due to the existence of an upstream trough: the importance of the position of the transitioning cyclone relative to the trough (phasing), the concepts of “phase locking” and associated resonant interaction [section 2b(1)], and the impact of ET on the upstream trough itself [section 2b(2)]. Section 2c introduces the idea of “preconditioning”: processes that occur before the onset of the actual ET and that promote interaction between the transitioning cyclone and the jet.

Fig. 2.
Fig. 2.

Overview of synoptic features and processes involved in Northern Hemispheric ET. Labels indicate relevant processes, starting with the section in which they are discussed. Transitioning cyclone presented by the black and white pictogram. The dark red line indicates axis of the undulating midlatitude jet stream separating stratospheric high PV air (dark gray; poleward) and tropospheric low PV air (light gray; equatorward), with the dashed line indicating an alternate configuration. The orange ellipse denotes the jet streak. The purple, semitransparent area signifies the forecast uncertainty for the downstream flow. The downstream cyclone is indicated by the “L” symbol and its associated fronts. The airplane symbol represents observation systems used for ET reconnaissance.

Citation: Monthly Weather Review 147, 4; 10.1175/MWR-D-17-0329.1

The midlatitude flow modifications introduced in section 2 often constitute the amplification or excitation of a Rossby wave packet (RWP; Wirth et al. 2018). Section 3 focuses on the downstream impacts of ET as mediated by RWP amplification or excitation. The modification of midlatitude RWPs by ET is discussed in section 3a. This subsection presents the mechanisms of downstream development during ET [section 3a(1)] before summarizing how RWP amplification during ET manifests in a climatological sense [section 3a(2)]. The contribution of ET to high-impact weather in downstream regions is the subject of section 3b.

Section 4 reviews predictability aspects (indicated by the semitransparent area enclosing potential alternative flow configurations in Fig. 2). Section 4a presents sources of forecast degradation during ET, whereas section 4b describes how forecast uncertainty associated with ET affects prediction downstream of ET. The potential impact of additional and targeted observations on the predictability associated with ET is presented in section 4c.

A summary and a presentation of current challenges and future directions for research are given in section 5.

2. Direct impacts on the midlatitude flow

During ET, the transitioning cyclone typically exerts a direct impact on the midlatitude flow, which manifests in a modification of the jet streak and the ridge–trough couplet immediately downstream of the transitioning cyclone. The processes associated with this impact (red labels in Fig. 2) are the subject of this section. J2003 discussed two mechanisms for this initial modification of the midlatitude flow. First, they hypothesized that the nonlinear-balanced circulation of the transitioning cyclone perturbs the gradient of potential vorticity (PV) associated with the midlatitude jet, thereby exciting RWPs and associated downstream development. The second mechanism occurs through diabatic PV modification and the presence of upper-tropospheric air with anticyclonic PV,1 originating from the TC outflow. This second mechanism was rather loosely defined by J2003 but has been described to enhance downstream ridging and jet streak formation and to promote anticyclonic wave breaking. More recent work has confirmed that both mechanisms operate and has clarified their respective roles. Furthermore, a third mechanism has been identified that is arguably the most important individual process: PV advection by the upper-tropospheric divergent outflow.

The amplification of the jet streak and of the ridge–trough couplet during ET is reviewed in section 2a. We discuss these processes in the context of wave amplification but stress that ET may also have a detrimental impact on downstream development. Section 2b discusses the large sensitivity that the TC–jet interaction exhibits to the relative position of the transitioning cyclone and the upstream trough (referred to as “phasing” hereafter). Section 2c introduces processes that impact the outcome of ET but occur before the actual ET commences. In this sense, these processes can be conceptually subsumed as preconditioning, a new concept that is not discussed by J2003.

a. Downstream ridge amplification, jet streak formation, and downstream trough amplification

The amplification (or generation) of a ridge–trough couplet and a jet streak downstream of ET are robust characteristics of the impact of ET on the midlatitude flow. These features have been found in idealized ET scenarios (Riemer et al. 2008; Riemer and Jones 2010), in numerical experiments and process-based analyses of real and composite cases (Agustí-Panareda et al. 2004; Harr and Dea 2009; Grams et al. 2011, 2013a; Griffin and Bosart 2014; Grams and Archambault 2016), in targeted observation studies (Chen and Pan 2010), and from ensemble (Torn 2010; Keller et al. 2014; Keller 2017) and climatological perspectives (Archambault et al. 2013, 2015; Torn and Hakim 2015; Quinting and Jones 2016).

Many of the studies cited above investigated ET from a PV framework, as proposed by J2003. In its most rigorous form, the PV framework considers the full PV budget of a region of interest and yields diabatic and advective PV tendencies [discussion of Eq. (4) by Teubler and Riemer (2016)]. The advective tendencies can be segregated further by carefully partitioning the full flow into distinct anomalies (e.g., Davis and Emanuel 1991). We here briefly introduce the terminology that is used throughout the paper to address the anomalies involved in ET (Figs. 3a,b; based on, e.g., Agustí-Panareda et al. 2004; Riemer et al. 2008). The flow attributed to the transitioning cyclone can be partitioned into three parts:

  1. The balanced (i.e., nondivergent) cyclonic circulation associated with the cyclonic PV tower and with low-level cyclonic PV anomalies, diabatically generated at the developing warm front.

  2. The balanced anticyclonic circulation associated with the anticyclonic PV anomaly of the upper-tropospheric outflow, consisting of air masses that have accumulated in the upper troposphere after having ascended in the presence of latent heat release from the lower troposphere (referred to “outflow anticyclone” in the remainder of the manuscript).

