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    Fig. 1.

    Buoy and mooring design at station 3.

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    Fig. 2.

    Tropical cyclones and observation stations.

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    Fig. 3.

    Variations in (a)–(c) temperature and (g)–(i) temperature anomalies before and after Kalmaegi above 150 m at the observation stations. Also shown are the averaged (d)–(f) vertical temperature and (j)–(l) temperature anomaly profiles before (black lines; averaged 24–48 h before the TC) and after (red lines; averaged for the second inertial period) Kalmaegi at the observation stations. In (a)–(c) and (g)–(i), the black solid lines are the depths that are 0.5°C lower than the sea surface temperature, the horizontal black dashed lines are the 21°C isotherms for stations 1 and 3 and 20°C for station 2, the vertical black lines show the arrival time of Kalmaegi (day 0), and the vertical brown dashed lines represent the range of the second inertial period after Kalmaegi.

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    Fig. 4.

    As in Fig. 3, but for (top) Rammasun, (middle) Sarika, and (bottom) Haima at the observation stations. The horizontal black dashed lines are the 21°C isotherms for Rammasun and 20°C isotherms for Sarika and Haima.

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    Fig. 5.

    (a)–(c)Sea surface shortwave radiation (black), longwave radiation (orange), sensible heat flux (blue), and latent heat flux (red) during Kalmaegi at S1, S2, and S3, respectively. (g)–(i) As in (a)–(c), but during Rammasun at S2, during Sarika at S4, and during Haima at S4, respectively. (d)–(f) Heat content (KJ cm−2) above 80 m (0–80 m; red), greater than 26°C [≥26°C, tropical cyclone heat potential (TCHP); orange], accumulated surface heat flux (blue, Qflux), and heat content within the mixed layer (ΔQinit + ΔQsub) during Kalmaegi at S1, S2, and S3, respectively. (j)–(l) As in (d)–(f), but during Rammasun at S2, during Sarika at S4, and during Haima at S4, respectively. Black dashed lines are zero lines. Accumulated surface heat flux values are the accumulated values of (a)–(c) and (g)–(i) regarding the change in water temperature, and the accumulated surface heat flux is set to 0 on day 0.

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    Fig. 6.

    (a) Surface fluxes (W m−2), (b) ocean temperature anomalies above 100 m (ΔT; °C), and (c) change in heat content (KJ cm−2) at station 1 before and after Kalmaegi. (d)–(f) As in (a)–(c) but for station 3 for Kalmaegi. (g)–(i) As in (a)–(c), but for station 2 for Rammasun. Day 0 is the time when TC is closest to the observation station. In (a), (d), and (g), the black line is the shortwave radiation, the orange line is the longwave radiation, the blue line is the sensible heat flux, and the red line is the latent heat flux. In (c), (f), and (i), the orange line is the heat anomaly of the water warmer than 26°C, which is also the tropical cyclone heat potential (TCHP), the red line is the heat anomaly in the initial mixed layer (ΔQinit), the blue line is the accumulated heat change caused by heat flux (Qflux, with 0 KJ cm−2 at day 0), the black solid line is ΔQinitQflux, and the black dashed line is the zero line. Anomalies are the original values minus their average from 24 to 48 h before the TC.

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    Fig. 7.

    (a)–(d) Temperature (T; °C), (e)–(h) temperature anomaly (ΔT; °C), and (i)–(l) net temperature anomaly (ΔTn; °C) before and after the tropical cyclones. Day 0 is the time when the tropical cyclone is closest to the observation station. The temperature anomaly is the temperature value minus its average from 24 to 48 h before the TC. The net temperature anomaly is the temperature anomaly low pass filtered that is lower than a 3-day period. Black arrows represent the propagation direction of the subsurface anomaly.

  • View in gallery
    Fig. 8.

    (a)–(f) Sketches of the vertical upper ocean temperature profiles modulated by (a) only mixing, (b) a combination of mixing and moderate upwelling, (c) a combination of mixing and downwelling, (d) a combination of mixing and strong upwelling that forces the mixed layer to be shallower than the initial mixed layer depth, (e) very strong upwelling with an outcropping isotherm so that only upwelling works, and (f) only downwelling. The term h0 is the initial mixed layer depth; hE and hT are the well-mixed layer thickness and transition layer thickness, respectively; and hU is the vertical movement of the isotherm by upwelling (positive value) or downwelling (negative value). The dashed (solid) line is the temperature profile before (after) TC forcing. The blue (red) shadow represents a cold (warm) anomaly. ΔQinit and ΔQmix are the heat changes in the initial mixed layer and forced mixed layer, respectively, and ΔQsub is the heat change in the subsurface between the initial and forced mixed layers. (g) Sketch of the effect of mixing and vertical advection (upwelling and downwelling) and the distribution of vertical temperature anomaly profiles in the across-track section of a tropical cyclone.

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    Fig. 9.

    Nondimensional SST change [ ΔSST=ΔSST/(h0Tz)] and mixing-induced near-surface cold anomaly ( ΔQcold) on the parameter space of the nondimensional transition layer thickness ( HT=hT/h0) and nondimensional well-mixed layer thickness ( HE=hE/h0). Note that ΔSST′ in (a) is on the condition when the nondimensional upwelling depth HU < 1 + HE + HT.

  • View in gallery
    Fig. 10.

    Sketches of the heat modulation after a TC. (a) Three-dimensional net heat transport by advection after a TC, with a sketch of heat transport in (b) upwelling and (c) downwelling regions. The brown dot in (a) represents the position of the TC center.

  • View in gallery
    Fig. A1.

    As in Fig. 3, but for filtered 3-day low-pass filtered temperature and temperature anomaly. The dashed red and black lines are the averaged of the raw (unfiltered) temperature profile before and after Kalmaegi, like the lines in Figs. 3d–f and 3j–l.

  • View in gallery
    Fig. A2.

    As in Fig. A1, but for Rammasun, Sarika, and Haima; the dashed red and black lines are the same as in Figs. 4d–f and 4j–l.

