Mineralogy and geochemistry of the carbonates in the Calatayud …hera.ugr.es/doi/15022092.pdf ·...

ELSEVIER Chemical Geology 130 (1996) 123-136

CHEMICAL GEOLOGY

INCLUDING

ISOTOPE GEOSCIENCE

Mineralogy and geochemistry of the carbonates in the Calatayud Basin (Zaragoza, Spain)

M.J. Mayayo ", B. Bauluz a, A. L6pez-Galindo u, J.M. Gonzfilez-L6pez a a ,~rea de Cristalografia y Mineralogla, Dep. Ciencias de la Tierra, Universidad de Zaragoza, Zaragoza, Spain

b Instituto Andaluz de Ciencias de la Tierra, CSIC-Universidad de Granada, Fac. Ciencias, Granada, Spain

Received 28 November 1994; accepted 30 October 1995

Abstract

The Calatayud Basin is an elongated NW-SE depression in northeastern Spain. It is filled with continental Miocene sediments resulting from erosion of Palaeozoic and Mesozoic rocks of the Castilian and Aragonese branches of the Iberian Range. Miocene series consist of lutitic, marly, carbonate and evaporitic levels characteristic of a short-lived salt lake or a low-salinity plays-lake depositional environment. The carbonates detected in the series were calcite, dolomite and magnesite. The calcites generally :;how micritic textures, anhedral to subhedral morphologies and contain from 0 to 1 mole% MgCO 3, and so can be considered "low-Mg calcites". Their mean isotopic values are ~lSo = -7.8%0 (PDB) and 813C = --7.7%0 indicating precipitation in isotopic equilibrium with meteoric water and an extensive contribution of CO 2 of organic origin to the total dissolved C, respectively. The calcites present Mg/Ca, Sr/Ca and Na/Ca molar ratios that suggest precipitation from Na-enriched meteoric water. Dolomites are non-stoichiometric (44.5-47.7 mole% MgCO3), disordered, micfitic and anhedral to subhedral; their mean isotopic values [SlsO = + 0.4%0 (PDB) and 813C = --3.8%0] indicate that they precipitated from more evolved water and with a higher contribution of atmospheric CO 2 to the total dissolved C, respectively, than those of the calcites. Moreover, they present Mg/Ca, Sr/Ca and Na/Ca molar ratios suggesting precipitation from water with similar composition to seawater. The isotopic composition of magnesite [81So = +4.6%0 (PDB) and 513C = -4.0%~], together with the higher concentration in Sr and Na, indicates that this phase probably precipitated from more evolved water. The isotopic values and trace-element contents of carbonates appear to confirm the mineralogical trend from calcite to dolomite and to magnesite corresponding to the progressive increase of evaporitic concentration as the water, in a hydrologically closed system, was subjected to more acutely arid conditions.

1. Geological setting

The Calatayud Basin (Zaragoza) is an elongated N W - S E depression > 130 km long and ~ 20 km wide, filled with Miocene lacustrine sediments. This basin constitutes the boundary between the two domains of the Iberian Cordillera known as the Castil- ian and Aragonese branches. These domains are made up of Palaeozoic and Mesozoic rocks and sediments and the Calatayud Basin is located on the

Palaeozoic rocks of the Ateca threshold. The Palaeo- zoic rocks consist of layers of quartzites, sandstones and dolostones intercalated in unequally consolidated slaty series with a tendency to monoclinal distribution around the whole periphery of the basin. These levels generally dip towards the SW, so that the quartzite crests face NE (Fig. 1).

From the start of sedimentation in this basin, the Calatayud-Daroca and Teruel-Montalb~m sub-basins were clearly differentiated. Both contain evaporitic

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124 M.J. Mayayo et al. / Chemical Geology 130 (1996) 123-136

sediments and are separated by the Daroca detrital- calcareous threshold. According to Julivert (1954), the basin is a tectonic trough generated by two or three orogenic phases during the Tertiary. These phases broke up the basement along longitudinal

faults, striking mainly NW-SE and NNE-SSW, and separated two raised marginal areas. The Tertiary movements probably began in the Aquitanian and have become active again more recently. Tectonic activity ceased at the end of the Turolian (Pontian),

PALEOZOIC

MIOCENE

¢ONeLOMERATES

CAItNONATES, MARLS o LUTIT[$

GYPSUM, LUTITES

P L I O C E N E

Q U A T E R N A R Y

F-] OlSCONOANT CONTACT

• " -~ CONCORDANT CONTACT "o I l l

- _ _ f

/

Fig. 1. Geological setting and location of profiles and boreholes (in circles).

M.J. Mayayo et al. / Chemical Geology 130 (1996) 123-136 125

which marks the maximum expansion of the Tertiary fill. This study centres on the Calatayud-Daroca sub-basin.

The Miocene fill derives from the erosion of the marginal Palaeozoic and Mesozoic reliefs and presents considerable horizontal and vertical variation of facies. Horizontally, these sediments are dis- tributed in more or less concentric rings. Coarse facies, such as breccias and slaty quartzite conglomerates, predominate on the periphery of the basin,

while sandstone, silt and, particularly, clay and marl facies are located in intermediate position, giving way to carbonate- and gypsum-beating facies towards the centre of the basin. Lateral facies changes are abrupt, changing from conglomerates to finer and evaporitic formations over a few hundred metres. This is indicative of an arid climate.

The lithological sequences filling the basin are characteristic of a short-lived salt lake or low-salinity playa-lake depositional environment, in which only

(S 41) GLOBAL SAMPLE Om.

