Understanding the Variation and Mechanisms of Tropical Cyclone Genesis Potential over the Western North Pacific during the Past 20 000 Years

Dubin Huan aNansen-Zhu International Research Centre, Institute of Atmospheric Physics, Chinese Academy of Sciences, Beijing, China
dUniversity of Chinese Academy of Sciences, Beijing, China

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Qing Yan aNansen-Zhu International Research Centre, Institute of Atmospheric Physics, Chinese Academy of Sciences, Beijing, China
bKey Laboratory of Meteorological Disaster, Collaborative Innovation Center on Forecast and Evaluation of Meteorological Disasters, Nanjing University of Information Science and Technology, Nanjing, China

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Ting Wei cState Key Laboratory of Severe Weather, Chinese Academy of Meteorological Sciences, Beijing, China

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Nanxuan Jiang aNansen-Zhu International Research Centre, Institute of Atmospheric Physics, Chinese Academy of Sciences, Beijing, China
dUniversity of Chinese Academy of Sciences, Beijing, China

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Abstract

To deepen our understanding of the behavior of tropical cyclones (TCs) over the western North Pacific (WNP) and its response to various external forcings, we investigate the variation of TC genesis potential over the WNP during the past 20 ka that experienced several large forcings. Using a set of transient simulations and a genesis potential index, our results indicate that TC genesis potential in the storm season shows an overall increase averaged over the WNP during the Heinrich Stadial 1 (HS1; ∼17–16 ka) relative to the Last Glacial Maximum. Subsequently, there is a sharp decrease in genesis potential during the Bølling–Allerød (BA; ∼13.5 ka) relative to the HS1, which is followed by an increase during the Younger Dryas (YD; ∼12.9–11.7 ka). During the Holocene, TC genesis potential shows a slight decrease during the early–middle Holocene (∼11.7–6 ka) and a significant increasing trend afterward. Further analysis shows that the contribution of each genesis factor to the genesis potential change varies greatly with time, which is in turn tied to changes in external forcings. The increased meltwater fluxes dominate the increased genesis potential during the HS1 and YD, whereas all external forcings (i.e., changes in meltwater, ice sheets, CO2, and insolation) contribute to genesis potential change during the BA. The long-term increasing trend from the mid-Holocene to the present is controlled by orbital insolation. These results may help improve our understanding on past TC activity and shed lights on TC response to various external forcings in a long-term future.

© 2023 American Meteorological Society. For information regarding reuse of this content and general copyright information, consult the AMS Copyright Policy (www.ametsoc.org/PUBSReuseLicenses).

Corresponding author: Qing Yan, yanqing@mail.iap.ac.cn

Abstract

To deepen our understanding of the behavior of tropical cyclones (TCs) over the western North Pacific (WNP) and its response to various external forcings, we investigate the variation of TC genesis potential over the WNP during the past 20 ka that experienced several large forcings. Using a set of transient simulations and a genesis potential index, our results indicate that TC genesis potential in the storm season shows an overall increase averaged over the WNP during the Heinrich Stadial 1 (HS1; ∼17–16 ka) relative to the Last Glacial Maximum. Subsequently, there is a sharp decrease in genesis potential during the Bølling–Allerød (BA; ∼13.5 ka) relative to the HS1, which is followed by an increase during the Younger Dryas (YD; ∼12.9–11.7 ka). During the Holocene, TC genesis potential shows a slight decrease during the early–middle Holocene (∼11.7–6 ka) and a significant increasing trend afterward. Further analysis shows that the contribution of each genesis factor to the genesis potential change varies greatly with time, which is in turn tied to changes in external forcings. The increased meltwater fluxes dominate the increased genesis potential during the HS1 and YD, whereas all external forcings (i.e., changes in meltwater, ice sheets, CO2, and insolation) contribute to genesis potential change during the BA. The long-term increasing trend from the mid-Holocene to the present is controlled by orbital insolation. These results may help improve our understanding on past TC activity and shed lights on TC response to various external forcings in a long-term future.

© 2023 American Meteorological Society. For information regarding reuse of this content and general copyright information, consult the AMS Copyright Policy (www.ametsoc.org/PUBSReuseLicenses).

Corresponding author: Qing Yan, yanqing@mail.iap.ac.cn

1. Introduction

The western North Pacific (WNP) is the most active ocean basin for tropical cyclone (TC) genesis on the planet, where about 26 TCs generate per year accounting for almost one-third of the global total (Chan 2005; Mei and Li 2022; Schreck et al. 2014). TCs threaten the economics, environment, and society of coastal areas and countries via inducing strong winds, heavy rain, and storm surges (e.g., Mei and Xie 2016; Tu et al. 2020). Moreover, TCs over the WNP play an important role in regulating precipitation, hot days, and extreme events in coastal areas of East Asia (Liu et al. 2021; Wang et al. 2020; Q. Zhang et al. 2018; Zhong et al. 2019). Given the important influences of TCs, extensive efforts have been made to understand their behavior and the associated dynamic mechanisms over the WNP. Observational-based studies help advance our knowledge on TC activity over the WNP under current warming. They, for example, suggested that TC activity over the WNP has become quieter since the late 1990s (Chang et al. 2021; W. Zhang et al. 2018; Zhao et al. 2018), manifested as an abrupt decrease in TC frequency, especially over the southeastern WNP (Chang et al. 2021; Zhao et al. 2018).

