96 (1992) 183-202
Elsevier Science Publishers B.V., Amsterdam 183
Formation of alunite, jarosite and hydrous iron oxides in a
hypersaline system: Lake Tyrrell, Victoria, Australia
D.T. Long a'~, N.E. Fegan a'~, J.D. McKee a, W.B. Lyons b'a, M.E.
Hines b and P.G. Macumber c
~Geological Sciences, Michigan State University, East Lansing, M148824, USA
blnstitute for Earth, Oceans. and Space, University of New Hampshire, Durham, NH 03824-3589, USA
CDepartment of Water Resources, Melbourne, Vic. 3000, Australia
(Accepted for publication August 30, 1991 )
Long, D.T., Fegan, N.E., McKee, J.D., Lyons, W.B_, Hines, M.E. and Macumber, P.G., 1992. Formation ofalunite, jarosite
and hydrous iron oxides in a hypersaline system: Lake Tyrrell, Victoria, Australia. In: W.B. Lyons, D.T. Long, A.L.
Herczeg amd M.E. Hines (Guest-Editors), The Geochemistry of Acid Groundwater Systems. Chem. Geol., 96:183-
(KAI3(SO4)2(OH)6) and jarosite
(KFe3(SO4)2(OH)6) are common weathering products ofaluminosilicates
and pyrite. Long-range transport of the constituents forming these minerals and the subsequent deposition of alunite and
jarosite in an evaporite setting have not been previously documented. Such conditions for the occurrence of alunite and
jarosite were investigated in a hypersaline system where acid groundwater enriched in K, AI, Fe (11I) and SO4 enter a salt-
playa lake. Sediment cores were studied by thin section, XRD and SEM-EDS. Groundwater and pore-water chemistries
were also analyzed. The results show that alunite and jarosite occur together or as individual layers and cements within the
top 20 cm of the sediments, where the groundwater is most concentrated due to evaporation. Hydrous Fe-oxides also occur
as cements or vein fillings with the alunile and jarosite, but are distributed throughout the 70-cm sediment column studied.
The results are consistent with a model in which alunite and jarosite precipitate as a result of the evaporation of water.
Thermodynamic modeling indicates that the pore water in the playa sediments maybe in equilibrium with alunite, jarosite
and hydrous Fe-oxides.
(KFe3 ( 504 ) 2 (OH) 6 ) form under
pH conditions and are often associated with
acid-mine drainage, weathering of sulfide ore
deposits, hydrothermal alteration, oxidation of
pyrite, or weathering in acid soils (Hemley et
al., 1969; van Breemen, 1973; Knight, 1977;
Nordstrom, 1982; Sullivan et al., 1986; Michel
correspondence should be directed to senior author.
~Department of Geological Sciences, California State
University-Hayward, Hayward, CA 94542, USA.
aHydrology/Hydrogeology Program, University of Ne-
vada, Mackay School of Mines, Reno, NV 89557, USA.
and van Everdingen, 1987). Alunite forms
during both hypogene and supergene pro-
cesses, whereas jarosite forms mainly during
the latter (Brown, 1971; Hladky and S!ansky,
1981; Stoffregen and Alpers, 1987; Rye et al.,
1989). Hypogene jarosite has been found to
occur at several hydrothermal deposits, but is
less common than hypogene alunite in zones
of acid-sulfate hydrothermal alteration (C. AI-
pers, pers. commun., 1989). In low-tempera-
ture environments the formation of these min-
erals has also been related to microbial
oxidation of iron and sulfur (Adams and Ha-
jek, 1978; Ivarson et al., 1979). The mecha-
nisms controlling the formation of alunite are
of interest because of its potential for seques-
0009-2541/92/$05.00 © 1992 Elsevier Science Publishers B.V. All rights reserved.
D,T. LONG ET AL.
tering A1 and SO4 in soils and its importance
as a source of Al for industry (Adams and Ra-
wajfih, 1977; Adams and Hajek, 1978; Ivarson
et al., 1979; Dutrizac and Jambor, 1987). Both
minerals have been studied as indicators of the
occurrence and potential mass of sulfide de-
posits (Bladh, 1982) and for their ability to
scavenge elements from solution (Ivarson et
al., 1979; Dutrizac and Jambor, 1987 ).
Alunite and jarosite comprise a family of
minerals with the general formula
AB3(XO4)E(OH)6, where A is a large cation
(e.g., K +, Na +, Rb +, NH4 +, Ag +, H30 +,
Ca 2+, Pb 2+, Ba 2+, Sr 2+, Ce 3+ ), B is an octa-
hedrally coordinated cation (e.g., AI s+, Fe 3÷,
Cu 2+, Zn 2+) and XO~4- is an anion (e.g.,
SO42-, PO43-, AsO 3-, CrO 2- ). Thus, the for-
mation of these minerals can affect the geo-
chemistry of many elements. The classifica-
tion of the minerals in the family is based on
whether AI 3+ or Fe 3+ is the dominant cation
in the B site, forming the alunite and jarosite
supergroups, respectively (Scott, 1987). Ex-
tensive solid solution between and within these
supergroups has been demonstrated experi-
mentally and by direct measurement of natu-
ral minerals (Brophy et al., 1962; Hartig et al.,
1984; Scott, 1987).
The source of acidity in solutions that pre-
cipitate alunite or jarosite can be the oxidation
of sulfide, particularly pyrite or H2S, precipi-
tation of hydrous Fe-oxides, and possibly large
inputs of acid rain into poorly buffered sys-
tems (Sass et al., 1965; Hemley et al., 1969;
Nordstrom, 1982; Mann, 1983). The oxida-
tion of pyrite also supplies Fe 3 ÷ and SO4. Ad-
ditional sources of SO4 can be acid-sulfate soil
water, acid rain and dissolution of sulfate min-
erals (van Breemen, 1973). The acid weath-
ering of clays and other aluminosilicate min-
erals mobilizes A1 for alunite formation as well
as K, which is a constituent of both alunite and
jarosite. K also may be derived from cation-
exchange reactions. Both SOn and K can origi-
nate from seawater and their concentration in
solution by evaporation could be related to the
formation of alunite and jarosite (Goldbery,
The principal mechanism for the formation
ofjarosite and alunite is thought to be by dis-
solution/re-precipitation reactions involving
pyrite and aluminosilicates. In some cases these
minerals precipitate after "local" migration
(on the order of I m) of the constituents (i.e.
K, A1, Fe, SO4) (Hemley et al., 1969). Long-
distance migration of the constituents prior to
their precipitation as alunite or jarosite min-
erals has not been previously documented, with
the exception of Michel and van Everdingen
( 1987 ), who have demonstrated such a case for
a jarosite deposit in shale. In this case jarosite
is precipitating from a groundwater seep. The
constituent elements originate from water-rock
reactions along the groundwater flow path.
In the southern half of the Australian conti-
nent, acid groundwater occurs over large areas
and migrates long distances (Macumber,
1983 ). This regional groundwater carries large
amounts of Fe, Si and A1 (Macumber, 1983;
Mann, 1983 ). In areas where this groundwater
discharges into playa lakes, evaporite minerals
form as the water evaporates (Long et al., 1992
in this issue ). Alunite, jarosite and hydrous Fe-
oxides are frequently found with the evaporite
minerals. The association of alunite and jaros-
ite with hypersaline environments in which
evaporite minerals are precipitating has been
little studied (Lock, 1988).
This study investigates the occurrence of al-
unite and jarosite in hypersaline sediments of
Lake Tyrrell, Australia. An evaporative origin
for these minerals is proposed. Their associa-
tion with Fe-oxides and the possibility of so-
lution-mineral equilibrium in the alunite-ja-
rosite and alunite-j arosite-Fe-oxide
assemblages are explored.
