2006-2011 Mission Kearney Foundation of Soil Science ...kearney.ucdavis.edu/NEW MISSION-LIVE... · Kearney Foundation of Soil Science: Understanding and Managing Soil-Ecosystem Functions

2006-2011 Mission Kearney Foundation of Soil Science: Understanding and Managing Soil-Ecosystem

Functions Across Spatial and Temporal Scales Final Report: 2007029, 1/1/2009-12/31/2009

Department of Land, Air and Water Resources, University of California Davis, CA 95616

*Principle Investigator

For more information contact Dr. Randy Southard ([email protected])

Fog Contributions to Pedogenesis and Hydrology in Pinus muricata Ecosystems on Santa Cruz Island

Randal Southard*, Julie Baker

Project Objectives

In the Mediterranean-type climates that occur over much of California, water is severely

limited during the summer months. Several researchers have studied fog as a supplemental

water source for vegetation in coastal areas (Dawson, 1998; Ingraham and Matthews, 1995).

The importance of fog as a water source in plant and ecosystem function varies depending

upon the relative availability of fog water and other water sources (rain or groundwater). In

contrast, the importance of fog water in soil formation in these locations is not known. This

research addressed four objectives related to the role of fog in soil-forming processes on Santa

Cruz Island:

(1) Characterize soil temperature and moisture regimes, determine depth of wetting of rain and

fog, and describe soil morphology and genesis on a lithosequence under pine forest.

(2) Identify precipitation inputs from rainfall versus fog drip for soils under Bishop pine (Pinus

muricata) canopy versus grassland.

(3) Characterize organic matter decomposition rates in combinations of soil climate/ hydrology

through litterbag decomposition experiments.

(4) Characterize isotope signatures of rainfall, fog, soil solution, and pedogenic phyllosilicates.

Approach and Procedures

We studied two pairs of soils on a litho-biosequence of chlorite schist and rhyolitic tuff/breccia

on Santa Cruz Island. Each pair contained a soil on a north-facing slope that was covered by

Bishop pine (Pinus muricata) canopy, and a soil on a south-facing slope that contained grass or

shrub vegetation. Various species of arboreal vegetation (Pinus radiata, Eucalyptus sp., Sequoia

sempervirens, Erica arborea) have been shown to intercept and collect fog water, as opposed to

open grass- or shrub-covered areas where no medium for interception is present (Ingraham and

Matthews, 1995; Dawson, 1998; Prada and da Silva, 2001). The canopy-covered and open sites

were chosen to provide a comparison of soil climate and pedogenic processes within and

between parent materials.

The study sites are located at the Santa Cruz Island Reserve. Santa Cruz Island, the

largest of the northern Channel Islands, is dominated by two east-west trending mountain ranges,

and harbors a wide range of parent materials and vegetation types (Figure 1). Local climate and

weather patterns produce summer fogs created by moisture-

Fog Contributions to Pedogenesis and Hydrology in Pinus muricata Ecosystems on Santa Cruz Island—Southard

Figure 1. Location of Santa Cruz Island (inset) off the California coast, with site locations, pine coverage, and selected geology (Dibblee, 2001).


laden marine air blown onto Santa Cruz Island by the prevailing northwest winds (Junak et al.,

1995). The study sites, known as Weather Station (WS) on the Santa Cruz Island schist, a

Mesozoic chlorite schist, and Sierra Blanca (SB) on the Blanca formation, a Miocene rhyolitic

tuff/breccia (Dibblee, 2001), were located about 2 km apart. The WS site was located at about

425 m elevation, whereas the SB site was at about 275 m elevation. Locations were chosen to

minimize differences in topography between the sites (Table 1). Two pedons were excavated at

each site location, with three additional auger holes at each pine site to characterize any soil

variability related to canopy cover and determine depth of fog wetting. Each profile was

described by morphologic horizon in the field, and samples collected from each horizon, air

dried, and sieved to separate the <2 mm fraction (Soil Survey Staff, 2004). Standard x-ray

diffraction and selective dissolution procedures were performed following the methods of the

Soil Survey Staff (2004), Whittig and Allardice (1986), and Jackson (1975).

