Soil evaporation is a key process in the water cycle and can be conveniently quantified using
Terrestrial ecosystems receive water from precipitation and subsequently release all or part of the water to the atmosphere through evapotranspiration. The evapotranspiration process consumes approximately 25 % of the incoming solar energy (Trenberth et al., 2009) and can be divided into two components, namely transpiration from plant leaves and evaporation from the soil surface. Soil evaporation varies from 10 % to 60 % of the total precipitation (Good et al., 2015; Oki and Kanae, 2006). Precise estimation of soil evaporative water loss relative to precipitation is critical for improving our knowledge of water budgets, plant water use efficiency, global ecosystem productivity, allocation of increasingly scarce water resources, and calibrating hydrological and climatic models (Kool et al., 2014; Oki and Kanae, 2006; Or et al., 2013; Or and Lehmann, 2019; Wang et al., 2014).
Water loss from soil progresses with air invasion into the soil in the order of large to small pores (Aminzadeh and Or, 2014; Lehmann and Or, 2009; Or et al., 2013). Soil pores can be divided into large, medium, and small pores. There is a minimum amount of small pore water at which liquid water in soil is still continuous or connected and below which liquid water is hydraulically disconnected, and vapor transport is the only way to further reduce water in soil. This water content is called the residual water content in the soil characteristic curve (Van Genuchten, 1980; Zhang et al., 2015). When large soil pores are filled with water, water in small pores does not participate in evaporation (Or and Lehmann, 2019; Zhang et al., 2015). Therefore, soil evaporation can be divided into three stages (Hillel, 1998; Or et al., 2013). Stage I is the evaporation front in the surface soil, and water in large and medium pores participates in evaporation, but larger pores are the primary contributors. With the progressive reduction of water in the larger pores, the evaporation rate gradually decreases. Stage II is the evaporation front still in the surface soil, but larger pores are filled with air, with water residing in the medium soil pores in the surface soil evaporates, and deep larger soil pores recharge the surface medium pores by capillary pull (Or and Lehmann, 2019), and the evaporation rate remains constant. Stage III is when the hydraulic connectivity between the surface medium pores and deep large pores breaks, such that the evaporation front recedes into the subsurface soil. Water in the surface small pores and water in medium pores on the evaporation front evaporates. The evaporation rate decreases to a low value (Or et al., 2013).
Furthermore, water in small pores and large pores may differ in isotopic compositions. As is well-known, pre-event soil water occupies the smallest pores. Depending on the rainfall amount and intensity, an event water may have three pathways. First, a subsequent small event water fills the empty small soil pores. Second, event water with small rates, but long duration, may also displace the pre-existing, saturated smaller pores with slow flow velocity (Beven and Germann, 1982; Brooks et al., 2010; Klaus et al., 2013; Sklash et al., 1996); in cases where the water flows into a relatively impermeable layer, the pre-event water in smaller pores may be forced into large pores due to the underlining hydraulic barriers (Si et al., 2017). Third, when the event water is large and intense, the event water preferentially enters large pores, bypassing the saturated small pores with large flow velocity (Beven and Germann, 1982; Booltink and Bouma, 1991; Kumar et al., 1997; Levy and Germann, 1988; Radolinski et al., 2021; Sprenger and Allen, 2020). Because the exchange rate between these two flow domains is small (Šimůnek and van Genuchten, 2008), small pores will lock the signature of first filling water. As the flow velocity is determined by the soil pore size, larger pores have greater hydraulic conductivity, and consequently, water residing in larger pores flows faster and, thus, drains first. Conversely, water residing in small pores drains last (Gerke and Van Genuchten, 1993; Phillips, 2010; Van Genuchten, 1980). Therefore, soil water in smaller pores has a longer residence time or memory (Sprenger et al., 2019b), while water in large pores generally have a short memory. This differing memory between large pore and smaller pores, due to the sequence of water infiltration and drainage, could introduce variability in the isotopic composition between soil pore spaces.
Additionally, due to seasonal, temperature, and amount effects of local precipitation events, there is strong temporal variation in the isotopic composition of precipitation (Kendall and McDonnell, 2012). As a result, precipitation events, differing in isotopic compositions, could recharge different soil pores, which may yield isotopic heterogeneities in soil pore spaces (Brooks et al., 2010; Goldsmith et al., 2012; Good et al., 2015). Isotopically, small-pore water may be similar to old precipitation, with large-pore water resembling new precipitation (Sprenger et al., 2019a, b).