On a longer time scale, the intense warming arising from increased greenhouse gases could induce accelerated melting of polar ice sheets (Muntjewerf et al. 2020; Shannon et al. 2019; Vizcaino et al. 2015; Yan et al. 2014), which in turn causes enhanced meltwater fluxes and may significantly regulate global climate (Bronselaer et al. 2018; Golledge et al. 2019; Schloesser et al. 2019). This highlights the important roles of ice sheets and the associated meltwater fluxes in the long-term future climate. Moreover, changes in orbital parameters can alter the incoming solar radiation at the top of the atmosphere, driving climate change at multimillennial time scales (Berger and Loutre 2002; Yi et al. 2018), as well as TC activity (Yan and Zhang 2017; Yan et al. 2019a). These external forcings are expected to vary considerably in the future on the geological time scale, but their roles have not been fully understood in current TC projections (e.g., Knutson et al. 2020). To better decipher the response of TCs over the WNP to various external forcings, we might look back into Earth’s history to explore how they might vary in the past.

Orbital insolation, greenhouse gases, meltwater fluxes, and continental ice sheets exhibited considerable changes during the past 20 ka. This period hence provides an opportunity to investigate how TCs over the WNP may respond to various external forcings and place current TC activity in a long-term paleoclimatic context. Geological evidence can help us understand the behavior of TCs in past climates, but the available reconstructions for TCs over the WNP were traced back only to the past several millennia (Chen et al. 2012; Wang et al. 2014, 2022; Woodruff et al. 2009; Yang et al. 2022; Zhou et al. 2019), which hampers our efforts to decipher the variation of TCs during the Last Glacial Maximum (LGM; ∼20 ka) and the last deglaciation (∼18–11 ka). Moreover, the sparsely distributed proxies might only document strong TCs, indicating that the variation of weak TCs may be neglected in reconstructions (Yan et al. 2015). Additionally, geological evidence is generally incapable of depicting large-scale environmental factors and atmosphere–ocean circulations, which is important for understanding the TC behavior.

Numerical simulation has been proven to be an effective tool to examine the response of TCs to various external forcings in the past, especially large-scale environmental factors that spawn storm genesis. Regarding the past 20 ka, existing modeling studies targeting TC activity over the WNP mainly focused on three periods: the LGM, mid-Holocene, and last millennium. During the LGM, Korty et al. (2012a) pointed out that despite the colder world, environmental conditions were more favorable for TC genesis over the WNP relative to the preindustrial era based on climate models from the Paleoclimate Modeling Intercomparison Project phase 2 (PMIP2), but TC genesis potential varied across ocean basins. Yoo et al. (2016) suggested that the number and mean intensity of TCs over the WNP during the LGM were not significantly different from those in modern times based on dynamical downscaling simulations. Furthermore, Lawton et al. (2021) pointed out that there was a decline in the frequency of TCs reaching category 4 or higher during the LGM, using a downscaling technique applied to three PMIP3 models. During the mid-Holocene (∼6 ka), in response to the higher summer orbital insolation, TC genesis potential generally decreased over the WNP relative to the preindustrial era, based on the PMIP2 simulations (Korty et al. 2012b). This was further confirmed by the results from PMIP3 models (Koh and Brierley 2015). Using the simulations from the EC-Earth model and a downscaling technique, Pausata et al. (2017) suggested that the greening of the Sahara and reduced dust loadings were favorable for TC genesis worldwide, but unfavorable over the majority of the WNP. With respect to the last millennium, Korty et al. (2012b) found that potential intensity exhibited no long-term trend over the WNP during the past millennium, based on Community Earth System Model (CESM) results, but it immediately reduced during the storm season following volcanic eruptions, confirmed by Yan et al. (2015). Similarly, Yan et al. (2018) used the CESM Last Millennium Ensemble outputs and indicated that tropical and Northern Hemisphere volcanic eruptions generally resulted in unfavorable conditions for TC genesis over the WNP, whereas the impacts of Southern Hemisphere eruptions were not significant. Besides, Yan et al. (2017) suggested that conditions were broadly conducive to TC genesis over the WNP during the Medieval Climate Anomaly (950–1200 AD) relative to the Little Ice Age (1600–1850 AD) based on PMIP3 simulations.

These modeling studies advance our knowledge of the variation of TC activity over the WNP during the LGM, mid-Holocene, and last millennium. However, transient evolution of the TCs over the WNP during the past 20 ka still remains unclear, especially over the last deglaciation when climatic fluctuations of large amplitude occurred. Moreover, the response of TCs over the WNP to several external forcings (e.g., meltwater fluxes and continental ice sheets) and the associated dynamic mechanisms have not been systematically investigated. Here, based on the transient simulations of climate evolution over the past 21 ka (TraCE-21ka), we investigate the potential evolution of TC genesis over the WNP during the past 20 ka via a genesis potential index. Furthermore, we identify the dominant forcing for TC changes and explore the underlying dynamic mechanisms. Such an investigation helps improve our knowledge of TC behavior over the WNP since the LGM and sheds light on the response of TCs to various external forcings in a long-term future.