Temperature and throughfall precipitation data were used to construct a simple water

budget for each site according to the Thornthwaite model (Thornthwaite, 1948; Thornthwaite

and Mather, 1955; Thornthwaite and Mather, 1957). Potential evapotranspiration (PE) was

calculated using monthly averages of measured ambient temperature data, Ta, in the

Thornthwaite equation, PE = 1.6 (10 Ta / I)a.

A litterbag experiment was performed over a period of three years in which both pine and

grass litter were placed at two site pairs. Each pair consisted of a forested site, which received

fog drip and had an isomesic soil temperature regime (mean annual soil temperature (MAST) of

14 °C, <6 °C change between seasons), and a grass site, which did not receive fog drip and had a

thermic (MAST 19 °C) or hyperthermic (MAST 22 °C) soil temperature regime.

Table 1. Santa Cruz Island site locations.

We measured stable isotope ( D and 18

O) values of throughfall, soil solution, and

pedogenic phyllosilicates at each site. Fog events occurred dominantly during summer months,

and were generally enriched in isotope values compared to rain events, which occurred mainly


during winter months. The soil <0.2 m fraction, believed to be authigenic (Tabor et al., 2002;

Baker, 2010), was analyzed for stable 18

O isotope composition by heating and oxidation with

BrF5 following the methods of Clayton and Mayeda (1963). Hydrogen isotope ratios were

determined by reduction of the sample to hydrogen gas using a zinc catalyst and subsequent

measurement on the mass spectrometer following the methods of Kendall and Coplen (1985).

Phyllosilicate samples from the A and Bt horizons of each soil were analyzed for elemental

composition on a Cameca SX 100 microprobe (Cameca, Gennevilliers, France).

Results and Discussion

Objective 1

Soils at the SB site were classified as coarse-loamy over clayey, mixed, semiactive,

isomesic Ultic Paleustalfs under pine canopy, and fine-loamy, mixed, active, thermic Ultic

Argixerolls under grass vegetation (Baker, 2010). At the WS site, pine forest soils were coarse-

loamy, mixed, superactive, isomesic Typic Haplustalfs, and grassland soils were fine-loamy,

mixed, superactive, hyperthermic Typic Haplustalfs (Baker, 2010). Both sites were

characterized by mountainous terrain with steep slopes (30-60%) and exposed ridge tops. Slopes

with north aspects tended to have pine and oak vegetation, while south-facing slopes were

dominated by grasses and shrubs. Vegetation at the pine sites was dominated by Bishop pine

(Pinus muricata), an endemic island species, while the grass sites contained a mixture of forbs

(Eriogonum arborescens and Eriogonum grande), perennial grasses (Nassella sp.), and annual

grasses (Bromus sp., Hordeum sp., Poa sp., Vulpia sp.).

Soils formed under pine canopy tended to be deeper, have greater total clay, lower pH

and base saturation than soils formed under grass or shrub vegetation (Table 2). Higher Feo/Fed

ratios indicate that iron oxides in the pine soils may be less crystalline than in grass soils,

corresponding to less drying of pine soils due to summer fog drip. On a total profile basis, total

pedogenic iron (Fed) and Feo/Fed ratios indicate that the pine soils are more weathered than the

grass soils, and that the schist soils are more weathered that the rhyolitic tuff/breccia soils (Table

3). Chlorite was not detected in the WS soils; either chlorite has completely weathered to

vermiculite in the silt and sand fractions, and to kaolinite and smectite in the fine clay fraction, or

it was not initially present. The apparent ratio of kaolinite to smectite, based on XRD peak

intensity, is highest in the upper horizons and decreases with depth. A similar pattern is found in

WS grass soils. The SB pine soils are dominated by quartz and feldspar in the silt and sand

fractions, and kaolinite in the fine clay fraction, with smectite in the lower horizons of the pedon.

The SB grass soils are dominated by smectite in the fine clay fraction, with minor amounts of

kaolinite. Smectites in both pine soils have higher d-spacings than in the grass soils; reduced

drying due to fog inputs and lower temperatures may produce more weathered, lower charge, or

less crystalline phyllosilicates than at the grass sites. Fog appears to control soil microclimates

on the island with profound effects on soil forming processes.


