A comparison of summertime water and CO
2
fluxes over
rangeland for well watered and drought conditions
Tilden P. Meyers
∗NOAA/ARL/Atmospheric Turbulence and Diffusion Division, 456 South Illinois Avenue, PO Box 2456, Oak Ridge, TN 37831-2456, USA
Received 27 January 2000; received in revised form 21 August 2000; accepted 23 August 2000
Abstract
Continuous measurements of the surface energy balance components (net radiation Rn, sensible heat flux H, latent heat
flux LE, ground heat flux G, and CO2fluxes began in early June of 1995 at the Little Washita Watershed, near Chickasha,
Oklahoma. A severe drought during 1998 provided a unique opportunity to evaluate the range of fluxes that can be expected during the summer period. Data obtained during four continuous summer periods were used to evaluate the year-to-year variability in summertime energy and CO2fluxes.
During the summer period (day 150–240), total evapotranspiration for non-drought years ranged from 224 to 273 mm with a mean and standard error of 253±12 mm. The mean and standard error of the net ecosystem exchange (NEE) rate of carbon dioxide for the same summer period was−120±36 g C/m2. In a year with severe drought (1998) total evapotranspiration for the summertime period was 145 mm. The lack of precipitation during this time resulted in total losses to the atmosphere of 155 g C/m2from soil respiration. © 2001 Elsevier Science B.V. All rights reserved.
Keywords: Sensible heat flux; Latent heat flux; Ground heat flux; Drought conditions; Carbon flux
1. Introduction
The sensitivity of regional and global scale (i.e., general circulation, or GCM) numerical models to the surface layer parameterizations of soil moisture and the surface energy balance components is well docu-mented (Troen and Mahrt, 1986; Meehl and Washing-ton, 1988; Sato et al., 1989; Atlas et al., 1993; Garratt, 1993). Characterization of the surface energy balance is necessary for determining the overall surface tem-perature, the surface flux of water vapor into the atmo-sphere, and the overall atmospheric heating rate. On short time scales, these processes determine the atmo-spheric stability, the height of the planetary boundary
∗Tel.:+1-423-576-1245; fax:+1-423-576-1327. E-mail address: [email protected] (T.P. Meyers).
layer (PBL) (Pan and Mahrt, 1987), and play a role in governing cloud formation and convective precipita-tion processes. On longer time scales, the local water balance could play a significant role in determining the capacity of the surface to fix carbon. For example, an accumulated deficit in the local water budget can potentially alter the sign and magnitude of the sea-sonal carbon budget, which can then have feedbacks into the evaporative processes. Ultimately, the short-and long-term carbon balance has the potential to be intimately linked to the local water balance.
To evaluate the impact of water deficits on the local surface energy balance and carbon fluxes on seasonal and annual time scales, long-term continuous mea-surements of the surface energy balance components and CO2fluxes were initiated in June of 1995 in the
Little Washita Watershed in south central Oklahoma
as a National Oceanic and Atmospheric Administra-tion (NOAA) contribuAdministra-tion to the GEWEX Continental Intercomparison Project (GCIP); (see Lawford, 1999). Data collected during the last 4 years have provided a unique opportunity to investigate the inter-annual vari-ability in summertime water and carbon balances for a rangeland location. In particular, a drought during the summer of 1998 (believed to be associated with the El Niño event) provided an opportunity to quan-tify extreme values for parameters related to the local water and carbon budgets.
2. Methodology
The turbulent fluxes of water vapor, sensible heat, and CO2 were measured using the eddy covariance
technique. Historically, the use of the eddy covari-ance method (Businger, 1986; Baldocchi et al., 1988) has been constrained to mainly short-term intensive field campaigns. However, improvements in instru-ment design, ruggedness, and stability over the past decade now allow for nearly continuous measure-ments of sensible and latent energy fluxes using the eddy covariance technique (Goulden et al., 1996; Grelle and Lindroth, 1996; Moore et al., 1996). With this technique, the average vertical turbulent eddy fluxes of sensible and latent heat and other scalars are determined as
w′χ′=
Pn
i=1(w− hwi)(χ− hχi)
n (1)
wherewis the vertical velocity component of the wind vector, andχ the scalar of interest (e.g., water vapor concentration). Here, the bracketed quantities denote an average or “mean” that is subtracted from the instantaneous values to obtain the fluctuating compo-nent. Average vertical turbulent fluxes (w′χ′) are com-puted in real time using a digital recursive filter (400 s time constant) for the determination of a running “mean”, which is subtracted from the instantaneous values to obtain the fluctuations from the mean. An averaging period of 30 min (denoted by the overbar) is used and is considered large enough for statistical confidence in the covariance quantity, but is also short enough to resolve the structure of the diurnal cycle.