2. Methodology

a. The TraCE-21ka simulations

TraCE-21ka simulations are based on the Community Climate System Model version 3 (CCSM3), a coupled atmosphere–ocean general circulation model. The TraCE-21ka includes a full-forcing simulation (TraCE-full), which is driven by realistic external forcings, including orbital insolation (ORB), greenhouse gases (CO2), meltwater fluxes (MWF), and continental ice sheets (ICE) (He 2011; Liu et al. 2009). Meantime, change in land–sea mask induced by the variations of ice sheets (e.g., the exposure of Sunda shelf during the past 20 ka) is considered in TraCE-21ka. Additionally, there are four single-forcing sensitivity experiments driven only by the change in the ORB (TraCE-ORB), CO2 (TraCE-CO2), MWF (TraCE-MWF), and ICE (TraCE-ICE). The simulations are all performed with T31_gx3v5 resolution. The horizontal resolution of the atmospheric and land model is about 3.75° × 3.75°. The ocean model is conducted at a longitudinal resolution of 3.6° and variable latitudinal resolution, with finer resolution near the equator (∼0.9°). The TraCE-21ka simulations have been proven to be capable of reproducing the major climatic characteristics during the past 21 ka, such as El Niño (Liu et al. 2014), East Asian monsoon (Wen et al. 2016; Wu et al. 2021), and midlatitude westerlies (Jiang et al. 2020). More details about the TraCE-21ka simulations are given in He (2011). It should be noted that we use summer and autumn seasonal mean in sensitivity experiments due to the lack of monthly outputs, and surface temperature is taken as the sea surface temperature (SST) because they are approximately equal over the ocean.

b. Large-scale TC genesis factors

The main factors for TC genesis include potential intensity, moisture in the middle troposphere, vertical wind shear, and low-level vorticity. Potential intensity is the theoretical maximum wind velocity of a TC, which depends on the atmospheric thermal structure. It is defined as (Bister and Emanuel 1998, 2002)
PI=CkCdSSTToTo(k0*k),
where SST is the sea surface temperature, To is the mean outflow temperature, k0 is the air saturation enthalpy at the sea surface, k is the enthalpy of a boundary layer air parcel, Ck is the surface enthalpy exchange coefficient, and Cd is the drag coefficient; (SSTTo)/To and k0k represents the effect of thermodynamic efficiency and thermodynamic disequilibrium, respectively. The latter is nonzero because the marine boundary is usually unsaturated (Korty et al. 2012b).
Sufficient middle-tropospheric moisture content favors the formation of TCs in that downdrafts will not overwhelm the moist inflow, which fuels the convection of a developing TC (Korty et al. 2012a). Here, moist entropy deficit χ is employed to measure the moisture content of the middle troposphere (Emanuel et al. 2008):
χ=s*sms0*s*,
where s0* and s* refer to the saturation moist entropies of the sea surface and free troposphere (600 hPa), respectively, and sm represents the moist entropy of the middle troposphere (600 hPa). The variation of χ is affected by relative humidity in the midtroposphere and vertical temperature contrast between sea surface and midtroposphere that contributes to the strength of surface heat fluxes (Yan and Zhang 2017; Yan et al. 2015, 2019b). A larger moist entropy deficit means less saturated midtroposphere and hence unfavorable for TC genesis.

Vertical wind shear affects the TC genesis by making the convection asymmetric and delivering comparatively dry air into it (Korty et al. 2012a). It is commonly defined as the magnitude of horizontal wind velocity difference between 200 and 850 hPa. Larger wind shear is generally unfavorable for TC genesis. Additionally, low-level (850 hPa) absolute vorticity is indispensable for the low-level convergence and aggregation process (Nolan et al. 2007).

Next, we employ a genesis potential index (GPI) to measure the TC genesis potential by combining the factors mentioned above. The GPI is defined as (Emanuel 2010; Korty et al. 2012a)
GPI=a[min(|η|,4×105)]3[max(PI35,0)]2χ4/3(25+VS)4,
where a is a normalizing coefficient, η is the absolute vorticity at 850 hPa, and VS is vertical wind shear between 200 and 850 hPa. GPI has been proven useful for TC research and widely used in past, present, and future climates (Camargo 2013; Cao et al. 2020; Korty et al. 2012a,b; Li et al. 2019; Yan and Zhang 2017; Yan et al. 2021). Preliminary evaluations indicate that compared with the observations from the International Best Track Archive for Climate Stewardship version 4 (Knapp et al. 2010), the GPI based on the TraCE-21ka results during 1951–90 reproduces the general pattern of modern TC genesis well (Fig. S1 in the online supplemental material).

c. Role of individual genesis factors

To identify the relative importance of each genesis factor for the variation of GPI from a key period to another, we recalculate the GPI, with interannual variations of three factors during a certain key period and the remaining factor varying interannually during the following period, and then the contribution of each factor can be quantified by the difference between the two GPIs (Camargo et al. 2007; Xu and Huang 2015). During the Holocene, we recalculate the GPI, but with three factors fixed at the climatology and the remaining one varying interannually, and then the impacts of each genesis factor on the trend of GPI can be compared (Camargo et al. 2007; Xu and Huang 2015).

d. Genesis-weighted mean

Observations and TraCE-21ka simulations show that the distribution of TC genesis is not uniform over the WNP, with more TC occurrences over the western side (Fig. S1). When analyzing the contribution of each genesis factor to the variation of GPI, we calculate the spatial average weighted by the modern spatial pattern of genesis potential in the TraCE-21ka simulations (i.e., during 1951–90) to emphasize the most prevalent regions of TC formation, following the methods of Held and Zhao (2011). The genesis-weighted mean is defined as
A=G(x,y)A(x,y)¯G(x,y)¯,
where A represents the studied variable, G refers to the distribution of modern TC genesis potential, and the overlines are the regional mean over the WNP (5°–35°N, 120°E–180°).

