Infiltration is the movement of water into the soil. This is possible,
because soil is not solid matter; instead it is a porous medium
comprising a matrix of solid granular particles and voids that may be
filled with air or water (Figure 21). Flow in a porous medium may be
unsaturated when some of the voids are occupied by air, or saturated
when all the voids are occupied by water. Considering the cross
section of a porous medium illustrated in Figure 21, the porosity is
defined as
total volume
n = volume of voids
The range of n for soils is approximately 0.25 to 0.75 depending
upon the soil texture. A part of the voids is occupied by water, and
the remainder by air. The volume occupied by water being measured
by the volumetric soil moisture content is defined as
total volume
θ = volume of water (7)
Hence 0 ≤ θ ≤ n; the soil moisture content is equal to the porosity
when the soil is saturated. Soil moisture content is also sometimes
characterized by the degree of saturation, defined as
Sd = θ/n (8)
The degree of saturation varies between 0 and 1.
average density of the mineral grains making up a soil
ρm= Mm/Vm (9)
where Mm is the mass and Vm the volume of the mineral grains. The
value of ρm is rarely measured, but is estimated based on the mineral
composition of the soil. A value of 2650 kg/m3, which is the density
of the mineral quartz, is often assumed. The bulk density, ρb, is the
dry density of the soil
where Vs is the total volume of the soil sample which is the sum of
the volume of the air, Va, liquid water, Vw, and mineral components,
Vm, of the soil respectively. In practice, bulk density is defined as the
mass of a volume of soil that has been dried for an extended period
(16 hr or longer) at 105 oC, divided by the original volume. The
porosity (6) is given by
Here Mswet and Msdry are the masses before and after drying,
respectively, and ρw is the density of water (1000 kg/m3). This
method for determining soil moisture is referred to as the gravimetric
method. In the field moisture content can be measured in a number
of other ways. Electrical resistance blocks use the inverse
relationship between water content and the electrical resistance of a
volume of porous material (e.g. gypsum, nylon or fiberglass) in
equilibrium with the soil. Neutron probe moisture meters are
combined sources and detectors of neutrons that are inserted into
Sunday, 1 January 2017
SOIL PROPERTIES
Hydrological Pathways involved in different runoff
generation processes. Infiltration excess pathways are shown in
red. Saturation excess and subsurface stormflow pathways are
shown in blue. Groundwater and baseflow pathways in black and
Evapotranspiration is green. (Courtesy of Mike Kirkby)
The general climatic regime controls the total volume of runoff in
any region through its effect on the water balance. In a broad sense,
over a time scale long enough that storage changes average out (are
negligible), and over a region large enough or with boundary defined
so that inflows (surface and subsurface) are negligible, the water
balance may be stated as
P = Q + E (1)
where P is the precipitation rate, Q the runoff rate, and E the
evapotranspiration rate. This equation indicates that the precipitation
input is disposed of either into runoff or evapotranspiration. In
general the climatic regime controls the overall proportioning. Here
groundwater recharge supplying baseflow is included in Q. Because
the quantities in equation (1) must be positive, this equation places
limits on the values of Q and E given any specific P. Both Q and E
are constrained to be less than P. This may be visualized in a
space
The domain of valid solutions is below the 1:1 line E=P. There is in
general an upper limit on the possible evapotranspiration, due to the
energy inputs required to evaporate water. This limit is related to the
solar radiation inputs as well as the capacity of the atmosphere to
transport evaporated water away from the surface (related to wind
and humidity). This limit has been denoted as Ep (potential
evapotranspiration) in Figure 14. The line E=Ep provides another
upper limit to the domain of valid solutions
precipitation is large (P >> Ep) water is not going to be limited at the
earth surface so E will approach Ep asymptotically for P tending to
infinity. Also, if precipitation is small (P << Ep) water is very limited
at the earth surface and may all evaporate, so E will approach the line
E=P asymptotically as P tends to 0 approaching the origin. These
constraints suggest a solution for the E versus P function of the form
indicated in Figure 14. In Figure 14 the axes have been scaled by Ep
to make them dimensionless. A nonlinear increase in Q with P as
P/Ep increases is suggested, and the index P/Ep serves as an index
of regional humidity or aridity, with P/Ep large (>1) in humid
regions and small (<1) in arid regions
Generalised dependence of Runoff Coefficient and
Style of Overland Flow on Arid-Humid scale and on Storm Rainfall
Intensities (Courtesy of Mike Kirkby)
