1. Introduction
Droughts are the world’s most damaging and pressing natural disasters (Keyantash and Dracup 2002; Svoboda et al. 2002; Romm 2011), causing tens of billions of dollars in global damages and collectively affecting more people than any other form of devastating climate-related hazard (Wilhite 2000). North China, where almost half of China’s population lives and most wheat and corn are grown, is facing a serious water crisis. Since the late 1990s severe and extreme droughts have frequently occurred, and drought-affected area has been increasing by 3.72% decade−1 in the past five decades, posing great challenges for regional sustainable development (Ma and Fu 2006; Yang et al. 2013; Yu et al. 2014; Zhang and Zhou 2015). In response, the South–North Water Transfer Project was put in place to alleviate this water crisis. However, this is far from solving the long-term arid condition of north China. A recent study documented that if climate continues to warm in the future, there is a high confidence level that drought over north China will continue to increase (Zhao and Dai 2015). Thus, it is of great importance to identify the drivers and dynamic mechanisms of north China drought in order to improve drought prediction.
During the past 60 years, frequent droughts observed in north China are the result of a combination of reduced precipitation and increased temperature (Zou et al. 2005; Xu et al. 2015; Chen and Sun 2015). This is associated with a pronounced interdecadal variability. A shift from wet to dry conditions is observed over north China from the late 1970s to present, linked to weakening East Asian summer monsoon (EASM) and hydrologically manifested as the “south flood–north drought” (Wang 2001; Yu and Zhou 2007; Zhou et al. 2009). Such interdecadal variability is believed to be driven by the phase transition of the Pacific decadal oscillation (PDO) from negative to positive around the end of 1970s (Li et al. 2010; Qian and Zhou 2014). In addition to internal climate variability, anthropogenic forcing may have also contributed to the long-term changes of north China precipitation (Polson et al. 2014; Jiang et al. 2013; Menon et al. 2002; Song et al. 2014; L. Zhang et al. 2017).
Interannual variability of north China precipitation is highly modulated by the East Asian monsoon variation. In general, summer precipitation deficiency over north China is closely related to suppressed Indian summer monsoon rainfall through the Silk Road teleconnection (Hu 1997; Wu et al. 2003; Enomoto et al. 2003; Zhang and Zhou 2012), which is a regional manifestation of a circumglobal teleconnection (CGT) pattern preferentially appearing in the summer preceding El Niño peak phase (Ding and Wang 2005; Ding et al. 2011). The south flood–north drought precipitation is also observed on the interannual time scale. It is dominated by a Pacific–Japan or East Asian–Pacific teleconnection triggered by anomalous anticyclonic circulation over the northwestern Pacific, which is forced by SST anomalies both locally and remotely from the Indian Ocean during El Niño decaying summer (Huang et al. 2012; Xie et al. 2009; Wu et al. 2009, 2010). Recent studies on the north China summer drought in 2014 and 2015 documented the importance of high-latitude circulation, such as summer Eurasian teleconnection pattern (Wang and He 2015; Wang et al. 2017).
So far much attention has been paid to the whole East Asian monsoon precipitation, which is dominated by Yangtze River valley precipitation variations. There are relatively few specific investigations on the north China drought, and the mechanisms behind the north China drought are not clear. The north China rainy season lasts from June to August, while dry spells are from autumn to the next spring. Before the rainy season arrives to north China, surface temperature warms up quickly in spring, enhancing the evapotranspiration and resulting in the continuous drying of soil moisture until early summer. Thus, the highest risk for onset of the north China drought is observed in spring and summer. A statistical analysis of the north China drought revealed that about 70% extreme north China drought events start in spring and persist into summer (Chen and Yang 2013). Investigation of the prolonged precipitation deficiency from spring to summer could provide useful guidance on when the north China drought begins in spring and what contributes to its persistence until summer. Therefore, this study focuses on the prolonged spring–summer drought events over north China and aims to answer the following questions: 1) What are the key large-scale circulations anomalies causing the prolonged precipitation deficiency over north China from spring to summer? 2) What drives the persistence of the key large-scale circulation anomalies?
