Geochimica et Cosmochimica Acta, Vol. 67, No. 11, pp. 1955–1972, 2003
Copyright © 2003 Elsevier Science Ltd
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0016-7037/03 $30.00 ⫹ .00
Pergamon
doi:10.1016/S0016-7037(02)01341-8
The transport of U- and Th-series nuclides in sandy confined aquifers
B. C. REYNOLDS,1,2,* G. J. WASSERBURG,1 and M. BASKARAN3
1
The Lunatic Asylum of the Charles Arms Laboratory, Division of Geological and Planetary Sciences, California Institute of Technology,
Pasadena, CA 91125 USA
2
Open University, Department of Earth Sciences, Walton Hall, Milton Keynes, Buckinghamshire MK7 6BT, UK
3
Department of Geology, Wayne State University, Detroit, MI 48202, USA
(Received June 6, 2002; revised 13 November 2002; accepted in revised form November 13, 2002)
Abstract—Abundances of 238U, 234U, 232Th, 226Ra, 228Ra, 224Ra, and 222Rn were measured in groundwaters
of the Ojo Alamo aquifer in northwest New Mexico. This is an arid area with annual precipitation of ⬃22 cm.
The purpose was to investigate the transport of U-Th series nuclides and their daughter products in an old,
slow-moving groundwater mass as a means of understanding water-rock interactions and to compare the
results with a temperate zone aquifer. It was found that 232Th is approximately at saturation and supports the
view of Tricca et al. (2001) that Th is precipitated irreversibly upon weathering, leaving surface coatings of
232
Th and 230Th on aquifer grains. Uranium in the aquifer waters has very high [234U/238U] ⬃ 9 and low 238U
concentrations. These levels can be explained by low weathering rates in the aquifer (w238U ⬃ 2 ⫻ 10⫺18 to
2 ⫻ 10⫺17s⫺1) using a continuous flow, water-rock interaction model. The Ra isotopes are roughly in secular
equilibrium despite their very different mean lifetimes. The 222Rn and 228Ra isotopes in the aquifer correspond
to ⬃10% of the net production rate of the bulk rock. This is interpreted to reflect an earlier formed irreversible
surface coating of Th that provides Ra and Rn to the aquifer waters. The surface waters that appear to be
feeding the aquifer have low [234U/238U] and high 238U concentrations. The flow model shows that it is not
possible to obtain the high [234U/238U] and low [238U] values in the aquifer from a source like the present
vadose zone input. It follows that the old aquifer waters studied cannot be fed by the present vadose zone input
unless they are greatly diluted with waters with very low U concentrations. If the present sampling of vadose
zone sources is representative of the present input, then this requires that there was a major change in water
input with much larger rainfall some several thousand years ago. This may represent a climatic change in the
Southwest. Copyright © 2003 Elsevier Science Ltd
lematic. We here consider the flow through sandstone. We note
that studies of the geochemistry of the groundwater from carbonate terrains (Banner et al., 1990) show it to be quite different. This is due to the dominant questions involving carbonate
solubility. Several studies have considered the aquifer transport
of U-Th series nuclides in sandy aquifers, but the interpretations of field data using theoretical models are limited (Ku et
al., 1992), and research has mainly focused on the interpretations of 234U/238U activity ratios in groundwaters without considering the whole decay series (Dickson and Davidson, 1985;
Frohlich and Gellermann, 1987) or the amounts of various
species of radon (Rn) released from the rocks to water during
the groundwater transport. A general problem that has long
been recognized is the high 222Rn content (far higher than the
content of its parent 226Ra) of groundwaters, which requires
that the products of ⬃5 to 10% of the total Rn production rate
in the host aquifer rocks must be accessible to the water. This
requires very small grains with very large surface area geometries (Semkow, 1990) or the possible existence of small microfissures or nanopores in the crystals to provide a leakage
path (Rama and Moore, 1984). However, these hypothesized
leakage paths are not evident from the low loss of argon from
irradiated samples (Krishnaswami and Seidemann, 1988).
Krishnaswami et al. (1982) calculated sorption reaction rate
constants and residence times of daughter nuclides by deducing
recoil inputs from 222Rn for a static system but did not consider
the transport and the effects of precipitation.
This study uses a transport model (Tricca et al., 2000, 2001)
to discuss the natural abundances of U, Th, Ra, and Rn nuclides
1. INTRODUCTION
The distribution of naturally occurring radionuclides can
provide important insights into the mobility of pollutant nuclides and the rates of natural geochemical processes. Modeling
of local mass balance can provide quantitative estimates of
transport rates and their retardation by physico-chemical processes (adsorption-desorption and dissolution-precipitation).
Processes occurring during water-rock interactions induce significant elemental and isotopic shifts between parent-daughter
activity ratios of the uranium (U) and thorium (Th) isotope
decay series, with half-lives as indicated:
238
U(4.5Ga) ␣3 234Th(24d) ␤3 234U(0.25Ma) ␣3 230Th(75ka)
␣ 226
3
Ra(1.6ka) ␣3 222Rn(3.8d)
232
␣ 228
Th 共14Ga兲 3
Ra(5.8a) ␤3 228Th(1.9a) ␣3 224Ra(3.7d)
The disequilibrium of the decay series are well suited to the
study of transport and physico-chemical processes because of
the diverse chemical properties of the elements and a wide
range of decay constants within each series. Although the
thermodynamic properties of U and Th have been extensively
studied in aqueous solutions (Langmuir, 1978; Langmuir and
Herman, 1980), the complex chemical compositions of natural
waters make the prediction of elemental behavior more prob-
* Author to whom correspondence
([email protected]).
should
be
addressed
1955
1956
B. C. Reynolds, G. J. Wasserburg, and M. Baskaran
in a large aquifer from an arid region to estimate the influences
of weathering, ␣-recoil, solubility, and surface reactivity (adsorption-desorption). This one-dimensional model considers
three phases: groundwater advecting along a flowline (without
dispersion), host rock comprised of typical arkose sandstone
mineralogy, and a surface layer coating some fraction of the
grains where reversible sorption and irreversible precipitation
takes place. The evolution of the nuclide activities along the
flow and aging of the water can be modeled and evaluated for
each nuclide using the measured activities in the aquifer. The
model predicts the relationship between the 238U content and
the supply rate of species “i” by weathering (wi) and the 234U
content and the fraction of nuclides ejected or leached into the
water after ␣ recoil in the rock for each species i (the recoil
fraction ␧i) and the evolution as a function of the flow path.
This model (Tricca et al., 2001) and the corresponding reference is subsequently be referred to as TWPB.
The model was previously applied to a relatively small,
unconfined aquifer at the Brookhaven National Laboratory
(BNL) on Long Island, USA (a temperate region), where activities in the groundwater were found to be dominated by
processes occurring in the vadose zone above the aquifer
(TWPB). These workers showed that Th is saturated in the
neutral pH waters at the BNL site. This saturation condition fed
by continuous precipitation of Th by an earlier phase of ongoing weathering is hypothesized to have precipitated Th onto
surface layers of mineral grains within the aquifer on nonexchangeable sites. The model predicts that the 230Th activity
in the surface layer will reach a steady-state value after ⬃1 Myr
but that the 232Th activity in the surface layer grows linearly
with time for the duration of the aquifer as long as the saturation condition obtains. The accumulation of 232Th and 230Th
onto mineral surfaces from such an earlier weathering episode
can provide the source for the high Rn contents typically found
in aquifer waters. According to the model, it is the Rn loss from
accumulated 230Th and 232Th on thin surface layers that may
typically provide a source for high Rn emanation and Ra in
rocks with high weathering rates. The direct observation of this
surface layer has not been made. It is considered that nanometer-sized pores in the radioactive minerals of the host rock are
not the source of the Rn as shown by Krishnaswami and
Seidemann (1988). The source of the Ra is considered to be
produced by decay of 230Th and 232Th in the precipitated
surface layer. In this study we will test the model in a distinct
environment where the climate is arid.
The retention of Th during weathering has been recently
recognized in the study of weathering products in riverwaters.
In contrast, uranium and radium have higher solubilities and are
considered to be much more easily leached and transported
relative to thorium during the chemical weathering of rocks
(Vigier et al., 2001). In the work by TWPB, it was shown that
the low Th concentration in the waters, if simply attributed to
weathering, would require 232Th weathering rates two or more
orders of magnitude lower than that of U. However, if the low
Th weathering rate is the result of a two-stage process of
weathering followed by irreversible precipitation on surfaces,
then the observations could be explained without a gross discrepancy in weathering rates. This is distinctive from the onestep, first-order kinetic leaching process used by others (Vigier
et al., 2001).
The present study was directed at water-rock interactions in
an arid environment to explore the applicability of the theoretical model under conditions of low groundwater flow. The Ojo
Alamo Aquifer of the San Juan Basin, near Farmington, New
Mexico, was selected for the study since the hydraulics of the
aquifer had been previously studied and groundwater ages had
been established using 14C dating, which was used to construct
a numerical flow model for the aquifer (Brimhall, 1973;
Tansey, 1984; Phillips et al., 1986, 1989). Data is, in general,
given in activity units (e.g., [234U]) or in activity ratios (e.g.,
[234U/238U] ⫽ [234U]/[238U]). In some cases, we use the notation ␦234U ⫽ ([234U/238U] ⫺ 1) ⫻ 103.
