Journal of Oceanography
Vol. 52, pp. 705 to 715. 1996
The Chemical States of Iron in Marine Sediments by Means of
Mössbauer Spectroscopy in Combination with Chemical Leachings
S.-Y. CHEN1,2, S. AMBE 1, N. TAKEMATSU 1 and F. AMBE 1
1The
Institute of Physical and Chemical Research (RIKEN), Wako-shi, Saitama 351-01, Japan
China Sea Institute of Oceanology, Academia Sinica, Guangzhou 510301, China
2South
(Received 11 October 1995; in revised form 29 February 1996; accepted 1 April 1996)
The chemical states of iron in near-shore and deep-sea sediments were investigated
by means of 57Fe Mössbauer spectroscopy in combination with selective and nonselective chemical leachings. As far as a limited number of the sediments analyzed
are concerned, Mössbauer spectra of near-shore sediments consist of high-spin
paramagnetic ferrous (δ = 1.13 mm/s, ∆Eq = 2.65 mm/s) and paramagnetic ferric
(δ = 0.35 mm/s, ∆Eq = 0.64 mm/s) components, while those of deep-sea sediments
are composed of high-spin paramagnetic ferrous, paramagnetic ferric and magnetic ferric (δ ~ 0.4 mm/s, H ~ 510 KG) components. The Fe2+/Fe3+ ratios of deepsea sediments are much smaller than those in near-shore sediments, while the total
contents of iron in the former are much higher than those in the latter. This is
principally due to the high contents of authigenic ferric oxides in deep-sea
sediments. Further, in the aluminosilicate fraction, the Fe2+/Fe3+ ratios of deep-sea
sediments are also smaller than those of near-shore sediments. This is probably
attributed to high contents of clay minerals and authigenic aluminosilicates in
deep-sea sediments relative to near-shore ones. The magnetic components in deepsea sediments are attributable to hematite, magnetite and/or maghemite.
1. Introduction
Marine sediments are mainly divided into lithogenous, hydrogenous and biogenous components (Goldberg, 1954). The lithogenous components are further divided into terrigenous and
authigenic aluminosilicates. Terrigenous aluminosilicates are supplied to the sea through the
atmosphere as airborne dusts and by rivers as suspended matter. They are composed of
chemically weathered aluminosilicates (clay minerals) and physically weathered ones (finegrained rock-forming minerals). The lithogenous components of terrigenous origin generally
contain a few percent of iron, and the ratio of Fe2+/Fe3+ reflects redox conditions of their formation. Whether or not authigenic clay mineral formation occurs to a significant extent in the
oceans is a question of some controversy, but chlorite, illite, glauconite, chamosite, etc. are
considered to be formed in the marine environment (Elderfield, 1976). Al-poor Fe-smectite
(nontronite) occurs in hydrothermal regions (Moorby and Cronan, 1983), and zeolite such as
phillipsite is formed from volcanic debris by halmyrolysis and from biogenous silica during
diagenesis (Elderfield, 1976).
The biogenic components are composed of calcium carbonate, opal and organic matter
which are formed by phytoplankton in overlying surface water. The contents of Fe in the
biogenous components are originally low (Turekian et al., 1973; Masuzawa et al., 1988) because
the concentration of iron in surface seawater is extremely low (Martin et al., 1989). The hydrogenous fraction of deep-sea sediments consists chiefly of iron-manganese oxides dispersed
706
S.-Y. Chen et al.
as micronodules in the sediments and present as coatings on sedimentary particles. Ironmanganese oxides are precipitated directly from seawater, from interstitial water during
diagenesis and from oxygenated hydrothermal fluids, and concentrate trace elements greatly.
Mössbauer spectroscopy is useful in identifying iron-bearing minerals as well as in
determining ferrous/ferric ratios and cation sites of iron-containing minerals. Several papers
have been published concerning Mössbauer spectroscopy of marine sediments and manganese
nodules (Johnston and Glasby, 1978, 1982; Minai et al., 1981; Thijs et al., 1981; Knedler and
Glasby, 1985; Chen et al., 1992). Minai et al. (1981) found latitudinal variations of ferrous and
ferric iron in marine sediments collected along 170°W. Abundance of ferrous iron shows a
minimum near the equatorial areas and that of ferric iron the reverse tendency. In accordance with
this, Thijs et al. (1981) reported that iron in the equatorial Pacific sediments is present almost in
the ferric state while about 15% of iron in S.E. Pacific sediments is in the ferrous state. Such
phenomena were attributed that the content of hydrogenous ferric oxides is high in the central
part of the ocean, while terrigenous aluminosilicates are less abundant as the distance from the
continents increases. Hydrogenous iron oxyhydroxides in deep-sea sediments are considered to
be the same in mineralogy as those in associated ferromanganese nodules because the chemical
composition of the former is similar to that of the latter (Piper, 1988). According to Thijs et al.