~, .~

L ~ , : ; ' ;

i t • /

. :_ ~;-;2-_-Z-_:,

7.~Z-----'-Z---3

~..a~.~-__~- - - :

38 I .

v/m, v////A I

M:

V//,/,, , / ~ z / / / / / / / / I

F////I//////////////I

I I

SILT FRACTION CLAY FRACTION

I I [ I I IIII I H~

t

I ~ - I II ~ I I

I , I : ~ I:==1

I I~ t f I:=11 I ~ I 1 I , , Io

V/ / / / / / / / / / I / / /4~r - -~ I I=III I=II

VzF::. I P--"III ~ II Fl//lll/I/lllll~i~ii: l I I I~ I " / ~ z ~ : : ~ i ~ i ~ i i : q I l ~ I F ~ " :'~--I

V/ / / / / / / / / / / / / ' I I I III IIII

LITHOLOGY GLOBAL SAMPLE CLAY MINERALS

L , . - T o . - o o , - , - I I ' ' L ' -

o o - - - - I I o ' - ' -

Fig. 2. L i tho log ica l sect ion and mine ra log ica l compos i t i on o f bo reho le $41. CM = c lay minera ls ; Qz = quartz; Fd = feldspars ; Ch = chlori te; Ka = kaolini te.


carbonates, sulphates and magnesian clays developed through precipitation of surface and subsurface water.

The main object of this study is the types of carbonate present in the different zones of this depositional environment. The textural and compositional characteristics of the carbonates, together with their trace elements and isotopic geochemistry, can help us to infer the nature of the solutions from which they precipitated. To this end, samples were selected

from the most carbonate-rich sections of profiles and boreholes from different sectors of the basin. Figs. 2 and 3 show the stratigraphic columns and mineralogical compositions of one borehole located in the lutitic plain zone and another in a more distal profile, respectively. The latter may very possibly have cor- responded to a short-lived salt lake zone.

A sepiolite deposit is located in the northeastern zone of the basin characterized by clays and marls. The mineralogical characteristics of this sepiolite

leom!C A) GLOBAL SAMPLE

L, I SILT FRACTION

F~'~'x-~y//xx> ;,~ >; I I====I I IIlilI

CLAY FRACTION

f

[

L,,,,

L Ill

~p;~//y////V.,',-',?/.?','/.~

I L x ~ ' ~ ~

L II

I ]

I III I Ill l 1

~/ / / / / / / / / / / / / / / )F=:~

t 7e ;~ ,I I

I IIII I

L, III t I]II I ~ I I ] [ .... -I: . . . . . . 4fl

I I

Om, LITHOLOGY

LIMESTONES ~ MARLS

DOLOSTONES ~ LUTITES

: _ I ~ MAQNESITE_ ~ ~ GYPSUM

GLOBAL SAMPLE

I I CALCITE ~ QI • Fd

DOLOMITE ~ C. M

MAeNIr$[TE ~ gYPSUM

CLAY MINERALS F"----] ILL ITE

SMECTITE

~ T ~ Ch. + Ko.

Fig. 3. Lithological section and mineralogical composition of profile CA. Legend as in Fig. 2.


were described by Ar~uzo et al. (1989). The informa- tion from the boreholes referred to here was pro- vided by the MYTA S.A. company, which mines this deposit.

2. Methods

was extracted after 20 min and again after 72 hr, as indicated by Waiters et al. (1972). The CO 2 was collected by freezing with liquid nitrogen and the analyses then were carried out in a SIRA-II mass spectrometer. The precision of the method is + 0.2%o for both carbon and oxygen.

The mineralogical composition of the powdered samples was determined by X-ray diffraction (XRD), using the Philips PW 1710 equipment, Cu-K~ radia- tion and an automatic slit. Analyses were carried out on both whole samples and their respective silt and clay fractions, the latller in air-dried oriented aggre- gates treated with ethylene glycol, dimethyl sulphox- ide and heated to 550°C. The reflective factors of Schultz (1964) and Barahona (1974) were used for the semi-quantitative analyses.

The molar percentage of MgCO 3 in the carbonates was determined by XRD, measuring the relative position of the (104) ~Lnd (101) reflections of carbonates and quartz, respectively; the latter being taken as standard (Goldsmith and Graf, 1958). The straight line regression of these authors was used for calcites and the equation by Lumsden (1979) for dolomites. The degree of ordering of dolomites was measured using the intensity ratio of the (105) and (110) reflections (Supko et al., 1974). The textures, grain size and morphologies of the carbonates were exam- ined by transmitted light microscopy and scanning electron microscopy (SEM) (ISM 6400) located at the Technical Service,,; of the University of Zaragoza.

The chemical analyses of the carbonates were carried out on the fraction dissolved in 1 N HC1, using inductively coupled plasma (Leeman PS 1000) at the University of Granada. The detection limits of elements are: A1, Fe, Na, Mg and Ca: 10 ppm; Sr: 5 ppm; Mn: 2 ppm; and[ Ba: 1 ppm.

The isotopic analyses were performed at the Sta- ble Isotopes Laboratory of the University of Sala- manca. The samples, 110-15 mg, were leached with 1 ml 100% pure H3PO 4 at 25°C. In the case of samples containing calci~e as the only carbonate, the resulting CO 2 was extracted 3 hr after the reaction, whereas when the samples contained pure dolomite and magnesite, the CO 2 was extracted after 72 hr, as described by McCrea (1950) and Craig (1957). In samples with mixed calcite and dolomite, the CO 2

3. Results

3.1. Mineralogical analysis by X-ray diffraction

Figs. 2 and 3 show the lithologies and mineralogical compositions of both the whole sample and the silt and clay fractions from borehole $41, which is representative of the transitional facies of the lutitic plain, and from profile CA, corresponding to the evaporitic facies of the short-lived salt lake.