2. SAMPLING AND ANALYTICAL PROCEDURES
This study focuses on the partially-confined Ojo Alamo
Aquifer of the San Juan Basin of northwestern New Mexico
(Fig. 1). The region has a continental, semiarid climate with
annual precipitation of ⬃22 cm, which is divided approximately between winter frontal storms from the Pacific and
summer monsoonal storms from the Gulf of Mexico. The
evapotranspiration yields a net input of 0.7 cm per year into the
groundwater. The land is sparsely populated and not cultivated,
with a vegetation cover type of north desert shrub. The mean
monthly temperature varies from 1 to 20 °C. The sampled
aquifer consists of an average 55-m thickness of fine- to coarsegrained alluvial arkosic sandstone deposited in the early Cenozoic and unconformably overlying the Cretaceous aquitard of
the Fruitland Formation and Kirtland Shale. It outcrops to the
south and west (see Fig. 1), and from this recharge area dips
gently northward. Toward the northeast, it is confined by the
conformably overlying Nacimiento Formation of shale and
siltstone, with thin interbedded sandstones. An extensive hydraulic study of the groundwater by Phillips et al. (1986, 1989)
has demonstrated that recharge is principally by water infiltration through the outcrop area and the subsequent water flow
northward is slow, typically at velocities of 1 m/yr (Brimhall,
1973; Phillips et al., 1989). The aquifer has porosity values on
the order of 20% and transmissivities in the range 0.5 to 8 ⫻
10⫺4 m2 s⫺1 (Brimhall, 1973; Phillips et al., 1989). Further
numerical modeling of the groundwater flow using hydraulic
heads and 14C activities enabled flow patterns within the aquifer to be estimated (Phillips et al., 1989), although variations
from those predicted by a simple piston flow model through the
aquifer were only slight (Tansey, 1984). Other studies of the
Ojo Alamo aquifer include measurements of the dissolved
noble gas concentrations for paleotemperature reconstructions
and deriving groundwater age constraints from excess 4He
accumulation (Phillips et al., 1986; Stute et al., 1995; Castro et
al., 2000). The groundwater ages derived from excess 4He were
made comparable to 14C ages using a model with exponential
decrease in the water velocity from 2 to 0.3 m/yr (Castro et al.,
2000) and a flow field that is not perpendicular to the measured
hydraulic head. From simple conservation of water, the modeled decrease in velocity would mean a large vertical flux of
water into the overlying Nacimento formation that has not been
demonstrated by field measurements. The groundwater ages
used in this study have been calculated using a constant velocity of 1.33 m/yr in the aquifer and estimating the flow distance
from the wellhead to the inferred recharge area where the Ojo
U, Th series in an aquifer or transport of U, Th series in an aquifer
Fig. 1. (A) Area of study showing the location of the sample sites. Inset shows location of study area (shaded) and the
main drainage pattern from the San Juan Mountains and continental divide feeding into the San Juan River. Major figure
shows sampling sites (open circles) with the location of the wells (W1–W10). Numbering is based on their distance from
the nearest southern outcrop of the Ojo Alamo sandstone (shaded), springs (Ojo Spring and Pete Spring), rock samples
(R1–R3), and the soil profile (Soil). River sample taken in Farmington. Seasonal river washes are shown as dashed lines
flowing to the San Juan River (solid line). The 6000-foot contour interval is shown (thin grey line) as well as the major roads
in the area (solid black lines). Estimated input source to sampled wells (by number) in the shaded outcrop of the Ojo Alamo
for samples indicated by X. The order of the outcrops from left to right correspond to well 2, Pete Springs wells 8, 6, 5,
9, 1, 4, 7, 3, and 10, respectively. (B) Cross section showing the Ojo Alamo aquifer from Phillips et al. (1989).
1957
1958
B. C. Reynolds, G. J. Wasserburg, and M. Baskaran
Alamo sandstone outcrops. We use 14C ages as a function of
flow path from Figure 13 in Phillips et al. (1989). Distances
were measured along mapped-out flowlines perpendicular to
the estimated hydraulic head (Tansey, 1984).
Groundwaters from the Ojo Alamo aquifer were collected in
February and June 2001 (12 samples). Sites were chosen to
cover the area to obtain waters from different recharge areas
and flow ages. The wells were equipped with large submersible
pumps able to draw several gallons per minute, except for the
springs and windmills (wells 1 and 2). Access to the wellheads
and help with the collection of several samples were given by
the Navajo Tribal Utility Authority (wells 3, 6, 7, and 8). Help
and access to submersible pumps was also given by the Carson
School (well 5), the Navajo Agricultural Products Industry
(NAPI) Feedyard (well 9), and a publicly available well at the
Brethren of Christ Mission School (well 4). Well 10 consisted
of a continuously flowing pumped well. The drawing up of the
water by the windmills may have affected the chemistry of the
waters, as equilibrium conditions of continuous flow might not
have been reached and gas-exchange may have taken place. A
riverwater sample was obtained from the San Juan River in
Farmington. Rock samples were taken from three places (see
Fig. 1): two in the recharge area (Rock #1 and 2) and one from
the cliffs near Farmington in the north (Rock #3). To investigate the vadose zone, a 1.5-m soil profile was taken from the
sides of a deeply cut aqueduct runoff channel (SP #1). Help in
the selection of a suitable site (Fig. 1) and access to the locality
was given by the Navajo Indian Irrigation Project (NIIP).
The pH and dissolved oxygen content were determined in the
field using a multiprobe flowcell setup, allowing for equilibrium to ensure that the measurements truly represented the
groundwater values. Only samples well 1 and 10 had visible
suspended particle loads, and all samples appeared colorless
and odorless. The concentrations of anions and cations in
filtered water samples (⬍ 0.45 microns) were determined
shortly after collection by Inner Mountain Labs, Farmington,
New Mexico. For analysis of long-lived isotopes (238U, 234U,
232
Th, and 230Th), both unfiltered and filtered water samples
were collected and were acidified (to pH ⫽ 2) in the field. The
filters used were 0.45-␮m cellulose filter cartridges that were
precleaned in 2 N HCl acid. Aliquots of each sample were
spiked with 236U, 233U, and 229Th tracers and left for 1 week to
allow thorough mixing of the spikes before carrying out extractions using iron coprecipitation. Uranium and thorium were
separated and measured by mass spectrometry following standard procedures (Chen et al., 1986). For analysis of short-lived
Ra and Th isotopes, 100-L samples were passed on line through
0.45-␮m prefilters and then through MnO-coated polypropylene filters. The MnO filters were ashed and measured using
a high-purity Ge Well gamma ray detector coupled to InSpector
at Wayne State University (Baskaran et al., 1993). Absolute
226
Ra activities were determined from 20 L of 0.45-␮m filtered
water, and passed through MnO-impregnated acrylic fibers that
have high Ra adsorption efficiency (Reid et al., 1979). The
222
Rn content of the waters was determined by a gas extraction
procedure. All separations and analyses for 222Rn were completed in the field. The Rn activities were measured by ␣
scintillation counting within a few hours after sample collection. The efficiency of the extraction was determined for the
procedure in the laboratory using a known 226Ra standard
(76%).
The three rock samples were analyzed for long-lived isotopes following crushing and powdering. Complete dissolution
of the rock minerals was achieved using combined HF, HClO4
acids and hot HNO3. U and Th were separated and measured
using the same methods as for the water samples.
A total of ⬃800 g of powdered material from the soil profile
was used for a water-leaching experiment. This consisted of a
combined sample of roughly 200 g of material for each of the
first 4 feet of the profile observed. This soil and rock sample
was mixed with ⬃500 g of ultrapure water and left for 24 h.
The water wash was filtered at 0.45 ␮m before analysis of U
and Th were made using the same method as for water samples.
The pH of the solution after leaching was unfortunately not
measured.
3. RESULTS
General chemical compositions of the water samples are
shown in Table 1. The suspended load (SPL) was determined
by weighing the dried filter before and after filtration. Total
dissolved solids were calculated from the chemical composition of the filtered waters. The cation/anion ratios varied between 0.96 and 1.06. The dominant ions in the groundwater are
generally Na⫹ and either HCO3⫺ or SO4⫺, with no systematic
relationship to each other. There is a relatively constant and
high level of total dissolved solids (TDS) throughout the aquifer (450 –770 mg/L corresponding to ions of 16 –25 meq/L).
The waters are slightly alkaline (measured pH varied between
7.1 and 9). Generally, major element data for unfiltered waters
(not shown) were indistinguishable from those of the corresponding 0.45-␮m filtered samples. The major ion chemistry is
principally derived from the dissolution of gypsum and
Na⫹/Ca2⫹ ion exchange that drives calcite into dissolution.
The low Cl⫺ concentrations and high pH exclude significant
NaCl or sulfide dissolution as a source of the ions. The TDS
appears to be roughly constant without any correlation with
distance. Comparison of the hydrochemical data from the same
sample localities from previous work spanning back over the
last 20 yr indicates only slight variations in the major ion
concentrations, although there may be significant variations in
the trace element composition of the groundwater (Tansey,
1984; Stute et al., 1995). There is also good agreement in the
major ion composition within the aquifer from the closest two
wells (5 and 6). Thus there appears to be relatively low temporal and spatial variations in the major ion chemistry. Generally, chloride is introduced into groundwaters through the seepage of rainwater, and there are no recognized sources for
chloride in the aquifer. We consider Cl to behave conservatively and that the chloride concentration in the aquifer reflects
the amount of evaporation that has taken place in the upper part
of the vadose zone. The Cl⫺ concentrations in rainwater in
New Mexico are ⬃0.25 mg/L (0.0076 meq/L), while general
Cl⫺ concentrations in the groundwater are around 8 mg/L (0.23
meq/L), which indicates that ⬃97% of the rainwater evaporates
before reaching the water table. The Cl concentrations for the
rainwater were generously provided by F. M. Phillips based on
a 5-yr monitoring of the Long Term Ecological Research Site
at the Sevilleta Wildlife Refuge. Thus, the annual groundwater
Table 1. General Characteristics of groundwater samples.