(1981), iron in deep-sea nodules occurs as a mixture of goethite (α-FeOOH) and either akaganéite
(β-FeOOH), lepidocrocite (γ-FeOOH) or poorly crystallized iron oxides. However, in order to
understand the cycle of iron itself and other minor elements in marine environments, it is
necessary to investigate whether ferrous and ferric iron in the sediments are present as
aluminosilicates or as hydrogenous iron oxyhydroxides.
In this study, the states of iron in deep-sea and near-shore sediments are investigated by
means of 57Fe Mössbauer spectroscopy using selective and non-selective chemical leachings,
together with X-ray diffraction.
2. Materials and Methods
In order to dissolve the hydrogenous fraction (mainly iron-manganese oxides) in sediments,
methods of selective and non-selective chemical leachings were examined using a ferromanganese
nodule, a hydrothermal iron oxide and a geostandard of basalt (JB-1a).
The selective leaching solutions tested are a mixture of 0.2M oxalic acid and 0.2M
ammonium oxalate (oxalate buffer: pH 3.0) (Schwertmann, 1964), and a mixture of 1.0M
hydroxylamine·hydrochloride and 0.175M Na-citrate (pH 5.0) (Robbins et al., 1984). Both of the
selective leaching solutions described above are reductive. Then, a chemical method using an
oxidizing reagent is necessary to check the effect of leaching solutions on the valency of Fe in
the fraction left after leaching. For oxidative chemical leaching, mixtures of dilute hydrochloric
acids and hydrogen peroxide (1.5%) were used. The leaching power of the mixtures was
regulated by the concentration of hydrochloric acid, which was varied between 0.1M and 2.0M.
The leaching was carried out at 80°C for 30 min with gentle stirring. The ferromanganese nodule
used is from the Penrhyn Basin (13°47.40′ S, 159°28.35′ W; 5162 m in depth) and the hydrothermal iron oxide is from the Okinawa Trough (31°05.5′ N, 139°53.2′ E; about 1650 m in depth).
The samples subjected to leaching and Mössbauer experiments were two air-dried nearshore and three air-dried deep-sea sediments, which are described in Table 1. The deep-sea
sediments were collected during the GH 80-1 cruise of R/V Hakurei-Maru, conducted by the
Geological Survey of Japan (Mizuno et al., 1982).
57 Fe Mössbauer spectra were obtained at room temperature using a 1024 channel constant
The Chemical States of Iron in Marine Sediments by Means of Mössbauer Spectroscopy
707
Table 1. Description of samples and the contents of elements.
Sample
Location
Depth
(m)
Mn
(%)
Fe
(%)
Co
(ppm)
Ni
(ppm)
Cu
(ppm)
3.40
9.6
25
36
92
T3: greenish-gray argillaceous mud, off Shikoku Island
1060
0.048
3.04
32°52.74′ N
133°53.32′ E
9.4
33
32
99
250
321
414
141
70
194
410
98
89
139
317
144
Stn. 5: greenish-gray argillaceous mud, Suruga Bay
ca. 600
0.039
34°44.52′ N
138°27.05′ E
B18: dark reddish brown zeolitic clay, Penrhyn Basin
5285
1.88
12°26.44′ S
157°57.20′ W
7.08
B26: brown to dark brown siliceous ooze, Central Pacific Basin
5351
0.524
3.41
3°16.42′ N
169°40.10′ W
B32: dark brown zeolitic clay, Mid-Pacific Mountains
5292
0.495
4.94
16°10.14′ N
179°19.82′ W
Zn
(ppm)
acceleration Mössbauer spectrometer equipped with a 740 MBq 57Co source in a Rh matrix. The
experimental spectra were fitted by a number of Lorentzian peaks, using a least-squares computer
program.