These regional water balance considerations based on climatic regime
serve as first order controls on the generation of runoff. However
data plotted in the form of Figure 14, shows scatter due to other
effects. Precipitation intensity is also important. Figure 15 shows the
interplay between humidity/aridity and precipitation intensity on
runoff processes and the runoff coefficient. In this figure, similar to
P/Ep the ratio E/Ep serves as a measure of aridity, with E/Ep
generation processes. Infiltration excess pathways are shown in
red. Saturation excess and subsurface stormflow pathways are
shown in blue. Groundwater and baseflow pathways in black and
Evapotranspiration is green. (Courtesy of Mike Kirkby)
The general climatic regime controls the total volume of runoff in
any region through its effect on the water balance. In a broad sense,
over a time scale long enough that storage changes average out (are
negligible), and over a region large enough or with boundary defined
so that inflows (surface and subsurface) are negligible, the water
balance may be stated as
P = Q + E (1)
where P is the precipitation rate, Q the runoff rate, and E the
evapotranspiration rate. This equation indicates that the precipitation
input is disposed of either into runoff or evapotranspiration. In
general the climatic regime controls the overall proportioning. Here
groundwater recharge supplying baseflow is included in Q. Because
the quantities in equation (1) must be positive, this equation places
limits on the values of Q and E given any specific P. Both Q and E
are constrained to be less than P. This may be visualized in a
space
The domain of valid solutions is below the 1:1 line E=P. There is in
general an upper limit on the possible evapotranspiration, due to the
energy inputs required to evaporate water. This limit is related to the
solar radiation inputs as well as the capacity of the atmosphere to
transport evaporated water away from the surface (related to wind
and humidity). This limit has been denoted as Ep (potential
evapotranspiration) in Figure 14. The line E=Ep provides another
upper limit to the domain of valid solutions
precipitation is large (P >> Ep) water is not going to be limited at the
earth surface so E will approach Ep asymptotically for P tending to
infinity. Also, if precipitation is small (P << Ep) water is very limited
at the earth surface and may all evaporate, so E will approach the line
E=P asymptotically as P tends to 0 approaching the origin. These
constraints suggest a solution for the E versus P function of the form
indicated in Figure 14. In Figure 14 the axes have been scaled by Ep
to make them dimensionless. A nonlinear increase in Q with P as
P/Ep increases is suggested, and the index P/Ep serves as an index
of regional humidity or aridity, with P/Ep large (>1) in humid
regions and small (<1) in arid regions
Generalised dependence of Runoff Coefficient and
Style of Overland Flow on Arid-Humid scale and on Storm Rainfall
Intensities (Courtesy of Mike Kirkby)
These regional water balance considerations based on climatic regime
serve as first order controls on the generation of runoff. However
data plotted in the form of Figure 14, shows scatter due to other
effects. Precipitation intensity is also important. Figure 15 shows the
interplay between humidity/aridity and precipitation intensity on
runoff processes and the runoff coefficient. In this figure, similar to
P/Ep the ratio E/Ep serves as a measure of aridity, with E/Ep
Surface Water Input
addition of a small amount of water can saturate this soil and cause
the water table to rise quite rapidly, resulting in subsurface stormflow,
surface saturation and saturation excess overland flow. The moisture
content in the capillary fringe can also be affected by the history of
wetting and drying of the soil, a phenomenon known as hysteresis.
When soil has been draining the moisture content tends to remain
above what it would be if it were filling at the same pressure. The
addition of a small amount of water can switch the soil from draining
to filling mode, enhancing the effect of the capillary fringe on the rise
of the water table and subsurface stormflow response. The capillary
fringe and hysteresis are discussed in more detail in Chapter 4.
The discussion thus far has focused on the main processes involved
in runoff generation on a hillslope. To complete the discussion on
runoff generation processes it is necessary to mention briefly some
other processes and factors involved. Interception of precipitation
by vegetation can play a significant role in reducing runoff, especially
in forested environments. Much intercepted water is eventually
evaporated back to the atmosphere (Figure 1). In some hydrologic
models, interception is sometimes modeled as an initial abstraction that
is subtracted from precipitation inputs before they are used in
infiltration or runoff calculations. In other hydrologic models
detailed representations of the interception, storage of water in the
canopy, throughfall or stem flow are used (e.g. Rutter et al., 1972).
