Clay Minerals (1995)30, 381-394
M I N E R A L O G Y A N D G E O C H E M I S T R Y OF D E V O N I A N
D E T R I T A L ROCKS FROM THE I B E R I A N R A N G E ( S P A I N )
B. B A U L U Z
LAZARO,
M . J . M A Y A Y O B U R I L L O , C. F E R N A N D E Z - N I E T O
AND J . M . G O N Z A L E Z L O P E Z
Departamento de Ciencias de la Tierra, Area de Cristalograffa y Mineralogfa, Universidad de Zaragoza, Pza San
Francisco s/n, 50.009 Zaragoza, Spain
(Received 9 December 1994; revised 5 June 1995)
ABSTRACT: Two profiles in Devonian marine deposits have been studied, consisting of pelites,
subgreywackes, greywackes and quartzites. Quartz and clay minerals are major components and
feldspar and calcite are minor ones. Phyllosilicates in the fine fractions are kaolinite and illite;
kaolinite has a high degree of ordering; illite is predominantly of a 1Md polytype, with low Na
content and poor crystallinity and has a phengitic composition in greywackes, whereas in pelites it is
muscovitic in composition. Both phyllosilicates may be inherited from a source area with intensive
weathering processes, although illite may also be a diagenetic phase. These mineral characteristics
indicate that the Devonian rocks did not reach the anchizone boundary in their post-depositional
evolution.
The chemical composition of pelites and subgreywackes reveals a high degree of chemical
maturity. Chondrite-normalized REE patterns indicate a higher degree of weathering of these
Devonian sediments than of Post-Archaean Australian Shales (PAAS), possibly as a consequence of
sedimentary recycling processes. The REE patterns of the Devonian rocks in addition to the high
Th/Sc, La/Sc and Th/Co ratios suggest a felsic composition of the primitive source area, probably a
K-rich granite.
GEOLOGICAL
SETTING
The Iberian Range is a mountain chain, trending
NW-SE in direction, located in the NE of Spain. It
is formed of Mesozoic and Palaeozoic sediments.
The Palaeozoic sediments consist of three units
with different structural and stratigraphic characteristics: the Herrera Unit, the Mesones Unit (both of
which belong to the Cantabrian Zone) and the
Badules Unit (a prolongation of the AsturianLeonese Zone) (Fig. 1) (Gozalo & Lifian, 1988).
The Devonian sediments crop out in the Herrera
Unit, where two profiles have been chosen as their
stratigraphic continuity and their litho- and biostratigraphy are well known. The Lower
Gedinnian-Upper Emsian and Frasnian-Famenian
periods are represented in these materials.
In the first profile, SE of Nogueras (Zaragoza),
the series is 1500 m thick and it is in sedimentary
continuity with the Silurian, representing the Lower
Devonian (Fig. 2). It consists mainly of pelites with
intercalations of greywackes, and it shows in its
lower part (Luesma Formation) some thick layers of
quartzites with ferruginous oolites and phosphorite
levels, whereas in the middle and upper parts
(Nogueras, Santa Cruz, and Mariposas Formations)
some carbonate layers appear (Cads & Gandl,
1967). The second profile studied is located in the
north of Tabuenca (Zaragoza) and is 1350 m thick,
spanning the Upper Devonian (Fig. 3). It consists
mainly of quartzites and pelites with intercalations
of greywackes, which compose the Rodanas,
Bolloncillos, Hoya and Huechaseca Formations
according to Gozalo (1990). The Middle Devonian
has not been sampled since no continuous
stratigraphic sequence is known.
The sedimentation in the basin during the
Devonian was completely marine and was nearly
95% siliciclastic. Its facies quite frequently present
vertical variation ranging from intertidal to pelagic
9 1995 The Mineralogical Society
382
B. Bauluz Lazaro et al.
N
J
I
HERRERA UNIT
MESON[S
I
UNIT
8AOULES UNIT
I PRECAMBRIAN
NOT INCLUDED
IN ANY U N I T
P U I 6 IdOR~NO
o
s
IO
19
Z Oxm,
MONTALBAN
Fie. 1. Geological setting of the Palaeozoic outcrops of the Iberian Range (Gozalo & Lifian, 1988), with the
location of the profiles studied. A: 'Nogueras' profile, B: 'Tabuenca' profile.
conditions, although they represent sedimentation in
a shallow shelf environment. Two source areas are
supposed to exist in the basin at this period: one of
them is located in the SW (Centro-Iberian Zone)
and the other one in the NE, Cantabro-Ebroian
Massif (Cads, 1988; Gozalo, 1990).
METHODS
Sixty samples from the detrital levels of each of the
two stratigraphic profiles cited have been studied.
Moreover, the fine grain size levels have been
sampled in more detail.
The mineralogical analyses of the whole samples
and their silt and clay fractions were performed by
X-ray diffraction (XRD) using a Philips PW 1710
diffractometer, Cu-Kct radiation, automatic divergence slit and diffracted-beam graphite monochromator. For semiquantitative estimation of
concentrations, the intensity factors of Schultz
(1964) and Barahona (1974) were used.
The crystallochemical parameters of the dominant phyllosilicates in the fine fractions of the
samples were determined as follows: the basal
spacings of illite were measured from the fifthorder basal reflection, on glycolated oriented
Mineralogy and geochemistry of Iberian Range Devonian rocks
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CLAY MINERALS
r
2/U-FRACTION
Fic. 2. 'Nogueras' schematic profile (Carls & Gandl, 1967) with mineralogy of average sample and clay fraction.