The concentrations of Mn, Fe, Co, Ni, Cu and Zn were measured by atomic absorption
spectroscopy. Their total contents were determined after decomposition of samples in Teflon
vessels under high pressure using a mixture of hydrofluoric, nitric and perchloric acids.
The mineralogy of the samples was obtained by X-ray diffraction (Cu-Kα, graphite
monochromator). The assignment of smectite (montmorillonite) was based on the shift of the
peak at 15 Å to 17 Å upon glycolation. Kaolinite was distinguished from chlorite by the chemical
resistance of the 7 Å peak of kaolinite to hot 6M HCl (Biscaye, 1965; Oinuma, 1968).
3. Results and Discussion
3.1 Mineralogy
Mineralogy of the sediments studied is given in Table 2. Clay mineral reflections are more
intense in deep-sea sediments (B18, B26 and B32) than in near-shore sediments (Stn. 5 and T3),
when normalized to the 4.26 Å (quartz) and 3.19 Å (feldspars) peaks. Further, background
reflections at the low 2θ angles corresponding to lattice distances longer than 10 Å are higher in
deep-sea sediments than in near-shore sediments. Of clay minerals, illite/mica and chlorite are
common constituents of the sediments. Smectite (montmorillonite) is less abundant in the
fraction smaller than 2 µm. Kaolinite is a minor constituent in these sediments (Table 2).
Of non-clay silicate minerals, quartz and feldspars are common, abundant constituents.
708
S.-Y. Chen et al.
Table 2. Mineralogy of the <2 µm fraction and bulk of sediments from near-shore and deep-sea environments.
Sample
Stn. 5
T3
B18
B26
B32
Sediment size
<2 µm
bulk
<2 µm
bulk
<2 µm
bulk
<2 µm
bulk
<2 µm
bulk
Nonclay minerals
Clay minerals
Qtz.
Felds.
Zeol.
S
M
?
M
S
?
W
M
S
M
M
?
M
M
M
Il./Mi.
Chl.
Smect.
Kaol.
S
M
S
M
S
W
S
M
S
S
M
M
M
M
M
W
M
M
M
M
tr.
?
tr.
?
M
?
M
W
M
W
tr.
?
tr.
?
W
?
W
?
W
W
Abbreviations: Qtz., quartz; Felds., feldspars; Zeol., zeolite; Il./Mi., illite and mica; Chl., chlorite;
Smect., smectite; Kaol., kaolinite; S, strong; M, medium; W, weak reflection intensity; tr., trace;
?, presence uncertain.
However, zeolite is the major constituent of the sample B18 and the intermediate one in the
sample B32. The zeolite is assigned to phillipsite by X-ray diffraction.
3.2 Comparison of the selective leachings with the non-selective one
Figure 1 shows the results of leaching experiments of the ferromanganese nodule and
hydrothermal iron oxide samples. As described before, the oxalate buffer method and the
hydroxylamine-citrate method were developed for selective dissolution of non-crystalline
hydrated ferric oxide and for selective leaching of iron-manganese oxides in marine sediments,
respectively, without attacking aluminosilicates to any great degree. In the case of the
ferromanganese nodule (Fig. 1A), the Fe leaching power of the oxalate buffer method (OB) is
comparable to that of the non-selective chemical leaching with 0.3M HCl-1.5% H2 O2. The
hydroxylamine-citrate solution (HC) leaches considerably less Fe than the oxalate buffer,
although it leaches Mn completely. Iron in the ferromanganese nodule is not readily extracted by
dilute hydrochloric acid because manganese oxides are not dissolved in dilute hydrochloric acid
without hydrogen peroxide. For hydrothermal iron oxide (Fig. 1B), the oxalate buffer method
(OB) and the non-selective leaching with 0.2M HCl-1.5% H2O2 show almost the same power but
the hydroxylamine-citrate method (HC) is not effective in leaching Fe. Dilute hydrochloric acid
can almost completely dissolve Fe in the hydrothermal iron oxide because the content of Mn in
it is less than 3%. In Table 3, the results of the selective and non-selective chemical leachings of
Fe for a geostandard of basalt are given together with those for near-shore sediments (Stn. 5 and
T3). Generally, the extent of Fe leaching is related to the acidity or pH of the leaching solutions.
