Introduction

Water is the most influential resource in ecosystems and societies in arid and semi-arid areas (Chen et al. 2013; Feng et al. 2016; Jaeger et al. 2017; Qi et al. 2017). In drylands, the spatio-temporal distribution and availability of water resources have been the major focuses of terrestrial ecosystem studies at local, regional and global levels (Groisman et al. 2017; Monier et al. 2017; Soja and Groisman 2018). In terrestrial ecosystems, hydrological cycling is largely motivated through energy transfer because of the close coupling of water and heat. Precipitation, via either rainfall or snow, is the largest flux term of water budget and can lead to changes in soil moisture and latent heat (LE), sensible heat (H), soil heat flux (G), and heat storage (S) by influencing plant transpiration and soil evaporation (Rodrigues et al. 2013). Knowledge about coupled ecosystem water cycling and energy balance can not only be applied to explain the integrated water-energy exchange between ground and atmosphere but also be used to address water scarcity and maintain ecosystem functions.

Rapid changes in land use and land cover and global warming have produced an increasing incidence and intensity of climatic extremes such as frequent and heavy precipitation events (Eade et al. 2012; IPCC 2013; Rey et al. 2017) and more intensive and prolonged drought. Changes in precipitation regimes contribute to higher uncertainty in ecosystem water-energy balance that requires further investigation of biomes and habitat. For instance, Minderlein and Menzel (2015) reported direct influences of precipitation on soil moisture, evapotranspiration (ET), and energy balance of shrub-grass ecosystems in the semi-arid areas of northern Mongolia. Nevertheless, there remains limited knowledge and consensus on the relationships between changes in precipitation and water-energy balance in grassland ecosystems.

Climate change, together with human disturbances, has had a large influence on ecological processes in the Eurasian Steppes, such as carbon assimilation and emission and energy and hydrological cycling. The meadow steppe appears to be the most sensitive ecosystem type in the grassland biome because of its relatively high biodiversity and high demands for water (Wang et al. 2007). Using a long-term research site of a meadow steppe in Northeast China, we explore the changes over time in water and energy fluxes in response to changes in climate, especially precipitation. Here, we focus particularly on the changes of major energy flux terms and the overall balances under different water availability. Based on 4 years of eddy covariance measurements, we first emphasize the distribution of energy and water fluxes and the possible ecological and climatological controls for a Songnen meadow steppe. Precipitation and its extremes, especially drought and transient waterlogging events, are examined for their roles in regulating the energy balance. Two major scientific questions are as follows: (1) What are the roles of shifting precipitation on water-energy balance of a meadow steppe? (2) What are the differences in how each energy flux term responds to spring/summer drought and transient waterlogging?

Materials and methods

Study site

This study was conducted at the Songnen Grassland Ecology Field Station of the Northeast Normal University (NENU) in Changling, Jilin, China (123° 30′ E, 44° 35′ N, 171 m a.s.l.). The site is representative of a temperate meadow steppe dominated by perennial grasses, including Leymus chinensis and Phragmitis communis. The study area has a temperate, semi-arid continental monsoonal climate, with a cold-dry spring and a warm-wet summer. The mean annual temperature is 5 °C, with the maximum and minimum temperatures being 39.2 °C and − 33.9 °C, respectively. The frost-free period is approximately 130 days to 165 days. The precipitation varied greatly within and between years, with an annual mean of 350 mm, 80% of which falls between June and August. The annual pan evaporation varies from 1200 to 1600 mm, which is approximately three to four times higher than the mean annual precipitation. The growing season runs from May through September. Main soil types in this area are alkaline soil, chernozem, and saline soil.

Flux measurements

To evaluate the energy and water exchanges between the land surface and the atmosphere, an open-path eddy-covariance (EC) flux tower was deployed in 2007 as part of the US-China Carbon Consortium (USCCC). A three-dimensional ultrasonic anemometer (CSAT3, Campbell Scientific Inc. (CSI)) and a CO2/H2O infrared gas analyzer (LI-7500, LI-COR) were mounted 2 m above the ground. Fluxes were calculated as the mean covariance of vertical wind speed fluctuation and the scalar fluctuation. Downward fluxes were indicated as negative and upward as positive. All raw data was sampled at 10 Hz and logged using a CR3000 data-logger (CSI). Footprint analysis using the flux source area model (FSAM) suggested a footprint in the prevalent wind direction (150–240°) extended to approximately 6.5 m to 77.8 m during unstable conditions (Monin-Obukhov length (L) < 0) and to about 11.3 m to 229.9 m during stable conditions (L > 0) at night, representing 90% of the total flux.