U.E. = Upper Emsian, Nog.Fm, = Nogueras Formation, M.Fm.= Mariposas Formation.
aggregates, using quartz as an internal standard and
a program which fixes the reflection maximum by
measuring intensities at regular intervals for equal
periods of time. The experimental conditions were:
0.005~ step-size between 42 and 46~ and a 1 s
measuring time at each point. The b0 values of illite
and kaolinite were measured from 060 reflections
on unoriented samples, using Si powder as an
internal standard with the following experimental
conditions: scan speed of 2~
time constant of 1
s and paper speed of 40 mm/~
Illite crystallinity
(width at half-maximum, I.K., Kubler, 1968) was
determined at the first-order reflection using a scan
speed of 2~
time constant of 1 s and paper
speed of 50 mm/~
The percentage of 2M1
p o l y t y p e illite was m e a s u r e d a c c o r d i n g to
Maxwell & Hower (1967). The degree of disorder
in kaolinite was determined by the Hinckley index
(1963) in the silt and clay fraction, and was
checked by the Stoch index (1974).
Chemical analyses of some of the phyllosilicates
were carried out by a CAMECA SX50 electron
B. Bauluz Lazaro et al.
384
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GLOBAL SAMPLE
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I I
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I:':':'.~
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llllIilmlli
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lllIIIlllIllllIll
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FIG. 3. 'Tabuenca' schematic profile (Gozalo, 1990) with mineralogy of average sample and clay fraction.
Boll.Fm. = BollonciUos Formation, H,Fm. = Huechaseca Formation. Legend in Fig. 2.
microprobe, at the Technical Services of the
University of Oviedo (Spain). Operating conditions
were: acceleration voltage 15 kV, probe current
5 nA and electron beam diameter 5 lira.
Chemical analyses of major and trace elements
were performed at the X-ray Assay Laboratories
(Ontario, Canada) by using the following techniques: major elements and Rb, Y, Ba, Zr, Sn and Nb
were quantified by X-ray fluorescence (XRF); Cr,
Cs, Hf, W, Th and U by neutron activation analysis
(NAA); V, Li, Sr, Co, Ni, Cu, Zn, and Sc by
inductively coupled plasma (ICP), and the rareearth elements (REE) by inductively coupled
plasma mass spectrometry (ICPMS).
RESULTS
Mineralogy
The schematic profiles of the two studied series
are reported in Figs. 2 and 3 where the mineral
composition of the global sample and clay fraction
Mineralogy and geochemistry of Iberian Range Devonian rocks
385
TABLE 1. Mean values (wt%) of minerals for different lithologies. Q = quartz, F = Feldspars,
Cal = calcite, Chl = chlorite, Clay Min. = clay minerals. Standard deviation in parentheses.
Pelites
(n = 45)
Greywackes
(n = 10)
Quartzarenites
(n = 5)
Q
F
Cal
Chl
Clay Min.
16.7
(8.4)
54:4
(13.8)
95.5
(3.2)
1.4
(1.9)
-
1.2
(3~8)
9.2
(8.4)
-
3.8
(5.3)
10
(13.6)
1.5
(2.1)
70
(3.7)
23.4
(11.1)
2.5
(1.7)
-
of some representative samples of the whole of
samples studied are shown.
The results of the mineralogical study by XRD
allow the distinction between three different
lithologic groups: pelites (>40% clay minerals),
greywackes (5-40%) and quartz-arenites (<5%).
This classification is accompanied by an increase in
the grain size. The average mineralogical composition of these three lithologic groups is listed in
Table 1, where chlorite is shown separately from
the other clay minerals. The major phases are clay
minerals as a whole and quartz, whereas calcite and
feldspar constitute minor components if detected. In
this table, a negative correlation between clay
minerals and quartz is noted.
The only phytlosilicates registered in the fine
fractions of the samples are iUite and kaolinite. The
mean values of the illite/kaolinite ratio and crystallochemical parameters of illites in the fine fractions
of greywackes and pelites are reported in Table 2.
Illite is the major phase in the fine fractions of
pelites and in the clay fraction of greywackes; the
illite/kaolinite ratio increases with decreasing grain
size. In the clay fractions the illitic component
consists of mixtures of illites and ISII interstratified,
with a smectite content always <15%. This has
been determined according to Srodofi (1984).
No trends have been observed in the variation of
the crystallinity index of illites throughout the
stratigraphic profile. In pelites, the more frequent
crystallinity indices range from 0.23 to 0.25~ in
the silt fraction and from 0.41 to 0.49~ in the clay
fraction, whereas in greywackes they range from
0.22 to 0.24~ in the silt fraction and from 0,30 to
0.35~ in the clay fraction.
The bo parameter of illites (Table 2) shows a
slight variation within any of the fine fractions of
pelites and also in their respective mean values; the
latter indicate muscovitic compositions, whereas in
greywackes the ranges of variation in the fractions
are wider and the respective mean values, very
similar to each other, indicate a clearly phengitic
composition (Guidotti, 1984).
The range of variability of the basal spacings of
illite is very similar to the range mentioned for the
bo values: there is a slight variability in the
fractions of pelites and a higher one in those of
the greywackes, with similar mean values in each
kind of rock. In this case the values are slightly
higher for pelites, although all of them indicate a
very low Na content.
The mean percentage of the 2M1 polytype
(Maxwell & Hower, 1967) is low, though greater
in greywackes than in pelites, and greater in the silt
fraction compared with that of the clay fraction in
both lithologies.