Well
Bottom
Altitude
m
Proposed
Flow1
Age ka
TDS2
mg/l
Ojo Spring
Pete Spring
River
Well #1
Well #2
Well #3
Well #4
Well #5
Well #6
Well #7
Well #8
Well #9
Well #10
2000
1805
1615
1893
1850
2009
1913
1855
1885
1920
1734
1729
1850
2000
1790
1615
1849
1812
1810
1839
1734
1748
1719
1643
1609
1792
0
0?
n/a
4.3
6.4
7.5
9.4
11.3
11.3
14.1
20.1
24.9
28.1
470
1150
200
630
650
770
580
530
550
450
590
—
470
1
Major Cations
pH
8.8
8.7
8.1
9.3
7.7
7.7
8.9
Major Anions
SPL3
mg/l
O24
mg/l
Na⫹
meq/l
K⫹
meq/l
Mg2⫹
meq/l
Ca2⫹
meq/l
HCO⫺
3
meq/l
CO2⫺
3
meq/l
Cl⫺
meq/1
SO 2⫺
4
meq/l
Sr
ppb
4.34
0.10
(3.93)
(2.11)
8.56
1.13
0.47
0.10
0.18
0.18
0.54
0.18
—
10.01
(3.12)
(3.16)
0.03
(3.22)
0.04
⬍0.01
0.01
0.04
—
⬍0.01
1.60
16.97
0.56
10.80
10.39
11.02
9.89
9.56
9.26
7.86
9.25
14.30
8.44
0.05
0.02
0.04
0.01
0.01
0.04
0.01
0.01
0.01
0.01
0.02
0.07
0.01
0.47
0.45
0.55
⬍0.01
⬍0.01
0.13
⬍0.01
⬍0.01
⬍0.01
⬍0.01
0.02
0.08
⬍0.01
5.36
1.29
2.38
0.05
0.04
1.00
0.04
0.02
0.04
0.03
0.36
1.55
0.03
2.56
6.95
1.6
4.57
3.11
6.59
5.12
4.88
3.84
3.29
4.94
—
4.58
⬍0.01
⬍0.01
⬍0.01
2.20
1.83
⬍0.01
1.10
1.76
2.20
2.20
⬍0.01
—
1.46
0.39
2.15
0.36
0.24
0.22
0.17
0.21
0.36
0.27
0.18
0.22
—
0.16
4.66
9.20
1.36
3.58
5.25
5.96
3.23
2.50
2.67
1.99
4.42
—
1.93
986
915
531
24
28
302
24
36
55
33
214
7115
28
87
Sr/86Sr
0.70829 ⫾ 4
0.70918 ⫾ 7
0.70957 ⫾ 6
0.70857 ⫾ 9
0.70887 ⫾ 6
0.70847 ⫾ 4
0.70890 ⫾ 6
0.70887 ⫾ 4
0.70871 ⫾ 8
0.71000 ⫾ 5
0.70883 ⫾ 5
0.70899 ⫾ 4
0.70930 ⫾ 5
Conc.
Factor5
55
303
51
34
31
24
30
51
38
25
31
23
Estimated flow distance is based on hyraulic flow lines from recharge area to well head from Tansey (1984).
TDS ⫽ Total Dissolved Solids.
SPL ⫽ Suspended Particulate Load (dry weight of particulates per liter ⬎0.45 ␮m).
4
O2 ⫽ measured dissolved oxygen concentrations. Parentheses indicate oxygen levels not measured from flowing well head. Major ions given in milk-equivalents per liter. Strontium (Sr) concentrations
given in ppb (␮g/l).
5
Concentraton factor is calculated as ratio of the measured Cl⫺ concentrations to the estimated rainfall concentration of 0.0071 meq/L.
2
3
U, Th series in an aquifer or transport of U, Th series in an aquifer
Location
Water
level
m
1959
1960
B. C. Reynolds, G. J. Wasserburg, and M. Baskaran
Table 2. Activities in the rock and soil profile.
[238U]
[232Th]
Location
ppm
dpm/kg
[234U]
dpm/kg
␦234U
[230Th]
dpm/kg
ppm
dpm/kg
[230Th/234U]
[230Th/232Th]
Rock 1 - Recharge
area
Rock 2 - Recharge
area
Rock 3 - Northern
edge
Soil Profile Leachate
Water ⬍0.45 ␮m
2.91
2172 ⫾ 13
2373 ⫾ 51
93 ⫾ 23
2651 ⫾ 27
8.04
1962 ⫾ 2
1.117 ⫾ 21
1.352 ⫾ 8
3.53
2631 ⫾ 5
2732 ⫾ 11
38 ⫾ 4
0.99
741 ⫾ 2
708 ⫾ 6
⫺45 ⫾ 8
1099 ⫾ 13
7.31
1785 ⫾ 1
1.553 ⫾ 9
0.616 ⫾ 19
0.04
0.316
0.439
392 ⫾ 5
0.01079
0.05
0.01231 ⫾ 2
0.0246
0.876 ⫾ 19
␦234U is the isotopic shift in permil deviation from secular equilibrium. ␦234U ⫽ ([234U/238U] ⫺ 1) ⫻ 103. [234U] and [238U] are the activities of
U and 238U in the water.
234
recharge is ⬃0.7-cm yr⫺1. This value corrects our original
erroneous estimate, which was much higher. This input can
supply the calculated subhorizontal groundwater flow of 133
cm/yr, given the 20% porosity and geometry of the aquifer
compared to the recharge area (F. M. Phillips, pers. comm.).
The local very high Cl⫺ concentrations from Pete Spring (76
mg/L) indicate that very large amounts of evaporation have
taken place in the input to the storage tank or from the water
tank before water collection (over 99% of the rainwater evaporated).
The dissolved oxygen concentrations pumped into the flowcell, excluding well 4, were very low (⬍ 0.08 mg/L or 1%
saturation), and so waters are strongly anoxic. Near-surface
waters from Ojo Spring and Pete’s Spring have dissolved
oxygen contents 30% above saturation. Elevated dissolved
oxygen concentrations in the samples from the windmills
(⬃45%) probably demonstrate that waters were not closed to
gas exchange whilst being drawn up and are indicated by
parentheses in Table 1.
The Sr concentrations of the groundwaters were highly variable from 24 ppb to 7.1 ppm but do not correlate with the
variations in the 87Sr/86Sr ratios, which varied from 0.7087 to
0.7100. Generally lower 87Sr/86Sr ratios correlate with higher
Ca/Sr ratios, which may be associated with gypsum dissolution.
The Sr concentrations of riverwater and springwater samples
are typically much higher than groundwater. The 87Sr/86Sr of
springwater is indistinguishable from the range of values we
found in groundwaters and in the riverwater.
The U and Th content of the rock samples is shown in Table
2. Previously measured U and Th concentrations are between
1.3 and 5 ppm and 2.4 and 6 ppm, respectively (Castro et al.,
2000). Both rock samples 1 and 2 are from the recharge zone
that shows heavy Fe-Mn staining. They have high U and Th
concentrations with activities of ⬃1800 to 2600 dpm/kg. Unless indicated otherwise, all data on radioactive nuclides are
given in activity units (dpm/kg). This is indicated by square
brackets. Rock sample 1 has daughter activities for both 234U
and 230Th in excess of unity with [234U/238U] and [230Th/234U]
of 1.093 and 1.117, respectively. However, for rock sample 3,
which is not from the recharge area and does not have heavy
mineral staining, the U concentration is much lower and the
[234U/238U] ⫽ 0.955. This corresponds to ␦234U ⫽ ⫺45‰.
Thus, the rock is deficient in 234U compared to 238U, suggesting significant enhanced 234U removal. This rock has Th con-
centrations similar to rocks 1 and 2, but the [230Th/232Th] ratio
is less than unity, roughly reflecting the low 238U/232Th ratio.
The high value of [230Th/234U] ⫽ 1.55 and [234U/238U] ⫽
0.955 in rock #3 can only be generated by an inherited 230Th
from precipitation of Th during weathering and some enhanced
loss of 234U.
Isotopic and abundance data for U and Th are given in Table
3. Measured 238U activities in both unfiltered and filtered
samples (⬍ 0.45 ␮m) from the aquifer vary from 15 ⫻ 10⫺3 to
156 ⫻ 10⫺3 dpm/kg (20 –210 ng/kg). Generally [238U] in the
unfiltered and filtered samples are the same within error; however, for wells 2, 9, and 10, [238U] in the unfiltered water is
somewhat higher compared to the filtered water. These wells
have higher suspended particulate concentrations and demonstrate that there must be U in the suspended particles of the
unfiltered samples. The [234U/238U] in the aquifer vary from
5.7 to 11.9 and so are all highly enriched in 234U. Generally,
[234U/238U] in the unfiltered and filtered samples are the same,
although well 9 has a significantly higher [234U/238U] in the
dissolved fraction. Given that there are variations in the 238U
activity by over an order of magnitude it is significant that the
[234U/238U] varies by less than a factor of 2 (see Fig. 2). In all
wells, [238U] does not increase with the TDS or correlate with
distance from the recharge area, as shown in Fig. 3.