e. Wave activity flux

To understand the remote impact of the change in ice sheets on large-scale environmental factors over the WNP, we utilize the wave activity flux to indicate the horizontal propagation of Rossby waves. The wave activity flux (W) is defined as (Takaya and Nakamura 2001)
W=12|U¯|[u¯(ψx2ψψxx)+υ¯(ψxψyψψxy)u¯(ψxψyψψxy)+υ¯(ψy2ψψyy)],
where U¯ represents the horizontal wind velocity background, ψ refers to the streamfunction, u¯ and υ¯ are climatological zonal and meridional wind velocities, respectively. Besides, primes refer to perturbations, and x and y denote the zonal and meridional gradients, respectively.

3. Evolution of GPI during the past 20 ka

During the colder LGM as seen in both the model and reconstructions (e.g., Korty et al. 2012a; Tierney et al. 2020; Waelbroeck et al. 2009) (Fig. 1a), TC genesis potential shifts northward during the storm season (JASO) relative to present, based on the GPI, with increased favorability for storm formation north of ∼15°N and decreased one to the south (Fig. 1a). The poleward shift is in agreement with previous studies based on PMIP models (Koh and Brierley 2015; Korty et al. 2012a). Subsequently, TC genesis potential over the WNP exhibits significant variations (Fig. 1b), especially during the last deglaciation. Briefly, environmental conditions are more favorable for TC genesis over the WNP during the Heinrich Stadial 1 (HS1; ∼17–16 ka) relative to the LGM (∼20 ka) based on the GPI. This is followed by a sharp decline in TC genesis potential during the Bølling–Allerød (BA; ∼13.5 ka) and then a weak rising during the Younger Dryas (YD; ∼12.9–11.7 ka). Afterward, TC genesis potential shows a slight decreasing trend over the WNP during the early–middle Holocene (∼11.7–6 ka), whereas an obvious increasing trend since the mid-Holocene (6 ka to present). Furthermore, our results indicate that the potential location of TC genesis tends to migrate northward and eastward during the HS1 relative to the LGM, and broadly shifts southward and westward subsequently (Figs. 1c,d). Next, we attempt to assess the contributions of individual genesis factors to the changes in GPI during these key periods and identify the dominant external forcing.

Fig. 1.
Fig. 1.

(a) Differences in GPI (shading) and SST (purple contours; units: °C) between the Last Glacial Maximum (LGM; 20–19 ka) and modern times during the storm season (JASO). Green lines represent modern main development region for TCs (GPI > 3 × 10−13). (b) Evolution of GPI over the western North Pacific (5°–35°N, 120°E–180°) during the past 20 ka (101-yr running average). The cold events (Heinrich Stadial 1 and Younger Dryas) and the warm event (Bølling–Allerød) are marked as the cyan and yellow vertical panels, respectively. (c) Zonal evolution of GPI anomaly averaged over 120°E–180° relative to the LGM. (d) Meridional evolution of GPI anomaly averaged over 5°–35°N relative to the LGM. The units of GPI are events m−2 season−1 10−13.

Citation: Journal of Climate 36, 10; 10.1175/JCLI-D-22-0638.1

a. Increased GPI during the LGM–HS1 transition

Although SST decreases over the majority of Northern Hemisphere during the HS1 relative to the LGM, it shows a slight increase over the WNP, broadly consistent with geological records on millennial time scale (Bolliet et al. 2011; Chen et al. 2010; Rosenthal et al. 2003; Stott et al. 2002; H. Yu et al. 2009) (Fig. 2a). Difference in GPI between the HS1 and the LGM shows a dipole pattern during the storm season over the WNP, with an increase north of ∼10°N and a decrease to the south (Fig. 2b), indicating a northward shift of the potential location of TC genesis during the HS1. Based on our sensitivity tests (see section 2c for details), decreased absolute vorticity plays a dominant role in the decreased GPI at the deep tropics (5°–10°N) during the HS1, whereas the reduction of moist entropy deficit is more important for the increased GPI at the relatively higher latitudes of the WNP (10°–35°N) (Figs. 2c,d). When averaged over the entire WNP, the increased GPI is mainly caused by larger potential intensity and lower moist entropy deficit, though absolute vorticity plays an opposite role (Figs. 2c,d).

Fig. 2.
Fig. 2.

(a) Differences in SST (units: °C) between the Heinrich Stadial 1 (HS1; 17–16 ka) and the LGM (20–19 ka). (b) Differences in GPI between the HS1 and the LGM. (c) Differences in GPI between the HS1 and the LGM (blue bars) resulting from changes in potential intensity (PI), moist entropy deficit (χ), vertical wind shear (VS), absolute vorticity (AV), and nonlinear effect (see section 2c for details). (d) Fractional changes in PI, χ, VS, and AV between the HS1 and the LGM (units: %). (e) Differences in GPI between the HS1 and the LGM during summer and autumn seasons (JJASON) in four sensitivity experiments. Nonfilled, dotted, and hatched bars show the regional mean over 5°–35°N, 120°E–180°; 5°–10°N, 120°E–180°; and 10°–35°N, 120°E–180°, respectively. Green stars represent the differences are significant at 95% significance level based on the Student’s t test. The units of GPI are events m−2 season−1 10−13 in (b) and (c) and 0.5 events m−2 month−1 10−13 in (e).