  3. The upper-tropospheric divergent outflow associated with latent heat release below.

Through “action at a distance,” all these features act on the midlatitude PV gradient, as hypothesized by J2003. Spatial integration of the associated PV tendencies over PV anomalies of interest (e.g., those that are associated with the downstream ridge) enables a quantitative assessment of the relative contribution of the advection through these processes to the overall evolution (exemplified in Fig. 3c; further discussed below).
Fig. 3.
Fig. 3.

PV anomalies involved in ET and their respective contribution to RWP modification near ET. The (a) 2D and (b) 3D schematics of the different flow features associated with ET. 1) Cyclonic circulation (thin orange arrows) associated with the cyclonic PV tower [orange TC symbol in (a), gray column in (b)] and with low-level cyclonic PV anomalies (small gray clouds and circulation) at the developing warm front. 2) Anticyclonic circulation associated with the anticyclonic PV anomaly of the upper-tropospheric outflow [white/gray cloud symbol in (a),(b) and thin blue arrows]. 3) Upper-tropospheric divergent outflow associated with latent heat release below (thin green arrows). The advective contribution of these flow features to the amplification of the downstream ridge and trough are given by the bold arrows in (a). The red contour and shading in (a),(b) are similar to Fig. 2. The lower layer in (b) exemplifies the developing warm front with colder air masses toward the pole. (c) Contributions to ridge amplification from the advection of potential temperature on the dynamic tropopause (a surrogate for upper-level PV advection) by the several flow features, integrated over a ridge for a 72-h forecast of an idealized scenario [in K km2 s−1; colors as in (a),(b)]. [Figure 8 from Riemer and Jones (2010), with modifications.]

Citation: Monthly Weather Review 147, 4; 10.1175/MWR-D-17-0329.1

Attributing the cyclonic and anticyclonic balanced circulations to the transitioning cyclone based on PV partition is justified by theory, but the attribution of the upper-tropospheric divergent outflow is not. In the context of ET, the literature agrees on the assumption that PV advection by upper-tropospheric divergent outflow is mostly associated with latent heat release within the transitioning cyclone below. In that sense, this PV advection is considered to be an indirect diabatic process. While the assumption has not yet been tested rigorously, a first test using the framework of the quasigeostrophic omega equation tends to support this assumption (Quinting and Jones 2016).

The formation of a jet streak can be considered from a PV perspective as manifestation of upper-tropospheric frontogenesis (Wandishin et al. 2000). During ET, jet streak formation is enhanced by the upper-tropospheric divergent outflow impinging on the large PV gradient associated with the midlatitude jet (Riemer and Jones 2010; Grams et al. 2013a; Archambault et al. 2013, 2015). Based on a quantitative analysis in an idealized ET scenario, this contribution is arguably as important as that of ongoing upper-tropospheric frontogenesis by the midlatitude dynamics (Riemer and Jones 2010). In addition, the outflow anticyclone constitutes a local elevation of the tropopause on the equatorward side of the jet, with respect to climatology. This elevation may locally increase the slope of the tropopause (i.e., strengthen the PV gradient on an isentrope intersecting the tropopause), thereby accelerating the jet locally and thus contributing to jet streak formation as well (Bosart 2003; Riemer and Jones 2010; Grams et al. 2013a). The latter mechanism has been discussed in a more general context and in a barotropic framework by Cunningham and Keyser (2000).

Complementary to the PV framework, ET can also be considered from a local eddy kinetic energy (Ke) perspective using the downstream baroclinic development paradigm [e.g., Eqs. (2.3) and (2.4) by Orlanski and Sheldon (1995)]. From that viewpoint, the amplification of the midlatitude ridge–trough couplet constitutes an increase in midlatitude Ke, and a jet streak appears as a local maximum in Ke, with the transitioning cyclone acting as an additional source of Ke. The notion that ET is a source of midlatitude Ke dates back to Palmén (1958). A number of more recent case studies further examined the processes underlying this source (Harr and Dea 2009; Keller et al. 2014; Quinting and Jones 2016; Keller 2017), and a consistent picture has emerged (depicted schematically in Fig. 4a and for a composite of real ET cases in Figs. 4b,c). Rising of warm air, mostly associated with latent heat release in the transitioning cyclone and along the baroclinic zone, releases Ke through the baroclinic conversion of eddy available potential energy (Figs. 4a,c). During ET, this Ke is redistributed via ageostrophic geopotential fluxes from the transitioning cyclone to the upstream end of the jet streak in the western flank of the ridge, evident by diverging fluxes emerging from the cyclone and converging fluxes in the western flank of the ridge (Figs. 4a,b).

Fig. 4.
Fig. 4.

Jet streak formation during ET from an energetics perspective. (a) Schematic representation, showing midlatitude jet (black line), developing Ke maxima (jet streak; gray ellipses), baroclinic conversion of Ke (clouds), ageostrophic geopotential flux (orange arrow), and its divergence (blue ellipses) and convergence (red ellipses). (b),(c) TC-relative composite of Ke budget for western North Pacific ETs, based on ERA-Interim for 1980–2010 [after Quinting and Jones (2016), their Figs. 12a,b]: vertically integrated Ke (shaded in 105 J m−2), 200-hPa geopotential (contours every 200 m2 s−2; thick black contour illustrates 11 800 m2 s−2), and (b) ageostrophic geopotential flux (vectors, reference vector in 106 W m−1; divergence as colored contours every 8 W m−2, divergence in blue, 0 W m−2 omitted) and (c) vertically integrated baroclinic conversion of Ke (red contours every 8 W m−2). The black box approximates the area that is captured by (b),(c). Composites are shown relative to the mean TC position.

