1. Introduction
Although the quest for improving tropical cyclone (TC) intensity forecasts has yielded some small, but consistent improvements (DeMaria et al. 2014), TC intensity prediction continues to struggle to accurately predict events of rapid intensification (RI; Gall et al. 2013; Kaplan et al. 2015; Rozoff et al. 2015). Part of this challenge exists because the rate of TC intensification depends on the interaction of processes that occur across a spectrum of spatial and temporal scales, ranging from stochastic, convective-scale processes (e.g., Judt and Chen 2016), to synoptic-scale environmental conditions, which may include external forcing for convection (e.g., Bosart et al. 2000; Kaplan et al. 2015; Fischer et al. 2017). Additionally, despite an improved understanding of the environmental conditions favorable for TC intensification, the environments of RI events have been shown to be, in large part, statistically similar to TCs that intensify more gradually (Hendricks et al. 2010). As a result, previous work has hypothesized internal dynamics primarily govern whether a given TC will undergo a period of RI, provided a favorable environment exists (Hendricks et al. 2010). The inherent decreased predictability related to such internal processes exacerbates the TC intensity prediction challenge (Judt and Chen 2016).
The intensity of mature TCs is also modulated by internal dynamics via eyewall replacement cycles (ERCs; Willoughby et al. 1982; Black and Willoughby 1992; Sitkowski et al. 2011; Kossin and DeMaria 2016). During a typical ERC, as the outer eyewall develops, the inner eyewall will contract toward the center of the TC, while the storm experiences a period of intensification (Sitkowski et al. 2011). As the outer eyewall intensifies, the inner eyewall begins to decay, resulting in a temporary period of weakening. This period of weakening can occur despite otherwise favorable environmental conditions. Eventually, the outer eyewall intensifies to a point where its tangential wind exceeds that of the inner eyewall, and the TC experiences a period of reintensification. The end result is a broader TC vortex, which can have important implications on potential TC impacts, such as increasing the area of damaging winds and storm surge (Irish et al. 2008; Lin et al. 2014).
Recent work has allowed for a better understanding of the characteristics of ERC events and, in turn, more accurate TC intensity forecasts. Sitkowski et al. (2011) created a climatology of 24 ERC events in the North Atlantic basin using aircraft reconnaissance observations. Building off this climatology, Kossin and Sitkowski (2012) and Kossin and DeMaria (2016) created statistical TC intensity prediction models containing parameters related to the presence of secondary eyewalls. Kossin and DeMaria (2016) found including such parameters can decrease short-term forecast (i.e., <24 h) errors of operational statistical models by over 50%, with the caveat of a relatively small sample size.
Despite the advances made in understanding and forecasting ERCs, the mechanisms by which a secondary eyewall forms in the first place are less clear, as multiple hypotheses have been put forth. Some of these hypotheses depend on processes internal to the TC. One potential mechanism for secondary eyewall formation (SEF) involves the presence of vortex Rossby waves (VRWs), which propagate along the TC’s negative radial vorticity gradient. Montgomery and Kallenbach (1997) linked the outward propagation, and eventual axisymmeterization, of spiral rainbands to VRWs. These outward propagating waves converged at a stagnation radius, where they acted to induce a secondary tangential wind maximum through wave–mean-flow interactions. Using high-resolution simulations of Hurricanes Katrina (2005) and Rita (2005), Abarca and Corbosiero (2011) analyzed convective evolutions consistent with VRW activity, with SEF occurring near the hypothesized VRW stagnation radius. Alternatively, Judt and Chen (2010) assessed a separate simulation of Rita and proposed the presence of a moat region prevented the outward propagation of VRWs. Because SEF was still simulated, Judt and Chen (2010) hypothesized VRWs are not a necessary ingredient for SEF. The importance of a TC’s radial vorticity profile for SEF was also emphasized by Terwey and Montgomery (2008), who argued SEF occurs via the axisymmeterization and upscale cascade of vorticity from rainband convection within a region associated with a weak radial gradient of vorticity, referred to as a “beta skirt.” Previous observational studies have also noted the presence of a beta skirt prior to SEF (e.g., Dougherty et al. 2018).
Other possible pathways to SEF involve negatively buoyant downdrafts, largely driven by stratiform precipitation in the downwind end of spiral rainbands. The evaporational cooling associated with stratiform precipitation can create a lower-tropospheric frontal zone between the eyewall and rainband region (Fang and Zhang 2012; Rozoff et al. 2012). This frontal zone facilitates the development of deep convection, which may axisymmeterize to form a secondary eyewall. Alternatively, the evaporational cooling of hydrometeors precipitating from the anvil exhaust of the primary eyewall have been hypothesized to lead a different type of top–down pathway to SEF (Tyner et al. 2018). Observational analyses of aircraft reconnaissance radar observations conducted by Didlake and Houze (2013b) and Didlake et al. (2018) identified instances of a mesoscale descending inflow jet, which ultimately terminated near the top of the TC boundary layer along the inner edge of a region of stratiform precipitation. Near the location where the descending inflow jet reached the entropy-rich TC boundary layer, a region of convergence and intense updrafts was observed, which preceded a local enhancement of the tangential wind field, a developing local secondary circulation, and SEF (Didlake et al. 2018).
Additional boundary layer processes have been hypothesized to be an integral part of SEF. Multiple studies have advocated for the importance of unbalanced dynamics in the generation of a secondary eyewall (Huang et al. 2012; Abarca and Montgomery 2013; Abarca et al. 2016; Huang et al. 2018). In this view, SEF begins with a broadening of the TC vortex and the associated tangential wind field. Thereafter, boundary layer inflow and convergence increase near the location of SEF, within a region of agradient forcing. As the inflow strengthens, the agradient force induces supergradient flow and ascent out of the boundary layer, within a thermodynamically favorable location for convection. Some debate exists, however, about the role of supergradient flow in SEF. Using three diagnostic boundary layer models, Kepert (2013) argued that a broad expansion of the tangential wind field does not necessarily cause a local frictional updraft. Additionally, Kepert (2013) concludes supergradient flow is a by-product of frictional convergence and is not critical to SEF. Instead, it is hypothesized SEF is the result of a feedback between an Ekman-like boundary layer response to a perturbation in the radial gradient of the gradient wind vorticity. Ascent out of the top of the boundary layer, near the location of the vorticity perturbation is presumed to cause a generation of additional vorticity through which the feedback process manifests. Alternatively, Montgomery et al. (2014a) argued agradient forcing in the boundary layer is an essential part of SEF and described multiple limitations of the Kepert (2013) study. In particular, Montgomery et al. (2014a) discovered the linearized diagnostic boundary layer model of Kepert (2013) poorly replicated the TC structure of two different full-physics, three-dimensional simulations of SEF. Kepert and Nolan (2014) provided a rebuttal to the concerns raised by Montgomery et al. (2014a), but also emphasized nonlinear boundary layer processes in SEF.
