Chemical and isotopic (87Sr/86Sr, 18O, D) constraints to the

Geochimica et Cosmochimica Acta, Vol. 65, No. 8, pp. 1259 –1275, 2001
Copyright © 2001 Elsevier Science Ltd
Printed in the USA. All rights reserved
0016-7037/01 $20.00 ⫹ .00
Pergamon
PII S0016-7037(00)00618-9
Chemical and isotopic (87Sr/86Sr, ␦18O, ␦D) constraints to the formation processes
of Red-Sea brines
M. C. PIERRET,1,*,† N. CLAUER,1 D. BOSCH,2 G. BLANC,3 and C. FRANCE-LANORD4
1
Centre de Géochimie de la Surface/Ecole et Observatoire des Sciences de la Terre, 1 rue Blessig, 67084 Strasbourg Cedex, France
2
Laboratoire de Tectonophysique, Place E. Bataillon, 34095 Montpellier Cedex, France
3
Département de Géologie et d’Océanographie, Avenue des facultés, 33405 Talence Cedex, France
4
Centre de Recherches Pétrographiques et Géochimiques, BP 20, 54501 Vandoeuvre-les-Nancy Cedex, France
(Received March 28, 2000; accepted in revised form October 13, 2000)
Abstract—About twenty deeps filled with hot brines and/or metalliferous sediments, are located along the
Red-Sea axis. These brines present a well-suited framework to study the hydrothermal activity in such a young
ocean. The present study outlines the results of a geochemical approach combining major-, trace-element and
isotopic (oxygen, hydrogen, strontium) analyses of brines in six of the deeps, to evaluate different processes
of brine formation and to compare the evolution of each deep. Important heterogeneities in temperature,
salinity, hydrographic structure and chemistry are recorded, each brine having its own characteristics. The
intensity of hydrothermal circulation varies among the deeps and ranges from being strong (Atlantis II and
Nereus) to weak (Port-Soudan) and even to negligible (Valdivia and Suakin) and it varies along the entire
Red-Sea axis. These observations do not favour a unique formational model for all of the brines. For example,
the brines of the Suakin deep appear to have been formed by an old sea water which dissolved evaporite beds,
without significant fluid circulation and hydrothermal input, while others such as Atlantis II or Nereus Deeps
appear to be dominated by hydrothermal influences. A striking feature is the absence of a relationship between
the position of the deeps along the axis and their evolutionary maturity. Copyright © 2001 Elsevier Science
Ltd
Zierenberg and Shanks, 1986). The available models propose
that sea water circulated and interacted with hot oceanic basalts, dissolved evaporitic salts during hydrothermal circulation, and due to their high density, the hydrothermal solutions
were trapped in the Atlantis-II depression.
The sediments and brines of the other deeps, such as those of
Suakin, Port-Soudan, Valdivia, Thetis and Nereus, were less
intensively sampled and studied and only a few comparisons
have been made among these different sites. Hydrothermalfluid circulations through the evaporite deposits seem to be the
main reason for the salty inputs into the deeps, inducing formation of brine layers at their bottoms. Alternatively, other
basins such as those of Bannock and Tyro in the eastern
Mediterranean Sea and that of Orca in the Gulf of Mexico,
contain brines which do not appear to result from hydrothermal
activities. The origin of these basins relates to subsurface salt
dissolution triggered by tectonic deformation (Camerlenghi,
1990). Our study addresses such questions. Are the brines from
different Red-Sea deeps formed the same way? Is the hydrothermal activity always necessary, and does it have the same
influence in each deep?
In mature oceanic ridges, such as the East-Pacific Rise and
the Mid-Atlantic Ridge, the hydrothermal fluids have different
salinities but never higher than 8% of NaCl (Edmond et al.,
1982; Von Damm et al., 1985; Von Damm, 1988; 1995), and
mineralizations occur as sulphide chimneys and mounds. The
difference between these ridge mineralizations and the hydrothermal mineralizations of Red-Sea deeps is largely due to the
occurrence of hot brine pools in the latter. The salty hydrothermal fluids of the Red Sea are trapped in the deeps and their
metals are precipitated as layers within the sediments. Moreover, no hydrothermal fluids have been directly sampled in the
1. INTRODUCTION
The Red Sea is a young ocean which evolved from a continental to an oceanic rift. The axial valley is characterised by
heterogeneous spreading rates along the axis. It differs from
mature oceanic ridges by having continuous oceanic crust in its
meridional part (between 16° and 19°N), while the axial trough
is discontinuous towards the north. Only a few isolated deeps
have a basaltic floor north of 19°N (Bonatti, 1985; Le Quentrec
and Sichler, 1991). These bathymetric depressions are characterised by the occurrence of hot brines and/or metalliferous
sediments and represent unique sites in the study of ocean-ridge
formation. In this study, we examine the formation processes of
the brines from six Red-Sea deeps.
Atlantis II Deep represents the largest deep of the Red Sea.
It contains the warmest and the saltiest brines, as well as the
most metal-enriched sediments. It is the most thoroughly studied deep, partly because of the economic interest in its metal
deposits. Characterisation and formation of the brine layers
have been studied since the 1960’s (Miller et al., 1966; Schoell
and Hartmann, 1973; 1978; Hartmann, 1980; Blanc and Anschutz, 1995; Anschutz and Blanc, 1996; Hartmann et al.,
1998a;b; Anschutz et al., 1999). Temperature and salinity have
been recorded since 1965. The occurrence of mineralised brines
and sediments in this deep has generally been related to hydrothermal activities and to leaching of the thick Miocene
evaporitic beds (Anschutz and Blanc, 1993a;b; 1995a;b; Blanc
et al., 1998; Degens and Ross, 1969; Pottorf and Barnes, 1983;
*Author to whom correspondence should be addressed.
†
Present address: Institut für Nukleare Entsorgung, Postfach 3540,
D-76021 Karlsruhe, Germany ([email protected]).
1259
1260
M. C. Pierret et al.
Fig. 1. Schematic map of the Red Sea with locations of the seven
studied deeps and corresponding CTD-rosettes numbers.
Red Sea, which implies that their chemical compositions and
temperatures could only be deduced. Ramboz et al. (1988)
measured temperature and salinity within fluid inclusions of
barite and anhydrite in the metaliferous sediments of the Atlantis II Deep and obtained a 400°C maximal temperature and
a maximum salinity of 330. These values should represent good
approximations for the hydrothermal fluids.
During the 1992 REDSED cruise aboard the R.V. Marion
Dufresne, sediments, waters and brines were for the first time
collected together in seven Red-Sea deeps (Suakin, PortSoudan, Valdivia, Chain, Atlantis II, Thetis and Nereus; Fig.
1). Analysis of these samples in the present study allowed
comparisons of the evolution of contemporaneous deeps. Physical, chemical and isotopic determinations were made to characterise the brines and to deduce their origins and formation
processes. The results address the following aims: (1) to evaluate the different formation processes of the brines occurring in
each deep; (2) to highlight the differences among the deeps; (3)
to quantify the impact of hydrothermal activity on each deep;
and (4) to deduce the general evolution of these deeps relatively
to their location along the spreading axis.
2. GEOLOGIC SETTING AND PREVIOUS STUDIES
The Red Sea is a narrow oceanic depression, about 2000 km
long, located between the Arabic and African plates. Its evo-
lution began with a rifting phase during Oligocene time (at
about 28 –32 Ma), followed by a magmatic and spreading
phase. Sea-floor spreading with formation of oceanic crust
lasted at least 7 Ma (Bohannon and Eittreim, 1991) with an
expansion rate of about 2 cm/yr. During the Miocene, high
evaporation rates induced formation of thick evaporite deposits.
These evaporites contributed to the formation of continental
brines in the Gulf of Suez (Issar et al., 1971; Rosenthal et al.,
1998) and to the brines of several Red-Sea depressions.
Discovery of the topographic depressions began in 1948
when scientists aboard the Swedish research vessel O.V. Albatross reported salinity and temperature anomalies relative to
normal sea water at 21°10⬘N and 38°09⬘E (corresponding to
Discovery Deep). These anomalies were the first evidence for
the occurrence of differentiated layers in the brines of the
depressions. Seismic transects correlated the salinity and temperature anomalies with topographic deeps located along the
axis (Charnock, 1964; Miller, 1964). In 1965, several studies
recorded temperatures of 56°C and salinities of 261 in Atlantis-II Deep. This deep was explored in detail during an international expedition in 1966 in which a stratified brine body and
the first oceanic metalliferous deposits of economic interest
were described (Miller et al., 1966; Degens and Ross, 1969).
Further investigations of the Red Sea resulted in the discovery
of about twenty deeps containing brine layers and/or metalliferous sediments between 18° and 26°N, along the active rift
zone (Bäcker and Schoell, 1972; Schoell, 1976).
Sea-floor spreading is often accompanied by hydrothermal
activity. In the Red Sea, sea water can interact with the oceanic
basalts and with the nearby sedimentary evaporites, black
shales enriched in metals, and biodetrital sediments. The dense
and saline hydrothermal fluids also fill the bathymetric depressions. Thus, trapping combined with double diffusive convection induces brine layering (Huppert and Turner, 1981; McDougall 1984a;b; Anschutz et al., 1998). Studies of fluid
inclusions in anhydrite of Atlantis II sediments suggest a supply of hot and salty fluids (Zierenberg and Shanks, 1983;
Ramboz et al., 1988). After trapping, the brines evolved via
different processes. Brines may be enriched by chemical dissolution, such as calcite or aragonite dissolution during settling
of biogenic particles in the water column. Moreover, different
chemical precipitations may take place into the brines (secondary carbonates, silicates or Fe-oxides; Pottorf and Barnes,
1983; Hartmann, 1985; Zierenberg and Shanks, 1986;
Butuzova et al., 1990; Anschutz and Blanc, 1993a;b; Anschutz
and Blanc, 1995a;b; Blanc et al., 1998).
Several studies have concluded, however, that not all RedSea brines were formed according to this model. For instance,
Schoell and Faber (1978) on the basis of oxygen and hydrogen
isotope ratios, concluded that the brines that were derived from
deep waters at various times are of three different types: (1)
isotopically unaltered present-day waters; (2) heavy palaeowaters from the last glacial maximum (18000 yr BP); and (3)
lighter palaeowaters from the last climatic optimum (8000 –
5000 yr BP). Zierenberg and Shanks (1986) suggested that the
brine of Valdivia Deep formed by karstic dissolution of Miocene evaporites without hydrothermal input, that the brine of
Suakin deep did not record high-temperature water/rock interactions, and that the lower brine of Atlantis II Deep reflects
exchanges at approximately 255°C. Monnin and Ramboz
Geochemical study of Red Sea brines
1261
Table 1. Hydrographic and CTD-rosette characteristics of the brines from the seven studied deeps.
Suakin
Port-Soudan
Chain
Atlantis II
Valdivia
Thetis
Nereus
Sea water
CTD-rosette
Depth (m)
Type of brine
B1
2800
B2
B5
B3, B4, B6
2600
2050
2120
B7
B8
B9
1673
1820
2460
Lower brine
Upper brine
1 brine
1 brine
LCL
UCL1
UCL2
UCL3
1 brine
1 brine
(1996) reported that anhydrite undersaturation seems to characterise brines with a hydrothermal signature. They suggested
that those of the Atlantis II, Discovery and Suakin deeps
contain a hydrothermal component, whereas the brine of
Valdivia Deep in equilibrium with anhydrite results only from
dissolution of Miocene evaporites. As has been discussed by
Monnin and Ramboz (1996) and Pierret et al. (2000) the origin
of the Suakin brine still remains a matter of debate. In summary, the chemical and isotopic characteristics of the Red-Sea
brines depend on parameters such as: (1) the supply of hydrothermal fluids; (2) the nature of the interactions between sea
water and either sedimentary deposits (evaporites, old sediments) or basalts; (3) the contribution of dissolved biodetrital
particles; and (4) chemical precipitation from brines.
We analysed the waters from seven different deeps comprising from South to North, Suakin, Port-Soudan, Atlantis II,
Chain, Valdivia, Thetis and Nereus (Fig. 1). The Suakin Deep
is about 30 km long and 9 km wide with an average depth of
2600 m; it consists of two subbasins (East and West) with the
bottom formed, at least locally, by basalt. Two brine layers
were described in this southernmost deep of the Red Sea
(Baumann et al., 1973). The Port-Soudan Deep is an elongated
graben structure (2836 m maximum depth; Bäcker et al., 1975)
located between Suakin and Atlantis II deeps (Fig. 1). It is a
single basin aligned in a NNW-SSE direction with a flat bottom
(Guennoc and Thisse, 1982), and is filled with the thickest
brine layer known in the Red Sea. The Atlantis II Deep is the
largest brine-filled basin of the Red Sea with a surface area of
about 52 km2 and a volume of about 4 km3 for the Lower
Convective Layer (LCL) brine (Hartmann, 1998a;b; Anschutz
and Blanc, 1996). The deep is a complicated graben structure
limited by antithetic tensional faults along its NE and SW
borders (Bäcker et al., 1975). It consists of four subbasins
separated by bathymetric highs that do not extend above the top
of the brine pool. The Chain Deep located to the South of
Atlantis II, consists of three basins (Blanc and Anschutz, 1995),
and is connected to Atlantis II at a depth of 1980 m (Anschutz
et al., 1999). The Valdivia Deep is located about 20 km to the
W-SW of Atlantis II on the western flank of the axial trough. It
consists of several small and shallow bathymetric depressions.