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Modulation of Upper Ocean Vertical Temperature Structure and Heat Content by a Fast-Moving Tropical Cyclone

Han ZhangaState Key Laboratory of Satellite Ocean Environment Dynamics, Second Institute of Oceanography, Ministry of Natural Resources, Hangzhou, China
bSouthern Marine Science and Engineering Guangdong Laboratory (Zhuhai), Zhuhai, China
cShanghai Typhoon Institute, China Meteorological Administration, Shanghai, China

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Abstract

The ocean temperature response to tropical cyclones (TCs) is important for TC development, local air–sea interactions, and the global air–sea heat budget and transport. The modulation of the upper ocean vertical temperature structure after a fast-moving TC was studied at the observation stations in the northern South China Sea, including TCs Kalmaegi (2014), Rammasun (2014), Sarika (2016), and Haima (2016). The upper ocean temperature and heat response to the TCs mainly depended on the combined effect of mixing and vertical advection. Mixing cooled the sea surface and warmed the subsurface, while upwelling (downwelling) reduced (increased) the subsurface warm anomaly and cooled (warmed) the deeper ocean. An ideal parameterization that depends on only the nondimensional mixing depth (HE ), nondimensional transition layer thickness (HT ), and nondimensional upwelling depth (HU ) was able to roughly reproduce sea surface temperature (SST) and upper ocean heat change. After TCs, the subsurface heat anomalies moved into the deeper ocean. The air–sea surface heat flux contributed little to the upper ocean temperature anomaly during the TC forcing stage and did not recover the surface ocean back to pre-TC conditions more than one and a half months after the TC. This work shows how upper ocean temperature and heat content varies by a TC, indicating that TC-induced mixing modulates the warm surface water into the subsurface, and TC-induced advection further modulates the warm water into the deeper ocean and influences the ocean heat budget.

Significance Statement

Tropical cyclones can cause a strong ocean response that modulates the upper ocean temperature structure and contributes to the local heat budget and transport. This manuscript shows how mixing and vertical advection modulate upper ocean temperature after four fast-moving tropical cyclones, and then gives a parameterization of how sea surface temperature and upper ocean heat change depend on the two mechanisms. The temperature anomalies can propagate into deeper ocean after the tropical cyclones, and sea surface heat flux is not important for upper ocean temperature response during a tropical cyclone. These results show how the upper ocean temperature responses to a tropical cyclone, and influences the local heat budget.

© 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: Han Zhang, zhanghan@sio.org.cn

Abstract

The ocean temperature response to tropical cyclones (TCs) is important for TC development, local air–sea interactions, and the global air–sea heat budget and transport. The modulation of the upper ocean vertical temperature structure after a fast-moving TC was studied at the observation stations in the northern South China Sea, including TCs Kalmaegi (2014), Rammasun (2014), Sarika (2016), and Haima (2016). The upper ocean temperature and heat response to the TCs mainly depended on the combined effect of mixing and vertical advection. Mixing cooled the sea surface and warmed the subsurface, while upwelling (downwelling) reduced (increased) the subsurface warm anomaly and cooled (warmed) the deeper ocean. An ideal parameterization that depends on only the nondimensional mixing depth (HE ), nondimensional transition layer thickness (HT ), and nondimensional upwelling depth (HU ) was able to roughly reproduce sea surface temperature (SST) and upper ocean heat change. After TCs, the subsurface heat anomalies moved into the deeper ocean. The air–sea surface heat flux contributed little to the upper ocean temperature anomaly during the TC forcing stage and did not recover the surface ocean back to pre-TC conditions more than one and a half months after the TC. This work shows how upper ocean temperature and heat content varies by a TC, indicating that TC-induced mixing modulates the warm surface water into the subsurface, and TC-induced advection further modulates the warm water into the deeper ocean and influences the ocean heat budget.

Significance Statement

Tropical cyclones can cause a strong ocean response that modulates the upper ocean temperature structure and contributes to the local heat budget and transport. This manuscript shows how mixing and vertical advection modulate upper ocean temperature after four fast-moving tropical cyclones, and then gives a parameterization of how sea surface temperature and upper ocean heat change depend on the two mechanisms. The temperature anomalies can propagate into deeper ocean after the tropical cyclones, and sea surface heat flux is not important for upper ocean temperature response during a tropical cyclone. These results show how the upper ocean temperature responses to a tropical cyclone, and influences the local heat budget.

© 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: Han Zhang, zhanghan@sio.org.cn

1. Introduction

A tropical cyclone (TC) is a strong natural synoptic phenomenon that occurs in the ocean. Because of the strong wind stress, TCs usually cause sea surface cooling in the range of about 1°–6°C (Chandra and Kumar 2021; Potter 2018; Chowdhury et al. 2020; Wang et al. 2016), sometimes even reaching ∼11°C, resulting in the reverse of the air–sea surface sensible and latent heat fluxes (Glenn et al. 2016). TC-induced sea surface cooling has increased by ∼0.05°C during the past 20–35 years (Da et al. 2021).

Understanding the upper ocean temperature response to a TC is important because it not only controls the TC–ocean feedback but also influences local and global ocean heat transport and circulation. TC intensity relates to upper ocean temperature variability, including sea surface temperature (SST), mixed layer depth, isothermal layer depth, 26°C isotherm, and TC heat potential (Busireddy et al. 2019; Potter and Rudzin 2021). Understanding the variation in SST and ocean stratification well improves the forecasting of TC intensity (Cione and Uhlhorn 2003; Domingues et al. 2021; Emanuel 2003; Huang et al. 2015; Jaimes et al. 2015); for example, SST cooling greater than 2.5°C is not conducive to TC strengthening (Emanuel 1999) and may even weaken a TC (Lin et al. 2008; Schade and Emanuel 1999). The sharpened subsurface vertical global temperature profile may contribute to a stronger ocean coupling (cooling) effect during the intensification of future TCs (Huang et al. 2015). After the TC, the TC-induced cold wake modulates local clouds and rainfall (Ma et al. 2020; Pasquero et al. 2021) and increases shortwave and longwave radiation (Zhang et al. 2021b).

TCs have recently been regarded as a contributor to the ocean heat budget and transport. Relative to SST cooling, some warm water is modulated into the subsurface by TC-induced mixing (heat pump effect), the surface cold anomaly recovers by the air–sea surface heat flux and the subsurface warm anomaly stays in the ocean, resulting in a net heat gain in the ocean (Cheng et al. 2015; Emanuel 2001; Mei et al. 2013); this process contributes to the local heat budget as well as the ocean transport of heat poleward (Korty et al. 2008; Liu et al. 2021) and equatorward (Jansen and Ferrari 2009), and thus strengthens subtropical gyre and meridional overturning circulations (Li and Sriver 2018). The TC-induced subsurface warm water parcels travel toward the equator and resurface in the eastern equatorial Pacific, which may be the cause of the permanent El Niño in the early Pliocene epoch (Fedorov et al. 2010). Moreover, TC-driven upwelling and ocean heat uptake are able to decrease the surface geostrophic velocity of the Kuroshio by ∼14 cm s−1 (16% of the mean velocity) (Park et al. 2021).