Meteorological measurements

Meteorological measurements were taken starting in May 2007. Net radiation (Rn) was measured using a four-component net radiometer (CNR-1, Kipp & Zonen) at 2 m above the ground. Photosynthetically active radiation (PAR), precipitation (P), wind speed (WS), and wind direction (WD) were measured by a quantum sensor (LI190SB, LI-COR), a tipping-bucket rain gauge (TE525MM, CSI), and a propeller anemometer (034B-L Met One Windset, CSI), respectively. Air temperature (Ta) and relative humidity (Rh) were measured at 2, 4, and 6 m with HMP45C probes (Vaisala). Soil temperature profiles (Ts) were measured at depths of 0.05, 0.1, and 0.3 m below the ground by thermistors (107-L, CSI). Soil heat flux (G) was measured at a depth of 0.05 m in three separate locations (HFT-3, REBS), and soil water content (SWC) was measured at a depth of 0.1 m by a time-domain reflectometry probe (CS616, CSI), which was buried horizontally in the soil. Additionally, soil water potential (SWP) was estimated using watermark probes at 0.1 and 0.3 m (257-L, CSI). All micrometeorological data were recorded and averaged or summed over a 30-min interval by a separate CR3000 data-logger.

Vegetation

Biometric measurements of aboveground biomass (AGB), vegetation height (Vt), and leaf area index (LAI) were conducted monthly during the growing season (May to September). Within a radius of 200 m around the EC tower, we placed 12 0.5 m2 quadrats for measuring Vt and LAI by a plant canopy analyzer (LAI-2000, LI-COR), while AGB was destructively sampled in 12 additional 0.5 m2 quadrats.

Data processing and gap filling

The EC flux data was processed using EdiRe. The planar fit rotation and Webb–Pearman–Leaning (WPL) density correction were applied to the flux calculations (Webb et al. 1980). Data from stable nocturnal periods were excluded, specifically when the friction velocity (u*) was < 0.15 m s−1. Anomalous or spurious data due to sensor malfunction, sensor maintenance, precipitation events, IRGA calibration, and power failure were also rejected. Consequently, 25.4%, 20.9%, 27.1%, and 30.6% of the data obtained during growing season in 2007, 2008, 2009, and 2010, respectively, were discarded. The introduced data gaps were filled with following the methods: for gaps of less than 2 h, linear interpolation was used by averaging the fluxes before and after the gaps; for larger data gaps, we used both the empirical relationships and look-up tables (Falge et al. 2001).

Calculation of soil heat storage

The Rn was partitioned into convective H for heating the atmosphere, LE for evaporating water from plants and the soil, G for heating the soil, and S for heat storage in the soil. Because the canopy and air heat storage are expected to be negligible in short canopies with minimal biomass, the canopy heat storage was not included in our energy balance algorithm:

$$ {R}_n=H+\mathrm{LE}+G+S $$
(1)

where soil heat storage (S) was calculated as follows:

$$ S={C}_s\frac{\Delta {T}_s}{\Delta t}d $$
(2)

where Ts is the average soil temperature (K) above heat flux plates, t is time (Δt = 1800 s), d is a depth between the heat flux plate and soil surface, and Cs is the soil heat capacity calculated from Shao et al. (2014):

$$ {C}_s={\rho}_b{C}_d+{\theta}_v{\rho}_w{C}_w $$
(3)

where ρb is soil bulk density, ρw is the density of water, Cd and Cw are the specific heat capacity of dry mineral soil (Cd = 890 J kg−1 K− 1) and soil water (Cw = 4190 J kg−1 K−1), respectively, and θv is volumetric SWC (%). The heat storage of air and organic matter in the soil were neglected due to their small quantities.