Other studies have attributed SEF to external, environmental forcings. For example, upper-tropospheric troughs have been hypothesized to initiate convection radially outward of the primary eyewall, which may cause SEF (Molinari and Vollaro 1989, 1990; Nong and Emanuel 2003; Leroux et al. 2013; Dai et al. 2017). Similarly, environmental vertical wind shear has been hypothesized to induce asymmetric rainband activity, which may axisymmeterize, resulting in SEF (Zhang et al. 2017; Dougherty et al. 2018).
Motivated by the important operational and research challenges RI, ERCs, and SEF pose, this study examines the evolution of an RI event in Hurricane Irma (2017) that features two rapidly evolving ERC events. The environmental, vortex, and convective-scale evolutions related to the RI event are examined using flight-level and tail Doppler aircraft reconnaissance observations, passive microwave satellite data, and model analyses of the environment. Here, we will show that the two ERC events are associated with different vortex evolutions and pathways to SEF, despite occurring within 24 h from another.
2. Datasets and methods
This study primarily utilized aircraft reconnaissance observations obtained from four Air Force WC-130 and five NOAA WP-3D (P3) missions into Hurricane Irma between 3 and 6 September 2017 (Table 1). In addition to the in situ measurements taken by the aircraft, each of the five P3 flights sampled the TC using both the C-band lower-fuselage (LFR) radar and the X-band Tail Doppler radar (TDR) onboard the aircraft. The LFR has the advantage of greater resiliency to attenuation and observations are stored at a greater temporal resolution than observations from the TDR. Alternatively, the dual-Doppler TDR is able to retrieve the three-dimensional kinematic structure of the TC vortex. Before the TDR observations were analyzed, they underwent an automated quality control and dealiasing procedure, which produces errors comparable to manual reprocessing (Rogers et al. 2012). The three-dimensional wind field was retrieved using a similar variational technique to those described in Gao et al. (1999) and Reasor et al. (2009). For each pass the aircraft made through the center of the storm, the observed wind field is solved onto a Cartesian grid with a horizontal resolution of 2 km and a vertical resolution of 0.5 km. Radar observations are available within a ~75–100-km wide swath, centered along the flight track, with some intermittent gaps due to either a lack of scatterers, quality control, or attenuation. The length of most radial legs flown through the TC were between 150 and 180 km.
Details of the Air Force WC-130 and NOAA P3 flights analyzed within this study. The maximum sustained 10-m wind (Vmax; m s−1) and minimum sea level pressure (p; hPa) are taken from the nearest best track entry. NOAA P3 flights are indicated by the boldface text.
Once the three-dimensional winds were obtained, the data were interpolated onto a TC-centered cylindrical grid, with a radial, azimuthal, and vertical grid spacing of 1 km, 1°, and 0.5 km, respectively, for each pass the aircraft made through the TC center. The TC center was identified following a technique similar to Reasor and Eastin (2012). In essence, an iterative process was employed, where the TC center was determined to be the point where the azimuthally averaged tangential wind was maximized within annuli of varying widths, ranging from 3 to 13 km, at a height of 2 km. Thereafter, the mean TC center latitude and longitude obtained from the multiple annuli examined were used to determine the “new” TC center.
3. Synoptic history of Hurricane Irma
Figure 1 shows the best track traces of both maximum sustained 10-m wind speed and minimum sea level pressure over the lifetime of Hurricane Irma. The easterly wave that spawned Hurricane Irma exited the African coast on 27 August 2017 (Cangialosi et al. 2018). As the disturbance gradually became better organized, tropical cyclogenesis was determined to occur around 0000 UTC 30 August. Six hours following the time of genesis, Irma reached tropical storm intensity and began to experience a period of RI, as the TC moved off to the west-northwest, across the eastern North Atlantic. Here, RI is defined following Kaplan and DeMaria (2003), which is equal to a 24-h increase in the maximum sustained 10-m wind ≥30 kt (1 kt ≈ 0.51 m s−1). During this RI event, Irma intensified by 65 kt in only 42 h—as highlighted by the shaded region in Fig. 1—and reached major hurricane intensity. Following this period of RI, Hurricane Irma encountered a region of cooler sea surface temperatures (SSTs; Fig. 2) and the storm’s intensity remained relatively steady for nearly 3 days. As the TC began to follow a more west-southwestward trajectory, Irma encountered increasing SSTs on 2 September (Fig. 2).
Best track evolution of maximum sustained 10-m winds (kt; blue line) and minimum surface pressure (hPa; orange line) associated with Hurricane Irma (2017). Time periods associated with RI are shaded in gray. The red vertical lines depict the times of two ERC events.
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
Path of Hurricane Irma beginning at 0000 UTC 30 Aug 2017 (black line; black circles denote the position of the TC every 24 h) and SST (°C; shaded). RI periods are represented by the magenta lines.
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
Beginning on 4 September 2017, Irma experienced a second RI period, intensifying to 155 kt by 1800 UTC 5 September (Fig. 1). The initiation of an RI event while a TC is at major hurricane intensity is a relatively rare event for storms in the North Atlantic basin, as Irma became the first TC to begin such an RI event since Hurricane Gilbert (1988), nearly 30 years earlier. Following Irma’s second RI event, the TC tracked off to the west-northwest (Fig. 2), maintaining category 5 hurricane intensity on the Saffir–Simpson Hurricane Wind Scale for approximately 60 h, which is the second longest consecutive period on record in the North Atlantic basin (Cangialosi et al. 2018). The path of Irma brought the storm through the northeastern Caribbean islands, the Turks and Caicaos, and the Bahamas, before making landfall in Cuba. Following landfall in Cuba, Irma began a turn to the northwest and then north, as the storm made additional landfalls in the Florida Keys and the Florida peninsula. In total, Hurricane Irma made seven landfalls, causing over 50 billion USD in damages and resulting in at least 47 direct deaths (Cangialosi et al. 2018).