Borehole $41 consists mainly of clayey deposits with intercalations of marly and carbonate layers. Clay minerals and dolomite predominate in the mineralogical composition, with calcite, quartz and feldspars as minor phases. The exception are a few calcite-rich levels. The main clay mineral phase is usually illite in association with smectite and very minor amounts of chlorite and kaolinite, although some levels have considerable concentrations of sepiolite. The latter is more abundant in other parts of the lutite plain, where it is at present being mined.

The profile CA is made up of gypsum-marl and gypsum-marl-lutite sequences in which evaporitic phases predominate. Gypsum is present either as a sole phase or together with small amounts of dolomite, magnesite, quartz and clay minerals. Small quantifies of gypsum were also detected in the marly and clayey layers. The carbonates that were detected in most of the series are dolomite and magnesite, occasionally forming almost monomineral layers. A limestone packet, consisting of almost pure calcite, outcrops at the top of the series. We did not observe the presence of efflorescences consisting of disori- ented fibrous hexahydrite crystals making up a non- consolidated network. Such features are known from some marly levels located between gypsum layers in zones sheltered from rainfall.

Table 1 shows the mineralogical composition, the molar MgCO 3 content of calcite and dolomite, the intensity ratio of the (015) and (110) reflections of


Table 1

Mineralogical and isotopic composit ions of analysed samples

Sample CM Qz Fd G Ca Do Mgs MgCO 3 l~oIs)/l~110) ~180 ~13C

(mole%) (%~ vs. SMOW) (%~ vs. PDB)

CA.27 0 1 0 0 99 0 0 0 - 5.9 - 7.8 CA.25 0 0 0 2 0 98 0 44.6 0.3 + 2.5 - 0.9 CA.21 0 0 0 0 0 0 100 + 4 . 6 - 4 . 0

$61.1 0 2 0 0 0 98 0 45.8 0.27 + 2 . 9 - 1.2

P1.L 0 2 0 0 0 98 0 47.5 0.33 + 0 . 8 - 5 . 1

P1.Y 0 5 0 0 0 95 0 47.9 0.32 + 0 . 9 - 5 . 0

P4.E (Ca) 14 2 0 0 5 79 0 0 0.0 - 6.5

P4.E (Do) 45.15 0.32 - 0 . 4 - 5 . 2

P5.D" 14 4 2 0 0 80 0 47.6 0.34 0.0 - 0.8

P5.C' 23 5 0 0 72 0 0 0.4 - 7.0 - 6.0

P7.25 0 2 5 0 0 93 0 47.5 0.48 - 5.4 - 1.6

$41.2 (Ca) 38 5 0 0 5 52 0 1 - 2.6 - 6.8

$41.2 (Do) 44.5 0.49 + 2.3 - 2.3 $41.11 0 0 0 0 100 0 0 0 - 8 . 7 - 9 . 2

$41.24 (Ca) 0 0 0 0 40 60 0 0 - 6.2 - 8.0

$41.24 (Do) 44.5 0.41 - 0.5 - 5.4 $64.10 0 0 0 0 0 100 0 47.7 0.3 + 1.7 - 3 . 6

$64.39 29 0 0 0 0 71 0 45.5 0.59 - 3.7 - 8.9

CM = clay minerals; Qz = quartz; Fd = feldspars; G = gypsum; Ca = calcite; Do = dolomite; Mgs = magnesite, l(ols)/Icj 10) = degree of

ordering of dolomites based on reflection intensity ratio (XRD).

Fig. 4. SEM micrographs of samples with compositions: (a) dolomite; (b) dolomite (95%) and quartz (5%); (c) calcite (72%), clay minerals (23%) and quartz (5%); and (d) dolomite (60%) and calcite (40%).


100

A

80

Ov

,=,, ~ 4 o

2o

o -11-1o-9 -8 -7 -6 -5 -4 -3 -2 -1 o 1 2 3 4 5 6

I~ 0 CALCITE

00

- B

-11-10-9 ~ 8 - 7 - 6 - 5 - 4 - 3 - 2 - 1 0 1 2 3, 4 5 6 7 8 9 10

( ~ e O DOLOMITE

100

80

a.., = so

~= 40

20

0 -11-10-9 -8 -7 -6 -5 -4 -3 -2 -1 0 1

( ~ ' 0 CALCITE

C

2 3 4 5 6

~oo D

-11-10-9-8 - 7 - 6 -5 -4 -3 -2 -1 0 1 2 3 4 5 6 7 8 9 10

"o ° ~ . ~

lOO

IV

E

\

-11-10-9 -8 -7 -6 45 -4 -8 -2 -1 0 1 2 8 4 S 6

~ " 0 CAI.CltE

100

oo

40

2o

- 1 1 - 1 0 - 9 - 8 - 7 - 6 - 5 - 4 - 3 - 2 - 1 0 1 2 3 4 5 6 7 8 9 10

Fig. 5. Isotopic compositions of water in equilibrium with carbonates: (A) micritic calcites; (B) micritic dolomites; micritic calcites (C) and dolomites (D) in the same: samples; sparitic calcite (E) filling veins in micritic dolomite (F). The 8]SO-values (SMOW) of water were obtained from the equations of Garlick (1974) and Anderson and Arthur (1983).

130 M.£ Mayayo et al. / Chemical Geology 130 (1996) 123-136

dolomite and the isotopic compositions of the carbonates in the selected samples.

The calcites generally have micritic textures, anhedral to subhedral morphologies (Fig. 4a and c), and their molar MgCO 3 content ranges from 0 to 1%. They are therefore "low-Mg calcites". Some- times (e.g., sample $41-24) the calcite is sparitic and fills veins and cracks in a micritic dolomite.

The dolomites are non-stoichiometric, with a molar MgCO 3 content of 44.5 to 47.7% and an average value of 45.25 _ 1.43%. They have low degrees of ordering, micritic textures and anhedral or subhedral morphologies (Fig. 4b and d).