Citation: Journal of Climate 36, 10; 10.1175/JCLI-D-22-0638.1

Next, we discuss the dominant external forcing for the variation of GPI over the WNP and the associated mechanisms. The sensitivity experiments show that the increase of GPI during the HS1 relative to the LGM is dominated by the enhanced injection of meltwater fluxes (Figs. 2e and 3e–h). Then, we investigate how the meltwater fluxes influence large-scale climate over the WNP and hence TC genesis potential during the LGM–HS1 transition. In response to the increased meltwater fluxes (Fig. 3c) and the associated weaker Atlantic meridional overturning circulation (AMOC) (He 2011; Liu et al. 2009), the SST increases over the tropical North Pacific (∼5°–20°N) via air–sea interactions (Kodama 1999; Lu and Dong 2008) and the variation of thermocline (L. Yu et al. 2009; Zhang and Delworth 2005) (Fig. 4a). The higher relative SST [defined as the difference between the local SST and tropical mean (∼20°S–20°N)] contributes to the increased potential intensity (Fig. S2a). Meanwhile, in response to warmer relative SST, precipitation is increased over the tropical North Pacific (∼10°–20°N), accompanied by anomalous ascending motion (Figs. 4b,c). This leads to increased relative humidity in the midtroposphere and hence reduced moist entropy deficit (Fig. 4c, Fig. S2b). Additionally, the enhanced precipitation results in larger diabatic heating, which triggers an anomalous cyclonic (anticyclonic) circulation in the lower (upper) troposphere over the northwestern flank of the heating source (∼10°–40°N) based on the Gill response (Gill 1980) (Figs. 4d,e). Owing to the anomalous circulations, vertical wind shear is strengthened and weakened over the western and eastern main development region (MDR), where modern TC generally forms (GPI > 3 × 10−13), respectively (Fig. S2c). Overall, the variations of TC genesis factors during the HS1 relative to the LGM are linked with the SST anomaly over the tropical North Pacific and the associated adjustment of atmospheric circulations, which is in turn largely caused by the increased meltwater fluxes.

Fig. 3.
Fig. 3.

(a)–(d) Evolution of external forcings used in TraCE simulations. (a) Orbital insolation at 20°N in summer (red line) and autumn (purple line) (units: W m−2). (b) Greenhouse gas concentration (units: ppm). (c) Meltwater fluxes in the Northern Hemisphere (units: meters of global sea level rise per 1000 years). (d) Ice volume in the Northern Hemisphere (units: 107 km3). (e)–(h) Evolution of GPI over the western North Pacific (5°–35°N, 120°E–180°) during summer and autumn seasons (JJASON) (11-point running average based on 10-yr mean) in the (a) TraCE-ORB, (b) TraCE-CO2, (c) TraCE-MWF, and (d) TraCE-ICE experiments. The cold events (Heinrich Stadial 1 and Younger Dryas) and the warm event (Bølling–Allerød) are marked as the cyan and yellow vertical panels, respectively. The units of GPI are events m−2 month−1 10−13.

Citation: Journal of Climate 36, 10; 10.1175/JCLI-D-22-0638.1

Fig. 4.
Fig. 4.

Differences in large-scale environmental conditions between the Heinrich Stadial 1 (17–16 ka) and the Last Glacial Maximum (19–18 ka) during summer and autumn seasons (JJASON) in the TraCE-MWF experiment. (a) Relative SST (units: °C). (b) Precipitation (shading; units: mm day−1). (c) Relative humidity at 600 hPa (shading; units: %) and vertical motion (dots; green represents upward motion and purple represents downward). (d) Wind field at 850 hPa (vectors; units: m s−1) and relative vorticity (shading; units: 10−6 s−1). (e) Wind field at 200 hPa (units: m s−1).

Citation: Journal of Climate 36, 10; 10.1175/JCLI-D-22-0638.1

Additionally, the GPI anomaly during the BA–YD transition exhibits a similar pattern to the LGM–HS1 transition (Fig. S3b), with decreased vertical wind shear playing an important role in the increased GPI (Figs. S3c,d). Similarly, increased meltwater fluxes dominate the variations of genesis factors and hence the increased GPI, via the aforementioned dynamic mechanisms. (Figs. S3e, S4, and S5).

b. Sharply decreased GPI during the BA

SST generally increases over the WNP during the BA compared with the HS1, though with two peaks and a valley in terms of temporal evolution (Fig. S6). The modeled overall SST warming is in accordance with proxies (Bolliet et al. 2011; Chen et al. 2010; Kubota et al. 2010; Rosenthal et al. 2003; Stott et al. 2002; Xu et al. 2021; H. Yu et al. 2009) (Fig. 5a). In contrast to the SST change, the evolution of GPI over the WNP shows a sharp decrease around 14.5 and 13.5 ka, and a sudden increase around 14 ka (Fig. 1b). Here, we focus on the relatively larger decrease in GPI around 13.5 ka and examine the underlying mechanisms. For the spatial pattern, the GPI shows a positive (negative) anomaly over the southwestern (northeastern) WNP during the BA relative to the HS1, indicating a southwestward shift of the potential location of TC genesis (Fig. 5b). The positive GPI anomaly over the southwestern WNP largely arises from the weakened vertical wind shear, whereas variations of the four TC genesis factors all contribute to the decreased GPI over the northeastern WNP (Figs. 5c,d). When averaged over the entire WNP, the decreased GPI is mainly caused by the decline of potential intensity and rising moist entropy deficit, though weakened vertical wind shear favors TC formation (Figs. 5c,d).