The remainder of this study will seek to understand the characteristics associated with, and leading up to, Irma’s second RI period, which occurred from 4 to 5 September (Fig. 1).
4. Environmental influences on Irma’s intensification
The evolution of the environmental vertical wind shear, moisture, SST, and ocean heat content leading up to, and during, Irma’s second RI period are provided in Fig. 3. Here, the environmental conditions are obtained from analyses via the Statistical Hurricane Intensity Prediction Scheme (SHIPS; DeMaria and Kaplan 1994). For reference, Irma’s intensity trace is provided in the dashed orange line. As seen in Fig. 3a, Hurricane Irma was located in an environment characterized by relatively favorable deep-layer (850–200 hPa) vertical wind shear magnitudes of less than 5 m s−1 at nearly every analysis time. The shear magnitudes are, generally, well below the climatological mean of all North Atlantic RI episodes of approximately 5 m s−1 (horizontal, red line in Fig. 3a). Early in Irma’s lifetime, including the first RI period, the vertical shear direction was out of the northeast (Fig. 3b). Following the completion of the first RI event on 1 September 2017, the shear direction shifted to a southerly direction, before rotating anticyclonically back toward a northeasterly direction near the start of Irma’s second period of RI on 4 September. Despite some oscillations, the vertical wind shear maintained a northerly component throughout Irma’s second RI event.
SHIPS analyses (blue lines) of the (a) 850–200-hPa vertical wind shear magnitude (m s−1), (b) 850–200-hPa vertical wind shear direction (meteorological direction), (c) 700–500-hPa layer-averaged relative humidity (%), (d) 850–700-hPa layer-averaged relative humidity (%), (e) sea surface temperature (°C), and (f) maximum potential intensity (m s−1). Analyses span from 1200 UTC 30 Aug to 0000 UTC 7 Sep 2017. For a given parameter, the mean value for all RI episodes occurring from 1989 to 2017 are depicted by the horizontal red line. The intensity of Hurricane Irma (m s−1) is shown in the dashed, orange line.
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
Figures 3c and 3d demonstrate that although Irma was originally located in a relatively moist environment compared to the mean of North Atlantic RI episodes, Irma encountered a drier environment near the end of the first RI period on 1 September, as measured by both 700–500-hPa (Fig. 3c) and 850–700-hPa relative humidity (Fig. 3d). The relative humidity in both layers maintained values 5%–10% less than the respective RI climatological means throughout Irma’s second RI period. The unfavorable decrease in lower- and midtropospheric humidity during this time, however, was offset by an increase in SSTs (Fig. 3e) and ocean heat content (Fig. 3f), beginning on 2 September and continuing through 6 September. By the beginning of Irma’s second RI event, SSTs had increased to 28°C and by the end of the RI event were near 29.5°C, which is well above the climatological mean for North Atlantic RI episodes (Fig. 3e). Ocean heat content displayed a particularly sharp increase in the 12–24 h preceding Irma’s second RI event, increasing from <20 kJ cm−2 to nearly 50 kJ cm−2 from 3 September to 4 September.
These results suggest Irma’s intensity evolution was closely linked to changes in the environment. Irma’s first RI period occurred in a low shear and relatively moist environment, despite being over marginal SSTs and low ocean heat content. As Irma traversed over cooler SSTs and into a drier environment, the TC experienced a plateau in intensity for approximately 3 days. It was not until SSTs and ocean heat content increased toward levels more typically associated with North Atlantic RI episodes that Irma underwent a second period of RI, before reaching its peak intensity.
Despite the strong relationship between the evolution of Irma’s intensity and its environment, it is important to understand the response of the convective and TC vortex structures during this time, as these features are what ultimately dictate the rate of TC intensity change.
5. Vortex evolution during RI
Irma’s second RI period was well-sampled by multiple aircraft reconnaissance missions, as seen by the flight paths of the Air Force WC-130 and NOAA P3 missions analyzed in this study (Fig. 4). NOAA P3 missions began sampling the core of Hurricane Irma around 2100 UTC 3 September, with subsequent missions flown every 12 h. Air Force missions were also conducted every 12 h, beginning on 4 September, with center-fix times staggered between the P3 fix times. Consequently, the inner core of Irma was sampled at a frequency of at least once every 6 h, beginning around 0900 UTC 4 September, near the start of the second RI event.
Flight patterns of all NOAA WP-3D (blue) and Air Force WC-130 (orange) aircraft reconnaissance missions analyzed within this study.
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
The evolution of the storm-relative, azimuthally averaged, tangential wind
Time series of azimuthally averaged tangential wind (
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
From Fig. 5, it is apparent that the maximum
The rate at which both ERCs complete is unusually rapid. In a climatological study of ERCs using flight-level aircraft reconnaissance observations, Sitkowski et al. (2011) determined that a typical ERC takes 36 h to complete. In the case of Irma, the secondary
The two ERCs noted in Fig. 5 are consistent with the convective evolution observed by passive microwave satellite imagery, shown in Fig. 6, which depicts storm-centered 85–91-GHz brightness temperatures from multiple sensors between 3 and 5 September. The 85–91-GHz band is particularly sensitive to frozen hydrometeors, with depressed brightness temperatures indicating a greater amount of “ice scattering,” and hence, increased deep-convective activity.
Storm-centered microwave overpasses displaying 85–91-GHz polarization corrected brightness temperatures (PCT; K; shaded) associated with Hurricane Irma. Red dashed radial-rings are spaced in 50-km increments. The overpasses occurred at (a) 1740 UTC 3 Sep, (b) 2136 UTC 3 Sep, (c) 0518 UTC 4 Sep, (d) 1010 UTC 4 Sep, (e) 2242 UTC 4 Sep, (f) 1059 UTC 5 Sep, (g) 1654 UTC 5 Sep, and (h) 2211 UTC 5 Sep 2017. The sensor and observing frequency of each overpass are given in the top-right corner of each panel.