It is worthwhile to notice that the oxalate buffer (pH 3.0) considerably extracts Fe from basalt,
although the buffer was expected to dissolve amorphous iron oxyhydroxides without attacking
aluminosilicates.
The Chemical States of Iron in Marine Sediments by Means of Mössbauer Spectroscopy
709
Fig. 1. The leaching behavior of Fe and Mn from a ferromanganese nodule and hydrothermal iron oxide
by selective and non-selective chemical leachings. OB and HC denote oxalate buffer and
(hydroxylamine·hydrochloride + Na-citrate) selective leachings, respectively.
Table 3. The percentage of leached iron to the total from a basalt (JB-1a) and near-shore sediments (Stn.
5 and T3) by selective and non-selective chemical leachings.
Samples
JB-1a
Stn. 5
T3
HC (%)
OB (%)
0.1M HCl (%)
0.2M HCl (%)
1.26
0.297
0.197
24.3
19.9
10.4
28.2
22.5
20.0
32.1
30.6
23.9
OB and HC denote oxalate buffer and (hydroxylamine·hydrochloride + Na-citrate) selective
leachings, respectively. 0.1M HCl and 0.2M HCl denote (0.1M HCl + 1.5% H2O2 ) and (0.2M HCl + 1.5%
H2 O2) non-selective chemical leachings, respectively.
710
S.-Y. Chen et al.
Fig. 2. The Fe2+/Fe3+ ratios and Fe contents of deep-sea (B18, B26 and B32) and near-shore sediments
(Stn. 5 and T3) before and after selective and non-selective chemical leachings. OB and HC denote
oxalate buffer and (hydroxylamine·hydrochloride + Na-citrate) selective leachings, respectively. The
HCl concentration of 0.0 designates untreated sediments.
The Chemical States of Iron in Marine Sediments by Means of Mössbauer Spectroscopy
711
3.3 Iron contents in the lithogenous and hydrogenous fractions of deep-sea and near-shore
sediments.
Figure 2 shows the variation of Fe-content in the sediment samples by different leaching
methods. As can be seen in the figure, the chemical leaching with 0.2M HCl-1.5% H2 O2 is
generally more powerful in leaching Fe than the oxalate buffer method for marine sediments, and
the hydroxylamine-citrate method is the weakest. As described above, the oxalate buffer
dissolves almost completely the ferromanganese nodule and hydrothermal iron oxide. Therefore,
iron in the fraction left after 0.2M HCl-1.5% H2O2 treatment is present in aluminosilicates, except
for that in magnetic iron components.
Iron contents in the lithogenous (aluminosilicates) fraction of the deep-sea clays (B18 and
B32) are about 4%, and decrease to about 2% after 2.0M HCl-1.5% H2 O2 treatment although
those in the hydrogenous fraction vary with sediments. The lithogenous iron content of the
siliceous ooze (B26) is about 2.5%, and decrease to about 1% after 2.0M HCl-1.5% H2O2
treatment. The leaching behavior of Fe in the lithogenous fraction of the deep-sea siliceous ooze
is similar to that of the deep-sea clays with respect to the HCl-H2 O2 extraction. Namely, the iron
contents in the lithogenous fraction of the deep-sea siliceous ooze and clays decrease to about
half after 2.0M HCl-1.5% H2O2 treatment. The low iron content of the lithogenous fraction in the
siliceous ooze is attributed to that biogenous opal contains little iron.
Lithogenous iron contents of near-shore sediments are lower than those of deep-sea clays.
This is probably attributed to that near-shore sediments contain more fine-grained rock-forming
minerals, as discussed below. From selective and non-selective chemical leachings, hydrogenous
iron contents in the sediments from Suruga Bay (Stn. 5) and off Shikoku Island (T3) are estimated
to be at most 1% and 0.5%, respectively. However, the iron measured as hydrogenous origin is
not necessarily located in the hydrogenous fraction of the sediments, because iron in basalt is
considerably extracted by the oxalate buffer method as described above.