Energy parameters

The energy balance ratio (EBR) was calculated as a ratio of the turbulent fluxes (LE + H) and the radiative energy fluxes (Rn–G–S):

$$ \mathrm{EBR}\kern0.5em =\kern0.5em \frac{\sum \left( LE+H\right)}{\sum \left({R}_n-G-S\right)} $$
(4)

Canopy conductance (gc) was calculated using the inverted form of the Penman–Monteith equation (Monteith and Unsworth 1990):

$$ {g}_c=\frac{1}{\frac{\uprho {C}_p}{\upgamma}\frac{D}{LE}+\left(\frac{\Delta}{\upgamma}\frac{H}{LE}-1\right)/{g}_a} $$
(5)

where ρ is air density, Cp is the specific heat of the air, Δ is the change of saturation vapor pressure with temperature, γ is the psychrometric constant, D is the vapor press deficit of air, and ga is the aerodynamic conductance of the air layer between the canopy and the flux measurement height. ga was calculated flowing Monteith and Unsworth (1990) as follows:

$$ {g}_a=\frac{1}{\frac{u}{u^{\ast 2}}+6.2{u}^{\ast -0.67}} $$
(6)

where u is wind speed and u* is friction velocity.

To calculate the Priestley-Taylor coefficient (LE/LEeq), the equilibrium LE flux (LEeq) was determined using (Priestley and Taylor 1972):

$$ {\mathrm{LE}}_{eq}\kern0.5em =\kern0.5em \frac{\Delta \left({R}_n-G-S\right)}{\Delta +\upgamma} $$
(7)

The LEeq is dependent on the Rn and temperature. Lower and higher values indicate evaporation rates that are lower and higher than the equilibrium rate, respectively (Wilson et al. 2002b). The sensitivity of evapotranspiration to stomatal control and the degree of aerodynamic coupling between vegetation and the atmosphere was expressed by the decoupling factor (Ω), obtained by the equation (Jarvis and McNaughton 1986; Runkle et al. 2014):

$$ \Omega =\frac{\Delta +\upgamma}{\Delta +\upgamma \left(1+\frac{{\mathrm{g}}_{\mathrm{a}}}{{\mathrm{g}}_{\mathrm{c}}}\right)} $$
(8)

The value of Ω ranges from 0 to 1. When it approaches 0, the vegetation and the atmosphere are fully aerodynamically coupled and the evapotranspiration proceeds at rates imposed by vapor pressure deficit; when it approaches 1, the vegetation and the atmosphere are completely aerodynamically decoupled and the evapotranspiration is controlled by Rn (Goldberg and Bernhofer 2001).

Results

Microclimate

The mean Ta in 2007–2010 was 8.23, 6.83, 5.64, and 7.91 °C, respectively (Table 1). For the growing season, the mean Ta in 2007 (21.26 °C) and 2010 (20.67 °C) were higher than that in 2008 (19.76 °C) and 2009 (19.88 °C). The maximum Ta occurred in July of 2007 and 2008, August of 2009, and June of 2010, while the minimum Ta occurred in January in all 4 years. For the non-growing seasons (October–April) of 2007–2008, 2008–2009, 2009–2010, and 2010–2011, the mean Ta was − 2.82, − 3.38, − 6.69, and − 5.62 °C, respectively. The Ts closely followed the changes of Ta (Fig. 1a). However, there was no significant difference between Ts (21.15 °C) and Ta (21.26 °C) in the growing season of 2007.

Table 1 Daily means of microclimatic variables and energy balance components during the growing season and non-growing season of 2007–2010 in Songnen meadow steppe
Full size table
Fig. 1
figure 1

Seasonal changes of major meteorological variables and energy fluxes during 2007–2010 in the meadow steppe: a daily mean air (Ta) and soil (Ts) temperatures, b net radiation (Rn), c latent heat (LE) and sensible heat (H), d soil heat flux (G), and e daily mean soil volumetric water content (SWC) and daily sum of rainfall (P)

Full size image

Precipitation amounts and distribution were significantly different during the 4 years. Annual precipitation from 2007 to 2010 was 207.9, 384.2, 281.5, and 265.6 mm, respectively. These years can be classified as dry (2007, 2009, and 2010) or wet years (2008) based on to the precipitation trending in the Songnen meadow steppe for the past 50 years (Fig. 2). The driest period was from DOY181 to DOY279 in 2007, with the first rain of > 3 mm. Precipitation or snowfall throughout the initial 6 months was minimal. In 2008, the meadow steppe received 85% more precipitation and experienced 18 more single precipitation pulses than in 2007. The first precipitation in 2008 (March 22nd) was 3 months earlier than in 2007, and the rainy season therein lasted longer and ended on November 15th. In 2009 and 2010, precipitation started in mid-April, and it ended in mid-October and mid-November, respectively. Total precipitation of March–May in dry year 2007 accounted for only 8.8% and 5.1% of that in 2009 and 2010, respectively, indicating that the meadow steppe experienced a serious spring drought in 2007. On the other hand, June–August precipitation was low in 2010 (90.7 mm) compared to that in 2007 (170 mm) and 2008 (231.9 mm). In the meadow steppe, transient waterlogging occurs usually after a single precipitation of > 30 mm, and the waterlogging condition often lasts for 2–3 days. Six transient waterlogging events were detected during the study period: DOY181, 190, 212 in 2007; DOY187, 214 in 2008; DOY125 in 2010.