Citation: Monthly Weather Review 147, 4; 10.1175/MWR-D-17-0329.1

The amplification of the ridge immediately downstream of ET was the focus of many of the studies cited at the beginning of this subsection. This ridge amplification may be vigorous enough to prevail over the ambient midlatitude Rossby wave dynamics. From the PV perspective, advection of anticyclonic PV by the upper-tropospheric divergent outflow makes a major contribution to ridge evolution and tends to dominate ridge amplification during the early phase of ET (Figs. 3a,b; Atallah and Bosart 2003; Riemer et al. 2008; Riemer and Jones 2010, 2014; Torn 2010; Archambault et al. 2013, 2015; Grams et al. 2013a; Lang and Martin 2013; Grams and Archambault 2016; Quinting and Jones 2016). In an idealized ET scenario, this process accounts, on average, for approximately 55% of the early-phase ridge amplification (from 36 to 72 h in Fig. 3c; Riemer and Jones 2010).

The source of the latent heat release, with which the upper-tropospheric divergent outflow and thus ridge building is associated, undergoes important changes during ET. Early during ET, ascent and associated latent heat release occurs mainly in the convection near the cyclone center (exemplified by trajectories depicted in Fig. 5a). At the same time, the cyclonic circulation of the transitioning cyclone advects warm and moist air masses out of the tropics toward the midlatitudes. When impinging on the baroclinic zone (Figs. 5a,b), these air masses begin to ascend slantwise along the sloping moist isentropes and convective bursts with associated latent heat release develop, usually poleward and east of the transitioning cyclone (Torn 2010; Grams et al. 2013a). During ET, these air masses become more stable, and saturated ascent becomes predominantly slantwise along the front (Fig. 5b), signifying the gradual evolution of a warm conveyor belt in the developing warm sector of the transitioning cyclone at later stages of ET (Agustí-Panareda et al. 2004; Evans and Hart 2008; Torn 2010; Grams et al. 2011, 2013a). The slantwise frontal ascent might still contain embedded convective cells. It thus seems plausible that the “elevator–escalator” perspective of Neiman et al. (1993), which describes ascent in the warm sector of extratropical cyclones as a combination of slantwise frontal upglide (escalator) and mesoconvective updrafts (elevator), also holds true for warm-frontal ascent during ET.

Fig. 5.
Fig. 5.

Lagrangian trajectories for the ET of Typhoon Jangmi showing ridge building and jet streak formation at (a) 0000 UTC 30 Sep and (b) 1200 UTC 1 Oct 2008. Shown are the 1.5-PVU surface (blue shading), 320-K surface of equivalent potential temperature (transparent gray shading), representing 3D baroclinic zone; |V| = 60 m s−1 (green shading), highlighting upper-level midlatitude jet; potential temperature at 990 hPa (shading at bottom; brown colors > 300 K, green ≈ 290 K) and geopotential height at 990 hPa (black contours; every 25 dam). Paths of representative trajectories [(a) 1200 UTC 28 Sep–1200 UTC 30 Sep 2008 and (b) 1200 UTC 30 Sep–1200 UTC 2 Oct 2008] colored by PV of air parcel moving along trajectory. Anticyclonic PV air (PV < 0.6 PVU) in gray shades and cyclonic PV values (PV > 0.6 PVU) in red shades (see legend in bottom right). [Figure 5 from Grams et al. (2013a).]

Citation: Monthly Weather Review 147, 4; 10.1175/MWR-D-17-0329.1

Advection of anticyclonic PV into the ridge by the cyclonic circulation associated with the transitioning cyclone, as hypothesized by J2003, is another contributor to ridge amplification (Fig. 3a; Riemer et al. 2008; Riemer and Jones 2010, 2014; Quinting and Jones 2016). This contribution to ridge amplification increases during ET as the cyclone moves closer to the midlatitude PV gradient, and it may become the dominant process for ridge building during the late stage of ET (Riemer et al. 2008; Riemer and Jones 2014).

In addition to the amplification of the downstream ridge, a transitioning cyclone directly amplifies the downstream trough through equatorward advection of cyclonic PV by the outflow anticyclone (Fig. 3a). This process has been observed in real cases and in idealized scenarios (Lazear and Morgan 2006; McTaggart-Cowan et al. 2006b; Riemer et al. 2008; Riemer and Jones 2010, 2014; Grams et al. 2013a,b; Grams and Blumer 2015; Grams and Archambault 2016). Furthermore, the presence of the outflow anticyclone implies an enhanced anticyclonic flow component in the region of the downstream trough and thereby indeed favors anticyclonic breaking of this trough (Riemer et al. 2008; Riemer and Jones 2010, 2014), as originally hypothesized by J2003. The advection of anticyclonic and cyclonic PV contribute equal to the direct amplification of the downstream ridge and trough, respectively, in the idealized scenario of Riemer and Jones (2014). Future studies, however, are needed to clarify the relative importance of both processes in the real atmosphere.

The PV and Ke frameworks provide complementary and broadly consistent views on the modification of the midlatitude flow by ET. Arguably, the Ke perspective provides a simpler general picture, whereas the strength of the PV perspective is apparent in the more detailed examination of individual processes. Note that individual terms in the respective budget equations of Ke and PV cannot be compared one to one. The two frameworks differ in particular in their interpretation of the secondary circulation associated with latent heat release. In the Ke perspective, the ascent associated with latent heat release is diagnosed as contributing to baroclinic conversion. The upper-tropospheric divergent outflow contributes to the ageostrophic geopotential flux. In an isentropic PV framework, (generalized) vertical motion is represented by diabatic heating, and hence diabatic PV modification is directly diagnosed. The upper-tropospheric divergent outflow is diagnosed as a separate process. More details on the differences between the two frameworks are provided by Teubler and Riemer (2016) and Wirth et al. (2018). Interpreting PV advection by the divergent flow as an indirect diabatic impact relies on the assumption that secondary circulations associated with dry, balanced dynamics are considerably smaller than those associated with latent heat release, at least near the transitioning cyclone. A rigorous test of this assumption is still missing. Arguably, connecting the PV framework with Lagrangian trajectory diagnostics yields the most comprehensive view on diabatic PV modification and cross-isentropic flow. The different manifestations of diabatic processes in the Ke and in the PV framework need to be kept in mind when interpreting the results.