3.4 Mössbauer spectra
Mössbauer spectra of Stn. 5 and B18 are shown as typical examples in Figs. 3A and B,
respectively. The spectra of near-shore sediments (Stn. 5 and T3) are composed of high-spin
paramagnetic ferrous (δ = 1.13 mm/s, ∆Eq = 2.65 mm/s) and paramagnetic ferric (δ = 0.35
mm/s, ∆Eq = 0.64 mm/s) components. On the other hand, the spectra of deep-sea sediments (B18,
B26 and B32) consist of high-spin paramagnetic ferrous, paramagnetic ferric and magnetic ferric
(δ ~ 0.4 mm/s, H ~ 510 KG) components. At the top of Fig. 3, the positions of the doublets of Fe3+
and Fe2+ components, and the sextet of the magnetic component are shown. The magnetic
component becomes distinguishable after selective and HCl-H2O2 leachings, as is evident from
Fig. 3B and given in Table 4. However, the intensity of the magnetic components decreased by
a treatment with 6M HCl-1.5% H2O2 .
Fe2+/Fe3+ ratios calculated from the Mössbauer spectra of the sediments before and after
selective and non-selective chemical leachings are shown in Fig. 2, together with their Fe
contents (in the calculated ratios, the magnetic ferric components are not included). In the deepsea sediments (B18, B26 and B32), Fe2+/Fe3+ ratios after selective leachings (OB and HC:
reductive leaching) are almost equal to those after the non-selective chemical leaching (0.2 M
HCl-1.5% H2 O2: oxidative leaching). This indicates that reductive and oxidative leachings do not
influence the Fe2+/Fe3+ ratios in the remaining phases after the leachings. In the case of near-shore
sediments (Stn. 5 and T3), the Fe2+/Fe3+ ratios in the sediments after selective leachings (OB and
HC) are lower than those in the untreated sediments. The reason is not clear, but the results should
712
S.-Y. Chen et al.
Fig. 3. The Mössbauer spectra of near-shore (Stn. 5) (A) and deep-sea (B18) (B) sediments at 296 K before
and after selective and non-selective chemical leachings. OB and HC denote oxalate buffer and
(hydroxylamine·hydrochloride + Na-citrate) selective leachings, respectively.
The Chemical States of Iron in Marine Sediments by Means of Mössbauer Spectroscopy
713
Table 4. The position and areal intensity of peaks in the Mössbauer spectra of the sample B18 before and
after oxalate buffer and 0.3M HCl-1.5% H2O2 treatments.
Peaks
1
2
3
4
5
6
7
8
9
10
(mm/s)
8.364
5.111
2.393
1.570
0.667
0.044
–0.196
–0.198
–4.230
–7.858
Areal intensity (%)
untreated
oxalate buffer
0.3M HCl-1.5% H 2 O2
1.26
1.12
4.69
0.99
41.93
41.93
4.69
0.99
1.12
1.26
3.27
2.69
9.38
2.13
32.54
32.54
9.38
2.13
2.69
3.27
3.56
3.55
10.84
2.10
29.96
29.96
10.84
2.10
3.55
3.56
Peaks 5 and 6 are the doublet of Fe3+ , peaks 3 and 7 are the doublet of Fe2+, and peaks 1, 2, 4, 8, 9
and 10 are the sextet of the magnetic component.
be the reverse if the Fe2+/Fe3+ ratios in the remaining phases are affected by the reductive
leachings (OB and HC).
In the deep-sea sediments (B18, B26 and B32), Fe2+/Fe3+ ratios in the untreated sediments
are smaller than those in the sediments after selective (OB and HC) and non-selective (0.2 M HCl1.5% H2O2 ) leachings. This is attributed to the presence of large amounts of hydrogenous ferric
oxides in deep-sea sediments. On the other hand, in the near-shore sediments, there is no distinct
difference in the Fe2+/Fe3+ ratios between untreated and treated (0.2 M HCl-1.5% H2 O2) ones.
The Fe2+/Fe3+ ratios in the lithogenous fraction of deep-sea sediments increase with the
concentration of HCl used in leaching treatments. This means the preferential leaching of Fe3+,
if the various aluminosilicates in the lithogenous fraction have the same Fe2+/Fe3+ ratio. This may
not be the case. The increase of Fe2+/Fe3+ ratios is probably attributed to the more ready leaching
of iron from Fe3+-rich aluminosilicates. What kind of aluminosilicates have low Fe2+/Fe3+ ratios
are not clear but those formed in situ in oxidizing deep-sea environments are the candidates. For
example, zeolites are formed in situ from volcanic ash and biogenic silica (Elderfield, 1976). Iron
in illite is considered to be predominantly ferric (Coey, 1975; Johnston and Glasby, 1982),
although the greater proportion of illite in marine sediments is of terrigenous origin (Windom,
1976). According to Johnston and Glasby (1982), the majority of Fe2+ in the lithogenous fraction
is attributable to chlorite, with a minor contribution from smectite. Generally speaking, Fe2+/Fe3+
ratios in physically weathered aluminosilicates (fine-grained rock-forming minerals) such as
feldspars are higher than those in chemically weathered ones (clay minerals) (Minai and
Tominaga, 1982). Therefore, iron is probably leached from different aluminosilicates with HCl
treatments of different concentrations.