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
Prior to the onset of RI, at 1740 UTC 3 September, Hurricane Irma displayed a ring of 85–91-GHz polarization-corrected brightness temperatures <250 K associated with the eyewall of the storm (Fig. 6a). Over the next 12 h, minimum brightness temperatures decreased and the eyewall contracted (Figs. 6a–c). Additionally, a convective band formed on the south side of the storm, which was downshear of the TC center according to SHIPS analyses (Fig. 3b). The presence of the downshear convective band is consistent with multiple previous studies that have found TC rainband activity occurs preferentially in the downshear part of the TC circulation (e.g., Willoughby et al. 1984; Corbosiero and Molinari 2002; Houze 2010; Hence and Houze 2012; Didlake and Houze 2013a). The azimuthal extent of the convective band on the south side of Irma increased with time, such that the rainband completely enclosed the eyewall by 1010 UTC 4 September (Fig. 6d). This time nearly coincides with the first observance of an outer
By the time of the next available overpass at 2224 UTC 4 September, the size of the eye had increased, with a single eyewall present (Fig. 6e). Such a convective evolution is consistent with previous ERC analyses (Black and Willoughby 1992; Kossin and Sitkowski 2009; Sitkowski et al. 2011; Didlake et al. 2018) and agrees with the larger RMW at flight-level shown in Fig. 5. In addition to the completion of ERC I, new convective bands developed on the northwest and southeast sides of the TC center (Fig. 6e). Because this time was only about 6 h prior to the first appearance of the secondary wind maximum associated with ERC II (Fig. 5), it is possible these convective bands were associated with secondary eyewall formation in ERC II; however, the relatively poor temporal resolution of microwave observations precludes a clear linkage. The following microwave overpass at 1059 UTC 5 September, over 12 h later and near the end of ERC II, revealed a nearly symmetric ring of PCT <200 K, indicative of intense convection (e.g., Mohr and Zipser 1996a,b; Mohr et al. 1996; Cecil and Zipser 1999). By the end of the RI period, minimum PCT continued to decrease to temperatures <180 K within a single, vigorous eyewall (Figs. 6g,h).
To understand how the three-dimensional TC vortex structure evolved during ERC I and ERC II, the following subsections will examine TDR observations from multiple P3 flights.
a. The evolution of ERC I
The development of the outer eyewall, as well as the completion of ERC I, can be viewed in the form of azimuthally averaged, radius–height sections of tangential wind and vertical velocity, as seen in Fig. 7. In this case, each panel displays the azimuthally averaged values for a given center pass. Beginning with the last pass of the 170903H1 flight, Fig. 7a displays the structure of Hurricane Irma prior to secondary eyewall formation, at 2358 UTC 3 September. At this time, a single, well-defined eyewall exists, with a RMW of 26 km at a height of 3 km and sloping outward to approximately 40 km at an altitude of 10 km. Just inside the RMW, azimuthally averaged ascent >0.5 m s−1 exists throughout much of the troposphere, with maximum azimuthally averaged ascent >1.5 m s−1 at a height of 12 km.
Azimuthally averaged tangential wind (
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
By the time of the first center pass of the 170904H1 mission at 0904 UTC 4 September, the TC vortex had intensified throughout the troposphere (Fig. 7b), consistent with the onset of RI in the best track beginning at 0600 UTC 4 September (Fig. 1). The RMW had contracted to a radius of 20 km at a height of 3 km, with the main eyewall ascent also shifting radially inward, remaining inward of the RMW. The first appearance of a secondary wind maximum was also observed around this time, indicated by an outward sloping local maximum in
During the next three passes, the secondary
During the next mission, at 2140 UTC 4 September, a pronounced, double
A view of the azimuthally averaged radial flow (
Azimuthally averaged radial wind (
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
Over the next three passes, the overturning circulation associated with the developing secondary eyewall progressively reached farther down, toward the surface (Figs. 8c–e). At the same time, boundary layer inflow that was originally feeding the primary eyewall became progressively weaker, until the inner eyewall began to collapse, with outflow noted throughout the lowest 2 km near the primary eyewall (Fig. 8e). Following the completion of ERC I, a more vigorous inflow layer became established, with structures consistent with previous observational studies of supergradient flow in the boundary layer (e.g., Montgomery et al. 2014b; Rogers et al. 2015; Abarca et al. 2016), as a region of outflow was located near the RMW, directly above the inflow region, between a height of 1–2.5 km (Fig. 8f). A secondary overturning circulation was observed above the boundary layer corner-flow region, characterized by inflow between approximately 3–7 km and outflow above 7 km, although these structures slanted upward with increasing radius. Consistent with the hypotheses of Smith et al. (2009), the midlevel inflow seen in Irma at this time was related to a broadening of the TC vortex (Fig. 7f), as relatively large angular momentum surfaces are advected inward above the boundary layer.
Despite the rapid completion of ERC I, the evolution of the TC vortex during the ERC displayed many similarities to the canonical ERC described in the literature (e.g., Sitkowski et al. 2011). Namely, an outer eyewall developed at a radial distance of 2–3 times the RMW, while a moat of subsidence formed between the eyewalls. With time, ascent in the inner eyewall weakened, followed by the total collapse of the inner eyewall, which was associated with a brief weakening of the maximum
b. The evolution of ERC II
The evolution of
As in Fig. 7, but for (a) the 2307 UTC 4 Sep pass, (b) the 0020 UTC 5 Sep center pass, (c) the 0149 UTC 5 Sep pass, (d) the 0943 UTC 5 Sep center pass, and (e) the 1110 UTC 5 Sep center pass.
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
The following pass through Irma 8 h later revealed a fairly complex vortex structure (Fig. 9d). Double
The double
The evolution of
As in Fig. 8, but for (a) the 2307 UTC 4 Sep pass, (b) the 0020 UTC 5 Sep center pass, (c) the 0149 UTC 5 Sep pass, (d) the 0943 UTC 5 Sep center pass, and (e) the 1110 UTC 5 Sep center pass.
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
At the time of the next pass through the TC 8 h later, a large increase in
Despite the formation of a secondary eyewall and completion of ERC II, maximum
The unique characteristics of ERC II, including the midtropospheric double
(a) Plan view of 3-km tangential velocity (m s−1) obtained from the 0943 UTC 5 Sep center pass. Gray radial rings are spaced in 25-km increments. (b) As in (a), but for radar reflectivity (dBZ). (c) A northwest–southeast vertical cross section of tangential velocity (m s−1) taken along the black line shown in (a). For perspective, the black triangle and square markers are shown. (d) As in (c), but for radar reflectivity (dBZ).