2. Study area
The geologic history, sedimentology and hy-
drogeochemistry of the Lake Tyrrell system has
been discussed by Teller et al. (1982), Mac-
ALUNITE, JAROSITE AND HYDROUS IRON OXIDES IN A HYPERSALINE SYSTEM 185
umber (1983, 1992 in this issue), Bowler
(1986), Bowler and Teller (1986), and Long
et al. (1992 in this issue); and will be only
briefly discussed here. Lake Tyrrell is a 26 × 7
km salt-playa lake located in the Murray Basin
of NW Victoria, Australia (Fig. 1 ). Lake Tyr-
rell became dry ~ 32 kyr BP, but is covered
with 1 O's of cm of water for ~ 3 months of the
year. The lake deposits are up to 5 m thick and
consist of clay, silt and sand covered by an
ephemeral halite-gypsum crust (30-60 cm).
Decaying algal mats occur locally on the lake
The lake beds are underlain by a large aqui-
fer system formed by the Parilla Sand. This
aquifer is Late Miocene to Pliocene in age. In
the study area the aquifer averages 60 m thick.
The quartz sand appears to have been depos-
ited in a marginal marine environment and
contains heavy minerals and organic matter.
The upper half of the Parilla is coarse-grained
sediment, while the sediment of the bottom
half is more fine grained. Most of the organic
matter and pyrite occur toward the base of the
aquifer, but this trend is not related to the
grain-size differences of the sand.
Within the study area the Parilla is a con-
fined to semi-confined aquifer. It is mostly
overlain by the Plio-Pleistocence lacustrine
Blanchetown Clay and underlain by the Upper
Oligocene-Lower Miocene marine Geera Clay.
The Blanchetown Clay, which is light to olive
grey with scattered sand lenses, is thickest (20
m) in the northern area of Lake Tyrrell and is
absent towards the south. In the southern area
the Tyrrell beds directly overlie the Parilla
Sand. The Geera Clay is dark grey to black,
~ 100 m thick and less permeable than the
Blanchetown Clay in the Lake Tyrrell region.
Two smaller playa lakes, Wahpool and Tim-
boram, occur east of Tyrrell. The regional
groundwater flow in the Parilla Sand is west-
ward (Fig. 1 ), interacting with the three playa
lakes as springs or as stream input. Lake Tim-
boram is at the highest elevation and is an open
or flow-through lake, whereas Lakes Tyrrell and
Wahpool are closed lakes. Groundwater enters
Lake Tyrrell from springs around the lake
88 Groundwoter Flow Line
S rlng •
" S cT-9
I "-- f2 ,AKE
,A ', ...... :_ ~ ~ t'~ J --t (. ") TIMBORAM
54 B2 groundwater ,N, I ~ J
/ ', ~g'
• ', ~-"~
~o26 6032 e2 5S-B ) ~ \
I, \ ..
~/- " Tyrrell Creek
o s Jo ,"- K) _.---.J ':
B 1 I I ~wn nf'~ TCST e
Groundwater Flow Line
Fig. 1. Study area location showing well distribution, general path of regional groundwater flow and zone of non-acid
186 D.T. LONG ET AL.
shore. A major groundwater divide occurs ~ 10
km west of Lake Tyrrell. Thus, groundwater
entering Lake Tyrrell along the western shore
has its origin in the regional groundwater flow
from the east (Fig. 1 ).
The groundwater in the Tyrrell system has
been characterized by salinity and location as
three major types (Macumber, 1983):
( 1 ) Regional groundwater from the Parilla
Sand which is similar in salinity to seawater
(3.5%). The water is oxic with pH < 4 and
enters Lake Tyrrell as seeps along the western
shore (Fig. 1 ).
(2) Brine which saturates the Parilla Sand
aquifer below Lake T~crrell and is formed by
reflux of evapo-concentrated water. This water
is more saline (25%) than regional groundwa-
ter. The water is suboxic to anoxic with near-
(3) Brine beneath Lakes Tyrrell and Wah-
pool which is also formed by evapo-concentra-
tion of regional water. The salinity of this brine
(12%) is about half that of the Tyrrell reflux
brine. This water is also suboxic to anoxic with
near-neutral pH and enters Lake Tyrrell from
springs along the eastern shore (Fig. 1 ).
Long et al. ( 1992 in this issue) conclude that
these three types of water masses have evolved
from a common parent with sea salt as the
source for the solutes. They suggest that pro-
cesses that occur early in the evolution of the
water such as calcrete and gypcrete formation
affect the chemical signature of the water. Aci-
dification of the water appears to happen early
in the geochemical evolution of the water and
is initiated by the oxidation of pyrite. Locally,
acid conditions are enhanced by the precipita-
tion of alunite, jarosite and hydrous Fe-oxides,
The wells established by Macumber ( 1983 )
were used to sample the three types of ground-
water discussed and are shown on Fig. 1. Mac-
umber ( 1983 ) put in a line of piezometers (TP
line) to sample the low-pH, oxic regional
groundwater entering the lake as springs on the
western shore (Fig. 2). These piezometers are
l m deep and cross the spring zone. Two of the
piezometers are outside the spring zone and
sample the Lake Tyrrell reflux brine (Fig. 2).
In order to characterize the spring zone water
further, a second piezometer line (NPZ line)
was established across the spring zone ~ 150
m north of the first line (Fig. 2 ). The chemical
data indicate that the Lake Tyrrell reflux brine
was not sampled by this new piezometer line.
Pore water in the spring zone was also col-
lected utilizing in situ pore-water samplers
called "sippers" (Fig. 2). These are basically
lysimeters made of Teflon ® with a porous Te-
flon ® sleeve (2-cm diameter) which can be
placed at various depths in the sediment
(Hines et al., 1992 in this issue). A vacuum
using an inert gas (N2) was drawn on the "sip-
per" and pore water was collected in situ. Sip-
pers were placed at each site in 0-3-, 4-9-, 10-
15- and 15-20-cm depth intervals. The data set
includes 89 samples from wells, piezometers
The wells and the piezometers were sampled
with either a Teflon ® bailer or a peristaltic
pump. The piezometers were pumped dry be-
fore sampling. Once retrieved, all samples for
o.. ,,,oL,.r. ,.o
Fig. 2. Spring zone with locations of piezometers and
ALUNITE, JAROSITE AND HYDROUS IRON OXIDES IN A HYPERSALINE SYSTEM 187
A1, Fe and SO4 analyses were maintained and
atmosphere in glove
bags. All samples were immediately filtered
through an acid-rinsed Teflon ® filtering sys-
tem using acid-rinsed Nuclepore ® 0.4-/~m fil-
ters for trace-metal and minor-element sam-
ples and Millipore ® 0.45-/1m filters for the
major elements. Except for the samples for al-
kalinity, SO4 and other anion determinations,
all samples were acidified to a pH < 2 using
ULTREX ® HNO3. Samples were stored in
acid-rinsed polyethylene bottles. These bottles
were rinsed thoroughly with deionized-dis-
tilled water before use. Samples for AI deter-
minations were handled in a similar manner to
the methods described by Mackin and Aller
(1984). These samples were acidified in the
laboratory 24 hr prior to A1 analysis in order to
minimize potential contamination from the
sample bottle and to maximize recovery of
monomeric A1 (J.E. Mackin, pers. commun,,
1987). The ULTREX ® HNO3 for acidifica-
tion of the AI samples had to be cleaned by sub-
boiling distillation in a Teflon ® evaporation
system. Lyons et al. (1992 in this issue) fur-
ther describe the clean procedures used in this
Chloride was determined by Mohr titration
and bromide was determined by colorimetry
using a modification of the chloramine-T
method proposed by Presley (1971) as de-
scribed by Long and Gudramovics (1982).
Other analytical methods and results for the
major- and minor-element determinations are
decribed by Long et al. (1992 in this issue).