Kaolinite has higher Hinckley indices than unity,
indicating that these are ordered kaolinites and,
furthermore, this parameter shows no systematic
variations throughout the stratigraphic profile. The
values of this parameter in pelites are slightly lower,
and with a lower variation range in the clay fraction
(1.68 + 0.45) than in the silt fraction (1.79 q- 0.61);
in greywackes, this index was only determined in the
silt fraction, showing slightly higher values than in
pelites (1.89 + 0.57). Moreover the Stoch index was
measured in some samples and it always had values
of <0.7, strengthening the argument that kaolinite is
an ordered phase. The bo parameter, determined only
in pelites, shows little variation, with very similar
mean values in the two fractions (8.934 -t- 0.006
in the silt fraction and 8.935 -I- 0.009 A in the clay
fraction).
Chlorite, which has not been detected in the fine
fractions, is mainly a 7 A phase, either of chamosite
or berthierine type, according to the study by XRD.
B. Bauluz Lazaro et al.
386
TABLE 2. Mean values (wt%) of illite/kaolinite ratio and crystallochemical parameters of illites for silt and clay
fractions. Silt Fr. = silt fraction, Clay Fr. = clay fraction, I1/Ka = illite/kaolinite, K.I. = Kubler Index, B.S. = basal
spacings. B.S. and bo in A, K.I. in ~ Standard deviation in parentheses.
Pelites
(n=45)
Greywackes
(n=10)
ll/Ka
K.I.
Silt fr.
bo
B.S.
%2M1
II/Ka
K.I.
Clay fr.
bo
B.S.
%2Mr
5.8
(7.9)
0.9
(1.2)
0.29
(0.14)
0.28
(0.06)
8.990
(0.009)
9.032
(0.042)
9.971
(0.009)
9.956
(0.026)
18
(10)
25
(5)
7.7
(12.3)
4.5
(5.1)
0.44
(0.09)
0.33
(0.10)
8.995
(0.009)
9.034
(0.043)
9.980
(0.009)
9.959
(0.032)
8
(10)
22
(6)
Study by petrographic microscope of the coarser
rocks has revealed that they display no homogeneous textures and have moderate sorting. The
framework consists mainly of subangular and
subrounded clasts of quartz and minor quantities
of feldspar; there are also kaolinites, white micas
and biotites, the latter being altered to chlorite;
these phyllosilicates are usually oriented parallel to
the bedding. As accessory constituents, zircon,
tourmaline and futile were noted. As minor
phases, different kinds of cements have been
observed as well, which are not homogeneous
throughout the profile: Fe-oxides (hematites),
chlorite, Fe-rich calcite, silica and kaolinite. The
Fe oxides are scattered irregularly in the rocks,
where there are even ferruginous levels, and they
can be considered as a very early cement, perhaps
synsedimentary. Chlorite cement fills the interstices
among subrounded quartz crystals, indicating that it
has probably been formed during the early
diagenetic stage, when compaction of the rocks
was not very high. This cement is frequently
associated with very Fe oxide rich zones. The
formation of authigenic quartz has been identified
at different stages, and, occasionally, pore-filling
kaolinite has been found which is possibly
diagenetic in origin.
Only quartz and kaolinite clasts were identified
in pelites due to the small size of the samples, the
kaolinite being parallel to the lamination.
Chemical analyses by electron microprobe
The results of electron microprobe analyses on
kaolinite and on chlorite crystals, the latter being
derived from altered biotite or as sandstone cement,
allow the deduction of the following mean
structural formulae:
kaolinite: Si4.06A13.920lo(OH)8(Ko.ol Nao.ol);
detrital chlorite: (Sis.77A12.E3)(A13.68Tio.olCro.olF e 6 . 1 2 M n o . o l M g 1 . 2 6 N i 0 , 0 1 ) O z o ( O H ) 1 6(Ko.02Cao.14Nao.02);
c h l o r i t e c e m e n t : (Sis.76A12.24)(A13.8zTio.19Fes.49Mg 1.o8)O2o(OH)16(Ko.o9Ca0.06Nao.os).
Geochemistry
The mean chemical analyses corresponding to
selected samples of pelites (DP), subgreywacke
pelites (DSGP) (this lithology is intermediate
between greywackes and pelites), and clay fractions
of pelites are reported in Table 3. One illite-rich
sample and other samples with similar contents of
illite and kaolinite have been differentiated among
the clay fractions in this table. Likewise, average
data of Post-Archaean Australian shales (PAAS)
(Taylor & McLennan, 1985) are included as a
reference. The mean values of some chemical ratios
of the same lithologies as in Table 3 are listed in
Table 4.
Major elements. The analysed rocks are noticeable by their high contents of the less mobile
elements, such as A1 and Ti. Compared with PAAS
shales, DP pelites contain greater amounts of
A1203, slightly more FeOt, MnO, TiO 2 and P205.
They have a similar K20 content and they are
depleted in SiO2, CaO, MgO and NazO. In
comparison with SGDP, DP pelites are enriched
in A1203 and K20, probably because of both the
scarcity of feldspars in the coarser grain rocks
(DSGP) and their lower contents in phyllosilicates.
Moreover, they have slightly higher percentages of
TiOz and MnO and lower percentages of SiO2,
CaO, MgO, NazO, P205 and particularly of FeOt
than the coarser grain size rocks.