Its main basin was sampled. The Thetis Deep discovered in
1972, 160 km northwest to Atlantis II Deep, is divided into
several subbasins. The main basin to the northwest is approximately 10 km long and 3 km wide with a depth of 1780 m. The
sediments of the Thetis Deep represent the second richest
Thickness (m)
22
29
123
42
2050 3 bottom
29
16
7
75
lack of brine
15
Temperature (°C)
Salinity
23.5°
23.3°
35.9°
45.3°
66.1°
55°
50°
46.5°
33.7°
147.5
145.5
214.5
270
270
155
118
90
242
29.9°
21.6°
223.5
40
metalliferous site in the Red Sea. The Nereus Deep situated
within the median valley, has an overall width of 12 km and a
length of 40 km. It is limited in the SW and NE by steep faults
(Fig. 1); Bäcker and Schoell, 1972; Bäcker et al., 1975; Bäcker,
1976; Bignell and Ali, 1976; Scholten et al., 1991).
3. SAMPLING AND METHODS
Water and brine samples were collected in the seven deeps described
above (Fig. 1). The results of different CTD-rosette characterizations
are summarised in Table 1.
The details of the sampling method have been described previously
(Blanc and Anschutz, 1995; Anschutz and Blanc, 1996). A BissettBerman modified Olliver CTD-rosette continuously recorded temperature, conductivity and hydrostatic pressure. Water and brines were
collected in ascending Niskin bottles (12 L capacity). The salinities
were obtained using a Goldberg optical refractometer. The pinger
position and the pressure recorded by the CTD-rosette allowed the
evaluation of the sampling depth based on sound-velocity determinations for sea water and different brines. The values of sound velocity
were extrapolated from Matthew’s tables (Matthew, 1939). The conductivity data together with temperature and salinity allowed reconstruction of a continuous salinity record with linear functions specific to
each environment. The water samples were filtered (0.45␮m pore size)
after collection, and subsequently acidified with HNO3.
The elemental concentrations were determined by atomic absorption
(Ca and Mg), emission spectrometry (Na and K), colorimetry (Si), ionic
chromatography (Cl and SO4), ICP-AES (Mn, Sr, Fe, Rb), and ICP-MS
(Li, Mo, Ba, U, Cr, Co, Cu, Zn and Pb). Reproducibility of the IAPSO
sea-water standard varied between 5 and 10% for these elemental
chemical determinations. The oxygen isotope compositions were analysed by CO2 micro-equilibration after 24 h. Hydrogen was extracted
by water reduction on uranium at 800°C (Friedman and Woodcock,
1957). The D/H ratios were measured on a VG 602D-type mass
spectrometer and 18O/16O ratios were determined using a triple collector VG Optima mass spectrometer. The ␦D and ␦18O relative to SMOW
(Standard Mean Ocean Water) are given in per mil (‰) values. The
analytical errors are ⫾2‰ and ⫾0.2‰ for the ␦D and ␦18O measurements, respectively. Corrections due to different salt concentrations and
CO2 equilibration with H2O were applied according to Gat and Confiantini (1981). Strontium was extracted using a cation-exchange procedure (type AG 50 ⫻ 12) according to Birck (1981) and was loaded
with a Ta activator on a single tungsten filament. The Sr isotopic data
were obtained on a VG Sector mass spectrometer. The NBS 987
standard was routinely measured and for 10 determinations made
during the course of the study, yielded an average 87Sr/86Sr ratio of
0.710256 ⫾ 0.000014 (2␴). Three CTD-rosette characterizations were
obtained in the Atlantis II Deep: B3 was located in the northern
passage, B4 in the eastern basin, and B6 in the west-southwestern
passage. Since they reflect identical stratification and chemical compositions, in what follows, mean composition values are given as an
average of the three CTD-rosettes of this deep.
1262
M. C. Pierret et al.
Table 2. Average concentrations of the brines from the seven studied Red-Sea deeps.
Suakin
pH
Alk (meq)
Na (1)
Cl (1)
K (2)
Ca (2)
Rb (3)
Mg (2)
So4 (2)
Fe (2)
Mn (2)
Li (3)
Ba (3)
Cu (3)
Zn (3)
7.85
0.9
2.48
2.76
38
56.1
8.2
72.2
33.3
0.007
0.48
45.7
1.1
0.5
2.6
PS
6.43
1.96
3.79
4.04
49
33.52
6.7
64.6
44
0.107
0.12
106
1.6
4.1
4.7
Atlantis II
Chain
Valdivia
Nereus
Thetis
5.21
1.11
4.88
5.37
69.9
146.3
25.3
32.8
6.7
1.4
1.52
563.4
15.3
3.8
41.6
6.48
0.26
4.81
5.17
66.8
143.3
25.1
36.8
8.5
0.02
0.71
429
5.3
6.9
9.9
6.21
2.64
4.13
4.59
52
25
3.1
95.3
72
0.13
0.06
98.1
0.5
⬍
⬍
7.43
0.44
3.74
4.35
80.2
237.7
31.1
80.4
7
0.06
0.94
168.7
4.5
12
18.4
8.1
2.46
0.54
0.64
12
12
1.5
63
33
⬍
⬍
31
0.06
⬍
⬍
(1) in mol/1; (2) in mmol/1 and (3) in ␮mol/1. P.S - Port-Soudan Deep. ⬍ ⫽ below detection limit.
4. RESULTS
Temperature and salinity were determined in the seven
deeps. The variations during the last 20 yr were published by
Anschutz et al. (1999). Four different stratified brine layers
were identified in the Atlantis II Deep; a Lower Convective
Layer (LCL) and three Upper Convective Layers (UCL1,
UCL2, UCL3; Blanc and Anschutz, 1995). Suakin Deep contains two layers of brines, Thetis Deep is filled with Red-Sea
bottom water, and the four other deeps are characterised by one
layer of brine. All brines are warmer and saltier than the bottom
water (Table 1). The temperatures range from 23.5°C in Suakin
to 66°C in the lower brine of Atlantis II and the salinities range
from 147.5 in Suakin to 270 in Atlantis II and Chain. The
Suakin brines differ from those of the other deeps by having
significantly lower temperatures and salinities. Each elemental
profile correlates with the temperature and the hydrographic
stratification (Pierret, 1998).
Differences in pH, major (anions and cations) and trace
elemental concentrations were observed among the different
brines (Table 2). The most basic brine is that of Suakin Deep
(pH ⫽ 7.85) and the most acidic one corresponds to the LCL of
Atlantis II (pH ⫽ 5). Nereus contains the brine which is the
most concentrated in K, Ca, Sr, and Rb and least concentrated
in Si, whereas the Atlantis II lower brine is the most enriched
in Si, Mn, Fe, Li, Ba and Zn. The lower brine of Suakin is the
least concentrated in Na, Cl, K, Fe (below detection limit) and
Li. The brine which is the most depleted in Ca, Sr, Mn, Rb, Ba,
Cu, and Zn, and the richest in Mg and SO4 is that of Valdivia
Deep (Table 2). Each brine is enriched in Na, Cl, Ca, Sr, Si, and
Rb relative to normal Red-Sea water, but the brines of Atlantis
II, Chain and Nereus deeps have low concentrations of SO4 and
those of Atlantis II and Chain contain less Mg. These large
differences suggest that the brines formed by different processes and that each deep operates as an isolated system allowing preservation of its chemical and physical characteristics.
Several bottom waters and brines were analysed for their
stable isotope compositions (oxygen and hydrogen; Table 3).
The isotope signatures of the bottom water plot into two areas
(Fig. 2) suggesting the presence of different water masses. The
differences may be partly explained by seasonal variations in
the exchange flow through the strait of Bab al Mandab connecting the Red Sea to the Gulf of Aden (Thompson, 1939;
Smeed, 1997), and by seasonal variations in temperature, circulation and current directions of the Red-Sea waters (Phillips,
1966). The highest ␦D was found for the lower brine of Suakin
(⫹22‰) and the lowest for the lower brine of Atlantis II
Table 3. ␦18O and ␦D ratios (‰/SMOW) of various waters and
brines in the Suakin, Port-Soudan, Chain, Atlantis II, Valdivia, Thetis
and Nereus deeps.
Deep
Atlantis II
Chain
Suakin
Port-Soudan
Valdivia
Thetis
Nereus
Type of
water
Depth
(m)
␦D (‰)
⫾2‰
␦18O (‰)
⫾0.2‰
LCL
LCL
LCL
LCL
LCL
LCL
Sea water
Sea water
B
Sea water
Sea water
B1
B1
B2
B2
Sea water
B
B
Sea water
Sea water
B
B
Sea water
Sea water
Sea water
Sea water
B
B
Sea water
Sea water
2087
2056
2101
2176
2056
Average
1944
1912
2090
1905
1799
2829
2808
1793
1778
1730
2643
2540
2443
2366
1600
1540
1520
1816
1677
986
2448
2443
2408
1990
6
6
7
8
8
13
12
13
10
10
9
21
20
20
16
7
11
12
9
8
9
9
11
12
8
10
11
10
9
10
1
1.1
1.1
1.1
1.1
1.8
1.9
1.8
1.1
1.7
1.8
1.9
1.9
2
2.3
2.1
2.4
2.5
2.6
2.7
2.2
2.2
2.2
2.1
2.2
3
3
2.1
2.1
(B - brine, B1 - lower brine in Suakin Deep, B2 ⫽ upper brine in
Suakin Deep and LCL ⫽ Lower Convective Layer in Atlantis II Deep).
Geochemical study of Red Sea brines
1263
BP; Schoell and Faber, 1978) fall along the line for ␦D and
␦18O of present-day Red-Sea surface water, given by the relation: ␦D ⫽ 6 ⫻ ␦18O (Craig, 1969; Fig. 2). The isotope
signatures of Red-Sea interstitial waters in Miocene to Pleistocene sediments, plot nearby (grey area in Fig. 2; Friedman and
Hardcastle, 1974; Lawrence, 1974). The figure outlines the fact
that the isotope signatures cannot be explained by simple
binary mixing of two end-members. In fact, different origins
and processes have to be considered: 1) the brines could be old
sea water which has (or has not) interacted with basalt and/or
with old sedimentary deposits; and 2) the isotopic fractionation
may have occurred during mineral precipitation and dissolution. These processes are discussed below.
The 87Sr/86Sr ratios range from 0.70913 ⫾ 0.00002 (2␴) in
the Thetis Deep filled with bottom sea water, to 0.70656 ⫾
0.00002 for the Nereus brine (Table 4). The Sr concentrations
range from 0.09 mmol/L (Thetis) to 0.63 mmol/L (Nereus),
which is in good agreement with the results obtained by Zierenberg and Shanks (1986) on brines of the Atlantis II, Valdivia
and Suakin deeps. Observed 87Sr/86Sr ratios plotted against
inverse Sr concentrations show a trend of low 87Sr/86Sr value in
Sr enriched brines such as those of Nereus and Atlantis II, and
high 87Sr/86Sr ratios at low Sr concentrations, such as those
occurring at Valdivia or Thetis (Fig. 3).
5. DISCUSSION
Fig. 2. ␦D versus ␦18O diagram with isotopic compositions of bottom
Red-Sea water and brines in the Atlantis II, Chain, Suakin, PortSoudan, Valdivia and Nereus deeps. The equation of Craig, 1969
relating surface Red-Sea water (␦D ⫽ 6 ⫻ ␦18O) and the domain of
interstitial waters in Miocene to Pleistocene Red-Sea sediments (1) is
calculated using the data of DSDP Leg 23 (Friedman and Hardcastle,
1974; Lawrence, 1974). IW ⫽ interstitial water; (2) Old sea-water
signature from Schoell and Faber (1978).
(⫹6‰). The ␦18O values range from ⫹3‰ for the Nereus brine
to ⫹1‰ for the lower brine of Atlantis II. Thus, Atlantis II has
the lightest water in oxygen and in hydrogen (Fig. 2). Two
palaeowaters corresponding to the last glacial maximum
(18000 yr BP), and to the last climatic optimum (8000 –5000 yr
Table 4.