For simplicity, some previous research assumed that TC-induced mixing (heat pump effect) dominated thermal modulation. Based on the heat pump assumption, Emanuel (2001) estimated the average column integrated ocean heat uptake by a heat pump to be ∼1.4 ± 0.7 PW during 1996, while others estimated it was ∼0.26–0.58 PW (Jansen et al. 2010; Mei et al. 2013; Sriver and Huber 2007; Sriver et al. 2008; Vincent et al. 2013). Because the total ocean heat transport is approximately 2.9 PW (Fasullo and Trenberth 2008), TCs may be an important contributor to it. A statistical analysis based on Argo data showed that the heat pump hypothesis seemed to hold for strong TCs (category ≥ 4) but not for relatively weak TCs (category ≤ 3), in which case subsurface warming was not detectable but near-surface cooling was still significant (Park et al. 2011). As TC-induced upwelling can devour the mixing-induced subsurface warm anomaly (Bueti et al. 2014; Lu et al. 2021; Price 1981; Vincent et al. 2013; Wang et al. 2020; Zhang et al. 2016), it merits further study of the TC-induced vertical temperature response.

The upper ocean temperature response to several TCs was observed by buoys and mooring arrays in the northern South China Sea, and its mechanisms were studied using numerical models (Wu et al. 2018; Zhang et al. 2016, 2018, 2019, 2020). This manuscript hopes to take an additional step and address the following two questions: 1) How do we best parameterize the contribution of upwelling and its accompanying downwelling on the upper ocean temperature profile and heat content? 2) How do the upper ocean temperature anomalies vary and recover after a TC? Section 2 describes the observation data, methods, and the four observed TCs. Section 3 studies the upper ocean temperature variations during and after TCs at the observation stations and analyzes the effects of mixing, vertical advection (upwelling/downwelling), and surface heat flux. Section 4 provides a one-dimensional parameterization method and gives the parameter space of SST and heat content change in the upper ocean based on mixing depth and vertical advection depth and verifies them by observation. Section 5 discusses the implications of this study. Section 6 presents the conclusions.

2. Data and methods

a. Observation and methods

The moored data were from the observation stations deployed in the northern South China Sea (Table 1) for Stations 1, 2, 3, and 4 (herein S1, S2, S3, and S4). The stations consist of both buoys and moorings (see Fig. 1 for the buoy and mooring design at S3 as an example). TCs Kalmaegi (2014), Rammasun (2014), Sarika (2016) and Haima (2016) were chosen for the validation of this study (Fig. 2) as they were close to our observation arrays, and the upper ocean temperature variation was measured by the temperature sensors on the stations. S1–S3 were deployed in 2014, capturing Kalmaegi and Rammasun; the SeaBird 37 recorders that observed temperature were at 20-m vertical intervals from the surface to ∼400 m and measurements were taken every 120 s. S2 (S1–S3) captured the temperature response to Rammasun (Kalmaegi); see Zhang et al. (2016, 2020) for more details of the stations. S4 was deployed in 2016, capturing TCs Sarika and Haima; the temperature sensors were at 5–20-m vertical intervals that were smaller (larger) at shallower (deeper) depths from the surface to ∼400 m on buoys and from ∼200 to 700 m, and these measurements were taken from every 1 s to every 300 s. The details of the observed temperature response refer to Zhang et al. (2016, 2018, 2019, 2020) and Wu et al. (2018, 2020).

Fig. 1.
Fig. 1.

Buoy and mooring design at station 3.

Citation: Journal of Physical Oceanography 53, 2; 10.1175/JPO-D-22-0132.1

Fig. 2.
Fig. 2.

Tropical cyclones and observation stations.

Citation: Journal of Physical Oceanography 53, 2; 10.1175/JPO-D-22-0132.1

Table 1

Location, water depth, and deployment of temperature sensors on buoys of the observation stations. Water depth of the observation stations was estimated by topography. S2, S4, S5, and S7 all consist of a buoy and mooring. Bold values indicate that the temperature sensor deployed at the depth was lost, has some error in data, or no data.

Table 1

This research used the Coupled Ocean Atmosphere Response Experiment 3.0 (COARE3.0) algorithm (Fairall et al. 2003) to calculate surface heat flux. The cloud cover data for shortwave radiation were from the U.S. National Centers for Environmental Prediction (NCEP) reanalysis products (http://www.esrl.noaa.gov/psd/data/gridded/data.ncep.reanalysis.html); air temperature, sea surface pressure, humidity, rainfall were from the in-site observations at S1–S4.

The ocean heat content is calculated by the following method:
Q=ρcpz1z2Tdz,
where Q is heat content, ρ is water density, cp is specific heat capacity for water, z1 and z2 are the depth range of water for the calculation of heat, and T is temperature. Take ρ as 1025 kg m−3 and cp as 4200 J kg−1 °C for simplicity.

In this study, the initial mixed layer depth (h0) is calculated by the temperature that is 0.5°C lower than the sea surface temperature before a TC. The post-TC well-mixed layer depth (h0 + hEhU) is the depth with highest vertical temperature gradient under h0. The term hE is the well-mixed layer thickness. The uplift distance by upwelling (hU) is the difference in the average depth of the 21°C (S1 and S3 for Kalmaegi and S2 for Rammasun) or 20°C (S2 for Kalmaegi and S4 for Sarika and Haima) isotherms before and after TCs and is positive for the uplift of the isotherm. The term hU is negative for downwelling. The transition layer thickness (hT) is taken as the thickness between h0 + hEhU and the temperature anomaly minimum under h0 + hEhU. The vertical temperature gradient for parameterization (Tz) is calculated by the average of the vertical temperature gradient of the initial temperature profile between h0 and h0 + hEhU.

b. Tropical cyclones

The TC tracks were from the China Meteorological Administration (CMA; http://tcdata.typhoon.org.cn/en/zjljsjj_zlhq.html) (Ying et al. 2014). The TC tracks from other best-track datasets (e.g., Joint Typhoon Warning Center and the Japan Meteorological Agency) were similar (Zhang et al. 2016, 2019, 2020). In Table 2, the radius of maximum wind (rmax) was obtained from the Joint Typhoon Warning Center (JTWC) dataset. The across-track distance of the station to TC (y) was calculated using the China Meteorological Administration (CMA) best-track data when the tropical cyclone was closest to our observation array. A positive (negative) distance corresponds to the right (left) side of the track. The values of Kalmaegi (2014) correspond to stations 1–3, the values of Rammasun (2014) correspond to station 2, and the values of Sarika (2016) and Haima (2016) correspond to station 4. The term τmax is the maximum wind stress of a TC.

Table 2

Information on tropical cyclones when closest to the observation array.

Table 2

As shown in Table 2, all four TCs were weak typhoons when traveling through our observation arrays, as the maximum wind speeds were from 36 to 42 m s−1, which were classified as category 1 under the Saffir–Simpson hurricane wind scale (see https://www.nhc.noaa.gov/aboutsshws.php). Note that the maximum sustained wind in the CMA dataset is 2-min averaged, while the Saffir–Simpson scale uses 1-min averaged wind, so the category may be relatively weak estimated by the CMA best-track data. Kalmaegi (Rammasun) was relatively fast (slow), with a translation speed of 8.80 m s−1 (4.68 m s−1). Sarika and Haima moved at a moderate speed, with translation speeds of 6.53 and 6.80 m s−1, respectively. All four TCs caused a significant near-inertial response in their lee, and the response biased to the right side of their tracks as their nondimensional translation speed (S) was greater than 1 (Zhang et al. 2020). The Rossby number for the mixed-layer current (R) was small (∼0.3) at the stations of Kalmaegi, Sarika, and Haima but large (R = 2.145) at S4 during Rammasun (mainly due to the shallow initial mixed layer depth than that of other TCs), which indicated that the effect of horizontal advection was not important at the stations during Kalmaegi, Sarika, and Haima but played a role at S2 during Rammasun.

