Hydrological controls on chemical weathering rates at the soil-bedrock interface Emmanuel J. Gabet Department of Geology, University of Montana, Missoula, Montana 59802, USA
Robin Edelman Department of Geology and Environmental Sciences, Northeastern Illinois University, Chicago, Illinois 60625, USA
Heiko Langner Department of Geology, University of Montana, Missoula, Montana 59802, USA
ABSTRACT Chemical weathering of bedrock is critical to maintaining terrestrial life, and climate, typically as manifested by precipitation, is often identified as having a first-order control on rates of chemical weathering. The ability of precipitation to dissolve rock, however, is modulated by the properties of the overlying soil that influence the contact time between water and minerals. Flume experiments were conducted to investigate the hydrological controls on rates of chemical weathering. Solute concentrations of runoff flowing across synthetic bedrock overlain by nonreactive pseudosoils of differing hydraulic conductivities were measured to examine the role of seepage velocity in influencing weathering rates on steep slopes. The results suggest that, where weathering is not limited by the supply of fresh minerals, weathering rates should increase with decreasing hydraulic conductivity. In addition, a mathematical relationship between hydraulic conductivity and chemical weathering on hillslopes is introduced to explore the hydrological controls on feldspar and calcite dissolution rates. The mathematical model supports the results from the experiments. Keywords: chemical weathering, soil production, dissolution, hydraulic conductivity, advectivedispersion equation, hillslope hydrology.
INTRODUCTION Chemical weathering of bedrock is an important geological process. For example, by providing habitat for the vast majority of terrestrial organisms, including the primary producers that transform inorganic compounds into highly energetic organic matter, the conversion of bedrock to soil is critical in maintaining terrestrial life. In addition, it has been proposed that chemical weathering of silicate rocks may influence the global climate by consuming and sequestering atmospheric CO2 (Berner, 1994). Finally, chemical weathering may have a dominant role in the development of topography, even in nonkarst landscapes (Mudd and Furbish, 2004). Numerous studies have examined the role of climate, particularly precipitation, in controlling chemical weathering rates (e.g., Riebe et al., 2004; West et al., 2005; White and Blum, 1995). However, because precipitation must pass through the soil mantle to reach fresh minerals, soil properties should also influence weathering rates. Gilbert (1877) recognized that the contact time between water and rock was a limiting factor in chemical weathering processes and proposed that bedrock weathering rates on hillslopes would be enhanced by a layer of soil that could retard
runoff. Carson and Kirkby (1972) extended Gilbert’s hypothesis and introduced a humped function whereby chemical weathering is low on bare bedrock, maximized under an intermediate soil depth, but then declines under thicker soils, where the high solute load of slowly circulating water inhibits net dissolution. No field studies have been specifically designed to test this hypothesis, but indirect evidence suggests that chemical weathering rates are lower on slopes mantled by thin soils and bare bedrock (Drever and Zobrist, 1992; Heimsath et al., 2001; Stallard, 1984). Although the humped function was presented as a relationship between soil depth and chemical weathering, its rationale is based primarily on the rate at which water flows through the soil (Carson and Kirkby, 1972). In soils with a high hydraulic conductivity (Ks), the water residence time is shorter and there should be fewer opportunities for mineral dissolution. In contrast, dissolution should be enhanced where the hydraulic conductivity is low and contact times between water and rock are longer. To test the proposition that chemical weathering rates should be inversely related to Ks, we built a physical model to simulate bedrock dissolution on a steep hillslope. Dissolution by runoff was simulated by
releasing a pulse of water over a reactive synthetic bedrock covered by nonreactive pseudosoils of differing Ks. These experiments were designed to mimic weathering processes along a hillslope where the hydrological contribution from upslope runoff is greater than from direction precipitation (i.e., at some distance from the divide). MATERIALS AND METHODS A Plexiglas flume (50 cm long by 11 cm wide) with an 18% slope was constructed to accommodate slabs of synthetic bedrock (Fig. 1). The synthetic bedrock was created by mixing gypsum powder (CaSO4·2H2O), table salt (NaCl), and water at a ratio of 1:2:0.56, respectively, and pouring the slurry into trays. The specific composition of the synthetic bedrock and the length of the flume were chosen because they allowed us to simulate a range of conditions, from interface limited to transport limited, over a short distance. Three sets of five experiments were performed. In each set, the first run consisted of releasing water over bare bedrock. In the four successive experiments, the synthetic bedrock was covered by a 6 cm layer of pseudosoil composed of well-sorted quartz sand or gravel with the following median grain diameters:
䉷 2006 Geological Society of America. For permission to copy, contact Copyright Permissions, GSA, or
[email protected]. Geology; December 2006; v. 34; no. 12; p. 1065–1068; doi: 10.1130/G23085A.1; 5 figures; Data Repository item 2006232.