The remainder of the paper is organized as follows. The data used in this paper are introduced in section 2. Section 3 presents the large-scale circulation anomalies responsible for the north China prolonged spring–summer drought (PSSD). In section 4, the role of air–sea interaction in maintaining the key large-scale circulation anomalies is examined. And finally, a summary and discussion are provided in section 5.
2. Data description
a. Observed and reanalysis datasets
Monthly observational and reanalysis datasets from 1960–2014 are used in this study, including 1) global land precipitation from the CRU Time Series, version 3.1 (TS3.1), datasets with a 0.5° × 0.5° resolution for the period 1961–2014 (Harris et al. 2014); 2) Precipitation Reconstruction (PREC) data for 1961–2014 provided by the NOAA/OAR/ESRL Physical Sciences Division (PSD) (Chen et al. 2002; available online at https://www.esrl.noaa.gov/psd/data/gridded/data.prec.html) that is constructed on a 2.5° latitude–longitude grid over the globe; 3) National Centers for Environmental Prediction (NCEP)–National Center for Atmospheric Research (NCAR) reanalysis for 1961–2014 (Kalnay et al. 1996); 4) global Hadley Centre Sea Ice and Sea Surface Temperature dataset (Rayner et al. 2003); and 5) Niño-3.4 index defined as the area-averaged SST anomalies over (5°N–5°S, 170°–150°W), compiled by NOAA/Climate Prediction Center (available online at http://www.cpc.ncep.noaa.gov/products/analysis_monitoring/ensostuff/ensoyears.shtml).
b. Model simulation description and ENSO events selection
To verify the large-scale drivers of north China PSSD, a 574-yr preindustrial control simulation forced by constant preindustrial external forcing is used here. The simulation was performed at the Met Office Hadley Centre with the Hadley Centre Global Environment Model, version 2, Earth System Model (HadGEM2-ES), a configuration of the Met Office’s Unified Model (MetUM). A full description about the model configurations can be found in Jones et al. (2011). We obtained the model output from the data archive for phase 5 of the Coupled Model Intercomparison Project (Taylor et al. 2009). The models using MetUM have shown several advantages in simulating and forecasting the East Asian climate. Zhang et al. (2012) documented that the Hadley Centre Global Atmosphere Model (HadGAM1) showed the highest fidelity in simulating the East Asian–northwestern Pacific summer climate. The natural variability modes of EASM simulated by the previous version of HadGEM2-ES, HadCM3, are comparable in magnitude and in temporal and spatial characteristics to those in observations (Lei et al. 2014). The reasonable simulation of EASM climate raises the fidelity of using this model simulation to verify the large-scale drivers of north China PSSD.
Given the model bias in simulating ENSO patterns, following Wu and Zhou (2016), we apply empirical orthogonal function (EOF) analysis to the monthly SST anomalies in the tropical Pacific (20°N–20°S, 120°E–80°W) to select ENSO events in model. To remove the potential influence of decadal variability, fluctuations with periods longer than 8 years are filtered out using a Lanczos filter (Duchon 1979) before EOF analysis. Then December–February (DJF) means of the obtained principal component (PC) time series are calculated as the ENSO index. The years with D(−1)JF mean PC values less than −0.3 standard deviation and DJF(+1) greater than 1.0 standard deviation are selected as ENSO events with La Niña transitioning to El Niño. The (−1) denotes the preceding year and (+1) denotes the subsequent year of La Niña transitioning to El Niño.
3. Key large-scale circulation anomalies responsible for north China PSSD
The climate mean precipitation and circulation in boreal spring [March–May (MAM)] and summer [June–August (JJA)] for 1961–2014 are first presented in Fig. 1. Climatologically, the total precipitation amount over north China (red box in Figs. 1a,b) shows a decrease from southeast to northwest both in spring and summer, with the maximum reaching 200 and 500 mm, respectively. The spring precipitation over north China accounts for 12%–20% of the total annual amount, while summer accounts for 50%–78%. This is distinguished from southern China, where the contribution of spring precipitation is comparable to summer. In spring (Fig. 1c), low-level westerly winds prevail over north China, southward to where it is dominated by southwesterly winds associated with the northwestern Pacific subtropical high (NWPSH; shading). At the upper level, a westerly jet core is seen along 32°N centered to the southeast of north China. In summer, there are prevailing southwesterlies in north China at the low level because of northward movement of NWPSH, transporting water vapor from the ocean to north China (Fig. 1d). The upper-level jet core moves 10° latitude northward correspondingly in summer with weaker magnitude relative to that in spring.