3.2. Isotopic composition of the carbonates

3.2.1. Oxygen isotopes The ~80-values of calcites vary from - 8 . 6 to

0.0%0 (PDB), although this wide range narrows con- siderably when only the same mineralogical assemblage is considered. For example, if calcite is the only carbonate found and its texture is micritic, the ~lSO-values range from - 8 . 6 to -5.9%0 (PDB), with an average value of - 7.8%0. If we assume that these calcites precipitated at normal surface temperatures (15-25°C), the ~180-values of water in isotopic equilibrium could have ranged between - 7 . 4 and - 5 . 1 % o (SMOW) (Anderson and Arthur, 1983) (Fig. 5A). In contrast, calcite associated in small percent- ages with dolomite (both fine-grained) has much heavier ~80-values (from - 2 . 6 to 0.0%o), with a mean value of -1.3%o. In this case, for the same temperature range as above, the ~80-values of water would have to be - 1.5 to + 0.7%0 (SMOW), that is enriched in 180 by ~ 6%0 (Fig. 5C). The dolomicrites associated with these calcites have ~180-values between - 0 . 4 and + 2.3%0, with a mean of + 1.9%o. Again, for the temperature range considered, the ~180-values of the water (cf. Garlick, 1974) would have to be - 1 . 5 and +0.6%0 (Fig. 5D). This suggests that the two carbonates of the assemblage may have formed contemporaneously. The samples in which the only carbonate is dolomite have ~80-values between - 5 . 4 and +2.9%0; the two heaviest values measured on samples from the central zones of the basin and/or from the portion of a zone nearest to the surface. The mean ~80-value is + 0.4%0 which, for the same temperature range, cor-

responds to dolomites precipitated in isotopic equilibrium with water whose ~180 would be - 3 . 0 to - 0.9%0 (Fig. 5B).

Microscopic examination of some of the dolomi- critic samples revealed veins and fractures filled with sparitic calcite. The ~180-value of the only such calcite analysed isotopically was -6.2%0, while the host dolomite had a value of -0.5%0 (Fig. 5E and F).

The 8180-value of +4.6%0 (PDB) detected in the magnesite sample is the highest recorded in the studied carbonates.

3.2.2. Carbon isotopes The ~3C-values of the carbonates analyzed also

cover a wide range ( - 9 . 2 to -0.8%o), typical of carbonates formed in continental environments, where decaying organic matter serves as a significant source of carbon for the dissolved inorganic carbon pool.

The lightest 8~3C-values are found in calcites where these are the only carbonate present, with values from - 9.2 to - 5.9%0, and a mean of - 7.7%0. The calcites associated with dolomicrites have a mean 8~3C of -6.6%0, and in the sparitic calcite filling fractures in dolomicrite the 813C-value is - 8.0%0.

The dolomicrites, on the other hand, have heavier 813C-values, from - 8 . 9 to -0.8%0, with a mean of - 3 . 8 % 0 . The heaviest values are almost equivalent to those typical of marine carbonates.

The ~13C of magnesite is -4.0%0, practically the same as the mean value for the dolomicrites.

3.3. Chemical analyses

Table 2 presents the results of the chemical analyses of the carbonate fraction of the selected samples. The variations in concentration of the elements analyzed reflect the diversity of their mineralogical compositions.

The statistically significant positive correlation between A1 and Fe ( r = 0.74) indicates that the concentration of Fe is mainly controlled by the contribution of the aluminosilicate fraction (clay minerals) dissolved during digestion of the samples. This is confirmed by the good correlation found between clay mineral content and A1 ( r = 0.90) and Fe ( r =

M.J. Mayayo et al. / Chemical Geology 130 (1996) 123-136

Table 2 Chemical analyses (in ppm) of carbonate fraction of studied samples

131

Sample A1 Fe Mn Na Mg Ca Ba Sr

$41.11 155 383 211 300 3,468 359,700 88 167 CA.27 145 133 27 327 2,946 361,200 21 302 $61.1 515 629 61 1,702 115,100 199,000 17 218 P1 .Y 652 699 195 861 116,900 238,900 102 178 CA.25 773 850 64 1,969 123,800 283,200 11 305 CA.21 127 74 95 722 222,500 18,130 7 47 P5.C 1,283 1,271 574 419 9,193 296,400 22 5,335 $41.24 297 1,915 11,340 567 85,910 209,900 205 98 P4.E 607 1,030 342 904 94,960 186,700 106 152 $41.2 1,939 3,091 249 1,270 86,390 1 64,600 48 191 P7.25 835 1,133 384 506 114,000 188,900 23 443 $64.39 1,088 3,329 2,925 647 108,200 203,300 16 89 P5.D 988 2,613 359 1,009 100,800 206,900 76 2,315 $64.10 793 2012 616 972 97,130 217,500 18 204

0.78), when the results of the XRD semi-quantitative mineralogical analyses are introduced into the correlation matrix. The chemical analyses can therefore be divided into two groups according to the Fe contents: one group of almost: exclusively carbonate samples (Fe < 900 ppm), and another containing appreciable amounts of clay minerals, quartz, feldspars, celes- tine, oxides, etc. (Fe > 1000 ppm) as well as carbonates as major components. This grouping is confirmed by the XRD mineralogical analyses of the whole samples.