Wind vector measurements made at experimental sites that are not perfectly flat can result in non-zero
vertical wind velocities measured from the “vertical” coordinate system of the measurement platform. At the end of an averaging period, vertical turbulent fluxes perpendicular to the mean horizontal wind (which generally follows the contour of the land surface) are obtained by mathematically rotating the coordinate system of the measurement frame of ref-erence (sonic anemometer) to obtain a zero mean vertical and transverse velocity (w=v =0). Details of this procedure are outlined by Businger (1986) and Baldocchi et al. (1988).
The three components of the wind vector were de-termined with a sonic anemometer (model R2, Gill In-struments, Hampshire, England). The stable long-term operational characteristics of this instrument and its ability to provide measurements during cold weather and light rain events as well as its low power consump-tion were important consideraconsump-tions in the selecconsump-tion of this anemometer (Yellard et al., 1994). The symmetric head design of the R2 with its slender support structure produces little flow distortion (Grelle and Lindroth, 1994) and is well suited for measurements in relatively flat and open sites with short vegetation. Fast response water vapor and CO2 concentration measurements
were made with an open-path, fast response infrared gas analyzer (Auble and Meyers, 1992). This sensor has been used extensively for flux measurements in coastal experiments (Crawford et al., 1993), and recent ARM (Doran et al., 1992) and BOREAS (Baldocchi and Meyers, 1998) experiments. In a recent evaluation of open- and closed-path sensors for water vapor and CO2 concentrations, Leuning and Judd (1996) found
that for the measurement of CO2, this sensor displayed
minimal cross-sensitivity to water vapor (see Leuning and Moncrieff, 1990). The gas analyzer is swapped out every 2 months with a recently calibrated sensor. Calibrations are done both before and after exposure in the field using CO2standards obtained from NOAA’s
Climate Monitoring and Diagnostics Laboratory (CMDL) and a chilled mirror that has certification by the National Institute of Standards and Technology (NIST). Changes in sensor calibration over a 2-month period for both H2O and CO2 have been observed
to be less than 5%. The sonic anemometer and the IRGA were placed about 3 m above ground level.
Table 1
Meterological variables measured at NOAA Energy Flux Monitoring Sites along with model number and manufacturer of instrumentation used
Meteorological variable Manufacturer Model number
Air temperature and RH Vaisala, Helsinki, Finland 50Y
Net radiation Radiation and Energy Balance Systems (REBS), Seattle, WA, USA Q∗7
Global radiation LI-COR, Lincoln, NE, USA LI-200 SB
Precipitation Texas Instruments, Dallas, TX, USA –
Wetness NOAA, Oak Ridge, TN USA –
Soil heat flux REBS, Seattle, WA, USA –
PAR LI-COR, Lincoln, NE, USA LI-190 SB
Atmospheric pressure Vaisala, Helsinki, Finland PTB101B
Surface temperature Everest, Fullerton, CA, USA 4000A
Soil temperature NOAA, Oak Ridge, TN, USA –
Soil moisture Vitel, Chantilly VA, USA Hydra
and direction, air temperature and relative humidity, precipitation, net radiation, incoming global radiation, incoming and reflected photosynthetically active ra-diation (PAR), barometric pressure, ground heat flux, surface wetness, and soil temperatures at six depths: 2, 4, 8, 16, 32, and 64 cm. A soil moisture sensor (Vitel) was installed before the 1997 summer season. This probe, which measures the dielectric constant of the soil, water, and air matrix, was placed in the mid-dle of a 10 cm soil layer. The soil moisture was then determined using the methodology outlined by Wang and Schmugge (1980). The surface temperature was measured with an infrared temperature sensor. These meteorological sensors (Table 1) are sampled every 2 s with a datalogger and multiplexor (CR21X, Campbell Scientific, Logan, UT), and averages are computed ev-ery 30 min, coincident with the eddy covariance data.
2.1. Data acquisition
A laptop computer was configured to perform three operations simultaneously. The priority task is to re-ceive data from the sonic anemometer, comprising the components of the wind vector, the speed of sound from which the virtual temperature can be derived and the digitized H2O and CO2 signals from the IRGA.