Mineralogy and geochemistry of Iberian Range Devonian rocks
387
TABLE 3. Mean values of chemical analyses of major (wt%) and trace (in p.p.m.) elements. DP = Devonian
pelites, DSGP = Devonian subgreywacke pelites, <2 Ilm Fr = <2 Ixm fraction, IL + KA = <2 lam fraction with
similar contents in illite and kaolinite, IL90% = <2 I.tm fraction with >90% in illite. Weight percent oxides
recalculated to 100% on a volatile-free and reduced Fe basis. Standard deviation in parentheses.
SiOz
AlzO3
TiOz
FeOt
MnO
CaO
MgO
NazO
KzO
P205
Cs
Ba
Rb
Sr
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Y
Th
U
Zr
Hf
Sn
Nb
W
Cr
V
Sc
Ni
Co
Cu
Zn
Li
DP
(n = 8)
DSGP
(n = 3)
PAAS
<2 lam Fr
(n = 8)
IL + KA
(n = 6)
IL90%
(n = l)
53.62 (1.60)
31.77 (2.83)
1.34 (0.08)
7.8 (1.64)
0.21 (0.43)
0.57 (0.12)
0.76 (0.12)
0.18 (0.16)
3.54 (1.65)
0.2 (0.07)
12 (5.2)
448 (357)
129 (69)
149 (40.1)
67.4 (4.9)
128.9 (14.3)
14.9 (1.9)
55 (9.6)
10.7 (2.9)
2.3 (0.6)
8.6 (2.5)
1.04 (0.38)
5.9 (1.9)
1 (0,3)
2.9 (0.6)
0.4 (0.1)
2.7 (0.3)
0.4 (0.1)
23 (6)
17 (1)
2.98 (0.7)
160 (48)
4 (1.5)
14 (3)
24 (4)
4 (0)
122 (15.8)
159 (17.5)
23 (2.5)
61 (19)
22 (9)
44 (39.4)
59 (27.7)
196 (86.7)
62.53 (1,72)
17.13 (4.55)
1.25 (0.48)
14.03 (5.61)
0.16 (0.13)
0.87 (0.28)
1.14 (0.29)
0.21 (0.22)
2.42 (1.65)
0.25 (0.24)
5 (2.7)
401 (346)
73 (44)
141 (59)
53,8 (22)
106.9 (42)
12.3 (4)
45.2 (12.8)
8.5 (1.4)
1.56 (0,3)
7.2 (2)
0.9 (0.4)
5.2 (2.2)
0.97 (0.44)
3.1 (1.3)
0.5 (0.2)
3.2 (1.2)
0.5 (0.2)
25 (11)
17 (8)
4.7 (1.8)
533 (292)
15.3 (8.4)
13 (2)
22 (6)
3 (2)
86.3 (14)
184 (48)
17.4 (2.1)
40 (8)
22 (9)
19 (15)
122 (126)
95 (28)
62.8
18.9
1
6.5
0.11
1.3
2.2
1.2
3.7
0.16
15
650
160
200
38
80
8.9
32
5.6
1.1
4.7
0.77
4.4
1
2.9
0.4
2.8
0.43
27
14.6
3.1
210
5
4
19
2.7
110
150
16
55
23
50
85
75
49.84 (1,32)
35.55 (2.2)
1.04 (0.08)
7.26 (1.27)
0.15 (0.25)
0.74 (0.11)
0.86 (0.11)
0.26 (0.19)
4.02 (1.7)
0.26 (0.09)
14.5 (6.4)
390 (164)
148 (71.3)
164 (50.6)
68.5 (7.3)
134.4 (19)
14.7 (1.7)
54.8 (8)
10.5 (2.1)
2.17 (0.5)
8.6 (2)
1.03 (0.3)
5.97 (1.7)
1.07 (0,3)
2.99 (0,6)
0.42 (0.1)
2.84 (0.31)
0.41 (0.1)
22 (5)
19 (2)
2.29 (0.4)
106 (26)
2.26 (0.6)
13 (2)
19 (2)
3 (2)
136 (16)
170 (18)
25.2 (2.3)
61 (15)
23 (9)
60.4 (33)
99.2 (29)
217.6 (95)
49.56 (1.35)
36.59 (0.66)
1.06 (0.08)
7.55 (1.29)
0.07 (0.05)
0.7 (0.09)
0.82 (0.07)
0.19 (0.08)
3.69 (0.76)
0.25 (0.06)
12.55 (5.7)
325.8 (99.5)
113 (30.6)
169 (51.6)
69.62 (7.7)
139.7 (18.2)
15.2 (1.55)
57 (5.99)
11.2 (1.6)
2.32 (0.44)
9 (1.6)
1.1 (0.25)
6.3 (1.53)
1.12 (0.26)
3.1 (0.57)
0.43 (0.08)
2.91 (0.33)
0.42 (0.05)
22.2 (4)
19 (2.4)
2.35 (0.42)
92.8 (11.1)
1.97 (0,19)
13 (2)
17.7 (1.4)
2,7 (1.37)
141.7 (13.3)
176 (15)
26.12 (1.3)
65 (13.4)
22.3 (5.6)
60.5 (33.8)
102.2 (33.5)
265 (41.4)
49.91
34.1
1
5,71
0.002
0.79
0.94
0.26
7.12
0,15
24.5
440
284
104
60.2
111
12.1
40.9
6.4
1.25
5
0.5
3.4
0.68
2.3
0.4
2.6
0.37
16
18
1.9
137
3.3
12
21
6
130
139
24.1
37
13
29.3
91.8
65
388
B. Bauluz Lazaro et al.
TABLE4. Mean values of some chemical ratios and the CIW index (Chemical Index Weathering) (Harnois, 1988)
and the CIA index (Chemical Index of Alteration) (Nesbitt & Yong, 1982). Legend in Table 3.