Table 2. Selected morphological and chemical properties of Santa Cruz Island soils.


Table 3. Selective dissolution of selected Santa Cruz Island soils.

Objective 2

Yearly average precipitation (rain and fog drip) measured as throughfall by a rain gauge

under pine canopy from 2006-2008 at the Weather Station pine site was 52.1 cm yr-1

, while the

Sierra Blanca pine site received 38.1 cm yr-1

for the same period (Table 4). The range of

precipitation at the WS pine site was 39.1 cm (2008) to 68.2 cm (2006), while the range at the

SB pine site was 24.7 cm (2007) to 50.9 cm (2008). By comparison, the WS grass site received

estimated rainfall (no fog) averaging 43.1 cm yr-1

with a range of 31.1 cm (2008) to 62.1 cm

(2006). The SB grass site received estimated rainfall averaging 35.3 cm yr-1

for the years 2006-

2008 with a low of 20.5 cm (2007) and high of 48.0 cm (2008). Monthly throughfall means

showed high variability at both sites, with the standard deviation approaching or exceeding

monthly totals in about half the year (Table 4). Fischer et al (2009) also observed high spatial

two near-average years and two years of below-normal precipitation compared to the long-term

record (Laughrin, 2009).

At the WS pine site, data from the 2007 and 2008 summers show that fog events

routinely increase soil moisture in the surface 5 cm, and large fog events can infiltrate beyond 10

cm (Figure 2). Field observations show many very fine and fine roots at these depths at both

sites, perhaps allowing the pine trees to maximize water uptake during summer fog events. At

the SB pine site (Figure 3), the relationship between fog events and soil moisture is not as

obvious, perhaps because of the smaller recorded volumes of throughfall. During many fog

events, the lowest probe, at 25 cm, actually shows a greater increase in moisture than the middle


two probes, at 15 and 20 cm. This may be due to the change in texture with depth; the surface

textures are sandy loams, which have a lower water holding capacity than the clay loams and

clays beneath them. Although the sandier soils should have a higher water potential than the

clayier soils with the addition of a similar amount of water, the moisture probes measured

volumetric water content, not water potential. Even though the water contents measured by the

ECH2O probes at the SB site only included the top of the estimated moisture control section, we

classified the soil moisture regime as ustic rather than udic as suggested by the nearby Theta

probes, given the spatial variability of soil properties and precipitation.

Although our throughfall data shows greater precipitation at the WS site than the SB site,

both the soil moisture probes indicate that the SB soils retain water at potentials above wilting

point for longer than the WS soils. This confirms Fischer et al. (2009) finding that an elevational

belt of maximum fog precipitation exists on Santa Cruz Island from about 200-400 m. The WS

soils, at 425 m, are above this belt and may experience a decreased frequency of fog events

compared to the SB soils, although there is high spatial variability due to topography and wind

conditions. We had only one rain gauge at each site due to budget constraints, so it is possible

that throughfall was underestimated at the SB site and is actually closer to Fischer and Still’s

(2007) reported amounts for site 7.

Table 4. Monthly average throughfall (rain and fog) collected under pine canopy.

Throughfall was collected from June 2005 to July 2009.


Figure 2. Soil moisture content at the Weather Station pine site, as measured by ECH2O probes.

Figure 3. Soil moisture at the Sierra Blanca pine site, as measured by ECH2O Missing data from Jun-Oct 2008 are due to data logger malfunction.


Potential evapotranspiration (PE) at each grass site reached a maximum at about twice

the value and a minimum at about half the value of its pine pair, corresponding to the greater

temperature fluctuations at the grass sites (Figure 4). Although winter rainfall at each site pair is

equal, summer precipitation at the grass sites was zero, creating a much larger water deficit (PE-

AE). Both of the grass sites show an estimated AE approaching zero by July, while water

utilization continues throughout the dry season at the pine sites until recharge begins in

November. According to this model, differences in available water for sites only about a

hundred meters apart indicate a severe deficit of water during the summer months at the grass

sites, but the addition of intercepted fog water at the pine sites provides an additional 10 mm (SB)

to 20 mm (WS) of water per month. The grass sites, which have lower AWC than the pine soils

due to depth, also have a water surplus at the end of the wet season. Modeled water balances for

the pine soils do not show a water surplus, but the Thornthwaite model does not account for any

reduction in solar radiation caused by fog or overcast. Fischer et al. (2009) found a 29%

reduction in water deficit (PE-AE), at the WS pine site (site 10 in their study), so it is possible

that the decreased PE due to fog could create a water surplus near the end of the wet season,

which would have implications for leaching and mineral weathering.