In the group consisting of almost pure carbonates, the relative concentrations of Ca and Mg distinguish between calcite, dolomite and magnesite. In the other group, with relatively high Fe contents, some sam- pies analyzed presented anomalous concentrations in some elements, e.g. P5-C and P5-D with high Sr contents, or $41-24 and $64-39, with high Mn contents and low Sr contents. The Sr anomalies in P5-C and P5-D can be explained by the presence of small quantities of celestite detected in the mineralogical analyses by XRD. "llae anomalies in Mn and Sr of

$41-24 and $64-39 could be attributed to the diagenetic effects produced by meteoric solutions which, in open systems, modify the chemical composition of carbonates in this way (Brand and Veizer, 1980), although the excess Mn may also be due to anoxic conditions in the basin or in the diagenetic system. Another possibility is that the Mn originated in the dissolved non-carbonate fraction of the samples. If this were the case, the dissolved phases supplying them must be the oxides detected in the microscopic examination, because higher concentration of this element is associated with a relatively high Fe content and a low A1 content ($41-24).

Table 3 presents the mean values of different elements in calcite, dolomite, magnesite and dolomite plus other phases, separated according to the criteria mentioned above. The data on samples with anomalous concentrations of an element have not been included in the last group. The calcites have a mean Mg content equivalent to a MgCO 3 molar percentage of 1.1%, which is very similar to that deduced from the XRD data and, therefore, falls within the range

Table 3 Mean values (wt%) of analysed elements in the carbonate fraction

Sample AI Fe Mn Na Mg Ca Ba Sr

Calcite 150 258 119 314 3,207 360,450 55 235 Dolomite 647 726 107 1,511 118,600 240,367 43 234 Magnesite 127 74 95 722 222,500 18,130 7 47 Dolomite + others 910 1,966 390 871 96,532 195,750 79 247


| Naf /

' l d 5 ' 10 5 ' 16" ' 1 0 '

(° M e / " C a ) w a~,,ai

Fig. 6. Mola r ratios o f solutions in equi l ibr ium with calcite. The

molar ratios of present-day seawater (SW), river water (RW) and shallow subsurface water (SSW) was taken from Majid (1983), Veizer (1983) and Machel (1986).

10 ~

"~ 10" R 0

- ___.r______.~ ~, =

- . - ' " • Fe

/ i r i i J i i

10 "5 10 "a 10" 10'

( 'Me/ 'Ca) w ~,tual

Fig. 7. Mola r ratios o f solutions in equi l ibr ium with dolomite.

Legend as in Fig. 6.

accepted for "low-Mg calcites" (Milliman, 1974; Brand, 1981). The dolomites present significantly higher A1 and Fe contents than the calcites and magnesite, which may indicate that their data also reflect some contribution of the non-carbonate phase of the samples. However, if we assume that the major elements Ca and Mg correspond to the carbonate, they are equivalent to a MgCO 3 molar content of 45% - - very similar to that deduced by XRD - - which suggests that the dolomites are deficient in Mg and are not stoichiometric. Naturally, the "dolomite plus other phases" group shows the highest A1, Fe and Mn contents.

Table 4 Mean values of molar ratios of carbonates

Calcite Dolomite Magnesite Dolomite+others

mFe/mca 5.1 21.7 29.3 73.3 ( X 10 4 )

mmn/mca 2.4 3.3 38.2 14.5 ()< 104 )

tuNa / i n C a 1.52 l l . 1 69.4 7.93 ( X 10 3 )

mMg / me= 1.47 82.7 2,023.4 81.9 (×10 2) msr / me= * 3 4.4 11.9 5.2 ( × 1 0 4 ) tuba / inCa 4.4 5.4 11 12 (XlO 5)

Table 4 shows the values of the molar concentrations of the elements normalized to the Ca molar concentrations. We observe a general progressive increase in the ratios of all the elements from calcite to dolomite and magnesite. Assuming that these trace elements are mainly found in the structures of the carbonates in substitution of Ca and Mg depending on their ionic radii (thus discounting the possibility that they occupy structural defects, are absorbed or form part of solid or liquid inclusions), these molar ratios and the appropriate solid-liquid distribution coefficients (Kinsman, 1969; Jacobson and Us- dowski, 1976; Kretz, 1982; Veizer, 1983) can be used to deduce the molar ratios corresponding to the solutions in equilibrium with the carbonates. Figs. 6 and 7 are graphic representations of these ratios for Mg, Sr, Na, Mn and Fe in calcite and dolomite, respectively. These are the elements normally used to indicate the chemistry of formation water repre- sented by the 1:1 lines in each diagram. By way of comparison, we have also projected the values corresponding to different types of present water, such as seawater (SW), river water (RW) and shallow subsurface water (SSW), as deduced from the data by Majid (1983), Veizer (1983) and Machel (1986). If the values of the 1:1 lines are accepted as first-order approximations to the compositions of Miocene water in the Calatayud Basin, their positions relative to the 1:1 lines of other waters can indicate the chem-

M.J. Mayayo et a l . / Chemical Geology 130 (1996) 123-136 133

istry of the water from which the carbonates precipitated.

4. Discussion and conclusions

The geochemical characteristics of trace elements and the isotopic ch~xacteristics of the carbonates analyzed describe the chemical evolution of the water from which they precipitated in the different zones of the basin.

In the mud-flats zone, the carbonates detected are calcite and dolomite, with clear predominance of the former. This is Mg-low calcite, with a mean MgCO 3 molar content of 1.1%, micritic texture and mean isotopic values of ~11tO = -7.2%0 (PDB) and ~13C =-6 .6%0. All these data suggest that these are primary calcites precipitated in equilibrium with isotopically light, meteoric water and with a strong organic contribution to the dissolved carbon reservoir, probably derived from C 3 plant decay (Ceding, 1984) and photosynthetic activity. Precipitation was therefore probably biologically induced. The variations observed in the isotopic compositions of C in these calcites show the relative influence of atmospheric CO 2 during precipitation. The correlation between the ~ 80 - and ~13C-values of the calcites ( r = 0.58) may indicate precipitation from water with relatively long residlence times in hydrologically closed systems, where the recharge- evaporation budget controls the isotopic evolution of the water, as against open lakes with short residence times for the water, in which the 8180 is relatively invariable and is typically associated with the isotopic value of the recharge water (Fontes and Gonflantini, 1967; Gonfiantini, 1986; Fritz et al., 1987; Talbot, 1990).