The major constituents of alunite, jar®site and
the hydrous Fe-oxides in the water samples
were analyzed as follows. Total Fe and Fe (II)
were determined by the Ferrozine ® technique
(St®®key, 1970; Gibbs, 1976; Lyons et al.,
1989) on filtered samples immediately after
collection. Sulfate was determined turbidi-
metrically (Hines et al., 1992 in this issue) and
K by atomic absorption spectrometry (AAS)
using a 1:210 dilution with a NaCI-LaCI3-HCI
matrix added. A1 was analyzed by AAS with
graphite furnace (AASG) using a 1:100 dilu-
tion. Standards for AI analysis were prepared
with Fe and C1 added in similar concentra-
tions to the sample being analyzed. It is as-
sumed that the dissolved AI obtained by this
technique is mainly monomeric. This was in
part confirmed by comparing the results from
the AASG technique to those obtained using
the A1 extraction technique of Barnes ( 1975 ).
This latter technique was used to selectively
extract monomeric AI from test samples. The
results show that similar A1 concentrations
were detected by both techniques.
To examine sediment mineralogy, 60-cm co-
res were taken from the spring zone along each
piezometer line using 9-cm-diameter polyvi-
nylchloride (PVC) pipe. Sediment was ex-
truded from the core and subsampled in the
field. Sediment mineralogy and specifically al-
unite, jar®site and hydrous Fe-oxides were
identified by X-ray diffraction (XRD) (Ri-
gaku ® Geigerflex D/Max-Ia) on bulk samples
using Cu-K, radiation. Alunite, jar®site and the
Fe-oxides were also examined visually and
chemically by scanning electron microscopy
(SEM) using energy-dispersive spectrometry
(EDS) (JEOL ® JSM-T20 with Tracor North-
ern ® EDS). During the sectioning of the sedi-
ment cores, selected sediments were dried and
vacuum impregnated with EPO-KWICK ®
epoxy cement for preparation of thin sections.
The chemical modeling program PHREEQE
(Parkhurst et al., 1980) was used to study the
potential for thermodynamic equilibrium
among alunite, jar®site and hydrous Fe-oxides
in the system. PHREEQE is a mass transfer/spe-
ciation code that uses the data base from
Truesdell and Jones (1974) and Ball et al.
( 1987 ) which has its origin in the Garrels and
Thompson (1962) model for the determina-
tion of activities of solutes in aqueous solu-
tions. We recognize that the ability of this
computer code to define activities at high ionic
strengths is debatable and that the Pitzer-based
(e.g., Pitzer, 1979) codes such as PHRQPITZ
(Plummer et al., 1988) are more suitable for
188 D.T LONG ET AL.
the study of high ionic strength solutions
(Harvie and Weare, 1980). However, Pitzer-
based codes are presently limited in their abil-
ity to describe metals such as Fe 3+ and AI 3+,
particularly in those systems involving oxida-
tion-reduction processes (Reardon and Be-
ckie, 1987) and were deemed not suitable for
4. Results and discussion
4. I. Occurrence of alunite, jarosite and
hydrous iron oxides
In addition to the sediment in the spring zone
along the western shore (Fig. 1 ), sediment was
also examined at the southern end of Lake
Tyrrell where Tyrrell Creek enters the lake and
at spring zones along the eastern shore. Of these
three areas, only in the western spring zone did
alunite, jarosite and hydrous Fe-oxides occur
together. The sediments in the spring zone
consist of bedded silts, sands, and some clays
overlain by a halite-gypsum crust. The crust is
thickest towards the center of the lake outside
the spring zone where it is > 40 cm thick at
piezometer TP0 (Fig. 2). Below the salt crust,
or top 2 cm of the sediment where the salt is
absent, there are bands of reduced organic
matter ~ 5 cm thick.
Within the spring zone, the near-surface sed-
iment occurs as thin layers (0. 1-5 cm) of silt
and some clay colored different shades of yel-
low, white and grey. Below these sediments are
layers of sand and clay, varying in thickness
between 10 and 50 cm, which are generally grey
with patches and streaks of red, yellow and
purple. The sands are compositionally mature,
containing generally in excess of 90% quartz.
In thin section, the quartz is mostly monocrys-
talline and nonundulose. The quartz grains
tend to be angular rather than rounded, so their
texture is not mature. In addition, quartz grains
coated with Fe-oxides are disseminated
throughout the sediments. These grains give the
appearance of feldspar in hand specimens of
the sediment. Few feldspars were actually
identified in the sediments from the spring
zone. Heavy minerals identified in the sedi-
ments include tourmaline, zircon, epidote and
dark opaques probably including magnetite, il-
menite, hematite and others. The clays identi-
fied are illite and kaolinite (Long et al., 1992
in this issue). Clay/silt layers in the lower sec-
tion of cores contain numerous fractures at
various angles to bedding. These fractures are
stained red and yellow and appear to have been
conduits for the upward flow of the oxidized
groundwater from the Parilla Sand. This ob-
servation is consistent with the hydrologic
model proposed for the spring zone system by
Macumber ( 1983, 1992 in this issue ).
Most of the red, purple and yellow streaks
and stains in the sediments are hydrous Fe-ox-
ides that occur either as discrete layers or len-
ses within the sediment or as intergranular ce-
ments. The yellow cement was identified by
XRD as goethite. This was the only Fe-oxide
conclusively identified by XRD; although mi-
nor amounts oflepidocrocite and hematite may
have been present. The red and purple streaks
were X-ray amorphous. Fig. 3 is a photomicro-
graph using crossed nicols of goethite cement
surrounding quartz grains and a microcline
grain. The microcline appears corroded, but the
timing of the corrosion is not known.
Like the Fe-oxides, alunite and jarosite oc-
cur as discrete layers within the sediment and
as cements. The alunite is white to greyish
white while the jarosite is a brilliant yellow.
Both minerals were frequently misidentified in
the field as gypsum and Fe-oxide, respectively.
When the XRD data of the alunite and jarosite
are compared to standard XRD data, these
minerals appear to be more like the K-rich end-
members than the Na-rich end-members (A1-
pers et al., 1989, 1992 in this issue). There-
fore, they would have the formulas of
respectively. The Na-enriched forms of these
minerals were not detected in the sediments.
The predominance of K over Na as the A cat-
ALUNITE, JAROSITE AND HYDROUS IRON OXIDES IN A HYPERSALINE SYSTEM 189
Fig. 3. Photomicrograph under crossed nicols ofgoethite cement surrounding a microcline grain (scale bar= 0.1 ram).
ion was qualitatively confirmed by EDS scans
(presented below). Alpers et al. ( 1992 in this
issue) provide data on unit cell dimensions of
alunite and jarosite based on powder XRD
which further confirm the dominance of the K-
Alunite and jarosite tend to occur in the top
20 cm of the sediment. This is in contrast to
the Fe-oxides, which occur at all depths within
the 70-cm length of the cores. Individual layers
of alunite and jarosite cements, usually < 5 cm,
are as thin as 0.3 cm. Neither these layers nor
the depth of occurrence of the alunite and ja-
rosite cements is laterally continuous through-
out the spring zone.
Although the regional groundwater entering
the spring zone are oxic (Macumber, 1983,
1992 in this issue; Long et al., 1992 in this is-
sue), the sediments through which the ground-
water flows in the spring zone do not appear to
have been completely oxidized. In addition to
the reduced sediment layer at or near the sur-
face, grey unoxidized sediment occurs
throughout the cores in layers from 0.3 to 7 cm
thick. Frequently, jarosite and alunite occur
immediately below these unoxidized sediment
Fig. 4 is a photomicrograph using plane-po-
larized light ofjarosite cementing quartz grains.
Periods of infilling of the jarosite cement are
seen as layering. Under crossed polarizers the
jarosite cement appears to be homogeneous.
No evidence of direct replacement of either
pyrite or aluminosilicates by jarosite was seen
in thin section. It was not possible to obtain a
thin section of the alunite suitable for photo-
microscopy. The alunite cement was not strong
enough to hold the sediment together during
thin section preparation and did not allow the
epoxy cement to penetrate the sediment enough
to bond the sediment and alunite. In addition,
part of the alunite cement disaggregated dur-
ing this preparation. XRD data showed the
presence of halite in the alunite cement, which
may have caused the apparent dissolution.