Although the water budget models show no surplus at the pine sites, the soils under pine

canopy are deep with well-developed argillic horizons, have low pH, and lack carbonates and

soluble salts (Baker, 2010). It is possible that much of the leaching that must have occurred to

create these soils took place during a prior climate with cooler and wetter conditions. Because

our sampling period captured only average and below-average precipitation years, it is also

possible that substantial leaching occurs during above-average precipitation years. The absence

of carbonates or soluble salts, which could accumulate without leaching, seems to confirm that

some leaching is occurring under the current climate.


Figure 4a. Weather Station pine site (WSP).

Figure 4b. Sierra Blanca pine site (SBP).


Figure 4c. Weather Station grass site (WSG).

Figure 4d. Sierra Blanca grass site (SBG). Figure 4. Water balance month numbers correspond to calendar months, starting with January. Estimated available water-holding capacity used in the calculation of actual evapotranspiration was 150 mm for the pine soils and 75 mm for the grass soils.


Objective 3

Decomposition rates of both pine and grass litter were highest at the pine sites,

suggesting that moisture is the most limiting factor in this environment (Figures 5 and 6). The

additional moisture received as fog drip at the pine sites accelerates decomposition processes

even though the pine sites are cooler than the grass sites. Decomposition rates of both litter types

decreased with time at the pine sites, while rates for both litters remained relatively constant at

the grass sites over the duration of the experiment. Grass litter decomposed faster than pine litter

at the grass sites, suggesting that litter composition and/or microbial community composition and

size regulate decomposition at the grass sites. Pine litter C:N ratios decreased with time, as did

C:N ratios of grass litter at pine sites (Figures 7 and 8). The C:N ratios of grass litter at grass

sites increased initially, then began to decrease over time. The C:N ratios of both litter types

were lowest at the pine sites throughout the experiment, consistent with the higher

decomposition rates at those sites. Litterbag decomposition rates for both pine and grass litter

were similar at the pine sites, indicating that climate, and specifically available moisture, not

organic matter composition, is the most limiting factor in this environment. Mineral weathering

may be accelerated under pine canopies where moisture from fog drip creates environments

favorable for increased decomposition and biological activity. Climate change scenarios that

decrease precipitation or alter the temporal distribution of fog drip may limit nutrient cycling and

availability, while scenarios that increase precipitation may accelerate organic matter

decomposition and deplete soil C stocks.

The additional water received as fog drip by soils under pine canopies may contribute to

an increase in leaching and mineral weathering compared to grass soils. Pine forest soils are

deeper, have more total clay, and lower pH and base saturation than grass soils (Baker, 2010).

Since many of the processes of mineral weathering are accelerated by microorganisms (Kostka et

al., 1999; Reith and McPhail, 2007), it is likely that the increased biological activity at the pine

sites occurs concurrently with increased mineral weathering. Although we did not measure

respiration, increased decomposition rates at the pine sites probably result in increased

production of CO2 and carbonic acid, leading to greater rates of mineral dissolution. It has been

widely reported in the literature that increased microbial and biological activity in the

rhizosphere contributes to greater mineral weathering and dissolution rates, especially in nutrient

limited environments (Hinsinger et al., 2005; Turpault et al., 2009; Uroz et al., 2009). Higher

total iron (Fed) and aluminum (Ald), measures of intensity of weathering (McFadden and

Hendricks, 1985), observed at the pine sites (Baker, 2010) indicate that mineral weathering has

progressed to a greater degree at the pine sites compared to the grass sites. In the Channel

Islands, fog appears to be linked to mineral weathering by creating environments favorable for

increased activity by microorganisms, which then accelerate weathering processes.