Fig. 5.
Fig. 5.

(a) Differences in SST (units: °C) between the Bølling–Allerød (BA; 14–13 ka) and the HS1 (17–16 ka). (b) Differences in GPI between the BA and the HS1. (c) Differences in GPI between the BA and the HS1 (blue bars) resulting from changes in PI, moist entropy deficit (χ), VS, AV, and nonlinear effect (see section 2c for details). (d) Fractional changes in PI, χ, VS, and AV between the BA and the HS1 (units: %). (e) Differences in GPI between the BA and the HS1 during summer and autumn seasons (JJASON) in four sensitivity experiments. Nonfilled, dotted, and hatched bars show the regional mean over 5°–35°N, 120°E–180°, region 1, and region 2, respectively. Green stars indicate that the differences are significant at 95% significance level based on the Student’s t test. The units of GPI are events m−2 season−1 10−13 in (b) and (c) and 0.5 events m−2 month−1 10−13 in (e).

Citation: Journal of Climate 36, 10; 10.1175/JCLI-D-22-0638.1

The impacts of external forcings on GPI change over the WNP from the HS1 to the BA are relatively complex (Fig. 5e). Higher insolation during the BA leads to a general increase (decrease) in GPI over the southwestern (northeastern) WNP relative to the HS1 (Figs. 3a and 5e). The rising CO2 gives rise to general increase of GPI over the WNP, especially the southwestern WNP (Figs. 3b and 5e). The decreased meltwater fluxes result in the decrease of GPI over the WNP and cause the location of TC genesis potential to shrink southwestward (Figs. 3c and 5e), whereas the lowering of ice sheets in North America generally reduces the GPI over the WNP (Figs. 3d and 5e). The effect of decreased meltwater fluxes on the genesis factors and GPI during the BA is just opposite to that during the HS1 (Figs. S7, S8). Therefore, we focus on the roles of ice sheets and CO2 in regulating the TC genesis potential over the WNP during the BA relative to the HS1, with the influence of orbital insolation discussed in section 3c.

In the TraCE-ICE experiment, the lowering of continental ice sheets in North America alters Earth’s orography, which changes the dynamic and thermal conditions over there, and hence leads to an eastward propagation of Rossby waves in the upper troposphere (e.g., Gao et al. 2020; Shi and Yan 2019) (Fig. 6a). As a result, there are negative geopotential height anomalies and anomalous cyclonic circulation in the upper troposphere over the WNP (Fig. 6b). The anomalous cyclonic circulation not only strengthens the wind shear over the eastern MDR (Fig. S9c), but also favors the convergence in the upper troposphere, which may lead to the descending motion over the WNP (Fig. 6c). This further contributes to reduced relative humidity in the middle troposphere and hence increased moist entropy deficit (Fig. S9b). Additionally, in response to the weakened AMOC (Shi and Yan 2019; Yan and Zhang 2017; Zhu et al. 2014), relative SST is decreased over the tropical ocean, especially the WNP (∼5°–20°N), leading to decreased potential intensity over the WNP (Fig. 6d, Fig. S9a). Overall, the lowering of continental ice sheets in North America regulates the TC genesis factors over the WNP through triggering an eastward propagation of Rossby waves and the effect of sea surface cooling induced by the weakening of AMOC.

Fig. 6.
Fig. 6.

Differences in large-scale environmental conditions between the Bølling–Allerød (14–13 ka) and the Heinrich Stadial 1 (17–16 ka) during summer and autumn seasons (JJASON) in the TraCE-ICE experiment. (a) Geopotential height with zonal mean removed (contours; units: gpm) and wave activity flux (vectors; units: m2 s−2) at 200 hPa. (b) Wind field at 200 hPa (units: m s−1). (c) Relative humidity at 600 hPa (shading; units: %) and vertical motion (dots; green represents upward motion and purple represents downward). (d) Relative SST (units: °C).

Citation: Journal of Climate 36, 10; 10.1175/JCLI-D-22-0638.1

In the TraCE-CO2 experiment, the rising CO2 leads to general warming in the troposphere, with stronger warming over the eastern tropical North Pacific (∼150°E–120°W) (Fig. 7a). The inhomogeneous warming leads to anomalous anticyclonic circulation over there, accompanied by anomalous upper-level easterlies at the deep tropics (Fig. 7b), which decreases the vertical wind shear over the southern WNP (∼5°–15°N) (Fig. S10c). The anomalous anticyclonic circulation further leads to anomalous ascending motion, which contributes to the increased midtropospheric relative humidity and hence smaller moist entropy deficit (Fig. 7c, Fig. S10b). Additionally, the larger relative SST over the southern WNP (∼5°–10°N) contributes to the increased potential intensity (Fig. 7d, Fig. S10a). Overall, the rising CO2 regulates the TC genesis factors and hence genesis potential via modulating the thermodynamic structure of atmosphere over the WNP and the associated variation of SST.

Fig. 7.
Fig. 7.