The 232Th activities in the filtered fractions of the riverwaters, springwaters, and groundwaters lie between 4 ⫻ 10⫺6 to
9 ⫻ 10⫺5 dpm/kg (14 –300 pg/kg) (Table 3). The [232Th] in the
unfiltered water samples are all higher, especially in well 10
and the riverwater, with activities between 7 ⫻ 10⫺5 and 7 ⫻
10⫺3 dpm/kg (0.3–30 ng/kg), as shown in Fig. 4. The estimated
solubility of Th in pure water is only ⬃10 pg/kg (2.4 ⫻ 10⫺6
dpm/kg) (Langmuir and Herman, 1980), although Th solubilities in natural waters can be increased greatly up to pH ⬃8 by
complexes with organic colloids and inorganic ligands such as
phosphate. Thorium forms a strong complex with CO32⫺ and
HCO3⫺. The residence time of thorium in Mono Lake was
found to be an order of magnitude higher than those in other
lakes and ocean water systems. Elements like Th have considerably enhanced solubility due to complexing with carbonate
and bicarbonate (Anderson et al., 1982). However, at pH values
much above 8, complexation should have little effect, as all
present Th4⫹ ions will be complexed by hydroxyl ions. The
measured [232Th] in the filtered waters are around the solubility
limit for thorianite with some enhancement presumably due to
U, Th series in an aquifer or transport of U, Th series in an aquifer
1961
Table 3. Activities of long-lived nuclides in the water samples.
Location
Ojo Spring-unfiltered
⬍0.45 ␮m
Pete Springunfiltered
⬍0.45 ␮m
River unfiltered
⬍0.45 ␮m
Well #1 unfiltered
⬍0.45 ␮m
Well #2 unfiltered
⬍0.45 ␮m
Well #3 unfiltered
⬍0.45 ␮m
Well #4 unfiltered
⬍0.45 ␮m
Well #5 unfiltered
0.45 ␮m
Well #6 unfiltered
⬍0.45 ␮m
Well #7 unfiltered
0.45 ␮m
Well #8 unfiltered
⬍0.45 ␮m
Well #9 unfiltered
⬍0.45 ␮m
Well #10 unfiltered
⬍0.45 ␮m
Water
Age*
ka
0.0
0?
4.3
6.4
7.5
9.4
11.3
11.3
14.1
20.1
24.9
28.1
[238U]
10⫺3 dpm/kg
[234U/238U]
Excessa
[234U]
10⫺3 dpm/kg
509 ⫾ 1
531 ⫾ 2
8500 ⫾ 150
1.703 ⫾ 2
1.709 ⫾ 2
1.534 ⫾ 7
357 ⫾ 2
377 ⫾ 3
4537 ⫾ 180
32.1 ⫾ 0.4
9.1 ⫾ 0.1
3.0 ⫾ 0.01
9055 ⫾ 192
650.4 ⫾ 1.4
634.0 ⫾ 4.1
76.2 ⫾ 0.2
73.7 ⫾ 0.2
42.7 ⫾ 0.1
37.6 ⫾ 0.1
15.3 ⫾ 0.1
14.9 ⫾ 0.1
156.3 ⫾ 4.9
154.4 ⫾ 0.5
107.6 ⫾ 0.8
108.1 ⫾ 0.4
—
121.6 ⫾ 0.5
14.6 ⫾ 0.1
14.7 ⫾ 0.1
77.1 ⫾ 0.1
77.6 ⫾ 1.0
30.4 ⫾ 0.1
21.4 ⫾ 0.1
84.0 ⫾ 1.0
79.8 ⫾ 0.5
1.511 ⫾ 5
1.479 ⫾ 1
1.503 ⫾ 2
8.011 ⫾ 17
7.999 ⫾ 22
8.881 ⫾ 29
8.899 ⫾ 18
7.269 ⫾ 104
7.456 ⫾ 24
7.290 ⫾ 34
7.210 ⫾ 15
—
10.695 ⫾ 18
—
9.839 ⫾ 20
11.772 ⫾ 65
11.874 ⫾ 55
8.405 ⫾ 22
8.390 ⫾ 51
5.735 ⫾ 35
5.922 ⫾ 23
10.291 ⫾ 627
10.687 ⫾ 133
4627 ⫾ 208
312 ⫾ 2
319 ⫾ 4
534 ⫾ 3
516 ⫾ 4
336 ⫾ 3
297 ⫾ 2
96 ⫾ 2
96 ⫾ 1
983 ⫾ 63
959 ⫾ 7
—
1048 ⫾ 8
—
1075 ⫾ 10
157 ⫾ 2
160 ⫾ 2
571 ⫾ 3
573 ⫾ 16
144 ⫾ 2
105 ⫾ 1
780 ⫾ 62
773 ⫾ 16
—
650 ⫾ 72
1.66 ⫾ 0.01
107.7 ⫾ 0.4
4.43 ⫾ 0.02
12.89 ⫾ 0.03
5.12 ⫾ 0.03
86.6 ⫾ 2.0
1.82 ⫾ 0.01
26.5 ⫾ 0.1
2.16 ⫾ 0.01
70.3 ⫾ 0.5
7.18 ⫾ 0.06
7.3 ⫾ 0.1
—
15.3 ⫾ 0.1
—
8.5 ⫾ 0.1
0.35 ⫾ 0.01
305.5 ⫾ 0.8
—
659.3 ⫾ 11.4
3.10 ⫾ 0.04
232
Th
10⫺5 dpm/kg
* Water ages calculated from the proposed flow distance using a constant flow velocity of 1.33 m/yr (4 ⫻ 10⫺6 cm/s).
Excess [234U] ⫻ {[234U/238U] ⫺ 1}[238U]
a
ligands. We conclude that the Th concentrations appear to be
controlled by the local solubility limit in each well. The Th
observations at this site are in full agreement with those found
by TWPB for groundwaters in the BNL site (see Fig. 4). We
have calculated the maximum concentration of 232Th in the
suspended load from the difference between the concentration
of unfiltered and filtered values and used the weights of SPL in
Table 1. In two cases the calculated concentrations are equal to
the values for the rocks (Table 3). Most of the values are a
factor of 10 or more below this. It follows that the excess 232Th
can easily be accounted for if some fraction (5% to 10% of the
SPL) was just transported soil particles. This suggests that there
is no great enhancement of Th precipitates (e.g., thorite) on
particles in the SPL. However, it is possible that soil/rock
particles are only a very small fraction of the SPL. In this case,
it might be argued that precipitated thorite is in the SPL. As it
is not possible to estimate the 238U concentrations, no arguments can be made for intrinsic Th enhancement relative to U,
which might be diagnostic of thorite precipitation. Measured
[232Th] in filtered water above ⬃10⫺5 dpm/kg (40pg/kg) are
explained by low pH, elevated ligand complexation, or poor
filtration of fine suspended particulate thorianite. Previously,
TWPB reported Th activities above 20 ⫻ 10⫺3 dpm/kg in
filtered waters (88ng/kg) and even higher activities in the
unfiltered waters from the Magothy aquifer, which underlies
the main glacial aquifer. This sample had very high Fe and Mn
concentrations and low O2. Tricca et al. (2001) inferred that the
high degree of mobilization of “insoluble” elements was related
to the low O2 content. This may be related to Th adsorption
onto suspended Fe-Mn colloids (⬍ 0.45 ␮m). There is no
evidence of enhanced Th concentration in the low O2 waters
studied here.
To determine the availability of labile U and Th, a wash of
the soil profile sample was carried out. This released relatively
high amounts of U and Th into the water (see Table 2), and the
measured activities were over 0.3 dpm/kg for 238U and above
0.012 dpm/kg for 232Th. The measured [234U] and [238U] are
quite similar to vadose zone water samples and had a ␦234U ⫽
⬃400‰. The 232Th concentrations are much higher than the
estimated solubility limit in all of the samples of waters collected in the field. This must reflect a large amount of Th
adsorbed onto a colloidal fraction. The [230Th/232Th] ratio in
this sample was ⬃0.88, which could reflect the activity ratio in
the bulk rock. The very low [230Th/234U] of ⬃0.025 indicates
there is substantial fractionation of U from Th isotopes, with U
being much more soluble. Due to the higher solubility of U in
natural waters, the 234U atom produced by 234Th recoil could
be leached with the water wash as opposed to the recoiled
230
Th atom, as thorium is less soluble in natural waters. This is
the same factor by which the 232Th/238U is decreased in the
leach over the value in the rock. For the typical concentration
in the rocks in the vadose zone (1000 dpm/kg), we find that a
relatively large fraction of the U (0.02%) and Th (0.001%) can
readily be dissolved in water. The U concentration and isotopic
composition of the leach is remarkably similar to the values
found for the riverwater and Ojo Spring samples.
All Ra isotopes are from the decay of parent Th isotopes. Ra
activities were only determined on the filtered (0.45 ␮m) frac-
1962
B. C. Reynolds, G. J. Wasserburg, and M. Baskaran
Fig. 2. Measured ␦234U and 238U activities for (A) the aquifer (filtered groundwater), river, and spring samples.
Springwaters are shown as diamonds and the riverwater as a triangle. For the groundwater, unfiltered samples are shown
as squares, and 0.45-␮m filtered water samples are shown as circles. Note that the aquifer samples are restricted to very low
238
U and very high ␦234U. Springs have high 238U and low ␦234U. (B) Data on filtered and unfiltered groundwaters only.