Figure 5. Mean decomposition rates and standard deviations of pine litter for Study 2. WS, Weather Station site; SB, Sierra Blanca site. Litterbags were installed in March 2006. Numbers in parentheses are mean decomposition rates for each sampling date.

Figure 6. Mean decomposition rates and standard deviations of grass litter for Study 2. WS, Weather Station site; SB, Sierra Blanca site. Litterbags were installed in March 2006. Numbers in parentheses are mean decomposition rates for each sampling date.


Figure 7. Mean C:N ratios and standard deviations of pine litter for Study 2. Ratio at time=0 is the initial composition of undecomposed plant material. WSP, Weather Station Pine; WSG Weather Station Grass; SBP, Sierra Blanca Pine; SBG Sierra Blanca Grass.

Figure 8. Mean C:N ratios and standard deviations of grass litter for Study 2. Ratio at time=0 is the initial composition of undecomposed plant material. WSP, Weather Station Pine; WSG Weather Station Grass; SBP, Sierra Blanca Pine; SBG Sierra Blanca Grass.


Objective 4

Stable oxygen (Figure 9) and hydrogen (Figure 10) isotope analyses of soil solution

samples indicate a depletion of the heavier isotopes in the rainwater compared to fog. Soil

solution samples from surface horizons measured during precipitation events show a shift in

isotope ratios to heavier values during fog events and lighter values during rain events, but

generally increase in value with depth as packets of water influenced by evaporation infiltrate

deeper into the profile. A plot of 18O versus D for the soil solution samples (Figure 11)

shows a decrease in slope from the Local soil water lines (LMWL), indicating that evaporation

has occurred, in particular at the grass sites.

Extracted soil solution isotope values indicate an approximation of piston flow of water

with each precipitation event (Figures 9 and 10). For example, the March 2006 rain event

brought lighter water into about the top 10 cm of the profile, while heavier water was pushed

down deeper into the profile. Similarly, the soil water profile from the May 2006 fog event

shows heavier water near the surface of the soil, while lighter water is pushed deeper into the

profile. Although evaporation may produce similar patterns of soil solution isotope ratios

(increase in values towards the bottom of the profile during wet periods, decrease in towards

the bottom of the profile during dry periods) (Hseih et al., 1998), samples were collected during

precipitation event to minimize evaporation effects. Soil solution in open areas appears to be

heavily influenced by evaporation, while soil solution under pine canopy did not show strong

evaporation effects, and the range of isotope values measured at the pine sites is much narrower

than those measured at the grass sites. Extracted soil solution from the grass sites shows an

enriched isotope signature over fog values, and plots off the meteoric water line, indicating that

fractionation during evaporation has occurred. In particular, the soil solution isotope

composition measured after rainfall events in December and January shows enriched values at

depth at the grass sites compared to the pine sites. This water, near the soil surface during the

warm summer months, likely indicates fractionation due to evaporation, then transport to deeper

horizons during subsequent rain events.

Local soil water lines for each site were constructed from measured soil solution 18

O

and D values (Table 5). Because this study focused on the formation of pedogenic

phyllosilicates, the composition of the soil solution with which the minerals formed in

equilibrium, rather than of precipitation, likely gives a better estimation of the mineral-water

isotope relationship. Grass sites, with LSWL slopes higher than those of the pine sites, showed a

strong influence of evaporation, and pine sites showed minimal evaporation effects, so soil water

samples were separated by site to give the best estimate of meteoric water composition in

equilibrium with pedogenic phyllosilicates. Estimated temperatures of formation are similar for

both grass sites (WSG and SBG); these sites are similar in terms of soil depth, climate, and water

contributions (i.e., no fog). Although we did not measure soil solution chemistry, pedogenic

minerals may be expected to crystallize during summer months as the soil dries and solutes are

concentrated in the soil solution (Furquim et al., 2008; Ryan and Huertas, 2009). The calculated

temperatures of formation may reflect this temporal preference for mineral formation with

increased surface temperatures. In the subsurface, temperature fluctuations are smaller and

temperatures lower than at the surface during the summer.