Under a binocular microscope, alunite, ja-
rosite and the Fe-oxides were clearly seen to
cement the quartz grains of the sediment. Both
alunite and jarosite cements encircled quartz
grains coated with Fe-oxides. More iron-coated
quartz grains were found within the jarosite
cement than in the alunite cement. This obser-
vation is consistent with the observation that
hydrous Fe-oxides t~:nd to occur more fre-
quently with the jarosite than with the alunite
in the sediments. Water was clearly seen to dis-
90 D.T. LONG ET AL.
Fig. 4. Photomicrograph under plane-polarized light ofjarosite cement
0. I mm).
Fig. 5_ SEM photomicrograph of alunite
aggregate the alunite cement, but not the jaros-
ite cement. This is consistent with the ob-
served association of halite with the alunite.
However, we cannot make the statement that
all alunite in the spring zone is associated with
halite whereas all jarosite is not, because we
have XRD data showing all three minerals in
4.2. Characterization of alunite, jarosite and
hydrous iron oxides
Figs. 5-8 show typical SEM images of alu-
nite, jarosite and the hydrous Fe-oxides. Alu-
nite has a sugary texture composed of fine-
grained, relatively equant crystals of < 1 /lm
cross-section (Fig. 5). Jarosite, on the other
ALUNITE, JAROSITE AND HYDROUS IRON OXIDES IN A HYPERSALINE SYSTEM 191
Fig. 6. SEM photomicrograph ofjarosite (scale bar= 2 #m).
Fig. 7. SEM photomicrograph of X-ray amorphous Fe-oxide (scale bar= 1/zm).
hand, tends to occur as hexagonal platelets
which are > 1/zm in cross-section and --, 0.2/zm
thick (Fig. 6). The X-ray amorphous Fe-ox-
ides (Fig. 7) occur as individual spheroids ~ 1
/zm in diameter which also coalesce into more
massive aggregates. This figure shows Fe-ox-
ides which have precipitated on quartz. Fig. 8
is a SEM image of halite. The EDS scans of the
flakes of material on the halite showed the
flakes to be composed ofAl, K and S (Fig. 9).
This indicates that the flakes are alunite and is
consistent with the observation of the close as-
192 D.T. LONG ET AL.
Fig. 8. SEM photomicrograph of halite with alunite flakes (scale bar= 2 ltm).
Fig. 9. EDS scans of minerals shown in Figs. 5-8: (a) al-
unite: (b) jarosite; (c) iron oxide; and (d) alunite flakes
peaks of Fe. It is not clear if these small peaks
reflect solid solution or mechanical mixing of
jarosite and alunite. This uncertainty has also
been noted by Alpers et al. (1992 in this is-
sue). In many cases, both alunite and jarosite
were detected by XRD in samples that ap-
peared monomineralic by macroscopic inspec-
tion. This might suggest that the jarosite and
alunite mainly occur as a mechanical mixture
rather than as a solid solution. However, more
work on the stoichiometry of these minerals
needs to be done before this problem can be
completely resolved. The color of jarosite-al-
unite mixtures is discussed by Alpers et al.
( 1992 in this issue).
4.3. Aqueous solution chemistry
sociation of alunite and halite from the thin
section and XRD results.
The EDS results (Fig. 9) clearly show that
alunite and jarosite are composed of the K end-
member. This was found to be the case for the
EDS scans of all the alunite and jarosite sam-
pies. Small peaks of Al were found in thejaros-
ire scans, and the alunite scans contained small
Table 1 summarizes the aqueous solution
chemistry in regional and spring zone ground-
water for the major constituents of alunite, ja-
rosite and hydrous Fe-oxides including pH; Br
is included because in provides a conservative
natural tracer. The regional groundwater in-
cludes only samples with pH < 5.5, the ap-
proximate division between the acid and near-
Concentrations (in mg 1-' ) of selected constituents
Al Fe(III) K Br SO4 pH
RGW SZW RGW SZW RGW SZW RGW SZW RGW SZW RGW SZW
X 16.3 68.2 8.91 3.74 279 769 151 460 5,442 19,838 3.99 3.47
M 2.4 41.0 5.43 1.94 240 691 110 398 3,084 19,035 3.75 3.56
Mo 2.4 10.2 2.77 1.62 240 1,050 70.1 362 4,150 3,285 3.70 3.60
GM 2.96 38.4 2.77 1.62 220 639 115 373 3,662 14,674 3.93 3.44
R 0.05-76.6 3.19-247 0.1-26.7 0.05-20.5 11.6-738 190-2,080 3.37-467 107-1,190 154-18,789 2,027-66,112 2.9-5.38 2.6-4.25
CF 0.05-7.2 1.6-18.3
RGW= regional groundwater, n = 35; SZW= spring zone ground- and pore waters, n = 23; X= mean; M= median; Mo = mode; GM = geometric mean;
R = range; CF = concentration factor relative to seawater.
194 D.T. LONG ET AL.
j~..d i ,
@ vQp-cur ve
n n n I
0.5 1.25 2 2.75 3.5
Log Br (mg/1)
Fig. 10. Iog,o CI vs. log,o Br. Seawater evaporation curve
from MacCaffrey et al. (1988).
neutral pH groundwater masses. The spring
zone chemistry combines the sipper and pi-
ezometer data within this zone (Fig. 2). Ex-
cept for Fe 3 +, the highest concentrations of the
constituents are found in the spring zone
groundwater where evaporation is occurring.
This indicates that evapo-concentration is in
part controlling the concentration of the con-
stituents. Fe(IlI) concentrations are approxi-
mately the same in both the regional and spring
zone groundwater, as are the pH-values.
A plot of Cl vs. Br (Fig. 10) shows that the
solute chemistry of the Lake Tyrrell ground-
water is similar to evaporated seawater (Long
et al., 1992 in this issue). Therefore, Br can be
used as an indicator of the degree to which a
water sample has been concentrated from a
seawater-like source. Concentration factors
(CF) are calculated as (Br in a sample ) / (Br
in seawater). These concentration factors are
listed in Table 1 along with the Br data.
Plots of the major constituents vs. Br are
shown in Fig. 11. Similar to C1, SO4 shows a
trend typical for evaporated seawater (Fig.
11 a). Chloride and sulfate geochemisty in the
Lake Tyrrell system has been shown to be con-
trolled by halite and gypsum precipitation, re-
spectively (Long et al., 1992 in this issue).
Potassium (Fig. 1 lb) also covaries with
bromide, but plots below the evaporative trend
e yap-cur ve
,,,,i,,o ...... , .... ,
S@O WO t Qr"
Q v£Tp -cur" ve
~,* * wells
• ~ s!ppers
i I I
2 2.5 3
Log Br (mg/1)
1.6 .. .
• • • Q-oo_
0.6 • ° • ol o•
-0.4 qbo •
L .... , ..... • • . I .... t
i .5 2 2.5 3 3.5
Log Br (rag/l)
Fig. 11. logno constituent vs. iog,o Br: (a) SO4; (b) K; (c) FeS+; and (d) AI. Seawater evaporation curves from Mac-
Caffrey et al. (1988).
ALUNITE, JAROSITE AND HYDROUS IRON OXIDES IN A HYPERSALINE SYSTEM 195
• dDqb •
• Qo •
• 1110 00
• • •
I .... t .... ] ......... I ..........
i . . . . i .... i .... 1,1..1.¸ . w .... w
• 0 ..'..
2 3 4 5
Fig. 12. Iogzo constituent vs. pH: (a)AI; and (b) Fe 3÷.
line for seawater with a slope of <1 (0.85).