K20/Na20
A1203/SIO2
AI203/Na20
Th/U
K/Cs
K/Rb
La/Sc
Th/Sc
La/Th
Ni/Co
Zr/Hf
Zr/Th
Yb/Hf
(La/Yb),
(La/Sm)n
(Gd/Yb)~
Eu/Eu*
REE
LREE
HREE
LREE/HREE
CIA
CIW
DP
DSGP
PAAS
<2 lira Fr
IL + KA
IL90%
19.7
0.59
176.5
5.7
2449
227.8
2.9
0.74
3.9
3
40
9.4
0.7
17.25
4.22
2.6
0.72
302.06
276.89
22.92
12.53
87
97
11.5
0.27
81.6
3.6
4018
275
3.09
0.98
3.16
1.8
34.8
31.3
0.21
12.21
3.94
1.86
0.62
249.9
226.7
21.6
11.3
81
92
3.08
0.3
15.75
4.7
2048
191.9
2.37
0.91
2.6
2.4
42
14.3
0.56
9.17
4.27
1.36
0.66
183
164.5
17.4
9.45
70
82
15.44
0.71
136.7
8.3
2301
225.4
2.7
0.75
3.6
2.7
46.9
5.6
1.3
16.3
4.11
2.45
0.7
308.42
282.93
23.33
12.13
86
97
16.6
0.74
192.6
8.1
2440
271
2.7
0.73
3.7
2.9
47.1
4.9
1.5
16.2
3.91
2.5
0.71
319.48
292.7
24.4
11.99
88
97
27.1
0.68
131.1
9.5
2412
208.1
2.5
0.75
3.3
2.8
41.5
7.6
0.8
15.65
5.92
1.56
0.68
247.1
230.6
15.25
15.12
79
95
The AIzO3/SiO2 ratio of DP rocks is almost twice
as high as that of PAAS and PSGD rocks; the
AlEOa/Na20 and K20/Na20 ratios of DP rocks are
also much higher, although they decrease from
DSPG to PAAS. The chemical index of alteration
(CIA) (Nesbitt & Young, 1982) and the chemical
index weathering (CIW) (Harnois, 1988) increase in
the order P A A S - D S G P - D P , the last values
corresponding with residual clays and indicating
the high degree of chemical maturity of these rocks.
The negative correlations between the pairs
AIzO3-CaO, (r = -0.68), A1203-FeOt (r = -0.80)
and AIzOz-Mg (r = -0.82) suggest that Ca, Fe and
Mg are not concentrated in dioctahedral clay
minerals but are distributed in other phases which
could be oxides, phosphates and/or trioctahedral
phyllosilicates, such as those suggested, for
instance, by the correlations P2Os-CaO (r = 0.60)
and Mg-SiO2 (r = 0.70).
The main chemical differences between pelites and
their clay fractions are enrichment in A1203, C a t ,
MgO and alkaline elements and the depletion in
Feet, Tit2 and MnO.
Trace elements. The large ion lithophile elements
Rb, Cs, Ba and Sr show a similar geochemical
behaviour, with increasing concentrations as both
the grain size of the rocks and their quartz contents
decrease. The DP pelites have lower concentrations
of these elements than PAAS. The correlations
between KzO and Rb, Cs and Ba are positive, with
correlation coefficients 0.98, 0.88 and 0.60,
respectively, suggesting that these elements are
preferentially fixed on clay minerals. Furthermore,
the decreasing trend in the K/Rb and K/Cs ratios
from DSGP to DP confirms this hypothesis. These
facts likewise reflect the relative concentration of
alkaline elements of larger ionic radius in the
weathering products, clay minerals, and, moreover,
the degree of chemical maturity noted above.
As far as the clay fractions of pelites are
concerned, Cs, Rb and Ba are selectively concentrated in the richest samples in illite.
The Th contents in the analysed samples are
15% higher than in PAAS, and the Th/U ratios
increase from DSGP to DP, which is probably
related to the higher mobility of U under the
Mineralogy and geochemistry of Iberian Range Devonian rocks
389
1000
+
Pelites
Subgreywacke
pelites
100
~
PAAS
O
10
La
I
I
!
I
I
I
I
I
I
I
I
I
I
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
FIG. 4. Chondrite-normalized REE patterns of pelites, subgreywacke pelites and PAAS. Chondrite values from
Taylor & McLennan (1985).
oxidizing conditions of weathering. Correlations
between Th and Zr, A1203 or P205 are not
significant, indicating that their distributions are
not controlled by zircon, clay minerals or
phosphates, but rather by a combination of these
minerals.
The DP pelites have higher levels of Cr, Sc and
Ni and similar contents of V and Co compared with
those of PAAS. They have a higher Cr, Sc, and Ni
content and a lower V content, if compared with
DSGP rocks. Linear correlations between A12Oa-Cr
(r = 0.93), A1203-Sc (r = 0.86) and AI203-Ni
( r = 0.61) show that these transition metals are
fixed by clay minerals, whereas the pairs V-Th (r =
0.74), U-V (r = 0.70) and V-Zr (r = 0.66) suggest
that V is mainly controlled by detrital heavy
minerals. In the clay fractions, there are also good
positive correlations between A1203 vs. Cr, Sc, Ni
and Li, suggesting that these elements are fixed
preferentially by kaolinite.