Long-term (1961–2014) averages of seasonal precipitation (contours; mm) and its fractional contribution to the annual total (shading; %) for (a) spring and (b) summer. The corresponding dynamical fields are shown for (c) spring and (d) summer, with SLP (shading; hPa), 200-hPa zonal winds (contours; m s−1), and 850-hPa winds (vectors; m s−1). The red box in (a),(b) marks north China.
Citation: Journal of Climate 31, 9; 10.1175/JCLI-D-17-0440.1
Precipitation anomaly is selected to measure drought intensity in this study, since precipitation deficiency is the main cause for agricultural failures and hydrological water shortages. The time series of regional mean precipitation anomalies for March–August averaged over north China are shown in Fig. 2a. As in Seager et al. (2015), the top 15% years with the most deficient precipitation are chosen as drought years. Eight drought years over north China are selected from the period of 1961–2014: 1965, 1968, 1972, 1986, 1997, 2001, 2006, and 2014. The drought intensity is largely contributed by the summer total precipitation owing to a much larger precipitation amount in summer than spring (Figs. 1a,b). To check whether deficient rainfall is observed in both spring and summer, the distributions of the spring and summer precipitation anomaly percentages composited for the eight drought years relative to seasonal climatology are shown in Figs. 2b,c. Significant below-normal precipitation can be observed in both seasons with the maximum reduction by 50% in spring and by 35% in summer. It demonstrates a prolonged precipitation deficiency from spring to summer over north China for the selected drought events.
(a) Time series of regionally averaged March–August precipitation anomalies (mm) over north China (33°–45°N, 105°–120°E). (b) Spring and (c) summer precipitation anomaly percentage relative to climate mean for the top 15 driest events. Dotted areas are statistically significant at the 10% level using the Student’s t test.
Citation: Journal of Climate 31, 9; 10.1175/JCLI-D-17-0440.1
The composite circulation and water vapor transport anomalies for the drought years are examined here to identify the dynamic processes of the north China PSSD. The anomalies of 850-hPa winds, sea level pressure (SLP), and water vapor transport are shown in Fig. 3. In the north China drought years, an anomalous cyclonic circulation over the northwestern Pacific persists from spring to summer, with significant northerly wind anomalies along the East Asian coast and westerly winds anomalies over the tropical North Pacific, demonstrating a weakened NWPSH. The northerly wind anomalies in the two seasons lead to continuous anomalous water vapor divergence over north China, and thus persistent reduction of warm and wet water vapor transport from the ocean to north China (Figs. 3c,d). The anomalous low-level winds are caused by weaker land–sea thermal contrast seen from the SLP anomalies distributions with a higher pressure anomaly over land paired with two negative anomaly centers over the northwestern Pacific, although only the negative anomaly over the tropical northwestern Pacific is statistically significant at the 10% level.
Composite (left) spring and (right) summer mean lower-level circulation anomalies and atmospheric moisture transports corresponding to the top 15 driest events: (a),(b) SLP anomalies (shading; hPa) and 850-hPa anomalous winds (vectors; m s−1) and (c),(d) anomalous vertical integrated water vapor flux convergence (shading; kg m−2 s−1) and water vapor transport (vectors; g m s−1). Only vectors statistically significant at the 10% level are shown. The dotted areas in (a),(b) and (c),(d) denote SLP and water vapor flux convergence anomalies, respectively, that are statistically significant at the 10% level.
Citation: Journal of Climate 31, 9; 10.1175/JCLI-D-17-0440.1
The spatial distributions of 200-hPa zonal winds, 200-hPa geopotential height, and 500-hPa vertical motion anomalies for the drought years are displayed in Fig. 4. The westerly winds at 200 hPa to the north of the jet core get weakened, but those to the south of jet core are enhanced from spring to summer in PSSD years (Figs. 4a,b), leading to a cyclonic circulation anomaly at the upper level that dynamically forces persistent anomalous descent motions over north China. On the other hand, the colder air temperature is associated with dry air masses, which also benefits the formation of drought (Lu et al. 2014). Since zonal wind at 200 hPa is determined by meridional temperature gradient, the geopotential height anomalies at 200 hPa that represent the average tropospheric temperature anomalies between 200 hPa and surface are shown in Figs. 4c,d. Significant cold anomalies along the jet core are observed in both seasons (Figs. 4c,d) although the climatological jet core shifts northward from spring to summer. It weakens the meridional gradient to the north of the jet core but enhances that to the south, resulting in the westerly wind anomalies as shown in Figs. 4a,b.