Figure 9a. Soil solution δ18O composition with depth at Weather Station Grass site.

Figure 9b. Soil solution δ18O composition with depth at Sierra Blanca Grass site.


Figure 9c. Soil solution δ18O composition with depth at Weather Station Pine site.

Figure 9d. Soil solution δ18O composition with depth at Sierra Blanca Pine site.


Figure 10a. Soil solution δD composition at Weather Station Grass site.

Figure 10b. Soil solution δD composition at Sierra Blanca Grass site.


Figure 10c. Soil solution δD composition at Weather Station Pine site.

Figure 10d. Soil solution δD composition at Sierra Blanca Pine site.


Figure 11. Isotope values of soil solution samples at all sites. WSP, Weather Station Pine; WSG, Weather Station Grass; SBP, Sierra Blanca Pine; SBG, Sierra Blanca Grass. See Table 5 for LSWL equations for each site.

Table 5. Estimated temperatures of crystallization for smectite-kaolinite mixtures assuming

equilibrium with meteoric water. Also given are estimates of soil water δ18O values assuming

equilibrium between phyllosilicates and meteoric water in the soil profile.


At the pine sites (WSP and SBP), the estimated temperatures of formation do not follow

the pattern established at the grass sites. At the WSP site, surface horizon temperatures are lower

than subsurface temperatures, and both temperatures are higher than the WSG site, a surprising

result given the measured climatic conditions at each site. Although organic matter was removed

prior to isolation of the <0.2 m fraction, both WSP and SBP A horizons had high organic matter

contents, and it is possible that some organic matter was protected from oxidation in

microaggregates. The inclusion of this organic matter in the sample for isotope analysis may

have altered the measured D value, introducing error to the temperature estimates.

Nonexchangeable soil organic matter D values are typically more depleted than our measured

phyllosilicate values, with reported del values ranging from -244 to -180 (Seki et al, 2010) and -

167 to -75 (Ruppenthal et al, 2010). In particular, at the SBP site, which had the lowest

measured surface phyllosilicate D value, contamination by organic matter could result in a

lower estimated meteoric water composition in equilibrium with the phyllosilicates at

crystallization, underestimating the input of fog. At the SBP site, both surface and subsurface

temperatures are similar, and lower than surface temperatures at the SBG site. This may be a

reflection of the isomesic soil climate at the SBP site, where seasonal variations in soil

temperatures are minimal. Since the SBP Bt horizon sample was contaminated with excess Na,

resulting in an imbalance of Al in the calculated smectite formula, and SBP A horizon smectite

chemical composition used in the calculation of fractionation factors and temperature for the Bt

horizon, it is also possible that this substitution resulted in the similar estimated temperatures for

both horizons.

In addition, an estimation of meteoric water 18

O in equilibrium with the phyllosilicates

was calculated from temperature estimates and mixture fractionation equations, using the

relationship 1000 ln mineral-water = 18

Omineral - 18

Owater (Table 5). The estimated water 18

O

values are within the range of measured soil solution values, and show that pedogenic minerals

in A horizons formed in equilibrium with consistently depleted waters compared to minerals in

Bt horizons. This may indicate that the isotope signature of rain is overwhelming the signature

of fog, resulting in the depleted ratio near the surface of the pedon. The lower surface

temperatures associated with fog events at the pine sites may also result in lower than expected

water 18

O values, underestimating the input of fog drip. This type of distribution may also be

encountered where evaporation has enriched surface soil solution during the dry season,

followed by displacement by wet season precipitation to the subsurface. Water budget estimates

for these soils indicate that any dry season soil solution may remain in the profile in normal to

dry precipitation years (Baker, 2010). At the grass sites, this is the likely mechanism for water

infiltration into the pedon, and may indicate that phyllosilicate formation is occurring dominantly

during the wet season. The depth profiles of soil solution isotope composition at each site

indicate that pine sites are not substantially affected by evaporation, since subsurface values do

not exceed surface values collected during fog events. At the pine sites, the reversal of the

pattern of higher temperatures in the surface compared to the subsurface that is seen at the grass

sites may indicate that some phyllosilicate crystallization from fog water is occurring during the

dry season, when temperatures are higher.