On the other hand, as deduced for water in equilibrium with calcite ( l : l line in Fig. 6), the Sr /Ca molar ratio coincides with that of river water, the Mg/Ca ratio falls between that of river water and shallow subsurface water, and the Na /Ca ratio is almost the same as seawater, these three elements being the most reliablte indicators of the chemistry of the solutions (Majid and Veizer, 1986). These data seem to indicate that these calcites were probably the first carbonates precipitated from the relatively diluted alkaline solutions originating from weathering of the Palaeozoic rocks and sediments on the basin

margin. The alkaline nature of the water is indicated by the fact that the country rocks contain no ancient pyritic or evaporitic slates to prevent the develop- ment of alkaline brine (Jones, 1966; Garrels and MacKenzie, 1967; Hardie and Eugster, 1970) and by the presence, in nearby zones, of sepiolite deposits with neoformed silex nodules. These require alkaline conditions for their precipitation. The processes that led to oversaturation and calcite precipitation could have been evaporitic concentration, degasification, mixing of water and temperature changes (Eugster and Hardie, 1978).

The dolomites analyzed are disordered, deficient in Mg, fine-grained and anhedral to subhedral. These crystal-chemical and morphological characteristics are typical of primary or early diagenetic dolomites formed at low temperature (Gunatilaka, 1990; Last, 1990; Last and De Deckker, 1990). Their mean isotopic values (8180 = +0.4%0 and 813C= - 3.8%o) are clearly heavier than those of the calcites and suggest that they precipitated from more intensely evaporated water. This hypothesis is supported by the value of the isotopic enrichment factor between calcite and dolomite, obtained from their corresponding mean ~80-values, e = - 7 . 6 % 0 , which is outside the - 6 to -3%o range generally accepted for cogenetic minerals (Fontes et al., 1970). The ~13C-values of the dolomites, heavier than those of the calcites, suggest a higher contribution of atmospheric CO 2 to the reservoir of dissolved carbon. The relative dispersion of the isotopic values, especially those of carbon, may indicate some type of mixture between evaporitic brine and meteoric water for the formation water (Coniglio et al., 1988).

The molar ratios deduced for the water in equilibrium with the dolomites coincide with that of seawater for Sr and Mg (Fig. 7), whereas that of Na is even higher. These data confirm what was suggested by the isotopic values as regards the higher degree of concentration attained by the solutions when the dolomite formed, as a result of longer residence times and/or more evaporitic conditions. These solutions must have had high Mg/Ca ratios, high alkalinity with predominance of CO 2- over HCO 3 ions and only small quantities of SO 2- in solution (Hsii, 1967; Baker and Kastner, 1981; Machel and Mountjoy, 1986; Last, 1990). Dolomite precipitation in the mud-flat zones of the basin was probably


caused by evaporitic pumping processes, or took place under a very shallow layer of water during periods of greater aridity.

The calcites occasionally found in association with the dolomites have geochemical and morphological characteristics that allow us to distinguish two gener- ations. The first of these is micritic to microsparitic calcite with a mean 818t-value of -1.3%o (PDB), clearly heavier than the other calcites. This suggests an enrichment factor for the associated calcite- dolomite pair of e = -3.2%0, within the accepted range for contemporaneous formation of both phases. The other calcite is sparitic and found as filling at fractures in the dolomite. This phase has a 818t-value (-6.2%0) that suggests precipitation from meteoric water. This calcite clearly formed later than the dolomite and its formation water was also very different.

In the central zones of the basin the marly and carbonate deposits are intercalated with gypsum- bearing beds that indicate more arid depositional conditions (Fig. 2). The most common gypsum facies is laminate, with grain-size gradation and rip- pies, that suggest direct precipitation in shallow lacustrine environments (Ortf, 1992). Less frequently gypsum is also present in nodular, lenticular and selenitic forms within the marly beds, and may have originated during diagenesis by evaporitic pumping. The carbonates in these deposits are dolomite and magnesite.

The dolomite has a 818t-value of + 2.5%0 (PDB), in isotopic equilibrium with waters of - 0 . 9 to

+ 1.2%o (SMOW) at temperatures of 15-25°C. These relatively heavy values indicate an increase in the rate of evaporation and/or the residence time of the formation water of the dolomite in comparison with those considered above. The 813C-value of -0.9%0 of this dolomite is almost identical to the typical values of marine carbonates and suggests that it precipitated in an arid environment lacking in plant remains (Lord et al., 1988) and with a high contribution of atmospheric CO 2 to the reservoir of dissolved carbon (Fontes et al., 1970; Deines, 1980; Anderson and Arthur, 1983). The values of the Sr /Ca and Na /Ca molar ratios of this dolomite are slightly higher than those of the dolomites described earlier, which confirms our observations on the evolution of the formation water as inferred from the isotopic compositions.