In conclusion, the hypothesis of J2003 that the balanced circulation of the transitioning cyclone perturbs the midlatitude PV gradient has been largely confirmed: the cyclonic circulation contributes to ridge building and the anticyclonic circulation of the outflow anomaly to trough amplification downstream. Arguably the largest individual contributor to ridge building, as well as jet streak formation, however, is the upper-tropospheric divergent outflow, which undergoes important changes during ET. While J2003 hypothesized on the role of upper-tropospheric air with anticyclonic PV,2 which is usually found within the divergent outflow, this important contribution of the divergent flow in modifying the midlatitude flow was not considered explicitly.

b. Interaction and phasing of transitioning cyclone with upstream trough

The interaction between the transitioning cyclone and the midlatitude flow, and thus the amplification of the downstream ridge and the formation of the jet streak, strongly depends on the relative spatial position of the transitioning cyclone and the upstream trough. The importance of this “phasing” as a major source of forecast uncertainty was identified by J2003. Through idealized (Ritchie and Elsberry 2003, 2007; Riemer et al. 2008; Riemer and Jones 2010; Scheck et al. 2011a,b) and real case studies (Grams et al. 2013b) and a climatological assessment (Archambault et al. 2013, 2015; Quinting and Jones 2016; Riboldi et al. 2019), it is now clear that phasing and the interaction of the transitioning cyclone with the upstream trough determines the downstream development during ET.

1) Phasing, phase locking, and resonant interaction

Phasing determines whether a transitioning cyclone moves into an area that is favorable to midlatitude cyclone development or not. Typically, the region ahead of an upper-tropospheric trough is considered favorable, as can be quantified by evaluating Petterssen development parameters3 (Petterssen and Smebye 1971). As described in Part I, TCs that track into such a region ahead of the trough reintensify as extratropical cyclones, which means that their phasing with the midlatitude flow is favorable. In contrast, TCs that miss this region of favorable conditions tend to decay after ET (Klein et al. 2002; Ritchie and Elsberry 2003, 2007; Grams et al. 2013b). Transitioning TCs that undergo reintensification as extratropical cyclones support stronger amplification of the downstream ridge through the processes explained in section 2a and may lead to strong downstream impacts (section 3; Archambault et al. 2013; Grams et al. 2013b). In contrast to phasing, the initial size and strength of the TC, or the initial amplitude of the upstream trough, are secondary to the intensity evolution in the extratropical phase of ET (Ritchie and Elsberry 2003) or the magnitude of the downstream impact (Quinting and Jones 2016; Riboldi et al. 2018, 2019).

This high sensitivity to phasing can be traced back to the existence of a bifurcation point in the steering flow near the tip of the upstream trough in a trough-relative frame of reference; that is, the full flow minus the phase speed of the trough (arrows in Fig. 6; Scheck et al. 2011b; Grams et al. 2013b; Riemer and Jones 2014). Near such a bifurcation point (dot in Fig. 6), small differences in the position of the transitioning cyclone lead to large differences in the subsequent cyclone track (black lines in Fig. 6). The cyclones either track northeastward and undergo ET or continue their westward movement without undergoing ET, which means that the highly sensitive behavior around the bifurcation point translates into a high sensitivity of whether or not the transitioning cyclone recurves (changes its motion component from westward to eastward relative to the trough), reintensifies, and potentially exerts a pronounced downstream impact (Grams et al. 2013b). A second bifurcation point near the tip of the downstream ridge (cross in Fig. 6) apparently distinguishes between transitions into either the northwest or the northeast circulation pattern introduced by Harr et al. (2000; see also Riemer and Jones 2014). Bifurcation points also exist for the potential interaction of the cyclone with upper-tropospheric cutoff lows. In this case, the interaction can be interpreted as vortex–vortex interaction, leading to the eventual merger or escape of the vortices [e.g., in the case of Hurricane Nadine (2012); Pantillon et al. 2016; Munsell et al. 2015].

Fig. 6.
Fig. 6.

Idealized scenario for steering flow topology during ET. Midtropospheric geopotential to illustrate the midlatitude wave pattern and winds in the frame of reference moving with this pattern (at 620 hPa; shaded and arrows, respectively) and different tracks of transitioning cyclones from sensitivity experiments (thin black lines). The bifurcation points associated with the upstream trough and the developing downstream ridge, deciding upon the development of a “no ET,” an “NW pattern,” or an “NE pattern” scenario, are marked by the dot and the cross, respectively. Dashed contours depict the streamlines that emanate from the bifurcation points. [Figure 11 from Riemer and Jones (2014).]