Previous studies on the brines of some Red-Sea deeps point
to varied geochemical properties, and the results obtained here
show important heterogeneities in the chemistries and isotopic
compositions of the brines. Several aspects of brine evolution
will be successively addressed in the discussion: 1) the formation processes including the contribution of basaltic Sr; 2) the
existence of secondary mechanisms which may modify the
brines; and 3) the origin and characteristics of each studied
brine. For deeps containing more than one brine layer, we were
especially interested in the data for the lower pools, as they are
thought to be the most illustrative of the origins of the brines.
5.1. Processes of Brine Formation: Hydrothermal
Alteration of Oceanic Crust and Dissolution of
Evaporites
The salinities of the studied Red-Sea brines (between 145
and 270; Table1) cannot only be due to hydrothermal sea
87
Sr/86Sr ratios and Sr concentrations (mmol/1) in the brines of the studied deeps. The two extreme cases are outlined in grey.
Deeps
Sample
Depth
(m)
Temperature
(°C)
[Sr]
(mmol/1)
87
Sr/86Sr
(⫾2␴ in 10⫺5)
Atlantis II
Atlantis II
Atlantis II
Atlantis II
Chain
Suakin
Suakin
Port-Soudan
Valdivia
Thetis
Nereus
B6-2 (LCL)
B6-6 (UCL1)
B4-7 (UCL 2)
B6-8 (UCL 3)
B5-1
B1-2
B1-6
B2-1
B7-3
B8-12
B9-1
2158
2042
2016
2007
2090
2829
2778
2643
1600
986
2453
66.4
55.8
50
38.2
45.4
23.5
23.3
35.9
33.7
21.6
29.9
0.479
0.225
0.194
0.179
0.468
0.189
0.185
0.243
0.167
0.092
0.625
0.70696 ⫾ 2
0.70747 ⫾ 2
0.70769 ⫾ 2
0.70782 ⫾ 1
0.70709 ⫾ 1
0.70801 ⫾ 1
0.70917 ⫾ 2
0.70850 ⫾ 2
0.70880 ⫾ 1
0.70913 ⫾ 2
0.70656 ⫾ 2
1264
M. C. Pierret et al.
almost zero (Edmond et al., 1982; Von Damm et al., 1985;
1988; 1995). The chemical mobilisation of various elements
depends on both the temperature and the depth of the circulation: SO4 is removed from sea water above 150°C (by anhydrite precipitation and then sulfate reduction), the removal of
sea water Mg is due to smectite formation at low temperatures
(⬍200°C) and chlorite crystallisation at higher temperatures.
Below 1 km depths at 300 to 400°C, transformation of sulfur in
rocks to H2S is accompanied by leaching of some metals (Fe,
Zn, Pb, Cu). Li and other alkalis (K, Rb, Ba) are leached from
the rocks at moderate to high temperatures (Honnorez et al.,
1983; Rosenbauer et al., 1983; Alt, 1995; Hannington et al.,
1995). In an active convection system, large volumes of sea
water circulate at low temperature through a thin layer of rock,
whereas only small amounts of sea water penetrate into the
deep rocks to react at higher temperature. In the Red Sea
system, thick evaporite layers constitute another type of rock
where sea water can react and reach very high salinities during
circulation. The main evaporite minerals are halite, anhydrite,
gypsum and sometimes magnesite, so that Ca, Na, Cl, Mg or
SO4 enrichments of brines relative to sea water may be due to
evaporite leachings.
Specific chemical ratios can also be used to identify a hydrothermal influence in Red Sea deeps, for instance the Fe/Mn
ratio is very different in hydrothermal fluids, hydrothermal
brines and non-hydrothermal brines (Table 5). Hydrothermal
fluids are more enriched in Fe than Mn relative to sea water.
The Zn/Cl ratio allows differentiation of Zn coming from sea
water/evaporite dissolution, and Zn resulting from interaction
with basalt (Table 5). Hydrothermal high temperature circulation of fluids with Cl involves an increase in some metals like
Zn in the fluids with a Cl concentrations close to that of sea
water (high Zn/Cl), whereas evaporite leaching includes a large
increase of Cl without clear Zn enrichment (low Zn/Cl).
The histograms in Figure 4 show that Atlantis II and Chain
deeps contain hydrothermal fluids which have interacted with
the oceanic crust at high temperatures (depletion in Mg and
SO4, enrichment in Li, Ba, Rb, K, Fe, Mn, Zn, Cu, high Fe/Mn
and Zn/Cl ratios). Nereus brine is depleted only in SO4 compared to sea water, but is enriched in Rb, K, L, Ba, Fe, Zn and
Cu, indicating a hydrothermal influence. The brines of the
Valdivia and Suakin deeps do not show clear hydrothermal
origins. The Fe/Mn ratio of Valdivia brine is high (of the same
order of magnitude as that of hydrothermal fluids), however,
Fig. 3. 87Sr/86Sr ratios versus 1/[Sr] (Sr in ppm) of the different
studied brines. Thetis Deep without brine can be considered to be
normal deep sea water characterised by a 87Sr/86Sr ratio of 0.70913 and
a Sr content of 0.092 mmol/L. The Sr isotopic system can be simplified
into a two end-member model (section 5.1.2.). The positions of the
end-member A (basaltic component) and of the two extreme endmembers (B) have been reported (point B⬘ corresponding to Nereus
brine and B⬙ to Valdivia brine). The B end-member is essentially
controled by evaporites and is located along the dotted line.
water/basalt interactions and require halite (evaporite) dissolution, as in the Salton-Sea hydrothermal system containing hypersaline metalliferous brines (McKibben et al., 1988). It is
proposed that the formation of the brines from the most studied
deep, Atlantis II, requires two main processes: hydrothermal
alteration of the oceanic crust by sea water and chemical
dissolution of thick Miocene evaporites (NaCl, CaSO4) by
circulating sea water. We have attempted to quantify these two
processes using different geochemical tracers.
5.1.1. Elemental concentrations
When sea water interacts with a basaltic rock at high temperature during hydrothermal circulation, the physical, chemical and isotopic characteristics of rock and waterboth change.
Compared to the original sea water, hydrothermal solutions at
midocean ridges (MOR) have higher temperature and elevated
concentrations of Li, K, Rb, Ba, Si, Ca, Sr, Mn, Fe, Zn, Cu,
lower pH and SO4 and Mg concentrations that decrease to
Table 5. Fe/Mn and Zn/Cl chemical ratios of different brines from Red-Sea (this study).
Ratios
Suakin
Port-Soudan
Fe/Mn
Zn/Cl
0.0145
0.93
0.885
1.24
Ratios
HF
(1)
Bannock
Fe/Mn
Zn/Cl
2.45
370
0.004
nd
Valdivia
2.24
1.15
(2)
Tyro
(2)
0.0165
nd
Nereus
Chain
Atlantis II
0.06
4.22
0.04
1.9
0.91
7.59
Orca
(3)
0.13
nd
Salton-Sea
(4)
1.1
178
Also reported for comparison are the Fe/Mn and Zn/Cl ratios for an average hydrothermal fluid (HF), several Mediterranean brines (Bannock and
Tyro), a brine from Orca basin and a hydrothermal brine from Salton Sea. (1; Edmond and Von Damm, 1985; Von Damm, 1995), (2: Saager et al.,
1993; De Lange et al., 1990; (3: Sackett et al., 1979), (4: McKibben and Williams, 1989; Thompson and Fournier, 1988). nd ⫽ not determined.
Geochemical study of Red Sea brines
Fig. 4. Histograms of pH (f), temperature (°C) (c) and some chemical elements (a,b, d, e, g, h, i, j) of brines of the Suakin,
Port-Soudan, Atlantis II lower layer, Chain, Valdivia, Thetis and Nereus deeps. Thetis Deep which does not contain brines,
represents generic deep Red-Sea water. The mean composition of hydrothermal fluid (from MOR) has been reported for
comparison, as well as the Fe/Mn ratios measured for the hydrothermal brine of the Salton Sea (McKibben and Williams,
1989), and non hydrothermal brines of the Orca basin (Sackett et al., 1979), and Bannock and Tyro basins (Saager et al.,
1993; De Lange et al., 1990).
1265
1266
M. C. Pierret et al.
Table 6. Values of the different contributions in the Sr isotopic budget.
Brines
[Sr] ppm
Nereus
Atlantis II
Chain B
Port-Soudan
Suakin
Valdivia
Thetis
46.81
34.97
34.17
18.35
14.92
12.40
8.04
87
Sr/86Sr
0.70658
0.70696
0.70709
0.70850
0.70801
0.70880
0.70913
[Sr] ppm End-MB
␣ (%)
␣sw (%)
32
27
25.1
17.2
13
12.1
/
13.7
7
7.9
0.9
1.5
0.2
/
33.6
25.1
24.6
13.2
10.7
8.9
/
The fourth column represents the Sr concentration of the variable B end-member corresponding to figure 3. The contribution of basaltic Sr (in %)
is calculated according to the equation: [Sr]s 䡠 (87Sr/86Sr)s ⫽ ␣ [Sr]A 䡠 (87Sr/86Sr)A ⫹ ␤ [Sr]B 䡠 (87Sr/86Sr)B, where ␣ ⫹ ␤ - 1. ␣ ⫽ contribution of
the A end-member (basalt) and ␤ ⫽ contribution of the B end-member (variable). The ␣sw factor is calculated with sea water as the end-member B.
See text for more explanation.
this ratio is due to very low Mn concentrations (0.06 mmol/L),
the lowest Mn content measured during the present study,
combined with 0.13 mmol/L of Fe (Table 2). For comparison,
Fe and Mn reach respectively 1.4 and 1.52 mmol/L, respectively 10 and 25 times more than the hydrothermal brine of the
Atlantis II Deep. The Fe/Mn ratio of the Valdivia brine is not
explained by hydrothermal activity, but more probably by
extreme Mn depletion of this deep. In view of the absence of
evidence for a hydrothermal component, the high salinities may
be simply due to submarine dissolution of exhumed evaporites
without deep circulation. Similar dissolution processes were
suggested for brine formation in the Orca Basin (Gulf of
Mexico; Addy and Behrens, 1980), and in the Bannock and
Tyro basins (eastern Mediterranean Sea; Camerlenghi, 1990).
The details of the processes in each deep are discussed below.
5.1.2. Sr isotopic compositions
The Sr in the brines may come from four major sources: sea
water, dissolution of biogenic carbonates, marine evaporites,
and basalts. The 87Sr/86Sr isotopic ratios of sea water (0.7091–
0.7092), biogenic carbonates (in equilibrium with sea water)
and the Pliocene-to-Pleistocene sedimentary units (0.7089;
Burke et al., 1982; Beets, 1992) are very similar. Sea water has
low Sr concentrations (0.09 mmol/L; 8 ppm). Old sedimentary
deposits, mainly consisting of biogenic carbonates, clay, and
silt (Supko et al., 1974), are less soluble than evaporites. The
Miocene evaporite units (mainly halite and anhydrite) have a
constant 87Sr/86Sr ratio at 0.70894 (Zierenberg and Shanks,
1986), but Sr concentrations that are highly variable (between
10 and 2000 ppm; Whitmarsh, 1974; Zierenberg and Shanks,
1986) because of variable proportions of anhydrite and halite.
Basaltic Sr in hydrothermal fluids decreases the 87Sr/86Sr ratio
relative to sea-water (e.g., Albarède et al., 1981; Vidal and
Clauer, 1981; Dosso et al., 1991). The Red-Sea basalts have
87
Sr/86Sr ratios ranging from 0.70257 to 0.70308 (with an
average of 0.7027) and an average Sr concentration of about
140 ppm (Altherr et al., 1988; Eissen et al., 1989; Bosch, 1990).
Variations of 87Sr/86Sr ratios between modern sea water
(0.70913) and the evaporites (0.70894) are small. In addition,
the Sr concentrations of evaporites (as large as 2000 ppm) are
significantly higher than sea water (8 –9 ppm). This is why the
mixing system may be simplified into a two-component mixing
model. To calculate the contribution of basaltic Sr to each
brine, the Sr isotopic system is simplified into this two end-
member model: one component (A) being the oceanic crust and
the second end-member (B) representing Sr from other sources
essentially reflecting evaporite dissolution (Fig. 3). Thus, the B
end-member has a constant 87Sr/86Sr ratio (the same as the
evaporite units: 0.70894), but variable Sr concentrations due to
(1) highly variable Sr concentrations in the evaporite units
(Whitmarsh, 1974; Zierenberg and Shanks, 1986), (2) the variable thickness of the Miocene evaporites in the Red Sea (Stoffers and Kühn, 1974), and (3) the different conditions of
evaporite dissolution depending on the temperature and chemical composition of the migrating fluids. The location of the
different B end-members fluctuates between the B⬘ and B⬙
positions (the B⬘ end-member representing the Nereus brine
and the B⬙ end-member the Valdivia brine) along the dotted
line in Figure 3. The contribution of the A and B end-members
can be calculated with the following relation (Faure, 1986)
(Table 6, Fig. 3, S ⫽ sample):
[Sr]S 䡠 (87Sr/86Sr)S ⫽ ␣ [Sr]A 䡠 (87Sr/86Sr)A ⫹ ␤ [Sr]B
䡠 (87Sr/86Sr)B
(1)
where ␣ ⫹ ␤ ⫽ 1, ␣ ⫽ basaltic Sr content and ␤ ⫽ Sr content
of the B end-member.