For its second task, the computer retrieves the stan-dard meteorological data from the CR21X datalogger every 30 min and appends the data to an existing file. In background (the third task), a terminate and stay resident (TSR) communications program is used to retrieve these data from the laptop computer via
a cellular phone about once every 2 days. The en-tire system is powered by a bank of nine 12 VDC deep cycle batteries that are charged by eight solar panels. The system uses approximately 3 A at 12 V, continuously.
3. Site description
3.1. Little Washita Watershed, Oklahoma
Fig. 1. Location of the NOAA flux tower site in Southern Great Plains of the United States in southwestern Oklahoma.
4. Results and discussion
The results presented and discussed here are from data collected during the summer periods beginning 1 June and ending 31 August for the 4 years 1995–1998 (hereafter denoted as Y95, Y96, Y97, and Y98). Data recovery rates for the summer periods Y96, Y97, and Y98 were 85, 99, and 94%, respectively. However, dur-ing the first year of operation (Y95), several periods during the summer had gaps in the eddy covariance data resulting is a data recovery rate of 65%. Gap pe-riods were typically 2–3 days. The longest gap period in 1995 (7 days) began day 192. Water and carbon fluxes during the gap periods for all years during the day were therefore estimated by first computing ratios of water and carbon fluxes to global radiation for peri-ods adjacent to the gap periperi-ods, then using these ratios
with the corresponding measurements of global radi-ation during the gap period to estimate the fluxes. The utility of these procedures was demonstrated by Brut-saert and Sugita (1992) during the first ISLSCP field experiment (FIFE). For filling in gaps during night-time periods, water fluxes were assumed to be zero and a mean CO2 flux of 0.2 mg CO2/(m2s) was
as-sumed. This nighttime carbon flux rate was obtained by averaging nighttime fluxes for the entire summer period.
4.1. Precipitation and evapotranspiration rates
Fig. 2. Accumulated precipitation (mm) in years 1995–1998, for summer season (day 150–240). 1998 was a drought year.
Y98, no measurable rainfall was recorded for over 50 days beginning around day 165. These highly wet and dry years provide a unique opportunity to examine up-per and lower limits of total accumulated evapotran-sipiration (ET) and associated carbon fluxes during the summer season. It is during this 90 day period of the year that ET comprises roughly 50% of the annual wa-ter loss, and the associated daily carbon accumulation rates are highest.
At the beginning of the summer period (day 155), the daily water loss from evaporative processes (aver-age of 10 days) for all years was about 3 mm per day (Fig. 3). By day 165, 15 days into the summer pe-riod, accumulated rainfall for Y95 and Y96 increased
Fig. 3. Seasonal cycle of 10-day average daily evapotranspiration at the Little Washita Watershed for years 1995–1998. 1998 was a drought year.
Fig. 4. Seasonal cycle of 10-day average ratio of actual evaporation to equilibrium evaporation at the Little Washita Watershed for years 1995–1998. 1998 was a drought year.
to 150 mm while only about 20 mm was received in Y97 and Y98. The daily ET rates for Y95, Y96, and Y97 remained near 3 mm per day till day 200. There was virtually no additional precipitation from day 150 to 200 for Y98 and the 10-day average daily ET rates had progressively dropped to 1 mm per day. Equilib-rium evaporation was calculated as
LEeq=
s
s+γ(Rn−G) (2)
where s is the slope of the saturation vapor pressure curve andγ the psychometric constant. LE/LEeq at
the beginning of the summer season (Fig. 4) for all years was comparable to earlier observations by Kus-tas et al. (1996) within the Washita Watershed. Except for Y98, all years maintained a LE/LEeq ratio of at
least 0.7 through day 200. With little rain and available soil moisture in Y98, LE/LEeqprogressively dropped
from 0.65 to a minimum of 0.2 around day 205. After day 205, the LE/LEeqratio for Y98 increased slightly
in response to a few minor precipitation events. For the remaining years, LE/LEeqdecreased slightly to 0.6
from day 200 to 240, although there was enough pre-cipitation to balance most of the evaporative losses for Y95 and Y96. Accumulated rainfall for Y97 from day 150 to 205 was only 80 mm and LE/LEeq decreased
to just below 0.5.
