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The DOI for this manuscript is doi: 10.1175/MWR-D-17-0329.1

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Keller, J., C. Grams, M. Riemer, H. Archambault, L. Bosart, J. Doyle, J. Evans, T. Galarneau, K. Griffin, P. Harr, N. Kitabatake, R. McTaggart-Cowan, F. Pantillon, J. Quinting, C. Reynolds, E. Ritchie, R. Torn, and F. Zhang, 2018: The Extratropical Transition of Tropical Cyclones Part II: Interaction with the midlatitude flow, downstream impacts, and implications for predictability. Mon. Wea. Rev. doi:10.1175/MWR-D-17-0329.1, in press.

3 The Extratropical Transition of Tropical Cyclones Part II: Interaction with the

4 midlatitude flow, downstream impacts, and implications for predictability

6 Julia H. Keller1 7 World Meteorological Organization, Geneva, Switzerland

8 Christian M. Grams 9 Institute for Atmospheric and Climate Science, ETH Zürich, Zürich, Switzerland and Institute of 10 Meteorology and Climate Research (IMK-TRO), Karlsruhe Institute of Technology, Karlsruhe, 11 Germany

12 Michael Riemer 13 Johannes Gutenberg-Universität Mainz, Mainz, Germany

14 Heather M. Archambault 15 NOAA/Geophysical Fluid Dynamics Laboratory, Princeton, New Jersey

16 Lance Bosart 17 Department of Atmospheric and Environmental Sciences, University at 18 Albany, State University of New York, Albany, New York

19 James D. Doyle 20 Naval Research Laboratory, Monterey, California

21 Jenni L. Evans 22 The Pennsylvania State University, University Park, Pennsylvania

23 Thomas J. Galarneau Jr. 24 The University of Arizona, Tucson, Arizona

25 Kyle Griffin 26 RiskPulse, Madison, Wisconsin

28 Corresponding Author: Julia H. Keller, World Meteorological Organization, Geneva, Switzerland, 29 email: [email protected]

1 Former affiliation: Deutscher Wetterdienst, Research & Development, Numerical Models, Offenbach, Germany

30 Patrick A. Harr 31 Naval Postgraduate School, Monterey, California

32 Naoko Kitabatake 33 Meteorological College, Kashiwa, Chiba, Japan

34 Ron McTaggart-Cowan 35 Numerical Weather Prediction Research Section, Environment and Climate Change Canada

36 Florian Pantillon 37 Institute of Meteorology and Climate Research (IMK-TRO), Karlsruhe Institute of Technology, 38 Karlsruhe, Germany

39 Julian Quinting2 40 School of Earth, Atmosphere and Environment, and ARC Centre of Excellence for Climate System 41 Science, Monash University, Clayton, Australia and Institute of Meteorology and Climate Research 42 (IMK-TRO), Karlsruhe Institute of Technology, Karlsruhe, Germany

43 Carolyn A. Reynolds 44 Naval Research Laboratory, Monterey, California

45 Elizabeth A. Ritchie 46 University of New South Wales, Canberra, Australia

47 Ryan D. Torn 48 Department of Atmospheric and Environmental Sciences, University at 49 Albany, State University of New York, Albany, New York

50 Fuqing Zhang 51 The Pennsylvania State University, University Park, Pennsylvania

52 Abstract

53 The extratropical transition (ET) of tropical cyclones often has an important impact on

54 the nature and predictability of the midlatitude flow. This review synthesizes the current

55 understanding of the dynamical and physical processes that govern this impact and highlights

56 the relationship of downstream development during ET to high-impact weather, with a focus

57 on downstream regions. It updates a previous review from 2003 and identifies new and

58 emerging challenges, and future research needs. First, the mechanisms through which the

59 transitioning cyclone impacts the midlatitude flow in its immediate vicinity is discussed. This

60 ‘direct impact’ manifests in the formation of a jet streak and the amplification of a ridge

61 directly downstream of the cyclone. This initial flow modification triggers or amplifies a

62 midlatitude Rossby wave packet, which disperses the impact of ET into downstream regions

63 (‘downstream impact’) and may contribute to the formation of high-impact weather. Details

64 are provided concerning the impact of ET on forecast uncertainty in downstream regions and

65 on the impact of observations on forecast skill. The sources and characteristics of the

66 following key features and processes that may determine the manifestation of the impact of

67 ET on the midlatitude flow are discussed: the upper-tropospheric divergent outflow, mainly

68 associated with latent heat release in the troposphere below, and the phasing between the

69 transitioning cyclone and the midlatitude wave pattern. Improving the representation of

70 diabatic processes during ET in models, and a climatological assessment of the ET’s impact

71 on downstream high-impact weather are examples for future research directions.

74 1. Introduction and motivation

75 Tropical cyclones (TC) that move poleward often interact with the midlatitude flow,

76 undergo profound structural changes, and transition into extratropical cyclones. This process

77 is known as extratropical transition (ET; Sekioka 1956; Palmén 1958). In recent years,

78 several ET cases were associated with extreme weather events, thus attracting the attention of

79 the general public. Hurricane Sandy (2012) inflicted widespread damage and severe

80 disruption of public life along the Northeast U.S. coast as it underwent ET (Blake et al. 2013;

81 Halverson and Rabenhorst 2013). Hurricane Gonzalo (2014), having undergone ET, tracked

82 across Europe and brought flooding and extreme winds to the Balkans (Brown 2015; Feser et

83 al. 2015). Extreme precipitation associated with Tropical Storm Etau (2015) during and after

84 its ET over Japan flooded areas north and east of Tokyo (AIR Worldwide 2015; Kitabatake et

85 al. 2017). In these examples, the high-impact weather was associated directly with the

86 transitioning cyclone. Such impacts, along with the structural evolution of the cyclone during

87 ET, are discussed in the first part of this review (Evans et al. 2017; referred to as Part I in the

88 remainder of this manuscript).

89 Extratropical transition may also lead to high-impact weather far downstream from

90 the actual cyclone. A prominent example for such a ‘downstream impact’ is provided by the

91 ET of Supertyphoon Nuri (2014) in the western North Pacific, displayed in Fig. 1. At the

92 onset of ET, Nuri moves poleward and starts to interact with the midlatitude flow (Fig. 1a).

93 This results in the formation of a jet streak (Fig. 1b) and a poleward deflection of the jet near

94 the transitioning cyclone in conjunction with the development of a ridge–trough couplet (Fig.

95 1b). At the same time a region of enhanced moisture flux—a so-called atmospheric river

96 (Zhu and Newell 1998)—forms ahead of the downstream trough. The ridge–trough couplet

97 continues to amplify, a new cyclone develops farther downstream, and the next downstream

98 ridge builds, which signifies the downstream propagation that arises from the initial local

99 changes in the jet near the site of ET (Fig. 1c). Meanwhile, Nuri reintensifies into a strong

100 extratropical cyclone and initiates cyclonic wave breaking over the western North Pacific

101 (Fig. 1c). Subsequently, the upper-level wave pattern amplifies further downstream,

102 establishing a high-amplitude ridge–trough couplet over North America. A heat wave

103 develops in the high-pressure conditions along the North American west coast, with highest

104 values occurring along the coast of California and over Alaska. The atmospheric river in the

105 western flank of the second downstream ridge (Fig. 1c) makes landfall in Alaska and British

106 Columbia, resulting in heavy precipitation (Fig. 1d). A cold-air outbreak occurs over the

107 continental and eastern North America. Further amplification of this pattern eventually leads

108 to a massive omega block over the west coast of North America and associated cold surges

109 and heavy snowfall in the continental and eastern U.S. (Bosart et al. 2015). Nuri is just one

110 example of the type of midlatitude flow modification due to ET. The processes acting during

111 such a midlatitude flow modification and the associated implications on downstream

112 extratropical regions are the subjects of this review.

113 Together with Part I, this review describes developments in our understanding of ET

114 since the first ET review by Jones et al. (2003). The review by Jones et al. (2003), referred to

115 as J2003 in the remainder of the manuscript, was motivated by the challenges that ET

116 typically poses to forecasters in terms of predicting the structural evolution of the

117 transitioning cyclone itself, and the high-impact weather that might be associated with it,

118 mostly in the immediate proximity of the storm. Since the publication of J2003 it has become

119 increasingly apparent that a forecasting challenge is also present for the region downstream

120 of ET because ET often leads to a basin-wide or even hemispheric reduction in the forecast

121 skill of numerical weather prediction (NWP) models. J2003 reviewed the then current

122 insights into ET and highlighted the need for a better understanding of the physical and

123 dynamic processes involved in ET, and their representation in NWP models. Since then, the

124 research community's understanding of the interactions that occur between a transitioning

125 cyclone and the midlatitude flow during ET has progressed considerably. The impact of ET

126 on the midlatitude flow configuration and on predictability both near the transitioning

127 cyclone and in downstream regions has now been quantified. These advancements motivate

128 this second part of the updated review, which synthesizes our current understanding, and

129 highlights open questions and current challenges, thus providing guidance for future research

130 activities.

131 The structure of the paper largely follows the sequence of processes involved in

132 downstream development during ET and is visualized in Fig. 2. The color of the labels in Fig.

133 2 indicates the section, whereas the index number refers to the subsection in which these

134 aspects are discussed. Section 2 reviews the impact of ET on the midlatitude flow in the

135 direct vicinity of the transitioning cyclone. The amplification of the downstream ridge, the

136 formation of a jet streak, and the amplification of the downstream trough are discussed in

137 section 2a because this material is key background information for the material that follows.

138 Section 2b introduces aspects that arise due to the existence of an upstream trough: the

139 importance of the position of the transitioning cyclone relative to the trough (“phasing”), the

140 concepts of “phase-locking” and associated resonant interaction (section 2b.1), as well as the

141 impact of ET on the upstream trough itself (section 2b.2). Section 2c introduces the idea of

142 “preconditioning”, in other words: processes that occur before the onset of the actual ET and

143 that promote interaction between the transitioning cyclone and the jet.

144 The midlatitude flow modifications introduced in section 2 often constitute the

145 amplification or excitation of a Rossby wave packet (RWP, Wirth et al. 2018). Section 3

146 focuses on the downstream impacts of ET as mediated by RWP amplification or excitation.

147 The modification of midlatitude RWPs by ET is discussed in section 3a. This subsection

148 presents the mechanisms of downstream development during ET (section 3a.1), before

149 summarizing how RWP amplification during ET manifests in a climatological sense (section

150 3a.2). The contribution of ET to high-impact weather in downstream regions is the subject of

151 section 3b.

152 Section 4 reviews predictability aspects (indicated by the semi-transparent area

153 enclosing potential alternative flow configurations in Fig. 2). Section 4a presents sources of

154 forecast degradation during ET, whereas section 4b describes how forecast uncertainty

155 associated with ET affects prediction downstream of ET. The potential impact of additional

156 and targeted observations on the predictability associated with ET is presented in section 4c.

157 A summary and a presentation of current challenges and future directions for research

158 are given in section 5.

159 2. Direct impacts on the midlatitude flow

160 During ET, the transitioning cyclone typically exerts a direct impact on the

161 midlatitude flow, which manifests in a modification of the jet streak and the ridge–trough

162 couplet immediately downstream of the transitioning cyclone. The processes associated with

163 this impact (red labels in Fig. 2) are the subject of this section. J2003 discussed two

164 mechanisms for this initial modification of the midlatitude flow. First, they hypothesized that

165 the nonlinear-balanced circulation of the transitioning cyclone perturbs the gradient of

166 potential vorticity (PV) associated with the midlatitude jet, thereby exciting RWPs and

167 associated downstream development. The second mechanism occurs through diabatic PV

168 modification and the presence of upper-tropospheric air with anticyclonic PV1, originating

169 from the TC outflow. This second mechanism was rather loosely defined by J2003, but has

170 been described to enhance downstream ridging and jet streak formation, and to promote

171 anticyclonic wave breaking. More recent work has confirmed that both mechanisms operate

1 We use the term anticyclonic PV to denote negative PV anomalies in the northern hemisphere and positive PV anomalies in the southern hemisphere. The term cyclonic PV is used accordingly.

172 and has clarified their respective roles. Furthermore, a third mechanism has been identified

173 that is arguably the most important individual process: PV advection by the upper-

174 tropospheric divergent outflow.

175 The amplification of the jet streak and of the ridge–trough couplet during ET is

176 reviewed in section 2a. We discuss these processes in the context of wave amplification but

177 stress that ET may also have a detrimental impact on downstream development. Section 2b

178 discusses the large sensitivity that the TC–jet interaction exhibits to the relative position of

179 the transitioning cyclone and the upstream trough (referred to as “phasing” hereafter). Section

180 2c introduces processes that impact the outcome of ET but occur before the actual ET

181 commences. In this sense, these processes can be conceptually subsumed as preconditioning,

182 a new concept that is not discussed in J2003.

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

184 amplification

185 The amplification (or generation) of a ridge–trough couplet and a jet streak

186 downstream of ET are robust characteristics of the impact of ET on the midlatitude flow.

187 These features have been found in idealized ET scenarios (Riemer et al. 2008; Riemer and

188 Jones 2010), numerical experiments and process-based analyses of real and composite cases

189 (Agustí-Panareda et al. 2004; Harr and Dea 2009; Grams et al. 2011, 2013a; Griffin and

190 Bosart 2014; Grams and Archambault 2016), targeted observation studies (Chen and Pan

191 2010), and from ensemble (Torn 2010; Keller et al. 2014; Keller 2017) and climatological

192 perspectives (Archambault et al. 2013, 2015; Torn and Hakim 2015; Quinting and Jones

193 2016).