The isotopic variations in the soil pore space could also result from mineral–water interaction, soil particle surface adsorption, and soil tension (Gaj et al., 2017a; Gaj and McDonnell, 2019; Oerter et al., 2014; Orlowski and Breuer, 2020; Thielemann et al., 2019).
Despite the recent progress in understanding evaporation processes and isotope partitioning in soil pore space, the latter, to the best of our knowledge, is not considered in the calculation of soil evaporative water loss in terms of the isotope-based method. The isotopic composition of bulk soil water, which is extracted by cryogenic vacuum distillation, containing all pore water, is still routinely used in evaporation calculations using the Craig–Gordon model (Allison and Barnes, 1983; Dubbert et al., 2013; Good et al., 2014; Robertson and Gazis, 2006; Sprenger et al., 2017). This might bias the evaporation estimates because of isotopic variation in pore space and the preference for larger-pore water by evaporation.
Therefore, we hypothesize that the isotopic composition in evaporating water (EW) is similar to that of water in larger pores but differs from that in bulk surface soil water (BW); thus, evaporative water loss based on isotope values in BW will be biased. The objectives of this study were to verify (1) whether isotopic compositions differ between EW and BW and (2) if the isotopic composition difference substantially biases the calculated evaporative water loss. This study may help improve our understanding of soil evaporation and ecohydrological processes.
The field experiment was conducted from June to September 2016 at
Huangjiabao village (34
A summer maize field (35 m long and 21 m wide) was selected for this study.
On 18 June 2016, maize seeds were sown in alternating row spaces of 70 cm
and 40 cm, with 30 cm seed intervals in each row. Seeds were planted at a
depth of 5 cm beneath the soil surface using a hole-sowing machine. On
26 August 2016, the field was irrigated with 30 mm water (
A randomized replication design was used to collect samples. To determine
the water isotopic composition in EW from the condensation water of the
evaporation vapor, we randomly selected three rectangular plots (40 cm long
and 30 cm wide) in the field. A channel of 3 cm deep was dug around the edge of the plot (Fig. 1). Subsequently, a piece of plastic film without holes (approximately 0.2 m
Photograph of a new plastic film cover and condensation water collection using a syringe
In addition, BW was obtained from 0–5 cm surface soil water (Wen et al.,
2016). The soil samples were collected using a soil auger every 3 d with
three replicates, and each was mixed well and separated into two subsamples, i.e., one for determining the soil gravimetric water content and the other for water
stable isotope analysis. The subsample for soil gravimetric water content
was stored in an aluminum box and oven-dried for 24 h at 105
A cryogenic vacuum distillation system (Li-2000; LICA United Technology
Limited, Beijing, China) with a pressure of approximately 0.2 Pa and a
heating temperature of 95
In total, five deep soil profiles were collected on 17 July 2016 (pre-precipitation), 3 August 2016 (10 d after precipitation; 10 DAP), 17 August 2016 (24 DAP), 1 September 2016 (6 d after irrigation; 6 DAI), and 16 September 2016 (21 DAI), with increments of 0–5, 5–10, 10–20, 20–30, 30–40, and 40–60 cm. These soil samples were used to measure soil texture (Dane and Topp, 2020), soil water content, and soil water isotopic composition. Furthermore, the lc excess of the soil water before the
Precipitation was collected during the entire growing season using three
rainfall collectors (Wang et al., 2010) in the experimental field. The amount of rainfall was determined by weighing using a balance. Subsequently, subsamples of these rainfall samples were transferred to 15 mL glass vials,
sealed immediately with Parafilm®, and placed in a refrigerator at 4
Hourly air and 0–5 cm soil temperature under the newly covered plastic film from 10 to 28 September 2016, were measured using an E-type thermocouple (Omega Engineering Inc., Norwalk, CT, USA) controlled by a CR1000 data logger (Campbell Scientific, Inc., Logan, UT, USA). The 0–5 cm field soil temperature was measured during the whole field season using an iButton® device (DS1921G; Maxim Integrated, San Jose, CA, USA) at a frequency of 1 h. The 0–5 cm soil temperature and air temperature under the plastic film are required to calculate the evaporation ratios, but these measurements were not available before 10 September 2016. To obtain these temperature values, a regression equation was established between the measured 0–5 cm soil temperature values under the newly covered plastic film and those without plastic film covering from 10 to 28 September 2016. We then used the equation to estimate 0–5 cm soil temperature under the newly covered plastic film before 10 September 2016, based on the iButton®-measured temperature of the 0–5 cm soil without the plastic film covering in the same period. Subsequently, another regression equation was obtained between air temperature and 0–5 cm soil temperature from 10 to 28 September 2016, both of which were under the newly covered plastic film. Then the air temperature under the newly covered plastic film before 10 September 2016 was estimated from the estimated 0–5 cm soil temperature under the newly covered plastic film. The regression equations are presented in the Supplement. Moreover, the hourly ambient air relative humidity was recorded by an automatic weather station (HOBO event logger; Onset Computer Corporation, Bourne, MA, USA) located 3 km away.