As in Fig. 3, but for the composite upper-level circulation anomalies with (a)–(d) superimposed climatological zonal winds (contours; m s−1) and (e),(f) midlevel vertical velocity anomalies: (a),(b) anomalous 200-hPa zonal winds (shading; m s−1), (c),(d) 200-hPa geopotential height anomalies (shading; m), and (e),(f) vertical velocity anomalies at 500 hPa (shading; hPa s−1). Dotted areas are statistically significant at the 10% level using the Student’s t test.
Citation: Journal of Climate 31, 9; 10.1175/JCLI-D-17-0440.1
The above analysis demonstrates that during the north China PSSD years, a cyclonic circulation anomaly over the northwestern Pacific lasts from spring to summer with reduced moisture transport to north China. This lower-level situation is aided by an East Asian tropospheric cooling along the upper-level westerly jet with dynamically forced anomalous descent above. Because of the persistence of the two key circulation anomalies, precipitation over north China remains deficient from spring to summer.
4. Air–sea interaction associated with the north China PSSD
a. Observed air–sea interaction and potential impact of ENSO phase transition
To reveal the reason for the long persistence of large-scale circulation anomalies of the north China PSSD, the composite sea surface temperature anomaly (SSTA), precipitation, and streamfunction anomalies at 850 hPa for the drought years are presented in Fig. 5. In spring, a large-scale meridional dipole anomaly in both 850- and 200-hPa winds is seen over the North Pacific, demonstrating an equivalent barotropic vertical structure. An anomalous low-level cyclonic circulation spans from the tropical to subtropical North Pacific centered at 20°–30°N, 160°–200°E and an anomalous anticyclonic circulation is seen over the high-latitude North Pacific centered at 50°–60°N, 180°. This pattern is coincidental with a negative phase of NPO, and the two key circulation anomalies in spring are set directly by a negative NPO phase (Linkin and Nigam 2008).
As in Fig. 3, but for (a),(b) SST anomalies (shading; °C) and 850-hPa wind anomalies; (c),(d) precipitation (shading; mm day−1) and 850-hPa streamfunction (contours; 10−6 m2 s−1); and (e),(f) geopotential height (shading; m) and 200-hPa wind (vectors; m s−1). The dotted areas denote that the shading is statistically significant at the 10% level when using the Student’s t test. The green and purple contours in (c),(d) indicate positive and negative streamfunction anomalies, respectively.
Citation: Journal of Climate 31, 9; 10.1175/JCLI-D-17-0440.1
The air–sea interaction plays a crucial role in sustaining the two key circulation anomalies in spring. First of all, NPO forcing alters the net heat flux over the North Pacific through changes in the surface wind speed and consequently leads to a meridional triple SSTA pattern over the North Pacific Ocean (Vimont et al. 2001, 2003a,b) in spring. Specifically, in the tropics, anomalous westerly winds around the southern cell weaken the trade winds over the central and eastern subtropical Pacific, reducing the wind speed and upward latent flux from ocean to air, thereby warming the underlying ocean (Fig. 5a). In contrast, the anomalous northerly winds in the subtropics lead to cold advection, enhance the local wind speed, and cool the subtropical SST. Consequently, a meridional triple SSTA pattern is seen over the North Pacific with negative SSTA along the Kuroshio Extension (20°–40°N, 120°–220°E) and positive over the tropics (0°–20°N, 150°–240°E) and high latitudes (north to 50°N) of the Pacific. Second, the SSTA in turn helps to sustain the cyclonic circulation over the midlatitude North Pacific as an adaptation of atmospheric circulation to vertical differential heating (Wu et al. 2011). The midlatitude SST cooling suppresses local deep convection with maximum cooling at the upper troposphere because the maximum deep convection heating usually occurs at the height between 300 and 400 hPa. In combination with the anomalous surface heating caused by stronger wind stress, the anomalous surface heating decreases with height. Because of Sverdrup balance, northerly wind anomalies over the midlatitude central North Pacific will be generated to balance the negative vorticity anomaly caused by the decrease in heating with altitude, maintaining the cyclonic circulation anomaly over the North Pacific from spring to summer.