Overall REE contents are greater in DP pelites
(~,REE = 302 p.p.m.) compared with those of
DSGP (250) and those of PAAS (183). Positive
correlation between A1203 and REE (r = 0.66)
indicates that overall the REE are mainly controlled
by clay minerals. Although their respective chondrite-normalized patterns (Taylor & McLennan,
1985) show similar general trends, they present
some differences (Fig. 4, Table 4). The fractionations, expressed by (La/Yb)n and (Gd/Yb)n ratios,
are almost twice as high in DP pelites as those from
PAAS and, despite these values decreasing strongly
in the DSGP rocks, they are still higher than those
for the reference rocks (PAAS).
Light rare earth elements (LREE) fractionation,
expressed by (La/Sm)n, is similar for the three kinds
of rocks, and a negative Eu anomaly is more
pronounced in the coarser grained rocks (DSGP).
From both the evidence that the finer grained rocks
are enriched in REE and the existence of correlations
such as A1203-(La/Yb), (r = 0.72), CIA-(Gd/Yb)n
(r = 0.86), PzOs-heavy rare earth element (HREE)
(r = 0.73) and Zr-Eu/Eu* (r = -0.93), the
importance of the chemical maturity of the rocks
can be claimed as a factor controlling the degree of
fractionation, as well as the relationship between the
mineral fractionation and the contents of HREE and
the size of Eu anomaly.
390
B. Bauluz Lazaro et al.
In the clay fractions, differences in the contents
and distribution of REE between illite-rich samples
and those which consist of mixtures of notable
proportions of illite and kaolinite can be note& The
REE content in illite has been deduced from a
sample with >90% of illite and that of kaolinite
from the average of the samples which have similar
illite and kaolinite contents, taking as reference the
deduced composition for illite. Kaolinites show
higher REE contents and (La/Yb)n and (Gd/Yb)n
ratios than those of illites (Fig. 5, Table 4).
The rest of the analysed trace elements show, as
major differences, that DSGP rocks contain greater
amounts of Zr and Hf, have a greater Zr/Th ratio
and a lesser Yb/Hf ratio than DP rocks, as well as
higher percentages of P2Os. These facts suggest
preferential accumulation of detrital heavy minerals
in the coarser grained rocks. It is observed that Zr,
Hf, Nb and W are positively correlated with K20 in
the clay fractions of pelites, suggesting their
preferential accumulation in illite-rich samples.
DISCUSSION
Mineralogy
The results of the mineralogical analyses confirm
the mainly detrital character of the materials
studied. Quartz constitutes the major mineral of
the coarser rocks; the morphologic and textural
characteristics of the majority of crystals indicate
that quartz is inherited from a near source area.
Moreover, the relatively high percentages of quartz
in quartz-arenites and greywackes suggest the
possibility that the nature of the primitive source
area was acidic. In these rocks, five kinds of cement
appear as components formed from post-sedimentary processes: Fe oxides, chlorite, Fe-rich calcite,
silica and kaolinite. The Fe oxides were probably a
very early cement, almost s y n s e d i m e n t a r y .
Subsequently, when the compaction undergone by
the sediments was still low, chlorite precipitated in
zones where there was a high availability of Fe (Fe
oxides), Si, A1 and Mg (biotites, Fe-rich micas,
kaolinites). In a more advanced stage of diagenesis,
calcite was precipitated. It is not possible to
establish the timing of quartz and kaolinite
formation in relation to the rest of the cements.
These authigenic phases are scarce in both
quartzites and pelites. In the greywackes they
represent <15% of the bulk rock.
With regard to the phytlosilicates, the phases
present in greywackes and pelites are illite,
kaolinite and chlorite, although chlorite does not
appear in the fine fraction due to its coarser grain
size; essentially, chlorite is formed by alteration of
biotite, and less commonly is formed as a cement.
The latter would form only in shallow conditions
and from transformation of ferruginous oolites or
detrital silicates (Velde, 1985).
The distribution of kaolinite and its ordered
nature suggests that it is mainly an inherited phase
(with the exception of the small quantities of porefilling kaolinite), probably originating from a source
area with intense weathering and good drainage
which eased the lixiviation of the most mobile
elements.
Illite is the only other phase present in the fine
fraction of the rocks studied. The crystallinity index
and basal spacing increase when both the grain size
of the rocks and fraction under consideration
decrease, as a result of a higher alteration degree
to which they have been subjected. Of the two
parameters, crystallinity index seems to be more
sensitive, since its variation is noticeable in the two
fractions of the same kind of rock as well as in the
comparison of the fine fractions of the two kinds of
rock studied, whereas basal spacing varies significantly only when the different rocks are compared.
The slighly higher percentages of 2M1 illites in
greywackes than in pelites, and also higher in the
silt fraction than in the clay fraction, indicate the
existence of detrital illite. Nevertheless, the
dominant polytype is 1Md which might have
originated by erosion and weathering of micas
from the source area (Yoder & Eugster, 1954;
Velde, 1965) or as an intermediate product from the
alteration of K-feldspar to kaolinite (Garrels &
Howard, 1959), and in both cases 1Md illite would
be inherited. It is more likely that it was a
diagenetic phase, produced by aggradation from
other phyllosilicates such as smectites or even
kaolinites.