Citation: Monthly Weather Review 147, 4; 10.1175/MWR-D-17-0329.1

In general, phasing evolves with time. There are processes, however, that promote a near-constant phasing over an extended period of time, referred to as phase locking. One such process is advection of midlatitude PV by the circulation associated with the outflow anticyclone (Fig. 3), which reduces the phase speed of the midlatitude Rossby wave and brings it closer to the translation speed of the transitioning cyclone (Riemer et al. 2008). In case of phase locking, the transitioning cyclone persistently amplifies the downstream ridge–trough couplet. In this sense, ET can be considered as a resonant interaction (Hodyss and Hendricks 2010; Scheck et al. 2011a,b), with the transitioning cyclone acting as a long-lived, local wave maker (Riemer et al. 2008) that moves in phase with the wave. Therefore, phase-locked configurations promote pronounced downstream impacts (Grams et al. 2013b; Riboldi et al. 2019) and favor strong reintensification of the transitioning cyclone after ET (Ritchie and Elsberry 2007). The local wave initiation and resonant interaction ideas imply that the transitioning cyclone constitutes an external forcing with persistent structure to the midlatitude wave. This idea is in marked contrast to traditional initial-value studies of baroclinic development, in which the initial perturbations are embedded in the midlatitude flow and are thus not an external forcing (e.g., Simmons and Hoskins 1979; Hakim 2000).

2) Evolution of the upstream trough

ET may also influence the upstream trough, which may experience modifications of its shape, meridional extension, and eventually break. These modifications influence phasing and thereby the overall flow evolution during ET.

The cyclonic circulation of the transitioning cyclone (Fig. 3a) impinging on the upstream trough may lead to trough amplification and/or thinning, as well as to a subsequent cyclonic wrap-up (McTaggart-Cowan et al. 2001; Agustí-Panareda et al. 2005; Riemer et al. 2008; Grams et al. 2011; Griffin and Bosart 2014; Riemer and Jones 2014; Quinting and Jones 2016). The upstream trough may further be modified by the upper-tropospheric divergent outflow, which might hinder the downstream propagation and cyclonic breaking of the trough. This hindering of downstream propagation may lead to trough thinning and the formation of a PV streamer (Agustí-Panareda et al. 2004; Grams et al. 2011; Pantillon et al. 2013a; Riemer and Jones 2014).

Interestingly, the observed impacts on the upstream trough during ET differ for different ocean basins and are sensitive to the large-scale midlatitude circulation pattern (J2003; Agustí-Panareda et al. 2005). Western North Pacific ETs tend to be associated with anticyclonic wave breaking and the formation of cutoff lows (i.e., the evolution follows the anticyclonic baroclinic life cycle paradigm; Davies et al. 1991; Thorncroft et al. 1993). Atlantic ETs tend to follow the cyclonic baroclinic life cycle with a cyclonic wrap-up of the trough and the formation of a broad and deep surface low (J2003 and references therein; Röbcke et al. 2004; Agustí-Panareda et al. 2004, 2005; Grams et al. 2011). The reasons for these differences in wave breaking and whether such large-scale circulation patterns associated with ET also exist in other ocean basins have not been investigated yet.

In conclusion, the relative position between the transitioning cyclone and the upstream trough (i.e., phasing) is crucial in determining the reintensification of the transitioning cyclone as an extratropical cyclone, the amplification of the downstream ridge–trough couplet, as well as the downstream impact of ET (in terms of RWP amplification). A reduction in the eastward propagation of the upstream trough by the divergent outflow and the cyclonic circulation of the TC, and a reduction of the phase speed of the RWP by the outflow anticyclone, may result in a phase-locked configuration. In this case, the transitioning cyclone and the upstream trough move in phase, and quasi-resonant interaction maximizes the amplification of the downstream ridge.

c. Preconditioning stage and predecessor rain events

The direct interaction between the transitioning cyclone and the midlatitude flow, as described above, might be preceded by processes that establish an extratropical environment that supports baroclinic development (preconditioning stage; Grams and Archambault 2016). J2003 mentioned the occasional occurrence of heavy precipitation events well poleward of the transitioning cyclone, which now have been established as so-called “predecessor rain events” (PRE; Cote 2007). Here, we consider the occurrence of PREs in the broader context of preconditioning.

Common to all processes involved in preconditioning is the poleward transport of warm and moist air of tropical origin (see Fig. 7 for the example of a PRE event). This transport can be facilitated if the transitioning cyclone recurves into a highly amplified wave pattern that yields a strong poleward steering flow (McTaggart-Cowan et al. 2006a,b). Alternatively, the poleward advection of tropical air masses may occur along the eastern side of a recurving TC and along the western flank of the subtropical high, showing the characteristics of a baroclinic moisture flux (McTaggart-Cowan et al. 2017).

Fig. 7.
Fig. 7.

Three-dimensional schematic depiction of the preconditioning stage with a PRE during western North Pacific ET anticyclonic PV air in the upper-level outflow of a TC and associated PRE as blue shading in the upper panel, and jet streak as green shading and 200-hPa waveguide as red contour separating high PV air (>3 PVU; orange shading) from lower PV air (<3 PVU; unshaded). Midlevel baroclinic zone as blue tilted surface. Trajectories of rapidly ascending air parcels as blue–red–blue lines, reflecting the diabatic PV modification of the parcels from low to high to low PVU, respectively. Mean sea level pressure (gray contours; every 8 hPa) and equivalent potential temperature (violet contours; 320 and 330 K) are indicated in the lower panel. [Figure 11a from Grams and Archambault (2016).]

Citation: Monthly Weather Review 147, 4; 10.1175/MWR-D-17-0329.1

During the preconditioning stage, exemplified for a PRE in Fig. 7, this tropical air impinges on the midlatitude baroclinic zone or experiences upper-tropospheric forcing for ascent ahead of the upstream trough (Fig. 7). The resulting ascent of the warm and moist tropical air mass may support extratropical cyclogenesis, the formation of a diabatic Rossby wave (Grams 2011, 2013; Riemer et al. 2014), or result in stationary heavy precipitation due to PREs well poleward of the transitioning cyclone. The upper-tropospheric divergent outflow associated with latent heat release in such a precursor weather system may initiate ridge building and jet acceleration (Fig. 7), similar to the transitioning cyclone itself (section 2a; Grams and Archambault 2016). Thus, prior to the onset of ET, these weather systems establish an extratropical environment that is characterized by an upstream trough and a downstream ridge, a flow configuration that may support the extratropical reintensification of the transitioning cyclone later (Fig. 7; Grams and Archambault 2016), and thus impacts the outcome of ET.