This corresponds to the extreme left-hand side of Fig. 2a.
tion due to the sampling method. Previous ultrafiltration studies
by TWPB demonstrated that Ra is carried mainly on colloids as
is expected for a surface-reactive element. Ra activities in the
sampled wells range from 0.12 to 2.3 dpm/kg for 228Ra, from
0.04 to 0.54 dpm/kg for 226Ra, and from 0.17 to 2.86 dpm/kg
for 224Ra (see Table 4). Despite the large variations in activity,
the [228Ra/226Ra] ratios are all between 1.7 and 4.0, and for the
daughter-parent [224Ra/228Ra] ratio (ignoring 228Th), the values
are between 0.77 to 1.65 (see Table 4). The variations in the
activity do not correlate with uranium activities or TDS but
U, Th series in an aquifer or transport of U, Th series in an aquifer
1963
Fig. 3. Measured activities in the aquifer waters vs. the estimated age for (A) 238U and (B) ␦234U. Unfiltered samples are
shown as squares, and 0.45-␮m filtered water samples are shown as circles. Well numbers as labeled.
appears to correlate with the concentration of Ca ions in the
groundwater (see Fig. 5). The samples with the higher Ra
activities are also the samples with the higher [224Ra/228Ra]
(see Fig. 6).
In contrast to the behavior of [228Ra/226Ra] and [224Ra/226Ra]
in the well waters, the ratio [226Ra/234U] ranges from 0.04 to
4.85. The ratio of [228Ra/232Th] ranges from 2.3 ⫻ 103 to
1.7 ⫻ 104 for the filtered samples. For the unfiltered samples,
1964
B. C. Reynolds, G. J. Wasserburg, and M. Baskaran
Fig. 4. Measured 232Th activities (on a logarithmic scale) vs. the groundwater age for groundwaters (unfiltered squares
and filtered circles). Also shown are filtered (dark grey) and unfiltered (light grey) Ojo Spring waters (diamonds) and
riverwater (triangles). Data for the filtered groundwater from the glacial aquifer from Tricca et al. (2001) are shown as X
for comparison. It can be seen that the Th concentrations are at a very low level. The highest [232Th] for a filtered sample
is 10⫺4 dpm/kg or 0.4 ppt.
the range in [228Ra/232Th] is 18.4 to 1.8 ⫻ 104. These results
are shown in Fig. 7. Note that 232Th decays directly to 228Ra,
while 234U 3 230Th 3 226Ra. It can be seen from Fig. 7 that
relative to the production rate in the water and suspended
particles, [226Ra/234U] is much lower than the production ratio
for six samples and is much greater for two samples. This
requires efficient 226Ra removal for the low ratios and an
additional source for 226Ra for the one high ratio. The
[228Ra/232Th] and [224Ra/232Th] are all in great excess of the
production ratio in the water. Note that the half-life of 224Ra is
only 3.66 d and that the observed 224Ra/228Ra ⬇ 1 for all
samples. 228Ra/224Ra and 224Ra/226Ra are approximately in
secular equilibrium (for normal Th/U ratios) and cannot be
provided by U or Th from the solution. This requires sites with
either rapid exchange rates for the radium isotopes or sites that
provide the Ra in the solution at or near the equilibrium value.
These results show that the Ra isotopes are far in excess of the
in situ production in the water.
The measured Rn activities vary from 55 to 485 dpm/kg as
shown in Table 4. These Rn activities are far higher than the
measured activities of other isotopes in the decay series, as was
observed by many previous workers. The lowest Rn activities
are from samples that have high dissolved oxygen concentrations, which suggests that there has been gas exchange with the
Table 4. Activities of short-lived nuclides in filtered groundwater (⬍0.45 ␮m).
Location
Well
Well
Well
Well
Well
Well
Well
Well
Well
#1
#2
#3
#4
#5
#6
#7
#8
#10
Water
Age
ka
[222Rn]*
dpm/kg
[226Ra]
10 dpm/kg
[228Ra]
10 dpm/kg
[224Ra]
10 dpm/kg
冤
4.3
6.4
7.5
9.4
11.3
11.3
14.1
20.1
28.1
(73)
(76)
309
(55)
485
103
282
365
381
41 ⫾ 14
538 ⫾ 26
210 ⫾ 7
43 ⫾ 6
74 ⫾ 15
330 ⫾ 22
388 ⫾ 27
80 ⫾ 8
115 ⫾ 50
2287 ⫾ 29
357 ⫾ 11
178 ⫾ 17
241 ⫾ 14
693 ⫾ 36
1536 ⫾ 22
121 ⫾ 32
—
2863 ⫾ 148
399 ⫾ 16
197 ⫾ 34
186 ⫾ 38
734 ⫾ 37
2534 ⫾ 151
165 ⫾ 16
—
1.25 ⫾ 0.07
1.12 ⫾ 0.06
1.11 ⫾ 0.22
0.77 ⫾ 0.16
1.06 ⫾ 0.08
1.65 ⫾ 0.10
1.36 ⫾ 0.38
⫺3
* Numbers in parentheses are likely lower limit of
222
⫺3
⫺3
224
228
Ra
Ra
冥
冤
226
228
Ra
Ra
冥
0.36
0.24
0.59
0.24
0.31
0.48
0.25
0.66
Rn as possibly due to water/gas exchange and loss of Rn while drawing the water sample.
U, Th series in an aquifer or transport of U, Th series in an aquifer
1965
Fig. 5. Activities of 224Ra (diamonds) and 228Ra (triangles) in aquifer waters plotted against Ca concentration. Note that
at high Ca levels, the Ra is also high.
air and hence loss of Rn gas while drawing the water. The
supposed air incorporation and Rn loss in well 4 is surprising,
as the water should have been pumped directly to the outlet.
Excluding the samples with possible loss of Rn gas (shown by
parentheses in Table 4, the Rn activities in all the samples lie
in a relatively narrow range of 103 to 485 dpm/kg. Since the
measured parent-daughter [226Ra/222Rn] in the water varies
from 1.7 ⫻ 10⫺3 to 8.9 ⫻ 10⫺5, this confirms that the Ra in the
water cannot be a significant source of Rn activity. If we
consider ⬃103 times the Ra in the water was adsorbed on
surfaces, then this would support the 222Rn. However, these
sources of Ra and Rn would correspond approximately to over
5% of the decay rate of U per kg of aquifer rock. This has been
the dilemma found by other workers from Rn and Ra studies.
4. MODELING GROUNDWATER ACTIVITIES
We will follow the model of TWPB for the radionuclide
transport model and will use exactly the same notation. The
reader is referred to that study for a full treatment of the
problem. The one-dimensional steady-state model is shown
schematically in Fig. 4 TWPB, and the model parameters
defined in Table 5 of that work. In the following, we use the
activity of species “i” as iAw. This is to distinguish the theoretical model parameters from the measured activities. In the
numerical calculations, we substitute the measured values [i]
for iAw.
4.1. The Vadose Zone
The activity of 238U in the water at depth x is from the
weathering of the host rock, while the 234U activity is the sum
of weathering and the ␣-recoil of 234Th from the rocks. Parameters in the vadose zone are indicated by primes. The activity in
the vadose water of species i is given by 共i A⬘w 兲 . For 238U, this
may be written:
238U
A⬘w ⫽
␳ r238UA⬘r w⬘238U x
s⬘ ␳ w q
v⬘
(1)
and for the ratio of the specific activities of 234U to 238U,
according to the model in the aquifer water may be written:
共 234UA⬘w/ 238UA⬘w兲 ⫽ 共w⬘234U ⫹ ␧⬘234Th␭ 234U兲/w⬘238U
(2)
Here w⬘i is the supply rate weathering of species i, v is the water
velocity in the vadose zone, q ⬅ n⬘/(1⫺n⬘), where n⬘ is the
porosity and s is the saturation index. ␳r and ␳w are densities of
rock and water, respectively, and x is the distance along the
macroscopic flowline. The recoil fraction is ␧⬘i, and ␭i is the
decay constant of species i. The excess activity in per mil units
is thus:
1966
B. C. Reynolds, G. J. Wasserburg, and M. Baskaran
Fig. 6. 224Ra vs. its parent 228Ra for aquifer samples. Note that five of the datapoints lie on a slope-1 line. However,
samples #3 and #8 show clear excesses of 224Ra above the equilibrium values.
冉
冊
␭ 234U 䡠 ␧⬘234Th
关 U/ U兴 ⫺ 1 ⫽
⫽ ␦ 234U ⫻ 10 ⫺3
w⬘238U
234
238
(3)
To estimate the weathering rate of U in the vadose zone, we
use Eqn. 3 and the value of [234U/238U] in Ojo Spring,
[234U/238U] ⫽ 1.709. This gives w⬘238U/␧⬘234Th ⫽ 1.3 ⫻ 10⫺13
s⫺1, which is essentially the same for the river and Ojo Spring
samples. For a recoil fraction ␧⬘234Th ⬃ 10⫺2, this would
correspond to a weathering rate w⬘238U ⬃ 1.3 ⫻ 10⫺15s⫺1.
