Download origin of iron-rich montmorillonite from the manganese nodule belt of

Clays and Clay Minerals, Vol. 27, No. 3, pp. 185-194, 1979.
ORIGIN OF IRON-RICH MONTMORILLONITE FROM THE
MANGANESE NODULE BELT OF THE
NORTH EQUATORIAL PACIFIC
JAMES R. H E I N , 1 H S U E H - W E N Y E H , z A N D E L A I N E A L E X A N D E R 1
U.S. Geological Survey, 345 Middlefield Road, Menlo Park, California 94025
2 Hawaii Institute of Geophysics, University of Hawaii, Honolulu, Hawaii 96822
Abstract--Clay minerals in the upper 50 cm of sediment that surround the Cu- and Ni-rich manganese
nodules in the North Equatorial Pacific form two fractions: terrigenous (mostly eolian) illite, chlorite, and
kaolinite, and authigenic smectite. Smectite increases with depth in box cores from 26 to 39% and from 53
to 66% in the easternmost and westernmost areas respectively, and with distance seaward from the Americas from 26 to 53% in surface deposits. The change in the amount of smectite relative to other clay minerals
is due to dilution by terrigenous debris; smectite probably forms at a uniform rate over much of the North
Pacific deep-sea floor. The 6018 value for the smectite is +29.6%0which suggests that it formed authigenically at a temperature characteristic of the deep-sea floor. The smectite is an Fe-rich montmorillonite that
probably forms by the low-temperature chemical combination of Fe hydroxides and silica. Silica is derived
from dissolution of biogenic debris, and the Fe hydroxide is from volcanic activity at the East Pacific Rise,
4000 to 5000 km to the east. AI in the authigenic montmorillonite may be derived from the dissolution of
large amounts of biogenic silica or from river-derived A1 that is adsorbed on Fe-Mn hydroxides in the
oceans. The Fe-montmorillonite contains relatively abundant Cu, Zn, and Mn and is of possible economic
importance as a source of these and other metals.
Key Words---Authigenic Formation, Deep-Sea Clay, Manganese Nodule Belt, Montmorillonite, Pacific
Sediments, Smectite.
INTRODUCTION
GENERAL GEOLOGIC SETTING
This investigation was undertaken as part of the
DOMES (Deep Ocean Mining and Environmental
Study) project designed to gain a more complete understanding of the formation of manganese nodules in
the North Equatorial Pacific. The study concentrated
on three geographic areas each measuring two degrees
on a side, centered at 9~
151~ (Area A), 12~
138.5~ (Area B), and 15~ 126~ (Area C) (Figure
1). Areas A, B, and C are within a larger area (bounded
by 5~ to 20~ latitude and 110~ to 180~ longitude)
that is of great commercial interest for the mining of
deep-sea nodules. Area A and C are about 2500 km
apart. Preliminary results of the DOMES geologic and
geochemical studies have been reported in Open-File
Reports of the U.S. Geological Survey compiled by
Bischoff (1976) and Piper (1977).
Within each 2~area several box cores measuring 50 •
50 cm were taken at each of several stations. Between
20 and 50 cm of sediment was penetrated and recovered
at each station. Stations 46-50 are in Area A, 51-55 are
in Area B, and 11-27 are in Area C. Every 2-cm interval
of the box core collected at Stations 47, 52, 16, 18B, 20,
and 25 was sampled; other cores were sampled at less
frequent intervals.
This report discusses the distribution of smectite, its
chemical and oxygen isotopic composition, and its
mode of formation. Conclusions are also drawn on the
possible significance of authigenic smectite to the geochemical balance of the oceans.
Areas A, B, and C are located between the Clarion
and Clipperton fracture zones, in abyssal-hill topography (Figure 1). Sediment overlying basaltic basement
is 150 to 200 m thick. The upper 50 cm of sediment sampled by DOMES consists of detrital and authigenic clay
minerals, quartz, feldspar, siliceous microfossils, locally abundant silicic volcanic glass, Fe and Mn hydroxides and oxides, microscopic Mn nodules, and
minor apatite (fish debris), zeolites, and opaque minerals (Hein et al., 1979; Piper et al., 1979). Based on
radiolarian biostratigraphy, the sediment is Quaternary, but contains minor to abundant reworked Miocene and Eocene microfossils (Quinterno and Theyer,
1979). Classified according to Deep-Sea Drilling Project
(DSDP) standards, DOMES sediments range from siliceous ooze to red clay. This is consistent with their
location within a narrow belt of siliceous biogenic deposits that occur between deposits of thick calcareous
ooze at the equator and thin, but extensive red clay in
the North Pacific. Twenty to 100% of the seafloor in the
DOMES areas is covered with Ni- and Cu-rich Mn nodules.
Copyright 9 1979, The Clay Minerals Society
METHODS
Clay mineralogy was determined on the <2-/zm size
fraction with a Norelco diffractometer using C u K a radiation, a curved crystal carbon monochrometer, and
a theta-compensating slit (Hein et al., 1976). Weighting
factors used in calculating percentages are those of Bis-
185
186
160 ~
150 o
140 ~
r
*
150"o
r
120 o
i
t
II 0 o
i00 o
~X~:.,I..
c
90 ~
.~
30 ~
Hawaiian
Islands
20 ~
...............