Direct precipitation onto a stream or water body also contributes to
runoff as indicated in Figure 6. This is important in areas where the
water surface is extensive, as with lakes, reservoirs and floodplains
that are flooded, because in these situations runoff generation is not
delayed by the usual hillslope processes.
The freezing state of the soil, in regions where freezing occurs, also
plays a role in runoff generation. Infiltration capacity is reduced due
to frozen ground, depending upon the soil moisture content at the
time of freezing.
Fire results in water repellency by soils which reduces infiltration
capacity. One cause for water repellency is chemicals released during
a fire that are absorbed in the soil, and can make it water repellent for
months to years following a fire. The heat from fire also removes the
thin films of irreducable water adhered to soil particles by capillary
forces, disconnecting potential flow paths. Penetration of water into
macropores following a fire is limited due to this effect. High
temperatures in deserts have the same effect, adding to the tendency
for infiltration capacities to be lower in arid regions making them
more subject to infiltration excess runoff generation processes. This
water repellency due to fire has been implicated in many floods
following severe bush or forest fires.
Many of the runoff generation processes described depend on the
soil moisture status of the soil. This is referred to as the antecedent
conditions. Between storms (surface water input events), processes of
evaporation, transpiration, percolation and drainage serve to set up
the soil moisture antecedent conditions. Runoff generation
mechanisms and processes therefore depend not only on conditions
during storms, but conditions in advance of storms and a complete
understanding or representation of all the land surface hydrologic
processes is required to quantify the generation of runoff.
Recognition of this has led to the development of continuous
simulation models, such as the National Weather Service Sacramento
soil moisture accounting model that keeps continuous track of the
state of different soil moisture components for the modeling of
runoff. Detailed presentation of these models is beyond the scope of
this module, although key ideas are reviewed at the end of this model.
The discussion above has reviewed, in a conceptual way many of the
processes and mechanisms involved in runoff generation. These can
be quite complex, and when efforts are made to perform quantitative
calculations the devil is in the details. Each watershed or hillslope is
different, with different topography, soils and physical properties.
The challenge for hydrologic modelers is to balance practical
simplifications with justifiable model complexity and the knowledge
that many specific physical properties required for detailed hydrologic
modeling are physically unknowable. Our understanding of runoff
generation involves the movement of water through soil pores and
macropores. These flows follow the physical laws governing fluid
flow (Navier Stokes equations) but we can never know in sufficient
detail the flow geometry to make use of fluid flow theory and
ultimately have to resort to simplifications or parameterizations of
the runoff generation processes. In the remainder of this module the
astute reader will note discrepancies between the physical
understanding given above and mathematical descriptions used to
perform practical calculations. The mathematical descriptions,
although frequently complex, incorporate significant simplifications
relative to the field based conceptual understanding of how runoff
processes work. This gap between field based and model based
representations makes the subject of rainfall – runoff processes a
fertile area for research to learn how to better model rainfall runoff
processes.
Figure 13 summarizes the main processes involved in runoff
generation, showing the interaction between infiltration excess,
saturation excess and groundwater flow pathways. Most rainfall
runoff models are organized around a representation similar to
Figure 13 involving partition of surface water input into infiltration
or overland flow, either due to infiltration excess or saturation excess.
Infiltrated water enters the soil regolith where it contributes to
interflow, percolates to deeper groundwater or is evaporated or
transpired back to the atmosphere. The quantity of water in the soil
affects the variable source area involved in the generation of
saturation overland flow. The deeper groundwater contributes to
baseflow and affects interflow through groundwater rise
the water table to rise quite rapidly, resulting in subsurface stormflow,
surface saturation and saturation excess overland flow. The moisture
content in the capillary fringe can also be affected by the history of
wetting and drying of the soil, a phenomenon known as hysteresis.
When soil has been draining the moisture content tends to remain
above what it would be if it were filling at the same pressure. The
addition of a small amount of water can switch the soil from draining
to filling mode, enhancing the effect of the capillary fringe on the rise
of the water table and subsurface stormflow response. The capillary
fringe and hysteresis are discussed in more detail in Chapter 4.