In the near-shore sediments, the Fe2+/Fe3+ ratios of the lithogenous fraction show a slight
increasing tendency with the concentration of HCl, and are much larger than those in deep-sea
sediments. The high Fe2+/Fe3+ ratios are attributable partly to the fact that physically weathered
714
S.-Y. Chen et al.
aluminosilicates (fine-grained rock-forming minerals) are more abundant in near-shore sediments than in deep-sea ones, and partly to the fact that the contents of authigenic aluminosilicates
in the former are lower than those in the latter.
Iron oxides and oxyhydroxides which may occur in marine environments are goethite (αFeOOH), akaganéite (β-FeOOH), lepidocrocite (γ-FeOOH), feroxyhite (δ′-FeOOH), ilumenite
(FeTiO3), hematite (α-Fe2 O3), magnetite (Fe3O4) and maghemite (γ-Fe2 O3) (Burns and Burns,
1981). Goethite (α-FeOOH) has a hyperfine field at room temperature. However, oxyhydroxides
(α-, β-, γ- and δ′-FeOOH) are of hydrogenous origin and dissolved by selective and non-selective
chemical leachings. Ilmenite has no magnetic components at room temperature (Murad and
Johnston, 1987). Therefore, the magnetic components remained after selective and non-selective
chemical leachings are probably hematite, magnetite and/or maghemite. This conclusion is not
inconsistent with the measured values for the Mössbauer parameters (δ ~ 0.4 mm/s and H ~ 510
KG). Hematite and maghemite have isomer shifts of 0.37 and 0.32 mm/s and hyperfine fields of
518 and 499 KG at 295 K, respectively. Maghemite is the completely oxidized end-member of
the magnetite-maghemite substitutional series. Magnetite has two sextet at 295 K (δ = 0.26
mm/s, H ~ 490 KG, ∆Eq = –0.02 mm/s) and (δ = 0.67 mm/s, H ~ 460 KG, ∆Eq = 0.00 mm/s)
(Murad and Johnston, 1987).
In brief, the low Fe2+/Fe3+ ratios of deep-sea sediments relative to near-shore ones are
attributed not only to the presence of hydrogenous ferric oxyhydroxides in the former but also
to the lower ratios of Fe2+/Fe3+ of the lithogenous fraction in the former than in the latter. The
magnetic components in deep-sea sediments are attributable to hematite, magnetite and/or
maghemite.
Acknowledgements
We thank Dr. A. Usui, Geological Survey of Japan, for donating a manganese nodule and
deep-sea sediments. We are likewise grateful to Dr. J. Naka, Japan Marine Science and
Technology Center, for donating a hydrothermal iron oxide. Thanks are also due to the
anonymous referees for improving the manuscript.
References
Biscaye, P. E. (1965): Mineralogy and sedimentation of recent deep-sea clay in the Atlantic Ocean and adjacent seas
and oceans. Bull. Geol. Soc. Amer., 76, 803–832.
Burns, R. G. and V. M. Burns (1981): Authigenic oxides. p. 875–914. In The Sea, Vol. 7, ed. by C. Emiliani, John
Wiley & Sons, New York.
Chen, S. Y., S. Ambe, G. L. Zhang, N. Takematsu and F. Ambe (1992): Mössbauer study on iron in marine sediments
and manganese nodules. Hyperfine Interactions, 70, 969–972.
Coey, J. M. D. (1975): Iron in a post glacial lake sediment core: a Mössbauer effect study. Geochim. Cosmochim.
Acta, 39, 401–415.
Elderfield, H. (1976): Hydrogenous materials in marine sediments; excluding manganese nodules. p. 137–215. In
Chemical Oceanography, Vol. 5, ed. by J. P. Riley and R. Chester, Academic Press, London.