3. Variation in vertical temperature in observations

a. Upper ocean temperature observations

TCs induced a significant upper ocean response at the observation stations (Figs. 3 and 4). There was background diurnal and semidiurnal tide and near-inertial response after the TCs, the temperature is averaged in one day before the TCs and one inertial period after the TCs to see the net temperature change (see Figs. 3d–f and 4j–l and Figs. 4d–f and 4j–l). The result of low-pass filtered temperature is similar; see Figs. A1 and A2 in the appendix for the 3-day low-pass filtered temperature before and after the TCs. After Kalmaegi, S1 and S3 (S2) show near-surface cold anomaly, subsurface cold (warm) anomaly and deeper ocean cold anomaly; see Figs. 3j and 3l (Fig. 3k). After Rammasun (Sarika and Haima), S2 (S4) shows a near-surface cold anomaly, subsurface warm anomaly, and deeper ocean warm anomaly.

Fig. 3.
Fig. 3.

Variations in (a)–(c) temperature and (g)–(i) temperature anomalies before and after Kalmaegi above 150 m at the observation stations. Also shown are the averaged (d)–(f) vertical temperature and (j)–(l) temperature anomaly profiles before (black lines; averaged 24–48 h before the TC) and after (red lines; averaged for the second inertial period) Kalmaegi at the observation stations. In (a)–(c) and (g)–(i), the black solid lines are the depths that are 0.5°C lower than the sea surface temperature, the horizontal black dashed lines are the 21°C isotherms for stations 1 and 3 and 20°C for station 2, the vertical black lines show the arrival time of Kalmaegi (day 0), and the vertical brown dashed lines represent the range of the second inertial period after Kalmaegi.

Citation: Journal of Physical Oceanography 53, 2; 10.1175/JPO-D-22-0132.1

Fig. 4.
Fig. 4.

As in Fig. 3, but for (top) Rammasun, (middle) Sarika, and (bottom) Haima at the observation stations. The horizontal black dashed lines are the 21°C isotherms for Rammasun and 20°C isotherms for Sarika and Haima.

Citation: Journal of Physical Oceanography 53, 2; 10.1175/JPO-D-22-0132.1

To further decompose the effects of mixing and vertical advection, a temperature 0.5°C lower than the SST is used as the variation in the well-mixed layer (h0 + hEhU), the depth of the 20° or 21°C isotherms as the variation in vertical advection (hU), and the depth between the well-mixed layer temperature anomaly minimum under the well-mixed layer as the transition layer thickness (hT); this method was similar to their decomposed depth or thickness in model simulations used in previous research (Zhang et al. 2016, 2018, 2019, 2020).

After Kalmaegi, S1–S3 all experienced net upwelling as the averaged isotherm uplifted below the mixed layer (Figs. 3d–f and Table 3). The deepened well-mixed layer quickly met the upwelling (Figs. 3a–c), with an uplift of the averaged isotherm under the forced mixed layer (Figs. 3d–f) and a negative heat anomaly in the forced mixed layer (ΔQinit + ΔQsub), which were −33.034, −21.566, and −26.011 KJ cm−2, respectively (Table 3). The heat anomaly above 150 m (ΔQ150) was also negative, which were −49.661, −28.259, and −48.431 KJ cm−2, respectively (Table 3). S1 and S3 were close to Kalmaegi’s track, and the upwelling was relatively strong and mixing was relatively weak, leading to only a net cold anomaly with no net warm anomaly in the upper ocean (Figs. 3j,l). S2 was relatively far from Kalmaegi’s track, and the upwelling was weak and mixing was strong, leading to subsurface warm anomaly (Fig. 3k).

Table 3

The averaged temperature response at the observation stations. The values in the table refer to the average temperature before and after the TCs, i.e., red solid lines in Figs. 3j–l and 4j–l as well as red dashed lines in Figs. A1j–l and A2j–l.

Table 3

The SST cold anomaly at S2 (−2.028°C) was also greater than that at S1 and S3 (−1.522° and −1.090°C) because the increased mixed layer depth at S2 (hE = 26.32 m, hT = 58.65 m) was greater than that at S1 and S3 (hE = 21.03 and 23.04 m, hT = 33.56 and 35.20 m).

Different from the stations after Kalmaegi, S2 and S4 experienced net downwelling after Rammasun, Sarika, and Haima (Fig. 4), and the deepened well-mixed layer quickly met the downwelling (Figs. 4a–c), with a downward push of the averaged isotherm under the forced mixed layer (Figs. 4d–f) and a positive heat anomaly in the forced mixed layer (ΔQinit + ΔQsub), which was −6.859, 4.513, and 17.197 KJ cm−2 respectively (Table 3). Heat content above 150 m was also positive (ΔQ150), with values of 4.043, 6.785, and 26.734 KJ cm−2 respectively (Table 3). The warm anomaly under the forced mixed layer was small at S4 after Sarika because the downwelling was weak (hU = −1.59 m). Haima further deepened the mixed layer slightly (Figs. 4c,f), but there were few near-surface cold anomalies (Figs. 4i,l) because Sarika had already mixed and uniformed the near-surface water. However, the transition layer was thicker and downwelling was stronger after Haima than after Sarika, resulting in greater and deeper subsurface warm anomalies (Figs. 4i,l, Table 3). Specifically, Sarika and Haima successively pushed some warm water into the subsurface water and deeper water.