194 Many of the studies cited above, investigated ET from a PV framework, as proposed

195 by J2003. In its most rigorous form, the PV framework considers the full PV budget of a

196 region of interest and yields diabatic and advective PV tendencies (discussion of Eq. 4 in

197 Teubler and Riemer 2016). The advective tendencies can be segregated further by carefully

198 partitioning the full flow into distinct anomalies (e.g., Davis and Emanuel 1991). We here

199 briefly introduce the terminology, which is used throughout the paper, to address the

200 anomalies involved in ET (Fig. 3a and 3b, based, e.g., on Agustí-Panareda et al. 2004;

201 Riemer et al. 2008). The flow attributed to the transitioning cyclone can be partitioned into

202 three parts: i) the balanced (i.e., non-divergent) cyclonic circulation associated with the

203 cyclonic PV tower and with low-level cyclonic PV anomalies, diabatically generated at the

204 developing warm front, ii) the balanced anticyclonic circulation associated with the

205 anticyclonic PV anomaly of the upper-tropospheric outflow, consisting of air masses that

206 have accumulated in the upper troposphere after having ascended in the presence of latent

207 heat release from the lower troposphere, (referred to “outflow anticyclone” in the remainder

208 of the manuscript), and iii) the upper-tropospheric divergent outflow associated with latent

209 heat release below. Through “action-at-a-distance”, all these features act on the midlatitude

210 PV gradient, as hypothesized by J2003. Spatial integration of the associated PV tendencies

211 over PV anomalies of interest (e.g., those that are associated with the downstream ridge),

212 enables a quantitative assessment of the relative contribution of the advection through these

213 processes to the overall evolution (exemplified in Fig. 3c, further discussed below).

214 Attributing the cyclonic and anticyclonic balanced circulations to the transitioning

215 cyclone based on PV partition is justified by theory, but the attribution of the upper-

216 tropospheric divergent outflow is not. In the context of ET, the literature agrees on the

217 assumption that PV advection by upper-tropospheric divergent outflow is mostly associated

218 with latent heat release within the transitioning cyclone below. In that sense, this PV

219 advection is considered to be an indirect diabatic process. While the assumption has not yet

220 been tested rigorously, a first test using the framework of the quasi-geostrophic omega

221 equation tends to support this assumption (Quinting and Jones 2016).

222 The formation of a jet streak can be considered, from a PV perspective, as

223 manifestation of upper-tropospheric frontogenesis (Wandishin et al. 2000). During ET, jet

224 streak formation is enhanced by the upper-tropospheric divergent outflow impinging on the

225 large PV gradient associated with the midlatitude jet (Riemer and Jones 2010; Grams et al.

226 2013a; Archambault et al. 2013, 2015). Based on a quantitative analysis in an idealized ET

227 scenario, this contribution is arguably as important as that of ongoing upper-tropospheric

228 frontogenesis by the midlatitude dynamics (Riemer and Jones 2010). In addition, the outflow

229 anticyclone constitutes a local elevation of the tropopause on the equatorward side of the jet,

230 with respect to climatology. This elevation may locally increase the slope of the tropopause

231 (i.e., strengthen the PV gradient on an isentrope intersecting the tropopause) thereby

232 accelerating the jet locally and thus contributing to jet streak formation as well (Bosart 2003;

233 Riemer and Jones 2010; Grams et al. 2013a). The latter mechanism has been discussed in a

234 more general context and in a barotropic framework by Cunningham and Keyser (2000).

235 Complementary to the PV framework, ET can also be considered from a local eddy

236 kinetic energy (Ke) perspective using the downstream baroclinic development paradigm (e.g.,

237 Eq. 2.3 and 2.4 in Orlanski and Sheldon 1995). From that viewpoint, the amplification of the

238 midlatitude ridge–trough couplet constitutes an increase in midlatitude Ke and a jet streak

239 appears as a local maximum in Ke, with the transitioning cyclone acting as an additional

240 source of Ke. The notion that ET is a source of midlatitude Ke dates back to Palmén (1958). A

241 number of more recent case studies further examined the processes underlying this source

242 (Harr and Dea 2009; Keller et al. 2014; Quinting and Jones 2016; Keller 2017) and a

243 consistent picture has emerged (depicted schematically in Fig. 4a, and for a composite of real

244 ET cases in Fig. 4b, c). Rising of warm air, mostly associated with latent heat release in the

245 transitioning cyclone and along the baroclinic zone, releases Ke through the baroclinic

246 conversion of eddy available potential energy (Fig. 4a, c. During ET, this Ke is redistributed

247 via ageostrophic geopotential fluxes from the transitioning cyclone to the upstream end of the

248 jet streak in the western flank of the ridge, evident by diverging fluxes emerging from the

249 cyclone and converging fluxes in the western flank of the ridge (Fig. 4a, b).

250 The amplification of the ridge immediately downstream of ET was the focus of many

251 of the studies cited at the beginning of this subsection. This ridge amplification may be

252 vigorous enough to prevail over the ambient midlatitude Rossby wave dynamics. From the

253 PV perspective, advection of anticyclonic PV by the upper-tropospheric divergent outflow

254 makes a major contribution to ridge evolution and tends to dominate ridge amplification

255 during the early phase of ET (Fig. 3a,b; Atallah and Bosart 2003; Riemer et al. 2008; Riemer

256 and Jones 2010, 2014; Torn 2010; Archambault et al. 2013, 2015; Grams et al. 2013a; Lang

257 and Martin 2013; Grams and Archambault 2016; Quinting and Jones 2016). In an idealized

258 ET scenario this process accounts on average for approximately 55% of the early-phase ridge

259 amplification (from 36–72-h in Fig. 3c; Riemer and Jones 2010).

260 The source of the latent heat release, with which the upper-tropospheric divergent

261 outflow and thus ridge building is associated, undergoes important changes during ET. Early

262 during ET, ascent and associated latent heat release occurs mainly in the convection near the

263 cyclone center (exemplified by trajectories depicted in Fig. 5a). At the same time, the

264 cyclonic circulation of the transitioning cyclone advects warm and moist air masses out of the

265 tropics toward the midlatitudes. When impinging on the baroclinic zone (Fig. 5a, b), these air

266 masses begin to ascend slantwise along the sloping moist isentropes and convective burst

267 with associated latent heat release develop, usually poleward and east of the transitioning

268 cyclone (Torn 2010; Grams et al. 2013a). During ET, these air masses become more stable

269 and saturated ascent becomes predominantly slantwise along the front (Fig. 5b), signifying

270 the gradual evolution of a warm conveyor belt in the developing warm sector of the

271 transitioning cyclone at later stages of ET (Agustí-Panareda et al. 2004; Evans and Hart 2008;

272 Torn 2010; Grams et al. 2011; Grams et al. 2013a). The slantwise frontal ascent might still

273 contain embedded convective cells. It thus seems plausible that the “elevator–escalator”

274 perspective of Neiman et al. (2003), which describes ascent in the warm sector of

275 extratropical cyclones as a combination of slantwise frontal upglide (escalator) and meso-

276 convective updrafts (elevator), also holds true for warm-frontal ascent during ET.

277 Advection of anticyclonic PV into the ridge by the cyclonic circulation associated

278 with the transitioning cyclone, as hypothesized by J2003, is another contributor to ridge

279 amplification (Fig. 3a; Riemer et al. 2008; Riemer and Jones 2010, 2014; Quinting and Jones

280 2016). This contribution to ridge amplification increases during ET as the cyclone moves

281 closer to the midlatitude PV gradient and it may become the dominant process for ridge

282 building during the late stage of ET (Riemer et al. 2008; Riemer and Jones 2014).

283 In addition to the amplification of the downstream ridge, a transitioning cyclone

284 directly amplifies the downstream trough, through equatorward advection of cyclonic PV by

285 the outflow anticyclone (Fig. 3a). This process has been observed in real cases and in

286 idealized scenarios (Lazear and Morgan 2006; McTaggart-Cowan et al. 2006b; Riemer et al.

287 2008; Riemer and Jones 2010, 2014; Grams et al. 2013a, b; Grams and Blumer 2015; Grams

288 and Archambault 2016). Furthermore, the presence of the outflow anticyclone implies an

289 enhanced anticyclonic flow component in the region of the downstream trough and thereby

290 indeed favors anticyclonic breaking of this trough (Riemer et al. 2008; Riemer and Jones

291 2010, 2014), as originally hypothesized by J2003. The advection of anticyclonic and cyclonic

292 PV contribute equal to the direct amplification of the downstream ridge and trough,

293 respectively, in the idealized scenario of Riemer and Jones (2014). Future studies, however,

294 are needed to clarify the relative importance of both processes in the real atmosphere.

295 The PV and Ke frameworks provide complementary and broadly consistent views on

296 the modification of the midlatitude flow by ET. Arguably, the Ke perspective provides a

297 simpler general picture, whereas the strength of the PV perspective is apparent in the more

298 detailed examination of individual processes. Note that individual terms in the respective

299 budget equations of Ke and PV cannot be compared one-to-one. The two frameworks differ in

300 particular in their interpretation of the secondary circulation associated with latent heat

301 release. In the Ke perspective, the ascent associated with latent heat release is diagnosed as

302 contributing to baroclinic conversion. The upper-tropospheric divergent outflow contributes

303 to the ageostrophic geopotential flux. In an isentropic PV framework, (generalized) vertical

304 motion is represented by diabatic heating and hence diabatic PV modification is directly

305 diagnosed. The upper-tropospheric divergent outflow is diagnosed as a separate process.

306 More details on the differences between the two frameworks is provided in Teubler and

307 Riemer (2016) and Wirth et al. (2018). Interpreting PV advection by the divergent flow as an

308 indirect diabatic impact relies on the assumption that secondary circulations associated with

309 dry, balanced dynamics are considerably smaller than those associated with latent heat

310 release, at least near the transitioning cyclone. A rigorous test of this assumption is still

311 missing. Arguably, connecting the PV framework with Lagrangian trajectory diagnostics

312 yields the most comprehensive view on diabatic PV modification and cross-isentropic flow.

313 The different manifestations of diabatic processes in the Ke and in the PV framework need to

314 be kept in mind when interpreting the results.

315 In conclusion, the hypothesis of J2003 that the balanced circulation of the

316 transitioning cyclone perturbs the midlatitude PV gradient has been largely confirmed: The

317 cyclonic circulation contributes to ridge building and the anticyclonic circulation of the

318 outflow anomaly to trough amplification downstream. Arguably the largest individual

319 contributor to ridge building, as well as jet streak formation, however, is the upper-

320 tropospheric divergent outflow, which undergoes important changes during ET. While J2003

321 hypothesized on the role of upper-tropospheric air with anticyclonic PV2, which is usually

322 found within the divergent outflow, this important contribution of the divergent flow in

323 modifying the midlatitude flow was not considered explicitly.

324 b. Interaction and phasing of transitioning cyclone with upstream trough

325 The interaction between the transitioning cyclone and the midlatitude flow, and thus the

326 amplification of the downstream ridge and the formation of the jet streak, strongly depend on

327 the relative spatial position of the transitioning cyclone and the upstream trough. The

328 importance of this “phasing” as a major source of forecast uncertainty was identified by

329 J2003. Through idealized (Ritchie and Elsberry 2003, 2007; Riemer et al. 2008; Riemer and

330 Jones 2010; Scheck et al. 2011a,b) and real case studies (Grams et al. 2013b), and a

331 climatological assessment (Archambault et al. 2013, 2015; Quinting and Jones 2016; Riboldi

332 et al. 2019), it is now clear that phasing and the interaction of the transitioning cyclone with

333 the upstream trough determines the downstream development during ET.

334 1) PHASING, PHASE LOCKING AND RESONANT INTERACTION

335 Phasing determines whether a transitioning cyclone moves into an area that is

336 favorable to midlatitude cyclone development or not. Typically, the region ahead of an upper-

337 tropospheric trough is considered favorable, as can be quantified by evaluating Petterssen

338 development parameters3 (Petterssen and Smebye, 1971). As described in Part I, TCs that

339 track into such a region ahead of the trough reintensify as extratropical cyclones, which

340 means that their phasing with the midlatitude flow is favorable. In contrast, TCs that miss this

341 region of favorable conditions tend to decay after ET (Klein et al. 2002; Ritchie and Elsberry

2 Direct diabatic PV modification, as diagnosed by the diabatic term in the PV equation, has been of secondary importance in all studies that have performed PV budget analysis and hence is not discussed in this review. 3 These are, for example, upper-tropospheric divergence, lower-tropospheric warm-air advection and midtropospheric cyclonic vorticity advection

342 2003, 2007; Grams et al. 2013b). Transitioning TCs that undergo reintensification as

343 extratropical cyclones support stronger amplification of the downstream ridge through the

344 processes explained in section 2.1, and may lead to strong downstream impacts (section 3;

345 Archambault et al. 2013; Grams et al. 2013b). In contrast to phasing, the initial size and

346 strength of the TC, or the initial amplitude of the upstream trough, do not determine the

347 intensity evolution in the extratropical phase of ET (Ritchie and Elsberry 2003) or the

348 magnitude of the downstream impact (Quinting and Jones 2016; Riboldi et al. 2018, 2019).

349 This high sensitivity to phasing can be traced back to the existence of a bifurcation

350 point in the steering flow near the tip of the upstream trough in a trough-relative frame of

351 reference, that is the full flow minus the phase speed of the trough (arrows in Fig. 6; Scheck

352 et al. 2011b; Grams et al. 2013b; Riemer and Jones 2014). Near such a bifurcation point (dot

353 in Fig. 6), small differences in the position of the transitioning cyclone lead to large

354 differences in the subsequent cyclone track (black lines in Fig. 6). The cyclones either track

355 northeastward and undergo ET, or continue their westward movement without undergoing

356 ET, which means that the highly sensitive behavior around the bifurcation point translates

357 into a high sensitivity of whether the transitioning cyclone recurves (changes its motion

358 component from westward to eastward relative to the trough), reintensifies, and potentially

359 exerts a pronounced downstream impact or not (Grams et al. 2013b). A second bifurcation

360 point near the tip of the downstream ridge (cross in Fig. 6) apparently distinguishes between

361 transitions into either the northwest or the northeast circulation pattern introduced by Harr et

362 al. (2000) (Riemer and Jones 2014). Bifurcation points also exist for the potential interaction

363 of the cyclone with upper-tropospheric cut-off lows. In this case, the interaction can be

364 interpreted as vortex–vortex interaction, leading to the eventual merger or escape of the

365 vortices (e.g., in the case of Hurricane Nadine (2012); Pantillon et al. 2016; Munsell et al.