The magnesite presents the heaviest 8~80-value ( + 4.6%0, PDB) of all those recorded and the Sr /Ca and Na /Ca molar ratios are much higher than those of dolomite. The isotopic fractionation factors as regards water are not yet well established, nor are the distribution coefficients for the trace elements, and so we cannot make calculations similar to those carried out for dolomite. However, the isotopic sig- nals and the geochemistry of trace elements in the magnesite suggest that it formed from even more concentrated solutions than those giving rise to dolomite and, therefore, in more intensely evaporitic conditions than the other carbonates. The positive correlation ( r = 0 . 8 8 ) existing between the mean isotopic values of O and C in the three carbonates

-10 0

-2

-4

-6

-8

-8 -6 -4 -2 0 2 4 6 I I I I t I I 0

A AA

[] .c

DE

D ~ M

/1 12

- 1 0 ~ I ~ ~ ~ t I

-10 -8 -6 -4 - 1 8 - -2 0 2 4 0 0

--2

.-4

--6

,-8

-10

13 Fig . 8. Plot o f ~ 8 0 - and 8 C-va lues o f carbonates . Squares = calci te; triangles = do lomi te ; rhombs = magnes i te . M e a n va lues are filled.


(Fig. 8) corresponds to the mineralogical transition from calcite to dolomite and magnesite produced by the gradual increase in the evaporitic concentration of the water. These results are consistent with long residence times of water in hydrologically closed systems, where evaporation and associated aridity produce progressively heavier isotopic compositions in comparison with those of recharge water (Talbot and Kelts, 1990).

It is probable that in these central zones of the basin the solutions were relatively enriched in Mg 2+ because of precipitation of gypsum and subsequent loss of SO 2- and Ca 2+ ions. They probably reached concentrations adequate for formation of dolomite as a first stage and the~a, as the Ca 2+ was consumed, magnesite. The pericxiical fluctuations due to sup- plies of recharge waLer and climatic variations produced the sequences of more or less evaporitic beds that the deposits form in these more central areas of the basin.

A massive limestone packet outcrops at the top of the series. Its isotopic values [8180 = -5 .9%, (PDB) and ~t3c = -7.8%0] indicate that the calcite precipitated from diluted meteoric water with a strong organic contribution to the dissolved carbon reservoir. The geochemistry of its trace elements confirms what the isotopic sig~aals suggest. All the data therefore seem to reveal an abrupt change in the chemistry of the water, probably related to a climatic change to clearly less arid conditions than those previously predominating.

Acknowledgements

We thank Dr. R.I. MacCandless of the English Philology Department of the University of Granada for his help on the la'anslation of the text, and two anonymous referees for their useful criticisms. This research was partially supported by projects PCB- 1192 (CONAI), Diputaci6n General de Arag6n, and AMB93-0794 (CICYT).(CA)

References

Anderson, T.F. and Arthur, M.A., 1983. Stable isotopes of oxygen and carbon and their .application to sedimentologic and pale- oenvironmental problems. In: M.A. Arthur, T.F. Anderson,

I.R. Kaplan, J. Veizer and L.S. Land (Editors), Stable Isotopes in Sedimentary Geology. Soc. Econ. Paleontol. Mineral., Short Course No. 10.

Arauzo, M.; Gonzalez Lopez, J.M. and Lopez Aguayo, F., 1989. Primeros datos sobre la mineralogia y genesis del yacimiento de sepiolita de Mara (prov. de Zaragoza). Bol. Soc. Esp. Min., 12: 329-340.

Baker, P.A. and Kastner, M., 1981. Constraints on the formation of sedimentary dolomite. Science, 213: 214-216.

Barahona, E., 1974. Arcillas de ladriller~a de la provincia de Granada: evaluaci6n de algunos ensayos de materias primas. Doctoral Thesis, University of Granada, Granada, 398 pp.

Brand, U., 1981. Mineralogy and chemistry of the Lower Pennsyl- vanian Kendrick fauna, eastern Kentucky, 1. Trace elements. Chem. Geol., 32: 1-16.

Brand, U. and Veizer, J., 1980. Chemical diagenesis of a multi- component carbonate system, 1. Trace elements. J. Sediment. Petrol., 50: 1219-1236.

Ceding, T.E., 1984. The stable isotopic composition of modern soil carbonate and its relationship to climate. Earth Planet. Sci. Lett., 71: 229-240.

Coniglio, M., James, N.P. and Aissaoui, D.M., 1988. Dolomitiza- tion of Miocene carbonates, Gulf of Suez, Egypt. J. Sediment. Petrol., 58: 100-119.

Craig, H., 1957. Isotopic standards for carbon and oxygen and correction factors for mass spectrometric analysis of carbon dioxide. Geochim. Cosmochim. Acta, 12: 133-149.

Deines, P., 1980. The isotopic composition of reduced organic carbon. In: P. Fritz and J.Ch. Fontes (Editors), Handbook of Environmental Isotope Geochemistry, 1. The Terrestrial Envi- ronment, A. Elsevier, Amsterdam, pp. 329-406.

Eugster, H.P. and Hardie, L.A., 1978. Saline lakes. In: A. Lerman (Editor), Lakes, Chemistry, Geology, Physics. Springer, New York, N.Y., pp. 237-293.

Fontes, J.Ch. and Gonfiantini, R., 1967. Comportement isotopique a u cours de r6vaporation de deux bassins sahariens. Earth Planet. Sci. Lett., 3: 258-266.

Fontes, J.Ch., Fritz, P. and L&olle, R., 1970. Composition isotopique, min6ralogique et gen~se des dolomies du Bassin de Paris. Geochim. Cosmochim. Acta, 34: 279-294.

Fritz, P., Morgan, A.V., Eicher, U. and McAndrews, J.H., 1987. Stable isotope, fossil Coleoptera and pollen stratigraphy in Late Quaternary sediments from Ontario and New York State. Palaeogeogr., Palaeoclimatol., Palaeoecol., 58: 183-202.

Garlick, G.D., 1974. The stable isotopes of oxygen, carbon, and hydrogen in the marine environment. In: E.D. Goldberg (Edi- tor), The Sea, Vol. 5. Wiley, New York, N.Y., pp. 393-425.