-
Clays and Clay Minerals
Hein, Yeh, and Alexander
;[
i - ON-E- -
"D
':
cD
10~
A O
_
.CL~~
IF
"_R
.TONFRACTUREzONE
......
~I
0o
~//
I000 kilometers
alt the equator
I
]
1
]
7
L
I
I0~
Figure 1. Location of DOMES Areas A, B, and C between the Clarion and Clipperton fracture zones.
caye (1965), that is, two times the sum of the area of the
chlorite and kaolinite peaks and four times the area of
the illite peak.
In general, authigenic smectite was isolated for
chemical and isotopic analysis by obtaining finer and
finer size fractions, but due to the limited amount of
sample available, not all samples were monomineralic
(see Table 4). All samples were treated with Morgan's
solution (sodium acetate + glacial acetic acid) and
HzO2 to remove calcite, organic matter, and Mn oxyhydroxides. These chemical treatments may also have
some influence on the exchangeable cations in the
smectite. The < 1-/zm size fraction was used for chemical analysis. Forty elements were scanned by the semiquantitative emission spectroscopic method of the U.S.
Geological Survey, the reliability of which can be estimated by comparing the spectroscopic values with the
values determined by atomic absorption (AA) spectroscopy on the <0.1-/zm size fraction of sample 47-10
(Table 2). Si was not determined by the emission spectroscopic technique, which is primarily used to determine minor and trace elements.
Oxygen isotopic compositions of the <l-/xm size
fractions (<0.1 ~m for sample 47-10) were determined
by the techniques of Yeh and Savin (1977). Isotope values are reported in the 8-notation and are presented relative to the Standard Mean Ocean Water (SMOW)
standard (Table 4).
RESULTS
Distribution of clay minerals
In Areas A and B clay minerals in the <2-/zm size
fraction are dominated by smectite; illite is of secondary importance, and kaolinite + chlorite is minor (Table 1). At stations 52, 53, and 54 in Area B, illite is equal
to smectite in abundance. In contrast, illite is dominant
(40-50%) in Area C except locally in the lower part of
most cores where smectite is the dominant clay min-
Table I. Lateral and temporal distribution of clay minerals
in DOMES cores.
Depth
(cm)
Surface
0-202
>20
Area C
%
Area B
%
Area A
%
Clay
mineraP
26
48
26
25
44
31
19
37
44
19
40
42
16
31
53
16
29
56
K+ C
I
S
K + C
I
S
23
38
39
18
36
46
13
22
66
K + C
I
S
Percentages are averaged for the cores in each area.
K = Kaolinite; I = Illite; S = Smectite; C = Chlorite.
=0-15 cm used for Area C.
Vol. 27, No. 3, 1979
Iron-rich montmorillonite from the Pacific manganese nodule belt
Table 2.
Sample ~
no.
187
Chemistry of DOMES smectites.
Si
%
AI
%
Ti
%
Fe
%
K
%
Ca
%
Na
%
Mg
%
P
%
Cu
ppm
Ni
ppm
Co
ppm
Mn
ppm
47-10
47-10 z
48-19
52-42
55-53
16B-49
18A-36
18B-37
nd
21.32
nd
nd
nd
nd
nd
nd
7.5
7.45
6.6
7.1
5.9
7.6
4.0
6.2
0.5
0.39
0.5
0.5
0.5
0.5
0.5
0.7
7
6.47
7
7
7
7
10
7
nd
1.86
nd
nd
nd
nd
nd
nd
0.2
0.63
0.3
0.2
0.2
0.2
2.0
0.2
3.0
2.06
3.0
3.0
2.0
2.0
1.5
1.5
2.0
3.14
3.0
3.0
3.0
2.0
2.0
2.0
------0.3
--
1500
1100
1500
1000
1500
1000
1000
700
150
500
150
70
150
100
100
100
100
200
70
50
70
70
30
50
2000
3300
2000
1000
2000
1500
3000
2000
Sample
no.
Zn
ppm
Ba
ppm
Zr
ppm
Cr
ppm
Pb
ppm
Sr
ppm
V
ppm
Li
ppm
B
ppm
Be
ppm
La
ppm
Nb
ppm
Sc
ppm
47-10
47-102
48-19
52-42
55-53
16B-49
18A-36
18B-37
300
1300
300
300
300
300
500
300
1500
500
1000
1000
1000
700
2000
1000
300
nd
500
500
500
500
700
300
30
nd
30
50
20
50
15
50
50
nd
50
70
50
70
100
70
70
nd
100
70
70
70
300
100
150
nd
150
150
100
200
200
150
500
nd
300
200
300
200
---
200
nd
200
300
200
300
100
300
5
nd
7
7
7
7
5
7
50
nd
70
30
50
30
70
30
-nd
15
15
15
15
20
15
50
nd
50
30
50
30
20
30
Sample
no,
Sn
ppm
Y
ppm
Ga
ppm
Yb
ppm
Pr
ppm
47-10
48-19
52-42
55-53
16B-49
18A-36
18B-37
15
20
15
15
15
---
150
t50
150
100
200
200
150
50
50
50
50
50
30
50
10
15
10
10
10
15
10
50
50
50
50
50
50
50
1 See Table 5 for purity of smectite analyzed.
2 Ignition loss is 18.43% at 1000~ <0,1-p.m size fraction analyzed by atomic absorption spectroscopy, all other samples
are the < 1.0-/zm fraction analyzed by emission spectroscopy.
nd, means not determined; - - , means below limit of detection.