Seasonal variation of SWC closely followed the changes in precipitation. The temporal change of SWC was clearly asymmetrical within any of the 4 years (Fig. 1e). A long dry period in 2007 resulted in severe water deficit in the spring and early summer when the SWC fell below 0.2 m3 m−3 in the upper soil layer. In contrast, frequent precipitation events during the summer of 2008 resulted in a consistently higher SWC. The highest SWC occurred in the early spring of 2009 and 2010, which declined sharply afterwards. SWC decreased to its minimum before July in 2010, followed by a summer drought. The SWC recovered in June of 2009 after a series of continuous precipitation events. A large precipitation event combined with several isolated precipitation events led to a temporary SWC peak during the 2010 summer compared to the previous summer. Interestingly, SWC at 10 cm depth did not drop below 0.4 m3 m−3 during the water-stressed periods in 2007 and 2009 (Fig. 2).

Fig. 2
figure 2

Long-term change in annual precipitation and its trend in Songnen meadow steppe during the past 49 years (1967 to 2015)

Full size image

A unimodal pattern of growing season AGB and LAI showed an overall higher AGB and earlier peak in 2008 than in the dry years (Fig. 3). Peak value of AGB in August 2007 was ~ 68% of that in July and August of 2009 and 2010. LAI was highest in 2008 and lowest in 2007.

Fig. 3
figure 3

Seasonal change of leaf area index (LAI) and aboveground biomass (AGB) with month in the meadow steppe during the four growing seasons of 2007–2010

Full size image

Energy balance closure

The energy balance enclosures using half-hourly and daily values produced reasonable results for all 4 years (Table 2). The slope of the enclosure was < 1 for all years, with an average value (± SD) of 0.75 (± 0.06), intercept of 10.48–21.17 W m−2, and R2 value of 0.84–0.98. For the growing season, the slope values were similar in 2008, 2009, and 2010. The average slope value increased to 0.85 (± 0.09), ranging from 0.74 to 0.93, when using the daily sums instead of the half-hourly data to calculate the regression coefficients of [LE + H] against [Rn–G–S].

Table 2 Energy enclosure based on half-hourly and daily sums of the growing seasons during 2007–2010
Full size table

The EBR provides an overall evaluation of energy closure at long temporal scale by averaging over random errors in the flux measurement. The mean values of EBR for the four growing seasons were 0.94 (2007), 0.83 (2008), 0.85 (2009), and 0.92 (2010), respectively, and increased to 1.00, 0.97, 0.91, and 0.93 at annual scale for the four consecutive years (Table 1).

Intra- and interannual variations in energy fluxes

The seasonal changes of daily Rn, LE, H, and G were similar (Fig. 1b–d). Rn showed maximum values of 15.67–20.86 MJ m−2 day−1 in July and the minimum values of from − 1.03–3.56 MJ m−2 day−1 in December among the 4 years. Seasonal changes in LE matched well with those of Rn, which gradually increased from May, peaked in July–August, and gradually decreased until soil was frozen in late October through February. During the spring drought in 2007, LE in March–May (3.51 MJ m−2 day−1) was 70% of the normal values (4.65–5.08 MJ m−2 day−1). LE during the peak growing season of 2010 (July–August) was 7.31 MJ m−2 day−1 and remained lower during 2007–2009 (~ 6.4 MJ m−2 day−1) despite the summer drought of 2010. Surprisingly, H showed a bimodal seasonal change, which increased at the beginning of a year, peaked in May, and began to decrease in June even though Rn continued its high values. Interestingly, H started to increase again in August and reached the second peak in September. As expected, G values were low, especially during the wintertime. At the beginning of March, G switched from negative to positive and continued to increase, reaching a maximum in May for 2007 (3.39 MJ m−2 day−1), 2009 (2.16 MJ m−2 day−1), and 2010 (2.15 MJ m−2 day−1), but in June of 2008 (2.10 MJ m−2 day−1).