Now consider the 238U activity directly to determine w⬘238U
using Eqn. 1. The 238U activity of Ojo Spring water is 0.53
dpm/kg and that of the rock in the vadose zone is ⬃2600
dpm/kg. For a vadose zone thickness of 103 cm, this requires
0.53⫽ (2(1⫺n⬘)/n⬘sv⬘)w⬘238U2.6⫻103⫻103, where we have
taken ␳ r / ␳ w • 2. For a porosity n ⬃ 0.2, we have w⬘238U ⬃ 1.3
⫻ 10⫺7n⬘sv⬘. To estimate n⬘sv⬘, we consider that total water
input per unit area into the aquifer is 0.7 cm per year (see
“Sampling and Analytical Procedures”) so that at steady
state, n⬘sv⬘ ⫽ 0.7 cm y⫺1 ⫽ 2.2 ⫻ 10⫺8cm s⫺1. Substituting
this in the above equation for wi yields w⬘238U ⫽ 2.9 ⫻
10⫺15s⫺1. The value of 103 cm for the depth of Ojo Spring
is certainly high. There is thus remarkable agreement in the
estimates of w⬘238U from the 238U concentrations and
[234U/238U] if the recoil fraction is ␧ ⬃ 10⫺2. The ready
availability of both 238U and high [234U/238U] in the vadose
zone is evident both in the riverwater and in the waterleaching experiment. These atoms then become available for
transport when flowing porewater is available. From the
recoil fraction it can be estimated that the effective grain size
of U-rich phases in the vadose zone is ⬃6 ␮m. However,
there is no evidence for ultrafine particles providing U in the
vadose zone. The readily soluble U atoms in the vadose zone
were most likely precipitated out of solution after weathering and are left on the surfaces due to evaporative losses.
The estimated weathering rate derived above from Eqn. 1
may be compared to estimates from solutes in rivers. Typical
chemical exhumation rates have been reported that are ⬃1
mm/kyr (Gaillardet et al., 1999), which corresponds to a
weathering removal rate of 3 ⫻ 10⫺14 ms⫺1. If this is
provided by weathering of a zone of 1-m thickness, this
corresponds to w ⬃ 3 ⫻ 10⫺14s⫺1, for a 10-cm zone, w ⬃3
⫻ 10⫺13s⫺1. It is not evident what surface zone thickness
provides the runoff or how directly the weathering rate of U
and Th, which are contained in accessory minerals, may be
related to bulk weathering (cf. Morton and Hawsworth,
1999). The agreement is quite good for these very independent methods. As noted by the reviewers, it appears that U
U, Th series in an aquifer or transport of U, Th series in an aquifer
1967
Fig. 7. [226Ra/234U] vs. [228Ra/232Th] for filtered (circles) and unfiltered (squares) aquifer water samples (NB logarithmic
scales). Well numbers are labeled. If these two daughter-parent systems were in equilibrium, the values would be unity. The
[226Ra/234U] values are not in equilibrium but only range from ⬃5 ⫻ 10⫺2 to 10. The [228Ra/232Th] range from ⬃10 to 106.
may be more effectively removed than Na. It is certainly not
to be expected that the “global” weathering rate should apply
here.
4.2. U in the Groundwater
The evolution of the U activities in the groundwater must
consider the initial input of U from the vadose zone. However,
a comparison of the 238U concentration and [234U/238U] values
for the aquifer samples with the springwater and riverwater
studied shows marked differences. The aquifer waters typically
have 238U concentrations of about an order of magnitude less
than that of either Ojo Spring or the riverwater. Further, the
[234U/238U] values of the aquifer waters are an order of magnitude greater than that of the springwater and riverwater (see
Fig. 2). As all data indicate that U is rather soluble and not
strongly surface reactive, it is not possible to consider that the
groundwaters are the result of the present vadose zone input
unless there is a funneling mechanism that feeds vadose zone
input into the aquifer with enormous dilution by U-free water
or a mechanism of U precipitation before entering the aquifer.
Further, to achieve the high [234U/238U] values in the aquifer,
the original vadose zone input to the present aquifer must have
been at much lower 238U concentrations (10⫺1 or less) than the
current vadose zone values. This follows the inverse relation-
ship between [234U/238U] (or ␦234U) and 238UAw, which has
been extensively discussed by TWPB.
The [234U/238U] values do not show any regular increase
with age but are all at high values, the lowest being at
[234U/238U] ⫽ 5.922. If we consider all of the [234U/238U]
values for the aquifer samples and assume that they are not
affected by any vadose input, then the values of w/␧234Th are
found to lie between 9.4 ⫻ 10⫺15 and 1.8 ⫻ 10⫺14s⫺1. For a
rather high value ␧ ⬃ 10⫺2, this implies w ⬃ 10⫺16s⫺1 and
hence low weathering rates throughout the aquifer. The 238U
activity in the waters is not correlated with the groundwater
age. Here, the age was calculated assuming a flow velocity of
1.33m/yr over the distance from the well to the aquifer input.
From consideration of the U concentrations, it is evident that
the samples studied do not represent sampling along welldefined flow lines.
The activity of 238U in the groundwaters according to Tricca
et al. (2001) should be governed by the equation:
238U
Aw ⫽
␳ rA rw 234U共1 ⫺ n兲
x ⫹ 238UA w0
␳ wnv
(4)
where 238UAw0 is the input from the vadose zone, and all other
parameters refer to the aquifer rock.
We next estimate the weathering rate from Eqn. 4 for U
assuming the flow model using the U concentration and then
1968
B. C. Reynolds, G. J. Wasserburg, and M. Baskaran
consider the vadose zone input that originally provided the
present aquifer water to be negligible. We take the two extremes with the highest and lowest U concentrations. This gives
w ⫽ 2.3 ⫻ 10⫺17s⫺1 and 1.5 ⫻ 10⫺18 s⫺1 for wells 4 and 7,
respectively. It can be seen that the inferred weathering rates
are very low. Using the typical value of w/␧234Th ⬃ 10⫺14 s⫺1
found from the ␦234U results for the aquifer waters and the
above weathering rates determined from the U concentration,
we find ␧234Th ⫽ 2 ⫻ 10⫺4 to 2 ⫻ 10⫺3. It follows that within
the aquifer, the weathering rates are very low and the recoil
fraction is low and typical for grain sizes of U-containing
minerals of ⬃30 to 100 ␮m.
4.3. Application of the Model to Thorium Activities
As shown in Fig. 4, the 232Th concentration is approximately
that which would be estimated from thorite saturation with
some relatively small enhancement. As the 232Th activities in
the filtered waters are close to the estimated solubility limit, we
conclude that 232Th is controlled by precipitation. The ratio
[232Th/238U] in the waters is typically ⬃10⫺4 for the filtered
samples and ⬃10⫺3 for unfiltered samples. The activity of 238U
and 232Th in the host rocks is approximately equal, so if the
chemical weathering rates of these two elements are roughly
similar (w238U ⬃ w232Th), the gross differences in their activities in the water must be generated by adsorption-precipitation
reactions. If the Th is precipitated out irreversibly at the solubility limit, then Th will build up on surface coatings. Application of this model requires the Th activity in the surface
coating (due to precipitation) to grow over time with the
continuous weathering of the host rocks in the aquifer. For the
total amount of 232Th deposited on such a surface coating, it
was shown by TWPB that:
␳ sc232ThA scS ␰ ⫽ ␳ rw 232Th232ThA rt
(5)
where 232ThAsc is the specific activity of the surface coating, S␰
is the volume of surface coating per unit mass of aquifer rock,
␳sc is the density of the surface coating, and t is the time
required to build up the deposited 232Th. To store the equivalent of ⬃10% of the 232Th activity in the rock in the surface
coating gives w232Tht ⫽ 10⫺1. For a weathering rate of w232Th
⬃ 10⫺13s⫺1, this requires a time of 3 ⫻ 104 years. The decay
product of 232Th is 228Ra, which then decays to 228Th. We
consider that the Ra is in solution and on the surface coating but
in exchangeable sites. This is distinct from the behavior of Th
isotopes, which are virtually all in the surface coating and not
in exchangeable sites. The daughter of 228Ra is 228Th (␶1/2 ⫽
1.9a), which will also be co-precipitated with the other Th.
The surface coating should thus have an inventory of 234Th
in steady state following the equation:
␭ 234Th␳ sc234ThA sc ⫽ ␳ r共w 234Th ⫹ ␧ 234Th␭ 234Th兲 234ThA r
(6)
The case for 230Th is different, as it will not be in steady state
due to its long mean life. The governing equation is:
␳ sc230ThA scS ␰ ⫽ ␳ r共w 230Th ⫹ ␭ 230Th␧ 230Th兲 238UAr•t.
The ratio of 230Th to
Eqn. 5 and 7:
232
(7)
Th in the surface coating is thus from
230Th
A sc/ 232ThA sc ⫽
共w 230Th ⫹ ␧ 230Th␭ 230Th兲 238UA r
w 232Th232ThA r
(8)
For high weathering rates (w230Th ⱖ 3 ⫻ 10⫺15s⫺1),
␭230Th␧230Th ⬍⬍ w230Th, it follows that the ratio of 230Th/232Th
in the surface coating should be:
230Th
A sc/ 232ThA sc ⬇ w 230Th/w 232Th238UA r/ 232ThA r
(9)
As the ratios on the right hand side are approximately unity, the
activity in the surface coating for these long-lived isotopes of
Th should be approximately equal for this case. This follows
the more detailed arguments presented by TWPB (see “Discussion and Conclusions”). Note, however, that for low weathering rates, the recoil contribution for 230Th in Eqn. 8 may
overwhelm the weathering term. We use these relationships in
discussing Ra and Rn in the following sections.
4.4. Application of the Model to Ra Activities
In the arguments given for Th, it was shown that 230Th is
precipitated out into non-exchangeable sites. The radium species as seen from the similar activities must certainly be on
exchangeable-reactive sites. As shown by several other workers, most of the Ra is on these sites and not in solution. We can
easily see this by first treating the problem without including
the Th in non-exchangeable sites.