This indicates that if seawater salts are the ma-
jor source for solutes in the system, then K is
removed early in the geochemical evolution of
the water (Long et al., 1992 in this issue). This
appears to be a typical behavior for K in hy-
persaline environments including deep basin
brines (Gudramovics, 1981; Long and Gud-
ramovics, 1982; Wilson and Long, 1986; Wil-
son, 1989). The removal of K is thought to be
via sorption by clays early in the evolution of
the water (Jones et al., 1977). If sorption of K
occurs only in the early stages of brine evolu-
tion, and saturation with K salts is not reached
during evaporation, then a plot of log~0 K vs.
IOgl0 Br will have a slope of one (Carpenter,
1978). Because the slope is < 1 in this case,
some process must continue to remove K from
solution during evaporative concentration.
The removal of K from the concentrated so-
lutions could be explained by continued sorp-
tion of K during evaporation. However, the ef-
ficiency of K sorption is decreased in
hypersaline systems, particularly when diva-
lent cations are present, so continued removal
of K by this means during evaporation would
be expected to be minor (Jones et al., 1977).
Mineral precipitation could also account for
the K removal. Graphical analysis and chemi-
cal modeling data indicate that these waters are
not saturated with salts of K typical of evapo-
rative environments (Long et al., 1992 in this
issue). Therefore, it is suggested that the pre-
cipitation of alunite and jarosite may in part
account for the removal of K from solution
during evaporation (Long et al., 1992 in this
The geochemistry of A1 and Fe is affected not
only by evaporation and mineral precipita-
tion, but also by pH and, in the case of Fe, Eh
controls. Consequently, a simple one-to-one
relationship between the logio concentration of
these elements and lOgl0 Br concentrations
would not be expected. This is the case for Fe 3+
(Fig. 1 lc) which shows no obvioius relation-
ship with Br. Aluminum concentration does
increase with increasing degree of evaporation
(Fig. 1 ld), but the scatter in the data indicates
that pH may also be a controlling factor. As
seen in Fig. 12a, AI concentration increases as
pH decreases. Fe (III) shows no clear relation-
ship with pH (Fig. 12b), suggesting that Fe
concentrations are controlled by a complex re-
lationship among pH, salinity and oxidation-
4.4. Chemical modeling
Because of the limitations of the computer
code and data set described earlier, with re-
gard to aq)aeous activity coefficeints in con-
centrted brines, disequilibrium indices of alu-
nite, jarosite and hydrous Fe-oxides could not
be directly calculated. Another approach to the
study of mineral-water equilibrium is to con-
struct activity-activity diagrams. For exam-
ple, Bladh (1982) studied theoretical solu-
tion-mineral equilibrium during the
weathering of sulfide minerals to produce ja-
rosite, alunite and clays. Using the computer
(Helgeson et al., 1970 ) for the cal-
culations, he plotted the changes in the loglo of
196 D.T. LONG ET AL.
the activities of Ala+/(H + )3 vs. Fe3+/(H + )3
during weathering. These changes could then
be compared to the alunite-jarosite phase
boundary which he defined with the reaction:
KAI3 (SO4)2 (OH)6 + 3Fe 3+ +
KFe3 (SO4)2(OH)6 + 3A13+ + 6H + (1)
This phase boundary has a slope of one and
water in equilibrium with both alunite and ja-
rosite would plot along it.
Because of the limitations discussed, activ-
ity-activity diagrams could not be constructed
for the study groundwater. Instead, a molal-
ity-molality plot of A13+/(H+) 3 vs. Fe3+/
(H ÷ )3 was constructed to detect a trend that
might indicate solution-mineral equilibrium.
This diagram is shown in Fig. 13 and includes
all samples for which both A1 and Fe 3÷ were
measured. The data appear to form a linear ar-
ray with a positive slope, but the slope is < 1.
This may be due to the fact that the data are
plotted as molalities rather than as activities or
that there is more than one data population in
Closer inspection of the data shows that in-
deed there are two trends on the diagram. The
change in slope occurs at a log l o [ Fe 3 + / (H + ) 3 ]
ratio of 7, which corresponds to a pH of ~ 4.0
(pH increases to the upper right on the dia-
gram). Samples which plot below this ratio
cluster along a line with a slope that is nearly
15 ........ N.~ ' " o'
12 ~X x
t • Regional
6 • • • /\Groundwater
O I . . , . . , . , , . . , . .
0 3 6 9 12 15
Log Fe+3/(H +)3
Fig. 13. logjo[Al3+/(H+) 3 ] vs. lOglo[Fe3+/(H+) 3 ] with
concentrations in molality.
one and have pH's < 4 (2.6-3.8). These data
are mainly from sipper and piezometer sam-
ples taken along the TP line (Fig. 2). Samples
which plot above this ratio are from the re-
gional groundwater with pH-value between 4.2
and 6.2 and plots below the trend line of the
first cluster. The trend of the data from the TP
line suggests possible solution equilibrium with
alunite and jarosite.
In order to test this hypothesis, PHREEQE
(Parkhurst et al., 1980) was used to predict to-
tal A13÷ and Fe 3÷ molalities and pH-values for
solutions in equilibrium with alunite and ja-
rosite. The range of molalities necessary to de-
fine a trend was calculated by varying the start-
ing pH of the modeled systems between 0.5 and
4.0. The resulting model pH-values were al-
ways higher than the starting pH's. The results
of the modeling are constrained by the choice
of equilibrium constants (Keq) for the reac-
tions and by the chemistry of the solution being
modeled. Table 2 is a summary of selected
constants and appropriate reactions. The val-
ues for amorphous Fe(OH)3 and goethite are
from the WATEQ-PHREEQE data set (Ball et al.,
1987; Parkhurst et al., 1987 ) and are generally
accepted as representative of these minerals
(D.K. Nordstrom, pers. commun., 1989).
However, for alunite and jarosite there are a
variety of equilibrium constants from which to
choose (Table 2).
The differences in equilibrium constants for
alunite and jarosite are due to: (a) differences
in the free energies assigned to these minerals
and in the thermodynamic data bases for
aqueous species used to calculate the equilib-
rium constants, or (b) differences in solubility
determinations or in their interpretation. For
example, using the free energies of alunite and
jarosite from Kashkay et al. (1975) and ther-
modynamic data for aqueous species from
either Naumov et al. (1971 ) or Robie et al.
( 1978 ), log Keq ranges from - 12.69 to - 9.23
for jarosite and from - 3.62 to - 1.22 for alu-
nit° (Table 2).
The equilibrium constants from Chapman et
ALUNITE, JAROSITE AND HYDROUS IRON OXIDES IN A HYPERSALINE SYSTEM
Equilibrium constants (Iogto) considered in this study
Reference Jarosite* J Alunite .2 Amorphous Goethite *a
Hladky and Slanksy ( 1981 ) -9.08 - 1.66
Bladh ( 1982 ) - 7.12 - 1.54
Ball et al. (1979) - 14.80
Truesdell and Jones (1974) - 1_346
Nordstrom et al. (1989)
Kashkay et al. (1975) *~ -9.23 - 1.87
Kashkay et al_ ( 1975 )'6 - 12.692 - 3.616
Chapman et al. (1983) - 9.21 - 1.22
*'KFej(SO4)2(OH)z+6H+~K+ + 3Fe3+ + 2SO24- +6H20.
*:AI3 (SO4)2 (OH)2 +6H+ ~K+ + 3A13+ +2SO4:- +6H20.
*JFe (OH)3 + 3H + ~Fe 3+ + 3H20.
*4FeOOH +3H+~Fe 3+ +2H20.
*SG°'s of alunite andjarosite from Kashkay et al. (1975), others from Naumov et al. ( 1971 ).
*6G°'s ofalunite and jarosite from Kashkay et al. (1975), others from Robie et al_ (1978).
al. ( 1983 ) were found to provide the best fit to
the data. These constants are based on data
from Naumov et al. ( 1971 ) and Kashkay et al.
( 1975 ) with modifications to the value for al-
unite from the work of Zotov (1971 ). These
data represent an internally consistent ther-
modynamic data base. D.K. Nordstrom (pers.
commun., 1989), for example, has found this
data base and the constants for jarosite pre-
sented by Chapman et al. ( 1983 ) to be better
in describing solution-mineral equilibria than
the other values listed in Table 2.