PREs are a particular type of preconditioning and have been studied extensively (e.g., Bosart and Carr 1978; Cote 2007; Stohl et al. 2008; Wang et al. 2009; Galarneau et al. 2010; Schumacher et al. 2011; Bosart et al. 2012; Byun and Lee 2012; Schumacher and Galarneau 2012; Cordeira et al. 2013; Baek et al. 2013; Moore et al. 2013; Parker et al. 2014; Bao et al. 2015; Galarneau 2015). PREs are regions of quasi-stationary convection and heavy precipitation that occur about 500–2000 km poleward of a recurving TC (Fig. 13 in Bosart and Carr 1978) and may develop in different synoptic-scale flow patterns (Moore et al. 2013). In general, PREs develop along a baroclinic zone when tropical air associated with the TC is ascending ahead of a trough and in the vicinity of a jet streak (Fig. 8). This results in heavy precipitation (Fig. 8; yellow–green ellipse) and associated diabatically enhanced upper-level outflow. Because of their preconditioning effect on the midlatitude flow, PREs may also influence the track of the transitioning cyclone (Galarneau 2015). About one-third of the North Atlantic TCs that made landfall in the United States between 1998 and 2006 produced at least one PRE (Cote 2007). Several PREs were associated with record-breaking amounts of precipitation (e.g., 500 mm in 48–72 h; Kitabatake 2002; Schumacher et al. 2011; Bosart et al. 2012). The heavy precipitation over Japan after the ET of Etau (2015), mentioned in the introduction, was superimposed on a PRE that developed well poleward of Typhoon Kilo (2015) at about the same time (Kitabatake et al. 2017). Furthermore, PREs may amplify the impact of the transitioning cyclone. The precipitation of the cyclone may impact the same regions that were affected by a PRE earlier, leading to exceptional flooding [recurrence frequency of 2000 years for the ET of Hurricane Erin (2007); Schumacher et al. 2011]. For Australia, a distinct impact has been observed. Enhanced ridge building over southeastern Australia due to PREs associated with recurving TCs at the Australian west coast can be important in the formation of heat waves, which in turn may favor bushfires (section 3b; Parker et al. 2013, 2014).

Fig. 8.
Fig. 8.

Conceptual model of the key synoptic-scale features during the occurrence of a PRE. Shown is the 200-hPa geopotential height (contours), the upper-tropospheric jet streak (gray shading), the midlatitude baroclinic zone [surface frontal structure with 925-hPa streamlines associated with warm (red) and cold (blue) air advection], and lower-tropospheric moisture flux along the eastern side of the recurving TC and the western flank of a subtropical high. The formation of the PRE is indicated by the yellow–green ellipse. [Fig. 20a of Moore et al. (2013).]

Citation: Monthly Weather Review 147, 4; 10.1175/MWR-D-17-0329.1

In summary, the poleward advection of tropical air masses prior to the actual ET may result in a preconditioning of the midlatitude flow, which may strongly impact the final outcome of the transition. Latent heat release and the associated upper-tropospheric divergent flow during extratropical cyclogenesis, the development of diabatic Rossby waves, or the formation of PREs may support the amplification of the upstream midlatitude trough and the first downstream ridge prior to ET.

3. Downstream impacts

The amplification of the first downstream ridge–trough couplet due to the processes elaborated in section 2 marks the initiation or modification of a midlatitude RWP (Fig. 1a; red contour). J2003 proposed that ET may excite Rossby waves on the upper-tropospheric PV gradient, which will disperse downstream by the mechanisms for downstream development of unstable baroclinic waves (Simmons and Hoskins 1979). Further, J2003 noted the importance of downstream development in the context of forecasting. The main focus of J2003 was on the amplification of the ridge–trough couplet directly downstream of ET. More recent work has investigated the processes that determine downstream development following the onset of ET beyond one wavelength [see Fig. 2 and blue labels for orientation; section 3a(1)] and identified a climatological signal of RWP development downstream of ET [section 3a(2)]. Furthermore, the development of high-impact weather in regions downstream of ET has been investigated more recently (section 3b).

a. Modification of midlatitude Rossby wave packets

1) Processes leading to modification of Rossby wave packets

The impact of ET is transmitted farther downstream by modifying the dispersion of RWPs (Riemer et al. 2008; Harr and Dea 2009; Riemer and Jones 2010, 2014; Grams et al. 2013b; Pantillon et al. 2013a; Griffin and Bosart 2014; Keller et al. 2014; Riemer et al. 2014; Archambault et al. 2015; Grams and Archambault 2016; Pryles and Ritchie 2016; Quinting and Jones 2016; Keller 2017). The concept of downstream baroclinic development (Orlanski and Sheldon 1995), introduced in section 2a, provides a succinct framework to describe this downstream propagation of the ET’s impact (Harr and Dea 2009; Keller et al. 2014; Keller 2017). The initial transmission of Ke from the transitioning cyclone into the Ke in the western flank of the first downstream ridge by ageostrophic geopotential fluxes and advection (referred to as total Ke flux) marks the initiation of downstream baroclinic development. Originating from this Ke maximum in the western flank of the trough, diverging and converging ageostrophic geopotential fluxes (Figs. 4a, 9a) and advection result in a total flux of Ke that is directed into downstream regions (Fig. 9b), leading to the further amplification of the RWP and its eastward propagation by group velocity. Baroclinic conversion in the remnants of the transitioning cyclone and along the baroclinic zone near the Ke maximum in the western flank, as well as in the downstream Ke maxima (Fig. 9c) and in possible downstream cyclones, further feed into the ongoing downstream baroclinic development (Orlanski and Sheldon 1995; Wirth et al. 2018).