366 2015).

367 In general, phasing evolves with time. There are processes, however, that promote a

368 near-constant phasing over an extended period of time, referred to as phase-locking. One

369 such process is advection of midlatitude PV by the circulation associated with the outflow

370 anticyclone (Fig. 3), which reduces the phase speed of the midlatitude Rossby wave and

371 brings it closer to the translation speed of the transitioning cyclone (Riemer et al. 2008). In

372 case of phase locking, the transitioning cyclone persistently amplifies the downstream ridge–

373 trough couplet. In this sense, ET can be considered as a resonant interaction (Hodyss and

374 Hendricks 2010; Scheck et al. 2011a, b) with the transitioning cyclone acting as a long-lived,

375 local wave-maker (Riemer et al. 2008) that moves in phase with the wave. Therefore, phase-

376 locked configurations promote pronounced downstream impacts (Grams et al. 2013b; Riboldi

377 et al. 2019) and favor strong reintensification of the transitioning cyclone after ET (Ritchie

378 and Elsberry 2007). The local wave initiation and resonant interaction ideas imply that the

379 transitioning cyclone constitutes an external forcing with persistent structure to the

380 midlatitude wave. This idea is in marked contrast to traditional initial-value studies of

381 baroclinic development, in which the initial perturbations are embedded in the midlatitude

382 flow and are thus not an external forcing (e.g., Simmons and Hoskins 1979; Hakim 2000).

383 2) EVOLUTION OF THE UPSTREAM TROUGH

384 ET may also influence the upstream trough, which may experience modifications of

385 its shape, meridional extension, and eventually break. These modifications influence phasing

386 and thereby the overall flow evolution during ET.

387 The cyclonic circulation of the transitioning cyclone (Fig. 3a) impinging on the

388 upstream trough may lead to trough amplification and/or thinning, as well as to a subsequent

389 cyclonic wrap-up (McTaggart-Cowan et al. 2001; Agustí-Panareda et al. 2005; Riemer et al.

390 2008; Grams et al. 2011; Griffin and Bosart 2014; Riemer and Jones 2014; Quinting and

391 Jones 2016). The upstream trough may further be modified by the upper-tropospheric

392 divergent outflow, which might hinder the downstream propagation and cyclonic breaking of

393 the trough. This hindering of downstream propagation may lead to trough thinning and the

394 formation of a PV streamer (Agustí-Panareda et al. 2004; Grams et al. 2011; Pantillon et al.

395 2013a; Riemer and Jones 2014).

396 Interestingly, the observed impacts on the upstream trough during ET differ for

397 different ocean basins and are sensitive to the large-scale midlatitude circulation pattern

398 (J2003; Agustí-Panareda et al. 2005). Western North Pacific ETs tend to be associated with

399 anticyclonic wave breaking and the formation of cut-off lows (i.e., the evolution follows the

400 anticyclonic baroclinic life-cycle paradigm; Davies et al. 1991; Thorncroft et al. 1993).

401 Atlantic ETs tend to follow the cyclonic baroclinic life-cycle with a cyclonic wrap-up of the

402 trough and the formation of a broad and deep surface low (J2003 and references therein;

403 Röbcke et al. 2004; Agustí-Panareda et al. 2004, 2005; Grams et al. 2011). The reasons for

404 these differences in wave breaking and whether such large-scale circulation patterns

405 associated with ET also exist in other ocean basins has not been investigated yet.

406 In conclusion, the relative position between the transitioning cyclone and the

407 upstream trough (i.e., phasing) is crucial in determining the reintensification of the

408 transitioning cyclone as an extratropical cyclone, the amplification of the downstream ridge–

409 trough couplet, as well as the downstream impact of ET (in terms of RWP amplification). A

410 reduction in the eastward propagation of the upstream trough by the divergent outflow and

411 the cyclonic circulation of the TC, and a reduction of the phase speed of the RWP by the

412 outflow anticyclone may result in a phase-locked configuration. In this case, the transitioning

413 cyclone and the upstream trough move in phase and quasi-resonant interaction maximizes the

414 amplification of the downstream ridge.

415 c. Preconditioning stage and predecessor rain events

416 The direct interaction between the transitioning cyclone and the midlatitude flow, as

417 described above, might be preceded by processes that establish an extratropical environment

418 that supports baroclinic development (“preconditioning stage”; Grams and Archambault

419 2016). J2003 mentioned the occasional occurrence of heavy precipitation events well

420 poleward of the transitioning cyclone, which now have been established as so-called

421 “predecessor rain events” (PRE; Cote 2007). Here, we consider the occurrence of PREs in the

422 broader context of preconditioning.

423 Common to all processes involved in preconditioning is the poleward transport of

424 warm and moist air of tropical origin (Fig. 7, for the example of a PRE event). This transport

425 can be facilitated if the transitioning cyclone recurves into a highly-amplified wave pattern

426 that yields a strong poleward steering flow (McTaggart-Cowan et al. 2006a, b). Alternatively,

427 the poleward advection of tropical air masses may occur along the eastern side of a recurving

428 TC and along the western flank of the subtropical high, showing the characteristics of a

429 baroclinic moisture flux (McTaggart-Cowan et al. 2017).

430 During the preconditioning stage, exemplified for a PRE in Fig. 7, this tropical air

431 impinges on the midlatitude baroclinic zone or experiences upper-tropospheric forcing for

432 ascent ahead of the upstream trough (Fig. 7). The resulting ascent of the warm and moist

433 tropical air mass may support extratropical cyclogenesis, the formation of a diabatic Rossby

434 wave (Grams 2011; Grams 2013; Riemer et al. 2014) or result in stationary heavy

435 precipitation due to PREs well poleward of the transitioning cyclone. The upper-tropospheric

436 divergent outflow associated with latent heat release in such a precursor weather system may

437 initiate ridge building and jet acceleration (Fig. 7), similar to the transitioning cyclone itself

438 (section 2a; Grams and Archambault 2016). Thus, prior to the onset of ET, these weather

439 systems establish an extratropical environment that is characterized by an upstream trough

440 and a downstream ridge, a flow configuration that may support the extratropical

441 reintensification of the transitioning cyclone later (Fig. 7; Grams and Archambault 2016), and

442 thus impacts the outcome of ET.

443 PREs are a particular type of preconditioning and have been studied extensively (e.g.;

444 Bosart and Carr 1978; Cote 2007; Stohl et al. 2008; Wang et al. 2009; Galarneau et al. 2010;

445 Schumacher et al. 2011; Bosart et al. 2012; Byun and Lee 2012; Schumacher and Galarneau

446 2012; Cordeira et al. 2013; Baek et al. 2013; Moore et al. 2013; Parker et al. 2014; Bao et al.

447 2015; Galarneau 2015). PREs are regions of quasi-stationary convection and heavy

448 precipitation that occur about 500–2000 km poleward of a recurving TC (Fig. 13 in Bosart

449 and Carr 1978) and may develop in different synoptic-scale flow patterns (Moore et al. 2013).

450 In general, PREs develop along a baroclinic zone when tropical air associated with the TC is

451 ascending ahead of a trough and in the vicinity of a jet streak (Fig. 8). This results in heavy

452 precipitation (Fig. 8; yellow-green ellipse) and associated diabatically enhanced upper-level

453 outflow. Due to their preconditioning effect on the midlatitude flow, PREs may also

454 influence the track of the transitioning cyclone (Galarneau 2015). About one-third of the

455 North Atlantic TCs that made landfall in the U.S. between 1998 and 2006 produced at least

456 one PRE (Cote 2007). Several PREs were associated with record-breaking amounts of

457 precipitation (e.g.; 500 mm in 48–72-h; Kitabatake 2002; Schumacher et al. 2011; Bosart et

458 al. 2012). The heavy precipitation over Japan after the ET of Etau (2015), mentioned in the

459 introduction, was superimposed on a PRE that developed well poleward of Typhoon Kilo

460 (2015) at about the same time (Kitabatake et al. 2017). Furthermore, PREs may amplify the

461 impact of the transitioning cyclone. The precipitation of the cyclone may impact the same

462 regions that were affected by a PRE earlier, leading to exceptional flooding (recurrence

463 frequency of 2000 years for the ET of Hurricane Erin (2007); Schumacher et al. 2011). For

464 Australia, a distinct impact has been observed. Enhanced ridge building over southeastern

465 Australia due to PREs associated with recurving TCs at the Australian west coast can be

466 important in the formation of heat waves, which in turn may favor bushfires (section 3b;

467 Parker et al. 2013, 2014).

468 In summary, the poleward advection of tropical air masses prior to the actual ET may

469 result in a preconditioning of the midlatitude flow, which may strongly impact the final

470 outcome of the transition. Latent heat release and the associated upper-tropospheric divergent

471 flow during extratropical cyclogenesis, the development of diabatic Rossby waves, or the

472 formation of PREs may support the amplification of the upstream midlatitude trough and the

473 first downstream ridge prior to ET.

474 3. Downstream impacts

475 The amplification of the first downstream ridge–trough couplet due to the processes

476 elaborated in section 2 marks the initiation or modification of a midlatitude RWP (Fig. 1a,

477 red contour). J2003 proposed that ET may excite Rossby waves on the upper-tropospheric PV

478 gradient, which will disperse downstream by the mechanisms for downstream development

479 of unstable baroclinic waves (Simmons and Hoskins 1979). Further, J2003 noted the

480 importance of downstream development in the context of forecasting. The main focus in

481 J2003 was on the amplification of the ridge–trough couplet directly downstream of ET. More

482 recent work has investigated the processes that determine downstream development

483 following the onset of ET beyond one wavelength (see Fig. 2 and blue labels for orientation;

484 section 3a.1) and identified a climatological signal of RWP development downstream of ET

485 (section 3a.2). Furthermore, the development of high-impact weather in regions downstream

486 of ET has been investigated more recently (section 3b).

487

488

489 a. Modification of midlatitude Rossby wave packets

490 1) PROCESSES LEADING TO MODIFICATION OF ROSSBY WAVE PACKETS

491 The impact of ET is transmitted further downstream by modifying the dispersion of

492 RWPs (Riemer et al. 2008; Harr and Dea 2009; Riemer and Jones 2010, 2014; Grams et al.

493 2013b; Pantillon et al. 2013a; Griffin et al. 2014; Keller et al. 2014; Riemer et al. 2014;

494 Archambault et al. 2015; Grams and Archambault 2016; Pryles and Ritchie 2016; Quinting

495 and Jones 2016; Keller 2017). The concept of downstream baroclinic development (Orlanski

496 and Sheldon 1995), introduced in section 2a, provides a succinct framework to describe this

497 downstream propagation of the ET’s impact (Harr and Dea 2009; Keller et al. 2014; Keller

498 2017). The initial transmission of Ke from the transitioning cyclone into the Ke in the western

499 flank of the first downstream ridge by ageostrophic geopotential fluxes and advection

500 (referred to as total Ke flux) marks the initiation of downstream baroclinic development.

501 Originating from this Ke maximum in the western flank of the trough, diverging and

502 converging ageostrophic geopotential fluxes (Fig. 4a, 9a) and advection result in a total flux

503 of Ke that is directed into downstream regions (Fig. 9b), leading to the further amplification

504 of the RWP and its eastward propagation by group velocity. Baroclinic conversion in the

505 remnants of the transitioning cyclone and along the baroclinic zone near the Ke maximum in

506 the western flank, as well as in the downstream Ke maxima (Fig. 9c) and in possible

507 downstream cyclones, further feed into the ongoing downstream baroclinic development

508 (Orlanski and Sheldon 1995; Wirth et al. 2018).

509 Consistent with this notion, the ET’s downstream impact is also sensitive to the

510 evolution of cyclones in the downstream region (downstream cyclones, Fig. 2), which are

511 main contributors to baroclinic conversion. Hence, a larger Rossby wave amplification near

512 ET does not necessarily lead to a more amplified RWP further downstream (Riemer and

513 Jones 2010, Pantillon et al. 2015). In other words, the feedback by downstream cyclone

514 development—including the associated moist processes (discussed below)—renders the

515 impact of ET on the downstream region highly nonlinear. Often, however, the generation or

516 amplification of RWPs near ET provides conditions conducive for downstream cyclone

517 development (Hoskins and Berrisford 1988; Agustí-Panareda et al. 2004, 2005; Riemer et al.

518 2008; Riemer and Jones 2010; Grams et al. 2013b; Pantillon et al. 2013b; Archambault et al.

519 2015; Grams and Archambault 2016) such that cyclone development tends to be faster and

520 stronger, thereby contributing to the amplification of RWPs downstream of ET.

521 The downstream impact is sensitive to characteristics of the midlatitude flow. For

522 instance, the generation of midlatitude RWPs, in general, depends on the configuration of the

523 midlatitude flow itself (Röthlisberger et al. 2016, 2018). An initially weaker upper-

524 tropospheric midlatitude jet is typically susceptible to a stronger meridional deflection during

525 ET than a strong jet and results in a more amplified RWP (Riemer et al. 2008; Riboldi et al.

526 2018). This is because phase-locking is more likely to be achieved with a weak jet and thus

527 the initial ridge building is more pronounced (section 2b.1). In contrast, a strong jet

528 immediately advects the anticyclonic PV air associated with the transitioning TC’s outflow

529 downstream and thus hinders ridge building and phase locking (see Riboldi et al. 2018 for a

530 detailed discussion).