The modeling was done assuming a solution
in equilibrium with gypsum, although the
presence or absence of gypsum made little dif-
ference in the output. Disequilibrium
[logm(IAP/Ksp)] with respect to halite was
varied in the models between -2.00 and 0.00,
which is the range calculated for the samples
in Fig. 11 (Long et al., 1992 in this issue). Be-
cause F is present in the groundwater (Dick-
son et al., 1988) and F is known to form rela-
tively stable aqueous complexes with both Fe
and A1 (Hem, 1968, 1985), solution-mineral
equilibrium was also studied as a function ofF
concentration. Fluoride was added to the
model solutions as NaF ranging in concentra-
tion from 1 to 20 ppm, the expected range for
these samples based on their similarity to sea-
water, degree of concentration and available
data (Dickson et al., 1988; Long et al., 1992 in
The result of the modeling is shown in Fig.
14 as a solid line for halite disequilibrium of
15 1 ................
12 /JAA . %[
• r :/ •
0 3 6 9 12 15
Log Fe +3/(H~ 3
Fig. 14. Phase relationship of alunite and jarosite pre-
dicted from PHREEQE. Concentrations are in molality,
Solid line is theorectical molal phase boundary between
alunite and jarosite in the presence of 6 ppm F- and a
disequilibrium index for halite of - 1. Dashed lines show
expected range in position of phase boundary in spring
zone because of variable F- concentrations and halite
saturation states. JAG and JAA points are the theoretical
equilibrium molal concentrations for jarosite-alunite-
goethite and jarosite-alunite-amorphous Fe(OH)~,
198 D.T. LONG ET AL.
1.00 and F concentrations of 6 ppm. This line
represents the molal phase boundary between
alunite and jarosite for an average sample from
the TP line. The dotted lines on either side of
this line represent the range in the position of
the boundary depending on halite disequili-
brium and F concentration. The samples from
the TP line cluster along this boundary, sup-
porting the hypothesis that the groundwater in
the spring zone is in equilibrium with alunite
and jarosite. The alunite-jarosite phase
boundary is plotted as a dotted line at
loglo[Fe3+/(H+) 3] ratios of >7, because
these minerals theoretically are not mutually
stable at pH's > 4 (Hladky and Slansky, 1981 ).
The data which plot higher than the ratio of 7
clearly do not cluster along the phase bound-
ary, which indicates that they are not con-
trolled by alunite-jarosite equilibrium. The re-
suits suggest that alunite-jarosite equilibrium
controls the geochemistry of AI and Fe in the
spring zone and that the spread of the data
along the phase boundary is due to differences
in pH among the samples.
However, the above models did not consider
the presence of hydrous Fe-oxides in the sedi-
ments and the potential for the solution to be
in equilibrium with alunite, jarosite and Fe-
oxide (Brown, 1971 ). Goethite and an amor-
phous Fe-oxide were identified in the sedi-
ments. Expected Fe3+/(H+) 3 and A13+/
(H ÷ )3 ratios in solutions in equilibrium with
alunite-jarosite-goethite (JAG) and alunite-
jarosite-amorphous Fe(OH)3 (JAA) equilib-
ria were also calculated. The results of these
calculations are shown in Fig. 14.
The equilibrium pH-values for the JAA so-
lutions were higher ( ~ 3.80) than those for the
JAG solutions (~ 1.30). Both showed little
change when the starting pH was varied so that
the JAG and JAA models plot as tight clusters
in Fig. 14 rather than as lines. The JAA cluster
plots at the high end the data on the phase
boundary. The JAG cluster plots on the phase
boundary, but below the data.
The failure of the field data to cluster at
either the JAG or the JAA point on the dia-
gram suggests that neither hydrous Fe-oxide is
dominant in controlling the solution chemis-
try. The spread of the data along the boundary
could be explained by variable mixtures of JAG
and JAA solutions. It could also be explained
by solutions being in equilibrium with alunite,
jarosite and several hydrous Fe-oxides, each
oxide having a different free energy. This might
be more realistic than mechanical mixtures of
JAG and JAA solutions. The JAA and possibly
JAG would then represent end-member equi-
libria in the system.
There is scatter in the data about the phase
boundary beyond the range attributable to the
variable F concentrations and disequilibrium
states of halite considered in the modeling.
Possible sources of data scatter include: (1)
clean and analytical techniques; (2) use of 0.4-
/~m filters rather than 0.1-/zm filters, which
might affect the Fe and A1 data; ( 3 ) model so-
lution chemistries may not reflect natural con-
ditions; and (4) incomplete characterization
of the alunite and jarosite stoichiometries. We
feel that our clean and analytical techniques
were the best possible for this study (cf. Lyons
et al., 1992 in this issue). Filtration through
0.1-#m filters was impossible under the cir-
cumstances of this study, so this factor will
have to remain as a possible cause of the scat-
ter. The calculated range in the phase bound-
ary in Fig. 14 does reflect the actual solution
conditions along the TP line. This is because
the calculated phase boundaries considered the
maximum expected variation in the concen-
trations of C1 and F, the two most important
ligands controlling Fe and AI concentrations,
respectively, at low pH.
Incomplete characterization of the alunite
andjarosite could be a major cause of the scat-
ter in the data. Although the XRD and EDS
data indicate that these minerals are close to
the K end-member, it is known that low-tem-
perature alunites and jarosites can have vari-
able amounts of hydronium
tution for K, as dicussed by Alpers et al. ( 1992
ALUNITE, JAROSITE AND HYDROUS IRON OXIDES IN A HYPERSALINE SYSTEM
in this issue). This substitution, which is not
detected by EDS, would of course affect the free
energies of these minerals, as would any solid
solution of Fe and A1 between the minerals.
Thus, resolution of the cause in the scatter
in the data requires better information on dis-
solved AI and Fe speciation as well as a better
understanding of the stoichiometry of alunite
and jarosite. Also, Pitzer's equations must be
employed in the chemical models for a more
rigorous test of solution-mineral equilibrium
in the alunite-jarosite-Fe-oxide system. How-
ever, PHREEQE (Parkhurst et al., 1980) did a
remarkable job in defining a phase boundary
consistent with the field data.
Alunite and jarosite were investigated in the
hypersaline sediments of Lake Tyrrell, Aus-
tralia. These minerals occur as layers and ce-
ments within the top 20 cm of the sediments.
The pore water in these sediments is highly
concentrated by evaporation and are precipi-
tating gypsum and halite. Alunite and jarosite
appear in many cases to occur together as a
mechanical mixture within the sediments.
Their chemistry approximates the K end-
member of the alunite and jarosite super-
groups. Alunite occurs as equant grains of < 1
lzm in cross-section, while jarosite occurs as
hexagonal platelets of > 1/zm in cross-section
and ~ 2/zm thick.
We conclude that the occurrence of alunite
and j arosite in the sediments of the spring zone
of Lake Tyrrell is consistent with a model in
which these minerals precipitate as a result of
evaporitic processes. The constituent ions are
not necessarily locally derived. Groundwater
enters the playa as acid solutions with high
concentrations of K, SO4, A1 and Fe 3+. Within
the playa, these ions are concentrated by evap-
oration. The geochemistry of A1 and Fe 3 +is also
affected by pH and in the case of Fe 3 +, oxida-
The field data from the spring zone area of
Lake Tyrrell plot along a molal equilibrium
boundary between alunite and jarosite and be-
tween equilibrium mixtures of alunite-jaros-
ite-goethite and alunite-jarosite-amorphous
Fe-oxide. The spread of the data along the
boundary can be explained by solutions of
varying pH in equilibrium with alunite and ja-
rosite and/or solutions in equilibrium with al-
unite, jarosite and different hydrous Fe-ox-
ides. In any case, the trend in the data when
compared to the results of the chemical mod-
eling is consistent with the hypothesis that the
groundwater in the spring zone is in equilib-
rium with alunite, jarosite, and possibly Fe-ox-
ides. This result is tentative because of the po-
tential limitations in the chemical modeling.