Fig. 9.
Fig. 9.

Downstream development during the ET of Typhoon Nabi (2005). Vertically integrated Ke (shaded; 105 J m−2) and 500-hPa heights (light gray contours; 60-m intervals) and (a) vertically integrated divergence (dashed) and convergence (solid) of the ageostrophic geopotential flux (contours; W m−2); (b) vertically integrated total Ke flux vectors (advection + dispersion; reference vector in lower right; 105 W m−1); and (c) vertically integrated baroclinic conversion (contours; W m−2). [Figures 3a–c from Harr and Dea (2009).]

Citation: Monthly Weather Review 147, 4; 10.1175/MWR-D-17-0329.1

Consistent with this notion, the ET’s downstream impact is also sensitive to the evolution of cyclones in the downstream region (downstream cyclones; Fig. 2), which are main contributors to baroclinic conversion. Hence, a larger Rossby wave amplification near ET does not necessarily lead to a more amplified RWP farther downstream (Riemer and Jones 2010; Pantillon et al. 2015). In other words, the feedback by downstream cyclone development—including the associated moist processes (discussed below)—renders the impact of ET on the downstream region highly nonlinear. Often, however, the generation or amplification of RWPs near ET provides conditions conducive for downstream cyclone development (Hoskins and Berrisford 1988; Agustí-Panareda et al. 2004, 2005; Riemer et al. 2008; Riemer and Jones 2010; Grams et al. 2013b; Pantillon et al. 2013b; Archambault et al. 2015; Grams and Archambault 2016) such that cyclone development tends to be faster and stronger, thereby contributing to the amplification of RWPs downstream of ET.

The downstream impact is sensitive to characteristics of the midlatitude flow. For instance, the generation of midlatitude RWPs, in general, depends on the configuration of the midlatitude flow itself (Röthlisberger et al. 2016, 2018). An initially weaker upper-tropospheric midlatitude jet is typically susceptible to a stronger meridional deflection during ET than a strong jet and results in a more amplified RWP (Riemer et al. 2008; Riboldi et al. 2018). This is because phase locking is more likely to be achieved with a weak jet, and thus the initial ridge building is more pronounced [section 2b(1)]. In contrast, a strong jet immediately advects the anticyclonic PV air associated with the transitioning TC’s outflow downstream and thus hinders ridge building and phase locking [see Riboldi et al. (2018) for a detailed discussion].

The downstream development associated with ET is also sensitive to moisture transport within the midlatitude flow (Riemer et al. 2008; Grams and Archambault 2016; Riboldi et al. 2018) in accordance with general Rossby wave dynamics (e.g., Gutowski et al. 1992; Teubler and Riemer 2016). Moisture transport toward the baroclinic zone by downstream cyclones, the accompanying latent heat release in ascending moist air masses, and the associated upper-tropospheric divergent outflow result in enhanced ridge building (similar to processes described in section 2a; Riemer et al. 2008; Grams and Archambault 2016). Increased availability of moisture in the downstream region therefore tends to increase the downstream impact of cyclones undergoing ET.

The sensitivity of downstream development during ET to jet configuration and midlatitude moisture dominates over the sensitivity to the initial size and intensity of the transitioning cyclone during its tropical stage. Hence, the downstream impact of ET is—climatologically speaking—relatively insensitive to the intensity and size of the transitioning cyclone during its tropical stage (Archambault et al. 2013; Quinting and Jones 2016; Riboldi et al. 2018). In the case of a midlatitude flow configuration that promotes RWP amplification (i.e., an initially relatively weak upper-tropospheric jet stream and the availability of low-level moisture), however, sensitivity of the downstream impact of ET to the characteristics of the transitioning cyclone emerges: initially stronger TCs often lead to more amplified RWPs (Riemer et al. 2008; Riemer and Jones 2010; Archambault et al. 2013, 2015; Grams and Archambault 2016). Furthermore, transitioning cyclones that reintensify more strongly during ET are associated with more amplified downstream RWPs (e.g., Archambault et al. 2013). Likewise, the strength and duration of baroclinic conversion of Ke within the transitioning cyclone determine the amount of additional Ke released by the transitioning cyclone that feeds the development of the RWP (Harr and Dea 2009; Keller et al. 2014; Keller 2017).

Favorable phasing is a prerequisite for the initiation of substantial downstream development [see section 2b(2)]. When favorable phasing occurs, the strength of the interaction between the midlatitude flow and the transitioning cyclone influences downstream development during ET, with strong interactions leading to more amplified RWPs than weak interactions. The strength of the interaction (i.e., interaction metric) can be approximately quantified by the upper-tropospheric advection of anticyclonic PV by the divergent outflow (Archambault et al. 2013). The divergent outflow (Fig. 10a) advects anticyclonic PV poleward (upper panel in Fig. 10a) and thereby enhances the PV gradient and deflects the jet stream, which results in jet streak formation (lower panel in Fig. 10a). Although the jet and the PV gradient might be initially weak, for western North Pacific strong interactions, a pronounced jet streak and downstream ridge evolve (Fig. 10b), whereas the jet remains weak and is less deflected poleward with weak interactions (Fig. 10c). Strong interactions typically lead to more amplified RWPs that reach North America (Fig. 10d), as compared to weak interactions (Fig. 10e), for which RWPs dissipate well prior to reaching North America (not shown). The interaction metric is in line with the more general concept that anticyclonic vorticity advection by the divergent wind acts as a Rossby wave source (Sardeshmukh and Hoskins 1988; Hodyss and Hendricks 2010).