Monthly LE value was higher than H value during the growing season in 2008 (Fig. 4). In dry years, the duration of higher LE over H appeared short (2 months) in 2007 (spring drought), mediate (3 months) in 2009, and long (4 months) in 2010, with different starting days in July, June, and May, respectively. For accumulated G, however, a positive G (22.72 MJ m−2 month−1) was observed in March 2008, which was − 2.00 MJ m−2 month−1 in 2009 and − 7.84 MJ m−2 month−1 in 2010. More importantly, G in May (67.91 MJ m−2 month−1) and June (132.91 MJ m−2 month−1) of 2007 was about twice the values of 2008 (35.21/71.93 MJ m−2 month−1), 2009 (39.85/63.93 MJ m−2 month−1), and 2010 (41.44/76.21 MJ m−2 month−1).

Fig. 4
figure 4

The monthly change of latent heat (LE), sensible heat (H), and soil heat flux (G) during 2007–2010 in the Songnen meadow steppe in northeastern China

Full size image

Sensible heat (H) was a major factor of Rn at annual scale, but LE was more influential during the growing seasons, with the mean LE/Rn ratio 0.46 to 0.53. LE/Rn values remained low during the non-growing season, with mean values in 0.12, 0.18, and 0.19 for 2007, 2008, and 2009, respectively (Table 1). The peak LE/Rn in 2007 arrived half a month later than that in 2010 (Fig. 5). Throughout the spring drought period in 2007, LE/Rn remained low until late summer. In contrast, LE/Rn in 2010 remained relatively high from the beginning of the growing season, when soil moisture was high due to the large precipitation input. The seasonal changes of H/Rn seemed to be a mirror image of LE/Rn in all 4 years. The lowest value of H/Rn was observed during the growing season when LE/Rn was the highest, while H/Rn began increasing in late September. The annual mean G/Rn was higher in 2007 (0.171) than in other years (0.0045–0.055) (Table 1). Although there was a higher G/Rn at the beginning of growing season (May–June) in 2007, no significant difference was found among LEs of the 4 years despite a severe summer drought in July–August of 2010.

Fig. 5
figure 5

Growing season fractions of latent heat, sensible heat, and soil heat flux to net radiation (LE/Rn, H/Rn, G/Rn, respectively) during 2007–2010

Full size image

Biophysical controls of energy partitioning

The Bowen ratio (β, H/LE) during the growing season varied from 0.21 to 4.63 (Fig. 6b). Due to a faster increase in H than that in LE during the spring, midday means of β over the growing season exhibited a concave shape, with β being > 1 during spring and autumn, but ~ 1 in summer. Over the growing season, midday means of β amounted to 1.39, 0.95, 1.29, and 1.04 for 2007–2010, respectively. The low LE and high H during the non-growing season resulted in a dramatic increase in the mean daytime β. Mean daily β varied from 0.02 to 12.85 over the 4-year study period, with the largest value found in the non-growing season of 2007 (DOY310), and the lowest in the growing season of 2010 (DOY233). The average of daily β during the spring drought period was 1.97, which was higher than that in the summer drought period (1.17). The gc value ranged from 0.02 to 19.67 mm s−1 during the growing season, with a peak in July of 2008, 2009, and 2010, but not in 2007, when the maximum gc was delayed to August (Fig. 6c). In 2007, LAI maintained a higher level after July while gc declined during the same time. During the growing season, the Ω value amounted a peak of 0.73 (DOY229), 0.81 (DOY214), 0.75 (DOY204), and 0.56 (DOY213), respectively, for the 4 years. Its seasonal changes appeared less pronounced compared to other measures of the fluxes, with the change in 2007 very different from that in other 3 years (Fig. 6).