Other elements scanned that were below the limits of detection in all samples are: Cd (15), Mo (5), TI (7), Ce (100), Nd
(50), Ag (1.5), Au (15), As (200), Bi (15), Pt (10), Sb (30), W (20), Te (700), U (300), Hf (100), Sm (100), Eu (2), Re (15), Ta
(50), Th (300), Ge (15), Pd (2), In (3); detection limits in brackets in ppm. Detection limits of Li and Sn are 200 and 5 ppm,
respectively.
eral. Similarly, at Stations 47 and 48 in A r e a A and at
Station 55 in A r e a B, smectite increases with depth in
the cores (Table 1); it makes up 90% of the clay minerals
in the l o w e r m o s t deposits at Station 48. At other stations in Areas A and B, the clay mineralogy does not
change with depth. Smectite not only increases with
depth below the seafloor but also increases in abundance with greater distance from the Americas, that is,
from A r e a C to Area A (Table 1). Smectite can, however, be locally abundant in any of the three areas.
Smectite chemistry and mineralogy
E x c e p t for sample 18A, little variation in the chemistry of the < 1-/zm size fractions was o b s e r v e d (Table
2). Sample 18A contains less AI and more Fe and Ca
than the other smectites. In all smectites, Cu, Ni, Mn,
and possibly Zn are relatively abundant. The Zn and Ba
concentrations for sample 47-10 obtained by A A and
emission spectroscopy are too different to be due simply to the difference in grain size of the samples and
suggest an analytical error for these two elements in the
emission spectroscopy technique (Table 2).
The structural formula for smectite (sample 47-10;
Table 3) was derived by the technique of Ross and Hendricks (1945). Si makes up most of the tetrahedral layer
(Si/A1 - 6.8), and AI, Fe, and Mg are primary constituents of the octahedral layer. All Fe was arbitrarily
considered as Fe a§ Other trace elements (Table 2),
such as Cu, Zn, Co, Mn, Pb, Ni, Cd, and Cr, can be
structurally bound in octahedral coordination or can be
present as exchangeable cations in clays (Counts et al.,
1973; Pronina and Varentsov, 1973; McBride, 1976;
K o p p e l m a n and Dillard, 1977). The exchangeable cations, h o w e v e r , are mainly Na, with some K and Ca
(Table 3). Because sample 47-10 may contain up to 10%
illite interlayers, or may contain minor illite contami-
188
Table 3. Atomic properties calculated for DOMES sample
47-10.
No. of ions
Tetrahedral
Octahedral
Interlayer
Clays and Clay Minerals
Hein, Yeh, and Alexander
Si
AI
Total
A1
Mg
Fe 3+
Ti
Mn
Zn
Cu
Ni
Total
3.484
0.516
4.000
0.7506
0.5921
0.5324
0.0367
0.0275
0.0090
0.0078
0.0037
1.9598
Na
K
Ca
Total
0.4085
0.2203
0.1441
0.7729
Octahedral + tetrahedral excess charge is -1.24.
(Ca.14K.22Na.41) (Ni.004Cu.01Zn.olMn.03
Ti.o4Mg.5~Fe3+53AI.Ts)(A1.52Si3.4s)
(O~o(OH)z)
Table 4. Oxygen isotopic composition of DOMES smectite,
split of same samples analyzed for chemistry.
Size
Sample
no.
Depth in
core (cm)
47-10
47-10
48-19
52-42
55-53
16B-49
16B-49
18A-36
18B-37
39 to 42
39 to 42
27 to 31
43 to 47
26 to 30
16 to 20
16 to 20
17 to 20
13 to 17
Smectite ~
s
(#.m)
80 TM
Smectite
(%)
d(001)
(.~)
<1.0
<0.1
<1.0
<1.0
<1.0
<1.0
>1.0
<1.0
<1.0
+27.6
+29.6
+26.9
+24.7
+28.3
+24.9
+20.4
+26.4
+22.7
95 to 100
98 to 100
98 to 100
65 to 75
94 to 100
65 to 75
-10
85 to 95
-46
16.9 S
16.7 MS
16.9 S
16.8 M
16.9 S
16.7 B
-17.7 VB
16.9 M
1 Percentage of the respective size fractions estimated
qualitatively from X-ray powder diffractograms.
2 Glycolatedsamples;S =sharp, M = moderate, B = broad.
16B which is a mixture of terrigenous, biogenic and authigenic material. In general, the variation in isotopic
compositions can be attributed to changes in the relative proportion of authigenic smectite to the detrital
minerals such as illite, kaolinite, quartz, and so on (Table 4); detrital minerals lower the ~O is values.
DISCUSSION
nation, some of the K may be present in illite rather
than as an exchangeable cation in the smectite. The
excess tetrahedral and octahedral charge (Table 3) is
large but is in the range for smectites (Weaver and Pollard, 1975). Smectites from the deep sea commonly
have larger layer charges (>0.8) than do those formed
on continents (e.g., Bischoff, 1972; Weaver and Pollard, 1975; Scheidegger and Stakes, 1977). The reason
for this difference is unknown; marine smectites may
be interlayered with amorphous oxyhydroxides, a condition that would change the structural formula and layer charge. Alternatively, a large number of minor elements (Table 2) that are not included in the layer charge
calculation may occupy the octahedral layer.