At all stations, the dominant terms controlling the air–sea heat flux during TCs were latent heat flux and shortwave radiation (solar radiation), as shown in Fig. 5. The latent heat flux increased because the intense TC wind strengthened evaporation and shortwave radiation decreased when the cloud of the TC covered (or partly covered) the sun. The ocean lost heat during the forced stage of the TCs because the decrease in shortwave radiation was associated with TC cloud cover and the increase in latent heat flux was associated with the strong TC wind. Surface heat flux (Qflux) were −1.555, −0.121, and −3.206 KJ cm−2 at S2, S4, and S5 during Kalmaegi, −0.702 KJ cm−2 at S4 during Sarika and 0.514 KJ cm−2 at S4 during Haima. However, the ocean seemed to gain heat at S4 during Rammasun (Qflux = 2.541 KJ cm−2) as S2 was far from Rammasun’s center, where TC clouds were coarse and wind was weak. Air–sea heat flux was responsible for the colder SST and inversion of the near-surface vertical temperature profiles with a colder SST than subsurface temperature at the stations (Figs. 3d–f and 4d,f). It also influenced the pre-TC vertical temperature, as the initial mixed layer at S2 before Rammasun was shallow (h0 = 9.74 m; Fig. 4d), and the pre-TC SST was colder than the subsurface temperature at S4 before Sarika (Fig. 4e), which was due to both the seasonal cycle of surface heat flux, which was in July (October), and the warming (cooling) of the sea surface at S2 (S4) before Rammasun (Sarika). Overall, the air–sea surface heat flux played a small role in SST but was not important for changes in the upper ocean heat. Note that the four TCs studied here were relatively strong TCs with a maximum wind speed (Vmax) ≥36 m s−1 and at least typhoon intensity, and the air–sea heat flux may be more important after weaker TCs (Potter et al. 2017; Zhang et al. 2021a).

Fig. 5.
Fig. 5.

(a)–(c)Sea surface shortwave radiation (black), longwave radiation (orange), sensible heat flux (blue), and latent heat flux (red) during Kalmaegi at S1, S2, and S3, respectively. (g)–(i) As in (a)–(c), but during Rammasun at S2, during Sarika at S4, and during Haima at S4, respectively. (d)–(f) Heat content (KJ cm−2) above 80 m (0–80 m; red), greater than 26°C [≥26°C, tropical cyclone heat potential (TCHP); orange], accumulated surface heat flux (blue, Qflux), and heat content within the mixed layer (ΔQinit + ΔQsub) during Kalmaegi at S1, S2, and S3, respectively. (j)–(l) As in (d)–(f), but during Rammasun at S2, during Sarika at S4, and during Haima at S4, respectively. Black dashed lines are zero lines. Accumulated surface heat flux values are the accumulated values of (a)–(c) and (g)–(i) regarding the change in water temperature, and the accumulated surface heat flux is set to 0 on day 0.

Citation: Journal of Physical Oceanography 53, 2; 10.1175/JPO-D-22-0132.1

In addition to mixing, vertical advection and surface heat flux, the temperature and heat change can be changed by horizontal advection. As mentioned before, the Rossby number for the mixed-layer current [ R=τmax/(ρ0hmixUf)] was relatively small at the stations during Kalmaegi, Sarika, and Haima (R ≈ 0.3) but relatively large (R = 2.145) at S4 during Rammasun, so horizontal advection was not important during the forced stage of Kalmaegi, Sarika, and Haima but played a role at S2 during Rammasun. Further validation and analysis of the processes and mechanisms during TCs using three-dimensional model simulations can be found in previous research (e.g., Zhang et al. 2016, 2018, 2019, 2020).

In brief, the upper ocean temperature response depended on the relative importance of mixing and vertical advection (upwelling and downwelling), surface heat flux modulated some of the SST before and after the TC, and horizontal advection was important when the Rossby number for the mixed-layer current (R) was large.

b. Long-term variation in the upper ocean temperature after a tropical cyclone

As a TC modulates the vertical temperature structure, mainly through heat pump and cold suction effects during its forced stage, a further unknown is how the temperature anomalies recover and vary after the TC.

It is generally recognized that the air–sea surface heat flux after a TC is responsible for the recovery of the SST and mixed layer anomalies (Emanuel 2001; Price et al. 1986, 2008; Zhang et al. 2021a). After Kalmaegi, the cooled initial mixed layer did not recover back to its initial condition for at least 47 days after the TC, with cold ΔQinit and cold SST anomalies (Figs. 6a–f). This result may be because Kalmaegi traveled through the observation array on 15 September, and the seasonal cycle was spanned from autumn to winter after Kalmaegi. The surface heat flux tends to cool the near-surface ocean. After Rammasun, the sea surface heat flux warmed and recovered ΔQinit, and the tropical cyclone heat potential (TCHP) returned to that before Rammasun in 5 days (Figs. 6h,i), but the near-surface ocean cooled again after that, which may be due to subsequent air–sea processes, such as TC Matmo (2014) from 22 to 23 July. Note that the strong cooling in the mixed layer near day 60 (nearly 15 September) at S4 was due to the effect of Kalmaegi. Specifically, the air–sea surface heat flux did not recover more than one and a half months after the TCs, although the seasonal cycle after Rammasun recovered the near-surface ocean temporarily. Except the air–sea surface heat flux, the background circulation of the South China Sea may lead to the near-surface temperature changes.

Fig. 6.
Fig. 6.

(a) Surface fluxes (W m−2), (b) ocean temperature anomalies above 100 m (ΔT; °C), and (c) change in heat content (KJ cm−2) at station 1 before and after Kalmaegi. (d)–(f) As in (a)–(c) but for station 3 for Kalmaegi. (g)–(i) As in (a)–(c), but for station 2 for Rammasun. Day 0 is the time when TC is closest to the observation station. In (a), (d), and (g), the black line is the shortwave radiation, the orange line is the longwave radiation, the blue line is the sensible heat flux, and the red line is the latent heat flux. In (c), (f), and (i), the orange line is the heat anomaly of the water warmer than 26°C, which is also the tropical cyclone heat potential (TCHP), the red line is the heat anomaly in the initial mixed layer (ΔQinit), the blue line is the accumulated heat change caused by heat flux (Qflux, with 0 KJ cm−2 at day 0), the black solid line is ΔQinitQflux, and the black dashed line is the zero line. Anomalies are the original values minus their average from 24 to 48 h before the TC.

Citation: Journal of Physical Oceanography 53, 2; 10.1175/JPO-D-22-0132.1

The post-TC geostrophic adjustment and background circulation may be responsible for the subsurface and deeper ocean temperature variations as the heat does not exchange with air directly. The TC-induced subsurface cold or warm anomalies tended to move into deeper depths with time (Figs. 7m–p). There seemed to be two paths for the downward propagation of subsurface anomalies. The downward movement of the main (minor) path was slow (relatively fast) and only dozens of meters (hundreds of meters) in 2–3 weeks. A model simulation several days after Kalmaegi (Zhang et al. 2018) showed that the main path was due to the recovery of subsurface anomalies by horizontal advection, and the minor path was due to the downward propagation of near-inertial waves. However, the background circulation may be more important for temperature and heat change 10 days or more after the TCs.

Fig. 7.
Fig. 7.

(a)–(d) Temperature (T; °C), (e)–(h) temperature anomaly (ΔT; °C), and (i)–(l) net temperature anomaly (ΔTn; °C) before and after the tropical cyclones. Day 0 is the time when the tropical cyclone is closest to the observation station. The temperature anomaly is the temperature value minus its average from 24 to 48 h before the TC. The net temperature anomaly is the temperature anomaly low pass filtered that is lower than a 3-day period. Black arrows represent the propagation direction of the subsurface anomaly.