The Nereus brine yields the highest contribution of basaltic
Sr (␣) at 13.7%. The ␣ values for the Atlantis II and Chain
brines range from 7 to 8%. This contribution is even smaller for
the Suakin (1.5%) and Port-Soudan (0.9%) brines and almost
insignificant for that of Valdivia (0.2%). We have also calculated the ␣SW factor taking sea water as the end-member B
(Table 6) as is the case for common sea water-basalt hydrothermal systems at oceanic ridges. In this case, ␣SW is greater
than in the previous case, but the evaporite Sr contribution has
not been integrated into the budget calculation.
Zierenberg and Shanks (1986) calculated the contribution of
basaltic Sr by isotope mass balance using only the isotopic
composition without concentrations:
XA 䡠 [87Sr/86Sr]A ⫹ XB 䡠 [87Sr/86Sr]B ⫽ [87Sr/86Sr]S
with XA⫹XB ⫽ 1
(2)
where XA represents the fraction of basaltic Sr, and XB represents the fraction of evaporitic Sr in the brine. The authors
indicated that approximately 31% the brine Sr is derived from
basalts in the lower brine of Atlantis II. This amount is higher
than that obtained here because calculations that do not take
Geochemical study of Red Sea brines
1267
Table 7. Values of the saturation index of three carbonates (calcite, aragonite and dolomite), two sulphates (anhydrite and gypsum) and halite, quartz
and amorphous silica in Red-Sea brines. NRSW ⫽ Normal Red Sea Water.
I ⫽ Q/K
Calcite
Aragonite
Dolomite
Gypsum
Anydrite
Halite
Quartz
Amorph.Si
NRSW
3.1
2
99.8
0.26
.148
4.13 䡠 10⫺3
0.148
7.16 䡠 10⫺3
Suakin
Port-Soudan
Valdivia
Nereus
Chain
Atlantis II
2.08
3.49
91.83
0.61
0.7
0.1
2.75
0.13
0.83
0.54
4.42
0.7
0.8
0.33
12.68
0.77
0.53
0.34
3.57
0.91
1.06
0.49
5
0.29
10
6.47
105.2
0.86
0.733
0.41
1.85
0.1
0.94
0.61
0.94
0.88
1.54
0.89
3.07
0.21
0.55
0.59
0.33
0.84
2.34
0.94
10
0.92
into account the concentrations of the different end-members
are less precise.
Volcanic glasses were observed in biodetrital Red Sea sediments (Schneider et al., 1976; Boger and Faure, 1976; Boger
et al., 1980) and in core sediments of the Suakin and Nereus
deeps (Bäcker et al., 1975; Jedwab et al., 1989; Bosch et al.,
1994; Pierret, 1998). Alteration and leaching of the basaltic
detritus may cause an increase of the contribution of the basaltic Sr in the brines. It will be shown below that the brines are
generally undersaturated with respect to amorphous silica (saturation indices of 0.13 and 0.10 in the brines of Suakin and
Nereus, Table 7). With a saturation index of 0.92 the brine of
Atlantis II Deep is the only one close to equilibrium, and no
volcanic detritus was observed in its sediments. This means that
the basaltic Sr of the Atlantis II brine is not supplied by
dissolved volcanic particles.
5.1.3. The oxygen and hydrogen isotope compositions
The ␦18O and ␦D values of the brines during hydrothermal
circulation seem to depend on water/rock interactions. The
differences between hydrothermal-fluid (HF) and sea-water
(SW) signatures (⌬HT-SW) due to fluid/basalt interaction at mid
oceanic ridge (MOR) are: ⌬HT-SW(␦D) from 0 to ⫹4 and
⌬HT-SW (␦18O) from 0 to ⫹2.5‰ (Bowers and Taylor, 1985;
Bowers, 1989; Böhlke et al., 1994; Shanks et al., 1995). During
migration sea water may also interact with the thick Miocene
evaporites consisting mainly of halite and/or anhydrite with
some units of black shales rich in pyrite and sphalerite, and also
with the Pliocene and Pleistocene sediments consisting of biogenic and detrital materials (Supko et al., 1974). Fluid-sediment
interactions cause isotopic evolution of water: ⌬(␦D) from 0 to
⫺3 and ⌬(␦18O) from 0 to ⫹2.5 ‰ (Bowers and Taylor, 1985;
Bowers, 1989; Böhlke et al., 1994; Shanks et al., 1995). The
oxygen and hydrogen signatures of the brines also strongly
depend on the signature of the original sea water. Schoell and
Faber (1978) estimated the oxygen and hydrogen isotope compositions of palaeoseawaters from the 18O-variations in foraminifera from Red-Sea sediments combined with the Craig
relation (␦D ⫽ 6 ⫻ ␦18O; Fig. 2). Our data show that the
bottom sea waters sampled in 1992 vary with the sampling
location (Table 3 and Fig. 2). These wide temporal and spatial
variations are a source of uncertainty in the determination of
the original oxygen and hydrogen isotope compositions of the
brines. However, hydrothermal activity started less than 20,000
yr ago (Bäcker and Richter, 1973) and it is reasonable to
assume that over 20,000 yr the isotope compositions of the
Red-Sea water have scattered around the Craig line, between
sea waters from two extreme climatic events (18000 yr BP and
8000 –5000 yr BP; Fig. 2). Except in the case of Suakin Deep,
which has a very peculiar signature, the brine data plot in a
relatively restricted area in the ␦D vs. ␦18O diagram (Fig. 2). A
hydrothermal origin implies an increase of the ␦D and ␦18O
values, and interactions between sea water and marine sediments can explain a decrease of ␦D. However, the oxygen and
hydrogen isotope compositions can also be modified by fractionations during secondary processes that occur after brines
are trapped in the deeps.
5.2. Chemical Processes in the Brines
When a brine is trapped in a submarine depression, different
processes can take place. Mineral dissolution and precipitation
may occur, e.g., biogenic carbonate dissolution, clay precipitation, etc. Such processes may involve oxygen and hydrogen
isotopic evolution. O’Neil et al. (1976) showed that the rates of
oxygen and hydrogen isotopic exchange may be more rapid in
saline waters, which means that the isotopic compositions of
brines may be modified during chemical reactions. The
fractionation factors for mineral-water (min-H2O) systems
(␣min-H2O ⫽ [1 ⫹ (␦min/1000)]/[1 ⫹ (␦H2O/1000)]) are ␣ ⬎ 1
for oxygen of clays and carbonates and ␣ ⬍ 1 for hydrogen of
clays (Savin and Epstein, 1970; O’Neil et al., 1969). Therefore,
chemical precipitation of clays and carbonates will tend to
deplete water in 18O and enrich water in D.
The saturation index (I) of several minerals has been studied
in the different available brines. It represents the ratio of the
ionic activity product (Q) of any mineral to its solubility
constant (Ks): I ⫽ Q/Ks. When I is equal or higher to 1, the
mineral can precipitate. Alternatively, the solution is undersaturated with respect to the mineral when I is smaller than 1, and
the mineral may dissolve. The EQPO program (Risacher,
1992), based on Pitzer’s model (Pitzer, 1979; 1984; Felmy and
Weare, 1986; Palaban and Pitzer, 1987; Greenberg and Moller,
1989; Spencer et al., 1990), has been used to calculate activity
coefficients (␥i) and activities (ai), and mineral-saturation indices from chemical compositions and temperatures of the brines.
The brines of Port-Soudan, Valdivia, Chain and Atlantis II
deeps are clearly undersaturated with respect to calcite and
aragonite, allowing dissolution of biogenic carbonates, whereas
preservation of skeletal carbonates is predicted in the Suakin
and Nereus deeps (Table 7). Since the brine in the Atlantis II
Deep is undersaturated with respect to the biogenic carbonates,
probably since the beginning of hydrothermal activity (An-
1268
M. C. Pierret et al.
schutz and Blanc, 1993b), we modeled its isotopic evolution
due to the dissolution of biogenic carbonate. Moreover, the
available data on Atlantis II Deep are the most complete
(history of the deep, volume and age of the brine). The biodetrital sedimentation rate is estimated to be 10 cm/1000 yr on
average in the Red Sea and constitutes approximately 65% of
biogenic carbonates (Bäcker, 1976; Blanc et al., 1998; Pierret,
1998). The surface of the Atlantis II Deep is 60 km2 and the
deed was filled by a stable acidic brine about 15,000 yr ago.
During 15,000 yr, 58.5 ⫻ 106 m3 of biogenic carbonates,
Oxy
representing Ncarb
⫽ 3.86 ⫻ 1012 mol of oxygen, may have
Oxy
been dissolved by 3.94 km3 of brines representing NBrine
⫽
12
262.4 ⫻ 10 mol of oxygen. The dependence of the evolution
of ␦18O of the water on the quantities of dissolved calcite is
given by:
Oxy
Oxy
Oxy
Oxy
␦18Ofinal ⫽ ␦18Oinitial ⫻ [NBrine
/(NBrine
⫹ NCaCO3
) NCaCO3
]
Oxy
Oxy
Oxy
Oxy
⫹ ␦18OCaCO3 ⫻ [NCaCO3
/(NBrine
⫹ Ncarb
) NCaCO3
]
⫽ ␦18Oinitial ⫹ 18O2 ,
with
␦18OCaCO3 ⫽ ⫹13.9‰/SMOW (O’Neil et al., 1969)
(3)
where 18O2 is the variation of ␦18O. The maximum value of
O2 is ⫹0.2 ‰ if all the biogenic carbonates were dissolved in
the Atlantis II Deep. Chemical precipitation of carbonates from
brine (chemical sedimentation) would reduce this slight change
(numerous secondary carbonates have been identified in the
Atlantis II sediments; Hofman et al., 1998; Pierret, 1998). Thus,
carbonates representing the most important mineral phase dissolved by the brine (the siliceous species were well preserved
in the Atlantis II sediments; Anschutz and Blanc, 1993), have
a limited influence on the isotopic compositions of the brine in
the Atlantis II Deep. The oxygen and hydrogen isotope signatures are therefore mainly due to the characteristics of the
fluid(s) forming the brines and secondary fractionations may be
considered as having only a minor influence.
Monnin and Ramboz (1996) calculated the anhydrite-saturation index in different Red-Sea brines. Using chemical data of
the brines collected in 1972, 1973 and 1977, the brine of Suakin
Deep was found to be undersaturated (I ⫽ 0.3– 0.4). Using
1977 data, anhydrite is approximately in equilibrium (I ⫽ 0.93)
with the brine of Valdivia Deep. In the case of the lower brine
of Atlantis II, these authors showed that anhydrite is not permanently in equilibrium with the brine; two periods of saturation (1966 and 1976) were followed by undersaturation periods.
They proposed a relation between anhydrite saturation and
hydrothermal activity. Since brines resulting only from evaporite dissolution must be saturation or nearby so with respect to
anhydrite, they suggested that undersaturated brines contain a
hydrothermal component. The 1992 data show that the Suakin
brine is always undersaturated, that the Valdivia brine is maintained at equilibrium with CaSO4 and that the lower brine of
Atlantis II Deep is in a saturation period (Table 7). However,
chemical evolution of a brine formed only by evaporite dissolution can come under modifications of the equilibrium with
anhydrite, such as sulfate reduction by bacteria and Ca precipitation as secondary carbonates. Thus, the undersaturation of
18
the Suakin brine with respect to anhydrite is not necessarily an
indication of hydrothermal activity.
5.3. Characteristics and Origins of the Different Red-Sea
Brines
5.3.1.Thetis
Thetis is the only deep of this study without brine; it is filled
with generic Red-Sea bottom water. The waters of this deep
(CTD-rosette B8) can, therefore, be considered to be the sea
water end-member.
5.3.2. Atlantis II
The Mg and SO4 concentrations of the Atlantis II brine are
lower than in sea water (Fig. 4e), although evaporites (mainly
consisting in NaCl and CaSO4) were leached as indicated by
elevated Na and Cl contents of the brine. Compared to sea
water, the Mg and SO4 depletions are typical for hydrothermal
fluids (Fig. 4e), and result from interactions between sea water
and the oceanic crust during hydrothermal circulation at temperatures above 250°C (cf. section 5.1). In addition, Li, Ba, Fe,
Zn, Cu, Mn, K and Rb which characterize hydrothermal fluids
are the most concentrated in the brine of Atlantis II (Fig.