531 The downstream development associated with ET is also sensitive to moisture

532 transport within the midlatitude flow (Riemer et al. 2008; Grams and Archambault 2016;

533 Riboldi et al. 2018) in accordance with general Rossby wave dynamics (e.g., Gutowski et al.

534 1992; Teubler and Riemer 2016). Moisture transport toward the baroclinic zone by

535 downstream cyclones, the accompanying latent heat release in ascending moist air masses,

536 and the associated upper-tropospheric divergent outflow result in enhanced ridge building

537 (similar to processes described in section 2a; Riemer et al. 2010; Grams and Archambault

538 2016). Increased availability of moisture in the downstream region therefore tends to increase

539 the downstream impact of cyclones undergoing ET.

540 The sensitivity of downstream development during ET to jet configuration and

541 midlatitude moisture dominates over the sensitivity to the initial size and intensity of the

542 transitioning cyclone during its tropical stage. Hence, the downstream impact of ET is —

543 climatologically speaking—relatively insensitive to the intensity and size of the transitioning

544 cyclone during its tropical stage (Archambault et al. 2013; Quinting and Jones 2016; Riboldi

545 et al. 2018). In the case of a midlatitude flow configuration that promotes RWP amplification

546 (i.e., an initially relatively weak upper-tropospheric jet stream and the availability of low-

547 level moisture), however, sensitivity of the downstream impact of ET to the characteristics of

548 the transitioning cyclone emerges: initially stronger TCs often lead to more amplified RWPs

549 (Riemer et al. 2008; Riemer and Jones 2010; Archambault et al. 2013, 2015; Grams and

550 Archambault 2016). Furthermore, transitioning cyclones that reintensify more strongly during

551 ET are associated with more amplified downstream RWPs (e.g., Archambault et al. 2013).

552 Likewise, the strength and duration of baroclinic conversion of Ke within the transitioning

553 cyclone determine the amount of additional Ke released by the transitioning cyclone that feeds

554 the development of the RWP (Harr and Dea 2009; Keller et al. 2014; Keller 2017).

555 Favorable phasing is a prerequisite for the initiation of substantial downstream

556 development (see section 2b.2). When favorable phasing occurs, the strength of the

557 interaction between the midlatitude flow and the transitioning cyclone influences downstream

558 development during ET, with strong interactions leading to more amplified RWPs than weak

559 interactions. The strength of the interaction (i.e., interaction metric) can be approximately

560 quantified by the upper-tropospheric advection of anticyclonic PV by the divergent outflow

561 (Archambault et al. 2013). The divergent outflow (Fig. 10a) advects anticyclonic PV

562 poleward (upper panel in Fig. 10a), and thereby enhances the PV gradient, deflects the jet

563 stream, which results in jet streak formation (lower panel in Fig. 10a). Although the jet and

564 the PV gradient might be initially weak, for western North Pacific strong interactions a

565 pronounced jet streak and downstream ridge evolve (Fig. 10b), whereas the jet remains weak

566 and is less deflected poleward with weak interactions (Fig. 10c). Strong interactions typically

567 lead to more amplified RWPs that reach North America (Fig. 10d) as compared to weak

568 interactions (Fig. 10e), for which RWPs dissipate well prior to reaching North America (not

569 shown). The interaction metric is in line with the more general concept that anticyclonic

570 vorticity advection by the divergent wind acts as a Rossby wave source (Sardeshmukh and

571 Hoskins 1988; Hodyss and Hendricks 2010).

572 The above-mentioned processes facilitate the amplification of RWPs, and thus

573 downstream development during ET. In cases where the transitioning cyclone interacts with

574 an already well amplified midlatitude RWP, ET may initiate Rossby wave breaking and can

575 thus be detrimental to downstream development (Riemer and Jones 2014).

576

577 2) A CLIMATOLOGICAL PERSPECTIVE ON RWP AMPLIFICATION DURING ET

578 Despite the large case-to-case variability and nonlinear interactions of the processes

579 that govern the downstream development initiated by ET, RWP amplification downstream of

580 ET reveals itself as a climatologically consistent feature in most ocean basins.

581 In the western North Pacific and South Indian Ocean, RWPs downstream of ET are

582 more amplified, and occur more frequently than in climatology. They are also more amplified

583 compared to RWPs associated with extratropical cyclones (Torn and Hakim 2015, Quinting

584 and Jones 2016). Between June and November, an enhancement of RWP frequency by up to

585 15% (Fig. 11a) becomes apparent downstream of ET across the western and central North

586 Pacific, as well as North America. For the same period, the South Indian Ocean RWP

587 frequency is enhanced by up to 18% (Fig. 11b).

588 The impact of ET on RWPs in the North Atlantic is less clear. Compared to RWPs

589 associated with extratropical cyclones, the RWPs downstream of ET in the North Atlantic

590 appear to be less amplified (Torn and Hakim 2015). Compared to climatology, however,

591 RWPs downstream of ET do not show significant differences in their amplitude (Quinting

592 and Jones 2016). These different results found for the North Atlantic basin might stem from

593 differences in the methods of these studies to detect RWPs and from the different sample

594 sizes. Furthermore, the lack of statistically significant differences from climatology in the

595 amplification of RWPs downstream of North Atlantic ET might result from the sensitivity to

596 the midlatitude jet structure. The North Atlantic jet is climatologically short and weak, and

597 thus prone to stronger amplification, but also to wave breaking, which may disrupt

598 downstream development (Wirth et al. 2018). The climatological results are confirmed by

599 case studies for the North Atlantic, which were not able to unambiguously attribute RWP

600 amplification in the North Atlantic to ET (e.g., Agustí-Panareda et al. 2004; McTaggart-

601 Cowan et al. 2001, 2003, 2004; Grams et al. 2011; Pantillon et al. 2015).

602 In summary, the downstream impact of ET can be interpreted as the local

603 modification of RWPs that then disperse this impact downstream. RWP amplification is more

604 likely if the midlatitude upper-tropospheric jet is initially relatively weak and enhanced low-

605 level moisture is available. In such a midlatitude flow susceptible to RWP amplification, and

606 if the transitioning cyclone is in favorable phasing with an upstream trough, cyclone

607 characteristics such as intensity and/or strength of the TC–midlatitude flow interaction further

608 modulate the downstream impact of ET. The downstream impact of ET manifests as a

609 climatologically consistent signal in RWP amplification downstream of the transitioning

610 cyclone in the western North Pacific and South Indian Ocean, whereas the climatological

611 signal in the North Atlantic might be masked by Rossby wave breaking initiated during ET.

612

613 b. Downstream high-impact weather

614 By triggering or amplifying midlatitude RWPs, ET may contribute to the development

615 of high-impact weather in downstream regions (e.g., Harr and Archambault 2016). The ET–

616 downstream high-impact weather relationship is due to the fact that strongly amplified

617 RWPs, in general, may result in blocking anticyclones (e.g., Nakamura et al. 1997; Renwick

618 and Revell 1999; Martius et al. 2013; Riboldi et al. 2019), establish atmospheric conditions

619 that are prone to strong cyclogenesis (e.g., Hoskins and Berrisford 1988), or favor PV

620 streamers and associated heavy precipitation (e.g., Martius et al. 2008). To date, the influence

621 of ET on high-impact weather in downstream regions, mediated by RWPs, has been

622 investigated via case studies. The findings of these case studies are presented here for each

623 ocean basin, but more research is needed to generalize and quantify this aspect of the ET’s

624 downstream impact and to improve the predictive capabilities for such events.

625 An enhanced frequency of blocking anticyclones downstream of ET has been reported

626 for the South Indian Ocean and the western North Pacific region (Riboldi et al. 2019). Over

627 the South Indian Ocean, such blocking anticyclones are, in general, a prerequisite for the

628 formation of southeast Australian heat waves (Quinting and Reeder 2017). The “pre–Black

629 Saturday heat wave” in 2009 caused devastating bushfires in southeast Australia and is a

630 prominent example, in which ET crucially affected the amplification of the blocking

631 anticyclone (Parker et al. 2013, 2014).

632 In the western North Pacific region, blocking anticyclones that result from the

633 amplification of the RWP during ET are suggested by Small et al. 2014 to be linked to the

634 peak in blocking frequency, found for September and October in the North Pacific. Such

635 blocking anticyclones have been associated with heat waves across western North America

636 and cold-air outbreaks and heavy precipitation in central and eastern North America (Grams

637 2011; Keller and Grams 2014; Bosart et al. 2015; Harr and Archambault 2016). An example

638 of such a development is the extratropical reintensification of Supertyphoon Nuri (2014)).

639 The midlatitude flow amplification during Nuri’s ET (cf. Fig. 1) resulted in the formation of a

640 major omega block along the west coast of North America (Fig. 12a), and several cold-air

641 outbreaks over continental North America (Bosart et al. 2015). A similar pattern has been

642 found for Typhoon Choi-Wan (2009) and other events, for which it could be shown that ET

643 did alter the location and severity of the heat-wave, cold-air outbreak and heavy precipitation,

644 whereas these events occurred in other places and to a weaker extend in a NWP forecast in

645 which the respective TCs had been removed (Fig. 12b; Grams et al. 2011; Keller and Grams

646 2014). In addition to the formation of blocking anticyclones in the eastern North Pacific, ETs

647 in the western North Pacific might contribute to the formation of Kona lows, which may

648 bring flash floods, high winds, and thunderstorms to Hawaii (Moore et al. 2008).

649 In the North Atlantic, several studies have documented that the amplification of the

650 first downstream ridge results in wave breaking and the formation of PV streamers and cut-

651 off lows over Europe (McTaggart-Cowan et al. 2007; Grams et al. 2011; Pantillon et al.

652 2013a; Grams and Blumer 2015; Pantillon et al. 2015). These PV streamers and cut-off lows

653 over Europe may affect the development of severe thunderstorms and heavy precipitation

654 (Pantillon et al. 2015; Grams and Blumer 2015), the formation of Mediterranean cyclones

655 (Pinto et al. 2001; Grams et al. 2011; Chaboureau et al. 2012; Pantillon et al. 2013a), or the

656 track and intensity of extratropical cyclones in the region (Agustí-Panareda et al. 2004; Hardy

657 et al. 2016, 2017). Furthermore, the PV streamers may also influence the development of

658 subsequent North Atlantic hurricanes (Galarneau et al. 2015).

659

660 4. Predictability

661 J2003 highlighted the often basin-wide reduction in NWP skill that may accompany

662 ET. During ET forecasters thus face the challenge to predict potential downstream high-

663 impact weather while forecast uncertainty is enhanced (purple labels and semi-transparent

664 shading in Fig. 2). As sources for the increased forecast uncertainty J2003 discussed

665 shortcomings of NWP models in the representation of moist processes and in capturing the

666 interaction between the small-scale TC and the large-scale midlatitude flow. The recent

667 findings on the important contribution of the upper-tropospheric divergent outflow to initial

668 ridge amplification, the sensitivity to phasing between the recurving TC and the upstream

669 trough, and the existence of the bifurcation point near the tip of the upstream trough support

670 these shortcomings identified by J2003. In this section we discuss the now better

671 understanding of the general intrinsic uncertainty associated with ET, and on how the

672 processes involved in ET impact forecast quality, both near the transitioning cyclone and in

673 downstream regions (sections 4a, b). Furthermore, we discuss the contribution of

674 observations to improve prediction during ET (section 4c).

675

676 a. Causes for forecast degradation in downstream regions during ET

677 1) INTERACTION WITH THE MIDLATITUDE FLOW AND PHASING

678 An incorrect representation of the phasing between the transitioning cyclone and the

679 upstream midlatitude trough in a forecast may lead to large position and intensity errors for

680 both the cyclone and the emerging midlatitude RWP (J2003). Thus, forecasts for ET are

681 particularly sensitive to the representation of the upstream midlatitude trough and the

682 transitioning cyclone (Kim and Jung 2009; Torn and Hakim 2009; Anwender et al. 2010;

683 Doyle et al. 2011; Pantillon et al. 2013b). As an example of this sensitivity, consider a

684 forecast initialized during the recurvature of Typhoon Shanshan (2006, Fig. 13). Singular

685 vector sensitivities pick up the transitioning cyclone, as well as the approaching upstream

686 trough, indicating that strongest error growth will be tied to the development of these two

687 flow features (Reynolds et al. 2009; Wu et al. 2009). The sensitivity to the upstream

688 midlatitude trough may become more dominant during and after recurvature of the

689 transitioning cyclone (Kim and Jung 2009).

690 The existence of the bifurcation point near the tip of the upstream trough (Fig. 6)

691 explains these sensitivities: Small differences on the order of 100 km in the TC track can

692 determine whether the cyclone starts to recurve in the trough-relative framework and

693 undergoes reintensification with potential downstream impact or decays (e.g., Grams et al.

694 2013b; Komaromi and Doyle 2018). Given that a track error of 100 km is quite typical in 48-

695 h forecasts over recent years (Lang et al. 2011; National Hurricane Center 2017), the trough-

696 cyclone phasing can thus be a substantial source for forecast errors.

697

698 2) DIABATIC PROCESSES

699 Another key source of forecast uncertainty is the representation of diabatic processes

700 in NWP models. The convective parameterization schemes employed in all global models,

701 and some regional models, might under-represent divergent outflow aloft, in part because the

702 divergent outflow is a grid-scale response to resolved precipitation processes, as well as

703 parameterized convective heating and moistening (Zadra et al. 2018). This error in the

704 representation of diabatic processes affects the correct simulation of initial ridge

705 amplification and RWP generation.

706 The importance of considering moist processes for predicting the midlatitude impact

707 of ET is corroborated from a singular vector perspective. Ensemble spread is significantly

708 larger in ET forecasts, for which the initial perturbations are obtained from moist singular

709 vectors4 (Buizza et al. 2003), as compared to ensemble forecasts constructed from dry

710 singular vectors (Lang et al. 2012). This corroborates the importance of considering moist

711 processes for predicting the midlatitude impact of ET. An increased ensemble spread is found

4 meaning that moist processes are considered when determining those flow features that are affected by strongest perturbation growth

712 both near of the transitioning cyclone and in downstream regions. Similar results are found

713 for increasing the horizontal resolution at which the singular vectors are calculated from

714 about 200 km to 80 km.