Because of the anomalous westerly winds associated with the south cell of NPO in spring, tropical SST continues to warm up and there is an onset of El Niño in summer (Fig. 5b). This lagged SSTA response to NPO is the so-called seasonal footprinting mechanism (SFM) (Vimont et al. 2001, 2003a,b). Forced by the warm equatorial SST anomalies in summer, convection over the central equatorial Pacific is enhanced and induces anomalous lower-level cyclonic circulation over the northwestern Pacific as a Gill–Matsuno response (Wu et al. 2009), maintaining the cyclonic anomalies over the North Pacific that has already existed in spring (Fig. 5d). Convection over the Maritime Continent and Indian summer monsoon region is suppressed due to weaker Walker circulation in El Niño summer (Fig. 5d). With a Rossby wave response to decreased heating over the Indian monsoon region, the CGT along the upper-level westerly jet stream over the African–Asian continent is triggered with reduced geopotential height (tropospheric cooling) over East Asia (Ding and Wang 2005; Ding et al. 2011). Therefore, precipitation is still below normal over north China in summer.
As documented by Vimont et al. (2001, 2003a,b) and Alexander et al. (2010), the fluctuations in NPO are most energetic and start to impact ocean in the preceding winter, and a La Niña cooling centered in the eastern tropical Pacific tends to create an NPO structure in winter that is favorable for SFM (Alexander et al. 2010; Park et al. 2013). To verify whether there is any precursor of NPO or La Niña in the preceding winter of a PSSD year, the corresponding circulation anomalies are further examined and shown in Fig. 6. Statistically significant meridional dipole SLP anomalies over the North Pacific are evident in the previous winter that demonstrates a negative NPO pattern (Figs. 6b,c). Besides the negative NPO pattern, significant cold SSTA dominates the equatorial eastern Pacific with negative SLP over the Maritime Continent and northwestern Pacific, consistent with the result of Alexander et al. (2010) and Park et al. (2013). With a Rossby wave response to enhanced convective heating forced by both the in situ ocean surface warming and the ascent forced remotely by the central Pacific cooling, another anomalous cyclonic circulation over the North Pacific is observed over the Philippine Sea (Figs. 6a,b) (Wang et al. 2000). The examination of SST evolution indicates that continuous spring–summer drought over north China tends to occur in the year when La Niña transits to El Niño. The associated air–sea interaction involved in the ENSO phase transition benefits the maintenance of the two key anomalous cyclonic circulations of the north China PSSD.
Composite circulation anomalies in the preceding winter of the top 15 driest years. (a) SST anomalies (shading; °C) and 850 hPa wind anomalies (vector; m s−1); (b) precipitation (shading; mm day−1) and 850 hPa streamfunction (contours; 10−6 m2 s−1); and (c) SLP anomalies (shading; hPa). The dotted areas denote that the shading is statistically significant at the 10% level when using the Student’s t test.
Citation: Journal of Climate 31, 9; 10.1175/JCLI-D-17-0440.1
To verify the relationship between ENSO phase transition and north China drought, the Niño-3.4 index evolution with individual drought events is examined as shown in Fig. 7a. All eight events are accompanied with negative Niño-3.4 below than −0.5°C in preceding winter, and seven cases (except 2001) show positive Niño-3.4 (exceeding 0.5°C) in the subsequent winter. Furthermore, six drought events (except 2006 and 2001) are preceded with a negative NPO in previous winter. To better illustrate the role of ENSO, lead–lag correlation coefficients between Niño-3.4 index and spring–summer mean north China precipitation are shown in Fig. 7b. The maximum positive correlation coefficient between them are observed in the winter before the drought years [D(−1)JF] exceeding the 5% significance level. Then the positive correlation coefficient is gradually replaced by a negative correlation that begins to pass the 5% significance test level in May–July (MJJ) and persists until the following winter (red line in Fig. 7b).