4b,g,h,i). The Fe/Mn and Zn/Cl ratios are also characteristic of
a hydrothermal origin (Table 5 and Fig. 4d,j). This brine has the
highest temperature and salinity and, as is typical for a hydrothermal origin it is also the most acidic (Fig. 4a,c,f). The stable
isotope signatures (average ␦18O ⫽ ⫹1.07‰ and ␦D ⫽
⫹7.1‰) are close to the ␦18O and ␦D values (⫹1.21 and
⫹7.4‰, respectively) obtained by Schoell and Faber (1978),
and also the values (⫹0.9 ‰ and ⫹6.25‰) obtained by Blanc
et al. (1995). The values clearly express the depletion in heavy
isotopes of oxygen and hydrogen compared to present-day
bottom sea water. Schoell and Faber (1978) explained these
stable isotopic characteristics by a supply of Red-Sea palaeowater of the last climatic optimum (8000 –5000 yr BP). The
authors proposed that high-temperature exchanges between
minerals and waters are improbable. Blanc et al. (1995) suggested that the brines resulted from mixing processes between
Red-Sea palaeowaters of pluvial and warm periods, interstitial
waters expelled from evaporitic sediments, and hydrothermal
inputs. Bowers (1989), Böhlke et al. (1994) and Shanks et al.
(1995) explained variations in oxygen and hydrogen isotopic
compositions as being due to water/basalt interactions during
hydrothermal circulation at MOR (section 5.1.3.) These fractionations systematically involve an increase in ␦18O of the
fluid. In this case, the initial sea water which circulated before
filling Atlantis II had a lower ␦18O value than the present-day
brine. Only certain interstitial waters from Miocene deposits
(which have variable values) have similar low ␦D and ␦18O
values (Friedmann and Hardcastle, 1974; Lawrence, 1974). The
brines may, therefore, result from mixing between sea water
and Miocene interstitial water which interacted with basalts
during hydrothermal circulation. The ␦D of the Atlantis II brine
is lower than the lowest sea-water ␦D from the last climatic
optimum (8000 –5000 yr BP). Only interactions between sea
water and old sediments can explain the ␦D measured signature.
The 87Sr/86Sr ratios of the brines are lower than that of sea
Geochemical study of Red Sea brines
water, which requires approximately a 7% addition of basaltic
Sr (Table 6). This Sr isotopic composition (87Sr/86Sr ⫽
0.70696) is similar to those reported by Zierenberg and Shanks
(1986) and Blanc et al. (1995) and can result from hydrothermal alteration of basalt by sea water. The relatively low value
of the basaltic contribution was explained by the occurrence of
a large quantity of Sr provided by the leaching of Miocene
anhydrite, consistent with the high concentrations of Na and Cl
from leaching of evaporites. Evaporite dissolution can, therefore, be envisaged during hydrothermal circulation. The dissolution of biogenic carbonates in the acidic brine, the dissolution
of Miocene anhydrite and the hydrothermal alteration would
result in a high Ca concentration. Thus, the Atlantis II brines
consist of a salty fluid reacting with basalt at high temperature
indicating significant hydrothermal activity.
5.3.3. Chain deep
The Chain Deep is adjacent and connected to Atlantis II to at
a depth of 1980 m. This depth is above the LCL/UCL1 transition zone in Atlantis II, and it does not allow a direct discharge of the lower brine from the top of Atlantis II into the
Chain basin. However, the isotopic signatures of the lower
Atlantis II and Chain brines are very close. The ␦18O and ␦D
values are of the same order of magnitude (Fig. 2 and Table 3)
and both 87Sr/86Sr ratios, 0.70696 ⫾ 0.00002 (Atlantis II) and
0.70709 ⫾ 0.00001 (Chain) and Sr concentrations (0.48
mmol/L and 0.47 mmol/L) are similar. Thus, in Figure 3, Chain
and Atlantis II lower brine data are very close to one another.
The salinity and K, Ca, Rb, Mg, SO4 concentrations are also
similar (Fig. 4a,b,e). The top of the LCL brine in Atlantis II and
the top of the Chain brine are at the same depth. Therefore, we
suggest the same origin for both: the brine in the Chain Deep
could have been derived from the lower brine of the Atlantis II
Deep. As direct communication above the seafloor cannot
occur because of the topography, it may have occurred beneath
the deeps along a fracture system. This hypothesis would also
explain why the top of the lower brine of Atlantis II Deep and
the top of the brine of Chain Deep are at the same depth.
The temperature and various elemental concentrations (Fe,
Mn, Zn) in the brine of Chain Deep are lower than in Atlantis
II Deep (Fig. 4h,i). Loss of heat may have occurred during
transport. Also, some chemical elements could have been fixed
in Atlantis II by chemical precipitation, or during transport.
Iron is less mobile than Mn in solution, and the Fe/Mn ratio is
actually lower in the Chain brine than in the Atlantis II lower
brine (Fig. 4d). Thus, the chemical and isotopic parameters
suggest the existence of subsurface connections between the
Atlantis II Deep and the Chain Deep with the consequence that
the brine of the Chain Deep may have resulted from an outflow
from the lower brine of Atlantis II but not by upper direct
discharge.
5.3.4. Nereus deep
The Nereus brine has the lowest 87Sr/86Sr ratio and the
highest Sr concentration of the studied brines. These parameters suggest the highest basaltic contribution (Table 6). The
SO4 concentration of the Nereus brine is lower than in sea
water represented by the Thetis reference (Fig. 4e), but that of
1269
Mg is higher. This distinction represents a significant difference
relative to the Atlantis II brine which has low concentrations of
SO4 and Mg relative to sea water. The enrichment in Mg
relative to sea water (Fig. 4e) may result from leaching of
evaporites. Actually, Mannheim (1974) identified Mg-rich layers in the Miocene evaporite deposits. The concentrations of Li,
Ba, K, Rb, Mn, Zn, and Cu, and the Zn/Cl ratio, support a
hydrothermal influence but with a smaller input than that in the
Atlantis II Deep (Fig. 4b,g,h,i,j). However, the contribution of
basaltic Sr in the Nereus brine (13.7%) appears to be higher
than in the Atlantis II brine (7%). This difference can be
explained by the occurrence of basaltic glasses in the sediments
of Nereus, as fragments of volcanic rocks were described in the
sediments of this deep (Bignell and Ali, 1976; Bonatti et al.,
1984; Jedwab et al., 1989; Bosch et al., 1994) but not in the
Atlantis II Deep. This volcanic-glass may have been dissolved
by the brines and may consequently have increased indirectly
the basaltic component of the brine. Indeed, the Nereus brine is
undersaturated with respect to amorphous silica (Table 7) and
such materials may be dissolved by the brines and increase both
the Si concentration and the basaltic Sr concentration. A precise
evaluation of the contribution of volcanic glasses dissolution is
difficult, but Sr input from dissolution of volcanic fragments
cannot be neglected in the budget. The temperature of the brine
is low (29.9°C; Table 1) and Nereus has the lowest thickness of
brine (15 m, Table 1). The low temperature may be explained
by a significant loss of heat at the sea water-brine interface due
to the high surface/volume ratio of the brine.
The Sr isotope composition, depletion in SO4, and enrichment of some other elements (Li, Ba, K, Mn, Zn, Cu) in
comparison with sea water suggest interactions between sea
water and oceanic crust above 150°C. In this case, the ␦D and
␦18O values of the initial sea water increased and the induced
variations are: ⌬(␦D) from 0 to ⫹4‰, ⌬(␦18O) from 0 to
⫹2.5‰ (see section 6.1.c). The metal-enriched sediments of
Nereus are Holocene (Bonatti et al., 1984), but it is impossible
to date the brine and to determine the age and isotope composition of the initial sea water (cf. section 5.1.3.). Also, the
hydrothermal circulations may be cyclical and initial sea waters
of different ages with varied isotopic signatures may have
interacted. However, the ␦D and ␦18O values of the Nereus
brine being higher than that of the bottom sea water (Fig. 2),
may be explained by interactions between sea water and oceanic crust.
The brine is saturated with respect to calcite and aragonite,
thus the ␦18O value did not increase because of carbonate
dissolution. The residence time of biogenic carbonates within
the Nereus brine is short because of the sedimentation rate and
no isotopic exchanges at equilibrium are possible. In addition,
the sediments contain beds of biogenic carbonates and in some
places, secondary carbonates such as siderite, rhodochrosite
and manganosiderite in low quantities. Small amounts of secondary clays were also identified in the Nereus sediments
(Bignell and Ali, 1976; Bonatti et al., 1984; Jedwab et al.,
1989). Precipitation of carbonates implies a decrease of the
␦18O values of water, while precipitation of clays implies a
decrease of the ␦18O values and an increase of the ␦D values.
However, such processes seem to have been negligible relative
to fractionation during sea water-basalt interactions. Thus, the
oxygen and hydrogen signatures are probably mainly due to
1270
M. C. Pierret et al.
interactions during migration of the hydrothermal fluids. In
summary, it may be assumed that the Nereus brine has been
formed from a fluid of hydrothermal origin with implied fixation of SO4 at a lower intensity than in the Atlantis II Deep
(lower Zn, Cu, Li, Ba concentrations in the Nereus brine).
Some chemical elements could have been brought to the brine
by dissolution of basaltic glass. Evaporites were dissolved
during the migration of hydrothermal sea water resulting in
high salinity and high Ca concentrations in the resulting brine
(Fig. 4a,b). The high Ca concentration implies that the leached
evaporite probably contained a high proportion of anhydrite.
5.3.5.Port-Soudan deep
The Port-Soudan Deep is filled with a brine having a temperature of 35.9°C and a salinity of 214.5 (Table 1) implying
evaporite dissolution. The brine is enriched in Mg and SO4 with
respect to sea water (Fig. 4e). This means that it did not mainly
consist in a high-temperature hydrothermal fluid, as was the
case for the Atlantis II and Nereus deeps. This brine is also
enriched in Li, Ba, Zn, Cu, K, Fe and Mn relative to sea water
(Fig. 4b,g,h,i), which is characteristic of a hydrothermal component. In addition, the Fe/Mn ratio is of the same order of
magnitude as in Atlantis II (Fig. 4d). These data imply that the
brine of the Port-Soudan Deep did not only result from sea
water dissolving evaporites. The contribution of basaltic Sr is
on the order of 1% (Fig. 3 and Table 6). No basaltic fragments
were identified in the sediment cores and the basaltic Sr in the
brine probably results from limited sea water-basalt exchange
during sea water circulation. In this case, it can be envisaged
that the Port-Soudan brine is characterised by a limited hydrothermal component with a limited sea water/basalt interaction.
The fluids mainly supplying the Port-Soudan Deep probably
did not reach a high temperature, and therefore did not penetrate the oceanic crust deeply.
5.3.6. Valdivia deep
The salinity of the brine in the Valdivia Deep (242) implies
dissolution of NaCl evaporites. We observed a clear enrichment
in SO4 and Mg relative to sea water (Fig. 4e), which excludes
the possibiliy that the brine was formed by a hydrothermal fluid
circulating in a basalt environment at high temperature. The
concentrations of Zn and Cu are belows detection limits, the
brine was the least concentrated in Mn, Rb, Ba, Ca and Sr
(Table 2), and Li and Fe concentrations are low. These elements are typically enriched in hydrothermal fluids. The high
Fe/Mn ratio of Valdivia brine is due to its low Mn content
(section 5.1.1.) but not hydrothermal activity. The content of
the basaltic Sr (0.2%; Table 6) in the brine is negligible and the
87
Sr/86Sr ratio (0.70880 ⫾ 0.00001) is close to the Sr isotopic
composition of evaporites (0.70894; Zierenberg and Shanks,
1986). Thus, elemental composition and Sr isotopic data preclude significant contribution of a hydrothermal fluid. The
Valdivia brine has the same oxygen and hydrogen isotopic
ratios as present-day bottom sea water and seems to be a recent
water without a clear hydrothermal input (Fig. 2). As for the
Nereus brine, the enrichment in Mg may result from leaching of
Mg-rich evaporites. Thus, the evaporitic units dissolved in this
case, consisted primarily of halite and magnesite and less
anhydrite.
Two processes can explain the formation of Valdivia brine.
Either it formed by sea water which dissolved an evaporitic
component during subsurface circulation, or as suggested by
Bäcker et al. (1975), the position of the Valdivia Deep on the
slope of the main trough, along a probable major transform
fracture zone, induced salt-dissolution influenced by tectonic
activity. In both cases, the hydrothermal influence was limited.