715 Another factor that affects perturbation growth in ensemble forecasts (assessed

716 through singular vectors and adjoints) is the baroclinicity of the midlatitude flow into which

717 the transitioning cyclone is moving, with stronger baroclinicity supporting stronger error

718 growth (Reynolds et al. 2009; Doyle et al. 2011). Furthermore, the use of explicit convection

719 may improve the forecasted cyclone track and phasing, compared to a simulation with

720 parameterized convection (Pantillon et al. 2013a). In that study, however, this effect results

721 from differences between explicit and parameterized convection along the associated RWP

722 rather than from improvements in the core of Helene. A strong sensitivity to moisture and

723 diabatic processes has also been found in previous studies (Riemer et al. 2008; Doyle et al.

724 2012; Grams and Archambault 2016), which further underscores the importance of diabatic

725 processes and associated rapid error growth downstream of ET events (e.g., Harr and Dea

726 2009; Hodyss and Hendricks 2010; Torn 2010, 2017; Archambault et al. 2015).

727 Research on the representation of diabatic processes in numerical models in the

728 context of ET has been limited to date. Investigations do exist, however, for the

729 representation of diabatic processes in extratropical cyclones (e.g., Davis et al. 1993;

730 Stoelinga et al. 1996). More recently, diabatic processes associated with mesoscale

731 convective systems and warm conveyor belts have been studied in the context of error

732 growth. Since these systems likewise modify RWPs by their upper-tropospheric divergent

733 outflow, it can be assumed that the findings of these investigations also hold true during ET.

734 The representation of diabatic processes associated with mesoscale convective

735 systems over the Great Plains has been identified as a source of short-term skill degradations

736 (“busts”) in ECMWF forecasts for Europe (Rodwell et al. 2013). In such cases, diabatic

737 processes act to decelerate the eastward progression of a synoptic-scale trough over the

738 Rocky Mountains, similar to the processes described in section 2.2. Errors in the

739 representation of these diabatic processes and their impact on the midlatitude flow may lead

740 to large phase errors in the representation of the downstream RWP. Another contribution to

741 ECMWF forecast errors stems from the representation of warm conveyor belt outflow

742 Substantial differences in PV generation, depending on the microphysical parameterization

743 used, may lead to variations in warm conveyor belt development and different positions of its

744 outflow (Joos and Wernli 2012, Joos and Forbes 2016). In addition, as warm conveyor belt

745 outflow is sensitive to environmental conditions, warm conveyor belt activity can amplify

746 initial condition error and project it on the large-scale circulation (e.g., Grams et al. 2018).

747 These differences in the outflow position and the associated tropopause structure translate

748 into differences in the midlatitude RWP and hence the synoptic development in downstream

749 regions (e.g., Dirren et al. 2003; Davies and Didone 2013; Joos and Forbes 2016; Lamberson

750 et al. 2016; Baumgart et al. 2018).

751 In summary, the skill of predicting the impact of a transitioning cyclone on the

752 midlatitude flow depends on the representation of the cyclone and the environmental features,

753 as well as their phasing. The ability of NWP models to properly forecast these flow features

754 and their phasing and interaction is, to a large extent, determined by the representation of

755 diabatic processes in the model. Small deviations in the simulation of these processes may

756 lead to a slightly different representation of these flow features, a positioning error for

757 example. Such an error may amplify rapidly due to the prevalence of a bifurcation point in

758 the steering flow, and can lead to large forecast errors in downstream regions.

759 b. Manifestation of midlatitude forecast uncertainty during ET

760 The decrease in forecast skill downstream of ET events was documented by J2003

761 (their Fig. 8) as a drop in the anomaly correlation for forecasts over the North Pacific. More

762 recent research has focused on investigating and characterizing the manifestation of this

763 forecast uncertainty, in particular in ensemble prediction systems.

764 Due to forecast uncertainty associated with the direct and downstream impact of ET,

765 ensemble standard deviation is typically increased downstream of the transitioning cyclone

766 (Fig. 14). A first increase is found at the onset of interaction between the transitioning

767 cyclone and the midlatitude flow in the direct vicinity of the cyclone, as also discussed in

768 section 6c of Part I (e.g., McTaggart-Cowan et al. 2006a; Munsell and Zhang 2014; Pantillon

769 et al. 2016; Torn et al. 2015). With increasing forecast lead time, the standard deviation

770 increases and spreads further downstream, often linked to the predicted position of the

771 transitioning cyclone. The increase in standard deviations varies for different ensemble

772 prediction systems (Fig. 14), due to differences in ensemble size, generation of initial

773 conditions and the different capabilities of models in simulating the processes associated with

774 ET (Harr et al. 2008; Anwender et al. 2008; Keller et al. 2011). Identifying which of the

775 models in Fig. 14 performs best in predicting the forecast uncertainty associated with ET,

776 hence, produces the best match between ensemble standard deviation and ensemble mean

777 RMS error, would require an investigation over many ET cases and has not been addressed

778 yet.

779 The increase in standard deviation is connected to the development of several forecast

780 scenarios across the members of one ensemble forecast, which can be revealed through a

781 cluster analysis. Besides providing information about the possible synoptic development for

782 forecasting purposes, these different forecast scenarios provide a means to study the

783 underlying physical and dynamic processes responsible for the different scenarios (Anwender

784 et al. 2008; Harr et al. 2008; Keller et al. 2011; Keller et al. 2014; Kowaleski et al. 2016).

785 Differences in phasing and the representation of diabatic processes affect the

786 predicted amplification of the first downstream ridge. Hence, largest differences in an

787 ensemble forecast are usually found in the crest and/or flanks of the first downstream ridge,

788 as depicted in Fig. 15 (e.g., Anwender et al. 2008; Harr et al. 2008; Keller et al. 2011, 2014;

789 Pantillon et al. 2016). This example has been derived from an empirical orthogonal function

790 analysis (EOF; e.g., Wilks 2011) applied to an ECMWF ensemble forecast for the potential

791 temperature at the tropopause. The strongest variability among the ensemble members,

792 described by the EOFs, is found in the flanks of the ridge (left; Fig. 15a, c) and the crests of

793 the ridge–trough couplet (right; Fig. 15a, c). Higher potential temperature is found in the

794 region of positive EOF signals (and lower potential temperature in regions of negative EOF

795 signals) for ensemble members that contribute positively to these EOF patterns. This leads to

796 an eastward tilt or shift in the downstream ridge and a stronger amplification of such

797 members (Fig. 15b) as compared to the ensemble mean. These shift- and amplitude patterns

798 (or a combination thereof) associated with the representation of the first downstream ridge

799 have been found in all the studies cited above. Hence, these patterns provide a robust signal

800 on how the forecast variability during ET affects the representation of the first downstream

801 ridge in ensemble forecasts.

802 Subsequently, the initial uncertainty then propagates farther downstream with the

803 group velocity of the midlatitude RWP in which the uncertainty is embedded (Harr et al.

804 2008; Anwender et al. 2008, 2010; Pantillon et al. 2013a; Grams et al. 2015). Without the

805 development of an RWP during ET (Quinting and Jones 2016), or for short forecast lead

806 times (Strickler et al. 2016), forecast uncertainty remains limited to the direct vicinity of the

807 transitioning cyclone, corroborating the essential role of the RWP in transmitting forecast

808 uncertainty into downstream regions. In addition to this downstream dispersion of forecast

809 uncertainty, the representation of diabatic processes within potentially developing

810 downstream cyclones may add additional uncertainty to the development of the midlatitude

811 RWP downstream of ET.

812 From a climatological standpoint, a statistically significant increase in forecast

813 uncertainty downstream is noted for western North Pacific, North Atlantic and South Indian

814 Ocean ET data (Aiyyer 2015; Quinting and Jones 2016; Torn 2017). Forecast uncertainty

815 associated with the downstream wave packet (measured in terms of normalized ensemble

816 spread, Fig. 16) tends to peak about 4–5 days after recurvature (2–3 days after ET) and

817 decreases to average conditions within 5–6 days, with the strongest increase observed in the

818 downstream troughs (Aiyyer 2015). The increase in spread results in a reduction of the

819 forecast skill horizon by about 2 days, mainly tied to the uncertainty in the amplification of

820 the first downstream ridge (Grams et al. 2015). The forecast uncertainty decreases when the

821 forecast is initialized closer to the completion of ET, with the phasing being already

822 developed. This suggests that predictive skill during ET is at least partly an initial value

823 problem (Harr et al. 2008; Anwender et al. 2008).

824 ET may not only affect midlatitude medium-range forecasts but may even deteriorate

825 the accuracy of sub-seasonal predictions. Although the transitioning cyclone was included in

826 the initial conditions (10-day average initial conditions), the National Center for

827 Environmental Prediction Climate Forecast System (NCEP-CFSv2) was not able to predict

828 the reconfiguration of the large-scale flow by Typhoon Nuri (Bosart et al. 2015).

829 In summary, the decrease in forecast skill associated with ET events initially

830 manifests as uncertainty in the prediction of the amplifying first downstream ridge.

831 Subsequently, the forecast uncertainty propagates downstream with the developing RWP,

832 while it may be further increased in the regions affected by developing downstream cyclones.

833 c. Impact of observations

834 As stated earlier, parts of the forecast uncertainty associated with ET might be tied to

835 an insufficient accuracy of initial conditions. J2003 noted that single observations may have a

836 strong influence on predicting ET and its downstream impact. They suggested further

837 investigation on how existing observations can be used in an optimum way and to exploit

838 new observational capabilities. Since J2003 appeared, additional targeted observations

839 (Majumdar 2016) have been gathered in field experiments, such as The Observing System

840 Research and Predictability Experiment (THORPEX; Parsons et al. 2017) Pacific Asian

841 Regional Campaign 2008 (T–PARC). These additional observations provided more detailed

842 insights in the processes of ET, as well as on the benefit of observations for ET forecasts.

843 The impact of observations on ET forecasts in the North Atlantic has been tested in

844 data-denial experiments with the ECMWF Integrated Forecasting System. In these

845 experiments, observations were removed either in sensitive regions (identified via singular

846 vectors as regions where errors grow most quickly) near ET, in sensitive regions in the

847 midlatitudes, or in randomly chosen regions (Cardinali et al. 2007; Chen and Pan 2010;

848 Anwender et al. 2012). Removal of observations in sensitive regions during an ET event

849 results in forecast degradations that are six times larger than degradations produced by

850 removal of observations that are randomly selected (Cardinali et al. 2007; Chen and Pan

851 2010). Compared to denying data in extratropical sensitive regions (SVout, Fig. 17a),

852 removing observations near ET (ETout, Fig. 17b), however, led to about the same magnitude

853 of forecast degradation (87% vs. 83% of degraded forecasts), measured in terms of the root

854 mean square difference for total energy (Anwender et al. 2012). Denying data near the

855 transitioning cyclone is, on average, more impactful than denying data in extratropical

856 sensitive regions for medium-range forecasts (Anwender et al. 2012). This also implies that

857 poorly observed transitioning cyclones yield, on average, larger forecast degradations than

858 unobserved extratropical sensitive regions. After completion of ET, however, larger

859 degradations are associated with denying data in extratropical sensitive regions.

860 Experiments with dropsonde data, gathered during T–PARC for the ET of Typhoons

861 Sinlaku and Jangmi corroborate these findings. Observations taken after recurvature toward

862 the completion of ET do not result in significant forecast improvements in the midlatitudes,

863 although the observations were taken in sensitive regions (Weissmann et al. 2011). In

864 contrast, significant forecast improvements, also in the midlatitudes, are found for

865 observations that were taken near the transitioning cyclone earlier in its life cycle.

866 Weissmann et al. (2011) explained this by the fact that observations during early stages of the

867 TC’s life cycle are typically taken in data-sparse regions, although the western North Pacific

868 midlatitudes have denser observation coverage and are thus better represented in the analysis

869 anyway.

870 In summary, a number of studies highlight the enhanced sensitivity to initial

871 conditions and increased potential for error growth during ET, both near the transitioning

872 cyclone and in the downstream midlatitudes. Although the results seem to be quite robust,

873 they are mainly based on case studies and case-to-case variability still needs to be assessed in

874 a systematic manner.

875

876 5. Conclusions and outlook

877 The ET of a tropical cyclone may modify the midlatitude flow and result in a basin-

878 wide reduction in forecast skill of NWP models, as summarized in the review paper by

879 J2003. Since J2003 appeared, the ET research community has worked toward a better

880 understanding of the interaction between a TC and the midlatitude flow. It has been

881 demonstrated that ET impacts the midlatitude flow such that a midlatitude RWP is initiated or

882 amplified. This RWP then spreads the impact of the ET downstream over a large

883 geographical region.

884 The first stage of this RWP amplification involves enhanced ridge building

885 immediately downstream of the transitioning cyclone, often accompanied by the development

886 of a jet streak, and is considered a direct impact (Fig. 2, red labels; section 2a). The

887 amplification of this ridge is a consequence of the favorable superposition of the dry

888 dynamics of the growing baroclinic wave and the diabatically enhanced upper-tropospheric

889 outflow associated with latent heat release. During the early stage of ET, this latent heat

890 release occurs primarily with the deep convection near the center of the cyclone. Later during

891 ET, the latent heat release that enhances the upper-tropospheric divergent outflow is

892 primarily tied to warm and moist air masses ascending slantwise along the baroclinic zone.

893 These air masses are advected poleward by the cyclonic circulation of the transitioning

894 cyclone and ascend as a warm conveyor belt along the baroclinic zone. In addition, the

895 cyclonic circulation of the transitioning cyclone advects anticyclonic PV into the ridge,

896 supporting its further amplification.