The discussion thus far has focused on the main processes involved
in runoff generation on a hillslope. To complete the discussion on
runoff generation processes it is necessary to mention briefly some
other processes and factors involved. Interception of precipitation
by vegetation can play a significant role in reducing runoff, especially
in forested environments. Much intercepted water is eventually
evaporated back to the atmosphere (Figure 1). In some hydrologic
models, interception is sometimes modeled as an initial abstraction that
is subtracted from precipitation inputs before they are used in
infiltration or runoff calculations. In other hydrologic models
detailed representations of the interception, storage of water in the
canopy, throughfall or stem flow are used (e.g. Rutter et al., 1972).
Direct precipitation onto a stream or water body also contributes to
runoff as indicated in Figure 6. This is important in areas where the
water surface is extensive, as with lakes, reservoirs and floodplains
that are flooded, because in these situations runoff generation is not
delayed by the usual hillslope processes.
The freezing state of the soil, in regions where freezing occurs, also
plays a role in runoff generation. Infiltration capacity is reduced due
to frozen ground, depending upon the soil moisture content at the
time of freezing.
Fire results in water repellency by soils which reduces infiltration
capacity. One cause for water repellency is chemicals released during
a fire that are absorbed in the soil, and can make it water repellent for
months to years following a fire. The heat from fire also removes the
thin films of irreducable water adhered to soil particles by capillary
forces, disconnecting potential flow paths. Penetration of water into
macropores following a fire is limited due to this effect. High
temperatures in deserts have the same effect, adding to the tendency
for infiltration capacities to be lower in arid regions making them
more subject to infiltration excess runoff generation processes. This
water repellency due to fire has been implicated in many floods
following severe bush or forest fires.
Many of the runoff generation processes described depend on the
soil moisture status of the soil. This is referred to as the antecedent
conditions. Between storms (surface water input events), processes of
evaporation, transpiration, percolation and drainage serve to set up
the soil moisture antecedent conditions. Runoff generation
mechanisms and processes therefore depend not only on conditions
during storms, but conditions in advance of storms and a complete
understanding or representation of all the land surface hydrologic
processes is required to quantify the generation of runoff.
Recognition of this has led to the development of continuous
simulation models, such as the National Weather Service Sacramento
soil moisture accounting model that keeps continuous track of the
state of different soil moisture components for the modeling of
runoff. Detailed presentation of these models is beyond the scope of
this module, although key ideas are reviewed at the end of this model.
The discussion above has reviewed, in a conceptual way many of the
processes and mechanisms involved in runoff generation. These can
be quite complex, and when efforts are made to perform quantitative
calculations the devil is in the details. Each watershed or hillslope is
different, with different topography, soils and physical properties.
The challenge for hydrologic modelers is to balance practical
simplifications with justifiable model complexity and the knowledge
that many specific physical properties required for detailed hydrologic
modeling are physically unknowable. Our understanding of runoff
generation involves the movement of water through soil pores and
macropores. These flows follow the physical laws governing fluid
flow (Navier Stokes equations) but we can never know in sufficient
detail the flow geometry to make use of fluid flow theory and
ultimately have to resort to simplifications or parameterizations of
the runoff generation processes. In the remainder of this module the
astute reader will note discrepancies between the physical
understanding given above and mathematical descriptions used to
perform practical calculations. The mathematical descriptions,
although frequently complex, incorporate significant simplifications
relative to the field based conceptual understanding of how runoff
processes work. This gap between field based and model based
representations makes the subject of rainfall – runoff processes a
fertile area for research to learn how to better model rainfall runoff
processes.
Figure 13 summarizes the main processes involved in runoff
generation, showing the interaction between infiltration excess,
saturation excess and groundwater flow pathways. Most rainfall
runoff models are organized around a representation similar to
Figure 13 involving partition of surface water input into infiltration
or overland flow, either due to infiltration excess or saturation excess.