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
Northwest–southeast vertical cross sections of both VT and reflectivity are shown in Figs. 11c and 11d. Within the cross sections, a double eyewall structure is evident, with the outer eyewall featuring vertically deeper maxima in both VT and reflectivity compared to the inner eyewall. The vertical cross sections indicate the lack of a pronounced moat at this time, as the two eyewalls appear to be in the process of merging together, particularly within the lowest 2 km of the troposphere. This result suggests the subsidence seen between the primary and secondary ascent maxima leading up to ERC II (Figs. 9c and 10c) may not have had sufficient time to generate a well-defined moat before the two eyewalls merged.
6. Potential mechanisms for secondary eyewall formation
a. SEF I: Examination of a mesoscale descending inflow jet
To understand the causes of the different vortex evolutions during ERC I and ERC II, potential physical mechanisms that led to the two SEF events need to be examined. Recent work by Didlake et al. (2018), who also analyzed TDR observations of an SEF event, determined shear-induced convective asymmetries were responsible for SEF. In the case of Irma, however, the vertical wind shear magnitude was ≤5 m s−1 during both ERC periods (Fig. 3a), which is typically not strong enough to produce robust convective asymmetries, especially in major hurricanes (Chen et al. 2006). Nevertheless, passive microwave observations of Irma detected asymmetric convection within the rainband region of the storm, especially in the times leading up to ERC I (Figs. 6b,c). Recall that this rainband activity was the most vigorous in the downshear side of the storm (Fig. 3b), consistent with the expected convective configuration for a TC in a more moderate shear regime (Corbosiero and Molinari 2002; Chen et al. 2006; DeHart et al. 2014). This raises the question of whether SHIPS analyses potentially underestimated the shear magnitude during the times of SEF.
Previous work has demonstrated the magnitude and direction of the vertical wind shear displays sensitivity to the area over which it is computed (Riemer et al. 2010; Reasor and Eastin 2012; Reasor et al. 2013). To test for potential sensitivities between the vertical wind shear and the domain size, a local vertical wind shear was computed using TDR observations following the methods of Reasor and Eastin (2012) and Reasor et al. (2013). Figure 12 depicts the zonal and meridional components of the local shear profiles, relative to the mean flow at a height of 2 km1 for each P3 mission into Irma. In general, the 2–9-km local shear magnitudes were between 5 and 9 m s−1, and thus, appreciably larger than the SHIPS analyses (Fig. 3a), which calculate shear using winds within a larger domain, that extends out to 500 km from the TC center. Additionally, the 2–9-km local shear directions were oriented 0°–90° to the left of the SHIPS shear direction. Active rainbands on the eastern side of Hurricane Irma on 5 September (Figs. 6f–h) are more consistent with the expected shear-relative convective structure obtained from the local shear estimates, which comprised a larger, westerly zonal-shear component, compared to SHIPS shear analyses.
Profiles of the 2–9 km local vertical wind shear (m s−1), relative to the mean flow at a height of 2 km, for flight 170903H1 (red circles), 170904H1 (orange triangles), 170904H2 (green squares), 170905H1 (blue diamonds), and 170905H2 (purple pentagons). Meridional shear values are displayed along the ordinate and zonal shear values displayed along the abscissa. Markers depict the 2–4, 2–6, and 2–8 km shear values. At each height, the profiles were computed using winds within 80 km of the TC center location.
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
Utilizing the 2–9-km local shear direction computed from the TDR, the downshear-left quadrant of the TC was examined to ascertain whether the SEF mechanisms documented by Didlake et al. (2018) were also found in Irma’s two SEF events. Figure 13 shows reflectivity, VT, VR, and divergence averaged within the downshear-left quadrant during the 0904 UTC 4 September center pass, near the time of the first SEF event (SEF I). Unlike the azimuthally averaged values at this time (Figs. 7b and 8b), the downshear-left portion of the inner eyewall displayed signs of weakening, with subsidence observed in the lowest 8 km of the troposphere, despite relatively large reflectivity values (Fig. 13a). Outward of this location, a broad, secondary ascent maximum was found within a 60–100-km radial band, with two smaller-scale ascent maxima embedded therein. Locations of locally greater ascent within the 60–100-km radial band were also associated with local maxima in VT in the midtroposphere. Similar to the findings of Didlake et al. (2018), the secondary ascent and VT maxima were collocated with the inner edge of a mesoscale descending inflow jet (Fig. 13c) and a band of lower-tropospheric convergence (Fig. 13d). Immediately above the leading edge of the inflow jet was a local maximum in outflow, indicative of a strengthening local secondary circulation and SEF, which was also seen in the SEF event examined by Didlake et al. (2018). Here, the formative secondary eyewall appeared to have acted as an inflow barrier to the primary eyewall, with minimal inflow between the primary and secondary eyewalls at heights between 0.5 and 8 km (Fig. 13c). Instead, the primary eyewall was linked to a secondary inflow jet existing above 10 km. The lack of lower–midtropospheric inflow associated with the primary eyewall likely explains the appearance of subsidence throughout this region, and signaled the upcoming demise of the primary eyewall during ERC I.
(a) Azimuthally averaged values of radar reflectivity (dBZ; shaded) and vertical velocity (black contours) for the southeastern quadrant of the storm as sampled during the 0904 UTC 4 Sep pass. Positive vertical velocity values are drawn every 0.5 m s−1 between 0.5 and 1.0 m s−1 and every 1.0 m s−1 for values greater than 1.0 m s−1. Negative values are dashed. (b) As in (a), but for tangential wind (VT; shaded; m s−1). (c) As in (a), but for radial wind (VR; shaded; m s−1). (d) As in (a), but for divergence (shaded; 10−4 s−1).
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
b. SEF II: Consistency with unbalanced boundary layer processes
Quadrant-averaged analyses of the downshear-left quadrant around the time of SEF prior to ERC II (SEF II), shown in Fig. 14, reveal a different TC structure compared to SEF I. During SEF II, the TC inner-core structure was characterized by two primary ascent maxima, spaced only 20–30 km apart. An area of subsidence was observed between the ascent maxima, collocated with the VT maximum (Fig. 14b). A secondary VT maximum was observed at the base of the outer ascent maximum, between heights of 1–3 km, but sloped radially outward above a height of 3 km. Unlike the VR structure seen in SEF I, the VR field associated with SEF II did not feature a prominent descending inflow jet (Fig. 14c). Instead, the lower-tropospheric inflow layer abruptly slowed at the interface of a local maximum in outflow associated with an apparent corner-flow region, similar to the azimuthally averaged analyses in Figs. 10a–c. A convergent boundary was observed at the interface of the inflow layer and corner-flow region, which fueled ascent into the formative outer eyewall (Fig. 14d). The secondary ascent maximum was also linked to a local overturning circulation, with a local maximum in outflow found extending outward from the top of the developing secondary eyewall (Fig. 14c).