The other parameter studied, bo, does not usually
undergo modifications during the sedimentary
processes, as reflected by the fact that the values
are practically identical in the two fractions of each
kind of rock. Consequently, it is used as an
indicator of the source area, especially that of the
silt fraction. However, in the Devonian materials
the differences between the bo of illites from
greywackes and pelites are very significant,
a/though the range of the variation in greywackes
Mineralogy and geochemistry of Iberian Range Devonian rocks
is so wide that it includes the values found in
pelites. This might be explained by the possibility
that the source areas of greywackes and pelites
were compositionally different, with more Fe-rich
micas in the provenance of the greywackes. Other
alternatives could be to consider that the distribution of these values is affected by the intensity of
the weathering undergone by the finer materials, the
pelites.
In relation to the diagenetic or metamorphic
evolution of these sediments, the presence of
notable quantities of kaolinite in the clay mineral
assemblages suggests that the diagenesis experienced by these Devonian rocks did not reach the
deep diagenesis phase. Kubler (1968) established
that the presence of kaolinite is possible up to what
he calls diagenetic zone III and disappears in the
last diagenetic zone, zone IV. As has been pointed
out in the literature, kaolinite disappears from the
clay mineral parageneses during the burial of the
sedimentary series due to different factors: combination with Mg from destabilized dolomite (Muffler
& White, 1969); combination with other phases to
produce illite and chlorite (Boles & Frank, 1979);
or mixed-layered minerals (Hower et al., 1976), or
by combination with quartz to produce pyrophyllite
(Winkler, 1964) near 275~ at 1 - 2 kbar total
pressure.
It has been observed that the temperature and
depth at which kaolinite disappears vary from one
sequence of rocks to another. This variability of
temperature-depth conditions supports the idea of a
non-unique cause of instability (Dunoyer de
Segonzac, 1969). The temperature (T) at which
kaolinite disappears has been reported to range
from 80~ (Dunoyer de Segonzac) to 200~
(Winkler, 1964).
The variability of kaolinite stability in the A12OaSiO2-H20 system has been the aim of numerous
experimental studies. Velde & Kornprobst (1969)
indicated that the transformation of kaolinite to
pyrophyllite and andalusite occurs at T ~ 300~ and
is practically independent of pressure, but such
conditions are never reached during diagenesis.
Sediments evolve by means of interaction with
interstitial solutions, and these are usually basic,
concentrated and rich in alkaline and alkaline-earth
elements. In this environment, the presence of K§
and Mg 1§ assures the illitization and chloritization
of clay materials (Long & Neglia, 1968) generating
the illite-chlorite facies of anchizone. The increase
in the K§ concentration in the solutions and in pH
391
strongly reduces the T of the kaolinite-mica
equilibrium, perhaps to near 100~
Therefore,
during deep diagenesis, saline solutions ensure the
systematic illitization of kaolinite since these
solutions reach and exceed 100~
Conversely, the other phyllosilicate that appears
associated with kaolinite in the fine fractions of
pelites is illite, and it is predominantly of 1Md
polytype and has a mean crystallinity index value of
0.44~
corresponding with the diagenetic zone of
Kisch (1990). These parameters seem to indicate
that the degree of post-depositional evolution of
these rocks has been scarce and probably lower
than the anchizone boundary.
Geochemistry
The geochemical characteristics of the analysed
rocks as a whole reveal their high degree of
chemical maturity. High values of the AlzO3/SiO2,
A1203/NaEO and K20/NaEO ratios and of the CIA
and CIW indices in both DP and DSGP rocks,
higher than those from PAAS, indicate the higher
degree of weathering which the materials have
undergone (Nesbitt et al., 1980; Wronkiewicz &
Condie, 1983). Likewise, the contents in mobile
(Ca, Mg, Sr, Na) and immobile (Ti) elements of DP
and DSGP are lower and higher, respectively, than
those from PAAS, probably as a result of
sedimentary recycling processes (chemically
mature source area).
The lower Rb, Cs and Ba contents in the
analysed rocks, in comparison to those of PAAS,
can be explained by their clay mineralogy, with
noticeable quantities of kaolinite, in addition to the
fact that these elements are preferably fixed by
illites, as their correlations with K20 reveal. These
percentages of kaolinite constitute an additional
evidence of the intensity of the weathering.
The small mobile trace elements, such as REE,
Th, Sc, Zr, Hf and Nb with very low water-solid
partition coefficients and very short residence times
in sea water, are almost wholly transferred into
clastic sedimentary rocks during the weathering and
transportation (Taylor & McLennan, 1985), and as a
result of this are used to determine the provenance.
The chondrite-normalized REE patterns of DP and
DSGP (Fig. 4), in addition to high Th/Sc, La/Sc and
Th/Co ratios and low Fe, Mg and Ni contents, are
all characteristics that suggest a felsic composition
of the primitive source area, probably K-rich granite
(Taylor & McLennan, 1985). In comparison to
392
B. Bauluz Lazaro et al.
1000
100
Kaolinites
0
9
Illites
1~I
1
I
I
I
I
I
I
I
I
I
I
I
I
I
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er TmYb Lu
FIG. 5. Chondrite-normalized REE patterns of illite and kaolinite. Chondrite values from Taylor & McLennan
(1985).