Potential relationship between ENSO and north China precipitation: (a) Niño-3.4 index from preceding winter to subsequent winter of individual drought years and (b) lead–lag correlation coefficient between Niño-3.4 index and north China precipitation averaged in March–August (red line), MAM (green line), and JJA (blue line). The gray lines denote the statistical threshold of the significance test at the 10% level.
Citation: Journal of Climate 31, 9; 10.1175/JCLI-D-17-0440.1
The lead–lag relationships of Niño-3.4 index with spring and summer north China precipitation are also examined separately as shown in Fig. 7b. Different from the total rainfall of spring and summer, the spring precipitation is positively correlated with Niño-3.4 index at a 3-month lag with the maximum correlation coefficient of 0.41, passing the threshold of the 5% significance level. Correlation falls below the 5% significance level during MAM, and no significant correlation is shown afterward (green line in Fig. 7b). In comparison, statistically significant negative correlation between Niño-3.4 index and summer precipitation is seen when the Niño-3.4 index leads by 2 months (−0.26) and reaches the maximum at 0 lag (−0.36), which persists until the subsequent winter. This lead–lag relation indicates that north China tends to be dry in the spring following a La Niña winter, while it may not last into summer without a concurrent development of El Niño in the subsequent seasons. And it is unnecessarily dry in the spring preceding an El Niño peak phase, although when summer precipitation is deficient.
To verify the importance of ENSO phase transition on the north China PSSD, a scatterplot between north China spring–summer rainfall and Niño-3.4 index in the preceding and following winter is presented in Fig. 8a. There are 11 years for 1961–2014 with negative Niño-3.4 in D(−1)JF and positive Niño-3.4 in DJF(+1) greater than 0.5°C. And eight of them (about three-quarters, except 1976, 1963, and 1975) are the north China PSSD years. We notice that the impact of ENSO phase transition on north China spring–summer rainfall is not symmetric, since extreme wet years in north China are not only seen in the year with El Niño transitioning to La Niña but also found in other years (such as 1979, 1990, and 1996). The scatterplot (Fig. 8b) shows that the spring precipitation deficiencies of the selected dry years are all below normal (besides 2014), although the precipitation anomaly magnitude is not very strong. There are 15 years with spring precipitation reduced by 20% or more for 1961–2014, and four-fifths of them follow the winters with negative Niño-3.4. It means that La Niña in the preceding winter favors less precipitation over north China in spring, causing drought onset in spring. However, the dry conditions over north China in spring are not always followed by less rainfall in summer. Seen from Fig. 8c, most of those years with a positive Niño-3.4 index in DJF(+1) witness continuous drying in summer. And seven-eighths of those years are the driest summers for 1961–2014. The top 15 driest summers are the same as the selected eight years based on total rainfall of spring and summer, further demonstrating the dominant role of summer rainfall in drought intensity. In summary, the north China drought tends to begin in the spring following a La Niña winter and to persist into summer when El Niño establishes afterward, thus causing persistent depleted water supplies to north China from spring to summer.
Scatterplots showing the connection of precipitation anomaly percentages (%) averaged over north China with ENSO. The x and y axes are the Niño-3.4 index (°C) in the preceding winter [D(−1)JF] and in the subsequent winter [DJF(+1)], respectively. (a) March–August, (b) MAM, and (c) JJA.
Citation: Journal of Climate 31, 9; 10.1175/JCLI-D-17-0440.1
b. Air–sea interaction when La Niña transitions to El Niño simulated by HadGEM2-ES
Because of the short observational record, only eight PSSD years of observations are selected in the study. To verify the impact of ENSO phase transition on the north China PSSD, a 574-yr control simulation from HadGEM2-ES is further examined. There are 33 years selected as years when La Niña transitions to El Niño in HadGEM2-ES, and 21 years (about 64%) witness the north China PSSD, 10% less than the observations (75%). The composite precipitation anomaly percentages in spring and summer for the years when La Niña transitions to El Niño are shown in Fig. 9. As in the observations, a prolonged deficient precipitation over north China is present from spring to summer, exceeding the 10% significance test level. The simulated maximum precipitation anomalies over north China are reduced by 10% in spring and 8% in summer, which are much weaker than the anomalies in observations. In addition, different from the observations, increased precipitation is simulated in summer over the northwestern corner of north China, which is insignificant at the 10% level. Interestingly, as in the observations, significant continuous below-normal precipitation anomalies over northwest China associated with north China is also captured by HadGEM2-ES.