897 This direct impact of ET crucially depends on the phasing between the transitioning

898 cyclone and the developing or already existing midlatitude wave pattern (section 2b). The

899 relative position of the transitioning cyclone to a bifurcation point near the tip of the trough in

900 the trough-relative frame of reference determines whether the transitioning cyclone enters a

901 region favorable for cyclone development. The most pronounced impact, in terms of ridge

902 amplification and downstream development, can be expected when the cyclone is located

903 ahead of and moves in phase with an upstream midlatitude trough. In such a phase-locked

904 configuration the cyclone is able to reintensify and continuously amplify the downstream

905 ridge–trough couplet. In other words, the transitioning cyclone acts as a local wave-maker

906 and the ET process can be interpreted as a resonant interaction.

907 The initial ridge building and the direct impact of ET may further be supported by the

908 so-called preconditioning, introduced in section 2c. Prior to ET, weather systems like warm

909 conveyor belts, predecessor rain events, or diabatic Rossby waves may precondition the

910 midlatitude flow. Through the poleward advection of warm and moist air masses or ridge

911 building ahead of ET, these systems create a midlatitude flow environment that supports a

912 potential reintensification of the transitioning cyclone and the initiation of a highly-amplified

913 RWP. The notion of preconditioning is a very recent one and its general importance for the

914 dynamics and predictability of ET needs further assessment.

915 To date, understanding of the direct impact of ET on the midlatitude flow is mostly

916 based on modeling studies or compositing approaches using model-based (re)analysis data.

917 Observational data providing information about the dynamics of the interaction and the role

918 of diabatic processes could elucidate how well model-derived results agree with observed ET

919 systems and how operational NWP systems perform in capturing the development. Data from

920 the recent North Atlantic Wave Guide and Downstream Impact Experiment (NAWDEX;

921 Schäfler et al. 2018), which took place in autumn 2016, could provide such valuable

922 observations. NAWDEX featured a unique set of high-resolution measurements for the ET of

923 Hurricane Karl both in clouds and cloud-free regions and that are not assimilated into models

924 on a routine basis. This provides the opportunity for studying diabatic processes during the

925 interaction of a transitioning cyclone with the midlatitude flow and the representation thereof

926 in models at a level of detail previously unavailable. The new generation of high-resolution

927 multi-spectral imagers, sounders, or scatterometers aboard satellites of the Global

928 Precipitation Measurement (GPM) mission, the Geostationary Operational Environmental

929 Satellite–R Series (GOES–R) or the HIMAWARI 8, are examples of observational resources

930 that may help improve the representation of ET in models on a routine basis.

931 The direct impact of ET propagates downstream, following RWP dynamics (Fig. 2,

932 blue labels; section 3). Depending on the phasing and the intensity of the interaction between

933 the transitioning cyclone and the midlatitude flow, the cyclone supplies additional eddy

934 kinetic energy to the midlatitude flow, supporting the amplification and downstream

935 propagation of the RWP, as described in section 3a.1. The development of the RWP,

936 however, also depends on the configuration of the midlatitude flow (section 3a). A weak jet is

937 susceptible to stronger RWP amplification as compared to a strong jet. Furthermore, the

938 availability of moisture in the downstream region, and the potential formation of downstream

939 cyclones and warm conveyor belts along the eastern flank of the downstream trough

940 influences the development of the downstream RWP. The diabatically enhanced upper-

941 tropospheric divergent outflow of such weather systems may support the further

942 amplification of the RWP in downstream regions through moist-baroclinic growth. Hence,

943 although a weak jet can be expected to yield a high-amplitude wave pattern downstream of

944 ET, the implied weak baroclinicity limits the development of downstream cyclones and, thus,

945 the positive feedback from moist-baroclinic growth. This raises the question about the

946 characteristics of an optimal jet that maximizes the downstream response. Beside the

947 amplification of a downstream RWP, ET may also initiate Rossby wave breaking, and thus

948 ultimately a de-amplification of the downstream midlatitude flow.

949 From a climatological perspective, ET events in the western North Pacific and the

950 South Indian Ocean are accompanied by an enhanced RWP activity in downstream regions

951 (section 3a.2). The findings for the North Atlantic are less clear, which might be tied to the

952 typically short and weak jet in this region, which is susceptive to Rossby wave breaking. Up

953 to now, the occurrence of Rossby wave breaking during ET has only been considered in a

954 few studies and its occurrence frequency has not been determined yet. Hence, a better

955 understanding of the effect of ET on Rossby wave breaking could help clarify the

956 climatological impact of ET in the North Atlantic.

957 The amplification of an RWP during ET often results in the development of high-

958 impact weather in downstream regions (section 3b). Heat waves and cold-air outbreaks may

959 develop when the amplified troughs and ridges become stationary. Strong cyclones, deep

960 convection, and heavy precipitation events developing on the eastern flank of the downstream

961 troughs may also be influenced by the downstream impact of ET. Thereby, ET may not

962 primarily be the trigger for the occurrence of such events, but it may at least alter their

963 position and intensity. A climatological assessment and a quantification of the ET’s

964 contribution to the formation of such weather events that goes beyond case studies would

965 complement current knowledge and could help to enhance prediction of such events.

966 The downstream impact of ET often leads to a degradation of predictability in

967 downstream regions (Fig. 2, purple labels; section 4), resulting in increased forecast

968 uncertainty particularly in the medium forecast range. As discussed in section 4a, the primary

969 sources for this increase in forecast uncertainty are shortcomings in the representation of

970 diabatic processes in numerical models and a high sensitivity of the subsequent evolution to

971 small changes in phasing. An insufficient representation of latent heat release or the position

972 of the cyclone with respect to the bifurcation point results in rapidly growing forecast errors

973 associated with the amplification of the first downstream ridge. These errors may then

974 propagate downstream with an RWP and may further grow due to nonlinearities in RWP

975 dynamics, in particular the contribution of diabatic processes to ridge amplification in

976 downstream regions. Typically, the strongest forecast uncertainty is found for the position

977 and amplitude of the downstream midlatitude RWP, which translates into uncertainty in the

978 geographical location and strength of associated weather systems and their impacts, as

979 discussed in section 4b. Observations may have a beneficial impact on forecast quality during

980 ET, as reviewed in section 4c. In particular, those observations taken near the transitioning

981 cyclone during early stages of ET reduce forecast errors. Observations taken in the

982 surrounding midlatitudes become equally important during later stages of the interaction

983 between the transitioning cyclone and the midlatitude flow. Given the possible occurrence of

984 high-impact weather downstream of ET, improved predictive capabilities during ET have

985 been, and still are, a major goal of ET research and require further work. A climatological

986 assessment is needed to better describe forecast degradation experienced during ET, for

987 example by using re-forecast datasets. This assessment should also consider the role of the

988 midlatitude flow configuration in propagating forecast errors into downstream regions and

989 how this differs across the ocean basins. The development of high-resolution, convection

990 permitting NWP models could provide a useful tool for capturing the diabatic processes

991 during ET. Moreover, this approach, when embedded in global models (e.g., via local grid

992 refinement), may also reduce forecast errors in the medium-range. Inherent uncertainty

993 associated with the representation of diabatic processes, however, may limit the predictability

994 of the downstream impact of ET. This motivates the need for improved probabilistic

995 prediction of downstream impacts associated with ET using ensemble prediction systems, and

996 additional research on the representation of model errors arising from diabatic processes

997 using techniques such as stochastic physics.

998 The research summarized in this review primarily focused on assessing the impact of

999 ET on the short- to medium-range forecast horizon. Preliminary results reveal a statistically

1000 significant correlation between monthly-mean values of selected teleconnection indices and

1001 ET event counts, as well as significant departures from climatology on the sub-seasonal to

1002 seasonal time scale in atmospheric field composites associated with ET events. The potential

1003 impact of ET on time scales beyond the medium-range calls for a further investigation of the

1004 factors that may impact ET on sub-seasonal to seasonal time scales, including persistent flow

1005 regimes and teleconnections, and on how this influences the predictability of ET events on

1006 these time scales. Attention should also be paid to ET-related modifications of the midlatitude

1007 flow configuration, including enhancements to poleward moisture transport, which may

1008 impact both the occurrence frequency and predictability of sub-seasonal regimes on basin- to

1009 hemispheric-length scales. The Sub-seasonal to Seasonal Project data base

1010 (www.s2sprediction.net; Vitart et al. 2017), which provides access to sub-seasonal to

1011 seasonal forecasts from eleven operational centers, could be a valuable resource for such

1012 investigations.

1013 On still longer time scales, the influence of a warming climate on the downstream

1014 impact of ET, in particular, is another aspect that deserves attention. Considering the

1015 important contribution of diabatic processes to the amplification of the downstream

1016 midlatitude RWP during ET, the increasing availability of water vapor in a changing climate

1017 suggests that an associated increase in latent heat release may strengthen impacts such as

1018 downstream ridge amplification. Research is needed to address this question and to explore

1019 how this extreme form of tropical–extratropical interaction could change with a changing

1020 climate.

1021

1022

1023 Acknowledgement

1024 The paper was jointly written by Julia H. Keller, Christian M. Grams, and Michael

1025 Riemer, based on contributions from all co-authors. We thank John Gyakum for his support

1026 in initiating this work, as well as Clark Evans for comments that helped to improve the

1027 manuscript. Well-thought-out and thorough comments by five anonymous reviewers and the

1028 editor David Schultz on a previous version of the manuscript provided very helpful guidance

1029 for improving the content and tangibility of this review. JHK made part of her contribution

1030 while she was employed at Deutscher Wetterdienst. The contribution by CMG was made

1031 while he held a Swiss National Science Foundation (SNSF) Ambizione fellowship under

1032 grant PZ00P2_148177/1 and finished under grant VH-NG-1243 of the Helmholtz

1033 Association. MR acknowledges support from the subproject “A4: Evolution and

1034 predictability of storm structure during extratropical transition of tropical cyclones” of the

1035 Transregional Collaborative Research Center SFB/TRR 165 “Waves to Weather” program

1036 funded by the German Science Foundation (DFG). Parts of the results presented in this

1037 review have been produced within the German PANDOWAE (FOR896) project, funded by

1038 DFG as a contribution to the WMO World Weather Research Programme THORPEX. Co-

1039 authors acknowledge the following support: Bosart – support from NSF (AGS-1240502 and

1040 AGS-1355960); Doyle and Reynolds – support of the Chief of Naval Research through the

1041 NRL Base Program PE 0601153N, and the ONR PE 0602435N; Quinting – support from the

1042 Australian Research Council Centre of Excellence for Climate System Science

1043 (CE110001028); Torn – support from NSF (ATM-1461753). Zhang – support from ONR

1044 (Grant N000140910526). This review was partly initiated at the World Meteorological

1045 Organization’s Eighth International Workshop on Tropical Cyclones, in 2014.

1046

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1455 Schumacher, R. S., T. J. Galarneau, and L. F. Bosart, 2011: Distant effects of a recurving 1456 tropical cyclone on rainfall in a midlatitude convective system: A high-impact 1457 predecessor rain event. Mon. Wea. Rev., 139, 650–667, 1458 https://doi.org/10.1175/2010MWR3453.1.

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1483 Teubler, F., and M. Riemer, 2016: Dynamics of Rossby wave packets in a quantitative 1484 potential vorticity–potential temperature framework. J. Atmos. Sci., 73, 1063–1081, 1485 https://doi.org/10.1175/JAS-D-15-0162.1.

1486 Thorncroft, C. D., B. J. Hoskins, and M. E. McIntyre, 1993: Two paradigms of baroclinic- 1487 wave life-cycle behaviour. Quart. J. Roy. Meteor. Soc., 119, 17–55, 1488 https://doi.org/10.1002/qj.49711950903.

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1490 extratropical transitions determined using ensemble-based sensitivity analysis. Mon. 1491 Wea. Rev., 137, 3388–3406, https://doi.org/10.1175/2009MWR2879.1.

1492 ——, 2010: Diagnosis of the downstream ridging associated with extratropical transition 1493 using short-term ensemble forecasts. J. Atmos. Sci., 67, 817–833, 1494 https://doi.org/10.1175/2009JAS3093.1.

1495 ——, and G. J. Hakim, 2015: Comparison of wave packets associated with extratropical 1496 transition and winter cyclones. Mon. Wea. Rev., 143, 1782–1803, 1497 https://doi.org/10.1175/MWR-D-14-00006.1.

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1501 ——, 2017: A comparison of the downstream predictability associated with ET and 1502 baroclinic cyclones. Mon. Wea. Rev., 145, 4651–4672, https://doi.org/10.1175/MWR- 1503 D-17-0083.1.

1504 Vitart, F., and Coauthors, 2016: The subseasonal to seasonal (S2S) prediction project 1505 database. Bull. Amer. Meteor. Soc., 98, 163–173, https://doi.org/10.1175/BAMS-D- 1506 16-0017.1.

1507 Wandishin, M. S., J. W. Nielsen-Gammon, and D. Keyser, 2000: A potential vorticity 1508 diagnostic approach to upper-level frontogenesis within a developing baroclinic wave. 1509 J. Atmos. Sci., 57, 3918–3938, https://doi.org/10.1175/1520- 1510 0469(2001)058<3918:APVDAT>2.0.CO;2.

1511 Wang, Y., Y. Wang, and H. Fudeyasu, 2009: The role of Typhoon Songda (2004) in 1512 producing distantly located heavy rainfall in Japan. Mon. Wea. Rev., 137, 3699–3716, 1513 https://doi.org/10.1175/2009MWR2933.1.

1514 Weissmann, M., F. Harnisch, C. Wu, P. Lin, Y. Ohta, K. Yamashita, Y. Kim, E. Jeon, T. 1515 Nakazawa, and S. Aberson, 2011: The influence of assimilating dropsonde data on 1516 typhoon track and midlatitude forecasts. Mon. Wea. Rev., 139, 908–920, 1517 https://doi.org/10.1175/2010MWR3377.1.