Garrels, R.M. and MacKenzie, F.T., 1967. Origin of the chemical composition of some spring and lakes. Am. Chem. Soc., Adv. Chem. Ser., pp. 222-242.

Goldsmith, J.R. and Graf, D.L., 1958. Structural and compositional variations in some naturally occurring dolomites. J. Geol., 66: 678-693.

Gonfiantini, R., 1986. Environmental isotopes in lake studies: In P. Fritz and J.Ch. Fontes (Editors), Handbook of Environmen- tal Isotope Geochemistry, 2. The Terrestrial Environment, B. Elsevier, Amsterdam, pp. 113-168.


Gunatilaka, A., 1990. Dolomite formation in coastal A1-Khiran, Kuwait Arabian Gulf - - a re-examination of the sabkha model. Sediment. Geol., 72: 35-53.

Hardie, L.A. and Eugster, H.P., 1970. The evolution of closed- basin brines. Mineral. Soc. Am., Spec. Pap., 3: 273-290.

Hsi], K.J., 1967. Chemistry of dolomite formation. In: G.V. Chilingar, H.J. Bissell and R.W. Fairbridge (Editors), Carbon- ate Rocks. Elsevier, Amsterdam, pp. 169-192.

Jacobson, R.L. and Usdowski, H.E., 1976. Partitioning of stron- tium between calcite and dolomite, and liquids. Contrib. Min- eral. Petrol., 59: 171-185.

Jones, B.F., 1966. Geochemical evolution of closed basin waters in the western Great Basin. N. Ohio Geol. Soc. 2nd Symp. on Salt, I: 181-200.

Julivert, M., 1954. Observaciones sobre la tectrnica de la depresirn de Calatayud. Arraona, pp. 1-18.

Kinsman, D.J.J., 1969. Interpretation of Sr 2+ concentrations in carbonate minerals and rocks. J. Sediment. Petrol., 139: 486- 508.

Kretz, R., 1982. A model for the distribution of trace elements between calcite and dolomite. Geochim. Cosmochim. Acta, 46: 1979-1981.

Last, W.M., 1990. Lacustrine dolomite - - an overview of modem, Holocene, and Pleistocene ocurrences. Earth-Sci. Rev., 27: 221-263.

Last, W.M. and De Deckker, P., 1990. Modem and Holocene carbonate sedimentology of two volcanic maar lakes in south- em Australia. Sedimentology, 37: 967-981.

Lord, B.K., Jones, L.M. and Faure, G., 1988. Evidence for the existence of the Gondwana ice sheet in the ISo depletion of carbonate rocks in the Permian formations of the Transantarc- tic Mountains. Chem. Geol. (Isot. Geosci. Sect.), 8: 163-172.

Lumsden, D.N., 1979. Discrepancy between thin section and X-ray estimates of dolomite in limestone. J. Sediment. Petrol., 49: 429-436.

Machel, H.G., 1986. Limestone diagenesis of Upper Devonian Nisku carbonates in the subsurface of central Alberta. Can. J. Earth. Sci., 23: 1804-1822.

Machel, H.G. and Mountjoy, E.W., 1986. Chemistry and environments of dolomitization - - a reappraisal. Earth-Sci. Rev., 23: 175-222.

Majid, A.H., 1983. Lithofacies, chemical diagenesis and dolomitization of the Tertiary carbonates of the Kirkuk oil field, Iraq. Ph.D. Thesis, University of Ottawa, Ottawa, Ont., 270 pp.

Majid, A.H. and Veizer, J., 1986. Deposition and chemical diagenesis of Tertiary carbonates, Kirkuk oil field, Iraq. Am. Assoc. Pet. Geol. Bull., 7: 898-913.

McCrea, J.M., 1950. On the isotopic chemistry of carbonates and a palaeotemperature scale. J. Chem. Phys., 18: 849-857.

Milliman, J.D., 1974. Marine Carbonates. Springer, Berlin, 375

pP. Ortl, F., 1992. Diagrnesis en las evaporitas continentales del

Terciario peninsular ib&ico, lII Congr. Geol. Esp. Simp., 1: 118-127.

Schultz, L.G., 1964. Quantitative interpretation of mineralogical composition from X-ray and chemical data for the Pierre Shale. U.S. Geol. Surv., Prof. Pap., 391-C: 1-31.

Supko, P.R.; Stoffers, P. and Coplen, T.B., 1974. Petrography and geochemistry of Red Sea dolomite. In: R.B. Whitmars, O.E. Weser, D.A. Ross (Editors), Initial Reports of the Deep Sea Drilling Project, Vol. 23. U.S. Gov. Print. Off., Washington, D.C., pp. 867-878.

Talbot, M.R., 1990. A review of the palaeohydrological interpretation of carbon and oxygen isotopic ratios in primary lacustrine carbonates. Chem. Geol. (Isot. Geosci. Sect.), 80: 261-279.

Talbot, M.R. and Kelts, K., 1990. Paleolimnological signatures from carbon and oxygen isotopic ratios in carbonates from organic carbon-rich lacustrine sediments. In: B.J. Katz (Editor), Lacustrine Basin Exploration - - Case Studies and Modem Analogs. Mere. Am. Assoc. Pet. Geol., 50: 99-112.

Veizer, J., 1983. Chemical diagenesis of carbonates: theory and application of trace element technique. In: M.A. Arthur, T.F. Anderson, I.R. Kaplan, J. Veizer and L.S. Land (Editors), Stable Isotopes in Sedimentary Geology. Soc. Econ. Paleontol. Mineral., Short Course No. 10.

Waiters, Jr., L.J., Claypool, G.G. and Choquette, P.W., 1972. Reaction rates and JSO variation for the carbonate-phosphoric acid preparation method. Geochim. Cosmochim. Acta., 36: 129-140.

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