Fig. 6
figure 6

The seasonal change of 5-day average midday LE/LEeq, Bowen ration (β), canopy conductance (gc), and decoupling coefficient (Ω) in four study years in the meadow steppe. Midday means were calculated for 10:00–15:00 h

Full size image

Discussion

Impact of precipitation intensity on water and energy balance

In this study, P and ET appeared to be more or less balanced in the wet year. Differing from previous reports on P vs ET in neighboring typical steppes on the Mongolia Plateau (Chen et al. 2009; Foken et al. 2006), the meadow steppe in our study showed an overall higher cumulative ET than cumulative P in dry years (2007, 2009, and 2010), suggesting there were other sources of water to enrich the soil (Fig. 7). The mean annual ET (409.5 mm) (Table 1) in the meadow steppe lies in the higher end of ET values (163~481 mm) for grasslands worldwide (Aires et al. 2008; Krishnan et al. 2012; Li et al. 2007; Ryu et al. 2008; Zha et al. 2010). However, ET (334.5 mm) in 2007—a spring drought—fell below the mean value. The ratio of ET/P varied between 1.05 and 1.71 for all study years (Table 1), and was higher than those in the Mediterranean grasslands (0.4–0.87) (Aires et al. 2008; Ryu et al. 2008), American semi-arid prairies (0.72–1.20) (Krishnan et al. 2012), northern temperate grasslands (0.98) (Hao et al. 2007; Wever et al. 2002), and alpine Kobresia meadows (0.60) (Li et al. 2013), but similar to those in the alpine wetland meadows (1.27) (Hu et al. 2009).

Fig. 7
figure 7

The annual and monthly accumulation of precipitation (P) (gray) and evapotranspiration (ET) (black) for the Songnen meadow steppe during 2007–2010

Full size image

Intensity of precipitation events had a crucial effect on water-energy balance and ecosystem recovery after drought (Cabral et al. 2010; Heisler-White et al. 2008; Knapp et al. 2008). Differences in seasonal ET/P in the meadow steppe suggested that non-growing season P from snow would supplement water resources required for the following growing season. The relative contribution of temporally isolated non-growing season precipitation events to water availability in the next growing season is considered a critical determinant of ecological processes (Eamus et al. 2001; Ma et al. 2013). Previous studies have indicated that freezing and thawing have important influences on ET and surface energy fluxes (Chen et al. 2011; Gu et al. 2005; Yao et al. 2008), by increasing SWC and ET from melted snow and frozen soil (Zhang et al. 2014; Zhang et al. 2003). Although cumulative P was lower than ET at annual scale, the monthly P could sometimes be higher than ET in the meadow steppe during the growing season. For example, heavy and frequent precipitation events in July of 2008 brought ~ 132.66 mm precipitation and 90.8 mm ET at our site, yielding a peak soil water content of 0.91. Similarly, although the maximum soil water content was equivalent to that of 2008, P was 86.5% lower and ET was 13.8% higher following the consecutive precipitation events in July 2009 compared to those in July 2008 (Figs. 1 and 7), suggesting a proportion of precipitation infiltrated to recharge groundwater in 2008. Heavy precipitation events in May 2010 (P = 105.1 mm, ET = 58.4 mm) alleviated summer drought in the growing season through water infiltration into deep soil under similar maximum soil water content (0.89 mm) in May 2009 (P = 14.5 mm, ET = 45.6 mm). Therefore, AGB between 2009 and the summer drought year of 2010 remained similar, likely because both the low infiltration rate in the alkali-saline soil of the meadow steppe and the limited precipitation input from short and sporadic showers could substantially reduce soil moisture. However, low ET/P in July 2007 (0.74), in July (0.68), and August (0.68) of 2008 and in May 2010 (0.56) (Fig. 7) indicated that infiltration might have happened after occasional and large precipitation events (Knapp et al. 2008; Thomey et al. 2011).

In our study, LE was apparently influenced by the process of temporary waterlogging caused by precipitation > 30 mm in a day (Fig. 8). LE increased because of low vegetation cover and high groundwater level at the beginning of the growing season, which is consistent with the findings of spring drought influence on evapotranspiration and water use efficiency in the same steppe (Dong et al. 2011). However, LE declined in response to the 3-day waterlogging at the growing season peak of 2007 and 2008, which could possibly be attributed to stomatal closure and reduction in both leaf photosynthesis and root respiration. Clearly, the severity of waterlogging effects on LE depend on the developmental stage of the grass.