Fig. 5. Seasonal cycle of soil moisture (upper 10 cm zone, volu-metric water content) for 1997 and 1998. 1998 was a drought year.
and Y98 was nearly equal at 400 mm. However, lit-tle precipitation (20 mm) was recorded between day 115 (15 April) and day 150 (30 May) for Y98 while 210 mm was observed for Y97 for the same time pe-riod, providing more available stored water for evap-otranspiration. The seasonal trends in soil moisture at 10 cm for Y97 and Y98 also depict the frequency and impact of the precipitation events (Fig. 5). For Y97, the volumetric water contentθv(in g H2O/cm3) was
above 0.20 more than 40% of the time. In Y98, θv
was above 0.20 for only 3% of the time.
The 10-day average air temperature at 2 m for all years peaked around day 210 but was highest in Y98, exceeding the other years by nearly 3◦C (Fig. 6).
Fig. 6. Seasonal cycle of 10-day average 2 m air temperature at the Little Washita Watershed for years 1995–1998. 1998 was a drought year.
Fig. 7. Seasonal cycle of 10-day average 64 cm soil temperature at the Little Washita Watershed for years 1995–1998. 1998 was a drought year.
Higher air temperatures during Y98 were reflected in the higher soil temperatures observed at the 64 cm depth which reached 25◦C, even though at the begin-ning of the summer season, the soil temperatures for Y95, Y96, and Y98 were nearly identical (Fig. 7). The maximum observed surface temperature, as measured from a downward looking infrared sensor, showed similar trends with seasonal maximums around day 210 (Fig. 8). The higher surface temperatures are consistent with the lower observed evaporation rates (Fig. 3). However, the differences between years is striking with Y98 maximum surface temperatures that were 12◦C above year Y96. Maximum surface
Fig. 9. Seasonal cycle of 10-day average daily net carbon fixation for the summertime period during 1995–1998. 1998 was a drought year.
temperatures for Y97, the second warmest summer, were 4◦C below the maximum observed surface tem-peratures in Y98. Visual observations in mid-July 1998 confirmed that little of the vegetation was green. Most of the grasses and weeds were brown from lack of precipitation.
The seasonal variation in the daytime and night-time carbon fluxes varied considerably from year to year. The largest daily average net ecosystem ex-change (NEE) rate (carbon fixation) observed for all four summer seasons was about−3 g C/m2 per day, with the largest fixation rates occurring in the early part of the summer period (Fig. 9). Although Y96 was the wettest of the four summers, daytime carbon fixation rates were not the highest since there was a much higher frequency of cloudy days and CO2
fixation rates were light-limited. For Y98, the range ecosystem was always a source of carbon for the day, with NEE rates ranging from 0 (daily uptake offset by nighttime losses) to nearly 4 g C/m2 per day by the end of the summer period. Nighttime respira-tion losses were computed by summing the fluxes from 18:00 LST to 06:00 LST the next morning with daytime fluxes summed for the remaining 12 h. The warm moist soils of Y96 also produced the largest nighttime fluxes with losses over 2.5 g C/m2 per day (Fig. 10). During the driest year Y98, maximum nighttime losses were just above 1 g C/m2 per day in the early summer and dropped to 0.5 g C/m2 per day by day 200, with slightly higher rates after a few minor precipitation events (Fig. 10).
Fig. 10. Seasonal cycle of 10-day average daily nighttime carbon fluxes for the summertime period during 1995–1998. 1998 was a drought year.
4.2. The surface energy balance and carbon fluxes
Using all data from the summer of 1998, closure of the energy balance is well within the combined uncer-tainty of the independent measurements of H, LE, Rn
and G, with a slope of 0.97 and offset of−16 W/m2 (Fig. 11).
Ensembles of the measured diurnal components of the surface energy balance (Rn, H, LE, and G) for
a non-stressed year (Y96) during mid-summer are
Fig. 12. Typical diurnal cycle of the components of the surface energy balance and CO2 flux at a grassland site in the Little Washita Watershed for non-stressed conditions.
