ESTIMATING THE MAGNITUDE OF COUPLED-FLOW EFFECTS IN THE ACTIVE LAYER AND UPPER PERMAFROST, BARROW, ALASKA U.S.A.

Transcription

1 ESTIMATING THE MAGNITUDE OF COUPLED-FLOW EFFECTS IN THE ACTIVE LAYER AND UPPER PERMAFROST, BARROW, ALASKA U.S.A. S. I. Outcalt 1, K. M. Hinkel 2, F. E. Nelson 3, L. L. Miller 4 1. Department of Geography, University of Cincinnati, Cincinnati, OH USA smout@compuserve.com 2. Department of Geography, University of Cincinnati, Cincinnati, OH USA Ken_Hinkel@compuserve.com 3. Department of Geography, University of Delaware, Newark, DE USA fnelson@udel.edu 4. Department of Geography, University of Cincinnati, Cincinnati, OH USA millll@ .uc.edu Abstract Upward migration of soil water and its eventual evaporation near the ground surface cools the active layer above permafrost during summer. The thermal magnitude of these effects was estimated using a numerical model based on conductive heat-transfer theory and incorporating fusion effects. The model utilizes precision measurements of soil temperature as initial and boundary conditions to simulate the evolution of the thermal profile over the annual cycle. The simulation indicates that upward water movement and near-surface evaporation cools the active layer several degrees during summer and depresses the mean annual ground temperature by about 0.4 C. Nonconductive cooling of the active layer also attenuates surface thermal variations and moderates the thermal signal that penetrates the upper permafrost. Because the ratio of the heat of evaporation to the heat of fusion is approximately seven, an increase in summer temperature may not cause dramatic permafrost degradation and surface subsidence at poorly drained sites. Introduction Soil moisture evaporation and upward migration of water are well known in unfrozen soil. Both entail coupled flow of mass and heat, the latter in latent or sensible form. However, quantifying the effect of these nonconductive heat-transfer processes on the soil temperature profile has proven problematic. One approach entails modeling the thermal evolution of the upper soil using a numerical model based on conductive heat-transfer theory and incorporating fusion effects. Previous observations and modeling of the midday thermal regime in moist peat above the thawing ice core of a palsa indicated that soil water from the thaw front flows upward and evaporates near the surface (Outcalt and Nelson, 1985). Although the surface temperature exceeded 40 C, the upward flowing meltwater maintained the soil temperature near 0 C at a depth of 5-10 cm. Near the surface, rapid temperature fluctuations indicated that the soil was desiccated. Vertical advection of cooler water and internal evaporation appear to have combined to depress soil temperatures in the desiccated region where water was evaporating to the atmosphere. Because evaporation entails consumption of energy, evaporative cooling slowed the advance of the thaw front below rates estimated by the model. A similar approach can be adopted to assess the seasonal and annual impact of coupled-flow processes. In this analysis, time series of measured temperature are used as model boundary conditions to simulate soil temperatures over time and with depth. The magnitude of coupled-flow effects can be estimated by examining the residuals created by subtracting simulated temperatures from observed temperatures. Data sets and model Two thermistor strings, separated laterally by about 50 m, have been recording hourly data at eight soil levels since 1993 at a site near Barrow, Alaska. Both are installed in reworked marine silts covered by a 5-cm thick organic mat. The first installation measures temperatures at depths of 0, 10, 20, 30, 45, 70, 95 and S. I. Outcalt, et al. 869