Infiltrated water enters the soil regolith where it contributes to
interflow, percolates to deeper groundwater or is evaporated or
transpired back to the atmosphere. The quantity of water in the soil
affects the variable source area involved in the generation of
saturation overland flow. The deeper groundwater contributes to
baseflow and affects interflow through groundwater rise
lateral flow
Lateral flow at the soil bedrock interface (Weiler and McDonnell,
2003) illustrated in Figure 11, occurs in steep terrain with relatively
thin soil cover and low permeability bedrock, where water moves to
depth rapidly along preferential infiltration pathways and perches at
the soil-bedrock interface. Since moisture content near the bedrock
interface is often close to saturated, the addition of only a small
amount of new water (rainfall or snowmelt) is required to produce
saturation at the soil-bedrock or soil-impeding layer interface. Rapid
lateral flow occurs at the permeability interface through the transient
saturated zone. Once rainfall inputs cease, there is a rapid dissipation
of positive pore water pressures and the system reverts back to a slow
drainage of matrix flow.
The processes involved in the generation of subsurface stormflow by
groundwater ridging are illustrated in Figure 12. An idealized cross
section of a valley with a straight hillslope is shown. In a simplified
situation with uniform soils the water table has an approximately
parabolic form, and soil moisture content decreases with increasing
height above the water table. The shaded areas represent graphs of
soil moisture at the base, middle and near the top of the hillslope (a)
before the onset of rainfall; (b) as an initial response to rainfall; and
(c) after continuing rainfall. Because (in a) before the onset of water
input the water table slopes gently towards the channel there will be a
slow flow of groundwater to maintain the baseflow of the stream.
With the onset of surface water input, water that infiltrates near the
base of the hillslope will quickly reach the water table and cause the
water table near the stream to rise, early in a storm. Further upslope
the soil is dryer and distance to the water table greater. It therefore
takes longer for infiltrating water to reach the water table and where
the water table is deep all the infiltrating water may go into storage in
the unsaturated zone and not reach the water table for many days
after the storm. The initial response to water input is therefore as
depicted in Figure 12b, where the water table has risen near the
stream but remained unchanged further upslope. The rising water
table near the stream causes an increase in the hydraulic gradient
between the groundwater and stream, and increased subsurface flow
into the stream results. This is subsurface stormflow, and is
frequently seen to be groundwater that has been displaced by the
infiltrating water, and is thus old or pre-storm water bearing the
chemical and isotopic signature of water in the hillslope prior to the
storm, which may be different from the chemical and isotopic
signature of overland flow from rainwater that has not infiltrated.
Measurement of chemical and isotopic signatures of stream water,
ground water and rain water is commonly used in hydrology as a way
of inferring hillslope flow pathways. After continuing rain (Figure
12c), the water table has risen to the surface over the lower part of
the hillslope and the saturated area is expanding uphill. Some water
emerges from this saturated area and runs down slope to the stream.
This is termed return flow. Direct precipitation onto the saturated
zone (DPS) forms saturation excess runoff as described above
Figure 12. Groundwater ridging subsurface stormflow processes
in an area of high infiltration rate. (redrawn following Water in
Environmental Planning, Dunne and Leopold, 1978)
Figure 12 illustrates a region just above the water table that was close
to saturation. This is known as the capillary fringe, and can play an
important role in runoff generation in certain situations. Capillary
forces due to the surface tension between water and soil particles act
to pull water into the soil matrix above the water table and maintain
the capillary fringe at moisture content very close to saturation.
2003) illustrated in Figure 11, occurs in steep terrain with relatively
thin soil cover and low permeability bedrock, where water moves to
depth rapidly along preferential infiltration pathways and perches at
the soil-bedrock interface. Since moisture content near the bedrock
interface is often close to saturated, the addition of only a small
amount of new water (rainfall or snowmelt) is required to produce
saturation at the soil-bedrock or soil-impeding layer interface. Rapid
lateral flow occurs at the permeability interface through the transient
saturated zone. Once rainfall inputs cease, there is a rapid dissipation
of positive pore water pressures and the system reverts back to a slow
drainage of matrix flow.
The processes involved in the generation of subsurface stormflow by
groundwater ridging are illustrated in Figure 12. An idealized cross
section of a valley with a straight hillslope is shown. In a simplified
situation with uniform soils the water table has an approximately
parabolic form, and soil moisture content decreases with increasing
height above the water table. The shaded areas represent graphs of
soil moisture at the base, middle and near the top of the hillslope (a)
before the onset of rainfall; (b) as an initial response to rainfall; and
(c) after continuing rainfall. Because (in a) before the onset of water
input the water table slopes gently towards the channel there will be a
slow flow of groundwater to maintain the baseflow of the stream.