As in Fig. 13, but for the 0149 UTC 5 Sep center pass.
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
The VR field during SEF II resembles the unbalanced boundary layer SEF mechanism proposed by Huang et al. (2012) and Abarca et al. (2016), among others, and suggests a different pathway to SEF compared to the first event less than 24 h earlier. Consistent with the unbalanced boundary layer pathway to SEF, a broadening of the VT field occurred (Figs. 9a–c) along with an increase in boundary layer convergence (Fig. 14d) near a local secondary ascent maximum, and a local VT maximum (Fig. 14b).
c. Evidence of vortex Rossby waves
It is possible additional processes contributed to either SEF I or SEF II. For example, previous studies have found SEF occurs via the axisymmeterization of outward-propagating VRWs near a stagnation radius (Montgomery and Kallenbach 1997; Wang 2002b; Qiu et al. 2010; Abarca and Corbosiero 2011). Potential VRW activity in Hurricane Irma was assessed through a combination of TDR and LFR analyses (Fig. 15). During the period near SEF I, Hurricane Irma was characterized by an elliptical eyewall, with an asymmetric maximum in both vorticity and reflectivity found within the southern eyewall and a spiral band extending anticyclonically outward from the southern eyewall at radii between approximately 20–30 km (Figs. 15a–c). Radially outward of these features, another spiral band was observed, beginning on the SE side of the vortex and trailing anticyclonically around the center of the storm toward the northeastern part of the vortex (Fig. 15c). A local minimum in reflectivity was observed immediately north of the inner eyewall, consistent with an azimuthal wavenumber-1 structure. A band of enhanced relative vorticity was again coincident with this outer spiral band of elevated reflectivity (Figs. 15a–c).
(a) Storm-centered, TDR-derived relative vorticity (ζ; 10−4 s−1) at a height of 2.5 km for 0904 UTC 4 Sep pass. (b) As in (a), but for the deviation of relative vorticity from the azimuthal mean (ζeddy; 10−4 s−1). (c) Storm-centered, composite LFR reflectivity (shaded; dBZ) using sweeps from 0902 to 0906 UTC 4 Sep. TDR-derived ζeddy is overlaid in the black contours at intervals of 2, 5, 10, 15, and 20 × 10−4 s−1. The dashed magenta circle represents the radius used in the diagram shown in Fig. 16a. (d)–(f) As in (a)–(c), but for the 2307 UTC 4 Sep pass. For (f), LFR reflectivity is composited using sweeps from 2305 to 2309 UTC 5 Sep. The dashed magenta circle represents the radius used in the diagram shown in Fig. 16b.
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
The period near, but prior to, SEF II also displayed overlapping banded structures of both reflectivity and vorticity (Figs. 15d–f). Here, a more prominent spiral band of reflectivity and vorticity was observed beginning on the eastern side of the eyewall and wrapped around the southern side of the TC. Closer inspection of the LFR at this time shows two smaller-scale spiral bands of reflectivity extending outward from the southern side of the eyewall, terminating at radii of 40–50 km from the TC center (Fig. 15f).
It is important to clarify that the TDR analyses (Figs. 15a,b,d,e) are constructed over a period of 30–60 min, while the plane traverses the TC vortex, whereas the LFR composites shown in Figs. 15c and 15f) use an averaging period of 4 min. Consequently, some small radial and azimuthal displacements in the locations of the banded vorticity and reflectivity structures may exist due to advective processes that occur while the TDR analyses are being constructed. Despite this caveat, the TDR–LFR synthesis exhibits general agreement in the locations of the banded structures and the TC eyewall. The spiral, banded structures of reflectivity and vorticity seen in Fig. 15 are consistent with prior work on VRWs (Montgomery and Kallenbach 1997; Reasor et al. 2000; Corbosiero et al. 2006; Moon and Nolan 2015; Guimond et al. 2019, manuscript submitted to J. Atmos. Sci.).
Figure 16 shows the evolution of the banded reflectivity structures immediately outward of the eyewall, as observed by the LFR. The brief time periods shown in Fig. 16 were selected to ensure the plane was near enough to the TC center to avoid appreciable attenuation of the radar beam. Figures 16a and 16b show time–azimuth sections of composite reflectivity taken along the radii encircled by the dashed magenta lines in Figs. 15c and 15f, respectively. If the aforementioned spiral bands of enhanced reflectivity and vorticity are indeed VRWs, they should exhibit characteristics of retrogression relative to the background vortex flow (Montgomery and Kallenbach 1997; Corbosiero et al. 2006).
(a) Time–azimuth diagram of LFR composite reflectivity (dBZ) observed during the 0904 UTC 4 Sep center pass, near the time of SEF I. Values are shown for a radius of 25 km from the TC center, corresponding to the dashed magenta line seen in Fig. 15c. (b) Time–azimuth diagram of LFR composite reflectivity (dBZ) observed during the 2307 UTC 4 Sep center pass, near the time of SEF II. Values are shown for a radius of 40 km from the TC center, corresponding to the dashed magenta line seen in Fig. 15f. In both panels, radar sweeps of reflectivity were averaged over a two-minute period centered along the time displayed on the ordinate of the diagram.