The Si, AI, Mg, and Fe atomic proportions show that
sample 47-10 is an Fe-rich montomrillonite. The (060)
X-ray spacing of most samples is very weak and broad
(from about 1.508 to 1.528/~) and cannot be used to
support the identification of a dioctahedral smectite
structure. The (001) peaks are, in general, sharp, of high
amplitude, and occur near 16.9 /~ when glycolated,
whereas the peak for sample 18A is very broad, weak,
and centered near 17.7/~ (Table 4). Higher order X-ray
spacings are very weak or absent.
Oxygen isotopic composition of smectite
8018 values range from +22.7 to +28.3 for the < l wm size fractions (Table 4). The isotopically heaviest
measurement (+29.6), however, is from the <0.1-/xm
fraction (pure smectite) of sample 47-10, and the lightest value (+ 20.4) is from the > 1-/~m fraction of sample
Abyssal sediment may be composed of terrigenous,
biogenic, volcanic, and authigenic debris. Also found
in abyssal sediment but not discussed much in the literature are amorphous Fe, AI, Mn, and Si hydroxides.
Amorphous debris enters the oceans from rivers (Si,
AI, Fe), from hydrothermal activity at oceanic spreading centers (Fe, Mn, Si), and from low-temperature
weathering of oceanic basaltic basement (Fe, Mn, Si,
A1). Most of these components are found in DOMES
cores, and the distribution and origin of the terrigenous,
volcanic, and biogenic debris is discussed by Hein et
al. (1979). The authigenic fraction consists of smectite
(discussed here), Mn nodules and micronodules, zeolites (phillipsite and clinoptilolite), and possibly apatite.
Source of clay minerals
DOMES illite, chlorite, and kaolinite are terrigenous,
probably mostly eolian in origin, whereas nearly all the
smectite is authigenic. K-Ar dating of illite isolated
from Holocene abyssal deposits shows that the illite is
continental and formed several hundred million years
ago (Hurley et al., 1959). It has also been shown that
illite abundance in the Pacific correlates with that of
quartz and that both correlate with the atmospheric circulation patterns. This implies that illite and quartz are
probably of mostly eolian origin (Heath, 1969; Windora, 1969). Dust collected from the atmosphere is rich
in quartz and illite (Windom, 1969). Oxygen and hydrogen isotopic compositions of deep-sea illite also
support a continental origin (Savin and Epstein, 1970b).
Illite probably will not form from smectite on the sea-
Vol. 27, No. 3, 1979
Iron-rich montmorillonite from the Pacific manganese nodule belt
floor as it does in thick sedimentary sequences on land,
even if the reaction is favored thermodynamically, because the reaction rate is too slow (Eberl and Hower,
1976). Griffin and Goldberg (1963) concluded that chlorite and kaolinite distributions in the Pacific indicated
their derivation from the continents. Therefore, the illite, chlorite, and kaolinite from DOMES cores are likely derived from the continents and are transported primarily by winds.
In contrast, abyssal smectite is formed in the marine
environment. Smectite is not an important part of atmospheric dust (Windom, 1969), nor is there a major
continental source for the Fe-rich smectite found in the
deep sea (Aoki et al., 1974). Griffin and Goldberg (1963)
suggested that smectite in the South Pacific formed
from the in situ alteration of volcanic debris. Apparently, only half of the smectite in South Atlantic deposits and a third of that in Indian Ocean sediment is
derived from the continents (Chester and Stoner, 1974).
Recently, Heath and Dymond (1977) found that smectite in the Bauer Basin (west of South America) formed
in the marine environment. By plotting the percentage
of smectite in DOMES samples versus the 8018 values,
it can be shown that the 8018 values of the <0. 1-/zm
fraction for each sample are about the same (+29.6%0
as that measured in sample 47-10 (Table 4). Thus,
DOMES smectite must have also formed in a deep-sea
environment (see below). Detrital smectites have 80 TM
values less than +22%o (Savin and Epstein, 1970a,
1970b; Yeh and Eslinger, 1979), whereas, the 8018 endmember value for authigenic smectite formed at 0~ in
the deep sea is +31%o (Yeh and Savin, 1977). Hence,
in contrast to the speculation of Hein and Jones (1977),
the isotopic data indicate that essentially all the smectite in the < 1-~m size fraction is authigenic, with terrigenous smectite making up less than about 5%. Smectite in the > 1-/zm fraction is likely to have had, but is
not limited to, an authigenic origin.
Formation o f D O M E S smectite
Four general mechanisms have been postulated for
the formation of authigenic smectite in the deep sea.
First, smectite may precipitate directly from hydrothermal fluids; for example, nontronite in the Red Sea
(Bischoff, 1972). Such smectite is rich in Fe and low in
AI and Mg (Table 5). Second, authigenic smectite may
precipitate from solutions at low temperatures into vesicles and fractures of basalt at oceanic spreading centers (Seyfried et al., 1976). Elements for this smectite
were derived from seawater or were leached from basalt by seawater. These smectites are rich in Fe and Mg
and have moderate AI contents (Table 5). Third, authigenic smectite may form by the alteration of volcanic
rock fragments and glass in the marine environment.