With the onset of surface water input, water that infiltrates near the
base of the hillslope will quickly reach the water table and cause the
water table near the stream to rise, early in a storm. Further upslope
the soil is dryer and distance to the water table greater. It therefore
takes longer for infiltrating water to reach the water table and where
the water table is deep all the infiltrating water may go into storage in
the unsaturated zone and not reach the water table for many days
after the storm. The initial response to water input is therefore as
depicted in Figure 12b, where the water table has risen near the
stream but remained unchanged further upslope. The rising water
table near the stream causes an increase in the hydraulic gradient
between the groundwater and stream, and increased subsurface flow
into the stream results. This is subsurface stormflow, and is
frequently seen to be groundwater that has been displaced by the
infiltrating water, and is thus old or pre-storm water bearing the
chemical and isotopic signature of water in the hillslope prior to the
storm, which may be different from the chemical and isotopic
signature of overland flow from rainwater that has not infiltrated.
Measurement of chemical and isotopic signatures of stream water,
ground water and rain water is commonly used in hydrology as a way
of inferring hillslope flow pathways. After continuing rain (Figure
12c), the water table has risen to the surface over the lower part of
the hillslope and the saturated area is expanding uphill. Some water
emerges from this saturated area and runs down slope to the stream.
This is termed return flow. Direct precipitation onto the saturated
zone (DPS) forms saturation excess runoff as described above
Figure 12. Groundwater ridging subsurface stormflow processes
in an area of high infiltration rate. (redrawn following Water in
Environmental Planning, Dunne and Leopold, 1978)
Figure 12 illustrates a region just above the water table that was close
to saturation. This is known as the capillary fringe, and can play an
important role in runoff generation in certain situations. Capillary
forces due to the surface tension between water and soil particles act
to pull water into the soil matrix above the water table and maintain
the capillary fringe at moisture content very close to saturation.
climate
The general climatic regime controls the total volume of runoff in
any region through its effect on the water balance. In a broad sense,
over a time scale long enough that storage changes average out (are
negligible), and over a region large enough or with boundary defined
so that inflows (surface and subsurface) are negligible, the water
balance may be stated as
P = Q + E (1)
where P is the precipitation rate, Q the runoff rate, and E the
evapotranspiration rate. This equation indicates that the precipitation
input is disposed of either into runoff or evapotranspiration. In
general the climatic regime controls the overall proportioning. Here
groundwater recharge supplying baseflow is included in Q. Because
the quantities in equation (1) must be positive, this equation places
limits on the values of Q and E given any specific P. Both Q and E
are constrained to be less than P. This may be visualized in a space
where E is plotted versus P
Thank you to Christina Bandaragoda and Yasir Kaheil, graduate
students at Utah State University who helped tremendously with
preparation of online material. Thank you to Mark Zachry and
Christine Hult in the English Department at Utah State University
for help on the pedagogy of online education. This module was
developed using database technology by 3GB Technologies
(http://www.3gb.com). David Brandon and Bill Reed at the
National Weather Service Colorado Basin River Forecast Center
collaborated on the production of this module.
This material was prepared by the Utah State University, Utah Water
Research Laboratory under a Subaward with the University
Corporation for Atmospheric Research (UCAR) under Cooperative
Agreement No. NA17WD2383 with the National Oceanic and
Atmospheric Administration (NOAA), U.S. Department of
Commerce (DOC). The statements, findings, conclusions, and
recommendations are those of the author and do not necessarily
reflect the views of NOAA, DOC or UCAR.
any region through its effect on the water balance. In a broad sense,
over a time scale long enough that storage changes average out (are
negligible), and over a region large enough or with boundary defined
so that inflows (surface and subsurface) are negligible, the water
balance may be stated as
P = Q + E (1)
where P is the precipitation rate, Q the runoff rate, and E the
evapotranspiration rate. This equation indicates that the precipitation
input is disposed of either into runoff or evapotranspiration. In
general the climatic regime controls the overall proportioning. Here
groundwater recharge supplying baseflow is included in Q. Because
the quantities in equation (1) must be positive, this equation places
limits on the values of Q and E given any specific P. Both Q and E
are constrained to be less than P. This may be visualized in a space
where E is plotted versus P
Thank you to Christina Bandaragoda and Yasir Kaheil, graduate
students at Utah State University who helped tremendously with
preparation of online material. Thank you to Mark Zachry and
Christine Hult in the English Department at Utah State University
for help on the pedagogy of online education. This module was
developed using database technology by 3GB Technologies
(http://www.3gb.com). David Brandon and Bill Reed at the
National Weather Service Colorado Basin River Forecast Center
collaborated on the production of this module.