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Figure 3. Mass of dissolved synthetic bedrock approaches maximum in runs where drainage time ~500 s (gray line highlights trend and is not regression). Dissolution rates decrease exponentially with drainage time, illustrating inhibitory effect of high solute concentrations in slowmoving flows. Regression equation and statistic apply to dissolution rate data.
Figure 1. Top: hillslope flume. Bottom: gate is pulled up, releasing water. Beaker at end of flume catches runoff. DI is deionized.
4.0 mm, 2.08 mm, 710 m, and 212 m. Although in nature soils may be highly reactive, we chose to use nonreactive soils to focus on weathering at the soil-bedrock interface. In addition, this experimental configuration simplifies the calculation of reactive surface area. Before each experiment, a fresh bedrock slab was mounted in the flume. Unless a bare bedrock run was scheduled, one of the four pseudosoils (rinsed with deionized [DI] water and dried) was emplaced over the bedrock to a depth of 6 cm; a minimum depth of 6 cm was necessary to contain the flow within the soil and prevent surface runoff. Immediately prior to the experiments with the sand soils, the sand was prewetted with a known amount of DI water from a spray bottle. During earlier experiments, we found that dry sand absorbed a large fraction of the flow, thus complicating the interpretation of the results. To minimize
Figure 2. Mass lost per unit area through dissolution increases with decreasing flow velocity.
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evaporative losses, a layer of plastic wrap was laid on top of the soil surface. At the start of each experiment, the gate was quickly pulled up to release 1 L of DI water stored behind it (Fig. 1). The runoff was collected at the outlet at timed intervals and its electrical conductivity (EC) and volume were measured. Typically, 10%–20% of the pulse water became stored in the soils (GSA Data Repository Table DR11). Because we were interested in the total amount of bedrock dissolved rather than only the solute flux, the soils were removed from the flume immediately after each experiment and the EC of the pore water was measured. Average seepage velocities were estimated from the runoff hydrograph. The Na and Ca concentrations in the runoff from one set of experiments were determined through inductively coupled plasma–atomic emission spectrometry to calibrate the EC readings. The Na concentrations were consistently ⬃10 times greater than Ca concentrations; in the results, we focus our analysis on the dissolution of NaCl. RESULTS From the 15 experiments, the result from 1 run was dismissed (run 10; Table DR1; see ootnote 1). The results from the other experiments show that the mass lost through dissolution declines with increasing seepage velocity (Fig. 2). Hydraulic conductivity controls the flow velocity and thus the contact time between runoff and minerals. A greater amount of synthetic bedrock is dissolved where the drainage time (i.e., the elapsed time for the water to move down the length of the flume) increases due to a decrease in Ks (Fig. 3). An average dissolution rate (D; g cm⫺2 s⫺1) was calculated for each run with D ⫽ 1GSA Data Repository item 2006232, Figure DR1 (flow pulse illustration) and Tables DR1 (experimental data) and DR2 (model parameters), is available online at www.geosociety.org/pubs/ ft2006.htm, or on request from editing@ geosociety.org or Documents Secretary, GSA, P.O. Box 9140, Boulder, CO 80301, USA.