1518 Wilks, 2011: Statistical Methods in the Atmospheric Sciences. Academic Press, 704pp.

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1522 Wu, C.-C., S.-G. Chen, J.-H. Chen, K.-H. Chou, and P.-H. Lin, 2009: Interaction of Typhoon 1523 Shanshan (2006) with the midlatitude trough from both adjoint-derived sensitivity 1524 steering vector and potential vorticity perspectives. Mon. Wea. Rev., 137, 852–862, 1525 https://doi.org/10.1175/2008MWR2585.1.

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1532

1533 LIST OF FIGURES

1534 Fig. 1: Downstream development during the ET of Supertyphoon Nuri (2014) at a) 0600 UTC 1535 5 Nov 2014, b) 1800 UTC 6 Nov 2014, c) 1800 UTC 7 Nov 2014, d) 0000 UTC 9 1536 Nov 2014. Panels shows IR Gridsat clouds (brightness temperature in °C as in color 1537 bar, Knapp et al. 2011), dynamical tropopause (2 PVU on the 330 K isentropic 1538 surface, red contour) and wind speed maxima highlighting jet streaks (wind speed on 1539 the 330 K isentropic surface as semi-transparent shading in yellow for 55 m s-1 and 1540 orange for 65 m s-1). Dynamical tropopause and wind speed are taken from ERA- 1541 Interim reanalyses. The TC symbol indicates the position of the transitioning cyclone 1542 (encircled during extratropical stage), the “L” the position of the downstream 1543 cyclone………..……………………………………………………………………….……..63

1544 Fig. 2: Overview of synoptic features and processes involved in northern hemispheric ET. 1545 Labels indicate relevant processes, starting with the section in which they are 1546 discussed. Transitioning cyclone presented by the black-and white pictogram. The 1547 dark red line indicates axis of the undulating midlatitude jet stream separating 1548 stratospheric high PV air (dark gray, poleward) and tropospheric low PV air (light 1549 gray, equatorward), with the dashed line indicating an alternate configuration. The 1550 light red ellipse denotes the jet streak. The purple, semi-transparent area signifies the 1551 forecast uncertainty for the downstream flow. The downstream cyclone is indicated by 1552 the “L”-symbol and its associated fronts. The airplane symbol represents observation 1553 systems used for ET reconnaissance………………………………………………………...64

1554 Fig. 3: PV anomalies involved in ET and their respective contribution to RWP modification 1555 near ET. a) 2D and b) 3D schematic of the different flow features associated with ET: 1556 i) cyclonic circulation (thin orange arrows) associated with the cyclonic PV tower 1557 (orange TC symbol in a), gray column in b)) and with low-level cyclonic PV 1558 anomalies (small gray clouds and circulation) at the developing warm front, ii) 1559 anticyclonic circulation associated with the anticyclonic PV anomaly of the upper- 1560 tropospheric outflow (white/gray cloud symbol in a)/b) and thin blue arrows) iii) the 1561 upper-tropospheric divergent outflow associated with latent heat release below (thin 1562 green arrows). The advective contribution of these flow features to the amplification 1563 of the downstream ridge and trough are given by the bold arrows in a). The red 1564 contour and shading in a) and b) are similar to Fig. 2. The lower layer in b) 1565 exemplifies the developing warm front with colder air masses toward the pole. c) 1566 Contributions to ridge amplification from the advection of potential temperature on 1567 the dynamic tropopause (a surrogate for upper-level PV advection) by the several flow 1568 features, integrated over a ridge for a 72-h forecast of an idealized scenario (in K km2 1569 s-1, colors as in a) and b)). Figure 8 from Riemer and Jones (2010), with 1570 modifications…………………………………………………………………………….…..65

1571 Fig. 4: Jet streak formation during ET from an energetics perspective. a) Schematic

1572 representation, showing midlatitude jet (black line), developing Ke maxima (jet streak)

1573 (gray ellipses), baroclinic conversion of Ke (clouds), ageostrophic geopotential flux 1574 (orange arrow) and its divergence (blue ellipses) and convergence (red ellipses). The 1575 black box approximates the area which is captured by panels b) and c). b), c) TC-

1576 relative composite of Ke budget for western North Pacific ETs, based on ERA-Interim 1577 reanalysis for 1980–2010 (after Quinting and Jones 2010, their Fig. 12a,b):vertically 5 −2 2 −2 1578 integrated Ke (shaded in 10 J m ), 200-hPa geopotential (contours every 200 m s , 1579 thick black contour illustrates 11 800 m2 s−2), and b) ageostrophic geopotential flux 1580 (vectors, reference vector in 106 W m−1; divergence as colored contours every 8 W 1581 m−2, divergence in blue, 0 W m-2 omitted) and c) vertically integrated baroclinic −2 1582 conversion of Ke (red contours every 8 W m ). Composites are shown relative to the 1583 mean TC position……………………………………………………………………...... 66

1584 Fig. 5: Lagrangian trajectories for the ET of Typhoon Jangmi showing ridge building and jet 1585 streak formation at (a) 0000 UTC 30 Sep 2008 and (b) 1200 UTC 1 Oct 2008. Shown 1586 are the 1.5 PVU PV surface (blue shading), 320 K surface of equivalent potential 1587 temperature (transparent gray shading), representing 3D baroclinic zone; |V| = 60 m s-1 1588 (green shading), highlighting upper-level midlatitude jet; potential temperature at 990- 1589 hPa (shading at bottom, brown colors >300 K, green ≈ 290 K) and geopotential height 1590 at 990-hPa (black contours, every 250 dam). Paths of representative trajectories 1591 (starting (a) 1200 UTC 28 Sep 2008 → ending 1200 UTC 30 Sep 2008 and (b) 1200 1592 UTC 30 Sep 2008 → ending 1200 UTC 2 Oct 2008) colored by PV of air parcel 1593 moving along trajectory. Anticyclonic-PV air (PV<0.6 PVU) gray shades, cyclonic 1594 PV values (PV>0.6 PVU) in red shades (see legend on bottom right). Figure 5 from 1595 Grams et al. (2013a)…………………………………………………………………………67

1596 Fig. 6: Idealized scenario for steering flow topology during ET. Midtropospheric geopotential 1597 to illustrate the midlatitude wave pattern and winds in the frame of reference moving 1598 with this pattern (at 620-hPa, shaded and arrows, respectively) and different tracks of 1599 transitioning cyclones from sensitivity experiments (thin black lines). The bifurcation 1600 points associated with the upstream trough and the developing downstream ridge, 1601 deciding upon the development of a “no ET”, a “NW pattern” or a “NE pattern” 1602 scenario, are marked by the dot and the cross, respectively. Dashed contours depict the 1603 streamlines that emanate from the bifurcation points. Figure 11 from Riemer and 1604 Jones (2014)…………………………………………………………………………………68

1605 Fig. 7: Three-dimensional schematic depiction of the preconditioning stage with a PRE 1606 during western North Pacific ET Anticyclonic PV air in the upper-level outflow of a 1607 TC and associated PRE as blue shading in the upper panel, jet streak as green shading 1608 and 200-hPa waveguide as red contour separating high PV air (>3 PVU; orange 1609 shading) from lower PV air (<3 PVU; unshaded). Mid-level baroclinic zone as blue 1610 tilted surface. Trajectories of rapidly ascending air parcels as blue-red-blue lines, 1611 reflecting the diabatic PV modification of the parcels from low to high to low PVU, 1612 respectively. Mean sea level pressure (gray contours; every 8 hPa) and equivalent 1613 potential temperature (violet contours; 320 and 330 K) are indicated in lower panel. 1614 Figure 11 from Grams and Archambault (2016)…………………………………………….69

1615 Fig. 8: Conceptual model of the key synoptic-scale features during the occurrence of a PRE. 1616 Shown is the 200-hPa geopotential height (contours), the upper-tropospheric jet streak 1617 (gray shading, the midlatitude baroclinic zone (surface frontal structure with 925-hPa 1618 streamlines associated with warm (red) and cold (blue) air advection. and lower-

1619 tropospheric moisture flux along the eastern side of the recurving TC and the western 1620 flank of a subtropical high. The formation of the PRE is indicated by the yellow-green 1621 ellipse (Fig. 20a of Moore et al. (2013))…………………………………………………….70

1622 Fig. 9: Downstream development during the ET of Typhoon Nabi (2005)). All panels: 5 -2 1623 Vertically integrated Ke (shaded; 10 J m ), 500-hPa heights (light gray contours, 60 1624 m intervals). a) vertically integrated divergence (dashed) and convergence (solid) of -1 1625 the ageostrophic geopotential flux (contours; W m ); b) vertically integrated total Ke 1626 flux vectors (advection + dispersion; reference vector in lower right, 105 W m-1), and 1627 500-hPa heights; c) baroclinic conversion (contours; W m-1). Figures 3c, a, b from 1628 Harr and Dea (2009)………………………………………………………………...………71

1629 Fig. 10: Interaction between a transitioning cyclone and the midlatitude flow expressed as 1630 advection of low PV air by the upper-level divergent outflow. a) Idealized 1631 representation of ridge amplification and jet streak intensification. Vectors represent 1632 the upper-tropospheric divergent outflow associated with the transitioning cyclone. 1633 Shading denotes anticyclonic PV advection by the divergent wind (Archambault et al. 1634 2013, their Fig. 4). Composite analyses of objectively defined b) strong and c) weak 1635 interactions at the time of maximum interaction. 500-hPa ascent (green, every 2x10-3 1636 hPa s-1, negative values only), total-column precipitable water (shaded according to 1637 grayscale, mm), 200-hPa PV (blue, every 1 PVU), irrotational wind (vectors, > 2 m s- 1638 1; purple vectors, > 8 m s-1), negative PV advection by the irrotational wind (dashed 1639 red, every 2 PVU day-1 starting at -2 PVU day-1), and total wind speed (shaded 1640 according to color bar, m s-1). The star denotes point of maximum interaction. The TC 1641 symbol denotes composite TC position. Downstream development of d) strong and e) 1642 weak interactions 36-h after time of maximum interaction as represented by 250-hPa 1643 meridional wind anomalies (shaded, m s-1; enclosed by black contours where 1644 significant at the 99% confidence level), PV (blue, every 1 PVU), and irrotational 1645 wind (vectors, >2 m s-1). Figures 8a, 8b, 5d and 6d, from Archambault et al. 2015………...72

1646 Fig. 11: Recurvature-relative composites of enhanced RWP frequency anomaly (shaded in %) 1647 (a) relative to June to November climatology for western North Pacific transitioning 1648 cyclones and (b) relative to December to April climatology for South Indian Ocean 1649 transitioning cyclones. Statistical significance at 95% confidence level hatched, mean 1650 track given by black like. Mean and range of recurvature longitudes indicated by white 1651 star and black bar, respectively. Data smoothed with a Gaussian filter. Figures 3a and 1652 5a from Quinting and Jones (2016)………………………………………………………….73

1653 Fig. 12: a) Illustration of omega block and high-impact weather downstream of Typhoon Nuri 1654 (2014) after Bosart et al. (2015). b) Downstream impact of Typhoon Choi-Wan 1655 (2009), based on PV surgery experiments where the storm has been removed from 1656 initial conditions (Keller and Grams 2014). Black items represent midlatitude flow 1657 features in the presence of ET, red items the evolution if ET influences were not 1658 present: 300-hPa geopotential height contour indicates upper-level waveguide (950- 1659 dam at 0000 UTC 22 Sep 2009), Arrows indicate shift of high-impact weather 1660 (precipitation, sunny and hot conditions, cold conditions) with symbol size 1661 representing magnitude……………………………………………………………………..74

1662 Fig. 13: Vertically integrated initial-time dry total energy singular vector sensitivities (shaded; 1663 values in color bar: J kg−1) with 500-hPa streamlines for TC Shanshan from 0000 UTC 1664 15 Sep 2006. Figure 3c from Reynolds et al. (2009).

1665 Fig.14: Increase in standard deviation of the 500-hPa geopotential height (in dam) in the 1666 Australian, the Canadian, the ECMWF and the THORPEX Interactive Grand Global 1667 Ensemble (TIGGE, Swinbank et al., 2015) Multimodel EPS for the ET of Hurricane 1668 Ike. Forecast initialized 0000 UTC 10 Sep 2008. TC position in ensemble members is 1669 marked by the black dots, best track 1670 position at ET time by the red dot. Figure 1 from Keller et al. (2011)……………………...75

1671 Fig. 15: a) Schematic of the shift and amplitude pattern of ensemble forecast uncertainty 1672 derived from the first two EOFs (thin solid and dashed lines) of potential temperature 1673 on the dynamical tropopause in ensemble members. The thick black line represents the 1674 strong potential temperature gradient on the dynamic tropopause in the midlatitudes. 1675 b) Synoptic patterns (shape of ridge) that result from the contribution to the variability 1676 patterns. c) EOF 1 (left, contours) and 2 (right, contours) for potential temperature at 2 1677 PVU (shaded in K) in an ECMWF ensemble forecast for Typhoon Maemi (2003). 1678 Values indicate percentage of total uncertainty captured by the respective EOF. 1679 Figures 9 and 10a, b from Anwender et al. (2008)………………………………………….76

1680 Fig. 16: Normalized ensemble spread of 500-hPa geopotential height as a function of forecast 1681 hour for NOAA’s 2nd generation global reforecasts initialized at recurvature time. Data 1682 covers all western North Pacific tropical cyclones from 1985-2013. Solid line shows 1683 mean, dashed line median, and the shaded region 25th–75th percentile range of the 1684 distribution. Statistically signiﬁcant values of the mean shown as thicker line. 1685 Hurricane symbol marks the time of recurvature, circle the median time of ET, and the 1686 thin vertical line the peak spread. Figure 3a from Aiyyer (2015)…………………………...77

1687 Fig. 17: Forecast degradation due to data denied in a) extratropical sensitive regions (SVout) 1688 and b) vicinity of the storm (ETout), expressed as root mean square difference total 1689 energy. Box-and-whiskers plot of the percentage impact over Europe (35–75° N, 10– 1690 30° E) for all denial cases. The 25 and the 75 quantile, median, and most extreme 1691 outliers are indicated by the box edges, red line and whiskers, respectively. Vertical 1692 dashed lines separate ET cases, vertical dotted lines indicate ET times. Adapted from 1693 Fig. 5 of Anwender et al. (2012)…………………………………………………………….78

1694

1695

1696 Fig. 1: Downstream development during the ET of Supertyphoon Nuri (2014) at a) 0600 UTC 1697 5 Nov 2014, b) 1800 UTC 6 Nov 2014, c) 1800 UTC 7 Nov 2014, d) 0000 UTC 9 Nov 2014. Panels 1698 shows IR Gridsat clouds (brightness temperature in °C as in color bar, Knapp et al. 2011), dynamical 1699 tropopause (2 PVU on the 330 K isentropic surface, red contour) and wind speed maxima highlighting 1700 jet streaks (wind speed on the 330 K isentropic surface as semi-transparent shading in yellow for 55 m 1701 s-1 and orange for 65 m s-1). Dynamical tropopause and wind speed are taken from ERA-Interim 1702 reanalyses. The TC symbol indicates the position of the transitioning cyclone (encircled during 1703 extratropical stage), the “L” the position of the downstream cyclone.