The calculated temperatures of mineral formation and associated meteoric waters in

equilibrium with the soil minerals indicate that the contribution of fog to mineral formation in

the Channel Islands may not be sufficient to distinguish a difference in isotope signatures of

pedogenic phyllosilicates formed in equilibrium with rain or fog water. However, the higher


estimated temperature of formation in the pine surface horizons may indicate some fog

influence. These soils contained a mixture of 1:1 and 2:1 phyllosilicates; if precipitation of

these minerals is temporally separated, the seasonal variation in isotope composition of the

precipitation may result in phyllosilicates with fractionations other than anticipated (Stern et

al., 1997). Nonetheless, pedogenic phyllosilicates appear to be capable of recording

systematic differences in temperature and meteoric water composition between surface and

subsurface soil horizons. For future study, a refinement of this technique may be useful for

detecting seasonal variation in phyllosilicate formation corresponding to meteoric water and

soil solution composition.

Conclusions

Soil moisture sensors, water balance models, and isotope signatures of precipitation

(Baker, 2010) indicate that water from fog drip infiltrates into the pedon. The fog water,

available at a time when soil temperatures are at a maximum, also contributes to increased

organic matter decomposition rates (Baker, 2010), possibly accelerating nutrient turnover and

mineral weathering and contributing to soil fertility. Historically, it is likely that the combination

of pine forest and foggy climate have aided in the formation of deeper, clayier soils than the

adjacent grasslands. Increased water and increased organic and carbonic acid production may

accelerate mineral weathering, leading to the well-developed argillic horizons observed in the

soils under pine canopy. While a change in weather patterns is unlikely to alter the soils already

in place over the short term, organic matter decomposition and nutrient cycling rates may

decrease if fog drip is reduced or shifts in temporal distribution, also threatening the pine forest

stands. It is clear that fog drip is an important part of the hydrologic cycle of the island pine

stands, but fog may also affect soil processes in less direct ways that are equally important to

pine tree growth and survival.

References

Baker, J.B. 2010. Fog contributions to pedogenesis and hydrology in Pinus muricata ecosystems

on Santa Cruz Island. Ph.D. diss. Univ. of California, Davis.

Dawson, T.E. 1998. Fog in the California redwood forest: ecosystem inputs and use by plants.

Oecologia 117:476-485.

Dibblee Jr., T.W. 2001. Geologic Map of Western Santa Cruz Island. Dibblee Geological

Foundation, Santa Barbara, CA.

Fischer, D.T., and C.J. Still. 2007. Evaluating patterns of fog water deposition and isotopic

composition on the California Channel Islands. Water Resour. Res. 43:W04420.

Fischer, D.T., C.J. Still, and A.P Williams. 2009. Significance of summer fog and overcast for

drought stress and ecological functioning of coastal California endemic plant species. J.

Biogeogr. 36:783-799.

Furquim, S.A.C., R.C. Graham, L. Barbiero, J.P.D. Neto, and V. Valles. 2008. Mineralogy and

genesis of smectites in an alkaline-saline environment of Pantanal wetland, Brazil. Clays

Clay Minerals 56:579-595.

Hinsinger, P., C. Plassard, and B. Jaillard. 2005. Rhizosphere: a new frontier for soil

biogeochemistry. J. Geochem. Explor. 88:210-213.


Hseih, J.C.C., O.A. Chadwick, E.F. Kelly, and S.M. Savin. 1998. Oxygen isotopic composition

of soil water: Quantifying evaporation and transpiration. Geoderma 82:269-293.

Ingraham, N.L. and R.A. Matthews. 1995. The importance of fog-drip water to vegetation: Point

Reyes Peninsula, California. J. Hydrology 164:269-285.

Jackson, M.L. 1975. Soil chemical analysis--advanced course. 2nd Ed. Published by the author,

Madison, WI.

Junak, S. T. Ayers, R. Scott, D. Wilken, and D. Young. 1995. A Flora of Santa Cruz Island.

Santa Barbara Botanic Garden: Santa Barbara, CA.

Kendall, C. and T.B. Coplen. 1985. Multisample conversion of water to hydrogen by zinc for

stable isotope determination. Anal. Chem. 57:1437-1440.