Nevertheless, the high salinity (S ⫽ 242) and the temperature
(33.7°C) determined for this brine suggest shallow sea-water
circulation before entering the Valdivia Deep.
5.3.7. Suakin deep
The Suakin brine is the least salty and has the lowest temperature of the Red-Sea brines (Table 1, Fig. 4a,c). The pH is
the least acidic and close to that of sea water (Fig. 4f). The K,
Fe, Li, and Si concentrations are the lowest, whereas the Mg
concentration is high (Fig. 4b,e,g,h). No sign of a clear hydrothermal origin was found in the chemical compositions. The
Suakin brine has the most atypical oxygen and hydrogen isotope signatures of this study. The ␦18O and ␦D values are
clearly different from those of sea water (past and present-day)
with ␦D distinctly enriched in deuterium (␦D of about ⫹20‰)
compared to the other brines. The brine’s isotopic data plots in
the field of the ancient interstitial waters (Fig. 2). We may
therefore adopt the idea proposed by Bath and Shackleton
(1984), Friedman et al. (1988) and France-Lanord and Sheppard (1992), that the isotopic oxygen and hydrogen compositions
of the interstitial waters in the ancient Red-Sea sediments
reflect the signature of bottom sea water variations trapped in
these sediments. Savin and Epstein (1970) and O’Neil and
Kharaka (1976) demonstrated that isotopic exchange between
minerals and interstitial waters does not occur in the temperature range of sedimentation and early diagenesis (T ⬍ 100°C).
The brine is saturated with respect to calcite and aragonite, thus
no oxygen variation (increase of ␦18O) occurred by carbonate
dissolution, as was the case for the Nereus brine. Alternatively,
the isotopic signature of oxygen and hydrogen can be explained
if the brine was derived from an old water which dissolved
exhumed evaporites without any other influence. The brine
could have formed similarly to the Bannock or Tyro basins
brines in the Mediterranean Sea (subsurface evaporite dissolution). As is the case for the Suakin brine, the temperature
differences between brines and bottom sea water in the Mediterranean basins are small (brine temperatures 1–2°C higher
than those of the bottom sea water; Boldrin and Rabitti, 1990).
Two layers of brine are observed in the Suakin Deep, but it is
possible to obtain a stratification of the brine with a simple
submarine dissolution of exhumed evaporites, as in the Bannock Basin which contains a double-layered brine (Bregant et
al., 1990). Anschutz et al. (1999) showed that an abnormal
salinity can persist in such depressions over thousands of years
without additional input of salt. In fact, the occurrence of
layered brines in any deep cannot be considered to be a record
of past or present hydrothermal activities.
The Fe/Mn ratio is low (0.014) with values in the interval of
the Bannock (0.004), Tyro (0.01) and Orca (0.13) brines (Fig.
4d). In the absence of hydrothermal inputs, the properties of Fe
Geochemical study of Red Sea brines
and Mn at the sea water/brine interface are dominated by
differences in the redox potential; as the solubility of Mn in
reduced environments is enhanced more than that of Fe. Saager
et al. (1993) showed that the particulate/dissolved ratios are
about 10⫺4 for Mn and 0.35 for Fe in the Bannock and Tyro
brines. Thus, the enrichment factors are about 2000 for Mn and
10 for Fe relative to sea water. In the Suakin brine, the Fe
concentration is very low (0.007 mmol/L) while the Mn concentration (0.48 mmol/L) is typical of a salty basin without
hydrothermal activity.
As in the Mediterranean basins, the high level of stagnation
prevented any mixing process other than molecular diffusion,
while the degradation of organic matter led to the development
of anoxic conditions (Bregant et al., 1990). Jorgensen (1983)
showed that eutrophication stimulates sulphate reduction and
accelerates the formation of H2S. Thus, SO4 concentrations in
brines result from a balance between dissolution of evaporitic
anhydrite and utilisation of SO2⫺
during oxidation of organic
4
matter. The mechanism of hydrogen sulphide generation is well
known in natural anoxic basins (Dryssen and Wedborg 1986;
Anderson et al., 1988; Luther et al., 1990) and can be represented, according to the Redfield et al.’s formulation (Redfield
et al., 1963), by:
2 CH2O ⫹ SO2⫺
N 2 HCO⫺
4
3 ⫹ H 2S
(4)
In the presence of sulphate-reducing bacteria, sulphate ions
replace oxygen as electron acceptors during oxidation of organic matter. In the Tyro Basin, the reduced sulphur (H2S and
R-SH) represent 5% of the sulphate concentration (Bregant et
al., 1990). This is the reason why the SO4/Ca ratio differs from
the anhydrite ratio and why the brine is not in equilibrium with
anhydrite (major constituent in evaporite). As Mg-enriched
evaporitic layers have been identified (Mannheim, 1974), the
Mg enrichment may result from evaporite dissolution, as suggested for the Valdivia brine. The brine is not depleted in Mg
or in SO4, which confirms the lack of high temperature basaltsea water interactions.
The Sr isotopic composition implies a contribution of 1.5%
of basaltic Sr (Fig. 3 and Table 6), appears to beconsistent with
a hydrothermal influence. However, as in the case of the Nereus
Deep, basaltic glass was found in the sediment cores, and the
brine of Suakin is undersaturated with respect to amorphous Si
(I ⫽ 0.13; Table 7). Thus, the basaltic Sr may be due to
dissolution of these volcanic fragments. We propose that the
brines of the Suakin Deep have been formed by old sea water
which dissolved exhumed evaporites without hydrothermal circulation.
6. IMPACT OF HYDROTHERMAL ACTIVITY IN
RED-SEA DEEPS
The brines of the different deeps record various origins, and
the extent of hydrothermal influences in each of them differ:
strong (Atlantis II and Nereus), weak (Port-Soudan) and negligible (Valdivia and Suakin). The occurrence of hydrothermal
fluids supplying deeps has clear consequences on the nature of
the sediments deposited in these deeps.
The Atlantis II Deep has the most metal-enriched sediments
with metal-sulphide units (Fe, Zn, Cu) and massive Fe/Mnoxide beds which represent the only deposit of Fe, Cu, Ag and
1271
Au of economic potential (Guennoc and Thisse, 1982; Mustafa
et al., 1984; Nawab, 1984). The Nereus Deep contains Fe-Mn
sediments (oxi-hydroxides, siderite, Mn-siderite, pyrite,
sphalerite, chalcopyrite), and pyrite, hematite and Mn- and
Fe-carbonates were identified in Port-Soudan cores (Jedwab et
al., 1989; Pierret, 1998). These sediments contain higher concentrations of trace metals (Co, Cr, Zn, Cu) than generic
Red-Sea sediments (Pierret, 1998). The chemical and mineralogical compositions of the sediments of Valdivia Deep are
close to normal biodetritic sediments of the Red Sea (Bäcker et
al., 1975; Pierret, 1998). Suakin Deep differs from the others;
its brine may have been formed without any hydrothermal
contribution. High seismic activity and many tectonic movements have been recorded in the southern part of the Red Sea
and the area of the Suakin Deep (Makris et al., 1991). This deep
could have been created by a tectonic event which fractured the
bottom of the sea and formed a depression. Evaporites may
have been exhumed after the tectonic deformation and the
dissolution of these evaporitic units by sea water could have
induced brine formation. No major transform fault was observed in the Suakin region by Pautot (1983) who suggested
that hydrothermally active deeps are preferentially located at
the junction between rift segments and transform faults. The
lack of evidence for a hydrothermal component in the Suakin
brine agrees well with the non hydrothermal origin proposed
for the framboı̈dal pyrite described in its sediments (Pierret et
al., 2000).
Hydrothermal circulation depends on a network of faults,
porosity of sediments, and the intensity of the geothermal
gradient. The fluid does not necessarily reach the oceanic crust
and the circulation may be shallow. Sea water and basalt may
interact at different temperatures according to the migration
depth of the sea water, strongly influencing the chemistry of the
fluid (fixation of sulphate and/or Mg, enrichment in Li, Ba, and
other metals). On the basis of Sr and Pb isotopic studies, Volker
et al. (1993) showed that the maturest oceanic crust (N-MORB
type) was located in the central part of the Red Sea, corresponding to the area of Atlantis II Deep. Eissen et al. (1989)
proposed that short-lived magmatic chambers (100 –1000 yr)
operate along the Red-Sea rift. Bogdanov et al. (1997) studied
the mineralogy and chemistry of sediment cores from the
spreading centre in the western Wodlark Basin of the western
Pacific Ocean. They showed that these centres are not synchronous along the spreading axis, each having a specific evolution.
Similary, except for the Chain Deep which is connected with
the Atlantis II Deep, the different deeps in the Red Sea may
have evolved independently.
7. CONCLUSIONS
The varied chemical and isotopic data obtained on brines
from different Red-Sea deeps suggest different mechanisms for
brine formation.
1. Some brines result from interactions with oceanic crust at
high temperatures during hydrothermal circulation. For instance, the brines of the Atlantis II and Chain deeps result
from hydrothermal fluids reacting with basaltic rocks at high
temperatures. It can be shown that the lower brine of the
Atlantis II Deep fills the Chain Deep next to it. The brine of
Nereus Deep is enriched in the same elements as the lower
1272
2.
3.
4.
5.
6.
M. C. Pierret et al.
brine of Atlantis II, has a lower 87Sr/86Sr ratio and it is
depleted only in SO4. Fluids supplying this deep have interacted with the oceanic crust at high temperatures, but the
intensity was probably lower than in the case of Atlantis II
Deep. These brines with basaltic Sr contributions between 7
and 13.7%, are the only ones that unequivocally indicate a
hydrothermal influence.
The brine of Port-Soudan Deep yields evidence for a slight
basalt interaction. Sea water probably circulated only superficially in the oceanic crust before discharging in this deep.
The brine of Valdivia Deep mainly results from evaporite
dissolution (especially halite) by recent sea water.
The brines of Suakin Deep having lower temperatures,
salinities, and Fe concentrations differ significantly from the
other brines. The oxygen and hydrogen isotopic signatures
are comparable to those of old interstitial waters. Suakin
Deep may have formed after tectonic activity, and the brines
seem to have resulted from old sea water which dissolved
exhumed evaporitic units without any hydrothermal contribution.
In summary, the formation of the Red-Sea brines involved
three major processes: leaching of Miocene evaporites, convective fluid circulation, and interaction with basaltic crust
and old sedimentary units. Varied combinations of these
processes allow distinction of the brines in the studied
deeps. This conclusion emphasizes the fact that hydrothermal activity is not identical along the whole Red-Sea axis; it
does not favour a unique formational model for the brines.
A striking feature of the brines is the lack of a relationship
between the position of the deeps along the axis and their
evolutionary maturity. Suakin is the closest deep to the
continuous oceanic crust (in the southern part of the Red
Sea), but its brines probably formed without hydrothermal
activity.
The different deeps appear to have evolved independently,
except for Chain Deep which is connected with Atlantis II
Deep. The highest hydrothermal activity seems to have been
in the area of the Atlantis II Deep.
Acknowledgments—We would like to thank sincerely, Dr. G. Faure
(The Ohio State University, USA), Dr. S. D. Scott (University of
Toronto, Canada), an anonymous reviewer and the Editor in charge of
the script for their very thoughtful comments and remarks and also for
their highly improving help in the English presentation. We are also
grateful to the chief scientist of the O.V. Marion-Dufresne, to its
captain and to its crew for completion of the 1992 REDSED expedition.
We thank F. Gauthier-Lafaye for discussions about oxygen and hydrogen isotopic results and F. Risacher for help and avaibility during
application of the EQP0 program. F. Chabaux is especially thanked for
his improving suggestion during the preparation of the script. Many
thanks are also due to Robert Rouault and Jean Samuel (Centre de
Géochimie de la Surface, Strasbourg) and to Etienne Devantibault
(Laboratoire de Tectonophysique, Montpellier) for analytical assistance during the course of the study. This work was funded by IFRTPTAAF and the French Marine Geosciences Committee INSU-95/1.
Associate editor: R. H. Byrne
REFERENCES
Addy S. K. and Behrens E. W. (1980) Time of accumulation of
hypersaline anoxic brine in Orca basin (Gulf of Mexico). Mar. Geol.
37, 241–252.
Albarède F., Michard A., Minster J. F., and Michard G. (1981) 87Sr/
86
Sr ratios in hydrothermal waters and deposits from the East Pacific
Rise at 21°N. Earth Planet. Sci. Lett. 55, 229 – 6.
Alt J. C. (1995) Subseafloor processes in Mid-Ocean Ridge hydrothermal systems. In Seafloor hydrothermal systems: Physical, chemical,
biological and geological interactions (ed. Geophysical Monograph
by the American Geophysical Union) pp. 85–114.