The equilibrium 226Ra activity according to the model (equation 24 TWPB) without Th precipitation is:
226Ra
A w⬁ ⫽
␳ r共1 ⫺ n兲 共␧ 226Ra ⫹ 0.5␧ 230Th兲 230Th
Ar
␳ wn
␹ 226Ra
(10)
where the term ␹226Ra is the ratio of the number of 226Ra atoms
in the surface coating divided by the number of atoms in the
water per unit volume of aquifer rock. Here, iAj is the activity
of species i in phase j (j ⫽ rock, water). The term ␧i is the
fraction of “i” nuclei that recoil from grains or surfaces into the
solution. For 230ThAr ⬃ 1800 dpm/kg, n ⬃ 0.2, ␳r ⬃ 2, this
yields:
226Ra
A w⬁ ⬇
2 ⫻ 0.8共␧ 226Ra ⫹ 0.5␧ 230Th兲
1.8 ⫻ 10 3
0.2 ␹ 226Ra
⬇
共␧ 226Ra ⫹ 0.5␧ 230Th兲 ⫻ 1.4 ⫻ 10 4
(11)
␹ 226Ra
For the sum of the recoil terms, we roughly estimate ⬃10⫺2.
Using a typical 226Ra concentration of ⬃10⫺1 dpm/kg, we
obtain ␹226Ra ⬃ 1.4 ⫻ 103. Considering the estimated parameters, this is in reasonable agreement with the estimate by
TWPB who give ␹226Ra ⬃ 700 for the exchangeable site model.
These considerations also apply to the other Ra isotopes and
confirm the well-established fact that Ra is highly surface
reactive. It appears that most of the Ra is adsorbed on reactive
surfaces. Note that this treatment does not include 226Rn
sources from irreversible precipitates on the surface layer (e.g.,
230
Th).
We now consider the Ra isotopes using a model of irreversible Th precipitation with ⬃1/10 of the 232Th in the aquifer
rock being stored on the surface layer, which requires w232Tht
⬃ 10⫺1. This is assumed to be from a first or earlier stage of
U, Th series in an aquifer or transport of U, Th series in an aquifer
1969
Fig. 8. [226Ra/228Ra] vs. [224Ra/228Ra]. Dashed vertical line marks secular equilibrium for [224Ra/228Ra] ⫽ 1. Although
Ra/228Ra are close to equilibrium and compatible with the model considering recoil effects (see section 4.4), the
depressed values of [226Ra/228Ra] require special consideration.
224
more rapid weathering when the rock minerals are initially
subject to chemical weathering. The weathering rate per second
is w232Th ⬇ 3 ⫻ 10⫺12/tkas⫺1, where tka is the time scale in
thousands of years. It follows from Eqn. 5 that 10 to 100 ka of
weathering of the aquifer rock would be sufficient to provide
this inventory (i.e., w232Th ⬃ 3 ⫻ 10⫺13 to 3 ⫻ 10⫺14s⫺1)). If
we now consider that the inventory of Ra isotopes on the
surface layer are mostly governed by the decay of 232Th, 230Th,
and 228Th contained in the surface layer, then:
Aw ⫽
␳ sc
关 Rak̂ ⫺1 ⫹ f 228Ra␰␭ 228Ra兴 232ThA *sc
␳ wRak̂ 1
(12)
Aw ⫽
␳ sc
关 Rak̂ ⫺1 ⫹ f 226Ra␰␭ 226Ra兴 230ThA *sc
␳ wRak̂ 1
(13)
Aw ⫽
␳ sc
关 Rak̂ ⫺1 ⫹ f 224Ra␰␭ 224Ra兴 232ThA *sc
␳ wRak̂ 1
(14)
228Ra
226Ra
224Ra
We assume that 228Th and 232Th are in secular equilibrium. The
terms Rak̂⫺1 and Rak̂1 are the rates of deposition (cm s⫺1) into
solution and from solution onto the surface layer. The term fiRa
is the fraction of iRa nuclides that are produced by decay of
their parent in the surface coating and are directly lost by recoil
to the water. The lifetime of 224Ra is much shorter than the
other Ra isotopes, and this term may be significant compared
with Rak̂⫺1. The terms f␰␭ for both 228Ra and 224Ra are negligible due to their longer lifetimes. The term representing the
ratio of atoms on the surface to those in solution (␹Ra) for this
version of the model is different from the value calculated
before without a precipitated Th source (cf. Eqn. 10) as the
dominant source is taken to be in the surface coating. In the
present case, we must now consider the precipitated Th isotopes as a source. This corresponds to ␹Ra ⬃ 2 ⫻ 104 compared
to the value as calculated before (i.e., ␹Ra ⬃ 103).
It is evident that the assumption of irreversible precipitation
of Th isotopes in a surface coating by an earlier stage of
weathering could provide the observed inventory of Rn and Ra.
This still requires substantial loss of U and Th from the source
that produced the surface layer, but it does not demand special
leakage paths for both Ra and Rn. We consider this to be a
plausible explanation of the observations. A detailed calculation of the parameters in this treatment could be done for every
Ra-Rn datapoint. The particular choice of the amounts on the
surface coatings and the parameters would change, but the
general qualitative and quantitative agreement would not be
altered.
The observed ratios of 224Ra/228Ra and 226Ra/228Ra are
informative (see Fig. 8). While four of the samples have
224
Ra/228Ra ⬇ 1, two of them have significantly higher values
(well 1 is 1.25 ⫾ 0.02; well 8 is 1.65 ⫾ 0.01). These would
require 共 f224Ra␰␭224Ra兲/Rak̂⫺1 ⬃ 0.2 to 0.7 (see Eqn. 15). This
may be readily obtained with a plausible choice of values for
the parameters (e.g., f ⬃ 10⫺2, ␰ ⬃ 10⫺4 cm, Rak̂⫺1 ⬃ 3 ⫻
10⫺10s⫺1). A much more difficult problem relates to the ratio
1970
B. C. Reynolds, G. J. Wasserburg, and M. Baskaran
226
Ra/228Ra. The 226Ra comes from the decay of 230Th in the
U series, while 228Ra is from the direct decay of 232Th. As
can be seen from Eqn. 13 and 14,
238
冊
(19)
Ar
⬇ 10共w 230Th/ ␭ 230Th ⫹ ␧ 230Th兲
Ar
(20)
230Th
A sc
⫽ 共w 230Th/ ␭ 230Th ⫹ ␧ 230Th兲 ⫻ 10
232Th
A sc
冉
238U
Ar
Ar
232Th
It can be seen that:
关 Rak̂ ⫺1 ⫹ f 226Ra␰␭ 226Ra兴 230Thsc
226
Raw/ 228Raw ⫽ Ra
关 k̂ ⫺1 ⫹ f 224Ra␰␭ 224Ra兴 232Thsc
226
(15)
228
Since the decay constants for Ra and Ra are over two
orders of magnitude smaller than that of 224Ra, it is not possible
to expect the second terms in the numerator and denominator to
be significant considering the estimate shown above for the
corresponding terms for 224Ra. Thus we would expect the
observed ratio to reflect 230Thsc/232Thsc ⬇ 238Ur/232Thr ⬇ 1.
We can find no changes in parameters that would explain the
low ratios observed (0.24 – 0.59). While we might appeal to an
original source with [238U/232Th] ⬃ 1/4, there is no evidence in
favor of such material in the aquifer rocks. If there is precipitation over sufficient time (⬎ ␶230Th), which is then stopped for
138 ka, then the 232Th is decreased by exp(⫺138 ka ⫻ 0.693).
This will give a decrease in 230Th used in Eqn. 15 and will give
a lower 226R/228Ra by that factor. A simple and intriguing
solution to this problem might be that the original weathering
that produced the surface coating of 230Th and 232Th ceased
⬃138 ka ago, which would give lower (230Th/232Th)sc. This
would result in a decrease of the 230Th without affecting the
232
Th. The 230Th would have decayed since the cessation of Th
precipitation and would match the observations. The remainder
of the discussion presented above would be unaffected. An
alternative model would be that the original weathering went
on for a total time of ⬃28 ka and then stopped without the
230
Th reaching a steady-state value. This would also give a low
230
Th, as it would not be in secular equilibrium. The latter case
cannot be in the regime of the present aquifer, which has very
low weathering rates. This follows from Eqn. 8 because:
冉
共 230Th/ 232Th兲 sc ⫽ 1 ⫹
冊
␭ 230Th␧ 230Th 230 232
共 Th/ Th兲 r
w 230Th
(16)
which
would
give
共230 Th/232 Th兲sc ⬃ 共3 to 30兲
230
232
⫻ 共 Th/ Th兲R which would violate the observed
226
Ra/228Ra ratios and would require extremely long times to
deposit the required amount of 230Th and 232Th. We infer that
the Th coatings were either produced at an earlier stage of
relatively rapid weathering (w ⬃ 10⫺13 to 10⫺14s⫺1), which
ceased roughly 130 ka before the present aquifer developed or
that the rapid weathering went on for 28 ka before the formation of the present aquifer waters.
An alternative interpretation is found if the original condition
of the surface coating of the aquifer had been at steady state
with precipitation of Th. In that case, the following relationships apply:
冉
␳ sc230ThA scS ␰ ⫽ ␳ r
230Th
232Th
w 230Th
⫹ ␧ 230Th
␭ 230Th
冊
238U
Ar
A sc
(w230Th/ ␭ 230Th ⫹ ␧ 230Th) 238U
⫽
Ar
A sc
w 232Th
(17)
(18)
If we again assume that the surface coating contains ⬃10% of
the total Th from the previous weathering, then:
冉
230Th
232Th
冊冒冉 冊
A sc
A sc
238U
232Th
For this ratio to be around unity requires that w230Th ⬃ 3 ⫻
10⫺14s⫺1 (i.e., a very reasonable, normal weathering rate). To
get a value of ((230ThA sc )/( 232ThA sc )) ⬃ 1/4 would correspond
to weathering rate of 7 ⫻ 10⫺15s⫺1. Considering uncertainties,
this is almost identical with the value of w⬘238U ⫽ 2 ⫻
10⫺14s⫺1 found for the vadose zone from the 238U concentration. It follows that an alternative explanation for the low
[224Ra/228Ra] values would be that the Th precipitation took
place over long times (t ⬎⬎ 1/␭230Th) and reached steady state
at a weathering rate of w ⬃ 7 ⫻ 10⫺15s⫺1. This would not
change the results for the daughters of 232Th. In this case, the
time restrictions on surficial weathering would be removed.