Kostka, J.E., J. Wu, K.H. Nealson, and J.W. Stucki. 1999. The impact of structural Fe(III)

reduction by bacteria on the surface chemistry of smectite clay minerals. Geochim.

Cosmochim. Acta. 63:3705-3713.

Laughrin, L. 2009. Santa Cruz Island Precipitation- Central Valley- 1904-2008. Santa Cruz

Island Reserve, University of California Natural Reserve System.

McFadden, L.D. and D.M. Hendricks. 1985. Changes in content and composition of pedogenic

iron oxyhydroxides in a chronosequence of soils in southern California. Quat. Res.

23:189-204.

Prada, S.N. and M.O. da Silva. 2001. Fog precipitation on the Island of Madeira (Portugal).

Environmental Geology 41:384-389.

Reith, F. and D.C. McPhail. 2007. Mobility and microbially mediated mobilization of gold and

arsenic in soils from two gold mines in semi-arid and tropical Australia. Geochim.

Cosmochim. Acta 71:1183-1196.

Ruppenthal, M., Y. Oelmann, and W. Wilcke. 2010. Isotope ratios of nonexchangeable hydrogen

in soils from different climate zones. Geoderma 155:231-241.

Ryan, P.C. and F.J. Huertas. 2009. The temporal evolution of pedogenic Fe-smectite to Fe-kaolin

via interstratified kaolin-smectite in a moist tropical soil chronosequence. Geoderma

151:1-15.

Seki, O., T. Nakasuka, H. Shibata, and K. Kawamura. 2010. A compound-specific n-alkane delta

C-13 and delta D approach for assessing source and delivery processes of terrestrial

organic matter within a forested watershed in northern Japan. Geochim. Cosmochim. Acta

74:599-613.

Soil Survey Staff. 2004. Soil survey laboratory methods manual. Soil Survey Investigations

Report No. 42. USDA-NRCS National Soil Survey Center, Lincoln, NE.

Tabor, N.J. and I.P. Montañez. 2002. Shifts in late Paleozoic atmospheric circulation over

western equatorial Pangaea: Insights from pedogenic mineral delta O-18 compositions.

Geology 30:1127-1130.

Thornthwaite, C.W. 1948. An approach toward a rational classification of climate. Geogr. Rev.

38:55-94.

Thornthwaite, C.W. and J.R. Mather. 1955. The water balance. Drexel Institute of Technology

Laboratory of Climatology: Centerton, NJ.

Thornthwaite, C.W. and J.R. Mather. 1957. Instructions and tables for computing potential

evapotranspiration and the water balance. Drexel Institute of Technology Laboratory of

Climatology: Centerton, NJ.

Turpault, M.P., C. Nys, and C. Calvaruso. 2009. Rhizosphere impact on the dissolution of test

minerals in a forest ecosystem. Geoderma 153:147-154.


Uroz, S., C. Calvaruso, M.P. Turpault, A. Sarniguet, W. de Boer, J.H.J. Leveau, and P. Frey-

Klett. 2009. Efficient mineral weathering is a functional trait of the bacterial genus

Collimonas. Soil Biol. Biochem. 41:2178-2186.

Whittig, L.D. and W.R. Allardice. 1986. X-ray diffraction techniques. p. 331-362. In A. Klute

(ed.) Methods of Soil Analysis, Part 1. Physical and Mineralogical Methods. 2nd Ed.

Agronomy Monograph no. 9. ASA-SSSA: Madison, WI.

This research was funded by the Kearney Foundation of Soil Science: Understanding and Managing Soil-Ecosystem Functions Across Spatial and Temporal Scales, 2006-2011 Mission (http://kearney.ucdavis.edu). The Kearney Foundation is an endowed research program created to encourage and support research in the fields of soil, plant nutrition, and water science within the Division of Agriculture and Natural Resources of the University of California

http://kearney.ucdavis.edu/

2006-2011 Mission Kearney Foundation of Soil Science ...kearney.ucdavis.edu/NEW MISSION-LIVE... · Kearney Foundation of Soil Science: Understanding and Managing Soil-Ecosystem Functions

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