2 120 cm; the resulting time series is used as the observed data in this analysis. Temperatures were simulated using a numerical model based on conductive properties and incorporating fusion effects. The methodology was used previously to explore the impact of internal evaporation over permafrost as compared to an adjacent non-permafrost seep site (Outcalt et al., 1997). The thermal properties of the soil were determined in a previous study conducted near the present site (McGaw et al., 1978), and the results are used in this analysis. Apparent thermal diffusivity (ATD) and measured temperature were fit to a polynomial to derive a temperature-dependent ATD value. ATD increased from a small value at 0 C to 6 x 10-7 m 2 s -1 at -6 C. For simplicity, the value of 6 x 10-7 m 2 s -1 was used above 0 C and below -6 C. The top and bottom observed temperature time series (at 0 and 120 cm) were used as boundary conditions and the model was initiated with the first thermal profile in the observed record. The internal node values were simulated using a fully implicit finite-difference model in which the new thermal vector is estimated by Gaussian elimination applied to the tri-diagional matrix and the current temperature vector (Press et al., 1989; Outcalt et al., 1997). The model makes a simple adjustment for unequal node spacing during calculation of the temperaturedependent Fourier modulus. The second thermistor string consists of eight thermistors at depths of 1, 8, 15, 22, 29, 50, 75 and 100 cm. To optimize the resolution of the thermistor measurements (0.015 C), measurements were restricted to temperatures above -22 C, making this data set inappropriate for numerical analysis of the annual cycle. However, in addition to temperature, measurements of soil voltage are made at the thermistor levels. The soil voltage was transformed to a surrogate of the soil water solute concentration (Hinkel and Outcalt, 1994; Hinkel et al., 1997), and this information was used to enhance interpretation of the results. Simulations and interpretation Figure 1 illustrates hourly temperature readings from the first installation over a complete annual cycle beginning 29 June Active-layer thickness at Barrow rarely exceeds 40 cm, so the lower probes are within the upper permafrost. In the active layer, four thermal regimes are generally discernible during the annual cycle; vertical dividing lines were added to delimit transitions between regimes. The "active layer" regime (AL) develops in summer when the thaw zone is thickening. Although conduction dominates heat transfer, coupled-flow processes can significantly affect the thermal evolution of the profile. The thawed soil is cooled by internal evaporation and the upward migration of colder water from the thaw front toward the evaporation front near the ground surface (Outcalt and Nelson, 1985). Infiltration of precipitation can also affect the Figure 1. Hourly time series of soil temperature at installation 1 for one year beginning on 29 June 1995 at Barrow, Alaska. 870 The 7th International Permafrost Conference

3 thermal regime significantly during this period (Hinkel et al., 1993; Hinkel and Nicholas, 1995). During autumn and early winter, the active layer cools and begins to refreeze. Often, the lower part of the unfrozen zone becomes isothermal near 0 C for several weeks or months. This is known as the zero-curtain effect (Sumgin et al., 1940; Outcalt et al., 1990), and the period during which this effect operates is referred to as the "zero curtain" (ZC) regime. Although heat conduction cannot occur across an isothermal zone, water and water vapour migrate vertically across the unfrozen layer in response to osmotic gradients (Outcalt et al., 1990), transporting both latent and sensible heat upward. This regime is dominated by nonconductive heat-transfer and phase-change effects. As the soil begins to freeze in mid to late winter, the "freezing" (FR) regime begins. Soil permeability decreases as pathways between pores are blocked by ice formation, and the movement of water and water vapour is restricted. This period is dominated by heat conduction. During spring snowmelt, meltwater typically infiltrates the soil. The downward transfer of mass and sensible heat alters the bulk soil thermal properties and disrupts the thermal field (Farouki, 1981; Hinkel et al., 1997). Rapid nonconductive warming of the frozen soil toward 0 C is characteristic of the "snowmelt" (SM) regime. As the soil thaws in early summer, the AL regime begins the cycle again. These regimes are based on the dominant heattransfer process, so the regime patterns are known from other analyses (Outcalt et al., 1997; Outcalt and Hinkel, 1992; 1996; Hinkel and Outcalt, 1993; 1995). However, none of these methods yields a good estimate of the thermal magnitude of the nonconductive effects. Figure 2 is a plot of hourly temperature residuals derived by subtracting simulated from observed temperature. Only traces from the middle of the probe stack, where residuals are maximized, are shown to avoid clutter. Large negative residuals (> 2 C ) are apparent during the AL-95 and AL-96 regimes at the 20, 30, and 45 cm probe levels; residuals tend to decrease in magnitude with depth. Negative residuals indicate that the soil is cooler than expected. Evaporation of soil water and upward advection of cool water from the thaw front are the likely mechanisms. Moderate positive residuals are observed during the ZC regime, indicating that the soil is warmer than predicted, and illustrates the effect of latent heat in retarding frost. Only during the winter FR regime, when the Figure 2. Time series of hourly residuals calculated by subtracting simulated from observed temperature for intermediate soil levels. S. I. Outcalt, et al. 871