The composite precipitation anomaly (%) for (a) spring and (b) summer composited for the years with La Niña transitioning to El Niño derived from the preindustrial control simulation of HadGEM2-ES. The dotted areas denote precipitation anomalies are statistically significant at the 10% level.
Citation: Journal of Climate 31, 9; 10.1175/JCLI-D-17-0440.1
The associated circulation anomalies from the HadGEM2-ES preindustrial control simulation are presented in Fig. 10. Generally speaking, the circulation anomalies associated with La Niña transitioning to El Niño simulated by HadGEM2-ES resemble observations but with weaker magnitude. A significant cold SSTA over the eastern tropical Pacific and a negative NPO are simulated in the preceding winter (Figs. 10a,d). Thus, there is a significant anticyclonic circulation over the high-latitude North Pacific and a cyclonic circulation anomaly over the midlatitude North Pacific (10°–40°N), with two cyclonic centers located at 20°N, 120°E and 30°N, 170°W. The former center is caused by the La Niña SSTA, while the latter is associated with the negative NPO phase. Those features are highly consistent with observations (Fig. 6).
The simulated evolution of SST (shading, top color bar; °C) and SLP (contours; contour interval 0.2 hPa) anomalies from (a) D(−1)JF, (b) MAM, and (c) JJA composited for the years with La Niña transitioning to El Niño simulated by the preindustrial control simulation of HadGEM2-ES. The dotted and hatched areas denote SST and SLP anomalies, respectively, are statistically significant at the 10% level. The negative SLP anomaly is shown as dashed contours in (a)–(c). (d)–(f) As in (a)–(c), but for the anomalies of geopotential height at 200 hPa (shading, bottom color bar; m) and winds at 850 hPa (vectors; m s−1), and the dotted areas denote geopotential height anomalies at 200 hPa are statistically significant at the 10% level.
Citation: Journal of Climate 31, 9; 10.1175/JCLI-D-17-0440.1
In spring (Figs. 10b,e), negative NPO phase persists, with a cyclonic circulation anomaly over midlatitude North Pacific and below-normal geopotential height at 200 hPa, which is statistically significant at the 10% level. Cold SSTAs over the eastern Pacific weaken and a triple SSTA pattern emerges, as in the observations, but with relatively weaker magnitude for the cold SSTA along 30°N. In summer (Figs. 10c,f), El Niño starts as a result of the SFM of NPO and sustains the cyclonic circulation anomaly over the midlatitude North Pacific. A low-level westerly anomaly is shown over the tropical Pacific Ocean, indicating weakened Walker circulation. As a result, the Indian subcontinent is dominated by an anomalous anticyclonic circulation (Fig. 10f), and Indian summer precipitation is suppressed (Fig. 9b), triggering the CGT with cooler tropospheric temperature (i.e., reduced 200-hPa geopotential height) over East Asia. The consistency between model simulations and observations further confirms the robust driving of the ENSO phase transition when accompanied by a negative NPO with the north China PSSD.
5. Summary and discussion
Drought over north China is prone to begin in spring and to persist until summer, showing the highest drought risk in these two seasons. Sustained precipitation deficiency from spring to summer dominates the whole annual water supply to north China, bringing great challenges to local agriculture growth and hydrological water management. In this study we explored the large-scale circulations responsible for north China drought prolonged from spring to summer and the associated air–sea interaction. The main finding is summarized in a conceptual plot shown in Fig. 11. This study identified two crucial circulation anomalies associated with PSSD over north China: an anomalous low-level cyclonic circulation spanning the midlatitude North Pacific (black vector in Fig. 11) and East Asian tropospheric temperature (TT) cooling along the upper-level westerly jet core (blue solid contour in Fig. 11). The low-level cyclonic circulation is related to a weaker NWPSH, which reduces warm and wet water transport from the adjacent ocean to north China. The TT cooling strengthens (weakens) the upper-level westerly winds south (north) of the jet core, thus leading to anomalous descent motion over north China.