Fig. 8
figure 8

The magnitudes of latent heat (LE) before, during (3-day), and after the transient waterlogging during 2007–2010

Full size image

Influence of drought on energy balance and partitioning

The Songnen meadow steppe exhibited the great intra- and interannual variability of energy partitioning into H and LE, which also differed greatly between the dry and wet years. The meadow steppe was LE-dominated when soil water content was > 0.3 m3 m−3, but switched to H-dominated status when soil water content fell below 0.3 m3 m−3 (Fig. 9). The timing of the shift from an H-dominated to a LE-dominated energy balance depended on the magnitude of the first precipitation event (> 30 mm day−1) or accumulated precipitation within five consecutive days over 30 mm, which enabled grass seed germination with SWC exceeding 0.3 m3 m−3. Similarly, the switch from a LE-dominated to an H-dominated ecosystem happened in concert with the start of biomass senescence when SWC fell below 0.3 m3 m−3. However, LE was dominant in energy balance even though SWC dropped to 0.3 m3 m−3 in the growing season of 2010 (since DOY200), which likely was due to the lag effect of a large precipitation event at early stage. In 2008–2010, the change in energy partitioning occurred in May–June. However, the spring drought led to a shift from H to LE occurring in mid-July in 2007. The increase in SWC and vegetation growth during the growing season altered solar heating of the surface via a decrease in albedo that likely was driven by the dark soil and vegetation surfaces after the precipitation events (Krishnan et al. 2012; Thompson et al. 2004).

Fig. 9
figure 9

Daily mean soil volumetric water content (SWC) and daily mean Bowen ration (β) in the meadow steppe during the four growing seasons of 2007–2010

Full size image

Interannual climate, especially the occurrence, duration, and extent of drought, can significantly reduce LE in this site. On the one hand, limited precipitation could cause a decrease in soil moisture and limit soil water recharge, thus decreasing LE. On the other hand, droughts with high temperature and radiation would result in a reduction in transpiration with lower stomatal conductance of foliage and canopy (Chen et al. 2009). For instance, spring drought in 2007 resulted in a higher LE over H in the beginning of growing season (May–June), but the summer drought of 2010 did not change the interannual relation of higher LE over H in June–August. The growth of plants was triggered by the monsoon precipitation for all grasses. Delayed and low LE in spring drought year appeared to be a direct consequence of slow grass germination and late starting of transpiration. The main reason for higher LE during summer drought meanwhile might be related to the soil water availability. Several large precipitation pulses at the beginning of the growing season could have replenished the soil profile to allow grass to reach a biomass peak despite the summer drought in 2010. The results suggested that LE was more sensitive to spring drought than to summer drought, and there was hardly a negative consequence of summer drought to LE in the meadow steppe.

Impacts of biotic and climate on LE were further examined for other flux terms and key parameters, including LE/LEeq, gc, and Ω. Evapotranspiration (i.e., latent heat) is mainly composed of plant transpiration (Tr), soil evaporation (Es), and the negligible proportion of canopy intercepted precipitation (I) in grasslands that can be inferred from LE/LEeq, gc, and Ω. LE/LEeq is conventionally used to estimate daily evapotranspiration (Wilson et al. 2002a). LE/LEeq of spring drought (0.53) and summer drought (0.61) in this study were typical events for the semi-arid grassland LE/LEeq (0.6–0.7) (Meyers 2001), indicating that the restriction of the water supply for LE over the meadow steppe was less compared to other grasslands. Decreased gc is usually attributed to low LAI during the growing season (Baldocchi et al. 2004) and decreasing stem-root hydraulic conductivity under conditions of low soil water content (Rodrigues et al. 2013). According to Dong et al. (2011), Es was significantly higher than Tr under spring drought conditions, during which Es released more than 85% of available soil water in May 2007. During the spring drought period, the low gc did not reflect the low transpiration rate. On the contrary, higher gc during summer drought may have been influenced by high LAI in 2007. Ω evaluated the response of LE to air humidity and net radiation changes (Jarvis and Mcnaughton 1986). Ω of growing seasons was higher in summer drought (0.34) than in spring drought (0.24), suggesting that LE was highly sensitive to decreasing air humidity. Moreover, similar Ω (0.33) in summer drought years and wet years further proved the significant effect of spring drought in increasing sensitivity of LE to low air humidity. In summary, although LE was not limited by water stress in the meadow steppe (4.58 MJ m−2 day−1 in spring drought and 5.43 MJ m−2 day−1 in summer drought), spring drought had an even stronger effect on LE than summer drought did.