A micro-lysimeter (Ding et al., 2013; Kool et al., 2014) replicated thrice and made of high-density polyethylene with a 10 cm in depth, 5.2 cm inner radius, and 3 mm thickness was used to obtain the soil evaporation amount. The micro-lysimeter was pushed into the soil surface between maize rows to retrieve an undisturbed soil sample. Subsequently, we sealed the bottom, weighed the micro-lysimeter, placed it back in the soil at the same level as the soil surface, and no other sensor was installed in the micro-lysimeter. After 2 d of evaporation, the lysimeter was weighed again. The mass difference was defined as the amount of soil evaporation. When evaporation occurs, unlike with soil outside the lysimeter, the soil within lysimeters is not replenished with water from deeper layers; thus, relative to soil outside the lysimeter, the soil water content within the lysimeters is generally smaller following continuous evaporation. Therefore, to represent the field soil conditions, the soil within the lysimeter was replaced every 4 d. In addition, after every rainfall or irrigation period, the inner soil was changed immediately.
All water samples were analyzed for
The results are reported in
The isotopic composition of EW was calculated using the condensation water
that adhered to the underside of the newly covered plastic film. We assumed
that the water vapor under the newly covered plastic film and above the
surface soil constitutes a closed system. Within the system, two equilibrium
fractionation processes are temperature dependent and occur independently, i.e., evaporation from surface soil water to air under the plastic film occurs during the day time (08:00 to 20:00 LT; Fig. 2), condensation from the water vapor under the plastic film to liquid water ensued at nighttime (20:00 to 08:00 LT), and the resulting dew (condensation water) adhered to the plastic film. The average temperatures from 08:00 to 20:00 LT and 20:00 to 08:00 LT on the day before water collection were used to calculate the equilibrium fractionation factor (
Temporal variation in temperature of soil under film, vapor under film, field soil, and ambient air during the study period.
Based on Eqs. (3) to (6) and Fig. 1b, the fractionation factors for the two
processes under the newly covered plastic film are expressed using Eqs. (7) and (8).
Combining Eqs. (7) and (8), we obtain the isotopic composition in the EW as follows:
For an open system (field soil condition; Fig. 1c), evaporation from surface
soil water to ambient air undergoes the following two processes: the equilibrium fractionation process from the surface soil to the saturated vapor layer above the soil surface and the kinetic fractionation process from the
saturated vapor layer to ambient air. The isotopic composition of evaporation vapor is controlled by the isotope values of the evaporating soil water and ambient vapor, equilibrium, and kinetic fractionations. The kinetic fractionation can be described by the enrichment factors (
The isotopic compositions of bulk soil water and evaporating water can be
used to evaporating soil water in the Craig–Gordon model (Eq. 14) to
calculate the isotope value of the evaporation vapor (
The dual-isotope plot of precipitation and 0–5 cm bulk soil water from 25 July to 25 August 2016 (period I). The regression line of precipitation represents the local meteoric water line.
A general linear model (GLM) was used to test if the regression lines for
isotopic composition/evaporative water loss of BW as a function of days
after precipitation/irrigation (DAP/I) differ from those of EW. GLM was also
used to compare the period I evaporative water loss derived from
Between the two large precipitation events on 24 July and 20 September 2016, there was no effective precipitation, except for an irrigation event of 30 mm on 26 August 2016 (Fig. 4a). Thus, two continuous evaporation periods can be identified, i.e., period I from 25 July to 25 August 2016 and period II from 27 August to 19 September 2016.
The amount of precipitation, irrigation, and 0–5 cm bulk soil water content
Soil water content in 0–5 cm reached field capacity (0.30 cm
The precipitation on 24 July 2016, had a
As expected, the
Temporal variation in
The evaporation line, defined as the change in water isotopes with
evaporation time in EW, was remarkably similar to that for BW (Fig. 5). For
example, in period II,
In period II,
In period I, we compared the mean
Temporal variation in deep soil water content,
The precipitation event on 24 July 2016 increased the soil water content
in the top 60 cm and decreased soil water
Similar to precipitation on 24 July 2016, the irrigation on 26 August 2016 increased the soil water content and decreased the
In period I, evaporative water loss (
Temporal variation in evaporative water loss (
During evaporation, light isotopes are preferentially evaporated, enriching
the residual liquid water in heavy isotopes (Mook and De Vries, 2000). This
could explain why, with increasing evaporation time,
For the latter, because there is negligible water input from the atmosphere
(both in vapor and liquid form), the only water input could be from the soil
below 5 cm. Indeed, because the evaporation amount was larger than the
0–5 cm soil water storage reduction (Sect. 3.1), the water below 5 cm must have moved upward as evaporation occurred. Consequently, due to evaporation, the order of the
No significant
Schematic of soil pore water partitioning during evaporation.