Conceptual plot of air–sea interaction in (a) MAM and (b) JJA on sustaining the crucial circulations associated with PSSD over north China. Dashed and solid lines denote the low- and upper-level circulation anomalies, respectively, whereas shading shows SSTAs.
Citation: Journal of Climate 31, 9; 10.1175/JCLI-D-17-0440.1
This study finds that the two crucial circulations in spring are directly caused by a negative phase of NPO that is generated in the previous winter with a La Niña–like equatorial Pacific cooling and is able to trigger the onset of El Niño in summer via the so-called seasonal footprinting mechanism. The fluctuations of NPO in spring cool the SST along the Kuroshio Extension and warm the tropical Pacific through surface heat fluxes (Fig. 11a). Consequently, El Niño begins in summer with suppressed Indian monsoon precipitation that triggers CGT along the upper-level westerly jet to sustain the East Asian tropospheric cooling (Fig. 11b). Enhanced precipitation over the central equatorial Pacific in summer further prolongs the anomalous lower-level cyclonic circulation over the northwestern Pacific through the Gill–Matsuno response.
Such a combination of the two crucial circulation patterns tends to occur when La Niña develops into El Niño, which could provide favorable connections between the seasonal footprinting mechanism and El Niño. The observed relationship between ENSO phase transition and north China PSSD is also shown in the long preindustrial control simulation of HadGEM2-ES, confirming the results found in this study. A further investigation of the observations reveals that four-fifths of the springs following La Niña winter show less precipitation during spring over north China despite no El Niño occurring in the subsequent winter, and the spring drought does not always prolong into summer. The summer with a concurrent El Niño in the subsequent seasons tends to show prolonged drying since spring. About three-quarters of the years with a La Niña transition to El Niño are present with the north China PSSD in observations, and seven-eighths of those years are extreme PSSD years.
In this study, the low-level anomalous cyclonic circulation over the North Pacific emerges in the previous winter, which is one season earlier than the PSSD and persists until summer via air–sea interaction. Although peaking El Niño phase is behind the drought in spring and summer, the negative NPO generated in the winter with La Niña plays the key role in forcing the onset of El Niño in summer through SFM. It provides a good predictor for the onset of El Niño and thus prolonged the north China PSSD. Hence, ENSO phase transition combined with a negative NPO can be seen as one of the important precursors of PSSD over north China.
Finally, our study demonstrated that ENSO phase transition from La Niña to El Niño is one of the important precursors of the north China PSSD; this does not necessarily mean that all north China PSSD events are caused by such an ENSO phase transition. In the observations, seven out of eight events support the ENSO phase transition scenario, implying other mechanisms may exist behind the PSSD. Since the short observational record does not allow for a further analysis, we selected the top 15 north China PSSD events in the 574 model years and examined the associated SSTA. No similar SSTAs as in the observations is found in the model results, implying another mechanism dominated the PSSD in the model simulation for these 15 events. In the monsoon community, some other mechanisms are proposed to affect north China drought, including the Eurasian spring snow (Yim et al. 2010); North Pacific processes (Y. Zhang et al. 2017); Tibetan Plateau heating (Zhang et al. 2018); joint impacts of Pacific SSTAs, Arctic sea ice anomalies, and warming over the European continent and Caspian Sea (Wang and He 2015); and the synergistic effect of El Niño evolution and Eurasian spring snow-cover reduction (Wang et al. 2017). The sample size from the HadGEM2-ES is 10 times larger than observations, and thus the long-term simulation should include more mechanisms. We hypothesized that in the 574-yr-long model simulation, the top 15 PSSD events may be dominated by some of the above mechanisms rather than the ENSO phase transition mechanism proposed here. How different mechanisms modulate the north China PSSD events at different decades warrants further investigation.
Acknowledgments
This work was jointly supported by the National Natural Science Foundation of China under Grants 41675076 and 41661144009 and Program of International S&T Cooperation under Grant 2016YFE0102400. Met Office staff is supported by the U.K.–China Research and Innovation Partnership Fund through the Met Office Climate Science for Service Partnership (CSSP) China as part of the Newton Fund.
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