Citation: Journal of Physical Oceanography 53, 2; 10.1175/JPO-D-22-0132.1

4. Parameterization of upper ocean vertical temperature and heat content change

a. Vertical temperature profile anomaly

If mixing is the only mechanism (Fig. 8a), the surface cold anomaly equals the subsurface warm anomaly, the upper ocean heat change is zero, and the heat change under the initial mixed layer (h0) is always positive; this is the “heat pump” condition as described and studied in previous works (Emanuel 2001; Mei et al. 2013). If the “cold suction” effect is added, the upwelling (downwelling) uplifts (downdrifts) the temperature profile and makes the mixed layer shallower (deeper). In so doing, upwelling (downwelling) will induce integral cooling (warming) in the vertical column. As mentioned by previous research (e.g., Lu et al. 2021; Price 1981; Zhang et al. 2018), vertical advection (i.e., upwelling and downwelling) usually delays mixing by approximately half a day, and the interaction between mixing and vertical advection is small during fast-moving TCs. If upwelling (downwelling) is moderate, the vertical temperature profile is a cold–warm–cold anomaly with more subsurface warm anomalies than surface cold anomalies, see Fig. 8b (Fig. 8c). If upwelling is strong, there will be no subsurface warm anomalies and only cold anomaly in the vertical temperature profile (Fig. 8d). The isotherm may outcrop if upwelling is too strong (Fig. 8e), but this condition is rare because there always exists a surface mixed layer after a TC, although it may be very shallow. On the other hand, there may be only a warm anomaly under h0 if downwelling is much stronger than mixing (Fig. 8f), and this condition was observed at S4 after Haima. Figure 8 presents the potential vertical temperature profile structures caused by the so-called heat pump (mixing) and cold suction (vertical advection) processes of a TC. Note that there is a transition layer (between h0 + hE and h0 + hE + hT) with the water being partially mixed that below the vertically uniform well-mixed layer (h0 + hE), according to Price et al. (1986).

Fig. 8.
Fig. 8.

(a)–(f) Sketches of the vertical upper ocean temperature profiles modulated by (a) only mixing, (b) a combination of mixing and moderate upwelling, (c) a combination of mixing and downwelling, (d) a combination of mixing and strong upwelling that forces the mixed layer to be shallower than the initial mixed layer depth, (e) very strong upwelling with an outcropping isotherm so that only upwelling works, and (f) only downwelling. The term h0 is the initial mixed layer depth; hE and hT are the well-mixed layer thickness and transition layer thickness, respectively; and hU is the vertical movement of the isotherm by upwelling (positive value) or downwelling (negative value). The dashed (solid) line is the temperature profile before (after) TC forcing. The blue (red) shadow represents a cold (warm) anomaly. ΔQinit and ΔQmix are the heat changes in the initial mixed layer and forced mixed layer, respectively, and ΔQsub is the heat change in the subsurface between the initial and forced mixed layers. (g) Sketch of the effect of mixing and vertical advection (upwelling and downwelling) and the distribution of vertical temperature anomaly profiles in the across-track section of a tropical cyclone.

Citation: Journal of Physical Oceanography 53, 2; 10.1175/JPO-D-22-0132.1

b. Relative positions of the mixing and vertical advection

Because of the relative positions of mixing, upwelling, and downwelling, the vertical temperature profiles in Figs. 8a–f can exist on different across-track positions (see Fig. 8g). In the Northern Hemisphere, TCs cause strong upwelling on the right-rear part of a TC and downwelling in front of the TC and on both sides of the upwelling zone (Zhang et al. 2018). The vertical current shear instability is usually responsible for the deepening of the mixed layer (Lu et al. 2021). The TC induced mixing biases to the right side of the TC track because of the TC wind-current resonance (Price 1981, 1983; Price et al. 1986) and upwelling biases to the right side of the TC track because of the resonance of pressure gradient and current (Price 1983). Note that the rightward bias of mixing is greater than upwelling. As a result, a typical upper ocean temperature response biases to the right side of the TC track in the Northern Hemisphere (e.g., Lu et al. 2021; Price 1981; Price et al. 1986, 1994; Zhang et al. 2016, 2018).

The vertical temperature anomaly after a TC in the upwelling zone is similar to Fig. 8b when upwelling is moderate (S2 during Kalmaegi; Fig. 3k), similar to Fig. 8d when upwelling is strong (S1 and S3 during Kalmaegi; Figs. 3j,l). The condition that upwelling is too strong and isotherm outcrops (Fig. 8e) is not captured by our observation. The downwelling zones are on both sides of the upwelling zone (Fig. 8g), and the vertical temperature anomaly is similar to Fig. 8c, which is the condition of S4 during Rammasun (Fig. 4j) or S4 during Sarika (Fig. 4k). The vertical temperature anomaly turns to only warm anomaly under initial mixed layer as shown in Fig. 8f if mixing is small or no mixing, which is the condition of S4 during Haima (Fig. 4l). There will also be a “pure mixing” position with no vertical advection between the upwelling and downwelling zones, with a heat pump vertical temperature structure, although it seems not to be captured by our observation. In the Southern Hemisphere, the upper ocean response will bias the left side of the TC track with a mirrored flip of the left and right sides of Fig. 8g.

c. Temperature and heat change by mixing

According to Fig. 8, assuming that the initial SST is T0, the vertical temperature gradient under the mixed layer is Tz, the deepening of the mixed layer by TC is hE, the transition layer is hT and the uplift distance of the temperature profile is hU; then, hU is positive (negative) for upwelling (downwelling). Assuming mixing immediately makes the temperature uniform in the mixed layer, the sea surface temperature cooling (T1T0) is ΔSST and ΔT = −ΔSST. The mixing-induced cold anomaly (ΔQcold) should equal the warm anomaly (ΔQwarm), thus:
ΔQcold=ΔTh0+ΔT22Tz=12(hEΔTTz)(hE+hTΔTTz)Tz=ΔQwarm.
Equation (2) is suitable for most cases, except when upwelling is very strong and isotherms outcrop:
ΔSST=(T1T0)=ΔT={hE(hE+hT)(2h0+2hE+hT)TzifhU<h0+hE+hT(hUh0)TzifhUh0+hE+hT.
Note that this ignores the interaction between mixing and vertical advection as it is weak after fast-moving TCs, so vertical advection only plays a role when upwelling is very strong and isotherms outcrop. However, the outcrop case is very rare especially after fast-moving TCs.
Then ΔSST can be nondimensionalized by Tz and h0:
ΔSS T=ΔSSTh0Tz={HE(HE+HT)(2+2HE+HT)ifHU<1+HE+HTHU+1ifHU1+HE+HT,
where
HE=hEh0,HT=hTh0,HU=hUh0;
HE is the nondimensional well-mixed layer thickness, HT is the nondimensional transition layer thickness, and HU is the nondimensional upwelling depth.
Combining Eq. (3) with Eq. (2), the mixing-induced cold anomaly (ΔQcold) and warm anomaly (ΔQwarm) are as follows:
ΔQcold=ΔQwarm=ρcp(ΔTh0ΔT22Tz)=ρcphE(hE+hT)(hE2+hEhT+4h02+4h0hE+2hTh0)2(2h0+2hE+hT)2Tz,
then the nondimensional mixing-induced cold anomaly ( ΔQcold) and warm anomaly ( ΔQwarm) are
ΔQcold=ΔQwarm=ΔQcoldρcph02Tz=HE(HE+HT)(4+4HE+HE2+2HT2+HEHT)2(2+2HE+HT)2.
Figure 9 shows the ΔSST′ and ΔQcold on the nondimensional space of HE and HT. Figure 8a only shows ΔSST′ when HU < 1 + HE + HT. TC induces only cold ΔSST', and its magnitude increases with the increases of HE and HT. Note that although HE alters ΔSST' in a greater extent when HE and HT, HT is usually several times greater than HE (e.g., 2.5–40 times in our observation; Table 3), so the well-mixed layer and transition layer are both important for SST and upper ocean thermal change in the real world.
Fig. 9.
Fig. 9.