This process produces moderate to high Fe and Mg
contents and higher Al-smectite than the first two
mechanisms and is the most commonly reported mech-
189
anism for the formation of smectite in basal marine deposits, in deep-sea ash beds, and in altered oceanic basaltic crust (Scheidegger and Stakes, 1977; VaUier and
Kidd, 1977; Hein and SchoU, 1978). Fourth, authigenic
smectite may form from the low-temperature combination of Fe oxyhydroxide and silica (Heath and Dymond, 1977). This reaction produces high-Fe, low-Al,
and probably low-Mg smectite when the clay forms
near where Fe oxyhydroxide is generated, and moderate- to high-Fe, moderate-A1, and low- to moderateMg clays at great distances from the Fe oxyhydroxide
source (Table 5; see discussion below).
A hydrothermal origin for DOMES smectite can be
eliminated because of the oxygen isotopic compositions. Using the smectite-water equilibrium equation of
Yeh and Savin (1977) and assuming 8018 of 0.0 for seawater, a maximum temperature of formation for
DOMES smectite is 7.0~ If as little as 2% terrigenous
quartz, chlorite, and illite contaminate the sample, the
temperature of formation would be near 0.0~ essentially the same as bottom water temperatures in
DOMES areas (0.8~176
(Gordon and Gerard, 1970).
Bischoff and Rosenbauer (1977) suggested that there is
a large component of metalliferous sediment at stations
18A and 18B that was produced by hydrothermal activity. The smectite examined at station 18B is from a
layer that occurs above the metalliferous sediment;
however, the smectite from station 18A included metalliferous deposits, and the oxygen isotope data indicate a low temperature of formation. A low-temperature, vesicle-filling origin for the smectite can be
eliminated because DOMES sites are not at an oceanic
spreading center.
The bulk chemistry of DOMES smectites does not
preclude their formation from the alteration of basaltic
volcanic debris, however, analyses of all 2-cm intervals
of many cores show little volcanic material. Also, the
relatively high percentage of some trace elements (Cu,
Zn, Mn, Co) is inconsistent with altered volcanic debris
as a source of the smectite. Petrographically, light-colored andesitic volcanic glass shards in DOMES cores
are very fresh (Hein et al., 1979) and possibly were
transported from Central and South America by winds.
Palagonite occurs in trace amounts in some cores and
is a significant component at only two DOMES stations-10 and 15--where it has been reworked from
nearby basement outcrops (Cook et al., 1977). Thus,
the formation of DOMES smectite by alteration of volcanic debris in the near-surface sediment is not probable (see further discussion on this point below).
The oxygen-isotope ratios suggest that the smectite
formed in the marine environment at temperatures typical of the deep-sea floor. Therefore, it is possible that
smectite formed by alteration of volcanic debris where
such debris is abundant and was then transported to the
DOMES sites by bottom currents. Basaltic volcanic
sediment is abundant near mid-Pacific volcanic island
Hein, Yeh, and Alexander
190
Table 5.
Reference
Clays and Clay Minerals
Elemental ratios for smectites that formed in the deep sea.
Ti/AI
Si/AI
Mg/AI
Ca/AI
Origin ~
0.87
2.50
0.05
0.13
2.86
--
0.42
0.50
0.08
0.50
4
~0.45
1.07
0.13
0.29
---
---
---
3
-20.00
--
15.33
0.64
0.32
1
16.10
--
18.93
--
--
4
0.76
2.05
2.61
0.56
0.03
0.03
0.30
0.38
3.13
3.46
3.55
2.34
0.64
1.84
0.42
1.40
0.04
0.07
0.07
0. l0
3
Drever (1976)
0.73
0.76 to 1.50
---
2.79
2.96 to 3.76
0.24
0.26 to 0.41
0.09
0.15 to 0.50
5
3
Scheidegger and Stakes (1977)
2.95
2.72
1.96 to 3.12
0.06
0.08
0.08
5.93
6.67
4.62 to 5.84
2.56
3.59
1.41 to 2.10
0.42
0.41
0.28 to 0.42
3
Banks (1972)
3.22
0.07
7.35
6.91
0.16
3
Seyfried et al. (1976)
4.23
0.05
7.77
3.80
0.06
This work,
sample 47-10
sample 18A
Hein and Scholl (1978)
Bischoff (1972)
Heath and Dymond (1977)
Melson and Thompson (1973)
Kastner (1976)
Fe/AI
-0.50
1.70
Aoki et al. (1974)
1.82 to 2.97
.
.
.
.
0.08 to 0.25
.
.
5.61 to 6.93
.
.
0.59 to 0.68
2
3
0.02 to 0.18
4
1 Origin of smectite as described in text: (1) hydrothermal precipitation; (2) low temperature precipitation at spreading
centers; (3) alteration of volcanic debris; (4) chemical combination of Fe oxyhydroxides and biogenic silica; (5) detrital
smectite for comparison.
chains; however, the major deep-sea circulation is from
south to north, with water-flow velocities of about 0.050.1 cm/sec (Knauss, 1962). Antarctic bottom water may
also enter the Equatorial East Pacific through passes in
the Line Island Ridge (Reid, 1969). These current directions suggest that deep-sea sediment surrounding
the north-south-trending volcanic island chains north
and southwest of the D O M E S sites are not likely
sources for DOMES smectite.