For large and intense precipitation events, event water preferentially
infiltrates into the empty large pores because of their high hydraulic
conductivity. The infiltrated water may partially or fully transfer to the
surrounding empty smaller pores, thus bypassing the small soil pores that
are filled with pre-event water at the point of water entry and along the
infiltration pathway (Beven and Germann, 1982; Booltink and Bouma, 1991;
Šimůnek and van Genuchten, 2008; Weiler and Naef, 2003; Zhang et
al., 2019). The bypass flow occurs universally (Lin, 2010) and has also been
reported in our experimental site at the Chinese Loess Plateau (Xiang et al.,
2018; Zhang et al., 2019). In our experiment, the precipitation event on
24 July 2016 was 31 mm, with the intensity of 10.3 mm h
On the other hand, at the end of the evaporation period, lc excess of 0–5 cm soil at 24 DAP, which had a lower soil water content than in period II, was still the smallest compared with deeper soil (Fig. 6d). Therefore, the evaporation front was in the surface soil during both periods. Accordingly, the evaporation in our experiment was in evaporation stages I or II, as indicated in the Introduction. During evaporation stages I and II, small-pore water does not evaporate (Or and Lehmann, 2019; Zhang et al., 2015), and larger-pore water is the primary source of water for evaporation (Lehmann and Or, 2009; Or et al., 2013).
Therefore, EW is mainly from larger-pore water, similar to the event water
in isotopic composition; BW contains EW and evaporation-insulated small-pore
water, similar to the pre-event water. Compared with pre-event water, event
water takes evaporation precedence. Therefore, the sequence of water in the
evaporation layer can be analogically summarized as adhering to a
“last-in-first-out” rule. Thus, when isotopic composition in the event
water was smaller than that in pre-event BW, such as
Furthermore, evaporative enrichment and loss of larger-pore water both
affect the temporal variation in
There was a significant difference in the isotopic composition between EW
and BW; however, the evaporative water loss derived from EW and BW did not
differ (
The event water was more enriched in heavy isotopes than pre-event soil
water, as shown by our
Larger-pore water, preferred by evaporation, also has a relatively higher matric potential and flows more rapidly and may, thus, be preferred by roots and dominate groundwater recharge (Sprenger et al., 2018). In other words, evaporation, transpiration, and groundwater preferentially tap the same pool of water, i.e., the water that resides in larger soil pores. This is inconsistent with Brooks et al. (2010), who separated soil water into the following two water worlds: mobile water, which eventually enters the stream, and tightly bound water, which is used by plants. In our study, soil water content was below field capacity, and thus, according to Brooks et al. (2010), all water in our soil is “tightly bound water”, including the large-pore water we discussed above. Therefore, in our study, the larger-pore water is still under the field capacity, i.e., the water that percolates into streams (groundwater) rather slowly and/or is adsorbed by plant roots, which has broad ecohydrological implications.
We performed an experiment in two continuous evaporation periods, namely a
relatively depleted water input in period I and a more enriched
The results showed that
The data that support the findings of this study are provided in the Supplement.
The supplement related to this article is available online at:
HW, JJ, BC, and BS designed the research, prepared and interpreted the data, and wrote the paper. MW offered constructive suggestions for the paper. HW and XM conducted the fieldwork.
The authors declare that they have no conflict of interest.
Publisher's note: Copernicus Publications remains neutral with regard to jurisdictional claims in published maps and institutional affiliations.
We thank Han Li, Wei Xiang, Eric Neil, and Huijie Li for the fruitful discussions.
This research has been supported by the National Natural Science Foundation of China (grant nos. 41630860 and 41371233), the Natural Sciences and Engineering Research Council of Canada (grant no. 11111111), the Major Scientific and Technological Innovation Projects of Shandong Key R & D Plan (grant no. 2019JZZY010710), and the China Scholarship Council (grant no. 201806300115).
This paper was edited by Natalie Orlowski and reviewed by two anonymous referees.