Goldberg, E. D. (1954): Marine geochemistry 1. Chemical scavengers of the sea. J. Geol., 62, 249–265.
Johnston, J. H. and G. P. Glasby (1978): The secondary iron oxidehydroxide mineralogy of some deep-sea and fossil
manganese nodules: A Mössbauer and X-ray study. Geochem. J., 12, 153–164.
Johnston, J. H. and G. P. Glasby (1982): A Mössbauer spectroscopic and X-ray diffraction study of the iron
mineralogy of some sediments from the Southwestern Pacific Basin. Mar. Chem., 11, 437–448.
Knedler, K. E. and G. P. Glasby (1985): Geochemistry and mineralogy of marine sediments from the Yasawa Trough
and Braemar Ridge, northwest of Viti Levu, Fiji. N.Z. J. Geol. Geophys., 28, 267–282.
Martin, J. H., R. M. Gordon, S. Fitzwater and W. W. Broenkow (1989): VERTEX: phytoplankton/iron studies in
the Gulf of Alaska. Deep-Sea Res., 36, 649–680.
The Chemical States of Iron in Marine Sediments by Means of Mössbauer Spectroscopy
715
Masuzawa, T., M. Koyama and M. Terazaki (1988): A regularity in trace element contents of marine zooplankton
species. Mar. Biol., 97, 587–591.
Minai, Y. and T. Tominaga (1982): Mössbauer analysis of iron (II) and iron (III) in geological reference materials.
Int. J. Appl. Radiat. Isot., 33, 513–515.
Minai, Y., T. Furuta, K. Kobayashi and T. Tominaga (1981): A Mössbauer study of deep sea sediments. Radiochem.
Radioanal. Letters, 48, 165–174.
Mizuno, A., S. Nakao, M. Joshima, O. Matsubayashi, Y. Okuda, K. Onodera, T. Saito, K. Tsurusaki and A. Usui
(1982): The GH80-1 cruise: general remarks. p. 1–35. In Regional Data of Marine Geology, Geophysics, and
Manganese Nodules: the Wake-Tahiti Transect in the Central Pacific, January–March 1980 (GH80-1 Cruise),
Cruise Report No. 18, ed. by A. Mizuno and S. Nakao, Geological Survey of Japan, Tsukuba-shi.
Moorby, S. A. and D. S. Cronan (1983): The geochemistry of hydrothermal and pelagic sediments from the
Galapagos Hydrothermal Mounds Field, D.S.D.P. Leg 70. Mineral. Mag., 47, 291–300.
Murad, E. and J. H. Johnston (1987): Iron oxides and oxyhydroxides. p. 507–582. In Mössbauer Spectroscopy Applied
to Inorganic Chemistry, ed. by G. J. Long, Plenum Press, New York and London.
Oinuma, K. (1968): Method of quantitative estimation of clay minerals in sediments by X-ray diffraction analysis.
Jour. Toyo Univ., General Education (Nat. Sci.), No. 10, 1–15.
Piper, D. Z. (1988): The metal oxide fraction of pelagic sediment in the equatorial North Pacific Ocean: A source
of metals in ferromanganese nodules. Geochim. Cosmochim. Acta, 52, 2127–2145.
Robbins, J. M., M. Lyle and G. R. Heath (1984): A sequential extraction procedure for partitioning elements among
co-existing phases in marine sediments. College of Oceanography, Oregon State University, Ref. 84-3,
Corvallis, 64 pp.
Schwertmann, U. (1964): Differenzierung der Eisenoxide des Bodens durch photochemische Extraktion mit sauer
Ammoniumoxalat-Lösung. Zeitschr. Pflanzenernähr. Düng. Bodenkunde, 105, 194–202.
Thijs, A., De Roy, E. F. Vansant, G. P. Glasby and T. Thijssen (1981): Mössbauer effect studies of iron in manganese
nodules and associated marine sediments in five areas in the equatorial and S.W. Pacific. Geochem. J., 15, 25–
37.
Turekian, K. K., A. Katz and L. Chan (1973): Trace element trapping in pteropod tests. Limnol. Oceanogr., 18, 240–
249.
Windom, H. L. (1976): Lithogenous material in marine sediments. p. 103–135. In Chemical Oceanography, Vol.
5, ed. by J. P. Riley and R. Chester, Academic Press, London.