PAAS, the higher overall R E E contents and the
higher (La/Yb)n and (Gd/Yb), ratios of the analysed
Devonian rocks are the main differences. These
differences are probably due to the higher degree of
weathering of their rocks (possibly as a consequence of sedimentary recycling). In fact, an
intense chemical weathering increases the R E E
fractionation because H R E E are preferentially
transported in solution (compared to LREE), since
H R E E form more stable carbonate, hydroxide and
even organic complexes (Balashov et al., 1964;
Ronov et al., 1967; Varshal et al., 1975; Cantrell &
Byrne, 1987). The good correlation between CIA
and (Gd/Yb)n (r = 0.86) in the Devonian rocks
would support this idea. Moreover, the clay
minerals produced by strong weathering are illite
and kaolinite, and these phases fix mainly LREE
(Nesbitt, 1979), as supported by the high (La/Yb)n
ratio in the clay fractions. In the chemical analyses
of these fractions, the effects of weathering and
sedimentary recycling in R E E are seen by
comparing their distributions in the illitic samples
and kaolinitic samples, assuming that the kaolinite
is the final product of stronger weathering (Fig. 5).
Kaolinite is notably enriched in REE, it has higher
(La/Yb)n ratios and a much higher (almost double)
(Gd/Yb), ratio than illite, and both phases show
LREE/HREE ratios which are higher than that of
PAAS.
The effect of sedimentary sorting is shown in the
lower overall R E E content, in the more pronounced
negative Eu-anomaly and in the lower (Gd/Yb)n
fractionation of the coarser grained rocks (DSGP)
compared to DP. The lower R E E content is
attributable to the quartz diluent effect, and the
other differences mentioned are attributable to the
preferential concentration of detrital heavy minerals
(zircon, monazite, apatite ...) in DSGP rocks, which
is supported by their Zr, Hf and P205 contents and
by the microscopic study. The good correlations
between P205 and H R E E and Zr and Eu/Eu*
explain that the differences between REE patterns
of DP and DSGP rocks are attributable to mineral
fractionation which produces a relative H R E E
enrichment in coarser grained rocks.
Analysed rocks show a slightly higher Th content
than those of PAAS, and this is consistent with a
source area of K-rich granitic composition (Taylor
& McLennan, 1985). The Th/U ratios in DP rocks
are higher than those of PAAS, whereas those from
PSGD rocks are lower. In each case these ratios are
in agreement with the contents registered in the
upper continental crust and sedimentary rock
averages (Taylor & McLennan, 1985), the higher
value displayed by DP rocks being explained by the
greater mobility of U under oxidizing conditions of
weathering. The Th/Sc and La/Sc ratios, particularly sensitive to average provenance composition
in terrigenous sediments (Taylor & McLennan,
1985; McLennan & Taylor, 1991; McLennan &
Hemming, 1992), are similar in DP and DSGP
rocks, and slightly higher than those of PAAS,
Mineralogy and geochemistry of Iberian Range Devonian rocks
Th
Hf
La
Th
Co
Sc
FIG. 6. Plot of average data of pelites (~A'), subgreywacke pelites (*), PAAS ('k), upper continental
crust (o), global continental crust (A) and oceanic
crust (ll) in the Th-Hf-Co and La-Th-Sc diagrams.
which suggests a more felsic composition of the
primitive source area, considering that sedimentary
recycling processes would tend to diminish the
differences (Taylor & McLennan, 1985). The fact
that these ratios increase sensitively in the coarsergrained DSGP rocks is possibly related to the
higher concentrations of detrital heavy minerals. In
the plot of average data of DP and DSGP in the
Th-Hf-Co and La-Th-Sc diagrams (Fig. 6) their
proximity to the dots corresponding to PAAS and
the upper continental crust can be observed,
although the position of DSGP deviates towards
the Th-Hf side, due to the above-mentioned mineral
fractionation. The plots of mean values of Devonian
detrital rocks suggest again that their primitive
source area has clearly been of a felsic composition;
this is also supported by the low concentration of
transition metals typically associated with mafic
compositions, such as Cr and Ni shown in these
rocks. A complementary possibility could be that
the differences noted are due to sedimentary
recycling processes.
CONCLUSIONS
Devonian detrital rocks from the Iberian Range
consist mainly of quartz and phyllosilicates, with
feldspar and calcite as minor phases. Illite and
kaolinite are the phyllosilicates encountered in fine
fractions. Illite has a low Na content, poor
crystallinity and a predominant 1Md polytype. Its
average composition is phengitic in greywackes and
muscovitic in pelites. Kaolinite has a high degree of
ordering. Both phyllosilicates were probably inher-
393
ited from a source area with intensive weathering
processes although illite might also have a
diagenetic origin. As far as the relation to
diagenesis or metamorphic degree undergone by
these rocks is concerned, the phyllosilicate assemblage suggests that they did not reach the lower
anchizone boundary.
It could also be claimed that high A1203/SIO2,
AI203/Na20, K20/Na20 ratios and CIW and CIA
indices indicate the high degree of weathering
undergone. This is supported by low Ca, Mg, St"
and Na contents and high Ti content compared to
PAAS. These chemical characteristics, in addition
to chondrite-normalized REE patterns, suggest a
high degree of maturity, possibly due to sedimentary recycling processes.
Chondrite-normalized REE patterns, high Th/Sc,
La/Sc and Th/Co ratios and low Fe, Mg and Ni
contents indicate that the primitive source area of
Devonian rocks studied was of felsic composition
(probably a K-rich granite).
ACKNOWLEDGMENTS
This work has been financed by Project PCB-1192
CONAI. The authors thank Dr De Caritat and an
anonymous referee for their suggestions, which helped
to improve the paper.
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