shown in Fig. 12. Although these ensembles were derived by averaging 3 days in which the atmospheric conditions were similar, the results are typical of non-stressed periods. Midday peaks of net radiation were near 650 W/m2, with nearly 50% of the net radi-ation partitioned into LE (about 300 W/m2), 35% into sensible heat flux (about 250 W/m2) and 15% into ground heat flux (100 W/m2). The evaporative fraction (LE/(Rn−G)≈0.46) is nearly constant and showed
no dependence on vapor pressure deficit, similar to several sites in the Washita Watershed reported by Kustas et al. (1996). The evaporative fraction is lower than what was observed by Verma et al. (1992) over a similar C4 grass ecosystem during mid-summer;
conversely the sensible heat fraction (H /Rn≈0.38)
is larger than the fractions observed by Verma et al. (1992). This in part could be attributed to active graz-ing that occurred throughout the summer, whereas the measurements of Verma et al. (1992) were made over ungrazed prairie. Over the diurnal cycle for this non-stressed period, LE slightly lagged behind net radiation, with H peaking earlier in the day and G peaking later around 14:00 LST. H peaked before noon LST while maximum G was around 14:00 LST, and actually exceeded H by 15:00 LST. The cycle of
LE was, however, in phase and proportional with the available energy (Rn−G), with no dependence on
vapor pressure deficit. The associated CO2flux was in
phase with both LE and net radiation with midday up-take rates peaking near−0.75 mg CO2/(m2s).
Night-time respiration losses were smaller in magnitude but were relatively constant (0.25 mg CO2/(m2s)).
For water stressed periods in Y98, surface en-ergy balance and carbon fluxes were very different (Fig. 13). The pattern of net radiation was similar to the net radiation in non-stressed years but the midday peak was about 50–70 W/m2 less than that for the non-stressed years. The difference can be reconciled with the increased outgoing longwave radiation during the stressed years (proportional to the fourth power of the surface temperatures, which were much higher in Y98) (Fig. 7). The sensible heat flux comprised the largest fraction of Rn (about 65%), with midday
val-ues that peaked near 400 W/m2. The second largest term was G, which accounted for nearly 25% of
Rn with midday values of 180 W/m2. LE comprised
the remaining 10% of Rn with midday fluxes near
Table 2
Seasonal (day 150–240) total precipitation, evaporation, net carbon ecosystem exchange, and average ratio of evaporation to equilib-rium evaporation
1998 67 145 155 0.39
75 W/m2. Soil moisture values in the upper 10 cm of the soil profile remained near 0.10.
With little active vegetation, the land surface was a source of CO2 to the atmosphere with midday
emis-sion rates near 0.25 mg/(m2s). Nighttime emission
rates were much smaller (0.05 mg/(m2s) with soil moisture levels near 0.1 (Fig. 5). The minimum car-bon flux appears to occur just after sunset. During this transition period, surface layer turbulence is usu-ally suppressed as the surface quickly begins to cool. Consequently, air-surface exchange rates are often spurious and are often relatively low.
For the entire summer season, total evaporation for non-drought years ranged from 224 to 273 mm (Table 1). For Y98, the drought year, total evaporative loss for the summer was only 145 mm, roughly 60% of the average losses for the other years. Daily evap-oration rates for the non-drought years (≈3 mm per day) are comparable to those found by Dugas et al. (1999) for native prairie in Texas. The amount of car-bon dioxide fixed each season was much more vari-able than the evaporation rates during the non-drought years, ranging from 41 to 196 g C/m2. The lack of precipitation in Y98 resulted in little to no carbon fix-ation from photosynthesis. This resulted in a net loss of carbon from the land surface to the atmosphere of 155 g/m2for the summer season (Table 2).
5. Conclusions
Summertime water and carbon fluxes over grass-land were determined from an eddy covariance sys-tem were measured over a 4-year time period from 1995 to 1998 with 1998 being a drought year with only 67 mm of precipitation for the 90 day summer period. For the non-drought years the average water
and net carbon fluxes (net fixation) were 253±12 mm and−118±36 g C/m2, respectively. The evaporative fraction (LE/Rn−G) was nearly constant (0.46) and
showed little dependence on vapor pressure deficit during non-drought periods. Similar to earlier mea-surements from the same region (Kustas et al., 1996), the LE/LEeqranged between 0.6 and 0.7. Peak daily
water losses (3 mm per day), however, are comparable to those reported by Kelliher et al. (1993) for similar grassland ecosystems. During the drought year (1998) summertime evaporative losses totaled 145 mm. The lack of precipitation coupled with high solar radia-tion resulted in surface or “skin” temperatures that exceeded 50◦C, resulting in lower values for net radiation because of the higher outgoing longwave ra-diation. The lack of rain also resulted in little carbon uptake via photosynthesis, changing the landscape from acting as a sink of carbon during non-drought years to a source of carbon dioxide with a loss of 155 g C/m2for the 90 day period from 1 June to 30 August.
Acknowledgements
I would like to acknowledge the support of Dr. Rick Lawford and Dr. John Leese of the NOAA Office of Global Programs for their support of this work.
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