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
Near the time of SEF I, a band of enhanced reflectivity at a radius of 25 km, just outside the eyewall, was observed to propagate cyclonically downwind (Fig. 16a). The band was first observed on the southwestern side of the vortex, with peak reflectivities traveling approximately 65° cyclonically downwind over the 18-min observing window. This corresponds to an azimuthal phase speed of approximately 26 m s−1. At this time,
Similar azimuthal propagation characteristics were observed near the time of SEF II (Fig. 16b). Here, multiple maxima in composite reflectivity were observed to propagate cyclonically around the southern side of the TC, consistent with the spiral bands extending from the eyewall of the TC as seen in Fig. 15f. Although some small differences in the azimuthal phase speed of these reflectivity maxima exist, even the most rapidly propagating maximum, beginning south-southwest of the TC center, orbits cyclonically nearly 60° in the 18-min interval. This corresponds to an azimuthal phase speed of approximately 39 m s−1, which is less than the background
The characteristics of the spiral, inner-core rainbands examined herein are consistent with theoretical work (Montgomery and Kallenbach 1997), numerical modeling studies (Wang 2002a,b; Qiu et al. 2010; Moon and Nolan 2015), and previous observational studies (Corbosiero et al. 2006) of VRW activity. Thus, we hypothesize the inner-core spiral bands observed near the times of SEF I and SEF II were associated with VRWs. In both SEF events examined here, it appears plausible the eventual stagnation and axisymmeterization of the ongoing VRW activity may have contributed to SEF. Although not shown here, estimates of the theoretical VRW stagnation radius, following methods similar to Reasor et al. (2000), suggest the stagnation radius resided close to the radial locations of both SEF I and SEF II.
7. Discussion
a. Different pathways to SEF
Despite occurring less than 24 h apart, the two SEF events in Hurricane Irma demonstrate that multiple mechanisms can contribute to the formation of a secondary eyewall. The characteristics associated with SEF I displayed many similarities to the SEF event in Hurricane Earl (2010) examined by Didlake et al. (2018), as a local maximum in both ascent and tangential wind developed in a region of enhanced lower-tropospheric convergence near the leading edge of the descending inflow jet, maximized in the downshear-left quadrant of the storm (Fig. 13). During SEF II, however, the same features were not observed in the downshear-left quadrant of Hurricane Irma. This result suggests a mesoscale descending inflow jet is a plausible, but not necessary ingredient for SEF.
In both SEF events, a broadening of the
Unfortunately, the lack of detailed observations of Hurricane Irma’s three-dimensional mass field prevents the direct computation of both the gradient wind as well as a potential agradient force. Furthermore, the vertical resolution of the TDR observations, particularly near the ocean surface, does not allow for a detailed analysis of the inflow structure within the boundary layer. Therefore, some uncertainty exists in regard to the role of unbalanced boundary layer processes—or lack thereof—in the two SEF events examined in this study. Despite these caveats, the TC vortex structure observed during SEF II is consistent with the vortex evolution described by Huang et al. (2012), Abarca et al. (2016), Huang et al. (2018), and others. Considering the observational limitations of this dataset, we cannot discount the potential for unbalanced boundary layer processes to have also contributed to SEF I.
Considering the aforementioned mechanisms and caveats, we hypothesize SEF I was initiated via a mesoscale descending inflow jet associated with a rainband region in the downshear-left quadrant of the storm. Similar to the hypotheses of Didlake and Houze (2013b) and Didlake et al. (2018), it is thought that boundary layer convergence near the leading edge of the descending inflow jet fueled a sustained updraft which generated a local VT maximum. The eventual axisymmeterization of this local VT maximum is hypothesized to have, at least in part, contributed to SEF I. SEF I may have been aided by the axisymmeterization of outward-propagating VRWs emanating away from the elliptical eyewall. Alternatively, we hypothesize SEF II was initiated through a combination of axisymmeterizing VRWs and enhanced lower-tropospheric convergence related to unbalanced boundary layer processes, both of which were found within a 40–50-km radial band and coincided with a secondary ascent maximum (Figs. 14 and 15d–f).
The hypothesized pathways to SEF likely depend on the convective nature of the TC. Specifically, the mesoscale descending inflow jet observed in SEF I is driven by evaporational cooling from stratiform precipitation (Didlake et al. 2018), while the presumed supergradient flow near the top of the boundary layer, the deeper lower-tropospheric inflow, and the more intense TC vortex observed around the time of SEF II are consistent with more intense convection. The relationship between the convective nature of the TC and SEF was evaluated in the form of contoured frequency by altitude diagrams (CFADs) of vertical velocity (VW), as sampled by the TDR (Fig. 17). Two prominent differences in the distributions of VW for SEF I (Fig. 17a) and SEF II (Fig. 17b) exist. The first difference is related to a shift in the mode of VW, at all heights, at the time of SEF II to slightly lower values compared to the distribution near the time of SEF I (Fig. 17c). The second difference is found at heights above 8–10 km, as SEF II is associated with a greater frequency of vertical velocities >2 m s−1 (Fig. 17c). The more robust updrafts in the upper troposphere observed near the time of SEF II indicate more vigorous, deep convective activity, when compared to SEF I. This result supports the hypothesis that the pathway to SEF displays some dependence to the convective nature of the TC. Although not shown here, similar CFAD analyses were performed using TDR observations from center passes just prior to the times of SEF I and SEF II, and yielded similar results. Nonetheless, some uncertainty remains since these CFADs were constructed using swaths of observations taken at a few different snapshots in time.
(a) Contoured frequency by altitude (CFAD) diagram of vertical velocity (VW; %) using values within 150 km of the TC center for the 0904 UTC 4 Sep center pass, near the time of SEF I. Values are binned in increments of 0.5 m s−1. (b) As in (a), but using the 0149 UTC 5 Sep center pass, near the time of SEF II. (c) The CFAD difference between (b) and (a). Here, positive values indicate the time near SEF II had a greater frequency of a given VW value than the time near SEF I.
Citation: Monthly Weather Review 148, 3; 10.1175/MWR-D-19-0185.1
b. A comparison of the TC vortex evolution during ERC I and ERC II
Following SEF, the two ERC events were also associated with differing vortex evolutions. ERC I displayed a more typical evolution, characterized by the decay of the inner eyewall while the outer eyewall intensified, which induced a temporary decrease in the azimuthally averaged lower-tropospheric winds (Fig. 5); however, the TC intensity continued to increase during this period according to best track estimates (Fig. 1). This discrepancy may be partially attributed to a temporal smoothing of the observations utilized within the best track, although the minimum central pressure measured by aircraft continued to fall between each center pass during this period.