Altherr R., Henjes-Kunst F., Puchelt H., and Baumann A. (1988)
Volcanic activity in the Red Sea axial trough-evidence for a large
mantle diapir ? Tectonophysics 150, 121–133.
Anderson L. G., Dryssen D., and Hall P. O. (1988) On the sulfur
chemistry of a super-anoxic fjord, Framvaren, South Norway. Mar.
Chem. 23, 283–293.
Anschutz P. and Blanc G. (1993a) L’histoire sédimentologique de la
fosse Atlantis II (mer Rouge). Les apports de la micropaléontologie.
C.R. Acad. Sci., Paris série II 317, 1303–1308.
Anschutz P. and Blanc G. (1993b) Le rapport NaCl/eau des boues
minéralisée de la fosse Atlantis II (mer Rouge). calcul de la teneur en
halite des sédiments et implication sur la paléotempérature du milieu.
C. R. Acad. Sci., Paris série II. 317, 1595–1600.
Anschutz P. and Blanc G. (1995a) Chemical mass balances in metalliferous deposits from the Atlantis II Deep, Red Sea. Geochim.
Cosmochim. Acta 59, 4205– 4218.
Anschutz P., Blanc G., and Stille P. (1995b) Origin of fluids and the
evolution of the Atlantis II deep hydrothermal system, Red Sea:
Strontium isotopic study. Geochim. Cosmochim. Acta 59, 4799 –
4808.
Anschutz P. and Blanc G. (1996) Heat and salt fluxes in the Atlantis II
Deep (Red Sea). Earth Planet. Sci. Lett. 142, 147–159.
Anschutz P., Blanc G., Chatin F., Geiller M., and Pierret M. C. (1999)
Hydrographic change during 20 years in the brine-filled basins of the
Red Sea. Deep Sea Res. 46, 1779 –1792.
Anschutz P., Turner J. S., and Blanc G. (1998) The development of
layering, the fluxes through double-diffusive interfaces, and the
location of hydrothermal sources of brines in the Atlantis II Deep;
Red Sea. J. Geophys. Res. C. 103, 27,809 –27,819.
Bäcker H. (1976) Fazies und Chemische Zusammensetzung rezenter
Ausfällungen aus Mineralquellen im Roten Meer. Geol. Jahrbuch.
D17, 151–172.
Bäcker H. and Richter H. (1973) Die rezente hydrothermal-sedimentäre
Lagestätte Atlantis II Tief im Roten Meer. Geol. Rundsch. 62,
697–741.
Bäcker H. and Schoell M. (1972) New deeps with brines and metalliferous sediments in the Red Sea. Nat. Phys. Sci. 240, 153–158.
Bäcker H., Lange K., and Richter M. (1975) Morphology of the Red
Sea central Graben between Subair Islands and Abul Kizan. Geol.
Jahrbuhr. D13, 79 –123.
Bath A. and Shackleton N. (1984) Oxygen and hydrogen istope studies
in sqeezed pore-waters, Deep Sea Drilling Project Leg 74, Hole
525B: Evidence for Mid-Miocene ocean isotopic change. Initials
reports of the DSDP. 74, 697– 699.
Baumann A., Richter H., and Schoell M. (1973) Suakin Deep: Brines
and hydrothermal sediments in the deepest part of the Red Sea. Geol.
Rundsch. 62, 684 – 697.
Beets C. J. (1992) Calibration of late Cainozoic marine strontium
isotope variations and its chronostratigraphic and geochemical applications. Thesis, Vrije Univeriteit, Amsterdam. 133 p.
Bignell R. D. and Ali S. S. (1976) Geochemistry and stratigraphy of
Nereus Deep, Red Sea. Geol. Jahrbuhr. D17, 173–186.
Blanc G. and Anschutz P. (1995) New stratification in the hydrothermal
brine system of the Atlantis II Deep, Red Sea. Geology 23, 543–546.
Blanc G., Boulègue J., and Michard A. (1995) Isotope compositions of
the Red Sea hydrothermal end-member. C.R. Académie des Sciences.
320, 1187–1193.
Blanc G., Anschutz P., and Pierret M. C. (1998) Metalliferous sedimentation in the Atlantis II Deep: A geochemical insight. In Sedimentation and Tectonics of Rift Basins: Red Sea-Gulf of Aden. (ed.
B. H. Purcer and D. W. J. Bosence) pp. 510 –524. Chapman et Hall.
Bogdanov Y. A., Gurvich E. G., Lisitzin A. P., Serova V. V., and
Gorbunova Z. N. (1997) Sediments of the active rift zone in the
western Woodlark basin and the development of hydrothermal and
volcanic activity. Mar. Geol. 142, 143–170.
Boger P. D. and Faure G. (1976) Systematic variations of sialic and
Geochemical study of Red Sea brines
volcanic detritus in piston cores from the Red Sea. Geochim. Cosmochim. Acta 40, 731–742.
Boger P. D., Boger J. L., and Faure G. (1980) Systematic variation of
87
Sr/86Sr ratios, Sr compositions, selected major-oxide concentrations, and mineral abundance in piston cores from the Red Sea.
Chem. Geol. 29, 13–38.
Bohannon R. G. and Eitreim S. L. (1991) Tectonic development of
passive continental margins of the southern and central Red Sea with
a comparison to Wilkes Land, Antartica. Tectonophysics 198, 129 –
154.
Boldrin A. and Rabitti S. (1990) Hydrography of the brines in the
Bannock and Tyro anoxic basins (eastern Mediterranean) Mar.
Chem. 31, 21–33.
Bonatti E., Colantoni P., Lucchini F., Rossi P. L., Taviani M., and
White J. (1984) Chemical and stable aspects of the Nereus Deep
(Red Sea) metal-enriched sedimentation. Mem. Soc. Geol. 27, 59 –
72.
Bonatti E. (1985) Punctiform initiation of seafloor spreading in the Red
Sea during transition from a continental to an oceanic rift. Nature
316, 33–37.
Bosch D. (1990) Evolution géochimique initiale et précose d’un rift:
Systématique isotopique Pb, Sr, Nd du diapir mantellique de Zabargad, de son encaissant gneissique et de son hydrothermalisme: Comparaison avec les sédiments métallifères et les MORB de la Mer
Rouge. Conséquences géodynamiques et métallogèniques. Thèse.
Montpellier II. 231 p.
Bosch D., Lancelot J., and Boulegue J. (1994) Sr, Nd and Pb isotope
constraints on the formation of the metalliferous sediments in the
Nereus Deep, Red Sea. Earth Planet. Sci. Lett. 123, 299 –315.
Böhlke J. K., and Shanks III W. C. (1994) Stable isotope study of
hydrothermal vents at Escanaba Trough: Observed and calculated
effects of sediments-seawater interaction. In Hydrothermal, and Biologic Studies at Escanaba Trough, Gorda Ridge, Offshore Northern
California (eds. J. L. Morton, R. A. Zierenberg, and C. A. Reiss)
U.S.G.S. Bull. 2022, 223–239.
Bowers T. S. and Taylor H. P. (1985) An integrated chemical and
stable isotopic model of the origin of mid-ocean ridge hot spring
systems. J. Geophys. Res. 90, 12,583–12,606.
Bowers T. S. (1989) Stable isotope signature of water-rock interaction
in mid-ocean ridge hydrothermal systems: Sulfur, oxygen, and hydrogen. J. Geophys. Res. 94, 5775–5786.
Bregant D., Catalano G., Civitarese G., and Luchetta A. (1990) Some
chemical characteristics of the brines in Babbock and Tyro Basins:
Salinity, sulfur compounds, Ca2⫹, F⫺, pH, As, PO3⫺
4 , SiO2, NH3.
Mar. Chem. 31, 35– 62.
Burke W. H., Denison R. E., Hetterington E. A., Koepnick R. B.,
Nelson H. F., and Otto J. B. (1982) Variation of sea water 87Sr/86Sr
throughout Phanerozoic time. Geol. 10, 516 –519.
Butuzova G. Y., Drits V. A., Morozov A. A., and Gorschkov A. I.
(1990) Processes of formation of iron-manganese oxyhydroxides in
the Atlantis II and Thetis Deep of the Red Sea. Spec. Publ. Int. Ass.
Sed. 11, 57–72.
Camerlenghi A. (1990) Anoxic basins of the eastern Mediterranean:
Geological framework. Mar. Chem. 31, 1–19.
Charnock H. (1964) Anomalous bottom water in the Red Sea. Nature
203, 591.
Craig H. (1969) Geochemistry and origin of Red Sea brines. In Hot
brines and recent heavy metal deposits in the Red Sea (eds. E. T.
Degens and D. A. Ross) pp. 208 –242. Springer-Verlag, New York.
Degens E. T. and Ross D. A. (1969) Hot brines and recent heavy metal
deposits in the Red Sea. Springer Verlag, New York., 600 p.
De Lange G. J., Middelburg J.J., Van Der Weijden C. H., Catalano
C. H., Luther G. W. III, Hydes D. J., Woittiez J. R. W., and
Klinkhammer G. P. (1990) Composition of anoxic hypersaline brines
in the Tyro and Bannock Basin, eastern Mediterranean. Mar. Chem.
31, 63– 88.
Dosso L., Hanan B. B., Bougault H., Schilling J. G., and Joron J. L.
(1991) Sr-Nd-Pb geochemical morphology between 10° and 17°N on
the Mid-Atlantic Ridge: A near MORB signature. Earth Planet. Sci.
Lett. 106, 29 – 43.
Dryssen D. and Wedborg M. (1986) Titration of dilphides and thiols in
natural waters. Anal. Chim. Acta 180, 473– 479.
1273
Edmond J. M., Von Damm K. L., McDuff R. E., and Measures C. I.
(1982) Chemistry of hot springs on the East Pacific Rise and their
effluent dispersal. Nature 297, 187–191.
Eissen J. P., Juteau T., Joron J. L., Dupré B., Humler E., and
Al’Mukhamedov A. (1989) Petrology and geochemistry of basalts
from the Red Sea axial rift at 18° North. J. Petrol. 30, 791– 839.
Faure G. (1986) Principles of isotope geology. Second Edition. John
Wiley and Sons. U.S.A. New-York, 589 p.
Felmy A. R. and Weare J. H. (1986) The prediction of borate mineral
equilibria in natural waters: Application to Searles Lake, California.
Geochim. Cosmochim. Acta 50, 2771–2783.
France-Lanord C. and Sheppard S. M. F. (1992) Hydrogen isotope
composition of pore waters and interlayer water in sediment from the
Central Western Pacific, leg 129. Proc. O.D.P., Sci. Res. 129, 295–
302.
Friedman I. and Woodcock A. H. (1957). Determination of deuteriumhydrogen ratios in Hawaiian waters. Tellus 9, 553–556.
Friedman I. and Hardcastle K. (1974) Deuterium in interstitial waters
from the Red Sea cores. Init. Rep. DSDP. 23, 969 –970.
Friedman I. and Hardcastle K. (1988) Deuterium in interstitial water
from deep-sea cores. J. Geophys. Res. 93, 8249 – 8263.
Gat J. R. and Confiantini R. (1981) Stable isotope hydrology. Deuterium and 18O in the water cycle. In International Atomic Energy
Agency Vienna.
Greenberg J. P. and Moller N. (1989) The prediction of mineral
solubilities in natural waters: A chemical equilibrium model fot the
Na-K-Ca-Cl-SO4-H2O system to high concentration from 0 to
250°C. Geochim. Cosmochim. Acta 53, 2503–2518.
Guennoc P. and Thisse Y. (1982) Genèse de l’ouverture de la mer
Rouge et des minéralisations des fosses axiales. Doc. Bourr. Rech.
Géol. Min. 51, 89. p.
Hannigton M. D., Jonasson I. R., Herzig P. M., and Petersen S. (1995)
Physical and chemical processes of seafloor mineralization at MidOcean Ridges. In Seafloor Hydrothermal Systems; Physical, Chemical, Biological and geological interactions. Geophysical Monograph 91, 115–157. Copyright 1995 by the American Geophysical
Union.
Hartmann M. (1980) Atlantis II deep geothermal brine system. Hydrographic situation in 1977 and changes since 1965. Deep Sea Res.
27A, 161–171.
Hartmann M. (1988) Atlantis II deep geothermal brine system. Chemical processes between hydrothermal brines and Red Sea deep water.
Mar. Geol. 64, 157–177.
Hartmann M., Scholten J. C., Stoffers P., and Wehner F. (1998a)
Hydrographic structure of brine-filled deeps in the Red Sea: New
results from the Shaban, Kebrit, Atlantis II and Discovery Deep.
Mar. Geol. 144, 311–330.