If the extensive DOMES clay deposits were formed
by low-temperature reaction of Fe oxyhydroxide and
silica in a deep ocean-floor environment, the source of
the Si, Al, and Fe must be discussed. Heath and Dymond (1977) proposed that for smectite in the Bauer
Basin, Fe in the form of Fe hydroxide or oxyhydroxide
was derived from hydrothermal activity at the East Pacific Rise (EPR) about 650 km to the west. Si was derived by the dissolution of siliceous microfossils.
We propose a similar origin for the Fe and Si in
DOMES samples. Because there are no local sources
for Fe, colloidal Fe oxyhydroxides must have been
transported 4000-5000 km from the EPR source. Volcanic cones are scarce in this part of the Pacific. The
Line and Hawaiian Islands and the Clarion fracture
zone, 500-800 km from D O M E S Area A, may have
provided some Fe, but only very minor amounts compared to that supplied by the EPR. The EPR is a welldocumented source for Fe and Mn hydroxides that are
apparently widely distributed adjacent to the ridge
(Bostrom, 1970; Lyle, 1976). The present study requires that volcanic Fe and Mn hydroxides produced
at the EPR are distributed for many thousands of kilometers. It is possible that oceanic spreading centers
(especially where spreading rates are high) provide
abundant colloidal Fe-Mn compounds that are distributed throughout the deep sea and become major sediment builders in areas where terrigenous and biogenic
sedimentation is minor.
Biogenic silica (5-50%) in DOMES cores decreases
in abundance with depth (Hein et al., 1979). Spicules
become progressively more abundant relative to other
siliceous microfossils with depth, suggesting that
wholesale dissolution of the more soluble, delicate
forms has taken place. Thus, biogenic silica was likely
available to combine with the Fe hydroxide derived
from the spreading centers to produce smectite. These
deposits probably cover extensive areas of the seafloor.
It is not clear at this time in what proportions the extensive smectite deposits in the South Pacific (Peterson
and Griffin, 1964) are due to the alteration of locally
derived basaltic volcanic debris versus the combination
of Fe hydroxide and silica. The problem is under investigation.
The origin of Al is not as easily understood, although
there are three possible sources. First, Al bound in or
adsorbed on biogenic silica will be released upon dis-
Voi. 27, No. 3, 1979
Iron-rich montmorillonite from the Pacific manganese nodule belt
solution. Second, A1 entering the oceans from rivers
may be deposited on the seafloor in the form of amorphous A1 hydroxides or it may be adsorbed on other
particles. Third, Al may be released during the alteration of minute volcanic dust particles in the sediment.
Hydes (1977) and MacKenzie et al. (1978) suggested
that Al is removed from seawater by living diatoms, and
Hurd (1973) believed that Al is adsorbed on frustules
of dead diatoms. In fact, it has been proposed
that smectite forms in the frustuies of living
diatoms (Van Bennekom and Van der Gaast,
1976). Sayles and Bischoff (1973) reported the occurrence of smectite pseudomorphs after radiolaria, and
Johnson (1976) proposed that authigenic smectite
formed in siliceous deposits from the silica released by
dissolution. F. T. MacKenzie and co-workers (personal
communication, 1978, Northwestern University) found
that A1 covaries with Si in pore waters of sediment containing biogenic silica. Their data suggest a release of
Al and Si by dissolution of diatoms and later uptake of
these elements by formation of an aluminosilicate. It is
likely that dissolution of siliceous biogenic debris provided some or perhaps most of the A1 for formation of
smectite, although other sources are also required.
Some of the Al may have been derived from Al hydroxides supplied by rivers. A1 may either have been
adsorbed on ferric hydroxide in the water column or,
less likely, it may have been deposited uncombined as
amorphous A1 hydroxide.
Fine volcanic dust in the sediment cannot be completely ruled out as a source for Al or for the smectite
in general. However, if volcanic dust were present, it
would be andesitic or more silicic (there is no source
for abundant basaltic dust) in composition and would
probably not alter within the short time period represented by the sediment. The coarser grained andesitic
glass is petrographically very fresh. Hein and Scholl
(1978) showed that deep-sea andesitic ash layers in the
Bering Sea take 3-5 million years to alter to smectite,
but fine dust should not take so long. Further, the chemistry of smectite commonly reflects the chemistry of the
parent ash (Kastner, 1976; Hein and Scholl, 1978), and
DOMES smectite more closely resembles basaltic rather than andesitic ash. From analysis of the bulk chemical composition of DOMES deposits, Bischoff et al.
(1979) concluded that Areas A and B contain virtually
no volcanic debris and Area C contains perhaps up to
10%.
The minor and trace elements (Table 2) are derived
primarily from seawater and biogenic debris. High Cu
and Zn values relative to Ni in the smectite is similar
to that found in nodules, where Cu and Zn are associated more with Fe phases and Ni with Mn phases.
Smectites, well known for their adsorptive properties,
probably scavenge trace elements from seawater. Fe
and Mn hydroxide also scavenge metals from seawater
191
(Hem, 1977; James and MacNaughton, 1977) and may
account for some trace metals in nodules and in clays.
SUMMARY AND CONCLUSIONS
DOMES clay minerals form two fractions: detrital
illite, chlorite, and kaolinite; and authigenic smectite.