Differences in large-scale environmental conditions between the Bølling–Allerød (14–13 ka) and the Heinrich Stadial 1 (17–16 ka) during summer and autumn seasons (JJASON) in the TraCE-CO2 experiment. (a) Tropospheric temperature averaged between 200 and 850 hPa (units: °C). (b) Wind field at 200 hPa (units: m s−1). (c) Relative humidity at 600 hPa (shading; units: %) and vertical motion (dots; green represents upward motion and purple represents downward). (d) Relative SST (units: °C).

Citation: Journal of Climate 36, 10; 10.1175/JCLI-D-22-0638.1

c. Two stage change in GPI during the Holocene

During the Holocene, TC genesis potential over the WNP experiences a slight decrease during the early–middle Holocene (∼11.7–6 ka), followed by an obvious upward trend from the mid-Holocene (6 ka) to the present (Fig. 1b). Here, we focus on the significant increasing trend during 6–0 ka and study how TC genesis potential over the WNP responds to changes in external forcings.

Since the mid-Holocene, SST shows an overall increasing trend over the WNP (Fig. 8a), while decreasing trend at higher latitudes in the Northern Hemisphere. Regarding the spatial distribution, GPI generally shows a decreasing (increasing) trend over the southwestern (northeastern) MDR from the mid-Holocene to the present, respectively (Fig. 8b), implying a northeastward shift of the potential location of TC genesis. The decreasing trend over the southwestern WNP mainly arises from enhanced vertical wind shear, whereas the increasing trend over the northeastern WNP is largely linked with reinforced potential intensity and weakened vertical wind shear (Figs. 8c,d). When averaged over the entire WNP, the increasing trend of GPI is mainly induced by increased potential intensity (Figs. 8c,d).

Fig. 8.
Fig. 8.

(a) Trends (per 10 ka) of SST (units: °C) from the mid-Holocene to the present (6–0 ka). (b) Trends (per 10 ka) of GPI from the mid-Holocene to the present (6–0 ka). (c) Trends (per 10 ka) of GPI from the mid-Holocene to the present (6–0 ka) (blue bars) resulting from changes in PI, moist entropy deficit (χ), VS, AV, and nonlinear effect (see section 2c for details). (d) Trends (per 10 ka) of PI (units: m s−1), χ (units: 10−2), VS (units: m s−1), and AV (units: 10−6 s−1) from the mid-Holocene to the present (6–0 ka). (e),(f) Trends [units: (10 ka)−1] of GPI from the mid-Holocene to the present (6–0 ka) in (e) summer and (f) autumn in four sensitivity experiments. Nonfilled, dotted, and hatched bars show the regional mean over 5°–35°N 120°E–180°, region 1, and region 2, respectively. The units of GPI are events m−2 season−1 10−13 in (b) and (c) and events m−2 month−1 10−13 in (e) and (f).

Citation: Journal of Climate 36, 10; 10.1175/JCLI-D-22-0638.1

Next, we try to identify the dominant external forcing for the increasing GPI and associated mechanisms from the mid-Holocene to the present. In the sensitivity experiments, the variation of GPI is dominated by the effect of orbital insolation, but the impacts of orbital insolation in summer and autumn are different (Figs. 3e–h). During the early–middle Holocene (∼11.7–6 ka), TC genesis potential increases in summer and decreases in autumn, respectively (Fig. 3e). Thus, given the offset between the two seasons, there is no significant trend in GPI during the storm season in the TraCE-full simulation (Fig. 1b). From the mid-Holocene to the present, the increasing trend is dominated by the role of orbital insolation in autumn (Figs. 3e and 8e,f), though there is a slight decrease in GPI in summer (Fig. 3e). Therefore, we mainly discuss how orbital insolation influences the genesis potential over the WNP in autumn from the mid-Holocene to the present.

In the TraCE-ORB experiment, tropospheric temperature shows an overall cooling trend in autumn, which might be attributed to lagged response to the decreased insolation in summer (Jiang et al. 2020) (Figs. 3a and 9a). However, the cooling is not spatially homogeneous, with weaker cooling over the central North Pacific (∼20°–30°N). This decreases the meridional temperature contrast over ∼10°–20°N and induces anomalous upper-level easterlies, which leads to increased (decreased) vertical wind shear over the southern (northern) MDR (Fig. 9b, Fig. S11c). Meantime, the relative SST generally shows a dipole pattern over the WNP, contributing to the decreased potential intensity over the southern WNP (∼5°–15°N) and increased one to the north (∼15°–30°N) (Fig. 9d, Fig. S11a). Overall, orbital insolation regulates the thermodynamic structure of atmosphere and SST over the WNP, which further influences the large-scale circulations and hence genesis factors.

Fig. 9.
Fig. 9.

Trends [unit: (10 ka)−1] of large-scale environmental conditions from the mid-Holocene to the present (6–0 ka) in autumn in the TraCE-ORB experiment. (a) Tropospheric temperature averaged between 200 and 850 hPa (units: °C). (b) Wind field at 200 hPa (units: m s−1). (c) Relative humidity at 600 hPa (shading; units: %) and vertical motion (dots; green represents upward motion and purple represents downward). (d) Relative SST (units: °C).