m/At, where m is the mass (g) of dissolved salt, A is the surface area of the bedrock (cm2), is porosity (included to account for available bedrock surface), and t is the drainage time (s). As the solute concentration in the pore water increases, the dissolution rate decreases exponentially (Fig. 3). Although there is a decline in the dissolution rate, the lower dissolution rates are compensated by the increases in contact time to yield the greater losses of mass observed in the runs with the finer pseudosoils. The exponential decline in dissolution rates, however, attests to the diminishing ability of low hydraulic conductivities to enhance chemical weathering rates. The dissolved mass approaches a maximum in the runs with a total contact time of ⬃500 s and only increases slightly as drainage times double (Fig. 3). The low dissolution rates for the runs with the fine pseudosoils suggest that the runoff is near saturation with Na, yet the average Na concentration for these runs, 22 g L⫺1, is much less than the saturation concentration, 360 g L⫺1 (Weast, 1985). This discrepancy suggests that the layer of water at the soilbedrock contact was not being significantly mixed with more dilute flow and that the rate of chemical diffusion from this layer to the upper layers was too slow to prevent the buildup of saturation concentrations. Alternatively, the low dissolution rates might have been due to a decrease in the supply of available salt; however, the relatively constant Na: Ca ratio indicates that the dissolution of the salt was not supply limited. MATHEMATICAL MODEL The role of Ks in controlling chemical weathering rates can be explored analytically. Assuming negligible dispersion, the advective-dispersive equation (ADE) for solute transport (Zheng and Bennet, 1995) is modified to account for dissolution: C AD C ⫽ ⫺u , t V x
(1)
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The net mass of mineral dissolved per unit area at a distance x from the drainage divide (Ex; M L⫺2) is the product of the dissolution rate and the contact time between water and rock. Because the dissolution rate decreases with time as the solute concentration increases, the mass of mineral dissolved is determined with: Ex ⫽
冕
x/u
D dt,
(3)
0
where x/u is the length of time necessary for all the water upslope to drain past x. Combining equations 2 and 3 yields: Ex ⫽
V A
冕
x/u
dC(t).
(4)
0
Figure 5. Dissolved mass of feldspar and calcite decrease with increasing seepage velocity according to equation 7. Note similarity between predicted calcite curve and flume data (Fig. 2); both calcite and salt dissolve quickly, and therefore approach equilibrium concentrations at relatively rapid seepage velocities.
In the simplest case, the dissolution rate can be expressed as (Lasaga, 1995): Figure 4. Illustration of control volume used to derive equation 1. Control volume is positioned such that its lower surface is soilbedrock interface. Shaded region of soil represents saturated layer. Input (i) of dissolved solids into control volume comes from upslope subsurface flow (uiCi) and dissolution of bedrock (D), where u is seepage velocity, C is solute concentration in soil water, and D is dissolution rate. Output (o) of dissolved solids is uoCo.
where C is concentration of dissolved weathering product (M L⫺3), t is time (T), A is surface area (L2) of the mineral, D is dissolution rate (M L⫺2 T⫺1), V is volume (L3) of water, u is seepage velocity (L T⫺1), and x is distance (L) (Fig. 4). Although the highly unsteady flow associated with a pulse of water precludes a simple analytical solution of the ADE to model chemical weathering rates for the flume experiments, equation 1 can be used to investigate hydrological controls on chemical weathering processes under more natural conditions. Assume a brief but intense rainstorm on a steep hillslope where the soil thickness and Ks are approximately uniform. Infiltration of the rain produces a wetting front that travels vertically down through the soil until it reaches the bedrock, at which point the subsurface flow begins to travel downslope along the soil-bedrock contact (Fig. DR1; see footnote 1). Because the water along the length of the hillslope reaches the bedrock at the same time and follows the same trajectory down the hillslope, the solute concentration in the flow is approximately uniform. With C/x 艐 0, equation 1 is recast as: D⫽
V dC . A dt
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(2)
[ ]
C(t) D⫽k 1⫺ , Ceq
(5)
where k is a dissolution rate constant (M L⫺2 T⫺1) and Ceq is the equilibrium concentration. Although this equation simplifies a complex process and it may not represent dissolution in all situations (e.g., Burch et al., 1993), there is evidence that it captures the essence of potassium feldspar and calcite dissolution kinetics (Drever and Clow, 1995). Combining equations 2 and 5 and integrating, the evolution of C through time can be expressed as
冦
[
C(t) ⫽ Ceq 1 ⫺ exp (⫺)
]冧
A k t . V Ceq