ERC II displayed an atypical vortex evolution. Although the inner eyewall began to decay during ERC II, the two eyewalls appeared to merge together before the inner eyewall could dissipate (Figs. 9 and 11). The merging of the eyewalls may stem from the smaller displacement between the primary and secondary eyewalls following SEF II, as well as the lack of a well-defined moat between the eyewalls (Zhu and Zhu 2015; Tyner et al. 2018). The continued intensification of the TC vortex during ERC II renders the event as a stark contrast to the ERC paradigm discussed within the literature (Sitkowski et al. 2011).
c. Multiscale nature of TC intensification
The RI of Hurricane Irma is a prime example of the complexities associated with the multiscale nature of TC intensification, as evidenced by the interaction of the environmental, vortex, and convective characteristics of the storm. For example, according to SHIPS analyses, Irma was located in an environment of favorable vertical wind shear magnitudes over a period of several days leading up the RI event (Fig. 3a); however, greater intensification rates were not realized until the TC moved over increasing SSTs and ocean heat content. As Irma encountered warmer waters, convective rainband activity also increased (Figs. 6a–c), and ERC I ensued. Contrary to best track estimates, airborne radar observations indicate the
A potential caveat to the aforementioned relationship between RI and the environment is found in the uncertainty of the estimates. In particular, airborne radar estimates of the local shear were, at times, around 5 m s−1 larger than operational model analyses (Fig. 12). Although a temporary decrease in the local shear was observed around the time of RI onset, estimates remained in the moderate (e.g., 5–10 m s−1) range throughout the remainder of the RI period. Considering previous work has shown TC intensification displays decreased predictability within such a shear regime (e.g., Zhang and Tao 2013), the results shown here emphasize the importance of accurate environmental observations when predicting TC intensity change.
8. Conclusions
The initiation of an RI event for a TC at major hurricane intensity is a rare event in the North Atlantic basin. This study examined the environmental and vortex-scale processes related to such an RI event through an observational analysis of Hurricane Irma. The evolution of the TC vortex during RI was examined using both in situ and TDR aircraft reconnaissance observations as well as microwave satellite data. The main findings of this study are highlighted below:
Two ERC episodes were observed during the RI of Hurricane Irma. The presence of multiple ERCs during a sustained RI event opposes the conceptual model within the literature that links ERCs to a period of weakening or, at the most, steady-state intensity (e.g., Sitkowski et al. 2011; Dougherty et al. 2018).
The two ERC events were hypothesized to feature different pathways to SEF. The first SEF event was linked to lower-tropospheric convergence along the leading edge of a descending inflow jet in the downshear-left quadrant of the TC. The second SEF event appeared to have originated from lower-tropospheric convergence associated with a presumed supergradient flow. In both events, it is plausible SEF was aided by the axisymmeterization of outward-propagating VRWs near the theoretical stagnation radius.
ERC I exhibited a relatively common vortex evolution, as the inner eyewall decayed at the expense of an intensifying outer eyewall. Although not reflected in the best track analysis, aircraft reconnaissance observations indicated a brief weakening of the maximum tangential wind throughout the lower troposphere. Alternatively, during ERC II, the two eyewalls appeared to merge together prior to the collapse of the inner eyewall, with no signs of the vortex weakening.
In each ERC, the time elapsed between the formation of a secondary wind maximum to the completion of the ERC was <12 h. This time period is significantly less than the mean ERC duration of 36 h (Sitkowski et al. 2011). Additionally, both ERCs featured a relatively small initial displacement between the primary and secondary eyewall (20–40 km), which is substantially less than the climatological mean of 71 km.
The possibility of an ERC to not induce a period of weakening in the TC has important implications for human impacts. In both ERCs analyzed here, a temporary expansion of the RMW was observed following the completion of the ERC event. A larger RMW would not only bring the most damaging winds over a larger area, but has also been shown to cause greater storm surge inundation (Irish et al. 2008; Lin et al. 2014). The TC intensity evolution observed during ERC II—and to some extent, ERC I—demonstrates a forecaster should not necessarily assume a period of steady-state intensity or weakening will commence following the appearance of a secondary eyewall.
A naturally arising question can be posed: How did Irma continue to experience RI during ERC II? We hypothesize the persistent ascent radially inward of the of primary eyewall was associated with the inward advection of relatively large angular momentum surfaces, along the lines of Montgomery and Smith (2014), Smith and Montgomery (2015), Smith and Montgomery (2016), and others. The persistence of this ascent was maintained by both a region of weak inflow seen above the corner-flow region near the top of the boundary layer, as well as the rapid pace of the ERC, which did not allow a significant amount of time for the inner eyewall to weaken prior to the eventual eyewall merger. Although a more quantitative analysis is ideal, the present study was limited by the intermittent sampling and the spatial resolution of the observational datasets. Future work will utilize representative high-resolution modeling simulations to gain a better understanding of the relationship between TC intensity change and the rapidly evolving vortex structure during the ERCs analyzed here.
The results of this study also raise other questions. For example, if ERC II was initiated by the convergence between the TC inflow layer and a corner-flow region, why do not all TCs that exhibit such supergradient flow experience SEF? Numerical simulations of Typhoon Sinlaku (2008) analyzed by Huang et al. (2012) suggest SEF occurs once boundary layer convergence associated with supergradient flow becomes increasingly persistent and exceeds a threshold value; however, it is challenging to test such a hypothesis using observations. In the case of Irma, ERC II was only detected due to the relatively high frequency of aircraft reconnaissance missions flown into the TC, as microwave satellite observations failed to detect the ERC. The results provided here not only underscore the utility of aircraft reconnaissance observations in studying TC intensity and structural changes, but also demonstrate the need for a more detailed observational network in the TC boundary layer. It is possible other TCs have exhibited vortex evolutions similar to ERC II, but were missed by satellite-based observing systems. The environmental conditions and predecessor vortex structures favoring a vortex evolution similar to ERC II need to be explored in more detail considering the TC intensity and inner-core size changes associated with such ERCs, which can compound societal impacts if the ERC precedes landfall.
Acknowledgments
The constructive and enlightening comments of Drs. Michael Montgomery, Frank Marks, Jun Zhang, and two anonymous reviewers improved the quality and clarity of this analysis. The authors thank John Kaplan for helpful discussions about environmental influences on Irma’s intensity evolution and Dr. Christopher Rozoff for providing the microwave brightness temperatures. Conversations with Drs. Anthony Didlake and Xiaomin Chen also provided valuable insight into this analysis. The first author was supported by the National Research Council Research Associateship Programs.
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The local shear was computed using winds within an 80-km radius relative to a TC center that was recalculated at every vertical level, following the same process described in section 2. Even so, the displacement between the TC center at the bottom and top of the 2–9-km shear layer was relatively small (<5 km) for each flight shown in Fig. 12.