This material was prepared by the Utah State University, Utah Water
Research Laboratory under a Subaward with the University
Corporation for Atmospheric Research (UCAR) under Cooperative
Agreement No. NA17WD2383 with the National Oceanic and
Atmospheric Administration (NOAA), U.S. Department of
Commerce (DOC). The statements, findings, conclusions, and
recommendations are those of the author and do not necessarily
reflect the views of NOAA, DOC or UCAR.
Soil Temperature Regime;
Soil
Temperature Regime;
In
soil taxonomy, the soil temperature regime is one of the ways of classifying
soils based on their mean annual temperatures.
Mean atmospheric temperatures are obtained to
estimate soil temperature regime (STR). Mean annual air temperature, mean
summer and mean winter temperatures have to be determined and these are
ultimately used to obtain Mean Annual
Soil Temperature (MAST).
The following
are the classes of temperature regimes;
·
Pergelic
STR:
The pergelic
soil temperature regime has mean annual soil temperatures of less than 0 °C at
50 cm below the surface. In this temperature regime, permafrost is
present.
·
Crylic
STR:
The cryic soil
temperature regime has mean annual soil temperatures of greater than 0 °C, but
less than 8 °C, with a difference between mean summer and mean winter soil
temperatures greater than 5 °C at 50 cm, and cold summer temperatures.
·
Frigid
STR:
The frigid soil
temperature regime has mean annual soil temperatures of greater than 0 °C, but
less than 8 °C, with a difference between mean summer and mean winter soil
temperatures greater than 5 °C at 50 cm below the surface, and warm
summer temperatures.
·
Mesic
STR:
The mesic soil
temperature regime has mean annual soil temperatures of 8 °C or more, but less
than 15 °C, and the difference between mean summer
and mean winter soil temperatures is greater than
5 °C at 50 cm below the surface.
·
Thermic
SRT:
The thermic soil
temperature regime has mean annual soil temperatures of 15° C or more, but less
than 22 °C; and a difference between mean summer and mean winter soil
temperatures of greater than 5 °C at 50 cm below the surface.
·
Hyperthermic
STR:
The hyperthermic
soil temperature regime has mean annual soil temperatures of 22 °C or more and
a difference between mean summer and mean winter soil temperatures of less than
5 °C at 50 cm below the surface.
·
Isofrigid
STR:
The isofrigid
soil temperature regime has mean annual soil temperatures of greater than
0 °C, but less than 8 °C, with a difference between mean summer and mean winter
soil temperatures of less than 5 °C at 50 cm. below the surface, and warm
summer temperatures.
·
Isohyperthemic
STR:
The
isohyperthermic soil temperature regime has mean annual soil temperatures of 22
°C or more and a difference between mean summer and mean winter soil
temperatures of less than 5 °C at 50 cm below the surface.
·
Isomesic STR:
The isomesic
soil temperature regime has a mean annual soil temperatures of 8 °C or more,
but a difference between mean summer and mean winter soil temperatures of less
than 5 °C at 50 cm below the
surface.
·
Isothermal
STR:
The isothermic
soil temperature regime that has mean annual soil temperatures of 15 °C or more
but, 5 °C difference between mean summer and mean winter soil temperatures at
50 cm. below the surface.
Importance of
Soil Temperature in Crop Production.
·
Microbiological
activities:
Extreme low and high temperature influence the
soil microbial population and rate of organic matter decompositionFriday, 23 December 2016
DO NOT PASS WITHOUT READING THIS
EHAT EACTLY HAPPENED IS AS FOLLOWS,
a young man you see on the photo was dating some oneelse wife,and as day went no one thought they will be caught.but unfortunately a day reached whereby they were caught and this was what happened,is really sad,please do not date some one else wife or husband it is dangerous.share to the world so as they can stop this kind of behaviour.
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