The proportion of individual surface energy balance component is a result of complex, long-term interactions between biogeochemical cycling, disturbance and climate, as well as short-term interactions of plant physiology and the dynamics of the atmospheric boundary layer (Aires et al. 2008; Baldocchi et al. 2004; Shao et al. 2012; Sun et al. 2010). The annual values of LE/Rn in our study (0.28–0.40) are comparable to the 0.31–0.35 that were reported for an annual grassland (Baldocchi et al. 2004), but lower than those in a tallgrass prairie (0.48–0.58) (Burba and Verma 2005) and in a Mediterranean C4 grassland (0.37–0.55) (Jongen et al. 2011). For the growing season, LE/Rn (0.49) was similar to that in an alpine meadow (Li et al. 2013), and higher than in temperate semi-arid grasslands in North America (Krishnan et al. 2012) and C3/C4 grassland in Southern Portugal (Aires et al. 2008), but less than the values reported in the Mongolian steppe (Li et al. 2007). When compared among the 4 years, the spring drought in 2007 caused the highest H/Rn and G/Rn but the lowest LE/Rn. The short-term water deficit at the initial stage of growing season could be partially responsible for the low LAI and AGB values. The cumulative available energy was 639.8 MJ m−2 in the extended drought period of 2010 (July–August), 6.6% lower than that of the same month in 2007. However, there was no obvious decrease in LE/Rn (Fig. 5), suggesting that the magnitude of LE was not constrained by the availability of energy but rather by the vegetation cover at dry period. As a result of earlier water inputs from precipitation, grass germinated earlier in the season, and the growing period was lengthened to maintain a stable and high AGB.

Soil heat flux (G) played an important role in energy partitioning at our site, which is similar to previous findings at other typical steppes (e.g., Shao et al. 2008). Soil heat flux tended to increase with Rn and H. A low canopy cover in the spring drought year jointly led to the highest G values (0.356 MJ m−2 day−1). Other than for the spring drought year, G/Rn was < 0.1 at the meadow steppe, which was lower than those reported in other grasslands (Aires et al. 2008) but similar to those at the Mongolian steppe and temperate semi-arid prairies (Krishnan et al. 2012; Li et al. 2007). More importantly, soil heat flux had wider diel variations in the dry year of 2007 than in other years (Fig. 10). The large seasonal fluctuations also reflected the seasonal variation in soil thermal conductivity, which were regulated by precipitation, soil moisture, and seasonal changes in vegetation coverage that directly determine value Rn (Heusinkveld et al. 2004; Rodrigues et al. 2013). A higher G was expected in a wet year (2008) because soil thermal conductivity increased with soil moisture (Hillel et al. 2005). However, the shading effect of vegetation cover likely offset the positive effect of soil thermal conductivity and resulted in lower G (0.69 MJ m−2 day−1). Finally, a higher S was found in spring drought year. A critical take-home message is that, if S is omitted from our analysis, G value at the topsoil may underestimate the surface heat flux by > 50% at our site (Ochsner et al. 2007; Shao et al. 2008).

Fig. 10
figure 10

Diurnal change of half-hourly soil heat flux (G) (closed circle) and soil heat storage (S) (open circle) during the growing seasons of 2007–2010

Full size image

The lack of energy closure is a common feature in most eddy flux sites. According to the FLUXNET study (Wilson et al. 2002a), the EBR for all vegetation types of site-years ranged from 0.53 to 0.99, with a mean of 0.79 ± 0.01. For the ChinaFLUX sites, the annual EBR ranged from 0.58 to 1.00, with a mean of 0.83 (Chen et al. 2011). In this study, the mean energy imbalance (0.89 ± 0.05) at meadow steppe is generally within, or higher than, the acceptable range reported within the FLUXNET and ChinaFLUX sites, which was higher compared with typical steppe 0.83 ± 0.07 (Chen et al. 2009).

Conclusions

Changes in precipitation resulted in frequent drought and waterlogging in meadow steppe of northeastern China. We measured the energy and water fluxes over a meadow steppe in Songnen Plain over a 4-year study period and found that spring and summer droughts resulted in a rapid shift in energy partitioning due to the change of water availability. However, spring drought produced more severe effects than summer drought. The rise of the groundwater level in our study region usually followed a large precipitation event or a series of continuous precipitation pulses that have profound influence on ecosystem water balance. Waterlogging influenced LE by enhancing its values during and after the waterlogged periods at the beginning of growing season in a dry year, but lowering them after the waterlogged periods in growing season. Disentangling the roles of different water table heights in controlling the energy balance, particularly LE remains a challenging research topic and should be prioritized in future studies.