Nondimensional SST change [ ΔSST=ΔSST/(h0Tz)] and mixing-induced near-surface cold anomaly ( ΔQcold) on the parameter space of the nondimensional transition layer thickness ( HT=hT/h0) and nondimensional well-mixed layer thickness ( HE=hE/h0). Note that ΔSST′ in (a) is on the condition when the nondimensional upwelling depth HU < 1 + HE + HT.

Citation: Journal of Physical Oceanography 53, 2; 10.1175/JPO-D-22-0132.1

According to the parameterization of SST and simplified vertical temperature profile in Fig. 8, heat change in initial mixed layer (ΔQinit) is as follows:
ΔQinit=ρcp{hE(hE+hT)(2h0+2hE+hT)h0TzifhU<hE+hThE(hE+hT)(2h0+2hE+hT)h0Tz12(hE2+3hEhThEhU+3hT2+hU2)TzifhE+hThU<h0+hE+hT(hU12h0)h0TzifhUh0+hE+hT.
Therefore, ΔQinit equals ΔSST if hU < hE + hT, and became colder if upwelling uplifts the forced mixed layer (h0 + hE + hThU) that shallower than initial mixed layer (h0 + hE + hT).

d. Temperature and heat change by vertical advection

As mentioned before, upwelling (downwelling) is able to upward (downward) push the isotherms. The heat content change by upwelling (ΔQup) when upwelling is not very strong:
ΔQup=12ρcp(hEΔTTz)(hE+hTΔTTz)Tz+12ρcp(hEhUΔTTz)(hE+hThUΔTTz)Tz.
If upwelling is very strong (Fig. 8e), the effect of mixing disappears, and then
ΔQup=ρcp{hE(hE+hT)hU(2h0+2hE+hT)Tz(hE+12hT12hU)hUTzifhU<h0+hE+hThU2Tz+12h02TzifhUh0+hE+hT.
Note that ΔQup is negative (positive) when hU is positive (negative), corresponding to upwelling (downwelling). Because of the heat conservation, heat content change by upwelling (ΔQup) equals to the heat that horizontal advected and modulated into deeper ocean (−ΔQdeep).
The heat change in subsurface (ΔQsup) equals to ΔQup − ΔQinit when upwelling is not strong and equals to 0 when uplifted isotherms shallower than initial mixed layer, then
ΔQsub=ρcp{hE(hE+hT)(h0hU)(2h0+2hE+hT)Tz(hE+12hT12hU)hUTzifhU<hE+hT0ifhUhE+hT.
Note that the heat change in the forced mixed layer (ΔQinit + ΔQsub) also equals ΔQup and −ΔQdeep.

e. Validation by observation

The observation data at the stations are used to check the one-dimensional parameterization model results. The model-estimated ΔSST and heat change (Table 4) are close to those in observations after Kalmaegi, Sarika, and Haima (Table 3). However, after Rammasun, the model estimated ΔSST (−1.642°C) and heat content (ΔQinit = −5.582 KJ cm−2; ΔQsub = 4.846 KJ cm−2) are warmer than observation (−3.682°C; −11.549 and 4.689 KJ cm−2). As shown in Table 1, the Rossby number for the mixed-layer current (R) was large (2.145), so the horizontal cold advection may be important at S2 after Rammasun, which is not considered in the one-dimensional parameterization model.

Table 4

Net temperature response at the observation stations by the one-dimensional theoretical model.

Table 4

Thus, the parameterizations are able to reproduce the SST and heat change in the upper ocean for fast-moving TCs when horizontal heat advection and air–sea heat flux is weak. Note that Table 4 uses one inertial period of average vertical movement of isotherms in the observation (hU) for the calculation of HU, and a simplified near-inertial movement of water. It also ignored the delay of vertical advection relative to mixing. The choice of mixed layer and vertical movement distance can improve the parameterization, but I do not pursue this issue further, as it would make the usage of the simplified model more complicated and require more observation data for validation.

5. Discussion

a. Effect of surface heat flux

Different from expectations, the SST, heat content in the initial mixed layer (Qinit) and TCHP did not recover back to pre-TC conditions after the four TCs at the observation stations. Thus, it was determined that a TC breaks the local air–sea heat balance, which does not return to pre-TC conditions, especially during the cooling phase of the seasonal cycle (e.g., from summer to winter). In contrast, the SST seems to have entered another air–sea heat balance. As TCs influence the subtropical ocean every year, their cumulative effect may influence the local air–sea heat budget. For example, subtropical ocean surfaces may be warmer without TCs. However, more observations and studies are needed to determine whether the insufficient surface and near-surface recovery are unique or universal.

b. Combined effect of mixing and advection on heat

Because of the heat conservation, although the subsurface was not as warm as the pure heat pump assumption, the cold suction effect actually modulated the subsurface warm anomalies deeper into the ocean than the forced layer. Thus, TC modulated warm surface water into the subsurface and deeper waters. In particular, the TC-induced combined effect of the heat pump and cold suction and the downward propagation of the temperature anomaly after a TC may be an important way for surface heat to transfer into the ocean interior (e.g., the propagation of surface global warming signals into the ocean interior and deep ocean). Note that the combination effect of heat pump and cold suction cause net cold heat change near the TC core region, as the cooling caused by upwelling is narrow near the TC track and warming by downwelling is extensive that on both sides of upwelling region. After TCs, some heat anomaly may stay in the ocean along with the geostrophic adjustment. Lu et al. (2020) and Park et al. (2021) did some pioneering work about the geostrophic response that perturbs the underlying ocean eddy after a TC, which may merit applying a similar method for the temperature and heat modulation by geostrophic adjustment after a TC in further studies.