Hartmann M., Scholten J. C., Stoffers P. (1998b) Hydrographic structure of brine-filled deeps in the Red Sea: Correction of Atlantis II
temperatures. Mar. Geol. 144, 331–332.
Hofman P., Schwark L., Brachert T., Badaut D., Rivière M., and Purcer
B. H. (1998) Sedimentation, organic geochemitry and diagenesis of
cores from the axail zone of the southern Red Sea: Relationships to
rift dynamics and climate. In Sedimentation and tectonics of rift
basins: Red Sea-Gulf of Aden. (eds. B. H. Prucer and D. W. J.
Bosence). chap. G2, pp. 479 –504. Chapman et Hall.
Honnorez J., Laverne C., Hubberten H., Emmermann R., and Muehlenbachs K. (1983) Alteration processes in layer 2 basalts from DSDP
Hole 504B, Costa Rica Rift. Init. Rep. DSDP 69, 509 –546.
Huppert E. H. and Turner J. S. (1981) Double-diffusive convection.
J. Fluid Mech. 106, 299 –329.
Issar A., Rosenthal E., Eckstein Y., and Bogoch R. (1971) Formation
waters, hot springs and mineralization phenomena along the eastern
shore of the Gulf of Suez. Bull. Int. Ass. Sci. Hydrol. 16, 25– 44.
Jedwab J., Blanc G., and Boulègue J. (1989) Vanadiferous minerals
from the Néréus Deep, Red Sea. Terra Nova 1, 188 –194.
Jorgensen B. B. (1983) Processes at the sediment-water interface. In
The major biogeochemical cycles and their interactions (eds. B.
Bolin and R. B. Cook), p. 477–509.
Lawrence J. R. (1974) Stable oxygen and carbon isotope variations in
the pore waters, carbonates and silicates, sites 225 and 228, Red Sea.
Init. Rep. DSDP. 23, 939 –942.
1274
M. C. Pierret et al.
Le Quentrec M. F. and Sichler B. (1991) 3-D inversion of deep tow
magnetic data on the Atlantis II Deep (Red Sea): Hydrothermal and
geodynamic interpretation. Tectonophys. 198, 421– 439.
Luther III G. W., Catalano G., De Lange G. J., and Woittiez J. R. W.
(1990) Reduced sulfur in the hypersaline anoxic basins of the Mediterranean Sea. Mar. Chem. 31, 137–152.
Manheim F. T. and Siems D. E. (1974) Chemical analyses of Red Sea
sediments. Init. Rep. D.S.D.P. 23, 923–938.
Matthews D. J. (1939) Tables of velocity of sound in pure water and
seawater in echo sounding and sound ranging. Hodrogr. Dep., Admiralty, London.
McDougall T. J. (1984a) Fluid dynamic implications for massive
sulfide deposits of hot saline fluid flowing into a submarine depression from below. Deep-Sea Res. 31, 145–170.
McDougall T. J. (1984b) Convective processes caused by a dense, hot
saline source flowing into a submarine depression from above.
Deep-Sea Res. 31, 1287–1309.
McKibben M. A., Andes J. P., and Williams A. E. (1988) Active ore
formation at a brine interface in metamorphosed deltaic lacustrine
sediments: The Salton Sea geothermal system, California. Econ.
Geol. 83, 511–523.
McKibben M. A. and Williams A. E. (1989) Metal speciation and
solubility in saline hydrothermal fluids: An empirical approach based
on geothermal brine data. Econ. Geol. 84, 1996 –2007.
Miller A. R. (1964) Highest salinity in the world ocean? Nature 203,
590.
Miller A. R., Densmore C. D., Degens E. T., Hathamay J. C., Manheim
F. T., McFarlin P. F., Pocklington H., and Jokela A. (1966) Hot
brines and recent iron deposits in deeps of the Red Sea. Geochim.
Cosmochim. Acta 42, 1103–1115.
Monnin C. and Ramboz C. (1996) The anhydrite saturation index of the
ponded brines and sediment pore waters of the Red Sea deeps. Chem.
Geol. 127, 141–159.
Mustafa H. E. Z., Nawab Z., Horn R., and Le Lann F. (1984) Economic
interest of hydrothermal deposits. The Atlantis II project. Proceedings of 2nd international seminar on offshore mineral resources.
Brest. 1984, 509 –539.
Nawab Z. (1984) Red Sea mining, a new area. Deep Sea Res. 31,
813– 824.
O’Neil J. R. and Kharaka Y. K. (1976) Hydrogen and oxygen isotope
exchange reactions between clay minerals and water. Geochim.
Cosmochim. Acta 40, 241–246.
O’Neil J. R., Clayton R. N. and Mayeda T. K. (1969) Oxygen isotope
fractionation in divalent metal carbonates. J. Chem. Phys. 51, 5547–
5558.
Palaban R. and Pitzer K. S. (1987) Thermodynamics of concentrated
electrolyte mixtures and the prediction of mineral solubilities to high
temperature for mixtures in the system Na-K-Mg-Cl-SO4-OH-H2O.
Geochim. Cosmochim. Acta 51, 2429 –2443.
Pautot G. (1983) Les fosses de la mer Rouge: Approche géomorphologique d’un stade initial d’ouverture océanique réalisée à l’aide
du Seabeam. Oceanologica Acta 6, 235–244.
Phillips O. M. (1966) On turbulent convective currents and circulation
of the Red Sea. Deep Sea Res. 13, 1148 –1160.
Pierret M. C. (1998) Les saumures et les sédiments de sept fosses de
l’axe de la Mer Rouge (19 –23°N). Etude minéralogique,
géochimique et isotopique. Rôle de l’hydrothermalisme. Thèse de
Doctorat. Université Louis Pasteur, Strasbourg. 383 p.
Pierret M. C., Blanc G., and Clauer N. (2000). Sur l’origine de la pyrite
framboı̈dale dans les sédiments de la fosse Suakin (Mer Rouge).
Compt. Rend. l’Acad. Sci. 330, 1– 8
Pitzer K. T. (1979) Theory: Ion interaction approach. In Activity coefficient in electrolyte solutions (ed. R. M. Pytkowicz.) Vol. 1, pp.157–
208. CRC Press, United States.
Pitzer K. S. (1984) Ionic fluids. J. Phys. Chem. 88, 2689 –2697.
Pottorf R. J. and Barnes H. L. (1983) Mineralogy, geochemistry and ore
genesis of hydrothermal sediments from the Atlantis II Deep, Red
Sea. Econ. Geol. Mon. 5, 198 –223.
Ramboz C., Oudin E., and Thisse Y. (1988) Geyser-type discharge in
Atlantis II deep, Red Sea: Evidence of boiling from fluid inclusions
in epigenetic anhydrite. Can. Mineral. 26, 765–786.
Redfield A. C., Ketchum B. H., and Richards F. A. (1963) The
influence of organisms on the composition of sea water. In The Sea
(ed. M. N. Hill) Chap. 2, pp. 26 –77. New York. Wiley-Interscience.
Risacher F. (1992) Géochimie des lacs salés et croûtes de sel de
l’altiplano bolivien. Sci. Géolog. 45, 135–219.
Rosenbauer R. J., Bischoff J. L., and Radtke A. S. (1983) Hydrothermal
alteration of graywacke and basalt by 4 molal NaCl. Econ. Geol. 78,
1701–1710.
Rosenthal A., Jones B. F., and Weinberger G. (1998) The chemical
evolution of Kurnub Group paleowater in the Sinai-Negev province-a mass balance approach. Appl. Geochem. 22, 1–17.
Saager P. M., Schijf J., and De Baar H. J. W. (1993) Trace-metal
distribution in sea water and anoxic brines in the eastern Mediterranean Sea. Geochim. Cosmochim. Acta 57, 1419 –1432.
Sackett W. M., Brooks J. M., Bernard B. B., Schwab C. R., Chung H.,
and Parker R. A. (1979) A carbon inventory for Orca basin brines
and sediments. Earth Planet. Sci. Lett. 44, 73– 81.
Savin S. M. and Epstein S. (1970) The oxygen and hydrogen isotope
geochemistry of clay minerals. Geochim. Cosmochim. Acta 31, 25–
42.
Schneider W., Koch G., and Lange G. (1976) Vulkanische Gläser in
sedimenten des Roten Meeres. Geolog. Jahr. D17, 107–134.
Schoell M. and Stahl W. (1972) The carbon isotopic composition and
the concentration of the dissolved inorganic carbon in the Atlantis II
deep brines (Red Sea). Earth Planet. Sci. Lett. 15, 206 –211.
Schoell M. and Hartmann M. (1973) Detailed temperature structure of
the hot brines in the Atlantis II deep area (Red Sea). Mar. Geol. 14,
1–14.
Schoell M. (1974) Valdivia VA01 Rotes Meer-Golf Von Aden. Hydrography II ⫹ III, Daten. Bundesanstalt für Bodenforschung. Hannover. 1063 pp.
Schoell M. and Faber E. (1978) New isotopic evidence for the origin of
Red Sea brine. Nature 275, 436 – 437.
Schoell M. and Hartmann M. (1978) Changing hydrothermal activity in
the Atlantis II deep geothermal system. Nature 274, 784 –785.
Scholten J. C., Stoffers P., Walters P., and Plüger W. (1991) Evidence
for episodic hydrothermal activity in the Red Sea from the composition and formation of hydrothermal sediments, Thetis Deep. Tectonophys. 190, 109 –117.
Shanks III W. C., Böhlke J. K., and Seal II R. R. (1995) Stable isotopes
in Mid-Ocean Ridge hydrothermal systems:interactions between fluids, minerals, and organisms. In Seafloor Hydrothermal Systems:
Physical, Chemical, Biological, and Geological interactions. (ed.
S. E. Humphris, R. A. Zierenberg, L. S. Mullineaux, and R. E.
Thomson) pp.194 –221. American Geophysical Union.
Smeed D. (1997) Seasonal variation of the flow in the strait of bab al
Mandab. Oceanolog. Acta 20, 773–781.
Spencer R. J., Moller N. and Weare J. H. (1990) The prediction of
mineral solubilities in natural waters: A chemical equilibrium model
for the Na-K-Ca-Mg-Cl-SO4-H2O system at temperature below
25°C. Geochim. Cosmichim. Acta 54, 575–590.
Stoffers P. and Kühn R. (1974) Red Sea evaporites: A petrographic and
geochemical study. Init. Rep. D.S.D.P. 23, 821– 847.
Supko P. R., Stoffers P., and Coplen T. B. (1974) Petrography and
geochemistry of Red Sea dolomite. Init. Rep. D.S.D.P. 23, 867– 878.
Thompson E. F. (1939) Chemical and physical investigations. The
exchange of water between the Red Sea and Gulf of Aden over the
“sill.” John Murray Expedition, 1933–34. Sci. Rep. 2, 83–103.
Thompson J. M. and Fournier R. O. (1988) Chemistry and Geothermometry of Brine Produced from the Salton Sea Scientific Drill
Hole, Imperial Valley, California. J. Geophys. Res. 13, 165–173.
Vidal P. and Clauer N. (1981) Pb and Sr isotopic systematics of some
basalts and sulfides from the East Pacific Rise at 21°N (project
RITA). Earth Planet. Sci. Lett. 55, 237–246.
Volker F., McCulloch M. T., and Alther R. (1993) Submarine basalts
from the Red Sea: New Pb, Sr and Nd isotopic data. Geophys. Res.
Lett. 20, 927–930.
Von Damm K. L., Edmond J. M., Measures C. I., and Grant B. (1985)
Chemistry of submarine hydrothermal solutions at Guyamas basin,
Gulf of California. Geochim. Cosmochim. Acta 49, 2221–2237.
Von Damm K. L. (1988) Systematics of postulated on submarine
hydrothermal solution chemistry. J. Geophys. Res. 93, 4551– 4561.
Von Damm K. L. (1995) Controls on the chemistry and tem-
Geochemical study of Red Sea brines
poral variability of seafloor hydrothermal fluids. In Seafloor
Hydrothermal Systems: Physical, Chemical, Biological,
and Geological Interactions (eds. S. E. Humphris, R. A.
Zierenberg, L. S. Mullineaux, and R. E. Thomson).
Geophysical Mongraph 91, 222–247. American Geophysical
Union.
1275
Zierenberg R. A. and Shanks III W. C. (1983) Mineralogy and geochemistry of epigenetic features in metalliferous sediment, Atlantis
II deep, Red Sea. Econ. Geol. 78, 57–72.
Zierenberg R. A. and Schanks III W. C. (1986) Isotopic contraints on
the origin of the Atlantis II, Suakin and Valdivia brines, Red Sea.
Geochim. Cosmochim. Acta 50, 2205–2214.

Download Report

Chemical and isotopic (87Sr/86Sr, 18O, D) constraints to the

Paperzz.com

Your Paperzz