The increase in smectite relative to the other clays with
depth in some cores is due to dilution of the authigenic
smectite by terrigenous debris in the upper deposits
(Hein et al., 1979). Authigenic smectite probably forms
at a relatively uniform rate over much of the North Pacific from the combination of Fe hydroxide produced
at the East Pacific Rise, Si and A1 released from biogenic silica, and perhpas from A1 that enters the ocean
from rivers. Harder (1978) synthesized iron-rich smectite at low temperature from hydroxides and silica. He
found that reducing conditions and the presence of Mg
in solution aid in the synthesis. DOMES smectite
formed near the sediment-water interface, where Mg
is abundant. The bottom waters, however, are highly
oxygenated, and, consequently, smectite may not form
until after some burial, perhaps to 10-20 cm where conditions are more reducing. Thus, silica released from
dissolution of biogenic debris in the upper few decimeters of sediment is being used very rapidly after production, probably by chemisorption on hydroxides
(production of smectite), or is being released into the
water column.
Volcanic debris is an unlikely source for DOMES
smectite because: (1) Basaltic volcanic debris has not
been found in DOMES deposits. Even if extensive alteration of the volcanic debris occurred within the short
time period represented by DOMES cores, one should
find at least traces of the parent basalt. (2) There is no
source of abundant basaltic glass or rock fragments in
or adjacent to the area of this study. (3) If volcanic dust,
too fine grained to be detected, were present it would
be andesitic or more silicic in composition and would
alter to smectite with a different chemical composition
than that determined for DOMES smectite. (4) More
volcanic debris would have to have been present in
DOMES deposits in order: to produce the large amount
of smectite found than is typically found in deep-sea
biogenic deposits of Quaternary age. (5) The trace element content of DOMES smectite is too high for volcanic debris to have been the parent material. (6) Phillipsite, commonly produced with smectite during
alteration of basaltic debris is rare in DOMES deposits.
Clinoptilolite, also uncommon in DOMES cores, can
form from siliceous biogenic debris rather than from
alteration of volcanic debris (Hein et al., 1978).
Authigenic smectite forms over an area at least 2700
km (E-W) by 700 km (N-S). Based on a conservative
mean clay content of 30% (Hein et al., 1979) of which
30% is smectite (Table l), authigenic smectite comprises about 10% of the sediment. In some areas it
192
Hein, Yeh, and Alexander
makes up more than 50%. Thus, there may be about
20 x 10a cubic meters of Cu-, Zn-, and Mn-containing
smectite in the upper 10 cm of sediment that connects
the three DOM ES areas. These deposits may offer a
long-range resource if areas where smectite, rich in
some of the uncommon and trace metals, can be delineated. Authigenic smectite must be considered in any
evaluation of the geochemical balance of the oceans,
especially in regard to Si, Fe, AI, Cu, Zn, Mn, and other
trace metals.
ACKNOWLEDGMENTS
Our gratitude to G. Ross Heath, University of Rhode
Island, J. I. Drever, University of Wyoming, and Tracy
L. Vallier, U.S. Geological Survey, for reviewing this
manuscript. C. Robin Ross and Jeanne Henning gave
technical assistance. Kam Leong provided the A A resuits, and Robert Mays and Chris Heropoulos the
emission spectroscopic data. Support for the isotopic
work is from N S F Grant DES71-00558-A03 to S. Epstein.
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Pe3mMe---F.rIHHHCTbIe MHHepa21bl a BepXHefi 50 CM HaCTH oca~IKOB, BMeI~amIRHX 6oraTb~e Cu H Ni
MapraHueahle KOHKpeUHH B ceBepnofi 3KBaTOpHa2IbHOfi qaCTH THxoro oKeana no~Ipa3~le3Lq}OTCa Ha ~se
qbparRnu: TeppnrenHb~e (a ocuoanoM ~O.rlOBble) HJI21HT, XJIOpHT H KaO~HHnT ~ ayTHreHHblfi CMeKTHT.
Co/Iepx~anHe CMeKTHTa ynena~naaeTcn c rny6enofi a ropo6qaTblX repHax OT 26 ~Io 39% n OT 53 ~O
66% COOTBeTCTBeHHO B CaMOf[ 8OCTOqHOfi H CaMOg 3anaJIHOg 3onax a C y~a~enHeM OT AMeparancrnx
KOHTHHeHTOBB CTOpOUyoKeana OT 26 nO 53% a HOBepXHOCTHblXOTJ10>KeHltflX.FIaMenenne conepxam~a
CMeKTnTa no OTHOmeRu~O K ~lpyrHM MuHepa.qaM o6n3aHo pa3y6ox~uaauu~ TeppHrenHbtMU oca~lKaMtt.