(6)
Combining equations 4 and 6 yields
冦
[
]冧
V A k x Ex ⫽ Ceq 1 ⫺ exp (⫺) , (7) A V Ceq u an expression for the mass lost by dissolution at any point along a hillslope as a function of seepage velocity. Using published values for Ceq and k for potassium feldspar and calcite, and reasonable values for the other parameters (Table DR2; see footnote 1), equation 7 predicts an increase in net dissolved mass with decreasing seepage velocity (Fig. 5), thus suggesting that the general results from the flume experiments may be applicable at the field scale. Therefore, under conditions of pulsed shallow subsurface flow, the hypothesized increase in chemical weathering rates with decreasing Ks (Carson and Kirkby, 1972) can be theoretically confirmed (Fig. 5). Note that this analytical solution does not account for weathering in the interludes between storms or the effects of soil
thickness on the dilution of flow at the soilbedrock interface; these elaborations are beyond the scope of this study. DISCUSSION AND CONCLUSION The experiments and the mathematical model support the hypothesis that chemical weathering rates are enhanced in soils with a lower Ks and longer soil water residence times. However, in order to isolate Ks as a controlling variable, the experiments and numerical model simplified the hydrological and weathering processes that might be observed in nature. First, because we modeled the bedrock (physically and mathematically) as a continuous surface, there was no flow within fractures. Where shallow subsurface flow is moving into and emerging from bedrock fractures (Anderson and Dietrich, 2001), the contact time may be essentially independent of Ks. Second, the uniform size of each pseudosoil created homogeneous flow conditions, whereas in nature, the development of argillic and calcic horizons, as well as the presence of macropores, will lead to heterogeneous flow (Stonestrom et al., 1998; White et al., 2005). The landscapes we seek to simulate are steep with high erosion rates and unlikely to exhibit strong horizonation; the importance of macropores is unknown but may be significant. Third, a fresh slab of bedrock was installed at the beginning of each experiment, simulating a landscape where chemical weathering is not limited by the exposure of fresh minerals. Although physical weathering may maintain a relatively constant supply of fresh minerals (Anderson, 2002; Gabet et al., 2003), on hillslopes where physical weathering is weak or infrequent, the dissolution rates will be limited by the availability of unweathered material (Oliva et al., 2003).
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With the caveats given, the experiments and the mathematical analysis presented here suggest that the seepage velocity of subsurface flow may exert a strong control on chemical weathering rates. The initial hypothesis linking soil thickness and chemical weathering implied an inverse relationship between soil thickness and hydraulic conductivity (Carson and Kirkby, 1972; Gilbert, 1877). For a given climate and parent material, the hydraulic conductivity of a soil may depend on a variety of factors that may or may not be related to soil thickness. For example, vegetation density may be important because plant roots bind fine particles, suggesting that soils that support a dense stand of vegetation may have a lower Ks than less vegetated soils (Drever and Zobrist, 1992). Alternatively, macropores created by the decay of plant roots may facilitate the movement of water through soil (Devitt and Smith, 2002). Bioturbation, particularly by fossorial animals, may affect soil hydrology by preventing horizonation and by maintaining a low soil bulk density (Gabet et al., 2003). Time may influence hydraulic conductivity. Older soils may have a lower hydraulic conductivity because they have had more time to accumulate clays and other small particles that might have weathered in situ or been blown in as dust (Reheis et al., 1995; White et al., 2005). Field investigations of soil chronosequences document an inverse relationship between soil age and Ks (Brooks and Richards, 1993; Lohse and Dietrich, 2005); however, the data are limited and the studies have only been performed on low-relief surfaces. In conclusion, the results from this study suggest that soil hydrological properties should play an important role in controlling rates of chemical weathering by acting as a filter between precipitation and the underlying bedrock. Results from a physical model and a mathematical model indicate that, given a constant supply of fresh minerals, the dissolution rate of bedrock should be dependent on the hydraulic conductivity of the soil mantle. ACKNOWLEDGMENTS Supplies and Edelman’s stipend were funded by the National Science Foundation. Thorough and helpful reviews were provided by S. Mudd and two anonymous reviewers. We thank K. Mehr and S. Brantley for discussions and advice.
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