1704

1705 Fig. 2: Overview of synoptic features and processes involved in northern hemispheric ET. 1706 Labels indicate relevant processes, starting with the section in which they are discussed. Transitioning 1707 cyclone presented by the black-and white pictogram. The dark red line indicates axis of the undulating 1708 midlatitude jet stream separating stratospheric high PV air (dark gray, poleward) and tropospheric low 1709 PV air (light gray, equatorward), with the dashed line indicating an alternate configuration. The light 1710 red ellipse denotes the jet streak. The purple, semi-transparent area signifies the forecast uncertainty 1711 for the downstream flow. The downstream cyclone is indicated by the “L”-symbol and its associated 1712 fronts. The airplane symbol represents observation systems used for ET reconnaissance.

1713

1714

1715 Fig. 3: PV anomalies involved in ET and their respective contribution to RWP modification 1716 near ET. a) 2D and b) 3D schematic of the different flow features associated with ET: i) cyclonic 1717 circulation (thin orange arrows) associated with the cyclonic PV tower (orange TC symbol in a), gray 1718 column in b)) and with low-level cyclonic PV anomalies (small gray clouds and circulation) at the 1719 developing warm front, ii) anticyclonic circulation associated with the anticyclonic PV anomaly of the 1720 upper-tropospheric outflow (white/gray cloud symbol in a)/b) and thin blue arrows) iii) the upper- 1721 tropospheric divergent outflow associated with latent heat release below (thin green arrows). The 1722 advective contribution of these flow features to the amplification of the downstream ridge and trough 1723 are given by the bold arrows in a). The red contour and shading in a) and b) are similar to Fig. 2. The 1724 lower layer in b) exemplifies the developing warm front with colder air masses toward the pole. c) 1725 Contributions to ridge amplification from the advection of potential temperature on the dynamic 1726 tropopause (a surrogate for upper-level PV advection) by the several flow features, integrated over a 1727 ridge for a 72-h forecast of an idealized scenario (in K km2 s-1, colors as in a) and b)). Figure 8 from 1728 Riemer and Jones (2010), with modifications.

1729

1730 Fig. 4: Jet streak formation during ET from an energetics perspective. a) Schematic

1731 representation, showing midlatitude jet (black line), developing Ke maxima (jet streak) (gray ellipses),

1732 baroclinic conversion of Ke (clouds), ageostrophic geopotential flux (orange arrow) and its divergence 1733 (blue ellipses) and convergence (red ellipses). The black box approximates the area which is captured

1734 by panels b) and c). b), c) TC-relative composite of Ke budget for western North Pacific ETs, based on 1735 ERA-Interim reanalysis for 1980–2010 (after Quinting and Jones 2010, their Fig. 12a,b): vertically 5 −2 2 −2 1736 integrated Ke (shaded in 10 J m ), 200-hPa geopotential (contours every 200 m s , thick black 1737 contour illustrates 11 800 m2 s−2), and b) ageostrophic geopotential flux (vectors, reference vector in 1738 106 W m−1; divergence as colored contours every 8 W m−2, divergence in blue, 0 W m-2 omitted) and 1739 c) vertically integrated baroclinic conversion of Ke (red contours every 8 W m−2). Composites are 1740 shown relative to the mean TC position.

1741

1742 Fig. 5: Lagrangian trajectories for the ET of Typhoon Jangmi showing ridge building and jet 1743 streak formation at (a) 0000 UTC 30 Sep 2008 and (b) 1200 UTC 1 Oct 2008. Shown are the 1.5 PVU 1744 PV surface (blue shading), 320 K surface of equivalent potential temperature (transparent gray 1745 shading), representing 3D baroclinic zone; |V| = 60 m s-1 (green shading), highlighting upper-level 1746 midlatitude jet; potential temperature at 990-hPa (shading at bottom, brown colors >300 K, green ≈ 1747 290 K) and geopotential height at 990-hPa (black contours, every 25 dam). Paths of representative 1748 trajectories (starting (a) 1200 UTC 28 Sep 2008 → ending 1200 UTC 30 Sep 2008 and (b) 1200 UTC 1749 30 Sep 2008 → ending 1200 UTC 2 Oct 2008) colored by PV of air parcel moving along trajectory. 1750 Anticyclonic-PV air (PV<0.6 PVU) gray shades, cyclonic PV values (PV>0.6 PVU) in red shades (see 1751 legend on bottom right). Figure 5 from Grams et al. (2013a).

1752

1753 Fig. 6: Idealized scenario for steering flow topology during ET. Midtropospheric geopotential 1754 to illustrate the midlatitude wave pattern and winds in the frame of reference moving with this pattern 1755 (at 620-hPa, shaded and arrows, respectively) and different tracks of transitioning cyclones from 1756 sensitivity experiments (thin black lines). The bifurcation points associated with the upstream trough 1757 and the developing downstream ridge, deciding upon the development of a “no ET”, a “NW pattern” 1758 or a “NE pattern” scenario, are marked by the dot and the cross, respectively. Dashed contours depict 1759 the streamlines that emanate from the bifurcation points. Figure 11 from Riemer and Jones (2014).

1760

1761 Fig. 7: Three-dimensional schematic depiction of the preconditioning stage with a PRE 1762 during western North Pacific ET Anticyclonic PV air in the upper-level outflow of a TC and 1763 associated PRE as blue shading in the upper panel, jet streak as green shading and 200-hPa waveguide 1764 as red contour separating high PV air (>3 PVU; orange shading) from lower PV air (<3 PVU; 1765 unshaded). Mid-level baroclinic zone as blue tilted surface. Trajectories of rapidly ascending air 1766 parcels as blue-red-blue lines, reflecting the diabatic PV modification of the parcels from low to high 1767 to low PVU, respectively. Mean sea level pressure (gray contours; every 8 hPa) and equivalent 1768 potential temperature (violet contours; 320 and 330 K) are indicated in the lower panel. Figure 11 1769 from Grams and Archambault (2016).

1770

1771

1772 Fig. 8: Conceptual model of the key synoptic-scale features during the occurrence of a PRE. 1773 Shown is the 200-hPa geopotential height (contours), the upper-tropospheric jet streak (gray shading, 1774 the midlatitude baroclinic zone (surface frontal structure with 925-hPa streamlines associated with 1775 warm (red) and cold (blue) air advection. and lower-tropospheric moisture flux along the eastern side 1776 of the recurving TC and the western flank of a subtropical high. The formation of the PRE is indicated 1777 by the yellow-green ellipse (Fig. 20a of Moore et al. (2013)).

1778

1779 Fig. 9: Downstream development during the ET of Typhoon Nabi (2005)). All panels: 5 -2 1780 Vertically integrated Ke (shaded; 10 J m ), 500-hPa heights (light gray contours, 60 m intervals). a) 1781 vertically integrated divergence (dashed) and convergence (solid) of the ageostrophic geopotential flux -1 1782 (contours; W m ); b) vertically integrated total Ke flux vectors (advection + dispersion; reference 1783 vector in lower right, 105 W m-1), and 500-hPa heights; c) baroclinic conversion (contours; W m-1). 1784 Figures 3c, a, b from Harr and Dea (2009).

1785

1786

1787 Fig. 10: Interaction between a transitioning cyclone and the midlatitude flow expressed as 1788 advection of low PV air by the upper-level divergent outflow. a) Idealized representation of ridge 1789 amplification and jet streak intensification. Vectors represent the upper-tropospheric divergent outflow 1790 associated with the transitioning cyclone. Shading denotes anticyclonic PV advection by the divergent 1791 wind (Archambault et al. 2013, their Fig. 4). Composite analyses of objectively defined b) strong and 1792 c) weak interactions at the time of maximum interaction. 500-hPa ascent (green, every 2x10-3 hPa s-1, 1793 negative values only), total-column precipitable water (shaded according to grayscale, mm), 200-hPa 1794 PV (blue, every 1 PVU), irrotational wind (vectors, > 2 m s-1; purple vectors, > 8 m s-1), negative PV 1795 advection by the irrotational wind (dashed red, every 2 PVU day-1 starting at -2 PVU day-1), and total 1796 wind speed (shaded according to color bar, m s-1). The star denotes point of maximum interaction. The 1797 TC symbol denotes composite TC position. Downstream development of d) strong and e) weak 1798 interactions 36-h after time of maximum interaction as represented by 250-hPa meridional wind 1799 anomalies (shaded, m s-1; enclosed by black contours where significant at the 99% confidence level), 1800 PV (blue, every 1 PVU), and irrotational wind (vectors, >2 m s-1). Figures 8a, 8b, 5d and 6d, from 1801 Archambault et al. 2015.

1802

1803 Fig. 11: Recurvature-relative composites of enhanced RWP frequency anomaly (shaded in %) 1804 (a) relative to June to November climatology for western North Pacific transitioning cyclones and (b) 1805 relative to December to April climatology for South Indian Ocean transitioning cyclones. Statistical 1806 significance at 95% confidence level hatched, mean track given by black like. Mean and range of 1807 recurvature longitudes indicated by white star and black bar, respectively. Data smoothed with a 1808 Gaussian filter. Figures 3a and 5a from Quinting and Jones (2016).

1809

1810

1811 Fig. 12: a) Illustration of omega block and high-impact weather downstream of Typhoon Nuri 1812 (2014) after Bosart et al. (2015). b) Downstream impact of Typhoon Choi-Wan (2009), based on PV 1813 surgery experiments where the storm has been removed from initial conditions (Keller and Grams 1814 2014). Black items represent midlatitude flow features in the presence of ET, red items the evolution if 1815 ET influences were not present: 300-hPa geopotential height contour indicates upper-level waveguide 1816 (950-dam at 0000 UTC 22 Sep 2009), Arrows indicate shift of high-impact weather (precipitation, 1817 sunny and hot conditions, cold conditions) with symbol size representing magnitude.

1818

1819

1820 Fig. 13: Vertically integrated initial-time dry total energy singular vector sensitivities (shaded; 1821 values in color bar: J kg−1) with 500-hPa streamlines for TC Shanshan from 0000 UTC 15 Sep 2006. 1822 Figure 3c from Reynolds et al. (2009).

1823

1824

1825 Fig.14: Increase in standard deviation of the 500-hPa geopotential height (in dam) in the 1826 Australian, the Canadian, the ECMWF and the THORPEX Interactive Grand Global Ensemble 1827 (TIGGE, Swinbank et al., 2015) Multimodel EPS for the ET of Hurricane Ike. Forecast initialized 1828 0000 UTC 10 Sep 2008. TC position in ensemble members is marked by the black dots, best track 1829 position at ET time by the red dot. Figure 1 from Keller et al. (2011).

1830

1831

1832 Fig. 15: a) Schematic of the shift and amplitude pattern of ensemble forecast uncertainty 1833 derived from the first two EOFs (thin solid and dashed lines) of potential temperature on the 1834 dynamical tropopause in ensemble members. The thick black line represents the strong potential 1835 temperature gradient on the dynamic tropopause in the midlatitudes. b) Synoptic patterns (shape of 1836 ridge) that result from the contribution to the variability patterns. c) EOF 1 (left, contours) and 2 (right, 1837 contours) for potential temperature at 2 PVU (shaded in K) in an ECMWF ensemble forecast for 1838 Typhoon Maemi (2003). Values indicate percentage of total uncertainty captured by the respective 1839 EOF. Figures 9 and 10a, b from Anwender et al. (2008).

1840

1841

1842 Fig. 16: Normalized ensemble spread of 500-hPa geopotential height as a function of forecast 1843 hour for NOAA’s 2nd generation global reforecasts initialized at recurvature time. Data covers all 1844 western North Pacific tropical cyclones from 1985-2013. Solid line shows mean, dashed line median, 1845 and the shaded region 25th–75th percentile range of the distribution. Statistically signiﬁcant values of 1846 the mean shown as thicker line. Hurricane symbol marks the time of recurvature, circle the median 1847 time of ET, and the thin vertical line the peak spread. Figure 3a from Aiyyer (2015).

1848

1849

1850 Fig. 17: Forecast degradation due to data denied in a) extratropical sensitive regions (SVout) 1851 and b) vicinity of the storm (ETout), expressed as root mean square difference total energy. Box-and- 1852 whiskers plot of the percentage impact over Europe (35–75° N, 10–30° E) for all denial cases. The 25 1853 and the 75 quantile, median, and most extreme outliers are indicated by the box edges, red line and 1854 whiskers, respectively. Vertical dashed lines separate ET cases, vertical dotted lines indicate ET times. 1855 Adapted from Fig. 5 of Anwender et al. (2012).