c. Control factors of the upper ocean temperature response

Generally, mixing depends on the intensity of TC wind stress, vertical advection depends on the intensity of the wind stress curl, and both depend on the upper ocean temperature structure. Previous research has shown that the upper ocean temperature response to TCs depends on the TC parameters (e.g., translation speed and wind speed) (Lu et al. 2021; Potter et al. 2017) and ocean conditions (e.g., SST, ocean stratification, and TC heat potential) (Busireddy et al. 2019; Potter and Rudzin 2021). Miyamoto et al. (2017) showed that the deepening of the mixed layer by a TC depends on a nondimensional number Co, which is the ratio of TC-induced velocity shear per depth in the initial mixed layer to the static stability beneath the mixed layer. Large-scale phenomena, such as El Niño–Southern Oscillation (ENSO), the Pacific decadal oscillation (PDO), and the North Atlantic Oscillation (NAO), can also modulate the upper ocean feedback to TCs through the influence of ocean stratification and TC-induced mixing (Cao et al. 2021).

The parameterization in section 4 show that we can estimate the change of signal-point upper ocean temperature profile and heat content after a TC with initial temperature profile, mixing depth (HE, HT), and upwelling depth (HU) after fast-moving TC with low Rossby number for the mixed-layer current. As HE, HT, and HU are controlled by TC parameters and ocean conditions, it is possible to estimate post-TC ocean single-point vertical temperature profile and heat content change based on tropical cyclone (or wind) characteristics and initial ocean temperature condition. In summary, further research on TC parameters, ocean conditions, and background processes that influence TC-induced SST and upper ocean heat responses is merited.

6. Conclusions

Typhoons Kalmaegi (2014), Rammasun (2014), Sarika (2016), and Haima (2016) traveled over our observation array consisting of buoys and moorings in the northern South China Sea and provided a valuable dataset that could be used to study the upper ocean temperature variation after TCs.

TCs cooled the SST at the observation stations, and the upper ocean temperature anomaly depended on the composition effects of mixing (heat pump) and vertical advection (cold suction). Mixing cooled the sea surface and warmed the subsurface, and upwelling (downwelling) cooled (warmed) the whole upper ocean. Subsurface temperature anomalies depended on the relative intensity of mixing and advection, which was mainly cold (warm) if upwelling was strong (weak), and downwelling could amplify mixing-induced subsurface warm anomalies. It is interesting that the subsurface warm anomaly could also amplify with little SST cooling, as mixing was weak and downwelling dominated at S4 during Haima. The upper ocean heat change was negative (positive) in the upwelling (downwelling) region. An SST and upper ocean heat change parameterizations that depend on only the nondimensional mixing depth (HE) and upwelling depth (HU) is found and used, and the SST and upper ocean heat change as well as their nondimensional parameter spaces could be approximately reproduced.

After TCs, the subsurface heat anomalies moved into the deeper ocean along with the recovery of subsurface anomalies and the downward propagation of near-inertial waves. Figure 10 provides sketches of the three-dimensional net heat transport (Fig. 10a) and the temperature modulation in the upwelling (Fig. 10b) or downwelling (Fig. 10c) region. During the TC-forced stage, the sea surface shortwave radiation decreased, and the latent heat flux increased, resulting in an air–sea heat negative heat flux. However, the air–sea heat flux contributed little to the upper ocean temperature anomaly relative to the heat pump and cold suction effects. After TCs, the SST and upper ocean temperature did not recover back to pre-TC conditions, which raises further questions about the net effect of TCs and air–sea heat flux on the sea surface and upper ocean.

Fig. 10.
Fig. 10.

Sketches of the heat modulation after a TC. (a) Three-dimensional net heat transport by advection after a TC, with a sketch of heat transport in (b) upwelling and (c) downwelling regions. The brown dot in (a) represents the position of the TC center.

Citation: Journal of Physical Oceanography 53, 2; 10.1175/JPO-D-22-0132.1

This work indicated that TCs could modulate surface heat into the subsurface by TC-induced mixing (heat pump), which was further modulated into the ocean interior and deeper ocean by TC-induced advection (cold suction). The subsurface anomalies could further propagate downward after TCs and influence local heat budget.

Acknowledgments.

This work was supported by the project supported by Southern Marine Science and Engineering Guangdong Laboratory (Zhuhai) (SML2021SP207), the National Natural Science Foundation of China (42176015) and the Scientific Research Fund of the Second Institute of Oceanography, MNR (QNYC2002), the Zhejiang Provincial Key Research and Development Project (2021C03186), the Innovation Group Project of Southern Marine Science and Engineering Guangdong Laboratory (Zhuhai) (311021001), the Shanghai Typhoon Research Foundation (TFJJ202111), and the Oceanic Interdisciplinary Program of Shanghai Jiao Tong University (SL2020MS032).

Data availability statement.

The TC tracks were from the China Meteorological Administration (CMA; http://tcdata.typhoon.org.cn/en/zjljsjj_zlhq.html). The cloud cover data for shortwave radiation were from the U.S. National Centers for Environmental Prediction (NCEP) reanalysis products (http://www.esrl.noaa.gov/psd/data/gridded/data.ncep.reanalysis.html). The observation data at the stations for this manuscript can be found on Mendeley Data (https://doi.org/10.17632/n4pj3c9nn6.1).

APPENDIX

Low-Pass Filtered Temperature

As the background diurnal and semidiurnal tide and TC-induced near-inertial waves influence the stations, it is worthy to remove the signals by low-pass filtering. Figs. A1 and A2 are 3-day low-pass filtered temperature and temperature anomalies at the stations. The variations of mixed layer and net updrift of isotherms are clear with the deepening of mixed layer and net upwelling (downwelling) at S1–S3 after Kalmaegi (S2 after Rammasun; S4 after Sarika and Haima); see Figs. A1a–c and A1j–l (Figs. A2a–c,j–l). The average temperature and temperature anomalies (Figs. A1d–f,j–l and A2d–f,j–l) are similar to that by raw data (Figs. 3d–f,j–l and 4d–f,j–l), so the study based on the average of raw data is acceptable.

Fig. A1.
Fig. A1.

As in Fig. 3, but for filtered 3-day low-pass filtered temperature and temperature anomaly. The dashed red and black lines are the averaged of the raw (unfiltered) temperature profile before and after Kalmaegi, like the lines in Figs. 3d–f and 3j–l.

Citation: Journal of Physical Oceanography 53, 2; 10.1175/JPO-D-22-0132.1

Fig. A2.
Fig. A2.

As in Fig. A1, but for Rammasun, Sarika, and Haima; the dashed red and black lines are the same as in Figs. 4d–f and 4j–l.

Citation: Journal of Physical Oceanography 53, 2; 10.1175/JPO-D-22-0132.1

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