However, considering these limitations, equi-
librium constants of 10 -~22 and
pear to best describe the hydrolysis reactions
of alunite and jarosite, respectively. This sup-
ports the results of Chapman et al, ( 1983 ) and
the work of Alpers et al. (1989) regarding the
formation ofjarosite in acid-mine drainages.
This work was supported by NSF Grant
EAR-12065 to D.T.L., W.Bi. and M.E.H. We
would like to thank Jean Long and Jane Matty
(Michigan State and Central Michigan Uni-
versities) for reviewing early versions of this
manuscript. Robert Lent (University of New
Hampshire) was instrumental in collecting and
the early processing of the field samples. Dis-
cussions with Charlie Alpers and Kirk Nords-
trom (U.S.G.S) were very important in our
understanding of the geochemical nature of al-
unite and jarosite. This work could not have
been done without the help of our Australian
colleagues. In particular we would like to thank
Andy Herczeg (CSIRO, Glen Osmond) and
Bruce Dickson (CSIRO, North Ryde).
Adams, F. and Hajek, B.F., 1978. Effects of solution sul-
fate, hydroxide, and potassium concentrations on the
200 D.T. LONG ET AL.
crystallization of alunite, basaluminite, and gibbsite
from dilute aluminum solutions. Soil Sci., 126: 169-
Adams, F. and Rawajfih, Z., 1977. Basaluminite and alu-
nite: a possible cause of sulfate retention by acid soils.
Soil Sci. Soc. Am. J., 41: 686-692.
Alpers, C.N., Nordstrom, D.K. and Ball, J.W., 1989. Sol-
ubility ofjarosite solid solutions precipitated from acid
mine waters, Iron Mountain, California, U.S.A. Sci.
Grol., 42: 281-298.
Alpers, C.N., Rye, R.O., Nordstrom, D.K., White, L.D.
and King, B.S., 1992. Chemical, crystallographic and
stable isotopic properties of alunite and jarosite from
acid-hypersaline Australian lakes. In: W.B. Lyons, D.T.
Long, A.L. Herczeg and M.E. Hines (Guest-Editors),
The Geochemistry of Acid Groundwater Systems.
Chem. Geol., 96:203-226 (this special issue).
Bail, J.W., Nordstrom, D.K. and Zachmann, D.W., 1987_
WATEQ2 -- A personal computer FORTRAN trans-
lation of the geochemical model WATEQ2 with re-
vised data base. U.S. Geol. Surv., Open-File Rep. 87-
50, 108 pp.
Barnes, R.B., 1975. The determination of specific forms
of aluminum in natural water. Chem. Geol., 15:177-
Bladh, K.W., 1982. The formation of goethite, jarosite,
and alunite during the weathering of sulfide-bearing
felsic rocks. Econ. Geol., 77: 176-184.
Bowler, J.M., 1986. Spatial variability and hydrologic
evolution of Australian lake basins: Analogue for
Pleistocene hydrologic change and evaporite forma-
tion. Palaeogeogr., Palaeoclimatol., Palaeoecol., 54:21-
Bowler, J.M. and Teller, J.T., 1986. Quaternary evaporite
and hydrological changes, Lake Tyrrell, north-west
Victoria. Aust. J. Earth Sci., 33: 43-63.
Brophy, G.P., Scott, E.S. and Snellgrove, R.A., 1962. Sul-
fate studies, II. Solid solution between alunite and ja-
rosite. Am. Mineral., 47:112-126.
Brown, J.B., 1971. Jarosite-goethite stabilities at 25 ° C, 1
atm. Miner. Deposita, 6: 245-252.
Carpenter, A.B., 1978. Origin and chemical evolution of
brines in sedimentary basins. Okla. Geol. Surv., Circ.,
Chapman, B.M., Jones, D.R. and Jung, R.F., 1983. Pro-
cesses controlling metal ion attenuation in acid mine
drainage streams. Geochim. Cosmochim. Acta, 47:
Dickson, B.L., Giblin, A.M. and Herczeg, A.L., 1988.
Geochemistry and radiochemistry of acid-saline waters
at Lake Tyrrell, Victoria. CSIRO (Commonw. Sci. Ind_
Res. Org. ), Inst. Miner. Energy Construct., Invest. Rep.
1771R, 47 pp_
Dutrizac, J.E. and Jambor, J.L., 1987. Behavior of cesium
and lithium during the precipitation of jarosite-type
compounds. Hydrometallurgy, 17:251-265.
Garrels, R.M. and Thompson, M.E., 1962. A chemical
model for seawater at 25 °C and one atmosphere total
pressure. Am. J. Sci., 260: 57-66.
Gibbs, C.R., 1976. Characterization and application of
Ferrozine iron reagent as a ferrous and total iron indi-
cator. Limnol. Oceanogr., 14: 357-367.
Goldbery, R., 1978. Early diagenetic, nonhydrothermal
Na-alunite in Jurassic flint clays, Makhtesh Ramon,
Israel. Geol. Soc. Am. Bull., 89: 687-698.
Goldbery, R., 1980. Early diagenetic, Na-alunites in Mio-
cene algal mat intertidal facies, Ras Sudar, Sinai. Se-
Gudramovics, R., 1981. A geochemical and hydrological
investigation of a modern coastal marine sabkha. Mas-
ter's Thesis, Michigan State University, East Lansing,
Hartig, C., Brand, P. and Bohmhammel, K., 1984. Fe-A1
lsomorphie und Strukturwasser in Kristallen von Ja-
rosit-Alunit-Typ. Z. Anorg. Allg. Chem., 508:159-164.
Harvie, C.E. and Weare, J.H., 1980. The prediction of
mineral solubilities in natural waters: the Na-K-Mg-
Ca-C1-SO4-H20 system from zero to high concentra-
tions at 25 degrees C. Geochim. Cosmochim. Acta, 44:
Heigeson, H.C., Brown, T.H., Nigrini, A. and Jones, T.A.,
1970. Calculation of mass transfer in geochemical pro-
cesses involving aqueous solutions. Geochim. Cos-
mochim. Acta, 34: 569-592.
Hem, J.D., 1968. Graphical methods of studies of aqueous
aluminum hydroxide, fluoride, and sulfate complexes.
U.S. Geol. Surv., Water-Supply Pap. 1827-B, 33 pp.
Hem, J.D., 1985. Study and interpretation of the chemi-
cal characteristics of natural water. U.S. Geol. Surv.,
Water-Supply Pap. 2254, 263 pp. (3rd ed. )
Hemley, J.J., Hostetler, P.B., Gude, A.J. and Mountjoy,
W.T., 1969. Some stability relations of alunite. Econ.
Geol., 64: 599-612.
Hines, M.E., Lyons, W.B., Lent, R.M. and Long, D.T.,
1992. Sedimentary biogeochemistry of an acidic, sa-
line groundwater discharge zone in Lake Tyrrell, Vic-
toria, Australia. In: W.B. Lyons, D.T. Long, A.L. Her-
czeg and M.E. Hines (Guest-Editors), The
Geochemistry of Acid Groundwater Systems. Chem.
Geol., 96:53-65 (this special issue).
Hladky, G. and Siansky, E., 1981. Stability of alunite
minerals in aqueous solutions at normal temperature
and pressure. Bull. Minrral., 104: 468-477.
Ivarson, K.C., Ross, G.J. and Mills, N.M., 1979. The mi-
crobiological formation of basic ferric sulfates, II.
Crystallization in presence of potassium, ammonium,
and sodium salts. J. Soil Sci. Soc. Am., 43: 908-912.
Jones, B.F., Eugster, H.P. and Rettig, S.L., 1977. Hydro-
geochemistry of the Lake Magadi Basin, Kenya. Geo-
chim. Cosmochim. Acta, 41: 53-72.