4 active layer. At this level, the soil is cooled about 0.4 C by evaporation. The soil below the layer experiencing evaporation is cooled by upward advection of cold water from the thaw front. These processes produce a cooler profile than expected from a purely conductive heat-transfer system. Figure 4 displays temperature and relative solute concentration profiles from the second installation. Plotted for noon on 12 August, the pattern is typical for a warm, sunny summer day. The thawed region above 40 cm is characterized by a higher solute concentration than the upper permafrost. This suggests that vertical advection of water from the base of the thawed zone, and evaporation of soil water near the surface, effectively pumps moisture to the atmosphere and increases the solute concentration of soil water during summer. The abrupt peak in solute concentration near the surface (7 cm) supports this interpretation. Figure 3. Observed and simulated mean annual ground temperature (MAGT) for all probe levels. volume of unfrozen water is minimal, does the conduction-based model reliably estimate the soil thermal profile, although nonconductive effects are still detectable. Observed and simulated mean annual ground temperatures (MAGT) are displayed in Figure 3. Since observations from the upper (0 cm) and lower (120 cm) probe levels were used as boundary conditions in the finite-difference model, the observed and simulated MAGTs are identical at these levels. Observed temperatures within the active layer and upper permafrost are cooler than simulated temperatures, and this is especially pronounced in the middle part (20 cm) of the Conclusions Evaporation of soil water, observed and modeled previously during a single summer day, persists throughout the summer at Barrow and cools the active layer by several degrees. Integrated over the annual cycle, internal evaporation depresses the mean annual temperature in the active layer by 0.4 C. Evaporation also cools the adjacent upper permafrost during summer. Evaporative cooling must, therefore, distort the surface thermal climate signal at depth in an irregular manner dependent on interannual variations in summer temperature, precipitation, and evaporation rates. It is possible that internal evaporative cooling will buffer both thaw deepening and warming of the upper permafrost against higher summer temperatures due to global or regional warming. Furthermore, widespread thermokarst might be mitigated at wet, poorly drained sites due to evaporative cooling. Acknowledgments This research was supported by NSF grants SES , OPP , and OPP to KMH and OPP to FEN. We are grateful to the Ukpeagvik Inupiat Corporation for administrative assistance and access to the Barrow Environmental Observatory. The paper benefited from the comments of M. Allard and two anonymous reviewers. Figure 4. Temperature and surrogate soil solute concentration at installation 2 at noon, 12 August The 7th International Permafrost Conference

5 References Farouki, O. T. (1981). Thermal Properties of Soils. CRREL Monograph 81-1, US Army Corps of Engineers, Cold Regions Research and Engineering Laboratory, Hanover, New Hampshire (136 pp). Hinkel, K. M. and Nicholas, J.R.J. (1995). Active layer thaw rate at a boreal forest site in central Alaska. Arctic and Alpine Research, 27, Hinkel, K. M. and Outcalt, S.I. (1993). Detection of nonconductive heat transport in soils using spectral analysis. Water Resources Research, 29, Hinkel, K. M. and Outcalt, S.I. (1994). Identification of heattransfer processes during soil cooling, freezing and thaw in Central Alaska. Permafrost and Periglacial Processes, 5, Hinkel, K. M. and Outcalt, S.I. (1995). The detection of heatmass transfer regime transitions in the active layer using fractal geometry. Cold Regions Science and Technology, 23, Hinkel, K. M., Outcalt, S.I. and Nelson, F.E. (1993). Near-surface summer heat-transfer regimes at adjacent permafrost and non-permafrost sites in central Alaska. In Proceedings, 6th International Conference on Permafrost. South China University of Technology Press, Beijing, pp Hinkel, K.M., Outcalt, S.I. and Taylor, A.E. (1997). Seasonal patterns of coupled flow in the active layer at three sites in northwest North America. Canadian Journal of Earth Sciences, 34, McGaw, R. W., Outcalt, S.I. and Ng, E. (1978). Thermal properties of wet tundra soils at Barrow, Alaska. In Proceedings, Third International Conference on Permafrost. National Research Council of Canada, Ottawa, pp Outcalt, S. I., and Hinkel, K.M. (1992). The fractal geometry of thermal and chemical time series from the active layer, Alaska. Permafrost and Periglacial Processes, 3, Outcalt, S. I. and Hinkel, K.M. (1996). Thermally-driven sorption, desorption, and moisture migration in the active layer in central Alaska. Physical Geography, 17, Outcalt, S. I., Hinkel, K.M., Miller, L.L. and Nelson. F.E. (1997). Modeling the magnitude and time dependence of nonconductive heat-transfer effects in taiga and tundra soils. In International Symposium on Physics, Chemistry, and Ecology of Seasonally Frozen Soils, Fairbanks, Alaska, June 10-12, CRREL Special Report 97-10, pp Outcalt, S. I. and Nelson, F.E. (1985). A model of near-surface coupled flow effects on the diurnal thermal regime in a peat-covered palsa. Archives of Meteorology, Geophysics, and Bioklimatology, Series A, 33, Outcalt, S. I., Nelson, F.E. and Hinkel, K.M. (1990). The zerocurtain effect: heat and mass transfer across an isothermal region in freezing soil. Water Resources Research, 26, Press, W. H., Flannery, B.P., Teukolsky, S.A. and Vetterling, W.T. (1989). Numerical Recipes in C: The Art of Scientific Computing. Cambridge University Press, Cambridge (735 pp). Sumgin, M. I., Kachurin, S.P., Tolstikhin, N.I. and Tumel', V.F. (1940). Obshchee Merzlotovedenie [General Permafrostology].: Academiia Nauk SSR, Moscow (240 pp). S. I. Outcalt, et al. 873