Citation: Journal of Climate 36, 10; 10.1175/JCLI-D-22-0638.1

4. Conclusions

Based on the TraCE-21ka simulations, we investigate the variation of TC genesis potential over the WNP during the past 20 ka and explore the associated physical mechanisms and dominant external forcings. Employing the GPI, we found that TC genesis potential exhibits fluctuations of large amplitude during the last deglaciation, with a long-term increasing trend since the mid-Holocene. The GPI experiences an increase during the HS1 relative to the LGM averaged over the WNP, with a positive (negative) anomaly over the northern (southern) WNP. This is followed by a substantial decrease of genesis potential during the BA, with more favorable conditions for TC formation shifting southwestward, which is opposite to the situation during the YD. During the early–middle Holocene, there is a slight decrease of GPI, while the genesis potential exhibits a significant increasing trend from the mid-Holocene to the present, with a northeastward shift of genesis location.

Moreover, the relative importance of each genesis factor and external forcings vary across periods. During the HS1 relative to the LGM, the overall increase in GPI is mainly attributed to increased potential intensity and reduced moist entropy deficit, with increased meltwater fluxes as the dominant forcing that modulates the large-scale environmental factors over the WNP via the SST variation induced by the weakened AMOC (Fig. 10a). During the BA relative to the HS1, the sharp decrease of GPI arises from lower potential intensity and larger moist entropy deficit. The decreased meltwater fluxes and the lowering of ice sheets dominate the decreased GPI, though the rising CO2 is conducive to the TC formation. The lowering of ice sheets in North America regulates the environmental factors over the WNP through triggering an eastward propagation of Rossby waves and the effect of sea surface cooling (Fig. 10b), whereas the rising CO2 influences genesis potential via modulating the thermodynamic structure of atmosphere and SST (Fig. 10c). From the mid-Holocene to the present, the long-term increasing trend is associated with increased potential intensity and weakened vertical wind shear, which is controlled by the orbital insolation (especially in autumn) via modulating the structure of air temperature and SST (Fig. 10d).

Fig. 10.
Fig. 10.

Schematic diagram for changes in large-scale environmental conditions over the western North Pacific (WNP) induced by (a) the increased meltwater fluxes, (b) the lowering of ice sheets, (c) the rising CO2, and (d) the decreased orbital insolation. The ocean is shaded by relative SST (RSST). (a) The higher RSST over the WNP induced by increased meltwater fluxes not only leads to increased PI, but also induces the anomalous ascending motion and decreased moist entropy deficit (χ) via increasing the relative humidity (RH) in the midtroposphere, which favors the TC formation. (b) The eastward propagation of Rossby waves induced by the lowering of ice sheets results in anomalous cyclonic circulation in the upper level over the WNP, leading to increased VS and anomalous descending motion, which is not conducive to the TC formation, together with the effect of lower RSST. (c) The rising CO2 modulates the air temperature and circulations, resulting in decreased VS and anomalous ascending motion, which favors the TC formation, together with the effect of higher RSST. (d) The decrease of orbital insolation alters the air temperature and circulations, leading to decreased VS, which favors the TC formation, together with the effect of higher RSST.

Citation: Journal of Climate 36, 10; 10.1175/JCLI-D-22-0638.1

As geological evidence for past TC activity can only be traced back to the past several millennia, our results are hard to constrain directly. However, our results are supported by previous studies based on PMIP models. As discussed in section 3, during the LGM, our results show that the potential location of TC genesis over the WNP shift northeastward relative to the preindustrial era (Figs. 1a,c,d), which is consistent with the findings based on PMIP2 and 3 models (Koh and Brierley 2015; Korty et al. 2012a). From the mid-Holocene to the present, our results show that TC genesis potential decreases (increases) over the southwestern (northeastern) MDR (Fig. 8b), in agreement with the results focusing on the mid-Holocene based on PMIP2 and 3 models (Koh and Brierley 2015; Korty et al. 2012b). Additionally, the anomaly patterns of GPI during the LGM and mid-Holocene share many similarities with the distributions of TC genesis and track density change based on the downscaling technique (Lawton et al. 2021; Pausata et al. 2017; Yoo et al. 2016). These results highlight the robustness of our findings.

As the GPI used here provides no information of TC track and landfalls, high-resolution regional climate models are needed to obtain more information on the TC behavior since the LGM, but this requires huge amounts of computational resources. Moreover, additional proxies for TC activity and other indirect records developed in the future could be used to evaluate our model results. Additionally, we simply divide the WNP into two regions based on GPI change to discuss the mechanisms, owing to the coarse resolution of data we use. However, it is important to further investigate the variation of TC activity over different subbasins of WNP with higher resolution simulations, as the dominating processes may be different (Mei and Li 2022). Nevertheless, our study helps to place the current TC activity over the WNP in a long-term paleoclimate context and advance our understanding on the response of TCs to various external forcings. We highlight the important roles of meltwater fluxes, ice sheets, and orbital insolation in regulating the TC genesis potential, which has not been fully considered in current TC projections.

Acknowledgments.

We sincerely thank Prof. C. M. Brierley and two anonymous reviewers for their very constructive and useful comments that helped to greatly improve our manuscript. This study is supported by National Natural Science Foundation of China (42022036, 41888101) and the Youth Innovation Promotion Association by CAS (2019080).

Data availability statement.

The TraCE-21ka simulations can be obtained from the Earth System Grid of the National Center for Atmospheric Research (https://www.earthsystemgrid.org/project/trace.html). The International Best Track Archive for Climate Stewardship version 4 can be obtained from NOAA National Centers for Environmental Information (https://www.ncdc.noaa.gov/ibtracs).

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