Kashkay, C.M_, Borovskaya, Yu.B. and Babazade, M.A.,
1975. Determination of Gf°29a of synthetic jarosite and
ALUNITE, JAROSITE AND HYDROUS IRON OXIDES IN A HYPERSALINE SYSTEM 201
its sulphate analogues, Geochem. Int., 12- 115-121.
Knight, J.E., 1977. A thermochemical study of alunite,
enargite, luzonite, and tennantite deposits. Econ. Geol.,
Lock, D.E., 1988, Alunite andjarosite formation in evap-
orative lakes of South Australia. SLEADS (Salt Lakes,
Evaporites, and Aeolian Deposits) 1988 Annu. Meet.,
Abstr., pp. 48-51.
Long, D.T. and Gudramovics, R_, 1982. Major element
geochemistry of brines from the wind-tidal fiat area,
Laguna Madre, Texas, J. Sediment. Petrol., 53: 797-
Long, D.T., Fegan, N,E., Lyons, W.B., Hines, M.E., Mac-
umber, P.G. and Giblin, A.M., 1992. Geochemistry of
acid brines: Lake Tyrrell, Victoria, Australia. In: W.B.
Lyons, D.T. Long, A.L. Herczeg and M.E. Hines
(Guest-Editors), The Geochemistry of Acid Ground-
water Systems. Chem. Geol., 96:33-52 (this special
Lyons, W.B., Chivas, A.R., Lent, R.M., Welsh, S., Kiss,
E, Mayewski, P.A., Long, D.T. and Carey, A.E., 1989.
Metal concentrations in surficial sediments from hy-
persaline lakes, Australia. Hydrobiologia, 197:13-22.
Lyons, W.B., Welch, S., Long, D.T., Hines, M.E., Giblin,
A.M., Carey, A.E, Macumber, P.G., Lent, R.M. and
Herczeg, A.L., 1992. The trace-metal geochemistry of
the Lake Tyrrell system brines (Victoria, Australia)_
In: W.B. Lyons, D_T. Long, A.L. Herczeg and M.E.
Hines (Guest-Editors), The Geochemistry of Acid
Groundwater Systems. Chem. Geol., 96:115-132 (this
MacCaffrey, M.A., Lazar, B. and Holland, H.D., 1988. The
evaporation path of seawater and the coprecipitation
of Br and K with halite, (Unpublished.)
Mackin, J.E. and Aller, R.C_, 1984. Dissolved AI in sedi-
ments and waters of the East China Sea: Implications
for authigenic mineral formation. Geochlm. Cosmo-
chim. Acta, 40: 218-297.
Macumber, P.B., 1983. Interactions between ground water
and surface systems in northern Victoria. Ph.D. Dis-
sertation, University of Melbourne, Melbourne, Vic.,
Macumber, P.G., 1992. Hydrological processes in the
Tyrrell Basin, southeastern Australia. In: W.B. Lyons,
D.T. Long, A.L Herczeg and M.E. Hines (Guest-Edi-
tors), The Geochemistry of Acid Groundwater Sys-
tems. Chem. Geol., 96:1-18 (this special issue).
Mann, A.W., 1983. Hydrogeochemistry and weathering
on the Yilgarn Block, W.A. -- Ferrolysis and heavy
metals in continental brines. Geochim. Cosmochim.
Acta, 47: 181-190.
Michel, F.A_ and van Everdingen, R.O., 1987. Formation
of a jarosite deposit on Cretaceous shales in the Fort
Norman area, Northwest Territories. Can. Mineral., 25:
Naumov, G.B., Ryzhenko, B.N. and Khodakovsky, I.L.,
1971. Handbook of Thermodynamic Data. Atomiz-
dat, Moscow, 373 pp. (translated by G.J. Soleimani,
U.S. Geol. Surv., NTIS/PB-226-722/AS, 1974; see es-
pecially pp. 18 I- 186).
Nordstrom, D.K., 1982. The effect of sulfate on alumi-
num concentrations in natural waters: some stability
relations in the system AI203-SO3-H20 at 298 K.
Geochim. Cosmochim. Acta, 46:681-692.
Parkhurst, D.L., Thorstenson, D.C. and Plummer, L.N.,
1980. PHREEQE -- A computer program for geo-
chemical calculations. U.S. Geol. Sum., Water-Re-
sour_ Invest. 80-96, 194 pp. (revised and reprinted Jan.
Pitzer, K.S., 1979. Theory: ion interaction approach. In:
R.M. Pytkowicz (Editor), Activity Coefficients in
Electrolyte Solutions, Vol. 1. CRC Press, Boca Raton,
Fla., pp. 157-208
Plummer, L.N., Parkhurst, D.L., Fleming, G.W. and
Dunkle, S.A., 1988. PHRQPITZ- A computer pro-
gram for geochemical calculations in brines. U.S. Geol.
Surv., Water-Resour. Invest. Rep. 88-4153, 310 pp.
Presley, F_B., 1971. Determination of selected minor and
major inorganic constituents_ In: E.L. Winterer, W.R.
Riedel, et al. (Editors), Initial Reports of the Deep Sea
Drilling Project, Vol. VII. U.S. Gov. Print_ Off-, Wash-
ington, D.C., pp. 1749-1755.
Reardon, E.J. and Beckie, R.D., 1987. Modeling chemical
equilibria of acid mine-drainage: the
H20 system. Geochim. Cosmochim. Acta, 51: 2355-
Robie, R.A., Hemingway, B.S. and Fisher, J.R., 1978.
Thermodynamic properties of minerals and related
substances at 298_ 15 K and 1 bar (105) pascals) pres-
sure and at higher temperatures. U.S. Geol. Surv., Bull.
1452, 456 pp.
Rye, R.O., Bethke, P.M. and Wasserman, M.D., 1989.
Diverse origins of alunite and acid-sulfate alteration:
stable isotope systematics. U.S. Geol. Surv., Open-File
Rep. 89-5, 18 pp.
Sass, E., Nathan, Y. and Nissenbaum, A., 1965. Mineral-
ogy of certain pyrite concretions from Israel and their
alteration products. Mineral. Mag., 35: 84-87.
Scott, K.M., 1987. Solid solution in, and classification of,
gossan-derived members of the alunite-jarosite fam-
ily, northwest Queensland, Australia. Am. Mineral., 72:
Stoffregen, R_E. and Alpers, C.N., 1987. Woodhouseite
and svanbergite in hydrothermai ore deposits: prod-
ucts of apatite destruction during advanced arglllic al-
teration. Can. Mineral., 25:201-211,
Stookey, L.L., 1970. Ferrozine -- a new spectrophoto-
metric reagent for iron. Anal. Chem., 42:779-781.
Sullivan, P.J., Mattigod, S.V. and Sobek, A.A., 1986. Dis-
solution of iron sulfates from pyritic coal waste. Envi-
ron. Sci. Technol., 20: 1013-1016.
Teller, J.T., Bowler, J_M. and Macumber, P_G., 1982.
202 D.T. LONG ET AL.
Modem sedimentation and hydrology in Lake Tyrrell,
Victoria. J. Geol. Soc. Aust., 29: 159-175.
Truesdell, A.H. and Jones, B.F., 1974. WATEQ- A com-
puter program for calculating chemical equilibria of
natural waters. J. Res. U.S. Geol. Surv., 2: 233-248.
van Breemen, N., 1973. Dissolved aluminum in acid sul-
fate soils and in acid mine waters_ Soil Sci. Soc. Am.
Proc., 37: 694-697.
Wilson, T.P., 1989_ Origin and geochemical evolution of
the Michigan Basin brine. Ph.D. Dissertation, Michi-
gan State University, East Lansing, Mich.
Wilson, T.P. and Long, D_T., 1986. Constraints on the
evolution of the Michigan Basin brines. Geol. Soc. Am.,
Abstr. Prog., 18: 791.
Zotov, A.V., 1971. Dependence of the composition of al-
unite on the temperature of its formation. Geochem.
Int., 8: 71-75_