CMeKTHT BO3MO)KHOorpa3yeTcn paanoMepHo na 6Oabtuefi cesepao~ qacTn TnxooKeaHcKoro rny6ogoBo:Iuoro :1Ha. 3naqen}te 80 ~a ~n~ CMeKTHTa +29,6%~, YTO npeano~araeT ero o6pa3oaanne ayTnrennTnqecKn npn TeMnepaType, xaparTepnofi ~ n rny6oKoao~moro aria. CMeKTHT npe~IcTaa~LaeT co6ofi
6oraTblfi Fe MOHTMOpH2121OI-IHT,KOTOpbII4BO3MO)KHOo6paayeTcn H3 UH31~OTeMuepaTypnhlX XHMHqeEKHX
coeaaHeuafi r~Ipoorncaoa F e u gpeMae3eMa. KpeMueaeM abl~ennexcn B pe3y.rlhTaTe pacTaopenHa
6HoreueTnqecKHx oca~lKoa, a FrUIpOOKHCJlhlFe o6paay~oTca a pe3y3IbTaTe ay,gannqecgofi aKTHBHOCTH
Ha BOCTOqHO-TuxooKeaHCKOMxpe6Te, 4000-5000 rM K BOCTOKy. A1 a ayTnreHnoM MOHTMOpH~OHHTe
MOr o6pa3oaaTbCa s pe3y~bTaTe pacTaopen~t~ 60~bU~HX ~O~HqeCTB 6aoreanoro rpeMne3eMa H~H
8 pe3y~lhTaTe a~lcop6u:aH AI peqnoro npoHcxo;~cj~ennn,a~Icop6apoaaHHoro a ogeanax rn~lpOOKHC~aMa
F e - Mn. Fe-MOHTMOpH2uIOHHT CO~lepTgnT OTnOCHTe.qhnO MnOrO Ca, Zn n Mn a BO3MO)KHO HMeeT
3KOHOMHqecKoe 3naqenne KaK HCTOqHHKS*THXH :lpyrnx MeTa~.rlOa.
Resiim~ Tonmineralien in den oberen 50 cm yon Sediment, welches die Cu- und Ni-reichen ManganNeste im nrrdlichen/iquatorialen Pazifik umgibt, formen zwei Fraktionen: terrigenes (meist fiolisches)
Illit, Chlorit und Kaolinit und authigenisches Smektit. Mit zunehmender Tiefe nimmt der Anteil des
Smektit in Kernen yon 26 bis 39% und von 53 bis 66% in den am weitesten 6stlich gelegenen beziehungsweise den am weitesten westlich gelegenen Gegenden zu und nimmt mit seew/irtigem Abstand von
Amerika von 26 bis 53% in Obertl~ichenablagerungen zu. Der Unterschied in der Menge des Smektiten
im Vergleich zu anderen Tonmineralien ist auf Verdiinnung mit terrigenem Schutt zur/ickzufiihren;
wahrscheinlich forint sich Smektit mit einheitlicher Geschwindigkeit fiber dem grrBten Tell der Tiefseesohle des Nord-Pazifik. Der 80 TM Weft f'fir die Smektite ist +29,6%o, was vorschl/igt, dab es
authigen geformt wurde, bei einer Temperatur, welche charakteristisch fiir die Tiefseesohle ist. Das
Smektit ist ein Fe-reiches Montmorillonit, dab sich wahrscheinlich dutch die chemische Reaktion yon
194
Hein, Yeh, and Alexander
Clays and Clay Minerals
Fe-Hydroxyden und Silika formt. Silika wird von der Aufl6sung von biogenischem Schutt abgeleitet
und das Eisenhydroxyd kommt von vulkanischer Aktivit/it in der Ost-Pazifik-H6he, 4000 bis 5000 km
6stlich. A1 im authigenischen Montmorillonit k6nnte von der Aufl6sung von grol3en Mengen von biogenischem Silika herstammen oder von vom Flu8 abgeleitetem AI, welches auf Fe-Mn Hydroxyden
im Ozean adsorbiert ist. Das Fe-Montmorillonit enth~ilt verh~iltnismS.13ig viel Cu, Zn und Mn und k6nnte
m6glicherweise wirtschaftliche Bedeutung erhalten als eine Quelle for diese Metalle.
R6sum6---Les min6raux argileux des 50 cm du dessus du s6diment entourant les nodules de manganese
riches en Cu- et en Ni dans l'Oc6an Pacifique 6quatorial Nord forment 2 fractions: l'illite, la chlorite, et
la kaolinite terrigineuses (surtout 6oliennes) et ia smectite authig6nique. La smectite dans des carottes
augmente proportionellement h la profondeur de 26 ~t 39% et de 53 h 66% dans les r6gions le plus
/~ l'est et les plus h l'ouest, respectivement, et elle augmente de 26 h 53% proportionellement /t la
distance des Am6riques darts les d6p6ts de surface. Le changement dans la quantit6 de smectite relative
aux autres min6raux argileux est dfi /~ la dilution par des d6bris terrigineux; la smectite est probablemerit form6e /~ une allure unfforme sur une grande partie du sol profond de l'Oc6an Pacifique Nord.
La valeur 8 0 t8 pour la smectite est +29,6 per rail ce qui sugg~re qu'elle est form6e authig6niquement
/~ une temp6rature caract6ristique du sol profond de l'oc6an. La smectite est une montmorillonite fiche
en Fe qui est probablement form6e par la combinaison chimique /i basse temperature d'hydroxides
F e e t de silice. La silice est d6riv6e de ia dissolution de d6bris biog6niques, et l'hydroxide Fe provient
de l'activit6 volcanique ~ l'East Pacific Rise, de 4000 ~ 5000 km A 1'est. AI dans la montmorillonite authig6nique peut ~tre d6riv6 de la dissolution de grandes quantit6s de silice biog6nique ou d'A1 d6rive de
rivi6res, adsorbe sur les hydroxides Fe-Mn dans les oc6ans. La montmorillonite-Fe contient assez
bien de Cu, Zn, et Mn et est possiblement d'importance 6conomique en tant que source de ces m6taux
et d'autres.