Fig. 10.
Fig. 10.

Interaction between a transitioning cyclone and the midlatitude flow expressed as advection of low PV air by the upper-level divergent outflow. (a) Idealized representation of ridge amplification and jet streak intensification. Vectors represent the upper-tropospheric divergent outflow associated with the transitioning cyclone. Shading denotes anticyclonic PV advection by the divergent wind (Archambault et al. 2013, their Fig. 4). Composite analyses of objectively defined (b) strong and (c) weak interactions at the time of maximum interaction. The 500-hPa ascent (green; every 2 × 10−3 hPa s−1, negative values only), total column precipitable water (shaded according to grayscale; mm), 200-hPa PV (blue; every 1 PVU), irrotational wind (vectors, > 2 m s−1; purple vectors, > 8 m s−1), negative PV advection by the irrotational wind (dashed red; every 2 PVU day−1 starting at −2 PVU day−1), and total wind speed (shaded according to color bar; m s−1). The star denotes point of maximum interaction. The TC symbol denotes composite TC position. Downstream development of (d) strong and (e) weak interactions 36 h after time of maximum interaction as represented by 250-hPa meridional wind anomalies (shaded; m s−1; enclosed by black contours where significant at the 99% confidence level), PV (blue; every 1 PVU), and irrotational wind (vectors; >2 m s−1). [Figures 8a,b, 5d, and 6d from Archambault et al. (2015).]

Citation: Monthly Weather Review 147, 4; 10.1175/MWR-D-17-0329.1

The abovementioned processes facilitate the amplification of RWPs and thus downstream development during ET. In cases where the transitioning cyclone interacts with an already-well-amplified midlatitude RWP, ET may initiate Rossby wave breaking and can thus be detrimental to downstream development (Riemer and Jones 2014).

2) A climatological perspective on RWP amplification during ET

Despite the large case-to-case variability and nonlinear interactions of the processes that govern the downstream development initiated by ET, RWP amplification downstream of ET reveals itself as a climatologically consistent feature in most ocean basins.

In the western North Pacific and south Indian Ocean, RWPs downstream of ET are more amplified and occur more frequently than in climatology. They are also more amplified than RWPs associated with extratropical cyclones (Torn and Hakim 2015, Quinting and Jones 2016). Between June and November, an enhancement of RWP frequency by up to 15% (Fig. 11a) becomes apparent downstream of ET across the western and central North Pacific, as well as North America. For the same period, the south Indian Ocean RWP frequency is enhanced by up to 18% (Fig. 11b).

Fig. 11.
Fig. 11.

Recurvature-relative composites of enhanced RWP frequency anomaly (shaded in %) (a) relative to June–November climatology for western North Pacific transitioning cyclones and (b) relative to December–April climatology for south Indian Ocean transitioning cyclones. Statistical significance at 95% confidence level hatched, mean track given by black line. Mean and range of recurvature longitudes indicated by white star and black bar, respectively. Data smoothed with a Gaussian filter. [Figures 3a and 5a from Quinting and Jones (2016).]

Citation: Monthly Weather Review 147, 4; 10.1175/MWR-D-17-0329.1

The impact of ET on RWPs in the North Atlantic is less clear. Compared to RWPs associated with extratropical cyclones, the RWPs downstream of ET in the North Atlantic appear to be less amplified (Torn and Hakim 2015). Compared to climatology, however, RWPs downstream of ET do not show significant differences in their amplitude (Quinting and Jones 2016). These different results found for the North Atlantic basin might stem from differences in the methods of these studies to detect RWPs and from the different sample sizes. Furthermore, the lack of statistically significant differences from climatology in the amplification of RWPs downstream of North Atlantic ET might result from the sensitivity to the midlatitude jet structure. The North Atlantic jet is climatologically short and weak and thus prone to stronger amplification, but also to wave breaking, which may disrupt downstream development (Wirth et al. 2018). The climatological results are confirmed by case studies for the North Atlantic, which were not able to unambiguously attribute RWP amplification in the North Atlantic to ET (e.g., Agustí-Panareda et al. 2004; McTaggart-Cowan et al. 2001, 2003, 2004; Grams et al. 2011; Pantillon et al. 2015).

In summary, the downstream impact of ET can be interpreted as the local modification of RWPs that then disperse this impact downstream. RWP amplification is more likely if the midlatitude upper-tropospheric jet is initially relatively weak and enhanced low-level moisture is available. In such a midlatitude flow susceptible to RWP amplification, and if the transitioning cyclone is in favorable phasing with an upstream trough, cyclone characteristics such as intensity and/or strength of the TC–midlatitude flow interaction further modulate the downstream impact of ET. The downstream impact of ET manifests as a climatologically consistent signal in RWP amplification downstream of the transitioning cyclone in the western North Pacific and south Indian Ocean, whereas the climatological signal in the North Atlantic might be masked by Rossby wave breaking initiated during ET.

b. Downstream high-impact weather

By triggering or amplifying midlatitude RWPs, ET may contribute to the development of high-impact weather in downstream regions (e.g., Harr and Archambault 2016). The relationship between ET and downstream high-impact weather exists because strongly amplified RWPs, in general, may result in blocking anticyclones (e.g., Nakamura et al. 1997; Renwick and Revell 1999; Martius et al. 2013; Riboldi et al. 2019), establish atmospheric conditions that are prone to strong cyclogenes