However, the strong contrast between the present vadose waters and the old groundwater still requires input of low U waters
into the aquifer.
5. DISCUSSION AND CONCLUSIONS
The Ojo Alamo aquifer, which is in an arid region, represents
a hydrologic system that is both old (up to 25 ka) and slow
moving. The isotopic composition of U in the groundwater is
highly enriched in 234U relative to the equilibrium value. The
typical value in the aquifer is ␦234U ⬃ 6000‰. These groundwaters also exhibit low 238U concentrations as a result of low
weathering rates within the aquifer. The U concentration is the
same in both the filtered and unfiltered water samples and
appears to be essentially completely in solution. The high
␦234U reflects the effects of a normal amount of 234Th recoil
and low weathering rates following the relationship:
␦ 234U ⫽
␭ 234U␧ 234Th
⫻ 10 3
w 238U
(21)
where w238U ⬃ 2 ⫻ 10⫺18 to 2 ⫻ 10⫺17s⫺1 and ␧234Th ⬃ 2 ⫻
10⫺4 to 2 ⫻ 10⫺3. In contrast to the aquifer waters, the present
vadose zone waters, springwaters, and riverwater have relatively low ␦234U ⬃ 500‰ and high 238U concentrations. This
is supported by laboratory leaches of vadose zone material with
water. It is not possible to evolve the groundwaters from the
present vadose zone input with high U concentrations unless
there is an almost complete removal (⬃90%) of U before
introduction into the aquifer. There is no evidence for such a
process of removal and storage of U in this system. It follows
that the present inventory of water in the Ojo Alamo aquifer
had to be derived under very different conditions. This requires
that the original vadose zone input responsible for providing
the present aquifer had to have low U concentrations, at least
one decade below the current vadose zone input. The temporal
change inferred for the vadose zone cannot be attributed to
anthropogenic processes, as the area has very little agricultural
or industrial activity (including U mining), and ore processing
is not associated with this area. If these assumptions are correct,
then we must conclude that there has been a natural change that
U, Th series in an aquifer or transport of U, Th series in an aquifer
occurred several thousand years ago in the nature of the vadose
waters. One possibility is that this difference is the result of a
climatic change where before ⬃5 ka ago there was much less
evapotranspiration and more rainfall. This could provide much
lower 238U concentrations in the vadose waters. This earlier
state would have had to persist for at least 10 ka to fill the
aquifer. The extent to which this hypothesis is consistent with
paleoclimatic records in the southwestern United States must
be evaluated. The observed effects on the U concentration and
[234U/238U] are clear. In recent papers, Polyak and Asmeron
(2001) and Polyak et al. (2001) proposed a wetter and cooler
late Holocene climate in the southwestern United States based
on a study of mites preserved in stalagmites. There is thus some
independent evidence for this proposal. However, we cannot
argue that the current net precipitation rate is inadequate to
provide the water in the Ojo Alamo over the past 104 years
The 232Th concentrations in filtered waters were found to be
approximately at the saturation level. This is in accord with the
results of Tricca et al. (2001) on an aquifer at the BNL Long
Island site. The precise value for saturation with ThO2 (Langmuir and Herman, 1980) depends on the concentration of other
ligands. The observed ratio 232Th/238U in filtered waters is
many orders of magnitude below the ratio in aquifer rocks. The
232
Th/238U ratio in unfiltered water samples is also far below
that in the rocks. It was also found that the ratio [228Ra/232Th]
in the waters is far below the ratio necessary to provide the
[228Ra] in the waters, requiring a predominant source for 228Ra
from the rock and surface coatings. These observations support
the idea that the Th concentration in natural waters is most
commonly in saturation equilibrium (TWPB). Any Th produced by weathering once the saturation condition is achieved
and maintained will then be irreversibly precipitated. In that
case, all models of transport are invalid if only surface coatings,
which are in reversible equilibrium and with exchangeable sites
for Th, are considered. The transport of Th and any of the
daughter elements must therefore be modified to include the
irreversible precipitation of Th isotopes. This greatly modifies
the equations governing a large number of nuclides (see
TWPB). The 232Th/238U ratio in waters is thus not simply
related to the weathering rates but must also involve the saturation condition. If 232Th/238U in water is not the ratio of the
weathering rates, then 关232 Th/238 U兴 ⫽ w232Th/w238U (see TWPB,
section 5). The [232Th] in the water is simply the 232Th of a
saturated solution. Changes in the chemical state of the aquifer
(including fO 2 and the presence of complexing ligands) may, of
course, drastically change the solubility limits and the role of
irreversible precipitation to reversible exchange with the solution.
If we first ignore the irreversible precipitation of Th and
consider only exchangeable sites on surface coatings (i.e.,
ignoring Th precipitation), it is found that almost all (99.9%) of
the Ra must reside in the coatings (i.e., ␧Ra ⬃ 103). This
calculation simply uses the flow model with weathering rates,
recoil, and the bulk-rock activity (cf. Eqn. 10 and 11) with
exchangeable sites for Ra. The Rn found in all aquifer water
samples requires a source that is far in excess of what would be
produced by recoil from and weathering of the aquifer rock. If
one considers 222Rn, then it is found that in the filtered waters
[222Rn/226Ra] ⬃ 10⫺3, which also requires that most of the Ra
(⬃99.9%) is on surface reactive sites. This is in accord with
1971
box model calculations done by other workers (cf. Ku et al.,
1992; Cowart and Burnett, 1994; Andrews et al., 1989; Krishnaswami et al., 1982). The radon isotopes, 228Ra and 224Ra, are
found to be roughly in equilibrium with small effects from
recoil enhancing 224Ra. The ratio [226Ra/228Ra] is low (⫻1/4)
compared to what would be expected for a source with
[238U/232Th] ⬃ 1. This requires that there be a shortfall in
226
Ra production (see below). Grossly, the Ra isotopes are
essentially well behaved with regard to normal rock sources.
The 222Rn and Ra concentrations each require sources corresponding to ⬃10% of the production in the bulk aquifer rock.
This would require special leakage paths both for Rn and Ra.
We now consider the results for 232Th and the conclusion
that it is saturated in the water and, hence, the role of nonexchangeable surface sites for Th. If saturated, then the 232Th
in a unit volume of aquifer is the amount in the rock plus the
amount precipitated on the surface of the grains in that volume,
the amount in the fluid being the saturation concentration times
the mass of water in the pores. This amount is negligible. The
amount of 232Th in the surface precipitate is proportional to the
weathering rate and the time duration of weathering. The activity of Th in the surface is 共␳r w232Th 232ThAr)t. As 230Th will
also be precipitated, then such a surface layer will contain both
isotopes. As shown by TWPB, a surface layer resulting from an
earlier stage of weathering could provide a source for both the
Ra and Rn.
It was noted above that [226Ra/228Ra] was below what was
expected from rock sources with [238U/232Th] ⬃ 1. This could
be achieved if the initial weathering that produced the precipitate Th stopped ⬃130 ka, allowing the 230Th (parent of 226Ra)
to be depleted by a factor of 4. Alternatively, it could be the
result of surficial weathering achieving steady-state deposition
of 230Th with a high weathering rate of w232Th ⬃ 7 ⫻
10⫺15s⫺1.
In summary, this study shows that the transport of U,Th
series nuclides in a slow-moving aquifer can be reasonably
described by a simple flow and reaction model. Physical insights into the hydrologic-chemical dynamics may be gained by
combining adequate datasets with consideration of theoretical
models. Further studies in aquifers with more detailed knowledge of hydrologic flow from the vadose zone inputs through
flow paths in the aquifer is required. Under the condition that
roughly 10% of the original Th was weathered in an earlier
weathering stage, then this would provide for the Ra and Rn in
the aquifer waters and surface coatings. As the coatings must be
very thin, a ready path for leakage of the rare gas Rn is
available and the release of Ra isotopes into the water can
easily take place. The near equilibrium of Ra isotopes is maintained despite their different lifetimes. The observations presented here are in strong support of this model.
Acknowledgments—This work was supported by DOE DE-FG0388ER13851. Caltech Division Contribution No. 8775(1086). Instruction in the arts of U-Th chemistry and mass spectrometry was kindly
given by Dr. J. Chen. We acknowledge the gracious aid in initial field
support from Fred Phillips and Doug McGhee (New Mexico Institute
of Technology) and Dale Worth from the Bureau of Land Management,
Farmington. In particular, Fred Phillips made critical contributions to
our understanding and to this report. Comments by the referees and the
associate editor on both substance and style were very helpful. For
kindly providing access and support, we thank members of the Navajo
1972
B. C. Reynolds, G. J. Wasserburg, and M. Baskaran
Agricultural Products Industry (NAPI) Feedyard, The Carson School,
the Navajo Indian Irrigation Project (NIIP), and especially the Navajo
Tribal Utility Authority (Shiprock). Field assistance by Sarah Trimble
of Wayne State University was of great help.
Associate editor: G. Helz
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